CHONDRULES

Chondrules are sub-millimetre spherical metal-sulphide-silicate objects which formed from the solar protoplanetary disk material, and as such provide an important record of the chronology and conditions of the solar system in pre-planetary times. Chondrules are a major constituent of chondritic meteorites; however, despite being recognised for over 200 years their origins remain enigmatic. This comprehensive review describes state-of-the-art research into chondrules, bringing together leading cosmochemists and astrophysicists to review the properties of chon- drules and their possible formation mechanisms based on careful observations of their chemis- try, mineralogy, petrology and isotopic composition, as well as laboratory experiments and theoretical modelling. Current and upcoming space missions returning material from chondritic asteroids and cometary bodies have invigorated research in this field, leading to new models and observations, and providing new insight into the conditions and timescales of the solar proto- planetary disk. Presenting the most recent advances, this book is an invaluable reference for researchers and graduate students interested in meteorites, asteroids, planetary accretion and solar system dynamics. sara s. russell is Head of Planetary Materials at the Natural History Museum in London, England, and a visiting professor at the Open University. She is a fellow of the Meteoritical Society and has been honoured with the eponymous asteroid 5497 Sararussell. She has been awarded the Antarctica Service Medal of the United States of America and the Bigsby Medal of the Geological Society. Her research interests include the formation of the solar system and the evolution of the Moon. harold c. connolly jr. is founding Chair and Professor in the Department of Geology, School of Earth and Environment, Rowan University in Glassboro, New Jersey. He is also a research associate at the American Museum of Natural History and was a special visiting professor at Hokkaido University in Sapporo, Japan. He has been awarded the Antarctica Service Medal of the United States of America, and in 2006 Asteroid 6761 Haroldconnolly 1981 EV19 was named in his honour. He is a co-investigator and Mission Sample Scientist for NASA’s New Frontiers 3 asteroid sample return mission OSIRIS-REX, and co-investigator for JAXA’s asteroid sample return mission, Hayabusa2. He is a fellow of the Meteoritical Society. His career has been devoted to understanding the formation and evolution of primitive planetary materials, chondritic meteorites, chondrule formation, the formation and dynamical evolution of asteroids and the origin of Earth-like planets. alexander n. krot is a researcher at the University of Hawai’iatMānoa, Honolulu, USA, from which he received the Regents’ Medal for Excellence in Research in 2004. He has also received a Humboldt Research Award and has been awarded the Antarctica Service Medal of the United States of America. He has been recognised by having both an asteroid (6169 Sashakrot Ex4) and a mineral (krotite) named in his honour. He is a fellow of the Meteoritical Society, by which he has been recently awarded the Leonard Medal. His research interests include astrophysical and cosmochemical problems related to the formation and early history of the solar system; chondritic meteorites and refractory inclusions; and isotope chronology. CAMBRIDGE PLANETARY SCIENCE

Series Editors:

Fran Bagenal, David Jewitt, Carl Murray, Jim Bell, Ralph Lorenz, Francis Nimmo, Sara Russell

Books in the Series: 1. Jupiter: The Planet, Satellites and Magnetosphere† Edited by Bagenal, Dowling and McKinnon 978-0-521-03545-3 † 2. Meteorites: A Petrologic, Chemical and Isotopic Synthesis Hutchison 978-0-521-03539-2 3. The Origin of Chondrules and Chondrites† Sears 978-1-107-40285-0 4. Planetary Rings† Esposito 978-1-107-40247-8 † 5. The Geology of Mars: Evidence from Earth-Based Analogs Edited by Chapman 978-0-521-20659-4 6. The Surface of Mars Carr 978-0-521-87201-0 7. Volcanism on Io: A Comparison with Earth Davies 978-0-521-85003-2 8. Mars: An Introduction to Its Interior, Surface and Atmosphere Barlow 978-0-521-85226-5 9. The Martian Surface: Composition, Mineralogy and Physical Properties Edited by Bell 978-0-521-86698-9 10. Planetary Crusts: Their Composition, Origin and Evolution† Taylor and McLennan 978-0-521-14201-4 † 11. Planetary Tectonics Edited by Watters and Schultz 978-0-521-74992-3 12. Protoplanetary Dust: Astrophysical and Cosmochemical Perspectives† Edited by Apai and Lauretta 978-0-521-51772-0 13. Planetary Surface Processes Melosh 978-0-521-51418-7 14. Titan: Interior, Surface, Atmosphere and Space Environment Edited by Müller-Wodarg, Griffith, Lellouch and Cravens 978-0-521-19992-6 15. Planetary Rings: A Post-Equinox View (Second edition) Esposito 978-1-107-02882-1 16. Planetesimals: Early Differentiation and Consequences for Planets Edited by Elkins-Tanton and Weiss 978-1-107-11848-5 17. Asteroids: Astronomical and Geological Bodies Burbine 978-1-107-09684-4 18. The Atmosphere and Climate of Mars Edited by Haberle, Clancy, Forget, Smith and Zurek 978-1-107-01618-7 19. Planetary Ring Systems Edited by Tiscareno and Murray 978-1-107-11382-4 20. Saturn in the 21st Century Edited by Baines, Flasar, Krupp and Stallard 978-1-107-10677-2 21. Mercury: The View after Messenger Edited by Solomon, Nittler and Anderson 978-1-107-15445-2 22. Chondrules: Records of Protoplanetary Disk Processes Edited by Russell, Connolly Jr. and Krot 978-1-108-41801-0

† Reissued as a paperback CHONDRULES Records of Protoplanetary Disk Processes

Edited by

SARA S. RUSSELL Natural History Museum, London

HAROLD C. CONNOLLY JR. Rowan University, New Jersey

ALEXANDER N. KROT University of Hawai’iatMānoa, Honolulu University Printing House, Cambridge CB2 8BS, United Kingdom One Liberty Plaza, 20th Floor, New York, NY 10006, USA 477 Williamstown Road, Port Melbourne, VIC 3207, Australia 314–321, 3rd Floor, Plot 3, Splendor Forum, Jasola District Centre, New Delhi – 110025, India 79 Anson Road, #06–04/06, Singapore 079906

Cambridge University Press is part of the University of Cambridge. It furthers the University’s mission by disseminating knowledge in the pursuit of education, learning, and research at the highest international levels of excellence.

www.cambridge.org Information on this title: www.cambridge.org/9781108418010 DOI: 10.1017/9781108284073 © Harold Connolly Jr., Alexander Krot and The Trustees of the Natural History Museum, London 2018 This publication is in copyright. Subject to statutory exception and to the provisions of relevant collective licensing agreements, no reproduction of any part may take place without the written permission of Cambridge University Press. First published 2018 Printed in the United Kingdom by TJ International Ltd. Padstow Cornwall A catalogue record for this publication is available from the British Library. Library of Congress Cataloging-in-Publication Data Names: Russell, Sara S. (Sara Samantha), 1966– editor. Title: Chondrules : records of protoplanetary disk processes / edited by Sara S. Russell, Natural History Museum, London, Harold C. Connolly Jr., Rowan University, New Jersey, and Alexander N. Krot, University of Hawaii, Manoa. Description: Cambridge, United Kingdom ; New York, NY : Cambridge University Press, 2018. | Series: Cambridge planetary science | Includes bibliographical references and index. Identifiers: LCCN 2017059979 | ISBN 9781108418010 (hardback : alk. paper) Subjects: LCSH: Chondrites (Meteorites) Classification: LCC QB758.5.C46 C456 2018 | DDC 549/.112–dc23 LC record available at https://lccn.loc.gov/2017059979

ISBN 978-1-108-41801-0 Hardback Cambridge University Press has no responsibility for the persistence or accuracy of URLs for external or third-party internet websites referred to in this publication and does not guarantee that any content on such websites is, or will remain, accurate or appropriate. Contents

List of Contributors page vii 1 Introduction 1 sara s. russell, harold c. connolly jr., and alexander n. krot

Part I Observations of Chondrules 9

2 Multiple Mechanisms of Transient Heating Events in the Protoplanetary Disk: Evidence from Precursors of Chondrules and Igneous Ca, Al-Rich Inclusions 11 alexander n. krot, kazuhide nagashima, guy libourel, and kelly e. miller

3 Thermal Histories of Chondrules: Petrologic Observations and Experimental Constraints 57 rhian h. jones, johan villeneuve, and guy libourel

4 Composition of Chondrules and Matrix and Their Complementary Relationship in Chondrites 91 dominik c. hezel, phil a. bland, herbert palme, emmanuel jacquet, and john bigolski

5 The Chondritic Assemblage: Complementarity Is Not a Required Hypothesis 122 brigitte zanda, e´ ric lewin, and munir humayun

6 Vapor–Melt Exchange: Constraints on Chondrite Formation Conditions and Processes 151 denton s. ebel, conel m. o’d. alexander, and guy libourel

7 Chondrules in Enstatite Chondrites 175 emmanuel jacquet, laurette piani, and michael k. weisberg

8 Oxygen Isotope Characteristics of Chondrules from Recent Studies by Secondary Ion Mass Spectrometry 196 travis j. tenner, takayuki ushikubo, daisuke nakashima, devin l. schrader, michael k. weisberg, makoto kimura, and noriko t. kita

v vi Table of Contents

9 26Al–26Mg Systematics of Chondrules 247 kazuhide nagashima, noriko t. kita, and tu-han luu

10 Tungsten Isotopes and the Origin of Chondrules and Chondrites 276 thorsten kleine, gerrit budde, jan l. hellmann, thomas s. kruijer, and christoph burkhardt

11 The Absolute Pb–Pb Isotope Ages of Chondrules: Insights into the Dynamics of the Solar Protoplanetary Disk 300 james n. connelly and martin bizzarro

12 Records of Magnetic Fields in the Chondrule Formation Environment 324 roger r. fu, benjamin p. weiss, devin l. schrader, and brandon c. johnson

Part II Possible Chondrule-Forming Mechanisms 341

13 Formation of Chondrules by Planetesimal Collisions 343 brandon c. johnson, fred j. ciesla, cornelis p. dullemond, and h. jay melosh

14 Making Chondrules by Splashing Molten Planetesimals: The Dirty Impact Plume Model 361 ian s. sanders and edward r. d. scott

15 Formation of Chondrules by Shock Waves 375 melissa a. morris and aaron c. boley

16 Evaluating Non-Shock, Non-Collisional Models for Chondrule Formation 400 alexander hubbard and denton s. ebel

17 Summary of Key Outcomes 428 harold c. connolly jr., alexander n. krot, and sara s. russell

Index 437 Colour plate section to be found between pages 246 and 247 Contributors

Conel M. O’D. Alexander Carnegie Institution of Washington, Washington DC, USA

John Bigolski City University of New York, New York, USA

Martin Bizzarro Natural History Museum of Denmark, Copenhagen, Denmark

Phil A. Bland Curtin University, Perth, Australia

Aaron C. Boley University of British Columbia, Vancouver, Canada

Gerrit Budde Institute for Planetology, University of Münster, Münster, Germany

Christoph Burkhardt Institute for Planetology, University of Münster, Münster, Germany

Fred J. Ciesla University of Chicago, Chicago, USA

James N. Connelly Natural History Museum of Denmark, Copenhagen, Denmark

Harold C. Connolly Jr. Rowan University, Glassboro, USA

Cornelis P. Dullemond Heidelberg University, Heidelberg, Germany

vii viii List of Contributors

Denton S. Ebel American Museum of Natural History, New York, USA

Roger R. Fu Harvard University, Cambridge, USA

Jan L. Hellmann Institute for Planetology, University of Münster, Münster, Germany

Dominik C. Hezel University of Cologne, Cologne, Germany

Alexander Hubbard American Museum of Natural History, New York, USA

Munir Humayun Florida State University, Tallahassee, USA

Emmanuel Jacquet National Museum of Natural History, Paris, France

Brandon C. Johnson Brown University, Providence, USA

Rhian H. Jones University of Manchester, Manchester, UK

Makoto Kimura Ibaraki University, Mito, Japan and National Institute of Polar Research, Tokyo, Japan

Noriko T. Kita University of Wisconsin–Madison, Madison, Wisconsin, USA

Thorsten Kleine Institute for Planetology, University of Münster, Münster, Germany

Alexander N. Krot University of Hawai’iatMānoa, Honolulu, USA List of Contributors ix

Thomas S. Kruijer Institute for Planetology, University of Münster, Münster, Germany and Lawrence Livermore National Laboratory, Livermore, USA

Éric Lewin Université Grenoble-Alpes and CNRS-INSU, ISTerre, Grenoble, France

Guy Libourel University of Côte d’Azur, Nice, France

Tutu H. Luu University of Bristol, Bristol, UK

H. Jay Melosh Perdue University, West Lafayette, USA

Kelly E. Miller University of Arizona, Tucson, USA

Melissa A. Morris State University of New York, Cortland, USA and Arizona State University, Tempe, USA

Kazuhide Nagashima University of Hawai’iatMānoa, Honolulu, USA

Daisuke Nakashima Tohoku University, Sendai, Japan

Herbert Palme Senckenberg Natural History Museum, Senckenberg, Germany

Laurette Piani Hokkaido University, Sapporo, Japan

Sara S. Russell Natural History Museum, London, UK x List of Contributors

Ian S. Sanders Trinity College, Dublin, Ireland

Devin L. Schrader Arizona State University, Tempe, USA

Edward R. D. Scott University of Hawai’iatMānoa, Honolulu, USA

Travis J. Tenner Los Alamos National Laboratory, Los Alamos, USA

Takayuki Ushikubo Japan Agency for Marine-Earth Science and Technology, Kochi, Japan

Johan Villeneuve Center for Petrographic and Geochemical Research, Vandoeuvre-lès-Nancy, France

Michael K. Weisberg Kingsborough Community College and The Graduate Center, City University of New York, USA and American Museum of Natural History, New York, USA

Benjamin P. Weiss Massachusetts Institute of Technology, Cambridge, USA

Brigitte Zanda IMPMC, Sorbonne Université, Muséum National d’Histoire Naturelle, CNRS and IMCCE, Observatoire de Paris, CNRS, Paris, France 1 Introduction

sara s. russell, harold c. connolly jr., and alexander n. krot

Abstract In this chapter, we review the history of chondrule research and introduce some of the basic concepts in the study of chondrules, including the classification of chondrites and the nomen- clature of chondrule types.

1.1 Introduction Chondrules are igneous-textured ferromagnesian silicate Æ Fe,Ni-metal Æ sulfides droplets that are ubiquitous in the majority of chondritic meteorites (chondrites). As survivors of the time when the early Sun was surrounded by a protoplanetary disk, their characteristics can provide clues to disk processes and planet formation. They formed from molten droplets, so their shape tends to be spherical, although they are also present as fragments in meteorites. A first-order observation of chondrules is that they have been melted and largely avoided evaporation, and therefore require conditions in which melts were stable. In this book, we present the latest mineralogical, petrologic, chemical, and isotopic observa- tions of chondrules that may place constraints on the environment in their formation regions in the protoplanetary disk. At the end of the book are chapters presenting models for chondrule formation. This book follows a workshop on the topic of Chondrules and the Protoplanetary Disk that was held on 27À28 February 2017 at the Natural History Museum in London, UK. We are grateful to the Natural History Museum, the Meteoritical Society, and the Royal Astronomical Society for their support of this workshop.

1.2 A Brief History of Chondrule Research These rounded inclusions were described in even the earliest recognized meteorite falls: in his description of the Benares meteorite fall in 1798, the English chemist Edward Howard remarked “Internally they consisted of a small number of spherical bodies, of a slate colour, embedded in a whitish gritty substance” (Howard, 1802). The word “chondrule” derives from the Ancient Greek χόνδρος or chondros, meaning grain. Gustav Rose (1798‒1873) coined the word “chondrite” in 1863, and soon after Tschermak began to call the rounded spherules within them “chondren,” which eventually became “chondrules” (see Connolly and Jones, 2017). An understanding that they were once molten droplets also came early; they were described by

1 2 Sara S. Russell, Harold C. Connolly Jr., and Alexander N. Krot

Henry Clifton Sorby as “drops of fiery rain” in 1877 during a talk at the then-new Natural History Museum in South Kensington, London. Chondrules have been the subject of research ever since. Merrill (1920) produced the first review of possible chondrule-forming mechanisms. Chondrules were the subject of a workshop in Houston in 1983, at which the chondrule was definitively defined and a book produced (King, 1983). At a later workshop in Albuquerque, New Mexico in 1993, definitive evidence was presented that chondrules formed in a flash heating event. This workshop resulted in a book on chondrules (Hewins, Jones, and Scott, Chondrules and the Protoplanetary Disk: Cambridge University Press, 1996). In this chapter, we cover some relevant basic concepts in meteoritics that may be useful for the nonexpert readers to make sense of succeeding chapters. An additional review of chondrule properties and some current chondrule-forming mechanisms can be found in Connolly and Jones (2017).

1.3 Primary Classification of Meteorites Meteorites are broadly divided into those that have been melted, and those which have remained unmelted since their accretion. Melted meteorites include irons, stony irons, and achondrites, or igneous-textured rocks. Specific chondrites have been identified as coming from asteroids, the Moon, and Mars. Some meteorites (called primitive achondrites) have been partly melted, and these may contain rare chondrules (e.g., Krot et al., 2014). The unmelted meteorites are called chondrites. Chondrites retain some of their primordial early solar system bulk chemical composition; within a factor of two they have a bulk composition similar to that of the Sun’s photosphere, excluding volatile gases (Palme et al., 2014). These meteorites are the hosts of most chondrules. Chondritic meteorites make up around 92 percent of our meteorite collections (Meteo- ritical Society, 2018). These are broadly divided into Ordinary + Rumuruti-like, Carbonaceous and Enstatite classes (Figure 1.1), based on differences in bulk chemical and oxygen isotopic compositions, along with the minor K and G grouplets (Weisberg et al., 2006, 2015). These meteorite classes are further subdivided into 13 groups that are classically presumed to have come from separate parent bodies. Recent work has suggested that may not always be the case, although the genetic relationship between groups is often controversial. For example, bulk chemical and oxygen isotopic compositions suggest the CVs and CKs may come from the same parent body (Greenwood et al., 2010), although bulk Cr isotopic compositions of CVs and CKs indicate that this may not be the case (Yin et al., 2017). Chondrites are composed of chondrules, refractory inclusions along with isolated metal and sulfide grains, all surrounded by a fine-grained matrix. The proportions of these components show very clear differences between classes (Weisberg et al., 2006). The size distribution of chondrules in each group is also very distinctive, to the extent that this characteristic can be used to assist classification. Chondrule sizes vary by almost an order of magnitude among chondrite classes and groups, from an average of 0.15 mm in COs to 1 mm in CVs (Jones, 2012). The classification of meteorites is described in more detail in Weisberg et al. (2006) and Krot et al. (2014). Introduction 3

Chemical Type

Ordinary Rumuritite Carbonaceous Kakangari Enstatite (OC) (RC) (CC) (KC) (EC)

H L LL CI CM CO CV CK CR CH CB EH EL

CR clan Petrological Type

123 4 5 6 7

Increasing metamorphism Melted Increasing aqueous alteration

Figure 1.1 A simplified classification scheme for chondrites. Kakangari chondrites (KC) define a grouplet; Rumuritite chondrites (RC) define a group.

1.4 Secondary Classification The primary classification of meteorites attempts to define groups that are clearly chemically and isotopically distinct and perhaps originate from separate parent bodies. In addition, meteor- ites have experienced geological processing while on their parent bodies. These include thermal and shock metamorphism, and aqueous alteration. Chondrites are designated a number (1–7) to describe this secondary processing (Van Schmus and Wood, 1967; Weisberg et al., 2006). Type 7s have been completely melted and, therefore, strictly speaking, are no longer chondrites. Types 2 and 1 have experienced increasing amounts of aqueous alteration, and types 3–6 have experienced increasing amounts of heating. Type 3 chondrites are, therefore, the meteorites that best retain their original petrographic features, mineralogy, mineral chemistry, and chemical compositions. Type 3s are further subdivided into types 3.0–3.9, indicating increasing amounts of parent body heating. Many chondrites have also experienced shock from asteroidal impacts. The extent of shock is highly variable, with some chondrites not displaying any evidence at all and some being melted or brecciated from the impact events. A shock classification scheme is best developed for the ordinary chondrites (Stöffler et al., 1991).

1.5 Tertiary Classification Some meteorites, especially those that are found (finds) rather than observed to fall (falls), experienced a further episode of changes while on the surface of the Earth. These processes have produced weathering effects, most notably oxidation of iron to form a series of iron oxides 4 Sara S. Russell, Harold C. Connolly Jr., and Alexander N. Krot and hydroxides (Bland et al., 2006) which has led to weathering schemes (e.g., Wlotzka, 1993). Additionally, some new minerals such as sulfates may form. For those interested in early solar system processes, it is important to recognize these effects so they can be disentangled from primordial early solar system processes.

1.6 Types of Chondrules 1.6.1 Textural Types

Chondrules exhibit a variety of textures that are a function of their initial composition and thermal histories. The most common textural type of chondrules is porphyritic. These are composed of euhedral to subhedral crystals embedded in a fine-grained or glassy mesostasis. The crystals can be olivine only (Porphyritic olivine (PO) chondrule), olivine plus pyroxene (POP) or, more rarely, except in enstatite chondrites, pyroxene only (PP). Another common texture for chondrules is barred olivine (BO). These chondrules are comprised of skeletal olivine crystals surrounded by mesostasis. The morphology of olivine is typically in the form of subparallel plates that are part of the same crystal (see Figure 3.1 for examples of different textures). Less common textures of chondrules include radial pyroxene (RP), cryptocrystalline (CC), and glassy (GC). Metal and sulfide grains are commonly found as blebs in chondrules, and tend to be concentrated around the outermost parts. In all cases, the mesostasis is composed of glass (in the lowest petrological type chondrites only) or a fine-grained mixture of crystals, most commonly high-Ca pyroxenes ((Ca,Mg,Fe)SiO3) and feldspar ((Na,Ca)(Si,Al)4O8). Chon- drules, especially porphyritic ones, often contain relict grains (see Chapter 2).

1.6.2 Chemical Types

Most chondrules from unaltered (type 3) chondrites are predominantly composed of olivine

(Mg, Fe)2SiO4 and low-Ca pyroxene (Mg, Fe)SiO3. Of these, the majority tend towards the magnesium end-member compositions (forsterite and enstatite) and are poor in volatile elements compared to solar compositions. These are called type I chondrules; they have a reduced mineralogy, and any iron within them is commonly in the form of metal blebs rather than combined in silicates. Some chondrules are more oxidized and contain ferromagnesian olivine and pyroxene (Fo > 10 mol% and En > 10 mol%, where Fo (at.%) = Mg/(Mg + Fe) Â 100 and En (at.%) = Mg/(Mg + Fe + Ca) Â 100); these are termed type II chondrules. Type I and type II chondrules coexist in most chondrite groups (with type I being always more abundant), but their proportions vary, with ordinary chondrites containing more type II chondrules than carbon- aceous or enstatite chondrites (Jones, 2012). These types are subdivided into type IA (silica- poor, olivine rich) and type IB (silica-rich, pyroxene-rich). Porphyritic chondrules are the predominant texture in most chondrites. However the exact proportions of textural types differs between meteorite groups, and, as discussed in Section 1.3, the size distribution also varies enormously between chondrite groups. Therefore, each group has sampled a unique reservoir, and models of the early solar system must account for this ability to separate reservoirs in space and/or time (Jones, 2012).

A minority of chondrules are Al-rich (>10 wt% Al2O3; Bischoff and Keil, 1983). Al-rich chondrules can contain anorthite, spinel, and/or glassy Al-rich mesostasis. They are found in all Introduction 5 chondrite groups, but are most abundant in CVs. Anorthite-rich chondrules (ARCs) are a subset of Al-rich chondrules (Kring and Holmen, 1988). ARCs may provide a genetic link between refractory inclusions and ferromagnesian chondrules (Krot and Keil, 2002). The related Na-rich chondrules, containing between 4 and 15 wt% Na2O, may have formed by melting together of Na-rich material and refractory-rich material (Ebert and Bischoff, 2016).

1.7 Refractory Inclusions Refractory inclusions are a rare component in most chondrite groups. There are two types of refractory inclusions: amoeboid olivine aggregates (AOAs) and calcium-aluminum-rich inclu- sions (CAIs). Amoeboid olivine aggregates are composed of Ca, Al-rich material enclosed by forsterite (Krot et al., 2004; note that olivine in AOAs is Ca-poor, see Sugiura et al., 2009). Calcium-aluminum-rich inclusions make up between 0 and 3 percent of chondrite groups (Hezel et al., 2008) and are composed of Ca, Al-rich oxides and silicates such as spinel (MgAl2O4), melilite ((Ca,Na)2(Al,Mg)(Si,Al)2O7), anorthite (CaAl2Si2O8), fassaite (Ca(Mg,Fe,Al)(Si, Al)2O6), and perovskite (CaTiO3) (MacPherson, 2014). CAIs can be either fluffy or compact in texture. Compact CAIs are igneous and formed by melting of solid precursors; they often show evidence for evaporation, suggesting that melts were not stable. Fluffy CAIs are aggre- gates of condensates from a hot gas. CAIs are very ancient and provide the time marker for the beginning of the solar system at 4567.3Æ0.16 Myr (Connelly et al., 2012).

1.8 Where Do Chondrites Come From? The texture of chondrites suggests they have never been melted, which points to their origin in a small minor body as a parent that has not had enough heat to melt. Asteroidal-sized bodies, therefore, are the likely parent of chondrites. In most cases, the exact body from which each meteorite comes from is unknown. In some instances, the orbital source of the meteorite can be determined from triangulating observations of fireballs (e.g., Brown et al., 2011), but this has only been successfully undertaken for a handful of meteorite falls. These orbital analyses typically indicate an origin from the main asteroid belt. The CI carbonaceous chondrite Orgueil has been suggested as coming from a comet from historical records of its fall (Gounelle et al., 2006). The Japanese Aerospace Exploration Agency (JAXA) Hayabusa mission provided the first proof that some chondrites originate in specific asteroids. This mission visited the 25143 Ito- kawa “S-type” asteroid in 2005 and brought fragments of this asteroid back to Earth in 2010. The returned material closely resembles LL5 ordinary chondrites, therefore providing a defini- tive link between S-type asteroids and the ordinary chondrites (Yurimoto et al., 2011).

References Bischoff, A., and Keil, K. (1983). Al-rich objects in ordinary chondrites: Related origin of carbonaceous and ordinary chondrites and their constituents. Geochimica et Cosmochi- mica Acta, 48, 693À709. Bland, P. A., Zolensky, M., Benedix, G., and Sephton, M. (2006) Weathering of chondritic meteorites. In D. Lauretta and H. McSween (Eds.), Meteorites and the Early Solar System II., 853À867. Tucson, AZ: University of Arizona Press. 6 Sara S. Russell, Harold C. Connolly Jr., and Alexander N. Krot

Brown, P., McCausland, P. J. A., Fries, M., et al. (2011). The fall of the Grimsby meteorite – I. Fireball dynamics and orbit from radar, video, and infrasound records. Meteoritics & Planetary Science, 46, 339–363. Connelly, J. N., Bizzarro, M., Krot, A. N., et al. (2012). The absolute chronology and thermal processing of solids in the solar protoplanetary disk. Science, 338, 651À655. Connolly, H. C. Jr., and Jones, R. (2017). Chondrules: The canonical and non-canonical view. Journal of Geophysical Research: Planets, 121, 1885À1899. Ebert, S., and Bischoff, A. (2016). Genetic relationship between Na-rich chondrules and Ca,Al- rich inclusions? – Formation of Na-rich chondrules by melting of refractory and volatile precursors in the Solar Nebula. Geochimica et Cosmochimica Acta, 177, 182À204. Friend, P., Hezel, D., and Mucerschi, D. (2016). The conditions of chondrule formation, Part II: Open system. Geochimica et Cosmochimica Acta, 173, 198À209. Gounelle, M., Spurny, P., and Bland, P. (2006). The orbit and atmospheric trajectory of the Orgueil meteorite from historical records. Meteoritics & Planetary Science, 41, 135À150. Greenwood, R. C., Franchi, I. A., Kearsley, A., and Alard, O. (2010). The relationship between CK and CV chondrites. Geochimica et Cosmochimica Acta, 74, 1684À1705. Hewins, R. H., Jones, R. H., and Scott, E. R. D. (Eds.). (1996). Chondrules and the Proto- planetary Disk. Cambridge, UK: Cambridge University Press. Hezel, D., Russell, S. S., Ross, A., and Kearsley, A. (2008). Modal abundances of CAIs: Implications for bulk chondrite elements and fractionations. Meteoritics and Planetary Science, 43, 1879–1894. Howard, E. (1802). Experiments and observations on certain stony and metalline substances, which at different times are said to have fallen on earth; also on various kinds of native iron. Philosophical Transactions of the Royal Society of London, 92, 168–212. Jones, R. H. (2012) Petrographic constraints on the diversity of chondrule reservoirs in the protoplanetary disk. Meteoritics & Planetary Science, 47, 1176–1190. King, A. (1983). Chondrules and Their Origins. Houston, TX: Lunar and Planetary Institute. Krot, A. N., Keil, K., Scott, E. R. D., Goodrich, C., and Weisberg, M. (2014). Classification of meteorites and their genetic relationships. In A. M. Davis (Ed.), Meteorites and Cosmo- chemical Processes. In H. D. Holland and K. K. Turekian (Eds.), Treatise on Geochemis- try (Second Edition). 1, 1À63. Oxford, UK: Elsevier. Krot, A. N., and Keil, K. (2002). Anorthite-rich chondrules in CR and CH carbonaceous chondrites: Genetic link between calcium-aluminium-rich inclusions and ferromagnesian chondrules. Meteoritics & Planetary Science, 37,91À111. Krot, A. N., Petaev, M. I., Russell, S. S., et al. (2004). Amoeboid olivine aggregates and related objects in carbonaceous chondrites: Records of nebular and asteroid processes. Chemie der Erde, 64, 185À239. Kring, D. A., and Holmen, B. A. (1988). Petrology of anorthite-rich chondrules in CV3 and CO3 chondrites. Meteoritics, 23, 282. Libourel, G., Krot, A. N., and Tissandier, L. (2006). Role of gas-melt interaction during chondrule formation. Earth and Planetary Science Letters, 251, 232–240. MacPherson, G. J. (2014) Calcium-aluminum-rich inclusions in chondritic meteorites. In A. M. Davis (Ed.), Meteorites and Cosmochemical Processes.InH.D.HollandandK.K.Turekian(Eds.), Treatise on Geochemistry (Second Edition), 1, 139‒179. Oxford, UK: Elsevier. Merrill, G. P. (1920). On chondrules and chondritic structure in meteorites. Proceedings of the National Academy of Sciences, 6, 449À472. Meteoritical Society, The. (2018). Meteoritical bulletin database. international society for meteoritics and planetary science. www.lpi.usra.edu/meteor/ Palme, H., Lodders, K., and Jones, A. (2014) Solar system abundances of the elements. In A. M. Davis (Ed.), Planets, Asteroids and Comets and the Solar System. In H. D. Holland and K. K. Turekian (Eds.), Treatise on Geochemistry (Second Edition), 2,15–36. Oxford, UK: Elsevier. Introduction 7

Stöffler, D., Keil, K., and Scott, E. R. D. (1991). Shock metamorphism of ordinary chondrite meteorites. Geochimica et Cosmochimica Acta, 55, 3845À3867. Sugiura, N., Petaev, M. I., Kimura, M., Miyazaki, A., and Hiyagon, H. (2009). Nebular history of amoeboid olivine aggregates. Meteoritics & Planetary Science, 44, 559–572. Van Schmus, W. R., and Wood, J. A. (1967). A chemical-petrologic classification for the chondritic meteorites. Geochimica et Cosmochimica Acta, 31, 747–754. Weisberg, M. K., Ebel, D. S., Nakashima, D., Kita, N. T., and Humayun, M. (2015). Petrology and geochemistry of chondrules and metal in NWA 5492 and GRO 95551: A new type of metal-rich chondrite. Geochimica et Cosmochimica Acta, 167, 269‒285. Weisberg, M. K., McCoy, T., and Krot, A. N. (2006), Systematics and Evaluation of Meteorite Classification. In D. Lauretta and H. McSween (Eds.), Meteorites and the Early Solar System II,19À52. Tucson, AZ: University of Arizona Press. Wlotzka, F. (1993). A weathering scale for the ordinary chondrites (abstract). Meteoritics, 28, 460. Yin, Q. -Z., Sanborn, M. E., and Ziegler, K. (2017). Testing the common source hypothesis for CV and CK chondrites (abstract). Lunar Planet. Sci. Conf., XLVIII, 1771. Yurimoto H., Abe K, Abe, M., et al. (2011). Oxygen isotopic composition of asteroidal materials returned from Itokawa by the Hayabusa mission. Science, 333, 1116À1119.

Part I Observations of Chondrules

2 Multiple Mechanisms of Transient Heating Events in the Protoplanetary Disk Evidence from Precursors of Chondrules and Igneous Ca, Al-Rich Inclusions alexander n. krot, kazuhide nagashima, guy libourel, and kelly e. miller

Abstract

In this chapter, we summarize our current knowledge of the mineralogy, petrography, oxygen- isotope compositions, and trace element abundances of precursors of chondrules and igneous Ca,Al-rich inclusions (CAIs), which provide important constraints on the mechanisms of transient heating events in the protoplanetary disk. We infer that porphyritic chondrules, the dominant textural type of chondrules in most chondrite groups, largely formed by incomplete melting of isotopically diverse solid precursors, including refractory inclusions (CAIs and amoeboid olivine aggregates (AOAs)), fragments of chondrules from earlier generations, and fine-grained matrix-like material during highly-localized transient heating events in dust-rich disk regions characterized by 16O-poor average compositions of dust (Δ17O~‒5‰ to +3‰). These observations preclude formation of the majority of porphyritic chondrules by splashing of differentiated planetesimals; instead, they are consistent with melting of dustballs during localized transient heating events, such as bow shocks and magnetized turbulence in the protoplanetary disk, and, possibly, by collisions between chondritic planetesimals. Like por- phyritic chondrules, igneous CAIs formed by incomplete melting of isotopically diverse solid precursors during localized transient heating events. These precursors, however, consisted exclusively of refractory inclusions, and the melting occurred in an 16O-rich gas (Δ17O~‒24‰) of approximately solar composition, most likely near the protosun. The U-corrected Pb–Pb absolute and Al–Mg relative chronologies of igneous CAIs in CV chondrites indicate that these melting events started contemporaneously with condensation of CAI precursors (4567.3 0.16 Ma) and lasted up to 0.3 Ma, providing evidence for the earliest transient heating events capable of melting refractory dustballs in the innermost part of the disk. There is no evidence that chondrules were among the precursors of igneous CAIs, which is consistent with an age gap between CAIs and chondrules. In contrast to typical (non–metal-rich) chondrites, the CB metal-rich carbonaceous chondrites contain exclusively magnesian nonporphyritic chon- drules formed during a single-stage event ~5 Ma after CV CAIs, most likely in an impact- generated gas–melt plume. Bulk chemical compositions of CB chondrules and equilibrium thermodynamic calculations suggest that at least one of the colliding bodies was differentiated. The uniformly 16O-depleted igneous CAIs in CB chondrites most likely formed by complete melting of preexisting refractory inclusions that was accompanied by gas–melt interaction in the plume. CH metal-rich carbonaceous chondrites represent a mixture of the CB-like materials

11 12 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

(magnesian skeletal olivine and cryptocrystalline chondrules and uniformly 16O-depleted igneous CAIs) formed in an impact plume and the typical chondritic materials (magnesian, ferroan, and Al- rich porphyritic chondrules, uniformly 16O-rich CAIs, and chondritic lithic clasts) that appear to have largely predated the impact plume event. We conclude that there are multiple mechanisms of transient heating events that operated in the protoplanetary disk during its entire lifetime and resulted in formation of chondrules and igneous CAIs.

2.1 Introduction

Based on bulk chemical and oxygen isotopic compositions, mineralogy, and petrography, 14 chondrite groups comprising four major classes are currently recognized: (1) carbonaceous – CI, CM, CR, CO, CV, CK, CH, and CB; (2) ordinary – H, L, and LL; (3) enstatite – EH and EL; and (4) Rumuruti – R; and K chondrite grouplet (Krot et al., 2014). Chondrules and associated Fe,Ni-metal, refractory inclusions [Ca,Al-rich inclusions (CAIs) and amoeboid olivine aggregates (AOAs)] and fine-grained matrix are the major components in most chondrites (Figure 2.1a–b; Scott and Krot, 2014). The important exceptions are CI chondrites, which virtually lack chondrules (Leshin et al., 1997), and metal-rich chondrites [CB, CH/CB, CH (Figure 2.1c‒f ), and two ungrouped meteorites – Grosvenor Mountains (GRO) 95551 and Northwest Africa (NWA) 5492] – which virtually lack fine-grained matrices (Krot et al., 2002; Weisberg et al., 2015). Chondrules, which are commonly composed of silicates Fe,Ni-metal Fe,Ni-sulfides, formed from molten or partially molten droplets that were freely floating in space before being accreted to the chondrite parent bodies. We note that igneous CAIs, the oldest solar system solids dated (Connelly et al., 2012), also satisfy this definition (see the following text) and, therefore, provide important constraints on the earliest transient heating events in the proto- planetary disk. The O-isotope compositions of chondrules and refractory inclusions in unme- tamorphosed chondrites (e.g., CR2‒3) are distinctly different (Figure 2.2), indicating that these chondritic components originated in isotopically distinct disk regions: 16O-poor planetary-like and 16O-rich solar-like, respectively. CAIs formed in a disk region, most likely near the protosun, that was exposed to irradiation by solar energetic particles, had high ambient temperature (at or above the condensation temperature of forsterite, >1,300 K), and had a dust-to-gas ratio that was approximately solar (1/100) (e.g., McKeegan et al., 2000; Scott and Krot, 2014; Sossi et al., 2017). They were subsequently transported radially away to the accretion regions of the chondritic and cometary planetesimals (e.g., Brownlee et al., 2006; Ciesla, 2010). In contrast, chondrules are thought to have formed in relatively cold (<1,000 K), dust-rich [dust-to-gas ratio ~(102–105) solar] regions of the protoplanetary disk (e.g., Cuzzi and Alexander, 2006; Alexander et al., 2008; Alexander and Ebel, 2012; Bischoff et al., 2017). Rubin (2010) pointed out many correlations among chondrule properties (textures, proportions of chondrules with igneous rims, average chondrule sizes, and so on) in different chondrite groups and used this as support for the idea that chondrules formed locally, at different heliocentric distances. It is not known, however, whether chondrule formation occurred only in the inner Solar System (i.e., inside Jupiter’s orbit) or took place in the outer disk (i.e., outside Jupiter’s orbit) as well (e.g., Walsh et al., 2011; Van Kooten et al., 2016; Budde et al., 2016a). Multiple Mechanisms of Transient Heating Events 13

Figure 2.1 (a‒c, e) Combined x-ray elemental maps in Mg (red), Ca (green) and Al (blue) and x-ray elemental maps in Ni (d, f ) of PCA 91082 (CR2), Tieschitz (H/L3), Hammadah al Hamra (HH) 237 (CB), and Isheyevo (CH/CB) chondrites. Typical chondrites consist of three major components: (1) chondrules 14 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

Figure 2.2 Oxygen–isotope compositions of refractory inclusions (CAIs and AOAs) and chondrules from CR carbonaceous chondrites. Individual refractory inclusions and chondrules are isotopically uniform. There are, however, significant differences in O-isotope compositions between these groups of objects, suggesting formation in isotopically distinct reservoirs. Refractory inclusions are 16O-enriched relative to chondrules, but 16O-depleted relative to the Sun. The inferred composition of the Sun is from McKeegan et al. (2011). The Terrestrial Fractionation (TF) and Carbonaceous Chondrite Anhydrous Mineral (CCAM) lines are shown for reference. Data from Krot et al. (2006a), Makide et al. (2009), Schrader et al. (2014), and Tenner et al. (2015).

Fine-grained matrices appear to be chemically and isotopically complementary to chondrules, suggesting their close genetic relationship – i.e., formation in the same disk region (e.g., Bland et al., 2005; Hezel and Palme, 2008, 2010; Palme et al., 2014, 2015; Ebel et al., 2016; Becker et al., 2015; Budde et al., 2016a, 2016b; Chapter 4). In contrast, Zanda et al. (2012; Chapter 5) advocate for the lack of chondrule-matrix complementarity and their origin in different disk regions. According to this hypothesis, chondrules formed near the protosun, were transported over large radial distances, and mixed with thermally unprocessed, primitive matrix in colder disk regions.

Caption for Figure 2.1 (cont.) + Fe,Ni-metal (Fe,Ni); (2) refractory inclusions (CAIs and AOAs); and (3) interstitial matrix (mx). The vast majority of chondrules have porphyritic olivine (PO), porphyritic pyroxene (PP) and porphyritic olivine-pyroxene (POP) textures; chondrules with radial pyroxene (RP) and barred olivine (BO) are less common. In contrast, the metal-rich CB, CH/CB, and CH carbonaceous chondrites lack matrix and contain abundant Fe,Ni-metal. CB chondrites contain exclusively magnesian nonporphyritic chondrules with skeletal olivine (SO) and cryptocrystalline (CC) textures, and rare refractory inclusions. CH chondrites contain chondrules with both porphyritic and nonporphyritic (SO and CC) textures. The CH/CB chondrite Isheyevo (Figures 2.1e and f) consists of the metal-rich, CB-like (left) and metal-poor, CH-like (right) lithologies with gradual boundaries between them. The origin of these lithologies is discussed by Morris, Garvie, and Knauth (2015). (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.) Multiple Mechanisms of Transient Heating Events 15

The prevalence of chondrules in minimally-altered chondrites suggests that chondrule formation was one of the most important processes that operated in the early Solar System. Among the currently discussed mechanisms of chondrule formation are (i) shock waves of different scale and nature (e.g., Desch et al., 2005; Morris et al., 2012; Morris and Boley, this volume), (ii) magnetized turbulence in the protoplanetary disk (McNally et al., 2013), and (iii) collisions between primitive or differentiated planetesimals (e.g., Krot et al., 2005a; Asphaug et al., 2011; Sanders and Scott, 2013, Chapter 14; Johnson et al., 2014, Chapter 13; Oulton et al., 2016). In this chapter, we review the recent data on the precursors of chondrules and igneous CAIs, mainly in carbonaceous chondrites, and discuss their significance for understanding the nature of transient heating events in the protoplanetary disk. Additional information on this topic can be found in several earlier reviews (e.g., Grossman et al., 1988; Hewins and Zanda, 2012; Connolly and Jones, 2016) and several chapters in this volume (Chapters 3–5, 7, and 8).

2.2 Porphyritic Chondrules and Their Precursors

Porphyritic chondrules are the dominant textural type of chondrules in most chondrite groups (e.g., Figure 2.1a–b); the only exceptions are metal-rich chondrites, which contain exclusively (in the case of CB chondrites) or a high proportion (in the cases of CH chondrites, GRO 95551, and NWA 5492) of nonporphyritic chondrules (Figure 2.1ce). Porphyritic chondrules crystal- lized from incomplete melts that retained several seed nuclei. Some of these nuclei can be recognized mineralogically, texturally or isotopically (oxygen) as relict grains (i.e., grains that did not crystallize from the host chondrule melt). Note that the open-system behavior of chondrule melts, which is supported by the mineralogy, O-isotope systematics, and bulk chemistry of chondrules (e.g., Libourel et al., 2006; Friend et al., 2016), may lead to disequilib- rium between the chemically and isotopically evolved chondrule melts and early crystallizing phenocrysts, which, as a result, may appear relict. Nonporphyritic (barred/skeletal olivine, radial pyroxene, and cryptocrystalline) chondrules crystallized from complete silicate melts and retained no mineralogical memory of chondrule precursors. However, in rare cases, the nature of precursors of nonporphyritic chondrules can be invoked from their bulk chemical compositions (e.g., Oulton et al., 2016).

2.2.1 Coarse-Grained Precursors of Porphyritic Chondrules

Based on textures, mineralogy, and chemical and O-isotope composition, various kinds of coarse (>10 µm) relict objects have been identified in porphyritic chondrules. These include refractory inclusions, chondrules from earlier generations and their fragments, and, possibly, fragments of thermally-processed planetesimals. Below, we summarize the major characteristics of these chondrule precursors and their significance for understanding chondrule formation.

2.2.1.1 Refractory Inclusions Refractory inclusions are mineralogically, chemically, and isotopically distinct from chondrules (e.g., MacPherson and Huss, 2005; Yurimoto et al., 2008; MacPherson, 2014), and can be easily 16 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller distinguished as relict grains in chondrules. However, reported relict CAIs and AOAs in chondrules are very rare (e.g., Bischoff and Keil, 1983a,1983b; Bischoff and Keil, 1984; Misawa and Fujita, 1994; Russell et al., 2000; Yurimoto and Wasson, 2002; Maruyama and Yurimoto, 2003; Russell et al., 2005; Krot et al., 2006b; Makide et al., 2009; Wakaki et al., 2013; Zhang et al., 2014; Nagashima et al., 2015a) and have not been systematically studied until recently (Krot et al., 2017a), possibly because refractory inclusions are a minor component (<3 vol.%) in most chondrite groups (Hezel and Palme, 2008). The most common relict minerals identified in chondrules, that are related to refractory inclusions, are spinel (MgAl2O4) and forsterite (Mg2SiO4). This may reflect the common presence of spinel in CAIs (MacPher- son, 2014) and its slower dissolution rate in silicate melts compared to melilite (åkermanite

(Ca2MgSi2O7) – gehlenite (Ca2Al2SiO7) solid solution) and Al,Ti-diopside, as well as the common presence of AOAs (~ 50% of all refractory inclusions) in most chondrite groups (Krot et al., 2004a). Detailed mineralogical, chemical and isotopic studies of the chondrules with relict refractory inclusions indicate that (i) porphyritic chondrules formed by incomplete melting of isotopically diverse solid precursors, (ii) refractory inclusions were among chondrule precur- sors, and (iii) porphyritic chondrules (at least some) postdate formation of refractory inclusions. These conclusions provide a possible explanation for the volatility-fractionated rare earth element (REE) patterns of some chondrules (e.g., Misawa and Nakamura, 1988), the CAI/ AOA-like O-isotope compositions of some chondrule olivines (Jones et al., 2004), and the origin of at least some Al-rich (bulk Al2O3 content > 10 wt%) chondrules (e.g., Bischoff and Keil, 1983a, 1983b, 1984; MacPherson and Huss, 2005; Zhang et al., 2014; Ebert and Bischoff, 2016). To investigate a large number of CAIs and chondrules in a relatively small number of polished sections of a chondrite group, Krot et al. (2017a) took advantage of the fine-grained nature of CH and CH/CB chondrite Isheyevo chondrites (hereafter CH chondrites) to systemat- ically study CAIs for possible effects of chondrule formation on their mineralogy and O-isotope compositions. Most CH CAIs are mineralogically and isotopically distinct from CAIs in other chondrite groups: (1) They are very refractory and composed of grossite (CaAl4O7), hibonite (CaAl12O19), perovskite (CaTiO3), gehlenitic melilite, and Al-rich pyroxene; anorthite (CaAl2- Si2O8) replacing melilite is rare (Bischoff et al., 1993; Kimura et al., 1993; Weber and Bischoff, 1994; Weber, Zinner, and Bischoff, 1995; Krot et al., 2008). (2) They show a bimodal 26 27 26 27 5 6 distribution of the inferred initial Al/ Al ratio [( Al/ Al)0], ~5 10 and <3 10 , 26 27 respectively; ~85% of CH CAIs have low ( Al/ Al)0. (3) They show a large range of 17 17 18 i i 16 i 16 Δ O values (= δ O – 0.52 δ O; where δ O=(O/ O)sample/( O/ O)SMOW 1) 1,000, i = 17 or 18; SMOW = Standard Mean Ocean Water), from 35‰ to 5‰; most individual CAIs, however, are isotopically uniform (Krot et al., 2008; Gounelle et al., 2009). In CH chondrites, Krot et al. (2017a) identified ~ 60 CAIs that were incompletely melted, most likely during formation of porphyritic chondrules. These include: (i) relict polymineralic CAIs in chondrules, (ii) CAIs surrounded by chondrule-like ferromagnesian silicate igneous rims, and (iii) objects intermediate between igneous Type C-like (anorthite-rich) CAIs and anorthite-rich chon- drules with clusters of relict spinel grains (Figures 2.3 and 2.4). Mineralogical and isotopic observations indicate that these CAIs were mixed with different proportions of ferromagnesian silicates and experienced incomplete melting and gas–melt interaction during chondrule forma- tion. These processes resulted in partial or complete dissolution of the CAI Wark-Lovering (WL) Multiple Mechanisms of Transient Heating Events 17

Figure 2.3 Backscattered electron (BSE) images of relict CAIs in porphyritic chondrules from CH chondrites. Minerals of relict refractory inclusions and host chondrules are labeled by black symbols in white boxes and by white symbols in black boxes, respectively. Relict CAIs in chondrules are dominated by grossite-rich, hibonite-rich, and spinel-rich types. They are typically surrounded by the layers of spinel and plagioclase. The spinel layers are remnants of Wark-Lovering (WL) rim layers; the plagioclase layers resulted from melting of meliliteAl-diopside and interaction with gaseous SiO. Numbers correspond to CAI-chondrule numbers listed in Figure 2.9. chd = chondrule; cpx = high-Ca pyroxene; gk = geikelite; grs = grossite; hib = hibonite; mel = melilite; mes = mesostasis; met = Fe, Ni-metal; ol = olivine; pl = plagioclase; pv = perovskite; px = low-Ca pyroxene; sp = spinel. 18 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

Figure 2.4 BSE images of melilite-rich CH CAIs melted to various degrees during formation of porphyritic chondrules with small addition of chondrule-like precursors (ferromagnesian silicatesFe,Ni-metal). Regions outlined in (a), (c), and (e) are shown in detail in (b), (d), and (f ), respectively. Melilite experienced melting and interaction with gaseous SiO, which resulted in its replacement by anorthite; spinel and hibonite largely avoided this melting. Wark-Lovering rims around CAIs experienced melting and O-isotope exchange with an 16O-depleted nebular gas. CAI minerals and chondrule minerals are labeled by black symbols in white boxes and by white symbols in black boxes, respectively. Numbers correspond to CAI-chondrule numbers listed in Figure 2.9. cpx = high-Ca pyroxene; hib = hibonite; mel = melilite; met = Fe,Ni-metal; ol = olivine; pl = plagioclase; pv = perovskite; px = low-Ca pyroxene; sp = spinel. Multiple Mechanisms of Transient Heating Events 19 rims in the host chondrule melts, replacement of melilite by Na-bearing plagioclase, dissolution and overgrowth of nearly end-member spinel by Cr- and Fe-bearing spinel, and O-isotope exchange of the melted portions of CAIs with the surrounding gas. Krot et al. (2017a) found that (1) only a small fraction (~5–10%) of CH CAIs experienced melting during chondrule formation; most CAIs show no evidence for being affected by chondrule melting events: they are surrounded by WL rims and preserved uniform O-isotope compositions. (2) The degree of CAI melting is highly variable and mineralogically controlled: grossite-rich, hibonite-rich, and spinel-rich CAIs preferentially survived chondrule melting and are commonly present as relict objects in porphy- ritic chondrules (Figure 2.3). Relict melilite-rich CAIs are rare, and often transformed into Type C-like CAIs/anorthite-rich chondrules with clusters of relict spinel grains (Figure 2.4). (3) Olivine and low-Ca pyroxene phenocrysts and plagioclase/mesostasis in the CAI-bearing chondrules show a larger range of Δ17O relative to those in the CH porphyritic chondrules without relict CAIs, probably due to dissolution of 16O-enriched relict CAIs (Figures 2.5 and 2.6). (4) The unmelted portions of relict CAIs are mineralogically and isotopically (oxygen and magnesium) similar to the CH CAIs apparently unaffected by chondrule melting (Figures 2.5 and 2.6). These observations indicate that CH porphyritic chondrules formed in a distinct disk region, where CH CAIs were present at the time of chondrule formation, but only a small fraction of these CAIs experienced melting. Because liquidus temperatures of typical (melilite-rich) CAIs are generally lower than those of magnesian porphyritic chondrules (Stolper, 1982; Hewins, 1997; Mendybaev et al., 2006), Krot et al. (2017a) conclude that CH porphyritic chondrules formed during localized heating events. Formation of metal-rich carbonaceous chondrites (CH, CB, and CR) in a distinct disk region has been recently inferred by Van Kooten et al. (2016) based on the whole-rock magnesium and chromium isotopic compositions.

2.2.1.2 Precursors of Al-Rich Chondrules in Ordinary Chondrites There are several varieties of Al-rich chondrules in ordinary and carbonaceous chondrites: Ca- Al-rich, (Ca,Na)-Al-rich, Na-Al-rich, and Na-Al-Cr-rich (e.g., Bischoff and Keil, 1983a, 1983b, 1984; Bischoff et al., 1989; Krot and Rubin, 1994; Russell et al., 2000; MacPherson and Huss, 2005; Ebert and Bischoff, 2016). There is a compositional continuum between Ca-Al- rich and Na-Al-rich chondrules (see Figure 3 in Ebert and Bischoff, 2016). Among these types,

Na-Al-rich chondrules (4‒15 wt% Na2O) are relatively common in ordinary chondrites of low (3.2‒3.8) petrologic types (Figure 2.7; see also Figures 1 and 2 in Ebert and Bischoff, 2016); they are virtually absent in carbonaceous chondrites. Although some enrichment in sodium of these chondrules could be explained by fluid-assisted thermal metamorphism on the chondrite parent bodies (e.g., Russell et al., 2000; Grossman and Brearley, 2005; Brearley and Krot, 2012), Ebert and Bischoff (2016) provided compelling evidence that this is not the case for the majority of Na-Al-rich chondrules composed of olivine, Al,Ti-rich pyroxene and spinel pheno- crysts embedded in an albite- and nepheline-normative glassy mesostasis: no gradient in sodium concentrations between chondrule edges and their interiors was detected. The Na-Al-rich chondrules have signatures of ultrarefractory, group II, and group III REE patterns. Based on these observations and major element chemical compositions, Ebert and Bischoff (2016) concluded that these chondrules formed by melting of CAIs and AOAs altered to various degrees in a nebular setting. This alteration resulted in replacement of melilite and anorthite by 20 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

Figure 2.5 Δ17O of CAIs incompletely melted during formation of CH porphyritic chondrules. These include (i) CAIs surrounded by chondrule-like ferromagnesian silicate igneous rims, (ii) relict polymineralic CAIs in chondrules, and (iii) objects intermediate between Type C-like CAIs and plagioclase-rich chondrules with clusters of relict spinel grains. Olivine and low-Ca pyroxene phenocrysts and plagioclase/mesostasis in the CAI-bearing chondrules show a larger range of Δ17O relative to those in the CH porphyritic chondrules without relict CAIs (see Figure 2.6). The unmelted portions of relict CAIs are isotopically similar to the CH CAIs surrounded by primordial WL rims and apparently unaffected by chondrule melting (see Figure 2.6). The terrestrial fractionation line (TF) is shown for reference. Numbers correspond to ID numbers of individual CAIs; some of them are shown in Figures 2.3 and 2.4. AB = addibischoffite; cpx = high-Ca pyroxene; grs = grossite; hib = hibonite; mel = melilite; ol = olivine; pl = plagioclase; pv = perovskite; px = low-Ca pyroxene. Data from Krot et al. (2016a). (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.) nepheline (NaAlSiO4) and sodalite (Na4Al3Si3O12Cl). A similar conclusion was reached by MacPherson and Huss (2005). (Note that some Na-Al-rich chondrules may have formed by splashing of mesostasis of incompletely crystallized chondrules (Bischoff et al., 1989; Ebert and Bischoff, 2016)). According to this hypothesis, high sodium content in chondrule melts was retained due to high partial pressure of sodium (10‒6–10‒3 bar) in the chondrule-forming regions that could have been created by evaporation of dust-rich nebular regions (e.g., Cuzzi and Alexander, 2006; Alexander et al., 2008) or by melting of planetary embryos that had a Na- rich atmosphere (Morris et al., 2012; Herbst and Greenwood, 2016). Although the presence of CAI/AOA-like materials among precursors of Al-rich chondrules is consistent with the observed systematic enrichment of their minerals in 16O compared to ferro- magnesian chondrules (Figure 2.8), the mixing of CAI/AOA-like and ferromagnesian chondrule Multiple Mechanisms of Transient Heating Events 21

Figure 2.6 Oxygen–isotope compositions of porphyritic chondrules and CAIs surrounded by primordial Wark-Lovering rims (unaffected by chondrule melting events) from the CH and CB metal-rich carbonaceous chondrites. Individual CAIs and their WL rims have similar oxygen isotopic compositions; there are, however, inter-CAI compositional differences. Most porphyritic chondrules analyzed are compositionally uniform. Several type II chondrules contain relict olivine grains (indicated by arrows) enriched in 16O compared to the host chondrule phenocrysts. The terrestrial fractionation line (TF) is shown for reference. Numbers correspond to ID numbers of individual CAIs. cpx = high-Ca pyroxene; grs = grossite; hib = hibonite; mel = melilite; ol = olivine; pv = perovskite; px = low-Ca pyroxene. Data from Krot et al. (2010, 2017b). (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.) precursors alone cannot explain these characteristics: (1) extensive gas–melt O-isotope exchange is required (Russell et al., 2000); (2) moreover, replacement of melilite and anorthite in CAIs and AOAs by nepheline and/or sodalite appears to be purely an asteroidal process (e.g., Tomeoka and Itoh, 2004; Brearley and Krot, 2012). Therefore, melting of altered refractory inclusions invoked by MacPherson and Huss (2005) and Ebert and Bischoff (2016) to explain the origin of Na-Al- rich chondrules is possible only if Na-Al-rich chondrules formed by melting of metasomatically altered asteroidal material (i.e., relatively late). If this is the case, the Na-Al-rich chondrules are expected to show no clear evidence for live 26Al. An alternative explanation would be an interaction between Na- and Si-rich gas and initially Na-free/poor Ca-Al-rich chondrule melt under high partial pressure of sodium (~10‒4 bar; Georges, Libourel, and Deloule, 2000; Matthieu, 2009; Matthieu et al., 2011). It was experimentally shown that under high sodium partial pressure, condensation of silica and sodium into Ca,Al-rich silicate melt could result in very high

Na2O content (up to 15 wt%) in the melt. Mathieu (2009) and Mathieu et al. (2011) have shown indeed that increasing polymerization in both Al2O3-free or Al2O3-bearing melts lead to a drastic 22 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

Figure 2.7 BSE images of a (Ca,Na)-Al-rich porphyritic chondrule from the LL3.2 chondrite Vicência. Regions outlined in (a) are shown in detail in (b) and (c). The chondrule is composed of coarse and skeletal crystals of high-Ca pyroxene (cpx), skeletal forsterite (ol), spinel (sp), and Na-rich glassy mesostasis (mes). increase in the sodium solubility in such silicate melts for a given sodium partial pressure in the gas phase; albite- and nepheline-normative glassy mesostasis in Al-rich chondrules being by definition close to fully polymerized melt. If this explanation is correct, Na-Al-rich chondrules are 26 27 expected to have similar ( Al/ Al)0 as ferromagnesian chondrules. Multiple Mechanisms of Transient Heating Events 23

Figure 2.8 Oxygen–isotope compositions of ferromagnesian (type I and type II) and Al-rich chondrules from ordinary chondrites. Data from Kita et al. (2010) and Russell et al. (2000).

2.2.1.3 Chondrules and Chondrule Fragments from Earlier Generations There are several pieces of evidence indicating that chondrules and chondrule fragments from earlier generations were one of the major coarse-grained chondrule precursors. These include (i) chondrule-like igneous rims around type I and type II chondrules (Figures 2.9 and 2.10), (ii) relict ferroan and/or “dusty” olivine grains in type I chondrules (Figure 2.11a), and (iii) coarse relict magnesian olivines in type II chondrules (Figure 2.11b). Igneous rims around chondrules provide a clear evidence for the repeatable nature and variable magnitude of transient heating events during chondrule formation (Wasson et al., 1994; Krot and Wasson, 1995; Rubin and Krot, 1996; Krot et al., 2004b). In most cases, igneous rims around type I and type II chondrules are compositionally similar to their host chondrules, suggesting repeatable melting events under similar redox conditions. Occasionally, type I chondrules in carbonaceous chondrites are surrounded by ferroan igneous rims compos- itionally similar to type II chondrules (Fig. 2.9), indicating melting under more oxidizing conditions. This process may have transformed some type I chondrules into type II chondrules, as was theoretically predicted by Grossman et al. (2012) and experimentally demonstrated by Villeneuve et al. (2015). Variable redox conditions may have also played an important role in the formation of chondrules in enstatite chondrites (Chapter 7). Chondrule pyroxenes are split into two categories: minor element-rich (MER; also called “red pyroxene” for its cathodoluminescence (CL) signature), and minor element-poor (MEP; also called “blue pyroxene”)(Simonetal., 2016). These different pyroxene populations may reflect different stages of a temporal sequence, with MER enstatite forming via solid-state reduction of enstatite and the subsequent condensation of MEP rims under slightly more oxidizing conditions (Weisberg et al., 1994; Simon et al., 2016). Weisberg et al. (1994) report a third kind of pyroxene in enstatite chondrite chondrules that is relatively FeO-rich (so-called “black pyroxene”) with Fs 6, which they argue is the precursor material for MER enstatite based on its ubiquitous 24 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

Figure 2.9 (a‒f ) BSE images of type I chondrules surrounded by type II chondrule-like igneous rims from the GRA (Graves Nunataks) 95229 (CR), Murchison (CM), and Y-81020 (CO) chondrites. Regions outlined in (a), (c), and (e) are shown in detail in (b), (d), and (f ), respectively. The igneous rim in GRA 95229 studied by ion microscopy contains abundant 16O-rich relict grains (see Figure 2.11). (g, h) BSE image (g) and x-ray elemental map in Fe (h) of a ferroan olivine chondrule reduced in type I chondrule melt. cpx = high-Ca pyroxene; mes = mesostasis; ol = olivine; px = low-Ca pyroxene; sf = sulfide. Multiple Mechanisms of Transient Heating Events 25

Figure 2.9 (cont.)

Figure 2.10 BSE images of type I chondrules surrounded by an igneous type II chondrule-like igneous rim from the CR chondrite GRA 95229.

(though minor) presence and similar minor element fractionation patterns. Thus, enstatite chondrite chondrule precursors appear to incorporate fragments formed under varying redox conditions. The relict nature of ferromagnesian olivines and pyroxenes in chondrules can be inferred from their textures, as well as their chemical and O-isotope compositions (e.g., Rambaldi, 1981; Rambaldi and Wasson, 1982; Rambaldi et al., 1983; Rubin and Krot, 1996; Jones, 1996; Jones and Danielson, 1997; Leroux et al., 2003; Kunihiro, Rubin, and Wasson, 2005; Kita et al., 2010; Hewins and Zanda, 2012; Hewins, Zanda, and Bendersky, 2012; Ushikubo et al., 2012; Schrader et al., 2013; Tenner et al., 2013, 2015, Chapter 8; Nagashima et al., 2013, 2016). Olivine and low-Ca pyroxene phenocrysts in individual chondrules from unmetamorphosed 26 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

Figure 2.11 BSE images of (a) type I chondrule with abundant relict dusty olivine grains from Semarkona and (b, c) type II chondrule with a relict AOA and a relict magnesium-rich olivine grains from DOM (Dominion Range) 08004. Region outlined in (b) is shown in detail in (c). The relict AOA consists of 16O- rich (Δ17O~‒25‰) forsteritic olivines with 120 triple junctions (indicated by white arrows in “c”). The coarse relict olivine is 16O-depleted (Δ17O~‒5‰) relative to the AOA, but 16O-enriched relative to olivine phenocrysts of the host chondrule (Δ17O~‒2‰). mes = mesostasis; met = Fe,Ni-metal; ol = olivine; px = low Ca pyroxene; sf = sulfide. Multiple Mechanisms of Transient Heating Events 27 chondrites have similar O-isotope compositions, suggesting crystallization from isotopically uniform melts (e.g., Chapter 8). Oxygen isotopic compositions of the relict magnesium-rich olivines and low-Ca pyroxenes in type II chondrules are distinctly different from those of the host chondrule phenocrysts, but are generally similar to a compositional range of phenocrysts in type I chondrules (Figure 2.12), consistent with formation of these relict grains by fragmentation of type I chondrules. Differences in oxygen isotopic compositions among magnesium-rich olivines in type I chondrules are widely used for distinguishing relict grains from chondrule phenocrysts (e.g., Ushikubo et al., 2012; Schrader et al., 2013; Tenner et al., 2013, 2015, Chapter 8). However, the nature of O-isotope heterogeneity in individual type I chondrules is not always clear. The apparent relict grains often have euhedral outlines and in BSE images are indistinguishable from chondrule phenocrysts, suggesting an overgrowth either of true relict olivine grains by olivine from the host chondrule melt or of chondrule phenocrysts that crystallized from an 16O-enriched chondrule melt prior to its significant exchange with the surrounding gas. As a result, the reported wide compositional ranges of chondrule “phenocrysts” (Δ17O from ‒5‰ to ‒25‰) within individual chondrules may represent either a mixture of 16O-rich relict olivine grains (fragments of chondrules and/or AOAs) and 16O-depleted olivine phenocrysts or an evolution of

Figure 2.12 Δ17O of relict ferromagnesian olivines in type I and type II chondrules from CR and CO chondrites. Outlined are compositional ranges of chondrule phenocrysts and refractory inclusions in CR and CO chondrites. Data from Schrader et al. (2013) and Tenner et al. (2013, 2015). Δ17O of most relict olivines are within the compositional ranges of chondrule phenocrysts, consistent with the relict grains being fragments of chondrules from earlier generations. Some relict olivine grains are significantly 16O- enriched relative to typical chondrules, suggesting genetic relationship with either 16O-enriched chondrules or refractory inclusions. Δ17O of olivine, low-Ca and high-Ca pyroxene, and plagioclase phenocrysts in two CR chondrules, #1 and #3, containing relict 16O-rich CAI-like spinel and anorthite are largely outside the compositional range of the CR chondrule phenocrysts. These observations provide evidence for a lack of complete O-isotope equilibrium between the chondrule melt and the nebular gas. Data from Makide et al. (2009). 28 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller the O-isotope composition of the chondrule melt (e.g., Jones et al., 2004). Olivine and pyroxene phenocrysts in Al-rich chondrules containing extensively melted relict CAIs are often 16O-enriched relative to phenocrysts in ferromagnesian chondrules (Figure 2.12), indicating a lack of complete O-isotope equilibrium between the chondrule melt and the surrounding nebular gas. In some cases, the outlines of possible relict forsteritic olivines in type I chondrules can be distinguished using CL and/or concentrations of minor elements (Ca, Al, Cr) (e.g., Steele, 1998; Weinbruch et al., 2000; Kita et al., 2010). Combined studies of chondrule olivines by CL, electron microprobe analysis, and ion microprobe analysis (SIMS) are required to understand the true nature of relict olivine grains in type I chondrules as well as to quantify their real proportions (Libourel et al., in preparation).

2.2.1.4 Fragments of Thermally Processed Planetesimals 2.2.1.4.1 Clasts with Granoblastic Textures Libourel and Krot (2006) described coarse- grained olivine-metal clasts inside type I porphyritic chondrules from the CV chondrite Vigar- ano (Figure 2.13). Olivine grains in these clasts show triple junctions with interfacial angles of ~120, indicative of granoblastic equilibrium textures requiring prolonged thermal annealing at high temperature. Since granoblastic textures cannot be produced by crystallization from chondrule melts or by thermal annealing in the chondrule-forming regions characterized by a relatively low ambient temperature, Libourel and Krot (2006) concluded that these clasts represent fragments of thermally processed planetesimals that experienced either strong thermal metamorphism or igneous differentiation. Subsequently, these fragments were incorporated into chondrule precursors and experienced melting and dissolution-crystallization in chondrule melts resulting in formation of “classical” porphyritic textures. [A similar conclusion was reached by Sokol et al. (2007) who described igneous-textured clasts (granitoids and andesitic fragments) in ordinary chondrite regolith breccias Adzhi-Bogdo (LL3‒6) and Study Butte (H3‒6)]. This hypothesis was tested by Whattam et al. (2008), who experimentally reproduced porphyritic textures by melting of thermally-annealed olivine clasts with granoblastic textures in a chondrule-like silicate melt. Due to the importance of these observations, detailed TEM (transmission electron micro- scopy) studies should be implemented to look for the occurrence of melt/glassy materials trapped at olivine-olivine grain boundaries. Our preliminary TEM study indeed reveals some material at grain boundaries in olivine triple junction clasts from Vigarano that could be interpreted either as residual or chondrule melt (Figure 2.13c). If the presence of chondrule melt along olivine grain boundaries is confirmed, then the textures of coarse-grained clasts are not truly granoblastic, and, could have resulted from crystallization of chondrule melt without subsequent annealing. More work is needed to resolve this issue.

2.2.1.4.2 Chromite-Rich Chondrules in Equilibrated Ordinary Chondrites: Formation by Melting of Metamorphosed Ordinary Chondrite-Like Materials Some chondrules in equilibrated ordinary chondrites (EOCs) of petrologic types 4‒6 contain a high abundance (up to 50 vol.%) of chromite and/or Cr-spinel grains associated with albitic plagioclase (Ramdohr, 1967; Krot et al., 1993). These chondrules have Al-rich bulk chemical compositions and very high Cr/Mg ratios [(180‒750) solar]. Equilibrium thermodynamic calculations Multiple Mechanisms of Transient Heating Events 29

Figure 2.13 (a) X-ray elemental map in Mg Kα and (b) BSE image of a magnesian porphyritic olivine- pyroxene (Type I) chondrule surrounded by a fine-grained matrix-like rim (FGR) from the CV carbonaceous chondrite Vigarano. Region outlined in (a) is shown in detail in (b). The chondrule contains a coarse-grained olivine Fe,Ni-metal fragment. Olivine grains in the fragment show triple junctions with interfacial angles of ~120 (arrows in (b)), indicative of granoblastic equilibrium textures. The granoblastic textures are not expected to have been produced by crystallization from rapidly cooling chondrule melts and require prolonged thermal annealing, most likely in an asteroidal setting (images from Libourel and Krot, 2006)). (c) TEM image at olivine grain boundary; location of TEM section extracted using focused ion beam (FIB) is indicated in (b). Notice the presence of unknown, possibly glassy material at the grain boundary as well as the remarkably low density of dislocations in the two adjacent olivine crystals. FGR = fine-grained rim; met = Fe,Ni-metal; ol = olivine; px = low-Ca pyroxene. 30 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller indicate that high Cr/Mg ratios observed in chromite-rich chondrules cannot be produced in the solar nebula and require unusual precursors (Krot et al., 1993). This is supported by the observations that no Cr-rich chondrules are known in the least metamorphosed ordinary chondrites of petrologic type 3.0; only small (<10 µm), rare chromite grains occur in type II chondrules. At the same time, chromite is a typical metamorphic mineral in EOCs, where it commonly associates with other metamorphic minerals, such as albitic plagioclase, phosphates, and ilmenite (Bunch et al., 1967; Snetsinger and Keil, 1969; Wlotzka, 2005). The origin of chromite-rich chondrules is controversial. Wlotzka (2005) suggested they resulted from Fe- and Cr-enrichment during thermal metamorphism of primary igneous spinel, which is observed in Al-rich chondrules from ordinary chondrites (Russell et al., 2000). In contrast, Krot and Rubin (1993) and Rubin (2003) suggested that chromite-rich chondrules formed by impact melting of metamorphic chromite-rich precursors. The second hypothesis appears to be consistent with (i) the common presence of fine-grained chromite-plagioclase- phosphate regions in heavily-shocked ordinary chondrites (Rubin, 2003), and (ii) the presence of a relict metamorphic fragment composed of chromite, ilmenite, albitic plagioclase, Fe,Ni- metal, and troilite inside a chromite-rich chondrule from the H3‒5 regolith breccia Magombedze (Figure 2.14). Oxygen–isotope measurements of Cr-spinel in chromite-rich chondrules using SIMS can potentially provide a test of both hypotheses. Spinel in the Al-rich chondrules from ordinary chondrites is 16O-enriched relative to normal ferromagnesian porphyritic chondrules, most likely reflecting CAI-like precursors (Russell et al., 2000; Kita et al., 2010). If Cr-spinels in chromite-rich chondrules resulted from Al + Mg $ Fe + Cr interdiffusion during thermal metamorphism (Wlotzka, 2005), these spinels could have preserved an 16O-rich isotopic signature of the CAI precursors. In contrast, if chromite-rich chondrules formed by melting of EOC precursors (Krot and Rubin, 1993; Rubin, 2003), the Cr-spinels are expected to have Δ17O values similar to those of bulk EOCs (Clayton et al., 1991).

2.2.2 Fine-Grained Precursors of Porphyritic Chondrules

Because of the high dissolutions rates of olivine in chondrule-like silicate melts (Soulié et al., 2012; Soulié, Libourel, and Tissandier 2017), relict 1‒10 µm-sized olivine grains would have survived only in chondrules that experienced a very small degree of melting. Chondrules in CV, CR, and ordinary chondrites are often surrounded by ferromagnesian silicate igneous rims, which are typically finer grained than the host chondrules (Rubin and Wasson, 1987; Krot and Wasson, 1995; Rubin and Krot, 1996; Krot et al., 2004b). These rims appear to have formed by melting of relatively fine-grained (<10 µm) solids that accreted on the surface of previously solidified chondrules (Krot and Wasson, 1995). Therefore, igneous chondrule rims could potentially provide important constraints on the nature of the fine-grained chondrule precursors. In situ O-isotope measurements of the relatively fine-grained chondrule igneous rims is challenging, but became possible with the development of two ion microscopes (the large- geometry Cameca ims-1270 and the ims-1280 secondary ion mass spectrometer (SIMS) combined with a stacked CMOS active pixel sensor (SCAPS) detector by H. Yurimoto’s group at Hokkaido University (Nagashima et al., 2001). Using this technique, Nagashima et al. (2013, 2016; Nagashima et al., 2014; Nagashima et al., 2015b) reported on O-isotope compositions of Multiple Mechanisms of Transient Heating Events 31

Figure 2.14 BSE images of a Cr-rich chondrule from Magombedze (H3‒5 regolith breccia). The boxed region from image (a) is shown enlarged in image (b). The chondrule consists of olivine (ol), euhedral chromium-spinel (Cr-sp) grains, and plagioclase-normative mesostasis (dark gray); it contains a relict fragment composed of anhedral chromite (chr), ilmenite (ilm), troilite (tr), kamacite (km), phosphate (ph), and albitic plagioclase (pl). The mineralogy and mineral chemistry of minerals in the fragment are similar to those in metamorphosed ordinary chondrites.

ferroan igneous rims around magnesian porphyritic (type I) chondrules in Acfer 094 (ungrouped), CO and CR carbonaceous chondrites. These rims are mineralogically and chem- ically similar to type II chondrules and composed mainly of ferroan olivines, Na-rich mesos- tasis, and sulfides (Figure 2.9). A similar ferroan igneous rim was also described around an AOA in a CO chondrite, suggesting that not only type I chondrules but also AOAs were reprocessed during type II chondrule-forming events (Nagashima et al., 2015a). These studies revealed the presence of abundant µm-sized relict 16O-rich olivine grains isotopically similar to those in CAIs and AOAs in the igneous chondrule rims (Figure 2.15). These observations indicate that (i) fine-grained chondrule precursors were isotopically heterogeneous, and (ii) some igneous rims around type I chondrules were melted under the redox conditions similar to those recorded by type II chondrules, suggesting that type I chondrules were among type II chondrule precursors (see also Villeneuve et al., 2015). 32 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

Figure 2.15 Oxygen–isotope compositions of a magnesium-rich porphyritic olivine-pyroxene chondrule surrounded by a ferroan olivine igneous rim from the CR chondrite GRA 95229 (see Figures 2.7a and b). Olivine and low-Ca pyroxene in the host chondrule are 16O-enriched relative to those in the igneous rim. The igneous rim contains abundant micron-sized 16O-rich relict olivines overgrown by ferroan olivines. Fe- ol = ferroan olivine; Fe-cpx = ferroan high-Ca pyroxene; Mg-ol = magnesian olivine; Mg-px = magnesian low-Ca pyroxene. Data from Nagashima et al. (2013).

2.2.3 Was Matrix Material among Chondrule Precursors?

The inferred presence of fine-grained material among chondrule precursors (see Section 2.2) and the proposed genetic relationship between chondrules and matrix based on their chemical and isotopic complementarity (e.g., Bland et al., 2005; Hezel and Palme, 2008, 2010; Palme et al., 2014, 2015; Becker et al., 2015; Budde et al., 2016a, 2016b; Chapter 4 Ebel et al., 2016) may imply that the fine-grained precursors could have been similar to matrix in primitive (unaltered and unmetamorphosed) chondrites. The presence of fine-grained matrix-like dust in the chondrule-forming regions is supported by the presence of microchondrules in fine-grained matrix-like rims surrounding chondrules (Krot and Rubin, 1996; Bigolski et al., 2016; Dobrica and Brearley, 2016). The microchondrules appear to have formed by melting of the host chondrule peripheries and to have subsequently become entrained in the surrounding dusty environment as the chondrules were accreting the fine-grained rims (Krot and Rubin, 1996). Nagashima et al. (2015b) reported on O-isotope compositions of chondrules, chondrule igneous rims, and matrix in the Kakangari (K) chondrite (Figure 2.16). It was found that (i) matrix olivines and low-Ca pyroxenes show a bimodal distribution of O-isotope compositions: 16O-rich solar-like and 16O-poor planetary-like. The 16O-rich olivines and low-Ca pyroxenes are similar to those of low-Ca pyroxene-bearing AOAs (Krot et al., 2005c) and most likely originated in the CAI/AOA-forming region. (ii) The 16O-poor matrix olivines and pyroxenes are similar to olivine and low-Ca pyroxene phenocrysts in Kakangari chondrules. (iii) Like Kakangari matrix, the chondrule igneous rims show a bi-modal distribution of O-isotope compositions: micron-sized 16O-rich olivine grains are overgrown by 16O-poor olivines and low-Ca pyroxenes. Based on these observations, Nagashima et al. (2015b) suggested that the Multiple Mechanisms of Transient Heating Events 33

Figure 2.16 (a) Combined x-ray elemental map in Mg (red), Ca (green), and Al (blue) of the Kakangari (K) chondrite. Region outlined in (a) is shown in detail in (d). (b, c) Oxygen–isotope compositions of olivine and low-Ca pyroxene in chondrules and matrix from Kakangari. There are two isotopically distinct populations of matrix grains, 16O-rich and 16O-poor. Chondrules are dominated by 16O-poor silicates, but contain relict 16O-rich olivines isotopically similar to those in the matrix. 34 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

Figure 2.16 (cont.) (d) Combined x-ray elemental map in Mg (red), Si (green) and Fe (blue), (e) BSE image, and (f, g) secondary ion images in “f” 27Al‒ and (g) δ18O of a porphyritic olivine-pyroxene chondrule and its coarse-grained igneous rim composed of magnesian olivine, low-Ca and high-Ca pyroxenes, and mesostasis. The rim contains abundant relict 16O-rich olivine grains (some of them are indicated by red arrows) overgrown by 16O-poor olivines. cpx = high-Ca pyroxene; mes = mesostasis; met = Fe,Ni-metal; ol = olivine; px = low-Ca pyroxene. Data from Nagashima et al. (2015b). (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.)

Kakangari chondrules and the 16O-poor matrix grains originated in the same gaseous reservoir, and the Kakangari matrix was probably one of the chondrule precursors. Formation of porphyritic chondrules by melting, evaporation, and condensation processes of clusters of fine-grained matrices and primordial chondrules (i.e., chondrules from earlier generations) has been experimentally studied in a Knudsen cell by Imae and Isobe (2017).

2.2.4 Maintaining Chondrule Diversity

One of the fundamental constraints imposed on the origin of chondrules is their overwhelming chemical, mineralogical, and isotopic diversity (for a review, see Jones, 2012). This may result from closed-system formation of chondrules from heterogeneous precursor materials (e.g., Hezel and Palme, 2007). Alternatively, in an open system, differences in precursor size (Friend Multiple Mechanisms of Transient Heating Events 35 et al., 2016), incorporation of multiple chondrule generations (see Section 2.1.2.), different cooling rates (Jacquet et al., 2015), and locally variable gas composition (e.g., Cohen et al., 2004) may result in chondrule diversity. Although fundamental similarities exist between chondrules in enstatite, ordinary, and carbon- aceous chondrites, petrographic and isotopic evidence suggests that each group samples a separate reservoir of chondrules (Jones, 2012; Miller et al., 2017; Chapter 7; Chapter 8). Given that 26Al ages for chondrules within a single chondrite span up to 1 Myr (Kita and Ushikubo, 2012; Chapter 9), and that radial transport was an important process in the protoplanetary disk (e.g., Westphal et al., 2009; Testi et al., 2014; Vacher et al., 2016) that may have acted on timescales shorter than 1 Myr (Cuzzi, Hogan, and Bottke, 2010), maintaining distinct chondrule populations is problematic (note that Goldberg et al., (2016) concluded that it strongly depends on the assumed turbulence parameter α). One possible solution is the early formation of planetesimals that acted as reservoirs for repeated cycles of chondrule formation by partially isolated precursor material (Weidenschilling, Marzari, and Hood, 1998). This model may explain O-isotope variations between chondrite groups, as well as the presence of type I and type II chondrules in all chondrite groups (Ruzicka, 2012), which have been experimentally reproduced from similar precursors under different redox conditions (Cohen et al., 2004; Villeneuve et al., 2015). Trace element data from ordinary (Jacquet et al., 2015) and carbonaceous chondrites (Jacquet et al., 2012) also support similar precursors for type I and type II chondrules, although carbonaceous chondrite precursors appear to be enriched in refractory elements (Jacquet et al., 2012). Future tests of this model will benefit from improved constraints on accretion and transport timescales, as well as more detailed modeling of the chemical and temperature evolution of impact plumes.

2.3 Precursors of Igneous CAIs

Igneous (compact Type A, Type B, and Type C) CAIs are among the oldest solar system solids dated using U-corrected Pb–Pb isotope systematics (Connelly et al., 2012; for the classification scheme CAIs see MacPherson, 2014). Assuming uniform distribution of 26Al in the protopla- 26 27 ‒5 netary disk at the canonical level [( Al/ Al)0 = 5.25 10 (Jacobsen et al., 2008; Larsen et al., 2011)], Al–Mg isotope systematics of igneous CAIs surrounded by WL rims indicate the CAI-melting events started at the very beginning of the Solar System and lasted for up to 0.3 Ma after formation of their precursors (Kita et al., 2013). We infer that igneous CAIs recorded the earliest transient heating events in the protoplanetary disk, and, therefore provide important constraints on the possible mechanisms of these events. Like porphyritic chondrules, igneous CAIs show mineralogical and isotopic evidence for multistage and often incomplete melting of chemically and isotopically diverse solid precursors during transient heating events (e.g., Aléon et al., 2007; Bullock et al., 2012; Ivanova et al., 2012, 2015; Krot et al., 2015). In contrast to porphyritic chondrules, igneous CAIs experienced melt evaporation and irradiation by solar energetic particles (e.g., MacPherson, 2014 and references therein; Sossi et al., 2017), suggesting they were melted in the CAI-forming region, i.e., near the protosun. The coexistence of isotopically similar igneous and non-igneous refractory inclusions in CV chondrites and the coexistence of igneous and nonigneous CAIs in individual AOAs (Krot et al., 2004a) suggest that the CAI melting events were highly localized in nature. 36 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

The precursors of igneous CAIs appear to have consisted exclusively of refractory inclusions from earlier generations 16O-rich forsterite-rich dust. For example, Bullock et al. (2012) concluded that forsterite-bearing Type B (FoB) CAIs formed by melting of aggregates composed of CAIs surrounded by 16O-rich forsterite condensates (Figure 2.17a). Ivanova et al. (2012, 2015) described the presence of ~25 relict CAIs inside a FoB CAI from

Figure 2.17 Combined x-ray elemental maps in Mg (red), Ca (green), and Al (blue) of the complex CV CAIs (a) SJ101 from Allende and (b) 3N from NWA 3118. Region outlined in (b) is shown in BSE in (c). (a) CAI SJ101 consists of the pyroxene-melilite-anorthite-spinel and forsterite-pyroxene lithologies surrounded by a forsterite-free pyroxene-anorthite-spinel mantle. The CAI formed by melting of a Type B-like CAI surrounded by forsterite-rich accretionary rim (after Bullock et al., 2012). (b, c) CAI 3N encloses of about 25 relict igneous CAIs of different types – Type B, Type C, compact Type A, and ultrarefractory (UR). It formed by melting of a dustball composed of multiple CAIs and forsterite-rich dust (after Ivanova et al., 2015). (d) Oxygen–isotope compositions of individual minerals in the relict ultrarefractory CAI 3N-24 and the host forsterite-bearing CAI 3N. The relict CAI is 16O-depleted relative to the host inclusion. Data from Ivanova et al. (2012). px = Al,Ti-diopside; an = anorthite; fo = forsterite; mel = melilite; sp = spinel; Zr,Sc-px = Zr-Sc-rich pyroxene; Zr,Sc,Y-ox = Zr-Sc-Y-rich oxides. (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.) Multiple Mechanisms of Transient Heating Events 37

Figure 2.18 Combined x-ray elemental maps in Mg (red), Ca (green), and Al (blue) overlaid on BSE images of a coarse-grained igneous CAI composed of Type A, B, and C portions from Vigarano (CV). The Type C mantle contains abundant chondrule fragments (indicated by yellow arrows in (c)). From MacPherson et al. (2012). (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.)

NWA 3118 (CV); one of these relict CAIs is an 16O-depleted ultrarefractory inclusion (Figure 2.17b‒d). There is no evidence that chondrules were among the precursors of igneous CAIs surrounded by WL rims (i.e., CAIs that avoided subsequent melting during chondrule formation; see Section 2.1.1). We note that several Type C and Type B CAIs experienced late- stage remelting with the addition of a small amount of 16O-poor ferromagnesian olivine and low-Ca pyroxene that was mineralogically and isotopically similar to chondrules (Itoh and Yurimoto, 2003; Krot et al., 2005b; MacPherson et al., 2012). Some of these ferromagnesian silicates survived as relict grains in the peripheral portions of the host CAIs (Figure 2.18), which make them appear to be chondrule-bearing CAIs (Itoh and Yurimoto, 2003). However, because this melting occurred in an 16O-poor gaseous reservoir and resulted in dissolution of the CAI WL rims, it must be related to chondrule formation (Krot et al., 2005b; MacPherson et al., 2012). Ivanova et al. (2014) discovered that >20% of cm-sized igneous CAIs in CV chondrites are bowl- and disk-shaped (Figure 2.19). This shape is not expected for crystallization of melt droplets in a low gravity field; nor can it be explained by postcrystallization shock deform- ation. Ivanova et al. (2014) suggested that the disk- and bowl-shaped igneous CAIs experienced aerodynamic heating, melting, and deformation by gas drag during their acceleration in the CAI- forming region by shock waves generated by x-ray flares. Liffman et al. (2016) explained this shape as the result of ejection of CAIs from the inner solar accretion disk via the centrifugal interaction between the solar magnetosphere and the inner disk rim.

2.4 Precursors of Nonporphyritic Chondrules in CB and CH Carbonaceous Chondrites

It is suggested that magnesian nonporphyritic chondrules [skeletal olivine (SO) and cryptocrys- talline (CC)], Fe,Ni-metal nodules and irregularly-shaped, often chemically-zoned Fe,Ni-metal grains in CB chondrites formed during a single-stage highly-energetic process 4562.49 0.21 38 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

Figure 2.19 (a) Combined x-ray elemental map in Ti (red), Ca (green), and Al (blue), and (b, c) BSE images of a plastically deformed coarse-grained igneous Compact Type A CAI from the CV chondrite NWA 3118. Regions outlined in (a) are shown in detail in (b) and (c). The hibonite-rich mantle is thicker on the convex side of the CAI than on the concave side, suggesting asymmetric heating of the convex and concave sides. cpx = Al,Ti-diopside; hib = hibonite; mel = melilite; sp = spinel. (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.) Multiple Mechanisms of Transient Heating Events 39

Ma, possibly in an impact generated gas–melt plume (Meibom et al., 2000; Campbell et al., 2001; Campbell et al., 2002; Krot et al., 2005a; Bollard et al., 2015; Fedkin et al., 2015; Oulton et al., 2016). According to this model, the SO chondrules and Fe,Ni-metal nodules represent the fraction of the plume that formed by melting of preexisting solids, while CC chondrules and irregularly-shaped Fe,Ni-metal grains condensed directly from the plume gas as liquids and solids, respectively. The nature of the CB chondrule precursors is poorly known. Based on the equilibrium condensation calculations from a high-temperature gaseous plume (Fedkin et al., 2015) and the bulk chemical compositions of the CB chondrules (Campbell et al., 2002; Oulton et al., 2016), Fedkin et al. (2015) concluded that the range of bulk major elemental compositions of the SO and CC chondrules could not be explained without invoking chemically differentiated colliding bodies (Figure 2.20). Krot et al. (2001a, 2001b) and Oulton et al. (2016) reported bulk REE abundances in chondrules from the Hammadah al Hamra 237 (CBb) and Gujba (CBa) chon- drites. Some of these chondrules show light REE (LREE) enrichment or depletion with a range of La/Sm ratio of (0.9–1.8) CI (Figure 2.21a). Oulton et al. (2016) concluded that these LREE variations are virtually impossible to achieve by processes involving vapor–liquid or vapor– solid exchange of REE in the plume, and must have been inherited from a differentiated target. The most distinctive evidence for inherited chemical differentiation is observed in highly refractory element (Sc, Zr, Nb, Hf, Ta, Th) systematics. The Gujba chondrules exhibit fractio- nations in Nb/Ta that correlate positively with Zr/Hf. These span the range known from lunar and Martian basalts, and exceed the range in Zr/Hf variation known from eucrites (Figure 2.21b). Variations of highly incompatible refractory elements (e.g., Th) against less incompatible elements (e.g., Zr, Sr, Sc) are not chondritic, but exhibit distinctly higher Th abundances requiring a differentiated crust to be admixed with depleted mantle in ratios that are biased to higher crust-to-mantle ratios than in a chondritic body. We note that collision between differentiated bodies cannot explain the 15N-rich bulk compos- itions (e.g., Ivanova et al., 2008) and the presence of CAIs in CB chondrites (Krot et al., 2017b); both seem to require reprocessing of some amount of chondritic material in the plume.

2.5 Recycling of CAIs in the CB Impact-Generated Gas–Melt Plume

Most CAIs and their WL rims in unmetamorphosed chondrites (e.g., CM2, CR2‒3, CO3.0) are uniformly 16O-rich (Δ17O~‒24‰) (Makide et al., 2009; Krot et al., 2017b). The mineralogy, petrology, and 16O-rich compositions of the WL rims suggest formation by evaporation/conden- sation, possible melting, and thermal annealing in the formation region of the host CAIs (e.g., Toppani et al., 2006; Matzel et al., 2013; Bodénan et al., 2014; Han et al., 2015; Krot et al., 2017b). Krot et al. (2017b) defined such rims as the primordial WL rims. As is discussed in Section 2.1.1, formation of porphyritic chondrules in most chondrite groups affected a relatively small fraction of CAIs, and resulted in melting, often incomplete, of the CAIs and their WL rims, as well as O-isotope exchange of the melted minerals with an 16O-poor nebular gas. In contrast to typical CAIs in carbonaceous chondrites, most CB CAIs (and about 10 percent of CH CAIs) are uniformly 16O-depleted igneous inclusions; Δ17O values between individual CAIs vary from ~ 15‰ to ~ 5‰ (Figures 2.22 and 2.23a). These CAIs have diverse Figure 2.20 (a) Curves showing the evolution of the calculated bulk composition (wt%) of the silicate fraction of the condensate in a plume produced by impact between a CR metal body and an H chondrite body at the indicated values of Ptot and Ni/H and Si/H enrichments relative to solar composition, compared to bulk analyses of BO and CC chondrules from CBb chondrites. Compositions in (a) and (b) are normalized to 100% (CaO + Al2O3) + MgO + SiO. Numbers on the curves are temperatures in K. The Multiple Mechanisms of Transient Heating Events 41

Figure 2.21 (a) CI-normalized REE abundances in Gujba chondrules. (b) Nb/Ta versus Zr/Hf in Gujba chondrules, lunar, Martian, and eucritic basalts. From Oulten et al. (2016).

Caption for Figure 2.20 (cont.) model curves miss most of bulk compositional data for CB chondrules. (b) Curves showing the evolution of the calculated bulk composition (wt%) of the silicate fraction of the condensate in plumes produced by impact vaporization of a mixture of CR core, CR residual mantle, CR crust, residual nebular gas and water ice, showing the effect of removal of the high-temperature condensate ‒3 assemblage that forms at 1,750 K. All cases shown are for Ptot =10 bar, Si/H and Ni/H enrichments of 300 and 3 104, respectively, relative to solar composition, and CR mantle/total silicate and water/total silicate weight ratios of 0.4 and 0.176, respectively. Curves are shown for removal of the indicated fractions of the early condensate. Note that an early condensate assemblage with almost the same composition as the BO chondrule with the highest CaO + Al2O3 content is predicted to form at 1,750 K, and that the compositions of almost all CC chondrule compositions are best fit at ~1,710–1,720 K after removal of 95–100 percent of that early condensate. From Fedkin et al. (2015). 42 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

Figure 2.22 (a, c) Combined x-ray elemental maps in Mg (red), Ca (green), and Al (blue), (b) x-ray elemental map in Ti, and (d‒f ) BSE images of the uniformly 16O-depleted igneous CAIs surrounded by igneous rims composed of Al-diopside and forsterite. Regions outlined in (c) and (e) are shown in detail in (d) and (f ), respectively. Numbers correspond to CAI numbers listed in Figure 2.21. cpx = high-Ca pyroxene; fo = forsterite; grs = grossite; hib = hibonite; mel = melilite; sp = spinel. (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.) Multiple Mechanisms of Transient Heating Events 43

Figure 2.23 (a) Oxygen–isotope compositions of magnesian nonporphyritic (SO and CC) chondrules and 16O-depleted (relative to typical CAIs carbonaceous chondrites having Δ17Oof~‒24‰) igneous CAIs surrounded by igneous rims from the CH and CB carbonaceous chondrites. Most chondrules have similar Δ17Oof~‒2.5‰; the CAIs are 16O-enriched relative to the chondrules. Individual CAIs and their igneous rims have similar O-isotope compositions; there are, however, inter-CAI compositional differences. It is suggested that magnesian CC and SO chondrules in CB and CH chondrites formed in a gas–melt plume generated by a collision between planetesimals, while the rimmed CAIs resulted from complete melting of preexisting CAIs followed by gas–melt interaction in the plume. The terrestrial fractionation line (TF) is shown for reference. Numbers correspond to numbers of individual inclusions; some of them are shown in Figure 2.20. (b) Oxygen–isotope compositions of magnesian porphyritic chondrules 1573–4-9 and 1–3-3 with coarse relict 16O-depleted spinels, and uniformly 16O-depleted spinel-rich igneous CAIs 1–3-4 and 1573–1-1 surrounded by igneous rims. These chondrules and CAIs are shown in Figure 2.22. Oxygen- isotope compositions of the relict spinel grains are similar to those of the uniformly 16O-depleted igneous CAIs in CB and CH chondrites. cpx = high-Ca pyroxene; fo = forsterite; grs = grossite; hib = hibonite; mel = melilite; ol = ferromagnesian olivine; pl = plagioclase; pv = perovskite; px = low-Ca pyroxene. Data from Krot et al. (2010, 2017a, 2017b); Krot, Nagashima, and Petaev (2012). (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.)

mineralogies (grossite-rich, hibonite-rich, melilite-rich, spinel-rich, and Al,Ti-diopside forsterite-rich), but are surrounded by mineralogically similar igneous rims composed of Al-diopside and Ca-rich (0.5‒1.4 wt% CaO) forsterite. The igneous rims and host CAIs have identical (within uncertainties of SIMS measurements) O-isotope compositions, suggesting that they crystallized from isotopically similar but chemically distinct melts. Krot et al. (2017b) suggested that the uniformly 16O-depleted igneous rims around uniformly 16O-depleted igneous CB and CH CAIs resulted from melting of preexisting refractory inclusions in an impact- generated gas–melt plume invoked as the origin of the CB and CH magnesian nonporphyritic chondrules. This melting was accompanied by gas–melt interaction and O-isotope exchange with an 16O-poor plume gas (Δ17O~2‰). Crystallization of the Si- and Mg-enriched CAI 44 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

Figure 2.24 BSE images of (a, b) relict 16O-depleted spinels surrounded by chondrule-like material composed of ferromagnesian olivine, low-Ca and high-Ca pyroxenes, mesostasis, and Fe,Ni-metal, and (c, d) uniformly 16O-depleted igneous CAIs surrounded by igneous Al-diopside-forsterite rims from the CH/CB chondrite Isheyevo and CH chondrite Acfer 214. Oxygen–isotope compositions of these objects are shown in Figure 2.21b. cpx = high-Ca pyroxene; mel = melilite; mes = plagioclase-normative mesostasis; ol = olivine; px = low-Ca pyroxene; sp = spinel.

peripheries resulted in formation of igneous Al-diopside-forsterite rims. As a result, these CAIs can be considered chondrules formed by complete melting of precursors composed largely of refractory inclusions. Several CH porphyritic chondrules contain 16O-depleted coarse relict spinel grains which are mineralogically and isotopically similar to the uniformly 16O-depleted CB-like igneous CAIs (Figures 2.23a and 2.23b, 2.24). If these relict spinels are indeed remnants of the CB CAIs melted in the impact plume, then these porphyritic chondrules must have postdated the plume event. Based on the lack of interchondrule matrix in CB chondrites, Krot et al. (2005a) suggested that the CB plume event occurred after nearly complete dissipation of dust in the protoplanetary disk. If this is the case, chondrules postdating the plume event may have formed by planetesimal collision(s). Multiple Mechanisms of Transient Heating Events 45 2.6 Discussion and Conclusions

1. Porphyritic chondrules are the dominant textural type of chondrules in most chondrite groups. The mineralogy, petrography, and isotope compositions of porphyritic chondrules suggest their formation by melting, often incomplete, of isotopically diverse solid precursors. The chondrule precursors included CAIs, AOAs, and chondrules from earlier generations, fine-grained matrix-like material, and, possibly, fragments of preexisting thermally pro- cessed planetesimals. The melting occurred during highly-localized transient heating events in different dust-rich regions of the protoplanetary disk characterized by 16O-poor average compositions of dust (Δ17O~‒5‰ to +3‰). Formation of chondrules in isotopically distinct regions of the protoplanetary disk is strongly supported by recent measurements of bulk Mg-, Ti- and Cr-isotope compositions of chondrules in ordinary, enstatite, and carbonaceous chondrites (e.g., Van Kooten et al., 2016; Bizzarro et al., 2017; Gerber et al., 2017). All together, these observations preclude formation of porphyritic chondrules by splashing of differentiated asteroids. Instead, they are consistent with melting of dustballs during loca- lized transient heating events, such as bow shocks, magnetized turbulence in the disk, and, possibly, by collisions between chondritic planetesimals (Chapters 13, 15, and 16). 2. Reprocessing of refractory inclusions during formation of porphyritic chondrules resulted in their melting to various degrees, destruction of WL rims, gas–melt interaction, replacement of melilite by Na-bearing plagioclase, oxygen–isotope exchange, and transformation into Type C-like igneous CAIs and anorthite-rich chondrules. AOAs extensively melted during chondrule formation could have been transformed in magnesian porphyritic chondrules. 3. Like porphyritic chondrules, igneous CAIs formed by melting, often incomplete, of iso- topically diverse solid precursors during localized transient heating events. In contrast to porphyritic chondrules, the precursors of igneous CAIs consisted exclusively of refractory inclusions from earlier generations 16O-rich forsterite-rich dust, and the melting occurred predominantly in an 16O-rich gas (Δ17O~‒24‰) of approximately solar composition, most likely near the protosun. The U-corrected Pb–Pb absolute and Al–Mg relative chronologies of igneous CAIs indicate that the CAI melting events started at the very beginning of the protoplanetary disk evolution and appear to have lasted up to 0.3 Ma after condensation of CAI precursors, providing clear evidence for very early transient heating events capable of melting refractory dustballs in the innermost part of the disk. There is no evidence that chondrules were among these precursors, consistent with an age gap between CAIs and chondrules (Chapter 9). 4. In contrast to typical (non–metal-rich) chondrites, the CB carbonaceous chondrites contain exclusively magnesian nonporphyritic (SO and CC) chondrules formed in an impact- generated gas–melt plume about 5 Ma after CV CAIs. Bulk chemical compositions of CB chondrules and equilibrium thermodynamic calculations suggest that the collision involved differentiated bodies. However, the 15N-rich bulk compositions of CB chondrites and the presence of CAIs in CB chondrites suggest that some amount of chondritic material must have been reprocessed in the plume as well. The uniformly 16O-depleted igneous CB CAIs most likely formed by complete melting of preexisting CAIs that was accompanied by gas- melt interaction and O-isotope exchange in the plume. 46 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

5. CH chondrites represent a mixture of the CB-like materials (magnesian SO and CC chon- drules and uniformly 16O-depleted igneous CAIs) formed in an impact plume and the typical chondritic materials (magnesian, ferroan, and Al-rich porphyritic chondrules, uniformly 16O-rich CAIs, and chondritic lithic clasts) that may have largely predated the impact plume event (see also Wasson and Kallemeyn, 1990). Some relict CAIs inside CH por- phyritic chondrules are texturally (igneous, spinel-rich) and isotopically (uniformly 16O-depleted) similar to the CB-like CAIs, which are interpreted as a result of complete melting and O-isotope exchange in the CB impact plume. If this is the case, these chondrules must have formed after the impact plume event, possibly by asteroidal collisions. We infer that CH chondrites contain multiple generations of chondrules formed by different mechanisms. 6. We conclude that there are multiple mechanisms of chondrule formation that operated during the entire lifetime of the protoplanetary disk.

Acknowledgments

The authors thank Drs. Edward R. D. Scott, Martin Bizzarro, and Harold C. Connolly Jr. for helpful discussions. We thank A. E. Rubin and an anonymous reviewer for constructive reviews. Handling of the manuscript by S. S. Russell is highly appreciated. This work is supported by the National Aeronautics and Space Administration (NASA) Emerging Worlds program, grants NNX15AH38G and NNX17AE22G (ANK, PI).

References

Abreu, N., and Brearley, A. J. (2010). Early solar system processes recorded in the matrices of two highly pristine CR3 carbonaceous chondrites, MET 00426 and QUE 99177. Geochim. Cosmochim. Acta, 74, 11461171. Aléon, J., El Goresy, A., and Zinner, A. (2007). Oxygen isotope heterogeneities in the earliest protosolar gas recorded in a meteoritic calcium–aluminum-rich inclusion. Earth Planet. Sci. Lett., 263, 114‒127. Alexander, C. M. O’D., Grossman, J. N., Ebel, D. S., and Ciesla, F. J. (2008). The formation conditions of chondrules and chondrites. Science, 320, 16171619. Alexander, C. M. O’D., and Ebel, D. S. (2012). Questions, questions: Can the contradictions between the petrologic, isotopic, thermodynamic, and astrophysical constraints on chon- drule formation be resolved? Meteorit. Planet. Sci., 47, 1157‒1175. Asphaug, E., Jutzi, M., and Movshovitz, M. (2011). Chondrule formation during planetesimal accretion. Earth Planet. Sci. Lett., 308, 369–379. Becker, M., Hezel, D. C., Schulz, T., Elfers, B. -M., and Münker, C. (2015). The age of CV chondrites from component specificHf–W systematics. Earth Planet. Sci. Lett., 432, 472‒482. Bigolski, J. N., Weisberg, M. K., Connolly, H. C., and Ebel, D. S. (2016). Microchondrules in three unequilibrated ordinary chondrites. Meteorit. Planet. Sci., 51, 235‒260. Bischoff, A., and Keil, K. (1983a). Ca-Al-rich chondrules and inclusions in ordinary chondrites. Nature, 303, 588592. Multiple Mechanisms of Transient Heating Events 47

Bischoff, A., and Keil, K. (1983b). Catalog of Al-Rich Chondrules, Inclusions, and Fragments in Ordinary Chondrites. Albuquerque, NM: Department of Geology and Institute of Meteoritics, University of New Mexico. Bischoff, A., and Keil, K. (1984). Al-rich objects in ordinary chondrites: Related origin of carbonaceous and ordinary chondrites and their constituents. Geochim. Cosmochim. Acta, 48, 693709. Bischoff, A., Palme, H., and Spettel, B. (1989). Al-rich chondrules from the Ybbsitz H4- chondrite: Evidence for formation by collision and splashing. Earth Planet. Sci. Lett., 93, 170‒180. Bischoff, A., Palme, H., Schultz, L., et al. (1993). Acfer 182 and paired samples, an iron-rich carbonaceous chondrite: Similarities with ALH85085 and relationship to CR chondrites. Geochim. Cosmochim. Acta, 57, 2631‒2648. Bischoff, A., Wurm, G., Chaussidon, M., et al. (2017). The Allende multicompound chondrule (ACC) – Chondrule formation in a local super-dense region of the early solar system. Meteorit. Planet. Sci., 52, 906‒924. Bizzarro, M., Wiellandt, D., Haugbølle, T., Nordlund, Å., and Connelly, J. N. (2017). Nucleo- synthetic diversity of chondrules – tracking disk mass transport and the formation of large- scale Solar System reservoirs. 27th Goldschmidt Conf., abstract. Bland, P. A., Alard, O., Benedix, G. K., et al. (2005). Volatile fractionation in the early solar system and chondrule/matrix complementarity. Proc. Nat. Acad. Sci., 102, 13755–13760. Bodénan, J.- D., Starkey, N. A., Russell, S. S., Wright, I. P., and Franchi, I. A. (2014). An oxygen isotope study of Wark-Lovering rims on Type A CAIs in primitive carbonaceous chondrites. Earth Planet. Sci. Lett., 401, 327–336. Bollard, J., Connelly, J., and Bizzarro, M. (2015). Pb-Pb dating of individual chondrules from the CBa chondrite Gujba: Assessment of the impact plume formation model. Meteorit. Planet. Sci., 50, 1197‒1216. Brearley, A. J. (1996). Nature of matrix in unequilibrated chondrites and its possible relation- ship to chondrules. In R. Hewins, R. Jones, and E. Scott (Eds.), Chondrules and the Protoplanetary Disk, 137–151. Cambridge, UK: Cambridge University Press. Brearley, A. J., and Krot, A. N. (2012). Metasomatism in the early solar system: The record from chondritic meteorites. In D. Harlov and H. Austrheim (Eds.), Metasomatism and the Chemical Transformation of Rock – Lecture Notes in Earth System Sciences, 659789. New York, NY: Springer. Brownlee, D., Tsou, P., Aléon, J., et al. (2006). Comet 81P/Wild 2 under a microscope. Science, 314, 17111716. Budde, G., Kleine, T., Kruijer, T. S., Burkhardt, C., and Metzler, K. (2016a). Tungsten isotopic constraints on the age and origin of chondrules. Proc. Nat. Acad. Sci., 113, 2886‒2891. Budde, G., Kleine T., Kruijer T. S., et al. (2016b). Isotopic complementarity of chondrules and matrix. 26th Goldschmidt Conference, abstract #2316. Bullock, E. S., MacPherson, G. J., Nagashima, K., et al. (2012). Forsterite-bearing type B refractory inclusions from CV3 chondrites: From aggregates to volatilized melt droplets. Meteorit. Planet. Sci., 47, 21282147. Bunch, T. E., Keil, K., and Snetsinger, K. G. (1967). Chromite composition in relation to chemistry and texture of ordinary chondrites. Geochim. Cosmochim. Acta, 31, 15691582. Campbell, A. J., Humayun, M., Meibom, A., Krot, A. N., and Keil, K. (2001). Origin of zoned metal in the QUE 94411 chondrite. Geochim. Cosmochim. Acta, 65, 163180. Campbell, A. J., Humayun, M., and Weisberg, M. K. (2002). Siderophile element constraints on the formation of metal in the metal-rich chondrites Bencubbin, Weatherford and Gujba. Geochim. Cosmochim. Acta, 66, 647660. 48 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

Cohen, B. A., Hewins, R. H., and Alexander, C. M. O’D. (2004). The formation of chondrules by open-system melting of nebular condensates. Geochim. Cosmochim. Acta, 68,1661‒1675. Connelly, J. N., Bizzarro, M., Krot, A. N., et al. (2012). The absolute chronology and thermal processing of solids in the solar protoplanetary disk. Science, 338, 651‒655. Connolly, H. C. Jr., and Jones, R. H. (2016). Chondrules: The canonical and noncanonical views. J. Geophys. Res.: Planets, 121, 1885‒1899. Ciesla, F. J. (2010). The distributions and ages of refractory objects in the solar nebula. Icarus, 208, 455467. Clayton, R. N., Mayeda, T. K., Olsen, E. J., and Goswami, J. N. (1991). Oxygen isotope studies of ordinary chondrites. Geochim. Cosmochim. Acta, 55, 23172337. Cuzzi, J. N., and Alexander, C. M. O’D. (2006). Chondrule formation in particle-rich nebular regions at least hundreds of kilometers across. Nature, 441, 483–485. Cuzzi, J. N., Hogan, R. C., and Bottke, W. F. (2010). Towards initial mass functions for asteroids and Kuiper Belt Objects. Icarus, 208, 518‒538. Desch, S. J., Ciesla, F. J., Hood, L. L., and Nakamoto, T. (2005) Heating of chondritic materials in Solar nebula shocks. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, 849‒872. ASP Conference Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Dobrica, E., and Brearley, A. J. (2016). Microchondrules in two unequilibrated ordinary chondrites: Evidence for formation by splattering from chondrules during stochastic collisions in the solar nebula. Meteorit. Planet. Sci., 51, 884905. Ebel, D. S., Brunner, C., Konrad, K., et al. (2016). Abundance, major element composition and size of components and matrix in CV, CO and Acfer 094 chondrites. Geochim. Cosmo- chim. Acta, 172, 322–356. Ebert, S., and Bischoff, A. (2016). Genetic relationship between Na-rich chondrules and Ca, Al- rich inclusions? – Formation of Na-rich chondrules by melting of refractory and volatile precursors in the solar nebula. Geochim. Cosmochim. Acta, 177, 182‒204. Fedkin, A. V., Grossman, L., Humayun, M., Simon, S. B., and Campbell, A. J. (2015). Condensates from vapor made by impacts between metal-, silicate-rich bodies: Comparison with metal and chondrules in CB chondrites. Geochim. Cosmochim. Acta, 164,236–261. Friend, P., Hezel, D. C., and Mucerschi, D. (2016). The conditions of chondrule formation, Part II: Open system. Geochim. Cosmochim. Acta, 173, 198‒209. Georges, P., Libourel, G., and Deloule, E. (2000). Experimental constraints on alkali condensa- tion in chondrule formation. Meteorit. Planet. Sci., 35, 1183‒1188. Gerber, S., Burkhardt, C., Budde, G., Metzler, K., and Kleine, T. (2017). Mixing and transport of dust in the early solar nebula as inferred from titanium isotope variations among chondrules. Astrophys. J., 841, L17. Goldberg, A. Z., Jacquet, E., and Owen, J. E. (2016). Chondrule transport in the early solar system. 79th Ann. Meteorit. Soc. Meet., abstract #6012. Gounelle, M., Krot, A. N., Nagashima, K., and Kearsley, A. (2009). Extreme 16O-enrichment in refractory inclusions from the Isheyevo meteorite: Implication for oxygen isotope com- position of the Sun. Astrophys. J., 698, L18L22. Greshake, A. (1997). The primitive matrix components of the unique carbonaceous chondrite Acfer 094: A TEM study. Geochim. Cosmochim. Acta, 61, 437–452. Grossman, J. N., and Brearley, A. J. (2005). The onset of metamorphism in ordinary and carbonaceous chondrites. Meteorit. Planet. Sci., 40,87‒122. Grossman, J. N., Rubin, A. E., Nagahara, H., and King, E. A. (1988). Properties of chondrules. In J. Kerridge and M. Matthews (Eds.), Meteorites and the Early Solar System, 619‒660. Tucson, AZ: University of Arizona Press. Grossman, L., Fedkin, A. V., and Simon, S. B. (2012). Formation of the first oxidized iron in the solar system. Meteorit. Planet. Sci., 75, 2160‒2169. Multiple Mechanisms of Transient Heating Events 49

Han, J., Keller, L. P., Needham, A. W., Messenger, S., and Simon, J. I. (2015). Microstructural investigation of a Wark-Lovering rim on a Vigarano CAI. 78th Ann. Meteorit. Soc. Meet., abstract #5243. Herbst, W., and Greenwood, J. P. (2016). A new mechanism for chondrule formation: Radiative heating by hot planetesimals. Icarus, 267, 364‒367. Hezel, D. C., and Palme, H. (2007). The conditions of chondrule formation, Part I: Closed system. Geochim. Cosmochim. Acta, 71, 4092‒4107. Hezel, D. C., and Palme, H. (2008). Constraints for chondrule formation from Ca-Al distribu- tion in carbonaceous chondrites. Earth Planet. Sci. Lett., 265, 716‒725. Hezel, D. C., and Palme, H. (2010). The chemical relationship between chondrules and matrix and the chondrule-matrix complementarity. Earth Planet. Sci. Lett., 294,85‒93. Hezel, D. C., Russell, S. S., Ross, A. J., and Kearsley, A. T. (2008). Modal abundances of CAIs: Implications for bulk chondrite element abundances and fractionations. Meteorit. Planet. Sci., 43, 18791894. Hewins, R. H. (1997). Chondrules. Ann. Rev. Earth Planet. Sci., 25,61‒83. Hewins, R. H., and Zanda, B. (2012). Chondrules: Precursors and interactions with the nebular gas. Meteorit. Planet. Sci., 47, 11201138. Hewins, R. H., Zanda, B., and Bendersky, C. (2012). Evaporation and recondensation of sodium in Semarkona Type II chondrules. Geochim. Cosmochim. Acta, 78,1‒17. Imae, N., and Isobe, H. (2017). An experimental study of chondrule formation from chondritic precursors via evaporation and condensation in Knudsen cell: Shock heating model of dust aggregates. Earth Planet. Sci. Lett., 473, 256‒268. Itoh, S., and Yurimoto, H. (2003). Contemporaneous formation of chondrules and refractory inclusions in the early solar system. Nature, 423, 728731. Ivanova, M. A., Kononkova, N. N., Greenwood, R. C., et al. (2008). The Isheyevo meteorite: Mineralogy, petrography, bulk chemistry, oxygen, nitrogen, carbon isotopic compositions and Ar-Ar ages. Meteorit. Planet. Sci., 43, 915‒941. Ivanova, M. A., Nagashima, K., Krot, A. N., and MacPherson, G. J. (2012). Calcium-alumi- num-rich inclusions with relict ultra-refractory inclusions rich in Zr, Y and Sc from Efremovka and North West Africa 3118 CV3 carbonaceous chondrites: Evidence for multistage formation in oxygen isotopic reservoirs of variable composition. Meteorit. Planet. Sci. , 47, 21072127. Ivanova, M. A., Lorenz, C. A., Shuvalov, V. V., et al. (2014). Plastically-deformed igneous calcium-aluminum-rich inclusions from CV carbonaceous chondrites: Clues to a nature of CAI melting events. LPSC #45, abstract #2166. Ivanova, M. A., Lorenz, C. A., Krot, A. N., and MacPherson, G. J. (2015). A com- pound Ca-, Al-rich inclusion from CV3 chondrite North West Africa 3118: Implications for understanding processes during CAI formation. Meteorit. Planet. Sci., 50, 1512‒1528. Jacobsen, B., Yin, Q.-Z., Moynier, F., et al. (2008). 26Al-26Mg and 207Pb-206Pb systematics of Allende CAIs: Canonical solar initial 26Al/27Al ratio reinstated. Earth Planet. Sci. Lett., 272, 353‒364. Jacquet, E. (2014). Transport of solids in protoplanetary disks: Comparing meteorites and astrophysical models. Comptes rendus – Geoscience, 346,312. Jacquet, E., Froman, S., and Gounelle, M. (2011). Radial transport of refractory inclusions and their preservation in the dead zone. Astron. Astrophys., 526, L8. Jacquet, E., Alard, O., and Gounelle, M. (2012). Chondrule trace element geochemistry at the mineral scale. Meteorit. Planet. Sci., 47, 1695‒1714. Jacquet, E., Alard, O., and Gounelle, M. (2015). Trace element geochemistry of ordinary chondrite chondrules: The type I/type II chondrule dichotomy. Geochim. Cosmochim. Acta, 155,47‒67. 50 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

Johnson, B. C., Minton, D. A., Melosh, H. J., and Zuber, M. T. (2014). Impact jetting as the origin of chondrules. Nature, 517, 339‒341. Jones, R. H. (1996). Relict grains in chondrules: Evidence for chondrule recycling. In R. Hewins, R. Jones, and E. Scott (Eds.), Chondrules and the Protoplanetary Disk, 163‒172. Cambridge, UK: Cambridge University Press. Jones, R. H. (2012). Petrographic constraints on the diversity of chondrule reservoirs in the protoplanetary disk. Meteorit. Planet. Sci., 47, 1176‒1190. Jones, R. H., and Danielson, L. R. (1997). A chondrule origin of dusty relict olivine in unequilibrated chondrites. Meteorit. Planet. Sci., 32, 753‒760. Jones, R. H., Leshin, L. A., Guan, Y., et al. (2004). Oxygen isotope heterogeneity in chondrules from the Mokoia CV3 carbonaceous chondrite. Geochim. Cosmochim. Acta, 68, 34232106. Jones, R. H., Grossman, J. N., and Rubin, A. E. (2005). Chemical, mineralogical and isotopic properties of chondrules: Clues to their origin. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, 251‒286. ASP Conf. Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Keil, K., Zucolotto, M. E., Krot, A. N., et al. (2015). The Vicência meteorite fall: A new unshocked (S1) weakly metamorphosed (3.2) LL chondrite. Meteorit. Planet. Sci., 50, 1089‒1111. Kimura, M., El Goresy, A., Palme, H., and Zinner, E. (1993). Ca-Al-rich inclusions in the unique chondrite ALH 85085 Petrology, chemistry and isotopic compositions. Geochim. Cosmochim. Acta, 57, 23292359. Kita, N. T., Nagahara, H., Tachibana, S., et al. (2010). High precision SIMS oxygen three isotope study of chondrules in LL3 chondrites: Role of ambient gas during chondrule formation. Geochim. Cosmochim. Acta, 74, 6610‒6635. Kita, N. T., and Ushikubo, T. (2012). Evolution of protoplanetary disk inferred from 26Al chronology of individual chondrules. Meteorit. Planet. Sci., 47, 1108‒1119. Kita, N. T., Yin, Q.- Z., MacPherson, G. J., et al. (2013). 26Al-26Mg isotope systematics of the first solids in the early solar system. Meteorit. Planet. Sci., 48, 13831400. Krot, A. N., and Rubin, A. E. (1993). Chromite-rich mafic silicate chondrules in ordinary chondrites: Formation by impact melting. Lunar Planet. Sci., 24, 827‒829. Krot, A. N., and Rubin, A. E. (1994). Glass-rich chondrules in ordinary chondrites. Meteorit. Planet. Sci., 29, 697‒707. Krot, A. N., and Wasson, J. T. (1995). Igneous rims on FeO-rich and FeO-poor chondrules in ordinary chondrites. Geochim. Cosmochim. Acta, 59, 4951‒4966. Krot, A. N., and Rubin, A. E. (1996). Microchondrule-bearing chondrule rims: Constraints on chondrule formation. In R. Hewins, R. Jones, and E. Scott (Eds.), Chondrules and the Protoplanetary Disk, 181‒184. Cambridge, UK: Cambridge University Press. Krot, A. N., Ivanova, M. A., and Wasson, J. T. (1993). The origin of chromitic chondrules and the volatility of Cr under a range of nebular conditions. Earth Planet. Sci. Letters, 119, 569‒584. Krot, A. N., McKeegan, K. D., Russell, S. S., et al. (2001a). Refractory Ca, Al-rich inclusions and Al-diopside-rich chondrules in the metal-rich chondrites Hammadah al Hamra 237 and QUE 94411. Meteorit. Planet. Sci., 36, 1189‒1217. Krot, A. N., Meibom, A., Russell, S. S., et al. (2001b). A new astrophysical setting for chondrule formation. Science, 291, 1776‒1779. Krot, A. N., Meibom, A., Weisberg, M. K., and Keil, K. (2002). The CR chondrite clan: Implications for early solar system processes. Meteorit. Planet. Sci., 37, 14511490. Krot, A. N., Petaev, M. I., Russell, S. S., et al. (2004a). Amoeboid olivine aggregates in carbonaceous chondrites: Records of nebular and asteroidal processes. Chem. Erde, 64, 185‒239. Multiple Mechanisms of Transient Heating Events 51

Krot, A. N., Libourel, G., Goodrich, C. A., and Petaev, M. I. (2004b). Silica-rich igneous rims around magnesian chondrules in CR carbonaceous chondrites: Evidence for fractional condensation during chondrule formation. Meteorit. Planet. Sci., 39, 1931‒1955. Krot, A. N., Amelin, Y., Cassen, P., and Meibom, A. (2005a). Young chondrules in CB chondrites from a giant impact in the early Solar System. Nature, 436, 989‒992. Krot, A. N., Yurimoto, H., Hutcheon, I. D., and MacPherson, G. J. (2005b). Relative chron- ology of CAI and chondrule formation: Evidence from chondrule-bearing igneous CAIs. Nature, 434, 998‒1001. Krot, A. N., Fagan, T. J., Yurimoto, H., and Petaev, M. I. (2005c). Origin of low-Ca pyroxene in amoeboid olivine aggregates: Evidence from oxygen isotopic compositions. Geochim. Cosmochim. Acta, 69, 1873‒1881. Krot, A. N., Libourel, G., and Chaussidon, M. (2006a). Oxygen isotope compositions of chondrules in CR chondrites. Geochim. Cosmochim. Acta, 70, 767‒779. Krot, A. N., McKeegan, K. D., Huss, G. R., et al. (2006b). Aluminum-magnesium and oxygen isotope study of relict Ca-Al-rich inclusions in chondrules. Astrophys. J., 629, 1227‒1237. Krot, A. N., Nagashima, K., Bizzarro, M., et al. (2008). Multiple generations of refractory inclusions in the metal-rich carbonaceous chondrites Acfer 182/214 and Isheyevo. Astrophys. J., 672, 713‒721. Krot, A. N., Nagashima, K., Yoshitake, M., and Yurimoto, H. (2010). Oxygen isotope compos- itions of chondrules from the metal-rich chondrites Isheyevo (CH/CBb), MAC 02675 (CBb) and QUE 94627 (CBb). Geochim. Cosmochim. Acta, 74, 21902211. Krot, A. N., Nagashima, K., and Petaev, M. I. (2012). Isotopically uniform, 16O-depleted calcium, aluminum-rich inclusions in CH and CB carbonaceous chondrites. Geochim. Cosmochim. Acta, 83, 159‒178. Krot, A. N., Keil, K., Goodrich, C., Weisberg, M. K., and Scott, E. R. D. (2014). Classification of meteorites. In A. M. Davis (Ed.), Meteorites and Cosmochemical Processes.InH.D. Holland and K. K. Turekian (Eds.), Treatise on Geochemistry (Second Edition), 1, 163. Oxford, UK: Elsevier. Krot, A. N., Nagashima, K., Ma, C., and Wasserburg, G. J. (2015). Forsterite-bearing Type B CAI with a relict eringaite-bearing ultra-refractory CAI. 78th Ann. Meteorit. Soc. Meet., abstract #5308. Krot, A. N., Nagashima, K., van Kooten, E. M. M., and Bizzarro, M. (2017a). Calcium- aluminum-rich inclusions recycled during formation of porphyritic chondrules from CH carbonaceous chondrites. Geochim. Cosmochim. Acta, 201, 185‒223. Krot, A. N., Nagashima, K., van Kooten, E. M. M., and Bizzarro, M. (2017b). High-temperature rims around calcium-aluminum-rich inclusions from the CR, CB and CH carbonaceous chondrites. Geochim. Cosmochim. Acta, 201, 155‒184. Kunihiro, T., Rubin, A. E., McKeegan, K. D., and Wasson, J. T. (2004). Oxygen-isotopic compositions of relict and host grains in chondrules in the Yamato 81020 CO3.0 chondrite. Geochim. Cosmochim. Acta, 68, 3599–3606. Kunihiro, T., Rubin, A. E., and Wasson, J. T. (2005). Oxygen-isotopic compositions of low- FeO relicts in high-FeO host chondrules in Acfer 094, a type 3.0 carbonaceous chondrite closely related to CM. Geochim. Cosmochim. Acta, 69, 3831–3840. Larsen, K., Trinquier, A., Paton, C., et al. (2011). Evidence for magnesium-isotope heterogen- eity in the solar protoplanetary disk. Astrophys. J., 735, L37L40. Leitch, C., and Smith, J. (1982). Petrography, mineral chemistry and origin of Type I enstatite chondrites. Geochim. Cosmochim. Acta, 46, 2083‒2097. Leroux, H., Libourel, G., Lemelle, L., and Guyot, F. (2003). Experimental study and TEM characterization of dusty olivines in chondrites: Evidence for formation by in situ reduc- tion. Meteorit. Planet. Sci., 38,81‒94. 52 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

Leshin, L. A., Rubin, A. E., and McKeegan, K. D. (1997). The oxygen isotopic composition of olivine and pyroxene from CI chondrites. Geochim. Cosmochim. Acta, 61, 835–845. Libourel, G., and Krot, A. N. (2006). Evidence for the presence of planetesimal material among the precursors of magnesian chondrules of nebular origin. Earth Planet. Sci. Lett., 254,1‒8. Libourel, G., Krot, A. N., and Tissandier, L. (2006). Role of gas-melt interaction during chondrule formation. Earth Planet. Sci. Lett., 251, 232‒240. Libourel, G., Michel, P., Delbo, M., et al. (2017). Primitive matter in the solar system. Icarus, 282, 375–379. Liffman, K., Cuello, N., and Paterson, D. A. (2016). A unified framework for producing CAI melting, Wark-Lovering rims and bowl-shaped CAIs. Monthly Notices of the Royal Astronomical Society, 462, 1137‒1163. MacPherson, G. J. (2014). Calcium-aluminum-rich inclusions in chondritic meteorites. In A. M. Davis (Ed.), Meteorites, Comets and Planets. In H. D. Holland and K. K. Turekian (Eds.), Treatise on Geochemistry (Second Edition), 1, 139‒179. Oxford, UK: Elsevier. MacPherson, G. J., and Huss, G. R. (2005). Petrogenesis of Al-rich chondrules: Evidence from bulk compositions and phase equilibria. Geochim. Cosmochim. Acta, 69, 3099–3127. MacPherson, G. J., Kita, N. T., Ushikubo, T., Bullock, E. S., and Davis, A. M. (2012). Well- resolved variations in the formation ages for Ca-Al-rich inclusions in the early Solar System. Earth Planet. Sci. Lett., 331,43‒54. Makide, K., Nagashima, K., Krot, A. N., et al. (2009). Oxygen and magnesium-isotope compositions of calciumaluminum-rich inclusions from CR2 carbonaceous chondrites. Geochim. Cosmochim. Acta, 73, 50185051. Maruyama, S., and Yurimoto, H. (2003). Relationship among O, Mg isotopes and the petrog- raphy of two spinel-bearing compound chondrules. Geochim. Cosmochim. Acta, 67, 3943–3957. Mathieu, R. (2009). Solubilite du sodium dans les silicates fondues. PhD thesis, INPL-CRPG Nancy. Mathieu, R., Libourel, G., Deloule, E., et al. (2011). Na2O-solubility in molten CaO–MgO– SiO2 system. Geochim. Cosmochim. Acta, 75, 608–628. Matzel, J. E. P., Simon, J. I., Hutcheon, I. D., et al. (2013). Oxygen isotope measurements of a rare Murchison Type A CAI and its rim. LPSC #44, abstract #2632. McKeegan, K. D., Chaussidon, M., and Robert, F. (2000). Incorporation of short-lived 10Be in a calcium–aluminum-rich inclusion from the Allende meteorite. Science, 289, 1334–1337. McKeegan, K. D., Kallio, A. P. A., Heber, V. S., et al. (2011). The oxygen isotopic composition of the Sun inferred from captured solar wind. Science, 332, 15281532. McNally, C. P., Hubbard, A., Mac Low, M.- M., et al. (2013). Mineral processing by short circuits in protoplanetary disks. Astrophys. J. Lett., 767, L2. Mendybaev, R. A., Richter, F. M., and Davis, A. M. (2006). Crystallization of melilite from CMAS-liquids and the formation of the melilite mantle of Type B1 CAIs: Experimental simulations. Geochim. Cosmochim. Acta, 70, 2622‒2642. Misawa, K. A., and Fujita, T. (1994). A relict refractory inclusion in a ferromagnesian chondrule from the Allende meteorite. Nature, 369, 163‒165. Meibom, A., Desch, S. J., Krot, A. N., et al. (2000). Large scale thermal events in the solar nebula: Evidence from FeNi metal grains in primitive meteorites. Science, 288, 839‒841. Miller, K. E., Lauretta, D. S., Connolly, H. C., et al. (2017). Formation of unequilibrated R chondrite chondrules and opaque phases. Geochim. Cosmochim. Acta, 209,24‒50. Misawa, K., and Nakamura, N. (1988). Highly fractionated rare-earth elements in ferromagne- sian chondrules from the Felix (CO3) meteorite. Nature, 334,47–50. Morris, M. A., Boley, A. C., Desch, S. J., and Athanassiadou, T. (2012). Chondrule formation in bow shocks around eccentric planetary embryos. Astrophys. J. Lett., 752, L27. Multiple Mechanisms of Transient Heating Events 53

Morris, M. A., Garvie, L. A. J., and Knauth, L. P. (2015). New insight into the Solar System’s transition disk phase provided by the metal-rich carbonaceous chondrite Isheyevo. Astrophys. J. Lett., 801, L22. Nagashima, K., Kunihiro, T., Takayanagi, I., et al. (2001). Output characteristics of stacked CMOS type pixel sensor for charged particles. Surf. Interface Anal., 31, 131‒137. Nagashima, K., Krot, A. N., Libourel, G., and Huss, G. R. (2013). Magnesian porphyritic chondrules surrounded by ferroan igneous rims from CR chondrite GRA 95229. LPSC #44, abstract #1780. Nagashima, K., Krot, A. N., and Libourel, G. (2014). Type I chondrules with ferroan igneous rims from Yamato 81020 CO3 and Acfer 094 ungrouped type 3 carbonaceous chondrites. 77th Ann. Meet. Meteorit. Soc., abstract #5424. Nagashima, K., Krot, A. N., and Park, C. (2015a). An amoeboid olivine aggregate surrounded by an igneous ferroan olivine-rich rim from CO3.0 chondrite DOM 08006. LPSC #46, abstract #2477. Nagashima, K., Krot, A. N., and Huss, G. R. (2015b). Oxygen-isotope compositions of chondrule silicates and matrix grains in Kakangari K-grouplet chondrite. Geochim. Cos- mochim. Acta, 151,49‒67. Nagashima, K., Krot, A., Libourel, G., and Schrader, D. (2016). 16O-rich olivine abundances in FeO-rich chondrules and their igneous rims from CR chondrites. 26th Goldschmidt Conference, abstract #2210. Oulton, J., Humayun, M., Fedkin, A., and Grossman, L. (2016). Chemical evidence for differentiation, evaporation and recondensation from silicate clasts in Gujba. Geochim. Cosmochim. Acta, 177, 254‒274. Palme, H., Spettel, B., and Hezel, D. C. (2014). Siderophile elements in chondrules of CV chondrites. Chem. Erde – Geochemistry, 74, 507–516. Palme, H., Hezel, D. C., and Ebel, D. S. (2015). The origin of chondrules: Constraints from matrix composition and matrix-chondrule complementarity. Earth Planet. Sci. Lett., 411, 11‒19. Rambaldi, E. R. (1981). Relict grains in chondrules. Nature, 293, 558‒561. Rambaldi, E. R., and Wasson, J. T. (1982). Fine, nickel-poor Fe-Ni grains in the olivine of unequilibrated ordinary chondrites. Geochim. Cosmochim. Acta, 46, 929‒939. Rambaldi, E. R., Rajan, R. S., Wang, D., and Housley, R. M. (1983). Evidence for relict grains in chondrules of Qingzhen, an E3 type enstatite chondrite. Earth Planet. Sci. Lett., 66,11‒24. Ramdohr, P. (1967). Chromite and chromite chondrules in meteorites. Geochim. Cosmochim. Acta, 31, 1961–1967. Rubin, A. E. (2003). Chromite-plagioclase assemblages as a new shock indicator: Implications for the shock and thermal histories of ordinary chondrites. Geochim. Cosmochim. Acta, 67, 2695–2709. Rubin, A. E. (2010). Physical properties of chondrules in different chondrite groups: Implica- tions for multiple melting events in dusty environments. Geochim. Cosmochim. Acta, 74, 4807–4828. Rubin, A. E., and Wasson, J. T. (1987). Chondrules, matrix and coarse-grained chondrule rims in the Allende meteorite – Origin, interrelationships and possible precursor components. Geochim. Cosmochim. Acta, 51, 1923–1937. Rubin, A. E., and Krot, A. N. (1996). Multiple heating of chondrules. In R. H. Hewins, R. H. Jones, and E. Scott (Eds.), Chondrules and the Protoplanetary Disk, 173‒180. Cambridge, UK: Cambridge University Press. Russell, S. S., MacPherson, G. J., Leshin, L. A., and McKeegan, K. D (2000). 16O enrichments in aluminum-rich chondrules from ordinary chondrites. Earth Planet. Sci. Lett., 184,57‒74. Russell, S. S., Krot, A. N., MacPherson, G. J., et al. (2005). Genetic relationship between refractory inclusions and chondrules. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), 54 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

Chondrites and the Protoplanetary Disk, 317353. ASP Conf. Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Ruzicka, A. (2012). Chondrule formation by repeated evaporative melting and condensation in collisional debris clouds around planetesimals. Meteorit. Planet. Sci., 47, 2218‒2236. Sanders, I. S., and Scott, E. R. D. (2013). The origin of chondrules and chondrites: Debris from low-velocity impacts between molten planetesimals? Meteorit. Planet. Sci., 47, 2170‒2192. Schrader, D. L., Connolly, H. C., Lauretta, D. S., et al. (2013). The formation and alteration of the Renazzo-like carbonaceous chondrites II: Linking O-isotope composition and oxida- tion state of chondrule olivine. Geochim. Cosmochim. Acta, 101, 302327. Schrader, D. L., Nagashima, K., Krot, A. N., Ogliore, R. C., and Hellebrand, E. (2014). Variations in the O-isotope composition of gas during the formation of chondrules from the CR chondrites. Geochim. Cosmochim. Acta, 132,5074. Scott, E. R. D., and Krot, A. N. (2014). Chondrites and their components. In A. M. Davis (Ed.), Meteorites and Cosmochemical Processes. In H. D. Holland and K. K. Turekian (Eds.), Treatise on Geochemistry (Second Edition), 1, 65137. Oxford, UK: Elsevier. Simon, S. B., Sutton, S. R., and Grossman, L. (2016). The valence and coordination of titanium in ordinary and enstatite chondrites. Geochim. Cosmochim. Acta, 189, 377‒390. Snetsinger, K. G., and Keil, K. (1969). Ilmenite in ordinary chondrites. Amer. Mineral., 54, 780‒786. Sokol, A. K., Bischoff, A., Marhas, K. K., Mezger, K., and Zinner, E. (2007). Late accretion and lithification of chondritic parent bodies: Mg isotope studies on fragments from primitive chondrites and chondritic breccias. Meteorit. Planet. Sci., 42, 1291‒1308. Sossi, P. A., Moynier, F., Chaussidon, M., et al. (2017). Early Solar System irradiation quantified by linked vanadium and beryllium isotope variations in meteorites. Nature Astronomy, 1, id. 0055. Soulié, C., Libourel, G., Tissandier, L., and Hiver, J.- M. (2012). Kinetics of olivine dissolution in chondrule melts: An experimental study. LPSC #43, abstract #1840. Soulié, C., Libourel, G., and Tissandier, L. (2017). Olivine dissolution in molten silicates: Applications to chondrule formation. Meteorit. Planet. Sci., 52, 225‒250. Steele, I. M. (1988). Primitive materials surviving in chondrites: Mineral grains. In J. Kerridge and M. Matthews (Eds.), Meteorites and the Early Solar System , 808‒818. Tucson, AZ: University of Arizona Press. Stolper, E. (1982). Crystallization sequence of Ca-Al-rich inclusions from Allende: An experi- mental study. Geochim. Cosmochim. Acta, 46, 21592180. Tenner, T. J., Ushikubo, T., Kurahashi, E., Kita, N. T., and Nagahara, H. (2013). Oxygen isotope systematics of chondrule phenocrysts from the CO3.0 chondrite Yamato 81020: Evidence for two distinct oxygen isotope reservoirs. Geochim. Cosmochim. Acta, 102, 226245. Tenner, T. J., Nakashima, D., Ushikubo, T., Kita, N. T., and Weisberg, M. K. (2015). Oxygen isotope ratios of FeO-poor chondrules in CR3 chondrites: Influence of dust enrichment and H2O during chondrule formation. Geochim. Cosmochim. Acta, 148, 228250. Testi, L., Birnstiel, T., Ricci, L., et al. (2014). Dust evolution in protoplanetary disks. In H. Beuther, R. Klessen, C. Dullemond, and T. Henning (Eds.), Protostars and Planets VI, 339‒361. Tucson, AZ: University of Arizona Press. Tomeoka, K., and Itoh, D. (2004). Sodium-metasomatism in chondrules in CO3 chondrites: Relationship to parent body thermal metamorphism. Meteorit. Planet. Sci., 39, 13591373. Toppani, A., Paque, J. M., Burnett, D. S., et al. (2006). Wark-Lovering rims at the nanometer scale: A transmission electron microscopy study. LPSC #37, abstract #2030. Multiple Mechanisms of Transient Heating Events 55

Ushikubo, T., Kimura, M., Kita, N. T., and Valley, J. W. (2012). Primordial oxygen isotope reservoirs of the solar nebula recorded in chondrules in Acfer 094 carbonaceous chondrite. Geochim. Cosmochim. Acta, 90, 242264. Vacher, L. G., Marrocchi, Y., Verdier-Paoletti, M. J., Villeneuve, J., and Gounelle, M. (2016). Inward radial mixing of interstellar water Ices in the solar protoplanetary disk. Astrophys. J. Lett., 827, L1. Van Kooten, E. M. M. E., Wielandt, D., Schiller, M., et al. (2016). Isotopic evidence for primordial molecular cloud material in metal-rich carbonaceous chondrites. Proc. Nat. Acad. Sci., 113, 2011‒2016. Villeneuve, J., Libourel, G., and Soulié, C. (2015). Relationships between type I and type II chondrules: Implications on chondrule formation processes. Geochim. Cosmochim. Acta., 160, 277–305. Wakaki, S., Itoh, S., Tanaki, T., and Yurimoto, H. (2013). Petrology, trace element abundances and oxygen isotopic compositions of a compound CAI-chondrule object from Allende. Geochim. Cosmochim. Acta., 102, 261–279. Walsh, K. J., Morbidelli, A., Raymond, S. N., O’Brien, D. P., and Mandell, A. M. (2011). A low mass for Mars from Jupiter’s early gas-driven migration. Nature, 475, 206209. Wasson, J. T., and Kallemeyn, G. W. (1990). Allan Hills 85085 – A subchondritic meteorite of mixed nebular and regolithic heritage. Earth Planet. Sci. Lett., 101, 148‒161. Wasson, J. T., Krot, A. N., Lee, M. S., and Rubin, A. E. (1994). Compound chondrules in ordinary chondrites: Evidence for multiple heating events and for large-scale heterogeni- eties in the nebula. Geochim. Cosmochim. Acta, 59, 1847‒1869. Wasson, J. T., and Rubin, A. E. (2009). Composition of matrix in the CR chondrite LAP 02342. Geochim. Cosmochim. Acta, 73, 14361460. Weber, D., and Bischoff, A. (1994). The occurrence of grossite (CaAl4O7) in chondrites. Geochim. Cosmochim. Acta, 18, 3855‒3877. Weber, D., Zinner, E., and Bischoff, A. (1995). Trace element abundances and magnesium, calcium, and titanium isotopic compositions of grossite-containing inclusions from the carbonaceous chondrite Acfer 182. Geochim. Cosmochim. Acta, 59, 803823. Weidenschilling, S., Marzari, F., and Hood, L. (1998). The origin of chondrules at jovian resonances. Science, 279, 681‒684. Weinbruch, S., Palme, H., and Spettel, B. (2000). Refractory forsterite in primitive meteorites: Condensates from the solar nebula? Meteorit. Planet. Sci., 35, 161171. Weisberg, M. K., Prinz, M., and Fogel, R. A. (1994). The evolution of enstatite and chondrules in unequilibrated enstatite chondrites: Evidence from iron-rich pyroxene. Meteoritics, 29, 362‒373. Weisberg, M. K., Ebel, D. S., Nakashima, D., Kita, N. T., and Humayun, M. (2015). Petrology and geochemistry of chondrules and metal in NWA 5492 and GRO 95551: A new type of metal-rich chondrite. Geochim. Cosmochim. Acta., 167, 269–285. Westphal, A. J., Fakra, S. C., Gainsforth, Z., et al. (2009). Mixing fraction of inner solar system material in Comet 81P/Wild2. Astrophys. J., 694, 18. Whattam, S. A., Hewins, R. H., Cohen, B. A., Seaton, N. C., and Prior, D. J. (2008). Granoblastic olivine aggregates in magnesian chondrules: Planetesimal fragments or thermally annealed solar nebula condensates? Earth Planet. Sci. Lett., 269, 200211. Wlotzka, F. (2005). Cr-spinel and chromite as petrogenetic indicators in ordinary chondrites: Equilibration temperatures of petrologic types 3.7 to 6. Meteorit. Planet. Sci., 40, 16731702. Yurimoto, H., and Wasson, J. T. (2002). Extremely rapid cooling of a carbonaceous-chondrite chondrule containing very 16O-rich olivine and a 26Mg-excess. Geochim. Cosmochim. Acta., 66, 4355–4363. 56 Alexander N. Krot, Kazuhide Nagashima, Guy Libourel, and Kelly E. Miller

Yurimoto, H., Krot, A. N., Choi, B. -G., et al. (2008). Oxygen isotopes of chondritic compon- ents. Rev. Mineral. Geochem., 68, 141187. Zanda, B., Humayun, M., and Hewins, R. H. (2012). Chemical composition of matrix and chondrules in carbonaceous chondrites: Implications for disk transport. LPSC #43, abstract #2413. Zhang, A. -C., Itoh, S., Sakamoto, N., Wang, R. -C., and Yurimoto, H. (2014). Origins of Al- rich chondrules: Clues from a compound Al-rich chondrule in the Dar al Gani 978 carbon- aceous chondrite. Geochim. Cosmochim. Acta., 130,78–92. 3 Thermal Histories of Chondrules Petrologic Observations and Experimental Constraints rhian h. jones, johan villeneuve, and guy libourel

Abstract

Thermal histories of chondrules can be deduced by studying the petrology and mineral chemistry of natural chondrules and their experimental analogs. Dynamic crystallization experi- ments have successfully reproduced chondrule textures, and provide general but broad con- straints on peak temperatures and cooling rates. Porphyritic textures result when a chondrule is heated to a maximum temperature close to, but below, its liquidus, and cooled at initial rates  between about 10 and 1,000 C/h. Typical liquidus temperatures for chondrules range from about 1,400–1,700 C. Nonporphyritic chondrules are produced when peak temperatures exceed the liquidus slightly (for barred/dendritic textures) and significantly (radiating textures) and chondrules cool at rates around 500–3,000 C/h. More quantitative constraints on cooling rates can be determined by considering growth and diffusion-related zoning in chondrule minerals. Results of such modeling are consistent with dynamic crystallization experiments. Rapid dissolution rates for relict olivine grains also indicate a limited time at high temperatures, and indicate fast cooling rates of hundreds to thousands of C/h, close to peak temperatures. Other cooling rate indicators include disequilibrium partition coefficients between minerals and chondrule glass, and consideration of chemical and isotopic diffusion between relict grains and their overgrowths. Interpretation of both these features is currently ambiguous. Several lines of evidence suggest that cooling rates decreased at lower temperatures, as the chondrule approached the solidus, to <50 C/h. These include slow cooling required to nucleate plagio- clase, cooling rates inferred from trace element diffusion profiles in metal grains, and exsolution microstructures in clinopyroxene. In contrast, clinoenstatite microstructures, the presence of chondrule glass, and dislocation densities in chondrule olivine appear to argue for rapid cooling (103–104 C/h) through the lower temperature regime, and textures in opaque (metal/sulfide) assemblages indicate cooling rates of hundreds of degrees per hour at subsolidus temperatures. Overall, thermal histories of chondrules can provide fundamental constraints for chondrule formation models. While high-temperature thermal histories are reasonably well constrained, there are currently some open questions about the nature of the cooling curve at lower temperatures. A better understanding of chondrule cooling rates at lower temperatures would help to discriminate between chondrule formation models that make quantitative predictions for thermal histories. Within a single chondrite, cooling rates may vary widely. It is also possible that the nature of cooling histories varies within a given population of chondrules. A statistical treatment of chondrule populations in which individual chondrules show distinct thermal

57 58 Rhian H. Jones, Johan Villeneuve, and Guy Libourel histories would help to make predictions about chondrule formation environments, and the diversity of processes that might be represented in a single chondrule-forming region.

3.1 Introduction

One of the most fundamental aspects of chondrule formation is that chondrules (or rather, their precursors) were heated to high temperatures capable of producing significant melting, and then cooled quickly enough to preserve a range of disequilibrium features. Any viable model for chondrule formation must be able to reproduce the thermal histories that chondrules record. In this chapter, we assess what is known about chondrule thermal histories, based on a combin- ation of petrologic observations of natural chondrules, and experimental studies of chondrule analogs. Petrologic observations of chondrules involve examination of what minerals are present, how those minerals are arranged (i.e. the chondrule texture), and chemical compos- itions of the minerals. Interpretation of petrologic features relies strongly on complementary experimental studies, which reproduce chondrule petrology under controlled conditions in the laboratory. The earliest petrologic studies of chondrites, using optical microscopy, recognized that chondrules must represent the solidified products of melted droplets: Sorby (1877) described chondrules as “drops of fiery rain”. Modern observations include detailed documentation of the chemical and isotopic compositions and zoning properties of chondrule minerals, using microbeam techniques including scanning electron microscopy (SEM), electron probe micro- analysis (EPMA), secondary ion mass spectrometry (SIMS), and laser-ablation inductively- coupled mass spectrometry (LA-ICP-MS) as well as microstructural information obtained using transmission electron microscopy (TEM). Experimental studies of chondrule analogs flourished in the post-Apollo years, particularly led by G. Lofgren and R. Hewins. In a typical dynamic crystallization experiment, starting material is prepared as a fine-grained mixture of oxide phases and/or mineral grains. The mixture is suspended on a wire loop in the hot zone of a one-atmosphere furnace and heated then cooled under controlled conditions. A mixture of reducing gases flows through the furnace to control the oxygen fugacity (i.e., availability of oxygen). There are many variable and controllable parameters, many of which have been examined in different studies so that a good overall understanding of chondrule thermal histories has emerged. In addition to these experiments, which primarily aim to match textures and zoning properties with natural chondrules, petrologic interpretations of chondrules rely on a large body of more general experimental petrology that includes measurements of partition coefficients and diffusion coefficients for major, minor, and trace elements. We do not make any attempt to review this literature here, but note that this work is a fundamentally important aspect of petrologic studies and interpretations, which we refer to throughout the following discussion. There are 15 different chondrite groups (e.g., Brearley and Jones, 1998), as well as many ungrouped chondrites. One of the complexities of chondrite studies is that the chondrule population differs among different chondrite groups (Rubin, 2010; Jones, 2012). This includes differences in properties such as chondrule size and abundance, as well as the distribution of different textural types of chondrules. In addition, for a given textural type of chondrule the petrologic details, such as composition of olivine and abundance of relict grains, vary among Thermal Histories of Chondrules 59 chondrite groups. Differences among chondrite groups can be attributed to differences in the chemical compositions of chondrule precursor assemblages, the local dust/gas ratio as well as compositions of dust and gas at the time of chondrule formation, and/or thermal histories. In this chapter, we aim to provide an overview of the general constraints on chondrule formation, and do not examine these group-scale differences in detail. However, it is important to remember that chondrule formation models must be able to account for variable localized conditions on a scale of chondrite groups, and that defining the nature of such heterogeneity is an important goal of petrologic and experimental studies. In this chapter, we first discuss different approaches to determining chondrule cooling rates, including experimental reproduction of chondrule textures, considerations of chemical zoning in mineral grains, dissolution rates of relict grains, pyroxene microstructures, and the kinetics of the glass transition. In the discussion, we evaluate how these parameters can be interpreted for several different end-member chondrule formation models. Finally, we discuss future research directions, and discuss how future experimental and chondrite observations could contribute towards evaluating and discriminating between different chondrule formation models.

3.2 Observations and Experiments Relevant to Interpreting Thermal Histories 3.2.1 Thermal Histories Determined from Textural Considerations

3.2.1.1 Overview of Chondrule Petrology The most basic description of chondrules classifies them according to their textures. Since chondrules are solidified droplets of molten material, their textures are indicative of the crystallization process and are described using the terms of igneous petrology. Chondrules exhibit a range of textures, which in order of increasing grain size, and decreasing aspect ratio, are described as cryptocrystalline (CC), radiating (R), barred or dendritic (B), and porphyritic (P) (see Figure 3.1). The dominant silicate mineralogy of chondrules consists of olivine [(Mg,

Fe)2SiO4] and pyroxene [(Ca,Mg,Fe)2Si2O6]: pyroxene typically consists of grains of low-Ca pyroxene with Wo < 2 mole % (where Wo = atomic Ca/(Ca+Mg+Fe)), commonly overgrown with narrow rims of high-Ca pyroxene with Wo > 30 mole % (e.g., Brearley and Jones, 1998). Barred textures are usually composed of bars of olivine (barred olivine, or BO chondrules), and radiating textures are usually composed of needles of pyroxene (radiating pyroxene, or RP chondrules). Porphyritic textures can include any ratio of olivine to pyroxene (PO = porphyritic olivine with >90% olivine, PP = porphyritic pyroxene with >90% pyroxene, POP = porphyritic olivine and pyroxene). A further chondrule classification refers to the initial (unequilibrated) FeO content of olivine and pyroxene (Brearley and Jones, 1998; see Figure 3.1): Chondrules which have Fa and Fs <10 mole %, are designated as type I, and those with Fa and Fs >10 mole %, are designated as type II (Fa = atomic Fe/(Fe+Mg) in olivine; Fs = atomic Fe/(Fe+Mg+Ca) in pyroxene). Most type I chondrules, particularly in carbonaceous and enstatite chondrites, have Fa < 2 mole % (e.g., Brearley and Jones, 1998). Within this classification, olivine-rich chondrules are desig- nated as type A and pyroxene-rich chondrules as type B: for example, a description as type IAB refers to an FeO-poor, porphyritic, olivine- and pyroxene-bearing chondrule. Type I and type II chondrules differ in several important petrologic respects, for example type I chondrules are 60 Rhian H. Jones, Johan Villeneuve, and Guy Libourel

Figure 3.1 Typical textures of natural chondrules from a variety of chondrites (backscattered electron images). Nonporphyritic chondrules include cryptocrystalline (CC), radiating and barred (or dendritic) textures. Radiating textures are commonly pyroxene (P)-rich (RP); barred and dendritic textures are commonly olivine (O)-rich (BO). Porphyritic (P) textures include olivine-rich (PO, or type A), olivine- and pyroxene-rich (POP, or type AB), and pyroxene-rich (PP, or type B). Chondrules with MgO-rich (Mg# >90) primary olivine and pyroxene compositions are designated type I, and chondrules with more FeO-rich (Mg# <90) primary olivine and pyroxene compositions are designated type II (Mg# = atomic Mg/(Mg + Fe)). Hence, for example, a chondrule with MgO-rich olivine and pyroxene is a type IAB, POP chondrule. Abbreviations: O = olivine, P = pyroxene, R = relict grain. Thermal Histories of Chondrules 61 highly reduced and can contain abundant Fe,Ni metal. Olivine and pyroxene grains in type II chondrules are often larger than those in type I chondrules, and commonly show significant compositional zoning arising from disequilibrium growth. Material interstitial to olivine and pyroxene is described as chondrule mesostasis. In unequi- librated chondrites, mesostasis in many chondrules is glassy, with a feldspathic composition. The glass represents the final liquid present during cooling, and commonly contains fine-grained elongate microcrystallites, which are interpreted as quench crystals. Mesostasis in unequili- brated chondrites can also be predominantly crystalline, typically consisting of an intergrowth of plagioclase (CaAl2Si2O8 – NaAlSi3O8) and high-Ca pyroxene with minor amounts of glass. Chondrules also contain an iron-nickel-sulfur component, containing minerals such as troilite, pyrrhotite, pentlandite, Fe,Ni metal, and magnetite. Within unequilibrated ordinary chondrite (UOC) type I chondrules, troilite generally occurs in association with Fe,Ni metal and oxidized minerals in opaque assemblages (hereafter OAs), so named because these minerals are opaque in transmitted light (Blum et al., 1989; Rubin et al., 1999; Tachibana and Huss, 2005). OAs commonly occur as spheroidal inclusions, which could result from the oxidation and sulfidation of Fe,Ni metal beads in the solar nebula (Lauretta et al., 1996; Lauretta et al., 1997) or form during secondary parent body alteration. Sulfides could also have been formed via the crystallization of Fe–Ni–S melts during chondrule cooling (e.g., Tachibana and Huss, 2005; Schrader, Davidson, and McCoy, 2016a). The presence of sulfur within metallic melts in chondrules has been interpreted as evidence that sulfur must have been present in the chondrule precursor material (Rubin et al., 1999). However, it has also been proposed that sulfides in chondrules formed through high-temperature processes of sulfur solubility within chondrule melts, followed by sulfide saturation (Marrocchi and Libourel, 2013; Marrocchi et al., 2016; Piani et al., 2016).

Minor chromite (FeCr2O4) is common in type II chondrules, and spinel (MgAl2O4) occurs rarely in type I chondrules. Silica (SiO2) occurs commonly in chondrules from enstatite (E) and CR chondrites, and rarely in ordinary (O) and other carbonaceous (C) groups. In addition to chondrules dominated by olivine and pyroxene, there are well-defined aluminium-rich chon- drules (ARC), which include plagioclase-rich chondrules (PRC) and glass-rich chondrules, as well as rare chondrules that contain unusually high abundances of chromites (see summaries in Brearley and Jones, 1998; Connolly and Desch, 2004).

3.2.1.2 Dynamic Crystallization Experiments That Have Reproduced Chondrule Textures Several textural features indicate that chondrules were largely molten at their highest tempera- tures, and that crystallization occurred during quite rapid cooling. Key textural features of rapid cooling include dendritic, elongate, and hopper morphologies of olivine and pyroxene grains, which result from disequilibrium crystal growth at significant degrees of undercooling (i.e., at temperatures below the equilibrium temperature for crystallization of the mineral phase). Many experimental studies have reproduced basic chondrule textures (Figure 3.2), and this has enabled us to place quantitative constraints on chondrule cooling rates. Earlier experiments have been summarized previously in several review papers (Lofgren, 1996; Desch and Connolly, 2002; Hewins et al., 2005; Desch et al., 2012), to which we refer the reader for further details. 62 Rhian H. Jones, Johan Villeneuve, and Guy Libourel

Figure 3.2 Examples of chondrule textures produced in dynamic crystallization experiments (back- scattered electron images). a) Experimental charge CH1–58, from Lofgren (1989) and Jones and Lofgren (1993): type II composition, cooled at 100 C/h, porphyritic olivine texture. The charge shows gravitational settling during the experiment (Up arrow points to top of charge): Lower zone (L) consists of fine-grained olivine, grown during the period the charge was held at the peak temperature; Central zone (C) consists of olivine grains that grew during the cooling period; Upper zone (U) is glassy and contains skeletal olivine that grew during the quench. b) Experimental charge Exp7 from Wick and Jones (2012): type IAB composition cooled at 25 C/h to 1,000 C, then 10 C/h to 800 C. Porphyritic olivine-pyroxene texture with high-Ca pyroxene overgrowths on low-Ca-pyroxene. c) Experimental charge CH1–61, from Jones and Lofgren (1993): type II composition, cooled at 100 C/h, then held at 1,200 C for 7 days before quenching. Porphyritic olivine-pyroxene texture. d) 3D tomographic images of partially resorbed olivine after heating experiments in a type I PO-like melt (c3 composition in Soulié et al., 2017) at 1531C for 366s (#E1) and 903s (#D7), and 1,465C for 598s (#G3) and 956s (#G4). Abbreviations: Ol = olivine, Pyx = Low-Ca pyroxene, Cpx = high-Ca pyroxene, Gl = glass.

One of the most important interpretations of chondrule textures is that they are largely controlled by the availability of nucleation sites for crystal growth as the melt cools (Lofgren and Russell, 1986; Radomsky and Hewins, 1990; Lofgren, 1996). Homogeneous nucleation occurs when nanometer-scale crystalline embryos within the melt grow to dimensions that exceed the critical radius for growth. In contrast, heterogeneous nucleation occurs on surfaces of existing grains, although these may be submicroscopic. Chondrule textures are most likely Thermal Histories of Chondrules 63 controlled by heterogeneous nucleation, which requires that solid material from the precursor assemblage is incompletely melted. Because of the importance of nucleation considerations, the relationship between texture and chondrule cooling rate is, to a certain extent, ambiguous. It is possible to produce different textures at the same cooling rate if the initial availability of nucleation sites is variable. Since availability of nucleation sites is a function of several variables, including precursor grain size, peak temperature, and time, textures alone do not provide unique thermal histories for specific textural types of chondrules, although they do provide general constraints. Figure 3.3a shows a cartoon generalizing the thermal history of radial, barred, and porphyritic chondrules, with thermal histories shown relative to the chon- drule liquidus. The liquidus temperature is a strong function of chondrule composition, and varies from around 1,400–1,700 C (e.g., Radomsky and Hewins, 1990). The cartoon in Figure 3.3a represents the typical case for most chondrule analogue experiments, in which chondrule precursors consist of fine-grained mineral grains, and the time spent at peak tempera- tures is relatively short (less than two hours). In general, porphyritic textures are produced when the peak temperature is below the liquidus, thus preserving multiple nucleation sites. Barred textures are produced when the peak temperature is slightly above the liquidus temperature, such that most, but not all, nucleation embryos are destroyed. Radial textures are produced from temperatures that clearly exceed the liquidus temperature, and in which few or no nucleation sites are available. These general statements are based on numerous experimental studies that have been described in detail in previous reviews (Lofgren, 1996; Desch and Connolly, 2002; Hewins et al., 2005; Desch et al., 2012). Ranges of cooling rates for each texture, given as labels on the cooling paths, give our best estimates based on the data shown in Figure 3.3b. In Figure 3.3b, we summarize the results of chondrule cooling rates that have been determined by reproducing different textures in dynamic crystallization experiments. The figure is similar to one drawn by Desch et al. (2012), but it differs from theirs because we interpret several of the experimental studies differently, based on our own evaluation of the papers included in the figure. As an example, our figure shows significant differences for cooling rates of nonporphyritic chondrules taken from Kennedy et al. (1993). We believe that significantly slower cooling rates shown by Desch et al. (down to 100 C/h and 1 C/h for BO and RP respectively) are because they misinterpreted Table 2 of Kennedy et al, in which headings of BO and RP refer to the bulk composition of the experimental charge, and not the texture of the resulting experiment. Overall, Figure 3.3b shows that there is a range of cooling rates for each textural type of chondrule, and that there is a general trend of increasing cooling rate from porphyritic to barred to radial textures. Most of the experiments represented in Figure 3.3b were cooled at single, linear, cooling rates from temperatures close to the liquidus to a specified quench temperature that was commonly significantly above the solidus (e.g., 1,200 C: Figure 3.2a). This means that cooling rates typically represent the high-temperature crystallization history of the chondrule. For type I porphyritic chondrules, experiments of Wick and Jones (2012: see Figures 3.2b and 3.3b #2) are an exception to this, as they examined slow and variable cooling rates to low temperatures (<1,000 C). Cooling rates for type I chondrule analog experiments are mostly in the range 10–500 C/h. Lower cooling rates (1–10 C/h) are more relevant to the lower temperature range (Wick and Jones, 2012): these conditions were found to be necessary in order to crystallize plagioclase, which is present in a limited but statistically significant number of type I chondrules. 64 Rhian H. Jones, Johan Villeneuve, and Guy Libourel

Temperature a

Liquidus in the range 1400 – 1700 °C

500-3000 1000-3000 °C/h °C/h

1-500 °C/h

(Assuming solid precursors) Time

10000

8 1 14 6 11 12 11 5 3 3 13 1 1000 4 4,9 4,9 10

10 7 13 6 8 100

2 Cooling rate, °C/h Cooling rate,

10

b 1 Type I Type II Barred Radial Porph Porph

Figure 3.3 Summary of chondrule heating and cooling conditions from experimental constraints. (a) Cartoon showing maximum (peak) temperature relative to the liquidus, for (from left to right) porphyritic, barred, and radial textures. The cartoon assumes solid precursors and short times (maximum 2 hours) at the peak temperature. The ranges of cooling rates given on the cooling paths are best estimates taken from panel (b). Small circles in the chondrules at peak temperatures represent nucleation embryos which may be submicroscopic. (b) Summary of chondrule cooling rate ranges from experiments that have reproduced chondrule textures. Bars show individual studies numbered as below. All experiments except for (2) had linear cooling rates; (2) used a sequence of decreasing cooling rates. The gray bar (5) refers to Thermal Histories of Chondrules 65

Several studies of type II porphyritic chondrule compositions give similar ranges of cooling rates to type I chondrules – i.e. 2–500 C/h (Figure 3.2c). In the work of Radomsky and Hewins (1990; Figure 3.3b #3), only one experiment with a cooling rate >100 C/h, an experiment at   1,000 C/h, gave a porphyritic texture: all others at >100 C/h had dendritic textures. Connolly and Hewins (1995; Figure 3.3b #10) produced a porphyritic texture at a cooling rate of 500 C/h when a large amount of fine-grained dust was puffed onto the molten chondrule. Experiments of Connolly and Hewins (1991; Figure 3.3b #4) and Connolly et al. (1998; Figure 3.3b #9) were all conducted at 500 C/h, but different textures were produced by varying bulk composition and grain sizes of precursor materials respectively. Type II chondrule textures were also produced by Nettles et al. (2006) in experiments that were heated to peak temperatures significantly below the liquidus (1,250–1,450 C) and cooled at rates of 10–1,000 C/h. Cooling rates for barred (or dendritic) textures from several studies are quite consistent, in  the range 500–3,000 C/h. For radiating textures, a wide range of cooling rates is shown in Figure 3.3b. Experiments of DeHart and Lofgren (1996; Figure 3.3b #1), and Hewins et al. (1981; Figure 3.3b #14) had relatively short dwell times (less than two hours) at peak temperatures. In contrast, experiments of Lofgren and Russell (1986; Figure 3.3b #6) and Kennedy et al. (1993; Figure 3.3b #11) had extended dwell times, above the liquidus, of 17 hours or more. As a result, all potential nucleation sites were destroyed and radiating textures  were produced over a wide range of cooling rates, down to 5 C/h. Connolly and Hewins (1995; Figure 3.3b #10) produced a radiating texture at a cooling rate of 500 C/h when a small amount of fine-grained dust was puffed onto the molten chondrule, thus seeding nucleation. Most of Figure 3.3b refers to experiments with bulk compositions that match ferromagnesian chondrules, in which olivine and pyroxene are the dominant silicate minerals. Dynamic crystal- lization experiments on aluminum-rich chondrule compositions, as well as a type C CAI compos- ition, were investigated by Tronche et al. (2007; Figure 3.3b #5). Porphyritic textures matching natural chondrules were produced at cooling rates from 10 – 1,000 C/h, with peak temperatures below liquidus temperatures. As for ferromagnesian chondrules, the main controlling factor on texture was the nucleation density, controlled by peak temperature relative to the liquidus.

3.2.2 Thermal Histories Constrained from Mineral Chemistry and Zoning in Olivine and Pyroxene

A promising approach to put more quantitative constrains on thermal history is the study of chemical and isotopic properties of chondrule components and their experimental analogs.

Caption for Figure 3.3 (cont.) aluminium-rich chondrules – all others are ferromagnesian chondrules. Stars (10) are experiments carried out at a cooling rate of 500 C/h, and dust was puffed onto the molten droplet as it cooled. Most experiments had dwell times at peak temperature of less than two hours. The exception is (6) which had dwell times of 17 hours. Experiments are numbered as follows: 1) DeHart and Lofgren, 1996; 2) Wick and Jones, 2012; 3) Radomsky and Hewins, 1990; 4) Connolly and Hewins, 1991; 5) Tronche, Hewins, and MacPherson, 2007; 6) Lofgren and Russell, 1986; 7) Jones and Lofgren, 1993; 8) Lofgren, 1989; 9) Connolly, Jones, and Hewins, 1998; 10) Connolly and Hewins, 1995; 11) Kennedy, Lofgren, and Wassenburg, 1993; 12) Lofgren and Lanier, 1990; 13) Tsuchiyama et al., 2004; 14) Hewins, Klein, and Fasano, 1981. 66 Rhian H. Jones, Johan Villeneuve, and Guy Libourel

During the last 30 years, tens of studies have been conducted on this topic (e.g., reviews by Lofgren, 1996; Hewins et al., 2005) that have provided strong inputs to chondrule formation models. They have also raised numerous unanswered questions as our knowledge of chondrules has increased. In the following we review constraints provided by this approach but also limitations and unknowns.

3.2.2.1 Growth Zoning Porphyritic olivine in type IIA and IIAB chondrules shows ubiquitous zoning patterns, with progressive enrichment in FeO as well as trace elements (Cr2O3, CaO, MnO) from core to rim. These patterns are interpreted as normal igneous zoning, produced during cooling while olivine crystallized in a closed system (e.g., Jones, 1990; Jones and Lofgren, 1993; Hewins et al., 2005). For instance, Jones (1990) showed that zoning in Fe-rich olivine in type IIA chondrules from Semarkona is consistent with closed-system fractional crystallization during a single thermal event with peak temperatures around 1,600 C and cooling rates ~1,000 C/h. Dynamic crystallization experiments (Lofgren and Russell, 1986, Jones and Lofgren, 1993; Rocha and Jones, 2012) and diffusion modeling (Miyamoto et al., 2009) succeeded in reproducing these observations for cooling rates ranging from 10–1,000 C/h, consistent with cooling rates inferred from textures (Figure 3.4). In order to explain large size olivine phenocrysts (up to 1,000 µm, e.g., Jones, 1990, 1996a) commonly observed in type II chondrules, relatively slow cooling rates, <50C/h, were usually proposed. This is in good agreement with dynamic experiments and also observations from olivine in terrestrial basalts (e.g., Welsch et al., 2013). These observations and experiments suggest that individual chondrules form during a single melting event and cooled slowly. In addition to normal zoning, olivine and pyroxene can show more complex zoning in type II chondrules. For instance, zoning of phosphorus is common in terrestrial and lunar magmatic olivine, as well as in type II chondrules (Milman-Barris et al., 2008; McCanta et al., 2008, 2016; Hewins et al., 2009; Wasson et al., 2014; Rubin et al., 2015; Villeneuve et al., 2015). Because P diffuses very slowly in olivine, it reveals details of the complex behavior of olivine during igneous processes experienced by chondrules, such as relict grains present in cores of pheno- crysts, annealing of skeletal cores, and oscillatory growth zoning (e.g., McCanta et al., 2008, 2016; Hewins et al., 2009; Villeneuve et al., 2015). Similarly, low-Ca pyroxene in type II chondrules also shows Fe oscillatory zoning (Watanabe et al., 1986; Jones, 1996a; Wasson et al., 2014). For some authors, this oscillatory zoning in FeO-rich olivine and low-Ca pyroxene hints at chondrule formation consisting of multiple brief heating events (Wasson et al., 2014; Rubin et al., 2015; Baecker et al., 2017). However, such a scenario is not relevant for explaining comparable oscillatory zoning in terrestrial magmatic olivine and pyroxene (e.g., Milman-Barris et al., 2008). Furthermore, experiments have successfully reproduced P zoning in olivine during single heating events at slow cooling rates (McCanta et al., 2008), and a melt–crystal interface effect that induces fluctuations in crystal growth with no link to the cooling rate is advocated to explain oscillatory zoning in olivine (McCanta et al., 2008; Milman-Barrat et al., 2008; Hewins et al., 2009). Similar melt–crystal interface effects were advocated to explain FeO oscillatory zoning in low-Ca pyroxene (e.g., Jones, 1996a and references therein). Hence, we argue that complex zoning in chondrule olivine and pyroxene is entirely consistent with a single-stage cooling model. Thermal Histories of Chondrules 67

Figure 3.4 Experimental and modeled zoning compared to natural igneous zoning measured in type II chondrule olivine from the ordinary chondrite Semarkona. a) FeO profiles in olivine from Semarkona and dynamic crystallization experiments at different cooling rates (Jones and Lofgren, 1993). b) FeO, Cr2O3, MnO, and CaO profiles from the same Semarkona olivine as a), black curves, and from experiments with a cooling rate of 30 C/h, gray curves (Rocha and Jones, 2012). c) and d) are modeled zoning that fit profiles measured in olivine (Miyamoto et al., 2009).

In some type II chondrules, FeO-rich olivine grains with normal zoning (FeO enrichment from the core to the edge) have widely varying grain sizes (from a few µm to hundreds of µm) and different Fa contents at the cores of small and large olivine grains (e.g., Wasson and Rubin, 2003). Based on these observations, these authors proposed, at odds with the single melting model, that the two sets of olivine could not have crystallized from the same melt at the same time, and thus that type II chondrules experienced at least two heating events. In that case, cores of large olivine grains are relicts and only the small olivine grains and the outer part of the large olivine grains crystallized during the last melting event, implying a brief second heating event. Similar to olivine, low-Ca pyroxene in type IIAB and IIB chondrules often shows normal zoning with enrichments in Fe and minor elements from core to rim. Cooling rates of ~100 C/h have been proposed based on zoning properties (e.g., Jones, 1996a). Since pyroxene crystallizes at lower temperature than olivine and is quite ubiquitous in chondrules, it can provide good 68 Rhian H. Jones, Johan Villeneuve, and Guy Libourel

 hints on the thermal history at temperatures below 1,500 C. However, to our knowledge, there is currently no experimental study that gives us quantitative information on thermal history experienced by pyroxene during its crystallization in chondrules.

3.2.2.2 Partition Coefficients It is well known that kinetic factors have an important effect on major and trace element partitioning between crystal and melt (e.g., Albarède and Bottinga, 1972): a fast cooling rate induces large deviations from the equilibrium values. Based on olivine, low-Ca-pyroxene, and glass assemblages produced over isothermal and dynamic experiments (cooling rates from  1–2,191 C/h), Kennedy et al. (1993) attempted to constrain the effect of kinetics on trace element partitioning. They observed that for dynamic experiments with cooling rates lower than 100 C/h, apparent partition coefficients (D*) did not depart from equilibrium partition coeffi- cients (Kd). In contrast, for experiments at higher cooling rates, D* of incompatible elements increased up to 3 orders of magnitude relative to Kd, although D* of compatible elements did not change significantly. Jones and Layne (1997) determined D* of trace elements (rare earth elements (REE) plus Sr, Y, and Zr) between pyroxene and mesostasis in chondrules from the Semarkona LL chondrite. For heavy rare earth elements (HREE) Sr, Y, and Zr, D* values show no significant differences from equilibrium values determined by Kennedy et al. (1993), while for light rare earth elements (LREE), D* values are higher than equilibrium values and HREE/ LREE ratios lower, in agreement with Alexander (1994). The authors suggested the importance of rapid cooling rates to explain such REE patterns. More recently, Jacquet et al. (2012, 2015) measured trace element abundances in olivine, low-Ca pyroxene, and mesostasis in type I and type II chondrules from CO and LL3 ordinary chondrites. From REE patterns and partitioning between olivine and mesostasis, they proposed that type I chondrules experienced a signifi- cant degree of batch (equilibrium) crystallization that implies cooling rates slower than 1–100 C/h, and probably < 10 C/h, while type II chondrules formed by fractional crystallization at higher cooling rates. In contrast, low-Ca pyroxene, which crystallized at a lower temperature  than olivine, seemed to have recorded rapid cooling rates of the order of 1,000 C/h, at odds with the thermal path usually inferred for chondrules from observations and experiments (e.g., Hewins et al., 2005: see previous text). However, the use of trace element partitioning between crystal and melt as a tool for constraining chondrule thermal histories remains uncertain, since (i) there are very few experiments that allow us to constrain precisely the impact of kinetics on partition coefficients, and (ii) the role of glass inclusions trapped in crystals is ambiguous since that would strongly affect measured D* (Kennedy et al., 1993; Jones and Layne, 1997). Studies on the olivine-melt Fe/Mg Kd or D* in chondrules show significant departures from equilibrium values of 0.3 (Roeder and Emslie, 1970), in highly unequilibrated ordinary and carbonaceous chondrites (Figure 3.5; Taylor and Cirlin, 1986; Jones and Scott, 1989; Jones, 1990; Lofgren and Le, 1998; Symes and Lofgren, 1999). Very low values of partitioning (down to 0.01) are commonly documented in type I PO chondrules of ordinary chondrites, while systematically higher values – often significantly higher than equilibrium (up to 0.9) – are observed in type II PO chondrules (Figure 3.5). BO chondrules, which are believed to have experienced higher cooling rates than porphyritic chondrules (Figure 3.3b), do not show Thermal Histories of Chondrules 69

Figure 3.5 Range of partition coefficient (Kd), or apparent partition coefficient (D*), for Fe–Mg between olivine and melt measured in various types of chondrules (black) and in experimental chondrule proxies (gray). The solid horizontal and bold line corresponds to the equilibrium value (0.3, Roeder and Emslie, 1970). The downward gray arrow indicates the direction of increasing cooling rates from 10 C/h–2,000 C/h (Symes and Lofgren, 1998; Nettles et al., 2006). Data are from Taylor and Cirlin (1986), Jones and Scott (1989), Jones (1990), Lofgren and Le (1998) and Symes and Lofgren (1998) for chondrules and from Wick and Jones (2012), Symes and Lofgren (1998) and Nettles et al. (2006) for experimental proxies. Except for Jones and Scott (1989) data, which are calculated from bulk olivine compositions, all the data correspond to olivine compositions as close as possible to the interface with the mesostasis, compared with bulk mesostasis compositions. Symes and Lofgren (1999) and Nettles et al. (2006) data correspond to Kds calculated for several individual olivine crystals from different chondrules while the others are average Kds for individual chondrules.

D* significantly different from type I PO chondrules (Figure 3.5; Taylor and Cirlin, 1986). D*  measured in experiments performed at low cooling rates (<25 C/h) that reproduced type I POP chondrule proxies give values between 0.04 and 0.11, lower than equilibrium and compatible with observations in type I PO chondrules (Figure 3.5; Wick and Jones, 2012). Based on experiments performed under open-system conditions, Villeneuve et al. (2015) pointed out that very low D* (<0.1), as measured in type I chondrules, can be easily reproduced if chondrules experienced chemical disequilibrium due to gas–melt interactions between crystals and melt during their formation, and thus might not be related to any kind of kinetic crystallization effect. In experimental type II PO chondrule proxies, measured D* seems to be more or less correlated with the cooling rate. Indeed, D* decreases from the equilibrium value to 0.10 for cooling rates increasing from 10C/h to 2,000 C/h (Figure 3.5; Symes and Lofgren, 1999; Nettles et al., 2006). However, partitioning values significantly higher than equilibrium values as observed in type II chondrules are not reproduced experimentally and thus cannot be explained by kinetic effects (Figure 3.5). Finally, it is worth noting that measured Kd or D* can be easily altered if minerals other than olivine, such as pyroxene, have also crystallized from the melt, as it is often the case in chondrules. 70 Rhian H. Jones, Johan Villeneuve, and Guy Libourel

3.2.2.3 Relict Grains A lot of chondrules contain relict grains that are not in chemical equilibrium with surrounding minerals and mesostasis, e.g., MgO-rich olivine grains in type II chondrules, dusty olivine in type I chondrules, relict olivine embedded in pyroxene crystals in POP and PP chondrules (e.g., Nagahara, 1981; Rambaldi, 1981; Miyamoto et al., 1986; Jones, 1996b; Wasson and Rubin, 2003; Jones and Carey, 2006; Ruzicka et al., 2007; Ruzicka et al., 2008; Chapter 2; see Figure 3.1). Relicts of MgO-rich olivine enclosed in FeO-rich olivine overgrowths in type II chondrules are very common with occurrences in >90% of chondrules in CO3.0 chondrites (Jones, 1992; Wasson and Rubin, 2003; Kunihiro et al., 2004), ~80% in Acfer 094 (Kunihiro et al., 2005) and >15% in type 3 ordinary chondrites (Jones, 1996b; Jones, 2012). The origin of these relicts remains uncertain, but they clearly come from reduced precursors that contain coarse grains (e.g., Jones, 1996b; Ruzicka, 2012; Hewins and Zanda, 2012). Also, MgO-rich relicts often show oxygen isotopic compositions that are 16O-rich compared to host FeO-rich olivine (e.g., Jones et al., 2000; Connolly and Huss, 2010; Kita et al., 2010; Libourel and Chaussidon, 2011; Rudraswami et al., 2011; Schrader et al., 2013; Tenner et al., 2013, 2015, Chapter 7; Ushikubo et al., 2012, 2013). Thus, objects like amoeboid olivine aggregates (AOAs) or type I-like chondrules can be proposed as the source of these relict grains (e.g., Jones, 1990, 1996b; Hewins and Zanda, 2012; Schrader et al., 2013, 2015; Villeneuve et al., 2015). The fact that olivine relicts are ubiquitous in chondrules and, as explained above (Section 3.2.1.2; Figure 3.3a), the necessity to preserve nuclei during the high-temperature phase of chondrule formation to obtain porphyritic textures, have led people to propose upper limits for peak tempera- tures, very fast cooling rates at high temperatures and/or very short heating histories (e.g., Hewins et al., 2005). Recently, an experimental study aimed at constraining precisely the dissolution rate of olivine, and thus the survival duration of relict olivine in chondrule-like melts, was conducted (Soulié et al., 2017; Figures 3.3d and 3.6). These experiments examined FeO-free compositions. Results of this study indicate that (i) forsterite dissolution is a very fast process from 5µm/min to 22µm/min at temperatures ranging 1,450–1,540 C and (ii) the dissolution rate strongly depends on the melt composition with decreasing rates for increasing MgO and SiO2 contents of the melts (Figure 3.6a). Empirical modeling based on forsterite saturation and viscosity of the melts succeeded at reproducing the forsterite dissolution rates as a function of peak temperatures, cooling rates, and chemical compositions (Figure 3.6b–c). It shows for instance that for a type I PO-like melt and peak temperatures of 1,500 C, cooling rates as high as 1,000–8,000 C/h are required to preserve olivine, while cooling rates are a bit lower, i.e. 300–3,000 C/h, for a type I PP-like melt (Figure 3.6b–c). In a chondrule, olivine dissolution may be slowed by the faster saturation of melt due to higher olivine/ melt ratios compared to the Soulié et al. (2017) experiments. However, as observed in type I POP chondrules, interactions with SiO gas enhance the dissolution of MgO-rich olivine (Soulié et al., 2017). Overall, these observations strongly suggest that porphyritic chondrules that contain olivine relict grains must have experienced cooling rates significantly higher than those traditionally inferred from dynamic crystallization experiments (Figure 3.3b). Another thermal constraint based on forsterite relicts comes from examining Fe–Mg inter- diffusion between the relict grain and either the melt or the fayalite-rich overgrowth. Thus, Villeneuve et al. (2015) showed in an experimental study that the survival of a forsterite relic of 50 µm in a type II PO-like chondrule would not exceed 10–20 minutes at 1,800 C and ~5 hours at 1,500 C. Chemical diffusive exchanges between relict forsterite grains and surrounding Thermal Histories of Chondrules 71

(a) Temperature (°C) 1400 1300 1200 1100 1000 900 5 5 0 0 m) m –5 4 –5 –10 Type IB –10 3 –15 –15 –20 –20 –25 2 –25 –30 –30 Type IA –35 –35 –40 1 Tpeak: 1445°C –40 –45 Cooling rate: 1000 K/h –45 Oliving radius variation ( variation Oliving radius –50 –50 0102030 Time (min) (b) Cooling rate (K/s) 0.001 0.01 0.1 1 2100 Resorbed length 1800 m 2000 50 m Type IA 40 mm 1700 30 mm 1900 20 mm 10 mm 1600 1800 1500 1700 1400 Peak temperature (°C) temperature Peak Peak temperature (K) temperature Peak 1600 1300 2 1 1600 1 10 100 1000 10000 Cooling rate (K/h) (c) Cooling rate (K/s) 0.001 0.01 0.1 1 2100 1800 Resorbed length 2000 50 mm Type IB 40 mm 1700 30 mm 1900 20 mm 10 mm 1600 1800 1500 1700 1400 Peak temperature (K) temperature Peak 1600 (°C) temperature Peak 1300 1500 1 10 100 1000 10000 Cooling rate (K/h)

Figure 3.6 Dissolution of olivine in chondrule-like melts as a function of the thermal regime. The figure is modified from Soulié et al. (2017). a) Evolution of the olivine dissolution distance during a linear cooling episode for various melt compositions. Composition 1 and 2 (in black and dark gray respectively) correspond to type IA mesostasis, while composition 3 and 4 (in black and dark gray) correspond to type IB mesostasis (i.e., increase of SiO2 content from type IA to type IB). Compositions 2 and 4 are derived from compositions 1 and 3 respectively, with 15 wt% enrichment in modal forsterite (i.e., increase of MgO content compared to compositions 1 and 3). These different starting compositions allow us to evaluate the contributions of viscosity and forsterite activity onto olivine dissolution rates (see Soulié et al. (2017) for details). b) and c) Empirical modeling of resorbed lengths (from 10 to 50 µm) as a function of the peak temperature and the cooling rate for the two types of melts. fayalite-rich overgrowths in type II chondrules have been studied during the last few years (e.g., Greeney and Ruzicka, 2004; Béjina et al., 2009, Hewins et al., 2009). Indeed, diffusion profiles for elements across the relict/overgrowth interface, such as Fe, Mg, Cr, Ca, or Mn, which have different diffusion rates, may provide powerful constraints on the cooling rates experienced by 72 Rhian H. Jones, Johan Villeneuve, and Guy Libourel type II chondrules. Thanks to the extensive study of diffusion in terrestrial minerals, empirical parameters for diffusion equations in olivine are well known and provide good inputs for diffusion modeling in olivine from chondrules (e.g., Jurewicz and Watson, 1988; Petry et al., 2004; Dohmen and Chakraborty, 2007; Coogan et al., 2005; Ito and Ganguly, 2006; Chakra- borty, 2010; Tachibana et al., 2013). However, a source of uncertainty comes from the fact that extrapolations from experimental conditions (temperature and/or oxygen fugacity) are often needed to match chondrule formation conditions. Cooling rates calculated by different authors from Fe, Mg, Cr, Ca, and Mn diffusion profiles in relicts from different ordinary chondrites scatter over several orders of magnitude between 1 and 104 C/h, with a majority  well above the 10–103 C/h range inferred from dynamic crystallization experiments of type II chondrules (Greeney and Ruzicka, 2004; Béjina et al., 2009; Hewins et al., 2009). Similarly, oxygen isotopic diffusion modelling between a forsterite relic and its overgrowth in a CO3.0  chondrite points toward a very high cooling rate, i.e., 105–106 C/h (Yurimoto and Wasson, 2002). These discrepancies between chondrules may be due to anisotropies in diffusion rates according to the crystalline structure (Hewins et al., 2009). However, this mismatch is absolutely irreconcilable when such scattering is observed within a single olivine grain, e.g., 340–18,500 C/h (Greeney and Ruzicka, 2004) or 700–3,000 C/h (Béjina et al., 2009). To illustrate this issue, Figure 3.7 shows the time needed to fitdiffusionprofiles for Fe and Cr from an olivine grain studied by Hewins et al. (2009) as a function of the linear cooling rate. As an example, for no cooling rate (i.e., an isothermal condition) and a temperature of 1,630 C, 2–5 minutes are needed to fit the Fe profile while 40–50 minutes areneededfortheCrprofile (Figure 3.7a, b, and d). The divergence becomes even worse when a cooling rate is considered. As illustrated in Figure 3.7d, (i) there is an important anisotropy of diffusion according to the crystal axis for both Fe and Cr (Tachibana et al., 2013, Ito and Ganguly, 2006) and (ii) uncertainties on experimental diffusion parameters can be quite large, e.g., diffusion for Cr along the c axis, illustrated by the thin dotted curve on Figure 3.7d (Ito and Ganguly, 2006). Obviously, such properties limit the use of chemical diffusion profiles in relict MgO-rich olivine to constrain the thermal history of chondrules. Furthermore, apart from cooling rate the gas-melt interactions during chondrule formation must be considered as a possible source for explaining diffusion profiles observed at the relict/ overgrowth interface (Villeneuve et al., 2015).

3.3 Constraints on Thermal Histories from Other Considerations

Along with textures and mineral chemistry of silicates, several other parameters can be used to interpret thermal histories of chondrules. We now describe constraints on the thermal history of chondrules that can be made from trace element diffusion profiles in metal grains, exsolution and microstructures in pyroxene, the glass transition, and dislocations in olivine.

3.3.1 Cu and Ga Diffusion Profiles of Metal Grains

Using laser ablation ICP-MS tracks that provided spatially resolved abundance distributions, Humayun (2012) and Chaumard et al. (2015) have shown that zoning profiles for Cu and Ga are Thermal Histories of Chondrules 73

Figure 3.7 Diffusion profiles of Fe and Cr between a MgO-rich olivine relict and its FeO-rich overgrowth. a) and b) show measured Fe and Cr profiles (dots) and example of modeled diffusion curve within a single type II chondrule olivine (white arrow on c). Data are from Hewins et al., 2009. Dark curves are examples of modeled diffusion profiles. d) Time needed to fit the profiles as a function of linear cooling rate and initial temperature of 1630 C. For Fe diffusion, the bold solid curve is based on Dohmen and Chakraborty (2007) diffusion parameters and the three bold dashed curves are based on Tachibana et al. (2013) diffusion parameters determined for the different crystal orientations indicated. Bold dotted curves are based on Ito and Ganguly (2006) diffusion parameters along a and c crystal axes for Cr diffusion. The thin dotted line that crosses the bold black curve is the fastest curve for Cr diffusion according to uncertainties reported by Ito and Ganguly (2006), for the c axis. observed in metal lumps in chondrules. When combined with precise diffusion coefficients (Righter et al., 2005), these zoning profiles provide constraints on the thermal history of chondrules as witnessed by metal (Figure 3.8). The cooling rates of metal lumps from the Acfer 097 CR2 chondrite were calculated to be    0.5–50 C/h for a temperature 1,200 Æ 100 C, with a maximum possible range of 0.1–400 C/h for temperatures of 927–1,527 C (Humayun, 2012; Figure 3.8). Similar results in the range of 1.2–86 C/h for temperature 1,300–1,400 C were also obtained using the same method from metal grains of type I chondrules in the Renazzo CR2 chondrite (Chaumard et al., 2015). These cooling rates are similar to, or systematically lower than, those determined from experimental studies of type I chondrule textures, 10–500 C/h (see previous text; Figure 3.3b) and broadly similar to cooling rates obtained from pyroxene exsolution in type I chondrules from carbon- aceous chondrites (Weinbruch and Müller, 1995; see Section 3.3.2). Humayun and coworkers 74 Rhian H. Jones, Johan Villeneuve, and Guy Libourel

Figure 3.8 Plot of cooling rates (C/h) against peak temperatures implied by different diffusion profiles of Cu, as a function of the characteristic length-scale of diffusion, a (µm). Each curve is a solution for a specific value of a, as labeled. Also shown is the melting point of Fe–Ni alloys (gray bar) that forms the ultimate limit on the peak temperature. The temperature range inferred from Cu-Ga thermometry (1,100–1,300 C) is shown as the gray solid box. Dynamic crystallization experiments for type I porphyritic chondrules yield cooling rates of 10–1,000 C/ h) (Radomsky and Hewins, 1990; Connolly and Hewins, 1995; DeHart and Lofgren, 1996) and peak temperatures of 1,600–1,700 C (gray mottled box). Figure adapted from Humayun (2012). assume thus that chondrule textures were established near the peak heating temperatures of chondrules (approximately 1,600–1,700 C), while the Cu and Ga diffusive profiles were established after solidification (T  1,227 C), consistent with nonlinear cooling, and consistent with nonlinear slow cooling rates (1–25 C/h) at low temperatures, from the dynamic crystal- lization experiments of Wick and Jones (2012) (see Figure 3.3b).

3.3.2 Exsolution Lamellae in Clinopyroxene

Lamellar exsolution microstructures in clinopyroxene have been used by several authors to constrain cooling rates of chondrules (Kitamura et al., 1983; Watanabe et al., 1985; Kitamura et al., 1986; Müller et al., 1995; Weinbruch and Müller, 1995; Weinbruch et al., 2001). In contrast to texture, mineral zoning, and trace element distributions, which yield cooling rates at temperatures between liquidus and solidus, the wavelength of exsolution lamellae in clinopyr- oxene can be used to infer cooling rates at subsolidus temperatures. Using an experimental approach, the coarsening of exsolution lamellae on (001) during cooling was modeled. Cooling rates between 7 and 25 C/h were obtained for Allende iron-poor granular olivine pyroxene chondrules (Weinbruch and Müller, 1995). Extension of this seminal work to iron-bearing clinopyroxenes has led the same group (Weinbruch et al., 2001) to obtain more accurate estimates of chondrule cooling rates. The subsolidus cooling rates of chondrules from various  H, L, LL, and CV chondrites are certainly below 50–60 C/h; in some cases they may have been as low as 0.1–0.2 C/h. These cooling rates are valid at temperatures between approximately 1,400 and 1,200 C for Fe-poor clinopyroxene and 1,100–900 C for Fe-rich clinopyroxene. From this clinopyroxene perspective, it thus seems that slow cooling of chondrules at subsolidus temperatures is a widespread phenomenon. Thermal Histories of Chondrules 75 3.3.3 Clinoenstatite Microstructure

Low-Ca pyroxene is the second most abundant crystalline phase in chondrules and provides additional constraints on thermal histories. Enstatite (i.e., MgO-rich, low-Ca pyroxene) shows three polymorphs: protoenstatite (orthorhombic), orthoenstatite (orthorhombic), and clinoensta- tite (monoclinic). Protoenstatite is the stable form at high temperature (>1,000 C). However, it is not quenchable: it is never observed at room temperature and spontaneously inverts to orthoenstatite, clinoenstatite, or a mixture of both during cooling, as soon as the temperature drops below 1,000 C. Experiments of Smyth (1974) made it possible to determine the conditions for inversion of proto- to clinoenstatite or orthoenstatite. He showed that the clino-/orthoenstatite ratios depend on the cooling rate. Only an instantaneous quenching (<1s) results in unmixed clinoenstatite, which is twinned. For slower cooling rates, a mixture of orthoenstatite and clinoenstatite in the form of microstructural intergrowths is obtained, in which the proportion of orthoenstatite increases gradually as the cooling rate decreases. Clinoenstatite, which is metastable at low temperature, is inverted to orthoenstatite when heated (as seen in metamorphosed chondrites). The work of Brearley and Jones (1993) developed a more quantitative experimental approach that links the proportions of two pyroxenes with cooling rate over a range of 2 C/h–10,000 C/h. Their results confirm those of Smyth (1974), that only cooling rates in the range of 10,000 C/h are able to produce pure clinoenstatite. In natural chondrules, where the polymorphism of enstatite has been studied extensively (e.g., Ashworth, 1980; Yasuda, Kitamura, and Morimoto, 1983; Watanabe et al., 1985), clinoenstatite has been shown to be abundant. It is generally accepted that it results from the inversion of protoenstatite during rapid cooling. Passage through the high-temperature poly- morph, protoenstatite, is attested by the observation of a network of fractures perpendicular to the <001> axis of the pyroxene (Figure 3.9a), resulting in the shortening that accompanies the transformation of orthoenstatite to clinoenstatite along this axis (protoenstatite c = 5.351 Å, clinoenstatite: c = 5.169 Å; Yasuda et al., 1983). An electron backscatter diffraction (EBSD) study of the Vigarano CV3 chondrite (Soulié, 2014) confirms the observations already made in other chondrites (Figure 3.9). The enstatite fracture network indicates the crystallization of pyroxenes at high temperature (>1,000 C). The proportion of clinoenstatite/orthoenstatite in porphyritic chondrules is very high, between 70 and 87 percent, and could possibly reach 100 percent in some PO chondrules (see also Weinbruch and Müller, 1995 in the case of Allende). Based on the experimental calibration of Brearley and Jones (1993), these proportions indicate highly variable cooling rates for porphy-    ritic chondrules between 250 C/h and 2,000 C/h and possibly as high as 10,000 C/h at a temperature around 1,000 C.

3.3.4 Presence of Glass and Critical Cooling Rate

From a general point of view, the most important transformation of a magma is its solidification due to cooling (i.e., the transition from a silicate melt to a rock). Solidified magmas may be crystalline, vitreous, or a mixture of glass and crystals. If the cooling rate is high enough to prevent crystallization, a magma can encompass the supercooling region without crystallization. In principle, the ability of a silicate liquid to persist in the amorphous state during cooling is 76 Rhian H. Jones, Johan Villeneuve, and Guy Libourel

Figure 3.9 Intergrowth of clinoenstatite (CEn) and orthoenstatite (OEn) in porphyritic chondrule 42 from Vigarano CV3 chondrite. a) BSE image showing the fracturing network in low-Ca pyroxene (arrows). b) EBSD image (Euler angles) showing polysynthetic twins in the clinoenstatite. Blue and green colours show CEn twin orientations. c) EBSD map of the entire chondrule. Clinoenstatite: yellow; orthoenstatite: gray; olivine: green; Fe–Ni metal: red. (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.) defined as glass-forming ability (Avramov et al., 2003; Cabral et al., 2003). The most common way to quantify the glass-forming ability is the critical cooling rate Rc, which is the minimum cooling rate at which a liquid can be frozen to a solid glass without crystallization, or with a percentage of crystals below 1 vol.% (Cabral et al., 2003). Glass-forming ability and the critical cooling rate Rc can be perceived in terms of two parameters, which reveal the ease or difficulty of a silicate melt to nucleate. Experimental data and theoretical models have shown that the glass-forming ability and the critical cooling rate of melts are a complex function of a) their chemical composition (e.g., SiO2 or bulk polymerization), which can be quantitatively esti- mated using the reduced glass transition parameter Trg = Tg/Tm (Tg, temperature of glass transition; Tm, temperature of melting), and b) the viscosity fragility concept (Cabral et al., 2003; Iezzi et al., 2009). For example, for terrestrial lavas, these data have shown that to vitrify a completely melted (silica-poor) basalt (from liquidus to superheated temperatures) a cooling rate of 360–1,800 C/h is not sufficient, while (silica-rich) rhyolitic glass can be obtained with cooling rates lower than 1–10 C/h. Intermediate subalkaline melts (basaltic andesites, andesites, and latites) are able to produce a glass with cooling rate values between those required for basalts and rhyolites (Iezzi et al., 2009). These general rules, being also valid for chondrule formation, suggest that the Thermal Histories of Chondrules 77 occurrence of glassy mesostases in many type I PO chondrules with composition as low as 45–50 2 3  wt% SiO2 suggests a high cooling rate (>10 –10 C/h) in the range of their glass transition temperature, around 680–880 C as calculated using the Dulong-Petit law. However, chondrule glass compositions vary widely, up to 73 wt% SiO2 in Semarkona chondrules (e.g., Brearley and Jones, 1998).

3.3.5 Dislocations in Olivine

Theoretically, dislocations in olivine are expected to develop during cooling of chondrules, as the transformation of liquid into glass is supposed to generate plastic strain within the crystal, due to the difference between the thermal expansion coefficients of glass and crystal. It has been shown that constant cooling experiments in the CaO–MgO-Al2O3-SiO2 (CMAS) system per- formed down to low temperature (900–800 C) provide samples in which the olivine crystalline phase displays a very high density of dislocations (see Figure 11 in Faure et al., 2003). In contrast, cooling experiments quenched above 1,000 C reveal olivines free of dislocations, even if performed at various degrees of undercooling (Faure et al., 2003). The unusual low density of dislocations observed in chondrule olivine crystals (see Figure 2.12) suggests quenching (i.e., high cooling rates) from high temperature, well above the glass transition temperature.

3.4 Discussion

From the wide variety of measurements discussed in Section 3.3, including observations of natural chondrules and experimental analogs, we can provide some general constraints on chondrule thermal histories. The presence of sulfur in chondrule precursors appears to constrain ambient temperatures in the chondrule-forming region to <400 C (Rubin et al., 1999), although this may be ambiguous (as discussed in Section 3.5). Peak temperatures in the chondrule-forming region are constrained broadly by the need to heat most chondrules, which have porphyritic textures, to temperatures close to, but not exceeding, the liquidus (Figure 3.3a).  However, because liquidus temperatures vary widely, from about 1,400–1,700 C, this con- straint is not very specific. Also, since nonporphyritic (barred and radial) textures require peak temperatures that exceed the liquidus, the upper limit on peak temperatures is somewhat poorly  defined. An upper limit of 1,800 C is reasonable, but higher peak temperatures – up to 2,000 C – cannot be excluded. A low-temperature limit is also difficult to constrain, as there is evidence for only limited or incipient partial melting in chondrules known as agglomeratic chondrules (Weisberg and Prinz, 1996; Ruzicka et al., 2012), corresponding to peak tempera- tures 1,200 C (e.g., Nettles et al., 2006). Determination of chondrule cooling rates is a more complex question, because it is often difficult to separate the many variables that can influence a particular observation, and because different cooling rates have been determined using different parameters, as discussed above. In the following discussion, we attempt to synthesize the interpretation of chondrule cooling rates, based on three different cooling models as illustrated in Figure 3.10. Model A is a simple model which involves single-stage, linear cooling from peak temperatures. Model B describes a 78 Rhian H. Jones, Johan Villeneuve, and Guy Libourel

Figure 3.10 Cartoon showing different pathways for chondrule thermal histories. The dashed gray line (Model A) indicates a linear cooling history from liquidus to low temperatures, with cooling rates ranging from 100–1,000 C/h. The black line (Model B) indicates a generalized thermal history determined for shock-wave models for chondrule formation (e.g., Desch et al., 2012), including a rapid rise from ambient to peak temperature, followed by decay of cooling rate over time, from cooling rates around 100–1,000 C/  h at temperatures close to the liquidus (TL) to cooling rates around 1–50 C/h at temperatures close to the glass transition temperature (Tg). The solid gray line (Model C) indicates an alternative cooling history, including a period of slow cooling of a few C/h at temperatures close to the liquidus followed by rapid cooling (1,000 – 10,000 C/h). For the two gray lines, no prepeak heating history is specified.

constantly decaying cooling rate, and is based on thermal histories predicted from shock wave models for chondrule formation (e.g., Desch et al., 2012; Chapter 15). Model C describes a model which is opposite from Model B, in that cooling rate is slow at high temperatures and then the chondrule is cooled rapidly through lower temperatures. By comparing these models, and assessing the evidence to support the cooling rate histories they infer, we hope to illustrate current uncertainties regarding chondrule cooling rates and key questions that need to be addressed in order to advance our understanding of this topic.

3.4.1 Continuous Linear Cooling Rates (Model A)

The simplest cooling rate model we can consider is one of linear cooling which extends at a constant rate from (super-) liquidus to subsolidus temperatures (Figure 3.10, Model A). Because of the simplicity of such a model, many studies of chondrules are based on this assumption. In terms of chondrule formation models, a continuous linear cooling rate would be particularly  valid if cooling rates are fast: a chondrule cooling at 1,000 C/h will cool from its peak temperature to below ambient temperature in less than two hours, and a small change in cooling rate (such as a decrease to hundreds of degrees per hour) over this interval would likely have Thermal Histories of Chondrules 79 negligible effects on chondrule properties. Our Model A is thus shown with cooling rates in the range of 100–1,000 C/h. Most dynamic crystallization experiments that reproduce chondrule textures have been conducted with continuous linear cooling rates, as well as relatively short durations (less than two hours) at peak temperatures (Figure 3.3b). Cooling rates of hundreds to thousands of degrees per hour are broadly consistent with both porphyritic and nonporphyritic chondrule textures, as well as features such as growth zoning in olivine grains from type II chondrules, dissolution rates of olivine precursor grains, diffusion profiles between relict grains and their overgrowths, and clinoenstatite microstructures. In addition, comparisons of metal/sulfide textures in chondrules (Tachibana et al., 2006; Schrader and Lauretta, 2010; Mori et al., 2016; Schrader, Fu, and Desch, 2016b) and experimental analogs (Tachibana et al., 2006) give subsolidus cooling rates around 100 C/h, indicating a fairly constant cooling rate and an  ambient temperature well below 400 C (Schrader et al., 2016b). The strength of a model that has linear cooling rates of 100–1,000 C/h is that all chondrule textures can be produced at these cooling rates, and differences between porphyritic and nonporphyritic chondrules can be explained by the peak temperature relative to the liquidus, which is a function of either chondrule composition, peak temperature, or both (Figure 3.3a). A linear cooling model with slower cooling rates is more problematic. Although slower  cooling rates (<100 C/h) are very reasonable for porphyritic textures, it is difficult to reproduce nonporphyritic textures under slow-cooled conditions. An experimental study that produced radial textures at slow cooling rates had a very long duration (17 hours) at peak temperatures (Lofgren and Russell, 1986; Figure 3.3b). An extended heating time at superliquidus tempera- tures is inconsistent with the constraints for porphyritic textures (i.e., short heating times at subliquidus temperatures). Thus, in order to call upon slow cooling rates throughout the chondrule-forming region, one would need widely varying conditions at peak temperatures. Although a continuous linear cooling model such as Model A has many strengths, several lines of evidence suggest that a more complex model is needed to fully explain all the observations. We now discuss Models B and C in Figure 3.10 to illustrate these complexities.

3.4.2 Nonlinear Cooling Rates, Constantly Decaying Cooling Rate (Model B)

One of the fundamental questions for a valid chondrule formation model is the manner in which the cooling path approaches the ambient temperature. Model B (Figure 3.10) is close to a thermal history inferred from shock wave models (e.g., Desch and Connolly, 2002; Morris and Desch, 2010; Desch et al., 2012 and refs therein), in which primordial dusty materials were flash heated   from ambient temperatures (<400 C) to above liquidus temperature (1,400–1,700 C) and then cooled at a rate slow enough to allow crystals to form. In these models, for example bow shocks driven by planetesimals or embryos on eccentric orbits, or large-scale shocks driven by gravitational instabilities (Desch et al., 2012; Morris et al., 2012, 2016), it is generally assumed that cooling rates from peak temperature are very high, i.e. in the range  103–104 C/h, and then slow down to  101–103 C/h in relation to the optical depth or the dust/gas ratio of the medium. For individual chondrules, heating duration from the peak temperature down to the closure of the system at the glass transition temperature (Tg) must have varied between several hours to several days. 80 Rhian H. Jones, Johan Villeneuve, and Guy Libourel

There is evidence from chondrule and experimental observations for a continuous decay of cooling rate through the cooling interval. Linear cooling rates of 100–1,000 C/h determined for reproduction of chondrule textures (see above) are largely dependent on features that are set at high temperatures, during the crystallization interval of olivine and pyroxene. These include olivine and pyroxene grain morphology, chemical compositions, growth zoning properties, disequilibrium apparent partition coefficients, and diffusion profiles between relict grains and  their overgrowths. However, several observations point to slower cooling rates (1–50 C/h) at lower temperatures. These include the need for slow cooling in order to crystallize plagioclase (Wick and Jones, 2012), exsolution microstructures in clinopyroxene (e.g., Weinbruch and Müller, 1995; Weinbruch et al., 2001), and the interpretation of diffusion-generated zoning profiles for trace elements in metal grains (Humayun, 2012; Chaumard et al., 2016), as discussed above. Although the presence of plagioclase is only required for a relatively limited subset of porphyritic chondrules (around 10% of type I chondrules in carbonaceous chondrites), the absence of plagioclase is not inconsistent with slow cooling. Experiments of Wick and Jones (2012) show that porphyritic textures, low values of apparent Kd for Fe–Mg, and the presence of glass, are features of experiments that were cooled slowly to low temperatures. One important feature of chondrule petrology that appears to be inconsistent with a non- linear decaying cooling curve is the high ratio of clinoenstatite / orthoenstatite observed in low- Ca pyroxene in type I chondrules, which indicates very high cooling rates in the range of thousands of degrees per hour (Fig. 3.9). If this cooling rate indicator is indeed a true reflection of cooling rate at temperatures around 1,000 C, it could potentially provide a robust and quantitative cooling rate constraint. However, at present it is difficult to reconcile this interpret- ation with other low-temperature cooling rate indicators. Further work on quantifying this thermometer and the temperature of the transition it records would be very valuable. Model B approximates the cooling curve predicted for a shock wave model (Desch et al., 2012; Wick and Jones, 2012). The cooling curves in these models approach very slow cooling rates (degrees per hour) at temperatures around 1,000 C. One important question is that it is not clear how ambient temperatures are achieved in shock wave models. If ambient temperatures were<400 C prior to chondrule heating, a very extended period of cooling is implied: cooling over 600 C at a rate of 1 C/h would take 600 hours, or approximately 25 days, and if cooling rates continue to decrease, this will be extended even more. Investigations of chondrules and experimental analogs with a focus on features that record processes in this low temperature range would be very helpful in defining chondrule thermal histories.

3.4.3 Nonlinear Cooling Rates, Two-Stage Thermal History (Model C)

Model C (Figure 3.10) proposes that porphyritic olivine chondrules, after periods of limited cooling as short as several tenths of minutes at high temperature, terminate their formation by fast cooling in the order of 103–104 C/h) or a quench, in order to prevent further crystallization and elemental diffusion. This model is based at high temperature on the behavior of olivine, i.e., the liquidus phase in most chondrules, and on the behavior of secondary phases (low- and high- Ca pyroxenes, glass) in the low temperature range. Model C suggests in fact that i) the main driving force for establishing the texture and the chemical gradients observed in chondrule olivines is not the cooling rate (dT/dt), but instead the Thermal Histories of Chondrules 81 chemical disequilibrium (dC/dt) occurring between chondrules and their gaseous environment, in which they formed, and that ii) the phase relationships and chemical compositions formed at high temperature are preserved (frozen) by the following quench or the fast cooling regime. As summarized in the previous text, this two-stage thermal history is supported by Fe–Mg inter- diffusion between relict and overlying olivines, the preservation of olivine grains in chondrule mesostasis which otherwise would have been dissolved according to their fast kinetics of dissolution, the preservation of glassy mesostases, and the high clinoenstatite/orthoenstatite ratio due to the inversion of protoenstatite during rapid cooling. The cooling curves from model C are significantly different from the ones predicted from shock heating (e.g., Morris and Desch, 2010; Desch and Connolly, 2012; Morris et al., 2012; Morris, Weidenschilling, and Desch, 2016). They are, on the other hand, very similar to cooling curves predicted by an alternative model stating that chondrules might have originated as a result of impacts or planetesimal collisions producing a splash of molten droplets that cooled down to become chondrules (Sanders and Taylor, 2005; Hevey and Sanders, 2006; Asphaug et al., 2011; Fedkin et al., 2012; Sanders and Scott, 2012; Fedkin and Grossman, 2013; Johnson et al., 2015). In this framework, the radiative cooling of a ballistically expanding spherical cloud of chondrule droplets ejected from an impact site has been recently modeled (Dullemond, Stammler, and Johansen, 2014; Dullemond et al., 2016). It was found that the temperature after the start of the expansion of the cloud remains constant for a time of the order of few tens of minutes and then drops rapidly. The time at which this temperature drop starts depends on the mass of the cloud, the expansion velocity and the size of the chondrule. The observation that, during the early subisothermal expansion phase, the density is so high that no large volatile losses are expected (Dullemond et al., 2016), is also consistent with the volatile behavior recorded in chondrules (e.g., Chapter 6).

3.5 Summary and Outlook for Future Research

As is clear from the previous discussion, there are many ways in which chondrule thermal histories can be investigated through petrology and experimental studies. Because some of these approaches appear to give contradictory interpretations, it is important to understand why that is the case, and to clarify what interpretations are well understood and what remains ambiguous. We can consider several parts of a chondrule thermal history. First, we need to know the ambient temperature in the chondrule-forming region. For many years now it has been agreed  that this is constrained to <391 C(<664 K ) by the presence of sulfur, which occurs in chondrule opaque assemblages (Rubin et al., 1999; Lodders, 2003). Since the 50-percent condensation temperature (Tc) for S is one of the lowest of all elements, this constraint has been considered to be robust. However, it is also possible that the sulfur present in chondrules is the result of open-system behavior when the chondrule interacted with surrounding gas (Tachi- bana and Huss, 2005; Schrader and Lauretta, 2010; Marrocchi and Libourel, 2013; Schrader et al., 2015, 2016a; Marrocchi et al., 2016; Piani et al., 2016), making the interpretation of ambient temperature somewhat ambiguous. It is unlikely that we will be able to confidently lower the estimated ambient temperature without, for example, strong arguments for the presence of elements such as Br, I, and Hg in chondrule precursors (condensation temperature 82 Rhian H. Jones, Johan Villeneuve, and Guy Libourel of 546, 535, and 252 K, respectively: Lodders, 2003) as well as confidence in the values of Tc for these elements, which are currently poorly known. The heating part of a chondrule’s thermal history is difficult to investigate. Several models have referred to “flash” heating of durations of minutes or less (e.g., Connolly et al., 1998), but it is unclear what quantitative constraints such experiments place on heating rates. Currently, our constraints on heating rates are included in constraints for heating durations at peak temperatures. For example, preservation of a relict grain that has rapid dissolution rates in a chondrule melt limits the duration of heating as well as the time spent at peak temperature. Also, constraints on heating rates and times spent at peak temperature can be made by the need to retain volatile elements such as sodium in chondrule melts in an open-system environment, with the caveat that in an open system such elements might recondense into a chondrule melt during cooling (see Chapter 6). Tachibana and Huss (2005) argued that heating rates must be >104–106   C/h in the temperature range of 1,000–1,200 C, in order to suppress the isotopic fractionation of sulfur from an Fe-S melt. Otherwise, we currently have few constraints on chondrule heating rates from ambient to peak temperatures. As discussed in Sections 3.2 and 3.3, peak temperatures for chondrule formation could range from 1,200–2,000 C. This range does not describe an uncertainty, but is a true reflection of the record preserved in chondrules. At the low-temperature end of the range, we observe agglom- eratic chondrules that apparently underwent only incipient or limited melting. At the upper end, we observe nonporphyritic chondrules with high liquidus temperatures, for which the peak temperatures clearly exceeded the liquidus. However, despite this wide range, we can make generalizations on chondrule peak temperatures for certain chondrite groups. For example, in carbonaceous chondrites, the majority of chondrules are porphyritic, and the majority of those are MgO-rich, type I chondrules that have high liquidus temperatures. Hence, we can generally state that most chondrules in carbonaceous chondrites were heated to peak temperatures around 1,600 C. Cooling rates are perhaps the most ambiguous part of chondrule thermal histories. Several lines of evidence point to short durations at peak temperatures, and rapid cooling rates (hundreds to thousands of C/h) at high temperatures, close to the peak temperature and in the initial cooling interval when crystallization of olivine and pyroxene largely takes place. For porphyritic textures, ambiguity in cooling rates arises because the same texture can be produced over a range of cooling rates (e.g., tens to hundreds of C/h). More quantitative measures of cooling rates also give a range of rates, leading to the conclusion that a range of cooling rates is actually represented in the chondrule population. A critical aspect of understanding cooling rates is the need to clarify the shape of the cooling curve as the chondrule cools (Figure 10). As discussed above, arguments can be made for a constantly decaying cooling curve (Model B), or for a rapid cooling event from high temperatures (Model C). One would expect a robust chondrule formation model to allow for cooling to approach ambient conditions (i.e., <400 C), although there are few observational constraints that this is necessary, or time limits on the subsolidus part of the cooling history. An improved understanding of chondrule cooling rates at low temperatures (<1,000 C) will be an important direction for future research. Promising avenues for such investigations include studies of trace elements in metal grains, microstructural investigations of minerals including pyroxene and olivine, studies of textures in opaque (metal/ sulfide) phases, investigation of the glass transition, and studies of the kinetics of crystallization Thermal Histories of Chondrules 83 of chondrule minerals at low temperatures. In all cases, a close relationship between observa- tions of natural chondrules and experimental analogs is essential, as has proved highly success- ful in the literature reviewed in this chapter. An important conclusion for chondrule formation models is that they should be able to accommodate a range of peak temperatures and thermal histories, as discussed in Chapter 15. Within a single chondrite, peak temperatures and cooling rates for individual chondrules may vary widely. It is also possible that the nature of cooling histories varies within a given population of chondrules (i.e. there may be a mix of chondrules with thermal histories that approximate Models A, B, and C of Figure 3.10). A statistical treatment of chondrule popula- tions in which individual chondrules show distinct thermal histories would help to make predictions about chondrule formation environments, and the diversity of processes that might be represented in a single chondrule-forming region.

Acknowledgments

The authors would like to thank D. Schrader for a helpful review that improved the manuscript.

References

Albarède, F., and Bottinga, Y. (1972). Kinetic disequilibrium in trace element partitioning between phenocrysts and host lava. Geochimica et Cosmochimica Acta, 36, 141–156. Alexander, C. M. O’D. (1994). Trace element distributions within ordinary chondrite chon- drules: Implications for chondrule formation conditions and precursors. Geochimica et Cosmochimica Acta, 58, 3451–3467. Asphaug, E., Jutzi, M., and Movshovitz, N. (2011). Chondrule formation during planetesimal accretion. Earth and Planetary Science Letters, 308, 369–379. Ashworth, J. R. (1980). Chondrite thermal histories: Clues from electron microscopy of orthopyroxene. Earth and Planetary Science Letters, 46, 167–177. Avramov, I., Zanotto, E. D., and Prado, M. O. (2003). Glass-forming ability versus stability of silicate glasses. II. Theoretical demonstration. Journal of Non-Crystalline Solids, 320, 9–20. Baecker, B., Rubin, A. E., and Wasson, J. T. (2017). Secondary melting events in Semarkona chondrules revealed by compositional zoning in low-Ca pyroxene. Geochimica et Cosmo- chimica Acta, 211, 256–279. Béjina, F., Sautter, V., and Jaoul, O. (2009). Cooling rate of chondrules in ordinary chondrites revisited by a new geospeedometer based on the compensation rule. Physics of the Earth and Planetary Interiors, 172,5–12. Blum, J. D., Wasserburg, G. J., Hutcheon, I. D., Beckett, J. R., and Stolper, E. M. (1989). Origin of opaque assemblages in C3V meteorites: Implications for nebular and planetary pro- cesses. Geochimica et Cosmochimica Acta, 53, 543–556. Brearley, A. J., and Jones, R. H. (1993). Chondrite thermal histories from low-Ca pyroxene microstructures: Autometamorphism versus prograde metamorphism revisited. Lunar and Planetary Science Conference, 24, 185–186. Brearley, A. J., and Jones, R. H. (1998). Chondritic meteorites. Reviews in Mineralogy and Geochemistry, 36, 3-1–3-398. Cabral, A. A., Cardoso, A. A. D., and Zanotto, E. D. (2003). Glass-forming ability versus stability of silicate glasses. I. Experimental test. Journal of Non-Crystalline Solids, 320, 1–8. 84 Rhian H. Jones, Johan Villeneuve, and Guy Libourel

Chakraborty, S. (2010). Diffusion coefficients in olivine, wadsleyite and ringwoodite. Reviews in Mineralogy and Geochemistry, 72, 603–639. Chaumard, N., Humayun, M., Zanda, B. and Hewins, R. H. (2015). Cooling rates of type I chondrules from the Renazzo CR2 chondrite: Implications for chondrule formation. Lunar and Planetary Science Conference #46, Abstract #1907. Connolly, H. C. Jr., and Desch, S. J. (2004). On the origin of the “kleine Kügelchen” called chondrules. Chemie der Erde, 64,95–125. Connolly, H. C. Jr., and Hewins, R. H. (1991). The influence of bulk composition and dynamic melting conditions on olivine chondrule textures. Geochimica et Cosmochimica Acta, 55, 2943–2950. Connolly, H. C. Jr., and Hewins, R. H. (1995). Chondrules as products of dust collisions with totally molten droplets within dust-rich nebular environment: An experimental investi- gation. Geochimica et Cosmochimica Acta, 59, 3231–3246. Connolly, H. C. Jr., and Huss, G. R. (2010). Compositional evolution of the protoplanetary disk: Oxygen isotopes of type-II chondrules from CR2 chondrites. Geochimica et Cosmo- chimica Acta, 74, 2473–2483. Connolly, H. C. Jr., Jones, B. D., and Hewins, R. H. (1998). The flash melting of chondrules: An experimental investigation into the melting history and physical nature of chondrule precursors Geochimica et Cosmochimica Acta, 62, 2725–2735. Coogan, L. A., Hain, A., Stahl, S., and Chakraborty, S. (2005). Experimental determination of the diffusion coefficient for calcium in olivine between 900 C and 1500 C. Geochimica et Cosmochimica Acta, 69, 3683–3694. Davidson, J., Schrader, D. L., Lauretta, D. S., et al. (2014). Petrology, geochemistry, stable isotopes, Raman spectroscopy, and presolar components of RBT 04133: A reduced CV3 carbonaceous chondrite. Meteoritics and Planetary Science 49, 2133–2151. DeHart, J. M., and Lofgren, G. E. (1996). Experimental studies of group A1 chondrules. Geochimica et Cosmochimica Acta, 60, 2233–2242. Desch, S. J., and Connolly Jr., H. C. (2002). A model of the thermal processing of particles in solar nebula shocks: Application to the cooling rate of chondrules. Meteoritics and Planetary Science, 37, 183–207. Desch, S. J., Morris, M. A., Connolly Jr., H. C., and Boss, A. P. (2012). The importance of experiments: Constraints on chondrule formation models. Meteoritics and Planetary Science, 47, 1139–1156. Dohmen, R., and Chakraborty, S. (2007). Fe–Mg diffusion in olivine II: point defect chemistry, change of diffusion mechanisms and a model for calculation of diffusion coefficient in natural olivine. Physics and Chemistry of Minerals, 34, 409–430. Dullemond, C. P., Stammler, S. M., and Johansen, A. (2014). Forming chondrules in impact splashes. I. Radiative cooling model. The Astrophysical Journal, 794, 91. Dullemond, C. P., Harsono, D., Stammler, S. M., and Johansen, A. (2016). Forming chondrules in impact splashes. II. Volatile retention. The Astrophysical Journal, 832, 91. Faure, F., Trolliard, G., and Soulestin, B. (2003). TEM investigation of forsterite dendrites. American Mineralogist, 88, 1241–1250. Fedkin, A. V., and Grossman, L. (2013). Vapor saturation of sodium: Key to unlocking the origin of chondrules. Geochimica et Cosmochimica Acta, 112, 226–250. Fedkin, A. V., Grossman, L., Ciesla, F. J., and Simon, S. B. (2012). Mineralogical and isotopic constraints on chondrule formation from shock wave thermal histories. Geochimica et Cosmochimica Acta, 87,81–116. Greeney, S., and Ruzicka, A. (2004). Relict forsterite in chondrules: Implications for cooling rates. Lunar and Planetary Science Conference #35, Abstract #1426. Hevey, P. J., and Sanders, I. S. (2006). A model for planetesimal meltdown by 26Al and its implications for meteorite parent bodies. Meteoritics and Planetary Science, 41, 95–106. Thermal Histories of Chondrules 85

Hewins, R. H., and Zanda, B. (2012). Chondrules: Precursors and interactions with the nebular gas. Meteoritics and Planetary Science, 47, 1120–1138. Hewins, R. H., Klein, L. C., and Fasano, B. V. (1981). Conditions of formation of pyroxene excentroradial chondrules. Proceedings of the 12th Lunar and Planetary Science Confer- ence, 448–450. Houston, TX: Lunar and Planetary Institute. Hewins, R. H., Connolly Jr., H. C., Lofgren, G. E., and Libourel, G. (2005). Experimental constraints on chondrule formation. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, 286–316. ASP Conference Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Hewins, R. H., Ganguly, J., and Mariani, E. (2009). Diffusion modeling of cooling rates of relict olivine in Semarkona chondrules. Lunar and Planetary Science Conference #40, Abstract #1513. Humayun, M. (2012). Chondrule cooling rates inferred from diffusive profiles in metal lumps from the Acfer 097 CR2 chondrite, Meteoritics and Planetary Science, 47, 1191–1208. Iezzi, G., Mollo, S., and Ventura, G. (2009). Solidification behaviour of natural silicate melts and volcanological implications. In N. Lewis and A. Moretti (Eds.), New Research on Volcanoes: Formation, Eruptions and Modeling, 127–151. Hauppauge, NY: Nova Science Publishers, Inc.. Ito, M., and Ganguly, J. (2006). Diffusion kinetics of Cr in olivine and 53Mn–53Cr thermochro- nology of early solar system objects. Geochimica et Cosmochimica Acta, 70,799–809. Jacquet, E., Alard, O., and Gounelle, M. (2012). Chondrule trace element geochemistry at the mineral scale. Meteoritics and Planetary Science, 47, 1695–1714. Jacquet, E., Alard, O., and Gounelle, M. (2015). Trace element geochemistry of ordinary chondrite chondrules: The type I/type II chondrule dichotomy. Geochimica et Cosmochi- mica Acta, 155,47–67. Johnson, B. C., Minton, D. A., Melosh, H. J., and Zuber, M. T. (2015). Impact jetting as the origin of chondrules. Nature, 517, 339–341. Jones, R. H. (1990). Petrology and mineralogy of type II, FeO-rich chondrules in Semarkona (LL3.0): Origin by closed-system fractional crystallization, with evidence for supercool- ing. Geochimica et Cosmochimica Acta, 54, 1785–1802. Jones, R. H. (1992). On the relationship between isolated and chondrule olivine grains in the carbonaceous chondrite ALHA77307. Geochimica et Cosmochimica Acta, 56, 467–482. Jones, R. H. (1996a). FeO-rich, porphyritic pyroxene chondrules in unequilibrated ordinary chondrites. Geochimica et Cosmochimica Acta, 60, 3115–3138. Jones, R. H. (1996b). Relict grains in chondrules: Evidence for chondrule recycling. In R. H. Hewins, R. H. Jones, and E. R. D. Scott (Eds.), Chondrules and the Protoplanetary Disk, 163–172. Cambridge, UK: Cambridge University Press. Jones, R. H. (2012). Petrographic constraints on the diversity of chondrule reservoirs in the protoplanetary disk. Meteoritics and Planetary Science, 47, 1176–1190. Jones, R. H., and Carey, E. R. (2006). Identification of relict forsterite grains in forsterite-rich chondrules from the Mokoia CV3 carbonaceous chondrite. American Mineralogist, 91, 1664–1674. Jones, R. H., and Layne, G. D. (1997). Trace element partitioning between pyroxene and melt in rapidly cooled chondrules. American Mineralogist, 82, 534–545. Jones, R. H., and Lofgren, G. E. (1993). A comparison of FeO-rich, porphyritic olivine chon- drules in unequilibrated chondrites and experimental analogues. Meteoritics, 28,213–221. Jones, R. H., and Scott, E. R. D. (1989). Petrology and thermal history of type IA chondrules in the Semarkona (LL3.0) chondrite. Proceedings of the 19th Lunar and Planetary Science Conference, 523–536. Houston, TX: Lunar and Planetary Institute. Jones, R. H., Saxton, J. M., Lyon, I. C., and Turner, G. (2000). Oxygen isotopic compositions of chondrule olivine and isolated olivine grains in the CO3 chondrite, ALHA77307. Meteori- tics and Planetary Science, 35, 849–857. 86 Rhian H. Jones, Johan Villeneuve, and Guy Libourel

Jurewicz, A. J. G., and Watson, E. B. (1988). Cations in olivine, Part 2: Diffusion in olivine xenocrysts, with applications to petrology and mineral physics. Contributions to Mineral- ogy and Petrology, 99, 186–201. Kennedy, A. K., Lofgren, G. E., and Wasserburg, G. J. (1993). An experimental study of trace element partitioning between olivine, orthopyroxene and melt in chondrules: Equilibrium values and kinetic effects. Earth and Planetary Science Letters, 115, 177–195. Kita, N. T., Nagahara, H., Tachibana, S., et al. (2010). High precision SIMS oxygen three isotope study of chondrules in LL3 chondrites: Role of ambient gas during chondrule formation. Geochimica et Cosmochimica Acta, 74, 6610–6635. Kitamura, M., Yasuda, M., Watanabe, S., and Morimoto, N. (1983). Cooling history of pyroxene chondrules in the Yamato-74191 chondrite (L3) – an electron microscopic study. Earth and Planetary Science Letters, 63, 189–201. Kitamura, M., Watanabe, S., and Morimoto, N. (1986). Pyroxene crystallization in chondrules– Autometamorphic evolution of chondrites. In Antarctic Meteorites XI, Papers presented to the 11th Symposium on Antarctic Meteorites, NIPR, Vol. 11, pp. 71–73. Kunihiro, T., Rubin, A. E., McKeegan, K., and Wasson, J. T. (2004). Oxygen-isotopic compositions of relict and host grains in chondrules in the Yamato 81020 CO3.0 chondrite. Geochimica et Cosmochimica Acta, 68, 3599–3606. Kunihiro, T., Rubin, A. E., and Wasson J. T. (2005). Oxygen isotopic compositions of low-FeO relicts in high-FeO host chondrules in Acfer 094, a type 3.0 carbonaceous chondrite closely related to CM. Geochimica et Cosmochimica Acta, 69, 3831–3840. Lauretta, D. S., Kremser, D. T., and Fegley Jr., B. (1996). The rate of iron sulphide formation in the solar nebula. Icarus, 122, 288–315. Lauretta, D. S., Lodders, K., and Fegley Jr., B. (1997). Experimental simulations of sulphide formation in the solar nebula. Science, 277, 358–360. Libourel, G., and Chaussidon, M. (2011). Oxygen isotopic constraints on the origin of Mg-rich olivines from chondritic meteorites. Earth and Planetary Science Letters, 301,9–21. Lodders, K. (2003). Solar system abundances and condensation temperatures of the elements. Astrophysical Journal, 591, 1220–1247. Lofgren, G. E. (1989). Dynamic crystallization of chondrule melts of porphyritic olivine compos- ition: Textures experimental and natural. Geochimica et Cosmochimica Acta, 53,461–470. Lofgren, G. E. (1996). A dynamic crystallization model for chondrule melts. In R. H. Hewins, R. H. Jones, and E. R. D. Scott (Eds.), Chondrules and the Protoplanetary Disk, 187–196. Cambridge, UK: Cambridge University Press. Lofgren, G. E., and Lanier, A. B. (1990). Dynamic crystallization study of barred olivine chondrules. Geochimica et Cosmochimica Acta, 54, 3537–3551. Lofgren, G. E., and Le, L. (1998). Partial melting of type 1 chondrule precursor aggregates: An experimental and petrographic study. Lunar and Planetary Science Conference #29, contribution #1441. Lofgren, G. E., and Russell, W. J. (1986). Dynamic crystallization of chondrule melts of porphyritic and radial pyroxene composition. Geochimica et Cosmochimica Acta, 50, 1715–1726. Marrocchi, Y., and Libourel, G. (2013). Sulfur and sulfides in chondrules. Geochimica et Cosmochimica Acta, 119, 117–136. Marrocchi, Y., Chaussidon, M., Piani, L., and Libourel, G. (2016). Early scattering of the solar protoplanetary disk recorded in meteoritic chondrules. Science Advances 2, e1601001. McCanta, M. C., Beckett, J. R., and Stolper, E. M. (2008). Zonation of phosphorus in olivine: Dynamic crystallization experiments and a study of chondrule olivine in unequilibrated ordinary chondrites. Lunar and Planetary Science Conference #39, Abstract #1807. McCanta, M. C., Beckett, J. R., and Stolper, E. M. (2016). Correlations and zoning patterns of phosphorus and chromium in olivine from H chondrites and the LL chondrite Semarkona. Meteoritics and Planetary Science, 51, 520–546. Thermal Histories of Chondrules 87

Milman-Barris, M. S., Beckett, J. R., Baker, M. B., et al. (2008). Zoning of phosphorus in igneous olivine. Contributions to Mineralogy and Petrology, 155, 739–765. Miyamoto, M., McKay, D. S., McKay, G. A., and Duke, M. B. (1986). Chemical zoning and homogenization of olivines in ordinary chondrites and implications for thermal histories of chondrules. Journal of Geophysical Research, 91, 12804–12816. Miyamoto, M., Mikouchi, T., and Jones, R. H. (2009). Cooling rates of porphyritic olivine chondrules in the Semarkona (LL3.00) ordinary chondrite: A model for diffusional equilibration of olivine during fractional crystallization. Meteoritics and Planetary Sci- ence, 44, 521–530. Mori, M., Tachibana, S., Piani, L., et al. (2016). Cooling Experiments of Fe-Fes Melts: A Cooling Speedometer of Chondrules. Goldschmidt Abstracts, 2016, 2147. Morris, M. A., and Desch, S. J. (2010). Thermal histories of chondrules in solar nebular shocks. Astrophysical Journal, 722, 1474–1494. Morris, M. A., Boley, A. C., Desch, S. J., and Athanassiadou, T. (2012). Chondrule formation in bow shocks around eccentric planetary embryos. Astrophysical Journal, 752,27–44. Morris, M. A., Weidenschilling, S. J., and Desch, S. J. (2016). The effect of multiple particle sizes on cooling rates of chondrules produced in large-scale shocks in the solar nebula. Meteoritics and Planetary Science, 51, 870–883. Müller, W. F., Weinbruch, S., Walter, R., and Muller-Beneke, G. (1995). Transmission electron microscopy of chondrule minerals in the Allende meteorite: Constraints on the thermal and deformational history of granular olivine-pyroxene chondrules. Planetary and Space Science, 43, 469–483. Nagahara, H. (1981). Evidence for secondary origin of chondrules. Nature, 292, 135–136. Nettles, J. W., Lofgren, G. E., Carlson, W. D., and McSween, H. Y. (2006). Extent of chondrule melting: Evaluation of experimental textures, nominal grain size, and convolution index. Meteoritics and Planetary Science, 41, 1059–1071. Petry, C., Chakraborty, S., and Palme, H. (2004). Experimental determination of Ni diffusion coefficients in olivine and their dependence on temperature, composition, oxygen fugacity, and crystallographic orientation. Geochimica et Cosmochimica Acta, 68, 4179–4188. Piani, L., Marrocchi, Y., Libourel, G., and Tissandier, L. (2016). Magmatic sulphides in the porphyritic chondrules of EH enstatite chondrites. Geochimica et Cosmochimica Acta, 195,84–99. Radomsky, P. M., and Hewins, R. H. (1990). Formation conditions of pyroxene-olivine and magnesian olivine chondrules. Geochimica et Cosmochimica Acta, 54, 3475–3490. Rambaldi, E. R. (1981). Relict grains in chondrules. Nature, 293, 558–561. Righter, K., Campbell, A. J., and Humayun, M. (2005). Diffusion of trace elements in FeNi metal: Applications to zoned metal grains in chondrites. Geochimica et Cosmochimica Acta, 69, 3145–3158. Rocha, S. E., and Jones, R. H. (2012). An experimental study of the conditions of type II chondrule formation in ordinary chondrites. Lunar and Planetary Science Conference #43, Abstract #2595. Roeder, P. L., and Emslie, R. F. (1970). Olivine-liquid equilibrium. Contributions to Mineral- ogy and Petrology, 29, 275–289. Rubin, A. E. (2010). Physical properties of chondrules in different chondrite groups: implica- tions for multiple melting events in dusty environments. Geochimica et Cosmochimica Acta, 74, 4807–4828. Rubin, A. E., Sailer, A. L., and Wasson, J. T. (1999). Troilite in the chondrules of type-3 ordinary chondrites: Implications for chondrule formation. Geochimica et Cosmochimica Acta, 63, 2281–2298. Rubin, A. E., Baecker, B., and Wasson, J. T. (2015). Overgrowth layers on olivine phenocrysts in high-FeO Semarkona chondrules revealed by P, Fe, and Cr X-ray maps: Evidence for multiple melting of chondrules. 78th Annual Meeting of the Meteoritical Society, Abstract #5033. 88 Rhian H. Jones, Johan Villeneuve, and Guy Libourel

Rudraswami, N. G., Ushikubo, T., Nakashima, D., and Kita, N. T. (2011). Oxygen isotope systematics of chondrules in the Allende CV3 chondrite: High precision ion microprobe studies. Geochimica et Cosmochimica Acta, 75, 7596–7611. Ruzicka, A. (2012). Chondrule formation by repeated evaporative melting and condensation in collisional debris clouds around planetesimals. Meteoritics and Planetary Science, 47, 2218–2236. Ruzicka, A., Hiyagon, H., Hutson, M., and Floss, C. (2007). Relict olivine, chondrule recycling, and the evolution of nebular oxygen reservoirs. Earth and Planetary Science Letters, 257, 274–289. Ruzicka, A., Floss, C., and Hutson, M. (2008). Relict olivine grains, chondrule recycling, and implications for the chemical, thermal, and mechanical processing of nebular materials. Geochimica et Cosmochimica Acta, 72, 5530–5557. Ruzicka, A., Floss, C., and Hutson, M. (2012). Agglomeratic olivine (AO) objects and type II chondrules in ordinary chondrites: Accretion and melting of dust to form ferroan chon- drules. Geochimica et Cosmochimica Acta, 76, 103–124. Sanders, I. S., and Taylor, G. J. (2005). Implications of 26Al in nebular dust: Formation of chondrules by disruption of molten planetesimals. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, 915–932. Astronomical Society of the Pacific Conference Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Sanders, I. S., and Scott, E. R. D. (2012). The origin of chondrules and chondrites: Debris from low-velocity impacts between molten planetesimals? Meteoritics and Planetary Science, 47, 2170–2192. Schrader, D. L., and Lauretta, D. S. (2010). High-temperature experimental analogs of primitive meteoric metal-sulfide-oxide assemblages. Geochimica et Cosmochimica Acta, 74, 1719–1733. Schrader, D. L., Connolly, H. C. Jr., Lauretta, D. S., et al. (2013). The formation and alteration of the Renazzo-like carbonaceous chondrites II: Linking O-isotope composition and oxidation state of chondrule olivine. Geochimica et Cosmochimica Acta, 101, 302–327. Schrader, D. L., Connolly, H. C. Jr., Lauretta, D. S., et al. (2015). The formation and alteration of the Renazzo-like carbonaceous chondrites III: Toward understanding the genesis of ferromagnesian chondrules. Meteoritics and Planetary Science, 50,15–50. Schrader, D. L., Davidson, J., and McCoy T. J. (2016a). Widespread evidence for high temperature formation of pentlandite in chondrites. Geochimica et Cosmochimica Acta, 189, 359–376. Schrader, D. L., Fu, R. R., and Desch, S. J. (2016b). Evaluating chondrule formation models and the protoplanetary disk background temperature with low-temperature, subsilicate solidus chondrule cooling rates. Lunar and Planetary Science Conference #47, Abstract #1180. Smyth, J. R. (1974). Experimental study on the polymorphism of enstatite. American Mineral- ogist, 59, 345–352. Sorby, H. (1877). On the structure and origin of meteorites. Nature, 15, 495–498. Soulié, C. (2014). Formation des chondres et relation avec leurs auréoles de matrice à grains fins. PhD thesis, Université de Lorraine. Soulié, C., Libourel, G., and Tissandier, L. (2017). Olivine dissolution in molten silicates: An experimental study with application to chondrule formation. Meteoritics and Planetary Science, 52, 225–250. Symes, S. J., and Lofgren, G. E. (1999). Distribution of FeO and MgO between olivine and melt in natural and experimental chondrules. Lunar and Planetary Science Conference #30, Abstract #1869. Tachibana, S., and Huss, G. R. (2005). Sulfur isotope composition of putative primary troilite in chondrules from Bishunpur and Semarkona. Geochimica et Cosmochimica Acta, 69, 3075–3097. Thermal Histories of Chondrules 89

Tachibana, S., Nagahara, H., and Mizuno, K. (2006). Constraints on cooling rates of chondrule from metal-troilite assemblages. Lunar and Planetary Science Conference #37, Abstract #2263. Tachibana, S., Tamada, S., Kawasaki, H., Ozawa, K. and Nagahara H. (2013). Interdiffusion of Mg-Fe in olivine at 1400–1600 C and 1 atm total pressure. Physics and Chemistry of Minerals, 40, 511–519. Taylor, L. A., and Cirlin, E. H. (1986). Olivine/melt Fe/Mg KD’s <0.3: Rapid cooling of olivine-rich chondrules. Lunar and Planetary Science, XVII,19–38. Tenner, T. J., Ushikubo, T., Kurahashi, E., Kita, N. T., and Nagahara, H. (2013). Oxygen isotope systematics of chondrule phenocrysts from the CO3.0 chondrite Yamato 81020: Evidence for two distinct oxygen isotope reservoirs. Geochimica et Cosmochimica Acta, 102, 226–245. Tenner, T. J., Nakashima, D., Ushikubo, T., Kita, N. T., and Weisberg, M. K. (2015). Oxygen isotope ratios of FeO-poor chondrules in CR3 chondrites: Influence of dust enrichment and H2O during chondrule formation. Geochimica et Cosmochimica Acta, 148, 228–250. Tronche, E. J., Hewins, R. H., and MacPherson, G. J. (2007). Formation conditions of aluminum-rich chondrules. Geochimica et Cosmochimica Acta, 71, 3361–3381. Tsuchiyama, A., Osada, Y., Nakano, T., and Uesugi, K. (2004). Experimental reproduction of classic barred olivine chondrules: Open-system behavior of chondrule formation. Geochi- mica et Cosmochimica Acta, 68, 653–672. Ushikubo, T., Kimura, M., Kita, N. T., and Valley J. W. (2012). Primordial oxygen isotope reservoirs of the solar nebula recorded in chondrules in Acfer 094 carbonaceous chondrite. Geochimica et Cosmochimica Acta, 76, 242–263. Ushikubo, T., Nakashima, D., Kimura, M., Tenner, T. J., and Kita, N. T. (2013). Contempor- aneous formation of chondrules in distinct oxygen isotope reservoirs. Geochimica et Cosmochimica Acta, 109, 280–295. Villeneuve, J., Libourel, G., and Soulié, C. (2015). Relationship between type I and type II chondrules: Implications on chondrule formation processes, Geochimica et Cosmochimica Acta, 160, 277–305. Wasson, J. T., and Rubin, A. E. (2003). Ubiquitous low-FeO relict grains in type II chondrules and limited overgrowths on phenocrysts following the final melting event. Geochimica et Cosmochimica Acta, 67, 2239–2250. Wasson, J. T., Baecker, B., and Rubin, A. E. (2014). Multiple, hierarchical heating of chon- drules and implications for cooling rates. Lunar and Planetary Science Conference #45, Abstract #2883. Watanabe, S., Kitamura, M., and Morimoto, N. (1985). A transmission electron microscope study of pyroxene chondrules in equilibrated L-group chondrites. Earth and Planetary Science Letters, 72,87–98. Watanabe, S., Kitamura, M., and Morimoto, N. (1986). Oscillatory zoning of pyroxenes in ALH- 77214 (L3). Papers Presented to the Eleventh Symposium on Antarctic Meteorites,74–75. Weinbruch, S., and Müller, W. F. (1995). Constraints on the cooling rates of chondrules from the microstructure of clinopyroxene and plagioclase. Geochimica et Cosmochimica Acta, 59, 3221–3230. Weinbruch, S., Müller, W. F., and Hewins, R. H. (2001). A transmission electron microscope study of exsolution and coarsening in iron-bearing clinopyroxene from synthetic analogues of chondrules. Meteoritics and Planetary Science, 36, 1237–1248. Weisberg, M. K., and Prinz, M. (1996). Agglomeratic chondrules, chondrule precursors, and incomplete melting. In R. H. Hewins, R. H. Jones, and E. R. D. Scott (Eds.), Chondrules and the Protoplanetary Disk, 119–127. Cambridge, UK: Cambridge University Press. Welsch, B., Faure, F., Famin, V., Baronnet, A., and Bachèlery, P. (2013). Dendritic crystalliza- tion: A single process for all the textures of olivine in basalts? Journal of Petrology, 54, 539–574. 90 Rhian H. Jones, Johan Villeneuve, and Guy Libourel

Wick, M., and Jones, R. H. (2012). Formation conditions of plagioclase-bearing type I chondrules in CO chondrites: A study of natural samples and experimental analogs. Geochimica et Cosmochimica Acta, 98, 140–159. Yasuda, M., Kitamura, M., and Morimoto, N. (1983). Electron microscopy of clinoenstatite from a boninite and a chondrite. Physics and Chemistry of Minerals, 9, 192–196. Yurimoto, H., and Wasson, J. T. (2002). Extremely rapid cooling of a carbonaceous-chondrite chondrule containing very O-16-rich olivine and a Mg-26-excess. Geochimica et Cosmo- chimica Acta, 66, 4355–4363. 4 Composition of Chondrules and Matrix and Their Complementary Relationship in Chondrites

dominik c. hezel, phil a. bland, herbert palme, emmanuel jacquet, and john bigolski

Abstract

Complementary chemical and isotopic relationships between chondrules and matrix have the potential to distinguish between categories of chondrule forming mechanisms, e.g., exclude all mechanisms that require different reservoirs for chondrules and matrix. The complementarity argument is, however, often misunderstood. Complementarity requires different average com- positions of an element or isotope ratio in each of the two major chondrite components chondrules and matrix, and a solar or CI chondritic bulk chondrite ratio of the considered elements or isotopes. For example, chondrules in carbonaceous chondrites typically have superchondritic Mg/Si ratios, while the matrix is subchondritic. Another example would be the Hf/W ratio, which is superchondritic in chondrules and subchondritic in matrix. We regard these ratios to be complementary in chondrules and matrix, because the bulk chondrite has solar Mg/Si and Hf/W ratios. In contrast, Al/Na ratios are also different in chondrules and matrix, but the bulk is not solar; therefore, Al/Na does not have a complementary relationship. A number of publications over the past decade have reported complementary relationships for many element pairs in different types of chondrites. Recently, isotopic complementarities have also been reported. A related, though different, argument can be made for volatile depletion patterns in chondrules and matrix, which can then also be considered as being complementary. The various models for chondrule formation require either that chondrules and matrix formed from a single (i.e., common) parental reservoir, or that chondrules and matrix formed in separate regions of the protoplanetary disk and were later mixed together. As chondrules and matrix have different compositions, mixing of these two components would result in a random bulk chondrite composition. The observation of complementary chondrule–matrix relationships together with a CI chondritic, element or isotope ratio is unlikely to be the result of a random mix of chondrules and matrix. It seems much more likely that chondrules and matrix formed in a single reservoir with initially CI chondritic element or isotope ratios. Incorporation of different minerals in chondrules and matrix together with volatile element depletion of the entire reservoir then resulted in chondrule-matrix complementarities and bulk chondrite volatile depletion. This excludes any chondrule formation mechanism that requires separate parental reservoirs for these components. Any chondrule forming mechanism must explain complementarity. Chondrules and matrix must have formed from a common reservoir.

91 92 Dominik C. Hezel, et al. 4.1 Introduction

Undifferentiated meteorites or chondrites are extraterrestrial materials that never experienced melting or differentiation and thus sampled primitive solar system material. The major element compositions of chondrites are close to the composition of the Sun for heavy, nonvolatile elements (e.g., Lodders, Palme, and Gail, 2009). Only one group of chondrites, the CI chondrites, fit within analytical uncertainties the solar photospheric abundances. Other types of chondrites deviate slightly from solar abundances, as described in more detail in Section 4.2. Chondrules are a major structural component of chondrites. These typically submillimeter-sized spherules were once partly or entirely molten at temperatures up to 2,000 K, and developed igneous textures during rapid cooling (Hewins, 1997). A second major component of chondrites is fine-grained matrix (typically <5 µm). Chondrules are dominated by high-temperature components, while matrix is a mixture of high-temperature and low-temperature phases. The protoplanetary disk was a highly dynamic system that stretched across tens of astro- nomical units. One category of chondrule formation models suggests that chondrules and matrix formed at different locations of the disk and one or both components were then transported to a common location where the parent body accreted. This category of chondrule formation model is typically referred to as ‘two-component models’ (e.g., Anders, 1964; Shu, Shang, and Lee, 1996; Zanda et al., 2006; Asphaug and Jutzi, 2011; Van Kooten et al., 2016). In contrast, models of a second category do not require transport and mixing of chondrules and matrix. For these models, both components formed in the same location and from a single chemical reservoir (e.g., Desch and Connolly, 2002; Morris and Desch, 2010; McNally et al., 2013). If it were possible to discriminate which of these two categories of models is correct, it would allow us to discard a large number of models, and focus on only the remaining category. Physical parameters that we extract from studying chondrules and matrix, such as tempera- ture, pressure, or chronology, contain no direct information on the location of chondrule or matrix formation and cannot be used to discriminate between the two categories of models. The combined suite of information about all individual components (chondrules, matrix, Ca-Al-rich inclusions (CAIs), opaque phases, etc.) in a chondrite, however, can provide insights on whether these components formed in a single or in different chemical reservoirs. For example, if the chondritic components formed in different reservoirs and were then randomly mixed into the parent body, bulk CI chondritic element or isotope ratios would be very unlikely. If the components formed in a single reservoir, a genetic (i.e., compositional) relationship between the components should be observed. Such a genetic relationship should in particular be present between chondrules and matrix, as these are the two most volumetrically abundant components (>95 vol.%). The relationship is best studied in the chondrites in which the relative abundances of chondrules and matrix are not very different (e.g., carbonaceous or Rumuruti chondrites). It has been known for a long time that bulk chondrule compositions are different from the matrix composition, chemically and isotopically (e.g., Clayton et al., 1991; Hezel et al., 2018a and references therein). Figure 4.1 illustrates how the combination of chondrules and matrix add up to the bulk chondrite composition, and what is, and what is not, a ‘complementary relationship’. To understand Figure 4.1 and complementarity, it is pivotal to note that: (i) in all cases, only element pairs are considered. Therefore, when talking about a chondrule, matrix, or Chondrule – Matrix Complementarity 93

(a) (b)

complementarity yes complementarity no chd chd bulk chondrite ratio Cl ratio

mixing line mixing line Cl ratio & bulk chondrite ratio bulk chondrite ratio bulk chondrite ratio ≠ Cl ratio element 2 element 2 = Cl ratio Cl mtx mtx Cl element 1 element 1

(c) (d)

complementarity yes complementarity no Cl Cl bulk chondrite = Cl bulk chondrite ≠ Cl

chd sample chd sample

mtx mtx

isotope isotope

Figure 4.1 Schematic illustration of the elemental (a) and isotopic (c) chondrule matrix complementarity: Chondrules and matrix need to be different, and the bulk chondrite needs to be CI chondritic. Although chondrules and matrix still satisfy mass balance if the bulk chondrite is not CI chondritic, this relationship is not called complementary (b,d). bulk chondrite composition, in almost all cases element ratios in chondrules, matrix, or the bulk chondrite are meant. This applies throughout the entire text. The absolute abundance of each element will typically not match the absolute CI abundances. This usage of the term ‘compos- ition’ is more obvious when talking about an isotopic composition of chondrules, or the bulk, which in this case also always means ratios of two isotopes. (ii) The contribution of CAIs and opaque phases to the bulk chondrite composition is negligible for most elements. Figures 4.1b and 4.1d illustrate how a bulk chondrite must have a random intermediate composition between chondrules and matrix when these have different compositions. Little can be learned from this trivial relationship. However, in Figures 4.1a and 4.1c, the bulk chondrite compositions are not random values between chondrules and matrix, but coincide with a unique value, which is the solar (i.e., CI chondritic) composition. Only in these cases is there a meaningful relationship that includes chondrules, matrix, and the bulk chondrite composition. Only these cases (i.e., Figures 4.1a and 4.1c) are designated as complementary relationships. These complementary relationships then imply the following: rather than having meteorite grains randomly mixed to coincidentally arrive at a CI chondritic element ratio, it is much more likely that all grains formed in a single location with initially CI chondritic element ratios, and the different compositions of chondrules and matrix were established during the formation of 94 Dominik C. Hezel, et al. these components. Complementarity, then, refers only to the observation that, although chon- drules and matrix have different element ratios, they sum up to a CI chondritic bulk (Figure 4.1). Therefore, any element pair (e.g., Si/Na or Ca/Mn) that has different ratios in chondrules and matrix, but a nonchondritic bulk chondrite ratio, is not complementary. Evidence for chondrule– matrix complementarity is strengthened if observed for (i) multiple element ratios, (ii) multiple meteorites from different groups, and (iii) not only chemical, but also isotopic ratios. A related, though different, argument for complementarity is based on volatile element depletion patterns in chondrules, matrix and the bulk chondrite. This will be discussed in more detail at the end of Section 4.5. Thus, complementarity becomes a criterion to discriminate between the two categories of chondrule formation models described above, and qualifies as one of the most critical con- straints to discriminate between these categories of models. In this chapter, we review the chemical, physical, and isotopic characteristics of bulk chondrites, their chondrules, their matrices in carbonaceous chondrites, and their genetic relationships and complementarities. We then discuss constraints that chondrule formation models must satisfy.

4.2 Bulk Compositions of Chondrites

The abundances of rock-forming elements in chondrites are similar to those in the Sun’s photosphere. Despite this similarity, there are large differences in petrography and mineralogy of chondrites resulting from thermal metamorphism and aqueous alteration on their parent bodies. These parent body processes should be distinguished from processes which occurred before the final assemblage of a parent body. To some extent, chondrites show bulk compos- itional variabilities, the most important of which are summarized in Figure 4.2. All data in this figure are normalized to Mg and the CI-composition. On the left side, the photospheric abundances are plotted. The dark band indicates the uncertainties in the photospheric abun- dances (approximately 20 percent), as listed by Palme et al. (2014a). The Si/Mg ratios of all chondrite groups are within the 20-percent band, except EH-chondrites. The Fe and refractory element (Al, Ca) variations are larger, but fit best with CI chondrites. Al and Ca are representa- tive of all refractory lithophile elements; the Fe fractionation reflects the behavior of siderophile elements. The volatile element Zn shows a good agreement with solar data only for CI chondrites. Zn is representative of many volatiles, most importantly S, which shows exactly the same behavior as Zn. The abundances of major and minor elements of the various chondrite groups are known within a few relative percent. For example, the average carbonaceous chondrite Si/Mg weight ratio of 22 samples analyzed by Wiik (1956) using wet chemical procedures is 1.102 0.034 (see compilation by Ahrens, 1965). Later, independent wet chemical analyses by Jarosewich (1990) gave 1.109 0.036 (19 samples), and Wolf and Palme (2001) found 1.102 0.025 (22 samples) using X-ray fluorescence (XRF). In these data sets, there is a small, barely resolvable decrease in the Si/Mg ratio of 3 to 4 relative percent, following the sequence of CI, CM, and CV chondrites. Allende (CV) bulk samples have, for example, an average Si/Mg ratio of 1.080 0.061 (Stracke et al., 2012) slightly below the CI-ratio of 1.12 0.05 Chondrule – Matrix Complementarity 95

2.0

1.8 Fe 1.6

1.4 Ca, Al Si 1.2

1.0

0.8

0.6 CI- and Mg-normalized Zn 0.4

0.2

0.0 SunCICMCVCOCKCRCHH L LLEHELR

Figure 4.2 Differences and similarities among the solar photosphere (Sun) and bulk chondrite element compositions. Data taken from: Wasson and Kallemeyn (1988); Wolf and Palme (2001). Red: Ca; Green: Al; Blue: Si; Red: Fe. (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.)

(Lodders et al., 2009) and in accord with a Si/Mg ratio of 1.077 for an average of a 4 kg Allende sample (Clarke et al., 1971). The solar photospheric Si/Mg ratio of 1.10 0.2 fits with the carbonaceous chondrite ratios, but has larger uncertainties (Lodders et al., 2009). All these data show that the bulk composition of carbonaceous chondrites is well defined, even if a group deviates slightly from CI values. In addition, Stracke et al. (2012) found that 0.5 g samples of Allende are sufficient to obtain chemically reproducible results. For Al, a few grams may be necessary, because of inhomogeneous distributions of CAIs. Since Allende is a relatively coarse-grained rock, the chemical homogeneity should be similar or even better for other groups of chondrites, as demonstrated by Jarosewich (1990), and more recently using X-ray powder diffraction for CV (Howard et al., 2011), CR, CM, and CI (King et al., 2015). Chondrites are not only characterized by approximately solar major element compositions, but they also have characteristic trace element patterns. The refractory trace elements are in most cases unfractionated. This implies flat CI-normalized rare earth element (REE) patterns and unfractionated highly siderophile element patterns (Ir, Os, Re, and so on). Also, the ratio of refractory lithophile to refractory siderophile elements is essentially constant among carbon- aceous chondrites. Minor variations in refractory element patterns may have been introduced by exotic components. Bulk Allende has a REE pattern influenced by the presence of fine-grained spinel-rich inclusions with so called group II REE patterns (see Stracke et al., 2012). Since bulk Fe contents are somewhat variable in chondrites, other siderophile elements, including refrac- tory siderophiles, will vary with bulk Fe, which affects the ratio of lithophile to siderophile elements. Volatile elements show particularly large variations in chondritic meteorites (Figure 4.3). They are in all cases depleted, never enriched relative to CI chondrites. Furthermore, their depletion is not dependent on the geochemical behavior; whether the element is siderophile, 96

Figure 4.3 Bulk chondrule, matrix, and chondrite element compositions of four different carbonaceous chondrites. Chondrules typically have superchondritic refractory and subchondritic volatile element concentrations. The opposite is observed in matrix. Figure taken from Hezel et al. (2018a). (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.) Chondrule – Matrix Complementarity 97 chalcophile, or lithophile, the depletion depends essentially on the condensation temperatures of the elements (Palme et al., 1988; Figure 4.3). In summary, element pairs with comparable cosmochemical behavior (i.e., 50% conden- sation temperatures) often have CI chondritic ratios in bulk carbonaceous chondrites, although the individual elements may have different geochemical behaviors. It should, however, be noted that element pairs with different cosmochemical behaviors can also have CI chondritic ratios. These observations are fundamental for complementarity (see Figure 4.1 and Sections 4.1 and 4.5). The processes responsible for bulk element fractionations (Figures 4.2 and 4.3) occurred in the pre-accretionary environment of the protoplanetary disk. These include the addition or loss of a refractory component and early-formed forsterite, metal addition or loss, and incomplete condensation of volatiles (e.g., Palme, Spettel, and Hezel, 2014b; see also Chapter 7).

4.3 Bulk Chondrule Characteristics of Chondrule Populations in Individual Chondrites

The chondrule population of each chondrite group has a unique and distinctive set of character- istics in terms of size, textural variation, and chemical as well as isotopic composition (Jones, 2012; Friedrich et al., 2015; Hezel et al., 2018a). Size distributions are typically unimodal and log-normal (cf. Friedrich et al., 2015; Friend et al., 2017). In a recent review, Hezel et al. (2018a) showed that chondrules in an individual chondrite have a significant spread in their bulk element compositions, while matrices are comparatively homogeneous. The bulk chondrule element distribution of chondrule populations is typically unimodal, rarely multimodal, and normal to log-normal (Hezel and Palme, 2007). Chondrule populations of individual chondrites also vary in their isotopic compositions. We summarize bulk chondrule stable isotope data in Table 4.1 and Figure 4.4. The number of reported data for a particular isotope ranges from only two chondrules up to several tens, or even hundreds of chondrules in the case of O isotopes. Still, the number of isotope data from individual chondrules or chondrule separates is limited. Reported relative deviations of chon- drules from standards for the light isotopes up to 18O (and including 41K) range between about –25 to +20‰, and between –134 to +164‰ in the case of 15N. Reported values for all other stable isotope systems are significantly lower, typically ranging between –2 and +1‰, with two chondrules reaching almost –4‰ (88Sr and 114Cd). Isotopes that plot on mass fractionation trends are most likely the result of mass dependent processes (solid lines in Figure 4.4), while isotopes that plot off their mass fractionation trends are the result of mixing of nucleosynthetic components (mass independent fractionation (not in the case of O); dashed lines in Figure 4.4). Only a few bulk chondrule isotope data exist for individual chondrite chondrule populations, and it is therefore difficult to determine reliable chondrule population distribution parameters. Chondrule populations of different chondrites, even within the same group (e.g., CV, CR, L, EH, and so on), are, however, not similar, but can differ either in their elemental and isotopic compositional spread, mean composition, or size distribution (see Jones, 2012; Friedrich et al., 2015; Hezel et al., 2018a). 98 Table 4.1 Isotope compositions of chondrules.

Isotope Variation Samples n Process Reference

δ7/6 Li 8.5 to +10‰ Allende, Semarkona, Bishunpur, 89 precursor variations, Seitz et al. (2012) Bremervörde, Bjurböle, Saratov (nucelosynthetic origin) δ7/6 Li +7.8, +10.7, +15.1‰ Px-Chd from Semarkona 3 precursor variations, Chaussidon and Robert (nucelosynthetic origin) (1998) δ18/16 O 25, +12‰ Krot et al. (2006) δ11/10 B 12.2, 10.9, 7.0‰ Px-Chd from Semarkona 3 precursor variations, Chaussidon and Robert (nucelosynthetic origin) (1998) 15/14 δ Nt 134 to 164‰ 6 Ordinary, 2 Carbonaceous, 68 precursor variations Das and Murty (2008) 2 Enstatite δ25/24 Mg 0.65 to +0.41‰ Murchison, Murray 23 parent body alteration and Bouvier et al. (2013) evaporation/recondensation during chd formation δ25/24 Mg +1.57 to +2.04‰ Allende 8 heterogeneous Galy et al. (2000) starting material δ25/24 Mg 0.32 to +0.80‰ Allende 18 Bizzarro et al. (2004) δ25/24 Mg 0.28 to +0.20 ppm; 0.09 to 2 CV (19), 3 CR (23) 42 Olsen et al. (2016) +0.28 ppm δ29/28 Si 0.06 to 0.17‰; (+0.58‰) Mokoia, Grosnaja 7 only one chd has a different Hezel et al. (2010) value δ29/28 Si 0.40 to 0.10‰ Murchison, Bjurböle 3 mineral-mineral fractionation Molini-Velsko et al. during chd formation (1986) δ29/28 Si 0.69 to +0.47‰ Parsa, Bjurböle, Chainpur, 42 Clayton et al. (1991) Dhajala, Weston ε34/32 S 0.068 to +0.089‰ Allende, Dhajala, EET99404, 5 Photochemical Irradiation Rai and Thiemens ALH85033 (2007) δ41/39 K 9.5 to +17.8‰ Semarkona 28 exchange with ambient gas Alexander and during cooling Grossman (2005) δ44/40 Ca +0.29 to +0.50‰ Allende 7 Amsellem et al. (2017) ε50/46 Ti 2.0 to +9.0 Allende, Kaidun, Murchison, 11 heterogeneities in the nebula Niemeyer (1988) Kakangari dust ε50/46 Ti 1.25 to +1.73 Allende, Dar al Gani 6 correlation 46Ti–50Ti, thermal Trinquier et al. (2009) processing μ54/52 Cr +1 to +201ppm; +121 to +172 2 CV (19), 3 CR (23) 42 different chd reservoirs Olsen et al. (2016) ppm ε54/52 Cr 0.48 to 0.08 Chainpur, Allende, Qingzhen 4 Trinquier et al. (2007) ε54/52 Cr 0.87 to 0.24 5 Connelly et al. (2012) δ56/54 Fe 1.33 to +0.65‰ Allende (9), Chainpur (3) 12 parent body alteration and Mullane et al. (2005) evaporation/recondensation during chd formation δ56/54 Fe 0.82 to +0.37‰ Allende (12), Mokoia (16), 34 evaporation/recondensation Hezel et al. (2010) Grosnaja (6) δ66/64(?) Zn +0.42‰ Mocs 1 Luck et al. (2005) δ66/64 Zn 0.45 to +0.18‰ Allende, Mokoia 9 Pringle et al. (2017) δ88/86 Sr 1.67 to +0.58‰ Allende 5 parent body alteration and Moynier et al. (2010) evaporation δ88/86 Sr 2.9 to 0.3‰ Allende (8), Richardton (1) 9 kinetic evaporation Patchett (1980) ε92/96 Mo +2.3 to +7.53‰ Allende 6 heterogeneous precursor Budde et al. (2016a) δ114/110 Cd 3.9 to 0.5‰ Allende 3 Wombacher (2008) ε183/184(6/4) W +0.05 to +2.05‰ Allende, Vigarano 7 heterogeneous precursor Becker et al. (2015) ε183/184(6/4) W +1.39 to +2.26‰ Allende 6 heterogeneous precursor Budde et al. (2016b) ε135/136(?) Ba +0.18 to +0.30‰ Allende 5 homogeneous precursor Budde et al. (2016a) 238/234 U 137.782 to 137.792 Allende 6 Brennecka et al. (2015)

n: number of chondrules for which bulk chondrule isotope compositions have been reported. 99 100 Dominik C. Hezel, et al.

–4 –3 –2 –1 0 1 2 3

135Ba dominant fractionation type 183 mass independent W mass dependent 92Mo 54Cr

50Ti 34S 114Cd 88Sr 66Zn 56Fe ‰ from a standard (d) 44Ca 29Si

25 Mg different Mg std 18O 41K 11 x10‰ from a standard (10d) B 7Li 15 x100‰ from a standard (100d) N –4 –3 –2 –1 0 1 2 3

Figure 4.4 Displayed are variations in bulk chondrule isotope compositions. Solid lines represent mass- dependent variations, and dashed lines nucleosynthetic variations. Data taken from the literature presented in Table 4.1.

4.4 Mineralogy of Chondrite Matrices

Although carbonaceous chondrites show significant group-to-group differences in matrix min- eralogy, the majority of those differences (e.g., matrix dominated by phyllosilicates in CMs and CRs, fayalitic olivine in CVs and COs) arguably arise from parent body processes (Krot et al., 1997a). Did different groups begin with a common matrix precursor mineralogy prior to parent body processing? The evidence suggests that they did. The most mineralogically primitive matrix is found in a small number of highly unequi- librated carbonaceous chondrites that appear to have experienced minimal parent body process- ing (e.g., Acfer 094 and ALH 77307). It is a mixture comprised of 10s–100s nm-sized materials, including forsteritic olivine, enstatic low-Ca pyroxene, amorphous silicates, Fe,Ni-sulphides, feldspar, metal with widely varying Ni contents, organics, minor element alloys (e.g., Cr-Fe alloy, and refractory metal nuggets), µm-scale CAIs, and presolar grains (Brearley, 1993; Greshake, 1997; Grossman and Brearley, 2005; Bland et al., 2007). Greshake (1997) concluded that the amorphous material formed by disequilibrium condensation. Crystalline phases were suggested to have formed either by condensation in the solar nebula, or by recrystallization from the amorphous material (Greshake, 1997). Nuth et al. (2005) also noted the primitive, unaltered nature of Acfer 094 and compared its matrix to anhydrous interplanetary dust particles (IDPs). Chondrule – Matrix Complementarity 101

Greshake (1997) and Nuth et al. (2005) proposed transient heating events (e.g., matrix proximal to a region experiencing chondrule formation) as a mechanism for producing amorphous solids and at least some of the crystalline material observed in Acfer 094 matrix. Bland et al. (2007) showed that a significant fraction of crystalline material in Acfer 094 has mineralogies consist- ent with equilibrium condensation modeling. This meteorite also contains an unusually high concentration of presolar grains, including presolar silicates (e.g., Nguyen and Zinner, 2004). Refractory metal nuggets (RMNs) are a trace component in chondrites. Some RMNs show evidence of condensation (e.g., Berg et al., 2009). In situ studies of RMN crystallography and chemistry (Daly et al., 2017a, 2017b), however, reveal a complex preaccretionary history for these materials with multiple reheating events. Although the matrices of most CR chondrites have been aqueously altered, mostly to phyllosilicates (serpentine), in addition to magnetite and minor calcite (Weisberg et al., 1993), rare examples (MET 00426 and QUE 99177) appear to represent comparatively unaltered end-members (Abreu and Brearley, 2010; Harju et al., 2014). In these meteorites, matrices are dominated by abundant amorphous Fe-rich silicate material, containing nanopar- ticles of Fe,Ni sulfides quite similar to Acfer 094 and ALH 77307 (Abreu and Brearley, 2010). Similar material has been observed in the relatively unaltered CM chondrite Paris (Leroux et al., 2015). The fact that the least altered CMs and CRs show similarities in matrix mineralogy with the most primitive carbonaceous chondrites suggests that all chondrites had similar matrix mineralogies prior to parent body alteration. Presolar grains predate any solar system processes. The survival of presolar grains indicates that subsequent processing at the midplane was inefficient, or that some small fraction of unprocessed material was added prior to accretion on the parent body. Hence, while all matrix precursors may have had a similar mineralogy in the various groups of chondritic meteorites, there is evidence that the compositions of matrices in the different groups of chondrites are to some extent compositionally and isotopically distinct. The most primitive members of a chondrite group are a highly variable and unequilibrated mixture of phases, apparently sourcing material that had formed or been processed in a variety of interstellar, circumstellar, and protoplanetary disk environments. Alteration produced the differences in matrix mineralogies presently observed in less primitive chondrites. A significant fraction of primitive matrices (e.g., amorphous silicate material and nanosulfide particles) may have formed in the solar nebula by rapid condensation following transient high-temperature events, such as during chondrule formation (e.g., Brearley, 1993; Greshake, 1997; Abreu and Brearley, 2010). Some appear to have formed by equilibrium condensation in the disk (Bland et al., 2007). Matrix was present during chondrule formation, but only a portion of the matrix might have experienced the high-temperature event that chondrules experienced.

4.5 The Chondrule-Matrix Complementarity

Bulk chondrites, in particular carbonaceous chondrites, have close to CI-chondritic (i.e., solar) compositions for a number of element and isotope ratios (see Section 4.2). Chondrules and matrix often have nonchondritic compositions for these element ratios and, in some cases, also for isotope ratios. We first consider element ratios. 102 Dominik C. Hezel, et al.

Figure 4.5 displays a number of element pairs for which bulk chondrites have solar ratios, as required for complementarity. Chondrules and matrix, however, have either super- or subchon- dritic element ratios. Their compositions are complementary relative to the CI chondritic bulk chondrite. As chondrules and matrix add up to typically >95 vol.% of the meteorite, it should in principle be possible to calculate the bulk chondrite composition from the modal abundances of chondrules and matrix, their densities and element concentrations. For example, the average Mg concentrations of Renazzo chondrules and matrix are 23.5 1 and 9 0.5 wt%; the average Si concentrations of Renazzo chondrules and matrix are 22.8 0.5 and 15.5 0.5 wt%; the chondrule and matrix densities are about 3.2 0.2 and 2.7 0.2 g/cm3; and chondrule and matrix modal abundances are 55 5 and 40 10 vol.% (Hezel et al., 2010 and references therein). This results in calculated bulk chondrite Mg/Si ratios between 0.83 and 0.99. The bulk Renazzo Mg/Si ratio of 0.91 (Mason and Wiik, 1962) is within this range. However, the large range of calculated Mg/Si ratios between 0.83 and 0.99 illustrates that sensible mass balance calculations are not possible, and complementary relationships need to be identified by measurements of chondrite components and bulk compositions. Complementarity cannot be tested using mass balance calculations, as attempted by Zanda et al. (Chapter 5), because of the large uncertainties in component modal abundances, as well as in chondrule and matrix compositions. Zanda et al. (Chapter 5) suggested that the complementary relationship of chondrules and matrix maybe an analytical artifact. This is immediately ruled out, as complementary relation- ships have been identified using various techniques: for example, the Mg/Si data of Figure 4.5a and 4.5b were obtained both with electron microprobe, but with different methods of data reduction (Hezel et al., 2010; Ebel et al., 2016); the Fe/Mg chondrule data of Figure 4.5c were obtained with Instrumental Neutron Activation Analysis (Jones and Schilk, 2009) and matrix data with the electron microprobe (Palme et al., 2015); all W/Hf data of Figure 4.5d were obtained using solution based mass spectrometry (Becker et al., 2015). Additional examples presented in Hezel et al. (2018a) unequivocally further demonstrate that complementary relationships are a typical characteristic of chondrules and matrix, and not an analytical artifact. A striking example for complementarity is the Fe/Mg ratio in components of CM chondrites: the fraction of chondrules in CM chondrites ranges from 20 to 40 vol.%, most of them are type I, FeO-poor chondrules. For example, the average FeO content of 11 chondrules in the CM meteorite El-Quss Abu Said is 2 wt% (including a conservative estimate of ~4 vol.%, (cf. Hezel et al., 2018b), opaque phases in chondrules, this concentration might rise to ~5 wt% FeO), typical of many CM chondrites (e.g., Hezel and Palme, 2010; Friend et al., 2018). Thus, a large fraction of a CM meteorite is almost free of Fe, but has high Mg contents (av. 43 wt% MgO in El-Quss Abu Said). The Fe/Mg ratio in bulk carbonaceous chondrites is around 1.5 to 2 (CI has an Fe/Mg ratio of 1.96), somewhat variable but much higher than the ratio of 0.06 (or 0.15, including opaques in chondrules) for the El-Quss Abu Said chondrules. A fraction of 20 vol.% chondrules in an CM chondrite would deliver half of the Mg of the bulk chondrite, but only a very small fraction of Fe. Thus, there must be a major fraction of the meteorite with high Fe and low Mg to compensate for the low Fe and high Mg in chondrules (see also Wood, 1967). This is matrix for most carbonaceous chondrites plus, in CM chondrites, a fine-grained alteration phase formerly designated PCP (poorly characterized phase), now often named TCI (tochelinite–cronstedtite intergrowth). As low-Fe olivine cannot be the result of metamorphic 103 Figure 4.5 Examples of element chondrule-matrix complementarities. Data taken from: a) Hezel et al., 2008; c) Palme et al., 2015; d) Becker et al., 2015; Budde et al., 2016a. Figure b) taken from Ebel et al., 2016. (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.) 104 Dominik C. Hezel, et al. redistribution, the conclusion is that neither chondrules nor matrix in CM chondrites have a composition that is, even in a broad sense, chondritic. In the case of El-Quss Abu Said, matrix has an Fe/Mg ratio of 3.3, compared to 0.06 (or 0.15, respectively) for chondrules and about 1.7 for the bulk meteorite. Both components, chondrules and matrix, must have formed from a single reservoir with chondritic composition. One may argue that type II, FeO-rich chondrules were initially much more abundant, and because FeO-rich chondrules are more susceptible to weathering they have not survived and ended up as matrix. This possibility is very unlikely. The Paris meteorite, the least altered CM chondrite, has an unusually high fraction of chondrules (46 vol.%), with only 1–2 vol.% FeO-rich, type II chondrules (Hewins et al., 2014). Even if there were initially another 20 vol.% of FeO-rich chondrules, the bulk Fe/Mg-ratio of the total chondrule fraction would still be far below the bulk meteorite ratio. Also, there are no chondrules with FeO contents between FeO-poor (type I) and FeO-rich (type II), which one would expect if type II chondrules had lost some of their iron to matrix. In addition, there is no correlation between the fraction of type II chondrules and the degree of weathering. The low fraction of type II chondrules is typical of CM chondrites, irrespective of the degree of alteration. The matrix composition of Paris has superchondritic Fe/Mg ratios, as required by mass balance and as determined by analyses of less altered and more altered zones in matrix (Hewins et al., 2014). In addition, it would be unexpected if in the least altered CM chondrite containing a few percent of metallic Fe, a large fraction of FeO-rich chondrules had lost their iron to matrix, given that there is little evidence of alteration of olivine (unlike metal or mesostasis). This is only one example to illustrate the argument of element complementarity. In past years, many elemental complementarities have been identified in most carbonaceous chondrite groups as well as in Rumuruti chondrites, and for many element pairs (Wood, 1963, 1985; Palme et al., 1992; Klerner and Palme, 1999a, 1999b, 2000; Klerner, 2001; Hezel and Palme, 2008, 2010; Ebel et al., 2008, 2016; Palme et al., 2014b, 2015, Friend et al., 2017b). Different element or isotope ratios in chondrules and matrix together with a CI chondritic bulk chondrite as demonstrated in Figure 4.1 is the basic concept of complementarity. It is, however, also possible to use volatile element depletion patterns in chondrules, matrix, and bulk chondrites to support the argument of complementarity: As we have seen in Section 4.2, bulk carbonaceous chondrites are always depleted in volatile elements relative to solar abundances – with the exception of CI chondrites – and element depletions are monotonically correlated with element 50-percent condensation temperatures (Figure 4.3). In bulk carbonaceous chondrites, volatile element contents decrease in the sequence CI > CM > CR > CV CO. Matrices have generally higher volatile element concentrations than the bulk chondrites, and the degree of volatile element depletion varies among chondrite groups (Figure 4.6). Bland et al. (2005) determined volatile element concentrations in matrices and found that contrary to the bulk chondrite, matrix volatile element depletion patterns only partly correlate with element

50-percent condensation temperatures (T50%, cond). In CV chondrites, for example, siderophile and chalcophile elements (e.g., As, Zn, Bi, Pb) are significantly more enriched in the matrix than lithophile elements (e.g., K, Rb, Cs), which show similar levels of depletion in matrix as in bulk (Figure 4.6). Chondrules then need to have the complementary volatile element pattern to the matrix (i.e., chondrules need to be depleted in siderophile and chalcophile elements and enriched in lithophile elements). Only then can the combination of matrix and chondrule volatile element patterns produce the observed monotonic, volatile element depletion pattern that Chondrule – Matrix Complementarity 105

Figure 4.6 Comparison of volatile element depletion patterns between CV and CM chondrites.

correlates with T50%, cond. It is highly unlikely that mixing of chondrules and matrix, each with different volatile element depletion patterns, results in a volatility-controlled depletion pattern. It is much more likely that chondrules and matrix formed together in a previously volatile element depleted reservoir. Then, lithophile elements were preferentially incorporated into chondrules and siderophile and chalcophile elements into matrix (see Section 4.8). As the high Na content of chondrule olivine requires high initial Na (e.g., Fedkin and Grossman, 2013 and references therein) and other alkalis in chondrule melts, aqueous alteration cannot be responsible for the high alkali contents in chondrules. As stated above, matrix in carbonaceous chondrites is not CI-like, and has different compos- itions in different groups. While matrix, chondrules and bulk follow similar trends, these are clearly distinct in various chondrites (e.g., CV and CM in Figure 4.6). The conclusion is that the various groups of carbonaceous chondrites represent reservoirs with different volatile element contents, and differences between these reservoirs are expressed in matrix and chondrules, as well as in bulk volatile element compositions. Each reservoir may have acquired its volatile element composition by incomplete condensation (e.g., Palme et al., 1988). Complementarity has in the past years primarily been investigated for element ratios. Recently, it has been reported that chondrules and matrix in Allende show complementarities of W and Mo isotopes (Becker et al., 2015; Budde et al., 2016a, 2016b), while the bulk chondrite shares essentially the same ratio as other meteorites (especially for W), thus enabling this to serve as a “solar” reference composition (Figure 4.7). These first reports of isotope complementarities further support the findings of element complementarities (see Chapter 10). 106 Dominik C. Hezel, et al.

Figure 4.7 Nucleosynthetic 183/184W isotope chondrule-matrix complementarity. 1Becker et al., 2015; 2Budde et al., 2016a.

4.6 Elemental Redistribution on Chondrite Parent Bodies

Brearley and Krot (2012) summarized the large body of studies on chondrite alteration with the statement that “fluids of some kind were present during thermal processing of virtually all chondrite classes.” This has also been concluded for petrologic type 3 chondrites and from the inverse perspective by Abreu and Brearley (2010), who stated that “only a few [type 3 chondrites] are now thought to be relatively pristine samples.” Alteration on meteorite parent bodies was wide- spread in the early solar system. One may be tempted to explain the complementary element ratios in chondrules and matrix by material exchange between these two components during alteration. Chondrite alteration has been studied in great detail (Brearley and Krot, 2012; Brearley, 2014). The focus of these studies was to understand the mineralogical changes produced by the alteration process. The amount of material redistributed among the chondrite components has not been quantified. Bland et al. (2009) studied the permeability of chondrites and demonstrated that chondrite alteration must be isochemical (i.e., the amount of material redistributed between chondrules and matrix must have been limited). This is also the conclusion of Abreu and Brearley (2011), who concluded that CV chondrite matrix mineralogy was the result of in situ parent body processes, while exchange of elements between chondrules and matrix was highly limited. Most chondrules in carbonaceous chondrites are type I, Mg-rich, Fe-poor, and are dominated by olivine and pyroxene. Matrix in carbonaceous chondrites is typically Mg-poor and Fe-rich (Hezel et al., 2018a; see Figure 4.3). Material redistribution between chondrules and matrix would increase Mg in matrix and decrease Mg in chondrules, with the inverse relationship for Fe. If the observed complementary Fe/Mg element ratios had been established by elemental redistribution, then chondrules should have been initially low in Mg and high in Fe, and gained Mg and lost Fe during alteration. At low temperatures, leaching of chondrules with an aqueous solution would have to remove Fe from chondrule minerals and deliver it to matrix. At the same time, Mg from matrix would have to be delivered to chondrules to make MgO-rich olivines. These processes make little sense, as at elevated temperatures thermodynamics require the homogenization of Fe and Mg concentrations. Therefore, any postulated redistribution of Chondrule – Matrix Complementarity 107

Mg-Fe between chondrules and matrix would require an even more pronounced initial Fe/Mg chondrule–matrix complementarity. Chondrule mesostases are frequently altered to various extents (Brearley, 2014 and refer- ences therein). Therefore, another possible explanation for the observed subchondritic Mg/Si ratio in chondrite matrices and the Mg/Si complementarity would be the redistribution of Si from the chondrules in the matrix. Mass balance calculations prohibit this explanation, as chondrule mesostases cannot release sufficient Si. Further, chondrite matrices have significantly less Si than chondrules (see Figure 4.3), which would be unexpected if chondrules had released major amounts of Si to the matrix. Both aspects were pointed out and calculated in more detail in Hezel et al. (2010) and Palme et al. (2015). Further, Bland et al. (2005) used the fluid-mobile Sr and Sr/Yb ratios in chondrules and matrix from Allende to argue that less than 1 percent of material was redistributed between chondrules and matrix. The two CV chondrites (Y-86751 and Allende) have Ca/Al complementarities between chondrules and matrix. However, in Y-86751 chondrules have superchondritic Ca/Al ratios, while in Allende chondrules have subchondritic Ca/Al ratios – and inverse for the matrices. If alteration was responsible, for example, for redistributing the fluid-mobile Ca, then the comple- mentarities should be the same in both chondrites, which is not the case (Hezel et al., 2008). More likely, spinel grains were incorporated in chondrules in the case of Allende, while these are abundant in the matrix of Y-86751 (Murakami and Ikeda, 1994). In addition, the two refractory elements Al and Ti are fluid-immobile, but display complementary relationship between chondrules and matrix (Ebel et al., 2016; Friend et al., 2017, 2018). Finally, the mass-independent isotope complementarities of 183W and Mo isotopes (Becker et al., 2015; Budde et al., 2016a, 2016b) cannot have been established by redistributing material between chondrules and matrix. Any elemental redistribution of W or Mo should have hom- ogenized all mass-independent isotope signatures between the components. This is in line with the earlier findings from Fe/Mg and Mg/Si complementarities that such redistribution is not possible. If some isotope redistribution occurred, the initial complementarities would have been even more pronounced. In summary, the observed complementarities cannot have been established by alteration. On the contrary, any elemental redistribution between chondrules and matrix during alteration would in many cases only lessen the initial complementarities.

4.7 The Origin of Complementarity

The complementary compositions of chondrules and matrix require significant material redistri- bution between these components before these accreted onto the parent body. Either by (i) the exchange of material between gas and chondrule melt during evaporation and recondensation (e.g., Tissandier et al., 2002; Hezel et al., 2003, 2010; Krot et al., 2004; Bland et al., 2005; Libourel et al., 2006; Friend et al., 2016; Chapter 6), with finer-grained, matrix-destined material possibly gaining proportionally more recondensed volatile elements than chondrules because of their large surface-to-volume ratio (Huss et al., 2005), or metal droplets separating from chondrules (Grossman and Wasson, 1984; Jacquet et al., 2013), or (ii) the preferential incorporation of (poly)mineral grains in either chondrules or matrix (Hezel et al., 2006, 2008; 108 Dominik C. Hezel, et al.

Palme et al., 2015; Budde et al., 2016a, 2016b). Each of these processes, as well as a combination of these, can produce the observed elemental complementarity within a single reservoir. The mass-independent W and Mo isotope complementarities could be explained by process (ii), which might also involve differential incorporation of presolar grains (e.g., Hubbard, 2016a, 2016b). In any case, elemental and isotopic complementarities require that the region initially had a CI chondritic composition for many element ratios, in particular the refractory and main elements. A relatively uniform (i.e., CI chondritic) chemical composition throughout the forming protoplanetary disc is indeed a natural expectation, given the limited relative variations in isotopic compositions observed in the present solar system. Subsequent bulk chemical fractionations – which were transparent for isotopic ratios – produced the nonsolar reservoirs (e.g., Palme et al., 2014b). Thus, CI chondritic ratios of many elements in a bulk chondrite are most likely the result of a CI chondritic parent region, rather than a coincidental mixture of chondrite components such as chondrules and matrix. Most chondrule populations have elemental and isotopic compositional distributions that are normal to log-normal (see Section 4.3). Such distributions result from processes that must have affected an entire chondrule population, e.g., due to variations of chondrule surface-to-volume ratios when material exchange happened with the surrounding gas during chondrule formation or statistical mixing of heterogeneous grains into chondrules (Chapter 14; Hezel and Palme, 2007; Hezel et al., 2010; Friend et al., 2016). Recently, Olsen et al. (2016) reported variable δ54Cr bulk chondrule compositions in the Allende CV chondrite. Also, δ50Ti might be variable, however, so far only four Allende chondrule 50Ti values have been reported (Trinquier et al., 2009). The δ54Cr, and, to an even greater extent, δ50Ti distributions are neither clearly (log)normal, nor multimodal. Olsen et al. (2016), however, interpret the δ54Cr distribution as being multimodal and suggest the individual peaks represent separate chondrule reservoirs, each with a different δ54Cr composition. Chon- drules from these different reservoirs were then mixed into the Allende parent body. However, the variable δ54Cr compositions might be better explained with a nugget effect of small, isotopically different grains added to the chondrule precursor aggregates, which all formed in a single reservoir (see Section 4.5 and Chapter 10). It was further suggested that δ54Cr might be in conflict with complementarity. This is not possible, as bulk Allende has a non-CI chondritic δ54Cr composition. It has been suggested that the olivine-rich, Mg/Si superchondritic chondrules might be the remnants of dunitic parent bodies (Libourel and Krot, 2007). A magmatic origin of chondrules has been suggested before (e.g., Hutchison, 1988; Bridges et al., 1995). Although this is probably true for rare granitic clasts (e.g., Ruzicka and Boynton, 1992; Bischoff et al., 1993), chondrules do not generally follow a magmatic fractionation trend, but rather follow mixing trends between olivine and pyroxene or SiO2, respectively (e.g., Hezel et al., 2006). Further, chondrules typically have flat, unfractionated REE patterns (e.g., Gerber, 2012), which is not expected, if chondrules had experienced magmatic processes inside a parent body. Several studies have recently examined chondrule formation by impact (e.g., Asphaug et al., 2011; Sanders and Scott, 2012; Johnson et al., 2015). In principle, a complementary relationship between millimeter-sized droplets solidified as chondrules and coejected finer or vaporized material could arise by chemical exchange if (i) the impacted target (or rather its surface – then Chondrule – Matrix Complementarity 109 excluding differentiation) had solar composition and (ii) the escaping plume representatively sampled its bulk chemistry. This, however, has so far not been demonstrated. In summary, the elemental and isotopic complementarities between chondrules and matrix, in cases when the bulk is CI chondritic, as well as the continuous elemental and isotopic distributions of chondrule populations, are best explained by preaccretionary material redistri- bution in a local region.

4.8 Physical Characteristics of Matrix Grains and Chondrule Rims

Microchondrules (40 µm) are a ubiquitous chondrule type within primitive chondrites (Bigolski et al., 2016; Dobricăand Brearley, 2016). As features that occur in unequilibrated ordinary chondrites, they are commonly found in the FeO-rich, fine-grained rims surrounding type I chondrules, yet are absent in fine-grained rims surrounding type II chondrules (e.g., Semarkona – LL3.00, NWA 5717 – ungrouped 3.05, Bishunpur – LL3.15; Bigolski et al., 2016). Microchondrules are also found as inclusions in matrix (e.g., Semarkona, MET 00526 – L3.05; Dobricăand Brearley, 2016). Microchondrules were probably produced from a combin- ation of the reheating of chondrule surfaces and the melting of dust within the immediate proximity of chondrules, and became entrained in the surrounding dust as the chondrules were accreting fine-grained rims (Krot et al., 1997b). Therefore, chondrules were not mixed into cold matrix from another region of the nebula; rather, chondrules, microchondrules, matrix, and fine- grained rims formed at the same time and in the same region. There are differing positions about the nature and origin of matrix. Some authors suggest that, based in part on the observation that the size frequency distribution (SFD) of matrix in some chondrites obeys a power law, matrix might be largely composed of fragmented chon- drules, (e.g., Alexander, Hutchison, and Barber, 1989). On a large scale, however, chondrule grain morphology in the most primitive chondrites is not consistent with fragmentation and comminution of chondrules to produce matrix. All matrix olivine grains studied in the sub- micron range (e.g., in Acfer 094; Greshake, 1997) have rounded or elongate morphologies and do not show typical fragmentation features. In addition, the massive small-scale mineralogic heterogeneity of the matrix is complex, which is contrary to the comparatively simple chondrule mineralogy. Notwithstanding possible cataclastic fracturing during parent body chondrule- fragmentation events (e.g., regolith gardening; Nelson and Rubin, 2002), the fine-scale miner- alogy of matrix makes forming this material from comminuted chondrules difficult. Further, matrix cannot be composed simply of fragmented chondrules. With respect to olivine compos- itions, some authors have suggested that most isolated grains in the matrices of CO3 chondrites are chondrule fragments (Jones, 1992). For Acfer 094, Greshake (1997) makes a strong case that they are not fragments, since chondrule and matrix olivines have a similar compositional peak, but matrix olivines do not show the broad compositional range of chondrule olivines. No matrix olivine with >1.9 wt% FeO was found, and in contrast to chondrule olivine, many matrix olivines contain appreciable MnO. In summary, petrographic and petrologic observations suggest that matrix and chondrules formed in the same region and were interacting with each other during multiple stages of chondrule formation. It is not possible to form large portions of matrix by fragmenting chondrules. 110 Dominik C. Hezel, et al. 4.9 Complementarity and Disk Transport

The considerable spreads in ages of chondrules in single chondrites suggested by Al–Mg and Pb–Pb systematics (e.g., Villeneuve et al., 2009; Connelly et al., 2012) are sometimes viewed as difficult to reconcile with the preservation of matrix/chondrule complementarity until accretion, and some authors have in fact argued for rapid accretion after chondrule formation on that basis (e.g., Budde et al., 2016a). Such statements have yet to be substantiated by theoretical modeling. Investigation of complementarity has only recently begun in the astrophysical community (Jacquet et al., 2012; Goldberg et al., 2015; Hubbard, 2016a, 2016b). Here we review the results obtained thus far. It is first necessary to define complementarity quantitatively for a given compositional parameter, say an elemental ratio r such as Mg/Si. Two criteria are mandatory for comple- mentarity: (i) the bulk chondrite needs to be close to solar (i.e., rblk/rCI ’ 1). However, the CI- normalized bulk ratio cannot in itself be the adequate measure of complementarity, for chon- drules and matrix could be both close to solar, in which case the near “solarness” of the bulk would not be informative. Therefore, (ii) chondrules and matrix also need to have compositions rchd and rblk significantly different from each other (i.e., from the bulk), so these have to differ significantly from solar. Specifically, their compositions should be less close to solar than the bulk is. Goldberg et al. (2015) quantified this by using the following complementarity parameter: , ζ ¼ rchd rblk rchd rCI rmtx rblk rmtx rCI

ζ was constructed as a “solar-normalized” matrix/chondrule ratio. That is, assuming one can mathematically construct an ideal mixture of the observed chondrules and matrix (that is, with the same rchd and rblk) with an exactly solar bulk, it can be expressed as: = ζ ¼ Xchd,measured Xmtx,measured Xchd,ideal=Xmtx,ideal

With Xchd and Xmtx signifying the chondrule and matrix fractions, respectively. │ζ-1│ is thus the relative deviation of the matrix/chondrule ratio from perfect complementarity in the observed chondrite. ζ is indeed of course exactly unity if rblk=rCI 6¼ rchd,mtx. More generally, the closer ζ is to 1, the more complementary the chondrite is (according to this definition), for the considered compositional parameter. For the ratio Mg/Si, Goldberg et al. (2015) calculated ζ values of 0.70–1.07 for carbonaceous chondrites from literature data and deemed ζ values satisfying │ζ-1│ < 0.3 to represent “acceptable” complementary relationships. The same calculations for ordinary and enstatite chondrites produce highly deviating ζ values, as these chondrites have clearly non-solar bulk Mg/Si ratios. Goldberg et al. (2015) modeled the aerodynamic redistribution of chondrules and dust grains (neglecting their aggregation; ‘dust’ in this context essentially means the material that is later forming the matrix) in a turbulent protoplanetary disk. For different epochs, they considered the average composition and petrographic makeup of different regions of the disk, assuming they represented that of the chondrites that would accrete there. This approach, it should be noted, ignores the possibility of size sorting upon accretion (e.g., during turbulent concentration; Cuzzi Chondrule – Matrix Complementarity 111 et al., 2001; Cuzzi, et al., 2010) which would have made the chondrule/dust ratio of the chondrite different from the average of the accretion region (the “accretion bias” of Jacquet et al., 2012) unless such size sorting operated on earlier formed representative matrix/chondrule assemblages (Jacquet, 2014). However, the very existence of complementarity argues against any significant size sorting at least for carbonaceous and Rumuruti chondrites, whose bulk composition would otherwise have been modified. Goldberg et al. (2015) considered essentially two starting conditions: (i) In the first case, the bulk disk composition was solar, and chondrules and dust formed in two separate broad regions. Then, chondrules and dust partly mixed. So an intermediate region arose where chondrules and dust were in “acceptable” complementary relationships. But this region hardly ever comprised a significant fraction of all chondrules in the disk (50 percent at most, for a short duration). Complementarity under those circumstances would evidently require delicate fine-tuning of the timing of accretion. (ii) In the second setup, chondrules and dust were placed in the same region and in exact complementary relationships. Since chondrules tend to drift sunward more rapidly than dust particles, they gradually fell out of complementary relationship with the ambient dust. Nevertheless, the time during which at least 50 percent of the chondrules in the simulation domain were still complementary to the co-located dust could be close to, or even longer than, the viscous timescale. The viscous timescale measures the time of turbulent radial mixing in the disk and is inversely proportional to the turbulence parameter α (Shakura and Sunyaev, 1973). This latter parameter, however, has order- of-magnitude uncertainties, and hence the viscous timescale is highly uncertain too. However, since magnetically dead zones are expected to be quiet (Gammie, 1996), the viscous timescale in the inner disk could be of the order of a million years. Chondrule populations could, therefore, be stored for prolonged times together with dust in their formation regions before accretion. Complementarity would be preserved, and is com- patible with significant time differences between chondrule formation events (i.e., ages) and late parent body agglomeration.

The reason underlying this robustness of complementarity is that chondrules and, a fortiori, matrix grains, are small objects. They are thus likely to be tightly coupled to the gas (i.e., have a small ( 1) solid/gas decoupling parameter S; Jacquet et al., 2012). As chondrules and dust then both follow the gas, they cannot significantly fractionate from each other within the viscous timescale. Thus, each parcel of chondrule+dust+gas will preserve its original (solar) compos- ition. This holds even if contributions from different sources are mixed together as a result of turbulence. Therefore, chondrules (and matrix grains) in a given chondrite could represent different chondrule-forming events and regions without contradicting complementarity (Cuzzi et al., 2005; Jacquet et al., 2012). Indeed, a mixture of individually complementary contribu- tions must itself be complementary overall; evidently solar + solar = solar. It should be noted that this argument is not specific to chondrule-forming sources. Obvi- ously, a parcel of “cold” CI chondritic dust that was never subjected to chondrule formation will be transported as such and can be mixed into a “processed” matrix (and chondrules) – and explain the presence of organic matter and presolar grains in chondrites – without affecting the overall complementarity (Jacquet et al., 2016). At the other temperature extreme, should a 112 Dominik C. Hezel, et al. parcel of originally entirely gaseous solar matter undergo sequential condensation and be transported in the same unbiased way envisioned in the previous text, the mix of condensates will retain complementarity (whatever the fraction thereof subsequently converted into chon- drules may be). In particular, CAIs may bear genetic relationships with their Al-depleted host (Hezel et al., 2008; Jacquet et al., 2012).

4.10 Summary and Conclusions

In Figure 4.8, we suggest a temperature-time framework for chondrule and matrix formation that explains the observations and processes presented in the previous text: CAIs are the oldest objects (e.g., Bouvier and Wadhwa, 2010; Connelly et al., 2012). These are most probably xenolithic material in chondrites (e.g., Pack et al., 2004; MacPherson et al., 2005; Hezel et al., 2008), which formed and were stored in various regions before incorporation into the meteorite parent bodies 2–4 million years later. Chondrule and matrix precursor material partly condensed in the protoplanetary disk and partly still was unprocessed interstellar material. Around this time, bulk chondrite volatile

Figure 4.8 Proposed sequence of events and processes in the protoplanetary disk. min: minutes; h: hours; comp. est. by e.g. incorp. of: complementarity established by e.g., incorporation of. (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.) Chondrule – Matrix Complementarity 113 depletion occurred to various degrees in different volumes of the protoplanetary disk, thereby forming chemically different reservoirs. These represent the parental reservoirs of the various meteorite groups. Also at this early time, presolar and/or C-rich material was added to these separate reservoirs. Further, during this stage, refractory material such as olivine, perovskite, or spinel preferentially agglomerated into material that subsequently became chondrule precursor aggregates. Additional preferential incorporation of grains such as metal/sulfide/and so on into matrix or chondrules can explain, for example, the complementary volatile element patterns of chondrules and matrix. During the still unknown chondrule-forming process, the chondrule precursor aggregates were heated above their liquidus, in cases up to 2,000 K and more. Chondrules exchanged material with the ambient gas during this high-temperature stage: for example, forsteritic olivine reacted with SiO(g) to produce low-Ca pyroxene rims, abundantly observed around chondrules (e.g., Friend et al., 2016). This material exchange between chon- drules and the surrounding gas (i.e., chondrules acting as open systems) may explain some of the bulk chondrule elemental variation. Thus, chondrite matrix represents a mix of condensed material from the chondrule formation process, as well as leftover precursor material that was present, but not heated during chondrule formation, and/or was later mixed in CI chondritic material. This initial mix might be called ‘protomatrix’ and was subsequently altered on the parent body to the matrix we now observe in chondrites (e.g., Abreu and Brearley, 2011; Brearley, 2014 and references therein). Complementarity mandates that all these processes happened in a common reservoir, but individual chondrites may contain contributions from repeated chondrule-forming events in these single reservoirs, as is evident from relict grains. However, no more than probably 2–5 cycles seem possible (Hezel and Palme, 2007; Baecker et al., 2017). High volatile element concentrations in chondrules – e.g., Na (Fedkin and Grossman, 2013 and references therein) – and stable isotope compositions of chondrules (e.g., Alexander et al., 2000; Hezel et al., 2010; Bouvier et al., 2013) require high ambient pressure significantly above the canonical 10–4 bar and/or high dust/gas ratios (e.g., Cuzzi and Alexander, 2006; Alexander et al., 2008) during chondrule formation. Finally, chondrules and matrix, together with variable amounts of CAIs formed the chondrite parent bodies. Alteration of the parent body might have redistributed elements and isotopes, but are not responsible for most of the observed complementarities. Minor differences in the different parental reservoirs during the sequence of all these processes probably resulted in chondrule populations with different size and compositional distributions, as well as different chondrule/matrix ratios and matrix compositions, leading to the conclusion of Jones (2012) that each chondrite represents a unique reservoir. Astrophysical models as the one presented in Section 4.9 can reconcile complementarity with observed age differences of chondrule formation. Chondrule formation could have happened within a reser- voir that contained all chondrite components for a prolonged period of time. Later addition of CI material with presolar grains would not disturb the observed complementarity and could not only explain the presence of presolar grains, but also potentially the distributions of bulk chondrule 54Cr compositions. From the observation of complementarity, we conclude a genetic relationship exists between chondrules and matrix. This complementary relationship requires the formation of chondrules and matrix from a common parental reservoir for each chondrite group. This finding supports the category of chondrule formation models in which chondrules and matrix form in the same 114 Dominik C. Hezel, et al. location of the protoplanetary disk, and excludes the category of models that require different locations for chondrules and matrix and later mixing of these components.

Acknowledgments

We thank K. Howard, A. Rubin, and A. N. Krot for their constructive reviews, which made this a much more concise contribution. We thank the organizers of this book and the insightful workshop at the Natural History Museum in London.

References

Abreu, N. M., and Brearley, A. J. (2010). Early solar system processes recorded in the matrices of two highly pristine CR3 carbonaceous chondrites, MET 00426 and QUE 99177. Geochimica et Cosmochimica Acta, 74, 1146–1171. Abreu, N. M., and Brearley, A. J. (2011). Deciphering the nebular and asteroidal record of silicates and organic material in the matrix of the reduced CV3 chondrite Vigarano. Meteoritics & Planetary Science, 46, 252–274. Ahrens, L. H. (1965). Observations on the Fe-Si-Mg relationship in chondrites. Geochimica et Cosmochimica Acta, 29, 801–806. Alexander, C. M. O’D., and Grossman, J. N. (2005). Alkali elemental and potassium isotopic compositions of Semarkona chondrules. Meteoritics & Planetary Science, 40, 541–556. Alexander, C. M. O’D., Hutchison, R., and Barber, D. J. (1989b). Origin of chondrule rims and interchondrule matrices in unequilibrated ordinary chondrites. Earth and Planetary Sci- ence Letters, 95, 187–207. Alexander, C. M. O’D., Grossman, J. N., Wang, J., et al. (2000). The lack of potassium- isotopic fractionation in Bishunpur chondrules. Meteoritics & Planetary Science, 35, 859–868. Alexander, C. M. O’D., Grossman, J. N., Ebel, D. S., and Ciesla, F. J. (2008). The formation conditions of chondrules and chondrites. Science, 320, 1617–1619. Amsellem, E., Moynier, F., Pringle, E. A., et al. (2017). Testing the chondrule-rich accretion model for planetary embryos using calcium isotopes. Earth and Planetary Science Letters, 469,75–83. Anders, E. (1964). Origin, age and composition of meteorites. Space Science Reviews, 3, 583–714. Asphaug, E., and Jutzi, M. (2011). Chondrule formation during planetesimal accretion. Earth and Planetary Science Letters, 308, 369–379. Baecker, B., Rubin A. E., and Wasson J. T. (2017). Secondary melting events in Semarkona chondrules revealed by compositional zoning in low-Ca pyroxene. Geochimica et Cosmo- chimica Acta, 211, 256–279 Becker, M., Hezel, D. C., Schulz, T., Elfers, B. -M., and Münker, C. (2015). The age of CV chondrites from component specificHf–W systematics. Earth and Planetary Science Letters, 432, 472–482. Berg T., Maul, J., Schönhense, G., et al. (2009). Direct evidence for condensation in the early solar system and implications for nebular cooling rates. The Astrophysical Journal, 702, L172–176. Bigolski, J. N., Weisberg, M. K., Connolly, H. C., and Ebel D. S. (2016). Microchondrules in three unequilibrated ordinary chondrites. Meteoritics & Planetary Science, 51, 235–260. Bischoff, A., Geiger, T., Palme, H., et al. (1993). Mineralogy, chemistry, and noble gas contents of Adzhi-Bogdo—an LL3–6 chondritic breccia with L- chondritic and granitoidal clasts. Meteoritics, 28, 570–578. Chondrule – Matrix Complementarity 115

Bizzarro, M., Baker, J. A., and Haack, H. (2004). Mg isotope evidence for contemporaneous formation of chondrules and refractory inclusions. Nature, 431, 275–278. Bland, P. A., Alard, O., Benedix, G. K., et al. (2005). Volatile fractionation in the early solar system and chondrule/matrix complementarity. Proceedings of the National Academy of Sciences, 102, 13755–13760. Bland, P. A., Stadermann, F. J., Floss, C., et al. (2007). A cornucopia of presolar and early solar system materials at the micrometer size range in primitive chondrite matrix. Meteoritics & Planetary Science, 42, 1417–1427. Bland, P. A., Jackson, M. D, Coker, R. F., et al. (2009). Why aqueous alteration in asteroids was isochemical: High porosity 6¼ high permeability. Earth and Planetary Science Letters, 287, 559–68. Bouvier, A., and Wadhwa, M. (2010). The age of the Solar System redefined by the oldest Pb– Pb age of a meteoritic inclusion. Nature Geoscience, 3, 637–641. Bouvier, A., Wadhwa, M., Simon, S. B., and Grossman, L. (2013). Magnesium isotopic fractionation in chondrules from the Murchison and Murray CM2 carbonaceous chon- drites. Meteoritics & Planetary Science, 48, 339–53. Brearley, A. J. (2014). Nebular Versus Parent Body Processing. In A. M. Davis (Ed.), Meteorites and Cosmochemical Processes. In H. D. Holland and K. K. Turekian (Eds.), Treatise on Geochemistry (Second Edition), 1, 309–334. Oxford, UK: Elsevier. Brearley, A. J. (1993). Matrix and fine-grained rims in the unequilibrated CO3 chondrite, ALHA 77307: Origins and evidence for diverse, primitive nebular dust components. Geochimica et Cosmochimica Acta, 57, 1521–1550. Brearley, A. J., and Krot, A. N. (2013). Metasomatism in the early solar system: The record from chondritic meteorites. In D. Harlov and H. Austrheim (Eds.), Metasomatism and the Chemical Transformation of Rock, 659–789. Lecture Notes in Earth System Sciences. Berlin and Heidelberg, Germany: Springer-Verlag. Brennecka, G. A., Budde, G., and Kleine, T. (2015). Uranium isotopic composition and absolute ages of Allende chondrules. Meteoritics & Planetary Science, 50, 1995–2002. Bridges, J. C., Franchi, I. A., Hutchinson, R., et al. (1995). Cristobalite-and tridymite-bearing clasts in Parnallee (LL3) and Farmington (L5). Meteoritics, 30, 715–727. Budde, G., Burkhardt, C., Brennecka, G. A., et al. (2016a). Molybdenum isotopic evidence for the origin of chondrules and a distinct genetic heritage of carbonaceous and non- carbonaceous meteorites. Earth and Planetary Science Letters, 454, 293–303. Budde, G., Kleine, T., Kruijer, T. S., Burkhardt, C., and Metzler, K. (2016b). Tungsten isotopic constraints on the age and origin of chondrules. Proceedings of the National Academy of Science, 113, 2886–2891. Chaussidon, M., and Robert, F. (1998). 7Li/6Li and 11B/10B variations in chondrules from the Semarkona unequilibrated chondrite. Earth and Planetary Science Letters, 164, 577–589. Clarke, R. S., Jarosewich, E., Mason, B., et al. (1971). The Allende, Mexico, meteorite shower. Smithsonian Contributions to the Earth Sciences, 5,1–53. Clayton, R. N., Mayeda, T. K., Goswami, J. N., and Olsen, E. J. (1991). Oxygen isotope studies of ordinary chondrites. Geochimica et Cosmochimica Acta, 55, 2317–2337. Connelly, J., Bizzarro, M., Krot, A., et al. (2012). The absolute chronology and thermal processing of solids in the solar protoplanetary disk. Science, 338, 651–655. Cuzzi, J. N., and Ciesla, F. J. (2005). Nebula evolution of thermally processed solids: Recon- ciling models and meteorites. In A. N. Krot, E. R. D. Scott, & B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk., 732–773. Astronomical Society of the Pacific Conference Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Cuzzi, J. N., and Alexander, C. M. O’.D. (2006). Chondrule formation in particle-rich nebular regions at least hundreds of kilometres across. Nature, 441, 483–485. 116 Dominik C. Hezel, et al.

Cuzzi, J. N., Hogan, R. C., Paque, J. M., and Dobrovolskis, A. R. (2001). Size-selective concentration of chondrules and other small particles in protoplanetary nebula turbulence. The Astrophysical Journal, 546, 496–508. Cuzzi, J. N., Hogan, R. C., and Bottke, W. F. (2010). Towards initial mass functions for asteroids and Kuiper Belt Objects. Icarus, 208, 518–538. Daly, L., Bland, P. A., Dyl, K. A., et al. (2017a). In situ analysis of refractory metal nuggets in carbonaceous chondrites. Geochimica et Cosmochimica Acta (in press) Daly, L., Bland, P. A., Dyl, K. A., et al. (2017b). Crystallography of refractory metal nuggets in carbonaceous chondrites: A transmission Kikuchi diffraction approach. Geochimica et Cosmochimica Acta (in press) Das, J. P., and Murty, S. V. S. (2008). Nitrogen isotopes in chondrules: Signatures of precursors and formation processes. Current Science, 94, 485–489. Desch, S. J., and Connolly, H. C. (2002). A model of the thermal processing of particles in solar nebula shocks: Application to the cooling rates of chondrules. Meteoritics & Planetary Science, 37, 183–207. Dobrica,̆ E., and Brearley, A. J. (2016). Microchondrules in two unequilibrated ordinary chondrites: Evidence for formation by splattering from chondrules during stochastic collisions in the solar nebula. Meteoritics & Planetary Science, 51, 884–905. Ebel, D. S., Weisberg, M. K., Hertz, J., and Campbell, A. J. (2008). Shape, metal abundance, chemistry, and origin of chondrules in the Renazzo (CR) Chondrite. Meteoritics & Planetary Science, 43, 1725–40. Ebel, D. S., Brunner, C., Konrad, K., et al. (2016). Abundance, major element composition and size of components and matrix in CV, CO and Acfer 094 chondrites. Geochimica et Cosmochimica Acta, 172, 322–56. Fedkin, A. V., and Grossman, L. (2013). Vapor saturation of sodium: Key to unlocking the origin of chondrules. Geochimica et Cosmochimica Acta, 112, 226–50. Friedrich, J. M., Weisberg, M. K., Ebel, D. S., et al. (2015). Chondrule size and related physical properties: A compilation and evaluation of current data across all meteorite groups. Chemie Der Erde – Geochemistry, 75, 419–43. Friend, P., Hezel, D. C., and Mucerschi, D. (2016). The conditions of chondrule formation, Part II: Open system. Geochimica et Cosmochimica Acta, 173, 198–209. Friend, P., Hezel, D. C., Barrat, J.-A., et al. (2017). Composition, petrology and chondrule- matrix complementarity of the recently discovered Jbilet Winselwan CM2 chondrite. Meteoritics & Planetary Science (in revisions). Friend, P., Hezel, D. C., Palme, H., Bischoff, A., and Gelissen, M. (2018). Complementary element relationships between chondrules and matrix in Rumuruti chondrites. Earth and Planetary Science Letters, 480,87–96. Galy, A., Young, E. D., Ash, R. D., and O’Nions, R. K. (2000). The Formation of chondrules at high gas pressures in the solar nebula. Science, 290, 1751–1753. Gammie, C. F. (1996). Layered accretion in T Tauri disks. The Astrophysical Journal, 457, 355–362. Gerber, M. (2012). Chondrule formation in the early Solar System: A combined ICP-MS, ICP- OES and petrological study. PhD thesis, University of Münster. Goldberg, A. Z., Owen, J. E., and Jacquet, E. (2015). Chondrule transport in protoplanetary discs. Monthly Notices of the Royal Astronomical Society, 452, 4054–4069. Gordon, S. G. (2009). The composition of components in primitive meteorites. PhD thesis, Imperial College, London. Greshake, A. (1997). The primitive matrix components of the unique carbonaceous chondrite Acfer 094: A TEM study. Geochimica et Cosmochimica Acta, 61, 437–452 Grossman, J. N., and Wasson, J. T. (1984). The origin and history of the metal and sulfide components of chondrules. Geochimica et Cosmochimica Acta, 49, 925–939. Grossman, J. N., and Brearley, A. J. (2005). The onset of metamorphism in ordinary and carbonaceous chondrites. Meteoritics & Planetary Science, 40,87–122. Chondrule – Matrix Complementarity 117

Harju, E. R., Rubin, A. E., Ahn, I., et al. (2014). Progressive aqueous alteration of CR carbonaceous chondrites. Geochimica et Cosmochimica Acta, 139, 267–292. Hewins, R. (1997). Chondrules. Annual Review of Earth and Planetary Science, 25,61–83. Hewins, R. H., Bourot-Denise, M., Zanda, B., et al. (2014). The Paris meteorite, the least altered CM chondrite so far. Geochimica et Cosmochimica Acta, 124, 190–222. Hezel, D. C., Harak, M., and Libourel, G. (2018a). What we know about elemental bulk chondrule and matrix compositions: Presenting the ChondriteDB database. Chemie der Erde – Geochemistry. doi:10.1016/j.chemer.2017.05.003 Hezel, D. C., Needham, A. W., Armytage, R., et al. (2010). A nebula setting as the origin for bulk chondrule Fe isotope variations in CV chondrites. Earth and Planetary Science Letters, 296, 423–433. Hezel, D. C., and Palme, H. (2007). The conditions of chondrule formation, Part I: Closed system. Geochimica et Cosmochimica Acta, 71, 4092–4107. Hezel, D. C., and Palme, H. (2008). Constraints for chondrule formation from Ca-Al distribution in carbonaceous chondrites. Earth and Planetary Science Letters, 265, 716–725. Hezel, D. C., and Palme, H. (2010). The chemical relationship between chondrules and matrix and the chondrule-matrix complementarity. Earth and Planetary Science Letters, 294, 85–93. Hezel, D.C., Palme, H., Brenker, F.E., and Nasdala, L. (2003). Evidence for fractional conden- sation and reprocessing at high temperatures in CH-chondrites. Meteoritics & Planetary Science, 38, 1199–1216. Hezel, D. C., Palme, H., Nasdala, L., and Brenker, F. E. (2006). Origin of SiO2-rich compon- ents in ordinary chondrites. Geochimica et Cosmochimica Acta, 70, 1548–1564. Hezel, D. C., Poole, G., Hoyes, J., et al. (2015). Fe and O isotope composition of meteorite fusion crusts: Possible natural analogues to chondrule formation? Meteoritics & Planetary Science, 50, 229–242. Hezel, D. C., Russell, S. S., Ross, A. J., and Kearsley, A. T. (2008). Modal abundances of CAIs: Implications for bulk chondrite element abundances and fractionations. Meteoritics & Planetary Science, 43, 1879–1894. Hezel, D. C., Wilden, J. S., Becker, D., et al. (2018b). Fe isotope composition of bulk chondrules from Murchison (CM2): Constraints for parent body alteration, nebula processes and chondrule-matrix complementarity. Earth & Planetary Science Letters, 490, 31–39. Howard, K. T., Alexander, C. M. O’D., Schrader, D. L., and Dyl, K. A. (2015). Classification of hydrous meteorites (CR, CM and C2 ungrouped) by phyllosilicate fraction: PSD-XRD modal mineralogy and planetesimal environments. Geochimica et Cosmochimica Acta, 149, 206–222. Howard, K. T., Benedix G. K., Bland P. A., and Cressey, G. (2009). Modal mineralogy of CM2 chondrites by PSD-XRD, Part 1: Total phyllosilicate abundance and the degree of aqueous alteration. Geochimica et Cosmochimica Acta, 73, 4579–4589. Howard, K. T., Benedix, G. K., Bland, P. A., and Cressey, G. (2010). Modal mineralogy of CV3 chondrites by PSD-XRD. Geochimica et Cosmochimica Acta, 74, 5084–5097. Howard, K. T., Benedix, G. K., Bland, P. A., and Cressey, G. (2011). Modal mineralogy of CM chondrites by PSD-XRD, Part 2: Degree, nature and settings of aqueous alteration. Geochimica et Cosmochimica Acta, 75, 2735–2751. Hubbard, A. (2016a). Ferromagnetism and particle collisions: Applications to protoplanetary disks and the meteoritical record. The Astrophysical Journal, 826, 152. Hubbard, A. (2016b). Partitioning tungsten between matrix precursors and chondrule precursors through relative settling. The Astrophysical Journal, 826, 151. Huss, G. R., Alexander, C. M. O’D., Palme, H., Bland, P. A., and Wasson, J. T. (2005). Genetic relationships between chondrules, fine-grained rims, and interchondrule matrix. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, 118 Dominik C. Hezel, et al.

701–731. Astronomical Society of the Pacific Conference Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Hutchison, R., Williams, C. T., Din, V. K., and Clayton, R. N. (1988). A planetary, H-group pebble in the Barwell, L6, unshocked chondritic meteorite. Earth and Planetary Science Letters, 90, 105–18. Jacquet, E. (2014). The quasi-universality of chondrule size as a constraint for chondrule formation models. Icarus, 223, 176–186. Jacquet, E., Gounelle, M., and Fromang, S. (2012). On the aerodynamic redistribution of chondrite components in protoplanetary disks. Icarus, 220, 162–173. Jacquet, E., Paulhiac-Pison, M., Alard, O., Kearsley, A. T., and Gounelle, M. (2013). Trace element geochemistry of CR chondrite metal. Meteoritics & Planetary Science, 48, 1981–1999. Jacquet, E., Barrat, J. -A., Beck, P., et al. (2016). Northwest Africa 5958: A weakly altered CM- related ungrouped chondrite, not a CI3. Meteoritics & Planetary Science, 51, 851–869. Jarosewich, E. (1990). Chemical analyses of meteorites: A compilation of stony and iron meteorites. Meteoritics, 25, 323–337. Jones, R. H. (1992). On the relationship between isolated and chondrule olivine grains in the carbonaceous chondrite ALHA77307. Geochimica et Cosmochimica Acta, 56, 467–482. Jones, R. H. (2012). Petrographic constraints on the diversity of chondrule reservoirs in the protoplanetary disk. Meteoritics & Planetary Science, 47, 1176–90. Jones, R. H., and Schilk, A. J. (2009). Chemistry, petrology and bulk oxygen isotope compos- itions of chondrules from the Mokoia CV3 carbonaceous chondrite. Geochimica et Cosmochimica Acta, 73, 5854–5883. King, A. J., Schofield, P. F., Howard, K. T., and Russell S. S. (2015). Modal mineralogy of CI and CI-like chondrites by X-ray diffraction. Geochimica et Cosmochimica Acta, 165, 148–160. Klerner, S. (2001). Materie im frühen Sonnensystem: Die Entstehung von Chondren, Matrix und refraktären Forsteriten. PhD Thesis, Universität zu Köln. Klerner, S., and Palme, H. (1999a). Origin of chondrules and matrix in carbonaceous chon- drites. 30th Lunar and Planetary Science Conference, abstract 1272. Klerner, S., and Palme, H. (1999b). Origin of chondrules and matrix in the Renazzo Meteorite (abstract). Meteoritics & Planetary Science, 34, (Supplement), A64. Klerner, S., and Palme, H. (2000). Large titanium/aluminium fractionation between chondrules and matrix in Renazzo and other carbonaceous chondrites (abstract). Meteoritics & Planetary Science, 35, 89. Krot, A. N., and Wasson, J. T. (1995). Igneous rims on low-FeO and high-FeO chondrules in ordinary chondrites. Geochimica et Cosmochimica Acta, 59, 4951–4966. Krot, A. N., Scott, E. R. D., and Zolensky, M. E. (1997a). Origin of fayalitic olivine rims and lath-shaped matrix olivine in the CV3 chondrite Allende and its dark inclusions. Meteori- tics, 32,31–49. Krot, A. N., Wasson, J. T., Rubin, A. E., Scott, E. R. D., and Keil, K. (1997b). Microchondrules in ordinary chondrites: Implications for chondrule formation. Geochimica et Cosmochi- mica Acta, 61, 463–475. Krot, A. N., Libourel, G., Goodrich, C., and Petaev, M. I. (2004). Silica-igneous rims around magnesian chondrules in CR carbonaceous chondrites: Evidence for fractional condensa- tion during chondrule formation. Meteoritics & Planetary Science, 39, 1931–1955. Krot, A. N., Yurimoto, H., McKeegan, K. D., et al. (2006). Oxygen isotopic compositions of chondrules: Implications for evolution of oxygen isotopic reservoirs in the inner solar nebula. Chemie Der Erde – Geochemistry, 66, 249–276. Leroux, H., Cuvillier, P., Zanda, B., and Hewins, R. H. (2015). GEMS-like material in the matrix of the Paris meteorite and the early stages of alteration of CM chondrites. Geochi- mica et Cosmochimica Acta, 170, 247–265. Chondrule – Matrix Complementarity 119

Libourel, G., and Krot, A. N. (2007). Evidence for the presence of planetesimal material among the precursors of magnesian chondrules of nebular origin. Earth and Planetary Science Letters, 254,1–8. Libourel, G., Krot, A., and Tissandier, L. (2006). Role of gas–melt interaction during chondrule formation. Earth and Planetary Science Letters, 251, 232–240. Lodders, K., Palme, H., and Gail, H. P. (2009). Abundances of the elements in the solar system. In J. E. Trümper (Ed.), Landolt-Börnstein, New Series, vol.VI/4B, 560–598. Berlin, Germany: Springer. Luck, J. -M., Othman, D. B., and Albarède, F. (2005). Zn and Cu isotopic variations in chondrites and iron meteorites: Early solar nebula reservoirs and parent-body processes. Geochimica et Cosmochimica Acta, 69, 5351–5363. MacPherson, G. J., Simon, S. B., Davis, A. M., Grossman, L., and Krot, A. N. (2005). Calcium- aluminum-rich inclusions: Major unanswered questions. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, 225–250. Astronomical Society of the Pacific Conference Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Mason, B., and Wiik, H. B. (1962). The Renazzo meteorite. American Museum Novitates, 2106,1–11. McNally, C. P., Hubbard, A., Mac Low, M. -M., Ebel, D. S., and D’Alessio, P. (2013). Mineral processing by short circuits in protoplanetary disks. The Astrophysical Journal, 767, L2–L7. Molini-Velsko, C., Mayeda, T. K., and Clayton, R. N. (1986). Isotopic composition of silicon in meteorites. Geochimica et Cosmochimica Acta, 50, 2719–2726. Moynier, F., Agranier, A., Hezel, D. C., and Bouvier, A. (2010). Sr stable isotope composition of Earth, the Moon, Mars, Vesta and meteorites. Earth and Planetary Science Letters, 300, 359–366. Morris, M. A., and Desch, S. J. (2010). Thermal histories of chondrules in solar nebula shocks. The Astrophysical Journal, 722, 1474–1494. Mullane, E., Russell, S. S., and Gounelle, M. (2005). Nebular and asteroidal modification of the iron isotope composition of chondritic components. Earth and Planetary Science Letters, 239, 203–218. Murakami, T., and Ikeda, Y. (1994). Petrology and mineralogy of the Yamato-86751 CV3 chondrite. Meteoritics, 29, 397–409. Nelson, V. E., and Rubin, A. E. (2002). Frequency distributions of chondrules and chondrule fragments in LL3 chondrites: Implications for parent-body fragmentation of chondrules. Meteoritics & Planetary Science, 37, 1361–1376. Nguyen, A., and Zinner, E. (2004). Discover of ancient silicate stardust in a meteorite. Science, 303, 1496–1499. Niemeyer, S. (1988). Titanium isotopic anomalies in chondrules from carbonaceous chondrites. Geochimica et Cosmochimica Acta, 52, 309–318. Nuth, J. A. III, Brearley, A. J., and Scott, E. R. D. (2005). Microcrystals and amorphous material in comets and primitive meteorites: Keys to understanding processes in the early Solar System. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, 675–700. Astronomical Society of the Pacific Conference Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Olsen, M. B., Wielandt, D., Schiller, M., Van Kooten, E. M. M. E., and Bizzarro, M. (2016). Magnesium and 54Cr isotope compositions of carbonaceous chondrite chondrules – insights into early disk processes. Geochimica et Cosmochimica Acta, 191, 118–38. Pack, A., Shelley, M. G., and Palme, H. (2004). Chondrules with peculiar REE patterns: Implications for solar nebular condensation at high C/O. Science, 303, 997–1000. Palme, H., Larimer, J. W., and Lipschutz, M. E. (1988). Moderately volatile elements. In J. F. Kerridge and M. S. Matthews (Eds.), Meteorites and the Early Solar System, 436–461. Tucson, AZ: University of Arizona Press. 120 Dominik C. Hezel, et al.

Palme, H., Spettel, B., Kurat, G., and Zinner, E. (1992). Origin of Allende chondrules. 23rd Lunar and Planetary Science Conference, abstract 1021. Palme, H., Lodders, K., and Jones, A. (2014a). Solar system abundances of the elements. In A. M. Davis (Ed.), Planets, Asteroids, Comets and the Solar System. In H. D. Holland and K. K. Turekian (Eds.), Treatise on Geochemistry (Second Edition), 2, 15–36. Oxford, UK: Elsevier. Palme, H., Spettel, B., and Hezel, D. C. (2014b). Siderophile elements in chondrules of CV- chondrites. Chemie der Erde, 74, 507–516. Palme, H., Hezel, D. C., and Ebel, D. S. (2015). The origin of chondrules: constraints from matrix composition and matrix-chondrule complementarity. Earth and Planetary Science Letters, 411,11–19. Pringle, E. A., Moynier, F., Beck, P., Paniello, R., and Hezel, D. C. (2017). The origin of volatile element depletion in early solar system material: Clues from Zn isotopes in chondrules. Earth and Planetary Science Letters, 468,62–71. Patchett, P. J. (1980). Sr isotopic fractionation in Allende chondrules: A reflection of solar nebular processes. Earth and Planetary Science Letters, 50, 181–88. Rai, V. K., and Thiemens, M. H. (2007). Mass independently fractionated sulfur components in chondrites. Geochimica et Cosmochimica Acta, 71, 1341–54. Rubin, A. E. (1984). Coarse-grained chondrule rims in type 3 chondrites. Geochimica et Cosmochimica Acta, 48, 1779–1789. Rubin, A. E., and Wasson, J. T. (1987). Chondrules, matrix and coarse-grained chondrule rims in the Allende meteorite: Origin, interrelationships and possible precursor components. Geochimica et Cosmochimica Acta, 51, 1923–1937. Ruzicka, A., and Boynton, W. V. (1992). A distinctive silica-rich, sodium-poor igneous clast in the Bovedy (L3) chondrite. Meteoritics, 27, 283. Sanders, I. S, and Scott, E. R. D. (2012). The origin of chondrules and chondrites: Debris from low-velocity impacts between molten planetesimals? Meteoritics & Planetary Science, 47, 2170–2192. Seitz, H. -M., Zipfel, J., Brey, G. P., and Ott, U. (2012). Lithium isotope compositions of chondrules, CAI and a dark inclusion from Allende and ordinary chondrites. Earth and Planetary Science Letters, 329,51–59. Shakura, N. I., and Sunyaev, R. A. (1973). Black holes in binary systems. Observational appearance. Astronomy & Astrophysics, 24, 337–355. Shu, F. H., Shang, H., and Lee, T. (1996). Toward an astrophysical theory of chondrites. Science, 271, 1545–1552. Stracke, A., Palme, H., Gellissen, M., et al. (2012) Refractory element fractionation in the Allende meteorite: Implications for solar nebula condensation and the chondritic compos- ition of planetary bodies. Geochimica et Cosmochimica Acta, 85, 114–141. Tissandier, L., Libourel, G., and Robert, F. (2002). Gas–melt interactions and their bearing on chondrule formation. Meteoritics & Planetary Science, 37, 1377–1389. Trinquier, A., Elliott, T., Ulfbeck, D., et al. (2009). Origin of nucleosynthetic isotope hetero- geneity in the solar protoplanetary disk. Science, 324, 374–76. Trinquier, A., Birck, J. L., and Allègre, C. J. (2007). Widespread 54Cr heterogeneity in the inner solar system. The Astrophysical Journal, 655, 1179–1185. Van Kooten, E. M. M. E., Wielandt, D., Schiller M., et al. (2016). Isotopic evidence for primordial molecular cloud material in metal-Rich carbonaceous chondrites. Proceedings of the National Academy of Sciences, 113, 2011–16. Villeneuve, J., Chaussidon, M., and Libourel, G. (2009). Homogeneous distribution of aluminum-26 in the solar system from the magnesium isotopic composition of chondrules. Science, 325, 985–988. Wasson, J. T., and Kallemeyn, G. W. (1988). Compositions of chondrites. Philosopical Transactions of the Royal Society, London A, 325, 535–544 Chondrule – Matrix Complementarity 121

Weisberg, M. K., Prinz, M., Clayton, R. N., and Mayeda, T. K. (1993). The CR (Renazzo-type) carbonaceous chondrite group and its implications. Geochimica et Cosmochimica Acta, 57, 1567–1586. Wiik, H. B. (1956). The chemical composition of some stony meteorites. Geochimica et Cosmochimica Acta, 9, 279–289. Wolf, D., and Palme, H. (2001). The solar system abundances of P and Ti and the nebular volatility of P. Meteoritics & Planetary Science, 36, 559–572. Wombacher, F., Rehkaemper, M., Mezger, K., Bischoff, A., and Muenker, C. (2008). Cadmium stable isotope cosmochemistry. Geochimica et Cosmochimica Acta, 72, 646–667. Wood, J. A. (1963). On the origin of chondrules and chondrites. Icarus, 2, 152–180. Wood, J. A. (1967). Olivine and pyroxene compositions in Type II carbonaceous chondrites. Geochimica et Cosmochimica Acta, 31, 2095–2108. Wood, J. A. (1985). Meteoritic constraints on processes in the solar nebula. In D. C. Black and M. S. Matthews (Eds.), Protostars and Planets II, 687–702. Tucson, AZ: University of Arizona Press. Zanda, B., Hewins, R. H., Bourot-Denise, M., Bland, P. A., and Albarède, F. (2006). Formation of solar nebula reservoirs by mixing chondritic components. Earth and Planetary Science Letters, 248, 650–660. 5 The Chondritic Assemblage Complementarity Is Not a Required Hypothesis brigitte zanda, e´ric lewin, and munir humayun

Abstract

Chondrules and matrices make up most of the bulk of chondrites. Chondrites have elemental compositions close to that of the Sun except for volatility related depletions and iron fraction- ation relative to silicon. Being igneous, chondrules are volatile depleted, while their host matrices are considered to have assembled mostly from low-temperature (volatile-rich) material. There is an outstanding controversy as to whether (i) chondrules and matrices formed inde- pendently and subsequent mixing of these components explains the departure of chondrite compositions from solar abundances for major elements, or whether (ii) their complementary patterns imply that the two dissimilar components were formed together from a reservoir with a solar composition and accreted without being separated to maintain this solar composition. The question around which this controversy centers has important implications for our understanding of the working of the protoplanetary disk, as complementarity between chondrules and matrices would rule out some chondrite formation models, such as those that imply an origin of these components from widely separated parts of the disk. In the present paper, we discuss literature data for chondrules, matrices, and bulk carbonaceous chondrites. We point out that, except for those of the CI group, all chondrites are fractio- nated with respect to the Sun for all major and minor elements across the condensation temperature range. We show that bulk chondrite compositions are better reproduced by adding a generic chondrule composition to the proper amount of CI-composition matrix, rather than by combining the in situ measured compositions of their chondrules and matrices. These results indicate that chondrule and matrix compositions in any given chondrite are not genetically related to one another. We also discuss the case of moderately volatile elements, for which a similarity of patterns between chondrules and matrices has been reported, and argue that this similarity was established as a result of exchange during alteration on the chondrite parent body and does not reflect a common nebular reservoir. Last, we contend that the recently discovered tungsten and molybdenum isotopic differences between chondrules and matrices argue in favor of distinct isotopic reservoirs for these chondritic components and hence against a genetic relationship between chondrules and their host matrices. A time interval between the formation of chondrules and of matrices and long-range relative transport of these components in the disk are, therefore, not ruled out by existing chemical and isotopic constraints.

122 The Chondritic Assemblage 123 5.1 Introduction

Most chondrites comprise components formed or processed at high temperature (chondrules and refractory inclusions) embedded in a volatile-rich fine-grained matrix assumed to be thermally unprocessed. Compared to other meteorites or rocks from planetary bodies, chondrites have bulk chemical compositions that resemble that of the Sun, within uncertainty of measurement of Sun’s photosphere. These compositions are qualified as “chondritic” or “solar” because the Sun accounts for more than 99 percent of the mass of the solar system. However, only CI chondrites (carbonaceous chondrites of Ivuna type) actually have the same composition as the Sun, except for the most volatile elements such as H, C, N, and O. The CI chondrite compositions are therefore widely used to estimate the abundances of almost all the elements in the solar system (e.g., Anders and Ebihara, 1982; Cameron, 1982; Anders and Grevesse, 1989; Allègre and Lewin, 2001; Lodders et al., 2009; Palme et al., 2014) and used as a reference for the composition of every other rock. Unlike other chondrites, CIs are essentially devoid of chondrules and other high temperature components and they have undergone the highest degree of aqueous alteration of all the meteorite classes (Scott and Krot, 2014). Destruction by aqueous alteration of the original high temperature components of CIs has been discussed, but the scarce grains that are observed lack alteration, contain preaccretional radiation tracks, and have distinct oxygen isotopic compos- itions (Reid et al., 1970, Goswami and Macdougall, 1983; Leshin et al., 1997; Frank et al., 2011). These observations suggest late addition of such grains to an altered body that was originally composed exclusively of matrix. The high H2O, C, N, and presolar grains concen- trations of CIs are consistent with this hypothesis, along with the additional observation that primitive carbonaceous chondrite matrices have similar concentrations of these species (Alexander, 2005; Zanda et al., 2009). All these findings suggest that the presence of high temperature components is responsible for the volatility-related fractionation of all other chondrite groups with respect to CIs (Anders, 1964; Wasson and Chou, 1974; Wai and Wasson, 1977; Palme et al., 1988; Humayun and Cassen, 2000; Palme, 2001; Bland et al., 2005; and Figure 5.1). Other chondrite groups also differ from CI chondrites in being fractionated in terms of Fe/Si ratio (Larimer and Wasson, 1988; Sears and Dodd, 1988; Palme and Lodders, 2009; Palme et al., 2014). Because terrestrial planets also differ from one another in terms of their volatile element budget and Fe/Si ratios (e.g., Palme, 2000), understanding the nature of elemental fractionation between chondrite groups may provide clues as to the genesis of these planets because similar mechanisms may have been at work. Chondrules and matrix may have formed independently and been mixed in random propor- tion in the different chondrite groups as suggested by Anders (1964). However, the discovery of chemical relations between chondrules and matrix within a given chondrite has led several authors (Wood, 1963, 1985; Bland et al., 2005; Hezel and Palme, 2010; Ebel et al., 2016) to suggest these components actually formed simultaneously from the same reservoirs, the com- positions of which would be chondritic (CI) in terms of major elements such as Si and Mg (Wood, 1963, 1985; Hezel and Palme, 2010) and to various degrees depleted in volatile elements depending on the chondrite group (Bland et al., 2005) as shown in Figure 5.2a. Because obtaining a reliable and representative bulk composition of chondrules within a chondrite is difficult, Bland et al. (2005) compared bulk and matrix compositions in carbonaceous 124 Brigitte Zanda, Éric Lewin, and Munir Humayun

Figure 5.1 Chondrite and Orgueil normalized concentrations of major and minor elements in the main carbonaceous chondrite groups. The right panel lists 50% equilibrium condensation temperature from Lodders (2003). Data from Wolf and Palme (2001) and Wasson and Kallemeyn (1988). After Palme (2001). (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.) chondrites (CCs). They showed that matrix and bulk mimic one another: if the bulk is highly depleted in volatiles (CVs, Vigarano-type carbonaceous chondrites), so is the matrix, but to a lesser extent. In contrast, when the bulk is moderately depleted (CMs, Mighei-type carbonaceous chondrites), then matrix is even less so. Bland et al. (2005) concluded this pattern indicated that chondrules and matrices had formed simultaneously from the same reservoir specific to each chondrite group. Looking at major and some minor elements, Hezel and Palme (2010) and additional work by this group (Palme et al., 1992; Hezel and Palme, 2008; Palme et al., 2015) argue that all CCs exhibit CI chondritic ratios of the major elements Si, Mg, and Fe and minor element Cr, and of the refractory elements Ca, Al, and Ti, but that each group has a specific matrix composition, with lower Mg/Si, Mg/Fe, Cr/Fe, and Ti/Al ratios than those of the bulk. For Ca/Al, they show that Allende has a superchondritic ratio in the matrix and Y-86751 a subchondritic ratio, while both CVs have chondritic bulk ratios (Hezel and Palme, 2008). These authors hence suggest that in each CC group, the composition and proportion of chondrules are exactly suited to those of their embedding matrix to add up to a chondritic bulk. They make the case that elemental distributions cannot result from parent body processes, and that chondrules and matrix must therefore have been formed simultaneously from a single reservoir of CI composition in terms of Mg, Fe, Si, and Cr ratios, and in terms of Ca, Al, and Ti ratios. However, Zanda et al. (2012) showed chemical redistribution between chondrules and matrix to have taken place in the course of parent body alteration in the Paris CM chondrite, which would explain the similarity in volatile depletion patterns for matrix and bulk in CCs. Matrix, originally CI in composition, would have lost some of its volatiles to the chondrules, resulting in a volatile depletion, moderate in the cases where the matrix comprises most of the rock (CMs) but much more pronounced in the cases where matrix is significantly less abundant than chondrules (CVs and COs – carbonaceous chondrites of the Ornans type). It is the aim of the present paper to discuss the case of the major elements based on data from Hezel and Palme (2010), subsequent The Chondritic Assemblage 125

Figure 5.2 Si and Orgueil normalized concentration plots as a function of 50% equilibrium condensation temperature from Lodders (2003). Orgueil from Palme et al. (2014). Except for panel (a) and matrix in 126 Brigitte Zanda, Éric Lewin, and Munir Humayun work by these authors (Palme et al., 2015) and additional literature data (see Table 5.1 for a full reference list). We will also discuss recent isotopic work by Budde et al. (2016a, 2016b), who found different nucleosynthetic isotopic signatures in W and Mo in chondrules and matrix in Allende and argue for isotopic complementarity between these components.

5.2 Methods

We mined the literature for comprehensive data on the chemical compositions of bulk carbon- aceous chondrites and of their main components, chondrules and matrix. Table 5.1 lists the data to which we had access, as well as matrix modal abundances also from the literature. This table shows that data for bulk rock, chondrule and matrix compositions together are only available for a handful of CCs: three CVs (Allende, Efremovka, Mokoia), one CO (Kainsaz) and one CR (Renazzo). In addition, we performed rastered beam electron microprobe (EMP) analyses of the Orgueil CI chondrite. These analyses were performed on the Cameca SX100 electron microp- robe at the Université Paris VI using 15keV, 10nA and a 10µm raster size, and quantification used the ZAF method. Sodium was analyzed first to prevent its loss, and both natural and synthetic standards were used.

5.3 Results

Table 5.2 lists all the chemical data used in the present study, except for individual chondrules or matrix points from the same source that were averaged. The data are expressed in wt% for the major elements (Mg, Fe, and Si) and some minor elements (Al, Ti, Ca, Cr, Mn, and Na).

Caption for Figure 5.2 (cont.) panel (b), all data and data sources are listed in Table 5.2. (a) the “ideal” CM chondrite as the basis for the “complementarity” hypothesis. It is fractionated in terms of moderately volatile and volatile elements (50% condensation T < 1,250K) as in Bland et al. (2005), but chondritic in terms of Mg, Fe, Si, and Cr ratios, and in terms Ca, Al, and Ti ratios as in Palme et al. (1992), Hezel and Palme (2008), Hezel and Palme (2010), Palme et al., (2015) and Ebel et al. (2016). (b) Solid line is the average chondrule of CM chondrite El-Quss Abu Saïd from Hezel and Palme (2010). Dotted line is the calculated composition of the complementary matrix to make up the “ideal” CM chondrite of panel (a), using a modal abundance of 70% for the matrix. Chondrule and matrix compositions smoothly follow the bulk symmetrically in the volatile element half of the diagram. In the major and refractory element half, the enrichments and depletions of the chondrule pattern impose symmetrical depletions and enrichments on the calculated matrix pattern. (c) Bulk chemical compositions of the 4 carbonaceous chondrites of different groups considered in the present paper as analyzed by wet chemistry. (d) Bulk chemical compositions of 3 of the 4 carbonaceous chondrites of different groups considered in the present paper as analyzed by XRF. (e) Bulk chemical compositions of 3 carbonaceous chondrites as calculated by using the composition of the weighted average Mokoia chondrule of Jones and Schilk (2009) and adding the listed amount of CI material (Palme et al., 2014) and metal containing 93 wt% Fe and 0.2 wt% Cr (see Table 5.2). The three compositions roughly correspond to those of Murchison, Kainsaz, and Efremovka. Kainsaz is a CO chondrite that contains an exceptionally high amount of metal (Diakonova et al., 1979); the matrix content of our model chondrite was adjusted to 50% to reproduce its abundance of volatile elements. Renazzo matrix abundance is very close to that of Efremovka. The corresponding pattern is not shown as it would be indistinguishable. Values are listed in Table 5.2. Table 5.1 Carbonaceous chondrites for which chemical data for bulk rock, matrix, and chondrules are available, and source references for these data. The last column lists matrix proportions and the references from which they were extracted.

Matrix Chondrules Matrix Bulk proportion CI Orgueil M&R77 (EMP); Z93 (100) (EMP); W&K88; A&G89; P&B93; W&P01; B+al12 (no Si); P +al14. Generic CI W&P01 (100)

CM Cold Bokkeveld M&R77 W&P01, W56 74.2 [M79] El Quss Abu Saïd H&P10 (11 chd) H&P10 (10 sel. pts) Mighei M&R77, Z93 W56 60.7 [M79] Murchison M&R77; (J90–WC) W&P01; J90 63.6 [M79] Murray M&R77, Z93 W&P01, W56 58.8 [M79] Nogoya M&R77, Z93 W&P01, W56 85.4 [M79] Average CM W&P01; W&K88

CV Allende K&I98 (av. chd) K&I98 (av. mtx); W&P01; J90, W56, H83 38.4 [M77b]; M&R77 53.7 [E+al16]* Bali M&R77 W&P01; J90 50.0 [M77b]; 42.6 127 [E+al16]* 128

Table 5.1 (cont.)

Matrix Chondrules Matrix Bulk proportion

Efremovka H&P10 (8 chd) H&P10 (10 sel. pts); W&P01; J90 35.7 [M77b] M&R77 Kaba K&I98 (13 chd) M&R77, Z93; K&I98 50.7 [M77b] (av. mtx); K01 Leoville K&I98 (av. chd) K&I98 (av. mtx); 35.1 [M77b]; M&R77 28.3 [E+al16]* Mokoïa K&I98 (13 ch); K&I98 (av. mtx); P+al15 W56 39.8 [M77b]; J&S09 (INAA 94 chd (150 pts, from K01); 50.7 [E+al16]* - no SiO2) M&R77 Vigarano M&R77, Z93; K01 W&P01, W56 34.5 [M77b]; 36.6 [E+al16]* Average CV W&P01; W&K88

CO Kainsaz H&P10 (10 chd) H&P10 (10 sel. pts); W&P01, D79 30.0 [M77a]; M&R77; G&B05. 29.9 [E+al16]* Ornans M&R77, R&W88 W&P01, W56 40.1 [M77a]; 41.8 [E+al16]* Average CO W&P01; W&K88

CR Al Rais M&R77 56.9 [M77b] LAP 02342 W&R09 (18 pts) MET 00426 A&B10 (5 pts) QUE 99177 A&B10 (3 pts) Renazzo H&P10 (9 chd); K01; H&P10 (10 sel. pts); M&W62 31.1 [M77b] P+al15 (73 chd) M&R77, Z93; K01; K&P99 (2 pts)

Acfer 94 W&R10 (10 pts) W&R10 (av. value)

Alha77307 B93

A&B10: Abreu and Brearley (2010); A&G89: Anders and Grevesse (1989); B+al12: Barrat et al. (2012); B93: Brearley (1993); D79: Diakonova et al., (1979); E+al16: Ebel et al. (2016); G&B05: Grossman and Brearley (2005); H&P10: Hezel and Palme (2010); H83: Haramura et al. (1983); J&S09: Jones and Schilk (2009); J90: Jarosewich (1990) + Jarosewich (2006); K&I98: Kimura and Ikeda (1998); K&P99: Kong and Palme (1999); K01: Klerner (2001); M&R77: McSween and Richardson (1977); M&W62: Mason and Wiik (1962); M77a: McSween (1977a); M77b: McSween (1977b); M79: McSween (1979); P&B93: Palme and Beer (1993); P+al14: Palme et al., (2014); P+al15: Palme et al. (2015); R&W88: Rubin and Wasson (1988); W&K88: Wasson and Kallemeyn (1988); W&P01: Wolf and Palme (2001); W&R09: Wasson and Rubin (2009); W&R10: Wasson and Rubin (2010); W56: Wiik (1956); Z93: Zolensky et al. (1993). av.: average; chd: chondrule; EMP: electron microprobe; INAA: instrumental neutron activation analysis; mtx: matrix; pts: points; sel.: selected; wc: wet chemistry; XRF: X-ray fluorescence. * Matrix proportions as listed in E+al16 includes lithic fragments, which are not included in M77a and M77b matrix proportions. To make the figures listed here more comparable, we estimated the proportions of lithic fragments based on (lithic + mineral fragments [M77a or M77b] – isolated olivines [E+al16]). The estimated proportions of lithic fragments were substracted from E+al16 figures. Except otherwise specified, all chondrule and matrix analyses are performed by EMP; bulk by wet chemistry or XRF. 129 130

Table 5.2 Summary of data used in the present study (wt% elt) and source references. Individual data points for chondrules and matrices are not listed, only their averages. Bulk compositions calculated based on chondrule and matrix compositions and proportions are also listed. For each of these calculations, the source for chondrule and matrix compositions is listed in column 2. Matrix proportions used for these calculations are listed in the table footnotes and “chondrule” proportions are calculated as (100-matrix volume %). This implies that “chondrule” stands for “high temperature fraction” (i.e., including refractory inclusions) as is done when comparing chondrules and matrices in references such as Bland et al. (2005), Hezel and Palme (2010), and Palme et al. (2015). Results from calculation adding various proportions (by weight) of CI material (Palme et al., 2014), weighted average Mokoia chondrule from Jones and Schilk (2009) and metal containing 93 wt% Fe and 0.2 wt% Cr are listed last (shown in Figure 5.2e).

Element Condensation T (L03) Al 1653 Ti 1582 Ca 1517 Mg 1336 Fe 1334 Si 1310 Cr 1296 Mn 1158 Na 958

Meteorite Reference{

CI Orgueil W&P01 0.83 0.05 0.91 9.60 18.69 10.69 0.26 0.19 Orgueil W&K88 0.86 0.04 0.92 9.70 18.20 10.50 0.27 0.19 Orgueil A&G89 0.87 0.04 0.90 9.53 18.51 10.67 0.27 0.20 0.49 Orgueil P&B93 0.86 0.04 0.95 9.61 18.23 10.68 0.27 0.19 Orgueil B+al12 0.79 0.04 0.84 9.42 19.52 10.52 0.26 0.19 0.48 Orgueil P+al14 0.84 0.04 0.91 9.54 18.66 10.70 0.26 0.19 0.50 Orgueil M&R77 (EMP) 1.22 0.02 0.11 11.64 14.07 13.88 0.29 0.12 0.26 Orgueil Z93 (EMP) 1.36 0.05 0.21 11.93 17.16 15.42 0.40 0.20 0.04 Orgueil This work (EMP) 1.18 0.05 0.43 11.37 16.80 14.75 0.38 0.22 0.42

CI average W&P01 W&P01 0.83 0.05 0.91 9.60 18.69 10.69 0.26 0.19

CM El-Quss Abu Said av. chondrule H&P10 0.77 0.14 1.46 25.99 1.61 23.55 0.36 0.10 0.02 El-Quss Abu Said matrix H&P10 1.65 0.05 1.98 8.12 27.90 12.06 0.23 0.18 0.56

Nogoya matrix M&R77 1.28 0.05 0.41 11.22 20.52 12.48 0.25 0.15 0.16 Nogoya matrix Z93 1.15 0.03 0.20 11.46 21.48 13.54 0.23 0.15 0.27 Nogoya bulk F1888 1.24 – 1.80 11.49 21.46 12.70 0.24 0.05 0.13 Nogoya bulk W&P01 1.17 0.06 1.25 11.16 19.82 12.60 0.28 0.17 Nogoya calculated bulk M&R77 mtx (H&P10 1.22 0.06 0.54 13.06 18.16 13.86 0.26 0.14 0.14 chd)

Murchison matrix M&R77 1.76 0.02 0.61 8.44 25.96 10.42 0.21 0.14 0.23 Murchison matrix Z93 1.57 0.04 0.54 9.20 25.99 12.73 0.25 0.15 0.27 Murchison bulk J90 1.14 0.08 1.35 12.02 22.13 13.59 0.33 0.15 0.18 Murchison bulk W&P01 1.16 0.07 1.28 12.14 21.40 13.48 0.31 0.17 Murchison calculated bulk M&R77 mtx (H&P10 1.40 0.06 0.91 14.76 17.19 15.15 0.27 0.13 0.15 chd) 64% CI + 4% Fe–Ni + 32% P+al14, J&S09* 1.28 0.08 1.31 12.37 19.83 13.84 0.29 0.16 0.44 Mokoia chd

CM average W&P01 W&P01 1.15 0.06 1.25 11.66 20.57 13.01 0.31 0.17

CV Efremovka av. chondrule H&P10 1.55 0.10 1.91 25.93 1.39 23.21 0.25 0.07 0.23 Efremovka matrix H&P10 1.11 0.04 1.28 11.73 29.72 15.59 0.31 0.26 0.24 Efremovka matrix M&R77 0.85 0.05 0.59 12.00 31.56 14.54 0.35 0.23 0.24 Efremovka bulk J90 1.78 0.10 1.90 14.90 22.49 16.04 0.38 0.15 0.19 Efremovka bulk W&P01 1.67 0.10 1.66 15.14 21.00 16.58 0.36 0.14 Efremovka calculated bulk H&P10 mtx & chd 1.38 0.08 1.67 20.49 12.24 20.29 0.28 0.14 0.23 38.3% CI + 5% Fe–Ni + 56.7% P+al14, J&S09* 1.63 0.11 1.63 14.75 19.18 16.49 0.31 0.14 0.40 Mokoia chd

Mokoia av. chondrule (INAA)* J&S09* 2.31 0.16 2.26 19.57 13.02 21.86* 0.35 0.11 0.36 Mokoia av. chondrule K&I98 2.3 0.1 2.0 21.5 6.9 19.0 0.5 0.1 0.3

131 Mokoia matrix K&I98 1.3 1.4 10.3 26.3 13.2 0.2 0.2 0.3 Mokoia matrix P+al15 1.61 0.10 1.35 11.01 28.19 13.52 0.23 0.18 0.30 132 Table 5.2 (cont.)

Element Condensation T (L03) Al 1653 Ti 1582 Ca 1517 Mg 1336 Fe 1334 Si 1310 Cr 1296 Mn 1158 Na 958

Meteorite Reference{ Mokoia matrix M&R77 1.86 0.04 1.71 10.19 21.61 12.39 0.21 0.08 0.20 Mokia bulk W56 1.33 0.06 1.83 14.46 24.05 15.61 0.36 0.15 0.38 Mokoia calculated bulk P+al15 mtx, J&S09 2.02 0.13 1.89 16.05 19.27 18.42* 0.30 0.14 0.34 chd* Allende av. chondrule K&I98 2.80 0.18 2.29 20.62 7.31 18.84 0.55 0.08 0.96 Separated chondrules J90 (WC) 2.95 0.16 2.90 20.71 8.47 19.57 0.32 0.11 0.61 Allende matrix K&I98 1.11 1.14 11.46 27.91 13.04 0.27 0.15 0.22 Allende matrix M&R77 1.22 0.05 1.69 12.18 24.80 13.09 0.26 0.16 0.16 Allende matrix Z93 0.89 0.05 0.16 12.45 28.65 13.82 0.34 0.15 0.04 Separated matrix J90 (WC) 1.62 0.08 1.91 12.92 26.96 15.48 0.38 0.17 0.33 Allende bulk J90 1.73 0.09 1.87 14.85 23.83 16.00 0.36 0.14 0.33 Allende bulk W&P01 1.72 0.09 1.87 14.94 23.60 15.90 0.35 0.15 Allende bulk H83 1.96 0.10 1.78 15.70 23.48 15.80 0.36 0.15 0.32 Allende bulk W56 1.82 0.08 1.84 15.47 23.16 15.94 0.40 0.15 0.35 Allende calculated bulk K&I98 mtx & chd 2.13 1.83 16.99 15.46 16.54 0.44 0.11 0.67 Allende calculated bulk M&R77 mtx, K&I98 2.18 0.13 2.05 17.28 14.23 16.56 0.43 0.11 0.65 chd CV3 average W&P01 1.62 0.09 1.76 14.72 22.55 15.80 0.35 0.14

CO Kainsaz av. chondrule H&P10 1.71 0.12 2.61 22.80 6.80 21.35 0.25 0.12 0.47 Kainsaz matrix H&P10 1.53 0.04 0.62 10.68 24.32 13.27 0.25 0.25 0.62 Kainsaz matrix M&R77 1.90 0.06 1.37 10.85 24.72 14.30 0.24 0.26 0.36 Kainsaz bulk D79 1.59 0.08 1.79 14.91 26.73 16.21 0.49 0.23 Kainsaz bulk W&P01 1.42 0.08 1.54 13.89 23.99 15.38 0.34 0.16 0.43 Kainsaz calculated bulk H&P10 mtx & chd 1.65 0.09 1.95 18.80 12.58 18.68 0.25 0.16 0.52 50% CI + 7% Fe–Ni + 43% P+al14, J&S09* 1.41 0.09 1.43 13.19 21.44 14.75 0.29 0.14 0.41 Mokoia chd

CO3 average W&P01 W&P01 1.37 0.08 1.53 13.77 23.98 15.41 0.34 0.16

CR Renazzo av. chondrule H&P10 1.65 0.11 2.18 23.35 2.04 23.86 0.50 0.16 0.07 Renazzo av. chondrule P+al15 1.20 0.09 1.35 22.50 5.72 22.27 0.49 0.22 Renazzo matrix H&P10 1.31 0.02 0.52 9.94 18.55 15.27 0.23 0.18 0.99 Renazzo matrix M&R77 1.41 0.04 0.62 9.53 18.89 14.68 0.24 0.26 0.86 Renazzo matrix Z93 1.62 0.04 0.44 10.03 21.35 16.18 0.21 0.16 0.58 Renazzo bulk M&W62 1.25 0.11 1.27 14.33 24.93 15.81 0.38 0.19 0.41 Renazzo calculated bulk H&P10 mtx & chd 1.53 0.08 1.60 18.64 7.84 20.84 0.41 0.17 0.39 35.1% CI + 9% Fe–Ni + 55% P+al14, J&S09* 1.59 0.10 1.59 14.29 22.20 15.97 0.31 0.13 0.38 Mokoia chd

A&G89: Anders and Grevesse (1989); B+al12: Barrat et al. (2012); D79: Diakonova et al., (1979); F1888: Friedheim (1888); H&P10: Hezel and Palme (2010); H83: Haramura et al. (1983); J&S09: Jones and Schilk (2009); J90: Jarosewich (1990) + Jarosewich (2006); K&I98: Kimura and Ikeda (1998); L03: Lodders, 2003; M&R77: McSween and Richardson (1977); M&W62: Mason and Wiik (1962); M77a: McSween (1977a); M77b: McSween (1977b); M79: McSween (1979); P&B93: Palme and Beer (1993); P+al14: Palme et al. (2014); P+al15: Palme et al. (2015); W&K88: Wasson & Kallemeyn (1988); W&P01: Wolf and Palme (2001); W56: Wiik (1956); Z93: Zolensky et al., (1993). adj.: matrix proportion in quoted reference has been adjusted so that sum of listed components for the chondrite ad up to 100%; chd: chondrule; mtx: matrix; WC: wet chemistry. { Matrix proportions used for calculated bulks: Nogoya, 87.5% [M79 adj.]; Murchison, 64% [M79 adj.]; Efremovka, 38.3% [M77b adj.]; Mokoia, 41.2% [M77b adj.]; Allende, 39.6% [M77b adj.]; Kainsaz, 33% [M77a adj.]; Renazzo, 35.1% [M77b adj.]. *weighted average of the data of Jones and Schilk (2009). A chondritic Si/Mg of 1.12 ratio was used. Except otherwise specified, all chondrule and matrix analyses are performed by EMP; bulk by wet chemistry or XRF. Condensation T is 50% equilibrium condensation temperature from Lodders (2003). 133 134 Brigitte Zanda, Éric Lewin, and Munir Humayun

Table 5.2 also lists source references and the bulk compositions calculated for each chondrite based on the compositions of their chondrules and matrix, the matrix modal abundance from Table 5.1 and the matching chondrule fraction. In doing so, we assume that modal abundances for chondrules and matrices are equivalent to weight fractions, which is probably not strictly true but should be unimportant in comparison to errors arising from the analytical methods (typically a few tens of percent on EMP data as discussed in the following text). When normalized, the plots that make up the figures are consistently normalized to Si and CI abundances from Palme et al. (2014), except for Figure 5.1 (normalized to Al, the most refractory element of the set) and for Figure 5.3a (normalized to Mg, as Si was not measured by Instrumental Neutron Activation Analysis, INAA).

5.3.1 A Comparison of Results Obtained with Different Analytical Methods Figure 5.3a compares the average bulk composition of chondrules in Mokoia as measured by INAA (Jones and Schilk, 2009 – we calculated the weighted average of their data) and by EMP (Kimura and Ikeda, 1998). Data from the two analytical methods are consistent within 20 percent except for Fe, for which the measurement by EMP is lower by a factor of 2 compared to INAA (Table 5.2), and for Cr, for which it is higher by a factor of 1.5. The difference in analytical results between in situ and bulk methods is even more striking when defocused beam

2 1. 8 Mokoia Cl Av. chondrule INAA 1. 6 Av. chondrule EMP 1. 4 1. 2 1 0.8 0.6 0.4 0.2 (a) 0 2.0 1. 8 Orgueil P+al14 (bulk) This work (EMP) 1. 6 1. 4 1. 2 1. 0 0.8 0.6 0.4 X/Si / in Orgueil X/Mg / in Orgueil 0.2 (b) 0 1650 1550 1450 1350 1250 1150 1050 950 Condensation Temperature AlTi Ca Mg Si Mn Na Fe Cr

Figure 5.3 A comparison of analytical results from EMP and bulk techniques in Si and Orgueil normalized concentration plots. (a) weighted average of Mokoïa chondrules analyzed by INAA (Jones and Shilk, 2009) and EMP average data (Kimura and Ikeda, 1998); (b) bulk Orgueil (Palme et al., 2014) and Orgueil analyzed by EMP with a rastered beam (this work). The Chondritic Assemblage 135

Figure 5.4 Mg and Si concentrations of Orgueil, analyzed by EMP with a rastered beam (open triangles). The average is shown by a large triangle, and the corresponding ratio by a thin black dotted line. The CI bulk concentrations as analyzed by wet chemistry or XRF is shown as a large gray circle and the corresponding ratio by a thick gray dotted line. When analyzed by broad beam EMP, Orgueil appears not to be chondritic. EMP data of matrices from Hezel and Palme (2010) are also plotted (dark stars on gray background) and their average as a larger point. Matrix ratios when analyzed with EMP do coincide with Orgueil EMP measured ratio. However, mineralogies differ and it is not possible to simply apply a correction based on the shift of Orgueil EMP measurements to matrix EMP measurements, hence a need for better quality in situ measurements (Zanetta et al., 2017). See Section 4.1 for a discussion of this discrepancy.

EMP measurements of Orgueil are compared with bulk analytical measurements (Tables 5.2 and 5.3; Figures 5.3b and 5.4). Bulk analytical measurements of Orgueil (lines 1–5 of Table 5.1 and Table 5.2) are internally consistent within a few percent (also see Figure 5.2c and 5.2d for other carbonaceous chondrites). EMP data (lines 6–8 of Tables 5.2 and 5.3) are also internally consistent, although they do exhibit more variability (up to a few tens of a percent). However, the differences between the bulk and the in situ measurements are much larger than the uncertainties (twice the spread of the bulk measurements) as shown from Figure 5.3b and Table 5.2, even where the major elements Si, Mg, and Fe are concerned (Mg/Si as determined by EMP is lower than the bulk measurement by 20 percent; Fe/Si by 40 percent). In fact, when analyzed with defocused beam EMP, Orgueil appears not to be chondritic. This observation shows that caution needs to be exercised when data obtained from different analytical methods are used as a basis for understanding chondritic compositions.

5.3.2 Chondrule and Matrix Compositions in Different Carbonaceous Chondrite Groups Figure 5.5 displays individual chondrule and matrix point compositions from Appendices A1 and A2 as published by Hezel and Palme (2010) for Efremovka (CV), Kainsaz (CO), El-Quss Abu Said (CM), and Renazzo (CR). Chondrule compositions as analyzed by EMP are almost uniformly depleted in Fe and in volatile elements (Cr, Mn, and Na), with the exception of the 136 Brigitte Zanda, Éric Lewin, and Munir Humayun

Table 5.3 A comparison of Si/Mg and Fe/Mg for Orgueil as measured in situ by EMP and by bulk methods (XRF, INAA and wet chemistry).

Si/Mg Fe/Mg

Electron probe: McSween & Richardson (1977) 1.19 1.21 Zolensky et al. (1993) 1.29 1.44 This work 1.30 1.48 Bulk analysis: Wolf and Palme (2001) 1.11 1.95 Wasson and Kallemeyn (1988) 1.08 1.88 Anders and Grevesse (1989) 1.12 1.94 Palme and Beer (1993) 1.11 1.90 Lodders et al. (2009) 1.12 1.96 Barrat et al. (2012) 1.12* 2.07

* Si/Mg ratio from Barrat et al. (2012) is derived from Lodders et al. (2009).

CO data. Refractory minor elements (Al, Ti, Ca) spread around the chondritic value. These patterns are fairly similar in all chondrite groups as shown by the bottom left panel of Figure 5.5, in which the averages for each group are plotted. They are also fairly similar to the “ideal bulk chondrite” as plotted in Figure 5.2a. Matrix patterns exhibit somewhat more spread, appearing enriched in Fe for El-Quss Abu Said, depleted in Fe for Renazzo, and straddling the chondritic Fe/Si value for Efremovka and Kainsaz. Matrix patterns are roughly chondritic for the volatile elements for El-Quss Abu Said and Kainsaz, depleted for Efremovka, and somewhat enriched (at least in Na) for Renazzo. They tend to be depleted in refractory minor elements, except for a peak in Ca in all analyses for El-Quss Abu Said and some analyses for Efremovka, a pattern most likely due to the presence of carbonates under the beam, possibly of terrestrial origin in the case of the El-Quss Abu Said (a find).

5.3.3 Reconstructing Bulk Chondrites from Their Chondrules and Matrices 5.3.3.1 Element Pairs: Si and Mg; Fe and Cr Hezel and Palme (2010) and Palme et al (2015) argue that Mg/Si and Cr/Fe are superchondritic in chondrules and subchondritic in matrices. Figure 5.6 displays their published data for Kainsaz and Renazzo as well as the CI bulk value (star) and ratio from Palme et al. (2014), after Lodders et al. (2009). Similar figures for El-Quss Abu Said and for Efremovka are not displayed here, as they do not differ significantly from Kainsaz and Renazzo for the Si-Mg pair and from Kainsaz for the Fe-Cr pair. These diagrams show that Mg/Si and Cr/Fe ratios of matrices are always close to bulk CI ratios, dispersed around it for Cr/Fe, and rather systematically lower than it for Mg/Si. Conversely, these ratios for chondrules exhibit more variability (standard deviation of ratio in chondrules is from 1.3 to 6.4 times greater than in matrices) and are quite systematically Figure 5.5 Si and Orgueil normalized concentration plots of individual CC chondrules and matrix points listed in Hezel and Palme (2010) as a function of 50% equilibrium condensation temperature from Lodders (2003). (a), (b): Efremovka; (c), (d): Kainsaz; (e), (f ): El-Quss Abu Saïd; (g), (h): Renazzo. (g) and (h) superimpose the averages for all these chondrites. Chondrules patterns exhibit a large range of variation, but their averages are fairly similar in the different carbonaceous chondrite representative of each group. Matrix patterns exhibit rather less variability within a given carbonaceous chondrite, but they tend to differ more from chondrite to chondrite. 138 Brigitte Zanda, Éric Lewin, and Munir Humayun

Figure 5.6 Mg-versus-Si and Cr-versus-Fe concentration data of Hezel and Palme (2010) for chondrules (circles) and matrices (diamonds) for Kainsaz and Renazzo. CI concentrations are shown with stars and the CI ratio with a line. The average for each group of concentrations can be computed (shown by a smaller black marker) and the relative amount of chondrules and matrix in a bulk with a CI ratio can be derived from the lever rule by measuring the distance between the chondrule and matrix averages and the intercept with the CI ratio line (small black star). The estimated relative amount of matrix is shown on each diagram and listed in Table 5.4 for all four carbonaceous chondrites considered in the present work. In Cr versus Fe (i) in Kainsaz: two chondrules fall below the CI line and skew the chondrule average towards CI; (ii) in Renazzo, as the average Cr/Fe ratio for the matrix is >CI ratio, the estimated amount of matrix is >100%. above bulk CI. However, for Mg/Si only, chondrule ratios are close to bulk CI, while for Cr/Fe, they are distinctly higher than bulk CI (mean values are higher than CI by a factor of 1.2 to 2.7 times their standard deviation). Chromium concentrations are undistinguishable in chondrules and in matrix, but Fe is comparatively very depleted in chondrules. This may only in part be attributed to the analytical technique: even if a factor of 2 is applied to bring the data in line with INAA measurements (see Section 5.3.1), Fe still remains depleted by a factor of 2 or more in chondrules compared to matrix. In such concentration diagrams, mixing is represented by straight lines and the lever rule may be applied in the framework of a two-component model such as the complementarity hypothesis of Hezel and Palme (2010), Palme et al. (2015), and Ebel et al. (2016). As long as the bulk chondrite has the CI ratio, the (Si, Mg) and (Cr, Fe) concentration pairs in chondrules and in matrix may then be used to estimate the relative (wt%) proportion of each of these compon- ents in the bulk mix. Results are displayed in Figure 5.6 and in Table 5.4. As average Si/Mg and The Chondritic Assemblage 139

Table 5.4 Wt% matrix in chondrite needed to generate a bulk CI Si/Mg or Cr/Fe ratio.

Si, Mg Fe, Cr Measured*

Efremovka 68 88 35.7 Kainsaz 73 82 30.0 El-Quss Abu Said 63 79 70{ Renazzo 32 103 31.1

* See Table 5.1 for references. The measured value is a modal abundance, hence expressed in vol%, not wt%. However, we hypothesize here that the density difference between chondrules, matrix, and bulk silicates is negligible compared to the analytical uncertainties. { Average value for CM chondrites, as there is no published measurement for El-Quss Abu Said.

Cr/Fe in the matrix are close to the bulk CI ratios, the estimated abundances of matrix are 60 wt%, much higher than the measured modal abundance, except probably for the El-Quss Abu Said CM (70 percent modal abundance of matrix is only an estimation). As the average Cr/Fe ratio of Renazzo matrix is almost equal to the bulk CI ratio but slightly above it, the estimated matrix abundance based on this element pair is 103 percent. On the other hand, the estimated matrix abundance of Renazzo based on the (Si, Mg) element pair is 32 percent, close to the 31.1 percent measured value (Table 5.1), because the average Si/Mg ratio of chondrules is also close to the CI bulk ratio (Figure 5.6).

5.3.3.2 Calculating Bulk Compositions Based on Matrix and Chondrule Compositions and Relative Proportions The bulk compositions of chondrites can be estimated based on the compositions of their matrices and chondrules as listed in Table 5.2 using the matrix modal abundances listed in Table 5.1. Results are listed in Table 5.2 and displayed in Figure 5.7 for Efremovka and Renazzo. For Efremovka, the calculated Si and CI-normalized element ratios differ from those measured by bulk chemical methods by about 20–40 percent depending on which bulk measurement is used, except for Mg (~10 percent excess in the calculated bulk) and Fe (50 percent depletion in the calculated bulk). As chondrules contribute very little to the Fe budget of the rock (Figure 5.7a), the difference between the estimated CI-normalized Fe/Si and the measured one (Figure 5.7b) cannot be attributed only to the chondrules being analyzed by EMP rather than by a bulk method such as INAA: even if the normalized Fe/Si ratio of the chondrule component is increased by a factor of 2 as discussed in the previous text (Section 5.3.1), the calculated normalized bulk Fe/Si ratio still falls short at about 50 percent of the measured one. This suggests that chondrules and matrix as measured in the present data set cannot make up the bulk composition of the rock and that there may be an analytical problem on Fe due to analyzing matrix with EMP (as suggested by Table 5.3), or that another component (Fe, Ni) contributes to the bulk, as will be discussed in Section 5.4. For Kainsaz, the patterns (not shown) are very similar to those of Efremovka. The agreement between calculated and measured Si and Orgueil normalized bulk compositions are better for most elements (10 percent) except for Mn (15–40 percent), Cr (40–55 percent) and Fe (57–59 percent) depending 140 Brigitte Zanda, Éric Lewin, and Munir Humayun

Figure 5.7 Si and Orgueil normalized plots of concentration data from Hezel and Palme (2010) for the average of chondrules and of matrix in (a) Efremovka (CV) and (c) Renazzo (CR). The calculated bulk compositions based on these chondrule and matrix compositions and on the modal abundances listed in Table 5.1 are shown as a light gray thick line. In (b) and (d), this calculated bulk composition is compared to the bulk chemical compositions of the two chondrites (Efremovka: bulk XRF data of Wolf and Palme, 2001; Renazzo: bulk wet chemical analysis of Mason and Wiik, 1962).

on which bulk composition is used (Wolf and Palme, 2001 or Diakonova et al., 1979). In the case of El-Quss Abu Said, there is no published bulk composition, but because of the similarities in chondrule compositions within a given group (see Section 5.3.2 and Figure 5.5) chondrules from El-Quss Abu Said can be used as a proxy for chondrules in other CMs for which bulk and matrix compositions as well as matrix modal abundances are available (see Table 5.1). Elemental abundance patterns (not shown) are similar to those for Kainsaz and Efremovka. For Renazzo, no bulk composition measured by X-ray fluorescence analysis is available, only a chemical composition published by Mason and Wiik (1962). The patterns are also very similar, except that the Fe discrepancy is even higher (76 percent) as can be seen from Figure 5.7d.

5.4 Discussion

5.4.1 Reconstructing Bulk Chondrite Compositions from Those of Their Primary Components As shown above in Section 5.3.3.1 and Table 5.4, matrix abundances derived from two different element pairs in several carbonaceous chondrites do not generally agree with one another and are inconsistent with measured modal abundances. In addition, bulk chondrite compositions cannot be satisfactorily reproduced from the addition of chondrules and matrix in the right The Chondritic Assemblage 141 proportion and their measured compositions (Section 5.3.3.2). Two main reasons can be invoked and are discussed in the following two subsections.

5.4.1.1 Comparing Data Obtained with Different Analytical Methods As shown in Section 5.3.1, in Tables 5.2 and 5.3, and in Figures 5.3 and 5.4, compositional data acquired by defocused beam EMP cannot be directly compared to data obtained by bulk analytical methods. As pointed out by Lux, Keil, and Taylor (1980), Jones and Scott (1989) and Warren (1997), this is mostly due to differences in the densities of phases under the beam. This is particularly true for phases that contain Fe, which are denser than the surrounding ones, resulting in an underestimation of Fe. Jones and Scott (1989) applied sulfide or Fe–Ni versus silicate corrections when measuring bulk chondrule compositions by defocused beam EMP. Indeed, Fe analyzed by EMP is lower than with other analytical methods: by 50 percent for Fe of Mokoia chondrules compared to INAA data and by 20 percent for Fe in Orgueil compared to bulk data. There is also an effect, albeit smaller, on the Si/Mg ratio. Figure 5.4 and Table 5.4 show that Si/Mg in Orgueil measured by EMP is 10–20 percent higher than by bulk methods. As the mineral assemblage in Orgueil is not identical to that of carbonaceous chondrite matrices, it is not known whether the difference between matrix compositions as measured by EMP and measured by bulk methods would be identical. Yet, it is interesting to note that Si/Mg of Orgueil as measured by EMP is identical to that of EMP-measured matrices of Hezel and Palme (2010). Hence, it is possible (but not demonstrated) that the Si/Mg ratio of carbonaceous chondrite matrices may be CI chondritic. This sheds doubt on the complementarity notion for Mg and Si as advocated by Hezel and Palme (2010) as it is based on the bulk chondrite being chondritic while the chondrules and the matrix sit respectively above and below the CI ratio line. However, in Figure 1 of their work, the bulk is analyzed by bulk chemical methods while the chondrules and matrix that are compared with it are analyzed by defocused beam EMP. If the Si/Mg ratio of the matrix were in fact chondritic as that of the bulk, it follows that chondrules should be too – a hypothesis that is not inconsistent with the dispersed data for chondrules shown in Figure 5.5.

5.4.1.2 A Model with Only Two Components Does Not Adequately Represent Chondrites Carbonaceous chondrites are not made exclusively of chondrules and matrix. In addition to these components, they also contain different proportions of refractory inclusions and opaque minerals (e.g., McSween, 1977a, 1977b, 1979; Ebel et al., 2016). Refractory inclusions account for ~10–12 percent of the volume of the rock in CVs and COs according to McSween (1977a, 1977b), or about half that amount according to Ebel et al. (2016) and Hezel et al. (2008). They contain ~20 wt% Al and Ca (Grossman et al., 2000), and hence might contribute by up to ~1 wt% Al and Ca to the total element budget of the rock. That figure is comparable to the bulk Al and Ca concentrations of CVs and COs as listed in Table 5.2, which would imply that a significant fraction of Al and Ca in these rocks is contained in refractory inclusions. It should, however, be observed that this is at odds with the results of Ebel et al. (2016), which indicate that most of these elements reside in the matrices. Neglecting the contribution of refractory inclusions in Ca, Al, and Ti inventories when calculating a bulk composition based only on chondrules and matrices may in part explain the discrepancy between calculated and measured bulk abundances for these elements. Similarly, neglecting Fe–Ni metal, troilite (FeS), and 142 Brigitte Zanda, Éric Lewin, and Munir Humayun magnetite (FeFe2O4), which together represent a few volume percent of the chondrites (McSween, 1977a, 1977b), or possibly less (Ebel et al., 2016), may also affect the calculated Fe budget.

5.4.1.3 Reconstructing the Bulk Composition of Chondrites The bulk compositions of carbonaceous chondrites for major and minor elements (Table 5.2 and Figure 5.2c and 5.2d) can be adequately reproduced by the simple mixing of one generic chondrule composition with the appropriate balance of CI composition matrix material with the addition of minor amounts of Cr-bearing Fe–Ni metal and refractory material. Results for such a calculation are shown in Table 5.2 and in Figure 5.2e. In this calculation, we used the composition of the weighted average of Mokoia chondrules from Jones and Schilk (2009) for the generic chondrule composition. As Si was not determined in the Jones and Schilk (2009) data, we used a chondritic ratio of 1.12 after Lodders et al. (2009). As Mokoia chondrules are fairly rich in Ca, Al, and Ti, there was no need to add refractory material to the present calculation. However, for calculations performed with the average Renazzo chondrule, a minor addition of CAI-like material was required. The results in Table 5.2 and Figure 5.2e compare satisfactorily to the bulk XRF analyses of Wolf and Palme (2001) shown in Figure 5.2d – better than the compositions calculated based on the observed chondrule and matrix compositions listed in Table 5.2 and displayed in Figure 5.7. This demonstrates that explaining the bulk compositions of carbonaceous chondrites in the different groups does not require that their chondrules and matrices had different and complementary compositions at the time of accretion.

5.4.2 Was Complementarity between Chondrules and Matrix a Required Element of Chondrule Formation and Chondrite Accretion? If chondrules and matrices had complementary elemental patterns adding together to create the solar ratios, complementarity would be a very significant constraint on models of chondrule origin, since the solar ratios are a large-scale chemical feature of our solar system. It would impose severe constraints on any chondrule forming mechanisms, for example, excluding models such as the X-wind model (Shu, Shang, and Lee, 1996) since chondrules and matrix would have to be generated or processed in the same nebular reservoir. Clarifying whether the constraint of complementarity is required is therefore of critical importance to our understand- ing of chondrule and chondrite formation.

5.4.2.1 Complementarity in the Volatile-Element Range In the previous text, we have concentrated on the major and refractory elements advocated by Hezel and Palme (2010), Palme et al. (2015) and Ebel et al. (2016) to exhibit complementary patterns adding up to CI between matrix and chondrules as shown in the left part of the diagrams in Figure 5.2a and 5.2b. Bland et al. (2005) offered another definition of complementarity, dealing with the moderately volatile and volatile elements. As mentioned in the introductory section of this chapter, they measured the bulk and matrix compositions of a set of carbonaceous chondrites and showed that these patterns “mimic” one another in the different groups. However, chondrule compositions were not measured and their compositions were not The Chondritic Assemblage 143 evaluated based on a mass balance calculation. We performed such calculations based on our own LA-ICP-MS (Laser Ablation Inductively Coupled Plasma Mass Spectrometry) analyses of bulk carbonaceous chondrites and their matrices, and showed that the chondrule compositions would be indistinguishable from one group to another (Zanda, Humayun, and Hewins, 2012). In addition, it is likely that exchange between chondrules and matrix has taken place as a result of parent body processes (Zanda et al., 2011a, 2011b): compared to the least altered zones, matrix in the moderately altered zones of the Paris CM breccia gained Ca, V, Fe, and the siderophile elements (Ni, Co, Mo, Ru, Rh, Pd, W, and so on) from the destruction of glass and metal grains in altered chondrules, while it lost Na, S, and the chalcophile elements (Zn, As, Cu, Sb, Te, Tl, Pb, and so on) due to the destruction of its FeS and the ubiquitous formation of PCPs (Poorly Characterized Phases) in the moderately altered zones. Except for Acfer 094, all the matrices analyzed by Bland et al. (2005) belong to chondrites which underwent some degree of parent- body transformation: the CM2s experienced more extensive alteration than the moderately altered zones of Paris (based on a comparison of PCP compositions – Bourot-Denise et al., 2010); while the COs and CVs experienced thermal metamorphism. Hence the findings by Bland et al. (2005) that matrix compositions mimic that of the bulk are likely to be due to parent body exchange: if the chondrite comprises mostly matrix (CMs), its bulk is only moderately depleted in volatile elements and the effects on matrix compositions of exchange with a comparatively small amount of chondrules will have been slight, whereas in chondrites con- taining a much larger fraction of chondrules (CVs and COs) the bulk is much more volatile depleted and the matrix compositions will have been more severely affected by such exchange.

5.4.2.2 Complementarity in the Major and Refractory Element Range As shown by Figure 5.1, chondrites are fractionated with respect to CIs over the whole range of condensation temperatures, excluding the refractory elements. Every single element of the 12 shown in the diagram is fractionated with respect to Al (the most refractory of the 12), except for Ti and Ca, which are also refractory (both have 50 percent equilibrium condensation temperatures, Tcond >1,500 K – Lodders, 2003). Even the most refractory element of the other nine considered (V, with Tcond = 1,429 K) is depleted by ~5 percent in CMs, ~10 percent in COs, and ~20 percent in CVs. Two points are noteworthy in this diagram. (a) Mg and Si, which have very similar Tcond (1,336 K and 1,310 K, respectively) closely follow one another, which explains why they are always in a CI ratio in carbonaceous bulk compositions. The same is true for Al, Ti, and Ca: the ratios between any two of these elements will be identical to that of CIs. However, except for Fe and Ni (see the following text), no other pair of elements can be in a chondritic ratio in all CCs as required by the complementarity hypothesis supported by Hezel and Palme (2010), Palme et al. (2015) and Ebel et al. (2016). CI ratios between two elements will occur in cases in one or two CC groups (e.g., Fe,Ni, and P in CMs and CVs, or Fe, Ni, and Cr in COs), but most element pairs will never be found to be complementary to one another. For example, P/Cr, Fe/Mg, and Si/Ca can only be below the CI ratio in any CC group considered. This is true of most element pairs except for those that have very similar volatility: all CCs relative to CIs are fractionated as a result of volatility-related processes to an extent that varies between chondrite groups and is in first approximation related to the amount of matrix they contain (Zanda et al., 2009). As Si and Mg have very similar volatilities, they follow one 144 Brigitte Zanda, Éric Lewin, and Munir Humayun another closely and rarely fractionate from one another as a result of evaporative processes. This makes it likely that both matrix and chondrules overall had a CI Mg/Si ratio at the time of accretion, as suggested in Section 5.4.1. (b) Another noteworthy point is the position of Fe and Ni in this diagram. These two elements closely follow one another, but they do not plot where they would be expected to, based on their volatility alone (i.e., above Mg (Ni) or between Mg and Si (Fe)). This suggests that another process was at work in the fractionation of Fe and Ni: either the redox conditions dramatically enhanced their volatility (this appears unlikely for Ni) or, more likely, there was physical separation of metal and silicates, a process that has long been known to have been at play in the genesis of chondrite compositions (e.g., Rambaldi and Wasson, 1981, 1984; Grossman and Wasson, 1985; Palme and Lodders, 2009; Palme et al., 2014).

5.4.2.3 Isotopic Complementarity Three nucleosynthetic processes, the s-process, the r-process, and the p-process, produce W and Mo with distinctive isotopic patterns (Burkhardt et al., 2011; Burkhardt and Schönbächler, 2015). Budde et al. (2016a, 2016b) found different nucleosynthetic isotopic signatures in W and Mo in chondrules (with an s-process deficit) and matrix (with an s-process excess) in the Allende CV3. They point out that such anomalies must result from the uneven distribution between chondrules and matrices of a presolar carrier enriched in s-process Mo and W nuclides creating an isotopic “complementarity.” They argue that this uneven distribution of nucleosynthetic anomalies was established during chondrule formation, and speculate that it may be related to metal-silicate fractionation (Budde et al., 2016b). However, no convincing mechanism is offered to explain how chondrule formation might have generated such an uneven distribution of the presolar carrier(s). They contend that presolar grains might have been separated between chondrules and matrix “according to their size or type (metal, silicates, oxides, sulfides),” or that small presolar grains (e.g., SiC) may have been “preferentially excluded” either from the “dust aggregates from which chondrules ultimately formed” or from the “melting that produced the chondrules” (incomplete melting). No actual chondrule formation mechanisms proposed so far have ever resulted in effects such as size or type sorting, or in incomplete melting followed by physical separation of the unmelted crystals. In addition, these ideas are totally at odds with the suggestion of Wood (1985), Hezel and Palme (2010) and Palme et al. (2015) that complementarity between chondrules and matrix was established because chondrule formation resulted in elemental losses that recondensed on surrounding dust (the complementary fraction) which, when added together, produce the bulk solar system composition. Budde et al. (2016a, 2016b) did not consider the idea that the difference in Mo and W isotopic signatures between chondrules and matrix indicate that these two components were derived from different isotopic reservoirs and, hence, rule out complementarity in the sense proposed by Hezel and Palme (2010). The reason for that must be the underlying assumption that chondrules (s-process depleted) and matrix (s-process enriched) add up to make a CI (solar) bulk analogous to that proposed by the elemental complementarity hypothesis. While that might be true for W isotope anomalies with an ε183W = 0 for bulk Allende, it does not appear to be applicable to Mo isotope anomalies. Budde et al. (2016b) reported two measurements of bulk Allende with significantly different Mo isotope compositions that plotted between the Mo The Chondritic Assemblage 145 anomalies of chondrule and matrix fractions, but which differ even more significantly from the Mo isotopic composition of bulk Allende reported by Burkhardt et al. (2011) that has a larger s- process Mo isotope deficit than the Allende chondrule fraction. Budde et al. (2016b) state that these results demonstrate “significant Mo isotope heterogeneity among bulk samples of Allende, even for samples prepared from powders of 40 g and 100 g material” and suggest that this must result from the uneven distribution of CAIs with highly anomalous Mo isotope compositions. These results show that the Mo isotopic composition of bulk Allende is not well constrained, that it is unlikely to be close to that of CIs as the available measurements are skewed towards s-process deficits, and that anomalous CAIs would need to be taken into account into the bulk Mo isotope budget of the rock in order to discuss isotopic comple- mentarity. It should also be observed here that oxygen isotopic signatures seem to indicate chondrules and matrices were derived from different isotopic reservoirs (Zanda et al., 2006) and that oxygen isotopic complementarity is ruled out as no chondrite is solar in Δ17O and no other chondrite group has the same oxygen isotope signature as CIs. Another concern is that parent-body effects may have modified the Mo and W isotopic signatures of chondrules and matrices. Allende has been classified as a >3.6 petrographic type (Bonal et al., 2006), and it has experienced fluid-assisted thermal metamorphism, resulting in Fe-alkali metasomatic alteration (e.g., Krot et al., 1998, 2006; Krot, Petaev, and Bland, 2004). Both Mo and W become very mobile during chemical alteration of Fe–Ni alloy under the oxidizing and sulfidizing conditions (Palme et al., 1994; Humayun, Simon, and Grossman,

2007) experienced by Allende CAIs. Scheelite (CaWO4), molybdenite (MoS2) and other alteration phases formed, and W-Mo were transported in veins crossing the CAI-matrix boundary (Campbell et al., 2003). The effect of these processes operating on Allende CAIs is to transfer isotopically anomalous Mo and W from the CAIs to the matrix creating an apparent isotopic difference between chondrules and matrix. However, nucleosynthetic isotope anomal- ies observed in CAIs are characterized by s-process isotope deficits, so that any preferential alteration of CAIs does not create the Mo s-process excess observed in the Allende matrix by Budde et al. (2016b). Interestingly, these authors dismissed such parent-body effects as a possible source of the nucleosynthetic anomalies on the grounds that “these would tend to decrease and not increase their magnitude as they potentially result in the redistribution and homogenization of isotopically heterogeneous materials (Yokoyama et al., 2011).” The differences in nucleosynthetic Mo and W isotopic signatures between chondrules and matrices hence appear unlikely to have been established as the result of the chondrule forming process operating on material derived from the same reservoir of an initially solar composition. Isotopic complementarity thus argues in favor of chondrules and matrices being derived from different isotopic reservoirs.

5.5 Conclusion

Except for CIs, chondrite compositions are fractionated with respect to the Sun in all elements across the condensation temperature range. This suggests condensation or evaporation pro- cesses played a role in establishing chondritic compositions, possibly as a result of chondrule formation. In addition, Fe (and Ni) are fractionated with respect to Si and Mg, which suggests 146 Brigitte Zanda, Éric Lewin, and Munir Humayun that physical separation of metal and silicates took place, possibly also as a result of chondrule formation. The fractionated major and minor element carbonaceous chondrite compositions are better reproduced by mixing an average chondrule composition with a CI matrix and Fe–Ni alloy containing 0.2 wt% Cr, than by combining the in situ measured compositions of their chondrules and matrices. This is likely to result from the difficulty in performing accurate in situ analyses of chondrules and matrices and attests to the need for analytical developments (Zanetta et al., 2017). Chondrule and matrix compositions are consequently not known well enough to argue that in any given chondrite they are genetically related to one another and accreted without being separated in order to maintain a solar composition, as suggested by Wood (1963, 1985), and later by Hezel and Palme (2008, 2010), Palme et al. (2015), and Ebel et al. (2016). Moreover, the similarity in patterns between bulk and matrices in the volatile element range described by Bland et al. (2005) is likely to have been established on the parent body rather than reflecting nebular processes in which chondrules and matrices from a given chondrite were formed together. Last, the W and Mo nucleosynthetic isotopic differences between chondrules and matrices measured by Budde et al. (2016a, 2016b) argue in favor of an origin from distinct isotopic reservoirs for these chondritic components rather than for their being formed together from the same reservoir of solar composition. Hence, we conclude that a genetic (comple- mentarity) relationship between chondrules and their host matrices is not a required hypothesis, which implies that a time interval between the formation of chondrules (close to the Sun?) and the accretion and transport of chondrules over large distances to be mixed in with the matrix in colder regions of the disk are not ruled out.

Acknowledgments

The authors thank the referees for their constructive reviews and H. C. Connolly for editorial handling. S. Russell, H.C. Connolly, and A. N. Krot are also gratefully acknowledged for convening the outstanding 2017 Chondrule Workshop in London and for organizing the present book.

References

Abreu, N. M., and Brearley, A. J. (2010). Early solar system processes recorded in the matrices of two highly pristine CR3 carbonaceous chondrites, MET 00426 and QUE 99177. Geochim. Cosmochim. Acta, 74(3), 1146–1171. Alexander, C. M. O’D. (2005). Re-examining the role of chondrules in producing the elemental fractionations in chondrites. Meteorit. Planet. Sci., 40(7), 943–965. Allegre,̀ C., Manhès, G., and Lewin, É. (2001). Chemical composition of the Earth and the volatility control on planetary genetics. Earth and Planet. Sci. Lett., 185,49–69. Anders, E. (1964). Origin, age, and composition of meteorites. Space Sci. Rev., 3(5–6), 583–714. Anders, E., and Ebihara, M. (1982). Solar-system abundances of the elements. Geochim. Cosmochim. Acta, 46, 2363–2380. Anders, E., and Grevesse, N. (1989). Abundances of the elements: Meteoritic and solar. Geochim. Cosmochim. Acta, 53, 197–214. Barrat, J. A., Zanda, B., Moynier, F., et al. (2012). Geochemistry of CI chondrites: Major and trace elements, and Cu and Zn isotopes. Geochim. Cosmochim. Acta, 83,79–92. The Chondritic Assemblage 147

Bland P. A., Alard O., Benedix G.K., et al. (2005). Volatile fractionation in the early solar system and chondrule/matrix complementarity. Proc. Nat. Acad. Sci., 102(39), 13755–13760. Bonal, L., Quirico, E., Bourot-Denise, M., and Montagnac, G. (2006). Determination of the petrologic type of CV3 chondrites by Raman spectroscopy of included organic matter. Geochim. Cosmochim. Acta, 70(7), 1849–1863. Bourot-Denise, M., Zanda, B., Marrocchi, Y., et al. (2010). Paris: The Slightly Altered, Slightly Metamorphosed CM that Bridges the Gap Between CMs and COs. In Lunar Planet. Sci. XLI, LPI Contribution No. 1533, p. 1683. Brearley, A. J. (1993). Matrix and fine-grained rims in the unequilibrated CO3 chondrite, ALHA77307: Origins and evidence for diverse, primitive nebular dust components. Geochim. Cosmochim. Acta, 57(7), 1521–1550. Budde, G., Kleine, T., Kruijer, T. S., Burkhardt, C., and Metzler, K. (2016a). Tungsten isotopic constraints on the age and origin of chondrules. Proc. Nat. Acad. Sci., 113(13), 2886–2891. Budde, G., Burkhardt, C., Brennecka, G. A., et al. (2016b). Molybdenum isotopic evidence for the origin of chondrules and a distinct genetic heritage of carbonaceous and non- carbonaceous meteorites. Earth Planet. Sci. Lett., 454, 293–303. Burkhardt, C., Kleine, T., Oberli, F., et al. (2011). Molybdenum isotope anomalies in meteor- ites: Constraints on solar nebula evolution and origin of the Earth. Earth Planet. Sci. Lett., 312(3–4), 390–400. Burkhardt, C., and Schönbächler, M. (2015). Intrinsic W nucleosynthetic isotop variations in carbonaceous chondrites: Implications for W nucleosynthesis and nebular vs. parent body processing of presolar materials. Geochim. Cosmochim. Acta, 165, 361–375. Cameron, A. G. W. (1982). Elemental and nuclidic abundances in the solar system. In C. A. Barnes, D. D. Clamon, and D. N. Schramm (Eds.), Essays in Nuclear Astrophysics,23–43. New York, NY: Cambridge University Press. Campbell, A. J., Simon, S. B., Humayun, M., Grossman, L. (2003). Chemical evolution of metal in refractory inclusions in CV3 chondrites. Geochim. Cosmochim. Acta, 67, 3119–3134. Diakonova, M. I., Kharitonova, V. I., and Iavnel, A. A. (1979). The Chemical Composition of Meteorites. (Book in Russian) Moscow; Russia: Izdatel’stvo Nauka, 68 p. Ebel, D. S., Brunner, C., Konrad, K., et al. (2016). Abundance, major element composition and size of components and matrix in CV, CO and Acfer 094 chondrites, Geochim. et Cosmochim. Acta, 172, 322–356. Frank, D., Zolensky, M., Martinez, J., et al. (2011). A CAI in the Ivuna CI1 Chondrite. In Lunar Planet. Sci. XLI, LPI Contribution No. 1608, p. 2785. Friedheim, C. (1888). Über die chemische Zusammensetzung der Meteoriten von Alfianello und Concepcion. Sitzber. Kgl. Preuss. Akad. Wiss. z. Berlin, 1, 345–367. Goswami, J. N., and Macdougall, J. D. (1983). Nuclear track and compositional studies of olivines in CI and CM chondrites. In LPSC XIII Part 2, A755–A764. Grossman, J. N., and Brearley, A. J. (2005). The onset of metamorphism in ordinary and carbonaceous chondrites. Meteorit. Planet. Sci., 40(1), 87–122. Grossman, J. N., and Wasson, J. T. (1985). The origin and history of the metal and sulfide components of chondrules. Geochim. Cosmochim. Acta, 49(4), 925–939. Grossman, L., Ebel, D. S., Simon, S. B., et al. (2000). Major element chemical and isotopic compositions of refractory inclusions in C3 chondrites: The separate roles of condensation and evaporation. Geochim. Cosmochim. Acta, 64(16), 2879–2894. Haramura, H., Kushiro, I., and Yanai, K. (1983). Chemical compositions of Antarctic meteor- ites I. Memoirs of NIPR, Special issue 30, 109–121. Hezel, D.C., and Palme, H. (2008). Constraints for chondrule formation from Ca–Al distribu- tion in carbonaceous chondrites. Earth Planet. Sci. Lett., 265(3–4), 716–725. 148 Brigitte Zanda, Éric Lewin, and Munir Humayun

Hezel, D. C., and Palme, H. (2010). The chemical relationship between chondrules and matrix and the chondrule matrix complementarity. Earth Planet. Sci. Lett., 294(1–2), 85–93. Hezel, D. C., Russell, S. S., Ross, A. J., and Kearsley, A. T. (2008). Modal abundances of CAIs: Implications for bulk chondrite element abundances and fractionations. Meteorit. Planet. Sci., 43(11), 1879–1894. Humayun, M., and Cassen, P. (2000). Processes determining the volatile abundances of the meteorites and terrestrial planets. In R.M. Canup and K. Righter (Eds.), Origin of the Earth and Moon,3–23. Tucson, AZ: University of Arizona Press. Humayun, M., Simon, S. B., and Grossman, L. (2007). Tungsten and hafnium distribution in calcium-aluminum inclusions (CAIs) from Allende and Efremovka. Geochim. Cosmochim. Acta, 71, 4609–4627. Jarosewich, E. (1990). Chemical analyses of meteorites: A compilation of stony and iron meteorite analyses. Meteorit. Planet. Sci., 25(4), 323–337. Jarosewich, E. (2006). Chemical analyses of meteorites at the Smithsonian Institution: An update. Meteorit. Planet. Sci., 41(9), pp. 1381–1382. Jones, R. H., and Schilk, A. J. (2009). Chemistry, petrology and bulk oxygen isotope compos- itions of chondrules from the Mokoia CV3 carbonaceous chondrite. Geochim. Cosmochim. Acta, 73(19), 5854–5883. Jones, R. H, and Scott, E.R.D (1989). Petrology and thermal history of type IA chondrules in the Semarkona (LL3.0) chondrite. In Proc. 19th Lunar Planet. Sci. Conf., pp. 523–536. Kimura, M., and Ikeda, Y. (1998). Hydrous and anhydrous alterations of chondrules in Kaba and Mokoia CV chondrites. Meteorit. Planet. Sci., 33(5), 1139–1146. Kong, P., and Palme, H. (1999). Compositional and genetic relationship between chondrules, chondrule rims, metal, and matrix in the Renazzo chondrite. Geochim. Cosmochim. Acta, 63(21), 3673–3682. Klerner, S. (2001). Materie im fruhen̈ Sonnensystem: Die entstehung von chondren, matrix, und refraktären forsteriten. Ph.D. thesis, Universität zu Köln. Krot, A. N., Hutcheon, I. D., Brearley, A. J., et al. (2006). Timescales and Settings for Alteration of Chondritic Meteorites. In D. S. Lauretta and H. Y. McSween Jr., Meteorites and the Early Solar System II, 525–553. Tucson, AZ: University of Arizona Press. Krot, A. N., Petaev, M. I., and Bland, P. A. (2004). Multiple formation mechanisms of ferrous olivine in CV3 carbonaceous chondrites during fluid-assisted metamorphism. Antarct. Meteorite Res., 17, 154–172. Krot, A. N., Petaev, M. I., Scott, E. R. D., et al. (1998). Progressive alteration in CV3 chondrites: More evidence for asteroidal alteration. Meteorit. Planet. Sci., 33, 1065–1085. Larimer, J. W., and Wasson, J. T. (1988). Siderophile element fractionation. In J. F. Kerridge and M. S. Matthews (Eds.), Meteorites and the Early Solar System, 416–435. Tucson, AZ: University of Arizona Press. Leshin, L. A., Rubin, A. E., and McKeegan, K. D. (1997). The oxygen isotopic composition of olivine and pyroxene from CI chondrites. Geochim. Cosmochim. Acta, 61, 835–845. Lodders, K. (2003). Solar System Abundances and Condensation Temperatures of the Elem- ents. Astrophys. J., 591(2), 1220–1247. Lodders, K., and Palme, H. (2009). Solar System Elemental Abundances in 2009. Suppl. to Meteorit. Planet. Sci., 44, A5154 Lodders, K., Palme, H., and Gail, H. -P. (2009). Abundances of the Elements in the Solar System. In J. E. Trumper̈ (Ed.), Landolt-Börnstein, New Series, Group VI/4B, 560–630. Berlin, Germany: Springer-Verlag. Lux, G., Keil, K., and Taylor, G. J. (1980). Metamorphism of the H-group chondrites - Implications from compositional and textural trends in chondrules. Geochim. Cosmochim. Acta, 44, 841–855. Mason, B. M., and Wiik, H. B. (1962). The Renazzo Meteorite, Novitates, 2106,1–11. The Chondritic Assemblage 149

McSween, H. Y., Jr. (1977a). Carbonaceous chondrites of the Ornans type – A metamorphic sequence. Geochim. Cosmochim. Acta, 41, 477–491. McSween, H. Y., Jr. (1977b). Petrographic variations among carbonaceous chondrites of the Vigarano type Geochim. Cosmochim. Acta, 41, 1777–1790. McSween, H. Y., Jr. (1979). Alteration in CM carbonaceous chondrites inferred from modal and chemical variations in matrix. Geochim. Cosmochim. Acta, 43, 1761–1770. McSween, H. Y., Jr., and Richardson, S. M. (1977). The composition of carbonaceous chondrite matrix. Geochim. Cosmochim. Acta, 41, 1145–1161. Palme, H. (2000). Are there chemical gradients in the inner Solar System? Space Sci. Rev., 92 (1–2), 237–262. Palme, H. (2001). Chemical and isotopic heterogeneity in protosolar matter in Origin and early evolution of solid matter in the Solar System. Roy. Soc. of London Phil. Trans A, 359 (1787), 2061–2075. Palme, H., and Beer, H. (1993). The composition of chondritic meteorites. In H. H. Voigt (Ed.), Landolt-Börnstein - Group VI/3A, 196–221. Berlin, Germany: Springer-Verlag. Palme, H., Hezel, D. C., and Ebel, D. S. (2015). The origin of chondrules: Constraints from matrix composition and matrix-chondrule complementarity. Earth Planet. Sci. Lett., 411, 11–19 Palme, H., Hutcheon, I. D., and Spettel, B. (1994). Composition and origin of refractory-metal- rich assemblages in a Ca,Al-rich Allende inclusion. Geochim. Cosmochim. Acta, 58, 495–513. Palme, H., Larimer, J. W., and Lipschutz, M. E. (1988). Siderophile element fractionation. In J. F. Kerridge and M. S. Matthews (Eds.), Meteorites and the Early Solar System, 436–461. Tucson, AZ: University of Arizona Press. Palme, H., and Lodders, K. (2009). Metal-silicate fractionation in carbonaceous chondrites. Suppl. to Meteorit. Planet. Sci., 44, A165. Palme, H., Lodders, K., and Jones, A. (2014). Solar System Abundances of the Elements. In A. M. Davis (Ed.), Planets, Asteroids, Comets and the Solar System. In H. D. Holland and K. K. Turekian (Eds.), Treatise on Geochemistry (Second Edition), 2, 15–36. Oxford, UK: Elsevier. Palme, H., Spettel, B., Kurat, G., and Zinner, E. (1992). Origin of Allende chondrules. In Lunar and Planet. Sci. XXIII, pp. 1021–1022. Rambaldi, E. R., and Wasson, J. T. (1981). Metal and associated phases in Bishunpur, a highly unequilibrated ordinary chondrite. Geochim. Cosmochim. Acta, 45(7), 1001–1015. Rambaldi, E. R., and Wasson, J. T. (1984). Metal and associated phases in Krymka and Chainpur: Nebular formational processes. Geochim. Cosmochim. Acta, 48(10), 1885–1897. Reid, A. M., Bass, M. N., Fujita, H., Kerridge, J. F., and Fredriksson, K. (1970). Olivine and pyroxene in the Orgueil meteorite. Geochim. Cosmochim. Acta, 34,1263–1255. Rubin, A. E., and Wasson, J. T. (1988). Chondrules and matrix in the Ornans CO3 meteorite: Possible precursor components. Geochim. Cosmochim. Acta, 52(2), 425–432. Scott, E. R. D., and Krot, A. N. (2014). Chondrites and their components. In A. M. Davis (Ed.), Meteorites and Cosmochemical Processes. In H. D. Holland and K. K. Turekian (Eds.), Treatise on Geochemistry (Second Edition), 1, 65–137. Oxford, UK: Elsevier. Sears, D. W. G., and Dodd, R. T. (1988). Overview and classification of meteorites. In J. F. Kerridge and M. S. Matthews (Eds.), Meteorites and the Early Solar System,3–31. Tucson, AZ: University of Arizona Press. Shu, F. H., Shang, H. and Lee, T. (1996). Toward an astrophysical theory of chondrites. Science, 271(5255), 1545–1552. Wai, C. M., and Wasson, J. T. (1977). Nebular condensation of moderately volatile elements and their abundances in ordinary chondrites. Earth Planet. Sci. Lett., 36,1–13. 150 Brigitte Zanda, Éric Lewin, and Munir Humayun

Warren, P. H. (1997). The unequal host-phase density effect in electron probe defocused beam analysis: an easily correctable problem. In Lunar Planet. Sci. XXVIII, p. 1406. Wasson, J. T., and Chou, C. -L. (1974). Fractionation of moderately volatile elements in ordinary chondrites. Meteoritics, 9,69–84. Wasson, J. T., and Kallemeyn, G. W. (1988). Compositions of chondrites. Phil. Trans. Royal Society of London, Series A, Mathematical and Physical Sciences, 325(1587), 535–544. Wasson, J. T., and Rubin, A. E. (2009). Composition of matrix in the CR chondrite LAP 02342. Geochim. Cosmochim. Acta, 73(5), 1436–1460. Wasson, J. T., and Rubin, A. E. (2010). Matrix and whole-rock fractionations in the Acfer 094 type 3.0 ungrouped carbonaceous chondrite. Meteorit. Planet. Sci., 45(1), 73–90. Wiik, H. B. (1956). The chemical composition of some stony meteorites. Geochim. Cosmochim. Acta, 9(5), 279–289. Wolf, D., and Palme, H. (2001). The solar system abundances of phosphorus and titanium and the nebular volatility of phosphorus. Meteorit. Planet. Sci., 36(4), 559–571. Wood, J. A. (1963). On the origin of chondrules and chondrites. Icarus, 2, 152–180. Wood, J. A. (1985). Meteoritic constraints on processes in the solar nebula. In D. C. Black and M. S. Matthews (Eds.), Protostars and Planets II, 687–702. Tucson, AZ: University of Arizona Press. Yokoyama, T., Alexander, C. M. O’D., and Walker, R. J. (2011). Assessment of nebular versus parent body processes on presolar components present in chondrites: Evidence from osmium isotopes. Earth Planet. Sci. Lett., 305(1–2), 115–123. Zanda, B., Hewins, R. H., Bourot-Denise, M., Bland, P. A., and Albarède, F. (2006). Formation of solar nebula reservoirs by mixing chondritic components. Earth Planet. Sci. Lett., 248, 650–660. Zanda, B., Humayun, M., Barrat, J. -A., Bourot-Denise, M., and Hewins, R. H. (2011a). Bulk and matrix composition of the Paris CM: Inferences on parent-body alteration and the origin of matrix-chondrule complementarity. In Lunar Planet. Sci. XLII, LPI Contribution No. 1608, p. 2040. Zanda, B., Humayun, M., Barrat, J. -A., Bourot-Denise, M., and Hewins, R. H. (2011b). Chemistry of carbonaceous chondrites matrices: Parent-body alteration and chondrule- matrix complementarity. 74th Annual Meeting of the Meteoritical Society, London. Suppl. to Meteoritics Planet. Sci., 46, A5358. Zanda, B., Humayun, M., and Hewins, R.H. (2012). Chemical composition of matrix and chondrules in carbonaceous chondrites: Implications for disk transport. In Lunar Planet. Sci. XLIII, LPI Contribution No. 1659, p. 2413. Zanda, B., Le Guillou, C., and Hewins, R. H. (2009). The relationship between chondrules and matrix in chondrites. Suppl. to Meteorit. Planet. Sci., 44, A5280. Zanetta, P.-M., Leroux, H., Le Guillou, C., et al. (2017). A new method for modal abundance, chemistry and density determination of fine grained matrices of primitive chondrites. Suppl. to Meteorit. Planet. Sci., 52, A6274. Zolensky, M., Barrett, R., and Browning, L. (1993). Mineralogy and composition of matrix and chondrule rims in carbonaceous chondrites. Geochim. Cosmochim. Acta, 57(13), 3123–3148. 6 Vapor–Melt Exchange Constraints on Chondrite Formation Conditions and Processes denton s. ebel, conel m. o’d. alexander, and guy libourel

Abstract

The bulk volatile contents of chondritic meteorites provide clues to their origins. Matrix and chondrules carry differing abundances of moderately volatile elements, with chondrules carrying a refractory signature. At the high temperatures of chondrule formation and the low pressures of the solar nebula, many elements, including Na and Fe, should have been volatile. Yet the evidence is that even at peak temperatures, at or near the liquidus, Na and Fe (as FeO and Fe-metal) were present in about their current abundances in molten chondrules. This seems to require very high solid densities during chondrule formation to prevent significant evaporation. Evaporation should also be accompanied by isotopic mass fractionation. Evidence from a wide range of isotopic systems indicates only slight isotopic mass fractionations of moderately volatile elements, further supporting high solid densities. However, olivine- rich, FeO-poor chondrules commonly have pyroxene-dominated outer zones that have been interpreted as the products of late condensation of SiO2 into chondrule melts. Late condensation of more refractory SiO2 is inconsistent with the apparent abundances of more volatile Na, FeO and Fe-metal in many chondrules. Despite significant recent experimental work bearing on this problem, the conditions under which chondrules behaved as open systems remain enigmatic.

6.1 Introduction

There has been a longstanding debate about whether chondrules behaved as open chemical systems, that is, gaining or losing material by exchange with surrounding H2-rich vapor during their formation (Wood, 1996; Hewins and Zanda, 2012; Connolly and Jones, 2016). Alterna- tively, chondrule chemical and isotopic compositions may primarily record the compositions of their precursors (Chapter 2), and they remained essentially closed systems upon melting and solidification. Experimental simulations of chondrule textures suggest that they were rapidly heated to near liquidus temperatures of 1,500–1,700 C for relatively brief periods and then cooled to solidus temperatures (1,000–1,200 C) at 10–1,000 C/hr (Hewins et al., 2005; Chapter 3). This implies formation timescales of hours to days. On these timescales, interactions between chondrules and gas would have been inevitable. Here we review the evidence for such interactions, note the constraints that they place on formation conditions, and highlight some unanswered questions.

151 152 Denton S. Ebel, Conel M. O’D. Alexander, and Guy Libourel 6.2 Clues from the Bulk Volatile Contents of Chondrites

A starting point in the discussion of volatile elements is the relationship between chondrules and the bulk volatile content of their host meteorites, and by extension their parent planetesimals. It has been concluded (Lodders et al., 2009) that deviations of the chondrite groups from CI composition can be understood, “at least in principle, by gas-solid fractionation processes”. Figure 6.1 illustrates the volatile element depletion of the bulk silicate Earth, known from the lithophile element abundances in mantle-derived rocks, compared to the volatile element depletions in LL and H ordinary chondrites (OC), CM, CR, CO, CV, CK, CH carbonaceous chondrites (CC), and EH enstatite chondrites. Elements are ordered according to their estimated 50 percent condensation temperatures (Lodders, 2003; Table 6.1). Abundances are normalized to the CI carbonaceous chondrite standard (Lodders et al., 2009) and to the refractory lithophile element Yb, as this is well measured in CI chondrites (Lodders et al., 2009), and, in contrast to normalization to Mg, reveals depletions in Fe, Mg, and Si. Furthermore, Yb is highly correlated with Al, often used as a proxy for refractory elements in general (Figure 6.1). Data for Mg, Fe, Si, Cr, Mn, Li, Cu, K, Ga, Na, Ge, Rb, Cs, Zn, Te, Sn, S, and Cd are plotted. Recent analytical results are consistent with the older data used in Figure 6.1. For example, most analyses of

Figure 6.1 Volatile trends in chondrites and silicate Earth. Bulk silicate Earth (BSE, filled diamonds), and CK (filled circles), LL (open diamonds), CV (open circles), H (open squares), CO (filled triangles), CR (open triangles), CM (filled squares) chondrite atomic abundances of selected elements, normalized to Yb and CI chondrite (dotted line at ‘1’) are plotted against their estimated 50% condensation temperatures (Lodders 2003). Earth data are from McDonough (2014); CI chondrite from Lodders et al. (2009); CK, LL, CV, H, CO, and CM from Wasson and Kallemeyn (1988); CR from Lodders and Fegley (1998). Trend lines for BSE and chondrites are calculated as described in text. Trends for CO and H are coincident. T (50%) for Rb and Cs are offset 20 K for clarity in this diagram. Vapor–Melt Exchange 153

Table 6.1 Elements, primary host phases, and 50% condensation temperatures (K; Lodders, 2003) used to construct Figure 6.1. Abbreviations: fo = forsterite, en = enstatite, fsp = feldspar.

Symbol Host T50%K

Al hibonite 1,653 Mg forsterite 1,336 Fe Fe alloy 1,334 Si fo + en 1,310 Cr Fe alloy 1,296 Mn fo + en 1,158 Li fo + en 1,142 Cu Fe alloy 1,037 K feldspar 1,006 Ga Fe alloy + fsp 968 Na feldspar 958 Ge Fe alloy 883 Rb feldspar 800 Cs feldspar 799 Zn fo + en 726 Te Fe alloy 709 Sn Fe alloy 704 S troilite 664 Cd en + troilite 652

Allende (CV3) by Makashima and Nakamura (2006) are within the standard deviation of 37 analyses by Stracke et al. (2012) for a majority of elements, and close to those of Jarosewich et al. (1987). The volatile trend lines were regressed using all these elements, except: Li in CR; Rb, Cs, and Cd in CK; and lithophiles Mg, Si, Cr, Mn, Li, K, Na, Rb, and Cs in H and LL. The regressions (Table 6.2, Figure 6.1) were constrained to pass through Yb. In the LL and H chondrites, the lithophile elements are enriched, but only weakly collinear. The EH and EL enstatite chondrites are omitted because their highly reduced nature indicates formation of many of their components in reducing conditions, which would strongly affect the condensation temperatures of moderately volatile elements (Lodders, 2003; Ebel and Alexander, 2011; Ebel and Sack, 2013). The depletion patterns in Figure 6.1 have been and continue to be the basis for extensive arguments about chondrule formation, summarized by Wood (1996). A two-component model mixing devolatilized chondrules and metal with CI-like matrix (e.g., Anders, 1964, 1977) would produce a step function, not continuous trends. Wood (1996) noted that a fractional condensa- tion model (e.g., Wai and Wasson, 1977; Wasson, 1977) involving “systematic withdrawal of gas containing uncondensed volatile elements as the protochondritic system cooled” could produce these trends, and that it is “a wonder” that such good correlations can be produced. 154 Denton S. Ebel, Conel M. O’D. Alexander, and Guy Libourel

Table 6.2 Results of the regressions computed for T(50%) versus depletion, producing the lines shown in Figure 6.1, with calculated r-squared. Fraction of matrix used in Figure 6.2 is noted (sources in text). * see text for elements included in regressions for H and LL.

Type Slope R2 Matrix

CI 0 0.998 CM 0.00068 0.98 0.7 CO 0.00100 0.98 0.419 CV 0.00107 0.96 0.4502 CR 0.00089 0.96 0.38 LL* 0.00115 0.99 0.25 H* 0.00102 0.94 0.23 CK 0.00119 0.97 0.4 CH 0.00109 0.77 0.05 BSE 0.00139 0.33

However, Wood also noted that for the elements with T (50% condensation) below about 700 K, the trends flatten. Indeed, these two models, and the fate of the missing volatiles, comprise 13–15 of the 15 “unresolved issues” in chondrule formation identified by Wood (1996). While many solutions have been proposed, their resolution based on meteoritic evidence remains elusive. Grossman (1996) attributed the fractionated bulk chondrite compositions in Figure 6.1 to losses or gains of at least six groups of elements. He concluded that these fractionations occurred “in chondrite-formation regions before chondrules” were formed. Similarly, Bland et al. (2005) found that although carbonaceous chondrites show monotonic volatile depletions in bulk, this pattern is not observed for matrix analyses. They attributed exchange of volatile elements between chondrules and matrix to thermal processing during subsequent chondrule formation, i.e., there is complementarity between matrix and chondrules (Bland et al., 2005; Palme et al., 2014; Ebel et al., 2016; Chapter 4). However, Zanda et al. (Chapter 5) contend that the matrices in the least altered carbonaceous chondrites are relatively unfractionated, and attribute the fractionations reported by Bland et al. (2005) to parent body processes. At present, the debate about whether chondrite matrices are dominated by a uniform, primitive CI-like material or material from the same reservoir from which chemically complementary chondrules formed remains unresolved, although isotopic evidence for the latter is accumulating (Chapter 10). There is a correlation between the volatile depletions of chondrites and their volume fractions of matrix. This correlation is discernable in the regressed slopes of the volatile trend lines in Figure 6.1, against the volume fraction of matrix (Figure 6.2; R2 = 0.975). Matrix fractions (Table 6.2) are from Weisberg, McCoy, and Krot (2006) and more recent measure- ments (Bayron et al., 2014, CR; Lobo et al., 2014, LL; and Ebel et al., 2016, CO and CV). The correlation is consistent with roughly CI abundances of the highly volatile elements, with T (50%) < ~750 K, in the matrix. In support of a very primitive, unheated component in matrix are the CI-like matrix-normalized abundances of presolar grains, volatile elements like Zn and organic C in the least metamorphosed chondrites (Huss, 1990; Huss and Lewis, 1995; Vapor–Melt Exchange 155

Figure 6.2 Relationship between slopes of volatile trends and matrix abundances (Table 6.2). Slopes (Table 6.2) are those of the lines in Figure 6.1. Sources of matrix data are described in the text.

Alexander, 2005; Davidson et al., 2014). Neither the presolar grains (or the noble gases they contain) nor the organic C would survive significant heating, let alone to temperatures required for chondrule formation (Hubbard and Ebel, 2015). This observed correlation between chondrule abundance and volatile depletion strongly suggests that chondrules are the carriers of volatile depletion in meteorite parent bodies. Furthermore, volatile element systematics suggest that planet formation and chondrule forma- tion are related processes. Although some objects in CB and/or CH chondrites may have impact origins, there are a great many arguments against such origins for the vast bulk of chondrules, including complementarity between matrix and chondrules (Bland et al., 2005; Palme et al., 2014; Chapter 4), among chondrules themselves (Ebel et al., 2008; Jones, 2012), and the ubiquity of chondrules among primitive materials (Taylor, Scott, and Kell, 1983; Grossman, 1988; Wood, 1996). It may be significant that bulk silicate Earth (BSE) plots at approximately “zero matrix” in Figure 6.2.

6.3 Measurements of Chondrules

At low pressures and chondrule liquidus temperatures (1,800–2,100 K) many elements are predicted to be volatile based on thermochemical calculations, and this has been confirmed by numerous experiments (e.g., Hashimoto, 1983; Davis et al., 1990; Floss et al., 1996;

Wang et al., 2001; Yu et al., 2003; Richter et al., 2011). The presence of H2 enhances evaporation rates above pressures of >10–6 bars (Nagahara and Ozawa, 1996; Kuroda and Hashimoto, 2002; Richter et al., 2002). For major elements, the order of decreasing volatility is S>alkalis>Fe>Si>Mg. Both theory and experiments indicate that the timescales for significant evaporation to occur at near peak chondrule formation temperatures would have been much shorter than the timescales for chondrule formation inferred from simulations of chondrule textures (Hewins et al., 2005). During free evaporation (i.e., no back reaction of a melt with evaporated material, and diffusion in the melt that is faster than evaporation), elemental fractionation is accompanied by isotopic fractionation with the melt becoming increasingly 156 Denton S. Ebel, Conel M. O’D. Alexander, and Guy Libourel enriched in heavy isotopes relative to lighter ones as the concentration of the element decreases (reviewed by Davis et al., 2005 and Day and Moynier, 2014). The relationship between heavy (α-1) isotope enrichment and element depletion follows the Rayleigh fractionation law, R = Rof , in which the ratio of isotopes (e.g., 66Zn/64Zn) remaining is, ideally, related to the initial isotope ratio (Ro) and the fraction (f ) of the element remaining (e.g., Zn) by the inverse square root of the mass ratios of the evaporating species (α, Davis and Richter, 2014). Hence, if chondrules formed in the low-pressure, low-density environment typically invoked for protoplanetary disks, they would be expected to show the characteristic Rayleigh fraction- ation behavior of evaporation in many elements. As summarized below, there have been multiple searches for evidence for evaporation in chondrules (Table 6.3). However, at present there is no unequivocal evidence for large degrees of evaporation.

6.3.1 Sulfur

Despite its volatility, one study of S isotopes in Semarkona (LL3.00) chondrules (Tachibana and Huss, 2005) found no evidence for evaporation, which would mass-dependently fractionate 32S from 34S, leaving heavy S in the melt. However, the petrologic evidence for S being present in chondrule melts during their formation is controversial. It is clear from experiments (Lauretta et al., 1997) that S will enter chondrules during parent body alteration/metamorphism, as observed in grades > 3.6 (Rubin et al., 1999), so careful study of unequilibrated chondrites is required. Hewins et al. (1997) argued that the most reduced (Type I) chondrules with coarsest phenocrysts lost more S than finer-grained, less thermally processed chondrules (see Zanda et al., 1994; Sears et al., 1996; Zanda, 2004). Rubin et al. (1999) found troilite (FeS) equally

Table 6.3 Isotopic fractionations observed in meteoritic chondrules.

T(50%) Maximum Element cond. fractionation References

Cd 642 chondrules are Wombacher et al. (2003, 2008) lighter than bulk S 664 < 1 ‰/amu Tachibana and Huss (2005) Zn 726 chondrules are Luck et al. (2005); Pringle et al. (2017) lighter than bulk K 1,006 < 1 ‰/amu Alexander et al. (2000); Alexander and Grossman (2005) Fe 1,334 < 1 ‰/amu Alexander and Wang (2001); Zhu et al. (2001); Kehm et al. (2003); Mullane et al. (2005); Poitrasson et al. (2005); Needham et al. (2009); Hezel et al. (2010); Wang (2013) Si 1,310 1.5 ‰/amu Molini-Velsko et al. (1986); Clayton et al. (1991); Georg et al. (2007); Hezel et al. (2010); Armytage (2011); Kühne et al. (2017) Mg 1,336 ~1 ‰/amu Esat and Taylor (1990); Galy et al. (2000); Young et al. (2002); Bouvier et al. (2013); Deng et al. (2017) Vapor–Melt Exchange 157 abundant in both highly reduced porphyritic and low-FeO chondrules in Semarkona (LL3.0), with no significant correlation between FeS content and grain size. The study of S is further hampered by the difficulty of accurately measuring the S content of Fe–Ni–S beads due to the heterogeneous exsolution of sulfides during crystallization and in the solid state. Despite this difficulty, Marrocchi and Libourel (2013) and Piani et al. (2016) have shown in CV and EH chondrites that sulfur concentrations and sulfide occurrence in chondrules obey high tempera- ture sulfur solubility and saturation laws, and that gas–melt interactions with high partial vapor pressures of sulfur could explain the co-saturation of low-Ca pyroxene and troilite (compare with Lehner et al., 2013). Nevertheless, it remains important to resolve whether the S in chondrules is primary and, if so, whether it retains isotopic evidence for evaporation because S can potentially place the strictest constraints on chondrule formation of any major element.

6.3.2 Zinc, Cadmium, and Copper

Zinc and Cd are even more volatile than S (Table 6.1). High-precision isotopic measurements of the five stable Zn isotopes in bulk carbonaceous and ordinary chondrites show that ratios such as 66Zn/64Zn are mass-dependently fractionated within chondritic materials (Luck et al., 2005; Moynier et al., 2009), but the Zn depletions in chondrites and their chondrules are not the result of evaporation (Figure 6.3; Moynier et al., 2017). Luck et al. (2005) measured whole

Figure 6.3 Three-isotope plot of δ68Zn versus δ66Zn showing average values for bulk chondrites reported by Luck et al. (2005), “L05” and Pringle et al. (2017), “P17”. “TL” refers to the Tagish Lake meteorite. Dashed line (slope 1.86) is a regression of Pringle et al. (2017) average values with the mean CI value of Barrat et al. (2012), at 0.46, 0.88. 158 Denton S. Ebel, Conel M. O’D. Alexander, and Guy Libourel rock 66Zn/64Zn, 67Zn/64Zn and 68Zn/64Zn in a large suite of ordinary and carbonaceous chondrites. Recently, Pringle et al. (2017) confirmed and extended the previous measure- ments. Both found that bulk chondrites become isotopically lighter with decreasing Zn content (Figure 6.1). Were Zn depletion in chondrites due to evaporation, the Zn remaining in the solids would be expected to be enriched in heavier isotopes. Luck et al. (2005) performed sequential leaching of Krymka (LL3.1), and Pringle at al. (2017) measured Zn isotopes in magnetic, sulfide, and silicate fractions separated from bulk chondrites Clovis (H3.6), GRA 95208 (H3.7) and ALH 90411 (L3.7). Sulfides were enriched in δ66Zn by ~0.65 ‰ relative to silicates in Krymka, the same as the mean enrichment found by Pringle et al. (2017). These authors also measured 66Zn/64Zn and 68Zn/64Zn in Allende matrix and in chondrules separated from CV3 chondrites Allende and Mokoia, finding that chondrules are isotopically lighter than bulk Allende, and matrix is heavier, consistent with the results for Krymka. Cadmium is a more volatile chalcophile metal than Zn (Figure 6.1; Wombacher et al., 2003, 2008). The δ114/110Cd compositions of bulk Allende, Murchison (CM2) and Orgueil (CI) are identical to Earth’s; however, ordinary and some enstatite chondrites show mass- dependent fractionations. Wombacher et al. (2008) found that Allende chondrules and Ca-, Al-rich inclusions (CAIs) were isotopically lighter in Cd than bulk, δ114/110Cd = 0.1 ‰. They also found that an Allende sample, heated at 1,100 C for 96 hours at the Ni-NiO oxygen buffer, became lighter, with δ114/110Cd = 0.9 ‰, likely due to the survival of refractory host minerals (i.e., in chondrules and CAIs). Russell et al. (2003) measured δ65Cu in NWA 801 (CR2) and reported that the chondrules are isotopically lighter than the bulk rock. Thus Cd and Cu isotopes appear to behave similarly to Zn isotopes; however, it must be noted that these elements, and Zn, are highly mobile and may be isotopically fractionated during parent body aqueous alteration (Walsh and Lipschutz, 1982; Palme et al., 1988; Friedrich, Wang et al., 2003). Luck et al. (2005) interpreted their measurements as due to interaction of Zn-depleted refractory components with vapor enriched in isotopically light Zn. Pringle et al. (2017) went further, suggesting that evaporative loss from sulfide that was enriched in the heavy isotopes of Zn from chondrite-forming reservoirs caused both Zn elemental depletions and enrichment in light Zn isotopes. A similar explanation may apply to chalcophile Cu and Cd. It is of particular interest that the differing depletions of chalcophile and lithophile elements in ordinary chon- drites (Figure 6.1) may be related to the isotopic fractionations of some of the chalcophile elements, as these authors suggest.

6.3.3 Alkali Metals: Potassium and Sodium

Sodium and K are the next most volatile of the major/minor elements in chondrules, and their abundances in chondrule mesostases can vary by over an order of magnitude. Sodium is monoisotopic, but K has two abundant stable isotopes and is an obvious target for isotopic searches for evidence of evaporation (Humayun and Clayton, 1995). Richter et al. (2011) showed that K diffusion in chondrule liquids is sufficiently rapid that there would be no suppression of isotopic fractionation by diffusion-limited evaporation. Thus, if there were free Vapor–Melt Exchange 159 evaporation of K, it should follow the expected Rayleigh fractionation law, as is observed in micrometeorites and cosmic spherules that were heated during atmospheric entry (Taylor et al., 2005). However, some caution must be taken when studying the alkalis because they can be very mobile in the glasses where they are concentrated, both during aqueous alteration and during thermal metamorphism. Thus, only chondrules from the most pristine chondrites should be studied, and even then considerable care is needed to only measure chondrules that do not appear to have exchanged with matrix (Huss et al., 2005). Secondary ion mass spectrometric (SIMS) measurements of K isotopic compositions in Semarkona (LL3.0) chondrules revealed no systematic isotopic fractionations consistent with Rayleigh fractionation (Alexander and Grossman, 2005). There were some isotopic fractiona- tions reported, but they were not systematic, and were attributed to instrumental artifacts associated with microcrystallites and fractures/holes in the mesostases that were measured. Alkali abundances are much lower in carbonaceous chondrite chondrules, suggesting that they could have large isotopic anomalies, but the low elemental abundances make the measurements very challenging. To date, no study of K isotopes in carbonaceous chondrite chondrules has been carried out. One potential explanation for the lack of K isotopic fractionations in Semarkona chondrules is that the K condensed into the chondrules at relatively low temperatures shortly before or even after final solidification. Indeed, this is what thermochemical models predict should happen under some conditions (Ebel and Grossman, 2000). The measured distribution coefficients between glass and clinopyroxene crystallites in the mesostases of chondrules show that Na, at least, was present in the chondrules prior to final solidification of the glass (Jones, 1990, 1996; Libourel et al., 2003; Alexander and Grossman, 2005). However, it is still possible that the alkalis entered the chondrules just before growth of the clinopyroxene crystallites. One way to test this explanation is to look for alkali zoning in chondrule phenocrysts. The expectation would be that Na contents would be very low throughout most of the phenocrysts’ growth, and rise very dramatically at their edges. Olivine is the most common phenocryst type, is normally the first phase to have begun crystallizing, and often exhibits igneous zoning in its major and minor elements. The olivine-melt distribution coefficient for Na is small (~0.003, Borisov et al., 2008). Nevertheless, it is high enough that Na concentrations in ordinary chondrite chondrule olivines can be measured by electron probe. Chondrules with porphyritic textures are thought to have approached their liquidus tempera- tures without destruction of all crystal nuclei so that crystal growth would have begun soon after the onset of cooling (Hewins et al., 2005). Thus, their phenocrysts should preserve a record of conditions throughout a chondrule’s cooling history. After carefully selecting only olivine-rich, porphyritic Semarkona chondrules with at least some interior regions of their mesostases that had not exchanged alkalis with the matrix, Alexander et al. (2008) measured the Na zoning profiles in their phenocrysts. Surprisingly, they found that, contrary to expectations, the phenocrysts had measurable levels of Na even in their cores and only modest increases toward their rims. Indeed, the zoning profiles were roughly consistent with essentially closed system crystallization. This was true for chondrules with wide ranges of Na and FeO contents and liquidus temperatures. One obvious concern is that the zoning profiles are the result of later diffusion of Na into the olivine. However, Na diffuses relatively slowly in olivine. The remarkable conclusion from these observations is that Na was present in Semarkona 160 Denton S. Ebel, Conel M. O’D. Alexander, and Guy Libourel chondrules at roughly the presently observed abundances throughout their crystallization. These measurements have been confirmed by several independent studies (Borisov et al., 2008; Kropf and Pack, 2008; Kropf et al., 2009; Hewins et al., 2012; Florentin et al., 2017), although there is some debate about exactly how much Na loss/gain (10–50 percent) during chondrule formation and cooling is allowable. Carbonaceous chondrite chondrules generally have much lower Na contents than ordinary chondrite chondrules, which is also reflected in the bulk chondrite compositions (Figure 6.1). Hence, the chondrules in carbonaceous chondrites potentially experienced much more evapor- ation, although the evidence from Zn, Cd, and Cu isotopes argues against evaporation. Unfor- tunately, the low Na contents also make the measurements of olivine phenocrysts much more challenging. The low concentrations make electron probe analyses impractical, and while in principle SIMS measurements do have the necessary sensitivity, ubiquitous Na surface contam- ination of thin sections has so far hindered any systematic studies.

6.3.4 Iron Oxide and Metal

Iron is the next most volatile major element after Na and K, whether it is present as FeO or Fe- metal (Tachibana et al., 2015). SIMS measurements of Fe isotopic compositions of olivine phenocrysts in ordinary chondrite chondrules with a wide range of FeO contents reveal no systematic fractionations that would be consistent with Rayleigh behavior although they should have been detectible if present (Alexander and Wang, 2001). Since then, higher precision bulk chondrule measurements of chondrules from a number of chondrites have found only relatively small mass fractionation effects (Table 6.3). These mass fractionation effects are much smaller than would be predicted if their range of Fe contents were the result of free evaporation. Since chondrules are a major component of chondrites, if they experienced significant Fe evaporation one might expect to see the isotopic signatures even in bulk chondrite measurements. However, as with the chondrules, bulk chondrites exhibit very subdued Fe isotopic fractionations (Zhu et al., 2001). Metallic iron is a common component of chondrules, particularly FeO-poor ones (type I). Texturally, Fe-metal seems to have been stable during chondrule formation. Villeneuve et al., (2015) showed by experiment that at high temperatures FeO-poor olivine (type IA)-like chondrules are readily oxidized to FeO-rich olivine (type IIA)-like chondrules, an argument for open-system addition requiring O-enriched vapor. However, Humayun et al. (2010, Figure 2) observed depletions of refractory siderophiles (e.g., W, Re, Os, Ir) relative to Fe, Co, and Ni in metal grains in the rims of chondrules in CR chondrites, compared to their abundances in metal grains in the chondrule interiors. They argued for the recondensation of the metal in rims from a refractory-depleted vapor formed during chondrule melting. Close variants of this hypothesis have also been proposed by others for CR chondrites (Zanda et al., 1994; Kong et al., 1999; Connolly et al., 2001; Humayun et al., 2002; Campbell et al., 2005). In contrast, Ebel et al. (2008, Figure 7) observed that interior metal grains in igneously zoned CR chondrules fall on the high temperature portions of Co/Fe versus Ni/Fe trends predicted by condensation calculations that include a nonideal metal solution model (Ebel and Grossman, 2000), while exterior metal grains had solar values. In these igneously zoned CR chondrules, Vapor–Melt Exchange 161 interior portions associated with high Co, Ni metal have nearly pure forsterite mineralogy, while outer parts are primarily enstatitic pyroxene (Hobart et al., 2015). They concluded that both silicate and metal chemistry are consistent with the melting of successive igneous rims of increasingly less refractory bulk compositions. In CR chondrites, stable melts with low FeO content in silicates require only modest enrichments in dust (Ebel and Grossman, 2000; Ebel, 2006, Plate 10). Similar major and trace element studies have yet to be conducted on meteorites from other chondrite groups.

6.3.5 Magnesium and Silicon

It is well documented that significant evaporation of Mg, Si, and O occurred in many Type B (melted) CAIs in CV3 chondrites, primarily Allende (L. Grossman et al., 2000, 2002; Richter et al., 2002, 2007). The melting temperatures of Type B CAIs (1,700–1,800 K) (1,800 K, e.g., Stolper, 1982; Stolper and Paque, 1986) are lower than those of FeO-poor Type I chondrules (1,900–2,100 K, e.g., Hewins and Radomsky, 1990; Alexander et al., 2008). Hence, these observations imply that if there were free evaporation of Mg, Si, and O during chondrule formation (from chondrule melts), there should be clear evidence for it in the Mg, Si, and O isotopes of chondrules. Galy et al. (2000), Young et al. (2002) and Young and Galy (2004) showed that whole- chondrules from Allende (CV3.6), Bjürbole (L/LL4), and Chainpur (LL3.4) contain slight 25Mg excesses that in Allende correlate with chondrule Mg/Al and size. These relations are consistent with small amounts of evaporation during chondrule formation. Recently, Deng et al. (2017) have improved the precision for measuring Mg-Al systematics by laser ablation inductively coupled plasma mass spectrometry (LA ICP-MS) and confirmed the presence of only small levels of Mg isotopic mass fractionations in Allende and Leoville (CV3) chondrules. Similarly small mass fractionation effects have been reported for Si isotopes in chondrules (Molini- Velsko et al., 1986; Clayton et al., 1991; Georg et al., 2007; Hezel et al., 2010). As with Fe, there are only very small Si isotope effects evident in bulk chondrites (Georg et al., 2007; Fitoussi et al., 2009).

6.4 Implications of the Lack of Evidence for Evaporation

If the range of elemental abundances observed in chondrules and reflected in whole rock compositions were produced by free evaporation, they should be accompanied by large and systematic isotopic fractions that are characteristic of Rayleigh-type behavior. The fact that such fractionations are not seen in any of the isotope systems that have been studied has potentially profound implications for conditions during chondrule formation. In addition, despite a wide range of Na contents in Semarkona chondrules, at least, the Na appears to have remained at roughly the observed abundances throughout chondrule cooling irrespective of chondrule type. Indeed, Zn, Cd, and Cu exhibit light isotope enrichments in chondrules relative to bulk meteorite. How can these observations be explained? Galy et al. (2000) suggested that chondrules formed with high partial pressures (~1 bar) of

H2 present. High pressures of H2 can reduce isotopic fractionation associated with evaporation 162 Denton S. Ebel, Conel M. O’D. Alexander, and Guy Libourel both by enhancing evaporation rates relative to diffusion rates in the melt, and by restricting diffusion rates in the gas so that there is more back reaction between the melt and the evaporated material (Ebel 2006, Plate 7). However, such high gas pressures are astrophysically unreason- able, making this explanation unlikely. Another suggestion is that chondrules were heated and cooled very rapidly, thereby prevent- ing significant evaporation (Wasson and Rubin, 2002). This explanation requires that the kinetics of melting and crystallization/solidification be much faster than those of evaporation. Rapid heating and cooling experiments are difficult to conduct, but Knudsen cell experiments (Imae and Isobe, 2017), and diffusion experiments do not support this explanation (Richter et al., 2011). However, there are natural “experiments” that can be used to test this idea. Micrometeorites and cosmic spherules are heated during atmospheric entry to a range of peak temperatures and cooled in a matter of seconds (e.g., Love and Brownlee, 1991). Despite these very short heating times, they show very large elemental and isotopic fractionations in K, Fe, O, Si, and Mg consistent with evaporative loss (Alexander et al., 2002; Taylor et al., 2005). The likely explanation is that the shorter the heating time the higher the degree of superheating required to achieve melting and, therefore, the faster the rates of evaporation. Thus, brief heating events are not the explanation for the lack of isotopic mass fractionation in chondrules.

6.4.1 High Solid Densities During Chondrule Formation

In the absence of plausible alternatives, the most likely explanation for the absence of system- atic isotopic fractionations in chondrules that are consistent with evaporation is that when chondrules formed they were stable melts that approached equilibrium with their surrounding gas. To generate an equilibrium vapor, there must have been some evaporation from chondrules and any other material present (e.g., dust, refractory inclusions, and so on). So there could have been some initial isotopic fractionation that was later all but erased by subsequent gas–melt re- equilibration. The extent of evaporation needed to generate the equilibrium vapor and the timescales needed to reach equilibrium will have depended on the density of solids (chondrule precursors, dust, and so on) present in the chondrule forming regions and the partial pressure of

H2 – the H2 increases the equilibrium vapor pressures of many species. The lower the solid densities in the chondrule forming regions, the greater the degree of evaporation that is required to generate the equilibrium vapor and the longer it takes to erase any isotopic fractionations. Thus, the elemental and isotopic compositions of chondrules can potentially be used to constrain their formation conditions if the equilibrium and kinetic factors that would have controlled the evolution of their compositions are understood. Equilibrium models indicate that to produce silicate melts, and to produce chondrule-like compositions by the time their solidi are reached, the solids in the formation regions must have been enriched in CI-like solids (i.e. dust) by factors of 15–100 (by numbers of atoms), relative to – solar, at a total pressure of Ptot =103 bars (Ebel and Grossman, 2000; Ebel, 2006, Plate 10). A total pressure of 10–3 bars is about the upper limit for astrophysical estimates of pressures in protoplanetary disks, even in shocks (D’Alessio et al., 2005). Higher enrichments are required – for lower Ptot. Enrichments >1,000 times solar even at 10 3 bars are needed if at near liquidus temperatures FeO-rich (type II) chondrules retained their current bulk FeO contents and metal Vapor–Melt Exchange 163 melts were stable in type I chondrules (Ebel and Grossman, 2000; Alexander, 2004; Alexander and Ebel, 2012; Tenner et al., 2015). Kinetic models of chondrule evaporation and gas–melt re- equilibration also suggest that solids densities must have been >1,000 at 10–3 bars or chondrule phenocrysts would have preserved isotopic mass fractionations in Fe, Si, O, and Mg that are not observed (Alexander, 2004; Fedkin and Grossman, 2013). However, the Na contents of chondrules provide the most stringent constraints. Even at – 1,000 times dust enrichments and Ptot =103 bars, at near liquidus temperatures Na and K will be entirely in the gas, and only recondense into the melt after most olivine and pyroxene has already crystalized. This is because the vapor pressures of Na alone must be > 10–3 bars for Na to condense into chondrules at near their liquidus temperatures (Lewis et al., 1993; Alexander and Ebel, 2012; Fedkin and Grossman, 2013). Thus, orders of magnitude higher solids enrichments are needed if chondrule melts maintained roughly their current Na contents at near liquidus temperatures (Alexander et al., 2008; Fedkin and Grossman, 2013). Furthermore, such solids enrichments would drive significant FeO condensation into chondrule silicates (Ebel and Grossman, 2000, Figure 8), yet high Na abundances are observed even in FeO-poor (type I) chondrules (Alexander et al., 2008; Kropf and Pack, 2008; Kropf et al., 2009). While Na contents provide the most stringent constraints at the moment, maintaining the observed FeO and Fe-metal contents at near-liquidus temperatures also require very high solids densities, which is a particularly important constraint for carbonaceous chondrite chondrules for which it is unclear whether the alkalis were retained at near peak temperatures. Alexander and Ebel (2012) summarized the constraints that can be placed on chondrule densities during formation from Na and Fe abundances. They also placed lower limits on chondrule number densities during chondrule formation, requiring compound chondrule and nonspherical chondrule abun- dances that are significantly higher than previously estimated (Ciesla and Hood, 2004; Rubin and Wasson, 2005). Recent petrologic evidence for so-called cluster chondrites (Metzler, 2012; Metzler and Pack, 2016; Bischoff et al., 2017) suggests that chondrule densities during forma- tion may have been much higher than these lower limits, and more in line with the density estimates based on Na.

6.5 Outstanding Problems

A model of chondrules as stable melts that formed in regions that had been enormously enriched in solids, relative to solar, explains many features of chondrules. However, the solids enrich- ments or absolute densities that are required by Na retention during ordinary chondrite chon- drule formation and Fe metal retention during ordinary and carbonaceous chondrite chondrule formation are much higher than current nebula models can explain (D’Alessio et al., 2005; Cuzzi, Hogan, and Shariff, 2008). Also, to melt the chondrules in such dense clumps requires a formation mechanism that can focus a large amount of energy into small regions, and no nebular model has been shown to be able to do this. Given the potential astrophysical implications, it is important to continue to test such models. There are, in fact, chemical and textural features that are not apparently entirely consistent with a high solids density model. For instance, as reviewed in the previous text, chondrules do retain small isotopic fractionations in many elements that are inconsistent with melts that fully 164 Denton S. Ebel, Conel M. O’D. Alexander, and Guy Libourel equilibrated with their surrounding vapor at all stages of formation. These fractionations could simply reflect incomplete re-equilibration after the initial phase of evaporation, and/or incom- plete equilibration between chondrules that had wide ranges of precursor compositions (Alexander and Ebel, 2012). The diversity of chondrule compositions indicates a wide range of precursors, and their diverse textures, even within single meteorites, require a broad range of thermal histories (Jones, Grossman, and Rubin, 2005; Ebel et al., 2016). Vapor from chondrules at the peripheries of a chondrule forming region would be able to diffuse out of the region before equilibrating with vapor. As a result, chondrules at the peripheries of a region would be expected to show significant isotopic fractionations, yet few chondrules (if any) do. This suggests the possibility of crudely estimating the typical size of a chondrule forming region. Assuming that fewer than 1 percent of chondrules exhibit significant isotopic fractionation associated with evaporation, and a spherical formation region with a homogeneous density of chondrules, then estimates of diffusion distances in the gas for typical estimates of chondrule formation timescales and conditions require that such “particle-vapor clump” regions must have been at least 150–6,000 km in radius (Cuzzi and Alexander, 2006). If they were any smaller, more than 1 percent of the chondrules would have been within one diffusion distance of the surface of the region. Chondrules, even within a single chondrite, exhibit a considerable diversity of compos- itions and textures that indicate a broad range of potential formation conditions (Jones et al., 2005). Perhaps most puzzling are compound chondrules composed of two or more distinct chondrule types that collided and stuck together while they were still hot and plastic. Since equilibration between chondrules involves diffusive exchange via the gas and diffusion distances in the gas on chondrule-forming timescales would have been much less than the sizes of the formation regions estimated above, it is possible that individual chondrule formation regions maintained numerous microenvironments that produced chondrules with different compositions (Cuzzi and Alexander, 2006). The development of turbulent mixing in such zoned chondrule forming regions is one possible explanation for multi-type com- pound chondrules. However, it also seems likely that the diversity of chondrule compos- itions and textures within chondrites requires multiple formation events with differing conditions.

6.5.1 Silica Metasomatism and the Sodium Paradox

Perhaps the most difficult observation to reconcile with formation of chondrules in dense clumps in which there was relatively little evaporation even of Na at near-liquidus tempera- tures is the presence of chondrules with pyroxene-rich outer zones around olivine-rich cores. The favored explanation for these rims is that volatilized SiO in the gas recondensed onto the chondrules creating silica enriched melts at their peripheries from which pyroxene crystal- lized. This process would seem to be entirely at odds with observations of Na and Fe metal being retained at near liquidus temperatures (Sections 6.3.3, 6.3.4), since Si is less volatile than Na or Fe. Examples of mineralogically zoned chondrules with Na-bearing olivine in their centers and low-Ca pyroxene in their outer zones (with poikilitic textures, olivine enclosed in low-Ca pyroxene) have been highlighted by several petrological surveys of different chondrite types Vapor–Melt Exchange 165

(Scott and Taylor, 1983; Tissandier et al., 2002; Krot et al., 2004; Libourel et al., 2006; Jones, 2012; Friend et al., 2016). The fraction of zoned chondrules in carbonaceous chondrites is estimated to be as high as 75 percent of all chondrules (Friend et al., 2016) demonstrating 1) the importance of this petrologic feature, and 2) the need to take account of this fact in chondrule formation models. Based on experiments that control the partial pressure of SiO gas above a crystallizing silicate melt (Tissandier et al., 2002; Kropf and Libourel, 2011), it has been suggested that gas– melt interactions may have played a major role during the formation of type I chondrules. According to these results, the formation of low-Ca pyroxene in a chondrule melt in which olivine is being dissolved: ðÞþðÞ¼ ðÞ Mg2SiO4 olivine SiO2 melt Mg2Si2O6 pyroxene (6.1) is buffered by exchange with the surrounding nebular gas:

SiO2ðÞ¼melt SiOðÞþ gas ½ O2ðÞgas : (6.2)

Increasing the activity of SiO2 in the melt promotes crystallization of low-Ca pyroxene (or silica) if saturation conditions are reached in the chondrule melt (Reaction 6.1). In this scenario (Libourel et al., 2006), the chondrule melt is considered to be an open system and its compos- ition, especially the SiO2 content, is controlled by (1) exchange with the surrounding gas, controlled by the SiO gas partial pressure, PSiO(gas) (Reaction 6.2), (2) the dissolution of precursor olivines (Kropf and Libourel, 2011; Soulié et al., 2017), and the crystallization of low-Ca pyroxenes (Tissandier et al., 2002; Chaussidon et al., 2008; Harju et al., 2015; Friend et al., 2016). The chemistry of the type I chondrules (PO, POP, PP) is thus dependent on

PSiO(gas) and/or the duration of gas–melt interactions at high temperatures.

6.5.2 Sodium Zoning in Mesostasis

A separate problem is the presence of Na in chondrule mesostasis, increasing toward chondrule rims. Grossman et al. (2002), and Alexander and Grossman (2005) concluded that this Na zoning is due to secondary exchange with matrix in chondrite parent bodies, based on correl- ations between Na and Cl and H in mesostasis. They noted that some chondrules are not zoned because they have fully equilibrated with Na in matrix. Others have observed that mesostasis compositions enriched in silica have higher

Na2O contents (Matsunami et al., 1993; Grossman et al., 2002; Libourel et al., 2003). Similarly, porphyritic pyroxene (PP) chondrules are systematically enriched in alkali and silica compared to porphyritic olivine pyroxene (POP) chondrules; porphyritic olivine (PO) chondrules being in comparison depleted in both Na and Si (Libourel et al. 2003; Berlin, 2010). Sodium diffusion in molten silicate is four orders of magnitude faster than diffusion of Si. This work leaves open the possibility of Na influx into chondrules prior to parent body accretion. Mathieu et al. (2008, 2011) have shown that Na solubility is primarily controlled by the silica content of the melt according to a bond species reaction of the type:

½1=2Si O Siþ½1=2Na O Na¼½Na O Si: (6.3) 166 Denton S. Ebel, Conel M. O’D. Alexander, and Guy Libourel

4+ Addition of Na inhibits the polymerization of [SiO4] tetrahedra in the melt. Reaction 6.3 clearly indicates that the more polymerized a melt, the higher its Na solubility. Condensation of

SiO into chondrule melts (i.e., increasing activity of SiO2(melt) and hence increasing Si-O-Si linkages of the melt) would enhance Na solubility. The very low Si diffusivity would induce a zonation of SiO2 in the melt, from rim to core. If the distribution of Na is controlled by the diffusion of SiO2 in the chondrule melt according to Reaction 6.3, then the heterogeneities in Na2O contents measured in chondrules could reflect variations of PSiO(gas) and/or different timing of the gas–melt interaction. Higher PSiO(gas) and/or longer exposure time would be required to produce PP chondrules. This scenario likewise could explain silica-bearing rims in CR chondrules (Krot et al., 2004), silica-bearing chondrules in enstatite chondrites (Lehner et al., 2013; Piani et al., 2016), and possibly silica-rich components in ordinary chondrites (Hezel et al., 2006). In the context of the particle-vapor clump model (Cuzzi and Alexander, 2006; Section 6.5.1), this proposed phenomenological model suggests variations of vapor composition or cooling time. Although it is consistent with multiple separate chondrules reservoirs, which are subjected to various conditions over extended time periods (Jones, 2012), this scenario must be tested by a systematic survey of silicon isotope systematics in chondrules.

6.6 Conclusions

The recent decade has seen a revolution in nontraditional stable isotope investigations enabled by LA-ICPMS methods (e.g., Young and Galy, 2004; Moynier et al., 2017). These studies have produced little evidence of evaporation of chondrule melts, and puzzling evidence of chalcophile element light isotopic enrichments in chondrules. Measurements of Na in olivine phenocrysts have prompted new questions about chondrule melt–vapor interactions, suggesting high partial pressures of Na due to heating of regions that were highly enriched in solids. New experiments demonstrate how late SiO condensation into chondrules could explain their olivine/pyroxene petrology and perhaps also Na zoning in chondrule mesostases, but does not explain why Fe- metal or FeO do not similarly enter chondrules. In detail, distributions of Na, SiO2, and FeO inside chondrules do not appear to fit prevailing models for chondrule formation. It might appear that with lots of data, there has been little progress (Wood, 2001). However, the abundance of new kinds of data bearing on vapor-melt exchange, while prompting new questions, also promises new, quantitative constraints on dynamical models for chondrule formation.

References

Alexander, C. M. O’D. (2004). Chemical equilibrium and kinetic constraints for chondrule and CAI formation conditions. Geochim. Cosmochim. Acta, 68, 3943–3969. Alexander, C. M. O’D. (2005). Re-examining the role of chondrules in producing the elemental fractionations in chondrites. Meteorit. Planet. Sci., 40, 943–965. Alexander, C. M. O’D., and Ebel, D. S. (2012). Questions, questions: Can the contradictions between the petrologic, isotopic, thermodynamic, and astrophysical constraints on chon- drule formation be resolved? Meteorit. Planet. Sci., 47, 1157–1175. Alexander, C. M. O’D., and Grossman, J. N. (2005). Alkali elemental and potassium isotopic compositions of Semarkona chondrules. Meteorit. Planet. Sci. 40, 541–556. Vapor–Melt Exchange 167

Alexander, C. M. O’D., and Wang, J. (2001). Iron isotopes in chondrules: Implications for the role of evaporation during chondrule formation. Meteorit. Planet. Sci. 36: 419–428. Alexander, C. M. O’D., Grossman, J. N., Wang, J., et al. (2000). The lack of potassium-isotopic fractionation in Bishunpur chondrules. Meteorit. Planet. Sci. 35, 859–868. Alexander, C. M. O’D., Taylor, S., Delaney, J. S., Ma, P., and Herzog, G. F. (2002). Mass- dependent fractionation of Mg, Si, and Fe isotopes in five stony cosmic spherules. Geochim. Cosmochim. Acta, 66, 173–183. Alexander, C. M. O’D., Grossman, J. N., Ebel, D. S., and Ciesla, F. J. (2008). The formation conditions of chondrules and chondrites. Science, 320, 1617–1619. Anders, E. (1964). Origin, age and composition of meteorites. Space Sci. Rev., 3, 583–714. Anders, E. (1977). Critique of “Nebular condensation of moderately volatile elements and their abundances in ordinary chondrites” by Chien M. Wai and John T. Wasson. Earth Planet. Sci. Lett., 36, 14–20. Armytage, R. M. G. (2011). The Silicon Isotopic Composition of Inner Solar System Materials. Ph.D. Dissertation. University of Oxford, 230pp. Barrat, J. A., Zanda, B., Moynier, F., et al. (2012). Geochemistry of CI chondrites: Major and trace elements, and Cu and Zn isotopes. Geochim. Cosmochim. Acta, 83, 79–92. Bayron, J. M., Erb, I. R., Ebel, D. S., Wallace, S., and Connolly, H. C. Jr. (2014). Modal abundances and chemistry of clasts in the Renazzo (CR2) chondrite by x-ray map analysis. LPSC #45, abstract 1225. Berlin, J. (2010). Mineralogy and Bulk Chemistry of Chondrules and Matrix in Petrologic Type 3 Chondrites: Implications for Early Solar System Processes. Ph.D. Dissertation. Free University, Berlin, 233pp. Bischoff, A., Wurm, G., Chaussidon, M., et al. (2017). The Allende multicompound chondrule (ACC) – Chondrule formation in a local super-dense region of the early solar system. Meteorit. Planet. Sci., 52, 906–924. Bland, P. A., Alard, O., Benedix, G. K., et al. (2005). Volatile fractionation in the early solar system and chondrule/matrix complementarity. Proc. Natl. Acad. Sci., 102, 13755–13760. Borisov, A., Pack, A., Kropf, A., and Palme, H. (2008). Partitioning of Na between olivine and melt: An experimental study with application to the formation of meteoritic Na2O-rich chondrule glass and refractory forsterite grains. Geochim. Cosmochim. Acta, 72, 5558–5573. Bouvier, A., Wadhwa, M., Simon, S. B., and Grossman, L. (2013). Magnesium isotopic fractionation in chondrules from the Murchison and Murray CM2 carbonaceous chon- drites. Meteorit. Planet. Sci., 48, 339–353. Chaussidon, M., Libourel, G., and Krot, A. N. (2008). Oxygen isotopic constraints on the origin of magnesian chondrules and on the gaseous reservoirs in the early Solar System. Geo- chim. Cosmochim. Acta, 72, 1924–19382. Ciesla, F. J., Lauretta, D. S., and Hood, L. L. (2004). The frequency of compound chondrules and implications for chondrule formation. Meteorit. Planet. Sci., 39, 531–544. Campbell, A. J., Zanda, B., Perron, C., Meibom, A., and Petaev, M. I. (2005). Origin and thermal history of Fe-Ni metal in primitive chondrites. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk,407–431. Astronomical Society of the Pacific Conference Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Clayton, R. N., Mayeda, T. K., Goswami, J. N., and Olsen, E. J. (1991). Oxygen isotope studies of ordinary chondrites. Geochim. Cosmochim. Acta, 55, 2317–2337. Connolly, H. C. Jr., Huss, G. R., and Wasserburg, G. J. (2001). On the formation of Fe-Ni metal in Renazzo-like carbonaceous chondrites. Geochim. Cosmochim. Acta, 65, 4567–4588. Connolly, H. C. Jr., and Jones, R. H. (2016). Chondrules: The canonical and noncanonical views. J. Geophys. Res. Planets, 121, 1885–1899. Cuzzi, J. N., and Alexander, C. M. O‘D. (2006). Chondrule formation in particle-rich nebular regions at least hundreds of kilometres across. Nature, 441, 483–485. 168 Denton S. Ebel, Conel M. O’D. Alexander, and Guy Libourel

Cuzzi, J. N., Hogan, R. C., and Shariff, K. (2008). Toward planetesimals: Dense chondrule clumps in the protoplanetary nebula. Astrophys. J., 687, 1432–1447. D’Alessio, P., Calvet, N., and Woolum, D. S. (2005). Thermal structure of protoplanetary disks. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, 353–372. Astronomical Society of the Pacific Conference Series, 341. San Francisco, CA: Astronomical Society of the Pacific., Davidson, J., Busemann, H., Nittler, L. R., et al. (2014). Abundances of presolar silicon carbide grains in primitive meteorites determined by NanoSIMS. Geochim. Cosmochim. Acta, 139, 248–266. Davis, A. M., Hashimoto, A., Clayton, R. N., and Mayeda, T. K. (1990). Isotope mass fractionation during evaporation of Mg2SiO4. Nature, 347, 655–658. Davis, A. M., Alexander, C. M. O‘D., Nagahara, H., and Richter, F. M. (2005). Evaporation and condensation during CAI and chondrule formation. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, 432–455. Astronomical Society of the Pacific Conference Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Davis, A. M., and Richter, F. M. (2014). Condensation and evaporation of solar system materials. In A. M. Davis (Ed.), Meteorites and Cosmochemical Processes.InH.D. Holland and K. Turekian (Eds.), Treatise on Geochemistry (Second Edition)., 1, 335–360. Oxford, UK: Elsevier. Day, J. M., and Moynier, F. (2014). Evaporative fractionation of volatile stable isotopes and their bearing on the origin of the Moon. Phil. Trans. Roy. Soc. London A., 372, 20130259. Deng, Z., Ebel, D. S., Gemma, M., Moynier, F., and Chaussidon, M. (2017). Contrasting Mg isotopic signatures in Leoville (CV3r) chondrules. Meteorit. Planet. Sci. Suppl. 52, abstract 6403. Ebel, D. S. (2006). Condensaton of rocky material in astrophysical environments. In D. Lauretta and H. Y. McSween Jr. (Eds.), Meteorites and the Early Solar System II, 253–277. Tucson, AZ: University of Arizona Press. Ebel, D. S., and Alexander, C. M. O’D. (2011). Equilibrium condensation from chondritic porous IDP enriched vapor: Implications for mercury and enstatite chondrite origins. Planet. Space. Sci., 59, 1888–1894. Ebel, D.S., and Grossman, L. (2000). Condensation in dust-enriched systems. Geochim. Cosmochim. Acta, 64, 339–366. Ebel, D. S., and Sack, R. O. (2013). Djerfisherite: Nebular source of refractory potassium. Contrib. Mineral. Petrol., 166, 923–934. Ebel, D. S., Weisberg, M. K., Hertz, J., and Campbell, A. J. (2008). Shape, metal abundance, chemistry and origin of chondrules in the Renazzo (CR) chondrite.Meteorit. Planet. Sci., 43: 1725–1740. Ebel, D. S., Brunner, C., Leftwich, K., et al. (2016). Abundance, composition and size of inclusions and matrix in CV and CO chondrites. Geochim. Cosmochim. Acta, 172, 322–356. Esat, T. M., and Taylor, S. R. (1990). Mg isotopic composition of chondrules from the unequilibrated ordinary chondrite Semarkona. Lunar Planet. Sci., 21, 333–334. Fedkin, A. V., and Grossman, L. (2013). Vapor saturation of sodium: Key to unlocking the origin of chondrules. Geochim. Cosmochim. Acta, 112, 226–250. Fitoussi, C., Bourdon, B., Kleine, T., Oberli, F., and Reynolds, B. C. (2009). Si isotope systematics of meteorites and terrestrial peridotites: Implications for Mg/Si fractionation in the solar nebula and for Si in the Earth’s core. Earth Planet. Sci. Lett., 287, 77–85. Florentin, L., Faure., F., Deloule, E., et al. (2017). Origin of Na in glass inclusions hosted in olivine from Allende CV3 and Jbilet Winselwan CM2: Implications for chondrule forma- tion. Earth Planet. Sci. Lett., 474, 160–171. Vapor–Melt Exchange 169

Floss, C., El Goresy, A., Zinner E., et al. (1996). Elemental and isotopic fractionations produced through evaporation of the Allende CV chondrite: Implications for the origin of the HAL- type hibonite inclusions. Geochim. Cosmochim. Acta, 60, 1975–1998. Friedrich, J. M., Wang, M. -S., and Lipschutz, M. E. (2003). Chemical studies of L chondrites. V: Compositional patterns for forty-nine trace elements in fourteen L4–6 and seven LL4–6 falls. Geochim. Cosmochim. Acta, 67, 2467–2479. Friend, P., Hezel, D. C., and Mucerschi, D. (2016). The conditions of chondrule formation, Part II: Open system. Geochim. Cosmochim. Acta, 173, 198–209. Galy, A., Young, E. D., Ash, R. D., and O’Nions, R. K. (2000). The formation of chondrules at high gas pressures in the solar nebula. Science, 290, 1751–1754. Georg, R. B., Halliday, A. N., Schauble, E. A., and Reynolds, B. C. (2007). Silicon in the Earth’s core. Nature, 447, 1102–1106. Grossman, J. N. (1988). Formation of chondrules. In J. F. Kerridge and M. S. Matthews (Eds.), Meteorites and the Early Solar System, 680–696. Tucson, AZ: University of Arizona Press. Grossman, J. N. (1996). Chemical fractionations of chondrites: Signatures of events before chondrule formation. In R. Hewins, R. H. Jones, and E. R. D. Scott (Eds.), Chondrules and the Protoplanetary Disk, 243–253. Cambridge, UK: Cambridge University Press. Grossman, J. N., Alexander, C. M. O’D., Wang, J., and Brearley, A. J. (2002). Zoned chondrules in Semarkona: Evidence for high-and low-temperature processing. Meteorit. Planet. Sci., 37, 49–74. Grossman, L., Ebel, D. S., Simon, S. B., et al. (2000). Major element chemical and isotopic compositions of refractory inclusions in C3 chondrites: The separate roles of condensation and evaporation. Geochim. Cosmochim. Acta, 64, 2879–2894. Grossman, L., Ebel, D. S., and Simon, S. B. (2002). Formation of refractory inclusions by evaporation of condensate precursors. Geochim. Cosmochim. Acta 66: 145–161. Hashimoto, A. (1983). Evaporation metamorphism in the early solar nebula – Evaporation experiments on the melt FeO-MgO-SiO2-CaO-Al2O3 and chemical fractionations of primitive materials. Geochem. J., 17, 111–145. Harju, E. R., Kohl, I. E., Rubin, A. E., and Young, E. D. (2015). Silicon isotope condensation in type IAB chondrules at near equilibrium conditions. LPSC #46, abstract 2658. Hewins, R. H., and Radomsky, P. M. (1990). Temperature conditions for chondrule formation. Meteoritics, 25, 309–318. Hewins, R. H., and Zanda, B. (2012). Chondrules: Precursors and interactions with the nebular gas. Meteorit. Planet. Sci., 47, 1120–1138. Hewins, R. H., Yu, Y., Zanda, B., and Bourot-Denise, M. (1997). Do nebular fractionations, evaporative losses, or both, influence chondrule compositions? Proc. NIPR Symp. Ant. Meteorit., 10, 275–298. Hewins, R. H., Connolly, H. C. Jr., Lofgren, G. E., and Libourel, G. (2005). Experimental constraints on chondrule formation. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the protoplanetary disk, 286–316. Astronomical Society of the Pacific Conference Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Hewins, R. H., Zanda, B., and Bendersky, C. (2012). Evaporation and recondensation of sodium in Semarkona Type II chondrules. Geochim. Cosmochim. Acta, 78, 1–17. Hezel, D. C., Palme, H., Nasdala, L., and Brenker, F. E. (2006). Origin of SiO2-rich components in ordinary chondrites. Geochim. Cosmochim. Acta, 70, 1548–1564. Hezel, D. C., Needham, A. W., Armytage, R., et al. (2010). A nebula setting as the origin for bulk chondrule Fe isotope variations in CV chondrites. Earth Planet. Sci. Lett., 296, 423–433. Hobart, K. K., Crapster-Pregont, E. J., and Ebel, D. S. (2015). Decoding the history of a layered chondrule through olivine grain orientation measurements using EBSD. LPSC #46, abstract 1978. 170 Denton S. Ebel, Conel M. O’D. Alexander, and Guy Libourel

Hubbard, A., and Ebel, D. S. (2015). Semarkona: Lessons for chondrule and chondrite forma- tion. Icarus, 245, 32–37. Humayun, M., and Clayton, R. N. (1995). Potassium isotope cosmochemistry: Genetic impli- cations of volatile element depletion. Geochim. Cosmochim. Acta, 59, 2131–2148. Humayun, M., Campbell, A. J., Zanda, B., and Bourot-Denise, M. (2002). Formation of Renazzo chondrule metal inferred from siderophile elements. LPSC #33, abstract 1965. Humayun, M., Connolly, H. C. Jr., Rubin, A. E., and Wasson, J. T. (2010). Elemental distribution in metal from the CR chondrites Acfer 059 and PCA 91082. LPSC #41, abstract 1840. Huss, G. R. (1990). Ubiquitous interstellar diamond and SiC in primitive chondrites: Abun- dances reflect metamorphism. Nature, 347, 159–162. Huss, G. R., and Lewis, R. S. (1995). Presolar diamond, SiC, and graphite in primitive chondrites: Abundances as a function of meteorite class and petrologic type. Geochim. Cosmochim. Acta, 59, 115–160. Huss, G. R., Alexander, C. M. O’D., Palme, P., Bland, P. A., and Wasson, J. T. (2005). Genetic relationships between chondrules, fine-grained rims, and interchondrule matrix. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, 701–731. Astronomical Society of the Pacific Conference Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Imae, N., and Isobe, H. (2017). An experimental study of chondrule formation from chondritic precursors via evaporation and condensation in Knudsen cell: Shock heating model of dust aggregates. Earth Planet. Sci. Lett., 473, 256–268. Jarosewich, E., Clarke, R. S., and Barrows, J. N. (1987). The Allende meteorite reference sample. Smithson. Contrib. Earth Sci., 27, 49 pp. Jones, R. H. (1990). Petrology and mineralogy of type II, FeO-rich chondrules in Semarkona (LL3.0): Origin by closed-system fractional crystallization, with evidence for supercool- ing. Geochim. Cosmochim. Acta, 54, 1785–1802. Jones, R. H. (1996). FeO-rich porphyritic pyroxene chondrules in unequilibrated ordinary chondrites. Geochim. Cosmochim. Acta, 60, 3115–3138. Jones, R. H. (2012). Petrographic constraints on the diversity of chondrule reservoirs in the protoplanetary disk. Meteorit. Planet. Sci., 47, 1176–1190. Jones, R. H., Grossman, J. N., and Rubin, A. E. (2005). Chemical, mineralogical and isotopic properties of chondrules: Clues to their origin. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk,251–285. Astronomical Society of the Pacific Conference Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Kehm, K., Hauri, E. H., Alexander, C. M. O’D., and Carlson, R. W. (2003). High precision iron isotope measurements of meteoritic material by cold plasma ICP-MS, Geochim. Cosmo- chim. Acta, 67, 2879–2891. Kong, P., Ebihara, M., and Palme, H. (1999). Distribution of siderophile elements in CR chondrites: Evidence for evaporation and recondensation during chondrule formation. Geochim. Cosmochim. Acta, 63, 2637–2652. Kropf, A., and Libourel, G. (2011). Gas-melt interaction experiments at high temperature and high SiO(g) partial pressure – Implication of melt composition to chondrule formation. LPSC #42, abstract 1160. Kropf, A., and Pack, A. (2008). Closed system behavior of chondrules – New constraints for the chondrule forming process. LPSC #39, abstract 2222. Kropf, A., Huss, G. R., Krot, A. N., and Pack, A. (2009). Closed system behavior of alkalis in type-1 chondrules – Understanding chondrules as igneous systems. LPSC #40, abstract 2464. Krot, A. N., Libourel, G., Goodrich, C. A., and Petaev, M. I. (2004). Silica-rich igneous rims around magnesian chondrules in CR carbonaceous chondrites: Evidence for condensation origin from fractionated nebular gas. Meteorit. Planet. Sci., 39, 1931–1955. Vapor–Melt Exchange 171

Kühne, P., Kadlag, Y., Tatzel, M., Frick, D. A., and Becker, H. (2017). Constraints on chondrule-matrix complementarity from silicon isotope compositions of components of the Allende CV3 meteorite. 80th Annual Meeting of the Meteoritical Society, abstract 6245. Kuroda, D., and Hashimoto, A. (2002). The reaction of forsterite with hydrogen – its apparent and real temperature dependence. Antarctic Meteorit. Res., 15, 152–164. Lauretta, D. S., Lodders, K., Fegley, B., and Kremser, D. T. (1997). The origin of sulfide- rimmed metal grains in ordinary chondrites. Earth Planet. Sci. Lett., 151, 289–301. Lehner, S. W., Petaev, M. I., Zolotov, M. Y., and Buseck, P. R. (2013). Formation of niningerite by silicate sulfidation in EH3 enstatite chondrites. Geochim. Cosmochim. Acta, 101, 34–56. Lewis, R. D., Lofgren, G. E., Franzen, H. F., and Windom, K. E. (1993). The effect of Na vapor on the Na content of chondrules. Meteoritics, 28, 622–628. Libourel, G., Krot, A. N., and Tissandier, L. (2003). Evidence for high temperature condensa- tion of moderately-volatile elements during chondrule formation. LPSC #34, abstract 1558. Libourel, G., Krot, A. N., and Tissandier, L. (2006). Role of gas–melt interaction during chondrule formation. Earth Planet. Sci. Lett., 251, 232–240. Lobo, A., Wallace, S., and Ebel, D. S. (2014). Modal abundances, chemistry, and sizes of clasts in the Semarkona (LL3.0) chondrite by x-ray map analysis. LPSC #45, abstract 1423. Lodders, K. (2003). Solar system abundances and condensation temperatures of the elements. Astrophys. J., 591, 1220–1247. Lodders, K., and Fegley, B. Jr. (1998). The Planetary Scientist’s Companion. New York, NY: Oxford University Press. Lodders, K., Palme, H., and Gail, H. P. (2009). Abundances of the elements in the solar system. In J. E. Trümper (Ed.), Landolt-Börnstein, New Series, Vol. VI/4B, 560–630. Berlin, Heidelberg, New York: Springer-Verlag. Love, S. G., and Brownlee, D. E. (1991). Heating and thermal transformation of micrometeor- oids entering the Earth’s atmosphere. Icarus, 89, 26–43. Luck, J. -M., Ben Othman, D., and Albarède, F. (2005). Zn and Cu isotopic variations in chondrites and iron meteorites: Early solar nebula reservoirs and parent-body processes. Geochim. Cosmochim. Acta, 69, 5351–5363. Makishima, A., and Nakamura, E. (2006). Determination of major, minor and trace elements in silicate samples by ICP-QMS and ICP-SFMS applying isotope dilution-internal standard- isation (ID-IS) and multi-stage internal standardisation. Geostand. Geoanalytical Res., 30, 245–271. Marrocchi, Y., and Libourel, G. (2013). Sulfur and sulfides in chondrules. Geochim. Cosmo- chim. Acta, 119, 117–136. Mathieu, R., Khedim, H., Libourel, G., et al. (2008). Control of alkali-metal oxide activity in molten silicates. J. Non-Cryst. Solids 354, 5079–5083. Mathieu, R., Libourel, G., Deloule, E., et al. (2011). Na2O solublity in CaO-MgO-SiO2 melts. Geochim. Cosmochim. Acta, 57, 608–628. Matsunami, S., Ninagawa, K., Nishimura, S., et al. (1993). Thermoluminescence and compos- itional zoning in the mesostasis of a Semarkona group A1 chondrule and new insights into the chondrule-forming process. Geochim. Cosmochim. Acta., 57, 2101–2110. McDonough, W. F. (2014). Compositional model for the Earth’s core. In R. W. Carlson (Ed.), The Mantle and Core. In H. D. Holland and K. K. Turekian (Eds.), Treatise on Geochem- istry (Second Edition), 3, 559–577. Oxford, UK: Elsevier. Metzler, K. (2012). Ultrarapid chondrite formation by hot chondrule accretion? Evidence from unequilibrated ordinary chondrites. Meteorit. Planet. Sci., 47, 2193–2217. Metzler, K., and Pack, A. (2016). Chemistry and oxygen isotopic composition of cluster chondrite clasts and their components in LL3 chondrites. Meteorit. Planet. Sci., 51, 276–302. 172 Denton S. Ebel, Conel M. O’D. Alexander, and Guy Libourel

Molini-Velsko, C., Mayeda, T. K., and Clayton, R. N. (1986). Isotopic composition of silicon in meteorites. Geochim. Cosmochim. Acta, 50, 2719–2726. Moynier, F., Dauphas, N., and Podosek, F. A. (2009). Search for 70Zn anomalies in meteorites. Astrophys. J. Lett., 700, L92–L95. Moynier, F., Vance, D., Fujii, T., and Savage, P. S. (2017). The isotope geochemistry of zinc and copper. Reviews in Mineralogy and Geochemistry, 82, 543–600. Washington, DC: Mineralogical Society of America. Mullane, E., Russell, S. S., and Gounelle, M. (2005). Nebular and asteroidal modification of the iron isotope composition of chondritic components. Earth Planet. Sci. Lett., 239, 203–218. Nagahara H., and Ozawa Z. (1996). Evaporation of forsterite in H2 gas. Geochim. Cosmochim. Acta, 60, 1445–1459. Needham, A. W., Porcelli, D., and Russell, S. S. (2009). An Fe isotope study of ordinary chondrites. Geochim. Cosmochim. Acta, 73, 7399–7413. Palme, H., Larimer, J. W., and Lipschutz, M. E. (1988). Moderately volatile elements. In J. F. Kerridge and M. S. Matthews (Eds.), Meteorites and the Early Solar System, 436–461. Tucson, AZ: University of Arizona Press. Palme, H., Hezel, D. C., and Ebel, D. S. (2014). Matrix chondrule relationship and the origin of chondrules. Earth Planet. Sci. Lett., 411: 11–19. Piani, L., Marrocchi, Y., Libourel, G., and Tissandier, L. (2016). Magmatic sulfides in the porphyritic chondrules of EH enstatite chondrites. Geochim. Cosmochim. Acta, 195, 84–99. Poitrasson, F., Levasseur, S., and Teutsch, N. (2005). Significance of iron isotope mineral fractionation in pallasites and iron meteorites for the core-mantle differentiation of terres- trial planets. Earth Planet. Sci. Lett., 234, 151–164. Pringle, E. A., Moynier, F., Beck, P., Paniello, R., and Hezel, D. C. (2017). The origin of volatile element depletion in early solar system material: Clues from Zn isotopes in chondrules. Earth Planet. Sci. Lett., 468, 62–71. Richter, F. M., Davis, A. M., Ebel, D. S., and Hashimoto, A. (2002). Elemental and isotopic fractionation of Type B calcium-, aluminum-rich inclusions: Experiments, theoretical considerations, and constraints on their evolution. Geochim. Cosmochim. Acta, 66, 521–540. Richter, F. M., Janney, P. E., Mendybaev, R. A., Davis, A. M., and Wadhwa, M. (2007). Elemental and isotopic fractionation of Type B CAI-like liquids by evaporation. Geochim. Cosmochim. Acta, 71, 5544–5564 Richter, F. M., Mendybaev, R. A., Christensen, J. N., Ebel, D., and Gaffney, A. (2011). Laboratory experiments bearing on the origin and evolution of olivine-rich chondrules. Meteorit. Planet. Sci., 46, 1152–1178. Rubin, A. E., and Wasson, J. T. (2005). Non-spherical lobate chondrules in CO3.0 Y-81020: General implications for the formation of low-FeO porphyritic chondrules in CO chon- drites. Geochim. Cosmochim. Acta, 69, 211–220. Rubin, A. E., Sailer, A. L., and Wasson, J. T. (1999). Troilite in the chondrules of type-3 ordinary chondrites: Implications for chondrule formation. Geochim. Cosmochim. Acta, 63, 2281–2298. Russell, S., Zhu, X., Guo, Y., et al. (2003). Copper and zinc isotope composition of Cr, CB and CH-like meteorites. EGS-AGU-EUG Joint Assembly, abstract #11331. Scott, E. R. D., and Taylor, G. J. (1983). Chondrules and other components in C, O, and E chondrites: Similarities in their properties and origins. Proc. Lunar Planet. Sci., 14, J. Geophys. Res. Suppl., 88, B275–B286. Sears, D. W. G., Huang, S., and Benoit, P. H. (1996). Open-system behaviour during chondrule formation. In R. Hewins, R. H. Jones, and E. R. D. Scott (Eds.), Chondrules and the Protoplanetary Disk, 221 –232. Cambridge, UK: Cambridge University Press. Vapor–Melt Exchange 173

Soulié, C., Libourel, G., and Tissandier, L. (2017). Olivine dissolution in molten silicates: An experimental study with application to chondrule formation. Meteorit. Planet. Sci., 46, 1–26. Stolper, E. (1982). Crystallization sequences of Ca-Al-rich inclusions from Allende: An experi- mental study. Geochim. Cosmochim. Acta, 52, 225–250. Stolper, E., and Paque, J. (1986). Crystallization sequences of Ca-Al-rich inclusions from Allende: The effects of cooling rate and maximum temperature. Geochim. Cosmochim. Acta, 50, 1785–1806. Stracke, A., Palme, H., Gellissen, M., et al. (2012). Refractory element fractionation in the Allende meteorite: Implications for solar nebula condensation and the chondritic compos- ition of planetary bodies. Geochim. Cosmochim. Acta, 85, 114–141. Tachibana, S., and Huss, G. R. (2005). Sulfur isotope composition of putative primary troilite in chondrules from Bishunpur and Semarkona. Geochim. Cosmochim. Acta, 69, 3075–3097. Tachibana, S., Nagahara, H., Ozawa, K., et al. (2015). Kinetic condensation and evaporation of metallic iron and implications for metallic iron dust formation. Astrophys. J., 736, 16–24. Taylor, G. J., Scott, E. R. D., and Keil, K. (1983). Cosmic setting for chondrule formation. In E. A. King (Ed.), Chondrules and their Origins, 262–278. Houston, TX: Lunar and Planetary Institute. Taylor, S., Alexander, C. M. O‘D., Delaney, J. S., et al. (2005). Isotopic fractionation of iron, potassium, and oxygen in stony cosmic spherules: Implications for heating histories and sources. Geochim. Cosmochim. Acta, 69, 2647–2662. Tenner, T. J., Nakashima, D., Ushikubo, T., Kita, N. T., and Weisberg, M. K. (2015). Oxygen isotope ratios of FeO-poor chondrules in CR3 chondrites: Influence of dust enrichment and H2O during chondrule formation. Geochim. Cosmochim. Acta, 148, 228–250. Tissandier, L., Libourel, G., and Robert, F. (2002). Gas-melt interactions and their bearing on chondrule formation. Meteorit. Planet. Sci., 37, 1377–1389. Villeneuve, J., Libourel, G., and Soulié, C. (2015). Relationships between type I and type II chondrules: Implications on chondrule formation processes. Geochim. Cosmochim. Acta, 160, 277–305. Walsh, T. M., and Lipschutz, M. E. (1982). Chemical studies of L chondrites–II. Shock-induced trace element mobilization. Geochim. Cosmochim. Acta, 46, 2491–2500. Wai, C. M., and Wasson, J. T. (1977). Nebular condensation of moderately volatile elements and their abundances in ordinary chondrites. Earth Planet. Sci. Lett., 36, 1–13. Wang, K. (2013). Iron Isotope Cosmochemistry. PhD Dissertation. Washington University, Saint Louis, 254pp. Wang, J., Davis, A. M., Clayton, R. N., Mayeda, T. K., and Hashimoto, A. (2001). Chemical and isotopic fractionation during the evaporation of the FeO-MgO-SiO2-CaO-Al2O3-TiO2 rare earth element melt system. Geochim. Cosmochim. Acta, 65, 479–494. Wasson, J. T. (1977). Reply to Edward Anders: A discussion of alternative models for explaining the distribution of moderately volatile elements in ordinary chondrites. Earth Planet. Sci. Lett., 36, 21–28. Wasson, J. T., and Kallemeyn, G. W. (1988). Composition of chondrites. Phil. Trans. Roy. Soc. London A, 325, 535–544. Wasson, J. T., and Rubin, A. E. (2002). Ubiquitous relict grains in type-II chondrules, narrow overgrowths, and chondrule cooling rates following the last melting event. LPSC #33, abstract 1141. Weisberg, M. K., McCoy, T. J., and Krot, A. N. (2006). Systematics and evaluation of meteorite classification. In D. Lauretta and H. Y. McSween Jr. (Eds.), Meteorites and the Early Solar System II,19–52. Tucson, AZ: University of Arizona Press. Wombacher, F., Rehkämper, M., Mezger, K., and Münker, C. (2003). Stable isotope compos- itions of cadmium in geological materials and meteorites determined by multiple-collector ICPMS. Geochim. Cosmochim. Acta, 67, 4639–4654. 174 Denton S. Ebel, Conel M. O’D. Alexander, and Guy Libourel

Wombacher, F., Rehkämper, M., Mezger, K., Bischoff, A., and Münker, C. (2008). Cadmium stable isotope cosmochemistry. Geochim. Cosmochim. Acta, 72, 646–667. Wood, J. A. (1996). Unresolved issues in the formation of chondrules and chondrites. In R. Hewins, R. H. Jones, and E. R. D. Scott (Eds.), Chondrules and the Protoplanetary Disk, 55–70. Cambridge, UK: Cambridge University Press. Wood, J.A. (2001). Chondrites: Tight-lipped witnesses to the beginning. unpublished Mazursky Lecture, 12 pp. Young, E. D., and Galy, A. (2004). The isotope geochemistry and cosmochemistry of magne- sium. Reviews in Mineralogy and Geochemistry, 55, 197–230. Washington, DC: Mineral- ogical Society of America. Young, E. D., Ash, R. D., Galy, A., and Belshaw, N. S. (2002). Mg isotope heterogeneity in the Allende meteorite measured by UV laser ablation-MC-ICPMS and comparisons with O isotopes. Geochim. Cosmochim. Acta, 66, 683–698. Yu, Y., Hewins, R. H., Alexander, C. M. O‘D., and Wang, J. (2003). Experimental study of evaporation and isotopic mass fractionation of potassium in silicate melts. Geochim. Cosmochim. Acta, 67, 773–786. Zanda, B. (2004). Chondrules. Earth Planet. Sci. Lett., 224, 1–17. Zanda, B., Bourot-Denise, M., Perron, C., and Hewins, R. H. (1994). Origin and metamorphic redistribution of silicon, chromium and phosphorus in the metal of chondrites. Science, 265, 1846–1849. Zhu, X. K., Guo, Y., O’Nions, R. K., Young, E. D., and Ash, R. D. (2001). Isotopic homogeneity of iron in the early solar nebula. Nature, 412, 311–312. 7 Chondrules in Enstatite Chondrites

emmanuel jacquet, laurette piani, and michael k. weisberg

Abstract

We review silicate chondrules and metal-sulfide nodules in unequilibrated enstatite chondrites (EH3 and EL3). Their unique mineral assemblages, with a wide diversity of opaque phases, nitrides, and nearly FeO-free enstatite, testify to exceptionally reduced conditions. While those have long been ascribed to a condensation sequence at supersolar C/O ratios, with the oldhamite- rich nodules among the earliest condensates, evidence for relatively oxidized local precursors suggests that their peculiarities may have been acquired during the chondrule-forming process itself. Silicate phases may have been then sulfidized in an O-poor and S-rich environment; whereas metal-sulfide nodules in EH3 chondrites could have originated in the silicate chondrules, those in EL3 may be impact products. The astrophysical setting (nebular or planetary) where such conditions were achieved, whether by depletion in water or enrichment in dry organics-silicate mixtures, is uncertain, but was most likely sited inside the snow line, consistent with the Earth-like oxygen isotopic signature of most EC silicates, with little data constraining its epoch yet.

7.1 Introduction

When thinking about the Chondrule Enigma, it is customary to concentrate on the common ferromagnesian, porphyritic chondrules found in ordinary or carbonaceous chondrites. Still, less usual chondrules (e.g., Al-rich, silica-rich, metallic...) can also offer their share of insight, for chondrule-forming scenarios should ideally account for them as well. There is a potential pitfall, however, in that nothing guarantees that all the objects we call “chondrules” formed by the same processes. For example, the currently widespread interpretation of the CB and CH chondrules as impact plume products (Chapter 2) does not necessarily hold for their more “mainstream” counterparts in other chondrite groups, which are very different petrographically and compos- itionally. Strangeness is an asset for the meteoriticist, but only within reasonable bounds. Chondrules in enstatite chondrites (EC) may provide such a golden mean. Their mineralogy is exceptionally reduced, with normally lithophile elements locked in unique sulfide (e.g., oldhamite, niningerite, alabandite) or nitride phases, high Si in metal and silica and FeO-free enstatite dominating over olivine. Yet the general textures and sizes of these chondrules are in the range seen in other chondrite clans1 (Jones, 2012), suggesting relatively similar geneses.

1 A clan is an association of several related chemical groups, whose parent bodies likely formed in a common reservoir. Ordinary chondrites (comprising H, L, and LL) and enstatite chondrites (EH and EL) are examples of chondrite clans.

175 176 Emmanuel Jacquet, Laurette Piani, and Michael K. Weisberg

Their isotopic signature is closer to that of the Earth and the Moon than any other chondrite group, suggesting that they represent material from the inner Solar System and not some aberrant exotic reservoir. As further evidence that they are not mere curiosities, enstatite chondrites represent at least two independent chemical groups (EH and EL, with the former more metal-rich and reduced than the latter) plus several anomalous samples presumably from yet other parent bodies (Weisberg and Kimura, 2012). The reason for the unusually reduced character of enstatite chondrite chondrules is not known. It may actually pertain not only to silicate chondrules, but also to the opaque (metal- sulfide) nodules common in EC which it is a matter of language convenience to also call chondrules or not depending on their genetic links to the former. A difficulty in the study of unequilibrated EC (Weisberg and Kimura, 2012) is their rarity, especially for observed falls (only three known, Qingzhen, Parsa, and Galim, all EH3s), whereas finds are prone to the easy weathering of their reduced mineralogy. Yet substantial progress has been accomplished in the past decade on EC chondrules, from the isotopic, chemical, and mineralogical point of view, with new insights on their astrophysical context of formation, and warrants a review of the field. In this chapter, we will describe the petrographic, chemical, and isotopic characteristics of chondrules, including opaque nodules. We will then discuss competing models for their reduced parageneses, in particular formation in a reduced condensation sequence or reduction of material during chondrule melting. Finally, we discuss the astrophysical setting that may have brought about the fractionations inferred for such environments, and its possible spatio-temporal location in the early solar system.

7.2 Petrography 7.2.1 Silicate Chondrules

Chondrules make up ~60–80 vol% of enstatite chondrites (Scott and Krot, 2014). Rubin and Grossman (1987) and Rubin (2000) determined an average chondrule diameter of about 220 µm for EH3 and 550 µm for EL3 chondrites. Schneider et al. (2002) reported similar sizes of 278 Æ 229 µm and 476 Æ 357 µm (average Æ standard deviation) for EH3 and EL3, respectively. The largest chondrules seem to be more often nonporphyritic in texture (Rubin and Grossman, 1987; Schneider et al., 2002). In the Sahara 97096 and Yamato 691 EH3 chondrites, Weisberg et al. (2011) found some chondrules up to 1 mm in diameter and one ~4 mm barred olivine chondrule. A broad correlation between chondrule and metal grain size in EH and EL chondrites may suggest aerodynamic sorting (Schneider et al., 1998). It is noteworthy that chondrules in E3 chondrites lack the fine-grained rims present in many other chondrites, and that compound chondrules are rare (Rubin, 2010). The chondrules in both EH3 and EL3 show a range of textures similar to those in ordinary and carbonaceous chondrites. However, enstatite (with generally <2 wt% FeO) is the major silicate phase in most E3 chondrules, with porphyritic pyroxene being the dominant chondrule texture (e.g., Figures 7.1a and b). In general, the chondrules contain enstatite Æ olivine (absent, or very minor, in many E3 chondrules) with a glassy mesostasis Æ Ca-rich pyroxene Æ minor silica (quartz, tridymite, cristobalite, and glass; Kimura et al., 2005). Olivine is commonly, but not exclusively, poikilitically enclosed in enstatite (Figure 7.1c). Although most of the Chondrules in Enstatite Chondrites 177

Figure 7.1 Mg-Ca-Al (red-green-blue) composite element map of the (a) Allan Hills (ALH) 81189 (,3) EH3 chondrite and (b) Queen Alexandra Range (QUE) 93351 (,6) EL3 chondrite showing chondrules of different textures and mineral abundances. In each image, the more common reds are dominantly enstatite, minor brightest reds are olivine and blues are feldspathic mesostases. Black regions are metal-sulfide nodules. (c) Example of a porphyritic chondrule from the ALH 81189 EH3 chondrite. The chondrule is dominantly enstatite (En) with poikilitic olivine (Ol), an albitic mesostasis (Mes) with Ca pyroxene (Ca- pyx) and minor Si-bearing Fe–Ni metal. (d) Example of a porphyritic chondrule from the MacAlpine Hill

(MAC) 88136 (,37) EL3 chondrite dominated by enstatite (Wo0.8Fs0.5) with minor mesostasis and pyroxene-hosted troilite mixed with daubréelite (Troil + Daub) enlarged in the insets. Troilite has 0.36 wt% Cr and 0.29 wt% Ti. (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.) chondrules in E3 chondrites may be characterized as type IB, they differ considerably from their counterparts in other chondrite groups (e.g., in the near pure endmember enstatite composition of their pyroxene or the occasional presence of silica). Most strikingly, the EC chondrules contain a wide range of opaque phases unique to this clan (e.g., Grossman et al., 1985; El Goresy et al., 1988; Ikeda, 1989b; Hsu, 1998; Rubin, 2010; Weisberg and Kimura, 2012; Lehner et al., 2013; Piani et al., 2016). In EH3 porphyritic chondrules, Cr-Ti-bearing troilite, Si-bearing Fe–Ni metal, oldhamite (CaS), and niningerite ([Mg,Fe,Mn]S) are the most abundant opaque phases, with average modal abundances of 2.1, 0.7, 0.5, and 0.4 vol%, respectively, for Sahara 97096 (Piani et al., 2016). Rare occurrences of minor caswellsilverite (NaCrS2), two Cr-sulfides (minerals A and B of Ramdohr, 1963), schreibersite ([Fe,Ni]3P), and perryite ([Fe,Ni]8[Si,P]3) have also been reported (Grossman 178 Emmanuel Jacquet, Laurette Piani, and Michael K. Weisberg et al., 1985; El Goresy et al., 1988; Piani et al., 2016). Troilite occurs both enclosed in low-Ca pyroxene and within the mesostasis associated with metal or other sulfides (Figure 7.1d). Oldhamite and niningerite are often found together with troilite in micrometer-sized assem- blages surrounded by mesostasis or connected to mesostasis channels between silicates (Piani et al., 2016). Niningerite is also frequently found in close association with low-Ca pyroxene, silica, and troilite in silica-rich areas of porphyritic chondrules (Piani et al., 2016) and abundant niningerite and troilite were reported in silica-rich chondrules (Lehner et al., 2013). The EL3 chondrite chondrules also contain sulfides (Rubin, 2010), typically oldhamite and alabandite ([Mn,Fe,Mg]S), with FeNi usually absent, but their textural and chemical properties and modal abundances have not been studied in detail. Olivine (sometimes with evidence of reduction) and FeO-rich (up to 10.2 wt% FeO) pyroxene grains are present in some chondrules of the least equilibrated E3 chondrites (Ram- baldi et al., 1983; Lusby et al., 1987; Weisberg et al., 1994), but are much less common than in ordinary or carbonaceous chondrite chondrules. No entire FeO-rich type II chondrules are present, but scarce ferroan pyroxene (up to 34 mol% ferrosilite) fragments (with silica2) and spherules have been described (Weisberg et al., 1994; Kimura et al., 2003; Jacquet et al., 2015). The rare olivine-rich porphyritic olivine/pyroxene (or olivine) chondrules in E3 chondrites (only 4 percent of EC chondrules; Jones, 2012) are similar to the type IA and IAB chondrules that are characteristic of ordinary and carbonaceous chondrites, but again may contain silica, Si-bearing metal, and sulfides that are not found in the latter.

7.2.2 Metal-Sulfide Nodules

In addition to silicate chondrules, the least equilibrated EL3 and EH3 chondrites contain abundant round to sub-round opaque assemblages ~50–800 µm in size (Weisberg and Kimura, 2012; Lehner et al., 2010), often concentrically layered in EH3 chondrites (where theymaymakeup30–40 vol% of the meteorite; Weisberg and Prinz, 1998; Lehner et al., 2014). Sometimes called “metal-sulfide chondrules” (e.g., Weisberg and Prinz, 1998), they will be henceforth referred to as “metal-sulfide nodules” (MSN; Ikeda, 1989a; Lehner et al., 2014). Fe-Ni metal (kamacite with ~2–3 wt% of Si in EH3; <1.4 wt% in EL3) and troilite are the most abundant phases in MSNs in both EL3 and EH3 chondrites and are often associated with a variety of other sulfides, schreibersite, perryite, graphite, and silicates. In the EH3 chondrite MSNs, the other sulfides include oldhamite, niningerite, djerfisher- ite ([K,Na]6[Fe,Ni,Cu]25S26Cl), caswellsilverite, daubréelite (FeCr2S4), and minor occur- rences of other Cu and Cr-sulfides (e.g., Kimura, 1988; Ikeda, 1989a; Lin and El Goresy, 2002). In the EL3 chondrite MSNs, these consist mostly of alabandite, daubréelite, sphaler- ite, and rare pendlandite ([Fe,Ni]9S8).Oldhamitewasreportedininclusionsinmetal(Lin and El Goresy, 2002), but may have been destroyed by terrestrial weathering in most EL3 chondrite samples.

2 Cristobalite and tridymite according to Raman spectra by E.J. on the Jacquet et al. (2015) objects. Chondrules in Enstatite Chondrites 179

Numerous silicates were reported in EH3 MSNs: low-Ca pyroxenes, silica (tridymite and cristobalite), roedderite ([Na,K]2[Mg,Fe]5Si12O30), albitic plagioclase (or albitic glass), and a porous amorphous silica (Keil, 1968; Kimura, 1988; Ikeda, 1989a; Hsu, 1998; Lehner et al., 2010). Silicates can be found as inclusions in kamacite, as intergrowths with sulfide and metal, or relatively continuous “garlands” (Figure 7.2; Weisberg and Prinz, 1998; Lehner et al., 2010). In most of the EL3 MSNs, euhedral to subhedral tabular or rod-shaped low-Ca pyroxenes form intergrowths with metal and sometimes sulfides (Weisberg et al., 1997; van Niekerk and Keil,

2011; El Goresy et al., 2017, Figure 7.2). Sinoite (Si2N2O) laths also sometimes occur associated with the low-Ca pyroxene (Horstmann et al., 2014, Figure 7.2) and in sinoite- graphite-oldhamite assemblages in an EL3 chondrite clast in the Almahata Sitta polymict breccia (Lin et al., 2011).

Figure 7.2 (a) MSNs in the EH3 chondrite Sahara 97096, and (b) and (c) metal-silicate intergrowths with enstatite and sinoite laths, minor schreibersite exsolution (black dotted line in (b)) and graphite inclusions in the EL3/4 MS-17 from the Breccia Almahatta Sitta (both from Horstmann et al., 2014) to which the reader is referred to for details; with permission from Elsevier). M = Fe–Ni metal, Tr = troilite, Schr = schreibersite, En = enstatite, Sil = silica, Perr = perryite, Nin = Niningerite, Oldh = oldhamite, Sin = sinoite, S = silicates. 180 Emmanuel Jacquet, Laurette Piani, and Michael K. Weisberg 7.3 Chemistry

Bulk silicate chondrule compositions in unequilibrated EC have been analyzed by instrumental neutron activation (Grossman et al., 1985), defocused/rastered electron microprobe (Ikeda, 1988; Schneider et al., 2002), and laser ablation inductively coupled plasma mass spectrometry (Lehner et al., 2014; Varela et al., 2015). While the element compositions of EC chondrules mostly overlap with chondrule populations from other chondrite clans (Jones, 2012), some differences are apparent. Some of the differences pertaining to lithophile elements merely reflect the bulk meteorites (the matrix being a minor component). These include enrichment in Si (e.g., as normalized to Mg) and volatile elements like Na and Cl, and impoverishment in refractory elements. A depletion relative to the whole rock in Ca, correlated with Eu and Se, seems to trace the loss of an oldhamite-rich component (Grossman et al., 1985), to which Varela et al. (2015) also ascribed negative (CI-normalized) anomalies in Nb, Ti, V, and Mn in pyroxene-rich chondrules. EC chondrules are also uniformly low in siderophile elements (e.g., Fe, Co, Ni, Ir, Au) which are tightly correlated (Grossman et al., 1985) owing to their concentration in the opaque phases. Negative Eu, Yb, and sometimes Sm anomalies are common (Lehner et al., 2014; see also leaching residue analyses by Barrat et al. (2014). If we turn to individual mineral compositions (reviewed by Brearley and Jones, 1998), it is remarkable that minor element (Fe included) compositions of silicates (olivine, pyroxene) differ little, qualitatively, from those of type I chondrules in ordinary chondrites, although a popula- tion of enstatite with blue (instead of red) cathodoluminescence shows concentrations of all minor elements below 0.1 wt% (Brearley and Jones, 1998; Schneider et al., 2002). This is true of their rare earth element (REE) patterns (Hsu and Crozaz, 1998; Jacquet et al., 2015) as well, although negative Eu anomalies are more common (Figure 7.3). In the most comprehensive

Figure 7.3 Representative rare earth element patterns of enstatite chondrite chondrule phases. Data sources are: Gannoun et al. (2011) for oldhamite (average of “type C-D” analyses in Sahara 97072); Crozaz and Lundberg (1995) for niningerite (average for Allan Hill 77156 (EH3/4)); Hsu and Crozaz (1998) for enstatite (here their representative “pattern II” and “pattern III” for “blue enstatite”); Jacquet et al. (2015) for olivine and mesostasis (averages of Sahara 97096 chondrules). (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.) Chondrules in Enstatite Chondrites 181

(SIMS-based) study on pyroxene trace element chemistry in EC chondrules to date, Hsu and Crozaz (1998) found various levels of overabundance in light REE (relative to equilibrium pyroxene-melt partitioning expectations) which they ascribe to glass contamination. This may explain the flat patterns reported for enstatite by other workers (Weisberg et al., 1994; Crozaz and Lundberg, 1995; Gannoun et al., 2011) and for ferroan pyroxene by Jacquet et al. (2015). Two porphyritic olivine-pyroxene (POP) chondrules as well as ferroan pyroxene spherules analyzed in Sahara 97096 exhibit negative REE pattern slopes which are as yet unexplained (Jacquet et al., 2015). The relative ordinariness of the silicates means that the bulk compos- itional peculiarities of EC chondrules are rather reflected in the modal mineralogy of the chondrules (e.g., the pyroxene/olivine ratio, sulfides) and the composition of the mesostasis (the main carrier of many incompatible elements). Compared to its ordinary chondrite counter- parts, the latter is indeed depleted in Ca, Eu, Yb, Mn (Jacquet et al., 2015), and enriched in S (0.2–0.5 wt%; Piani et al., 2016), Cl (0–2 wt%; Grossman et al., 1985; Piani et al., 2016) and alkalis such as Na – more so in EH3 (5–13 wt% Na2O; Grossman et al., 1985) than in EL3 (0–4 wt%; Schneider et al., 2002) although Manzari (2010) finds up to 9 wt% in Elephant Moraine 902299 (EL3). Oldhamite, whether inside or outside silicate chondrules, is typically enriched by 1–2 orders of magnitude in REE relative to CI, with frequent positive anomalies in Eu and Yb in EH3 and negative Eu anomalies in EL3 chondrites (Larimer and Ganapathy, 1987; Crozaz and Lundberg, 1995; Gannoun et al., 2011; El Goresy et al., 2017; Figure 7.3). Niningerite shows nearly chondritic heavy REE contents with monotonic depletion in light REE (Crozaz and Lundberg, 1995; Gannoun et al., 2011; Figure 7.3).

7.4 Isotopic Composition

Oxygen isotope ratios of chondrule olivine and pyroxene in EH3 chondrite (and the anomalous Lewis Cliff 87223) show a wide range of values as large as 10 ‰ in both δ18O and δ17O. Most chondrules plot along the terrestrial fractionation (TF) line on a 3-isotope diagram (Figure 7.4a). However, some chondrules overlap the OC field and some extend toward more 16O-rich compositions. Tanaka and Nakamura (2017) interpreted the range of whole-chondrule O isotopic compositions they measured (including a silica-rich chondrule with a record-high δ18O = 6.846 ‰) by varying mixtures between a carbonaceous chondrite-like precursor and nebular gas in the chondrule-forming environments. Most of the FeO-rich pyroxene in E3 chondrites plot on the TF line similar to enstatite, suggesting it formed locally in the EC region, indicating variable oxidation/reduction conditions within the E3 chondrule-forming region (Kimura et al., 2003; Weisberg et al., 2011). Olivine in some E3 chondrules shows a wide range of Δ17O values (À4 ‰ to +2 ‰) and forms a distinct mixing line approximately parallel to, but displaced from, the carbonaceous chondrite anhydrous mineral (CCAM) line (Figure 7.4a). Weisberg et al. (2011) referred to this as the “enstatite chondrite mixing line,” but it is similar to the “primitive chondrule minerals line” defined by chondrules from the Acfer 094 chondrite (Ushikubo et al., 2012). Weisberg et al. (2011) showed that some chondrules in EH3 chondrites have oxygen isotope heterogeneity among their minerals. These are interpreted to indicate incomplete melting of the chondrules, survival of minerals from previous generations of chondrules, and chondrule 182 Emmanuel Jacquet, Laurette Piani, and Michael K. Weisberg

Figure 7.4 (a) Oxygen 3-isotope diagram showing SIMS data for chondrules in E3 (EH3 and the anomalous EC Lewis Cliff 87223), LL3 and C3 chondrites. Also shown is the range of chondrule compositions in R3 chondrites and the Allende CV3 chondrite and the terrestrial fractionation (TF), carbonaceous chondrite anhydrous mineral (CCAM) and enstatite chondrite mixing (ECM) lines. Chondrules from E3 chondrites have compositions that overlap those from LL3 and R3 chondrites but their distribution appears to be unique with most data points clustered around the TF line. The E3 chondrules also fall along a line referred to as the ECM. The ECM is similar and possibly the same as the primitive chondrite mixing (PCM) line defined by chondrules in the primitive chondrite Acfer 094. (b) Δ17O versus δ18O for chondrules in E3 chondrites compared to those of chondrules in LL3 and CR3 chondrites. Data are from Weisberg et al. (2011); Kita et al. (2010); Nakashima et al. (2010); Rudraswami et al. (2011); Tenner et al. (2013); Tenner et al. (2015). (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.) Chondrules in Enstatite Chondrites 183 recycling. Weisberg et al. (2011) found in particular one chondrule with a relict grain with a high, R chondrite-like Δ17O value, suggesting limited admixtures of materials from other reservoirs. From their oxygen isotope values, the chondrules in E chondrites seem to represent a distinct population formed in a local nebular reservoir different from chondrules in other groups (see e.g., Chapter 8) but most closely related to ordinary and R chondrites. This is also shown by their Δ17O values, which overlap with the range of compositions from LL3 and R3 chondrules. However, the majority of E3 chondrules have values close to 0 (i.e., Earth–Moon). In compari- son, most LL3 and R3 chondrules have higher Δ17O values while those in carbonaceous chondrites generally have negative Δ17O values (Figure 7.4b). The isotopic composition of elements other than oxygen (e.g., Ca, Ti, Cr, Ni, Mo, Ru, W, Nd; see Burkhardt et al. (2017) and references therein) has been measured for bulk enstatite chondrites and generally found indistinguishable from the Earth’s. There are, however, few such data on individual chondrules. Gerber et al. (2017) found that the Ti isotope compositions of EC chondrules were fairly uniform, around the terrestrial value, distinct from that of the (equally uniform) ordinary chondrite chondrules. This contrasts with the range exhibited in single carbonaceous chondrites, attributed by them to variable admixtures in their precursors of CAI-like material, presumably lacking in the enstatite chondrite reservoir.

7.5 A Special Condensation Sequence …

Clearly, the reduced mineralogy of enstatite chondrite chondrules (MSN included) sets them apart from other chondrite clans, but at which stage did their specificities arise? Was it the formation of their precursors – presumably some nebular condensation sequence – or/and the chondrule melting process itself? Much of the past literature has envisioned the first option (Grossman et al., 2008; Lehner et al., 2013 and references therein). Some of it, in fact, envisioned chondrules in general, and those of EC in particular, as direct condensates (e.g., Blander et al., 2009; Varela et al., 2015), although support for this scenario has lately decreased, given the instability of liquids at the low total pressures expected in average protoplanetary disks (Ebel and Grossman, 2000), the presence of relict grains and prevalence of porphyritic textures pointing to solid precursors (Hewins et al., 2005) or the few-Ma age spread of chondrules (Connelly et al., 2012; Kita et al., 2013), indicating repeated, short-lived formation episodes well beyond the “hot (inner) solar nebula” epoch. It is nonetheless also conceivable that chondrule-forming episodes essentially preserved the composition of the EC chondrule precursors, e.g., because of short timescales which would have averted devolatilization (espe- cially in sulfur; Grossman et al., 1985) and for which Rubin (2010) finds supporting evidence in the smallness and lack of dust mantle of EC chondrules (in contradistinction to CV and CR type I chondrules, for example). The EC metal-sulfide nodules could also be nebular condensates (Rambaldi et al., 1986; Larimer and Ganapathy, 1987; El Goresy et al., 1988; Kimura, 1988; Ikeda, 1989a; Lin and El Goresy, 2002; Lehner et al., 2010), more or less affected by remelting and/or sulfidation events, which may account for reverse zoning in some niningerite crystals (i.e., FeS increasing outward; Skinner and Luce, 1971; El Goresy et al., 1988). Which condensation conditions would be required specifically? It has long been known that a condensation sequence of a solar mixture cannot account for the mineralogy of EC, in 184 Emmanuel Jacquet, Laurette Piani, and Michael K. Weisberg particular the dominance of pyroxene over olivine (even at low temperatures or pressures; Pignatale et al., 2016) or their array of sulfides – unless, perhaps, one assumes very high pressures (e.g., 1 bar) and iron supersaturation (Blander, 1971; Herndon and Suess, 1976; Blander et al., 2009). The stability of the reduced phases specific to EC is, however, possible if the C/O ratio is increased over the solar value of 0.5 (Larimer, 1968; Baedecker and Wasson, 1975; Larimer and Bartholomay, 1979). This is because the stable species CO hereby locks an increasing fraction of the oxygen, hence lower oxygen fugacities. However, C/O ratios above ~1 would entail the condensation of phases such as carbides, graphite, and nitrides, which are scarce in ECs (Larimer and Bartholomay, 1979; Grossman et al., 2008); even though the then higher-temperature condensation of oldhamite would allow it to concentrate REE at the observed levels (Lodders and Fegley, 1993), albeit with Eu and Yb anomalies of the wrong sign (negative) compared to observations (Gannoun et al., 2011). The relatively modest enhancements of Si contents in kamacite compared to carbonaceous chondrite metal also favor more moderate C/O ratios, perhaps around 0.8 (Grossman et al., 2008). The subsolar Mg/Si ratios and refractory element abundances (e.g., Al/Si ratio) of EC chondrules (and whole-rocks), common to other noncarbonaceous chondrites, require fractiona- tions among nonvolatile species too. Baedecker and Wasson (1975) and Hutson and Ruzicka (2000) postulated the loss of an olivine-rich refractory condensate. Petaev and Wood (1998) obtained this effect in a variant of fractional condensation calculation where 0.7–1.5 percent of the solids equilibrating with the gas were isolated from it after each degree of cooling. Lehner et al. (2013) proposed Mg/Si fractionation of enstatite chondrites was associated with the volatility of Mg in niningerite, which would, however, require another mechanism for the other noncarbonaceous chondrite clans (Jacquet et al., 2015).

7.6 … or Special Chondrule-Forming Conditions?

Evidence is, however, mounting that EC chondrule precursors may actually have been less than remarkable. First, there is no evidence for a particular CAI population: The rare CAIs in E3 chondrites are petrographically and isotopically similar to those in other chondrite groups; in particular, their oxygen isotopic compositions, even though statistically indistinguishable from the enstatite chondrite mixing line (Weisberg et al., 2011), are very 16O-rich and plot on the CCAM line, similar to CAIs in carbonaceous chondrites (Guan et al., 2000a; Guan et al., 2000b; Fagan et al., 2001; Lin et al., 2003). Second, the FeO-rich pyroxene and most relict olivine grains, with their fairly unexceptional chemistries (see Section 7.3), share the same O isotope signature as the bulk EC (see Section 7.4); thus, they cannot be mere interlopers from other reservoirs, and more likely represent a population of EC chondrule precursors (or perhaps, an early stage of the chondrule-forming process itself ). Third, Simon et al. (2016) found that 10–70 percent of the Ti in E3 chondrules was tetravalent, although it is predicted to be largely trivalent for oxygen fugacities equal to or lower than solar, indicating relatively oxidized precursors whose Ti valence was not efficiently reset. As to metal-sulfide nodules, the patterns of siderophile elements measured in the metal of EH and EL MSNs do not correspond to those expected for a condensate, but favor metal crystallization from a melt (Horstmann et al., 2014). Silicate laths intergrown with metal or sulfides in EL3 MSNs have also been interpreted as the results of melt crystallization, with van Chondrules in Enstatite Chondrites 185

Niekerk and Keil (2011) suggesting impact on the parent body as the cause of the melting. Following Weisberg et al. (1997), Horstmann et al. (2014), noting the absence of textural indications for in situ melting (absence of melt pockets, FeS veins, or intergrowths of the MSN phases within the matrix) or significant REE redistribution between silicates and sulfides (Barrat et al., 2014), preferred impact prior to accretion. However, El Goresy et al. (2017) argue against an igneous origin of the MSN on the basis of the delicate zoning (e.g., in C isotopes) of constituent grains or lack of quench textures. Still, euhedral pyroxene/metal intergrowths reminiscent of those are present in EH impact melt breccias like Abee and absent in unmelted EH3s (Rubin and Scott, 1997), although their exact formation mechanism (as melted target material? or direct (liquid?) condensates from an impact plume? and so on) is unknown. As to MSN in EH3 chondrites, the presence in oldhamite of numerous inclusions and intergrowths of silicate phases, often with euhedral shapes, also argues for an igneous origin (Hsu, 1998). A further clue may be provided by their complementarity with silicate chondrules (Lehner et al., 2014; Ebel et al., 2015), in particular for siderophile elements, Ca (all depleted in the chondrules) and rare earth elements (with opposite Eu and Yb anomalies; Figure 7.3). The petrographic continuum between silicate-bearing metal-sulfide nodules to sulfide-bearing silicate chondrules (Lehner et al., 2014) also suggests a genetic link. Specifically, McCoy et al. (1999) suggested that the MSN were lost as immiscible droplets from the chondrules, in the same way opaques were likely expelled from molten chondrules in other groups, whether through surface tension or inertial acceleration effects (Grossman and Wasson, 1985; Campbell et al., 2005; Jacquet et al., 2013), and may have further evolved by interaction with the gas. This has also been recently suggested for “sulfide chondrules” in unequilibrated R chondrites (Miller et al., 2017), and could be entertained for the MSN in LL3 chondrites (Weisberg et al., 2016) as well. Jacquet et al. (2015) proposed that igneous partitioning between oldhamite and the liquid chondrule mesostasis may explain the high REE content of oldhamite (as suggested for aubritic oldhamite by Dickinson and McCoy (1997)) as well as its frequent positive Eu and Yb anomalies if these elements were in divalent form. This would, of course, also account for the complementary negative Eu and Yb anomalies and Ca depletion of chondrule mesostases. Interestingly, a few ordinary chondrite chondrules present negative Eu, Yb, and sometimes Sm anomalies (Pack et al., 2004; Ebert and Bischoff, 2016; Metzler and Pack, 2016) and are all CaO-poor (<0.8 wt%), suggesting that they, or at least their precursors, underwent similar fractionations. While Metzler and Pack (2016) objected that loss of oldhamite (light (L) REE-enriched) would have incurred LREE depletion which are not observed in the chondrules, (i) the LREE enrichment of oldhamite is in fact moderate and likely further alleviated by accompanying phases such as (LREE-depleted) niningerite and (ii) the EC chondrule mesostases, in lacking the marginal LREE enrichment seen in other chondrite clans, do evidence some degree of relative LREE depletion (see Figure 7.3). It is noteworthy that the EC chondrule mesostases’ Eu and Yb anomalies are not reflected in the silicates, suggesting that the latter were already crystallized (whether formed at the beginning of the high-temperature episode in question or as relict of an earlier episode; Jacquet et al., 2015) when the opaque assemblages formed. Indeed, experimental heating in closed system of the Indarch EH4 chondrite indicates that the metal-sulfide components completely melt at   1,000 C(McCoyetal.,1999)andnosulfide (solid or molten) remains above 1400 Cinthe presence of a silicate melt (Fogel et al., 1996; Fogel, 1998), suggesting moderate temperatures at this stage. 186 Emmanuel Jacquet, Laurette Piani, and Michael K. Weisberg

What would the origin of sulfides in EC chondrules then be? The absence of niningerite and oldhamite enclosed in low-Ca pyroxene of the Sahara 97096 porphyritic chondrules indicates that these sulfides are not preserved from the chondrule precursors, unlike (possibly) olivine or Fe–Ni metal grains (Piani et al., 2016). In order to avoid a rapid loss of sulfur from the molten chondrules (Yu et al., 1996; Uesugi et al., 2005), a high ambient partial pressure of sulfur is necessary (e.g., Marrocchi and Libourel, 2013) – such as has long been suggested to explain the opaque paragenesis of enstatite chondrites (e.g., Fleet and MacRae, 1987; El Goresy et al., 1988; Zhang et al., 1995; Lin and El Goresy, 2002; Lehner et al., 2013). The stability of metal-sulfide mixtures, as a buffer of the S2 fugacity, has been used by Lehner et al. (2013) to suggest S/H ratios 2,500–10,000 times higher than solar. Then, the sulfidation of silicates (olivine and pyroxene) by a S-rich gas at high temperatures (> 1,000 C) could result in the formation of Fe- , Ca- , or Mg-sulfides (Rubin, 1983; Fleet and MacRae, 1987; Lehner et al., 2013). Piani et al. (2016) noted that under the temperatures required for the sulfidation of solid silicates, with the chondrules being partially molten, the high partial pressure of sulfur should have interacted with the silicate melt. Thus sulfur would have dissolved in the silicate melt (hence the high S content of the mesostasis) and eventually saturated, forming immiscible sulfide melts (e.g., Fincham and Richardson, 1954; McCoy et al., 1999). The involvement of melt in the formation of niningerite may explain the large amounts of coexisting silica as well (Piani et al., 2016). Another question is the origin of the high pyroxene/olivine ratio in EC chondrules compared to other chondrite clans. It may to some extent be a mechanical result of the bulk Mg/Si fractionation of the EC reservoir, whatever its origin (see page 184). It is, however, noteworthy that olivine in chondrules from other clans is predicted to dissolve within minutes in a melt of the composition measured in the mesostasis (Soulié et al., 2017). Libourel et al. (2006) suggested that the melt was enriched in Si by influx of SiO from the ambient gas, promoting pyroxene crystallization at the expense of olivine, and only a rapid cooling rate at that stage saved olivine (often rounded and partly resorbed) from wholesale dissolution. Then the lower olivine content in EC chondrules may perhaps reflect longer exposures to a SiO-rich gas and/or higher partial pressures of SiO compared to other chondrule-forming reservoirs. This has yet to be investigated. In sum, it is conceivable that the specificities of EC chondrules (MSN included) were acquired from processing of relatively (chemically) “normal” precursors in unusually S (and other volatile-)-rich and O-poor environments. Then, the conditions of the condensation models discussed in Section 7.5 may actually pertain to the EC chondrule-forming regions. Indeed, thermodynamic equilibrium (to the extent it was approached by the chondrules) predictions are independent of the history of the reservoir. Whatever that may be, we have yet to find an astrophysical context for these reducing conditions.

7.7 Fractionation Mechanisms

Although Grossman et al. (2008) entertain the possibility of primordial heterogeneities inherited from the presolar cloud, the weakness of the observed isotopic nucleosynthetic anomalies, even in refractory inclusions (Birck, 2004), leaves little prospect of significant relative chemical Chondrules in Enstatite Chondrites 187 anomalies at disk formation. One thus has to rely on dynamical processes within the proto- planetary disk to explain the nonsolar characteristics of the enstatite chondrite reservoir(s).

One way to increase the C/O ratio (or reduce H2O/H2) is a loss of rock and (if condensed) ice through gas–solid interaction. This could be effected by, for example, settling to the midplane (Ikeda, 1989b; Krijt et al., 2016), or radial drift sunward (see simulations by Ciesla and Cuzzi, 2006; Estrada, Cuzzi, and Morgan, 2016). Another possibility would be efficient accretion of water and rock diffused beyond the snow line – a “cold finger”–, which, if more rapid than inward radial transport, could deplete the inner disk in water by orders of magnitude (Ciesla and Cuzzi, 2006). Perhaps the S enrichment seen by enstatite chondrites would have been achieved beyond the troilite condensation front by similar “cold trapping” of troilite there as well (Pasek et al., 2005). Yet Grossman et al. (2008) object that these different oxygen-depleted regions would be even more depleted in the other condensable elements, and thus unsuitable to form EC components. Still, radial drift of early coarse refractory condensates would be a way to achieve the nonvolatile element fractionations of EC (Larimer and Wasson, 1988; Jacquet et al., 2012). A quite different scenario toward the required reducing conditions would be a regional enhancement,by2–3 orders of magnitude, in a dry mixture of rock and organic matter (some being still preserved in EC; see Piani et al. (2012) and references therein), e.g., an analog of “Chondritic Porous” (CP) interplanetary dust particles (Wood and Hashimoto, 1993; Petaev and Wood, 2001; Ebel and Alexander, 2005). Indeed, even though the global C/O ratio may not necessarily be supersolar, the condensation of silicates would drastically deplete the gas in oxygen. This may also be a means to yield the sulfur enhancements proposed by Lehner et al. (2013). This requires a process which can efficiently fractionate water from the other condens- able elements, most likely inside the snow line where the former is gaseous and the latter solid. Radial drift appears unable to produce enrichments above one order of magnitude in the inner disk (Ciesla and Cuzzi, 2006; Estrada et al., 2016) and would likely include water from the outer disk (Jacquet and Robert, 2013). One may thus be left with local processes such as turbulent concentration (Cuzzi et al., 2001) or vertical settling, the effect of which seems to be limited to one order of magnitude each in current models (Jacquet et al., 2012), although disk astrophysics admittedly remains quite uncertain. It is, however, not necessary to restrict attention to “background disk” (“nebular”) environ- ments. The needed enrichments in condensable elements may indeed come about in a planetary setting (Piani et al., 2016), e.g., an impact plume (Fedkin and Grossman, 2013) on a reduced (chemically EC-like?) parent body (such as already inferred for metal-silicate intergrowths in EL3 chondrites; e.g., Horstmann et al., 2014). This, however, has yet to be investigated in detail from the modeling point of view in the specific context of enstatite chondrites.

7.8 Enstatite Chondrite Chondrules in Space and Time

However mysterious their context of formation remains, can we infer anything about where and when chondrules in enstatite chondrites formed? An origin inside the snow line, as suggested just above, would be consistent with the lack of aqueous alteration in enstatite chondrites. Asteroids most spectroscopically consistent with enstatite chondrites, namely those of the 188 Emmanuel Jacquet, Laurette Piani, and Michael K. Weisberg

E type, are most prevalent in the inner and middle parts of the main belt (DeMeo and Carry, 2014). It is nonetheless unclear whether they lie significantly sunward of their S-type counter- parts, for the most conspicuous E asteroid population, among the Hungarias (inside 2 AU), may largely constitute a single Hirayama family (Milani et al., 2010), that is one initial parent body only. Yet, the putative mineralogy of Mercury is notoriously reminiscent of enstatite chondrites (McCoy and Nittler, 2014), suggesting similar building blocks. Another hint at an inner Solar System provenance is the isotopic similarity in many elements of enstatite chondrites (and their chondrules) of both chemical groups to the Earth and the Moon. With the aubrites and the brachinite-like ungrouped achondrite Northwest Africa 5363, they may represent an “inner disk (isotopically) uniform reservoir” (Dauphas et al., 2014). Radial homogenization would indeed be more rapid at smaller heliocentric distances, given the shorter spatial scales and higher turbulent diffusivity. The great degree of isotopic homogenization in the EC reservoir may also indicate a formation epoch later than other chondrite clans. This was also proposed by Jacquet et al. (2012) to allow for the Mg/Si and refractory element fractionation. Perhaps the relatively high proportion of type 3 in enstatite chondrites (especially EH) compared to ordinary chondrites (Scott and Krot, 2014) may be traced back to then-weaker heat sources in their parent bodies at the time of their accretion. Direct radiochronological data on chondrules, while consistent with a late formation, are, however, scarce: Individual I-Xe dating of EH3 chondrite chondrules yields ages of 4,561–4,564 Ma (Whitby et al., 2002); Guan et al. (2006) found evidence of in situ decay of 26Al in only one out of 13 Al-rich objects in unequilibrated enstatite chondrites, but suggested disturbances by incipient metamorphism; and Trieloff et al. (2013) dated sphalerites in Sahara 97096 (EH3) by Mn–Cr at 4562.7Æ0.5 Ma.

7.9 Outlook

At the close of this tour of enstatite chondrite chondrule literature, we as a community are actually far from the end of the journey regarding the understanding of these objects. Still, with the progress accomplished in recent years, we may be able to ask better questions. So let us list some as concluding thoughts:

– Can we test further the possible link between silicate chondrules and the metal-sulfide nodules in enstatite chondrites? – What are the petrographic characteristics of sulfides in EL3 chondrite chondrules and their relationships with the silicates as compared to EH3? – Under which redox conditions can experimental partition coefficients between oldhamite (and other sulfides) and silicate melts reproduce observations, especially as to the REE signatures? Is a condensate origin still possible for such signatures considering the observed overall mineralogy of enstatite chondrites? – Why is olivine so rare in EC chondrules? Was it dissolved following Si influx in the chondrules during formation, and if so, what would the effect on their O isotopic composition be? – Is it possible to increase the (water-free) condensable/gas ratio by 2–3 orders of magnitude above solar in the protoplanetary disk, as suggested by the S-rich composition of EC chondrules? Chondrules in Enstatite Chondrites 189

– Could such conditions be achieved in a planetary environment (e.g., an impact plume on a parent body with reduced mineralogy)? – Did EH and EL silicate chondrules and metal-sulfide nodules form by similar processes? – How do EC chondrules fit in the absolute and relative chronology of chondrule formation in other clans? – What are the parent bodies of EC chondrites? Where and when did they accrete?

Acknowledgments

We thank the anonymous referee and the A.E. Dr. Connolly for their reviews, which improved the precision of the manuscript.

References

Baedecker, P. A., and Wasson, J. T. (1975). Elemental fractionations among enstatite chon- drites. Geochimica et Cosmochimica Acta, 39, 735–765. Barrat, J. A., Zanda, B., Jambon, A., and Bollinger, C. (2014). The lithophile trace elements in enstatite chondrites. Geochimica et Cosmochimica Acta, 128,71–94. Birck, J. (2004). An overview of isotopic anomalies in extraterrestrial materials and their nucleosynthetic heritage. Reviews in Mineralogy and Geochemistry, 55,25–64. Washing- ton, D.C.: Mineralogical Society of America. Blander, M. (1971). The constrained equilibrium theory: Sulphide phases in meteorites. Geo- chimica et Cosmochimica Acta, 35,61–76. Blander, M., Pelton, A. D., and Jung, I. -H. (2009). A condensation model for the formation of chondrules in enstatite chondrites. Meteoritics and Planetary Science 44, 531–543. Brearley, A. J., and Jones, R. H. (1998). Chondritic meteorites. In J. J. Papike (Ed.), Planetary Materials. Reviews in Mineralogy and Geochemistry, 36, 3-1–3-398. Washington, D.C.: Mineralogical Society of America. Burkhardt C., Dauphas N., Tang H., et al. (2017). In search of the Earth-forming reservoir: Mineralogical, chemical and isotopic characterizations of the ungrouped achondrite NWA 5363/NWA 5400 and selected chondrites. Meteoritics and Planetary Science, 52, 807–826. Campbell, A. J., Zanda, B., Perron, C., Meibom, A., and Petaev, M. I. (2005). Origin and Thermal History of Fe-Ni Metal in Primitive Chondrites. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, 407–431. Astronomical Society of the Pacific Conference Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Ciesla, F. J., and Cuzzi, J. N. (2006). The evolution of the water distribution in a viscous protoplanetary disk. Icarus, 181, 178–204. Connelly, J., Bizzarro, M., Krot, A., et al. (2012). The absolute chronology and thermal processing of solids in the solar protoplanetary disk. Science, 338, 651. Crozaz, G., and Lundberg, L. L. (1995). The origin of oldhamite in unequilibrated enstatite chondrites. Geochimica et Cosmochimica Acta, 59, 3817–3831. Cuzzi, J. N., Hogan, R. C., Paque, J. M., and Dobrovolskis, A. R. (2001). Size-selective Concentration of Chondrules and Other Small Particles in Protoplanetary Nebula Turbu- lence. The Astrophysical Journal, 546, 496–508. Dauphas, N., Burkhardt, C., Warren, P. H. , and Teng, F. -Z. (2014). Geochemical arguments for an Earth-like Moon-forming impactor. Philosophical Transactions of the Royal Society A, 372(2024). 20130244. 190 Emmanuel Jacquet, Laurette Piani, and Michael K. Weisberg

DeMeo, F. E., and Carry, B. (2014). Solar System evolution from compositional mapping of the asteroid belt. Nature, 505, 629–634. Dickinson, T. L., and McCoy, T. J. (1997). Experimental REE partitioning in oldhamite: Implications for the igneous origin of aubritic oldhamite. Meteoritics and Planetary Science, 32, 395–412. Ebel, D. S., and Alexander, C. M. O’D. (2005). Condensation from cluster-IDP enriched vapor inside the snow line: Implications for mercury, asteroids, and enstatite chondrites. In S. Mackwell and E. Stansbery (Eds.), 36th Lunar and Planetary Science Conference, #1797. Ebel, D. S., Boyet, M., Hammouda, T., et al. (2015). Complementary rare earth element abundances in enstatite and oldhamite in EH3 chondrites. In 46th Lunar and Planetary Science Conference, #1832. Ebel, D. S., and Grossman, L. (2000). Condensation in dust-enriched systems. Geochimica et Cosmochimica Acta, 64, 339–366. Ebert, S., and Bischoff, A. (2016). Genetic relationship between Na-rich chondrules and Ca,Al- rich inclusions? – Formation of Na-rich chondrules by melting of refractory and volatile precursors in the solar nebula. Geochimica et Cosmochimica Acta, 177, 182–204. El Goresy, A., Lin, Y., Miyahara, M., et al. (2017). Origin of EL3 chondrites: Evidence for variable C/O ratios during their course of formation–A state of the art scrutiny. Meteoritics and Planetary Science, 52, 781–806. El Goresy, A., Yabuki, H., Ehlers, K., Woolum, D., and Pernicka, E. (1988). Qingzhen and Yamato-691: A tentative alphabet for the EH chondrites. Antarctic Meteorite Research, 1, 65. Estrada, P. R., Cuzzi, J. N., and Morgan, D. A. (2016). Global modeling of nebulae with particle growth, drift, and evaporation fronts. I. methodology and typical results. The Astrophysical Journal, 818, 200. Fagan, T. J., McKeegan, K. D., Krot, A. N., and Keil, K. (2001). Calcium-aluminum-rich inclusions in enstatite chondrites (II): Oxygen isotopes. Meteoritics and Planetary Science, 36, 223–230. Fedkin, A., and Grossman, L. (2013). Vapor saturation of sodium: Key to unlocking the origin of chondrules. Geochimica et Cosmochimica Acta, 112, 225–250. Fincham, C. J. B., and Richardson, F. D. (1954). The behaviour of sulphur in silicate and aluminate melts. Proceedings of the Royal Society of London, Series A, 223,40–62. Fleet, M. E., and MacRae, N. D. (1987). Sulfidation of Mg-rich olivine and the stability of niningerite in enstatite chondrites. Geochimica et Cosmochimica Acta, 51, 1511–1521. Fogel, R. A. (1998). High-sulfur/low-iron silicate melts: Low-oxygen-fugacity phenomena of importance to aubrite formation. Meteoritics and Planetary Science Supplement, 33, A52. Fogel, R. A., Weisberg, M. K., and Prinz, M. (1996). The solubility of CaS in aubrite silicate melts. In 27th Lunar and Planetary Science Conference, p. 371. Gannoun, A., Boyet, M., El Goresy, A., and Devouard, B. (2011). REE and actinide micro- distribution in Sahara 97072 and ALHA77295 EH3 chondrites: A combined cosmochem- ical and petrologic investigation. Geochimica et Cosmochimica Acta, 75, 3269–3289. Gerber, S., Burkhardt, C., Budde, G., Metzler, K., Kleine, T. (2017). Mixing and transport of dust in the early solar nebula as inferred from titanium isotope variations among chondrules. The Astrophysical Journal Letters, 841, L17. Grossman, J. N., Rubin, A. E., Rambaldi, E. R., Rajan, R. S., and Wasson, J. T. (1985). Chondrules in the Qingzhen type-3 enstatite chondrite: Possible precursor components and comparison to ordinary chondrite chondrules. Geochimica et Cosmochimica Acta, 49, 1781–1795. Grossman, J. N., and Wasson, J. T. (1985). The origin and history of the metal and sulfide components of chondrules. Geochimica et Cosmochimica Acta, 49, 925–939. Chondrules in Enstatite Chondrites 191

Grossman, L., Beckett, J. R., Fedkin, A. V., Simon, S. B., and Ciesla, F. J. (2008). Redox conditions in the solar nebula: Observational, experimental and theoretical constraints. In G. J. MacPherson (Ed.), Oxygen in the Solar System. Reviews in Mineralogy and Geo- chemistry, 68, 93–140. Washington, D.C.: Mineralogical Society of America. Guan, Y., Huss, G. R., Leshin, L. A., MacPherson, G. J., and McKeegan, K. D. (2006). Oxygen isotope and 26Al-26Mg systematics of aluminum-rich chondrules from unequilibrated enstatite chondrites. Meteoritics and Planetary Science, 41,33–47. Guan, Y., Huss, G. R., MacPherson, G. J., and Wasserburg, G. J. (2000a). Calcium-aluminum- rich inclusions from enstatite chondrites: indigenous or foreign? Science, 289, 1330–1333. Guan, Y., McKeegan, K. D., and MacPherson, G. J. (2000b). Oxygen isotopes in calcium- aluminum-rich inclusions from enstatite chondrites: New evidence for a single CAI source in the solar nebula. Earth and Planetary Science Letters, 181, 271–277. Herndon, J. M., and Suess, H. E. (1976). Can enstatite meteorites form from a nebula of solar composition. Geochimica et Cosmochimica Acta, 40, 395–399. Hewins, R. H., Connolly H. C. Jr, Lofgren, G. E., and Libourel, G. (2005). Experimental constraints on chondrule formation. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, 286–316. Astronomical Society of the Pacific Conference Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Horstmann, M., Humayun, M., and Bischoff, A. (2014). Clues to the origin of metal in Almahata Sitta EL and EH chondrites and implications for primitive E chondrite thermal histories. Geochimica et Cosmochimica Acta, 140, 720–744. Hsu, W. (1998). Geochemical and petrographic studies of oldhamite, diopside, and roedderite in enstatite meteorites. Meteoritics and Planetary Science, 33, 291–301. Hsu, W., and Crozaz, G. (1998). Mineral chemistry and the origin of enstatite in unequilibrated enstatite chondrites. Geochimica et Cosmochimica Acta, 62, 1993–2004. Hutson, M., and Ruzicka, A. (2000). A multi-step model for the origin of E3 (enstatite) chondrites. Meteoritics and Planetary Science, 35, 601–608. Ikeda, Y. (1988). Petrochemical study of the Yamato-691 enstatite chondrite (EH3) I: Major element chemical compositions of chondrules and inclusions. Proceedings of the NIPR Symposium on Antarctic Meteorite Research, 2, 147–165. Ikeda, Y. (1989a). Petrochemical study of the Yamato-691 enstatite chondrite (E3) III: Descriptions and mineral compositions of chondrules. Antarctic Meteorite Research, 2, 75–108. Ikeda, Y. (1989b). Petrochemical study of the Yamato-691 enstatite chondrite (E3) V: Com- parison of major element chemistries of chondrules and inclusions in Y-691 with those in ordinary and carbonaceous chondrites. Antarctic Meteorite Research 2, 147–165. Jacquet, E., Alard, O., and Gounelle, M. (2015). The formation conditions of enstatite chon- drites: Insights from trace element geochemistry of olivine-bearing chondrules in Sahara 97096 (EH3). Meteoritics and Planetary Science, 50, 1624–1642. Jacquet, E., Gounelle, M., and Fromang, S. (2012). On the aerodynamic redistribution of chondrite components in protoplanetary disks. Icarus, 220, 162–173. Jacquet, E., Paulhiac-Pison, M., Alard, O., and Kearsley, A. (2013). Trace element geochemis- try of CR chondrite metal. Meteoritics and Planetary Science, 48, 1981–1999. Jacquet, E., and Robert, F. (2013). Water transport in protoplanetary disks and the hydrogen isotopic composition of chondrites. Icarus, 223, 722–732. Jones, R. H. (2012). Petrographic constraints on the diversity of chondrule reservoirs in the protoplanetary disk. Meteoritics and Planetary Science, 47, 1176–1190. Keil, K. (1968). Mineralogical and chemical relationships among enstatite chondrites. Journal of Geophysical Research, 73, 6945–6976. Kimura, M. (1988). Origin of opaque minerals in an unequilibrated enstatite chondrite Yamato- 691, Antarctic Meteorite Research, 1,51–64. 192 Emmanuel Jacquet, Laurette Piani, and Michael K. Weisberg

Kimura, M., Hiyagon, H., Lin, Y., and Weisberg, M. K. (2003). FeO-rich silicates in the Sahara 97159 (EH3) enstatite chondrite: Mineralogy, oxygen isotopic compositions, and origin. Meteoritics and Planetary Science, 38, 389–398. Kimura, M., Weisberg, M. K., Lin, Y., et al. (2005). Thermal history of the enstatite chondrites from silica polymorphs. Meteoritics and Planetary Science, 40, 855–868. Kita, N. T., Nagahara, H., Tachibana, S., et al. (2010). High precision SIMS oxygen three isotope study of chondrules in LL3 chondrites: Role of ambient gas during chondrule formation. Geochimica et Cosmochimica Acta, 74, 6610–6635. Kita, N. T., Yin, Q. -Z., MacPherson, G. J., et al. (2013). 26Al-26Mg isotope systematics of the first solids in the early Solar System. Meteoritics and Planetary Science, 48, 1383–1400. Krijt, S., Ciesla, F. J., and Bergin, E. A. (2016). Tracing water vapor and ice during dust growth. The Astrophysical Journal, 833, 285–298. Larimer, J. W. (1968). An experimental investigation of oldhamite, CaS; and the petrologic significance of oldhamite in meteorites. Geochimica et Cosmochimica Acta, 32, 965–982. Larimer, J. W., and Bartholomay, M. (1979). The role of carbon and oxygen in cosmic gases – Some applications to the chemistry and mineralogy of enstatite chondrites. Geochimica et Cosmochimica Acta, 43, 1455–1466. Larimer, J. W., and Ganapathy, R. (1987). The trace element chemistry of CaS in enstatite chondrites and some implications regarding its origin. Earth and Planetary Science Letters, 84, 123–134. Larimer, J. W., and Wasson, J. T. (1988). Refractory lithophile elements. In J. F. Kerridge and M. S. Matthews (Eds.), Meteorites and the Early Solar System, 394–415. Tucson, AZ: University of Arizona Press. Lehner, S. W., Buseck, P. R., and McDonough, W. F. (2010). Origin of kamacite, schreibersite, and perryite in metal-sulfide nodules of the enstatite chondrite Sahara 97072 (EH3). Meteoritics and Planetary Science, 45, 289–303. Lehner, S. W., McDonough, W. F., and NéMeth, P. (2014). EH3 matrix mineralogy with major and trace element composition compared to chondrules. Meteoritics and Planetary Sci- ence, 49, 2219–2240. Lehner, S. W., Petaev, M. I., Zolotov, M. Y., and Buseck, P. R. (2013). Formation of niningerite by silicate sulfidation in EH3 enstatite chondrites. Geochimica et Cosmochimica Acta, 101, 34–56. Libourel, G., Krot, A. N., and Tissandier, L. (2006). Role of gas-melt interaction during chondrule formation. Earth and Planetary Science Letters, 251, 232–240. Lin, Y., and El Goresy, A. (2002). A comparative study of opaque phases in Qingzhen (EH3) and MacAlpine Hills 88136 (EL3): Representatives of EH and EL parent bodies. Meteori- tics and Planetary Science, 37, 577–599. Lin, Y., El Goresy, A., Boyet, M., et al. (2011). Earliest solid condensates consisting of the assemblage Oldhamite, Sinoite, Graphite and excess 36S in Lawrencite from Almahata Sitta MS-17 EL3 chondrite fragment. In Workshop on Formation of the First Solids in the Solar System, LPI Contributions, 1639, 9040. Lin, Y., Kimura, M., Hiyagon, H., and Monoi, A. (2003). Unusually abundant refractory inclusions from Sahara 97159 (EH3): A comparative study with other groups of chon- drites. Geochimica et Cosmochimica Acta, 67, 4935–4948. Lodders, K., and Fegley, B. (1993). Lanthanide and actinide chemistry at high C/O ratios in the solar nebula. Earth and Planetary Science Letters, 117, 125–145. Lusby, D., Scott, E. R. D., and Keil, K. (1987). Ubiquitous high-FeO silicates in enstatite chondrites. Journal of Geophysical Research, 92, 679. Manzari, P. (2010). Investigation of Enstatite Chondrites: Mineralogical and Chemical Features of EH3 and EL3 Chondrules. PhD thesis, Dipartimento Geomineralogico – Università degli Studi di Bari, Italy. Chondrules in Enstatite Chondrites 193

Marrocchi, Y., and Libourel, G. (2013). Sulfur and sulfides in chondrules. Geochimica et Cosmochimica Acta, 119, 117–136. McCoy, T. J., Dickinson, T. L., and Lofgren, G. E. (1999). Partial melting of the Indarch (EH4) Meteorite: A textural, chemical and phase relations view of melting and melt migration. Meteoritics and Planetary Science, 34, 735–746. McCoy, T. J., and Nittler, L. R. (2014). Mercury. In A. M. Davis (Ed.), Planets, Asteroids, Comets and the Solar System. In H. D. Holland and K. K. Turekian (Eds.), Treatise on Geochemistry (Second Edition), 1, 119–126. Amsterdam, Netherlands: Elsevier. Metzler, K., and Pack, A. (2016). Chemistry and oxygen isotopic composition of cluster chondrite clasts and their components in LL3 chondrites. Meteoritics and Planetary Science, 51, 276–302. Milani, A., Kneževic, Z., Novakovic, B., and Cellino, A. (2010). Dynamics of the Hungaria asteroids. Icarus, 207, 769–794. Miller, K. E., Lauretta, D. S., Connolly, H. C. Jr., et al. (2017). Formation of unequilibrated R chondrite chondrules and opaque phases. Geochimica et Cosmochimica Acta, 209, 24–50. Nakashima, D., Kimura, M., Yamada, K., et al. (2010). Study of chondrules in CH chondrites I: Oxygen isotope ratios of chondrules. Meteoritics and Planetary Science Supplement, 73, 5288. Pack, A., Shelley, J. M. G., and Palme, H. (2004). Chondrules with peculiar REE patterns: Implications for solar nebular condensation at high C/O. Science, 303, 997–1000. Pasek M. A., Milsom J. A., Ciesla F. J., et al. (2005). Sulfur chemistry with time-varying oxygen abundance during Solar System formation. Icarus, 175,1–14. Petaev, M. I., and Wood, J. A. (1998). The condensation with partial isolation model of condensation in the solar nebula. Meteoritics and Planetary Science, 33, 1123–1137. Petaev, M. I., and Wood, J. A. (2001). Condensation in Fractionated Nebular Systems. II. Formation of Enstatite Chondrites in Dust-enriched Nebular Reservoirs. Meteoritics and Planetary Science Supplement, 36, A162. Piani, L., Marrocchi, Y., Libourel, G., and Tissandier, L. (2016). Magmatic sulfides in the porphyritic chondrules of EH enstatite chondrites. Geochimica et Cosmochimica Acta, 195,84–99. Piani, L., Robert, F., Beyssac, O., et al. (2012). Structure, composition, and location of organic matter in the enstatite chondrite Sahara 97096 (EH3). Meteoritics and Planetary Science, 47,8–29. Pignatale, F. C., Liffman, K., Maddison, S. T., and Brooks, G. (2016). 2D condensation model for the inner Solar Nebula: an enstatite-rich environment. Monthly Notices of the Royal Astronomical Society, 457, 1359–1370. Rambaldi, E. R., Rajan, R. S., Housley, R. M., and Wang, D. (1986). Gallium-bearing sphalerite in a metal-sulfide nodule of the Qingzhen (EH3) chondrite. Meteoritics, 21,23–31. Rambaldi, E. R., Rajan, R. S., Wang, D., and Housley, R. M. (1983). Evidence for RELICT grains in chondrules of Qingzhen, an E3 type enstatite chondrite. Earth and Planetary Science Letters, 66,11–24. Ramdohr, P. (1963). Opaque minerals in stony meteorites. Journal of Geophysical Research, 68, 2011–2036. Rubin, A. E. (1983). The Adhi Kot breccia and implications for the origin of chondrules and silica-rich clasts in enstatite chondrites. Earth and Planetary Science Letters, 64, 201–212. Rubin, A. E. (2000). Petrologic, geochemical and experimental constraints on models of chondrule formation. Earth Science Reviews, 50,3–27. Rubin, A. E. (2010). Physical properties of chondrules in different chondrite groups: Implica- tions for multiple melting events in dusty environments. Geochimica et Cosmochimica Acta, 74, 4807–4828. 194 Emmanuel Jacquet, Laurette Piani, and Michael K. Weisberg

Rubin, A. E., and Grossman, J. N. (1987). Size-frequency distributions of EH3 chondrules. Meteoritics, 22, 237–251. Rubin, A. E., and Scott, W. R. D. (1997). Abee and related EH chondrite impact-melt breccias. Geochimica et Cosmochimica Acta, 61, 425–435. Rudraswami, N. G., Ushikubo, T., Nakashima, D., and Kita, N. T. (2011). Oxygen isotope systematics of chondrules in the Allende CV3 chondrite: High precision ion microprobe studies. Geochimica et Cosmochimica Acta, 75, 7596–7611. Schneider, D. M., Akridge, D. G., and Sears, D. W. G. (1998). Size Distribution of Metal Grains and Chondrules in Enstatite Chondrites. Meteoritics and Planetary Science Supplement, 33, 136. Schneider, D. M., Symes, S. J. K., Benoit, P. H., and Sears, D. W. G. (2002). Properties of chondrules in EL3 chondrites, comparison with EH3 chondrites, and the implications for the formation of enstatite chondrites. Meteoritics and Planetary Science, 37, 1401–1416. Scott, E. R. D., and Krot, A. N. (2014). Chondrites and their components. In A. M. Davis (Ed.), Meteorites and Cosmochemical Processes.InH.D.HollandandK.K.Turekian(Eds.), Treatise on Geochemistry (Second Edition), 1, 65–137. Amsterdam, Netherlands: Elsevier. Simon, S. B., Sutton, S. R., and Grossman, L. (2016). The valence and coordination of titanium in ordinary and enstatite chondrites. Geochimica et Cosmochimica Acta, 189, 377–390. Skinner, B., and Luce, F. (1971). Solid solution of the type (Ca, Mg, Mn, Fe)S and their use as geothermometers for the enstatite chondrites. American Mineralogist, 56, 1269–1296. Soulié, C., Libourel, G., and Tissandier, L. (2017). Olivine dissolution in molten silicates: An experimental study with application to chondrule formation. Meteoritics and Planetary Science, 52, 225–250. Tanaka, R., and Nakamura, E. (2017). Silicate-SiO reaction in a protoplanetary disk recorded by oxygen isotopes in chondrules. Nature Astronomy, 1, 0137. Tenner, T. J., Nakashima, D., Ushikubo, T., Kita, N. T., and Weisberg, M. K. (2015). Oxygen isotope ratios of FeO-poor chondrules in CR3 chondrites: Influence of dust enrichment and H2O during chondrule formation. Geochimica et Cosmochimica Acta, 148, 228–250. Tenner, T. J., Ushikubo, T., Kurahashi, E., Kita, N. T., and Nagahara, H. (2013). Oxygen isotope systematics of chondrule phenocrysts from the CO3.0 chondrite Yamato 81020: Evidence for two distinct oxygen isotope reservoirs. Geochimica et Cosmochimica Acta, 102, 226–245. Trieloff, M., Storck, J. -C., Mostefaoui, S., et al. (2013). A precise 53Mn-53Cr age of sphalerites from the primitive EH3 chondrite sahara 97158. Meteoritics and Planetary Science Supplement, 76, 5251. Uesugi, M., Akaki, T., Sekiya, M., and Nakamura, T. (2005). Motion of iron sulfide inclusions inside a shock-melted chondrule. Meteoritics and Planetary Science, 40, 1103. Ushikubo, T., Kimura, M., Kita, N. T., and Valley, J. W. (2012). Primordial oxygen isotope reservoirs of the solar nebula recorded in chondrules in Acfer 094 carbonaceous chondrite. Geochimica et Cosmochimica Acta, 90, 242–264. van Niekerk, D., and Keil, K. (2011). Metal/sulfide-silicate intergrowth textures in EL3 meteorites: Origin by impact melting on the EL parent body. Meteoritics and Planetary Science, 46, 1484–1497. Varela, M. E., Sylvester, P., Brandstätter, F., and Engler, A. (2015). Nonporphyritic chondrules and chondrule fragments in enstatite chondrites: Insights into their origin and secondary processing. Meteoritics and Planetary Science, 50, 1338–1361. Weisberg, M. K., Ebel, D. S., Bigolski, J. N., and Friedrich, J. M. (2016). Metal-sulfide nodules in enstatite and ordinary chondrites. 79th Annual Meeting of the Meteoritical Society, 1921, 6549. Weisberg, M. K., Ebel, D. S., Connolly, H. C., Kita, N. T., and Ushikubo, T. (2011). Petrology and oxygen isotope compositions of chondrules in E3 chondrites. Geochimica et Cosmo- chimica Acta, 75, 6556–6569. Chondrules in Enstatite Chondrites 195

Weisberg, M. K., Fogel, R. A., and Prinz, M. (1997). Kamacite-enstatite intergrowths in enstatite chondrites. In 28th Lunar and Planetary Science Conference, p. 523. Weisberg, M. K., and Prinz, M. (1998). Sahara 97096: A highly primitive EH3 chondrite with layered sulfide-metal-rich chondrules. In 29th Lunar and Planetary Science Conference, #1741. Weisberg, M. K., Prinz, M., and Fogel, R. A. (1994). The evolution of enstatite and chondrules in unequilibrated enstatite chondrites: Evidence from iron-rich pyroxene. Meteoritics, 29, 362–373. Weisberg, M., and Kimura, M. (2012). The unequilibrated enstatite chondrites. Chemie der Erde, 72, 101–115. Whitby, J. A., Gilmour, J. D., Turner, G., Prinz, M., and Ash, R. D. (2002). Iodine-Xenon dating of chondrules from the Qingzhen and Kota Kota enstatite chondrites. Geochimica et Cosmochimica Acta, 66, 347–359. Wood, J. A., and Hashimoto, A. (1993). Mineral equilibrium in fractionated nebular systems. Geochimica et Cosmochimica Acta, 57, 2377–2388. Yu, Y., Hewins, R. H., and Zanda, B. (1996). Sodium and sulfur in chondrules: Heating time and cooling curves. In R. H. Hewins, R. H. Jones, and E. R. D. Scott (Eds.), Chondrules and the Protoplanetary Disk, 213–219. Cambridge, UK: Cambridge University Press. Zhang, Y., Benoit, P. H., and Sears, D. W. G. (1995). The classification and complex thermal history of the enstatite chondrites. Journal of Geophysical Research, 100, 9417–9438. 8 Oxygen Isotope Characteristics of Chondrules from Recent Studies by Secondary Ion Mass Spectrometry

travis j. tenner, takayuki ushikubo, daisuke nakashima, devin l. schrader, michael k. weisberg, makoto kimura, and noriko t. kita

Abstract

We review recent chondrule oxygen isotope studies by secondary ion mass spectrometry (SIMS). We discuss primary O-isotope fractionation characteristics of chondrule phases, and how they are used to garner information related to the physicochemical environment from which they formed. This includes high temperature gas–melt interactions, sampling of common precursors among different chondrite types, and how precursor compositions influenced redox states during chondrule formation. We also explore how primary O-isotope ratios of chondrule phases are disturbed by secondary alteration.

8.1 Introduction

Oxygen isotope ratios of chondrules provide critical information related to the protoplanetary disk. This includes, but is not limited to, chondrule precursor characteristics (Chapter 2), dust- to-gas ratios, and high-temperature gas–melt interactions (Chapter 6). The spherulitic morphol- ogies and igneous textures of chondrules indicate formation by rapid heating and crystallization (e.g., Hewins and Radomsky, 1990), and they dominate the assemblage of unequilibrated chondrites (20–80 percent of their volume; Scott et al., 1996). As such, chondrules represent widespread thermal processing of early solar system materials, including CO-rich gas, silicate dust, and H2O ice. These precursors likely had different O-isotope ratios, which influenced (1) the value(s) within a chondrule and (2) the O-isotope variability observed among chondrule populations. Oxygen has three isotopes, 16O (99.757 atom %), 17O (0.038 atom %), and 18O (0.205 atom %) (Rosman and Taylor, 1998), which measurably fractionated during chondrule formation. However, primary O-isotope characteristics of chondrules may have been disturbed by secondary parent body processing, and determining if and how this occurred is challenging. To this extent, in situ analysis by secondary ion mass spectrometry (SIMS) is useful, because O-isotope ratios of chondrule phases are individually interrogated, with typical spot analysis diameters of 2–15 µm. Such investigations complement bulk chondrule O-isotope data (e.g., Clayton, 1993), as they allow for better understanding primary and secondary characteristics of chondrules and their related processes. Here, we highlight recent SIMS studies of chondrules and discuss implications regarding their formation, building upon previous summaries (e.g., Krot et al., 2006a; Yurimoto et al., 2008).

196 Oxygen Isotope Characteristics of Chondrules 197 8.2 Background 8.2.1 Delta Notation

SIMS O-isotope ratios are commonly reported in delta notation, or parts-per-thousand fractionation relative to Vienna Standard Mean Ocean Water (VSMOW; Baertschi, 1976): 17,18 17,18 16 18 δ O(‰) = [(Rsample/RVSMOW) – 1] 1,000, where R = O/ O. SIMS δ O values are calibrated with VSMOW-scale matrix-matched standards. SIMS δ17O values of unknowns are calculated by assuming bracketing standard δ17O values equal 0.52 δ18O (details: Appendix EA1; Tenner et al., 2013). Data are reported on an oxygen three-isotope plot, or δ17O versus δ18O. Data are also reported as Δ17O(=δ17O – 0.52 δ18O), approximating the difference in δ17O relative to the terrestrial fractionation line.

8.2.2 Mass-Dependent and Mass-Independent O-Isotope Fractionation

During mass-dependent fractionation, O-isotopes change with a δ17O versus δ18O slope of ~0.52, according to the mass differences of 17O and 18O, relative to 16O. The terrestrial fractionation line is largely defined by kinetic O-isotope fractionation of meteoric waters generated by preferential uptake of (1) light isotopes during evaporation, and (2) heavy isotopes during condensation. Regarding chondrules, their O-isotopes may also have fractionated mass- dependently by high-temperature evaporation/condensation, at tens-of-per-mil levels (e.g., Alexander, 2004). However, such signatures could have been partially or totally erased by gas–melt O-isotope exchange during chondrule formation (e.g., Yu et al., 1995; Alexander, 2004; Nagahara and Ozawa, 2012). High-temperature equilibrium fractionation can also occur, which is small among minerals (‰) but appreciable (‰-level) between silicates and ambient gas (Onuma et al., 1972; Kita et al., 2010; Javoy et al., 2012). Regarding carbonaceous chondrites, a remarkable feature is that O-isotopes of their com- ponents, including chondrules, Ca, Al-rich inclusions (CAIs), and unaltered matrix minerals are mass-independent fractionated (MIF). These materials plot along a slope-1 line with δ17O and δ18O values of –75‰ to +200‰, meaning they significantly differ in 16O relative abundances (Figure 8.1). Traditionally, two regression lines are used to represent primary MIF, the carbon- aceous chondrite anhydrous mineral (CCAM) line (Clayton et al., 1977) and the Young and Russell (1998) line, which are based on gas-source mass spectrometry measurements of CAI minerals from CV chondrites. Several origins for early solar system MIF are proposed, (e.g., Thiemens and Heidenreich, 1983; Dominguez, 2010), and the most currently favored hypoth- esis, called self-shielding, involves isotope selective photodissociation of CO gas by stellar ultraviolet radiation. The premise of CO self-shielding is that, because C16O is much more abundant than C17O and C18O, its dissociating photons could have been dampened within the core of the protosolar molecular cloud (Yurimoto and Kuramoto, 2004) and/or at the surface of the protoplanetary disk (Lyons and Young, 2005). Thus, C17O and C18O would have preferen- tially dissociated in equal amounts at greater column depths, producing a zone enriched in 17O, 18O, and C16O. Such characteristics are observed in nearby protosolar systems (e.g., Ando et al., 2002; Sheffer et al., 2002), where self-shielding is likely preserved through sequestration of 17 18 O and O into H2O ice in outer disk regions (Young, 2007). This is an important consider- ation, as ~1/4 of oxygen in the early solar system resided in H2O, half was associated with CO, 198 Travis J. Tenner, et al.

200 Acfer 094 COS 150

100 slope ~1 TF (slope 0.52)

O (‰) 50

17 matrix IOM δ

0 CC chondrules

reprocessed CAI minerals -50 fine-grained Sun refractory inclusions -100 -100 -50 0 50 100 150 200

δ18O (‰)

Figure 8.1 Oxygen three-isotope plot showing (1) the mass-dependent fractionated terrestrial fractionation line established from rocks and waters (e.g., Clayton et al., 1977; Clayton and Mayeda, 1983; Luz and Barkan, 2010); and (2) the slope ~1 mass-independent fractionated mixing line established by carbonaceous chondrite materials. Data: fine-grained refractory inclusions: Bodénan et al. (2014) and Ushikubo et al. (2017); reprocessed CAI minerals: Clayton et al. (1977); Clayton (1993); carbonaceous chondrite (CC) chondrules: Krot et al. (2006a); matrix insoluble organic matter (IOM): Hashizume et al. (2011) (average of areas #1 and #2 in their study); in situ SIMS analyses of Acfer 094 matrix cosmic symplectites (COS), inferred to represent O isotopes of primordial H2O: Sakamoto et al. (2007); the Sun, estimated from solar wind measurements (McKeegan et al., 2011). and ~1/4 was associated with silicate oxides, based on early solar system elemental abundances (Allende Prieto et al., 2001; 2002; Lodders, 2003); here we assume Fe, Ni, and S existed as metals/sulfides at the reducing conditions of solar gas (log fO2 < IW – 6; Grossman et al., 2008). Yurimoto and Kuramoto (2004) predict that, for typical protosolar clouds, the 16O-poor 17 18 16 H2O reservoir would have a δ O δ O of +100‰ to +250‰, and the O-rich CO gas reservoir would have a δ17O δ18Oof‒60‰ to ‒400‰, relative to the bulk value. Evidence for primordial heavy isotope enriched H2O ice comes from Fe-O-S-bearing cosmic symplectites within Acfer 094 (δ17O δ18O: +180‰; Sakamoto et al., 2007; Seto et al., 2008). Evidence for 16O-rich CO-bearing gas comes from a few CAIs and a chondrule (Kobayashi et al., 2003; Gounelle et al., 2009; Bodénan et al., 2014) that have more 16O-rich compositions than the estimated value of the Sun from solar wind measurements (δ17O δ18O: ~–59‰; McKeegan 16 et al., 2011). Timescales for the transport of O-poor H2O ice from the outer disk to the inner rocky planet-forming region are estimated at 105–106 years (Young, 2007). Chondrules, which postdate CAIs by 1–3 Myr based on Al–Mg isotope systematics (e.g. Kita and Ushikubo, 2012; Chapter 9; also see Chapter 11 for an alternative explanation of chondrule chronology), have relatively intermediate O-isotope ratios (δ17O δ18O: ‒15‰ to +15‰; Figure 8.1). As chondrules are predicted to have formed at high dust-to-gas ratios (e.g., Ebel and Grossman, Oxygen Isotope Characteristics of Chondrules 199

2000; Fedkin and Grossman 2006), it is reasonable to hypothesize that their O-isotope ratios reflect the early solar system silicate dust component (e.g., Kita et al., 2010; Krot et al., 2010a; Ushikubo et al., 2012; Kööp et al., 2016), and that they also represent mixing and/or evolution of the earliest 16O-poor and 16O-rich reservoirs (e.g., Connolly and Huss, 2010; Schrader et al., 2013; 2014a; Tenner et al., 2015). It is worth mentioning that if the O-isotope ratio of the Sun is the bulk solar system value, and if respective O-isotopes of cosmic symplectites and chondrules represent those of early solar system H2O ice and silicate reservoirs, a simple mass balance incorporating the proportions of each reservoir mentioned previously predicts the CO-rich gas reservoir would have a δ17O δ18O near –200‰. Such 16O-rich materials in meteoritic materials have not currently been discovered. However, a small number of refractory inclusions investigated by Gounelle et al. (2009) with δ17O δ18O: ~–67‰, as well as a single magnesian cryptocrystalline CH chondrite chondrule with δ17O δ18O: –75‰ (found by Kobayashi, Imai, and Yurimoto, 2003) indicate sampling of a 16O-enriched reservoir compared to the Sun, 16 perhaps depleted in O-poor H2O ice. In oxygen three-isotope plots throughout this chapter, we label the terrestrial fractionation, Young and Russell, and carbonaceous chondrite anhydrous mineral lines as TF, Y & R, and CCAM, respectively.

8.2.3 O-Isotope Exchange During Secondary Processing

O-isotope exchange may be recorded if chondrules experienced secondary processing during parent body aqueous/thermal metamorphism. Such exchange is driven by oxygen diffusion when chondrule phases are heated, and rates are phase-specific, based on experiments (e.g., Farver, 2010; Figure 8.2). We note experimental data are lacking below 600 C, meaning extrapolations are required to estimate effects at timescales of parent body metamorphism. In Figure 8.2 we show this as sustained temperatures required for oxygen to diffuse 10 µm2 per 1 Myr, where it can be seen that olivine, pyroxene, and spinel require sustained temperatures of 730 C–910 C to achieve 10 µm2 of O-diffusion per 1 Myr. As such, primary O-isotope ratios of these phases are largely unaffected by secondary processing, because estimated peak metamorphic temperatures are not this high for unequilibrated meteorites (e.g., Huss et al., 2006). Plagioclase and glass have faster O-diffusion rates, corresponding to 10 µm2 of O-diffusion per 1 Myr at 270 C–500 C in a dry system (Figure 8.2a), and at 150 C–230 C in a wet system (Figure 8.2b). Therefore, primary O-isotope ratios of chondrule plagioclase and glass are more susceptible to disturbance by parent body aqueous/thermal metamorphism, as discussed later.

8.3 SIMS Interphase O-Isotope Characteristics of Chondrules 8.3.1 Primary Interphase O-Isotope Characteristics from Acfer 094 Chondrules

Acfer 094 is an ungrouped carbonaceous chondrite and is classified as petrologic type 3.00 (Kimura et al., 2008). This means it experienced a very low extent of thermal metamorphism and aqueous alteration, and is arguably the most pristine chondrite from which O-isotope ratios 200 Travis J. Tenner, et al.

1600 800 600 400 200 (°C)

-16 (a) oxygen diffusion An Glass (dry) -18

Sp -20

HPx -22 2 Ol 10 µm Ab per 1 Myr

/s) -24 2

Ab (b) oxygen diffusion -16 Glass log D (m (wet) Ol

-18 An

-20 HPx

-22

-24

4 8 12 16 20 24 10000/T (K)

Figure 8.2 (a) Dry and (b) wet phase O-diffusion rates as a function of temperature. Dashed lines are extrapolations allowing for estimating temperatures required for 10 µm2 per 1 Myr. Data sources: Merigoux et al. (1968); Giletti et al. (1978); Gerard and Jaoul (1989); Ryerson et al. (1989); Matthews et al. (1994); Ryerson and McKeegan (1994); Pacaud et al. (1999); Behrens et al. (2004; 2007).

of chondrules have been studied. As such, we begin by highlighting Acfer 094 chondrule data because they serve as a logical benchmark for primary chondrule O-isotope signatures.

8.3.1.1 Primary O-Isotope Homogeneity, and Host and Relict O-Isotope Signatures Ushikubo et al. (2012) determined interphase O-isotope characteristics of Acfer 094 chondrules by SIMS. They investigated 42 chondrules, of which 29 are type I (Mg# > 90, where Mg# = mol.% MgO/[MgO + FeO] of constituent olivine and/or low-Ca pyroxene), 12 are type II (Mg#

< 90), and one is Al-rich (>10 wt% bulk Al2O3; Bischoff and Keil, 1984). They measured O-isotopes from coexisting olivine, low-Ca pyroxene, high-Ca pyroxene, plagioclase, and glass. Per chondrule, O-isotopes of coexisting pyroxene, plagioclase, and glass are homogeneous, with 2SD uncertainties ~< 1‰ in δ17O and δ18O (e.g., Figure 8.3a and 8.3b). Regarding chondrule olivine, most O-isotope data also equal those of coexisting pyroxene, plagioclase, and glass (e.g., Figure 8.3a and 8.3b), except for relict olivines described below. Per chondrule, observations that multiple mineral phases match those of coexisting glass indicate O-isotope ratios of Acfer 094 chondrules remained constant as they cooled from a molten state. Specific- ally, olivine was likely the first phase to crystallize at high temperatures, followed by low-Ca pyroxene, high-Ca pyroxene, plagioclase, and glass, as chondrules cooled (e.g., Figure 1 of 2 ch95 (a) 0 ch95 TF CCAM -2

-4 Y & R Ol LPx -6 HPx Plag -8 Glass -10 -8 -6 -4 -2 0 2 4

-2 ch56 (b) -4

-6 ch56 -8

-10

-12

-14 -10 -8 -6 -4 -2 0 O (‰) ch63 17 5 (c) ch63 0

-5 1 -10 2 Relict Ol

-15 -15 -10 -5 0 5

ch64 4 (d) 2 ch64

0

-2

-4

-6 -4 -2 0 2 4 6 18O (‰)

Figure 8.3 Examples of SIMS interphase O-isotope data from Acfer 094 chondrules (Ushikubo et al., 2012). Uncertainties are the 2SD of standard bracketing analyses. Data include 15 µm and 3 µm diameter spot analyses (not to scale in SEM images), which have δ17,18O uncertainties of ~0.4‰ and ~1‰ for large and small spots, respectively. Ol = olivine LPx = low-Ca pyroxene; HPx = high-Ca pyroxene; Plag = plagioclase. Ch95 (a) and ch56 (b) are type I porphyritic olivine-pyroxene (POP) chondrules. In ch63 (c), a type II porphyritic olivine (PO) chondrule, there are relict olivine grains with FeO-poor cores (labeled 1 and 2). In ch64 (d), a type I porphyritic pyroxene (PP) chondrule, there are relict olivine grains with dusty textures (see also supplement EA3 from Ushikubo et al., 2012 for images of dusty olivine in G64). Scalebars: 100 μm. 202 Travis J. Tenner, et al.

Ustunisik et al., 2014). Therefore, O-isotope ratios of such homogeneous phase data represent the “host” chondrule O-isotope ratio, defined as that of the final chondrule melt. For reference, the final chondrule melt is defined as any melt from which crystalline phases originated during the final chondrule-forming event. Ushikubo et al. (2012) calculate host chondrule O-isotope ratios from phase data within three standard deviation (3SD) analytical uncertainties, relative to the averaged Δ17O. O-isotope ratios of olivine grains within some Acfer 094 chondrules exhibit heterogeneity. In particular, within type II chondrules, olivine grains with FeO-poor cores are 16O-enriched relative to coexisting FeO-rich olivine and glass (Figure 8.3c), similar to observations from Kunihiro et al., (2005). Ushikubo et al. (2012) also found a type I chondrule containing olivine with a “dusty” texture of fine-grained Fe metal, and that such olivine is 16O-poor relative to coexisting phases (Figure 8.3d). Together, these are important results because of the relationship between O-isotope ratios of such olivines and their FeO/Fe metal textures. The term “relict” is commonly used to describe olivine grains which have FeO-poor cores and FeO-rich rims, as well as dusty olivines with fine-grained Fe metal (e.g., Nagahara, 1981; Rambaldi, 1981; Jones, 1992), because such textures could not have formed by a single crystallization event from a liquid. Instead, such grains likely formed in a prior heating and crystallization event, were incorporated into later chondrule precursors, and remained partially or fully intact during final chondrule formation. Olivine has a high propensity to survive as relicts because it has a higher melting temperature than other chondrule minerals (Ustunisik et al., 2014). Olivine also has a slow O-diffusion rate (e.g., Figure 8.2), meaning relict O-isotope signatures of olivines would have remained largely intact during the final chondrule-forming event. However, because olivine dissolution into silicate chondrule melt is rapid (i.e., microns per minute; Soulié et al., 2017), relict olivines probably only survived in chondrules that experienced low degrees of melting and/or short heating durations above solidus temperatures. Importantly, the link between non-host O-isotope ratios of chondrule olivines with FeO zoning and dusty textures confirms relict signatures can be determined from O-isotope ratios. This is relevant because Ushikubo et al. (2012) demonstrate many olivines with relict O-isotope signatures show no obvious major element signatures and/or textural features, similar to data from Jones et al. (2004). This suggests many relict olivines equilibrated FeO and MgO with the final chondrule melt, even though their O-isotopes did not. Supporting this suggestion are 1 atm. experimental results showing Fe–Mg diffusion rates in olivine are several orders of magnitude faster than those for oxygen, over a wide range of temperature and fO2 (e.g., Dohmen et al., 2007; Chakraborty, 2010). Ushikubo et al. (2012) defined relict olivines as those with Δ17O values outside 3SD of averaged homogeneous chondrule phase Δ17O. In addition to host and relict features, Ushikubo et al. (2012) found four Acfer 094 chondrules with scattered interphase O-isotope data, meaning none cluster together within 3SD of their averaged Δ17O (e.g., Figure 8.4). Thus, neither host O-isotope ratios nor relict olivine signatures can be identified. Ushikubo et al. (2012) simply call these chondrules “heterogeneous,” in terms of their O-isotopes. Of the 42 chondrules studied by Ushikubo et al. (2012), 38 have multiple O-isotope data per chondrule that are homogeneous (i.e. Δ17O 2SD 1‰), allowing for calculating host chondrule values. Four of 42 chondrules are heterogeneous in their mineral/glass O-isotope ratios, and 15 of 42 contain relict olivine. Oxygen Isotope Characteristics of Chondrules 203

ch68 -4 TF Y & R -6

-8 ch68 CCAM O (‰) -10 17 δ -12 olivine × 4 δ17O –28.6 to –45.6‰; -14 δ18O –26.5 to –43.4‰ -16 -12 -8 -4 0 δ18O (‰)

Figure 8.4 Acfer 094 chondrule G68 (type I POP) (Ushikubo et al., 2012), an example of a chondrule with heterogeneous interphase O-isotope data, meaning they do not overlap within analytical uncertainties. Symbols, lines, and uncertainties are the same as in Figure 8.3. Four olivine spot data are significantly 16O- rich (red arrow); their range of δ17O and δ18O values are given. Scalebar: 100 μm.

8.3.1.2 The Primitive Chondrule Mineral (PCM) Line Acfer 094 chondrule mineral O-isotope data plot between the CCAM and Y & R lines (e.g., Figures. 8.3 and 8.4). Based on the petrologic type 3.00 nature of Acfer 094, Ushikubo et al. (2012) defined a new mixing line called the primitive chondrule mineral (PCM) line (Figure 8.5a–d). Such data are well-constrained along a slope ~1 line with δ18O between 44‰ and +4‰,defining the regression δ17O = (0.987 0.013) δ18O – (2.70 0.11). In oxygen three- isotope plots throughout this chapter, we label the primitive chondrule line as PCM. The PCM line is significant for several reasons. It intersects the terrestrial fractionation line at δ18O: 5.8 0.4‰, which coincides with the terrestrial mantle (δ18O: 5.5‰; Eiler, 2001) and lunar rocks (Weichert et al., 2001; Spicuzza et al., 2007), as well as multiple achondrites (Clayton, 2003; Franchi, 2008; Greenwood et al., 2017). O-isotopes of 16O-rich refractory inclusions from Acfer 094 are also consistent with the PCM line (e.g., Figure 8.5a), as are those of extremely 16O-poor Acfer 094 cosmic symplectites (e.g., Figure 8.5d). Together, these observations are most consistent with a slope-1 fractionation line produced by CO photodisso- ciation (e.g., Shi et al., 2011; Gao et al., 2013), and are not as consistent with galactic chemical evolution, where slope-1 O-isotope fractionation lines are not uniquely produced (Lugaro et al., 2012). Based on analyses of unaltered Allende CAI minerals, Young and Russell (1998) suggest the majority of early solar system materials, which are enriched in δ18O compared to their mixing line, underwent varying degrees of mass-dependent fractionation/exchange, relative to a pri- mary O-isotope reservoir. For example, Allende whole rock data (e.g., Clayton et al., 1977) may be systematically shifted right in δ18O because of mass-dependent heavy isotope enrichment of chondritic materials during parent body aqueous alteration (e.g., Bridges et al., 1999; Young et al., 2002; Bullock et al., 2012; Doyle et al., 2015; Krot and Nagashima, 2016). Additionally, bulk and matrix separate CR chondrite O-isotope ratios show that those with low-degree 204 Travis J. Tenner, et al.

-42 Acfer 094 Acfer 094 fine-grained -25 Chd. Ol refractory Chd. Px inclusions -44 -30

-46 -35

-48 (a) (b) -40 -48 -46 -44 -42 -40 -40 -35 -30 -25 O (‰)

17 5 200 δ Acfer 094 COS 0 190

-5 TF 180 -10 Y & R

170 -15 CCAM (c) (d) PCM -20 160 -15 -10 -5 0 5 160 170 180 190 200 δ18O (‰)

Figure 8.5 Primitive chondrule mineral (PCM) O-isotope mixing line (solid line in all panels), defined by chondrule phenocryst data from Ushikubo et al. (2012). Uncertainties are the 2SD of standard bracketing measurements. From their study, only high-precision, 15 µm-sized SIMS spot data from olivine and pyroxene phenocrysts were used to establish the PCM line. Panels (a)–(d) traverse the 16O-rich to 16O- poor regions of the PCM line. Panel (a) includes Acfer 094 fine-grained refractory inclusion data from Ushikubo et al. (2017). Panel (d) includes in situ SIMS analyses of Fe-O-S-bearing cosmic symplectites from Sakamoto et al. (2007), which are hypothesized to be the reaction products of FeS bearing materials 16 with primordial O-poor H2O. Shown for comparison are TF, Young and Russell, and CCAM (Clayton et al., 1977) mixing lines (dashed and dotted lines, respectively, in all panels). aqueous alteration plot near the Y & R line, while progressively altered CR chondrites are increasingly enriched in δ18O, toward 16O-poor water (Schrader et al., 2011; 2014b). In contrast, many chondrule minerals (especially pyroxene, olivine, and spinel) from CV, CR and several other chondrites show no obvious evidence of mass-dependent fractionation/exchange, and their SIMS data plot on the PCM line (as discussed later).

8.3.2 SIMS Interphase O-Isotope Characteristics of Other Chondrite Chondrules

8.3.2.1 O-Isotope Homogeneity of Chondrule Pyroxene Per chondrule, the vast majority of pyroxene phenocrysts have homogeneous O-isotopes (e.g., Figures 8.3a, 8.3b, and 8.3d), and this characteristic includes many chondrites with such data. Including the following studies: Chaussidon et al. (2008) (CV3 and CR2), Kita et al. (2010) Oxygen Isotope Characteristics of Chondrules 205

(LL3), Krot et al. (2010b) (CBb, CH/CBb), Rudraswami et al. (2011) (Allende CV3), Weisberg et al. (2011) (enstatite or E), Ushikubo et al. (2012) (Acfer 094 ungr, C3), Nagashima, Krot, and Huss (2015) (Kakangari or K), Miller et al. (2017) (Rumuruti or R), Weisberg et al. (2015) (Grosvenor 95551-Northwest Africa 5492, or G), Tenner et al. (2013; 2015; 2017) (CO3, CR2, Yamato (Y) 82094 ungr, C3.2, respectively), and Schrader et al. (2014a; 2017) (CR2), there are 728 SIMS spot analyses representing multiple pyroxene data (n 2) on a per-chondrule basis, coming from 186 chondrules (Table 8.1). Of the spot data, 665, or 91 percent, represent homogeneous pyroxene data under the criterion that, when averaged, their per-chondrule pyroxene Δ17O 2SD is 1‰. Within the bounds of such spot data, this corresponds to 90 percent (167 of 186) of pyroxene-bearing chondrules having homogeneous pyroxene O- isotope ratios. This indicates pyroxene accurately records host chondrule O-isotope ratios. Supporting this hypothesis is that O-diffusion in pyroxene is sufficiently slow (e.g., Figure 8.2), meaning their primary O-isotope signatures were unlikely disturbed during parent body thermal metamorphism. Pyroxene also has a lower melting temperature than olivine, so it is less likely to have survived multiple heating events as relicts.

8.3.2.2 Host and Relict Signatures of Chondrule Olivine Recent studies from multiple chondrites, including LL3 (Kita et al., 2010), enstatite (Weisberg et al., 2011), K (Nagashima et al., 2015), R (Miller et al., 2017), G (Weisberg et al., 2015), CV3 (Rudraswami et al., 2011), CO3 (Tenner et al., 2013), CR2 (Tenner et al., 2015; Schrader et al., 2017), and Y-82094 (ungr. C3.2) (Tenner et al., 2017) find that the vast majority of chondrules have host olivines with δ18O and δ17O matching those of coexisting pyroxene, similar to Acfer 094 chondrule characteristics (Figure 8.6). Their data are fit by the following regressions: 18 18 2 17 17 2 δ Opyroxene = 0.999 δ Oolivine (R : 0.964); and δ Opyroxene = 1.004 δ Oolivine (R : 0.986). These results indicate that, excluding relict grains, olivines and pyroxenes were co- magmatic, inheriting the same O-isotope ratio as the final chondrule melt. Although O-isotopes of coexisting chondrule olivine and pyroxene generally agree (exclud- ing relicts), arguments against their co-magmatic origin exist. This is driven by observations that many type I chondrules have olivine-rich interiors and low-Ca pyroxene-rich peripheries (e.g., Grossman et al., 2002; Friend et al., 2016). This arrangement has been experimentally produced by condensation of SiO-rich gas onto chondrule analogues as they cooled (e.g., Tissandier et al., 2002). Further, Chaussidon et al. (2008) found that, among CV3 and CR2 chondrites, olivines are systematically enriched in 16O relative to coexisting pyroxenes per chondrule (Figure 8.6, dashed lines), although subsequent studies show comparative O-isotope homogeneity among CV3, CR2 and other chondrite chondrules (Figure 8.6). Nonetheless, Chaussidon et al. (2008) hypothesize that most olivines within type I chondrules are relicts, and that pyroxene O-isotope signatures were established during reactions involving olivine dissolution by addition of SiO2 to chondrule melt from ambient gas. Finally, Marrocchi and Libourel (2013) hypothesize that an association between low-Ca pyroxene and sulfides among CV3 chondrite chondrules indi- cates crystallization after olivine and after SiO from ambient gas condensed onto chondrule surfaces. While possible, this finding is not a unique solution because sulfides were immis- cible liquids in the chondrule melt during olivine and pyroxene crystallization (as they have relatively low crystallization temperatures of <990 C; Schrader et al., 2015), meaning they Table 8.1. Compilation of SIMS pyroxene O-isotope spot data from recent studies 206 GRO 95551- CBb, Ungrouped Enstatite Kakangari Rumuruti NWA 5492 Chondrite type CV3 CO3 CR2 CH/CBb C LL3 (E) (K) (R) (G)

Source [1] [2] [3] [4] [5] [6] [7] [8](a) [9] [10] [11] [12] [13] [14] [15] Totals Total px data 20 36 67 11 194 7 12 6 105 107 57 24 28 8 46 728 when n 2 data per chondrule # of chondrules 8 11 19 3 41 3 5 3 27 26 3 8 6 3 20 186 represented Homogeneous 18 36 67 2 187 7 12 6 80 107 55 16 28 5 39 665 px spot data (per chondrule Δ17O 2SD 1‰) # of chondrules 7 11 19 1 39 3 5 3 20 26 2 6 6 2 17 167 represented Heterogeneous 20097000 25028 0 3 3 59 px spot data (per chondrule Δ17O 2SD 1‰) # of chondrules 10022000 7012 0 1 3 19 represented

(a) only porphyritic chondrule data were considered; cryptocrystalline chondrule data were not considered because there were only single measurements per chondrule. Sources: [1] Chaussidon et al. (2008); [2] Rudraswami et al. (2011); [3] Tenner et al. (2013); [4] Chaussidon et al. (2008); [5] Tenner et al. (2015); [6] Schrader et al. (2014a); [7] Schrader et al. (2017); [8] Krot et al. (2010b); [9] Ushikubo et al. (2012); [10] Tenner et al. (2017); [11] Kita et al. (2010); [12] Weisberg et al. (2011); [13] Nagashima et al. (2015); [14] Miller et al. (2017); [15] Weisberg et al. (2015) Oxygen Isotope Characteristics of Chondrules 207

15 15

10 LPx 10 Al-rich LPx 5 HPx 5 δ 0 0 17 O Px O Px

18 -5 -5 δ -10 -10

-15 -15

-20 -20 -20 -10 0 10 -20 -10 0 10 δ18O Ol δ17O Ol

Figure 8.6 Comparisons of coexisting chondrule olivine and pyroxene O-isotope data from recent studies of ordinary, enstatite, Rumuruti (R), Grosvenor 95551-Northwest Africa 5492-like (G), Kakangari (K), and carbonaceous chondrites (sources: Kita et al., 2010; Krot et al., 2010b; Rudraswami et al., 2011; Weisberg et al., 2011; 2015; Ushikubo et al., 2012; Tenner et al., 2013; 2015; 2017; Nagashima et al., 2015; Miller et al., 2017; Schrader et al., 2017). LPx = low-Ca pyroxene; HPx = high-Ca pyroxene. Data exclude relict grains identified in the studies. Each datum is a comparison of averaged olivine versus pyroxene data within a single chondrule, where uncertainties are 2SD. The majority of coexisting chondrule olivine and pyroxene exhibit a 1 to 1 relationship for δ18O and δ17O (solid trend lines), Also shown are regression lines from Chaussidon et al. (2008) (dashed lines), which represent fits to their coexisting olivine and pyroxene data. formed only after olivine and pyroxene. For example, CR chondrite chondrules show different associations in which pyroxene-rich peripheries among type I chondrules are sulfide-absent, while sulfides are ubiquitous in olivine-dominated type II chondrules (e.g., Schrader et al., 2015). In order to explain O-isotope homogeneity among single chondrules with olivine-rich cores and low-Ca pyroxene-rich peripheries, it seems likely that chondrule melts behaved as open systems during chondrule formation, but that the surrounding gas and chondrule melts had similar O-isotope ratios. As such, the ambient gas surrounding a given chondrule may simply have consisted of evaporated material from the chondrule melt (including SiO), which then returned to the system by condensation (e.g., Nagahara et al., 2008). This is a reasonable hypothesis when considering the high dust-to-gas ratios invoked during chondrule formation (e.g., Ebel and Grossman, 2000). Although kinetic mass-dependent O-isotope fractionation would have occurred during this process (as suggested by olivines that are ~1‰ higher in δ18O than coexisting phases among CO and CR chondrite chondrules; Tenner et al., 2013; 2015), thermodynamic modeling (Alexander, 2004) shows that tens-of-per-mil level fractiona- tions are erased by gas–melt isotope exchange at cooling rates consistent with experiments that produce type I chondrule assemblages (e.g., 5–20 C/hr.; Wick and Jones, 2012). In this scenario, it is possible, although not required, that precursors contained relict olivines with O- isotopes that did not exchange during chondrule formation. Another way that olivine and pyroxene O-isotope homogeneity has been explained is that (1) precursors consisted predomin- antly of relict olivines; and that (2) such grains were fully or partially resorbed as chondrule melts gained silica from ambient gas condensation, because olivine would fall out of equilib- rium with the system (e.g., Marrocchi and Chaussidon, 2015; Soulié et al., 2017). In this 208 Travis J. Tenner, et al. scenario, host olivines and pyroxenes would have co-crystallized from the silica-enriched, yet pyroxene-undersaturated melt, recording uniform O-isotope ratios at dust to gas proportions >10 (e.g., Figure 8 from Marrocchi and Chaussidon, 2015; note that Ebel and Grossman, 2,000 ‒ predict dust to gas ratios >12.5 are necessary to form chondrules at a Ptot of 10 3 bar); incompletely resorbed precursor olivines would retain their relict O-isotope signatures. In addition to chondrules with coexisting olivine and pyroxene O-isotope data, others have only olivine O-isotope data available. Therefore, determining host versus relict olivine signatures from such chondrules requires different evaluation criteria. Textural evidence is helpful in this regard, such as forsterite zoning and dusty olivines, which provide convincing evidence for a relict origin (e.g., Kunihiro et al., 2004; 2005; Ruzicka et al., 2007). Additionally, imaging by cathodoluminescence (CL) can reveal the presence of refractory relict olivine (e.g., Pack et al., 2004; Kita et al., 2010). Regarding barred olivine chondrules studied by SIMS, (e.g., Connolly and Huss, 2010; Rudraswami et al., 2011; Ushikubo et al., 2012; Schrader et al., 2013; 2014a; Tenner et al., 2013; 2015), olivine data are homogeneous per chondrule (Δ17O2SD 1‰), indicating they represent host O-isotope ratios. This finding is somewhat expected, as BO textures indicate complete melting and rapid olivine crystallization (e.g., Hewins et al., 2005). Recent studies provide large chondrule datasets where host and relict O-isotope signatures are defined in a similar manner as the Acfer 094 data from Ushikubo et al. (2012). This allows for evaluating proportions of host versus relict olivine grains within various chon- drites (Table 8.2). For example, some chondrites have few relict olivine-bearing chondrules, such as CV3 (1 of 24; Rudraswami et al., 2011), CR2 (11 of 77; Connolly and Huss, 2010;

Schrader et al., 2013; 2014a; 2017; Tenner et al., 2015), CBb,CH/CBb (1 of 11; Krot et al., 2010b), LL3 (3 of 31; Kita et al., 2010), K chondrites (1 of 7; Nagashima et al., 2015), and G chondrites (0 of 17; Weisberg et al., 2015). In contrast, other chondrites have chondrules with larger proportions of relict olivine-bearing chondrules, such as CO3 (13 of 26; Tenner et al., 2013), Acfer 094 (ungr. C3) (15 of 34; Ushikubo et al., 2012), Y-82094 (ungr. C3) (16 of 30; Tenner et al., 2017), enstatite chondrites (4 of 12; Weisberg et al., 2011), and R chondrites (2 of 6; Miller et al., 2017). Previous studies of CO chondrites (e.g., Jones et al., 2000; Kunihiro et al., 2004) and Acfer 094 (e.g., Kunihiro et al., 2005) show similar proportions of relict olivine-bearing chondrules. If considering all olivine data from Table 8.2, the majority represent host chondrule O-isotope ratios (e.g., 62.5–100 percent, depending on chondrite type). Note there may be some uncertainty in these statistics, as chondrule datasets may not accurately represent proportions of all chondrule types within a given chondrite. O-isotope values of relict olivine can help to reveal their origins. If relict olivines have O-isotope ratios within the range of host values of a given chondrite, they likely originated locally within the accretion region (e.g., Figure 8.7a). In contrast, several SIMS studies report relict olivine grains that are significantly 16O-rich relative to their coexisting chondrule phases (Figure 8.7b). Possible origins include amoeboid olivine aggregates (AOAs), as well as pristine and reprocessed CAIs, and the spread in relict olivine values could represent mixtures of refractory and host materials analyzed by SIMS. Among such chondrules, the O-isotope variability among their olivine and pyroxene phenocrysts tracks chondrule melt-gas inter- actions. Relict olivines that are significantly 16O-poor compared to ranges of host chondrule values within chondrites are rarely observed. Table 8.2 Compilation of host and relict chondrule olivine SIMS O-isotope data from recent studies. Excludes heterogeneous chondrules reported in studies.

GRO CBb, Enstatite Kakangari Rumuruti 95551-NWA Chondrite Type CV3 CO3 CR2 CH/CBb Ungrouped C LL3 (E) (K) (R) 5492 (G) Source [1](c) [2] [3–7] [8](d) [9](e,f) [10] [11](g) [12] [13] [14](h) [15]

Total chondrules with 24 26 77 11 34 30 31 12 7 6 17 olivine data(a)(b) Total relict olivine- 1 13 11 1 15 16 3 4 1 2 0 bearing chondrules(a)(b) %Relict olivine-bearing 4 50 14 9 44 53 10 33 14 33 0 chondrules Total host olivine spot 65 68 278 23 111 99 92 20 33 18 39 data Total relict olivine spot 1 32 14 1 35 43 6 12 1 5 0 data % Host olivine spot data 98.5 68.0 95.2 95.8 76.0 69.7 93.9 62.5 97.1 78.3 100

(a) Only chondrules with 2 or more spot measurements of olivine and/or pyroxene were considered. (b) Designations do not include other phases, such as spinel, plagioclase, or glass that were also measured. (c) Two heterogeneous chondrules from this study have four 16O-rich olivine grains (Δ17O: 11 to 17‰) interpreted as relicts. (d) Only porphyritic chondrule data were considered; cryptocrystalline chondrules have only single measurments. (e) Only spot data from Electronic Appendices used to determine host and relict O-isotope ratios were considered. (f ) Two heterogeneous chondrules from this study have seven 16O-rich olivine grains (Δ17O: 14.5 to 23‰) interpreted as relicts. (g) Only spot data from Electronic Appendices used to determine host and relict O-isotope ratios were considered. (h) One heterogeneous chondrule from this study has three olivine grains (Δ17O: 1.5 to 3.6‰) interpreted as relicts. Sources: [1] Rudraswami et al. (2011); [2] Tenner et al. (2013); [3] Connolly and Huss (2010); [4–6] Schrader et al. (2013; 2014b; 2017); [7] Tenner

209 et al. (2015);[8] Krot et al. (2010b); [9] Ushikubo et al. (2012); [10] Tenner et al. (2017); [11] Kita et al. (2010); [12] Weisberg et al. (2011); [13] Nagashima et al. (2015); [14] Miller et al. (2017); [15] Weisberg et al. (2015). 210 Travis J. Tenner, et al.

Y-81020. Y12 Relict Ol 0 -20 LPx TF fine-grained Host Ol refractory Relict Ol inclusions -4 -30 Y & R O (‰) 17 δ -8 Y25 -40 PCM LPx reprocessed Host Ol a CAI minerals b -12 Relict Ol -50 CCAM -12 -8 -4 0 4 -50 -40 -30 -20 δ18O (‰)

Figure 8.7 Examples of relict olivine (Ol) SIMS O-isotope characteristics. LPx = low-Ca pyroxene. Reported 2SD uncertainties are shown. (a) Relict olivines that plot between host data from two Yamato (Y) 81020 (CO3.05) chondrules, suggesting transport and mixing of multiply-processed olivine grains between chondrule-forming O-isotope reservoirs. (b) Relict chondrule olivine grains from various studies that are significantly 16O-rich. Fine grained refractory inclusion data (CAI and AOA): Bodénan et al. (2014) and Ushikubo et al. (2017). Reprocessed CAI mineral data: Clayton et al. (1977) and Clayton (1993). Relict olivine data: Rudraswami et al. (2011); Ushikubo et al. (2012); Tenner et al. (2013); Zhang et al. (2014); Nagashima et al. (2015); Schrader et al. (2017); Tenner et al. (2017).

8.3.2.3 Host and Relict Spinel Signatures Spinel is uncommon in chondrules, and is instead typically found within CAIs. Spinel is observed in carbonaceous and ordinary chondrites, within porphyritic and barred olivine chondrules (e.g., Maruyama et al., 1999; Kimura et al., 2006; Ma et al., 2008; Rudraswami et al., 2011) and in Al-rich chondrules (e.g., Maruyama et al., 1999; Guan et al., 2006; Krot et al., 2006b; Zhang et al., 2014; Tenner et al., 2017). In Allende (CV3.6), Rudraswami et al. (2011) found spinels with O-isotope ratios equaling coexisting olivine and low-Ca pyroxene in three chondrules (e.g., Figure 8.8a). They interpret these spinels crystallized from the final chondrule melt, therefore representing host O-isotope ratios. Kimura et al. (2006) and Jiang et al. (2015) also report chondrule spinels in ordinary chondrite chondrules with similar Δ17Oas coexisting olivine and pyroxene, and with igneous textures consistent with crystallization from the final chondrule melt. If considering (1) the high melting temperature of spinel compared to other chondrule minerals (Figure 1 of Ustunisik et al., 2014); and (2) the slow O-diffusion rate in spinel (Figure 8.2), it could be expected at least some chondrule spinels are unmelted relicts of CAI- like precursors (Krot et al., 2006c; Makide et al., 2009; Krot et al., 2017). Indeed, O-isotope ratios of some chondrule spinels are significantly 16O-rich relative to their coexisting phases, approaching values of pristine and reprocessed CAI minerals (e.g., Figure 8.8b). Relict spinels that are significantly 16O-poor relative to those of coexisting host chondrule phases are not observed. Oxygen Isotope Characteristics of Chondrules 211

-2 Allende -10 Relict Sp TF All-1-Ch-9 fine-grained -4 LPx -20 refractory Ol inclusions Host -30 -6 Sp O (‰) 17

δ Y & R CCAM -40 -8 reprocessed CAI minerals a PCM b -50 -10 -8 -6 -4 -2 0 -50 -40 -30 -20 -10 δ18O (‰)

Figure 8.8 Examples of host and relict spinel (Sp) SIMS O-isotope data. Uncertainties are the reported 2SD values. Ol = olivine; LPx = low-Ca pyroxene. (a) Host spinel datum from Allende (CV3.6) chondrule All-1-Ch-9 (Rudraswami et al., 2011) that matches O-isotopes of coexisting olivine and low-Ca pyroxene. (b) Relict spinel data. Sources: Maruyama et al. (1999); Krot et al. (2006b); Zhang et al. (2014); Tenner et al. (2017). Fine grained refractory inclusion data: Bodénan et al. (2014) and Ushikubo et al. (2017). Re-processed CAI mineral data: Clayton et al. (1977) and Clayton (1993).

8.3.2.4 Host, Relict, and Secondary Alteration O-Isotope Signatures of Chondrule Plagioclase O-isotope data from chondrule plagioclase are relatively sparse, as it is often associated with mesostasis (exceptions are Al-rich chondrules, where plagioclase is prevalent). For primitive chondrites, such as Acfer 094, CR2 and Y-82094 (ungr. C3.2), chondrule plagioclase have δ17O and δ18O equaling those of coexisting pyroxene at the per-mil level (e.g., Tenner et al., 2015; 2017; Schrader et al., 2017) (Figure 8.9a). A small number of plagioclase from CR2 Al-rich chondrules are significantly 16O-rich relative to coexisting pyroxene (Krot et al., 2006b; Schrader et al., 2017), and are interpreted as relicts, based on characteristics indicating refrac- tory materials within precursor components. Overall, the O-isotope agreement of coexisting plagioclase and pyroxene among Acfer 094, CR2, and Y-82094 chondrules indicates plagio- clase O-isotope ratios (1) were undisturbed by parent body thermal metamorphism; and (2) represent host chondrule values. Furthermore, such plagioclase show no evidence of replace- ment by nepheline, a reaction which likely occurred during secondary processing (e.g., Kimura and Ikeda, 1997; Ichimura et al., 2017). Unlike Acfer 094, CR2 and Y-82094 chondrites, plagioclase O-isotope data from other chondrites indicate disturbance by parent body thermal metamorphism. Rudraswami et al. (2011) report two chondrules from Allende (CV3.6) that have significantly 16O-poor plagioclase relative to coexisting pyroxenes, plotting right of the CCAM line (Figure 8.9b). Importantly, Allende is estimated to have experienced metamorphic temperatures of 340–600 C (Huss and Lewis, 1994; Brearley, 1997). Therefore, O-isotopes of 10 µm-sized Allende chondrule plagio- clase grains could have equilibrated with fluid on the parent body in 0.1–10 million years, based on plagioclase O-diffusion rates (Figure 8.2). Indeed, the Δ17O of Allende chondrule 212 Travis J. Tenner, et al.

2 2

a b 0 0

-2 TF -2

-4 -4 Y & R O (‰) Allende CV3.6

17 QUE 99177 3658-Ch17 δ -6 -6 PCM (CR2.8) ch13 Ol Ol LPx -8 LPx HPx -8 CCAM Host Plag Altered Plag -10 -10 -6 -4 -2 0 2 4 6 -6 -4 -2 0 2 4 6 δ18O (‰)

Figure 8.9 Examples of chondrule plagioclase (Plag) with host (a) and disturbed (b) O-isotope characteristics. Ol = olivine; LPx = low-Ca pyroxene; HPx = high-Ca pyroxene. Uncertainties are the reported 2SD. Each plot shows individual SIMS O-isotope measurements collected per chondrule. Data in (a) are from a Queen Alexandra Range (QUE) 99177 chondrule (ch13; Schrader et al., 2017); QUE 99177 is characterized as a pristine, petrologic type 2.8 CR chondrite (Harju et al., 2014; Howard et al., 2015). Data in (b) are from Allende (CV3.6) chondrule 3658-ch17 from Rudraswami et al. (2011). Allende experienced peak metamorphic parent body temperatures of 340–600 C (Huss and Lewis, 1994; Brearley, 1997), which likely disturbed primary O-isotope ratios of plagioclase, but would not have disturbed primary O-isotope ratios of olivine and pyroxene (e.g., Figure 8.2). plagioclase (2.6‰) is consistent with CV3 (including Allende) magnetite and fayalite (1‰ to 3‰; Choi et al., 1997; 2000; Doyle et al., 2015), where magnetite and fayalite are interpreted as reaction products that involved water on the parent body (e.g., Krot et al., 1997; 1998; Davidson et al., 2014). Furthermore, the difference in δ18O between Allende chondrule plagio- clase (δ18O: +5%; Rudraswami et al., 2011) and Allende magnetite (δ18O mean averages: 18 –2.6‰ to –7.1‰; Choi et al., 1997) (i.e., Δ Oplagioclase-magnetite: 7.6‰–12.1‰) is consistent with equilibrium anorthite-magnetite fractionation calibrations (Clayton and Kieffer, 1991) 18 from 340 to 600 C(Δ Oanorthite-magnetite: 5.3‰–10.0‰), where magnetite has the lower δ18O value. Additionally, Chaussidon et al. (2008) report chondrule plagioclase data from CV3 chondrites that are 16O-poor relative to their coexisting phases, and suggest the possibility their O-isotopes were disturbed by parent body thermal metamorphism. When comparing Grove Mountains (GRV) 022410 (H4), GRV 052722 (H3.7) and Julesberg (L3.6) chondrites, Jiang et al. (2015) found plagioclase within Al-rich chondrules has the same O-isotope ratio as nepheline. Further, both phases are higher in δ18O along a mass-dependent fractionation line relative to chondrule pyroxene, olivine, and spinel. They suggest plagioclase and nepheline exchanged O-isotopes with parent body fluid during thermal metamorphism, with nepheline representing complete reaction of plagioclase and Na-bearing aqueous fluid. Finally, Zhang et al. (2014) report 16O-poor anorthite (Δ17O: –3‰) relative to enstatite, olivine, and sapphirine (Δ17O: –7‰) within a sapphirine-bearing Al-rich chondrule (SARC) from Dar al Gani 978 (ungr C). They attribute the anorthite signature to exchange with 16O-poor fluid during thermal metamorphism (850–950 K; Zhang and Yurimoto, 2013); further evidence for thermal Oxygen Isotope Characteristics of Chondrules 213 metamorphism in the SARC includes high-temperature FeO-MgO exchange among chondrule enstatite, olivine, and spinel (e.g., Kimura and Ikeda, 1995).

8.3.2.5 Evidence of Secondary Alteration from Chondrule Glass O-Isotopes There are relatively few O-isotope data from chondrule glass. Although Acfer 094 chondrules have host glass O-isotopes equaling those of coexisting host olivine and pyroxene (Figure 8.3b– 8.3d), attesting to its 3.00 petrologic type, O-isotopes of LL3 chondrule glass indicate it is susceptible to O-isotope disturbance during low-degree aqueous/thermal metamorphism (e.g., Figure 8.2). Kita et al. (2010) report that, per chondrule, glasses from Krymka, Bishunpur and Semarkona (petrologic types 3.2, 3.15 and 3.01, respectively; Grossman and Brearley, 2005; Kimura et al., 2008) are generally enriched in δ17O and δ18O, relative to coexisting olivine and pyroxene data (Figure 8.10). These data plot along a slope ~0.8 line on an oxygen three-isotope plot (dashed line, Figure 8.10), suggesting O-isotope exchange with a 16O-poor fluid. In particular, Semarkona chondrule 36 glass from Kita et al. (2010) plots on a fractionation line defined by Semarkona magnetite (Choi et al., 1998; Doyle et al., 2015), which is considered the product of parent body aqueous alteration (e.g., Krot et al., 1997). As such, Kita et al. (2010) interpret that LL3 chondrule glass exchanged O-isotopes with Δ17O~5‰ aqueous fluid on the parent body, and that greater O-isotope deviations from chondrule olivine and pyroxene data

20 CH36 Glass Glass

15 Semarkona Magnetite

10 O (‰)

17 TF δ

5

LL3 Chondrule 0 Ol & Px data

-5 0 5 10 15 20 δ18O (‰)

Figure 8.10 SIMS O-isotope data from LL3 chondrite chondrule glass (Kita et al., 2010; chondrules B4, B26, K10, CH11, CH33, CH52, CH4, CH36, and K27). Uncertainties: 2SD of bracketing standard measurements. Several glass data are enriched in δ17O and δ18O, relative to those of chondrule olivine and pyroxene (dark gray shaded region), and they plot along a slope 0.8 line (dashed line). This indicates primary O-isotope ratios of chondrule glasses were disturbed to various extents during low-temperature parent body alteration. Furthermore, Semarkona CH36 glass plots on the fractionation line produced by Semarkona magnetite (Choi et al., 1998; Doyle et al., 2015), inferred to represent that of water that oxidized metal during parent body alteration. The relatively fast O-diffusion rate of glass (e.g., Figure 8.2) means even during mild low-temperature aqueous alteration, chondrule primary glass O-isotope characteristics are susceptible to disturbance. 214 Travis J. Tenner, et al.

(e.g., Figure 8.10) correspond to greater extends of exchange. Jiang et al. (2015) also found chondrule glass with similar characteristics from higher petrologic type ordinary chondrites. Specifically, O-isotopes from L3.6, H3.7, and H4 chondrule glass generally match those of nepheline interpreted to have inherited the O-isotope ratio of aqueous fluid when it replaced plagioclase. Finally, Chaussidon et al. (2008) report glass O-isotope data from Vigarano chondrule 477–2-Ch9 (CV3.1–3.4; Bonal et al., 2007), which are enriched in δ17Oby ~7–11‰, and in δ18O by ~15–20‰, relative to coexisting olivine and pyroxene. These glass data plot right of the CCAM line, suggesting disturbance by parent body alteration.

8.3.2.6 Chondrules With Heterogeneous O-Isotope Ratios Chondrules with heterogeneous SIMS O-isotope data are uncommon. For clarity, heteroge- neous chondrules lack any O-isotope data clustering within analytical uncertainties (e.g., Figure 8.4), meaning their host O-isotope ratios cannot be determined. Thus, heterogeneous chondrules may represent precursor grains that were only marginally melted, or were sintered, preserving their O-isotope variability. Using published and unpublished data from 330 chon- drules, representing 9 chondrite types, Kita et al. (2016) calculate a maximum of 20 percent for CV, but only 0–10 percent from other chondrites, have heterogeneous O-isotope ratios. Excluding relict olivine, Kita et al. (2016) define heterogeneous chondrules as those with four or more SIMS analyses of olivine and/or low-Ca pyroxene, and of which their per-chondrule Δ17O 2SD exceeds 1‰.

8.4 Ranges of Chondrule O-Isotope Ratios in Various Chondrites 8.4.1 Ordinary Chondrite Chondrules

Kita et al. (2010) determined O-isotope characteristics of 36 chondrules from LL chondrites Semarkona (LL3.01), Bishunpur (LL3.15), and Krymka (LL3.2). Olivine and pyroxene data, which define host chondrule O-isotope ratios, parallel the terrestrial fractionation line but plot slightly above it (Δ17O: 0.4‰–2.2‰; Figure 8.11a). A key finding is that, among type I chondrules, refractory element-rich porphyritic olivine (PO) chondrules tend to have the lowest δ18O values, volatile element-enriched porphyritic pyroxene (PP) chondrules tend to have the highest δ18O values, and porphyritic olivine-pyroxene (POP) chondrules plot inter- mediately in δ18O (Figure 8.11b). This trend covers a ~6‰ range in δ18O. Kita et al. (2010) attribute this relationship to an environment with a low dust-to-gas ratio (~100 times solar), where evaporation and condensation of precursors invoked O-isotope and Si fractionation between chondrule melts and ambient gas. Specifically, Kita et al. (2010) hypothesize that, opposite of Rayleigh fractionation, ambient gas favors heavy O-isotopes during evaporation of chondrules with olivine-pyroxene compositions, based on thermodynamic calculations (e.g., Richet et al., 1977; Clayton and Kieffer, 1991). If true, light O-isotope enrichment of chondrule melts would have accompanied increased Mg/Si through multiple heating/evaporation steps in which ambient gas was removed from the system (see Figure 13 from Kita et al., 2010). In turn, condensation of such ambient gas enriched in heavy O-isotopes and Si would explain LL3 type I PP chondrule signatures. Oxygen Isotope Characteristics of Chondrules 215

LL3 chondrite chondrule data

CH61 Ol 6 Type I PO 4 B47 Host Avg. Type I POP Y & R CH26 Host Avg. Type II PP 0 CH26 Rel. Ol 4 Type II -4 O (‰)

17 -8 δ 2 CH44 Type I PO TF CCAM -12 Blue CL Host Ol Host Glass 0 PCM (a)-16 (b) Red CL Rel. Ol

0 2 4 6 -16 -12 -8 -4 0 4 8 δ18O (‰)

Figure 8.11 SIMS LL3 chondrite chondrule data from the study of Kita et al. (2010). (a) Host chondrule O-isotope ratios, where each datum corresponds to the average of homogeneous olivine and pyroxene data per chondrule (uncertainties: 2SD). (b) 16O-rich chondrule data. Individual spot data are shown for CH61, as well as relict olivine data from CH26, with uncertainties shown as the 2SD of bracketing standard measurements. Averages of multiple spot analyses and 2SD uncertainties are shown for B47, CH26 (excluding relict olivine), and red CL olivine, blue CL olivine, and glass in CH44.

Type II chondrules from LL3 chondrites have a limited range of O-isotope ratios, with an average δ18O of 4.5 0.4‰, and an average Δ17O of 0.5 0.6‰ (Figures. 8.11a and 8.11b). Kita et al. (2010) ascribe these characteristics to an environment with higher dust-to-gas ratios

(~1,000–10,000 times solar) and a higher abundance of H2O ice (~1/10 that of CI chondrites) that generated more oxidizing conditions to form type II chondrules. The higher partial pressure of this system would have suppressed evaporation and condensation, and the presence of

H2O would have aided in homogenizing chondrule O-isotope ratios due to efficient exchange between chondrule melt and H2O vapor (e.g., Yu et al., 1995). Kita et al. (2010) report four chondrules with 16O-rich olivine (type I PO CH44 and CH61, type I POP B47, and type II POP CH26), plotting mainly between the Young and Russell and CCAM lines (Figure 8.11c). In particular, type I PO chondrule CH44 is unusual because olivines show cathodoluminescence (CL) images associated with 16O-poor relict cores and 16O-rich overgrowths. Collectively, the 16O-rich signatures of these chondrules indicate the ordinary chondrite accretion region contained at least some precursors from a carbonaceous chondrite-like reservoir. Jiang et al. (2015) studied seven Al-rich chondrules from L3.6, H3.7, and H4 chondrites. Although plagioclase and glass data are most likely disturbed, chondrule pyroxene and olivine are interpreted as having primary O-isotope ratios. Several olivines and pyroxenes have lower δ18O values (0‰ to 6‰) than those from Kita et al. (2010), plotting between the terrestrial fractionation line and the Young and Russell lines. Based on these characteristics Jiang et al. hypothesize the Al-rich chondrules contained a mixture of type C CAI and ferromagnesian 216 Travis J. Tenner, et al. chondrule-like precursors (e.g., Russell et al., 2000), and required multiple heating and exten- sive open-system exchange with 16O-poor gas.

Libourel and Chaussidon (2011) report Fo9299 olivine data from eight Semarkona chon- drules. Five of the eight have δ18O and δ17O (3.8–6.9‰, and 2.0–3.9‰, respectively) that overlap with data from Kita et al. (2010). The other three chondrules have lower δ18O and δ17O(–1.6 to –6.6‰, and –2.2 to –4.9‰, respectively) consistent with host data from Jiang et al. (2015).

8.4.2 Enstatite (E) Chondrite Chondrules

Enstatite chondrite chondrules studied by Weisberg et al. (2011) show an appreciable O-isotope range. Many chondrule olivines, enstatites, and FeO-enriched pyroxenes (2–10 wt% FeO) plot on the terrestrial fractionation line, with δ18Oof3‰–6‰, and overlap with whole-rock enstatite chondrite data (Figure 8.12a). However, several enstatite grains have O-isotopes consistent with those from ordinary chondrites, and two relict olivine data plot within the range of R chondrites (Figure 8.12a). Further, several olivine grains from two chondrules plot below the terrestrial fractionation line, forming a slope ~1 line with δ18O and δ17O values extending to ~–5‰ and ~–7‰, respectively (Figures 8.12a and 8.12b). Weisberg et al. (2011) refer to this line as the Enstatite Chondrule Mineral (ECM) line, but note it coincides with the PCM line. Overall, while most phenocrysts from enstatite chondrite chondrules have O-isotope ratios consistent with formation from a reservoir on the terrestrial fractionation line, at least some precursor

Enstatite chondrite chondrule minerals

6 Ol RC 0 En Fe 4 Px Y & R -2 TF PCM

O (‰) OC

17 -4

δ CCAM 2 EC -6 (a) (b) 0 -8 2 4 6 8 -6 -4 -2 0 2 δ18O (‰)

Figure 8.12 SIMS enstatite chondrite chondrule mineral data from Weisberg et al. (2011). Ol = olivine; En = enstatite; Px = pyroxene. In panel (a) enstatite chondrite (EC) whole rock data (oval) are from Clayton and Mayeda (1984) and Weisberg et al. (1995); the oval representing ordinary chondrite whole rocks (OC) is based on data from Clayton et al. (1991), and the oval representing R chondrite whole rocks (RC) is based on data from Weisberg et al. (1991), Bischoff et al. (1994), Kallemeyn et al. (1996), Bischoff et al. (2011), and Isa et al. (2014). Panel (b) shows olivine data that are comparatively 16O-rich and plot on/near the PCM line. Oxygen Isotope Characteristics of Chondrules 217

O-isotopes are consistent with those from other chondrites. Further details of enstatite chondrite chondrules are provided in Chapter 7.

8.4.3 Rumuruti (R) Chondrite Chondrules

SIMS O-isotope data from Rumuruti (R) chondrite chondrules are reported by Miller et al. (2017) from Mount Prestrud (PRE) 95404, which hosts type 3.2 material. Type I and type II chondrules are present, and their Mg#’s(74–98.5, with one chondrule having an Mg# of 58.6) are similar to those from LL3 chondrite chondrules (71–99.6; Kita et al., 2010). This suggests similar fO2 ranges during R and LL chondrite chondrule formation (1–3.9 log units below the iron-wüstite buffer), based on metal-silicate phase equilibria (e.g., Ebel and Grossman, 2000). O-isotope ratios of PRE 95404 chondrules (δ18Oandδ17O: 0‰–4.5‰) also overlap with LL3 chondrite chondrule data from Kita et al. (2010), and are distinct from bulk R chondrite O-isotope ratios (Figure 8.13). As previously suggested (e.g., Greenwood et al., 2000; Isa et al., 2011; Kita et al., 2015), these results indicate most R and LL chondrite chondrules formed in the same O-isotope environment. Three relict olivine grains plot below host chondrule data, between the PCM and CCAM lines, and a matrix refractory forsterite grain is relatively 16O-rich (δ18Oandδ17O: –5.3‰), plotting near the Young and Russell line (Figure 8.13). Based on the difference between bulk and chondrule O-isotope ratios, Miller et al. (2017) suggest the R chondrite matrix is relatively 16O-poor, reflecting signatures of original silicate and/

R chondrite chondrules

6 Type I RC 4 Type II

2 LL3 chondrite 0 chondrules O (‰) 17 -2 Y & R TF CCAM -4 PCM Relict Ol -6 Refrac. Ol

-8 -8 -6 -4 -2 0 2 4 6 d18O (‰)

Figure 8.13 SIMS Rumuruti (R) chondrite chondrule data from Miller et al. (2017). Each datum is the average of spot measurements per chondrule (uncertainties: 2SD), excluding relict olivine (Ol) grains. The gray oval represents the range of LL3 chondrite chondrule data from Kita et al. (2010) and the black oval represents R chondrite whole rock data (RC) from Weisberg et al. (1991), Bischoff et al. (1994), Kallemeyn et al. (1996), Bischoff et al. (2011), and Isa et al. (2014). 218 Travis J. Tenner, et al.

or ice grains. Using (1) the average R chondrite chondrule/matrix ratio (0.79); (2) their averaged chondrule δ18Oandδ17Oof2.3‰ and 2.1‰, respectively; and (3) a bulk R chondrite δ18Oand δ17Oof5.1‰ and 5.4‰, respectively (Bischoff et al., 2011), Miller et al. (2017) estimate a bulk R chondrite matrix δ18Oandδ17Oof7.1‰ and 7.8‰ (Δ17O: 4.1‰), respectively.

8.4.4 Grosvenor Mountains 95551 and Northwest Africa 5492 (G) Chondrite Chondrules

Weisberg et al. (2015) investigated O-isotopes of chondrules from Grosvenor Mountains (GRO) 95551 and Northwest Africa (NWA) 5492. These chondrites are metal-rich, but are not related to CH or CB chondrites, based on their siderophile elemental abundances, average metal compositions, and O-isotope characteristics. Further, they differ from CB chondrites because their silicates are more reduced, and because their metals and sulfides are not intergrown. As such, Weisberg et al. (2015) propose GRO 95551 and NWA 5492 belong to a unique chondrite set, which they call G chondrites. Currently, only these two chondrites have this designation. Like R chondrite chondrules, O-isotope ratios of most G chondrite chondrules plot within the range of LL3 chondrite chondrule data from Kita et al. (2010), with one chondrule that is consistent with R chondrite whole rock data (Figure 8.14). However, G chondrite chondrules

G chondrite chondrules 8

FeO-poor R chondrites 6 FeO-rich

4 LL3 chondrite chondrules O (‰)

17 2 δ

0 TF

-2 Y & R PCM CCAM

-2 0 2 4 6 8 δ18O (‰)

Figure 8.14 GRO 95551 and NWA 5492-like (G) chondrite chondrule SIMS data from Weisberg et al. (2015). Each datum is the average of spot measurements per chondrule (uncertainties: 2SD). Chondrules consist of 28 porphyritic (27 FeO-poor, 1 FeO-rich), 2 barred olivine, 2 enstatite-metal angular fragments, 1 radial pyroxene, 1 Al-rich, and 1 Fe-rich pyroxene lithic fragment. The gray oval represents the range of LL3 chondrite chondrule data from Kita et al. (2010) and the black oval represents R chondrite whole rock data from Weisberg et al. (1991), Bischoff et al. (1994), Kallemeyn et al. (1996), Bischoff et al. (2011), and Isa et al. (2014). Oxygen Isotope Characteristics of Chondrules 219 differ from LL3 and R chondrite chondrules because their silicates are predominantly magnes- ian (Mg#’s: 97.0–99.8), with only 3 of 35 chondrules studied having FeO-rich pyroxenes with Mg#’s ranging from 60 to 91. This suggests a greater proportion of G chondrite chondrules formed at reducing conditions, based on metal-silicate phase equilibria. Nonetheless, Weisberg et al. (2015) point out the considerable overlap in O-isotope ratios between G, LL3, and R chondrite chondrules, as well as enstatite chondrite chondrules (e.g., Figures 8.11–8.14), indicates mixing between these chondrule and parent body forming environments. As ordinary, E, R, and G chondrites are relatively dry, and have isotopic similarities to the Earth and Moon, Weisberg et al. (2015) suggest they represent materials from the inner solar system.

8.4.5 Kakangari (K) Chondrite Chondrules

SIMS O-isotope ratios of chondrule and matrix olivine and pyroxene in Kakangari (K) chondrites were investigated by Nagashima et al. (2015). Olivines have forsterite values of 95–97 and pyroxenes have enstatite values of 90–96. This uniformity is likely related to partial equilibration of MgO, FeO, and MnO during parent body thermal metamorphism, as Kakangari has taenite grains with Ni zoning profiles consistent with cooling between 1 and 10 C/Myrat~500–600 C(e.g., Berlin, 2009). Further, the presence of apatite around chondrule peripheries is a feature of petrologic type >3.5 chondrites (e.g., Huss et al., 2006). However, it is unlikely primary O-isotope ratios of olivine and pyroxene were disturbed because of their slow O-diffusion rates (Figure 8.2). Among Kakangari chondrules, most olivines and pyroxenes plot on the terrestrial fractionation line, with δ18Ofrom0‰ to 4‰. These data generally overlap bulk Kakangari chondrule, matrix, and whole rock data (Prinz et al., 1989; Weisberg et al., 1996). However, a chondrule olivine grain within a coarse igneous rim is extremely 16O-rich, plotting near AOA data with a δ17Oof–45.2‰ and a δ18Oof–46.4‰ (Figure 8.15a). Matrix olivines and pyroxenes have nearly identical O-isotope

K chondrite chondrule minerals K chondrite matrix minerals 4 10 (a) (b) 0 2 -10

0 -20

O (‰) TF -40

17 -30

δ AOA -2 Y & R -45 -40 LPx -50 Ol PCM -50 -4 CCAM -50 -45 -40 -60 -2 0 2 4 6 -60 -50 -40 -30 -20 -10 0 10 δ18O (‰)

Figure 8.15 Kakangari (K) chondrite chondrule mineral and AOA data (a) and matrix mineral data (b) from Nagashima et al. (2015). LPx = low-Ca pyroxene; Ol = olivine. Uncertainties are the reported 2SD. 220 Travis J. Tenner, et al. characteristics as those from chondrules, with some plotting on the TF line, and others having extremely 16O-rich values (Figure 8.15b). This indicates Kakangari chondrules and matrix are related, meaning (1) matrix materials were likely a component of chondrule precursors; and/or (2) matrix minerals experienced similar thermal processing as chondrule precursors.

8.4.6 Carbonaceous Chondrite Chondrules

8.4.6.1 Acfer 094 and CO3 Chondrite Chondrules Acfer 094 (ungr. C3.00) and CO chondrite chondrules have near-identical O-isotope character- istics, largely defined by a bimodal distribution of their host values (Figures 8.16a and 8.16b, top panels). Most type I chondrules from Acfer 094 and Yamato (Y) 81020 (CO3.05) are relatively 16O-rich, with Mg#’s > 97 and Δ17O near 5‰, while Mg# < 97 type I chondrules and type II chondrules are relatively 16O-poor, with Δ17O near 2‰. In both chondrites, there are relict olivines with Δ17O near 5‰ and 2‰ within some chondrules (Figures 8.16a and 8.16b, bottom panels), indicating transport of materials between these O-isotope environments. Tenner et al. (2013) also interpret three low-Ca pyroxene grains from Y-81020 as relicts, with Δ17O values of –3.6‰ to –5.0‰. Significantly 16O-rich relict olivines are also found in Acfer 094 and CO chondrites (Figures 8.16a and 8.16b, bottom panels), suggesting a refractory origin. Four Acfer 094 chondrules have heterogeneous O-isotope ratios, and their phase data cover the range of host and relict olivine values.

8.4.6.2 CV3 Chondrite Chondrules Chaussidon et al. (2008) and Libourel and Chaussidon (2011) report O-isotope ratios of chondrule phases from Allende, Vigarano, Mokoia, and Efremovka, and Rudraswami et al. (2011) investigated O-isotopes of chondrule phases from Allende. Their data overlap (1) each other; and (2) bulk chondrule data on an oxygen three isotope plot (Figure 8.17a–8.17c), covering a host δ18O range of ~15‰–~+5‰ along the PCM line. Per-chondrule olivine and pyroxene O-isotope data from Rudraswami et al. (2011) generally overlap, allowing for determining host chondrule O-isotope ratios, while corresponding data from Chaussidon et al. (2008) are systematically displaced in δ17O and δ18O (e.g., Figure 8.6; top panels). As such, we show the range of individual spot data from Chaussidon et al. (2008) in Figure 8.17b. Libourel and Chaussidon (2011) report data from olivine grains in FeO-poor chondrules, and their averaged spot data per chondrule are shown in Figure 8.17b. Based on host data from Rudraswami et al. (2011), which exclude plagioclase with disturbed O-isotope ratios (e.g., Figure 8.9b), FeO-poor Allende chondrules (n =19)haveΔ17O values of 5.3‰ 0.6‰ (2SD), and five porphyritic chondrules (FeO-rich and FeO-poor) have Δ17O between 2‰ and 3‰. In addition, averaged olivine Δ17O from FeO-poor CV3 chondrules reported by Libourel and Chaussidon (2011) range from 4.4‰ to 7.3‰.These values are similar to those found among Acfer 094 and CO chondrite chondrules (Figure 8.16). Four FeO-poor BO chondrules investigated by Rudraswami et al. (2011) have host Δ17Oranging from 5‰ to 0‰. These values overlap with host O-isotope data from porphyritic chondrules, suggesting both textures were produced in the same environment. These data also indicate not all CV3 BO chondrules plot on the terrestrial fractionation line, as had been suggested from previous Oxygen Isotope Characteristics of Chondrules 221

5 (a) Acfer 094. 5 (b) Y-81020. CO3.05 ungr. C3.00 0 0

-5 TF -5

-10 -10

PCM Host Type I Host Type I -15 Host Type II -15 Host Type II

-15 -10 -5 0 5 10 -15 -10 -5 0 5 10 O (‰) 17 δ 10 10 Acfer 094 Y-81020 0 Relict Ol 0 Relict Ol Heterogeneous Relict LPx -10 Spot Data -10

-20 -20

-30 -30

-40 -40

-50 -50 -50 -40 -30 -20 -10 0 10 -50 -40 -30 -20 -10 0 10 δ18O (‰)

Figure 8.16 SIMS host chondrule O-isotope data from Acfer 094 (ungr. C3.00) and Yamato (Y) 81020 (CO3.05) chondrites (a & b, respectively, top panels) and their corresponding relict grain and heterogeneous chondrule spot data (a & b, respectively, bottom panels). Ol = olivine; LPx = low-Ca pyroxene. For host chondrules, each datum is the average of homogeneous spot data for one chondrule, showing 2SD uncertainties. For relict grain and heterogeneous spot data the uncertainties are the 2SD of bracketing standard measurements. Data sources: Acfer 094: Ushikubo et al. (2012); Y-81020: Tenner et al. (2013). bulk studies mainly of FeO-rich BO chondrules (e.g., Clayton and Mayeda, 1983). Nine of the 36 chondrules studied by Rudraswami et al. (2011) have olivine data with heterogeneous O-isotope ratios (Figure 8.17a, bottom panel). Some of these data are interpreted as having refractory origins, due to their low Δ17O(downto18‰). However, most heterogeneous chon- drule phenocryst data overlap the range of host chondrule values, suggesting common origins.

8.4.6.3 CR2 Chondrite Chondrules Recent SIMS O-isotope data from CR2 chondrite chondrules come from Krot, Libourel, and Chaussidon (2006b), Libourel and Chaussidon (2011), Connolly and Huss (2010), Schrader et al. (2013; 2014a; 2017), and Tenner et al. (2015). Their data overlap, plotting along the PCM line with host chondrule δ18O values of ~5‰– ~+7‰ (e.g., Figure 8.18). 222 Travis J. Tenner, et al.

CV chondrite chondrule data

5 (a) Rudraswami 5 (b) SIMS 0 et al. (2011) spot data 0 -5 -5 TF -10 -10 C et al. 08.Ol -15 Host Type I -15 C et al. PCM Host Type II 08.Px -20 Host -20 L&C 11. FeO-poor BO Ol -25 -25 -20 -10 0 10 -20 -10 0 10 O (‰) 17

δ 5 5 Rudraswami (c) bulk 0 et al. (2011) 0 chondrule data -5 -10 -5

-15 -10 -20 -15 -25 Relict Ol Clayton et al. 83 BO Heterogeneous -20 Clayton et al. 83 porph -30 Chondrule Spot Rubin et al. 90 porph Data Jones et al. 04 -35 -25 -30 -25 -20 -15 -10 -5 0 5 10 -20 -10 0 10 δ18O (‰)

Figure 8.17 CV3 chondrite chondrule O-isotope data by SIMS (panels a & b) and by bulk analysis (panel c). (a) top panel: Host O-isotope data from Allende chondrules (Rudraswami et al., 2011). Each datum represents the average of homogeneous SIMS spot data for one chondrule, showing 2SD uncertainties. bottom panel: SIMS spot data from relict olivine and from nine heterogeneous Allende chondrules from Rudraswami et al. (2011). Uncertainties are the 2SD of bracketing standard measurements (b) SIMS data from Chaussidon et al. (2008) and from Libourel and Chaussidon (2011) (L&C 11). Collectively, the data come from Allende, Vigarano, Efremovka, and Mokoia. Data from Chaussidon et al. (2008) are individual olivine (Ol) and pyroxene (Px) data from their electronic annex, along with their reported 2SD. Regarding Libourel and Chaussidon (2011), each datum is the average of olivine spot measurements for one chondrule, showing 2SD uncertainties. (c) bulk chondrule O-isotope data. Each datum represents one chondrule. Data from Clayton and Mayeda (1983) and Rubin et al. (1990) are from Allende, and data from Jones et al. (2004) are from Mokoia.

Connolly and Huss (2010) investigated O-isotopes of type II CR2 chondrite chondrules, which comprise ~4 percent of the chondrule population (i.e., ~96 percent of chondrules have Mg#’s > 90; Schrader et al., 2011; 2015). The type II data are generally 16O-poor (i.e., Δ17O: 2.0‰– +1.1‰) relative to type I host chondrule data from Krot et al. (2006b) and Libourel and Chaussidon (2011) (Figure 8.18a). More recent investigations by Schrader et al. Oxygen Isotope Characteristics of Chondrules 223

CR chondrite chondrules

5 5 (a) K (2006b); (b) Schrader et al. L & C (2011) (2013; 2014a; 0 (type I) 0 2017)

TF -5 -5

Connolly & Huss (2010) -10 PCM -10 (type II)

-10 -5 0 5 10 -10 -5 0 5 10 O (‰) 17 δ

5 (c) Schrader 5 (d) Tenner et al. (2014a) et al. (2015)

0 0

-5 -5 Host Type I Host Type II Barred Olivine Relict Ol -10 -10 Heterogen. Chondrules Spot Data

-10 -5 0 5 10 -10 -5 0 5 10 δ18O (‰)

Figure 8.18 SIMS chondrule O-isotope data from CR chondrite chondrules. Each host datum represents the average of homogeneous spot data measured within a single chondrule, showing 2SD uncertainties. Uncertainties for relict olivine (Ol) grains and spot data from heterogeneous chondrules are the reported 2SD. Unless otherwise noted, data come from porphyritic chondrules. A small number of host type I data are from chondrules that could also be defined as Al-rich. In panel (a), type I chondrule data are from Krot et al. (2006b) and Libourel and Chaussidon (2011). With regard to (b), there are two data from Schrader et al. (2017) that plot outside of the panel parameters, a relict olivine grain and a relict plagioclase grain that are significantly 16O-rich. In panel (d) one of the host type I chondrules has a BO texture. Regarding data from Schrader et al. (2014a; 2017), calculated host isotope ratios omit spot analyses of silica and mixed phases, as corrections for SIMS instrumental biases may be incorrect.

(2013; 2014a; 2017) and Tenner et al. (2015) determined host O-isotope ratios of type I and type II CR2 chondrite chondrules (Figure 8.18b–8.18d). Like the previous studies, they find type I chondrules are predominantly 16O-rich relative to type II chondrules, with host δ18O values extending to 5‰ on the PCM line. Schrader et al. (2014a) show this type I versus type II distinction among eleven barred olivine chondrules (Figure 8.18c) and their host O-isotope ratios overlap with those from porphyritic chondrules (e.g., Figures 8.18a, 8.18b, 8.18d). As BO chondrules likely experienced more complete melting than porphyritic chondrules, their 224 Travis J. Tenner, et al.

O-isotope ratios should decidedly reflect the bulk of their precursors. Based on this, Schrader et al. (2014a) suggest the overlap between host BO and porphyritic chondrule O-isotope ratios means (1) they formed in the same general precursor environment; and (2) that the O-isotopes of olivines from such porphyritic chondrules are not relict in origin. Among 100+ CR2 chondrules investigated in the aforementioned studies, there is a high- degree of O-isotope homogeneity per chondrule, excluding relict grains. Only four chondrules are identified as isotopically heterogeneous. Few relict olivine grains are found in CR2 chondrite chondrules, and when identified, the majority have O-isotope ratios plotting within the range of host chondrule values (Figure 8.18), indicating common origins.

8.4.6.4 Yamato 82094 (Ungrouped C3.2) Chondrules Tenner et al. (2017) measured SIMS O-isotope ratios of 35 chondrules from Yamato (Y) 82094 (ungr. C3.2). Per chondrule, all chondrules with multiple pyroxene data (21 of 21) exhibit O-isotope homogeneity of this phase; all chondrules with multiple coexisting pyroxene plus plagioclase data (17 of 17) show that these two phases have homogeneous O-isotope ratios per chondrule. Among 25 chondrules with olivine plus pyroxene data, 21 have at least one olivine datum that matches O-isotopes of coexisting pyroxene data. The O-isotope homogeneity among these phases indicates they crystallized from the final chondrule melt, defining host O-isotope ratios. Host O-isotope ratios of Y-82094 chondrules plot within 0.7‰ of the PCM line, with δ18O ranging from ~11‰ to ~4‰ (Figure 8.19a), similar to Acfer 094, CO, CV, and CR chondrite chondrules (Figures 8.16–8.18). Among Mg# > 90 chondrules no systematic

Y-82094 (ungr. C3.2) chondrules

a LL3 0 b chondrite 0 chondrules -10 TF -5 -20 O (‰)

17 Host δ Type I -10 Host -30 Al-rich Relict Ol PCM Host Relict Sp -40 -15 Type II

-10 -5 0 5 -40 -30 -20 -10 0 δ18O (‰)

Figure 8.19 O-isotope ratios of Yamato (Y) 82094 chondrules (Tenner et al., 2017). (a) host O-isotope ratios of Y-82094 chondrules. Each datum represents the average of homogeneous SIMS O-isotope spot data per chondrule, showing 2SD uncertainties. Type I and Al-rich chondrules are FeO-poor (Mg# > 90). The dashed oval represents the range of host LL3 chondrule O-isotope ratios reported by Kita et al. (2010). (b) relict olivine (Ol) and spinel (Sp) SIMS spot data from Y-82094 chondrules. Uncertainties are the 2SD of bracketing standard measurements. Oxygen Isotope Characteristics of Chondrules 225 differences exist when comparing ranges of type I porphyritic and Al-rich host O-isotope ratios. Three type II chondrules plot slightly above the PCM line and near the terrestrial fractionation line (Δ17O: ~0.1‰). Their O-isotope ratios, as well as their olivine MnO (wt%) versus forsterite content (e.g., Figure 6 from Tenner et al., 2017), are consistent with LL3 chondrite chondrules. Seventeen chondrules have relict olivine and/or spinel grains. Most relict olivines plot within the range of host Y-82094 chondrule values, indicating common origins (e.g., Figure 8.19b). However, some relict olivines, and all relict spinels, are significantly 16O-rich when compared to host Y-82094 chondrule values, suggesting a refractory origin.

8.4.6.5 CH, CH/CBb and CBb Chondrite Chondrules

Krot et al. (2010b) investigated chondrules from the CH/CBb chondrite Isheyevo and from CBb chondrites MacAlpine Hills (MAC) 02675 and Queen Alexandra Range (QUE) 94627. Twenty porphyritic chondrules from Isheyevo (11 type I, 9 type II), and 51 cryptocrystalline chondrules from Isheyevo, MAC 02675 and QUE 94627 (43 magnesian, 6 magnesian inclusions inside chemically-zoned Fe,Ni-metal condensates, and 2 ferroan) were studied. Regarding magnesian cryptocrystalline chondrule O-isotopes, data from the three chondrites are indistinguishable, plotting on the PCM line with δ18Oof–1.1‰–2.9‰ (Figure 8.20a). The two Isheyevo ferroan cryptocrystalline chondrules plot slightly above the terrestrial fraction- ation line (Δ17O: 1.1‰ and 1.5‰, respectively), with higher δ18O (3.7‰ and 6.7‰, respect- ively) than magnesian cryptocrystalline chondrules. The O-isotope uniformity of magnesian cryptocrystalline chondrules suggests CH and CB chondrites are related, and their textures indicate complete melting and rapid crystallization (Chapter 2). Based on these characteristics, Krot et al. (2010b) hypothesize such chondrules formed as single-stage gas–melt condensates by asteroidal impact. Support for this hypothesis comes from magnesian CBb chondrite chondrules enriched in light Mg isotopes (Gounelle et al., 2007), and relatively young chondrule ages that are inconsistent with formation by nebular shock waves (Krot et al., 2005). Porphyritic Isheyevo chondrules occupy a larger range of host δ18O along the PCM line (~–1‰ to ~11.5‰; Figure 8.20b). Two porphyritic chondrules were identified by Krot et al. (2010b) as heterogeneous; their spot data, all from olivine grains, occupy a large δ18O range along the PCM line (–34‰– +0.5‰) with one grain identified as a relict (Figure 8.20c). The 16O-rich nature of some heterogeneous chondrule data indicate refractory origins. Nine of twenty Isheyevo porphyritic chondrules have host O-isotope ratios plotting above the terrestrial fractionation line, including seven of eleven type I chondrules (Figure 8.20b). As such, no systematic differences exist when comparing Isheyevo type I and type II chondrule O-isotope ratios. These observations of host data differ relative to chondrules from other carbonaceous chondrites (e.g., Figures 8.16–8.19), in which (1) the vast majority of host O-isotope ratios plot below the terrestrial fractionation line; and (2) type I chondrules are generally 16O-enriched relative to type II chondrules. These differences, plus the common presence of 26Al-poor relict CAIs in chondrules that are mineralogically and isotopically similar to host meteorite CAIs (e.g., Krot et al., 2007), led Krot et al. (2010b) to conclude that Isheyevo and other CH porphyritic chondrules formed in a different environment or at a different time than other carbonaceous chondrite chondrules. Nakashima et al. (2011) also investigated SIMS O-isotope systematics of magnesian cryptocrystalline chondrules from CH chondrite Sayh al Uhaymir (SaU) 290. δ17O and 226 Travis J. Tenner, et al.

Isheyevo (CH/CBb), MAC 02675 (CBb), QUE 94627 (CBb) and SaU 290 (CH) chondrite chondrule data 15 15 Magnesian Cryptocrystalline Isheyevo 10 Isheyevo 10 Porphyritic PCM MAC 02675 Host QUE 94627 Type I 5 TF 5 Host Type II 0 0 (a) Krot et -5 al. (2010) -5 Isheyevo (b) Krot et -10 Ferroan Cyptocrystalline -10 al. (2010)

-10 -5 0 5 10 15 -10 -5 0 5 10 15 O (‰)

17 15 δ 10 Magnesian Cryptocrystalline 5 (c) Krot et 10 SaU 290 0 al. (2010) -5 5 -10 -15 0 -20 -25 Isheyevo porphyritic -5 -30 Relict Ol (d) Nakashima -35 heterogeneous spot data -10 et al. (2011) -40 -40 -35 -30 -25 -20 -15 -10 -5 0 5 10 15 -10 -5 0 5 10 15 δ18O (‰)

Figure 8.20 O-isotope ratios of Isheyevo (CH/CBb), MacAlpine Hills (MAC) 02675 (CBb) and Queen Alexandra Range (QUE) 94627 (CBb) chondrules from Krot et al. (2010b), and Sayh al Uhaymir (SaU) 290 (CH) chondrite chondrules from Nakashima et al. (2011). Uncertainties are 2SD. (a) magnesian (Mg# > 90) and ferroan (Mg# < 90) cryptocrystalline chondrule spot data from Krot et al. (2010b). Three chondrules have two spot measurements and the rest of the chondrules have a single spot measurement. MAC 02675 and QUE 94627 data include those from cryptocrystalline silicate inclusions inside chemically zoned Fe,Ni-metal condensates (n = 3 and 3 inclusions, respectively). (b) porphyritic chondrule data from Isheyevo (Krot et al., 2010b). Each datum represents the single spot measurement or average of multiple spot measurements from one chondrule. (c) olivine (Ol) spot data from two heterogeneous chondrules and a relict grain, from Isheyevo, as identified from Krot et al. (2010b). (d) data from SaU 290 (CH) magnesian cryptocrystalline chondrules from Nakashima et al. (2011). Each datum represents the average of three SIMS spot analyses per chondrule.

δ18O were determined in eight of the nine chondrules studied, and 6 of the 8 overlap with

CH/CBb and CBb magnesian cryptocrystalline chondrule data from Krot et al. (2010b) (Figure 8.20d). Nakashima et al. interpret that these chondrules formed in the same environment as CH/CBb and CBb magnesian cryptocrystalline chondrules. However, two SaU 290 magnesian Oxygen Isotope Characteristics of Chondrules 227 cryptocrystalline chondrules significantly deviate from the other data (Figure 8.20d), plotting at the 16O-rich and 16O-poor ends of the host Isheyevo porphyritic chondrule data (e.g., Figure 8.20b); additionally, chondrule 09 from Nakashima et al. (2011) has Δ17O values ranging from –5.2‰ to –7.4‰ (but no reported δ17O and δ18O data), and differs from the majority of magnesian cryptocrystalline chondrule data with Δ17O near –2‰. As such, Nakashima et al. (2011) suggest at least some magnesian cryptocrystalline chondrules formed in a similar environment as that which formed type I Isheyevo porphyritic chondrules.

8.5 Mg# and O-Isotope Relationships of Chondrules: Links to Physicochemical Conditions of Their Formation Environment 8.5.1 Mg# Versus Δ17O: Qualitative Insights

Chondrule Mg#’s are valuable because they document redox states of chondrule formation, based on metal-silicate phase equilibria. Qualitatively, decreases in chondrule Mg# correspond to more oxidized formation conditions (e.g., Kring, 1988; Ebel and Grossman, 2000). In turn, redox conditions were likely controlled by proportions of dust, gas, and H2O ice within chondrule-forming environments (e.g., Ebel and Grossman, 2000; Fedkin and Grossman, 2006; Grossman et al. 2008; Fedkin and Grossman, 2016). Qualitatively, increased dust-to- gas ratios and increased ice abundances in chondrule precursors correspond to more oxidizing chondrule-forming conditions. As such, chondrule Mg#’s can constrain these regional charac- teristics, especially when combined with corresponding chondrule O-isotope ratios. Most LL3, E, R, G, and K chondrite chondrules have a narrow range of host Δ17O, respectively, plotting on or parallel to the TF line (Figures 8.11–8.15). However, they have an appreciable Mg# range (~60–100; Kita et al., 2010; Weisberg et al., 2011; 2015; Nagashima et al., 2015; Miller et al., 2017), suggesting variable redox conditions within respective chondrule- forming environments, potentially due to significant ranges of dust-to-gas ratios. The relatively constant Δ17O values of LL3, E, R, G, and K chondrite chondrules suggests bulk O-isotope ratios of respective precursor dusts and gases were constant. However, observed mass-dependent O-isotope fractionation among LL3, E, R, G, and K chondrite chondrules indicates evaporation/ condensation processes and gas–melt interactions, especially at low dust-to-gas ratios. Regarding carbonaceous chondrites, several have type II chondrules that are systematically 16O-poor relative to coexisting type I chondrules, plotting along the PCM line (e.g., Figures

8.16–8.19); CH, CH/CBb and CBb chondrite chondrules do not exhibit this type I/type II relationship (e.g., Figure 8.20), suggesting they formed in a different environment relative to other carbonaceous chondrite chondrules. In more detail, chondrules from Acfer 094, CO, CV, CR, and Y-82094 chondrites generally show increases in Δ17O with decreasing chondrule Mg# (Figure 8.21). As such, Connolly and Huss (2010) were among the first to hypothesize that, in addition to increased dust-to-gas ratios, such a trend originated from addition of an oxidizing 16 agent to high-Mg# chondrule precursors, which they inferred to be O-poor H2O ice (e.g., Figure 8.1). This hypothesis is supported by observations of forsteritic relict olivine grains in some type II chondrules that are 16O-rich and have similar values as host type I chondrules (e.g., Kunihiro et al. 2004; 2005; Connolly and Huss, 2010; Ushikubo et al., 2012; Tenner et al., 2013; Schrader et al., 2013). 228 Travis J. Tenner, et al.

f log O2 – log IW CI dust/gas ratio 2 1 -3.5 -3-2.5 -2 -1 50 100 200 400 2000 0 a b –1 –2 –3 –4 –5 –6 Acfer 094 –7 ungr. C3 –8 CV3 ox CO3 –9 CV3 red O (‰)

17 1 cd Δ 0 –1 –2 –3 –4 –5 –6 –7 –8 –9 CR2 Y-82094 100 99 98 97 96 95 94 90 80 70 60 50 40 99 98 97 96 95 94 90 80 70 60 50 40 30 chondrule Mg#

Figure 8.21 Chondrule Mg#’s versus host Δ17O from various carbonaceous chondrites. Δ17O uncertainties are 2SD, and Mg# uncertainties are those reported from studies. In some datasets it is assumed reported olivine forsterite contents represent the chondrule Mg#. Oxygen fugacities and CI dust-to-gas ratios are calculated using Table 8.3 and Eqn. 8.1 (page 229). (a) Acfer 094 (Ushikubo et al., 2012) and Yamato 81020 CO3.05 (Tenner et al., 2013) data. (b) Allende chondrule data from Rudraswami et al. (2011) (CV3 ox), from which chondrule Mg#’s were calculated using only low-Ca pyroxene data, and reduced CV3 chondrule data from Efremovka and Vigarano (Libourel and Chaussidon, 2011). (c) CR2 chondrite chondrule data from Connolly and Huss (2010), Schrader et al. (2013; 2014a; 2017) and Tenner et al. (2015). (d) Yamato (Y) 82094 (ungr. C3.2) chondrule data from Tenner et al. (2017).

8.5.2 Quantifying the Links Between Chondrule Mg#, Oxygen Fugacity, and Precursor Assemblages

Based on the Mg# versus Δ17O relationships among chondrite chondrules, the influence of precursor components on redox conditions during chondrule formation can be quantified. Links between chondrule Mg#, oxygen fugacity, and dust and H2O ice enrichment are illustrated in Figure 8.22, using CI dust and solar gas compositions (Table 8.3) as examples. Oxygen fugacity is strongly controlled by ratios of hydrogen, carbon, and oxygen in the precursor assemblage (e.g., Eqn. 24 from Grossman et al., 2008), as these elements are dominant in solar gas, CI dust, and

H2O ice (e.g., Table 8.3). Relative to gas of a solar composition (atomic H/O: 2041; Table 8.3), increased proportions of H2O ice (atomic H/O: 2) and/or CI dust (atomic H/O: 1.1; e.g., Table 8.3) diminish the bulk atomic H/O of a chondrule precursor assemblage, making it more oxidizing Oxygen Isotope Characteristics of Chondrules 229

IW [1] 1000x IW–1 [2] 125x 500x [3] 1000x IW–2 "icy" 500x 125x IW–3 75x 100x 50x 100x "icy" IW–4 Solar Solar IW–5 gas relative to the Fe-wüstite buffer

2 factor of CI dust added to O

f gas of a Solar compostion IW–6

log a b IW–7 1 10 100 1000 95 90 85 80 75 70 65 60 55 atomic H/O ratio of chondrule Mg# precursor assemblage

Figure 8.22 (a) atomic H/O versus oxygen fugacity (log fO2 relative to the Fe-wüstite [IW] buffer) of various assemblages. For a dust enrichment of n, a bulk assemblage is calculated by adding (n – 1) units of dust to gas. References [1–3]: Ebel and Grossman (2000), Fedkin and Grossman (2006), and Grossman et al. (2008), respectively. In Ebel and Grossman (2000) and Fedkin and Grossman (2006), dust of a CI chondritic composition was employed. “Icy” dust from Fedkin and Grossman (2006) assumes 25.5 wt%

H2O. Ebel and Grossman (2000) use the Solar gas composition from Anders and Grevesse (1989), while Fedkin and Grossman (2006) and Grossman et al. (2008) employ a modified Solar gas composition, using C and O atomic abundances from Allende Prieto et al. (2001; 2002). “Icy” Solar gas from Grossman et al. (2008) assumes 9 times the amount of water already present in the inner Solar nebula (based on the Ciesla and Cuzzi (2006) model), corresponding to 6.62 107 atoms of oxygen and 2.80 107 atoms of hydrogen (normalized to 1 106 atoms of Si). Open symbols are raw oxygen fugacities, while closed symbols are data corrected to carbon-free systems. The carbon correction is based on a fit to Solar gas data from Figure 10 of Grossman et al. (2008) at 1,480 K. Collectively the fit to the closed symbols is Eqn. 8.1. (b) chondrule Mg# versus oxygen fugacity. Data points are taken from Figure 8a from Ebel and Grossman (2000) at 1480K, the temperature at which >98 percent of Fe condensed for any dust-to-gas ratio above 100. It is assumed olivine fayalite compositions match chondrule Mg#’s. For a given chondrule Mg#, the corresponding oxygen fugacity is calculated from the H/O and C/O ratio of the dust enriched assemblage, using the fit to the data shown in panel (a) (Eqn. 8.1). The fit to the data points in panel (b) is Eqn. 8.2.

(note that, in part, the atomic H/O abundance of CI dust is low due to an H2O concentration of 18.1 wt%). Atomic C/O ratios of solar gas and CI dust are not as different relative to one another

(0.5 and 0.1 respectively; Table 8.3), and so C/O does not influence fO2 as strongly from changes to the dust-to-gas ratio of an assemblage. Effects of H/O and C/O on redox conditions imposed by an assemblage are shown in Figure 8.22a, where log fO2 is calculated relative the iron-wüstite (Fe-FeO or IW) buffer. For reference, Figure 4 from Ebel and Grossman (2000) shows that log fO2 versus temperature curves of dust-enriched assemblages are concentric with the IW buffer from 1,200 to 2,400K. Based on an empirical fit to thermodynamic equilibrium model data (e.g., Figure 8.22a), the oxygen fugacity of a chondrule precursor assemblage is calculated as:

log f O2 log IW ¼0:7740 lnðÞ H=O 0:0217 þ 1:2007 lnðÞ 1 C=O (8.1) 230 Travis J. Tenner, et al.

Table 8.3 Atomic abundances of solar gas and CI dust used in the mass- balance model from Tenner et al. (2015)

Solar Gas CI Dust

H 2.79E+10 8.26E+06(a) He 2.72E+09 O 1.37E+07 7.63E+06 C 6.85E+06 7.56E+05 N 3.13E+06 5.98E+04 Mg 1.07E+06 1.07E+06 Si 1.00E+06 1.00E+06 Fe 9.00E+05 9.00E+05 S 4.46E+05 5.15E+05 Al 8.49E+04 8.49E+04 Ca 6.11E+04 6.11E+04 Na 5.74E+04 5.74E+04 Ni 4.93E+04 4.93E+04 Cr 1.35E+04 1.35E+04 P 1.04E+04 1.04E+04 Mn 9.55E+03 9.55E+03 K 3.77E+03 3.77E+03 Ti 2.40E+03 2.40E+03 Co 2.25E+03 2.25E+03 H/O 2041 1.08 C/O 0.50 0.10

* Values are from Anders and Grevesse (1989), and Allende Prieto et al. (2001; 2002). Abundances are normalized to 1E+06 atoms of Si. (a) Value calculated from Tenner et al. (2015) to account for available oxygen

that would exist as H2O ice. See text for details.

Using Eqn. 8.1, redox conditions imposed by a given assemblage are used to estimate the Mg# of a chondrule that would be formed. This is accomplished by combining (1) estimated CI dust enrichments that form chondrules with various Mg#’s (e.g., Figure 8a from Ebel and Grossman, 2000); and (2) oxygen fugacities imposed at each CI dust enrichment factor, using Eqn. 8.1. Results are shown in Figure 8.22b, and the empirical fit to the data is:

chondrule Mg# ¼ 100 expðÞðÞ½þ log f O2 log IW 3:444 =0:6649 (8.2) Ebel and Grossman (2000) predict CI dust-to-gas ratios of at least 12.5 are necessary to stabilize a silicate liquid at 10–3 bar, and ratios of 100 to 1,000 are commonly inferred as conditions that formed most chondrules. Similarly, Schrader et al. (2013) calculated the differences in H2O/H2 and oxygen fugacity during formation of type I and type II CR chondrite chondrules to be ~10–100 times Solar (IW –4.0 to –2.3) and ~230–740 times Solar (IW –1.6 to –0.8), respectively. Therefore, the bulk O-isotope composition of dust in chondrule precur- sors most likely dictated host chondrule values. Oxygen Isotope Characteristics of Chondrules 231

For chondrites with relatively uniform host Δ17O values per chondrule (e.g., LL3, E, R, G, and K) their chondrule Mg# ranges may simply reflect variable dust-to gas ratios in respective environments. For example, chondrule Mg#’s of 98, 88, and 77 correspond to CI dust-to-gas ratios of 100, 500, and 1,000, respectively (Figure 8.22b). As chondrule Mg#’s indicate dust is by far the dominant component of chondrule precursors, it suggests bulk precursor dusts from LL3, E, R, G, and K chondrites have O-isotope ratios that differ from those of most carbon- aceous chondrite chondrule precursors (Figures 8.11–8.20).

8.5.2.1 Accounting For Chondrule Precursor Dusts With Variable O-Isotopes As mentioned in Section 8.5.1, host chondrule Δ17O values from several carbonaceous chon- drite types change with chondrule Mg# (e.g., Figure 8.21), suggesting that, along with variable dust-to-gas ratios, their precursor dust O-isotope ratios differed. In particular, Tenner et al. (2015) found that among FeO-poor CR chondrite chondrules, there is a monotonic increase in host chondrule Δ17O, from ~–6‰ to ~–1‰, as chondrule Mg#’s decrease from ~99.2 to ~94

(Figure 8.23). This implies an increase in fO2 by approximately one log unit (e.g., Eqn. 8.2) corresponded to an increase in the Δ17O of chondrule precursors by ~5‰ along the PCM line.

From this relationship, it is possible to estimate H2O ice proportions and dust-to-gas-ratios of chondrule precursors that could produce the data trend from Tenner et al. (2015). This is accomplished by mass balance, where chondrule precursors are split into four components:

(1) Solar gas; (2) silicate in the dust; (3) organic material in the dust; and (4) H2O ice in the dust: ÀÁÀÁ Δ17 ¼ : Δ17 þ : Δ17 chondrule O ÀÁfrac OSolar gas OSolar gas frac OsilicateÀÁ in dust Osilicate in dust 17 17 þ frac:Oorganics in dust Δ Oorganics in dust þ frac:OH2O in dust Δ OH2O in dust (8.3) The mass balance uses compositions of Solar gas and CI dust given in Table 8.3. Note that for any given dust-to-gas ratio, as well as H2O enhancement, the Mg# of a chondrule is calculated from the atomic H/O and C/O of the assemblage (e.g., Eqns. 8.1 and 8.2). As part of the mass balance, component proportions of oxygen in the dust must be defined. Using atomic abundances of Mg, Si, Al, Ca, Na, Cr, P, N, K, and Ti in CI dust (e.g., Table 8.3) and assigning oxygen to each as the following: MgO, SiO2,Al2O3, CaO, Na2O, Cr2O3,P2O5, MnO, K2O, and TiO2, 44 percent of oxygen is constrained to silicate in the dust. It is assumed Fe, S, Ni, and Co existed as metal and sulfide at ambient conditions. For oxygen associated with organic material in the dust, the atomic abundance of carbon in CI dust is combined with the inorganic matter C/O measured in Orgueil and Murchison (i.e., 5; Binet et al., 2002; Remusat et al., 2007), and this corresponds to 2 percent of oxygen in the dust. The remaining 54 percent of oxygen is assigned to H2O ice in CI dust. In the mass balance, the proportion of H2O ice is allowed to deviate relative to that in CI dust, in order to evaluate its influence on chondrule Δ17O (see Table 8.3 from Tenner et al., 2015). O-isotopes of components in the mass balance (Eqn. 8.3) are defined as follows. For Solar gas, the estimated bulk solar system value is employed (Δ17O: –28.4‰; McKeegan et al., 2011). For silicate in the dust, a Δ17Oof–5.9‰ is assigned, which is the value of the most 16O-rich chondrule measured by Tenner et al. (2015), and has an Mg# of 99.1. Thus, it most likely 17 represents the Δ O of chondrule precursors that were the most reduced and H2O-free. For 232 Travis J. Tenner, et al.

CR2.6-2.8 chondrite host chondrule data f log O2 – log IW –3.5 –3 –2.5 –2 –1

2 2500 dust-to-gas ratio 1000 200 300 500 100 CI dust 0 50 75% ice

50% ice -2 O (‰)

17 25% ice Δ

-4

anhydrous CI dust -6

100 99 98 97 96 95 94 93 90 80 70 60 50 40 chondrule Mg#

Figure 8.23 CR2.6–2.8 chondrite chondrule Mg# versus host Δ17O data from Tenner et al. (2015), and accompanying mass balance model. Δ17O uncertainties are 2SD, and Mg# uncertainties are the range of measured values within a chondrule. The composition of dust with a specified abundance of H2O ice (relative to the atomic abundance in CI dust; right side of the panel) is used to determine the oxygen fugacity imposed by an assemblage, and the corresponding chondrule Mg#, for a given dust-to-gas ratio (see Table 8.3, page 230; Eqn. 8.1, page 229; Eqn. 8.2, page 230). Fractions of oxygen from each component in the mass balance (e.g., Eqn. 8.3) are also determined for a given assemblage, and when combined with the Δ17O values assigned to each source the bulk Δ17O of a chondrule is determined. In the 17 model, Δ O values of Solar gas, silicate in the dust, organic matter in the dust, and H2O in the dust are –28.4‰, –5.9‰, +11.3‰, and +5.1‰, respectively (see Eqn. 8.3, page 231). organic material in the dust, a Δ17O value of +11.3‰ is used, which is the average of measurements from CR2 chondrite Yamato 793495 (Hashizume et al., 2011). For H2O ice in the dust, the Δ17Oisdefined by the mass balance (Eqn. 8.3), by assuming (1) type II chondrules formed at high dust-to-gas ratios, where the O-isotope contribution from solar gas was negli- gible; (2) type II CR chondrite chondrule precursor dust was CI-like in composition; and (3) the bulk Δ17O of such dust is the same as those of host type II chondrules from Tenner et al. (2015), or +0.4‰ (Figure 8.23). With these constraints, as well as the aforementioned fractions of oxygen and Δ17O values of other components in the mass balance, the predicted Δ17Oof

H2O ice in the dust is +5.1‰. This value is intermediate relative to that of primordial H2O ice inferred from cosmic symplectites (Δ17O: ~+80‰; Sakamoto et al., 2007), and to that of

H2O that participated in parent body alteration of chondritic materials (e.g., magnetite and fayalite Δ17O values of +1‰ to ‒3‰; Choi et al., 1997; Choi et al., 2000; Doyle et al., 2015).

The mass balance assumes complete O-isotope exchange of chondrule melt with H2O ice that Oxygen Isotope Characteristics of Chondrules 233 evaporated into ambient gas during chondrule heating. This assumption generally agrees with exchange experiments (e.g., Yu et al., 1995; Di Rocco and Pack, 2015), where high-temperature O-isotope exchange achieves completion/near-completion in tens of minutes to several hours, depending on fO2. If desired, the mass balance could evaluate scenarios of inefficient exchange, 17 where qualitatively the Δ OofH2O ice would increase as a function of decreasing exchange efficiency, in order to explain the Δ17O versus Mg# relationship among CR chondrite chon- drules (see Figure 10b from Schrader et al. (2013) and related discussion). Results of the mass balance are shown in Figure 8.23. Thick semi-horizontal lines represent enrichment of solar gas with dust containing a specified percentage of H2O ice, relative to the atomic abundance within CI dust. Thin semi-vertical lines show various dust to gas ratios, and 17 exhibit curvature with increasing Δ O due to an increasing proportion of H2O ice within dust, making the assemblage more oxidizing. The mass balance predicts most FeO-poor CR chon- drite chondrules formed at dust-to-gas ratios between 100 and 200, and that H2O ice abundances within chondrule precursor dusts range from anhydrous to ~75 percent of the atomic abundance within CI dust. These parameters are within constraints predicted by dynamic models of the protoplanetary disk (e.g., Cassen, 2001; Cuzzi et al., 2001; Ciesla and Cuzzi, 2006). Regarding type II chondrules from Tenner et al. (2015), the mass balance predicts formation at significantly higher dust enrichments (~2,500), from dusts with CI chondritic H2O abundances. Such high dust enrichments are difficult to achieve by dynamic models of the protoplanetary disk, and may in part explain the low proportion of type II chondrules in CR and other carbonaceous chondrites (<1 percent–25 percent; Krot et al., 2002; Kunihiro et al., 2005; Schrader et al., 2011; 2015; Kimura et al., 2014; Scott and Krot, 2014). It is worth mentioning that Acfer 094, CO, CV, CR, and Y-82094 chondrites have a large contingent of chondrules with Mg#’s greater than 98, and with Δ17O values between 4‰ and 6‰ (Figure 8.21). This suggests a common O-isotope environment was sampled by multiple carbonaceous chondrites, which was relatively low in its dust-to-gas ratio, highly reducing and largely devoid of H2O (e.g., Figure 8.23). This may have been the dominant chondrule-forming environment when considering the high proportion of type I chondrules among carbonaceous chondrites (70–99 percent; Scott and Taylor, 1983; Grossman et al., 1988; Weisberg et al., 1993; Newton et al., 1995; Kunihiro et al., 2005; Kimura et al., 2014). Admittedly, this conclusion poses a challenge if carbonaceous chondrites formed beyond Jupiter and past the snow line. Evidence for this arrangement comes from plots of ε54Cr versus ε50Ti and ε54Cr versus Δ17O, as they show a bimodality among carbonaceous and non-carbonaceous meteorites (e.g., Warren, 2011; Gerber et al., 2017). Further, the Grand-Tack hydrodynamic model of Walsh et al. (2011) predicts that S-type (stony) planetesimals were sequestered to the inner solar system, while C-type (carbonaceous) planetesimals were constrained to the outer solar system, due to inward-outward migration of Jupiter. If true, it would be difficult to understand how the majority of carbonaceous chondrite chondrules formed while avoiding H2O. One possibility is that Δ17O: 4‰ to 6‰ chondrule precursors contained an appreciable amount of 16O-poor 16 H2O ice, but that the silicate component was more O-rich than that modeled by Tenner et al. (2015), perhaps due to the incorporation of refractory material. For instance, by changing the Δ17O of the silicate component in the dust from –5.9‰ to –10‰ in the Tenner et al. (2015) mass balance model, it predicts that Mg#: 99, Δ17O: 4‰ to 6‰ chondrules originated from dusts with 30 percent–60 percent of the atomic CI chondritic abundance of H2O, at dust to gas 234 Travis J. Tenner, et al. ratios between 50 and 100; similar dust to gas ratios (10 to 100) are predicted by Schrader et al. (2013) with respect to type I chondrule formation. Potentially complicating this scenario, however, are examples of ordinary chondrites containing chondrules with carbonaceous chondrite-like characteristics (e.g., Kita et al., 2010) as well as carbonaceous chondrites with ordinary chondrite-like chondrules (e.g., Tenner et al., 2017). This suggests OC and CC chondrule precursors were not completely isolated from each other in the protoplanetary disk. Relative to Mg# > 98 chondrules, Acfer 094, CO, CR, CV, and Y-82094 chondrules with Mg#’s below 98 generally have higher Δ17O, with values between ~3‰ and ~+1‰ (Figure 8.21). The Mg# range of such chondrules with O-isotope data is large, extending to below 40. Collectively this indicates FeO-rich chondrules from these chondrites derived from 16 precursors with increased proportions of O-poor H2O ice, as well as much higher, and highly variable dust-to-gas ratios, relative to the FeO-poor chondrule-forming environment. It is conceivable such precursors consisted of dust that produced FeO-poor chondrules from the 17 16 Δ O “5‰” O-isotope reservoir, plus additional O-poor H2O ice (e.g., Figure 8.23). This environment was likely minor, because the relative proportion of FeO-rich chondrules in carbonaceous chondrites is low. The inference that several carbonaceous chondrites sampled precursors from a common reduced and high Mg# Δ17O 5‰ reservoir is in general agreement with the hypothesis from Libourel and Chaussidon (2011), in that distinct modes of O-isotope ratios are found among chondrules. Their explanation for this observation is that chondrule precursors derive from broken apart planetesimals and that each isotope mode represents a distinct planetesimal. While possible, the variability of dust-to-gas ratios, along with different abundances of 16O-poor

H2O ice within precursors (e.g., Figs. 8.21 and 8.23), can also explain O-isotope variability among chondrules.

8.5.3 Insights Regarding Chondrule-Forming Mechanisms

When evaluating how chondrules formed, many constraints must be satisfied, including peak temperatures and cooling rates, multiple heating events, volatile retention, and the time period over which chondrules formed. The most popular chondrule-forming mechanisms involve shock heating (e.g., Desch et al., 2012; Morris et al., 2016; Chapter 15) and planetesimal impacts (e.g., Krot et al. 2005; Asphaug et al., 2011; Sanders and Scott, 2012; Chapter 14; Johnson et al., 2015; Chapter 13). Shock heating explains several chondrule characteristics, including heating and cooling histories, chondrule-matrix complementarity, and chondrule formation time intervals (e.g., Desch et al., 2012; Chapter 15). Additionally, if shocks passed through disk regions with variable dust, gas, and ice abundances, it would readily explain

O-isotope and fO2 variabilities observed among chondrules. Regarding impacts, Johnson et al. (2015) show that impactors with diameters of 5,000 km down to 100 km yield realistic cooling – – rates (10 K h 1–1,000 K h 1, respectively). In particular, large impactors explain FeO-poor chondrule textures (cooling rates <25 K h–1; Wick and Jones, 2012), while smaller impactors explain FeO-rich chondrule textures (cooling rates of 1,000 K h–1; Lofgren and Le, 2000). Additionally, Johnson et al. (2015) predict higher mass (larger diameter) impactors were located closer to the Sun than lower mass (smaller diameter) impactors, producing chondrules Oxygen Isotope Characteristics of Chondrules 235

10,000 to 4 million years after CAIs. As such, this scenario is consistent with FeO-poor chondrule formation in regions deficient in H2O ice, and FeO-rich chondrule formation in an ice-enhanced environment. However, it is currently unclear if impact models can explain the observed fO2 and O-isotope diversity recorded in chondrules. In addition, the common presence of relict grains in chondrules is inconsistent with formation by splashing of melt during molten planetesimal collisions. As summarized in Krot and Nagashima (2017), chondrules likely formed by both shock heating (e.g., porphyritic) and by impacts (e.g., non-porphyritic chon- drules in metal-rich chondrites).

8.6 Conclusions

Investigations of chondrule O-isotope characteristics by SIMS reveal the physicochemical constraints in which chondrules were processed. Interphase comparisons allow for evaluating primary chondrule O-isotope signatures, as well as the extent of disturbance by parent body thermal metamorphism. Chondrules from the petrologic type 3.00 chondrite Acfer 094 exhibit O-isotope homogeneity among coexisting phases pyroxene, olivine, plagioclase, and glass. This indicates primary O-isotope signatures of chondrules remained constant as minerals crystallized from the host chondrule melt. As chondrules likely formed in an open system, based on textures indicative of gas–melt interactions (e.g., olivine-rich cores surrounded by pyroxene-rich peripheries, particularly among type I chondrules), this suggests (1) the O-isotope ratios of a given chondrule melt and its surrounding gas were similar; and that (2) gas–melt O-isotope exchange effectively erased evidence of kinetic mass-dependent fractionation. Many chondrules from various chondrites contain relict olivine and spinel grains with different O-isotope ratios relative to host chondrule values. This attests to the high melting temperatures of these minerals, as well as their slow oxygen diffusion rates. Most relict grains have signatures within the range of host chondrule O-isotope ratios for a given chondrite, suggesting common origins and localized transport. However, some relict grains are significantly 16O-rich, suggest- ing refractory (e.g., CAI or AOA) origins. This inference is consistent with formation of refractory inclusions before chondrules. Overall, the presence of relict grains in chondrules (which survived despite fast dissolution rates in silicate melts), as well as a small percentage of chondrules with heterogeneous O-isotope ratios (interpreted as representing sintering of precur- sors with variable O-isotopes), indicates many chondrules experienced incomplete melting. This simple observation may preclude models in which chondrules can only form by complete melting. Per chondrule, approximately 90 percent of pyroxene-bearing chondrules have homo- geneous pyroxene O-isotope data. This indicates pyroxene is an accurate recorder of host chondrule values. In addition, the majority of measured chondrule olivines (62.5–100 percent, depending on chondrite type) have O-isotopes matching those of coexisting pyroxene, or are otherwise homogeneous when pyroxene comparisons are not available. This suggests the majority of chondrule olivines crystallized during the final chondrule-forming event, represent- ing host chondrule O-isotope ratios. O-isotope measurements of chondrule plagioclase and glass reveal each phase is sensitive to O-isotope disturbance. While plagioclase is found to have O-isotope ratios that match coexisting pyroxene and olivine (per chondrule) in some chondrites (e.g., Acfer 094 ungr. C3.00; CR2.6–2.8; Y-82094 ungr. C3.2), it has higher δ18O when compared to coexisting pyroxene 236 Travis J. Tenner, et al. and olivine in chondrites that experienced higher extents of thermal metamorphism (e.g., Allende CV3.6; Efremovka CV3.1–3.4; L & H >3.6; Dar al Gani 978 ungr. C>3.5). Disturbed plagioclase O-isotope signatures are often linked to those of other alteration phases, such as magnetite or nepheline, which have O-isotope ratios inferred to represent the fluid they equilibrated with during thermal metamorphism. Similar disturbed O-isotope characteristics are found in chondrule glass, extending to even less thermally metamorphosed chondrites, such as Semarkona (LL3.01). Ranges of host chondrule O-isotope ratios within a chondrite further reveal characteristics about the chondrule-forming environment. For example, mass-dependent O-isotope fraction- ation is related to bulk chondrule Mg/Si ratios in LL3 chondrites, indicating evaporation and condensation processes within the chondrule-forming environment. O-isotope ratios of enstatite (E), Rumuruti (R), and Grosvenor 95551-Northwest Africa 5492-like (G) chondrite chondrules overlap with each other and with those of LL3 chondrite chondrules, suggesting each chondrite accretion region sampled the same O-isotope reservoir. In Kakangari chondrites, the similarity in O-isotope modes among its chondrules and matrix suggest they are genetically related.

Among CH, CH/CBb, and CBb chondrites, O-isotope ratios of most cryptocrystalline chon- drules are tightly clustered, supporting formation by asteroidal impact. O-isotope ratios of some chondrules also suggest they were transported from one chondrite accretion region to another. For example, among enstatite chondrites, while most chondrules follow the terrestrial fraction- ation line, others have O-isotopes consistent with R-chondrites, ordinary chondrites, and carbonaceous chondrites. Carbonaceous chondrite chondrules plot along the mass-independent fractionated primitive chondrule mineral (PCM) line, established from Acfer 094 chondrule O-isotope data. In addition, most carbonaceous chondrites exhibit a relationship in which decreasing chondrule Mg#’s correspond to heavy O-isotope enrichment (i.e., Δ17O increases). Increased proportions of heavy-isotope enriched H2O ice in chondrule precursors, along with increasing dust-to-gas ratios, could have produced such a trend.

Acknowledgments

The authors thank reviewers Yves Marrocchi, Kazuhide Nagashima, and Associate Editor Alexander Krot for constructive comments that improved the content of this chapter. Through Los Alamos National Laboratory, this document is approved for unlimited release under LA-UR-17–27650.

References

Alexander, C. M. O’D. (2004). Chemical equilibrium and kinetic constraints for chondrule and CAI formation conditions. Geochim. Cosmochim. Acta 68, 3942–3969. Alexander, C. M. O’D., Barber, D. J., and Hutchison, R. (1989). The microstructure of Semarkona and Bishunpur. Geochim. Cosmochim. Acta 53, 3045–3057. Allende Prieto, C., Lambert, D. L., and Asplund, M. (2001). The forbidden abundance of oxygen in the Sun. Astrophys. J. 556, L63–L66. Allende Prieto, C., Lambert, D. L., and Asplund, M. (2002). A reappraisal of the Solar photospheric C/O ratio. Astrophys. J. 573, L137–L140. Oxygen Isotope Characteristics of Chondrules 237

Anders, E., and Grevesse, N. (1989). Abundances of the elements: Meteoritic and solar. Geochim. Cosmochim. Acta 53, 197–214. Ando, M., Nagata, T., Sato, S., et al. (2002). Near-infrared and CO (J =1–0) observations of photodissociation regions in M17. Astrophys. J. 574, 187–197. Asphaug, E., Jutzi, M., and Movshovitz, N. (2011). Chondrule formation during planetesimal accretion. Earth Planet. Sci. Lett. 308, 369–379. Baertschi, P. (1976). Absolute 18O content of standard mean ocean water. Earth Planet. Sci. Lett. 31, 341–344. Behrens, H., Zhang, Y. X., Leschik, et al. (2007). Molecular H2O as carrier for oxygen diffusion in hydrous silicate melts. Earth Planet. Sci. Lett. 254,69–76. Behrens, H., Zhang, Y. X., and Xu, Z. G. (2004). H2O diffusion in dacitic and andesitic melts. Geochim. Cosmochim. Acta 68, 5139–5160. Berlin, J. (2009). Mineralogy and bulk chemistry of chondrules and matrix in petrologic type 3 chondrites: Implications for early solar system processes. Ph.D. thesis, The University of New Mexico. Binet, L., Gourier, D., Derenne, S., and Robert, F. (2002). Heterogeneous distribution of paramagnetic radicals in insoluble organic matter from the Orgueil and Murchison meteor- ites. Geochim. Cosmochim. Acta 66, 4177–4186. Bischoff, A., Geiger, T., Palme, et al. (1994). Acfer 217-A new member of the Rumuruti chondrite group (R). Meteoritics 29, 264–274. Bischoff, A., and Keil, K. (1984). Al-rich objects in ordinary chondrites: Related origin of carbonaceous and ordinary chondrites and their constituents. Geochim. Cosmochim. Acta 48, 693–709. Bischoff, A., Vogel, N., and Roszjar, J. (2011). The Rumuruti chondrite group. Chemie der Erde 71, 101–133. Bodénan, J. -D., Starkey, N. A., Russell, S. S., Wright, I. P., and Franchi, I. A. (2014). An oxygen isotope study of Wark-Lovering rims on type A CAIs in primitive carbonaceous chondrites. Earth Planet. Sci. Lett. 401, 327–336. Bonal, L., Bourot-Denise, M., Quirico, E., Montagnac, G., and Lewin, E. (2007). Organic matter and metamorphic history of CO chondrites. Geochim. Cosmochim. Acta 71, 1605–1623. Brearley, A. J. (1997). Disordered biopyriboles, amphibole, and talc in the Allende meteorite: Products of nebular or parent body aqueous alteration? Science 276, 1103–1105. Bridges, J. C., Franchi, I. A., Sexton, A. S., and Pillinger, C. T. (1999). Mineralogical controls on the oxygen isotopic compositions of UOCs. Geochim. Cosmochim. Acta 63, 945–951. Bullock, E. S., MacPherson, G. J., Nagashima, K., et al. (2012). Forsterite-bearing type B refractory inclusions from CV3 chondrites: From aggregates to volatilized melt droplets. Meteorit. Planet. Sci. 47, 2128–2147. Cassen, P. (2001). Nebular thermal evolution and the properties of primitive planetary materials. Meteorit. Planet. Sci. 36, 671–700. Chakraborty, S. (2010). Diffusion coefficients in olivine, wadsleyite, and ringwoodite. In Y. Zhang and D. J. Cherniak (Eds.), Reviews in Mineralogy and Geochemistry 72, 603–639. Washington D.C.: Mineralogical Society of America. Chaussidon, M., Libourel, G., and Krot, A. N. (2008). Oxygen isotopic constraints on the origin of magnesian chondrules and on the gaseous reservoirs in the early Solar System. Geo- chim. Cosmochim. Acta 72, 1924–1938. Choi, B. -G., Krot, A. N., and Wasson, J. T. (2000). Oxygen isotopes in magnetite and fayalite in CV chondrites Kaba and Mokoia. Meteorit. Planet. Sci. 35, 1239–1248. Choi, B. -G., McKeegan, K. D., Krot, A. N., and Wasson, J. T. (1998). Extreme oxygen- isotope compositions in magnetite from unequilibrated ordinary chondrites. Nature 392, 577–579. 238 Travis J. Tenner, et al.

Choi, B. -G., McKeegan, K. D., Leshin, L. A., and Wasson, J. T. (1997). Origin of magnetite in oxidized CV chondrites: In situ measurement of oxygen isotope compositions of Allende magnetite and olivine. Earth Planet. Sci. Lett. 146, 337–349. Ciesla, F. J., and Cuzzi, J. N. (2006). The evolution of the water distribution in a viscous protoplanetary disk. Icarus 181, 178–204. Clayton, R. N. (1993). Oxygen isotopes in meteorites. Annu. Rev. Earth Planet. Sci. 21,115–149. Clayton, R. N. (2003). Oxygen isotopes in meteorites. In A.M. Davis (Ed.), Meteorites, Comets, and Planets. In H. D. Holland and K. K. Turekian (Eds.), Treatise on Geochemistry (First Edition) 1, 129–142. Oxford, UK: Elsevier-Pergamon. Clayton, R. N., Grossman, L., and Mayeda, T. K. (1973). A component of primitive nuclear composition in carbonaceous meteorites. Science 182, 485–488. Clayton, R. N., and Kieffer, S. W. (1991). Oxygen isotopic thermometer calibrations. In H. P. Taylor Jr., J. R. O’Neil, and I. R. Kaplan (Eds.), Stable Isotope Geochemistry: A Tribute to Samuel Epstein,3–10. Special Publication, 3. San Antonio, TX: Geochemical Society. Clayton, R. N., and Mayeda, T. K. (1983). Oxygen isotopes in eucrites, shergottites, nakhlites, and chassignites. Earth Planet. Sci. Lett. 62,1–6. Clayton, R. N., and Mayeda, T. K. (1984). Oxygen isotopic compositions of enstatite chondrites and aubrites. J. Geophys. Res. 89, C245–C249. Clayton, R. N., Mayeda, T. K., Goswami, J. N., and Olsen, E. J. (1991). Oxygen isotope studies of ordinary chondrites. Geochim. Cosmochim. Acta 55, 2317–2337. Clayton, R. N., Onuma, N., Grossman, L., and Mayeda T. K. (1977). Distribution of the pre- solar component in Allende and other carbonaceous chondrites. Earth Planet. Sci. Lett. 34, 209–224. Connolly, H. C. Jr., and Huss, G. R. (2010). Compositional evolution of the protoplanetary disk: Oxygen isotopes of type II chondrules from CR2 chondrites. Geochim. Cosmochim. Acta 74, 2473–2483. Cuzzi, J. N., Hogan, R. C., Paque, J. M., and Dobrovolskis, A. R. (2001). Size-selective concentration of chondrules and other small particles in protoplanetary nebula turbulence. Astrophys. J. 546, 496–508. Davidson, J., Krot, A. N., Nagashima, K., Hellebrand, E., and Lauretta, D. S. (2014). Oxygen isotope and chemical compositions of magnetite and olivine in the anomalous CK3 Watson 002 and ungrouped Asuka-881595 carbonaceous chondrites: effects of parent body ther- mal metamorphism. Meteorit. Planet. Sci . 49, 1456–1474. Desch,S.J.,Morris,M.A.,Connolly,H.C.Jr.,andBoss,A.P.(2012).Theimportanceof experiments: Constraints on chondrule formation models. Meteorit. Planet. Sci. 47,1139–1156. Di Rocco, T., and Pack, A. (2015). Triple oxygen isotope exchange between chondrule melt and water vapor: an experimental study. Geochim. Cosmochim. Acta 55, 2317–2337. Dohmen, R., Becker, H. -W., and Chakraborty, S. (2007). Fe-Mg diffusion in olivine I: Experimental determination between 700 and 1200 C as a function of composition, crystal orientation and oxygen fugacity. Phys. Chem. Minerals 34, 389–407. Dominguez, G. (2010). A heterogeneous chemical origin for the 16O-enriched and 16O-depleted reservoirs of the early solar system. Astrophys. J. Lett. 713, L59–L63. Doyle, P. M., Jogo, K., Nagashima, K., et al. (2015). Early aqueous activity on the ordinary and carbonaceous chondrite parent bodies recorded by fayalite. Nat. Commun. 6,1–10. Ebel, D. S., and Grossman, L. (2000). Condensation in dust-enriched systems. Geochim. Cosmochim. Acta 64, 339–366. Eiler, J. M. (2001). Oxygen isotope variations of basaltic lavas and upper mantle rocks. In J. W. Valley and D. R. Cole (Eds.), Reviews in Mineralogy and Geochemistry 43, 319–364. Washington, D.C.: Mineralogical Society of America. Farver, J. R. (2010). Oxygen and hydrogen diffusion in minerals. In Y. Zhang and D. J. Cherniak (Eds.), Reviews in Mineralogy and Geochemistry 72, 447–507. Washington D.C.: Mineralogical Society of America. Oxygen Isotope Characteristics of Chondrules 239

Fedkin, A. V., and Grossman, L. (2006). The fayalite content of chondritic olivine: Obstacle to understanding the condensation of rocky material. In D. S. Lauretta and H. Y. McSween (Eds.), Meteorites and the Early Solar System II, 279–294. Tucson, AZ: University of Arizona Press. Fedkin, A. V., and Grossman, L. (2016). Effects of dust enrichment on oxygen fugacity of cosmic gases. Meteorit. Planet. Sci. 51, 843–850. Fegley, B. (2000). Kinetics of gas-grain reactions in the solar nebula. Space Sci. Rev. 92, 177–200. Franchi, I. A. (2008). Oxygen isotopes in asteroidal materials. In G. J. MacPherson (Ed.), Reviews in Mineralogy and Geochemistry 68, 345–397. Washington D.C.: Mineralogical Society of America. Friend, P., Hezel, D. C., and Mucerschi, D. (2016). The conditions of chondrule formation, Part II: Open system. Geochim. Cosmochim. Acta 173, 198–209. Ganguly, J., and Tazzoli, V. (1994). Fe2+-Mg interdiffusion in orthopyroxene: Retrieval from the data on intracrystalline exchange reaction. Am. Mineral. 79, 930–937. Gao, H., Song, Y., Chang, Y. -C., et al. (2013). Branching ratio measurements for vacuum ultraviolet photodissociation of 12C16O. J. Phys. Chem. 117, 6185–6195. Gerard, O., and Jaoul, O. (1989). Oxygen diffusion in San Carlos olivine. J. Geophys. Res. 94, 4119–4128. Gerber, S., Burkhardt, C., Budde, G., Metzler, K., and Kleine, T. (2017). Mixing and transport of dust in the early solar nebula as inferred from titanium isotope variations among chondrules. Astrophys. J. Lett. 841, L17 (7 pp). Giletti, B. J., Semet, M. P., and Yund, R. A. (1978). Studies in diffusion-III. Oxygen in feldspars: An ion microprobe determination. Geochim. Cosmochim. Acta 42,45–57. Gounelle, M., Young, E. D., Shahar, A., Tonui, E., and Kearsley, A. (2007). Magnesium isotopic constraints on the origin of CBb chondrites. Earth Planet. Sci. Lett. 256, 521–533. Gounelle, M., Krot, A. N., Nagashima, K., and Kearsley, A. (2009). Extreme 16O enrichment in calcium-aluminum-rich inclusions from the Isheyevo (CH/CB) chondrite. Astrophys. J. 698, L18–L22. Greenwood, J. P., Rubin, A. E., and Wasson, J. T. (2000). Oxygen isotopes in R-chondrite magnetite and olivine: Links between R chondrites and ordinary chondrites. Geochim. Cosmochim. Acta 64, 3897–3911. Greenwood, R. C., Burbine, T. H., Miller, M. F., and Franchi, I. A. (2017). Melting and differentiation of early-formed asteroids: The perspective from high-precision oxygen isotope studies. Chemie der Erde 77,1–43. Grossman, J. N., Alexander, C. M. O’D., Wang, J., and Brearley, A. J. (2002). Zoned chondrules in Semarkona: Evidence for high- and low-temperature processing. Meteorit. Planet. Sci. 37,49–73. Grossman, J. N., and Brearley, A. J. (2005). The onset of metamorphism in ordinary and carbonaceous chondrites. Meteorit. Planet. Sci. 40,87–122. Grossman, J. N., Rubin, A. E., Nagahara, H., and King, E. A. (1988). Properties of chondrules. In J. Kerridge and M. Shapley Matthews (Eds.), Meteorites and the Early Solar System, 619–659. Tucson, AZ: University of Arizona Press. Grossman, L., Beckett, J. R., Fedkin, A. V., Simon, S. B., and Ciesla, F. J. (2008). Redox conditions in the solar nebula: Observational, experimental, and theoretical constraints. In G. J. MacPherson (Ed.), Reviews in Mineralogy and Geochemistry 68, 93–140. Washing- ton D.C.: Mineralogical Society of America. Guan, Y., Huss, G. R., Leshin, L. A., MacPherson, G. J., and McKeegan, K. D. (2006). Oxygen isotope and 26Al-26Mg systematics of aluminum-rich chondrules from unequilibrated enstatite chondrites. Meteorit. Planet. Sci. 41,33–47. Harju, E. R., Rubin, A. E., Ahn, et al. (2014). Progressive alteration of CR carbonaceous chondrites. Geochim. Cosmochim. Acta 139, 267–292. 240 Travis J. Tenner, et al.

Hashizume, K., Takahata, N., Naraoka, H., and Sano, Y. (2011). Extreme oxygen isotope anomaly with a solar origin detected in meteoritic organics. Nature Geosci. 4, 165–168. Hewins, R. H., Connolly, H. C. Jr., Lofgren, G. E., and Libourel, G. (2005). Experimental constraints on chondrule formation. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrules and the Protoplanetary Disk, 286–316. ASP Conference Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Hewins, R. H., and Radomsky, P. M. (1990). Temperature conditions for chondrule formation. Meteoritics 25, 309–318. Howard, K. T., Alexander, C. M. O’D., Schrader, D. L., and Dyl, K. A. (2015). Classification of hydrous meteorites (CR, CM and C2 ungrouped) by phyllosilicate fraction: PSD-XRD modal mineralogy and planetesimal environments. Geochim. Cosmochim. Acta 149, 206–222. Huss, G. R., and Lewis, R. S. (1994). Noble gases in presolar diamonds II: Component abundances reflect thermal processing. Meteoritics 29, 811–829. Huss, G. R., Rubin, A. E., and Grossman, J. N. (2006). Thermal metamorphism in chondrites. In D. S. Lauretta and H. Y. McSween (Eds.), Meteorites and the Early Solar System II, 567–586. Tucson, AZ: University of Arizona Press. Ichimura, S., Seto, Y., and Tomeoka, K. (2017). Nepheline formation in chondrite parent bodies: Verification through experiments. Geochim. et Cosmochim. Acta 210, 114–131. Isa, J., Rubin, A. E., Marin-Carbonne, J., McKeegan, K. D., and Wasson, J. T. (2011). Oxygen- isotopic compositions of R-chondrite chondrules. LPSC XLII, #2623. Isa, J., Rubin, A. E., and Wasson, J. T. (2014). R-chondrite bulk-chemical compositions and diverse oxides: Implications for parent-body processes. Geochim. Cosmochim. Acta 124, 131–151. Javoy, M., Balan, E., Méhut, M., Blanchard, M., and Lazzeri, M. (2012). First-principles investigation of equilibrium isotopic fractionation of O- and Si-isotopes between refractory solids and gases in the solar nebula. Earth Planet. Sci. Lett. 319–320, 118–127. Jiang, Y., Hsu, W., Guan, Y., and Wang, Y. (2015). In situ SIMS oxygen isotope analyses: Evidence for the formation of aluminum-rich chondrules from ordinary chondrites. Science China 58, 1748–1757. Johnson, B. C., Minton, D. A., Melosh, H. J., and Zuber, M. T. (2015). Impact jetting as the origin of chondrules. Nature 517, 339–341. Jones, R. H. (1992). On the relationship between isolated and chondrule olivine grains in the carbonaceous chondrite ALHA77307. Geochim. Cosmochim. Acta 56, 467–482. Jones, R. H., Leshin, L. A., Guan, et al. (2004). Oxygen isotope heterogeneity in chondrules from the Mokoia CV3 carbonaceous chondrite. Geochim. Cosmochim. Acta 68, 3423–3438. Jones, R. H., Saxton, J. M., Lyon, I. C., and Turner, G. (2000). Oxygen isotopes in chondrule olivine and isolated olivine grains from the CO3 chondrite Allan Hills A77307. Meteorit. Planet. Sci. 35, 849–857. Kallemeyn, G. W., Rubin, A. E., and Wasson, J. T. (1996). The compositional classification of chondrites: VII. The R chondrite group. Geochim. Cosmochim. Acta 60, 2243–2256. Kimura, M., Barrat, J. A., Weisberg, et al. (2014). Petrology and bulk chemistry of Yamato- 82094, a new type of carbonaceous chondrite. Meteorit. Planet. Sci. 49, 346–357. Kimura, M., Grossman, J. N., and Weisberg, M. K. (2008). Fe-Ni metal in primitive chondrites: Indicators of classification and metamorphic conditions for ordinary and CO chondrites. Meteorit. Planet. Sci. 43, 1161–1177. Kimura, M., and Ikeda, Y. (1995). Anhydrous alteration of Allende chondrules in the solar nebula II: Alkali-Ca exchange reactions and formation of nepheline, sodalite and Ca-rich phases in chondrules. Proc. NIPR Symp. Antarctic Meteorites 8, 123–138. Kimura, M., and Ikeda, Y. (1997). Comparative study of anhydrous alteration of chondrules in reduced and oxidized CV chondrites. Antarctic Meteorite Res. 10, 191–202. Oxygen Isotope Characteristics of Chondrules 241

Kimura, M., Nakajima, H., Hiyagon, H., and Weisberg, M. K. (2006). Spinel group minerals in LL3.00–6 chondrites: Indicators of nebular and parent body processes. Geochim. Cosmo- chim. Acta 70, 5634–5650. Kita, N. T., and Ushikubo, T. (2012). Evolution of protoplanetary disk inferred from 26Al chronology of individual chondrules. Meteorit. Planet. Sci. 47, 1108–1119. Kita, N. T., Nagahara, H., Tachibana, et al. (2010). High precision SIMS oxygen three isotope study of chondrules in LL3 chondrites: Role of ambient gas during chondrule formation. Geochim. Cosmochim. Acta 74, 6610–6635. Kita, N. T., Tenner, T. J., Defouilloy, et al. (2015). Oxygen isotope systematics of chondrules in R3 clasts: A genetic link to ordinary chondrites. LPSC XLVI, #2053. Kita, N. T., Tenner, T. J., Ushikubo, et al. (2016). Internal homogeneity of oxygen isotope ratios in chondrules. LPSC XLVII, #2375. Kobayashi, S., Imai, H., and Yurimoto, H. (2003). New extreme 16O-rich reservoir in the early solar system. Geochem. J. 37, 663–669. Kööp, L., Davis, A. M., Nakashima, et al. (2016). A link between oxygen, calcium and titanium isotopes in 26Al-poor hibonite-rich CAIs from Murchison and implications for the hetero- geneity of dust reservoirs in the solar nebula. Geochim. Cosmochim. Acta 189,70–95. Kring, D. A. (1988). The petrology of meteoritic chondrules: Evidence for fluctuating condi- tions in the solar nebula. PhD thesis, Harvard University. Krot, A. N., Amelin, Y., Cassen, P., and Meibom, A. (2005). Young chondrules in CB chondrites from a giant impact in the early Solar System. Nature 436, 989–992. Krot, A. N., Hutcheon, I. D., and Keil, K. (2002). Plagioclase-rich chondrules in the reduced CV chondrites: Evidence for complex formation history and genetic links between cal- cium-aluminum-rich inclusions and ferromagnesian chondrules. Meteorit. Planet. Sci. 37, 155–182. Krot, A. N., Libourel, G., and Chaussidon, M. (2006b). Oxygen isotope compositions of chondrules in CR chondrites. Geochim. Cosmochim. Acta 70, 767–779. Krot, A. N., McKeegan, K. D., Huss, et al. (2006c). Aluminum-magnesium and oxygen isotope study of relict Ca-Al-rich inclusions in chondrules. Astrohys. J. 639, 1227–1237. Krot, A. N., and Nagashima, K. (2016). Evidence for oxygen-isotope exchange in chondrules and refractory inclusions during fluid-rock interaction on the CV chondrite parent body. Meteorit. Planet. Sci. 51 (Suppl.), A392 (abstr.). Krot, A. N., and Nagashima, K. (2017). Constraints on mechanisms of chondrule formation from chondrule precursors and chronology of transient heating events in the protoplanetary disk. Geochem. J. 51,45–68. Krot, A. N., Nagashima, K., Ciesla, F. J., et al. (2010a). Oxygen isotopic composition of the Sun and mean osygen isotopic composition of the protosolar silicate dust: Evidence from refractory inclusions. Astrophys. J. 713, 1159–1166. Krot, A. N., Nagashima, K., Huss, G. R., et al. (2007). Relict refractory inclusions in magnesian porphyritic chondrules from the CH and CH/CB carbonaceous chondrites. Meteorit. Planet. Sci. 42 (Suppl.), A90 (abstr.). Krot, A. N., Nagashima, K., van Kooten, E. M. M., and Bizzarro, M. (2017). Calcium- aluminum-rich inclusions recycled during formation of porphyritic chondrules from CH carbonaceous chondrites. Geochim. Cosmochim. Acta 201, 185–223. Krot, A. N., Nagashima, K., Yoshitake, M., and Yurimoto, H. (2010b). Oxygen isotopic compositions of chondrules from the metal-rich chondrites Isheyevo (CH/CBb), MAC 02675 (CBb) and QUE 94627 (CBb). Geochim. Cosmochim. Acta 74, 2190–2211. Krot, A. N., Petaev, M. I., Scott, E. R., et al. (1998). Progressive alteration in CV3 chondrites: More evidence for asteroidal alteration. Meteorit. Planet. Sci. 33, 1065–1085. Krot, A. N., Yurimoto, H., McKeegan, K. D., et al. (2006a). Oxygen isotopic compositions of chondrules: Implication for evolution of oxygen isotopic reservoirs in the early solar nebula. Chemie der Erde 66, 249–276. 242 Travis J. Tenner, et al.

Krot, A. N., Zolensky, M. E., Wasson, J. T., et al. (1997). Carbide-magnetite-bearing type 3 ordinary chondrites. Geochim. Cosmochim. Acta 61, 219–237. Kunihiro, T., Rubin, A. E., McKeegan, K. D., and Wasson, J. T. (2004). Oxygen-isotopic compositions of relict and host grains in chondrules in the Yamato 81020 CO3.0 chondrite. Geochim. Cosmochim. Acta 68, 3599–3606. Kunihiro, T., Rubin, A. E., and Wasson, J. T. (2005). Oxygen-isotopic composition of low-FeO relicts in high FeO-host chondrules in Acfer 094, a type 3.0 carbonaceous chondrite closely related to CM. Geochim. Cosmochim. Acta 69, 3831–3840. Libourel, G., and Chaussidon, M. (2011). Oxygen isotopic constraints on the origin of Mg-rich olivines from chondritic meteorites. Earth Planet. Sci. Lett. 301,9–21. Lodders, K. (2003). Solar system abundances and condensation temperatures of the elements. Astrophys. J. 591, 1220–1247. Lofgren, G. E. and Le, L. (2000). Experimental evidence for a partial melting origin for most porphyritic chondrules. LPSC XXXI, #1809. Lugaro, M., Liffman, K., Ireland, T. R., and Maddison, S. T. (2012). Can galactic chemical evolution explain the oxygen isotopic variations in the solar system? Astrophys, J. 759, 1. Luz, B., and Barkan, E. (2010). Variations of 17O/16O and 18O/16O in meteoritic waters. Geochim. Cosmochim. Acta 74, 6276–6286. Lyons, J. R., and Young, E. D. (2005). CO self-shielding as the origin of oxygen isotope anomalies in the early solar nebula. Nature 435, 317–320. Ma, C., Beckett, J. R., Connolly, H. C. Jr., and Rossman, G. R. (2008). Aluminous spinels in ferromagnesian chondrules from Allende. LPSC XXXIX, #2030. Makide, K., Nagashima, J., Krot, A. N., et al. (2009). Oxygen- and magnesium-isotope compositions of calcium-aluminum-rich inclusions from CR2 carbonaceous chondrites. Geochim. Cosmochim. Acta 73, 5018–5050. Marrocchi, Y., and Chaussidon, M. (2015). A systematic for oxygen isotopic variation in meteoritic chondrules. Earth Planet. Sci. Lett. 430, 308–315. Marrocchi, Y., and Libourel, G. (2013). Sulfur and sulfides in chondrules. Geochim. Cosmo- chim. Acta 119, 117–136. Maruyama, S., Yurimoto, H., and Sueno, S. (1999). Oxygen isotope evidence regarding the formation of spinel-bearing chondrules. Earth Planet. Sci. Lett. 169, 165–171. Matthews, A., Palin, J. M., Epstein, S., and Stolper, E. M. (1994). Experimental study of 18 16 O/ O partitioning between crystalline albite, albitic glass, and CO2 gas. Geochim. Cosmochim. Acta 58, 5255–5266. McKeegan, K. D., Kallio, A. P. A., Heber, V. S., et al. (2011). The oxygen isotopic composition of the Sun inferred from captured solar wind. Science 332, 1528–1532. Merigoux, H. (1968). Etude de la mobilite d l’oxygen dans the feldspaths alcalins. Bull. Soc. Fr Mineral. Crystallogr. 91,51–64. Miller, K. E., Lauretta, D. S., Connolly, H. C. Jr., et al. (2017). Formation of unequilibrated R chondrite chondrules and opaque phases. Geochim. Cosmochim. Acta 209,24–50. Miller, M. F., Franchi, I. A., Sexton, A. S., and Pillinger, C. T. (1999). High precision δ17O isotope measurements of oxygen from silicates and other oxides: methods and applications. Rapid Comm. Mass Spectrom. 13, 1211–1217. Morris, M. A., Weidenschilling, S. J., and Desch, S. J. (2016). The effect of multiple particle sizes on cooling rates produced in large-scale shocks in the solar nebula. Meteorit. Planet. Sci. 51, 870–883. Nagahara, H. (1981). Evidence for secondary origin of chondrules. Nature 292, 135–136. Nagahara, H., Kita, N. T., Ozawa, K., and Morishita, Y. (2008). Condensation of major elements during chondrule formation and its implication to the origin of chondrules. Geochim. Cosmochim. Acta 72, 1442–1465. Nagahara, H., and Ozawa, K. (2012). The role of exchange reactions in oxygen isotope fractionation during CAI and chondrule formation. Meteorit. Planet. Sci. 47, 1209–1228. Oxygen Isotope Characteristics of Chondrules 243

Nagashima, K., Krot, A. N., and Huss, G. R. (2015). Oxygen-isotope compositions of chon- drule phenocrysts and matrix grains in Kakangari K-grouplet chondrite: Implication to a chondrule-matrix genetic relationship. Geochim. Cosmochim. Acta 151,49–67. Nakashima, D., Ushikubo, T., Gowda, R. N., et al. (2011). Ion microprobe analyses of oxygen three-isotope ratios of chondrules from the Sayh al Uhaymir 290 CH chondrite using a multiple-hole disk. Meteorit. Planet. Sci. 46, 857–874. Newton, J., Bischoff, A., Arden, J. W., et al. (1995). Acfer 094, a uniquely primitive carbon- aceous chondrite from the Sahara. Meteoritics 30,47–56. Onuma, N, Clayton, R. N., and Mayeda, T. K. (1972). Oxygen isotope geothermometer. Geochim. Cosmochim. Acta 36, 169–188. Pacaud, L., Ingrin, J., and Jaoul, O. (1999). High-temperature diffusion of oxygen in synthetic diopside measured by nuclear reaction analysis. Miner. Mag. 63, 673–686. Pack, A., Yurimoto, H., and Palme, H. (2004). Petrographic and oxygen-isotopic study of refractory forsterites from R-chondrite Dar al Gani 013 (R3.5–6), unequilibrated ordinary and carbonaceous chondrites. Geochim. Cosmochim. Acta 68, 1135–1157. Prinz, M., Weisberg, M. K., Neru, C. E., et al. (1989). Petrologic and stable isotope study of the Kakangari (K-group) chondrite: Chondrules, matrix, and CAI’s. Lunar Planet. Sci. 20, 870–871 (abstr.). Rambaldi, E. R. (1981). Relict grains in chondrules. Nature 293, 558–561. Remusat, L., Robert, F., and Derenne, S. (2007). The insoluble organic matter in carbonaceous chondrites: Chemical structure, isotopic composition and origin. C.R. Geosci. 339, 895–906. Richet, P., Bottinga, Y., and Javoy, M. (1977). Review of hydrogen, carbon, nitrogen, oxygen, sulfur, and chlorine stable isotope fractionation among gaseous molecules. Annu. Rev. Earth Planet. Sci. 5,65–110. Rosman, J. J. R., and Taylor, P. D. P. (1998). Isotopic compositions of the elements 1997. Pure Appl. Chem. 70, 217–236. Rubin, A. E., Wasson, J. T., Clayton, R. N., and Mayeda, T. K. (1990). Oxygen isotopes in chondrules and coarse-grained chondrule rims from the Allende meteorite. Earth Planet. Sci. Lett. 96, 247–255. Rudraswami, N. G., Ushikubo, T., Nakashima, D, and Kita, N. T. (2011). Oxygen isotope systematics of chondrules in the Allende CV3 chondrite: High precision ion microprobe studies. Geochim. Cosmochim. Acta 75, 7596–7611. Russell, S. S., MacPherson, G. J., Leshin, L. A., and McKeegan, K. D. (2000). 16O enrichments in aluminum-rich chondrules from ordinary chondrites. Earth Planet. Sci. Lett. 184,57–74. Ruzicka, A., Hiyagon, H., Hutson, M., and Floss, C. (2007). Relict olivine, chondrule recycling, and the evolution of nebular reservoirs. Earth Planet. Sci. Lett. 257, 274–289. Ryerson, F. J., Durham, W. B., Cherniak, D. J., and Lanford, W. A. (1989). Oxygen diffusion in olivine: Effect of oxygen fugacity and implications for creep. J. Geophys. Res. 94, 4105–4118. Ryerson, F. J., and McKeegan, K. D. (1994). Determination of oxygen self-diffusion in akermanite, anorthite, diopside, and spinel: Implications for oxygen isotopic anomalies and the thermal histories of Ca-Al rich inclusions. Geochim. Cosmochim. Acta 58, 3713–3734. Sakamoto, N., Seto, Y., Itoh, S., et al. (2007). Remnants of the early solar system water enriched in heavy oxygen isotopes. Science 317, 231–233. Sanders, I. S., and Scott, E. R. D. (2012). The origin of chondrules and chondrites: Debris from low-velocity impacts between molten planetesimals? Meteorit. Planet. Sci. 47, 2170–2192. Schrader, D. L., Connolly, H. C. Jr., Lauretta, D. S., et al. (2013). The formation and alteration of the Renazzo-like carbonaceous chondrites II: Linking O-isotope composition and oxidation state of chondrule olivine. Geochim. Cosmochim. Acta 101, 302–327. 244 Travis J. Tenner, et al.

Schrader, D. L., Connolly, H. C. Jr., Lauretta, D. S., et al. (2015). The formation and alteration of the Renazzo-like carbonaceous chondrites III: Towards understanding the genesis of ferromagnesian chondrules. Meteorit. Planet. Sci. 50,15–50. Schrader, D. L., Davidson, J., Greenwood, R. C., Franchi, I. A., and Gibson, J. M. (2014b). A water-ice rich minor body from the early Solar System: The CR chondrite parent asteroid. Earth Planet. Sci. Lett. 407,48–60. Schrader, D. L., Franchi, I. A., Connolly, H. C. Jr., et al. (2011). The formation and alteration of the Renazzo-like carbonaceous chondrites I: Implications of bulk-oxygen isotopic com- position. Geochim. Cosmochim. Acta 75, 308–325. Schrader, D. L., Nagashima, D., Krot, A. N., Ogliore, R. C., and Hellebrand, E. (2014a). Variations in the O-isotope composition of gas during the formation of chondrules from CR chondrites. Geochim. Cosmochim. Acta 132,50–74. Schrader, D. L., Nagashima, K., Krot, A. N., et al. (2017). Distribution of 26Al in the CR chondrite chondrule-forming region of the protoplanetary disk. Geochim. Cosmochim. Acta. 201, 275–302. Scott, E. R. D., and Krot, A. N. (2014). Chondrites and their components. In A. M. Davis (Ed.), Meteorites, Comets and Planets. In H. D. Holland and K. K. Turekian (Eds.), Treatise on Geochemistry (Second Edition), 1, 143–200. Oxford, UK: Elsevier. Scott, E. R. D., Love, S. G., and Krot, A. N. (1996). Formation of chondrules and chondrites in the protoplanetary nebula. In R. H. Hewins, R. H. Jones, and E. R. D. Scott (Eds.), Chondrules and the Protoplanetary Disk,87–96. Cambridge, UK: Cambridge University Press. Scott, E. R. D., and Taylor, J. G. (1983). Chondrules and other components in C, O, and E chondrites: Similarities in their properties and origins. Proceedings, 14th Lunar and Planetary Science Conference, Part 1. J. Geophys. Res. (Suppl.) 88, B275–B286. Seto, Y., Sakamoto, N., Fujino, K., et al. (2008). Mineralogical characterization of a unique material having heavy oxygen isotope anomaly in matrix of the primitive carbonaceous chondrite Acfer 094. Geochim. Cosmochim. Acta 72, 2723–2734. Sheffer, Y., Lambert, D. L., and Federman, S. R. (2002). Ultraviolet detection of interstellar 12C17O and the CO isotopometric ratios toward X Persei. Astrophys. J. 547, L171–L174. Shi, X., Yin, Q. -Z., Luo, Z., Huang, H., and Ng, C. -Y. (2011). Testing “self-shielding” model with laboratory experiment for the oxygen isotope evolution in the early solar nebula – A progress report. LPSC XLII #2705. Soulié, C., Libourel, G., and Tissandier, L. (2017). Olivine dissolution in molten silicates: An experimental study with application to chondrule formation. Meteorit. Planet. Sci. 52, 225–250. Spicuzza, M. J., Day, J. M. D., Taylor, L. A., and Valley, J. W. (2007). Oxygen isotope constraints on the origin and differentiation of the Moon. Earth Planet. Sci. Lett. 253, 254–265. Tachibana, S., Nagahara, H., Mostefaoui, S., and Kita, N. T. (2003). Correlation between relative ages inferred from 26Al and bulk compositions of ferromagnesian chondrules in least equilibrated ordinary chondrites. Meteorit. Planet. Sci. 38, 939–962. Tenner, T. J., Ushikubo, T., Kurahashi, E., Kita, N. T., and Nagahara, H. (2013). Oxygen isotope systematics of chondrule phenocrysts from the CO3.0 chondrite Yamato 81020: Evidence for two distinct oxygen isotope reservoirs. Geochim. Cosmochim. Acta 102, 226–245. Tenner, T. J., Nakashima, D., Ushikubo, T., Kita, N. T., and Weisberg, M. K. (2015). Oxygen isotope ratios of FeO-poor chondrules in CR3 chondrites: Influence of dust enrichment and H2O during chondrule formation. Geochim. Cosmochim. Acta 148, 228–250. Tenner, T. J., Kimura, M., and Kita, N. T. (2017). Oxygen isotope characteristics of chondrules from the Yamato-82094 ungrouped carbonaceous chondrite: Further evidence for common Oxygen Isotope Characteristics of Chondrules 245

O-isotope environments sampled among carbonaceous chondrites. Meteorit. Planet. Sci. 52, 268–294. Thiemens, M. H., and Heidenreich, J. E. (1983). The mass-independent fractionation of oxygen: A novel isotope effect and its possible cosmochemical implications. Science 219, 1073–1075. Tissandier, L., Libourel, G., and Robert, F. (2002). Gas–melt interactions and their bearing on chondrule formation. Meteorit. Planet. Sci. 37, 1377–1389. Ushikubo, T., Kimura, M., Kita, N. T., and Valley, J. W. (2012). Primordial oxygen isotope reservoirs of the solar nebula recorded in chondrules in Acfer 094 carbonaceous chondrite. Geochim. Cosmochim. Acta 90, 242–264. Ushikubo, T., Tenner, T. J., Hiyagon, H., and Kita, N. T. (2017). A long duration of the 16O-rich reservoirs in the solar nebula, as recorded in fine-grained refractory inclusions from the least metamorphosed carbonaceous chondrites. Geochim. Cosmochim. Acta 201, 103–122. Ustunisik, G., Ebel, D. S., Walker, D., and Boesenberg, J. S. (2014). Experimental investigation of condensation predictions for dust-enriched systems. Geochim. Cosmochim. Acta 142, 27–38. Walsh, K. J., Morbidelli, A., Raymond, S. N., O’Brien, D. P., and Mandell, A. M. (2011). A low mass for Mars from Jupiter’s early gas-driven migration. Nature 475, 206–209. Warren, P. H. (2011). Stable-isotopic anomalies and the accretionary assemblage of the Earth and Mars: A subordinate role for carbonaceous chondrites. Earth Planet. Sci. Lett. 311, 93–100. Weichert, U., Halliday, A. N., Lee, D. -C., et al. (2001). Oxygen isotopes and the Moon- forming giant impact. Science 294, 345–348. Weisberg, M. K., Boesenberg, J. S., Kozhusko, G., et al. (1995). EH3 and EL3 chondrites: A petrologic-oxygen isotopic study. Lunar. Planet. Sci. XXVI, 1481–1482. Weisberg, M. K., Ebel, D. S., Connolly, H. C. Jr., Kita, N. T., and Ushikubo, T. (2011). Petrology and oxygen isotope compositions of chondrules in E3 chondrites. Geochim. Cosmochim. Acta 75, 6556–6569. Weisberg, M. K., Ebel, D. S., Nakashima, D., Kita, N. T., and Humayun, M. (2015). Petrology and geochemistry of chondrules and metal in NWA 5492 and GRO 95551: A new type of metal-rich chondrite. Geochim. Cosmochim. Acta 167, 269–285. Weisberg, M. K., Prinz, M., Clayton, R. N., and Mayeda, T. K. (1993). The CR (Renazzo-type) carbonaceous chondrite group and its implications. Geochim. Cosmochim. Acta 57, 1567–1586. Weisberg, M. K., Prinz, M., Clayton, R. N., et al. (1996). The K (Kakangari) chondrite grouplet. Geochim. Cosmochim. Acta 60, 4253–4263. Weisberg, M. K., Prinz, M., Kojima, H., et al. (1991). The Carlisle Lakes-type chondrites: A new grouplet with high Δ17O and evidence for nebular oxidation. Geochim. Cosmochim. Acta 55, 2657–2669. Wick, M. J., and Jones, R. H. (2012). Formation conditions of plagioclase-bearing type I chondrules in CO chondrites: A study of natural samples and experimental analogs. Geochim. Cosmochim. Acta 98, 140–159. Young, E. D. (2007). Time-dependent oxygen isotopic effects of CO self shielding across the solar protoplanetary disk. Earth Planet. Sci. Lett. 262, 468–483. Young, E. D., Ash, R. D., Galy, A., and Belshaw, N. S. (2002). Mg isotope heterogeneity in the Allende meteorite measured by UV laser ablation-MC-ICPMS and comparisons with O- isotopes. Geochim. Cosmochim. Acta 66, 683–698. Young, E. D., and Russell, S. S. (1998). Oxygen reservoirs in the early solar nebula inferred from an Allende CAI. Science 282, 452–455. Yu, Y., Hewins, R. H., Clayton, R. C., and Mayeda, T. K. (1995). Experimental study of high temperature oxygen isotope exchange during chondrule formation. Geochim. Cosmochim. Acta 59, 2095–2104. 246 Travis J. Tenner, et al.

Yurimoto, H., Krot, A. N., Choi, B. -G., et al. (2008). Oxygen isotopes of chondritic compon- ents. In G. J. MacPherson (Ed.), Reviews in Mineralogy and Geochemistry, 68, 141–186. Washington D.C.: Mineralogical Society of America. Yurimoto, H., and Kuramoto, K. (2004). Molecular cloud origin for the oxygen isotope heterogeneity in the solar system. Science 305, 1763–1766. Zhang, A. -C., Itoh, S., Sakamoto, N., Wang, R. -C., and Yurimoto, H. (2014). Origins of Al- rich chondrules: Clues from a compound Al-rich chondrule in the Dar al Gani 978 carbon- aceous chondrite. Geochim. Cosmochim. Acta 130,78–92. Zhang, A. -C., and Yurimoto, H. (2013). Petrography and mineralogy of the ungrouped type 3 carbonaceous chondrite Dar al Gani 978. Meteorit. Planet. Sci. 48, 1651–1677. 9 26Al–26Mg Systematics of Chondrules

kazuhide nagashima, noriko t. kita, and tu-han luu

Abstract

The 26Al–26Mg systematics of chondrules from ordinary and carbonaceous chondrites and 26 27 26 27 their implications are reviewed. The initial Al/ Al ratios [( Al/ Al)0] based on in situ analyses of chondrules from the least metamorphosed chondrites range from unresolved ‒5 26 =27 Internal from zero to ~1.2 10 and thus no chondrules have Al Al 0 ratios corresponding to the canonical level (~5.2 10‒5) recorded by CAIs. Assuming homogeneous distribution of 26Al in the protoplanetary disk at the canonical level, these observations suggest chondrule formation started ~1.5 million years after CAIs and lasted over a few million years. The 26Al–26Mg systematics of bulk chondrules could have recorded 26 =27 Bulk – Al Al 0 ratios of chondrule precursors and may suggest that Al Mg fractionation recorded by chondrule precursors started contemporaneously with CAIs and lasted over ~1.5 million years. The comparisons of formation ages of different meteorites and their components have been made with 26Al–26Mg, 182Hf–182W, and 206Pb–207Pb systematics. While the ages determined by 26Al–26Mg and 182Hf–182W systematics are generally consistent, those determined by 26Al–26Mg and 206Pb–207Pb systematics are largely inconsistent. The homogeneous versus heterogeneous distribution of 26Al in the protoplanetary disk remains controversial.

9.1 Introduction

26 26 Al is a short-lived radionuclide which decays to Mg with a half-life (t1/2) of ~0.71 million years (Norris et al., 1983; Nishiizumi et al., 2004). The presence of now-extinct 26Al in the early solar system was inferred from excesses of daughter isotope 26Mg (26Mg*) among Ca- and Al- rich inclusions (CAIs), the oldest solar system objects (e.g., Lee et al., 1976). A correlation of excesses of 26Mg with Al/Mg ratio in constituent minerals is considered to be indicative of in situ decay of 26Al (i.e., an isochron) and allows an initial 26Al/27Al ratio to be inferred for when these minerals crystallized. Unlike radiometric dating using long-lived nuclides that provides ages in an absolute timescale (hereafter called as “absolute” age), such as the U–Pb chronom- eter, Al–Mg dating provides a relative chronology: the inferred 26Al/27Al ratios of multiple objects measure the time difference among them by assuming their 26Al/27Al ratios decrease

247 248 Kazuhide Nagashima, Noriko T. Kita, and Tu-Han Luu

26 26 27 26 27 with time according to the Al half-life. When Al/ Al of a sample [( Al/ Al)sample]is 26 27 26 27 compared with that of a reference Al/ Al ratio [( Al/ Al)ref], a “relative” age (Δt) of the sample is calculated with a following equation: 2 3 26Al 1 6 27Al 7 Δt ¼ ln 4ref 5 λ 26Al 27 Al sample

7 26 where λ (~9.8 10 ) is the decay constant of Al, defined as λ = ln(2)/t1/2. If an absolute age of the reference is known, the derived relative age of the sample can be converted to an absolute age. In the past ~40 years, Al–Mg systematics has been analyzed in a large number of CAIs. MacPherson et al. (1995) summarized initial 26Al/27Al ratios reported in CAIs by that time and suggested the canonical 26Al/27Al ratio of ~5 105 represented the 26Al/27Al ratio at the time when most CAIs formed. This conclusion was further strengthened by more recent high precision bulk analyses of CAIs using multi-collector inductively coupled plasma source mass spectrometry (MC-ICPMS) by several groups (e.g., Bizzarro et al., 2004; Jacobsen et al., 2008; Larsen et al., 2011), and precisely merged to ~5.2 105. In addition to the bulk analyses, in situ analyses of unmelted CAIs from CV (Vigarano-type) chondrites using secondary ion mass spectrometry (SIMS) are also consistent with the canonical ratio (MacPherson et al., 2010, 2012). The canonical 26Al/27Al ratio has been assumed to represent an initial 26Al/27Al ratio of the entire solar system – i.e., the initial 26Al/27Al ratio was homogeneously distributed through- out the protoplanetary disk, and has been used as the reference 26Al/27Al ratio to obtain relative ages of various early solar system objects. With the assumption of homogeneous distribution of 26Al in the protoplanetary disk, the 26Al–26Mg system plays an important role for dating of chondrules as it provides high-time- resolution chronology (1 million years). This is because of the short half-life of 26Al, and because chondrules often contain Al-rich phases suitable for measuring the 26Al–26Mg system. The 26Al–26Mg ages of chondrules suggest the majority of chondrules formed ~13 million years after CAIs, and, therefore, can potentially provide important cosmochemical constraints such as a lower limit on the lifetime of the protoplanetary disk, duration of chondrule-forming and chondrite accretion processes (e.g., Kita and Ushikubo, 2012 and references therein). However, the use of 26Al–26Mg systematics to date chondrules has been challenged by Larsen et al. (2011), who suggest there is large-scale heterogeneity of 26Al/27Al in the protoplanetary disk, inferred by combining high-precision 26Al–26Mg measurements and 54Cr isotopic com- positions of bulk chondrites. At the same time, recent 182Hf–182W relative age determinations of multiple meteoritic samples show a good agreement with the 26Al–26Mg ages, supporting the homogeneous distribution of 26Al/27Al (Kruijer et al., 2014). Since then, there have been a large number of analyses and debates on homogeneous versus heterogeneous distribution of 26Al in the protoplanetary disk, and the ability of 26Al to date chondrules is at stake. In this chapter, we summarize recent progresses and issues on 26Al–26Mg systematics of chondrules, especially from the last ~5 years (the latest review on this topic was published by Kita and Ushikubo (2012)), and discuss how they would impact on the understanding of chondrule forming events. 26Al–26Mg Systematics of Chondrules 249 9.2 Initial 26Al/27Al Abundances in Chondrules Inferred from Internal Isochrons

Most chondrules consist of ferromagnesian silicates (olivine and pyroxene), Fe-Ni metal, and mesostasis composed of feldspathic glass and high-Ca pyroxene crystallites. In many chon- drules from carbonaceous chondrites, crystalline plagioclase ((Na,Ca)(Si,Al)4O8) is present. Plagioclase is generally anorthitic (CaAl2Si2O8) in FeO-poor, type I chondrules and is often more albitic (NaAlSi3O8) in FeO-rich, type II chondrules. These plagioclases and some feldspathic glass have low concentrations of MgO (1 wt%), corresponding to high 27Al/24Mg ratios (30). Since we look for an excess of 26Mg (26Mg*) due to in situ decay of 26Al, and 26 Mg* depends on Al/Mg ratio, plagioclase and feldspathic glass are good candidates to detect 26 26 27 26 =27 Internal Mg* and to determine initial Al/ Al ratios [ Al Al 0 ] recorded in chondrules. However, in general, these phases in chondrules are small (less than tens of microns), and, therefore, difficult to be isolated and measured with the conventional mass spectrometry such as MC-ICPMS. Instead, they can be measured in situ using SIMS, also known as ion microprobe. SIMS uses high-energy ions (i.e., primary ions) to excavate materials from sample surface with leaving a hole (i.e., sputtering). Among those sputtered by primary ions, less than 10 percent of Al and Mg atoms are ionized and extracted into mass spectrometer to analyze for elemental and isotopic compositions. Because primary ion beam can be focused from tens of microns to submicron, SIMS allows measurement of plagioclase and glass in chondrules without having any contribution from adjacent minerals. Indeed, the first detection of 26Mg* in a plagioclase- bearing chondrule from the Semarkona (LL3.00) (Grossman and Brearley, 2005; LL3.01 proposed by Kimura et al., 2008) chondrite by Hutcheon and Hutchison (1989) was made with a small geometry magnetic sector SIMS instrument. Plagioclase grains in type I chondrules show a narrow spread in their Al/Mg ratios, typically ranging from ~30 to 60 (e.g., Kita and Ushikubo, 2012 and references therein), and higher Al/Mg ratios are often found in anorthitic plagioclase from Al-rich chondrules and in albitic plagioclase from type II chondrules (e.g., Rudraswami et al., 2008; Nagashima et al., 2014), due to variable MgO contents introduced by igneous fractionation (e.g., Kurahashi et al., 2008). These Al/Mg variations, however, are often insufficient to construct a precise internal isochron. As a result, low Al/Mg phases such as olivine and pyroxene are measured, and an internal isochron is determined from low Al/Mg phase(s) and plagioclase and/or feldspathic glass within a single chondrule. The slope of the 26 =27 Internal isochron corresponds to Al Al 0 at the time of the last melting of the chondrule precursor(s) as long as low Al/Mg ferromagnesian silicate(s) and plagioclase and/or glass formed from a same chondrule melt. Examples of internal isochrons obtained through analyses of plagioclases and olivine and/or pyroxene using the SIMS technique are shown in Figure 9.1.

9.2.1 Disturbed/Undisturbed Al–Mg Systematics in CAIs and Chondrules: Towards Determination of Representative Initial 26Al/27Al Ratios in Chondrules

Since the first detection of 26Mg* in a chondrule from Semarkona, many chondrules from unequilibrated ordinary chondrites (UOCs) and carbonaceous chondrites (CCs) with different petrologic subtypes ranging from type 2 to >3.6 have been analyzed using SIMS, and their 26 =27 Internal Al Al 0 were obtained from internal isochrons (e.g., Kita and Ushikubo (2012) and 250

Figure 9.1 26Al–26Mg systematics of chondrules from the least metamorphosed chondrites. BSE images (a, c, e) and 26Al–26Mg evolution diagrams (b, d, f ) of chondrules from CR2 (El Djouf 001), CV3.1 (Kaba), and CH3 (Acfer 214) chondrites, respectively. The chondrules in (a) and (e) are Al-rich chondrules, and that in (c) is a type I (magnesian) chondrule. Aluminum and magnesium isotopic compositions were obtained from crystalline plagioclase, olivine, and/or low-Ca 26 26 pyroxene using secondary ion mass spectrometry (SIMS). SIMS measurement spots (~5–10 µm size) are indicated by white arrows. Excesses of Mg (δ Mg*) 26 26 27 26 =27 Internal are correlated with their Al/Mg ratios, indicative of in situ decay of Al. Their inferred initial Al/ Al ratios [ Al Al 0 ] are also shown in (b, d, and f ). Note the data for CH3 are only from plagioclase and the regression line is forced through the origin. Data are from Nagashima et al. (2014, 2017); Krot et al. (2014b). cord: cordierite; mgt: magnetite; ol: olivine; pl: plagioclase; px: low-Ca pyroxene; sp: spinel. 26Al–26Mg Systematics of Chondrules 251 references therein). These studies show that most chondrules from the least metamorphosed chondrites (petrologic types <3.1) have 26Mg excesses at levels detectable by ion microprobe analysis (e.g., Kita et al., 2000; Mostefaoui et al., 2002). In contrast, 26Al–26Mg systematics in chondrules from chondrites of higher petrologic types (>3.2) often do not show resolvable 26Mg* (e.g., Huss et al., 2001). This systematic difference could be explained by redistribution of Al and Mg isotopes during thermal metamorphism in higher petrologic type samples. The disturbance of 26Al–26Mg systematics has been also observed in primary anorthite grains in CAIs from metamorphosed chondrites. MacPherson et al. (1995) compiled the existing 26Al–26Mg data on Al-rich objects from chondrites and on achondrites. They demonstrated that initial 26Al/27Al ratios recorded in CAIs have a bimodal distribution, with one peak at ~5 105 and another peak at ~0. The former corresponds to the canonical 26Al/27Al ratio that may represent the initial abundance of 26Al in the early Solar System at the time of CAI formation. 26 27 On the other hand, rare CAIs with ( Al/ Al)0 of ~0 such as FUN CAIs (Fractionated and Unidentified Nuclear isotope effects) (Wasserburg et al., 1977) and PLACs (platy hibonite crystals) (Ireland, 1990) were interpreted as either forming with low 26Al abundance or formed 26 27 with the canonical ( Al/ Al)0 but reprocessed several million years later either in the solar nebula or in an asteroidal setting. The majority of “normal” (i.e., non-FUN) CAIs with low 26 27 ( Al/ Al)0 ratios show petrologic and mineralogical evidence for thermal metamorphism and/ 26 27 or metasomatic alteration, suggesting that their low ( Al/ Al)0 are a result of disturbance of Al–Mg isotope systematics by redistribution of these isotopes. The redistribution of Mg isotopes during thermal metamorphism is supported by Mg-self-diffusion rate in anorthitic plagioclase. LaTourette and Wasserburg (1998) reported Mg-isotope self-diffusion coefficients in anorthite and showed that metamorphic temperature of 500‒600 C could cause Mg-isotope exchange within ~100 µm-sized anorthite crystal in ~1 million years. The model calculations by Yurimoto et al. (2000) and Ito and Messenger (2010) also demonstrated that Al–Mg isotope systematics of anorthite grains in CV CAIs could have been disturbed by redistribution of magnesium isotopes by diffusion at ~400‒600 C in ~1 million years during thermal metamorph- ism. Since the Mg diffusion rate in albitic plagioclase is much faster than in anorthite (Van Orman et al., 2014), Al–Mg isotope systematics in type II chondrules that contain albitic plagioclase are more susceptible to redistribution of Mg. Several recent studies on Al–Mg system in chondrules from CV (Vigarano-type) chondrites also show disturbance related to parent body thermal metamorphism. The inferred 26 =27 Internal Al Al 0 ratios reported in chondrules from CV chondrites (both reduced (CVred) 5 and Allende-like oxidized (CVoxA)) of petrologic subtype >3.1 range from ~1.5 10 to <2 106; about 40 percent of them show no resolvable excess of 26Mg* (Srinivasan et al., 2000; Hutcheon et al., 2009; Sano et al., 2014). In contrast, all chondrules from Kaba (Bali-like 26 =27 Internal oxidized CV3.1 chondrite (CV3.1oxB)) have Al Al 0 resolved from zero and fall within a narrow range, (4.8 2.1) 106 (Nagashima et al., 2017). In addition to petrologic subtypes, careful petrologic investigation of plagioclase grains in chondrules could help to identify disturbed/undisturbed Al–Mg systematics in chondrules. Figure 9.2 shows two 26 26 examples of disturbed Al– Mg systematics in chondrules from Yamato 980145 (CVoxA) and Efremovka (CVred). While these chondrites are classified as weakly metamorphosed chondrites (subtypes 3.1‒3.4, Bonal et al., 2006; Komatsu et al., 2014), some plagioclase grains contain nepheline laths replacing primary plagioclase, indicating the chondrule could have experienced relatively high temperature metasomatic alteration (>400 C), as estimated for 252 Kazuhide Nagashima, Noriko T. Kita, and Tu-Han Luu

Figure 9.2 26Al–26Mg systematics of weakly metamorphosed chondrules. BSE images (a, c) and 26Al–26Mg evolution diagrams (b, d) of type I and Al-rich chondrules from CV3s, Yamato 980145 (CVoxA), and Efremovka (CVred), respectively. These chondrules contain nepheline replacing plagioclase, indicating these chondrules experienced thermal metamorphism on their asteroid. Their 26Al–26Mg systematics do not show correlated excesses of 26Mg (δ26Mg*) with their Al/Mg ratios, most likely due to disturbance of Al–Mg systematics in these chondrules during metamorphism. Data are from Nagashima et al. (2017) and Sano et al. (2014). M: iron sulfide; ne: nepheline.

26 the CVoxA (Bonal et al., 2006; Huss et al., 2006). Accordingly, the lack of Mg* in these – chondrules probably resulted from the disturbance of Al Mg systematics during this alteration. 26 =27 Internal Nakashima et al. (2016) observed systematic differences in Al Al 0 of chondrules between unaltered and altered domains of the reduced CV breccia RBT 04143. The excess silica components induced into anorthite along with relatively higher MgO concentrations (0.5‒1wt%) are observed in anorthite from the least metamorphosed chondrites and indicate anorthite crystal- lization under high temperature, while they were erased systematically in mildly metamorphosed chondrites (Tenner et al., 2014). Collectively it is important to investigate chondrules from chondrites of very low petrologic types (<3.1) in order to obtain undisturbed 26 =27 Internal 26 27 Al Al 0 , representing Al/ Al abundances when chondrules formed. It is poorly understood whether low temperature aqueous alteration has affected 26Al–26Mg systematics in chondrules. In carbonaceous chondrites that experienced extensive aqueous alter- ation (e.g., CM, Mighei-type), glassy mesostases are often completely replaced by phyllosilicates, and, therefore, primordial 26Al–26Mg systematics cannot be obtained. Plagioclase grains that have survived in weakly aqueously altered chondrites, such as most CR2 (Renazzo-type) chondrites, appear to be unaffected, and could retain undisturbed 26Al–26Mg systematics. With diffusion coefficients of Mg self-diffusion in plagioclase obtained under dry conditions (LaTourrette and 26Al–26Mg Systematics of Chondrules 253

Wasserburg, 1997; Faak et al., 2013; van Orman et al., 2014), it is impossible to disturb Mg-isotopes in plagioclase under low temperature aqueous alteration experienced by CRs (<100 C) and even Kaba (<300 C). Cherniak (2010) demonstrated that water has little effect on Na–K interdiffusion in alkali feldspars, but affects the CaAl–NaSi interdiffusion in plagioclase. It is concluded that migration rate of tetrahedral cations may depend on water. Although Mg self-diffusion in plagioclase under wet conditions has not been measured yet, because Mg does not diffuse through tetrahedral sites in plagioclase, it is plausible that Mg diffusion in plagioclase is not enhanced under hydrothermal conditions. Tenner et al. 26 –26 (2013), Nagashima et al. (2014), and Schrader et al. (2017) measured Al Mg systematics in 26 =27 Internal chondrules from CR2 3 chondrites, and found very similar distributions of Al Al 0 ratios from CR3.0s to those of weakly altered CR2s. These data also support that 26Al–26Mg systematics in chondrule plagioclase are undisturbed in weakly altered chondrites. We suggest 26Al–26Mg systematics obtained from crystalline plagioclase grains in chondrules from weakly aqueously altered chondrites most likely represent 26Al/27Al abundances when chondrules crystallized.

9.2.2 Initial 26Al/27Al Ratios in Chondrules from the Least Metamorphosed Chondrites 26 –26 Kita and Ushikubo (2012) summarized the existing data on Al Mg systematics in chondrules 26 =27 Internal from type 3.0 chondrites and discussed representative Al Al 0 for chondrules in the least metamorphosed chondrites. They showed that there is a spread in the initial 26Al/27Al ratios among chondrite groups (Figure 9.3). Chondrules from LL3.00 (Semarkona), CO3.05 (Yamato 26 =27 Internal 81020), and Acfer 094 (ungrouped C3.00) have Al Al 0 ratios corresponding to (7.9 3.0) 106,(6.1 1.9) 106, and (5.7 1.5) 106 (average 1SD), respectively. The relative probability density plots in Figure 9.3b also show that LL3.00 chondrules appear to have 26 =27 Internal slightly higher Al Al 0 compared to those of CO3.05 and Acfer 094. To minimize inter-laboratory biases, Kita and Ushikubo (2012) used only the data obtained with ims-1270/ 1280 at the Geological Survey in Japan and at the University of Wisconsin–Madison. The 26 =27 Internal Al Al 0 ratios for the LL3.00, CO3.05, and Acfer 094 chondrules they reported are (7.1 1.7) 106,(6.7 2.0) 106, and (6.6 1.5) 106, respectively. Those selected by Kita and Ushikubo (2012) and datasets compiled here are similar, with possibly higher Internal 26Al=27Al ratios in UOC chondrules compared to other two groups. 0 26 =27 Internal Since Kita and Ushikubo (2012) paper, representative Al Al 0 have been obtained for chondrules from CR23 (Tenner et al., 2013; Nagashima et al., 2014; Schrader et al., 2017), CH3 (Krot et al., 2014b), and CV3.1 (Nagashima et al., 2017) chondrites. The CR chondrites used in these studies do not include extensively altered CR2s, and none of the chondrules show any clear evidence for aqueous alteration or thermal metamorphism. CH3 chondrites appear to have avoided thermal metamorphism: their peak metamorphic temperature was estimated 300 C (Reisener et al., 2000). The CVoxB3.1 chondrite Kaba experienced extensive metaso- matic alteration, but in contrast to Allende (CVoxA3.6), largely avoided subsequent thermal metamorphism (Krot et al., 1997; Kimura and Ikeda, 1998; Bonal et al., 2006). As discussed in Krot et al. (1998a) and Zolotov et al. (2006), the secondary mineral assemblages in Kaba are consistent with their formation by aqueous alteration under relatively low temperature, below 254 Kazuhide Nagashima, Noriko T. Kita, and Tu-Han Luu

(a)

(b)

26 27 26 =27 Internal Figure 9.3 (a) Initial Al/ Al ratios [ Al Al 0 ] obtained via in situ analysis of chondrules from the least metamorphosed ordinary (LL3.00) and carbonaceous (CO3.05, Acfer 094 (ungrouped type 3.00), CR23, CH3, and CV3.1) chondrites. Also shown are relative ages of the chondrules calculated from 26 27 26 27 5 their ( Al/ Al)0 ratios and the canonical ( Al/ Al)0 ratio of 5.2 10for CV CAIs (Jacobsen et al., 2008; 26 =27 Internal Larsen et al., 2011). (b) Relative probability density plots for Al Al 0 shown in (a). The 26 27 ( Al/ Al)0 ratios among the chondrite groups appear to show systematic differences, LL CO A094 CV > CR CH. Data for LL3.00 (Semarkona) are from Hutcheon and Huchison (1989), Kita et al. (2000), McKeegan et al. (2000), Huss et al. (2001), Rudraswami et al. (2008), and Villeneuve et al. (2009); for CO3.05 (Yamato 81020) from Yurimoto and Wasson (2002), Kunihiro et al. (2004), and Kurahashi et al. (2008); for Acfer 094 (ungrouped type 3.00 carbonaceous chondrite) from Ushikubo et al. (2013); for CR2‒3 chondrites from Kurahashi et al. (2008), Schrader et al. (2017), Tenner et al. (2013), and Nagashima et al. (2014); for CH3 chondrites from Krot et al. (2014b); for CV3.1 (Kaba) from Nagashima et al. (2017).

300 C. Therefore, these chondrules meet the criteria discussed in Section 9.2.1, and their 26 –26 26 =27 Internal Al Mg systematics are likely undisturbed. Note the Al Al 0 ratios in CH3 chondrules reported by Krot et al. (2014b) are model isochrons in which regression lines are forced through the origin (i.e., olivine or pyroxene has not been measured). 26Al–26Mg Systematics of Chondrules 255

26 27 The initial Al/ Al ratios from CR2 3 and CH3 chondrules are distinctly different from 26 =27 Internal 6 those of UOCs and COs: while a few chondrules have Al Al 0 ratios of ~6 10 , comparable to those of LL3.00, CO3.05, and Acfer 094, most of the data are systematically lower than ~3 10 6 (and often unresolved from 0); >60 percent of CR23 chondrules and 26 ~80 percent of CH3 chondrules have no resolvable Mg* excesses (Figure 9.3). The very 26 =27 Internal similar distributions of chondrule Al Al 0 ratios between CR2 3s and CH3 are consistent with the fact that these two chondrite groups appear to be closely related (e.g., Weisberg et al., 1995; Krot et al., 2002). 26 =27 Internal 6 The Al Al 0 ratios of Kaba chondrules, (4.8 1.1) 10 , are similar to those of LL3.00, CO3.05, and Acfer 094, and systematically higher than those of CR and CH chon- 26 =27 Internal drules. The Al Al 0 ratios of the porphyritic chondrules from the chondrite groups > decrease in the order UOCs COsAcfer 094 CVs CRs CHs. 26 =27 Internal In addition to the differences in Al Al 0 among chondrite groups, each chondrite 26 =27 Internal group may also have several populations of chondrules with different Al Al 0 ratios. Kita and Ushikubo (2012) pointed out that variability among chondrules is very small for 26 27 CO3.0 and Acfer 094, which contain a few chondrules with marginally different ( Al/ Al)0 ratios. A few studies investigated the presence of distinct populations of chondrule 26 =27 Internal Al Al 0 ratios within several chondrite groups (UOCs, COs, CRs, and Acfer 094) using statistical approach (e.g., Villeneuve et al., 2009; Schrader et al., 2017). They found that each chondrite group contains more than one population of chondrules. Some of these differences may be linked to gas/solid fractionation for Mg, Si, and other volatile elements among LL3 chondrules (Mostefaoui et al., 2002; Tachibana et al., 2003), but also to oxygen-fugacity and oxygen–isotope compositions in chondrule forming environments for CR3 chondrules (Tenner et al., 2013). If, after an epoch of CAI formation characterized by heterogeneous distribution of 26Al/27Al (Krot et al., 2009 and references therein), 26Al/27Al was homogenized in the proto- planetary disk at the canonical level (26Al/27Al ~ 5.2 105, Jacobsen et al., 2008), the 26 =27 Internal Al Al 0 ratios recorded by chondrules correspond to their formation ages (a vertical scale in the right of Figure 9.3a). If this is the case, almost all chondrules measured to date for Al–Mg isotope systematics formed after ~1.5 million years, postdating the time indicated from canonical 26Al/27Al ratios of CAIs. Most chondrules from UOCs, COs, and CVs formed between 1.5 and 3 million years, while many chondrules from CRs and CHs formed more than 3 million years after canonical CAIs. The 26Al–26Mg ages of chondrules indicate the > chondrule formation lasted at least 2 million years, most likely 3 million years. In this scenario, 26 =27 Internal the distinct populations of Al Al 0 ratios within a single chondrite group suggest each chondrite group contains chondrules formed at different times (i.e., multiple generations of chondrules).

9.3 26Al/27Al Abundances for Chondrule Precursors Inferred from Bulk Chondrules

Internal 26Al–26Mg isochrons from minerals in chondrules date the last melting event, and, therefore, provide little clue on the age of chondrule precursors. In contrast, 26Al–26Mg systematics of bulk chondrules measured using MC-ICPMS with much higher precision (~ppm to tens of ppm) than internal 26Al–26Mg isochrons together with an assumed initial 256 Kazuhide Nagashima, Noriko T. Kita, and Tu-Han Luu

26Mg/24Mg ratio are being used to constrain model ages of chondrule precursors. Bulk chon- drule measurements are undertaken by either separating chondrules from the host meteorites (e.g., Galy et al., 2000; Luu et al., 2015) or by sampling of materials from chondrules exposed on meteorite section using microdrilling technique (e.g., Bizzarro et al., 2004). Galy et al. (2000) first reported excesses of 26Mg* in bulk chondrules, followed by Bizzarro et al. (2004), 26 who showed most bulk chondrules from Allende CV3 chondrite have excesses of Mg* 26 =27 Bulk ‒5 corresponding to Al Al 0 ~(3 6) 10 (Figure 9.4). The authors of these two papers argue that bulk chondrule isochrons are less susceptible to the effects of postcrystallization 26 redistribution ofMg* amongst different minerals than internal isochrons, and interpreted the 26 =27 Bulk inferred bulk Al Al 0 as more reliable chondrule formation ages than the in situ measured internal 26Al–26Mg isochrons. Bizzarro et al. (2004) suggested that Allende chondrule formation began contemporaneously with the formation of CAIs, and continued for at least 1.4 26 =27 Bulk million years. However, the interpretation that Al Al 0 ratios correspond to formation ages of chondrules has been questioned, and it is now generally accepted that they likely correspond to the formation time of chondrule precursors, not chondrule crystallization ages (e.g., Kita et al., 2005; Krot et al., 2009, Luu et al., 2015). Based on the data reported by Bizzarro et al. (2004), some chondrule precursors can be as old as CAIs, while others have been produced over a period of a few million years in the protoplanetary disk. Recently Luu et al. (2015) reported 26Al–26Mg systematics of bulk chondrules from Allende. These chondrules, including Al-rich chondrules, appear to form a single correlation line in µ26Mg* (µ-notation: per 106 deviations from the reference value) versus 27Al/24Mg diagram (Figure 9.4). Luu et al. (2015) interpreted that the line indicates a bulk chondrule isochron 26 =27 Bulk 5 corresponding to Al Al 0 ratio of (1.2 0.2) 10 . The isochron is clearly distinct from the canonical 26Al/27Al ratio (~5.2 105) obtained from measurements of bulk CAIs (Jacobsen et al., 2008; Larsen et al., 2011) and many bulk chondrule data from Bizzarro et al. (2004). Together with the high precision bulk chondrule data from Galy et al. (2000), Bizzarro et al. (2004), and Villeneuve et al. (2009), Luu et al. (2015) suggested that some chondrule precursors might have started forming contemporaneously with the formation of CAIs precursors 26 27 – (with the canonical Al/ Al ratio) but this process of Al Mg fractionation in chondrule precur- 26 =27 Bulk 5 sors had likely ended up by the time Al Al 0 decreased to ~1.2 10 , this inferred ratio the authors call “minimum” bulk-chondrule isochron. With an assumption that 26Al/27Al was homogeneously distributed throughout the protoplanetary disk at the canonical level, the min- imum bulk-chondrule isochron corresponds to 1.5 0.2 million years after the formation of CAI precursors, and may reflect the timing of when condensation of chondrule precursors ceased or the timing of when chondrule precursors were separated from a nebular reservoir (Luu et al., 2015). 26 27 The difference of ( Al/ Al)0 between the bulk and in situ chondrule data may correspond to a time difference between the precursor formation and the last melting of chondrules. Luu et al. 26 –26 (2015) attempted to obtain such time differences by measuring in situ Al Mg systematics 26 =27 Internal from the chondrules in which the bulk chondrule data were acquired. The Al Al 0 ratios 5 5 from in situ measurements range from (0.350.57) 10 to (2.6 1.53) 10 .Mostofthe 26 =27 Internal individual chondrules measured have a Al Al 0 ratio lower than that of the bulk isochron, which is expected as the formation of chondrules cannot predate the formation of their precursors. Combining bulk and in situ data actually shows that melting of precursors to form chondrules took place from 0 to ~ 2 million years after the formation of their precursors. The few 26Al–26Mg Systematics of Chondrules 257

Figure 9.4 26Al–26Mg systematics of bulk chondrules from CV and CR chondrites. Magnesium isotopic compositions were obtained with multi-collector inductively coupled plasma mass spectrometry 26 27 (MC-ICPMS) and shown asµ-notation (i.e., parts per million) relative to standards. The initial Al/ Al 26 =27 Bulk 5 ratios of bulk chondrules [ Al Al 0 ] show a large spread ranging from ~0 to ~5 10 . The bulk chondrule isochron likely represents formation ages of chondrule precursors. Luu et al. (2015) obtained a 26 =27 Bulk 5 correlation line from most of the chondrules, corresponding to Al Al 0 ~1.2 10 (dotted-line), and interpreted as the minimum bulk chondrule isochron which corresponds to the time that major production of chondrule precursors stopped in the protoplanetary disk. Data from Olsen et al. (2016) have a spread ranging from ~1 105 to ~2 105. Enlarged view for a low Al/Mg region is shown in (b). The canonical 26 27 5 26 ( Al/ Al)0 of 5.2 10 with initial Mg-isotope compositions (µ Mg*0)of40 ppm (Jacobsen et al., 2008) and 16 ppm (Larsen et al., 2011) are also shown in (b) as dotted-dashed and solid lines, respectively. 258 Kazuhide Nagashima, Noriko T. Kita, and Tu-Han Luu exceptions reported in Luu et al. (2015) might correspond to chondrules having a higher bulk 26Al/27Al ratio than the minimum bulk chondrule isochron of 1.2 105. Additional studies on 26Al–26Mg systematics of bulk chondrules from carbonaceous (CV, CR, and CH/CB) and ordinary chondrules have been reported by several groups (Claydon et al., 2014, 2015; van Kooten et al., 2016; Olsen et al., 2016; Hayakawa et al., 2017). The bulk chondrule 26Al–26Mg systematics from CV chondrites appear to show large variations. This is in contrast to the in situ measured 26Al–26Mg systematics of chondrules from Semarkona (LL3.00), Yamato 81020 (CO3.05), and CR23s that have narrow ranges of the inferred 26 27 ( Al/ Al)0 and show general consistency among data obtained by different laboratories. About half of chondrules measured from CV3 chondrites Allende, Vigarano, and Mokoia by Claydon 26 =27 Bulk 5 et al. (2014, 2015) are consistent with Al Al 0 of (1.8 0.2) 10 , slightly higher than the (1.2 0.2) 105 obtained by Luu et al. (2015), while the rest have lower ratios or do not have resolvable 26Mg*. The bulk chondrules from CV3 chondrites Vigarano and North West Africa (NWA) 3118 reported by Olsen et al. (2016) are characterized by systematic deficits in µ26Mg* (Figure 9.4b) relative to the solar value of 4.5 ( 1.0) ppm defined by CI chondrites (Larsen et al., 2011). These µ26Mg* values, together with the initial Mg-isotope composition (~ 16 ppm; Larsen et al., 2011) obtained by bulk isochron with CAIs and AOAs 26 =27 Bulk 5 (amoeboid olivine aggregates) are consistent with model Al Al 0 ratios of ~0.8 10 5 26 =27 Bulk 5 5 to ~2.7 10 (note they correspond to Al Al 0 ratios of ~2 10 to ~7 10 with the initial Mg-isotope composition of ~40 ppm reported by Jacobsen et al. (2008)). All CV3 chondrites experienced metasomatic alteration to varying degrees, and chondrules in them could have exchanged some elements and isotopes not only within chondrules but also with surrounding matrix materials (i.e., experienced open-system alteration on their parent asteroid (e.g., Brearley and Krot, 2012)). The Fe–Mg exchange often observed in chondrule olivines from CV3 chondrites can modify bulk chondrule Al/Mg ratios and Mg-isotope exchange between chondrule constituents and matrix materials can modify bulk chondrule 26 =27 Bulk Mg-isotope compositions. These processes produce a lower Al Al 0 ratio than the original one, and thus open-system alteration could be responsible in part for some variations observed in bulk 26Al–26Mg systematics in CV chondrules observed by Luu et al. (2015), Claydon et al. (2015), and Olsen et al. (2016). However, these processes do not seem to explain Allende chondrules measured by Bizzarro et al. (2004) which have much higher 26 =27 Bulk Al Al 0 than those obtained from less-metamorphosed CV chondrules by Olsen et al. (2016). One explanation could be that some of the bulk chondrules used in Bizzarro et al. (2004) are Al-rich chondrules which often contain recycled CAI materials (e.g., Krot et al., 2009), and thus some 26Mg* might be inherited from CAI materials.

9.4 Homogeneous/Heterogeneous Distribution of 26Al/27Al in the Protoplanetary Disk

While it is generally accepted that the CAI-forming region was largely characterized by initial 26Al/27Al ratio of the canonical value (~5.2 10 5) inferred from high precision 26Al–26Mg systematics of bulk CAIs from multiple laboratories (e.g., Jacobsen et al., 2008; Larsen et al., 2011), the homogeneous versus heterogeneous distribution of 26Al in the early solar system 26Al–26Mg Systematics of Chondrules 259 throughout the protoplanetary disk (i.e., in regions where chondrules formed) is highly contro- versial (e.g., Kita et al., 2013). There are several studies both in favor of and against the homogeneous 26Al distribution within the protoplanetary disk. Villeneuve et al. (2009) determined initial 26Al/27Al ratios and their initial Mg-isotope 26 24 composition [( Mg/ Mg)0] in chondrules from Semarkona and found that the all ferromagne- sian chondrules plot on the growth curve expected from the canonical 26Al/27Al ratio of 5.2 5 26 10 and initial µ Mg*0 of –40 29 ppm inferred from bulk CAI measurements by Jacobsen et al. (2008), indicating that precursors of these chondrule were derived from the same reservoir. Therefore, Villeneuve et al. (2009) concluded that 26Al was homogeneously distributed within the protoplanetary disk characterized by the canonical 26Al/27Al ratio. In contrast, Larsen et al. (2011) proposed large 26Al heterogeneity in the protoplanetary disk. They measured 26Al–26Mg systematics of bulk CAIs and AOAs and confirmed the canonical 26Al/27Al ratio but obtained 26 an initial Mg-isotopic ratio, µ Mg*0 = 15.9 1.4 ppm, which is at a face value much higher than that indicated by Jacobsen et al. (2008). They also measured bulk meteorites, including CI carbonaceous, enstatite, and ordinary chondrites, and found that model isochrons of these 26 chondrites calculated together with the µ Mg*0 of –16 ppm are inconsistent with the canonical 26Al/27Al ratio but have 26Al/27Al ratios as low as ~2 105. These authors suggested large- scale heterogeneity may have existed throughout the inner solar system. This interpretation has been questioned by Wasserburg et al. (2012) and Kita et al. (2013) who suggested that the 26 µ Mg*0 of 16 ppm is not accurate due to inclusion of AOAs in the regression, which might have formed later than CAIs. If the regression line is calculated for the bulk CAIs without AOAs, the bulk µ26Mg* of chondrites are consistent with uniform distribution of 26Al in the disk at the canonical level. On the other hand, we note that Villeneuve’s approach does not 26 27 26 5 distinguish the two sets of initial Al/ Al ratio and µ Mg*0, [5.2 10 and 40 ppm] versus [~2 105 and 16 ppm] as the growth curves calculated from these two sets are very similar and are consistent with the 26Al–26Mg systematics of bulk ferromagnesian chondrules reported by Villeneuve et al. (2009). More recently, van Kooten et al. (2016) reported that whole-rock, bulk chondrules, and bulk lithic clasts from CR and CH/CB chondrites have systematic deficits in 26Mg* ranging from ~ –18 to ~ –2 ppm compared to ~4.5 ppm of CI chondrite (Figure 9.4b). The bulk 26Al–26Mg systematics of more chondrules from CR and CV chondrites subsequently reported by Olsen et al. (2016) show that most of them have deficits in 26Mg* and the µ26Mg* values from CR chondrules are systematically lower than those of CV chondrules (Figure 9.4b). These authors discussed that the majority of bulk chondrules with 27Al/24Mg ratios close to the solar value (0.101) are apparently incompatible with the hypothesis of 26Al homogeneity with the canonical 26 26 Al abundance with initial µ Mg*0 of –16 ppm (Larsen et al., 2011) and require reduced and heterogeneous distribution of 26Al in the protoplanetary disk, with an assumption that the chondrules did not experience a complex history involving multiple Al/Mg fractionation events. These two studies (van Kooten et al., 2016; Olsen et al., 2016) together with those reported by Larsen et al. (2011) identified two trend lines in the µ54Cr versus µ26Mg* diagram (Figure 9.5). Larsen et al. (2011) interpreted the correlation line through CAIs, chondrites (CI, CM, ordinary, enstatite, and R), achondrites (ureilite, acapulcoite, angrite), and Martian meteorites reflects preferential thermal processing of infalling molecular cloud materials, including 26Al-rich or 54Cr anomaly carriers, resulting in an enrichment in the abundance of 26Al and 54Cr in the gas 260 Kazuhide Nagashima, Noriko T. Kita, and Tu-Han Luu

Figure 9.5 µ54Cr versus µ26Mg* plot for bulk chondrites and chondrules from carbonaceous, ordinary, enstatite chondrites, angrite, ureilite, and Martian meteorites (from Larsen et al., 2011 and van Kooten et al., 2016). The correlation line defined from most chondrites, achondrites, and CAIs is assumed to have resulted from preferential sublimation of thermally unstable and isotopically anomalous carriers and subsequent fractionation of gas from remaining dust, creating isotopically anomalous reservoirs. In contrast, chondrules and clasts from CR, CH, and CB chondrites are off from the correlation line and making a trend extending to a predicted composition of 26Al-free and thermally unprocessed molecular cloud material, indicated by a light-gray box (van Kooten et al., 2016). The trend was interpreted as a mixing array between two end members defined by the compositions of primordial molecular cloud matter and 26Al-polluted, thermally processed inner solar system material, akin to that of bulk CM chondrites. The 26Mg* deficits in CR, CH, and CB chondrite components may reflect depletions in the initial 26Al/27Al value of their precursors relative to the value defined by CI chondrites, requiring a contribution from a 26Al-poor reservoir different from most inner solar system bodies. phase relative to residual midplane dust, analogous to that proposed by Trinquier et al. (2009). Because µ54Cr and µ26Mg* are positively correlated, the apparent linearity of the array suggests that the individual carrier phases of 26Al and 54Cr have comparable thermal properties. In addition, van Kooten et al. (2016) identified the other trend line based on an array of CR and CH/CB chondrules and lithic clasts, which extends to the hypothetical μ54Cr and μ26Mg* compositions of the 26Al-free and thermally unprocessed molecular cloud material from the line defined by the other meteorites (Figure 9.5). They interpreted the array reflects the binary mixing of two distinct reservoirs, an 26Al-free primordial molecular cloud component and thermally processed inner solar system materials similar to CM chondrites, and suggested that CR and CH chondrites and their components formed in the outer solar system, where high fractions of primordial molecular cloud material are expected to survive. Luu et al. (2016) reported similar bulk Mg-isotope compositions in CR and other chondrites. However, they showed that with all high precision data on bulk chondrites combined, there remains at best only a marginal correlation between μ26Mg* and μ54Cr. Instead, they pointed out that the variability in μ26Mg* for bulk chondrites could easily be explained by differences in their bulk Al/Mg and 26Al–26Mg Systematics of Chondrules 261

26 27 26 27 a uniform, canonical ( Al/ Al)0. Thus the original evidence cited for ( Al/ Al)0 heterogen- eity can be readily interpreted in another way. Moreover, the interpretation of 26Mg* deficits as a reduced and heterogeneous distribution of 26Al/27Al requires an assumption that the chondrules did not experience Al/Mg fractionation. As pointed by Kita et al. (2013), this assumption may not be valid. The chondrule forming process was likely open-system as gas–melt interaction played a major role in the evolution of mineralogy, bulk chemical, and isotopic compositions (e.g., Tissandier et al., 2002; Libourel et al., 2006; Nagahara et al., 2008). The chondrule precursors included CAIs, AOAs, chondrules of earlier generations, fine-grained matrix-like material, and, possibly fragments of preexisting thermally processed planetesimals (e.g., Krot and Nagashima, 2017 and references therein). In the case of CV chondrules, the exchange of elements during parent body metamorphism could also modify Al/Mg ratios. Magnesium-isotope heterogeneity can be an alternative explanation for the 26Mg* deficits, although the lack of large anomalies in Si-isotopic compositions in bulk solar system reservoirs suggests limited Mg heterogeneity, as Mg and Si are predicted to be synthesized by similar nucleosynthetic processes (Pringle et al., 2013). We note Si-isotope heterogeneity among the solar system objects is less than 15 ppm (Pringle et al., 2013), and, therefore, does not completely rule out the possibility of Mg-isotope heterogeneity, as the variation observed by Olsen et al. (2016) is also less than 15 ppm. Instead of investigating only 26Al–26Mg systematics to constrain 26Al homogeneity/hetero- geneity in the protoplanetary disk, the comparisons of ages determined using three radiometric dating methods (26Al–26Mg, 182Hf–182W, and U-isotope ratio corrected 206Pb–207Pb) have been made for angrites and chondrules from different chondrite groups. U-isotope corrected 206Pb–207Pb dating utilizes long-lived radionuclides of 235U and 238U and provides an absolute age, which does not require the assumption of initial homogeneous distribution of the parent nuclides. 182Hf most likely originated from steady-state galactic stellar nucleosynthesis prior to the formation of the protosolar molecular cloud; therefore, its homogenous distribution in the protoplanetary disk is a very reasonable assumption (e.g., Holst et al., 2013; Lugaro et al., 2014). Therefore, ages determined using these two dating methods compared to those from 26Al–26Mg systematics would provide important constraints on 26Al homogeneity/heterogen- eity in the protoplanetary disk. Figure 9.6 shows relative ages determined for two quenched angrites, D’Orbigny and Sahara 99555, using 26Al–26Mg, U-isotope ratio corrected 207Pb–206Pb, and 182Hf–182W systematics (Amelin, 2008a, 2008b; Spivak-Birndorf et al., 2009; Schiller et al., 2010, 2015; Kleine et al., 2012). The 207Pb–206Pb ages originally reported by Amelin (2008a, 2008b) are corrected with U-isotope ratios from Brennecka and Wadhwa (2011) and Connelly et al. (2012). These ages are calculated relative to those of CV CAIs (see Kita et al. (2013) for an alternative age comparison anchoring to angrites); U-corrected Pb–Pb: 4,567.30 0.16 million years (Con- nelly et al., 2012); Al–Mg: 5.2 105 (Jacobsen et al., 2008; Larsen et al., 2011); Hf–W: 1.02 104 (Kruijer et al., 2014). The relative ages of the angrites from Al–Mg systematics are generally consistent among the three data sets (Spivak-Birndorf et al., 2009; Schiller et al., 2010; 2015) and range from ~4.7 to ~5.1 million years after the CV CAIs. These ages and ranges are in good agreement with those from 182Hf–182W relative ages of ~4.6 to ~5.1 million years with ~600,000-year errors (Kleine et al., 2012). In contrast, Pb–Pb ages of the angrites 262 Kazuhide Nagashima, Noriko T. Kita, and Tu-Han Luu

Figure 9.6 Relative chronology of quenched angrites from 26Al–26Mg, U-isotope ratio corrected 207Pb–206Pb, and 182Hf–182W systematics. The relative ages are calculated with respect to those of CV CAIs (U-corrected Pb–Pb: 4,567.300.16 million years (Connelly et al., 2012); Al–Mg: 5.2 105 (Jacobsen et al., 2008; Larsen et al., 2011); Hf–W: 1.02 104 (Kruijer et al., 2014)). Symbols with and without “3” are for D’Orbigny and Sahara 99555, respectively. Al–Mg data are from Spivak-Birndorf et al. (2009), Schiller et al. (2010), Schiller et al. (2015). Pb–Pb data are from Amelin (2008a, 2008b) with U-isotope ratio correction using those from Brennecka and Wadhwa (2011) and Connelly et al. (2012). Hf–W data are from Kleine et al. (2012). Note the Al–Mg relative ages calculated with the reduced 26Al (~1.6 105) abundance for angrite parent body suggested by Larsen et al. (2011) are consistent with their Pb–Pb ages and inconsistent with the Hf–W ages. after the CV CAIs are ~3.7 and ~3.9 million years with ~300,000-year uncertainty and distinctly different from the Al–Mg ages. There is no consensus to interpret the consistency/inconsistency among the ages from the three different chronometers. Schiller et al. (2015) argued that only the Pb–Pb ages are correct and thus the discrepancy between Pb–Pb and Al–Mg ages are due to a reduced abundance of 26Al in the protoplanetary disk, possibly with an initial 26Al/27Al ratio of ~1.3 105. They also speculated that the relative Hf–W ages may not be accurate due to the requirement of secondary corrections for nucleosynthetic effects on W isotopes in CAIs. On the other hand, Kruijer et al. (2014) argued the consistency between Hf–W and Al–Mg ages as an evidence for the homogeneous distribution of 26Al/27Al and pointed out that the alternative CAI Pb–Pb age of 4,567.94 0.31 million years reported by Bouvier et al. (2011) makes ages of angrites from the three chronometers consistent. The similar comparisons have been made for chondrules from different chondrite groups (Figure 9.7). The relative ages are calculated relative to the CV CAI ages listed in the previous text. In addition to the in situ Al–Mg ages shown also in Figure 9.3, the bulk chondrule Al–Mg age from CB metal-rich chondrite (Olsen et al., 2013) is also shown. The relative ages of chondrules from the three chronometers are available for CV and CR carbonaceous chondrites. The U-isotope ratio corrected Pb–Pb ages have been also reported for ordinary and CB chondrules. 26Al–26Mg Systematics of Chondrules 263

(a) UOC

CV

CR -5 =5.2×10

0 CB Al) 27 CO Al/

Al-Mg 26 ( A094 Pb-Pb CH Hf-W

-1 0 12345 6 Time after CV CAIs (Myr)

(b) UOC

CV

CR -5

CB =1.5×10 0 Al)

27 CO Al/

26 A094 ( CH

-1 0123456 Time after CV CAIs (Myr)

Figure 9.7 Relative chronology of chondrules from 26Al–26Mg, U-isotope ratio corrected 207Pb–206Pb, 182 182 26 27 and Hf– W systematics. The relative ages are calculated with respect to the canonical ( Al/ Al)0 of 264 Kazuhide Nagashima, Noriko T. Kita, and Tu-Han Luu

For CV chondrules, the 26Al–26Mg systematics of eight Kaba chondrules have (4.8 2.1) 106 as an average and 2 standard deviation, corresponding to 2.4 (+0.6/0.4) million years after the CV CAIs. There are two very different 182Hf–182WagesofCVchondrules reported by Becker et al. (2015) and Budde et al. (2016b). Becker et al. (2015) determined very old ages based on chondrule separates and other components from Allende and Vigarano, ranging from 1.2 to 0.2 million years with large uncertainty (~3 million years). On the other hand, Budde et al. (2016b) obtained 2.2 0.8 million years after CV CAIs as 182Hf–182W ages of CV chondrules from chondrule separates containing hundreds to thou- sands of chondrules and their fragments and matrix materials, and pointed that Becker et al. (2015) ages are compromised by analytical artifacts on measured W isotope compositions, probably caused by organic interferences on 183W. If this is the case, the Hf–W chondrule age of 2.2 0.8 million years after CV CAIs is in excellent agreement with the Al–Mg age of 2.4 (+0.6/0.4) million years. In contrast, Pb–Pb ages of CV chondrules reported by several studies show a large spread ranging from ~0 to ~3 million years after CAIs, based on individual chondrule ages and pooled data from multiple chondrules. There appears to be a difference between the different methods. The pooled chondrule ages from Amelin and Krot (2007), Connelly et al. (2008), Connelly and Bizzarro (2009) with U-isotope ratio correction made by Brennecka et al. (2015) are consistent with each other within their uncertainties and also consistent with the Al–Mg and Hf–W ages. In contrast, the individual chondrule ages reported by Connelly et al. (2012) are systematically older than the pooled chondrule data and thus inconsistent with the Al–Mg and Hf–Wages. There is a similar data set in CR23 chondrules. The Al–Mg ages of CR chondrules are systematically younger than chondrules from other chondrite groups, except for CB chondrites (Olsen et al., 2013; Nagashima et al., 2014; Schrader et al., 2017). Their simple average and 26 27 2 standard deviations of ( Al/ Al)0 ratios (i.e., including potential multiple generations of chondrules) correspond to 4.1 (+∞/1.7) million years after the CAIs. The relative mean Hf–W age of chondrules from four different CR2 chondrites is 3.7 0.6 million years (Budde et al., 2017), in good agreement with the Al–Mg age. The pooled Pb–Pb age of six CR chondrules (Amelin et al., 2002 with U-isotope correction by Schrader et al., 2017) is 3.7 0.6 million years, in excellent agreement with the Al–Mg and Hf–W ages, while the Pb–Pb ages of individual chondrules have a large spread over ~4 million years (Connelly et al., 2012; Bollard et al., 2014; Bollard et al., 2015) and are largely inconsistent with the Al–Mg and Hf–W ages. Pooled Pb–Pb ages reflect average ages of multiple chondrules included in measurements (e.g.,

Caption for Figure 9.7 (cont.) 5.2 105 and those of CV CAIs in (a) (see Figure 9.6 for data sources). The Al–Mg ages are shown as averages of the data shown in Figure 9.3 and two standard deviation of them. In (b), chondrule relative ages from Al–Mg systematics are calculated with a reduced 26Al abundance of 1.5 105.Pb–Pb data with “3” are based on pooled analyses of multiple chondrules. Pb–Pb ages are from Amelin and Krot (2007), Connelly et al. (2008), Connelly and Bizzarro (2009) with U-isotope ratio correction made by Brennecka et al. (2015), Connelly et al. (2012), Bollard et al. (2014, 2015), Amelin et al. (2002) with U-isotope correction made by Schrader et al. (2016), Huyskens et al. (2016). Hf–W data are from Becker et al. (2015) and Budde et al. (2016b; 2017). Al–Mg data for Hammadah al Hamra 237 CB chondrules are from Olsen et al. (2013); see Figure 9.3 for other Al–Mg data sources. 26Al–26Mg Systematics of Chondrules 265

Connelly et al., 2017). The individual chondrule ages have a large spread and are older than those of the pooled chondrule ages. This may suggest that the old chondrules, such as those with ages similar to the CAIs, are rare and therefore nonrepresentative. Alternatively, the pooled age may not represent an true average of chondrules but may be dominated by a single chondrule that has the largest radiogenic Pb, or the difference might be due to an analytical artifact if the individual chondrule Pb–Pb ages are compromised by the significant amounts of common Pb (Kita et al., 2015). In ordinary chondrite chondrules, the range of Al–Mg ages is relatively narrow (1.8 (+0.7/ 0.4) million years), while the Pb–Pb ages of individual chondrules have a large spread over ~4 million years (Connelly et al., 2012; Bollard et al., 2014; Bollard et al., 2015). In Figure 9.7b, relative Al–Mg ages of the chondrules are shown with respect to initial 26Al/27Al ratio of 1.5 105 as an alternative initial 26Al/27Al ratio in the protoplanetary disk suggested by Larsen et al. (2011), Schiller et al. (2015), and Olsen et al. (2016). This normalization makes all Al–Mg ages older by ~1.3 million years and less consistent with those of Hf–W ages in CV and CR chondrules. The narrow ranges of the Al–Mg ages still contrast with the ~4-million-year spread observed by Pb–Pb ages of individual chondrules. The comparisons made in Figure 9.7 are based on different chondrules independently measured by these three chronometers (i.e., not a comparison of ages from a same chondrule). Bollard et al. (2015) measured Pb–Pb ages and Al–Mg systematics in the same chondrules 26 27 from Allende and NWA 5697 (L3) chondrites, and found that the ( Al/ Al)0 ratios are much lower than those expected from their Pb–Pb ages and the canonical 26Al abundance, support- ing the reduced abundance of 26Al in the disk regions where chondrules originated. We note, however, it is difficult to obtain reliable in situ 26Al–26Mg systematics from Allende chondrules. Considering the data and interpretations described in the previous text, it appears we have not yet reached a consensus whether 26Al/27Al was homogeneously distributed within the protoplanetary disk or not. The good agreement between Al–Mg and Hf–W ages of both angrites and chondrules relative to CAIs supports the homogeneous distribution of 26Al/27Al, while the inconsistency between Al–Mg and Pb–Pb ages of angrites and chondrules and large 26Mg* deficits in the bulk chondrule data are potentially against the homogeneous distribution of 26Al/27Al. Even if the Pb–Pb ages of individual chondrules are correct and 26Al/27Al abundance in the protoplanetary disk was low as ~1.5 10–5, no chondrules in UOCs and CV chondrites were formed at ~0 or later than 3 million years, and therefore the inconsistency – – between Al Mg and Pb Pb systematics remains (Figure 9.7b). We conclude that it is not clear 26 =27 Internal yet what the distinct Al Al 0 ratios within each chondrite group indicate: they could be due to spatial heterogeneity of 26Al/27Al abundances in a chondrule-forming region and/or reflect the presence of multiple generations of chondrules formed at different times.

9.5 Implication for Thermal History of Parent Asteroids

Despite the possibility of heterogeneous distribution of 26Al/27Al in the protoplanetary disk, the 26 27 ( Al/ Al)0 ratios recorded in chondrules still provides important constraints on the thermal 266 Kazuhide Nagashima, Noriko T. Kita, and Tu-Han Luu

Figure 9.8 Initial 26Al/27Al ratios inferred from in situ measurements of chondrules from the least metamorphosed chondrites and estimated 26Al/27Al ratios when parent asteroids accreted based on thermal modeling of the asteroids (solid lines and gray boxes from Doyle et al. (2015) and Sugiura and Fujiya (2014), respectively).

history of their parent asteroids. Even though the duration of accretion of chondrite parent bodies is poorly understood, formation of chondrules must have predated accretion of their 26 27 parent asteroids. Therefore, the ( Al/ Al)0 ratios in chondrules from an individual chondrite group provide an upper limit on 26Al/27Al abundances available for heating up their parent 26 26 27 asteroid due to decay of Al. Figure 9.8 shows ( Al/ Al)0 ratios recorded in chondrules 26 27 from different groups with ( Al/ Al)0 ratios predicted by thermal modeling calculations 26 27 (Sugiura and Fujiya, 2014; Doyle et al., 2015). The ( Al/ Al)0 ratios from the models were determined from 26Al abundance required for an asteroid of 50 km in radius to achieve a peak metamorphic temperature, determined for each chondrite group, at least in the center of the asteroid. The peak metamorphic temperatures of the ordinary, CO, CV, and CR parent bodies are estimated to be ~950 C, ~500–600 C, ~500–600 C, and ~100 C, respectively (e.g., Huss et al., 2006 and references therein; Cody et al., 2008; Jilly et al., 2014). The predicted 26 27 26 27 ( Al/ Al)0 ratios are generally comparable to or slightly lower than the inferred ( Al/ Al)0 of chondrules measured in each chondrite group, except for CR and CH chondrites, suggest- ing a rapid accretion of parent asteroid after the latest chondrule-forming event. This is consistent with theoretical models of accretion through turbulent concentration (Cuzzi et al., 2010) and streaming and gravitational instability in a dusty midplane (Youdin and Shu, 2002; Johansen and Klahr, 2011), and the high dust/gas ratios inferred for the regions where chondrules formed (e.g., Alexander et al., 2008). It may also explain the unique chemical, isotopic, and mineralogical characteristics of known chondrite groups (e.g., Krot et al., 2014a), and the W and Mo isotopic complementarity of chondrules and matrix in Allende 26 =27 Internal (Budde et al., 2016a, 2016b). Many CR and CH chondrule Al Al 0 ratios are lower 26Al–26Mg Systematics of Chondrules 267 than those predicted by the models. This might be due to the higher peak temperatures (~240 C and ~260 C for CR and CH parent asteroids, respectively, Busemann et al., 2007) used for the model by Sugiura and Fujiya (2014). The paleomagnetic records in the Allende and Kaba chondrites are interpreted as evi- dence for the existence of molten core in the CV asteroid (Elkins-Tanton et al., 2011; Carporzen et al., 2011; Gattacceca et al., 2013). To reach melting temperature (~1,250 C), aCV-likeasteroid>50 km in radius must have accreted within the first 1.5 million years of the solar system formation, which, assuming uniform distribution of 26Al/27Al in the 26 27 > 5 protoplanetary disk, corresponds to ( Al/ Al)0 of 1.2 10 (e.g., Elkins-Tanton 26 =27 Internal 26 27 et al., 2011). The Al Al 0 ratios of Kaba chondrules suggest that Al/ Al available for the CV asteroid was <5 106 and therefore too low to melt the CV asteroid even if we assume its instantaneous accretion, right after chondrule formation. This contra- dicts the existence of a molten core in the CV asteroid and appears to support an alternative interpretation of paleomagnetic records in CV chondrites as being due to the short-duration magnetic field generated by other mechanisms, e.g., impacts (Bland et al., 2014) or by a young Sun (Tarduno et al., 2016). Alternatively, the data might indicate that the chondritic materials accreted on an earlier formed differentiated body, as hypothesized by Elkins- Tanton et al. (2011) and Sanders and Scott (2012).

9.6 Summary and Unresolved Questions

The 26Al–26Mg systematics of chondrules inferred from in situ analyses of chondrule minerals could constrain 26Al/27Al ratios when chondrules crystallized. However, only a handful of chondrules likely have preserved 26Al/27Al ratios undisturbed because metasomatic alteration/ thermal metamorphism experienced by chondritic asteroids could modify Al and Mg isotopic compositions in some mineral phases. Therefore, it is important to limit existing 26Al–26Mg data in chondrules to those from the least metamorphosed chondrites. The possible best data sets 26 27 from the existing ( Al/ Al)0 ratios in chondrules are from LL3.00, CO3.05, Acfer 094 (ungrouped C3.00), CV3.1oxB, mildly aqueously altered CR23, and CH3 chondrites. Their 5 26 27 values range from unresolved to ~1.2 10 , and thus no chondrules have ( Al/ Al)0 ratios 5 26 27 corresponding to the canonical level (~5.2 10 ). The ( Al/ Al)0 ratios of chondrules among the different chondrite groups appear to have different populations and decrease in the order of UOCs COs Acfer 094 CVs > CRs CHs. As it has been thought, with an assumption of homogeneous distribution of 26Al/27Al in the protoplanetary disk, the 26Al–26Mg systematics of chondrules suggests chondrule formation started ~1.5 million years after CAIs and lasted over a few million years. 26 26 26 27 The Al– Mg systematics of bulk chondrules could have recorded ( Al/ Al)0 ratios of chondrule precursors, instead of the last melting event that chondrules experienced. The existing data of the bulk chondrules are controversial; some suggest formation of chondrule precursors started contemporaneously with CAIs and lasted over ~1.5 million years, while some are consistent with their formation with reduced 26Al/27Al abundance of around ~(12) 10 5, suggesting heterogeneous distribution of 26Al/27Al in chondrule forming regions. Since the 268 Kazuhide Nagashima, Noriko T. Kita, and Tu-Han Luu latter interpretation is only from a single laboratory, studies of 26Al–26Mg systematics of bulk chondrules should be evaluated with the same analytical precision by other groups in order to confirm/reject the data supporting the heterogeneous distribution of 26Al/27Al. The correct interpretation may require considering complex Al/Mg fractionation during and post chondrule formation, possible Mg-isotope anomalies, and partial re-equilibration of Mg-isotopes due to interaction of chondrule melts with nebular gas. Therefore, careful complementary studies of petrography, mineralogy, and possibly other isotope systems (e.g., O, Si, and Cr-isotopes) of the chondrules should also be undertaken. In addition, as model 26Al/27Al ratios of bulk chondrules 26 rely on knowledge of the initial Mg-isotope composition (µ Mg*0) of the solar system, it is 26 important to determine with high precision whether solar system µ Mg*0 is –40 ppm or –16 ppm or another value. The homogeneous/heterogeneous distribution of 26Al/27Al in the protoplanetary disk was also assessed by comparisons of formation ages of CAIs, chondrules, and angrites determined by Al– Mg, Pb–Pb and Hf–W isotope systematics, and there are no agreements among them. There is a general consistency among the ages determined by Al–Mg and Hf–W systematics, supporting homogeneous 26Al/27Al distribution. In contrast, the ages determined by Al–Mg and Pb–Pb systematics are largely inconsistent as per chondrite group the range of chondrule ages determined by Al–Mg is only ~12 million years, while the Pb–Pb ages have a spread over ~4 million years, supporting 26Al heterogeneity. Since the Hf–W ages of chondrules are averages of thousands of chondrules, these ages represent dominant populations of chondrules but do not rule out the presence of old chondrules that Pb–Pb ages suggest. Therefore, the age comparison of individual chondrules determined by Al–Mg and Pb–Pb systematics may provide more critical evaluation. While the Al–Mg systematics of chondrules has been determined by many different groups, both Hf–WandPb–Pb ages of chondrules are from only a couple of laboratories. It is necessary to have independent evaluations of Pb–Pb and Hf–W ages by different groups before concluding about a homogenous/heterogeneous distribution of 26Al/27Al in the protoplanetary disk. Regardless the possibility of 26Al/27Al heterogeneity, initial 26Al/27Al abundances recorded in chondrules are still useful to constrain thermal history of chondritic asteroids. The values and 26 27 narrow ranges of ( Al/ Al)0 ratios of chondrules in each chondrite group are consistent with those required to explain thermal histories of parent asteroids, suggesting a rapid accretion of parent asteroid after the latest chondrule-forming event.

Acknowledgments

We thank Johan Villeneuve, Shogo Tachibana, and Alexander Krot for their constructive reviews, and the Editors for handling the manuscript. T-H Luu is funded by a grant from the European Research Council (ERC grant 321209 - ISONEB).

References

Alexander, C. M. O’D., Grossman, J. N., Ebel, D. S., and Ciesla, F. J. (2008). The formation conditions of chondrules and chondrites. Science, 320, 1617–1619. 26Al–26Mg Systematics of Chondrules 269

Amelin, Y. (2008a). U-Pb ages of angrites. Geochim. Cosmochim. Acta, 72, 221–232. Amelin, Y. (2008b). The U-Pb systematics of angrite Sahara 99555. Geochim. Cosmochim. Acta, 72, 4874–4885. Amelin, Y., Krot, A. N., Hutcheon, I. D., and Ulyanov, A. A. (2002). Lead isotopic ages of chondrules and calcium-aluminum-rich inclusions. Science, 297, 1678–1683. Amelin, Y., and Krot, A. N. (2007). Pb isotopic age of the Allende chondrules. Meteorit. Planet. Sci., 42, 1321–1335. Becker, M., Hezel, D. C., Schulz, T., Elfers, B.-M., and Münker, C. (2015). Formation timescales of CV chondrites from component specificHf–W systematics. Earth Planet Sci. Lett., 432, 472–482. Bizzarro, M., Baker, J. A., and Haack, H. (2004). Contemporaneous formation of chondrules and refractory inclusions. Nature, 431, 275–278. Bland, P. A., Collins, G. S., Davison, T. M., et al. (2014). Pressure–temperature evolution of primordial Solar System solids during impact- induced compaction. Nat. Commun., 5, 5451. Bollard, J., Connelly, J. N., and Bizzarro, M. (2014). The absolute chronology of the early Solar System revisited. Meteorit. Planet. Sci., 77, abstract #5234. Bollard, J., Kawasaki, N., Sakamoto, N., et al. (2015). Early disk dynamics inferred from isotope systematics of individual chondrules. Meteorit. Planet. Sci., 78, abstract #5211. Bonal, L., Quirico, E., Bourot-Denise, M., and Montagnac, G. (2006). Determination of the petrologic type of CV3 chondrites by Raman spectroscopy of included organic matter. Geochim. Cosmochim. Acta, 70, 1849–1863. Bouvier, A., Brennecka, G. A., and Wadhwa, M. (2011). Absolute chronology of the first solids in the Solar System. Workshop on Formation of the First Solids in the Solar System abstract #9054. Budde, G., Burkhardt, C., Brennecka, G. A., et al. (2016a). Molybdenum isotopic evidence for the origin of chondrules and a distinct genetic heritage of carbonaceous and non- carbonaceous meteorites. Earth Planet Sci. Lett., 454, 293–303. Budde, G., Kleine, T., Kruijer, T. S., Burkhardt, C., and Metzler, K. (2016b). Tungsten isotopic constraints on the age and origin of chondrules. Proc. Natl. Acad. Sci. USA, 113, 2886–2891. Budde, G., Kruijer, T. S., and Kleine, T. (2017). Hafnium-tungsten chronology of CR chon- drites. Lunar Planet. Sci., 48, abstract #1886. Busemann, H., Alexander, C. M. O’D., and Nittler, L. R. (2007). Characterization of insoluble organic matter in primitive meteorites by microRaman spectroscopy. Meteorit. Planet. Sci., 42, 1387–1416. Brearley, A. J., and Krot, A. N. (2012). Metasomatism in the early Solar System: The record from chondritic meteorites. In D. E. Harlov and H. Austrheim (Eds.), Metasomatism and the Chemical Transformation of Rock – Lecture Notes in Earth System Sciences, 659–789. New York NY: Springer-Verlag. Brennecka, G. A., Budde, G., and Kleine, T. (2015). Uranium isotopic composition and absolute ages of Allende chondrules. Meteorit. Planet Sci., 50, 1995–2002. Brennecka, G. A., and Wadhwa, M. (2011). Uranium isotope compositions of mineral separates from a single refractory inclusion. 74th Meteorit. Soc. Meet., abstract #5030. Carporzen, L., Weiss, B. P., Elkins-Tanton, L. T., et al. (2011). Magnetic evidence for a partially differentiated carbonaceous chondrite parent body. Proc. Natl. Acad. Sci. USA, 108, 6386–6389. Cherniak, D. J. (2010). Cation diffusion in feldspars. Rev. Min. Geochem., 72, 691–733. Claydon, J. L., Lai, Y.-J., Coath, C. D., et al. (2014). The Al-Mg system in chondrules: A two- pronged approach. Meteorit. Planet. Sci., 77, abstract #5164. 270 Kazuhide Nagashima, Noriko T. Kita, and Tu-Han Luu

Claydon, J. L., Elliott, T., Coath, C. D., et al. (2015). A chondrule from the Mokoia (CV3) chondrite with anomalously low 26Mg*: Evidence for a multi-stage history. Meteorit. Planet. Sci., 78, abstract #5250. Cody, G. D., Alexander, C. M. O’D., Yabuta, H., et al. (2008). Organic thermometry for chondritic parent bodies. Earth Planet. Sci. Lett., 272, 446–455. Connelly, J. N., Amelin, Y., Krot, A. N., and Bizzarro, M. (2008). Chronology of the Solar System’s oldest solids. Astrophys. J. Lett., 675, L121. Connelly, J. N., and Bizzarro, M. (2009). Pb–Pb dating of chondrules from CV chondrites by progressive dissolution. Chem. Geol., 259, 143–151. Connelly, J. N., Bizzarro, M., Krot, A. N., et al. (2012). The absolute chronology and thermal processing of solids in the solar protoplanetary disk. Science, 338, 651–655. Connelly, J. N., Bollard, J., and Bizzarro, M. (2017). Pb-Pb chronometry and the early Solar System. Geochim. Cosmochim. Acta, 201, 345–363. Cuzzi, J. N., Hogan, R. C., and Bottke, W. F. (2010). Towards initial mass functions for asteroids and Kuiper Belt Objects. Icarus, 208, 518–538. Doyle, P. M., Jogo, K., Nagashima, K., et al. (2015). Early aqueous activity on the ordinary and carbonaceous chondrite parent bodies recorded by fayalite. Nat. Commun., 6, 1–10. Elkins-Tanton, L. T., Weiss, B. P., and Zuber, M. T. (2011). Chondrites as samples of differentiated planetesimals. Earth Planet. Sci. Lett., 305,1–10. Faak, K., Chakraborty, S., and Coogan, L. A. (2013). Mg in plagioclase: Experimental calibra- tion of a new geothermometer and diffusion coefficients. Geochim. Cosmochim. Acta, 123, 195–217. Galy, A., Young, E. D., Ash, R. D., and O’nions, R. K. (2000). The formation of chondrules at high gas pressures in the solar nebula. Science, 290, 1751–1753. Gattacceca, J., Weiss, B. P., Gounelle, M., Lima, E. A., and Rochette, P. (2013). More evidence for a partially differentiated CV parent body from the meteorite Kaba. Lunar Planet. Sci., 44, abstract #1721. Grossman, J. N., and Brearley, A. J. (2005). The onset of metamorphism in ordinary and carbonaceous chondrites. Meteorit. Planet. Sci., 40,87–122. Hayakawa, A., Fukuda, K., Iizuka, T., and Hiyagon, H. (2017). High precision magnesium isotopic measurements for CV chondrite CAIs and LL3.15 chondrite chondrules. Lunar Planet. Sci., 48, abstract #1923. Holst, J. C., Olsen, M. B., Paton, C., et al. (2013). 182Hf–182W age dating of a 26Al-poor inclusion and implications for the origin of short-lived radioisotopes in the early Solar System. Proc. Natl. Acad. Sci. USA, 110, 8819–8823. Huss, G. R., MacPherson, G. J., Wasserburg, G. J., Russell, S. S., and Srinivasan, G. (2001). Aluminum-26 in calcium-aluminum-rich inclusions and chondrules from unequilibrated ordinary chondrites. Meteorit. Planet. Sci., 36, 975–997. Huss, G. R., Rubin, A. E., and Grossman, J. N. (2006). Thermal Metamorphism in Chondrites. In D. Lauretta and H. McSween (Eds.), Meteorites and the Early Solar System II, 567–586. Tucson, AZ: University of Arizona Press. Hutcheon, I. D., and Hutchison, R. (1989). Evidence from the Semarkona ordinary chondrite for 26Al heating of small planets. Nature, 337, 238–241. Hutcheon, I. D., Marhas, K. K., Krot, A. N., Goswami, J. N., and Jones, R. H. (2009). 26Al in plagioclase-rich chondrules in carbonaceous chondrites: Evidence for an extended duration of chondrule formation. Geochim. Cosmochim. Acta, 73, 5080–5099. Huyskens, M. H., Yin, Q. -Z., Sanborn, M. E., et al. (2016). Possible uranium isotopic heterogeneity in Allende chondrules and its impact on Pb-Pb ages: A first case of U and Pb isotopes from a single chondrule. Lunar Planet. Sci., 47, abstract #2727. 26Al–26Mg Systematics of Chondrules 271

Ireland, T. R. (1990). Presolar isotopic and chemical signatures in hibonite-bearing refractory inclusions from the Murchison carbonaceous chondrite. Geochim. Cosmochim. Acta, 54, 3219–3237. Ito, M., and Messenger, S. (2010). Thermal metamorphic history of a Ca, Al-rich inclusion constrained by high spatial resolution Mg isotopic measurements with NanoSIMS 50L. Meteorit. Planet. Sci., 45, 583–595. Jacobsen, B., Yin, Q.-Z., Moynier, F., et al. (2008). 26Al-26Mg and 207Pb-206Pb systematics of Allende CAIs: Canonical solar initial 26Al/27Al ratio reinstated. Earth Planet. Sci. Lett., 272, 353–364. Jilly, C. E., Huss, G. R., Nagashima, K., and Schrader, D. L. (2014). Oxygen isotopes and geothermometry of secondary minerals in CR chondrite. Meteorit. Planet. Sci., 77, abstract #5395. Johansen, A., and Klahr, H. (2011). Planetesimal formation through streaming and gravitational instabilities. Earth, Moon, and Planets, 108,39–43. Kimura, M., and Ikeda, Y. (1998). Hydrous and anhydrous alterations of chondrules in Kaba and Mokoia CV chondrites. Meteorit. Planet. Sci., 33, 1139–1146. Kimura, M., Grossman, J. N., and Weisberg, M. K. (2008). Fe-Ni metal in primitive chondrites: Indicators of classification and metamorphic conditions for ordinary and CO chondrites. Meteorit. Planet. Sci., 43, 1161–1177. Kita, N. T., Huss, G. R., Tachibana, S., et al. (2005). Constraints on the origin of chondrules and CAIs from short-lived and long-lived radionucleides. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, 558–588. ASP Conf. Ser. 341. San Francisco, CA: Astronomical Society of the Pacific. Kita, N. T., Nagahara, H., Togashi, S., and Morshita, Y. (2000). A short duration of chondrule formation in the solar nebula: Evidence from 26Al in Semarkona ferromagnesian chon- drules. Geochim. Cosmochim. Acta, 64, 3913–3922. Kita, N. T., Tenner, T. J., Ushikubo, T., et al. (2015). Why do U-Pb ages of chondrules and CAIs have more spread than their 26Al ages? Meteorit. Planet. Sci., 78, abstract #5360. Kita, N. T., and Ushikubo, T. (2012). Evolution of protoplanetary disk inferred from 26Al chronology of individual chondrules. Meteorit. Planet. Sci., 47, 1108–1119. Kita, N. T., Yin, Q. -Z., MacPherson, G. J., et al. (2013). 26Al-26Mg isotope systematics of the first solids in the early Solar System. Meteorit. Planet. Sci., 48, 1383–1400. Kleine, T., Hans, U., Irving, A. J., and Bourdon, B. (2012). Chronology of the angrite parent body and implications for core formation in protoplanets. Geochim. Cosmochim. Acta, 84, 186–203. Kleine, T., Mezger, K., Palme, H., Scherer, E., and Münker, C. (2005). Early core formation in asteroids and late accretion of chondrite parent bodies: Evidence from 182Hf-182W in CAIs, metal-rich chondrites, and iron meteorites. Geochim. Cosmochim. Acta, 69, 5805–5818. Komatsu, M., Fagan, T. J., Mikouchi, T., and Yamaguchi, A. (2014). Alteration sequence of CV3 chondrites: Matrix textures and Raman spectroscopy. 5th Symp. Polar Science, National Inst. Polar Res. Abstract #370. Krot, A. N., Amelin, Y., Bland, P., et al. (2009). Origin and chronology of chondritic compon- ents: A review. Geochim. Cosmochim. Acta, 73, 4963–4997. Krot, A. N., Anders, M., Weisberg, M. K., and Keil, K. (2002). Invited Review: The CR chondrite clan: Implications for early Solar System processes. Meteorit. Planet. Sci., 37, 1451–1490. Krot, A. N., and Hutcheon, I. D. (1997). Highly oxidized and metamorphosed chondritic or igneous (?) clasts in the CV3 carbonaceous chondrite Mokoia: Excavated material from the interior of the CV3 asteroid or previously unsampled asteroid. Lunar Planet. Sci., 28, abstract #767. 272 Kazuhide Nagashima, Noriko T. Kita, and Tu-Han Luu

Krot, A. N., and Keil, K. (2002). Anorthite-rich chondrules in CR and CH carbonaceous chondrites: Genetic link between Ca,Al-rich inclusions and ferromagnesian chondrules. Meteorit. Planet. Sci., 37,91–111. Krot, A. N., Keil, K., Goodrich, C., Weisberg, M. K., and Scott, E. R. D. (2014a). Classification of meteorites. In A. M. Davis (Ed.), Meteorites and Cosmochemical Processes.InH.D. Holland and K. K. Turekian (Eds.), Treatise on Geochemistry,1,1–63. Oxford, UK: Elsevier. Krot, A. N., and Nagashima, K. (2017). Constraints on mechanisms of chondrule formation from chondrule precursors and chronology of transient heating events in the protoplanetary disk. Geochem. J., 51,45–68. Krot, A. N., Nagashima, K., Bizzarro, M. (2014b). Aluminum-magnesium isotope systematics of porphyritic chondrules and plagioclase fragments in CH carbonaceous chondrites. Lunar Planet. Sci., 45, abstract #2142. Krot, A. N., Nagashima, K., Bizzarro, M., et al. (2008). Multiple generations of refractory inclusions in the metal-rich carbonaceous chondrites Acfer 182/214 and Isheyevo. Astro- phys. J. 672, 713–721. Krot, A. N., Petaev, M. I., Scott, E. R. D., et al. (1998a). Progressive alteration in CV3 chondrites: More evidence for asteroidal alteration. Meteorit. Planet. Sci., 33, 1065–1085. Krot, A. N., Petaev, M. I., Zolensky, M. E., et al. (1998b). Secondary calcium-iron-rich minerals in the Bali-like and Allende-like oxidized CV3 chondrites and Allende dark inclusions. Meteorit. Planet. Sci., 33, 623–645. Kruijer, T. S., Kleine, T., Fischer-Godde, M., Burkhardt, C., and Wieler, R. (2014). Nucleosyn- thetic W isotope anomalies and the Hf–W chronometry of Ca–Al-rich inclusions. Earth Planet. Sci. Lett., 403, 317–327. Kunihiro, T., Rubin, A. E., McKeegan, K. D., and Wasson, J. T. (2004). Initial 26Al/27Al in carbonaceous-chondrite chondrules: Too little 26Al to melt asteroids. Geochim. Cosmo- chim. Acta, 68, 2947–2957. Kurahashi, E., Kita, N. T., Nagahara, H., and Morishita, Y. (2008). Al-26-Mg-26 systematics and petrological study of chondrules in CR chondrites. Geochim. Cosmochim. Acta, 72, abstract #504. Kurahashi, E., Kita, N. T., Nagahara, H., and Morishita, Y. (2008). 26Al-26Mg systematics of chondrules in a primitive CO chondrite. Geochim. Cosmochim. Acta, 72, 3865–3882. Larsen, K. K., Trinquier, A., Paton, C., et al. (2011). Evidence for magnesium isotope heterogeneity in the solar protoplanetary disk. Astrophys. J., 735, L37. LaTourrette, T., and Wasserburg, G. J. (1998). Mg diffusion in anorthite: Implications for the formation of early Solar System planetesimals. Earth Planet. Sci. Lett., 158,91–108. Lee, T., Papanastassiou, D. A., and Wasserburg, G. J. (1976). Demonstration of 26Mg excess in Allende and evidence for 26Al. Geophys. Res. Lett., 3 (1), 41–44. Libourel, G., Krot, A. N., and Tissandier, L. (2006). Role of gas-melt interaction during chondrule formation. Earth Planet. Sci. Lett., 251, 232–240. Lugaro, M., Heger, A., Osrin, D., et al. (2014). Stellar origin of the 182Hf cosmochronometer and the presolar history of Solar System matter. Science, 345, 650–653. Luu, T. -H., Young, E. D., Gounelle, M., and Chaussidon, M. (2015). Short time interval for condensation of high-temperature silicates in the solar accretion disk. Proc. Natl. Acad. Sci. USA, 112, 1298–1303. Luu, T. -H., Hin, R. C., Coath, C. D., and Elliott, T. (2016). High precision mg-isotope measurements of bulk chondrites and the homogeneity of 26Al in the solar nebula. Meteorit. Planet. Sci., 79, abstract #6485. MacPherson, G. J., Bullock, E. S., Janney, P. E., et al. (2010). Early solar nebula condensates with canonical, not supracanonical, initial 26Al/27Al ratios. Astrophys. J., 711 L117–L121. 26Al–26Mg Systematics of Chondrules 273

MacPherson, G. J., Davis, A. M., and Zinner, E. K. (1995). The Distribution of 26Al in the early Solar-System – A reappraisal. Meteoritics, 30, 365–386. MacPherson, G. J., Kita, N. T., Ushikubo, T., Bullock, E. S., and Davis, A. M. (2012). Well- resolved variations in the formation ages for Ca-Al-rich inclusions in the early Solar System. Earth Planet. Sci. Lett., 331,43–54. McKeegan, K. D., Greenwood, J. P., Leshin, L. A., and Cosarinsky, M. (2000). Abundance of 26Al in ferromagnesian chondrules of unequilibrated ordinary chondrites. Lunar Planet. Sci., 31, abstract #2009. Makide, K., Nagashima, K., Krot, A. N., et al. (2009). Oxygen- and magnesium-isotope compositions of calcium-aluminum-rich inclusions from CR2 carbonaceous chondrites. Geochim. Cosmochim. Acta, 73, 5018–5050. Mostefaoui, S., Kita, N. T., Togashi, S., et al. (2002). The relative formation ages of ferro- magnesian chondrules inferred from their initial aluminum-26/aluminum-27 ratios. Meteorit. Planet. Sci., 37, 421–438. Nagahara, H., Kita, N. T., Ozawa, K., and Morishita, Y. (2008). Condensation of major elements during chondrule formation and its implication to the origin of chondrules. Geochim. Cosmochim. Acta, 72, 1442–1465. Nagashima, K., Krot, A. N., and Huss, G. R. (2014). 26Al in chondrules from CR2 chondrites. Geochem. J., 48, 561–570. Nagashima, K., Krot, A. N., and Komatsu, M. (2017). 26Al-26Mg systematics in chondrules from Kaba and Yamato 980145 CV3 carbonaceous chondrites. Geochim. Cosmochim. Acta, 201, 303–319. Nakashima, D., Ishida, H., Tenner, T. J., Kita, N. T., and Nakamura, T. (2016). Al-Mg chronology of chondrules in the RBT 04143 CV3 chondrite. Goldschmidt Abstracts, 2016 #2236. Nishiizumi, K. (2004). Preparation of 26Al AMS standards. Nucl. Instr. Meth. Phys. Res. B., 223–224, 388–392. Norris, T. L., Gancarz, A. J., Rokop, D. J., and Thomas, K. W. (1983). Half-life of 26Al. J. Geophys. Res., 88, B331–B333. Olsen, M. B., Schiller, M., Krot, A. N., and Bizzarro, M. (2013). Magnesium isotope evidence for single stage formation of CB chondrules by colliding planetesimals. Astrophys. J., 776, L1. Olsen, M. B., Wielandt, D., Schiller, M., van Kooten, E. M. M. E., and Bizzarro, M. (2016). Magnesium and 54Cr isotope compositions of carbonaceous chondrite chondrules – Insights into early disk processes. Geochim. Cosmochim. Acta, 191, 118–138. Pringle, E. A., Savage, P. S., Jackson, M. G., Barrat, J.-A., and Moynier, F. (2013). Si isotope homogeneity of the solar nebula. Astrophys. J., 779, L123. Reisener, R., Meibom, A., Krot, A. N., Goldstein, H. I., and Keil, K. (2000). Microstructure of condensate Fe-Ni metal particles in the CH chondrite PAT 91546. Lunar Planet. Sci. 31, abstract #1445. Rudraswami, N. G., Goswami, J. N., Chattopadhyay, B., Sengupta, S. K., and Thapliyal, A. P. (2008). Al-26 records in chondrules from unequilibrated ordinary chondrites: II. Duration of chondrule formation and parent body thermal metamorphism. Earth Planet. Sci. Lett., 274,93–102. Sanders, I. S., and Scott, E. R. D. (2012). The origin of chondrules and chondrites: Debris from low-velocity impacts between molten planetesimals? Meteorit. Planet. Sci., 47, 2170–2192. Sano, Y., Takada, M., Takahata, N., Fujiya, W., and Sugiura, N. (2014). Ion microprobe Al-Mg dating of single plagioclase grains in an Efremovka chondrule. Geochem. J., 48, 133–144. Schrader, D. L., Nagashima, K., Krot, A. N., et al. (2017). Distribution of 26Al in the CR chondrite chondrule-forming region of the protoplanetary disk. Geochim. Cosmochim. Acta, 201, 275–302. 274 Kazuhide Nagashima, Noriko T. Kita, and Tu-Han Luu

Schiller, M., Baker, J. A., and Bizzarro, M. (2010). 26Al–26Mg dating of asteroidal magmatism in the young Solar System. Geochim. Cosmochim. Acta, 74, 4844–4864. Schiller, M., Connelly, J. N., Aslaug, G. C., Mikouchi, T., and Bizzarro, M. (2015). Early accretion of protoplanets inferred from a reduced inner Solar System 26Al inventory. Earth Planet. Sci. Lett., 420,45–54. Spivak-Birndorf, L., Wadhwa, M., and Janney, P. (2009). 26Al–26Mg systematics in D’Orbigny and Sahara 99555 angrites: Implications for high-resolution chronology using extinct chronometers. Geochim. Cosmochim. Acta, 73, 5202–5211. Srinivasan, G., Huss, G. R., and Wasserburg, G. J. (2000). A petrographic, chemical, and isotopic study of calcium-aluminum-rich inclusions and aluminum-rich chondrules from the Axtell (CV3) chondrite. Meteorit. Planet. Sci., 35, 1333–1354. Sugiura, N. and Fujiya, W. (2014). Correlated accretion ages and epsilon Cr-54 of meteorite parent bodies and the evolution of the solar nebula. Meteorit. Planet. Sci., 49, 772–787. Tachibana, S., Nagahara, H., Mostefaoui, S., and Kita, N. T. (2003). Correlation between relative ages inferred from 26Al and bulk compositions of ferromagnesian chondrules in least equilibrated ordinary chondrites. Meteorit. Planet. Sci., 38, 939–962. Tarduno, J. A., O’Brien, T. M., and Smirnov, A. V. (2016). Does the magnetization of CV meteorites record a parent body core dynamo? Lunar Planet. Sci., 47, abstract #2609. Tenner, T. J., Ushikubo, T., Nakashima, D., Kita, N. T., and Weisberg, M. K. (2013). 26Al in chondrules from the CR3.0 chondrite Queen Alexandra Range 99177: A link with O isotopes. Lunar Planet. Sci., 44, abstract #2010. Tenner, T. J., Ushikubo, T., Nakashima, D., et al. (2014). Silica excess in anorthitic plagioclase from type 3.00 chondrite chondrules: Evidence for retaining primary 26Al-26Mg systemat- ics. Lunar Planet. Sci., 45, abstract #1187. Tissandier, L., Libourel, G., and Robert, F. (2002). Gas-melt interactions and their bearing on chondrule formation. Meteorit. Planet. Sci., 37, 1377–1389. Trinquier, A., Elliott, T., Ulfbeck, D., et al. (2009). Origin of nucleosynthetic isotope hetero- geneity in the solar protoplanetary disk. Science, 324, 374–376. Ushikubo, T., Nakashima, D., Kimura, M., Tenner, T. J., and Kita, N. T. (2013). Contempor- aneous formation of chondrules in distinct oxygen isotope reservoirs. Geochim. Cosmo- chim. Acta, 109, 280–295. van Kooten, E. M. M. E., Wielandt, D., Schiller, M., et al. (2016). Isotopic evidence for primordial molecular cloud material in metal-rich carbonaceous chondrites. Proc. Natl. Acad. Sci. USA, 113, 2011–2016. Van Orman, J. A., Cherniak, D. J., and Kita, N. T. (2014). Magnesium diffusion in plagioclase: Dependence on composition, and implications for thermal resetting of the 26Al–26Mg early Solar System chronometer. Earth Planet. Sci. Lett., 385,79–88. Villeneuve, J., Chaussidon, M., and Libourel, G. (2009). Homogeneous distribution of Al-26 in the Solar System from the Mg isotopic composition of chondrules. Science, 325, 985–988. Wasserburg, G. J., Lee, T., and Papanastassiou, D. A. (1977). Correlated O and Mg isotopic anomalies in Allende inclusions: II. Magnesium. Geophys. Res. Lett., 4, 299–302. Wasserburg, G. J., Wimpenny, J., and Yin, Q. -Z. (2012). Mg isotopic heterogeneity, Al-Mg isochrons, and canonical 26Al/27Al in the early Solar System. Meteorit. Planet. Sci., 47, 1980–1997. Weisberg, M. K., Prinz, M., Clayton, R. N., et al. (1995). The CR chondrite clan. Proc. NIPR Symposium, 8,11–32. Weisberg, M. K., Prinz, M., Clayton, R. N., and Mayeda, T. K. (1997). CV3 chondrites: Three subgroups, not two. Meteorit. Planet. Sci., 32, abstract #139. Youdin, A. N., and Shu, F. (2002). Planetesimal formation by gravitational instability. Astro- phys. J., 580, 494–505. 26Al–26Mg Systematics of Chondrules 275

Yurimoto, H., Koike, O., Nagahara, H., Morioka, M., and Nagasawa, H. (2000). Heterogeneous distribution of Mg isotopes in anorthite single crystal from Type-B CAIs in Allende meteorite. Lunar Planet. Sci., 31, abstract #1593. Yurimoto, H., and Wasson, J. T. (2002). Extremely rapid cooling of a carbonaceous-chondrite chondrule containing very 16O-rich olivine and a 26Mg-excess. Geochim. Cosmochim. Acta, 66, 4355–4363. Zolotov M. Y., Mironenko M. V., and Shock E. L. (2006). Thermodynamic constraints on fayalite formation on parent bodies of chondrites. Meteorit. Planet. Sci. 41, 1775–1796. 10 Tungsten Isotopes and the Origin of Chondrules and Chondrites

thorsten kleine, gerrit budde, jan l. hellmann, thomas s. kruijer, and christoph burkhardt

Abstract

Chondrules and matrix from carbonaceous chondrites exhibit complementary nucleosynthetic W isotope anomalies that result from the depletion of a metallic s-process carrier in the chondrules, and the enrichment of this carrier in the matrix. The complementarity is difficult to reconcile with an origin of chondrules in protoplanetary impacts and also with models in which chondrules and matrix formed independently of each other in distinct regions of the disk. Instead, the complementarity indicates that chondrules formed by localized melting of dust aggregates in the solar nebula. The Hf–W ages for metal-silicate fractionation in CV and CR chondrites are 2.2 0.8 Ma and 3.6 0.6 Ma after formation of Ca-Al-rich inclusions, and are indistinguishable from Al–Mg ages for CV and CR chondrules. The good agreement between these ages strongly suggests that 26Al was homogeneously distributed in the solar protoplane- tary disk and that therefore Al–Mg ages are chronologically meaningful. The concordant Al–Mg and Hf–W ages reveal that chondrule formation (as dated by Al–Mg) was associated with metal- silicate fractionation (as dated by Hf–W), both within a given chondrite but also among the different subgroups of ordinary chondrites. These age data indicate that chondrules from a given chondrite group formed in a narrow time interval of <1 Ma, and that chondrule formation and chondrite accretion were closely linked in time and space. The rapid accretion of chondrules into a chondrite parent body is consistent with the isotopic complementarity, which requires that neither chondrules nor matrix were lost prior to chondrite accretion. Combined, these observations suggest that chondrule formation was an important step in the accretion of planetesimals.

10.1 Introduction

Chondrules are the major component of most chondrites, and judging by their abundance, must have formed by one of the most common processes in the early solar system. The nature and timing of this process remain poorly understood, however. The different formation hypotheses that have been proposed for the origin of chondrules can be divided into two general categories (e.g., Connolly and Desch, 2004): those in which chondrules formed by melting of dust in the solar protoplanetary disk (the solar nebula) and those in which chondrules formed during protoplanetary impacts. Distinguishing between these distinct models is important not only for understanding the origin of chondrules, but also for assessing the role chondrule formation might have played in planet formation. If chondrules formed by melting of dust aggregates

276 Tungsten Isotopes and the Origin of Chondrules and Chondrites 277 within the disk, followed by rapid accretion into a parent body, their formation may have been important for mediating the accumulation of dust into planetesimals. If, however, chondrules formed by protoplanetary impacts, then they would merely be a by-product of planet formation. Studies aimed at identifying the chondrule-forming process typically focus on reproducing the properties of chondrules, but often ignore the other major components of chondrites, namely fine-grained matrix and metal. However, these three components seem to be genetically linked. For instance, chondrules and matrix in carbonaceous chondrites have complementary chemical compositions, suggesting that chondrules and matrix from a given chondrite group formed from a single reservoir of precursor dust (Bland et al., 2005; Hezel and Palme, 2008; Hezel and Palme, 2010; Palme et al. 2015; Ebel et al., 2016; Chapter 4). Further, metal in primitive chondrites predominantly seems to have been reprocessed or newly formed during chondrule formation (Zanda et al., 1994; Connolly et al. 2001; Campbell et al., 2005), meaning that chondrules and metal are probably also genetically linked. Thus, any successful model for the origin of chondrules must not only identify the chondrule-forming process, but must also account for the formation of matrix and metal, and the subsequent accretion of all three components to a chondrite parent body. In this chapter, we review how the W isotopic compositions of chondrules, matrix, and metal from primitive (unmetamorphosed) chondrites help to shape models of chondrule formation and chondrite accretion. In particular, we will discuss how the W isotopic data can distinguish between an impact and solar nebula origin of chondrules, and review the Hf–W chronology of chondrule formation and metal-silicate fractionation in chondrites. Finally, we compare the Hf-W ages to results obtained from other chronometers, with the ultimate goal to establish a coherent timescale of chondrule formation and chondrite parent body accretion.

10.2 Tungsten Isotope Systematics

Tungsten isotope variations can be nucleosynthetic, radiogenic, or cosmogenic in origin (Kleine and Walker, 2017). Cosmogenic W isotope variations result from the interaction with cosmic rays and predominantly occur in iron meteorites and lunar samples. For chondrites and their components, cosmogenic effects on W isotopes are not important and will therefore not be discussed here. Nucleosynthetic W isotope variations arise through the heterogeneous distribu- tion of isotopically distinct presolar materials, and are powerful tracers to assess genetic links between meteoritic and planetary materials. Radiogenic W isotope variations result from the decay of short-lived 182Hf to 182W with a half-life of 8.9 million years (Ma) and form the basis of the Hf–W chronometer to date early solar system processes.

10.2.1 Nucleosynthetic W Isotope Variations

The heterogeneous distribution of isotopically distinct presolar materials leads to correlated variations of 182W/184W and 183W/184W (Burkhardt et al., 2012; Figure 10.1). Such nucleosyn- thetic W isotope variations can be used to assess genetic links between meteoritic materials, because the heterogeneous distribution of presolar material between two genetically linked components leads to complementary W isotope anomalies relative to the bulk material. 278 Thorsten Kleine, et al.

2 2

182 1 Hf decay 1 182Hf decay

0 0

–1 –1 s-process deficit W (6/4) W (6/3) s-process excess 182 182 e –2 e –2 s-process deficit –3 182 –3 182 s-process excess Initial e W of the Initial e W of the solar system solar system –4 –4

–1–0.50 0.5 1 –1 –0.5 00.51 e183W (6/4) e184W (6/3)

Figure 10.1 Schematic illustration of W isotope variations. Tungsten isotope ratios are shown in ε-units, corresponding to the deviation from terrestrial standard values in part per 10,000. (6/4) and (6/3) stand for normalization of the W isotope ratios to a fixed 186W/184Wor186W/183W, respectively.

Nucleosynthetic isotope variations predominantly affect W isotope ratios involving 184W, and so measurements of 183W/184W (when internally normalized to 186W/184W) are ideally suited to quantify the extent of nucleosynthetic W isotope heterogeneity. By contrast, the 182W/183W ratio (when internally normalized to 186W/183W) shows only small nucleosynthetic variations (Figure 10.1). This is because 184W has the largest contribution from the s-process of stellar nucleosynthesis, while 182W, 183W and 186W have similar s-process contributions (Arlandini et al., 1999; Qin et al., 2008b; Burkhardt et al., 2012). Thus, the heterogeneous distribution of s-process matter affects 182W, 183W, and 186W almost equally, resulting in only small changes of the 182W/183W ratio. These small changes can be reliably corrected, meaning that the presence of nucleosynthetic W isotope anomalies does not compromise the use of the 182Hf–182W chronometer.

10.2.2 182Hf–182W Chronometer

The 182Hf–182W chronometer can be used for dating processes that occurred in the first ~60 Ma of the Solar System (Kleine et al., 2009; Kleine and Walker, 2017). There are several properties that distinguish the Hf–W system from other chronometers typically utilized for dating chon- drule formation, such as the 26Al–26Mg and 207Pb–206Pb systems. First, the Hf–W system does not directly provide crystallization ages for individual chondrules, but instead provides the timing of Hf–W fractionation among chondrite components. As Hf is lithophile, but W siderophile, the dominant process fractionating Hf from W is metal-silicate fractionation. Chondrules are typically enriched in refractory lithophile and depleted in siderophile elements, whereas matrix displays the complementary depletions and enrichments (e.g., Palme et al., 2015). Moreover, metal in primitive chondrites predominantly seems to have formed or was reprocessed during chondrule formation (e.g., Campbell et al., 2005). These observations combined reveal that chondrules, matrix, and metal have different Hf/W ratios, most likely as a result of metal-silicate fractionation during chondrule formation, making it possible to date chondrule formation using the Hf–W system. Tungsten Isotopes and the Origin of Chondrules and Chondrites 279

Another important property of the Hf–W system is its high closure temperature of typically 800–900 C (Kleine et al., 2008), making it more robust against thermal resetting than the Al–Mg and Pb–Pb systems (e.g., the closure temperature for Al–Mg in anorthite is <600 C; Van Orman et al., 2014). Finally, 182Hf was homogeneously distributed within the solar protoplanetary disk, and so unlike for 26Al there is no doubt that variations in initial 182Hf/180Hf ratios reflect time differences rather than distinct formation locations within the disk (Kleine et al., 2009). AHf–W isochron for chondrites is typically defined by data for chondrule or silicate fractions and matrix or metal. The age of a sample is then obtained from the slope of the isochron, which provides the 182Hf/180Hf at the time of Hf–W closure. The Hf–W ages are expressed as the time elapsed since formation of Ca-Al-rich inclusions (CAIs), generally thought to represent the first solids formed in the solar system: ! 1 182Hf=180Hf t ¼ ln sample CAI λ 182 =180 Hf HfCAI where the initial 182Hf/180Hf of CAIs is (1.018 0.043) 10‒4,asdefined by a Hf–W isochron ‒ for CAIs (Burkhardt et al., 2008; Kruijer et al., 2014a), and λ = 0.078 0.002 Ma 1 (Vock- enhuber et al., 2004). In summary, W isotope variations among chondrules and other components of primitive chondrites can be used not only for dating the formation of chondrite components, but also for assessing potential genetic links among them. This combination of constraining both the genetic heritage and age of a sample makes W isotopes a uniquely powerful tool for investigating the origin of chondrules and chondrites.

10.3 Isotopic Complementarity and the Origin of Chondrules 10.3.1 Complementary Nucleosynthetic Isotope Anomalies in Chondrules and Matrix

Chondrule and matrix samples from the Allende CV3 chondrite exhibit nucleosynthetic 183W anomalies that are indicative of an s-process deficit (or r-excess) in chondrules and a complementary s-excess (or r-deficit) in matrix (Figure 10.2) (Budde et al., 2016b). The same Allende chondrules and matrix samples also exhibit complementary Mo isotope anomalies1 (Figure 10.3) (Budde et al., 2016a). Of note, the Mo and W isotope anomalies are broadly correlated, suggesting strongly that they result from the heterogeneous distribution of the same presolar carrier. Unlike for W, the Mo isotopic data can be used to distinguish between s- and r-process heterogeneity and, as such, make it possible to determine whether a presolar carrier is enriched in matrix and depleted in chondrules or vice versa. As the Mo isotope variations

1 Complementarity is strictly defined as complementary compositions of two components relative to a well-defined bulk composition (see Chapter 4). The prerequisite of a well-defined bulk composition is met for 183W, because bulk meteorites as well as the Earth, Mars, and Moon have very similar 183W compositions, in contrast to the large 183W anomalies observed for chondrules and matrix. For Mo isotopes, the situation is different because bulk meteorites exhibit large Mo isotope variations. Thus, although the Mo isotope anomalies in chondrules and matrix reflect the complementary depletion and enrichment of a single s-process carrier, these anomalies are not complementary relative to a common bulk composition. 280 Thorsten Kleine, et al.

Earth -deficit s-excess Moon s Mars Bulk Angrites planetary bodies Eucrites Ureilites Magmatic irons Chondrites Matrix Chondrules

Allende Bulk

–3–2 –1 0 1 2 3 e183W

Figure 10.2 Complementary ε183W compositions for Allende chondrules and matrix. Bulk meteorites show no or only small ε183W anomalies. W isotopic data from Budde et al. (2016b) and references therein.

Allende CV3 chondrite r-process excess 4 Chondrules Bulk Matrix s-process 2 deficit Mo 95 e

0

–2 –4–2 0 246 e94Mo

Figure 10.3 Complementary Mo isotopic anomalies for Allende chondrules and matrix. Data are given in ε-units, as the parts per 10,000 deviations of 95Mo/96Mo and 94Mo/96Mo ratios from terrestrial standard values. The chondrule and matrix samples plot along a mixing line (dashed line) between presumed s-process enriched and depleted end member compositions; the heterogeneous distribution of r-process matter would have resulted in different Mo isotope anomalies, as indicated by the dotted line. Data from Budde et al. (2016a). among chondrules and matrix reflect the heterogeneous distribution of an s-process carrier (Figure 10.3), these data therefore indicate that chondrules are depleted, whereas matrix is enriched in this carrier (Budde et al., 2016a). By contrast, the heterogeneous distribution of an r-process carrier would have led to different Mo isotope signatures, most notably larger 95Mo anomalies for a given 94Mo anomaly (Figure 10.3). It is noteworthy that the different chondrule separates display variable Mo and W isotope anomalies (Figures 10.2 and 10.3), indicating variable depletions of the s-process carrier in the chondrules. As each of these separates consists of several hundreds to thousands of chondrules, the isotopic variability among the chondrule separates also implies that there are even larger Mo and W isotope anomalies in individual chondrules. Tungsten Isotopes and the Origin of Chondrules and Chondrites 281

Unlike for W and Mo, the chondrule and matrix samples do not show complementary isotope anomalies for Ba (Budde et al., 2016a) and Ti (Gerber et al., 2017). Thus, isotopic comple- mentarity only seems to exist for siderophile (Mo, W), but not lithophile (Ba, Ti), elements. This, and the observation that the Mo (and W) isotope anomalies in different chondrule separates are correlated with magnetic susceptibility (and hence initial metal content), indicates that the nucleosynthetic isotope variations result from the heterogeneous distribution of a metallic presolar carrier (Budde et al., 2016a). Thus, Allende chondrules are variably depleted in a presolar metal enriched in s-process Mo and W nuclides, and this same carrier is enriched in the matrix. As discussed in Budde et al. (2016a, 2016b), several lines of evidence indicate that the Mo and W isotope anomalies are not the result of parent body alteration. First, the isotope anomalies cannot reflect the liberation and redistribution of Mo and W from isotopically anomalous CAIs, because most CAIs have smaller 183W anomalies than chondrules and because CAIs only exhibit 183W excesses but no deficits (Burkhardt et al., 2008; Kruijer et al., 2014a). As such, neither the large 183W excesses in chondrules nor the 183Wdeficits in matrix can be explained in this manner. Moreover, whereas the isotope anomalies in most CAIs result from an excess in r- process nuclides, the anomalies in chondrules and matrix reflect heterogeneity in an s-process carrier (note that r- and s-process heterogeneity results in very different Mo isotope patterns). Thus, the isotope anomalies observed in CAIs are of different origin than those observed in chondrules and matrix. Second, chondrules are characterized by a deficit in s-process isotopes, meaning that to account for this composition by parent body alteration would require the removal of a presolar s-process carrier from the solidified chondrules. However, as igneous objects, chondrules contain no presolar grains. Thus, the complementary isotope anomalies in chondrules and matrix cannot be the result of parent body processes, but must have been established during the formation of these two components.

10.3.2 Genetic Link between Chondrules and Matrix

The isotopic complementarity of chondrules and matrix implies that both components are genetically linked and must therefore have formed from the same reservoir of dust. This conclusion arises from several observations. First, the isotope anomalies reflect the heteroge- neous distribution of a single s-process carrier, which is most easily explained by a single process in which chondrules became depleted, and matrix enriched in the very same presolar s-process carrier. Second, in spite of the large 183W anomalies in chondrules and matrix, bulk meteorites, including Allende, show no or only very small 183W anomalies. Thus, chondrules and matrix are mixed in exactly the right proportions so that no 183W anomaly is produced in the bulk chondrite. It would take extraordinary circumstances to accomplish this by random mixing of chondrules and matrix that formed independently of each other. Instead, formation of chondrules and matrix with their large and complementary 183W anomalies from a single reservoir having the average 183W composition of the inner solar system (as defined by bulk meteorites, Earth, and Mars) naturally leads to a bulk chondrite with no 183W anomaly, provided that neither chondrules nor matrix were lost (Figure 10.2). Finally, the nucleosynthetic W (and Mo) isotope anomalies of chondrules and matrix are larger than those observed for any bulk meteorite (Figure 10.2). Moreover, the matrix is 282 Thorsten Kleine, et al. characterized by an excess in s-process nuclides, unlike the compositions observed for any bulk meteoritic or planetary material. As such, the nucleosynthetic isotope anomalies in chondrules and matrix cannot represent signatures of distinct formation regions in the disk2, but must reflect processes that occurred during chondrule formation itself.

10.3.3 Implications for the Origin of Chondrules

The isotopic complementarity of chondrules and matrix is a fundamental observation that is key for constraining the origin of chondrules. Any successful model for chondrule formation must therefore provide (i) a mechanism for generating the complementary isotope anomalies, (ii)a process by which chondrules and matrix derive from the same reservoir of dust, and (iii)a process by which chondrules and matrix were accreted into a parent body in their original proportions, that is in the correct proportion to not produce a 183W anomaly in the bulk chondrite. Below we evaluate to what extent existing models for the origin of chondrules are consistent with these requirements. Impact models for the origin of chondrules have great difficulty in meeting the constraints provided by the isotopic complementarity. In its original form, the impact model posits that chondrules originated as ‘splashes’ expelled by the collision of molten planetesimals (Zook, 1981; Asphaug et al., 2011; Sanders and Scott, 2012). Support for this model comes from the observation that the parent bodies of differentiated meteorites accreted earlier than the chondrite parent bodies (Kleine et al., 2005; Scherstén et al., 2006; Qin et al., 2008a; Kruijer et al., 2014b 2017), implying that differentiated and probably still molten bodies were present at the time of chondrule formation. However, chondrules exhibit larger nucleosynthetic Mo and W isotope anomalies than observed for bulk, differentiated meteorites (Figure 10.2). Thus, even though differentiated bodies were present, melt splashes from these bodies cannot represent the material from which chondrules ultimately formed. Moreover, the splashing model does not provide a mechanism for producing nucleosynthetic isotope anomalies in the matrix. In this model, matrix either represents primordial dust from the solar nebula or derives from the regolith of the colliding planetesimals (e.g., Sanders and Scott, 2012). However, in both cases the matrix would not be expected to have nucleosynthetic isotope anomalies relative to bulk meteorites and would also not be expected to exhibit complementary isotope anomalies relative to chondrules. In a recent modification of the splashing model, now termed the ‘dirty plume’ model (Chapter 14), it is proposed that within the impact plume the melt droplets mixed with primitive dust (containing presolar grains), and that during this mixing an s-process carrier was preferentially incorporated into the matrix and not included into the melt droplets that then became chondrules. In this model, the 183W excesses in the chondrules would reflect the contamination of isotopically normal melt droplets with primitive dust that is characterized by large 183W excesses. Such large excesses could in principle exist in presolar grains enriched in r-process nuclides. As argued above, however, the nucleosynthetic isotope anomalies in the

2 Note that Zanda et al. (Chapter 5) suggest that the isotopic complementarity indicates that chondrules and matrix formed in distinct areas of the disk. As outlined here, however, this cannot be correct. Tungsten Isotopes and the Origin of Chondrules and Chondrites 283 chondrules do not reflect addition of an r-process carrier but result from the removal of an s-process carrier. Moreover, mass balance indicates that 183W excesses of dust grains resulting from the removal of an s-process carrier would not be very large because only a very small fraction of the bulk W is removed. Consequently, large amounts of s-process- depleted dust grains would need to be added to the melt droplets to produce the 183W excesses in the final chondrules. Adding such large amounts of material to the melt droplets, while at the same time specifically excluding an s-process carrier from this material, seems highly unlikely. Other versions of the impact model invoke collisions of planetesimals that are either undifferentiated or maintained a primitive, unmelted crust (e.g., Chapter 13). In this model, chondrules originate from the melt produced by impact jetting. This model is, nonetheless, also difficult to reconcile with the constraints from isotopic complementarity. First, there is no a priori reason why the impact melting should have produced melt droplets that are variably depleted in a presolar s-process carrier. Second, and most importantly, these melt droplets would have to be remixed with primitive dust (i.e., the matrix) that contained the comple- mentary proportion of this s-process carrier to yield the original and nearly constant W isotopic composition of bulk chondrites and inner solar system planets (Figure 10.2). To meet this requirement, the melt droplets that ended up as chondrules in the final chondrite, and the source material from which these melt droplets originally derived (i.e., the material that became the matrix), must be sampled in an unbiased way, because otherwise the final chondrule-matrix mixture would not have the 183W composition of bulk inner solar system planets. However, it seems quite unlikely that the original proportion of chondrules and matrix could be retained throughout the impact melting and subsequent ejection and accretion of material. This example, as well as the example of the dirty plume model discussed above, illustrate that attempting to account for the isotopic complementarity of chondrules and matrix adds substantial complexity to impact models. The x-wind model for the origin of chondrules (Shu et al., 1996) (or any other model that invokes transport of chondrules over wide distances in the disk) also is difficult to reconcile with the isotopic complementarity of chondrules and matrix. In this model, chondrules originated near the Sun and were then transported outwards by several astronomical units (AU) to fall back onto the protoplanetary disk, where they were mixed with primitive dust (i.e., the matrix). However, the random mixing of chondrules and matrix predicted in the x- wind model is inconsistent with the observation that in spite of the large 183Wanomaliesin chondrules and matrix, bulk chondrites have no significant 183W anomalies (Figure 10.2). Moreover, the x-wind model provides no mechanism for generating isotope anomalies in chondrules and matrix, because these anomalies are unlike those observed for bulk meteorites and, hence, not a signature of distinct formation regions in the disk (see above). Of note, nucleosynthetic isotope anomalies in bulk meteorites reveal that material formed at greater heliocentric distance exhibits a larger s-process deficit than material formed closer to the Sun (Burkhardt et al., 2011; Render et al., 2017). In the x-windmodel,matrixshouldtherefore exhibit an s-process deficit compared to chondrules. However, matrix is characterized by an excess, not a deficit, in s-process matter. In contrast to the x-wind and impact models, an origin of chondrules by localized melting of dust aggregates in the solar nebula (e.g., Connolly and Desch, 2004) can readily account for the 284 Thorsten Kleine, et al. isotopic complementarity observed. This is because such models provide viable mechanisms for generating the complementary distribution of presolar grains between chondrules and matrix, and they also allow for keeping matrix and chondrules together. For instance, the comple- mentary isotope anomalies could have been produced by metal-silicate fractionation during chondrule formation, which may have resulted in the preferential incorporation of a presolar metal enriched in s-process matter into the chondrite matrix. Alternative options include the sorting of dust grains according to their size or magnetic properties. For instance, the enhanced sticking properties of magnetized Fe-metal bearing dust grains could lead to their settling to the midplane of the disk, preferentially removing metal from the precursor dust of chondrules (Hubbard, 2016a, 2016b). Finally, the melting temperature of refractory metal grains may have been at least partially below the liquidus temperature of some chondrules, meaning that some metal grains might not have been melted and, as such, might have become part of the matrix instead of being incorporated into the chondrules (Budde et al., 2016a, 2016b). Thus, although the exact mechanism remains uncertain, formation of chondrules by localized melting events in the disk can account for the observed isotopic complementarity of chondrules and matrix. The formation of chondrules from a spatially restricted reservoir of solar nebula dust also provides a mechanism for keeping chondrules and matrix together. If chondrules and matrix from a given chondrite originated from the same reservoir of dust, and provided that no material was lost from this reservoir after chondrule formation and before chondrite accretion, then the bulk chondrite would naturally have the same isotopic composition as the original reservoir from which the chondrules and matrix derive. A corollary of this is that to avoid loss of either chondrules or matrix, there probably was a strong temporal link between chondrule formation and chondrite accretion (see Section 10.4). Alexander and Ebel (2012) argued that in a turbulent solar nebula, chondrules would rapidly mix on timescales of <1Ma and on this basis inferred that the distinct physical and chemical properties of chondrules from a given chondrite group could not be preserved, unless chondrules rapidly accreted into planetesimals. Nevertheless, Jacquet et al. (2012) argued that immediate chondrite accretion following chondrule formation may not be necessary to preserve chondrule-matrix comple- mentarity, provided that solids and gas are tightly coupled to each other. Under these conditions, more complex accretion histories might be possible, in which chondrites consist of several distinct dust populations (each consisting of chondrules and matrix) that could, in principle, have formed over an extended period of time (Goldberg et al., 2015). However, these models provide no precise estimate for the timescale over which chondrule-matrix complementarity could be preserved, preventing significant tightening of constraints on a possible time gap between chondrule formation and chondrite accretion in these models. We, therefore, stress that by far the most straightforward way of preserving complementarity is the immediate accretion of chondrules and matrix into a parent body. In this case, no special conditions need to be invoked to preserve distinct dust populations for an extended period of time. The derivation of chondrules and matrix from a single reservoir of dust raises the question of whether this implies complete chemical or thermal processing of matrix during chondrule formation. Evidently this was not the case because the matrix contains presolar grains and thus has, at least in part, never been heated significantly (e.g., Huss et al., 2003). Moreover, Tungsten Isotopes and the Origin of Chondrules and Chondrites 285

Alexander (2005) argued that the matrices of all chondrites are dominated by CI-like material, also implying that matrix (at least significant parts of it) had not been significantly modified during chondrule formation. However, the complementarity does not require thermal processing of matrix; it only requires that the material that was removed from the chondrules (or the chondrule precursors) was incorporated into the matrix and was not lost. Thus, the evidence that matrix escaped thermal processing does not contradict the inference of complementarity that chondrules and matrix derive from a common reservoir of solar nebula dust. In summary, of the different models proposed for the origin of chondrules, only models in which chondrules are the result of localized melting events in the solar protoplanetary disk are consistent with the isotopic complementarity of chondrules and matrix. Other models, such as the x-wind model as well as the splashing and impact jetting models, cannot account for some or all of the constraints derived from the isotopic complementarity.

10.3.4 Significance of Isotope Anomalies in Single Chondrules

One important implication of chondrule-matrix complementarity is that chondrules from a given chondrite group should all derive from spatially restricted areas of the disk. However, variations in 54Cr isotopic compositions of individual chondrules from CV chondrites have been inter- preted to reflect formation at different orbital distances, followed by transport of chondrules across the disk before accretion to the CV parent body (Olsen et al., 2016). This interpretation is based on the observation that the range of 54Cr compositions observed among CV chondrules is similar to the range of 54Cr compositions for bulk meteorites. Obviously, such a disk-wide transport of chondrules is inconsistent with the formation of chondrules in localized regions as inferred from the isotopic complementarity. However, it is important to note that the 54Cr data cannot be used to distinguish between transport of chondrules and transport of chondrule precursors. As such, the 54Cr signatures do not provide evidence that chondrules themselves formed in different regions of the disk; they only show that chondrules formed from isotopically heterogeneous precursors. This heterogeneity may have arisen locally within the disk, or may reflect transport of isotopically anomalous dust grains across the disk, followed by incorporation of these dust grains into chondrule precursors, which then later melted to form chondrules. Direct evidence for such transport processes comes from Ti isotopic data for chondrules. As for 54Cr, individual CV chondrules also display a wide range of 50Ti isotopic compositions. Gerber et al. (2017) demonstrated that the 50Ti variations reflect the variable admixture of CAI- like and 50Ti-enriched material to the chondrules. The 50Ti variations, therefore, reflect a ‘nugget’ effect arising from the incorporation of isotopically anomalous grains into chondrule precursors, rather than providing a unique signature of distinct formation locations in the disk. The variable 54Cr compositions of CV chondrules are thus most likely also due to a nugget effect, reflecting the incorporation of 54Cr-rich dust grains into chondrule precursors (Gerber et al., 2017). As such there is no need to invoke disk wide transport of chondrules to account for their variable 54Cr compositions, and the 54Cr (and 50Ti) anomalies in CV chondrules are consistent with formation in localized regions of the disk, as required by the isotopic comple- mentarity of CV chondrules and matrix. 286 Thorsten Kleine, et al. 10.4 Hf–W Chronology of CV and CR Chondrules 10.4.1 182Hf–182W Ages

The Hf–W system can be used to date chondrules, because chondrule formation and associated metal-silicate fractionation led to substantial Hf–W fractionation among the constituents of primi- tive chondrites. For instance, in the CV3 chondrite Allende, chondrules have elevated Hf/W ratios, while matrix is characterized by correspondingly low Hf/W ratios (Becker et al., 2015; Budde et al., 2016b). In CR chondrites, chondrules typically contain metal and for these samples the major Hf–W fractionation therefore is between silicates and metal (Budde et al., 2018). The Hf–W system cannot easily be applied to single chondrules, because the W concen- trations in chondrules are too low for precise isotope analysis (chondrules contain ~90 ng/g W, but ~30 ng W are needed for precise W isotopic analysis). The Hf–W data are, therefore, typically obtained on pooled chondrule separates, each consisting of hundreds to thousands of chondrules or chondrule fragments (Budde et al., 2016b, 2018). To date, there are two precise Hf–W ages for primitive carbonaceous chondrites. Chondrules and matrix samples from the CV3 chondrite Allende define a Hf–W isochron that corresponds to an age of 2.2 0.8 Ma after CAI formation at 4,567.3 Ma (Figure 10.4a) (Budde et al., 2016b), and metal and silicate fractions from four CR2 chondrites defineaHf–Wisochronwithanageof3.6 0.6 Ma after CAI formation (Figure 10.4b) (Budde et al., 2018). Of note, an isochron regression using only data for silicate and metal fractions separated from CR chondrules (Figure 10.4c) yields the same age as the combined isochron for bulk metal and silicate fractions. This is consistent with the idea that the major metal-silicate fractionation in CR chondrites was associated with chondrule formation (e.g., Kong et al., 1999; Campbell et al., 2005; Jacquet et al., 2013). It has been suggested that the Hf–W ages are not chronologically meaningful, because the very presence of nucleosynthetic W isotope anomalies would mean that there had been no W isotopic equilibration among the components used to construct the Hf–W isochrons (Chapter 11). However, the presence of nucleosynthetic isotope anomalies by no means excludes or requires that chondrules, matrix, and metal were not in Hf–W isotopic equilibrium. This is because the nucleosynthetic isotope anomalies arise from the heterogeneous distribution of small presolar grains with very large isotopic anomalies. As such, the presolar grains contain a negligible amount of the bulk W present in chondrules and matrix, meaning that the heterogeneous distribution of presolar grains has no effect on the chemical fractionation and equilibration of Hf and W between chondrules and matrix (or metal). In other words, the nucleosynthetic W isotope anomalies simply reflect the complementary depletion and enrich- ment of presolar grains between two components that otherwise were in isotopic equilibrium. Moreover, as noted above (sect. 10.2.1), nucleosynthetic W isotope anomalies are very small for the 182W/183W ratio normalized to 186W/183W. Thus, using these data in the isochron regression makes it possible to assess whether there had been any primordial W isotopic heterogeneity unrelated to the nucleosynthetic W isotope variations. Of note, the 182W/183W data (normalized to 186W/183W) provide precise isochrons for both Allende chondrules and matrix as well as for CR metal and silicates. This demonstrates that there was no initial W isotopic heterogeneity (other than the nucleosynthetic variations) among the components used to define the isochrons. Thus, the Hf–W data define meaningful isochrons. 2 (a) Allende CV3 chondrite 182Hf/180Hf = (8.6 ± 0.4) × 10-5 1 i e182 Wi = –3.34 ± 0.13 t =2.2±0.8 Ma 0

W –1 182 e

–2

Chondrules –3 Matrix Bulk –4 012345 180Hf/184W 4 (b) CR2 chondrites 182 180 -5 Hf/ Hfi = (7.7 ± 0.2) × 10 e182 2 Wi = –3.13 ± 0.05 t = 3.6 ± 0.6 Ma

W 0 182 e

Metal –2 Metal + silicates Bulk Chondrule silicates –4 02468 180Hf/184W 4 (c) CR2 chondrules 182 180 -5 Hf/ Hfi = (7.6 ± 0.4) × 10 e182W = –3.15 ± 0.11 2 i t =3.7±0.8 Ma

W 0 182 e

–2 Bulk CR chondrites Chondrule silicates Chondrule metal –4 02468 180Hf/184W

Figure 10.4 Hf–W isochrons for CV and CR chondrites. Ages calculated using the initial 182Hf/180Hf obtained from the isochrons and equation 1. (a) Allende chondrules and matrix. Data from Budde et al. (2016b). (b) Metal and silicate fractions from four different CR chondrites, and (c) metal and silicate fractions from chondrules of CR chondrites. Data for CR chondrites from Budde et al. (2018). 288 Thorsten Kleine, et al. 10.4.2 Comparison to 26Al–26Mg Ages

The Hf–W ages for CV and CR chondrules are in very good agreement with Al–Mg ages reported for these two groups of chondrites. For CV chondrites, the Al–Mg system in chondrules might be partially or wholly reset during parent body alteration (Alexander and Ebel, 2012), but Nagashima et al. (2017) argued that for the least metamorphosed CV chondrite Kaba, the Al–Mg system has not been disturbed. For Kaba chondrules, these authors reported an initial 26Al/27Al = (4.8 2.1) 10‒6, corresponding to an Al–Mg age of 2.4 0.4 Ma after CAI formation [using 26Al/27Al = (5.23 0.13) 10‒5 for CAIs; Jacobsen et al., 2008]. This Al–Mg age for Kaba chondrules is in excellent agreement with the Hf–W age of 2.2 0.8 Ma for Allende chondrules and matrix. Schrader et al. (2017) showed that the initial 26Al/27Al of CR chondrules is significantly lower than those in chondrules from other chondrite groups, and also showed that this lower initial 26Al/27Al does not reflect disturbance and resetting by parent body processes. These authors reported a weighted average initial 26Al/27Al = (1.33 0.29) 10‒6 for CR chondrules, corresponding to an Al–Mg age of 3.8 0.2 Ma after CAI formation. Again, this age is in excellent agreement with the Hf–W age for CR chondrules of 3.6 0.6 Ma after CAIs. The good agreement of Hf–W and Al–Mg ages for CV and CR chondrules not only demonstrates that these formation ages for chondrules are robust, but also has important implications for assessing the distribution of 26Al in the solar nebula. Although the Al–Mg system can provide very precise ages for early Solar System objects, differences in inferred initial 26Al/27Al ratios could potentially also reflect the heterogeneous distribution of 26Al in the solar nebula, rather than different formation times. For instance, on the basis of correlated 54Cr and 26Mg anomalies in bulk meteorites and CAI, Larsen et al. (2011) argued for a high degree of 26Al/27Al heterogeneity of up to ~80 percent, in which case the Al–Mg system could not be used as a chronometer. However, the 26Mg variations more likely reflect small nucleosynthetic Mg isotope anomalies and not large-scale 26Al/27Al heterogeneity (Kruijer et al., 2014a). Neverthe- less, for a reliable chronological interpretation of Al–Mg data it is important to independently evaluate the level of 26Al/27Al homogeneity/heterogeneity. To assess the chronological significance of Al–Mg ages it is useful to compare them to ages obtained from other chronometers, such as the Hf–W system. Such a comparison can be made for bulk CAIs, CV, and CR chondrules as well as angrites (basaltic achondrites), because for all these samples precise Al–Mg and Hf–W data are available. Moreover, these samples formed over almost the entire lifetime of 26Al (i.e., within the first ~5 Ma of the Solar System), making it possible to evaluate possible heterogeneities in 26Al/27Al at different times of Solar System evolution. Figure 10.5 illustrates that the Al–Mg and Hf–W ages of the aforementioned samples are concordant. The good agreement between ages of four different samples that formed at different times and in different areas of the protoplanetary disk provides strong evidence for a homogeneous distribution of 26Al at the scale of bulk CAIs and bulk meteorites. Nevertheless, owing to the much shorter half-life of 26Al compared to 182Hf, small 26Al/27Al heterogeneities may still exist that could not be resolved by age comparison to the Hf–W system. Taking the scatter around the correlation in Figure 10.5 as a measure of the permissible degree of 26Al/27Al heterogeneity limits it to smaller than ~10–20 percent (Budde et al., 2017). Note that this value only provides a maximum estimate and by no means indicates or requires an actual heterogen- eity in the distribution of 26Al. As such, these data demonstrate that Al–Mg ages for chondrules and other samples have chronological significance. Tungsten Isotopes and the Origin of Chondrules and Chondrites 289

6 Angrites 5

4 CR chondrules 3 CV chondrules

2 Al-Mg age (Ma) 1 Slope = 1.03 ± 0.12 0 CAI

0123456 Hf-W age (Ma)

Figure 10.5 Comparison of Al–Mg and Hf–W ages for different meteorites. All ages are calculated relative to CAIs, but excluding the uncertainty on the initial 182Hf/180Hf and 26Al/27Al ratios of the CAIs. This was done to use CAIs as an independent data point in the age comparison. For data sources, see Budde et al. (2018).

10.4.3 Comparison to 207Pb–206Pb Ages

The Hf–W ages for the CV and CR chondrules are also in good agreement with the average Pb– Pb ages reported for chondrules from these two chondrite groups. For multi-chondrule fractions from Allende, Pb–Pb ages of 4,565.6 1.0 Ma (Amelin and Krot, 2007), 4,564.5 0.5 Ma (Connelly et al., 2008) and 4,564.3 0.8 Ma (Connelly and Bizzarro, 2009) were reported. Note that these ages were corrected for age bias induced by differences in U isotopic compos- itions, as summarized in Brennecka et al. (2015). These Pb–Pb ages correspond to age differences relative to CAIs of 1.7 1.0 Ma, 2.8 0.5 Ma and 3.0 0.8 Ma and as such are in good agreement with the Hf–W age of 2.2 0.8 Ma of Allende chondrules and matrix. Amelin et al. (2002) reported a Pb–Pb age for a multi-chondrule fraction from a CR chondrite, which after correction for the U isotopic composition of CR chondrites, is 4,563.6 0.6 Ma or 3.7 0.6 Ma after CAI formation (Schrader et al., 2017). Again, this age is in very good agreement with the Hf–W age of 3.6 0.6 Ma for metal-silicate fractionation in CR chondrites. Thus, for multichondrule fractions of CV and CR chondrules there is very good agreement between the Hf–W and Pb–Pb ages. Matters become more complicated when Pb–Pb ages for individual chondrules are also considered. Connelly et al. (2012) reported Pb–Pb ages of 4,567.3 0.4 Ma (i.e., 0.0 0.4 Ma after CAIs) and 4,566.2 0.6 Ma (i.e., 1.1 0.6 Ma after CAIs) for two chondrules from the CV chondrite Allende. Of note, both ages are older than the aforementioned Pb–Pb ages determined for multi-chondrule separates from Allende. More recently, Bollard et al. (2017) reported Pb–Pb ages for chondrules from ordinary and CR chondrites, some of which also seem to be as old as CAI. On this basis, both Connelly et al. (2012) and Bollard et al. (2017) argued that chondrules formed continuously over at least the first ~3 Ma of the Solar System. 290 Thorsten Kleine, et al.

Such an extended formation period for chondrules from a single chondrite group is not apparent in the Al–Mg ages for chondrules, however. For instance, chondrules from the CV chondrite Kaba all have indistinguishable Al–Mg ages, with a mean age of 2.4 0.4 Ma after CAIs (Nagashima et al., 2017). As noted above, Kaba is one of the least metamorphosed CV chondrites and has been classified as petrologic type 3.1 (Bonal et al., 2006); as such the Al–Mg systematics of chondrules from Kaba should have remained largely undisturbed as a result of parent body metamorphism. For other, more strongly metamorphosed CV chondrites (e.g., Allende, petrologic type 3.6; Bonal et al., 2006) the Al–Mg chondrule ages are more variable and, compared to chondrules from Kaba, extend to younger apparent ages. However, this range of ages reflects disturbance of the Al–Mg system during parent body metamorphism and not an extended period of chondrule formation (Nagashima et al., 2017). The duration of CV chon- drule formation as derived from Al–Mg ages can therefore only be reliably deduced from the least metamorphosed sample (in this case, Kaba). As such, the Al–Mg data for Kaba indicate that CV chondrules formed within a time interval of <1 Ma. The Al–Mg age distributions of individual chondrules from other chondrite groups show that in general the majority of chondrules from a given group formed in a narrow time interval. For instance, although Villeneuve et al. (2009) identified several peaks in the Al–Mg age distribu- tion of chondrules from primitive ordinary chondrites, ~80 percent of these chondrules formed in a time interval between ~1.5 and ~2.8 Ma after CAIs. Likewise, Schrader et al. (2017) observed that ~90 percent of chondrules from CR chondrites formed at ~3.5 Ma after CAIs. It is noteworthy that the true range of formation ages for chondrules from a given chondrite might be smaller than the reported range of Al–Mg ages, because the Al–Mg system might be partially disturbed by parent body processes (Alexander and Ebel, 2012) and because the observed range in ages at least in part reflects analytical uncertainties and not a true range in formation ages. Thus, taken together, the Al–Mg data indicate that the vast majority of chondrules from a given chondrite formed in a narrow time interval of <1 Ma. The discrepancy between this short interval of <1 Ma obtained from Al–Mg data and the extended period of ~3 Ma deduced from the Pb–Pb ages cannot reflect a heterogeneous distribution of 26Al at the scale of individual chondrules. This is because 26Al/27Al heterogen- eity among the precursor material of individual chondrules would lead to an apparent range in ages and not to one consistent age. The discrepancy between the Al–Mg and Pb–Pb ages also cannot reflect a reduced initial 26Al/27Al of bulk chondrites. For instance, to account for the mismatch of Pb–Pb and Al–Mg ages for chondrules from ordinary and CR chondrites, Bollard et al. (2017) argued that at the time of CAI formation chondrule precursors had a lower initial 26Al/27Al of ~1 10–5 than CAIs themselves (26Al/27Al ~5 10–5). This reduced initial 26Al/27Al would make the chondrule Al–Mg ages appear ~2 Ma too young. However, this possibility is effectively being ruled out by the good agreement of Al–Mg and Hf–W ages of CV and CR chondrules, angrites and CAIs (see above and Figure 10.5), demonstrating that the initial 26Al/27Al ratios of these samples are consistent with radioactive decay from a common Solar System initial value as measured in CAIs. Thus, the mismatch between Al–Mg and Pb–Pb ages for individual chondrules cannot be accounted for by a heterogeneous distribution of 26Al, meaning that the Al–Mg data support neither the range of Pb–Pb ages observed for chondrules for a given chondrite group, nor the presence of chondrules as old as CAI. Tungsten Isotopes and the Origin of Chondrules and Chondrites 291

The observation that some CV chondrules have apparent Pb–Pb ages as old as CAI is also difficult to reconcile with other chronological data for CV chondrites. This is because if there were a significant fraction of CV chondrules as old as CAI, then there must also be a significant fraction of much younger chondrules with ages of around ~4–5 Ma after CAIs. This is necessary to account for the Pb–Pb and Hf–W ages of between ~2 and ~3 Ma after CAIs obtained for multi-chondrule separates from Allende. However, there are no CV chondrules with reported formation ages as young as ~4–5 Ma after CAIs. Moreover, such young formation ages of CV chondrules would be inconsistent with independent constraints on the timing of parent body accretion and alteration. Secondary fayalites in CV chondrites, which formed during aqueous alteration on the parent body, have a 53Mn–53Cr age of 4.2 0.8 Ma after CAIs (Doyle et al., 2015; Jogo et al., 2017). As chondrule formation obviously must predate aqueous alteration on the parent body, the Mn–Cr age of secondary fayalites effectively rules out that CV chondrules could have formed as late as 4–5 Ma after CAIs. Moreover, to facilitate heating and formation of aqueous fluids that ultimately lead to alteration, there must be a time gap between chondrule formation (and hence chondrite accretion) and parent body alteration. Using thermal modeling of asteroids heated internally by 26Al decay, it has been shown that to account for the formation conditions and ages of secondary fayalites, the CV parent body must have accreted between ~2.5 Ma and ~3.3 Ma after CAI formation (Doyle et al., 2015; Jogo et al., 2017). On this basis, there should be no CV chondrules that formed later than ~3 Ma after CAIs and, consequently, only a very minor fraction if any of CV chondrules can be as old as CAIs. Finally, the formation of CV chondrules over an extended period of ~3 Ma as inferred from Pb–Pb chronometry is difficult to reconcile with the isotopic complementarity of chondrules and matrix. As argued above, this complementarity is most easily explained by rapid accretion of both chondrules and matrix into a chondrite parent body. This rapid accretion would also allow for the preservation of distinct chondrule populations, which is necessary to account for the distinct chemical and physical properties of chondrules from a given chondrite group (Alexan- der and Ebel, 2012). In summary, the isotopic complementarity of chondrules and matrix, the narrow range of Al– Mg ages of <1 Ma for chondrules from a given chondrite group, the good agreement of Al–Mg and Hf–W ages for chondrules, and the chronology of aqueous alteration on the CV chondrite parent body combined with constraints on the time of parent body accretion, all indicate that the vast majority of chondrules for a single chondrite group formed in a narrow interval of time and that chondrule formation and chondrite accretion were about coeval. By contrast, the prolonged interval of chondrule formation inferred from Pb–Pb chronometry is inconsistent with all these constraints. As such, the variable Pb–Pb ages reported for chondrules from a given chondrite group either do not reflect true differences in chondrule formation ages, or, alternatively, the Pb–Pb ages have been obtained on a set of chondrules that represents only a very minor fraction of the entire population of chondrules.

10.5 Hf–W Chronology of Ordinary Chondrites

Ordinary chondrites contain abundant metal and as such can be well dated using the Hf–W system through metal-silicate isochrons. To date, the Hf–W system has been applied to several 292 Thorsten Kleine, et al. equilibrated ordinary chondrites of petrologic types 4 to 6 (Kleine et al., 2008; Hellmann et al., 2017). As the Hf–W closure temperature is similar to the estimated peak metamorphic temperatures of ordinary chondrites, the Hf–W ages closely approximate the time of the thermal peak for each sample (Kleine et al., 2008). As such, the Hf–W ages provide insights into the high-temperature thermal evolution of ordinary chondrites, information that is critical for assessing the heat source for metamorphism and the internal structure of ordinary chondrite parent bodies. The Hf–W ages for type 4 ordinary chondrites typically are ~3–4 Ma, while type 5 and 6 chondrites have ages of between ~6 and ~12 Ma, post CAI formation (Figure 10.6). This range in ages is consistent with the expected timescale of thermal metamorphism in asteroids that accreted at ~2 Ma after CAI formation and were internally heated by decay of 26Al (Trieloff et al., 2003; Kleine et al., 2008; Henke et al., 2012). Of note, this timescale of parent body accretion is consistent with Al–Mg ages for chondrules from primitive ordinary chondrites (Kita et al., 2000; Rudraswami and Goswami, 2007; Villeneuve et al., 2009), indicating that

(a) Type 6

Type 5

LL Type 4 L H

0 2 4 6 8 10 12 14 Time after CAIs (Ma)

(b) Type 6 CAI initial

Type 5

LL Type 4 L H

–3.8 –3.4 –3–3.2 –2.8 –2.6 –2.4 e182W

Figure 10.6 (a) Hf–W ages and (b) initial ε182W of ordinary chondrites. Note that for a given petrologic type, there are no systematic differences between the Hf–W ages for H, L, and LL chondrites. By contrast, metal from type 5 and 6 LL chondrites always has higher ε182W compared to H and L metals of the same petrologic type. Data from Hellmann et al. (2017). Tungsten Isotopes and the Origin of Chondrules and Chondrites 293 there was a close temporal link between chondrule formation and chondrite accretion (see also Section 10.6). The H, L, and LL subgroups of ordinary chondrites are distinguished based on different metal contents, most likely as a result of metal-silicate fractionation prior to parent body accretion. Of note, whereas for all three subgroups metal samples from type 4 samples have similar and comparatively low ε182W (providing the initial 182W composition at the time of Hf–W closure), samples of higher petrologic types (type 5 and 6) exhibit higher and also more variable initial ε182W values (Figure 10.6). As samples of a given petrologic type from the H, L, and LL subgroups have similar Hf–W ages, their different initial ε182W cannot reflect different times of Hf–W closure. Instead they reflect distinct Hf/W ratios of the H, L, and LL parent bodies, which over time led to the observed variations in initial ε182W. The highest initial ε182W are observed for type 5 and 6 LL chondrites, while H chondrites of similar age have much lower initial ε182W (Figure 10.6). The H chondrites therefore evolved with the lowest and LL chondrites with the highest Hf/W. Hellmann et al. (2017) demonstrated that these distinct Hf/W ratios were established between 1.5 and 3 Ma after CAI formation. A fractionation earlier than ~1.5 Ma can be excluded because in this case the type 4 ordinary chondrites would already have had distinct initial ε182W values, and because the back projection of the ε182W composition of type 4 metal from L and LL chondrites would result in ε182W values below the solar system initial value defined by CAIs. A fractionation later than ~3 Ma can also be excluded because the Hf–W fractionation among the different subgroups cannot postdate the onset of thermal metamorphism in each of the bodies, as given by the Hf–W ages for type 4 chondrites (Figure 10.6). The metal-silicate fractionation among the different subgroups of ordinary chondrites at 1.5–3 Ma after CAIs was about coeval with formation of ordinary chondrite chondrules at ~2 Ma, consistent with the idea that chondrule formation was associated with metal-silicate fractionation. It is noteworthy that H chondrites have the same Fe/Si (and hence metal-to- silicate) ratio as CI chondrites (Palme et al., 2014), implying that the Hf/W ratio of H chondrites reflects the average ratio prior to metal-silicate fractionation among the ordinary chondrite subgroups. If this is correct, then metal must have been removed from L and LL chondrites, to accommodate their higher Hf/W. Of note, the aerodynamic properties of metal in ordinary chondrites are most similar to those of chondrules from H chondrites (Jacquet, 2014), suggest- ing that aerodynamic sorting did not significantly affect the chondrule-to-metal ratio in H chondrites. By contrast, in L and LL chondrites, chondrules might have been fractionated from metal grains. Thus, the distinct metal-silicate ratios of the ordinary chondrite subgroups probably reflect varying degrees of metal loss, possibly by aerodynamic sorting from H chondrite-like precursors.

10.6 Chondrules and the Accretion of Planetesimals

Chondrite parent bodies underwent different thermal histories, ranging from high temperature thermal metamorphism in ordinary chondrite parent bodies to the much lower temperature aqueous alteration typical for many carbonaceous chondrites (e.g., Huss et al., 2006; Krot et al., 2006). Figure 10.7 shows that there is an inverse correlation between chondrule ages and peak temperatures inferred for each chondrite parent body. For instance, chondrules from ordinary 294 Thorsten Kleine, et al.

1400 Fe-FeS eutectic

OC 1000 CV CO 600 Temperature (K) Temperature CR

200 12345 Time after Solar System formation (Ma)

Figure 10.7 Peak temperature reached inside chondrite parent bodies versus chondrule ages. Whereas the ages for OC and CO chondrules are Al–Mg ages (Kurahashi et al., 2008; Villeneuve et al., 2009), the ages shown for CV and CR chondrules are Hf–W ages. As discussed in the text, there is very good agreement between the Hf–W and Al–Mg ages for these chondrites. The dashed line is the maximum temperature increase through decay of 26Al, calculated by assuming the Al concentration of 1.167 wt. % of H chondrites (Wasson and Kallemeyn, 1988). These calculations only serve to illustrate that the observed thermal histories of chondrites are consistent with the effects expected from heating through 26Al decay at different times of parent body accretion. In reality, the thermal evolution of chondrite parent bodies is more complex. Estimated peak temperatures for the different chondrite groups taken from Scott (2007). The gray box shows the temperature field in which melting will occur, indicating that bodies accreted within the first ~2 Ma of the solar system underwent at least partial melting. chondrites are about ~1–2 Ma older than chondrules from CR chondrites, and peak metamorphic temperatures estimated for ordinary chondrite parent bodies were much higher (up to ~1,200 K) than those for the CR chondrite parent body (lower than ~500 K). This inverse correlation between chondrule ages and chondrite peak temperatures provides several important observations. First, such a correlation is expected only if the chondrule ages closely approximate the time of parent body accretion, defining the time at which heating inside the bodies started. Of note, this is consistent with the chronology of chondrule formation and independent constraints on the time of chondrite parent body accretion. For instance, based on Al–Mg chronometry, CV chondrules formed at 2.4 0.4 Ma after CAIs, consistent with an ~2.5–3.3 Ma accretion age of the CV parent body derived from Mn–Cr dating of fayalites combined with thermal modeling (see Section 10.4). Similarly, the chronology of the thermal metamorphism of ordinary chondrites combined with thermal modeling provides an accretion age of ~2 Ma after CAIs for the ordinary chondrite parent bodies. Again, this is consistent with Al–Mg ages for chondrules from the most primitive ordinary chondrites (see Section 10.5). This close temporal link between chondrule formation and chon- drite accretion is also consistent with the isotopic complementarity of chondrules and matrix (see Section 10.3). Another important observation from Figure 10.7 is that the decay of 26Al must have been the dominant heat source determining the thermal history of chondrite parent bodies. This is because 26Al is the only heat source whose power strongly decreases as a function of accretion time. This suggests that, for instance, CR chondrites predominantly underwent aqueous alteration on their parent body not only because they contained water ice, but also because the CR chondrite parent body, owing to its late accretion, contained too little 26Al for significant thermal metamorphism. Of note, the inverse correlation shown in Figure 10.7 in principle also Tungsten Isotopes and the Origin of Chondrules and Chondrites 295 extends to differentiated meteorites. These derive from bodies that underwent large-scale melting and chemical differentiation, and, based on Hf–W chronometry, accreted within the first ~1 Ma of the Solar System (Kruijer et al., 2014b, 2017; Kleine and Walker, 2017). At this early time, 26Al was sufficiently abundant to cause complete melting and efficient metal-silicate differentiation (Figure 10.7). A final important observation from Figure 10.7 is that the oldest chondrule population (i.e., those from ordinary chondrites) formed at exactly the time when 26Al had decayed to a level at which it can no longer cause melting and differentiation of planetesimals. This is unlikely to be coincidence but rather suggests that planetesimals that formed earlier than the ordinary chon- drite parent bodies also contained chondrules initially, but that these chondrules were destroyed later during subsequent melting and differentiation. Thus, chondrule formation probably did not only commence at ~2 Ma after solar system formation, but started as soon as the earliest planetesimals formed. A corollary of this observation is that planetesimals in general might have contained chondrules, implying that chondrule formation was closely linked to, and might have even been a critical step towards, making planetesimals.

10.7 Conclusions

Tungsten isotope studies on chondrites and chondrite components have led to several major constraints that are key for understanding the origin of chondrules and chondrites: Chondrules and matrix (or metal) have complementary nucleosynthetic isotope anomalies, indicating that there is a strong genetic link between these three components. The isotopic complementarity reflects the uneven distribution of a metallic presolar carrier between chondrules and matrix (or metal), and as such most likely results from metal-silicate fractionation during chondrule formation. The complementarity is difficult to reconcile with an origin of chondrules by protoplanetary impacts and also with models in which chondrules and matrix formed in different regions of the protoplanetary disk (e.g., x-wind model). As such, the complementarity indicates that chondrules formed by localized melting of dust aggregates in the solar nebula. The isotopic complementarity is best explained by a rapid accretion of chondrules and matrix into a parent body. This is consistent with Hf–W and Al–Mg evidence that chondrules from a given chondrite group formed in a narrow time interval of <1 Ma, and implies a strong temporal link between chondrule formation and chondrite accretion. Such a link is also evident from the inverse correlation of chondrule ages and peak metamorphic temperatures reached inside chondrite parent bodies. Metal-silicate fractionation among subgroups of ordinary chondrites was about coeval to chondrule formation and may have occurred by aerodynamic sorting of chondrules (or their precursors) and metal grains. This provides further evidence that chondrule formation was associated with metal-silicate fractionation. CR chondrules formed significantly later than other chondrules (except for CB chondrules, which likely formed by a different formation mechanism than other chondrules; Chapter 2), implying that chondrule formation and accretion of planetesimals occurred for at least ~3–4 Ma after solar system formation. 296 Thorsten Kleine, et al.

The good agreement of Hf–W and Al–Mg ages for several early solar system objects indicates a homogeneous distribution of 26Al in the solar nebula, and limits any possible 26Al/27Al heterogeneity to less than 10–20 percent. The Al–Mg system can therefore be used as a precise chronometer for early Solar System processes. Decay of 26Al was the dominant heat source determining the thermal history of planetesimals, including the parent bodies of both differentiated meteorites and chondrites.

Acknowledgments

We thank Emmanuel Jacquet and Graeme Poole for constructive reviews that have greatly improved the chapter. We also thank Sara Russell, Sasha Krot and Harold Connolly for organising the chondrule workshop in London, and for their assistance and support in the production of this chapter and the entire book. Funding by the Deutsche Forschungsge- meinschaft (DFG) is gratefully acknowledged.

References

Alexander, C. M. O’D. (2005). From supernovae to planets: The view from meteorites and interplanetary dust particles. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, 972–1002. Astronomical Society of the Pacific Conference Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Alexander, C. M. O.’D., and Ebel, D. S. (2012). Questions, questions: Can the contradictions between the petrologic, isotopic, thermodynamic, and astrophysical constraints on chon- drule formation be resolved? Meteoritics & Planetary Science, 47, 1157–1175. Amelin, Y., and Krot, A. N. (2007). Pb isotopic ages of the Allende chondrules. Meteoritics & Planetary Science, 42, 1321–1335. Amelin, Y., Krot, A. N., Hutcheon, I. D., and Ulyanov, A. A. (2002). Lead isotopic ages of chondrules and calcium-aluminum-rich inclusions. Science, 297, 1678–1683. Arlandini, C., Käppeler, F., and Wisshak, K. (1999). Neutron capture in low-mass asymptotic giant branch stars: Cross sections and abundance signatures. Astrophysical Journal, 525, 886–900. Asphaug, E., Jutzi, M., and Movshovitz, M. (2011). Chondrule formation during planetesimal accretion. Earth and Planetary Science Letters, 308, 369–379. Becker, M., Hezel, D. C., Schulz, T., Elfers, B. -M., and Münker, C. (2015). Formation timescales of CV chondrites from component specificHf–W systematics. Earth and Planetary Science Letters, 432, 472–482. Bland, P. A., Alard, O., Benedix, G. K., et al. (2005). Volatile fractionation in the early solar system and chondrule/matrix complementarity. Proceedings of the National Academy of Sciences of the United States of America, 102, 13755–13760. Bollard, J., Connelly, J. N., Whitehouse, M. J., et al. (2017). Early formation of planetary building blocks inferred from Pb isotopic ages of chondrules. Science Advances, 3,(8),e1700407. Bonal, L., Bourot-Denise, M., Quirico, E., and Montagnac, G. (2006). Determination of the petrologic type of CV3 chondrites by Raman spectroscopy of included organic matter. Geochimica et Cosmochimica Acta, 70, 1849–1863. Brennecka, G. A., Budde, G., and Kleine, T., (2015). Uranium isotopic composition and absolute ages of Allende chondrules. Meteoritics & Planetary Science, 50, 1995–2002. Budde, G., Burkhardt, C., Brennecka, G. A., et al. (2016a). Molybdenum isotopic evidence for the origin of chondrules and a distinct genetic heritage of carbonaceous and non- carbonaceous meteorites. Earth and Planetary Science Letters, 454, 293–303. Tungsten Isotopes and the Origin of Chondrules and Chondrites 297

Budde, G., Kleine, T., Kruijer, T. S., Burkhardt, C., and Metzler, K. (2016b). Tungsten isotopic constraints on the age and origin of chondrules. Proceedings of the National Academy of Sciences, 113, 2886–2891. Budde, G., Kruijer, T. S., and Kleine, T. (2018). Hf–W chronology of CR chondrites: Implica- tions for the timescales of chondrule formation and the distribution of 26Al in the solar nebula. Geochimica et Cosmochimica Acta, 222, 284–304. Burkhardt, C., Kleine, T., Dauphas, N., and Wieler, R. (2012). Nucleosynthetic tungsten isotope anomalies in acid leachates of the Murchison chondrite: Implications for Hf–W chronom- etry. The Astrophysical Journal Letters, 753, L6. Burkhardt, C., Kleine, T., Oberli, F., et al. (2011). Molybdenum isotope anomalies in meteor- ites: Constraints on solar nebula evolution and origin of the Earth. Earth and Planetary Science Letters, 312, 390–400. Burkhardt, C., Kleine, T., Palme, H., et al. (2008). Hf–W mineral isochron for Ca,Al-rich inclusions: Age of the solar system and the timing of core formation in planetesimals. Geochimica et Cosmochimica Acta, 72, 6177–6197. Campbell, A. J., Zanda, B., Perron, C., et al. (2005). Origin and thermal history of Fe-Ni metal in primitive chondrites. In A. N. Krot (Ed.), Chondrites and the Protoplanetary Disk. pp. 407–431. Astronomical Society of the Pacific Conference Series, 341. San Francisco, CA: Astronomical Society of the Pacific. Connelly, J. N., Amelin, Y., Krot, A. N., and Bizzarro, M. (2008). Chronology of the solar system’s oldest solids. The Astrophysical Journal Letters, 675, L121. Connelly, J. N., and Bizzarro, M. (2009). Pb–Pb dating of chondrules from CV chondrites by progressive dissolution. Chemical Geology, 259, 143–151. Connelly, J. N., Bizzarro, M., Krot, A. N., et al. (2012). The Absolute Chronology and Thermal Processing of Solids in the Solar Protoplanetary Disk. Science, 338, 651–655. Connolly, H. C., and Desch, S. J. (2004). On the origin of the “kleine Kügelchen” called Chondrules. Chemie der Erde – Geochemistry, 64,95–125. Connolly, H. C., Huss, G. R., and Wasserburg, G. J. (2001). On the formation of Fe-Ni metal in Renazzo-like carbonaceous chondrites. Geochimica et Cosmochimica Acta, 65, 4567–4588. Doyle, P. M., Jogo, K., Nagashima, K., et al. (2015). Early aqueous activity on the ordinary and carbonaceous chondrite parent bodies recorded by fayalite. Nature Communications, 6,7444. Ebel, D. S., Brunner, C., Konrad, K., et al. (2016). Abundance, major element composition and size of components and matrix in CV, CO and Acfer 094 chondrites. Geochimica et Cosmochimica Acta, 172, 322–356. Gerber, S., Burkhardt, C., Budde, G., Metzler, K., and Kleine, T. (2017). Mixing and transport of dust in the early solar nebula as inferred from titanium isotope variations among chondrules. Astrophysical Journal Letters, 841, L17. Hellmann, J. L., Kruijer, T. S., and Kleine, T. (2017). Constraining the timescale of solar nebula metal-silicate fractionation using Hf–W chronometry of ordinary chondrites. 48th Lunar and Planetary Science Conference, abstract #2046. Henke, S., Gail, H. P., Trieloff, M., Schwarz, W. H., and Kleine, T. (2012). Thermal history modelling of the H chondrite parent body. Astronomy & Astrophysics, 545, A45. Hezel, D. C., and Palme, H. (2010). The chemical relationship between chondrules and matrix and the chondrule matrix complementarity. Earth and Planetary Science Letters, 294,85–93. Hezel, D. C., and Palme, H. (2008). Constraints for chondrule formation from Ca–Al distribu- tion in carbonaceous chondrites. Earth and Planetary Science Letters, 265, 716–725. Hubbard, A. (2016a). Partitioning tungsten between matrix precursors and chondrule precursors through relative settling. The Astrophysical Journal, 826, 151. Hubbard, A. (2016b). Ferromagnetism and particle collisions: Applications to protoplanetary disks and the meteoritical record. The Astrophysical Journal, 826, 152. Huss, G. R., Meshik, A. P., Smith, J. B., and Hohenberg, C. M. (2003). Presolar diamond, silicon carbide, and graphite in carbonaceous chondrites: Implications for thermal process- ing in the solar nebula. Geochimica et Cosmochimica Acta, 67, 4823–4848. 298 Thorsten Kleine, et al.

Huss, G. R., Rubin, A. E., and Grossman, J. N. (2006). Thermal metamorphism in chondrites. In D. S. Lauretta, and H. Y. McSween (Eds.), Meteorites and the Early Solar System II, 567–586. Tucson, AZ: University of Arizona Press. Jacobsen, B., Yin, Q. -Z., Moynier, F., et al. (2008). 26Al-26Mg and 207Pb-206Pb systematics of Allende CAIs: Canonical solar initial 26Al/27Al ratio reinstated. Earth and Planetary Science Letters, 272, 353–364. Jacquet, E. (2014). Transport of solids in protoplanetary disks: Comparing meteorites and astrophysical models. Comptes Rendus – Geoscience, 346,3–12. Jacquet, E., Paulhiac-Pison, M., Alard, O., et al. (2013). Trace element geochemistry of CR chondrite metal. Meteoritics & Planetary Science, 48, 1981–1999. Jacquet, E., Gounelle, M., and Fromang, S. (2012). On the aerodynamic redistribution of chondrite components in protoplanetary disks. Icarus, 220, 162–173. Jogo, K., Nakamura, T., Ito, M., et al. (2017). Mn–Cr ages and formation conditions of fayalite in CV3 carbonaceous chondrites: Constraints on the accretion ages of chondritic asteroids. Geochimica et Cosmochimica Acta, 199,58–74. Johnson, B. C., Minton, D. A., Melosh, H. J., and Zuber, M. T. (2015). Impact jetting as the origin of chondrules. Nature, 517, 339–341. Kita, N. T., Nagahara, H., Togashi, S., and Morishita, Y. (2000). A short duration of chondrule formation in the solar nebula: Evidence from 26Al in Semarkona ferromagnesian chon- drules. Geochimica et Cosmochimica Acta, 64, 3913–3922. Kita, N. T., and Ushikubo, T. (2012). Evolution of protoplanetary disk inferred from 26Al chronology of individual chondrules. Meteoritics & Planetary Science, 47, 1108–1119. Kleine, T., Mezger, K., Palme, H., Scherer, E., and Münker, C. (2005). Early core formation in asteroids and late accretion of chondrite parent bodies: Evidence from 182Hf-182WinCAIs, metal-rich chondrites and iron meteorites. Geochimica et Cosmochimica Acta, 69, 5805–5818. Kleine, T., Touboul, M., Bourdon, B., et al. (2009). Hf–W chronology of the accretion and early evolution of asteroids and terrestrial planets. Geochimica et Cosmochimica Acta, 73, 5150–5188. Kleine, T., Touboul, M., Van Orman, J. A., et al. (2008). Hf–W thermochronometry: Closure temperature and constraints on the accretion and cooling history of the H chondrite parent body. Earth and Planetary Science Letters, 270, 106–118. Kleine, T., and Walker, R. J. (2017). Tungsten isotopes in planets. Annual Review of Earth and Planetary Sciences, 45, 389–417. Kong, P., Ebihara, M. and Palme, H. (1999). Distribution of siderophile elements in CR chondrites: Evidence for evaporation and recondensation during chondrule formation. Geochimica et Cosmochimica Acta, 63, 2637–2652. Krot, A. N., Hutcheon, I. D., Brearley, A. J., Pravdivtseva, O., and Petaev, M. I. (2006). Timescales and settings for alteration of chondritic meteorites. In D. S. Lauretta and H. Y. McSween (Eds.), Meteorites and the Early Solar System II, 525–554. Tucson, AZ: University of Arizona Press. Kruijer, T. S., Kleine, T., Fischer-Godde, M., Burkhardt, C., and Wieler, R. (2014a). Nucleo- synthetic W isotope anomalies and the Hf–W chronometry of Ca-Al-rich inclusions. Earth and Planetary Science Letters, 403, 317–327. Kruijer, T. S., Touboul, M., Fischer-Godde, M., et al. (2014b). Protracted core formation and rapid accretion of protoplanets. Science, 344, 1150–1154. Kruijer, T. S., Burkhardt, C., Budde, G., and Kleine, T. (2017). Age of Jupiter inferred from the distinct genetics and formation times of meteorites. Proceedings of the National Academy of Sciences, 114, 6712–6716. Kurahashi, E., Kita, N. T., Nagahara, H., and Morishita, Y. (2008). 26Al-26Mg systematics of chondrules in a primitive CO chondrite. Geochimica et Cosmochimica Acta, 72, 3865–3882. Larsen, K. K., Trinquier, A., Paton, C., et al. (2011). Evidence for magnesium isotope heterogeneity in the solar protoplanetary disk. The Astrophysical Journal Letters, L37, L37. Nagashima, K., Krot, A. N., and Komatsu, M. (2017). 26Al–26Mg systematics in chondrules from Kaba and Yamato 980145 CV3 carbonaceous chondrites. Geochimica et Cosmochi- mica Acta, 201, 303–319. Tungsten Isotopes and the Origin of Chondrules and Chondrites 299

Olsen, M. B., Wielandt, D., Schiller, M., Van Kooten, E. M. M. E., and Bizzarro, M. (2016). Magnesium and 54Cr isotope compositions of carbonaceous chondrite chondrules – Insights into early disk processes. Geochimica et Cosmochimica Acta, 191, 118–138. Palme, H., Hezel, D. C., and Ebel, D. S. (2015). The origin of chondrules: Constraints from matrix composition and matrix-chondrule complementarity. Earth and Planetary Science Letters, 411,11–19. Palme, H., Lodders, K., and Jones, A. (2014). Solar System Abundances of the Elements. In H. D. Holland and K. K. Turekian (Eds.), Treatise on Geochemistry (Second Edition), 15–36. Oxford: Elsevier. Qin, L., Dauphas, N., Wadhwa, M., Masarik, J., and Janney, P. E. (2008a). Rapid accretion and differentiation of iron meteorite parent bodies inferred from 182Hf-182W chronometry and thermal modeling. Earth and Planetary Science Letters, 273,94–104. Qin, L. P., Dauphas, N., Wadhwa, M., et al. (2008b). Tungsten nuclear anomalies in planet- esimal cores. Astrophysical Journal, 674, 1234–1241. Render, J., Fischer-Gödde, M., Burkhardt, C., and Kleine, T. (2017). The cosmic molybdenum- neodymium isotope correlation and the building material of the Earth. Geochemical Perspectives Letters, 3, 170–178. Rudraswami, N. G., and Goswami, J. N. (2007). 26Al in chondrules from unequilibrated L chondrites: Onset and duration of chondrule formation in the early solar system. Earth and Planetary Science Letters, 257, 231–244. Sanders, I. S., and Scott, E. R. D. (2012). The origin of chondrules and chondrites: Debris from low-velocity impacts between molten planetesimals? Meteoritics & Planetary Science, 47, 2170–2192. Scherstén, A., Elliott, T., Hawkesworth, C., Russell, S. S., and Masarik, J. (2006). Hf–W evidence for rapid differentiation of iron meteorite parent bodies. Earth and Planetary Science Letters, 241, 530–542. Schrader, D. L., Nagashima, K., Krot, A. N., et al. (2017). Distribution of 26Al in the CR chondrite chondrule-forming region of the protoplanetary disk. Geochimica et Cosmochi- mica Acta, 201, 275–302. Scott, E. R. D. (2007). Chondrites and the protoplanetary disk. Annual Review of Earth and Planetary Sciences, 35, 577–620. Shu, F. H., Shang, H., and Lee, T. (1996). Toward an astrophysical theory of chondrites. Science, 271, 1545–1552. Trieloff, M., Jessberger, E. K., Herrwerth, I., et al. (2003). Structure and thermal history of the H-chondrite parent asteroid revealed by thermochronometry. Nature, 422, 502–506. Trinquier, A., Elliott, T., Ulfbeck, D., et al. (2009). Origin of nucleosynthetic isotope hetero- geneity in the solar protoplanetary disk. Science, 324, 374–376. Van Orman, J. A., Cherniak, D. J., and Kita, N. (2014). Magnesium diffusion in plagioclase: Dependence on composition, and implications for thermal resetting of the 26Al-26Mg early solar system chronometer. Earth and Planetary Science Letters, 385,79–88. Villeneuve, J., Chaussidon, M., and Libourel, G. (2009). Homogeneous distribution of Al-26 in the Solar System from the Mg isotopic composition of chondrules. Science, 325, 985–988. Vockenhuber, C., Oberli, F., Bichler, M., et al. (2004). New half-life measurement of 182Hf: Improved chronometer for the early solar system. Physical Review Letters, 93, article # 172501. Wasson, J. T., and Kallemeyn, G. W. (1988). Compositions of chondrites. Philosophical Transactions of the Royal Society of London Series A-Mathematical Physical and Engin- eering Sciences, 325, 535–544. Zanda, B., Bourot-Denise, M., Perron, C., and Hewins, R. H. (1994). Origin and metamorphic redistribution of silicon, chromium, and phosphorus in the metal of chondrites. Science, 265, 1846–1849. Zook, H. A. (1981). On a new model for the generation of chondrules. Lunar and Planetary Science Conference, XII, 1242–1244. Abstract. 11 The Absolute Pb–Pb Isotope Ages of Chondrules Insights into the Dynamics of the Solar Protoplanetary Disk james n. connelly and martin bizzarro

Abstract

The parent nuclides 238U and 235U decay to 206Pb and 207Pb, respectively, with half-lives that makes this system uniquely suited to define the temporal framework of the solar protoplanetary disk, including the timing and duration of chondrule formation. Lead isotope data for 22 indi- vidual nebular chondrules indicate that the oldest chondrules formed contemporaneously with CAIs and that chondrules were recycled for ~4 Myr within the protoplanetary disk. Integrating the initial Pb isotopic compositions and ages of these individually-dated chondrules reveals that they appear to have formed in two distinct epochs. A primary phase of chondrule production occurred within 1 Myr of the formation of the Sun during the most energetic phase of the protoplanetary disk when mass accretion rates were highest. This epoch of primary chondrule production transitioned into a phase dominated by the reworking of existing chon- drules, which lasted for the remainder of the protoplanetary disk’s lifetime. Such a model is consistent with a transition from heating by shock waves related to gravitational instabilities during the more energetic first 1 Myr to heating by bow shocks around early formed planetesimals and planetary embyros. The age of chondrules from the CB meteorite Gujba formed from a vapor–melt plume caused by impacting planetary embyros indicates that the solar protoplanetary disk had dissipated within 4.5 Myr. The Pb–Pb ages require that any appearance of chemical or isotopic complementarity between matrix and chondrules does not imply rapid chondrule formation and accretion or that matrix and chondrules in a single chondrite group have a strict cogenetic relationship. In this view, inferences about the range of ages for chondrule formation based on a 182Hf–182W decay method and the assumption of cogenetically-formed matrix and chondrules cannot be meaningful. Finally, the preponderance of chondrules (>50%) having formed in the first 1 Myr of the protoplanetary disk lifetime is consistent with models of early, efficient growth of planetary embryos by pebble accretion.

The distribution of U-corrected Pb–Pb ages for chondrules is inconsistent with ages for chondrules based on the short-lived 26Al–26Mg system that implies a +1-Myr gap between the formation of CAIs and the first chondrules. We attribute the discrepancy between Pb–Pb and 26Al–26Mg ages to an incorrect assumption of homogeneously distributed 26Al relative to 27Al between the CAI- and chondrule-forming regions. Establishing precise Pb–Pb and 26Al–26Mg ages by the internal isochron method for the same individual chondrules would provide another

300 The Absolute Pb–Pb Isotope Ages of Chondrules 301 important test of homogeneity of 26Al relative to 27Al. In addition, higher-precision in situ measurements of younger chondrules that currently have no resolvable excess of 26Mg* would help test whether the 4-Myr duration of chondrule formation defined by the Pb–Pb ages is also recorded by the 26Al–26Mg system. Collectively, these two proposals may help solve the apparent discordance between the Pb–Pb and 26Al–26Mg systems that currently impedes progress of chondrule research within the cosmochemical community.

11.1 Introduction

Chondritic meteorites are sedimentary rocks that accreted during the first few million years of solar system formation and are chemically similar to the nonvolatile composition of the Sun (Scott and Krot, 2003). Having never experienced parent body melting, these primitive meteorites provide direct information on the formation and earliest evolution of the Sun and its protoplanetary disk. The major constituent of chondrites are chondrules, millimeter- sized inclusions that were once molten in the protoplanetary disk and accumulated in the disk midplane together with several other kinds of particles, including low temperature components. Chondrules are mainly composed of olivine and pyroxene, which crystallized within minutes to days at temperatures between ~1,800 and ~1,300 K (Scott, 2007). These silicate minerals are also the main constituents of the fine-grained matrix that mantles chondrules and other chondritic inclusions and fills the space between them in chondrites (Scott and Krot, 2005). Recent simulations indicate that the main growth of asteroids can result from the gas-drag–assisted accretion of chondrules (Johansen et al., 2015), a process analogous to pebble accretion. In these models, the largest planetesimals of a population with a characteristic radius of ~100 km undergo runaway accretion of chondrules, forming Mars-sized planetary embryos within a timescale of ~3 Myr. If chondrules represent the building blocks of planetary embryos and, by extension, terrestrial planets, understanding their chronology and formation mechanism(s) is critical to determine the point at which conditions became favorable for forming planetary bodies during the early evolution of the solar system. The chronology of chondrule formation is historically based on the short-lived 26Al- to-26Mg decay system (26Al decays to 26Mg with a half-life of 0.705 Myr; Norris et al., 1983). Assuming that 26Al/27Al ratio was uniform within the protoplanetary disk with the canonical value of ~5 10–5 commonly observed in the solar system’s first solids – calcium-aluminium-rich inclusions, CAIs – the 26Al–26Mg systematics of chondrules sug- gest that these objects formed >1 Myr after CAIs and rapidly accreted into chondrite parent bodies together with matrix in discrete events during the lifetime of the disk (Krot et al., 2009). This view is consistent with chondrule formation restricted to the high-density regions of the solar protoplanetary disk and the existence of a chemical complementarity between chondrules and coexisting matrix in chondrite meteorites (Hezel and Palme, 2010; Palme et al., 2015; Ebel et al., 2016). However, a number of studies have suggested that the initial 26Al/27Al ratio of disk material was lower than the canonical abundance based on CAIs such that the age gap between CAIs and chondrules is an artifact of the heterogeneous distribution of 26Al and has no chronological significance (Larsen et al., 2011; Schiller et al., 2015a, 2015b). 302 James N. Connelly and Martin Bizzarro

Of the various radiometric clocks, U-corrected Pb–Pb dating is the only method that can provide a high-resolution chronology of the first 10 Myr of the solar system that is not based on an assumption of a homogeneously distributed parent nuclide. It is based on two isotopes of U that decay in a chain to stable Pb isotopes, namely 235Uto207Pb with a half-life of ~0.7 Gyr and 238Uto206Pb with a half-life of ~4 Gyr. Using this approach, it has recently been demonstrated by Connelly et al. (2012) and Bollard et al. (2017) that chondrule formation started contemporaneously with the condensation of CAIs at 4,567.3 0.16 Myr and lasted ~4 Myr. Moreover, variability in the titanium and chromium stable isotope compositions of chondrules from individual chondrites suggests that these objects or their precursors were formed in distinct regions of the protoplanetary disk and subsequently transported to the accretion regions of their respective parent bodies (Trinquier et al., 2009; Olsen et al., 2016). These Pb–Pb ages combined with the Ti and Cr data are at odds with the traditional view of a short formation history for chondrule populations from individual chondrites based on chondrule-matrix complementarity, as well as the timescales and style of chondrite parent body accretion (Bizzarro et al., 2017). In this contribution, we review the current state-of-the-art data with respect to the absolute Pb–Pb chronology of individual chondrules from various chondrite groups and discuss how these data can be used to provide novel insights into the thermal and chemical evolution of the solar protoplanetary disk.

11.2 Uranium-Corrected Pb–Pb Dating

The U–Pb Decay System: Lead has four naturally occurring isotopes, 204Pb, 206Pb, 207Pb, and 208Pb, that were inherited from the molecular cloud in relative abundances that have been estimated using U-free troilite from the Type IAB Canyon Diablo iron meteorite (Tatsumoto et al., 1973). All Pb isotopes except 204Pb have increased their abundances through the radioactive decay of 238U, 235U, and 232Th. Only 204Pb preserves its primordial abundances such that the radiogenic ingrowth of Pb isotopes is most conveniently expressed as 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb ratios that evolved according to the general equation: 20x 20x 23z 232 Pb ¼ Pb þ U , Th λt 204 204 204 e 1 , (11.1) Pb today Pb t Pb today where t represents the time before present and λ represents the decay constant of the parent isotope. In objects with only initial Pb (Pbi) and radiogenic Pb (Pbr) present, this permits the theoretical possibility of calculating three independent ages based on the 235U–207Pb, 238U–206Pb, and 232Th–208Pb. Only the two U isotopes have decay constants large enough to provide ages for chondrules with errors of less than 1 Myr. Precleaning of all chondrules is required to remove ubiquitous modern terrestrial Pb contamination, an action that fractionates U/Pb ratios, rendering Equation 11.1 ineffective for determining ages. To circumvent the uncertainty of the U/Pb ratio, we employ the dual decay of the U–Pb system that results in unique radiogenic 207Pb/206Pb ratios for different times in the past, by combining equations 11.2 and 11.3: 207 207 235 Pb Pb U λ t ¼ þ ð 1 Þ 204 204 204 e 1 , (11.2) Pb today Pb t Pb today The Absolute Pb–Pb Isotope Ages of Chondrules 303 206 206 238 Pb Pb U λ t ¼ þ ð 2 Þ 204 204 204 e 1 , (11.3) Pb today Pb t Pb today to derive: 2 3 207Pb 207Pb λ t 6 204Pb 204Pb 7 235U e 1 1 4today t5 ¼ 206 206 λ t , (11.4) Pb Pb 238U e 2 1 204 204 today Pb today Pb t

207 206 207 206 where the left side of the equation equals the radiogenic ( Pb/ Pb) ratio [( Pb/ Pb)r]so that Equation 11.4 can be reduced to: λ t 207Pb 235U e 1 1 ¼ (11.5) 206 238 λ2t Pb r U today e 1 Equation 11.5 forms the basis for determining ages based on radiogenic 207Pb/206Pb ratios, including those based on the “inverse Pb–Pb diagram” that plots 207Pb/206Pb versus 204Pb/206Pb ratios (Figure 11.1). In this construct, the y-intercept of the array corresponds to the Pb isotope composition that lacks 204Pb and must, therefore, equal the radiogenic 207Pb/206Pb ratio. This approach works most effectively for highly radiogenic systems where at least some points fall close to the y-axis, thus minimizing the extrapolation to (and uncertainties of ) the y-intercept. All chondrules dated by the Pb–Pb method employ the inverse Pb–Pb method such that the discussion on the following pages concentrates on this approach. 207 206 Defining a Pb–Pb age based on the ( Pb/ Pb)r ratio requires that any initial Pb incorpor- ated into different phases had a single Pb isotopic composition, that the system remained closed, and that the 238U/235U ratio is known for each individual sample (but not the U/Pb ratio). It is

Figure 11.1 An inverse Pb–Pb diagram showing the results for a single chondrule after a single stage evolution with an elevated µ-value. In this case, back projection of the isochron passes through the solar system initial Pb composition. Modified from Connelly et al. (2017). 304 James N. Connelly and Martin Bizzarro important to note that these requirements are common to all chronometric methods based on radiogenic isotopes and are, therefore, not unique to the Pb–Pb method of dating. Whether an individual sample adheres to these requirements can be tested by evaluating the linearity of the data used to calculate ages. In some rare cases, the sample may not contain any initial Pb such 207 206 that a pure ( Pb/ Pb)r component is directly measured. The widespread addition of tetraethyl Pb to gasoline for over 50 years has contributed to the near-ubiquitous contamination of terrestrial Pb in meteorites. As such, this component of Pb in meteorites must be recognized and, in the most general cases, be removed from the sample for meteoritic Pb isotopic measurements to provide accurate age information (see also Connelly et al., 2017). Multistage Pb–Pb Evolution: The discussion in the preceding section describes a simple single-stage evolution of uranogenically-produced 207Pb and 206Pb after last closure (corres- ponding to the last melting event) from an initial Pb isotopic composition. The initial Pb isotopic composition will depend on the prehistory of the chondrule. If a chondrule or its precursor material never acquired a higher 238U/204Pb (µ) than the solar value of ~0.15, then the initial Pb isotopic composition will not become measurably different from the primordial Pb composition if the last closure occurred within the lifetime of the protoplanetary disk. As such, the projection of an isochron with a solar premelting µ-value should intersect the solar system initial value. In contrast, a chondrule that acquired an early, elevated µ-value through Pb devolatilization will have an evolved Pb isotopic composition commensurate with its µ-value and the length of time between acquiring the elevated µ-value and last closure. Figure 11.2 shows a two-stage evolution where a chondrule acquires a µ-value of 250 at t0 and accumulates

Figure 11.2 An inverse Pb–Pb diagram showing a theoretical example of a two-stage Pb evolution for a chondrule. The first stage evolves for 2 Myr with a µ-value of 250, at which time the chondrule is fully SS 2Myr melted and the intial Pb (Pbi ) in Pb-bearing phase(s) and radiogenic Pb (Pbr ) in U-bearing phase(s) are homogenized and incorporated into Pb-bearing phase(s). The second stage evolves with a µ-value of 2Myr 250 from 2 Myr after t0 with the reset initial Pb (Pbi ) until today, with the second stage radiogenic Pb today (Pbr ) accumulated in the U-bearing phase. The Absolute Pb–Pb Isotope Ages of Chondrules 305 radiogenic Pb for 2 Myr, at which time the sample is rehomogenized during the last melting event. This mixture represents an evolved isotopic composition (relative to the primordial values) that serves as the initial Pb composition for the second-stage evolution from 2 Myr after t0 until today. The inferred initial Pb composition of individual chondrules is used in the following text to evaluate the prehistories of chondrules. 238U/235U Composition of the Early Solar System: Brennecka et al. (2010) showed that the 238U/235U ratio historically used to calculate Pb–Pb ages (137.88; Steiger and Jäger, 1977) is not invariant in CAIs. However, several more recent studies have concluded that significant U isotope variability appears to be limited to the CAI-forming reservoir. Analyses of U isotopic compositions of the eucrite Juvinas, the unclassified basaltic meteorite NWA 2976, the enstatite chondrite Sahara 97159, the CI chondrite Ivuna, the CB chondrite Gujba, three chondrules from CV3 Allende, and a bulk sample of L3 ordinary chondrite NWA 5697 returned overlapping compositions with an average value of 137.786 0.013, which is inferred to represent the solar 238U/235U composition (Connelly et al., 2012). It is important to note that this solar 238U/235U value is based on high-precision measurements of chondrites that formed in different disk regions based on their distinct nucleosynthetic heritages. For example, it includes a metal-rich chondrite (Gujba) that is thought to have accreted beyond the orbits of the gas giants (Van Kooten et al., 2016), a CI chondrite that formed beyond the H2O ice-line and samples of primitive (ordinary chondrite) as well as differentiated (eucrite and angrite) asteroids believed to have accreted in the inner solar system. The solar 238U/235U composition suggested by Connelly et al. (2012) is in agreement with the analyses of six angrites by Brennecka and Wadhwa (2012) who reported an average value of 137.780 0.021. Moreover, Brennecka et al., (2015) reported the U isotopic compositions of pooled chondrules for Allende as well as bulk Allende and matrix from Allende that were overlapping with an average value of 137.786 0.004. Finally, Bollard et al. (2017) report U isotopic compositions for seven individual chondrules from the ordinary chondrite NWA 5697 that overlap the solar 238U/235U value of 137.786 0.013 (Connelly et al., 2012). Collectively, these results support two important conclusions about the U isotopic composition of the early solar system. First, the analyses of these diverse materials, including CAI-free chondrites, chondrules, and several differentiated bodies, strongly supports homogeneity of U in the solar system outside of the CAI-forming region with a present-day value of 137.786 0.013. Second, fractionation of U during high- temperature thermal and magmatic processing has not been documented in these studies within the analytical uncertainties. There are two studies that report U isotopic composition for early Solar System materials outside of this limit, namely Goldmann et al. (2015) and Kaltenbach (2012). As discussed in more detail in Connelly et al. (2017), this variability can be attributed to secondary processing, unrepresentative sampling or analyses of small amounts of U. Avoiding these issues, all studies to date report U isotopic compositions for a range of chondrites and achondrites that overlap and, therefore, support a present-day bulk solar system value of 137.786 0.013. For the sake of dating chondrules by the Pb–Pb method, it is clear from Figure 11.3 that all U isotopic measurements of bulk and individual chondrules overlap this value. While these studies support the use of a solar U isotope composition of 137.786 0.013 to calculate Pb–Pb ages of meteoritic material, it must be emphasized that the U isotopic composition should be verified whenever there is sufficient sample size. 306 James N. Connelly and Martin Bizzarro

Figure 11.3 238U/235U ratios of individual chondrules from the CV chondrite Allende (individual chondrules by Connelly et al., 2012; and pooled chondrules by Brennecka et al., 2015) and ordinary chondrite NWA5697 (Bollard et al., 2017). All analyses are consistent with a 238U/235U ratio of 137.786 0.013 as documented by Connelly et al. (2012) for the present-day solar value and shown in the gray band. The larger uncertainties on the analyses of individual chondrules reflect the smaller amounts of U in these analyses compared to the larger Allende chondrules.

11.3 Pb–Pb Ages of Chondrules

Early attempts to date chondrules by the Pb–Pb method involved analyzing individual pieces of different chondrules obtained from a lightly-crushed CR chondrite Acfer 059 (Amelin et al., 2002) or pooling chondrules from the CV chondrite Allende (Amelin and Krot, 2007; Connelly and Bizzarro, 2009). Amelin et al. (2002) reported an age of 4,563.66 Ma, while Amelin and Krot (2007) and Connelly and Bizzarro (2009) obtained ages of 4,565.6 1.0 Ma and 4,564.32 0.81 Ma, respectively (all ages adjusted for a U isotopic composition of 137.786 after Connelly et al., 2012). That individual pieces of CR chondrules fall on a single isochron is consistent with the analyzed chondrules having formed within a narrow time frame at 4,563.66 Ma. The ages of 4,565.6 1.0 Ma and 4,564.32 0.81 Ma for pooled CV chondrules overlap with each other and most likely represent the average age of formation of the chondrules analyzed in these studies, but they do not constrain the duration of their formation. There are now 22 individual nebular chondrules dated by the Pb–Pb method (Figure 11.4) with measured 238U/235U ratios for six of these objects (Connelly et al., 2012; Bollard et al., The Absolute Pb–Pb Isotope Ages of Chondrules 307

Figure 11.4 Pb–Pb dates for individual chondrules from NWA 5697 (L3.10), NWA 6043 (CR2), NWA 7655 (CR2), and Allende (CV3) as reported in Bollard et al. (2017). The timing of CAI formation is accepted to be 4,567.30 0.16 Ma (Connelly et al., 2012). The CB chondrules are interpreted as having formed from colliding planetesimals at 4562.49 0.21 Ma (Bollard et al., 2015a). Modified from Bollard et al. (2017).

2017). The dated nebular chondrules come from CV (Allende), ordinary (NWA 5697), and CR (NWA 6043 and 7655) chondrites and comprise all major petrographic classes, including both porphyritic and nonporphyritic texture types as well as FeO-rich and FeO-poor varieties (see Bollard et al., 2017 for full descriptions). The six chondrules from ordinary chondrite NWA 5697 with measured 238U/235U ratios define ages that range from 4,567.61 0.54 Ma to 4,564.65 0.46 Ma. This age range is comparable to that of 4,567.57 0.56 Ma to 4,563.24 0.62 Ma defined by all remaining chondrules for which the solar 238U/235U value is used to calculate the Pb–Pb ages (Figure 11.4). Using the 238U/235U of the bulk host chondrites or, alternatively, that for multichondrule analyses of the same meteorites returns ages that are within 100,000 years of that obtained using the solar 238U/235U value of 137.786 0.013 (Bollard et al., 2017) and, hence, well within the uncertainties of the final ages reported. These data convincingly demonstrate that the production and melting of chondrules began contemporaneously with CAI formation and continued for ~4 Myr (Figure 11.4). In addition, there are four chondrules from the CB chondrite Gujba that are interpreted to have been generated in an impact plume (Krot et al., 2005a; Bollard et al., 2015a). The CB chondrites are metal-rich meteorites with characteristics that sharply distinguish them from other chondrite groups. Citing their unusual chemical and petrologic features and a young formation age of bulk chondrules from the CB chondrite Gujba, Krot et al. (2005a) interpreted these chondrules to have formed late in the evolution of the protoplanetary disk during an impact between planetesimals after the disk was mainly cleared of gas and dust. Pb–Pb analyses of fragments of four individual chondrules defined a weighted mean age of 4,562.49 0.21 Myr (Bollard et al., 2015a), consistent with their late origin from the vapor–melt impact plume generated by colliding planetesimals. 308 James N. Connelly and Martin Bizzarro

Linking Pb–Pb Ages to the Last Chondrule-Melting Event: To understand the signifi- cance of Pb–Pb ages of chondrules, the Pb isotopic compositions of the individual mineral phases for two chondrules from the NWA 7655 CR chondrite and the NWA 5697 ordinary chondrite were measured using in situ methods to determine the host phase of U (Bollard et al., 2017). This study established that the radiogenic Pb (and hence U) resides in the fine-grained chondrule mesostasis, whereas nonradiogenic Pb is predominantly hosted by sulfides in the mesostasis. This result is consistent with partition coefficient data predicting that U typically behaves as an incompatible element in most silicates (Blundy and Wood, 2003) as well as the systematics of the sequential acid dissolution approach used to define the Pb–Pb isochron ages (Connelly and Bizzarro, 2009). In detail, the fractions with the most radiogenic Pb isotope compositions are typically released with the first use of HF acid, which preferentially dissolves nonrefractory silicate phases. Therefore, it is clear that the Pb–Pb isochron ages of individual chondrules reflect crystallization of the mesostasis that can be directly related to the last chondrule-melting event.

11.3.1 Comparing Pb–Pb Ages to Other Radiogenic Isotope Systems 26Al–26Mg Systematics of Chondrules: Apart from Pb–Pb dating, the only other method utilized to provide insights into the timing of crystallization of individual chondrules via the internal isochron approach is based on the 26Al-to-26Mg short-lived decay system (see Chapter 9 for a review of this method). However, the validity of 26Al–26Mg relative ages is based on the assumption that the 26Al/27Al ratio was homogeneous in the solar protoplanetary disk with an initial value of ~5 10–5 at the time of CAI formation. Attempts to provide a chronology of chondrule formation based on this system indicate that chondrule formation began ~1.5–2.0 Myr after CAI formation (Krot et al., 2009). In the following text we discuss possible explan- ations to account for the inconsistency between ages of individual chondrules derived by the Al–Mg and U–Pb systems. One possibility is that the age variability or, alternatively, the preponderance of old ages inferred from Pb–Pb dating may be due to mass-independent fractionation caused by the progressive step-leaching dissolution technique (Amelin et al., 2009) utilized in Connelly et al., (2012) and Bollard et al. (2017). Three lines of evidence indicate that this is not the case. First, independent estimates of the Pb–Pb age for the SAH99555 angrite using the step-leaching dissolution technique (Connelly et al., 2008) and a more traditional mineral separation approach (Amelin, 2008) return ages concordant within 280,000 405,000 years. Second, independent estimates for the timing of CAI formation using the different techniques mentioned above (Amelin et al., 2010; Connelly et al., 2012) yield ages that are within 140,000 505,000 years. In both cases, the potential offset between the two techniques is within the typical uncertainties of individual chondrule ages. Third, using the step-leaching dissolution technique, Bollard et al. (2015a) report Pb–Pb ages for four individual chondrules from the Gujba metal-rich chondrite that are identical within 340,000 years, in accordance with their formation from a vapor–melt plume produced by a giant impact (Krot et al., 2005a). Collectively, these observations suggest that the published Pb–Pb dates summarized here for individual chondrules are accurate to within their stated uncertainties. The Absolute Pb–Pb Isotope Ages of Chondrules 309

Assuming that the Pb–Pb dates are accurate, a possible explanation for the observed age offset between the Pb–Pb and 26Al–26Mg ages is that the assumption of initial 26Al/27Al disk homogeneity is incorrect. Indeed, several recent studies have documented a reduced initial 26Al/27Al ratio in solids that accreted to form protoplanets (Larsen et al., 2011; Schiller et al., 2015a), which would translate into younger 26Al–26Mg ages relative to the Pb–Pb dates. In Figure 11.5, we show the ages of NWA 5697 chondrules inferred from Pb–Pb dating relative to 26Al–26Mg ages of chondrules from the most primitive ordinary chondrites, assuming that the initial inner disk 26Al/27Al inventory was ~1.5 10–5 (Schiller et al., 2015a). Under this assumption, the two dating methods return comparable age distributions and age range of chondrule formation, consistent with the hypothesis of a reduced inner disk 26Al inventory relative to the canonical abundance. We note that preliminary results reporting the 26Al–26Mg

A 7 U-corrected Pb-Pb ages

6 NWA 5697 (L3.10)

5

4

3

2

Number of chondrules 1

0 0 1234 B 25 26Al-26Mg ages

Semarkona (LL3.00) 20 LEW86134 (L3.00) QUE97008 (L3.05) 15 Bishunpur (L3.10)

10

5 Number of chondrules

0 01234 Time after CAI formation (Myr)

Figure 11.5 Histograms depicting the absolute and relative ages of chondrules from unequilibrated ordinary chondrites of low petrologic type (3.1) based on internal isochron relationships. (a) Pb–Pb dates of chondrules from the NWA 5697 ordinary chondrite (Bollard et al., 2017). (b) 26Al–26Mg ages of chondrules from Semarkona, LEW 86134, QUE97008 and Bishunpur ordinary chondrites (Huss et al., 2001; Mostefaoui et al., 2002; Rudraswami and Goswami, 2007). The relative 26Al–26Mg ages are calculated assuming that the precursor material from which these chondrules formed had a reduced initial 26Al/27Al value corresponding to ~1.5 10–5 (Schiller et al., 2015a). Three chondrules record initial 26Al/27Al slightly higher than 1.5 105 but are within analytical uncertainty of this estimate and, hence, have been assigned a T = 0 formation age for simplicity. Modified from Bollard et al. (2017). 310 James N. Connelly and Martin Bizzarro and Pb–Pb ages of the same individual chondrules are consistent with the proposal of a reduced inner disk 26Al inventory (Bizzarro et al., 2014; Bollard et al., 2015b). Bizzarro et al. (2004) reported 26Al–26Mg systematics of individual bulk chondrules from the Allende CV chondrite and suggested that a fraction of these objects formed with the canonical 26Al/27Al value of ~5 10–5, suggesting coeval formation of chondrules and CAIs. This observation is this at odds with the proposal of a reduced inner solar system 26Al in disk material, including chondrules. However, using refined methods for ultra-high-precision Mg isotope measurements (Bizzarro et al., 2011), a recent study (Olsen et al., 2016) of bulk individual chondrules from the Efremovka and NWA 3118 CV chondrites did not find excesses comparable to that reported by Bizzarro et al. (2004). A possible explanation to account for this discrepancy is that the chondrules with resolvable excess 26Mg* reported in Bizzarro et al. (2004) have incorporated refractory, 26Mg*-rich CAI material. Although this possibility was discarded based on the stable 25Mg isotope composition of chondrules, knowledge of the Mg stable isotope variability in CAIs was limited at the time of publication of the Bizzarro et al. (2004) report. Reassessment of this possibility based on published stable Mg isotope compos- itions of CAIs (i.e., Larsen et al., 2011) suggest that admixing of CAI-like material in chondrule precursors can account for the excess 26Mg* reported in Bizzarro et al. (2004). Lastly, it has been inferred that metal-rich chondrites, including CR chondrules, formed from an 26Al-poor reservoir located beyond the orbits of the gas giant planets (Olsen et al., 2016; Van Kooten et al., 2016). More specifically, the 26Mg* and 54Cr isotope compositions of CR chondrules require significant amounts (25–50 percent) of primordial 26Al-free molecular cloud matter in their precursor material. Accepting the Pb–Pb age distribution of CR chondrules at face value predicts that many of these objects will lack evidence for live 26Al. These predictions are in line with recent 26Al–26Mg systematics of CR chondrules, which record initial 26Al/27Al values of typically less than 3 10–6 and more than 50 percent of the objects investigated having no evidence of live 26Al at the time of their crystallization (Nagashima et al., 2014; Schrader et al., 2017). These observations, together with the clear evidence for 26Al heterogen- eity at the time of CAI formation (Holst et al., 2013; Thrane et al., 2008), emphasize that the 26Al–26Mg system cannot provide an accurate chronology of chondrule formation and, hence, disk processes. 182Hf–182W Systematics of Chondrules: Budde et al. (2016a) attempted to use the short- 182 182 lived Hf– W decay system (t1/2 ~ 8.9 Myr) to establish the timing and duration of chondrule formation in the CV chondrite Allende (see Chapter 10 for a review of this method). Since single chondrules have insufficient W to measure the isotopic composition accurately enough to constrain an isochron, they pooled 100s to 1,000s of chondrules together for single analyses. Six analyses of pooled chondrules were combined with three matrix and two bulk Allende analyses to define an array in 182W/184W versus 180Hf/184W space that they interpreted to represent an isochron. Furthermore, they noted an apparent isotopic complementarity between the chondrules and matrix reflected in the nucleosynthetic inventory of 183W. Using the Hf–W decay system, they define an age of 2.2 0.8 Ma after CAI formation as their preferred age of chondrule formation in this meteorite. They conclude that this represents an average age of the “vast majority of Allende chondrules” and that they formed “within a narrow time interval of less than 0.6 Ma” (the analytical scatter on the array). However, in the following text we show that the existence of variable W nucleosynthetic anomalies among chondrules and The Absolute Pb–Pb Isotope Ages of Chondrules 311 matrix demonstrates that the fundamental assumption of isotopic equilibrium that underpins the isochron method is violated. For the conclusion of Budde et al. (2016a) to be correct requires that the 182W/184W ratio fully equilibrated between matrix and chondrules at the time of Hf/W fractionation. As explored in Figure 11.6, there are two end-member scenarios for the evolution of the W isotope composition during Hf/W fractionation that they associate with chondrule formation. In the first scenario, if a low Hf/W phase is transferred from chondrule to matrix with its full unequilibrated W inventory intact (i.e., not equilibrated with the rest of the chondrule), the compositions of chondrules and matrix will move in opposite directions along a sloping line at the time of chondrule formation (Figure 11.6a). If the timing of last Hf/W fractionation was t0, the final isochron slope will reflect the age of the solar system (assuming no other alteration events affecting these reservoirs before or after chondrule formation). In the second scenario, W in the chondrules is first internally equilibrated so that chondrules have a homogeneous

Figure 11.6 End-member scenarios for the redistribution of Hf and W during chondrule formation. a) Shows an initial condition for the dust at t0 for the solar system (1) that evolves to the time of chondrule formation (2) when a W-rich phase is transferred to the matrix without it having first equilibrated with the rest of the chondrule and (3) the final isochron today. b) Shows an initial condition for the dust at t0 for the solar system (1) that evolves to the time of chondrule formation (2) when W is homogenized within chondrules before a W-rich phase is transferred to the matrix and (3) the final isochron today. 312 James N. Connelly and Martin Bizzarro

182W/184W composition before W is preferentially lost to the matrix (Figure 11.6b), resulting in both matrix and chondrules having the same 182W/184W ratio, but with different Hf/W ratios, provided the matrix and all chondrules had the same bulk initial 182W/184W and Hf/W ratios prior to chondrule formation. Only this case generates a final isochron slope corresponding to the correct age of chondrule formation. However, the nucleosynthetic 183W/184W differences between matrix and chondrules indi- cate that full W isotopic equilibration could not have occurred within the chondrules before the fractionation of Hf–W. This lack of isotopic equilibration is also demonstrated by the different Mo isotope signatures between matrix and chondrules (Budde et al., 2016b). If Mo- and W-rich pre-solar metal grains in the matrix are indeed the presolar carrier of nucleosynthetic W isotope anomalies, as proposed by Budde et al. (2016a and b), the conditions for an isochronous relationship between matix and chondrules are violated and the ages cannot reflect the timing or duration of chondrule formation. In contrast, the 1.77 1.55 Myr (after t0) age corresponding to the array defined by the six bulk chondrule analyses may provide an accurate estimate of the average age, but not the duration, of chondrule formation. This result is fully consistent with the ages of Allende chondrules defined by the Pb–Pb method.

11.4 Implications of Chondrule Pb–Pb Ages for Understanding Disk Processes

Mass Transport and Storage: Figure 11.7 shows the Pb–Pb age distribution of chondrules, which infers a progressive reduction in chondrule production rate through time. Approximately 50 percent of these chondrules formed within the first Myr of the solar protoplanetary disk, implying that chondrule formation was more efficient in early times. The preponderance of ancient chondrules is unexpected, given that the residence time of millimeter-sized solids in evolving disks is predicted to be extremely short relative to the typical lifetimes of

Figure 11.7 Histogram of Pb–Pb ages of individual chondrules from Bollard et al. (2017) showing that over 50 percent of the dated chondrules have a formation age within 1 Myr of CAI formation at 4,567.30 0.16 Ma (Connelly et al., 2012). Modified from Bollard et al. (2017). The Absolute Pb–Pb Isotope Ages of Chondrules 313 protoplanetary disks due to aerodynamic drag (Weidenschilling, 1977). Although mechanisms exist to limit the inward drift of millimeter-sized solids such as dust trapping, these operate on timescales typically shorter than the age variability reported here for chondrules (Testi et al., 2014). Thus, the presence of an ancient component and the significant age variability amongst chondrules from individual chondrites (Figure 11.4) require outward mass transport and/or storage of chondrules during the lifetime of the disk. Recycling of early-formed refractory solids such as CAIs during chondrule-forming events has been observed in a number of primitive chondrites (Itoh and Yurimoto, 2003; Krot et al., 2005b, 2017). Together with petrological evidence suggesting that many chondrules experienced multiple melting events (Krot and Wasson, 1995; Krot et al., 2004), these observations raise the possibility that younger chondrule populations predominantly reflect remelting and, hence, recycling of chondrules formed at earlier times. This possibility can be tested by considering the initial Pb isotopic composition of each individual chondrule in light of its age of last closure. Evidence for Early Formation and Protracted Recycling of Chondrules: It is well documented that the brief heating events resulting in chondrule melting leads to evaporative Pb loss and, therefore, enhancement of the 238U/204Pb ratio (μ) relative to the solar value of ~0.15 given the refractory nature of U (Connelly et al., 2012). This is in line with the μ-values published for individual chondrules that range from ~2 to ~183 (Bollard et al., 2017), which corresponds to up to ~99.94 percent of Pb loss relative to solar abundances. As described in the previous text, an early-formed chondrule that acquires a relatively high μ-value will accumulate measurable amounts of radiogenic Pb relative to the primordial levels over the lifetime of the protoplanetary disk. Remelting of such a chondrule at later times will result in an evolved initial Pb isotope composition relative to a chondrule that only experienced a single melting event. Thus, chondrules with protracted complex thermal histories involving more than one melting episode are expected to record evolved initial Pb isotopic compositions relative to objects formed from precursors with near-solar μ-values. Figure 11.8 shows the back projection of the regressions for individual chondrules with well-defined isochron relationships, which enables the assessment of their initial Pb isotope compositions. Collectively, the chondrules show variable initial Pb isotope compositions that correspond to a range of ~120 ε-units in the 207Pb/206Pb ratio, with most chondrules recording evolved compositions compared to the most primitive chondrules (chondrules 2-C1 and C1 of Bollard et al., 2017). A component of the elevated 207Pb/206Pb ratio could also reflect mass-dependent heavy isotope enrichment associated with evaporative loss of Pb. To assess this possibility, Bollard et al. (2017) measured the isotope composition of Zn in a subset of chondrules, an element with comparable volatility as Pb (Lodders, 2003). The δ66Zn values range from 0.26 0.05 to 2.20 0.05‰, representing light isotope enrichment that is not expected during simple evaporation, which typically results in heavy Zn isotope composition (Moynier et al., 2009). Thus, the light and limited variability of Zn isotope fractionation in these chondrules (~2‰) confirms that the range of the 207Pb/206Pb (~12‰) ratios is not a product of mass-dependent fractionation. Importantly, Figure 11.9 shows that the initial Pb isotope compositions of chondrules are positively correlated with their crystallization ages. Whereas only chondrules formed in the first million years of the disk evolution record primitive compositions, chondrules formed at later times show progressively more evolved initial Pb isotope compositions. The progressively more evolved initial Pb isotope compositions and the lack of primitive initial Pb 314 James N. Connelly and Martin Bizzarro

Figure 11.8 Initial Pb isotopic compositions of individual chondrules. The initial Pb isotope compositions are defined by the intersection of the individual isochrons and a Pb evolution array anchored to the solar 207 206 system’s primordial Pb isotopic composition of Tatsumoto et al. (1973). The ε( Pb/ Pb)initial values shown in Figure 11.7 reflect the offset from the two least evolved chondrules shown in the blue box. Individual chondrule data points have been displaced to the left and right of the array for clarity. Modified from Bollard et al. (2017).

207 206 207 206 Figure 11.9 Initial Pb/ Pb ratios and age variation diagram. The ε( Pb/ Pb)initial values are reported in deviations per 10,000 (ε-unit) from the composition defined by the least evolved chondrules 207 206 shown in Figure 11.7. The blue box reflects the range of ε( Pb/ Pb)initial compositions of the >1 Myr chondrules back-calculated to 4,566.8 Ma. Modified from Bollard et al. (2017). The Absolute Pb–Pb Isotope Ages of Chondrules 315 recorded by younger chondrules suggest that these formed from precursors having already accumulated radiogenic Pb, which is indicative of material characterized by an elevated μ-value. The evolving initial Pb isotope compositions of chondrules during the lifetime of the disk are most easily understood as reflecting recycling and, hence, remelting of preexisting chondrules. Thermally-processed precursors may represent chondrule fragments or, alternatively, entire chondrules having experienced more than one chondrule-forming event. Dynamical Evolution of the Solar Protoplanetary Disk: Unlike CAIs that formed as fine- grained condensates near the proto-Sun during collapse of the presolar molecular cloud core, the majority of chondrules are thought to be products of flash heating of dust aggregates in different disk regions, possibly caused by shock wave heating (Ciesla and Hood, 2002). As shocks require the existence of a gaseous disk, the absolute ages of nebular chondrules provide a means of assessing the lifetime of the protoplanetary disk, including the timing of its establishment relative to the proto-Sun. It is significant that six chondrules define absolute ages identical to that of CAIs within analytical uncertainty, defining a population with a weighted mean age of 4,567.41 0.19 Ma (Bollard et al., 2017). It is inferred from both astronomical observations and numerical simulations of young stars and their disks that protoplanetary disks are formed shortly after star formation during the deeply-embedded stage (Tobin et al., 2015; Tomida et al., 2015). The age uncertainty associated with the ancient chondrule population mentioned in the previous text indicates that these objects formed at most ~138,000 years after CAIs. This defines an upper age limit for the establishment of the protoplanetary disk after CAI formation and, by extension, collapse of the proto-Sun. It is inferred that the disk was largely cleared of dust and gas (Morris et al., 2015) at the time of formation of the impact-generated chondrules from the Gujba metal-rich chondrite at 4562.49 0.21 Ma (Bollard et al., 2015a). Combined with the age and uncertainty of the youngest nebular chondrule from Bollard et al. (2017), this defines the minimum and maximum lifetimes of ~3.4 and ~4.5 Myr, respectively, for the active phase of the solar protoplanetary disk. We note that these timescales are in keeping with astronomical observations of young stars and their disks (Evans et al., 2009) and, importantly, indicate that the formation of a disk amiable to the production of asteroidal bodies and planetary embryos occurred shortly after collapse of the molecular cloud and formation of the proto-Sun. The emerging absolute age data and initial Pb isotope compositions of individual chondrules may provide critical insights into the accretion history and thermal processing of dust in the protoplanetary disk. It is well understood that mass accretion to the protostar occurs from the surrounding envelope via a circumstellar disk structure during the early stages of low mass star formation (Williams and Cieza, 2011). From an astronomical perspective, this epoch represents the deeply-embedded phase of star formation that only lasts for the first 500,000 years of the disk lifetime. During this time, fresh, volatile-rich envelope material is accreted to the disk and, thus, is available to participate in the formation of early solar system solids. It is apparent from the integrated dataset presented in Bollard et al. (2017) that chondrules formed within the first ~1 Myr of the disk lifetime record primitive initial Pb isotope compos- itions consistent with incorporation of thermally-unprocessed material. In contrast, younger chondrules with evolved compositions require limited or no admixing of primordial dust to their precursors. If correct, the chondrules may be recording the transition between two distinct accretionary regimes during the early evolution of the protoplanetary disk. The first ~1 Myr reflecting the main epoch of accretion of envelope material to the disk, which represents the 316 James N. Connelly and Martin Bizzarro regime of primary chondrule production. Conversely, the >1 Myr regime may reflect an epoch dominated by the transport and recycling of disk material, including remelting of chondrules formed during the first epoch. The limited evidence for admixing of primordial dust in >1 Myr chondrules indicates that the envelope of accreting material surrounding the proto-Sun was largely cleared by that time.

11.5 Chondrule Formation and Recycling Mechanisms

It is inferred that chondrules formed by the melting of isotopically diverse solid precursors in dust-rich regions of the protoplanetary disk during repeated, localised transient heating events (Scott, 2007). The thermal histories of chondrules, as well as the high solid densities required for their formation (Alexander et al., 2008), have been used to argue for shock wave heating as the primary source of energy for the thermal processing and melting of chondrule precursors. The 26Mg* and 54Cr systematics of CR chondrules suggest that these objects formed in a reservoir distinct from chondrules from other chondrite groups, a reservoir possibly located beyond the orbits of the gas giant planets (Olsen et al., 2016; Van Kooten et al., 2016). As such, these data require an efficient heat source resulting in chondrule formation both in the inner and outer disk. Viable mechanisms for producing chondrule-like objects in gaseous disks include shocks produced by disk gravitational instability (Boss and Durisen, 2005; Boley and Durisen, 2008) and eccentrically orbiting planetesimals and planetary embryos (Hood et al., 2009; Morris et al., 2012; Boley et al., 2013). Bollard et al. (2017) suggest that the primary chondrule production occurred during the deeply-embedded stage of the proto-Sun, which is characterized by significant accretion of envelope material to the disk. The most efficient source of shocks during this epoch is shock fronts associated with spiral arms generated in a gravitationally unstable disk (Boley and Durisen, 2008; Pérez et al., 2016). This shock regime is modeled to be highly efficient in the inner disk region and possibly extend to ~10 astronomical units (Boley and Durisen, 2008; Desch et al., 2012), which provides an efficient mechanism for the global thermal processing of disk solids in early times. As mass accretion to the disk decreases, the resulting disk mass is thought to be too low to sustain gravitational instabilities such that a different shock source(s) may be required to remelt chondrules from ~1 to ~4 Myr after proto- Sun collapse. Bow shocks resulting from planetesimals and planetary embryos traveling on eccentric orbits provide a possible source of heating for the thermal processing of solids in the >1 Myr regime. Thus, it is apparent that different heat sources are likely required to satisfy the Pb–Pb isotope ages of individual chondrules, which point to a multiplicity of chondrule-forming and remelting mechanisms.

11.6 Chondrule Age Variability and the Chondrule-Matrix Complementarity

The bulk chemistry of chondrites is defined by the two major components, chondrules and matrix. Numerous studies have investigated the apparent chemical relationship between chon- drules and matrix in individual chondrite groups (Hezel and Palme, 2010; Palme et al., 2015; Ebel et al., 2016 and Chapter 4). These studies have concluded that the average compositions of chondrules and matrix are typically different for a number of elements in an individual The Absolute Pb–Pb Isotope Ages of Chondrules 317 chondrite, whereas the bulk composition, which reflects mixtures of chondrules and matrix, has approximately solar elemental abundances. This so-called chondrule-matrix complementarity has been used to argue for a genetic link between these two components and, therefore, formation from a single reservoir. Given the short residence time of solids in a protoplanetary disk due to gas drag (Weidenschilling, 1977), this model predicts that all chondrules from a single chondrite should be in isotopic equilibrium and have the same age. Thus, the chondrule- matrix complementarity requires that chondrule formation and asteroidal accretion are intim- ately linked. However, the observed 54Cr variability suggests that chondrules from individual chondrite groups formed from isotopically diverse precursor materials in different regions of the protoplanetary disk and were subsequently transported to the accretion regions of their respect- ive parent bodies (Olsen et al., 2016). This is consistent with the presence of age variability of ~4 Myr between chondrules from individual chondrites (Connelly et al., 2012), which requires transport and/or storage. At face value, these data appear inconsistent with the concept of chondrule-matrix complementarity as originally envisaged, namely that all chondrules from an individual chondrite are genetically related to the coexisting matrix (see also Chapter 5). Recent models of evolving viscous disks, however, suggest that a complementary relation- ship between chondrules and dust can be preserved for long timescales, provided that the decoupling between chondrules and gas is limited (Goldberg et al., 2015). In these models, various chondrule populations remained in complementarity such that the bulk contribution from each source is chemically solar and, thus, so is the final mixture. However, these experi- ments assume that the main transport mechanism of chondrules occurs through outward diffusion via the disk midplane. In disk models where outward transport of material is associated with stellar outflows (Shu et al., 1996; Hu, 2010), the coarse-grained dust component (i.e., CAIs and chondrules) is not expected to be efficiently coupled to the gas and, thus, it is unclear how complementarity can be preserved. It is possible that the observed chondrule- matrix complementarity is an expression of the generic process of chondrule formation and does not reflect a strict genetic link. In this view, the matrix comprises a complement related to the chondrule formation process (Alexander, 2005) such that the bulk composition of the matrix is shifted from its starting composition and, thus, appears complementary to a chondrule compos- ition. This does not require matrix and chondrules to be genetically linked in an individual chondrite, but merely that some matrix has experienced earlier chondrule formation events. If correct, fractions of the matrix in a particular chondrite may be complementary to chondrule populations in other chondritic meteorites. Likewise, the apparent isotopic complementarity for siderophile elements such as W and Mo (Budde et al., 2016a, 2016b) between chondrules and matrix may reflect the selective destruction and/or removal of isotopically anomalous metal phases during chondrule formation.

11.7 Implications for Mechanism and Style of Asteroidal Accretion

The protracted timescales for the formation of chondrules inferred from the U-corrected Pb–Pb dating method and 26Al–26Mg dating are inconsistent with a model where chondrule formation and planetesimal accretion are linked temporally and spatially. Indeed, the ~4 Myr formation interval for chondrules from various chondrite groups indicates prolonged accretion timescales 318 James N. Connelly and Martin Bizzarro for chondritic parent bodies and a possible decoupling of chondrule formation and asteroidal accretion in space and time. Recent numerical simulations suggest that the formation of asteroidal bodies and planetary embryos may be a two-step process, where first generation planetesimal seeds of approximately 50 km diameter form rapidly via streaming instabilities (Johansen et al., 2007; Lambrechts and Johansen, 2012) followed by the protracted gas-drag– assisted accretion of chondrules during the lifetime of the protoplanetary disk (Johansen et al., 2015). The gas-drag–assisted accretion of chondrules onto planetesimals is a process analogous to pebble accretion, which is the accretion of centimeter- to meter-sized particles loosely bound to the gas onto planetesimals seeds (Bitsch et al., 2015). In these simulations, the largest planetesimals of a population with a characteristic radius of approximately 100 km undergo runaway accretion of chondrules forming Mars-sized planetary embryos within a timescale of ~3–5 Myr. This timescale is in agreement with the timing of formation and differentiation of Mars inferred from 182Hf–182W chronology (Dauphas and Pourmand, 2011). An abundance of chondrules in the first million years of disk evolution as indicated by the Pb–Pb isotope ages suggests that the precursor material required to drive efficient and rapid formation of planetary objects via chondrule accretion was available at early times. This is consistent with recent astronomical observations of young protoplanetary disks (Carrasco-Gonzalez et al., 2016) that suggest the accumulation of sizeable planetary objects over timescales comparable to that inferred for the primary epoch of chondrule production in the solar system. Finally, a model of continuous asteroidal accretion during the lifetime of the protoplane- tary disk (Larsen et al., 2016) has important implications for the thermal evolution of asteroidal bodies, given that the accretion process is completed beyond the time when 26Al can provide enough energy to induce heating and differentiation (Larsen et al., 2011). Thus, protracted asteroidal accretion predicts the existence of partially differentiated aster- oidal bodies, namely onion-shell structured bodies with differentiated interiors consisting of silicate mantles and metallic cores surrounded by unmelted chondritic crusts (Weiss and Elkins-Tanton, 2013). Although controversial, this proposal is apparently supported by the discovery of remnant magnetism in chondritic meteorites suggesting the existence of a dynamo field (Carporzen et al., 2011), which can only occur through the establishment of a convecting metallic core.

11.8 Summary and Future Work

U-corrected Pb–Pb dating is the only approach that can establish a chronological framework for understanding the formation history of individual disk solids that is not based on an a priori assumption of a homogeneously-distributed parent nuclide. Using this method, it is now established that the production and melting of nebular chondrules began contemporaneously with CAI formation and continued for ~4 Myr. Combined with the absolute Pb–Pb ages of individual chondrules, the initial Pb isotope compositions of these objects provide insights into the thermal history of their precursors. Collectively, these data indicate that the primary production of chondrules was restricted to the first million years after formation of the Sun, and that these existing chondrules were recycled for the remaining lifetime of the protoplanetary disk. This is consistent with an epoch of primary chondrule formation during the early The Absolute Pb–Pb Isotope Ages of Chondrules 319 high-mass accretion phase of the protoplanetary disk that transitions into a longer period of chondrule reworking during a more quiescent phase of mass accretion. Protracted chondrule formation and reworking during the disk lifetime suggest multiple heating sources that may be linked to the evolving dynamics of the protoplanetary disk. The emerging picture of protracted chondrule formation and recycling based on the U- corrected Pb–Pb dating method is at odds with the implications of the perceived genetic complementarity between chondrules and the host matrix in chondrite meteorites. In the current view, chondrule formation and their subsequent accretion onto chondrite parent bodies must have all occurred rapidly in a localized part of the disk. Problems with the model of chondrule- matrix complementarity may already be apparent from the inferred range of 26Al–26Mg ages of chondrules from individual chondrite groups that indicates >1 Myr formation interval, irre- spective of whether the ages are accurate relative to CAI formation. This point alone casts serious doubts on attempts to constrain the chronology of chondrule formation using the 182Hf–182W systematics of multiple chondrule and matrix fractions, an approach that requires contemporaneous formation of all chondrules and a genetic link between chondrules and matrix. It appears that a significant amount of work is required to resolve the current discrepancy that exists between the absolute and relative chronometers and, by extension, evaluate the concept of chondrule-matrix complementarity. One avenue of research is determining 26Al–26Mg and U- corrected Pb–Pb ages for individual chondrules from primitive chondrites. This will provide an accurate determination of the initial 26Al/27Al ratios in chondrule forming regions, which will allow us to test the assumption of initial 26Al disk homogeneity. Moreover, it is also critical to determine the full duration of the chondrule formation in various primitive chondrites based on the 26Al–26Mg system. If the Pb–Pb dates of chondrules are indeed accurate to their stated uncer- tainty, ultra-high-precision 26Al–26Mg systematics of individual chondrules from primitive chon- drites using the internal isochron approach should return an age distribution that is comparable to that inferred from the Pb–Pb dates within uncertainty for the same meteorite. If correct, this would undermine the concept of a genetic complementarity between chondrule and matrix and, thus, solve an important conundrum that currently plagues the cosmochemical community.

References

Alexander, C. M. O’D. (2005). Re-examining the role of chondrules in producing the elemental fractionations in chondrites. Meteorit. Planet. Sci., 40, 943–965. Alexander, C. M. O’D., Grossman, J. N., Ebel, D. S., and Ciesla, F. J. (2008). The formation conditions of chondrules and chondrites. Science, 320, 1617–1619. Amelin, Y. (2008). The U–Pb systematics of angrite Sahara 99555. Geochim. Cosmochim. Acta, 72, 4874–4885. Amelin, Y., Connelly, J. N., Zartman, R. E., et al. (2009). Modern U-Pb chronometry of meteorites: Advancing to higher time resolution reveals new problems. Geochim. Cosmo- chim. Acta, 73, 5212–5223. Amelin, Y., Kaltenbach, A., Iizuka, T., et al. (2010). U–Pb chronology of the Solar System’s oldest solids with variable 238U/235U. Earth Planet. Sci. Lett., 300, 343–350. Amelin, Y., and Krot, A. N. (2007). Pb isotopic age of the Allende chondrules. Meteorit. Planet. Sci., 42, 1321–1335. Amelin, Y., Krot, A. N., Hutcheon, I. D., and Ulyanov, A. A. (2002). Lead isotopic ages of chondrules and calcium-aluminum-rich inclusions. Science, 297, 1678–1683. 320 James N. Connelly and Martin Bizzarro

Becker, M., Hezel, D. C., Schulz, T., Elfers, B. -M., and Münker, C. (2015). Formation timescales of CV chondrites from component specificHf–W systematics. Earth Planet. Sci. Lett., 432, 472–482. Bitsch, B., Lambrechts, M., and Johansen, A. (2015). The growth of planets by pebble accretion in evolving protoplanetary discs. Astron. Astrophys., 582, A112. Bizzarro, M., Baker, J. A., and Haack, H. (2004). Mg isotope evidence for contemporaneous formation of chondrules and refractory inclusions. Nature, 431, 275–278. Bizzarro, M., Paton, C., Larsen, K., et al. (2011). High-precision Mg-isotope measurements of terrestrial and extraterrestrial material by HR-MC- ICPMS – Implications for the relative and absolute Mg-isotope composition of the bulk silicate Earth. J. Anal. At. Spectrom., 26, 565–577. Bizzarro, M., Connelly, J.N., and Krot, A.N. (2017). Chondrules – Ubiquitous chondritic solids tracking the evolution of the solar protoplanetary disk. In M. Pessah and O. Gressel (Eds.), Formation, Evolution, and Dynamics of Young Solar Systems, 161–195. Cham, Switzer- land: Springer International. Bizzarro, M., Olsen, M., Itoh, S., et al. (2014). Evidence for a reduced initial abundance of 26Al in chondrule forming regions and implications for the accretion timescales of protoplanets. Meteorit. Planet. Sci, 49 (Suppl.), Abstract #A43. Blundy, J., and Wood, B. (2003). Mineral-melt partitioning of uranium, thorium and their daughters. Rev. Mineral. Geochem., 52, 59–123. Boley, A. C., and Durisen, R. H. (2008). Gravitational instabilities, chondrule formation, and the FU Orionis phenomenon. Astrophys. J., 685, 1193–1209. Boley, A. C., Morris, M. A., and Desch, S. J. (2013). High-temperature processing of solids through solar nebular bow shocks: 3D radiation hydrodynamics simulations with particles. Astrophys. J., 776, 101. Bollard, J., Connelly, J. N., and Bizzarro, M. (2015a). Pb–Pb dating of individual chondrules from the CBa chondrite Gujba: Assessment of the impact plume formation model. Meteorit. Planet. Sci., 50, 1197–1216. Bollard, J., Kawaski, N., Sakamoto, N., et al. (2015b). Early disk dynamics inferred from isotope systematics of individual chondrules. Meteorit. Planet. Sci., 50 (Suppl.), Abstract #5211. Bollard. J., Connelly, J. N., Whitehouse, M. J., et al. (2017). Early formation of planetary building blocks inferred from Pb isotopic ages of chondrules. Sci. Adv., 3 (8), e1700407. Boss, A. P., and Durisen, R. H. (2005). Chondrule-forming Shock Fronts in the Solar Nebula: A Possible Unified Scenario for Planet and Chondrite Formation. Astrophys. J. Lett., 621, L137–L140. Brennecka, G. A., Weyer, S., Wadhwa, M., et al. (2010). 238U/235U variations in meteorites: extant 247Cm and implications for Pb–Pb dating. Science, 327, 449–451. Brennecka, G. A., Budde, G., and Kleine, T. (2015). Uranium isotopic composition and absolute ages of Allende chondrules. Meteorit. Planet. Sci., 50, 1995–2002. Brennecka, G. A., and Wadhwa, M. (2012). Uranium isotope compositions of the basaltic angrite meteorites and the chronological implications for the early Solar System. Proc. Natl. Acad. Sci., 109, 9299–9303. Budde, G., Burkhardt, C., Brennecka, G. A., et al. (2016b). Molybdenum isotopic evidence for the origin of chondrules and a distinct genetic heritage of carbonaceous and non- carbonaceous meteorites. Earth Planet. Sci. Lett., 454, 293–303. Budde, G., Kleine, T., Kruijer, T. S., Burkhardt, C., and Metzler, K. (2016a). Tungsten isotopic constraints on the age and origin of chondrules. Proc. Natl. Acad. Sci., 113, 2886–2891. Carporzen, L., Weiss, B. P., Elkins-Tanton, L. T., et al. (2011). Magnetic evidence for a partially differentiated carbonaceous chondrite parent body. Proc. Natl. Acad. Sci., 108, 6386–6389. Carrasco-Gonzalez, C., Henning, T., Chandler, C. J., et al. (2016). The VLA view of the HL Tau Disk - Disk Mass, Grain Evolution, and Early Planet Formation. Astrophys. J., 821, L16. The Absolute Pb–Pb Isotope Ages of Chondrules 321

Ciesla, F. J., and Hood, L. L. (2002). The nebular shock wave model for chondrule formation: Shock processing in a particle -gas suspension. Icarus, 158, 281–293. Connelly, J. N., Bizzarro, M., Thrane, K., and Baker, J. A. (2008). The Pb–Pb age of Angrite SAH99555 revisited. Geochim. Cosmochim. Acta, 72, 4813–4824. Connelly, J. N., and Bizzarro, M. (2009). Pb–Pb dating of chondrules from CV chondrites by progressive dissolution. Chemical Geology, 259, 143–151. Connelly, J. N., Bizzarro, M., Krot, A. N., et al. (2012). The absolute chronology and thermal processing of solids in the solar protoplanetary disk. Science, 338, 651–655. Connelly, J. N., Bollard, J., and Bizzarro, M. (2017). Pb–Pb chronometry and the early Solar System. Geochim. Cosmochim. Acta, 201, 345–363. Dauphas, N., and Pourmand, A. (2011). Hf-W-Th evidence for rapid growth of Mars and its status as a planetary embryo. Nature 473, 489–492. Desch, S. J., Morris, M. A., Connolly, H. C., and Boss, A. P. (2012). The importance of experi- ments: Constraints on chondrule formation models. Meteorit. Planet. Sci., 47, 1139–1156. Ebel, D. S., Brunner, C., Konrad, K., et al. (2016). Abundance, major element composition and size of components and matrix in CV, CO and Acfer 094 chondrites. Geochim. Cosmo- chim. Acta, 172, 322–356. Evans, N. J. II, Dunham, M. M., Jørgensen, J. K., et al. (2009). The Spitzer c2d Legacy results: star formation rates and efficiencies; evolution and lifetimes. Astrophys. J. Suppl. Ser., 181, 321–350. Goldberg, A. Z., Owen, J. E., and Jacquet, E. (2015). Chondrule transport in protoplanetary discs. Mon. Not. R. Astron. Soc., 452, 4054–4069. Goldmann, A., Brennecka, G., Noordmann, J., Weyer, S., and Wadhwa, M. (2015). The uranium isotopic composition of the Earth and the Solar System. Geochim. Cosmochim. Acta, 148, 145–158. Hezel, D. C., and Palme, H. (2010). The chemical relationship between chondrules and matrix and the chondrule matrix complementarity. Earth Planet. Sci. Lett., 294, 85–93. Holst, J. C., Olsen, M. B., Paton, C., et al. (2013). 182Hf-182W age dating of a 26Al-poor inclusion and implications for the origin of short-lived radioisotopes in the early Solar System. Proc. Natl. Acad. Sci., 110, 8819–8823. Hood, L. L., Ciesla, F. J., Artemieva, N. A., Marzari, F., and Weidenschilling, S. J. (2009). Nebular shock waves generated by planetesimals passing through Jovian resonances: Possible sites for chondrule formation. Meteorit. Planet. Sci., 44, 327–342. Hu, R. (2010). Transport of the First Rocks of the Solar System by X-winds. Astrophys. J., 725, 1421–1428. Huss, G. R., MacPherson, G. J., Wasserburg, G. J., Russell, S. S., and Srinivasan, G. (2001). Aluminum-26 in calcium-aluminum-rich inclusions and chondrules from unequilibrated ordinary chondrites. Meteorit. Planet. Sci., 36, 975–997. Itoh, S., and Yurimoto, H. (2003). Contemporaneous formation of chondrules and refractory inclusions in the early Solar System. Nature, 423, 728–731. Johansen, A., Low, M. -M. M., Lacerda, P., and Bizzarro, M. (2015). Growth of asteroids, planetary embryos, and Kuiper belt objects by chondrule accretion. Sci. Adv., 1, e1500109–e1500109. Johansen, A., Oishi, J. S., Low, M. -M. M., et al. (2007). Rapid planetesimal formation in turbulent circumstellar disks. Nature, 448, 1022–1025. Kaltenbach, A. (2012). Uranium isotopic analyses of terrestrial and extraterrestrial samples. Ph. D. Thesis, University of Otago, Dunedin, New Zealand, 174 p. Krot, A. N., and Wasson, J. T. (1995). Igneous rims on low-FeO and high-FeO chondrules in ordinary chondrites. Geochim. Cosmochim. Acta, 59, 4951–4966. Krot, A. N., Libourel, G., Goodrich, C. A., and Petaev, M. I. (2004). Silica-rich igneous rims around magnesian chondrules in CR carbonaceous chondrites: Evidence for condensation origin from fractionated nebular gas. Meteorit. Planet. Sci. Arch., 39, 1931–1955. 322 James N. Connelly and Martin Bizzarro

Krot, A. N., Amelin, Y., Cassen, P., and Meibom, A. (2005a). Young chondrules in CB chondrites from a giant impact in the early Solar System. Nature, 436, 989–992. Krot, A. N., Yurimoto, H., Hutcheon, I. D., and MacPherson, G. J. (2005b). Chronology of the early Solar System from chondrule-bearing calcium-aluminium-rich inclusions. Nature, 434, 998–1001. Krot, A. N., Amelin, Y., Bland, P., et al. (2009). Origin and chronology of chondritic compon- ents: A review. Geochim. Cosmochim. Acta, 73, 4963–4997. Krot, A. N., Nagashima, K., van Kooten, E. M. M., and Bizzarro, M. (2017). Calcium- aluminum-rich inclusions recycled during formation of porphyritic chondrules from CH carbonaceous chondrites. Geochim. Cosmochim. Acta, 201, 185–223. Lambrechts, M., and Johansen, A. (2012). Rapid growth of gas-giant cores by pebble accretion. A. Astron. Astrophys., 544, A32. Larsen, K. K., Trinquier, A., Paton, C., et al. (2011). Evidence for magnesium isotope heterogeneity in the solar protoplanetary disk. Astrophys. J. Lett., 735, L37. Larsen, K. K., Schiller, M., and Bizzarro, M. (2016). Accretion timescales and style of asteroidal differentiation in an 26Al-poor protoplanetary disk. Geochim. Cosmochim. Acta, 176, 295–315. Lodders, K. (2003). Solar System Abundances and condensation temperatures of the elements. Astrophys. J., 591, 1220–1247. Morris, M. A., Boley, A. C., Desch, S. J., and Athanassiadou, T. (2012). Chondrule formation in bow shocks around eccentric planetary embryos. Astrophys. J., 752, 27. Morris, M. A., Garvie, L. A. J., and Knauth, L. P. (2015). New insight into the Solar System’s transition disk phase provided by the unusual meteorite Isheyevo. Astrophys. J., 801, L22. Mostefaoui, S., Kita, N. T., Togashi, S., et al. (2002). The relative formation ages of ferro- magnesian chondrules inferred from their initial aluminum-26/aluminum-27 ratios. Meteorit. Planet. Sci., 37, 421–438. Moynier, F., Beck, P., Jourdan, F., et al. (2009). Isotopic fractionation of zinc in tektites. Earth Planet. Sci. Lett., 277, 482–489. Nagashima, K., Krot, A. N., and Huss, G. R. (2014). 26Al in chondrules from CR2 chondrites. Geochem. J. 48, 561–570. Norris, T. L., Gancarz, A. J., Rokop, D. J., and Thomas, K. W. (1983). Half-life of 26Al. J. Geophys. Res.: Solid Earth, 88, B331–B333. Olsen, M. B., Schiller, M. Krot, A. N., and Bizzarro M. (2013). Magnesium isotope evidence for single stage formation of CB chondrules by colliding planetesimals. Astrophys. J. Lett., 776, L1. Olsen, M. B., Wielandt, D., Schiller, M., Van Kooten, E. M. M. E. and Bizzarro, M. (2016). Magnesium and 54Cr isotope compositions of carbonaceous chondrite chondrules – Insights into early disk processes. Geochim. Cosmochim. Acta, 191, 118–138. Palme, H., Hezel, D. C., and Ebel, D. S. (2015). The origin of chondrules: Constraints from matrix composition and matrix-chondrule complementarity. Earth Planet. Sci. Lett., 411, 11–19. Pérez, L. M., Carpenter, J. M., Andrews, S. M., et al. (2016). Spiral density waves in a young protoplanetary disk. Science, 353, 1519–1521. Rudraswami, N. G., and Goswami, J. N. (2007). 26Al in chondrules from unequilibrated L chondrites: Onset and duration of chondrule formation in the early solar system. Earth Planet. Sci. Lett., 257, 231–244. Schiller, M., Connelly, J. N., Glad, A. C., Mikouchi, T., and Bizzarro, M. (2015a). Early accretion of protoplanets inferred from a reduced inner solar system 26Al inventory. Earth Planet. Sci. Lett., 420, 45–54. Schiller, M., Paton, C., and Bizzarro, M. (2015b). Evidence for nucleosynthetic enrichment of the protosolar molecular cloud core by multiple supernova events. Geochim. Cosmochim. Acta, 149, 88–102. The Absolute Pb–Pb Isotope Ages of Chondrules 323

Schrader, D. L., Nagashima, K., Krot, A. N., et al. (2017). Distribution of 26Al in the CR chondrite chondrule-forming region of the protoplanetary disk. Geochim. Cosmochim. Acta, 201, 275–302. Scott, E. R. D., and Krot, A. N. (2003). Chondrites and their Components. In A. M. Davis (Ed.), Meteorites, Comets and Planets. In H. D. Holland and K. K. Turekian (Eds.), Treatise on Geochemistry (First Edition), 1, 144–200. Oxford, UK: Elsevier. Scott, E. R. D. (2007). Chondrites and the Protoplanetary Disk. Annu. Rev. Earth Planet. Sci., 35, 577–620. Scott, E. R. D., and Krot, A. N. (2005). Thermal processing of silicate dust in the solar nebula: clues from primitive chondrite matrices. Astrophys. J., 623, 571. Shu, F. H., Shang, H., and Lee, T. (1996). Toward an astrophysical theory of chondrites. Science, 271, 1545–1552. Steiger, R. H., and Jäger, E. (1977). Subcommission on geochronology: Convention on use of decay constants in geo- and cosmochronology. Earth Planet. Sci. Lett., 36, 359–362. Tatsumoto, M., Knight, R. J., and Allegre, C. J. (1973). Time differences in the formation of meteorites as determined from the ratio of Lead-207 to Lead-206. Science, 180, 1279–1283. Testi, L., Birnstiel, T., Ricci, L., et al. (2014). In H. Beuther, R. S. Klessen, C. P. Dullemond, and T. Henning (Eds.), Protostars and Planets VI, 339–361. Tucson, AZ: University of Arizona Press. Thrane, K., Kazuhide, N., Krot, A. N., and Bizzarro M. (2008). Discovery of a new FUN CAI from a CV carbonaceous chondrite: Evidence for multistage thermal processing in the protoplanetary disk. Astrophys. J. Lett., 680, L141–L144. Tobin, J. J., Looney, L. W., Wilner, D. J., et al. (2015). A sub-arcsecond survey toward Class 0 protostars in Perseus: Searching for signatures of protostellar disks. Astrophys. J., 805, 125. Tomida, K., Okuzumi, S., and Machida, M. N. (2015). Radiation magnetohydrodynamic simulations of protostellar collapse: Nonideal magnetohydrodynamic effects and early formation of circumstellar disks. Astrophys. J., 801, 117. Trinquier, A., Elliott, T., Ulfbeck, D., et al. (2009). Origin of nucleosynthetic isotope hetero- geneity in the solar protoplanetary disk. Science, 324, 374–376. Van Kooten, E. M. M. E., Wielandt, D., Schiller, M., et al. (2016). Isotopic evidence for primordial molecular cloud material in metal-rich carbonaceous chondrites. Proc. Natl. Acad. Sci., 113, 2011–2016. Weidenschilling, S. J. (1977). Aerodynamics of solid bodies in the solar nebula. Mon. Not. Roy. Astron. Soc., 180, 57–70. Weiss, B. P., and Elkins-Tanton, L. T. (2013). Differentiated planetesimals and the parent bodies of chondrites. Annu. Rev. Earth Planet. Sci., 41, 529–560. Williams, J. P., and Cieza, L. A. (2011). Protoplanetary disks and their evolution. Annu. Rev. Astron. Astrophys., 49, 67–117. 12 Records of Magnetic Fields in the Chondrule Formation Environment

roger r. fu, benjamin p. weiss, devin l. schrader, and brandon c. johnson

Abstract

Chondrules contain ferromagnetic minerals that may retain a record of the magnetic field environments in which they cooled. Paleomagnetic experiments on separated chondrules can potentially reveal the presence of remanent magnetization from the time of chondrule formation. The existence of such a magnetization places quantitative bounds on the frequency of inter- chondrule collisions, while the intensity of magnetization may be used to infer the strength of nebular magnetic fields and thereby constrain the mechanism of chondrule formation. Recent advances in laboratory instrumentation and techniques have permitted the isolation of nebular remanent magnetization in chondrules, providing the potential basis to probe the formation environments of chondrules from a range of chondrite classes.

12.1 Introduction

Chondrules are millimeter-scale, nearly spherical igneous inclusions found in chondritic meteorites (chondrites). Their abundance in chondrites and early crystallization within several million years (My) of the formation of calcium aluminum-rich inclusions (CAIs) suggests that they may have played an important role in the formation of solid bodies in the solar system [e.g., Villeneuve et al. (2009); Kita and Ushikubo (2012); Schrader et al. (2017)]. However, funda- mental questions about the origin of chondrules remain unanswered, including the mechanism of formation and the spatial distribution of chondrule formation regions in the solar nebula. Paleomagnetism is a technique that takes advantage of naturally occurring ferromagnetic minerals in rocks to infer the presence and intensity of ancient magnetic fields. Although only a handful of paleomagnetic studies to date have measured separated, individual chondrules, infor- mation about magnetic fields during chondrule formation may yield valuable information about their formation environment. Interactions with the partially ionized nebular gas caused the embed- ded magnetic fields to vary in time and position in the solar nebula (Balbus, 2003; Wardle, 2007; Bai, 2016), permitting knowledge of the magnetic field to constrain potentially the timing and location of chondrule formation. Furthermore, proposed chondrule formation models predict a range of magnetic field strengths during the formation event, implying that paleomagnetic infer- ences of paleofield intensity may help to identify the mechanism or mechanisms of chondrule formation. Finally, although beyond the scope of this chapter, paleomagnetic experiments on

324 Records of Magnetic Fields in the Chondrule Formation Environment 325 chondrules may reveal the strength of background nebular magnetic fields, which likely played a significant role in disk evolution and planet formation (Fu et al., 2014b). In this chapter, we begin by outlining the potential contributions of paleomagnetism to questions about chondrule formation. We then turn to a review of existing paleomagnetic studies on individual chondrules, highlighting the significant challenges posed by sample selection and by the measurement of small, weakly magnetized samples.

12.2 The Potential of Chondrule Paleomagnetism 12.2.1 Implications of a Preserved Magnetization

Paleomagnetic constraints on the chondrule formation environment come from two observables related to the chondrules’ natural remanent magnetization (NRM), the semipermanent align- ment of electron spins within ferromagnetic materials imparted by past magnetic fields. First, the existence of any identifiable, linear component of NRM places requirements on the stability of the magnetic field in the chondrule’s frame of reference (Figure 12.1). Second, the intensity of chondrule NRM may be used to infer the strength of background magnetic fields during the formation process. In both cases, the recorders of ancient magnetic fields within chondrules are inclusions of ferromagnetic minerals such as magnetite, pyrrhotite, and FeNi alloys including kamacite, martensite, taenite, and tetrataenite (Uehara et al., 2011). Among these, FeNi alloys and, in some chondrule groups, pyrrhotite are most likely to have a nebular origin and record the magnetic field during chondrule formation (Krot et al., 1998; Schrader et al., 2016a). Similar to igneous rocks on Earth and other solar system bodies, the form of NRM carried by

Figure 12.1 Schematic summary of potential constraints on chondrule formation derived from paleomagnetic experiments. Abbreviation Tc refers to Curie temperature, which is 780 C for the native Fe mineral kamacite. 326 Roger R. Fu, et al. ferromagnetic grains in chondrules may be a thermoremanent magnetization (TRM) if their cooling at temperatures below the Curie point (700–780 C for typical FeNi compositions) took place in a directionally stable background magnetic field (Swartzendruber et al., 1991; Kimura et al., 2008). We refer to such a magnetization as a primary natural remanent magnetization (primary NRM), meaning that its acquisition occurred during the formation of the chondrule and distinguishing it from secondary NRM acquired during thermal and/or aqueous alteration events in the chondrule’s subsequent history. Laboratory demagnetization of separated chondrules via heating or room temperature alter- nating field (AF) protocols are designed to remove secondary overprints and isolate the primary component of magnetization, if any. For a primary NRM to be properly recognized as a component of magnetization, the primary NRM must display a consistent direction during demagnetization after removal of any secondary magnetization components (Tauxe, 2010). During thermal demagnetization, the temperatures over which the observed sample magnetiza- tion trends in a consistent direction is known as the blocking temperature range of that magnetization component. For a chondrule to carry such an identifiable magnetization compon- ent, the background magnetic field must have remained constant with respect to the chondrule or, for a rotating chondrule, the chondrule’s angular momentum axis during cooling over the blocking temperature range. In the latter case, a spinning or precessing chondrule is expected to record a component of magnetization parallel to its angular momentum axis (Fu et al., 2014b; Takac and Kletetschka, 2015). The identification of a primary NRM in a population of chondrules therefore implies that these specimens must have avoided significant reorientation of their angular momentum vectors as they cooled through the observed range of blocking temperatures for the primary NRM, which can be quantified using stepwise thermal demagnetization. The initial rotation rate of chondrules during cooling above the solidus was likely in excess of 100 rad s1 based on their three-dimensional shapes (Tsuchiyama et al., 2003; Miura et al., 2008). Under minimum mass solar nebula conditions [MMSN; (Hayashi, 1981)], the frictional decay of angular velocity for a 1-mm diameter chondrule takes place over the course of several hours at 1 AU and over 103 hours at 10 AU (Fu and Weiss, 2012). Other solar nebula models predict gas densities and nebular temperatures that vary by a factor of ~2 and ~4 from the MMSN, respectively, while gas density may have declined by more than an order of magnitude within 3–4 My after CAIs (Bitsch et al., 2015; Wang et al., 2017). As such, the spin-down timescales calculated above are likely to represent lower bounds during the time period of peak chondrule formation between 2 and 3 My after CAIs. Given chondrule cooling rates of 10–1,000 Ch1 (Hewins et al., 2005; Tachibana et al., 2006; Schrader et al., 2016b), these spin-down rates imply that chondrules are expected to be spinning during most or all of the cooling process assuming background nebular gas densities, especially in the outer solar system. Given these rotation rates, random collisions with gas molecules during the spin-down process impart negligible changes to the rotation state, equivalent to ~0.01 rad s1 for a 1 mm diameter chondrule at 1 AU (Fu et al., 2014b). Therefore, a spinning chondrule cooling in the gas environment of the solar nebula is expected to retain a nearly constant angular momentum vector orientation, permitting the acquisition of a primary NRM if a significant magnetic field is present. Enhancement in nebular gas density of several orders of magnitude would not affect this expectation in the outer solar system at 2–3 My after CAIs. On the other hand, early formation Records of Magnetic Fields in the Chondrule Formation Environment 327 of chondrules in the inner solar system (<2 AU) combined with local enhancement of gas densities during chondrule formation may result in chondrules that de-spin completely before sufficient cooling to record ambient magnetic fields. In such cases, the orientation of the chondrule in space may undergo Brownian motion due to collisions with gas molecules, resulting in non-unidirectional magnetization. On the other hand, collisions with other chondrules would dramatically change the angular momentum of a cooling chondrule. If such collisions occurred frequently during the cooling process below the Curie temperature, then chondrules would not be able to record an identifi- able primary NRM. The observation of a primary NRM in a chondrule would therefore imply an upper bound on the probability of interchondrule collisions during the interval of cooling corresponding to TRM acquisition. For a collection of equal-sized spherical particles with radius rc moving with mutual velocity vc, the expected number of collisions (N) experienced 2 by a single particle in time Δt is given by N ¼ 4πr vcnΔt, where n is the number density of particles (Ciesla et al., 2004). The parameter Δt may be estimated by combining the temperature interval of remanence acquisition observed from thermal demagnetization and subsilicate solidus cooling rates derived from petrographic observations. To constrain cooling rates at temperatures below the 780 C Curie point of kamacite, we use the texture and compositions of sulfides in some minimally altered chondrite groups. The morphology, texture, and equilibration temperature of these phases indicate formation at high temperatures during chondrule cooling and rule out a low temperature origin on the parent asteroids (Schrader et al., 2015, 2016a, 2016b). The existence of a primary NRM in a population of chondrules, which constrain N, therefore constrains the number density and velocity of chondrules (Figure 12.1), analogous to observations of compound chondrule frequency. Finally, even in the case of a consistent rotation axis, chondrules would not record a primary NRM if the background magnetic field direction were unstable during the cooling time window. Therefore, the detection of a primary NRM implies that the local magnetic field remained constant in direction for at least the several hours during which the chondrule cooled through the block temperature range.

12.2.2 Implications of the Paleointensity

If and when a primary NRM is identified in a chondrule, further paleomagnetic experiments may be able to infer the paleointensity, or the strength of the ancient magnetic field during TRM acquisition (Figure 12.1). The most prevalent paleointensity protocols impart an artificial remanent magnetization in a known laboratory field using a variety of both high temperature and room temperature techniques (Tauxe, 2010). By comparing the strengths of the natural and artificial remanence magnetizations, the approximate strength of the original natural magnetic field may be deduced. Similar to other extraterrestrial samples, ferromagnetic phases in chon- drules often experience severe chemical alteration during laboratory heating. As such, paleoin- tensity estimates on single chondrules are typically based on room temperature methods that are uncertain by a 2-standard deviation factor of 3–5 [see discussion in Weiss and Tikoo (2014)]. In the best-case scenario where laboratory cooling experiments are available as calibrations for room temperature paleointensities (Lappe et al., 2013), this uncertainty may be reduced to a 328 Roger R. Fu, et al. factor of ~1.3. As a second source of uncertainty, chondrule rotation during cooling implies that the recorded paleointensity would be attenuated by a factor cos θ, where θ is the angle between the rotation axis and the magnetic field. Although the factor cos θ is unknown for any one chondrule, the mean value across a population of randomly oriented chondrules approaches 0.5 in the limit of a large sample size (Fu et al., 2014b). The experimental and sample selection challenges of chondrule paleomagnetism limit the practical sample size, leading to uncertainty on the multiplier used to correct for chondrule rotation. Together, these uncertainties imply that chondrule-derived paleointensities are accurate to an order of magnitude in the absence of calibrations against a laboratory TRM and to a factor of 1.5 in the best-case scenario where such experiments are available (Fu et al., 2014b). With these limitations in mind, paleomagnetic constraints on field intensity during chondrule formation may be able to distinguish between different formation mechanisms that predict significantly different field intensities. Fortunately, existing models indeed place chondrule formation in broadly disparate magnetic field environments. The x-wind model of chondrule formation places the site of chondrule formation between approximately 0.1 and 0.5 AU from the sun, resulting in exposure to strong magnetic fields. Chondrules traveling in the magnetocentrifugal winds invoked by the x-wind model experience relatively constant magnetic fields of 400–800 µT during cooling between the 780 C Curie point of kamacite and 300 C, assuming chondrule formation during the embedded phase of the protosun (Shu et al., 1996, 1997). These field values drop to 80–160 µT for chondrule formation during the revealed phase. Models that place chondrule formation in the protoplanetary disk itself make chondrule paleointensity predictions that depend on the local nebular magnetic field strength. The variation of nebular field strength with orbital radius is complex and model-dependent. Within 5–10 AU of the sun, higher concentrations of dust grains and higher optical depths at the disk midplane may lead to the formation of a “dead zone,” in which magnetic field strengths are suppressed due to low gas ionization levels (Gammie, 1996; Turner and Sano, 2008; Bai, 2011). However, recent simulations that more fully incorporate nonideal magnetohydrodynamical (MHD) phe- nomena including the Hall effect suggest that significant magnetic fields may persist in the dead zone [Figure 12.2; (Bai, 2014; Lesur et al., 2014)]. At the same time, the regions of anomalous low or high gas density, which may affect local magnetic field strength, have been observed in other protoplanetary disks and may have formed with or without the presence of planets (van der Marel et al., 2013; Zhu, Stone, and Rafikov, 2013; Flock et al., 2015). However, the advection of magnetic fields into these regions may result in minimal change in the magnetic field strength even in zones of lower gas density (Wang et al., 2017). Finally, the inferred dissipation of the nebular gas at least in specific regions of the disk by ~3–4 My after CAIs (Bitsch et al., 2015; Wang et al., 2017) implies that the strength of nebular magnetic fields may have evolved rapidly on 1 My timescales. As such, accurate geochronology of chondrules, preferably on the same specimens being analyzed for paleomagnetism, is critical for the interpretation of paleomagnetic results. Despite these uncertainties, an order-of-magnitude theoretical estimate of background nebu- lar field strengths may be derived under the assumption that magnetic fields led to the bulk of angular momentum transport at all radii in the protosolar disk. In this scenario, the magnetic field strength is a function of the assumed mechanism of transport and accretion rate and Records of Magnetic Fields in the Chondrule Formation Environment 329

Figure 12.2 Model predictions for background nebular magnetic field intensities. Points at 5, 15, and 30 AU denote the approximate maximum midplane field intensities observed in numerical MHD simulations that fully include nonideal MHD effects (Bai, 2015). Because the modeled field strengths fluctuate in time, these peak values may be regarded as upper bounds on the field strength. Solid line denotes the magnetic –8 -1 field intensities required to produce a mass accretion rate of 10 M☉ y assuming accretion by coupling of the radial and azimuthal components of the magnetic field [(Bai and Goodman, 2009); Equation 16]. Dashed line shows the magnetic field intensities for the same accretion rate but assuming accretion by radial-vertical magnetic field coupling [(Bai and Goodman, 2009); Equation 7]. Dotted gray line denotes the minimum, nonfluctuating net vertical magnetic field strength needed to produce the above accretion rate according to numerical MHD simulations (Bai, 2015). This value may be regarded as a lower bound on the midplane field intensity. Physical field strengths from numerical simulations are computed by combining Equation 14 of Bai (2015) and the definition of the plasma β parameter for a MMSN scenario. decreases with radius following a power-law [Figure 12.2; (Wardle, 2007; Bai and Goodman, – 2009)]. Assuming a typical sun-like protostellar accretion rate of 10 8 solar masses per year -1 (M☉ y ), the background nebular field strength at 1 AU is between 5 and 100 µT, which declines to 0.02–1 µT at 30 AU. These field strengths are consistent with estimates assuming a turbulent, one-dimensional disk (Guilet and Ogilvie, 2014) and with numerical MHD simulations (Bai, 2014) assuming net vertical fields of ~10 µT at 1 AU. Although the value of the net vertical field is not known directly, a ~10 µT value is in agreement with estimates from models of magnetic field diffusion in the protoplanetary disk (Desch and Mouschovias, 2001; Okuzumi et al., 2014). The magnetic field intensity potentially recorded in chondrules formed in the protoplanetary disk is a function of these background magnetic field intensities as well as any modification due to the chondrule formation mechanism itself. The degree of field amplification by local hydrodynamic processes depends on the magnetic diffusivity of the gas and the size of the region of interest. For the impact formation model of chondrules, the spatial scale of the chondrule-forming jets is 105 km. The presence of dust particles leads to very low ionization fractions (<10–18) for nebular gas near the disk midplane, corresponding to high magnetic diffusivities on the order of 1018–1020 m2 s1 (Wardle, 2007). The 780 C temperatures in which magnetic fields may be recorded by chondrules are unlikely to lead to significant thermal ionization (McNally et al., 2013), while the enhanced solids concentration may lead to even higher magnetic diffusivities. Under these conditions, magnetic fields are expected to diffuse 330 Roger R. Fu, et al. across the thickness of chondrule-forming outflows in less than 1 s (Asphaug et al., 2011; Johnson et al., 2015). Chondrules formed in planetesimal collisions would therefore experience no enhancement of the background nebular magnetic field, predicting recorded paleointensities of 0.02–100 µT, depending on the location in the protoplanetary disk (Figure 12.2). In the case of the nebular shock model of chondrule formation, compression of the nebular gas over potentially large spatial scales may prevent the rapid diffusion of magnetic fields, resulting in enhancement of the nebular magnetic field by a factor of approximately 10 (Desch and Connolly, 2002). As such, predicted recorded paleointensities for chondrules formed via this mechanism are in the 100–1,000 µT range in the innermost disk around 1 AU, overlapping with fields expected for the x-wind model. Meanwhile, shock-enhanced field strengths in and beyond the asteroid belt are between 1 and 100 µT. Two other hypothesized mechanisms of chondrule formation, the magnetic reconnection flare and short circuit instability models (Levy and Araki, 1989; McNally et al., 2013), predict potentially very high local magnetic field strengths of 500 and 104 µT, respectively, during the peak heating phase of chondrule formation. However, these strong magnetic fields may decay significantly by the time temperature drops below the 780 C Curie temperature of kamacite. More detailed modeling of the simultaneous decay of magnetic field strength and temperature is required to produce quantitative paleointensity predictions for chondrules formed in these mechanisms.

12.3 Previous Paleomagnetic Experiments on Chondrules

Several basic properties of chondrules pose formidable challenges to their paleomagnetic characterization. The small size and nearly spherical morphology of chondrules have created difficulties for the accurate measurement of their magnetic moment and for the preparation of mutually oriented samples. As importantly, the complex, protracted, and often poorly known geological histories of chondrules in the solar nebula, on their asteroidal parent bodies, and on Earth result in a wide range of secondary processes that may have potentially resulted in partial or complete remagnetization. As a result of these challenges, few robust paleomagnetic meas- urements of individual chondrules have been obtained, and even fewer among these can provide meaningful constraints on the chondrule formation process itself. A critical experiment used to determine the origin of chondrule NRM is the paleomagnetic conglomerate test. In this analysis, consistent magnetization directions found in multiple, mutually oriented chondrules indicates that the magnetization postdates the assembly of the meteorite and therefore is not a primary record of magnetic fields in the solar nebula during chondrule formation. The paleomagnetic conglomerate test can only be performed on chon- drules whose mutual orientations are maintained through the extraction process. The earliest paleomagnetic experiments on separated chondrules focused on unoriented specimens from the Allende CV3 carbonaceous chondrite (Lanoix et al., 1978). Paleointensity experiments using a heating-based protocol suggested very high paleointensites of 1,600 µT over a blocking temperature range of 350–500 C. However, because the chondrules in these experiments lacked mutual orientation, the authors could not rule out a late remagnetization event during processing on the parent body or due to exposure to strong artificial magnetic fields on Earth. Indeed, the very high inferred paleointensities are suggestive of contamination due to Records of Magnetic Fields in the Chondrule Formation Environment 331 hand magnet magnetic fields (Wasilewski, 1981; Weiss et al., 2010). As further evidence that the chondrule paleointensities obtained by Lanoix et al. (1978) were the result of anomalous hand magnet contamination, a follow-up study of mutually oriented Allende chondrules by the same research group did not find such high paleointensities in magnetization over the same blocking temperature range (Sugiura et al., 1979). Instead, the later study obtained a paleoin- tensity of ~100 µT on a different component of magnetization blocked below 300 C. Because the direction of this lower temperature magnetization was consistent among different chon- drules and with matrix samples, it must have been acquired after the accretion of the parent body and cannot constrain preaccretional fields during chondrule formation (Carporzen et al., 2011). As such, early published nebular field intensities of 100–>1,000 µT based on the Lanoix et al. (1978) work [e.g., Levy and Sonett (1978)] should be disregarded. Furthermore the ~100 µT paleointensity of CV parent body magnetic fields obtained by Sugiura et al. (1979) has been mistakenly cited by later studies as corresponding to the nebular magnetic field [e.g., Levy and Araki (1989); Stepinski (1992); Shu et al. (1996); Nübold and Glassmeier (2000); Johansen (2009)]. Finally, other early paleointensities derived from unoriented chondrules, such as the values of up to ~1,000 µT obtained from the Allende, Allegan (H5), Bjurböle (L/LL4), and Chainpur (LL3.4) chondrites, were also likely influenced by contamination from exposure to hand magnets (Wasilewski and O’Bryan, 1994). To identify similar contamination by hand magnets, future studies of meteorite paleomagnetism should evaluate the ratio of NRM to isothermal remanent magnetization (IRM), which is anomalously high for samples exposed to strong fields (Weiss et al., 2010). Furthermore, the measurement of mutually oriented samples that include fusion crust material would aid in the confident detection of hand magnet-induced overprints, which affect the fusion crust and are expected to be heterogeneous in strength and direction (Weiss and Elkins-Tanton, 2013). Later studies of mutually oriented chondrules from Allende have revealed ostensible evi- dence that a primary NRM recording chondrule formation magnetic fields may be preserved. Sugiura et al. (1979) and Sugiura and Strangway (1985) showed that a high temperature component of magnetization blocked above ~300 C in Allende chondrules passes the paleo- magnetic conglomerate test (i.e., random magnetization directions suggesting preaccretional origin). Motivated by these findings, Fu et al. (2014a) performed additional experiments demonstrating that, while chondrules separated from Allende indeed exhibit distinct magnetiza- tion directions, subsamples of chondrules also showed randomized magnetizations (Figure 12.3a). These random magnetization directions within a single chondrule sum up to an overall random direction for the chondrule, thereby explaining the positive conglomerate test but invalidating the inference that the magnetization predates accretion. Such heterogeneity at <1-mm length scales is not consistent with a primary NRM acquired during chondrule cooling. At the same time, high-resolution magnetic mapping shows that magnetization in Allende chondrules is associated closely with altered mesostases [Figure 12.4; Glenn et al. (2017)], which formed after parent body accretion. These lines of evidence indicate that the non- unidirectional magnetization in Allende chondrules was acquired in the presence of a weak magnetic field during or after aqueous alteration on the CV parent body. The misidentification of a parent body magnetic record in Allende chondrules as a primary nebular NRM highlights the importance of measuring subsamples of individual chondrules to 332 Roger R. Fu, et al.

Figure 12.3 Equal area stereonet projections showing contrasting magnetization distributions in chondrules and chondrule sub-samples from the (a) Allende and (b) Semarkona meteorites. Data points represent the direction of high temperature or high coercivity magnetization of the indicated sample plotted using a Lambert equal-area projection where hollow (solid) symbols correspond to the upper (lower) hemispheres. Samples from each meteorite are mutually oriented. Note that subsamples of Allende chondrules show random directions, while dusty olivine-bearing chondrules from Semarkona have internally consistent directions. avoid the possibility of a false-positive conglomerate test resulting from locally randomized magnetizations. As importantly, careful identification of the mineralogical basis of magnetiza- tion is necessary to distinguish a primary NRM from the product of later remagnetization. Motivated in part by the latter requirement, recent experiments on chondrules from the Semarkona LL3.00 ordinary chondrite have focused on dusty olivine grains, which contain micrometer to submicrometer grains of Fe-metal exsolved from Fe-bearing olivine during primary cooling of the chondrules under relatively reducing conditions (Jones and Danielson, 1997; Leroux et al., 2003). Because these Fe-metal grains are situated within the host olivine, they are effectively shielded from the aqueous alteration that has severely affected the matrix and chondrules of most weakly metamorphosed chondrites of petrologic type 3 and lower (Uehara and Nakamura, 2006). As a further advantage of studying dusty olivine-bearing chondrules, experiments on artificially synthesized dusty olivines have permitted calibrations of room-temperature paleointensity protocols with accuracies of ~30 percent (Lappe et al., 2013), which is a dramatic improvement over typical uncertainties of several hundred percent associated with these methods (see Section 12.2.2). To test the primary origin of dusty olivine magnetization in Semarkona, Fu et al. (2014b) performed a conglomerate test using eight chondrules, and, furthermore, divided two chon- drules into two subsamples each to test the internal unidirectionality of magnetization. Five of the seven dusty olivine-bearing chondrules subjected to AF demagnetization carried identifi- able, origin-trending, high-coercivity components. Such magnetizations are likely to have a Records of Magnetic Fields in the Chondrule Formation Environment 333

Figure 12.4 Optical and magnetic microscopy of chondrules from the Allende and Semarkona meteorites. (a) Reflected light optical photomicrograph showing a barred olivine chondrule in Allende with altered mesostasis regions highlighted. (b) Quantum diamond microscope (QDM) map of magnetic fields due to a near-saturation isothermal remanent magnetization (IRM) in the upward, in-plane direction. Note the correspondence between magnetic sources and altered mesostasis. (c) Transmitted light photomicrograph of a dusty olivine-bearing chondrule from Semarkona with dusty olivine regions highlighted. (d) QDM map of magnetic fields due to an out-of-plane IRM. Note the spatial correspondence between magnetic sources and dusty olivines, which were formed during chondrule formation and permit the preservation of a primary NRM. Images in panels (a) and (b) are after Glenn et al. (2017). Images in panels (c) and (d) are after Fu et al. (2014b). (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.) primary origin because their high-coercivity carrier grains are difficult to remagnetize, while their origin-trending directions imply that no higher-coercivity, and therefore higher-stability, magnetization exists in the sample. Meanwhile, thermal demagnetization of one dusty olivine- bearing chondrule isolated a high temperature component of magnetization that is blocked between 400 and 600 C. Given that the cooling timescale in the laboratory (~5 min) is much 334 Roger R. Fu, et al. shorter than the timescale of chondrule formation (several hours), these laboratory unblocking temperatures correspond to temperatures during natural chondrule cooling of 360 C and 580 C, respectively (Garrick-Bethell and Weiss, 2010). Such a temperature range is fully consistent with a primary NRM carried by Fe-metal that was subsequently partially demagnet- ized during parent body metamorphism at ~200 C for ~1 My. Similar to the Allende chondrules discussed above, the magnetization directions of individual chondrules are randomly distributed (Figure 12.3b). However, subsamples of single Semarkona chondrules showed consistent magnetization directions. This critical piece of evidence strongly suggests that dusty olivine- bearing chondrules in Semarkona, unlike chondrules in Allende, preserve a primary NRM recorded in the solar nebula during chondrule cooling. Room temperature paleointensity experiments on five dusty olivine-bearing chondrules from Semarkona yielded a raw mean of 27 µT, which resulted in a final chondrule cooling field intensity estimate of 54 21 µT, assuming continuous rotation of chondrules around a single axis during cooling. The formation location of our measured chondrules in the protoplanetary disk is poorly known and depends on the extent of radial migration both before and after accretion. Assuming Semarkona chondrules were formed in the protoplanetary disk at between 2 and 3 AU corresponding to the modern asteroid belt, this paleointensity is consistent with both the impact and shock wave models for chondrule formation but is likely too low to be compatible with the x-wind model. The existence of inferred primary NRM in six out of eight Semarkona dusty olivines suggests that approximately 75 percent of chondrules avoided reorientation of their spin axes due to interchondrule collisions during cooling across the 360–580 C unblocking temperature range. We therefore estimate the number density of chondrules during formation (see Section 12.2.1) assuming a collision probability of 0.25, a cooling rate of ~100 Ch–1 inferred for the 600–400 C cooling interval (Schrader et al., 2016b), and uniform chondrule diameters of 0.5 mm correspond- ing to the dusty olivine-bearing specimens. Estimates of mutual velocities among chondrules vary widely between approximately 0.001 and 1 m s–1 (Cuzzi and Hogan, 2003; Carrera et al., 2015), where the higher value is a mean mutual velocity due to high levels of nebular turbulence (α parameter of 10–2) and the lower value is based on a maximum likelihood analysis of particles with similar mean mutual velocities, but considering only particles located close to each other. Higher concentrations of solids in the chondrule formation region may have resulted in correlated motions and even lower mutual velocities. Future models of chondrule mutual velocities that consider in detail the feedback between particle and gas motion [e.g., (Johnson et al., 2015)] are required to create more robust estimates of chondrule collision rates. Adopting the collision probability inferred from chondrule paleomagnetism measurements, the lower and upper bounds on mutual velocity cited above correspond to chondrule number densities of 4 104 and 40 m3, respectively, as the chondrules cooled through the 360 to ~580 C interval. The higher number densities are consistent with the 103–104 m3 values at temperatures above the silicate solidus based on the abundance of sodium in olivine (Alexander and Ebel, 2012). Meanwhile, the lowest estimates of chondrule number density derived above would imply a decrease in density by at least two orders of magnitude during cooling, suggesting rapid expansion of the chondrule formation region. The lower number density estimates presented above would be consistent with those predicted in the nebular shock model of chondrule formation for chondrule enrichment ratios Records of Magnetic Fields in the Chondrule Formation Environment 335 of greater than 1800 relative to the background solid density (Desch and Connolly, 2002). However, such high enrichment ratios would imply heating of the chondrules to >2,000 K, which is inconsistent with the preservation of relict dusty olivine grains. As such, the Semarkona dusty olivine-bearing chondrules likely did not form via the shock mechanism. Alternatively, the lack of an identifiable high-coercivity magnetization in two chondrules may be due to other factors such as the lack of sufficient high-coercivity ferromagnetic grains or the alignment of the chondrule spin axes nearly perpendicular to the background field. In this case, more than 75 percent of chondrules may have avoided collisions during cooling below the Curie temperature, implying lower solid enrichments. Future paleomagnetic analyses of a greater number of chondrules, particularly using thermal demagnetization, are required for more robust constraints on the fraction of chondrule that carry a unidirectional magnetization. For the impact jetting model of chondrule formation, estimates of inter-chondrule velocities may be complicated by the effect of correlated particle motion. The computation by Johnson et al. (2015) predicts 4 collisions per hour for each chondrule during the late cooling stage corresponding to our sampled temperature range, which appears inconsistent with the acquisi- tion of a primary NRM in most dusty olivine-bearing chondrules in Semarkona. However, the collision rate and number density estimates in Johnson et al. (2015) were based on modeling jetted material as a disk that expands only in the radial direction. A more realistic model that includes vertical expansion of this disk would likely reduce this collision rate at late times. The paleomagnetic measurement of a greater number of chondrules with primary ferromag- netic mineralogy and a history of minimal secondary alteration will be necessary to confirm and expand the results discussed here. Studying chondrules from a range of meteorite classes will expand the temporal and spatial coverage of the paleomagnetic record. Although the most reliable record of magnetic fields during chondrule formation so far is derived from dusty olivine chondrules from a single meteorite, the youngest Al–Mg ages of Semarkona chondrules differ from the oldest at the >4σ level (Kita et al., 2013), implying that the eight chondrules studied in Fu et al. (2014b) likely sample a number of chondrule formation events over a time span of at least several hundred thousand years. Other chondrite groups, such as the Renazzo- like carbonaceous chondrites, are expected to yield significantly younger chondrules (Schrader et al., 2017), potentially constraining changes in the nebular magnetic field strength through time. The formation of sampled chondrules at multiple ages implies that the sampled space is in excess of the volume of a single chondrule formation event. However, given broad uncertainties in the degree of planetesimal migration within the protoplanetary disk, the spatial coverage of chondrules from single and multiple chondrule groups remains highly uncertain.

12.4 Conclusions and Future Directions

Chondrule paleomagnetism provides unique constraints on the formation mechanism of chon- drules as well as on the properties of their cooling environment. Early attempts to measure chondrule magnetization were hindered by the challenges of maintaining mutual orientation, inadequate characterization of the chondrules’ ferromagnetic mineralogy, and the lack of suitably sensitive and high–spatial-resolution magnetometers. Advances in paleomagnetic laboratory protocols now permit the mutual orientation of chondrules with better than 5 accuracy (Fu et al., 2014b; Shah et al., 2017), while new magnetometers such as the SQUID 336 Roger R. Fu, et al.

– microscope can detect magnetizations as weak as 5 10 14 Am2 (Fu et al., 2017), which are more than a thousand times weaker than those of the first separated chondrules measured from Allende. At the same time, the quantum diamond microscope (QDM; Figure 12.4c) is capable of mapping the distribution of magnetization at ~1 µm spatial resolution, permitting the identification of ferromagnetic carriers and helping to determine the origin of magnetization. Detailed information has also become available for the alteration and metamorphic history of primitive chondrites [e.g., Alexander et al., (1989), Brearley and Krot (2012), Schrader et al. (2015)]. Combined with high-resolution compositional and magnetic mapping, these insights permit the robust identification of primary ferromagnetic phases in chondrules and the isolation of magnetization recording nebular magnetic fields. Although few published studies so far have used the full complement of these techniques to reach strong conclusions about the origin and paleointensity of single chondrule magnetizations, studies in the near future may be expected to improve the sample statistics of already studied samples, such as Semarkona, and to produce results from a broader range of chondrite groups. At the same time, similar paleomagnetic experiments on nonchondrule inclusions such as igneous CAIs may also extent the coverage of paleomagnetic information to the formation region of these objects. In time, these experiments on chondrules with different formation times may reveal new clues to the mechanism of chondrule formation and the evolution of chondrule production regions through time.

Acknowledgments

We thank X. -N. Bai, A. J. Brearley, J. Gattacceca, A. Johansen, and A. N. Krot for helpful discussions that improved the manuscript. R. R. F. was supported in part by a Postdoctoral Fellowship from the Lamont-Doherty Earth Observatory. B. P. W. thanks the NASA Emerging Worlds (grant NNX15AH72G), the NASA Solar System Exploration and Research Virtual Institute (grant NNA14AB01A), the U. S. Rosetta program, and Thomas F. Peterson Jr. for their generous support.

References

Alexander, C. M. O’D., Barber, D. J., and Hutchison, R. (1989). The microstructure of Semarkona and Bishunpur. Geochim. Cosmochim. Acta 53, 3045–3057. Alexander, C. M. O’D., and Ebel, D. S. (2012). Questions, questions: Can the contradictions between the petrologic, isotopic, thermodynamic, and astrophysical constraints on chon- drule formation be resolved? Meteorit. Planet. Sci. 47, 1157–1175. Asphaug, E., Jutzi, M., and Movshovitz, N. (2011). Chondrule formation during planetesimal accretion. Earth Planet. Sci. Lett. 308, 369–379. Bai, X. -N. (2016). Towards a global evolutionary model of protoplanetary disks. Astrophys. J. 821, 80. Bai, X. -N. (2015). Hall effect controlled gas dynamics in protoplanetary disks. II. Full 3D simulations toward the outer disk. Astrophys. J. 798, 84. Bai, X. -N. (2014). Hall-effect-controlled gas dynamics in protoplanetary disks. I. Wind solutions at the inner disk. Astrophys. J. 791, 137. Bai, X. -N. (2011). Magnetorotational-instability-driven accretion in protoplanetary disks. Astrophys. J. 739, 1–19. Records of Magnetic Fields in the Chondrule Formation Environment 337

Bai, X. -N., and Goodman, J. (2009). Heat and dust in active layers of protostellar disks. Astrophys. J. 701, 737–755. Balbus, S. A. (2003). Enhanced angular momentum transport in accretion disks. Annu. Rev. Astron. Astrophys. 41, 555–597. Bitsch, B., Johansen, A., Lambrechts, M., and Morbidelli, A. (2015). The structure of proto- planetary discs around evolving young stars. Astron. Astrophys. 575, A28. Brearley, A. J., and Krot, A. N. (2012). Metasomatism in the early solar system: The record from chondritic meteorites, in: D.E. Harlov and H. Austrheim (Eds.), Metasomatism and the Chemical Transformation of Rock., 659–789. Berlin: Springer-Verlag. Carporzen, L., Weiss, B. P., Elkins-Tanton, L. T., et al. (2011). Magnetic evidence for a partially differentiated carbonaceous chondrite parent body. Proc. Natl. Acad. Sci. USA 108, 6386–6389. Carrera, D., Johansen, A., and Davies, M. B. (2015). How to form planetesimals from mm-sized chondrules and chondrule aggregates. Astron. Astrophys. 579, A43. Ciesla, F. J., Lauretta, D. S., and Hood, L. L. (2004). The frequency of compound chondrules and implications of chondrule formation. Meteorit. Planet. Sci. 39, 531–544. Cuzzi, J. N., and Hogan, R. C. (2003). Blowing in the wind I. Velocities of chondrule-sized particles in a turbulent protoplanetary nebula. Icarus 164, 127–138. Desch, S. J., and Connolly, H. C. (2002). A model of the thermal processing of particles in solar nebula shocks: Application to the cooling rates of chondrules. Meteorit. Planet. Sci. 37, 183–207. Desch, S. J., and Mouschovias, T. C. (2001). The magnetic decoupling of star formation. Astrophys. J. 550, 314–333. Flock, M., Ruge, J. P., Dzyurkevich, N., et al. (2015). Gaps, rings, and non-axisymmetric structures in protoplanetary disks – From simulations to ALMA observations. Astron. Astrophys. 574, A68. Fu, R. R., Lima, E. A., and Weiss, B. P. (2014a). No nebular magnetization in the Allende CV carbonaceous chondrite. Earth Planet. Sci. Lett. 404, 54–66. Fu, R. R., and Weiss, B. P. (2012). Detrital remanent magnetization in the solar nebula. J. Geophys. Res. 117, E02003. Fu, R. R., Weiss, B. P., Lima, E. A., et al. (2014b). Solar nebula magnetic fields recorded in the Semarkona meteorite. Science. 346, 1089–1092. Fu, R. R., Weiss, B. P., Lima, E. A., et al. (2017). Evaluating the paleomagnetic potential of single zircon crystals using Bishop Tuff zircons. Earth Planet. Sci. Lett. 458, 1–13. Gammie, C. F. (1996). Layered accretion in T Tauri disks. Astrophys. J. 457, 355–362. Garrick-Bethell, I., and Weiss, B. P. (2010). Kamacite blocking temperatures and applications to lunar magnetism. Earth Planet. Sci. Lett. 294, 1–7. Glenn, D. R., Fu, R. R., Kehayias, P., et al. (2017). Micrometer-scale magnetic imaging of geological samples using quantum diamond microscopy. Geochem. Geophys. Geosyst. 18, 3254–3267. Guilet, J., and Ogilvie, G. I. (2014). Global evolution of the magnetic field in a thin disc and its consequences for protoplanetary systems. Mon. Not. R. Astr. Soc. 441, 852–868. Hayashi, C. (1981). Structure of the solar nebula, growth and decay of magnetic fields and effects of magnetic and turbulent viscosities on the nebula. Sup. Prog. Theor. Phys.70,35–53. Hewins, R. H., Connolly, H. C., Lofgren, G. E., and Libourel, G. (2005). Experimental constraints on chondrule formation. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk. ASP Conference Series, 341, 286–316. San Francisco, CA: Astronomical Society of the Pacific. Johansen, A. (2009). The role of magnetic fields for planetary formation. In K. G. Strassmeier, A. G. Kosovichev, and J. E. Beckman (Eds.), Cosmic Magnetic Fields: From Planets, to Stars and Galaxies. Proc. IAU Symp. 259, 119–128. Cambridge, UK: Cambridge Univer- sity Press. 338 Roger R. Fu, et al.

Johnson, B. C., Minton, D. A., Melosh, H. J., and Zuber, M. T. (2015). Impact jetting as the origin of chondrules. Nature 517, 339–341. Jones, R. H., and Danielson, L. R. (1997). A chondrule origin for dusty relict olivine in unequilibrated chondrites. Meteorit. Planet. Sci. 32, 753–760. Kimura, M., Grossman, J. N., and Weisberg, M. K. (2008). Fe-Ni metal in primitive chondrites: Indicators of classification and metamorphic conditions for ordinary and CO chondrites. Meteorit. Planet. Sci. 43, 1161–1177. Kita, N. T., and Ushikubo, T. (2012). Evolution of protoplanetary disk inferred from 26Al chronology of individual chondrules. Meteorit. Planet. Sci. 47, 1108–1109. Kita, N. T., Yin, Q. -Z., MacPherson, G. J., et al. (2013). 26Al-26Mg isotope systematics of the first solids in the early solar system. Meteorit. Planet. Sci. 48, 1–18. Krot, A. N., Petaev, M. I., Scott, E. R. D., et al. (1998). Progressive alteration in CV3 chondrites: More evidence for asteroidal alteration. Meteorit. Planet. Sci. 33, 1065–1085. Lanoix, M., Strangway, D. W., and Pearce, G. W. (1978). The primordial magnetic field preserved in chondrules of the Allende meteorite. Geophys. Res. Lett.5,73–76. Lappe, S. -C. L. L., Harrison, R. J., Feinberg, J. M., and Muxworthy, A. (2013). Comparison and calibration of nonheating paleointensity methods: A case study using dusty olivine. Geochem. Geophys. Geosyst. 14, 1–16. Leroux, H., Libourel, G., Lemelle, L., and Guyot, F. (2003). Experimental study and TEM characterization of dusty olivines in chondrites: Evidence for formation by in situ reduc- tion. Meteorit. Planet. Sci. 38, 81–94. Lesur, G., Kunz, M. W., and Fromang, S. (2014). Thanatology in protoplanetary discs: The combined influence of Ohmic, Hall, and ambipolar diffusion on dead zones. Astron. Astrophys. 566, A56. Levy, E. H., and Araki, S. (1989). Magnetic reconnection flares in the protoplanetary nebula and the possible origin of meteorite chondrules. Icarus 81, 74–91. Levy, E. H., and Sonett, C. P. (1978). Meteorite magnetism and early solar system magnetic fields. In T. Gehrels (Ed.), Protostars and Planets: Studies of Star Formation and of the Origin of the Solar System., 516–532. Tucson, AZ: University of Arizona Press. McNally, C. P., Hubbard, A., Mac Low, M. -M., Ebel, D. S., and D’Alessio, P. (2013). Mineral processing by short circuits in protoplanetary disks. Astrophys. J. 767, L2. Miura, H., Nakamoto, T., and Doi, M. (2008). Origin of three-dimensional shapes of chon- drules: I. Hydrodynamics simulations of rotating droplet exposed to high-velocity rarefied gas flow. Icarus 197, 269–281. Nübold, H., and Glassmeier, K. -H. (2000). Accretional remanence of magnetized dust in the solar nebula. Icarus 144, 149–159. Okuzumi, S., Takeuchi, T., and Muto, T. (2014). Radial transport of large-scale magnetic fields in accretion disks. I. Steady solutions and an upper limit on the vertical field strength. Astrophys. J. 785, 127. Schrader, D. L., Connolly, H. C., Lauretta, D. S., et al. (2015). The formation and alteration of the Renazzo-like carbonaceous chondrites III: Toward understanding the genesis of ferro- magnesian chondrules. Meteor. Planet. Sci. 50, 15–50. Schrader, D. L., Davidson, J., and McCoy, T. J. (2016a). Widespread evidence for high-temperature formation of pentlandite in chondrites. Geochim. Cosmochim. Acta 189, 359–376. Schrader, D. L., Fu, R. R., and Desch, S. J. (2016b). Evaluating chondrule formation models and the protoplanetary disk background temperature with low-temperature, sub-silicate solidus chondrule cooling rates. Lunar Planet. Sci. Conf. XLVII, 1180. Schrader, D. L., Nagashima, K., Krot, A. N., et al. (2017). Distribution of 26Al in the CR chondrite chondrule-forming region of the protoplanetary disk. Geochim. Cosmochim. Acta 201, 375–302. Shah, J., Bates, H. C., Muxworthy, A. R., et al. (2017). Long-lived magnetism on chondrite parent bodies. Earth Planet. Sci. Lett. 475, 106–118. Records of Magnetic Fields in the Chondrule Formation Environment 339

Shu, F. H., Shang, H., Glassgold, A. E., and Lee, T. (1997). X-rays and fluctuating x-winds from protostars. Science. 277, 1475–1479. Shu, F. H., Shang, H., and Lee, T. (1996). Toward an astrophysical theory of chondrites. Science. 271, 1545–1552. Stepinski, T. F. (1992). Generation of dynamo magnetic fields in the primordial solar nebula. Icarus 97, 130–141. Sugiura, N., Lanoix, M., Strangway, D. W., (1979). Magnetic fields of the solar nebula as recorded in chondrules from the Allende meteorite. Phys. Earth Planet. Inter. 20, 342–349. Sugiura, N., and Strangway, D. W. (1985). NRM directions around a centimeter-sized dark inclusion in Allende. Proc. Lunar Planet. Sci. Conf. XV, C729–C738. Swartzendruber, L. J., Itkin, V. P., and Alcock, C. B. (1991). The Fe-Ni (iron-nickel) system. J. Phase Equilibria 12, 288–312. Tachibana, S., Nagahara, H., and Mizuno, K. (2006). Constraints on cooling rates of chondrules from metal-troilite assemblages. Lunar Planet. Sci. Conf. XXXVII. Takac, M., and Kletetschka, G. (2015). Meteorite movement during deceleration studied analogically with magnetic remanence in the bullet. In AGU Fall Meeting. Abstract # GP43B-1244. Tauxe, L. (2010). Essentials of Paleomagnetism. Berkeley, CA: University of California Press. Tsuchiyama, A., Shigeyoshi, R., Kawabata, T., et al. (2003). Three-dimensional structures of chondrules and their high-speed rotation. Lunar Planet. Sci. Conf. XXXIV. Turner, N. J., and Sano, T. (2008). Dead zone accretion flows in protostellar disks. Astrophys. J. Lett. 679, L131–L134. Uehara, M., Gattacceca, J., Leroux, H., Jacob, D., and van der Beek, C. J. (2011). Magnetic microstructures of metal grains in equilibrated ordinary chondrites and implications of paleomagnetism of meteorites. Earth Planet. Sci. Lett. 306, 241–252. Uehara, M., and Nakamura, N. (2006). Experimental constraints on magnetic stability of chondrules and the paleomagnetic significance of dusty olivines. Earth Planet. Sci. Lett. 250, 292–305. van der Marel, N., van Dishoeck, E. F., Bruderer, S., et al. (2013). A major asymmetric dust trap in a transition disk. Science. 340, 1199–1202. Villeneuve, J., Chaussidon, M., and Libourel, G. (2009). Homogeneous distribution of 26Al in the Solar System from the Mg isotopic composition of chondrules. Science. 325, 985–988. Wang, H., Weiss, B. P., Bai, X. -N., et al. (2017). Lifetime of the solar nebula constrained by meteorite paleomagnetism. Science. 355, 623–627. Wardle, M. (2007). Magnetic fields in protoplanetary disks. Astrophys. Sp. Sci. 311, 35–45. Wasilewski, P. (1981). New magnetic results from Allende C3(V). Phys. Earth Planet. Inter. 26, 134–148. Wasilewski, P. J., and O’Bryan, M. V. (1994). Chondrule magnetic properties. Lunar Planet. Sci. Conf. XXV, 1467. Weiss, B. P., and Elkins-Tanton, L. T. (2013). Differentiated planetesimals and the parent bodies of chondrites. Annu. Rev. Earth Planet. Sci. 41, 21. Weiss, B. P., Gattacceca, J., Stanley, S., Rochette, P., and Christensen, U. R. (2010). Paleo- magnetic records of meteorites and early planetesimal differentiation. Sp. Sci. Rev. 152, 341–390. Weiss, B. P., and Tikoo, S. M. (2014). The lunar dynamo. Science. 346, 1246753–1. Zhu, Z., Stone, J. M., and Rafikov, R. R. (2013). Low-mass planets in protoplanetary disks with net vertical magnetic fields: The planetary wake and gap opening. Astrophys. J. 768, 143.

Part II Possible Chondrule-Forming Mechanisms

13 Formation of Chondrules by Planetesimal Collisions

brandon c. johnson, fred j. ciesla, cornelis p. dullemond, and h. jay melosh

Abstract

Chondrules are the millimeter-scale previously molten droplets found in chondritic meteorites. These pervasive yet enigmatic particles hint at energetic processes at work in the nascent solar system. Chondrules and chondrites are well studied and many of the details about their compositions, ages, and thermal histories are well known. Without the proper context of a formation mechanism, however, we can only imagine what chondrules may reveal about the processes at work in the early solar system. In this chapter, we explore the hypothesis that chondrules were formed by impacts between growing planetary embryos. Specifically, we focus on shock heating associated with accretionary impacts as a means for melting chondrule precursor material. Although we discuss previous work on impact origin for chondrules, much of this chapter focuses on a new incarnation of this old idea, the impact jetting model. We explore the predictions of this model and its implications for our understanding of early solar system history and meteoritics. Throughout the chapter, we discuss potential issues and uncer- tainties with the model while identifying avenues for further development and testing of the impact origin hypothesis.

13.1 Introduction

An impact origin for chondrules was hypothesized as early as 1952, with chondrites formed in the high temperature region produced at the center of the impact (Urey, 1952). As subsequent laboratory studies and measurements became more comprehensive, however, it seemed impos- sible to reconcile an impact origin for chondrules with these observations and associated constraints as petrologic and chemical data seemed at odds with this idea. For example, impact ejecta on the Earth and Moon are dominated by brecciated rock and less abundant impact melt is in the form of glassy spherules, yet chondrites and other meteorites lack brecciated rock and glassy or cryptocrystalline chondrules, which has been used to argue against an impact origin (Taylor et al., 1982). As a result, chondrule formation mechanisms that occur in pre-accretionary environments – that is, an environment where chondrule precursors (dust aggregates) have not been accreted into planetesimals – have been strongly favored over postaccretionary environments, where the precursor material originates in early planetesimals. Shock waves in nebular gas are examples of preaccretionary mechanisms for which quantitative models have been developed, enabling

343 344 Brandon C. Johnson, et al. detailed comparisons between theoretical predictions and chondrule and chondrite properties (Desch et al., 2012). For details on the shock wave model and the comparison with the meteoritic record, see Chapter 15. Recent measurements of the volatile content of chondrule olivine, however, have led to serious questions about all mechanisms occurring in a low pressure preaccretion environment as expected under canonical solar nebula conditions (Alexander et al., 2008; Alexander and Ebel, 2012; Fedkin and Grossman, 2013). The higher-than-expected volatile contents suggest that chondrules formed in regions with some combination of high dust enrichments and high pressures to suppress evaporation and isotopic fractionation of volatiles at the time when chondrules were molten (Alexander et al., 2008; Alexander and Ebel, 2012; Fedkin and Grossman, 2013). Dust enrichment is a measure of the abundance of solids in a system relative to a system of solar composition (defined as a dust enrichment of 1). More precisely, a dust enrichment of n means the system has a bulk composition described by adding (n – 1) units of CI dust to a solar composition (Ebel and Grossman, 2000). Fedkin and Grossman (2013) found that at pressures of 100 Pa (10–3 bar), which are on the high end expected for the solar nebula (e.g., Ruden and Pollack, 1991), dust enrichments larger than 106 are needed to explain the composition of chondrule olivine. These dust enrichments are >104 times more than the maximum values estimated by some authors (e.g., Cassen, 2001). The chondrules found in CB chondrites represent a notable exception to the general preference for formation in a preaccretion environment. The CB chondrites are the youngest among known chondrite groups and formed ~4.8 0.3 Ma after formation of Ca, Al-rich inclusions (CAIs), the earliest solar system solids dated (Krot et al., 2005; Connelly et al., 2012; Bollard et al., 2015). The late formation time, chondrule ages consistent with formation in a single event, and the presence of chemically zoned Fe,Ni metal grains that condensed directly from vapor enriched in dust by factors of >107 (Campbell et al., 2002), suggest that the CB chondrites are the product of an impact generated gas-melt plume (Krot et al., 2005). Compared to other chondrites, CB chondrites are highly-depleted in volatiles (Weisberg et al., 2001) and Fe,Ni metal grains show evidence of isotopic fractionation (Alexander and Hewins, 2004; Richter et al., 2014). As chondrules appear to have formed in environments with very high concentrations of solids, and CB chondrites are accepted as the product of collisions in the early solar system, this has renewed interest in understanding whether the other chondrites and chondrules may also be consistent with formation in by impacts (Sanders and Scott, 2012; Dullemond et al., 2014; Johnson et al., 2015). Unlike most other chondrule formation mechanisms, high-velocity impacts are accepted as common events during the early epoch of planet formation (e.g., Davison et al., 2013). Recent studies have challenged some of the other objections raised for an impact origin for chondrules. Many argued against an impact origin for chondrules on the basis that planetesimals would only accrete once chondrules formed. Chondrules may have formed at nearly the same time as CAIs and continued to be produced for the next few million years (Scott, 2007; Connelly et al., 2012). Isotopic analyses of magmatic iron meteorites, however, suggest that the parent bodies of magmatic iron meteorites greater than 10–100 km in diameter likely accreted ~0.1–0.3 Myr after the formation of CAIs, indicating planetesimals had formed before most chondrules (Kruijer et al., 2014). Once such objects formed, they would perturb one another’s orbits and experience a number of collisions, which would in turn create planetary Formation of Chondrules by Planetesimal Collisions 345 embryos roughly the size of the Moon or Mars all within ~0.1–1 Myr (Chambers, 2004; Chiang and Youdin, 2010; Johansen et al., 2014). Thus, as early as 0.1 Myr after the formation of CAIs, we anticipate planetesimals were accreting onto planetary embryos as large as the Moon or Mars. The conclusion that planetary embryos existed at the time of chondrule formation is further supported by isotopic measurements indicating Mars itself accreted on a timescale of 1.8 Myr and was larger than the Moon for a majority of the time chondrules were forming (Dauphas and Pourmand, 2011). Thus, it appears chondrule formation must have been occurring in an environment where planetesimals and their collisions were abundant.

13.2 Impact Jetting

Chondrules are thought to form in a gas rich environment as without abundant gas, chondrules cannot be concentrated to form planetesimals. This conclusion is consistent with astronomical observations suggesting the typical lifetime of a protoplanetary disk (i.e., when nascent stars are surrounded by a gas-rich disks) is a few Myr (Evans et al., 2009), comparable to the estimated duration of chondrule formation (Bollard et al., 2017). In addition to facilitating efficient accretion, the gas effectively damps eccentricities and inclinations of planetesimals, ultimately leading to relatively low impact velocities dominated by the mutual escape velocity of the target and projectile (Hasegawa et al., 2015; Johnson et al., 2015). Once large bodies form, including Moon- and Mars-sized embryos, impact velocities of a few to several km/s may occur. Even higher impact velocities can develop as Jupiter grows and begins to excite the orbits of bodies near resonances (Weidenschilling et al., 1998). Impact experiments and numerical models, however, suggest velocities exceeding ~10 km/s are required to produce appreciable impact melt (Pierazzo et al., 1997). Even for high velocity impacts, ejecta found around terrestrial impact craters are overwhelmingly dominated by cold solid fragments, which are not seen in the meteorite record (Taylor et al., 1982). For these reasons, it has been difficult to imagine how impacts could produce the large number of chondrules seen in chondrites, especially the ordinary chondrites, which are 60–80 percent chondrules by volume and represent more than 80 percent of meteorite falls (Scott and Krot, 2003).1 If the target and impactor are molten, questions about the abundance of cold solid ejecta and how much melt low velocity impacts can create are circumvented. If heating by short-lived 26Al has melted much of the interior of the target and projectile, a collision could efficiently splash out molten material (Asphaug et al., 2011; Sanders and Scott, 2012). However, material ejected by low-velocity collisions would create meter-scale blobs of melt instead of the millimeter-sized spherules that are needed to form chondrules (Johnson et al., 2015). In addition, the partially molten target bodies would likely be differentiated. Thus, it is not clear that this model could produce chondrules with primitive compositions (Taylor et al., 1982). For a more detailed description of molten planetesimal splashing models, see Chapter 14. Impact jetting, however, may provide a means of reconciling an impact origin for chondrules with observational constraints (Johnson et al., 2015). Jetting is an extreme process that occurs

1 It should be remembered, however, that it is unclear how representative the meteorites we collect are of the bodies in the asteroid belt, and thus high abundances in the meteorite record is not necessarily indicative of abundance in the asteroid belt (Scott and Krot, 2003). 346 Brandon C. Johnson, et al. very early in an impact when the projectile is still coming into contact with the target. During this short time, a small amount of material is shocked to high pressures and temperatures as it is squirted out from the region where the projectile is coming into contact with the target. This material is typically ejected at velocities exceeding the impact velocity. A schematic of this process is shown in Figure 13.1. Our understanding of the process greatly benefits from the fact that impact jetting is observed in laboratory scale impacts (Kurosawa et al., 2015). Jetting is the result of the physics of oblique shockwaves. The initiation of jetting is determined by the angle described by the impactor and target, α, as shown in Figure 13.2. Below a critical angle αcr, which depends on material type and impact velocity, unshocked material moves through a single shockwave attached to (i.e., originates from and moves along with) the collision

Figure 13.1 Schematic representation of the formation of chondrules and protomatrix by impact jetting. Early in the impact a small amount (a few percent of a projectile mass) of primitive crustal material is ejected above the escape velocity of the target body. Some of this material is partially melted and breaks up becoming melt droplets. The rest of the material is only lightly shocked and may become protomatrix material. As the jetted material moves outward it radiatively cools. The vast majority of impact ejecta are bound to the target body.

Figure 13.2 Schematic showing how the angle α is defined for the vertical impact of a spherical impactor on a flat target. The thick black curve labeled S shows the extent of the shockwave and vimp is the impact velocity. The gray region is material that is already shocked. The darker gray dots represent a locus of points where the projectile is just coming into contact with the target. We refer to this as the collision point. After Johnson et al. (2014).2

2 Reprinted from B. C. Johnson, T. J. Bowling, and H. J. Melosh (2014), Jetting during vertical impacts of spherical projectiles. Icarus, Vol 238, Pages 13–22. Copyright (2014), with permission from Elsevier. Formation of Chondrules by Planetesimal Collisions 347 point and no jet forms. The strength of this shockwave increases with α to turn material through the angle α/2. This shock is always stronger than that produced by a planar impact at the same velocity. Above αcr a single shock of arbitrary strength is incapable of turning material through the angle α/2; the shock becomes detached from the collision point and a rarefaction accelerates material, which squirts out from the collision point at a velocity exceeding the impact velocity. During the impact of a spherical projectile, the angle (α) between the spherical projectile and target increases with time such that αcr is inevitably exceeded and jetting occurs (Figure 13.2). In addition to being ejected at speeds greater than the impact velocity, this jetted material is highly shocked. We refer the interested reader to Walsh et al., (1953) for a comprehensive description of the physics of jetting. Many of the details of impact jetting and their relation to chondrule formation remain unexplored. As we describe in Section 13.4, other than initial and final state, the thermodynamic path of jetted material is uncertain. Kieffer (1975) originally suggested that chondrules are the product of jetting during the collision of millimeter- to centimeter-sized objects. However, it is unlikely that small objects would collide at the few km/s needed to cause melting. Further, because jetting involves only a small fraction of mass involved in the impact, a few percent of a projectile mass, much of the mass involved in these impacts would be unmelted. This is true regardless of the size of the impactor and target. If the target bodies are Moon- to Mars-sized and impactors are 10–1,000 km in diameter, Johnson et al. (2015) argue that many of the potential issues with an impact-jetting origin for chondrules can be resolved. Johnson et al. (2015) show that jetting will eject approximately 1 percent of a projectile mass of impact melt when impact velocities exceed 2.5 km/s typical of impacts onto Moon and Mars-sized planetary embryos. The 1 percent of a projectile mass of material is a mix of both projectile and target material. This estimate is independent of the size of the target body except when the impactor and target size are comparable and target curvature increases the jetting efficiency (Johnson et al., 2014). While this jetted material will be ejected above escape velocity of the target body, much of the normal solid ejecta will be gravitationally bound and reaccreted. Thus, the target body acts as a sieve sorting melt rich material from the bulk of the impact ejecta (Figure 13.1). Only the high-velocity material not bound to the target body will accrete onto existing planetesimals or potentially accrete to form completely new planetesimals. This offers an explanation of how impacts can produce chondrites that are dominated by chondrules rather than cold solid fragments. Given that jetting would allow the types of collisions expected in the early solar system to produce a significant amount of melt that is launched away from the colliding bodies, it is worthwhile to consider the fate of this melt and its relationship to meteoritic materials. In the following sections, we critically discuss how the impact-jetting model compares to constraints on chondrule sizes, abundance, cooling rates, volatile content, isotopic fractionation, primitive compositions, and complementarity. We will also discuss how the process relates to the formation of chondrites.

13.3 Chondrule Sizes

Chondrules are typically submillimeter- to centimeter in size (e.g., Friedrich et al., 2015). Although the comparison is imperfect, if chondrules are produced by impacts, their closest 348 Brandon C. Johnson, et al. terrestrial analogs are the impact spherules found in distal impact ejecta layers (Glass and Simonson, 2012). The most familiar of these layers is the K-Pg boundary layer, which marks the end of the reign of the dinosaurs (Schulte et al., 2010). This nearly global layer of impact spherules is the result of the expansion of a massive impact vapor plume, and the spherules in this layer likely condensed directly from vapor (Johnson and Melosh, 2012a, 2012b). The textures of chondrules are not indicative of direct nucleation and condensation from vapor (Lofgren, 1989) and indeed impact velocities in the nascent solar system were much too low to produce significant vaporization (Johnson and Melosh, 2012a; Johnson et al., 2015). In this framework, chondrules would likely be more closely related to impact spherules found in more proximal material originally contained in the less highly shocked ejecta curtain (Schulte et al., 2010; Glass and Simonson, 2012; Johnson and Melosh, 2014). These spherules likely formed by impact melt breaking up and forming droplets, rather than condensing directly from vapor (Schulte et al., 2010; Glass and Simonson, 2012; Johnson and Melosh, 2014). It is not clear that any ejecta deposits corresponding to jetted material have been found. At the 20 km/s impact velocities typical on Earth (Minton and Malhotra, 2010), jetting is inefficient in part due to the increase in critical angle with increasing impact velocity (Johnson et al., 2014) and jetted material would be vaporized. This may indicate that for terrestrial impact ejecta, jetted material is impossible to distinguish from ejecta produced in the impact vapor plume. Despite the noted differences between chondrules and terrestrial impact spherules, further comparison and study of terrestrial impact spherules may provide a better understanding of the origin of chondrules. Terrestrial spherule layers may also act as ground truth for models of chondrule formation by impacts. Johnson and Melosh (2014) determined how the typical size of melt droplets depend on impactor size, impact velocity, and ejection velocity. The model is based on the balance of drag forces associated with acceleration of a mixture of droplets and vapor and surface tension forces. For conditions appropriate for terrestrial ejecta layers, the model produces estimates that agree with observed spherule sizes (Johnson and Melosh, 2014). Applying this model to impact jetted chondrules suggests that 10–1,000 km-scale impactors will produce millimeter-scale chondrules. This model only provides rough estimates of typical droplet size. The largest uncertainties likely come from estimates of the rate of acceleration, a, of the two-fluid mixture. – The droplet size d / a 1/2, so even a factor of 10 error in acceleration would change our estimates by a factor of ~3. Other sources of uncertainty are the critical Weber number, Wec, and σ the droplet surface tensionpffiffiffiffiffiffiffiffiffiffiffi, however, we expect these parameters are better constrained than a and droplet size d / Wecσ (Johnson et al., 2014). In addition to refining size estimates, detailed models of droplet breakup and size evolution may help us understand how chondrule size distributions are established and if any post impact size sorting of chondrules is required to explain chondrule sizes observed in meteorites (Friedrich et al., 2015).

13.4 Chondrule Abundance

Approximately 80 percent of all meteorites found on Earth contain chondrules (Scott, 2007). Although there is some indication that the current-day meteorite record may not be representa- tive of the composition of the asteroid belt (e.g., Schmitz et al., 2016), the overwhelming Formation of Chondrules by Planetesimal Collisions 349 abundance of chondrules in the meteorite record suggests that chondrules and chondrites are a major, if not dominant, fraction of the asteroid belt. Thus, we expect that a successful model of chondrule formation should be able to produce a current-day asteroid belt that includes a large fraction of chondrules by mass. To estimate the total abundance of chondrules produced by impacts, we need estimates of how many chondrules a given impact can produce. Because we assume that material ejected below the escape velocity of the target body does not produce chondrules, we need to know the ejection velocity of the material as well as its thermal state. Shock physics codes, also known as hydrocodes, allow estimates of the amount of partially melted material ejected at a given velocity to be made (e.g., Johnson et al., 2015). The impact models of Johnson et al. (2015) suggest that when impact velocities exceed 2.5 km/s the amount of potentially chondrule- forming mass created by a given impact is on the order of 1 percent of the mass of the projectile. Further, to estimate the total abundance of chondrules produced during the first few Myr of planetary accretion, we also need to know how many collisions there are, what velocities these impacts occur at, and the sizes of the targets and impactors involved. This has been investigated using both a Monte Carlo accretion model (Johnson et al., 2015) and a semianalytic accretion model (Hasegawa et al., 2015). Johnson et al. (2015) couple results of how the ejected chondrule mass depends on impact velocity and escape velocity of the target body to their accretion models. This is in contrast to Hasegawa et al. (2015), who simply assume any impact occurring at a velocity greater than 2.5 km/s produces 1 percent of a projectile mass of chondrules. Both models suggest that the impacts of smaller bodies onto growing planetary embryos (Moon- to Mars-sized bodies) produce the majority of chondrules. Modeling accretion in a nebula three times as massive as the minimum mass solar nebula (MMSN), Johnson et al. (2015) found that 3.3 1022 kg of chondrules form between 2 and 3 astronomical units (AU) from the Sun during the first 5 Myr of accretion, while Hasegawa et al. (2015) find 1023 kg of chondrule are formed in the same region for the similar model conditions; considering the uncertainties associated with these models and the differences between them, their results are in good agreement. To compare our model results to the chondrule abundance implied by the meteorite record, we must first make some estimate of the primordial chondrule abundance required to match the abundance of chondrules in the meteorite record. The mass of the current-day asteroid belt is thought to be depleted by a factor of ~1,000 times from the early asteroid belt based on a MMSN (Weidenschilling, 2011). Because most of the primordial mass is thought to be concentrated in large Moon- to Mars-mass bodies, a numerical depletion of 10–100 times (meaning 90–99 percent of the bodies are removed from the main belt either by being accreted by growing protoplanets, accreted by the sun, or ejected from the solar system) can explain a ~1,000 times depletion in mass (Weidenschilling, 2011). Hood et al. (2009) suggest typical chondritic material contain ~1/3 chondrules by mass. Johnson et al. (2015) assume the current-day asteroid belt is composed of at most ~1/3 chondrules by mass and suggest that producing 3–30 times the main belt mass (10–100 for the 3 MMSN) in chondrules is sufficient to explain their current prevalence in the meteorite record. Accretion models suggest that for a disk with three times the mass of the MMSN, more than 11–33 main belt masses of chondrules could be produced in the main belt region during the first few Myr of accretion (Johnson et al., 2015; Hasegawa et al., 2015). In addition to depletion of the primordial asteroid belt, it may be important to consider the amount of chondritic material that is transported to the asteroid belt (located between 2 and 350 Brandon C. Johnson, et al.

3 AU from the sun) from another location. The accretion models of Johnson et al. (2015) suggest that 13 times more chondrules form in the inner solar system (within 2 AU of the sun) than in the asteroid belt for the 3 MMSN model. Dynamical models suggest that the main belt may be composed of 10–30 percent inner solar system materials that formed at smaller heliocentric distances and then were scattered to their current locations (O’Brien et al., 2007). The most extreme example of material transport to the main belt is the so called ‘Grand-Tack’(Walsh et al., 2011). In this model, after its formation Jupiter migrates into the inner solar system. Jupiter is then captured into a resonance by Saturn and migrates out to its current position. This migration produces a main belt that has a large fraction of inner solar system material (Walsh et al., 2011). The model also suggests that C-type asteroids are outer solar system bodies implanted in the main asteroid belt by this process. If giant planet migration occurred after most chondrules formed (Johnson et al., 2016; Levison et al., 2016) this would require that chondrules also formed in the outer solar system. Recent modeling of the formation of the giant planets through pebble accretion show that Moon- to Mars-sized bodies would form in 1–5 105 years in the outer solar system (Levison et al., 2015). Thus, chondrule forming impacts may be expected in the outer solar system too, which may be consistent with recent work indicating an outer solar system origin for some chondrites (Warren, 2011; Van Kooten et al., 2016; Kruijer et al., 2017). Future work using more complex direct accretion simulations capable of modeling pebble accretion like LIPAD (Levison et al., 2012) may provide the more robust estimates of the abundance of chondrules that impacts can create and the locations where chondrules were forming. Where, when, and how many chondrules were forming is clearly dependent on the dynamics of accretion and history of the early solar system. It is well known that most impacts are oblique, occurring with a typical impact angle of 45 degrees. Johnson et al. (2015) only modeled vertical impacts to estimate the amount of potentially chondrule-forming material a given impact would create. Vickery (1993) estimated that very oblique impacts could jet 10 times more material than vertical impacts with similar conditions; if oblique impacts produce significantly more chondrule forming material, then the estimates of the total abundance of chondrules made during planetary accretion could change considerably. The estimates of Vickery (1993) are based on application of thin plate theory. Recent experiments that resolve impact jetting using ultra-fast imaging show that jet velocities are typically about half the velocity expected from simple theory of impacting thin plates (Kurosawa et al., 2015). Numerical models also exhibit this lower-than-expected jet velocity (Johnson et al., 2014). Although this was originally attributed to limited model resolution (Johnson et al., 2014), the experimental study suggests this lower velocity is physical and indicates impact jetted material has a different thermodynamic path than that implied by thin plate theory (Kurosawa et al., 2015). Moreover, the numerical modeling of Johnson et al. (2014) demonstrates that application of thin plate theory to impact jetting greatly overestimates the total jetted mass. Thus, conclusions that are based on application of thin plate theory to the process of impact jetting should be viewed with skepticism. Gaining a robust understanding of the thermodynamics of jetting is crucial to further test the hypothesis that chondrules are produced by impact jetting, with a focus on the processing of materials in oblique impacts. Despite this caveat, we suspect that oblique impacts will produce more jetted material than head-on collisions. However, without direct simulation using high resolution fully Formation of Chondrules by Planetesimal Collisions 351 three-dimensional models, it is unclear how the impact angle will change the amount of potentially chondrule-forming material produced by a given impact. It is also possible that oblique impacts will jet impact melt at a lower threshold impact velocity. Another area for refinement is to consider the curvature and topography of the target body. Johnson et al. (2014) found that impacts of equal sized bodies produce more than five times the amount of material jetted during a similar impact into a flat target. This is the result of the symmetry of the impact and the fact that the critical angle for jetting is reached much earlier than during an impact into a flat target (Johnson et al., 2014). Inclusion of oblique impacts and target curvature could change estimates of overall chondrule abundance significantly, perhaps by an order of magnitude or more.

13.5 Cooling Rates

The cooling rate of an individual chondrule in empty space is 107 Kh1, while chondrule textures imply cooling rates of ~10–1,000 K h1 (Desch et al., 2012). There are multitudes of melt droplets in an impact-produced jet or plume and as a droplet radiatively cools, it will absorb radiation from other nearby cooling chondrules, prolonging the cooling timescale. Initially chondrule number densities in the plume will exceed 108 m3 and the plume will be optically thick, only radiating from its exposed surface (Dullemond et al., 2014). The problem of a radiatively cooling plume of chondrules is well suited for simple radiative transfer models (Johnson et al., 2015) and for a spherically expanding plume an analytical approximation exists (Dullemond et al., 2014). Both Dullemond et al. (2014) and Johnson et al. (2015) assume the chondrules in the plume have initial temperatures of 2,000 K. Johnson et al. (2015) found that expanding jets, modeled as one-dimensional slabs with densities that evolve as the jets, produced by 100–1,000 km diameter impactors, expand outward will cool at typical rates of 10–1,000 K h1. In contrast to the jet-like geometry modeled by Johnson et al. (2015), Dullemond et al. (2014) consider the geometry of a spherically expanding plume. In both models, the interior of the plume begins cooling later than the outer parts of the plume but typical cooling rates are relatively insensitive to location in the plume, except for very near the outer edge of the plume where cooling rates are much higher (Johnson et al., 2015; Dullemond et al., 2014). For expansion velocities of a few km/s, appropriate for jetted material (Johnson et al., 2015), Dullemond et al. (2014) find the initial scale of the expanding plume must be 10–1,000 km to produce expected chondrule cooling rates. This is somewhat larger than the initial 2–20 km thickness of jets described by Johnson et al. (2015). At late times the radial dimension of the either geometry is r vt, where v is the expansion velocity and t is the time. – The density of material for the jet-like geometry scales is r 2, where the spherically expanding plume has a density that scales as r–3. Thus, the spherically expanding plume will have a lower opacity than the jet-like geometry consistent with the finding that spherically expanding plumes cool faster than a more appropriate jet-like geometry. Recent laboratory experiments suggest that chondrule melts remained at relatively high temperatures for tens of minutes to hours before cooling at high rates (>103 K/hr) (Villeneuve et al., 2015). This is the predicted cooling behaviors of impact produced plumes (Dullemond et al., 2014; Johnson et al., 2015). For example, the temperature of a jet produced by a 100-km 352 Brandon C. Johnson, et al. diameter impactor will remain relatively constant for ~40 minutes before the cooling rate increases to a few thousand K/hr (Johnson et al., 2015).

13.6 Volatile Content and Lack of Isotopic Fractionation

Although chondrules are generally depleted in volatiles as one might expect for ‘drops of fiery rain’ (Sorby, 1877), recent observations show that chondrule olivine is more volatile-rich than one might expect for flash-heated droplets cooling in free space, where vaporization of volatile species would be expected. These observations of the volatile content of chondrule olivine indicate that chondrules formed in highly dust-enriched environments (Alexander et al., 2008; Fedkin and Grossman, 2013). Sodium is one example of a volatile element that would otherwise be lost from chondrules, as it vaporizes at temperatures well below the peak temperatures to which chondrules were exposed. That sodium is found in these olivines means that this vaporization was suppressed somehow. One possibility is that the chondrules formed in such a dust-rich region that while each chondrule lost some sodium, the resulting partial pressure from all the chondrules was large enough to saturate the surrounding gas and prevent further vaporization. Another possibility is that the total pressure around the chondrule was so high that the sodium vapor remained trapped around the chondrule, acting as if the entire gas was saturated, and that there was no subsequent motion of the chondrule with respect to this “confined” gas. The measured sodium content of chondrule olivine requires total pressures and/or dust enrichments orders of magnitude above expected values for the solar nebula (Fedkin and Grossman, 2013). Jetted material quickly reverts to the background pressure of the solar nebula, suggesting that dust enrichment is the major factor determining the volatile content of chondrule olivine. Simple estimates of the evolution of chondrule number density in impact- jetted material, required for cooling rate calculations, suggest that dust enrichments exceed 106 for times longer than it takes for the chondrules to cool to their solidus temperature (Johnson et al., 2015). This indicates that impact jetting can explain the higher-than-expected volatile content of chondrule olivine. In addition to providing further support of this claim, recent calculations by Dullemond et al. (2016) provide insight into the time history of volatile evaporation and condensation in an expanding plume. Their work shows that after an initial phase of evaporation, volatiles can condense back onto chondrules as the plume expands and cools (Dullemond et al., 2016). Although chondrules are relatively depleted in volatiles, they are not depleted in lighter isotopes of these volatile species (Davis et al., 2005). This suggests that chondrules were in equilibrium with their degassed volatiles and formed in regions larger than 150–6000 km in radius (Cuzzi and Alexander, 2006). The model of Johnson et al. (2015) suggests that the outer, faster-moving portion of the jet moves ~6 km/s. After escaping the gravity well of the target 2 ¼ 2 2 body, the outer part of the plume will move with v∞ vej vesc. Assuming an escape velocity of 2.5–5 km/s (appropriate for Moon- to Mars-sized bodies) and a cooling time of hours to days, this velocity implies that the chondrules fill regions 104–105 km across and maintain number densities greater than 103 m3 as they cool through their solidus. This simple calculation shows that the impact jetting model is consistent with the lack of significant isotopic fractionation. Cuzzi and Alexander (2006) suggest that variations in chondrule number density and cooling Formation of Chondrules by Planetesimal Collisions 353 rates may explain some of the variation of chondrule compositions. Chondrule number densities and cooling rates vary significantly within an impact produced plume (e.g., Johnson et al., 2015). More detailed calculations that account for these variations in chondrule number density and cooling rates may help explain some of the variation of chondrule compositions.

13.7 Primitive Composition and Complementarity

CI chondrites have compositions very similar to the solar photosphere (e.g., Palme et al., 2014). Variations in the composition of other chondrite groups are related to condensation temperature (Palme et al., 2014). The compositions of chondrules and matrix material are often described as complementary, where chondrules and surrounding matrix material have distinct compositions but the bulk meteorite has a composition similar to CI chondrites (e.g., Hezel and Palme, 2010). This complementarity may be explained if chondrules are depleted in volatiles while the surrounding matrix material is enriched (Bland et al., 2005). This complementarity seems to indicate that chondrules and matrix material are not physically separated and that the volatiles released by melted chondrules condense onto protomatrix material. The impact jetting model of chondrule formation suggests that lightly shocked (minimally heated) protomatrix is ejected along with partially melted chondrule-forming material (Johnson et al., 2015). As we discuss next, this aspect of the model may be consistent with complementarity. For more on chondrule matrix complementarity, see Chapter 4. Johnson et al. (2015) suggest the lightly shocked solid material ejected above the target bodies’ escape velocity represent protomatrix. This material would be cold and have properties (e.g., grain size and composition) determined by the source material. The relative abundance of protomatrix to chondrule-forming material ejected during any given impact is dependent on impact velocity, vimp, and the escape velocity, vesc, of the target body (Johnson et al., 2015). As vimp/vesc increases, ejected material can transition from being dominated by molten material to being dominated by solids (protomatrix). Volatiles may preferentially condense onto protoma- trix, which has a larger surface area and is initially relatively cold compared to the slowly cooling molten chondrules. The protomatrix, however, will still be exposed to the thermal radiation from hot chondrules and could be melted or evaporated. If the cloud of chondrules and fine-grained solids expands quickly enough, then some vapor never makes it back onto the ejected material (Dullemond et al., 2016). This vapor will mix into the nebular gas and subsequently condense out (over a much longer timescale) onto the solids in the solar nebula: in particular, those with the largest surface area, which is typically the fine-grained dust in the nebula. If the free-floating chondrules and dust subsequently form a new body or accrete onto existing bodies, the final composition of chondrules and matrix will be complementary (pro- vided the chondrules and the dust do not dynamically separate over large scales in the disk) (Dullemond et al., 2016). Further modeling with radiative transfer models coupled with simple condensation models like those of Dullemond et al. (2016) could prove useful for better understanding of how chondrule-matrix complementarity may be achieved. A mechanism that can explain chondrule-matrix complementarity in volatiles does not readily explain comple- mentarity seen in more refractory components like Ca/Al, Al/Ti, or Hf/W (Palme et al., 2014; Becker et al., 2015). 354 Brandon C. Johnson, et al.

In addition to complementarity in refractory elements, recent detailed measurements show that chondrules and matrix are also complementary in tungsten isotopes (Becker et al., 2015; Budde et al., 2016). Budde et al. (2016) suggest that this complementarity is explained by preferential incorporation of presolar grains in matrix material and argue that impact jetting is inconsistent with these observations. This conclusion is based on the assumption that jetted material is completely melted; however, this need not be the case. Impact velocities are typically low in the early solar nebula, and potential chondrule-forming impacts modeled by Johnson et al. (2015) occur at velocities of only a few km/s. At these velocities, much of the chondrule- forming material is partially melted. Indeed, the threshold for potentially chondrule-forming material used by Johnson et al. (2015) is material has bulk temperatures above the solidus after decompression from high pressure. Note also that shock heating in real material is heterogenous (e.g., Bland et al., 2014). Thus, the mechanism proposed by Budde et al. (2016) to explain isotopic complementarity (i.e., that refractory presolar grains remain unmelted and are prefer- entially incorporated in matrix material) is as viable for impact jetting as it is for formation of chondrules in any model that leaves some component of solids unmelted. The details of how refractory grains would preferentially get into the matrix remain unexplored for any model. A better understanding of the thermodynamics and microphysical processes occurring in the impact jet could shed light on the details of these processes. Another possibility is that presolar grains are physically separated from the chondrule formation process (i.e., more likely to be freely floating in the nebula) and are later incorporated into matrix. To achieve complementar- ity, the combination of matrix and chondrules that go into a chondrite must combine to produce a chondrite with a solar composition. This would require that the region where chondrules and matrix are located somehow acts as a closed system over the temporal and spatial scales relevant to chondrite formation. For the jetting model to be consistent with the relatively primitive nature of chondrites and chondrule-matrix complementarity, the jetted material must have a chondritic composition before impact. Magmatic iron meteorites suggest that many bodies were differentiated or differentiating at the time of chondrule formation (Kruijer et al., 2014). Paleomagnetism studies of Allende and thermal models suggest that even differentiated bodies may have had undifferen- tiated chondritic crusts, the thickness of which depends on the size of the body, time of accretion, and abundance of 26Al (Carporzen et al., 2011; Elkins-Tanton et al., 2011; Sanders and Scott, 2012). More detailed modeling of magma buoyancy suggests that relatively thick undifferentiated chondritic crusts may be preserved even on large parent bodies for bulk compositions typical of CV and CM chondrites (Fu and Elkins-Tanton, 2014). Impact models suggest that the high-velocity jet is sourced from near the surface of the impactor and target (Johnson et al., 2014, 2015). Thus, the jetting model is consistent with the primitive nature of chondrites if most bodies had primitive crusts. This suggests that at the time chondrules were forming, bodies with basaltic crusts like Vesta (Jutzi et al., 2013) were rare. Analysis of zircons from eucrite meteorites suggests that Vesta’s basaltic crust was built over ~35 Myr starting approximately 2 Myr after the formation of CAIs, with ages peaking at ~4 Myr after CAI (Roszjar et al., 2016). Thus, it is conceivable that Vesta itself had a primitive crust at the time chondrules were forming. Detailed estimates of the provenance depth of potentially chondrule-forming material based on three-dimensional numerical impact models may pro- vide predictions of the minimum thickness of these required pristine crusts. Additionally, Formation of Chondrules by Planetesimal Collisions 355 coupling these estimates with accretion histories may provide predictive constraints regarding the rate at which bodies accrete relatively pristine material and the rate at which basaltic crusts are produced.

13.8 Chondrules to Chondrites

Chondrules are inevitably accreted along with surrounding matrix material to form chondritic material, which is ultimately found on Earth as chondritic meteorites (Scott, 2007). Understand- ing how chondrites are produced may provide valuable constraints for chondrule formation models. Chondrules formed during the first few Myr after the formation of CAIs (Scott, 2007). Although the frequency of impacts on planetary embryos decreases rapidly with time after the nebular gas has been removed (Davison et al., 2013), large impacts certainly continued to occur after chondrule formation had apparently ceased. Johnson et al. (2015) argue that without nebular gas to damp out the relative velocity of chondrules, they will grind to dust and eventually be blown out of the solar system by radiation pressure (e.g., Johnson et al., 2012). This implies that these late impacts did not produce chondritic material. Johnson et al. (2015) suggest that for their model to produce enough chondritic material to be consistent with the abundance of chondrites in the meteorite record, this material must be preferentially accreted onto smaller bodies, which dominate the surface area of accreting bodies. Alternatively, this jetted material could produce a new generation of planetesimals. Preliminary work by Hasegawa et al. (2016) suggests that the efficiency of accretion of chondrules onto chondrite parent bodies depends on the strength of nebular magnetic fields and the size of the bodies that are accreting chondrules. They find that for some estimates of nebular magnetic field strength (Fu et al., 2014) planetesimals smaller than 1,000 km in diameter can efficiently accrete chondrule bearing chondritic crusts. This accreted material may be compacted and heated by subsequent impacts (Bland et al., 2014) or in some cases jetted to make a second generation of chondrules.

13.9 Conclusions

Our understanding of planet formation suggests that planets are built through a series of mergers and impacts. Accretion models imply that the impact velocities necessary to jet partially melted material are an expected outcome of the accretion process (Hasegawa et al., 2015; Johnson et al., 2015). Other models indicate this jetted material would be efficiently accreted onto smaller bodies (<1,000 km in diameter) (Hasegawa et al., 2016). Thus, in contrast to some other models for chondrule formation, the potentially chondrule-forming impact jetting process was very likely occurring during the millions of years when chondrules were formed. Thus, if chondrules formed via jetting, they can be considered a natural consequence of planet forma- tion. If impact jetting is not responsible for chondrule formation, and no other signs of this process are identified, this suggests that such collisions did not occur. By extension, this would imply that Moon-sized bodies never formed near the region where chondritic meteorite parent bodies formed. Such a conclusion is difficult to reconcile with our current conception of planetary accretion and is contradicted by the isotopically constrained formation timescale of 356 Brandon C. Johnson, et al.

Mars (Dauphas and Pourmand, 2011), both of which suggest the formation of embryos occurred readily in the early solar system. An impact origin for chondrules is also potentially more universal than other suggested formation mechanisms. CB chondrites are the youngest among known chondrite groups (Krot et al., 2007), and are thought to have formed in an impact generated gas-melt plume ~4.8 0.3 Ma after formation of CAIs (Krot et al., 2005; Bollard et al., 2015). The formation of CB chondrites requires high enough impact velocities to partially vaporize iron. The major differ- ence between the origin of most chondrules and those in CB chondrites may be the result of some process that leads to increased impact velocities such as the gas-driven migration of the giant planets (Johnson et al., 2016). This comparison demonstrates that an impact origin for chondrules can provide valuable information about the dynamic processes at work in the early solar system. Perhaps some of the differences between other chondrite types can be explained in a similar way. Clearly, more detailed modeling is needed to further test the impact-jetting model against observations. However, simple comparisons show that an impact origin for chondrules is broadly consistent with many chondrule constraints, including chondrule sizes, cooling rates, overall abundance in the meteorite record, lack of isotopic fractionation, and the volatile content of chondrule olivine. Much work remains to be done to better understand the process of impact jetting in relation to chondrule and chondrite formation. Better understanding of the details of chondrule formation by impact jetting may lead to deeper insights and connection of meteorite observations and the processes that were at work in the early solar system.

Acknowledgments

We thank D. Hezel for thorough review of this chapter. FJC would like to acknowledge funding from NASA’s Planetary Geology and Geophysics research program.

References

Alexander, C. M. O’D., and Ebel, D. S. (2012). Questions, questions: Can the contradictions between the petrologic, isotopic, thermodynamic, and astrophysical constraints on chon- drule formation be resolved? Meteoritics & Planetary Science 47, 1157–1175. Alexander, C. M. O’D., Grossman, J. N., Ebel, D. S., and Ciesla, F. J. (2008). The formation conditions of chondrules and chondrites. Science 320, 1617–1619. Alexander C. M. O’D., and Hewins, R. H. (2004). Mass fractionation of Fe and Ni isotopes in metal in Hammadah al Hamra 237 (abstract). Meteoritics & Planetary Science 39, A13. Asphaug, E., Jutzi, M., and Movshovitz, N. (2011). Chondrule formation during planetesimal accretion. Earth and Planetary Science Letters 308, 369–379. Becker, M., Hezel, D. C., Schulz, T., Elfers, B., and Münker, C. (2015). Formation timescales of CV chondrites from component specificHf–W systematics. Earth and Planetary Science Letters, 432, 472–482. Bland, P. A., Alard, O., Benedix, G. K., et al. (2005). Volatile fractionation in the early solar system and chondrule/matrix complementarity. Proceedings of the National Academy of Sciences 102, 13755–13760. Bland, P. A., Collins, G. S., Davison, T. M., et al. (2014). Pressure–temperature evolution of primordial solar system solids during impact-induced compaction. Nature Communica- tions 5, 5451. Formation of Chondrules by Planetesimal Collisions 357

Bollard, J., Connelly, J. N., and Bizzarro, M. (2015). Pb-Pb dating of individual chondrules from the CB achondrite Gujba: Assessment of the impact plume formation model. Meteoritics & Planetary Science 50, 1197–1216. Bollard, J., Connelly, J. N., Whitehouse, M. J., et al. (2017). Early formation of planetary building blocks inferred from Pb isotopic ages of chondrules. Science Advances 3, e1700407. Budde, G., Kleine, T., Kruijer, T. S., Burkhardt, C., and Metzler, K. (2016). Tungsten isotopic constraints on the age and origin of chondrules. Proceedings of the National Academy of Sciences 113, 2886–2891. Campbell, A. J., Humayun, M., and Weisberg, M. K. (2002). Siderophile element constraints on the formation of metal in the metal-rich chondrites Bencubbin, Weatherford, and Gujba. Geochimica et Cosmochimica Acta 66, 647–660. Carporzen, L., Weiss, B. P., Elkins-Tanton, L. T., et al. (2011). Magnetic evidence for a partially differentiated carbonaceous chondrite parent body. Proceedings of the National Academy of Sciences 108, 6386–6389. Cassen, P. (2001). Nebular thermal evolution and the properties of primitive planetary materials. Meteoritics & Planetary Science, 36, 671–700. Chambers, J. E. (2004). Planetary accretion in the inner Solar System. Earth and Planetary Science Letters 223, 241–252. Chiang, E., and Youdin, A. N. (2010). Forming planetesimals in solar and extrasolar nebulae. Annual Review of Earth and Planetary Sciences 38, 493–522. Connelly, J. N., Bizzarro, M., Krot, A. N., et al. (2012). The absolute chronology and thermal processing of solids in the solar protoplanetary disk. Science 338, 651–655. doi:10.1126/ science.1226919. Cuzzi, J. N., and Alexander, C. M. O’D. (2006). Chondrule formation in particle-rich nebular regions at least hundreds of kilometres across. Nature 441, 483–485. Dauphas, N., and Pourmand, A. (2011). Hf-W-Th evidence for rapid growth of Mars and its status as a planetary embryo. Nature 473, 489–492. Davis, A. M., Alexander, C. M., Nagahara, H., and Richter, F. M. (2005). Evaporation and condensation during CAI and chondrule formation. Chondrites and the Protoplanetary Disk 341, 432–455. Davison, T. M., O’Brien, D. P., Ciesla, F. J., and Collins, G. S. (2013). The early impact histories of meteorite parent bodies. Meteoritics & Planetary Science 48, 1894–1918. Desch, S. J., Morris, M. A., Connolly, H. C., and Boss, A. P. (2012). The importance of experiments: Constraints on chondrule formation models. Meteoritics & Planetary Science 47, 1139–1156. Dullemond, C. P., Stammler, S. M., and Johansen, A. (2014). Forming chondrules in impact splashes. I. Radiative cooling model. The Astrophysical Journal 794,91–12. Dullemond, C. P., Harsono, D., Stammler, S. M., and Johansen, A. (2016). Forming chondrules in impact splashes. II. Volatile retention. The Astrophysical Journal 832,91–19. Ebel, D. S., and Grossman, L. (2000). Condensation in dust-enriched systems. Geochimica et Cosmochimica Acta 64, 339–366. Elkins-Tanton, L. T., Weiss, B. P., and Zuber, M. T. (2011). Chondrites as samples of differentiated planetesimals. Earth and Planetary Science Letters 305,1–10. Evans, N. J. I., Dunham, M. M., Jørgensen, J. K., et al. (2009). The Spitzer c2d legacy results: Star-formation rates and efficiencies; evolution and lifetimes. The Astrophysical Journal Supplement 181, 321–350. Fedkin, A. V., and Grossman, L. (2013). Vapor saturation of sodium: Key to unlocking the origin of chondrules. Geochimica et Cosmochimica Acta 112, 226–250. Friedrich, J. M., Weisberg, M. K., Ebel, D. S., and Biltz, A. E. (2015). Chondrule size and related physical properties: A compilation and evaluation of current data across all meteorite groups. Chemie der Erde - Geochemistry 75, 419–443. 358 Brandon C. Johnson, et al.

Fu, R. R., and Elkins-Tanton, L. T. (2014). The fate of magmas in planetesimals and the retention of primitive chondritic crusts. Earth and Planetary Science Letters 390, 128–137. Fu, R. R., Weiss, B. P., Lima, E. A., et al. (2014). Solar nebula magnetic fields recorded in the Semarkona meteorite. Science 346, 1089–1092. Glass, B. P., and Simonson, B. M. (2012). Distal impact ejecta layers: spherules and more. Elements 8,43–48. Hasegawa, Y., Turner, N. J., Masiero, J., et al. (2016). Forming chondrites in a solar nebula with magnetically induced turbulence. The Astrophysical Journal Letters 820, L12. Hasegawa, Y., Wakita, S., Matsumoto, Y., and Oshino, S. (2015). Chondrule formation via impact jetting triggered by planetary accretion. The Astrophysical Journal 816,1–14. Hezel, D. C., and Palme, H. (2010). The chemical relationship between chondrules and matrix and the chondrule matrix complementarity. Earth and Planetary Science Letters 294,85–93. Hood, L. L., Ciesla, F. J., Artemieva, N. A., Marzari, F., and Weidenschilling, S. J. (2009). Nebular shock waves generated by planetesimals passing through Jovian resonances: Possible sites for chondrule formation. Meteoritics & Planetary Science 44, 327–342. Johansen, A., Blum, J., Tanaka, H., et al. (2014). The multifaceted planetesimal formation process. In H. Beuther, R. S. Klessen, C. P. Dullemond, and T. Henning (Eds.), Protostars and Planets VI, 547–570. Tucson, AZ: University of Arizona Press. Johnson, B. C., Bowling, T. J., and Melosh, H. J. (2014). Jetting during vertical impacts of spherical projectiles. Icarus 238,13–22. Johnson, B. C., Lisse, C. M., Chen, C. H., et al. (2012). A self-consistent model of the circumstellar debris created by a giant hypervelocity impact in the HD 172555 system. The Astrophysical Journal 761, 45. Johnson, B. C., and Melosh, H. J. (2014). Formation of melt droplets, melt fragments, and accretionary impact lapilli during a hypervelocity impact. Icarus 228, 347–363. Johnson, B. C., and Melosh, H. J. (2012a). Formation of spherules in impact produced vapor plumes. Icarus 217, 416–430. Johnson, B. C., and Melosh, H. J. (2012b). Impact spherules as a record of an ancient heavy bombardment of Earth. Nature 485,75–77. Johnson, B. C., Minton, D. A., Melosh, H. J., and Zuber, M. T. (2015). Impact jetting as the origin of chondrules. Nature 517, 339–341. Johnson, B. C., Walsh, K. J., Minton, D. A., Krot, A. N., and Levison, H. F. (2016). Timing of the formation and migration of giant planets as constrained by CB chondrites. Science Advances 2, e1601658. Jutzi, M., Asphaug, E., Gillet, P., Barrat, J. A., and Benz, W. (2013). The structure of the asteroid 4 Vesta as revealed by models of planet-scale collisions. Nature 494, 207–210. Kieffer, S. W. (1975). Droplet Chondrules. Science 189, 333–340. Krot, A. N., Amelin, Y., Cassen, P., and Meibom, A. (2005). Young chondrules in CB chondrites from a giant impact in the early Solar System. Nature 436, 989–992. Krot, A. N., Keil, K., Scott, E. R. D., Goodrich, C. A. and Weisberg, M. K. (2007). Classifica- tion of meteorites and their genetic relationships. In H. Holland and K. Turekian (Eds.), Treatise on Geochemistry (Second Edition), 1, 1–63. Oxford, UK: Elsevier. Kruijer, T. S., Touboul, M., Fischer-Godde, M., et al. (2014). Protracted core formation and rapid accretion of protoplanets. Science 344, 1150–1154. Kruijer, T. S., Burkhardt, C., Budde, G., and Kleine, T. (2017). Age of Jupiter inferred from the distinct genetics and formation times of meteorites. Proceedings of the National Academy of Sciences 312, 201704461–5. Kurosawa, K., Nagaoka, Y., Senshu, H., et al. (2015). Dynamics of hypervelocity jetting during oblique impacts of spherical projectiles investigated via ultrafast imaging. Journal of Geophysical Research: Planets 120, 1237–1251. Levison, H. F., Duncan, M. J., and Thommes, E. (2012). A Lagrangian Integrator for Planetary Accretion and Dynamics (LIPAD). The Astronomical Journal 144, 119. Formation of Chondrules by Planetesimal Collisions 359

Levison, H. F., Kretke, K. A., and Duncan, M. J. (2015). Growing the gas-giant planets by the gradual accumulation of pebbles. Nature 524, 322–324. Lofgren, G. (1989). Dynamic cyrstallization of chondrule melts of porphyritic olivine com- position: Textures experimental and natural. Geochimica et Cosmochimica Acta 53, 461–470. Minton, D. A., and Malhotra, R. (2010). Dynamical erosion of the asteroid belt and implications for large impacts in the inner Solar System. Icarus 207, 744–757. O’Brien, D. P., Morbidelli, A., and Bottke, W. F. (2007). The primordial excitation and clearing of the asteroid belt—Revisited. Icarus 191, 434–452. Palme, H., Lodders, K., and Jones, A. (2014). Solar system abundances of the elements. In H. Holland and K. Turekian (Eds.), Treatise on Geochemistry (Second Edition), 2, 15–36. Oxford, UK: Elsevier. Pierazzo, E., Vickery, A. M., and Melosh, H. J. (1997). A reevaluation of impact melt production. Icarus 127, 408–423. Richter, F. M., Huss, G. R., and Mendybaev, R. A. (2014). Iron and nickel isotopic fractionation across metal grains from three CBb meteorites. Lunar Planet. Sci. Conf. XLV, #1346. Roszjar, J., Whitehouse, M. J., Srinivasan, G., et al. (2016). Prolonged magmatism on 4 Vesta inferred from Hf–W analyses of eucrite zircon. Earth and Planetary Science Letters 452, 216–226. Ruden, S. P., and Pollack, J. B. (1991). The dynamical evolution of the protosolar nebula. The Astrophysical Journal, 375, 740–760. Sanders, I. S., and Scott, E. R. D. (2012). The origin of chondrules and chondrites: Debris from low-velocity impacts between molten planetesimals? Meteoritics & Planetary Science 47, 2170–2192. Schmitz, B., Yin, Q. Z., Sanborn, M. E., et al. (2016). A new type of solar-system material recovered from Ordovician marine limestone. Nature Communications 7, ncomms11851. Schulte, P., Alegret, L., Arenillas, I., et al. (2010). The Chicxulub asteroid impact and mass extinction at the Cretaceous-Paleogene boundary. Science 327, 1214–1218. Scott, E. R. D. (2007). Chondrites and the protoplanetary disk. Annual Review of Earth and Planetary Sciences 35, 577–620. Scott, E. R. D., and Krot, A. N. (2003). Chondrites and their components. In H. Holland and K. Turekian (Eds.), Treatise on Geochemistry (Second Edition), 1, 65 –137. Oxford, UK: Elsevier. Sorby, H. C. (1877). On the structure and origin of meteorites. Nature 15, 495–498. Taylor, G. J., Scott, E. R. D., and Keil, K. (1982). Cosmic setting for chondrule formation. Abstracts of Papers Presented to the Conference on Chrondrules and Their Origins, 493, 58. Houston, TX: Lunar and Planetary Institute. Urey, H. C. (1952). Chemical fractionation in the meteorites and the abundance of the elements. Geochimica et Cosmochimica Acta 2, 269–282. Van Kooten, E. M. M. E., Wielandt, D., Schiller, M., et al. (2016). Isotopic evidence for primordial molecular cloud material in metal-rich carbonaceous chondrites. Proceedings of the National Academy of Sciences 113, 2011–2016. Vickery, A. M. (1993). The theory of jetting: Application to the origin of tektites. Icarus 105, 441–453. Villeneuve, J., Libourel, G., and Soulié, C. (2015). Relationships between type I and type II chondrules: Implications on chondrule formation processes. Geochimica et Cosmochimica Acta 160, 277–305. Walsh, J. M., Shreffler, R. G., and Willig, F. J. (1953). Limiting conditions for jet formation in high velocity collisions. Journal of Applied Physics 24, 349. Walsh, K. J., Morbidelli, A., Raymond, S. N., and O’Brien, D. P. (2011). A low mass for Mars from Jupiter’s early gas-driven migration. Nature 475, 206–209. 360 Brandon C. Johnson, et al.

Warren, P. H. (2011). Stable-isotopic anomalies and the accretionary assemblage of the Earth and Mars: A subordinate role for carbonaceous chondrites. Earth and Planetary Science Letters 311,93–100. Weidenschilling, S. J. (2011). Initial sizes of planetesimals and accretion of the asteroids. Icarus 214, 671–684. Weidenschilling, S. J., Marzari, F., and Hood, L. L. (1998). The origin of chondrules at jovian resonances. Science 279, 681–684. Weisberg, M. K., Prinz, M., Clayton, R. N., et al. (2001). A new metal-rich chondrite grouplet. Meteoritics & Planetary Science 36, 401–418. 14 Making Chondrules by Splashing Molten Planetesimals The Dirty Impact Plume Model ian s. sanders and edward r. d. scott

Abstract

The antiquity of iron meteorites and the inferred early intense heating by the decay of 26Al suggest that many planetesimals were molten beneath a thin insulating cap at the same time as chondrules were being made. As those planetesimals were colliding and merging, it seems inevitable that impact plumes of droplets from their liquid interiors would have been launched into space and cooled to form chondrules. We call the process splashing; it is quite distinct from making droplets by jetting during hypervelocity impacts. Evidence both for the existence of molten planetesimals, and for the cooling of chondrules within a plume setting, is strong and growing. Detailed petrographic and isotopic features of chondrules, particularly in carbon- aceous chondrites (that probably formed beyond the orbit of Jupiter), suggest that the chondrule plume would have been ‘dirty’ and the otherwise uniform droplets would have been contamin- ated with earlier-formed dust and larger grains from a variety of sources. The contamination possibly accounts for relict grains, for the spread of oxygen isotopes along the primitive chondrule mineral (PCM) line in carbonaceous chondrites, and for the newly recognized nucleosynthetic isotopic complementarity between chondrules and matrix in Allende.

14.1 Introduction and History

The splashing hypothesis for chondrule formation posits that most chondrules are frozen droplets of spray that cooled within impact plumes from the collision of planetesimals that were already molten beneath a thin, insulating carapace. It differs from the proposal that chondrules are formed by jetting during hypervelocity impacts between larger solid bodies (Johnson et al., 2015 and Chapter 13). With splashing, the energy for melting is radioactivity; it comes almost entirely from the decay of short-lived 26Al and collisions are ‘messy and slow’ (Asphaug et al., 2011). With jetting, the heat for melting is supplied by the huge kinetic energy of impact. We reviewed the splashing hypothesis in detail about five years ago (Sanders and Scott, 2012), so we focus our attention here on new evidence that has come to light in the past five years. Before doing that we give a summary of the historical background to splashing, and the evidence that supports it. Splashing was proposed by Herbert Zook (Zook, 1980, 1981) and also, at about the same time, by Helmut Wänke and co-workers (Wänke et al., 1981; Wänke et al., 1984). However, it

361 362 Ian S. Sanders and Edward R. D. Scott did not catch on. One possible reason for this is that so-called relict grains had just been reported in chondrules (Nagahara, 1981; Rambaldi, 1981). These are unequilibrated xenocrysts of olivine or orthopyroxene in porphyritic chondrules. They were interpreted as unmelted residual grains in the context of the conventional view that chondrules began as dust balls – free-floating clumps of dusty precursor fragments – which were melted by an intense external heat source such as nebular shock waves, lightning, or energetic flares. Another reason for the lack of enthusiasm for splashing was that the features seen in chondrules and chondrites were deemed to be at odds with the then-prevailing ideas on the products of collisions between planetesimals, whether molten or solid (Taylor et al., 1983). There was also a strong sense that, because chondrites have primitive chemistry, and because they contain calcium-aluminium-rich inclu- sions (CAIs) which are the oldest solar system solids, they must be samples of the first generation of planetesimals. In this case, the chondrules within them must predate the existence of planetesimals. So, the splashing model was relegated to a quiet backwater, far from mainstream thinking, and the conventional view that chondrules were produced by the ‘flash melting’ of clumps of nebular dust went largely unchallenged throughout the 1980s and 1990s. The conventional view remains the most popular view of chondrule formation to this day despite appeals by influential scientists like Jerry Wasserburg and Gunter Lugmair to take a closer look at the splashing model (Chen et al., 1998; LaTourrette and Wasserburg, 1998; Lugmair and Shukolyukov, 2001) Seeds of doubt regarding the conventional view were sown in our minds by a report on a 2-mm–sized fragment of microgabbro in Parnallee (LL3) by Kennedy et al. (1992).* This fragment shows that Parnallee’s chondritic parent body was not assembled at the very beginning of the solar system, as conventional thinking would imply. Instead, it shows that accretion happened later, after the planetesimal from which the microgabbro was derived had accreted, become heated, acquired an igneous crust, and been fragmented. The first 26Al–26Mg age of a chondrule was published by Hutcheon and Hutchison (1989). As more ages followed during the 1990s, it became apparent that most chondrules are probably about 2 Myr younger than CAIs, and that the parent bodies of chondrites (such as Parnallee) must therefore have accreted even later than this. The late accretion of chondrites is consistent with the widely accepted idea that 26Al (half- life ~ 0.72 Myr) caused planetesimal melting. Assuming that at the start of the solar system 26Al /27Al was everywhere close to 5 Â 10–5 (the so-called canonical value), thermal modeling (e.g., Hevey and Sanders, 2006) suggests that planetesimals could only have become totally melted and yielded differentiated products if they accreted within the first 1–1.5 Myr. The corollary is that the planetesimals that supply chondritic meteorites, since they did not melt, must have accreted after about 2–2.5 Myr, by which time 26Al had largely decayed away. This view was reinforced by the results of 182Wdeficit dating of iron meteorites, which indicate that the parent bodies of most iron meteorites had accreted, melted, and segregated well before 2 Myr (e.g., Kleine et al., 2005; Kruijer et al., 2017). The effect of 26Al heating on early planetesimal interiors is illustrated in Figure 14.1. It shows whether (and when) complete melting, partial melting, or no melting occurs as a function of the time of accretion. It relates to planetesimals of any radius at a depth greater than roughly 10 km. It shows that the interior of a body accreting from dry dust right at the beginning (A) would have had 6.4 kJ of radioactive energy in each gram, and would have become fully molten The Dirty Impact Plume Model 363

8

6.4 6 1.6 kJ/g is the energy needed to melt completely, from cold, the insulated interior

Al in ‘dry’ Al in of a primitive planetesimal. 26 4 About 10 km of insulation 3.2 is effective. primitive dust (kJ/g) primitive

Energy in 2 1.6 0.8

0

solid D 0.8

250 K partial

Al at C melt 26 1.6 1425 K

solidus 1850 K instantaneous accretionliquidus

Energy in 3.2 B

time of accretion (kJ/g) convecting magma ocean

A 6.4 012 34 5 Time (Myr after CAIs)

Figure 14.1 The upper diagram is an estimate of the exponential decline of 26Al in kilojoules of stored radioactive energy in one gram of dry solar system dust with ~1.2 wt% Al and canonical 26Al/27Al during the first five million years. The lower diagram shows how this declining energy probably affected the insulated interiors (deeper than about 10 km) of planetesimals that accreted between 0 and 2.5 Myr after the beginning. Four examples, A to D, with accretion one half-life apart, are described in the text. after just 300 kyr. In case (A) continued, intense heating is presumed to have led to strong thermal convection in the liquid, with consequent reduction of the thickness of the insulating cover to perhaps 1 km or less, by melting upwards from below. A body accreting after one half- life (B), starting with 3.2 kJ/g, would have taken another 0.75 Myr to heat up and melt completely in its insulated interior. A body accreting after two half-lives (C), with 1.6 kJ/g, would have had just enough 26Al for its interior to melt totally, but it would not have done so until about 5 Myr after the beginning, and the insulating carapace would probably have remained thick and rigid. A body accreting after three half-lives, at about 2.2 Myr from the beginning (D), with 0.8 kJ/g, would not even have begun to melt. Most chondrites appear to come from parent bodies that followed a path like (D). Primitive achondrites (like ureilites) might be products of path (C), and fully differentiated meteorites presumably come from parent bodies that accreted early, and took paths like (A) or (B). The chronology of meteorites and their components, recently reviewed by Kleine and Wadhwa (2017), is remarkably consistent with the timing of events predicted in Figure 14.1. What is the relevance of Figure 14.1 for chondrule formation? The answer is that it paves the way for understanding splashing. In cases A and B, planetesimal collisions after the time of melting would potentially have led to splashing. One can imagine that from about 300 kyr onwards many planetesimals in the protoplanetary disk were substantially, if not completely, molten beneath a thin insulating cap. As those planetesimals collided and merged to make larger 364 Ian S. Sanders and Edward R. D. Scott bodies, in the inevitable early stages of planet building, it is easy to imagine how some of their liquid contents would have been launched as huge plumes of molten droplets that cooled to become chondrules. However, the potential for splashing would, conceivably, have diminished over the next two or three million years because, with planetesimal cooling, the insulating cap would steadily have become thicker and stronger, as depicted in Figure 2 of Sanders and Scott (2012). Another important point shown by Figure 14.1 is that chondrules formed (by any mechanism) in the first 1.5 Myr would probably be scarce, as is observed, because they would still have been very radioactive and, following fairly prompt reaccretion, they would mostly have been destroyed by melting inside their new host planetesimals. The review by Sanders and Scott (2012), and earlier accounts of the splashing model by Sanders (1996), Sanders and Taylor (2005) and Asphaug et al. (2011) show how splashing can be reconciled with many features of chondrules and chondrites. The model has the potential to make chondrules with ‘primitive’ silicate chemistry, at near-liquidus tempera- tures, and in very large numbers, at just the right time. Also, the heat source for melting is well understood. Perhaps the most compelling case for splashing is the discovery that Na is present in chondrule olivine at a level which implies that cooling and crystallization occurred when chondrules were in large, densely-populated swarms (Alexander et al., 2008). Hewins et al., (2012) suggested that a plausible setting would be an impact plume. Grossman et al., (2012) reached the same conclusion, but based on the argument that chondrules must have been densely spaced in an oxidizing vapor to account for the presence of iron in silicates. It had earlier been recognized that a plume setting for chondrule cooling was consistent with the incidence of compound chondrules, and with the retarded cooling rate inferred for chondrules, over hours rather than seconds. Also, the resurgence of the concept of ‘hot accretion’ based on chondrules that indent one another (Metzler, 2012) is entirely consistent with rapid reaccre- tion of chondrules to a body within a hot plume, before they had time to solidify. Sanders and Scott (2012) pointed out that the plume they envisage may have contained, in addition to molten droplets, a great deal of dust, with variously sized grains. Here, to emphasize the importance of dust in the plume, we suggest that the splashing model may alternatively be named the dirty plume model. The dust could have several possible sources. It could have come from the outermost, unconsolidated carapace (the regolith) on one or both of the colliding pair of planetesimals. It could have come from an unmelted planetesimal colliding with a molten one, where the lack of melting might have been due to late accretion or to an abundance of accreted ice. Another possibility is that one of the colliding planetesimals was still actively accreting and was enveloped in its own, inwardly spiraling stream of dense nebular dust. In all these cases the dust would largely end up as matrix, or as fine-grained rims on chondrules, but some of the larger grains, possibly derived from the fragmentation of older chondrules or older planetesimals, or from early condensates in the nebula, could have been captured in flight by the molten droplets and incompletely dissolved up, giving rise to so-called relict grains (see Chapter 2). Sanders and Scott (2012) emphasized that all chondrules, even those produced by splashing, were once freely floating objects, and therefore products of a nebular setting. A plume is simply a local, very dense, and ephemeral part of the nebula, and the fact that its contents come from planetary bodies is incidental. In this light, splashing should be seen as simply another possible way of the formation of molten droplets in the nebula. The Dirty Impact Plume Model 365

Sanders and Scott (2012) invoke total, or near total, melting to account for the primitive (olivine-rich) composition of most chondrules (see their Figures 2 and 3). However, the likelihood of magma oceans with primitive compositions is deemed improbable by some researchers, who believe that early-formed partial melt (basaltic magma) would have made its way rapidly upwards, carrying with it the 26Al heat source, and thus forestalling further melting in the deeper interior. In support of primitive magma oceans, Sanders and Scott (2012) note that the very high magmatic temperature implied by sulfur-poor iron meteorites of the IVB group would entail total, or near total, melting of the silicate mantle. We add here that even if basaltic liquid did rise, it would probably have formed an extensive (planetesimal-wide?) shallow intrusion, as described by Wilson and Keil (2012). In this case, any late accreting debris, or any small merging planetesimals, would probably have become mingled and digested in the radioactive, heat-producing basaltic liquid, and rendered it more magnesian and primitive (olivine-rich) and appropriate for chondrule compositions.

14.2 General Developments Since 2012 That Bear on the Splashing Model

Since our 2012 review, an important development has been the growing volume of data that supports the isotopic dichotomy of meteorites. Every meteorite, whether primitive or differenti- ated, is now widely believed to come from one of just two giant reservoirs in the disk, labeled carbonaceous and noncarbonaceous, and thought to have existed, respectively, beyond and inside the orbit of Jupiter. While Warren (2011) brought the dichotomy to the attention of a wide audience, the unique isotopic properties of carbonaceous chondrites were earlier recog- nized by Trinquier et al., (2007), and by Niemeyer (1985). Most recently, the dichotomy has been extended to include many iron meteorites (Kruijer et al., 2017), which shows that the two reservoirs remained isolated almost from the start of the solar system (when iron meteorite parent bodies accreted) until about four million years later (when the youngest chondrites accreted). The significance of isotopic dichotomy for chondrule formation is that the process, splashing or otherwise, would have been operating both within and beyond the orbit of Jupiter. Importantly for splashing, since large-scale planetesimal melting and production of iron cores took place in both reservoirs, then the formation of chondrules by splashing is plausible in both reservoirs. Differences in oxygen isotopic compositions between chondrules in carbonaceous chondrites, and those in ordinary and R-chondrites are widely recognized (e.g., Kita et al., 2010, 2015). Recently titanium isotopes in chondrules from carbonaceous chondrites were found to differ from those in non-carbonaceous chondrites (Gerber et al., 2017), implying that, as one might expect, chondrules were probably formed in the same reservoir as their host meteorite. An ongoing debate since 2012 has concerned the homogeneity, or otherwise, of 26Al/27Al in the disk. Some have argued that in parts of the disk where chondrules were formed 26Al/27Al – was variable and much lower than the canonical value of about 5 Â 10 5 (Larsen et al., 2011; Schiller et al., 2015). The inferred subcanonical 26Al/27Al would barely have provided enough energy to cause planetesimal melting by 3 Myr, so the splashing model, which requires abundant molten planetesimal interiors much earlier than this, would be ruled out. However, Budde et al. (2018) published new Hf-W ages of CR chondrules, and combined them with 366 Ian S. Sanders and Edward R. D. Scott existing Hf-W and Al–Mg data to show that the initial 26Al/27Al and the initial 182Hf/180Hf of CV CAIs, CV chondrules, CR chondrules, and angrites are all remarkably concordant, and consistent with the disk having initial canonical 26Al/27Al (Æ 10%) throughout. These results suggest that the Pb–Pb ages of angrites are discordant. Clearly, early solar system chronology has some outstanding problems to be resolved. Meanwhile, with the jury still out, we favor the results of Budde et al. (2017) and assume initially canonical 26Al/27Al where the first planet- esimals accreted. New general support for the idea that chondrules cooled in a plume comes from both theoretical and experimental studies. Firstly, thermal modeling of an expanding plume by Dullemond et al. (2014) shows that chondrules would remain more or less at the same high temperature for as long as they were in an optically thick part of the plume, and they would cool very rapidly once the plume around them became optically thin. Dullemond et al. (2016) extended their thermal model to include the behavior of volatile elements in an expanding plume. They showed that Na, for example, would partly evaporate, then later recondense, in a way that is consistent with the interpretation of Na in Semarkona chondrules by Hewins et al. (2012). Secondly, late rapid cooling of chondrules (1,000–8,000 K per hour), in line with the prediction of Dullemond et al. (2014) has been inferred by Soulié et al. (2017) based on experimentally observed rapid dissolution rates of olivine in chondrule melts and the preser- vation of relict olivine grains. Villeneuve et al. (2015) similarly invoke late quenching to account for the preservation of glassy mesostases, and envisage chondrules cooling in an impact plume. Theoretical, petrographic, and experimental studies have also been directed to understand- ing the interiors of planetesimals as they melt, and as they cool. Faure et al. (2012, 2017) infer from glass inclusions inside phenocrysts of magnesian olivine in chondrules that the pheno- crysts grew within a magma ocean rather than within cooling chondrule droplets. This is in line with our own inference that certain large or oscillatory zoned phenocrysts in chondrules grew in a magma ocean. Also, Florentin et al. (2017) favor the crystallization of magnesian olivine phenocrysts in magma oceans based on high levels of sodium in glass inclusions. With regard to metal segregation, numerical modeling by Lichtenberg et al. (2016, 2018) suggests that metal drained downwards rapidly as the planetesimals heated up and melted. Since some chondrules contain metal, these authors hint that splashing took place at a critical stage of partial melting, before metal had drained away. Their results, however, are not in agreement with those of Kiefer and Mittlefehldt (2017) who suggest that metal would segregate much more slowly. A minor correction to our 2012 review concerns a correlation between the Al–Mg ages of chondrules and their Mg/Si ratios (Tachibana et al., 2003). We linked it to protracted crystal- liquid fractionation in magma oceans. Our discussion is now immaterial because further data have shown that the putative correlation does not exist (Villeneuve et al., 2012).

14.3 Evidence Since 2012 That is Claimed to Refute the Splashing Model

Connelly et al. (2012) published precise Pb–Pb ages for chondrules, showing some to be as old as CAIs. They reasoned that these old chondrules could not be products of impact splashing The Dirty Impact Plume Model 367 because there would not have been enough time for planetesimal melting. Their claim, however, is not strictly valid. The time needed for total planetesimal melting at the very beginning was only about 300 kyr (Figure 14.1) assuming cold accretion. Even less time would have been needed if nebular temperatures were higher and the first crop of planetesimals accreted hot. When errors in the ages of chondrules and the ages of CAIs are taken into account, there is enough latitude for planetesimal melting to have occurred before the earliest chondrules were produced. An ongoing criticism of splashing concerns so-called chemical complementarity in carbon- aceous chondrites, where neither chondrules nor matrix have primitive, solar compositions, but when combined together the bulk composition is solar. Palme et al. (2015) reiterate their earlier arguments that chondrule formation by splashing is inconsistent with complementarity. They infer that complementarity in a given meteorite points to a single reservoir of primitive (chondritic) dusty material, and that chondrules with a nonprimitive composition are produced from some of the dusty material (presumably by ‘flash melting’), then promptly reunited with the complementary unmelted fraction of the dust (with a different composition) during accre- tion, thus preserving the original primitive composition of the reservoir in the bulk meteorite. Our concerns with this interpretation are firstly that we have difficulty in imagining a flash melting process that will selectively melt only part of a dusty reservoir while the unmelted residue remains close at hand then quickly remixes with the melted fraction, and secondly that the scenario seems at odds with the presence in a single carbonaceous chondrite of chondrules that have differing ages or show isotopic diversity (e.g., Olsen et al., 2016; Gerber et al., 2017). We suggest that since each carbonaceous chondrite group has its own distinctive features, then perhaps the reservoir for each parent body comprised a narrow annulus of orbiting material that included matrix dust and splash-produced chondrules from several impacts within the reservoir, with significant recycling of material between evolving planetesimals but no overall change to the bulk chemistry of the annular reservoir. On a separate note, Zanda et al. (Chapter 5) question the status of complementarity, showing that the chemical analysis of matrix by electron microp- robe can lead to erroneous results. Nagashima et al. (2015) and Schrader et al. (2017) argue that relict grains, and the diverse oxygen isotopic compositions of chondrules in some meteorites, are inconsistent with the splashing model. These authors envisage splashing as capable only of producing huge numbers of isotopically identical droplets. However, the dirty plume scenario we have described means that the droplets can become very variable, depending on the nature and extent of their contamination by trapped grains. The abundance of 16O-rich grains in the matrix of Kakangari (Nagashima et al., 2015) is evidence that the ambient dust in the disk contained many such grains. In this light, we wonder whether the spread of oxygen isotopic compositions of pristine carbonaceous chondrite chondrules along the primitive chondrule mineral (PCM) line (Chapter 8; Figure 14.2) results from mixing between chondrule droplets, with compositions at the top end of the range, and fragments of CAIs and amoeboid olivine aggregates (AOAs) in the dust whose oxygen isotopic compositions started a long way down the slope-1 line, close to the position inferred for the Sun. We envisage appreciable dissolution of trapped grains in the molten droplets prior to cooling, with only remnants surviving as relict grains. Also on the subject of oxygen isotopes in chondrules, Bischoff et al. (2017) describe a remarkable multistage compound chondrule from Allende in which the individual chondrule 368 Ian S. Sanders and Edward R. D. Scott

10

PCM 5 Bulk Earth

0(‰) 0 17 CC TFL chondrules –5 to AOA and CAI fragments

–10 –10 –5 0 5 10 180(‰)

Figure 14.2 Oxygen three-isotope diagram modified from Tenner et al. (2015) showing typical distribution of pristine carbonaceous chondrite chondrules along the so-called PCM slope-1 line. We suggest the linear trend is a result of contamination of original chondrule droplets at the top end of the distribution by AOA and CAI fragments in a dusty plume. caps have a diversity of compositions. These authors claim that their observation rules out the splashing model, believing that splashing would create a suite of identical droplets, which would make all the chondrule caps the same. Again, the possibility that chondrule compositions can change by contamination within a dirty plume counters this argument. In the Allende multicompound chondrule, incidentally, the chondrule cap compositions fall along the ‘heavy’ end of the CCAM line, raising the possibility that the isotopic spread may be due to partial isotopic exchange with ‘heavy’ ambient gas, possibly polluted by evaporation of ‘heavy’ nebular ice. We suggest that this evaporation and contamination could have occurred in a plume environment though it might also have resulted from alteration within the parent body. Finally, perhaps the most serious challenge to the splashing model since 2012 comes from a recent and remarkable discovery of nucleosynthetic isotopic complementarity between chon- drules and matrix in Allende (CV3). Budde et al. (2016a) reported that ε183W in Allende chondrules is positive, with values between +1.5 and +2.5 in six different chondrule separates, and that in three separates of the matrix ε183W is negative with values between –0.5 and –1.5. In bulk Allende ε183W is zero, the same as ε183W for other bulk meteorites, both primitive and differentiated, and for the Earth. They interpreted the complementarity as resulting from an excess of s-process tungsten in the matrix, and a corresponding deficit in the chondrules. They claimed that these observations rule out the splashing model because molten planetesimals, like bulk meteorites, must have had ε183W = 0, so droplets splashed from the molten planetesimals must also have had ε183W = 0. Moreover, chondrules formed by splashing, they suggest, would be isotopically identical, yet the Allende chondrules have substantial variation in ε183W from one separate to another. The Dirty Impact Plume Model 369

They proposed that in Allende the chondrules and matrix were formed together from a uniform primitive dusty reservoir (ε183W = 0), which means that the chondrule-forming mechanism involved the partitioning of nuclides from the s-process carrier phase preferentially into the matrix, leaving a deficit in the dust that melted to become chondrules. They also inferred that, in order to retain a value of zero in the mixture of chondrules and matrix, the chondrules and matrix must have stayed in close proximity and accreted together very soon after their joint formation. Budde et al. (2016b), using the same Allende chondrule and matrix separates as were used to measure tungsten isotopes, discovered similar complementarity in the isotopes of molybdenum, with an excess of s-process Mo in the matrix, and a deficit in the chondrules. They drew the same conclusions as for tungsten: (1) that chondrule formation by splashing is ruled out and (2) the chondrule forming process must start with a uniform primitive reservoir and lead to partitioning of nuclides from an s-process presolar carrier phase preferentially into the matrix, leaving the chondrules with a deficit in these nuclides. We find these new observations intriguing, and admit that they challenge the splashing model, but we are not convinced that they rule it out. In fact, we find it hard to envisage how any chondrule-forming mechanism, splashing or otherwise, could yield molten droplets whose W and Mo isotopic abundances deviate from zero. Budde et al. (2016b) noted that chondrules with less metal have larger deficits of s-process W and Mo, and invoked metal-silicate fractionation during chondrule formation. The physical mechanism they invoke is not clear to us, but we assume they envisage W and Mo from the s-process carrier(s) being partitioned into metal during melting, and that liquid metal physically migrates, to varying degrees, from the chondrules into the dust (which becomes the matrix) before it can equilibrate isotopically with the silicate liquid. We speculate on a possible alternative mechanism. We suggest that newly made molten chondrules with no initial W and Mo anomalies were immediately mixed with ambient dust in such a way that s-process W and Mo in the dust were scavenged from their presolar carrier(s) by molten metal (hot condensates?) and mostly ended up in the matrix, while other dust grains (silicates? with, on average, a slight, complementary depletion of s-process W and Mo) were engulfed in the droplets and digested. This alternative mechanism could, we suppose, have taken place in various cosmic settings. It is illustrated graphically for tungsten in Figure 14.3 in the context of a dirty impact plume. We speculate, as we speculated about chemical complementarity, that the grains in the plume would have included old, recycled chondrules as well as pristine presolar dust, and that the bulk rock values of zero for ε183W and ε92Mo can be the result of mixing and interaction of chondrules and dust in several separate impact and reaccretion cycles in a closed reservoir occupying a narrow annulus in the disk; they do not need to imply a single episode of chondrule production followed quickly by accretion in a small closed reservoir.

14.4 Conclusions

The possibility that chondrules were produced in impact plumes created during low-speed mergers between planetesimals that were already molten (beneath a thin insulating carapace) 370 Ian S. Sanders and Edward R. D. Scott

Before plume mixing

average W in average W in average W in s-enriched grains all grains s-depleted grains

newly splashed droplet

?-10 0 +10? ε183W

After plume mixing

average W in average W in residual purged contaminated dust chondrule

-4 -2 0 +2 +4 ε183W

Figure 14.3 Basis for a possible reconciliation between the splashing model and the nucleosynthetic complementarity of tungsten between chondrules and matrix in Allende, CV3. Before mixing (top diagram) neither the newly splashed droplets nor the dust in the dirty plume has a tungsten isotopic anomaly, but the dust contains unprocessed presolar grains (solid rectangles) with an excess of s-process W (negative anomalies) and other grains (open triangles) which, to compensate, must have, on average, a positive W anomaly. The values of the anomalies in ε183W shown (approximately +10 and À10) are chosen arbitrarily for illustrative purposes only. In reality, it is likely that the amount of s-process carrier in the dust would be miniscule and the anomaly in its s-process W enormously negative. After mixing (lower diagram, which is what is observed in Allende), any shift to the right in chondrule composition following digestion of dust grains (silicate?) with positive ε183W would result in a complementary shift to the left in the residual dust that would later become matrix.

remains viable. Perhaps the strongest argument in its favor is that no other nebular setting seems capable of generating the very high concentrations of chondrule droplets inferred from the levels of Na in chondrule phenocrysts and the presence of iron in chondrule silicates. A collision plume environment for chondrule formation is supported by recent numerical modeling and also by experimental and petrographic evidence that cooling was initially slow and then very fast (consistent with an expanding plume changing from being optically thick to optically thin). However, much more work in the difficult area of plume simulation is needed if splashing is to become more widely accepted. While turbulently-mixed, planetesimal-wide spheres of magma are postulated as the source of chondrules, it is not yet clear that such huge magma oceans ever existed. It is possible that early melt segregation forestalled total meltdown. However, the very high temperatures needed for sulfur-poor iron liquids would have led to total, or near-total, silicate melting. Some glassy inclusions in phenocrysts in chondrules suggest growth of the grains within a large body of magma, and not within cooling droplets. The Dirty Impact Plume Model 371

Recent observations of isotopic dichotomy in iron meteorites show that early planetesimal meltdown happened not only in the inner solar system but also beyond the orbit of Jupiter in a huge carbonaceous meteorite reservoir. This implies that splashing could have occurred in the carbonaceous chondrite reservoir. Here the abundant dust may have contaminated molten chondrule droplets in a dirty plume setting, so accounting for the oxygen isotopic spread down the PCM line, the presence of so-called relict grains in carbonaceous chondrite chondrules, and also (perhaps) accounting for the nucleosynthetic isotopic complementarity in W and Mo seen in Allende. The splashing model might be renamed the “dirty plume” model. With the possibility of mixing between molten droplets and presolar dust, larger mineral grains, recycled chondrules and refractory inclusions, we suggest that splashing offers a scenario which encompasses more dimensions of early solar system evolution than most other chondrule-forming mechanisms. As a final note, we do not claim that splashing is the source of all chondrules, but we suspect that most chondrules were formed in this way, if only because of the likely prevalence of colliding molten bodies in the disk at the same time as chondrules were being produced. We believe that molten droplets were also formed in the disk by other mechanisms, including high- velocity impacts, igneous fire fountaining, and frictional melting during rapid deceleration in atmospheres since these are all well-known processes on the Earth or the Moon. We wonder, though, whether ‘flash melting’ of dust clumps, other than (perhaps) by irradiation within a plume, led to significant numbers of chondrules.

Dedication

We dedicate this chapter to the memory of two of the authors of Kennedy et al. (1992). Stuart Agrell, who found the microgabbro fragment in Parnallee, was a remarkable scientist and was an encouraging and supportive mentor to each of us as PhD students in Cambridge at the time of the Apollo landings. Robert Hutchison was a friend and an inspiration, never shy to say, in his inimitable Glaswegian accent, that he believed chondrule formation was associated with impacts on planetary bodies and had little to do with ‘flash melting.’

Acknowledgments

We thank Thorsten Kleine and Noēl Chaumard for their constructive comments on an early draft of this chapter, which led to significant improvements.

References

Alexander, C. M. O’D., Grossman, J. N., Ebel, D. S., and Ciesla, F. J. (2008). The formation conditions of chondrules and chondrites. Science, 320, 1617–1619. Asphaug, E., Jutzi, M., and Movshovitz, M. (2011). Chondrule formation during planetesimal accretion. Earth and Planetary Science Letters 308, 369–379. Bischoff, A., Wurm, G., Chaussidon, M., et al. (2017). The Allende multicompound chondrule (ACC) – Chondrule formation in a local super-dense region of the early solar system. Meteoritics and Planetary Science 52, 906–924. 372 Ian S. Sanders and Edward R. D. Scott

Budde, G., Kleine, T., Kruijer, T. S., Burkhardt, C., and Metzler, K. (2016a). Tungsten isotopic constraints on the age and origin of chondrules. Proceedings of the National Academy of Sciences 113, 2886–2891. Budde G., Burkhardt C., Brennecks G. A., et al. (2016b). Molybdenum isotopic evidence for the origin of chondrules and a distinct genetic heritage of carbonaceous and non-carbonaceous meteorites. Earth and Planetary Science Letters 454, 293–303. Budde, G., Kruijer, T. S., and Kleine, T. (2018). Hafnium-tungsten chronology of CR chon- drites: Implications for the timescales of chondrule formation and the distribution of 26Al in the solar nebula. Geochimica et Cosmochimica Acta 222, 284–304. Chen, J. H., Papanastassiou, D. A., and Wasserburg, G. J. (1998). Re-Os systematics in chondrites and the fractionation of the platinum group elements in the early solar system. Geochimica et Cosmochimica Acta 62, 3379–3392. Connelly J. N., Bizzarro M., Krot A. N., et al. (2012). The absolute chronology and thermal processing of solids in the solar protoplanetary disk. Science 338, 651–655. Dullemond, C. P., Stammler, S. M., and Johansen, A. (2014). Forming chondrules in impact splashes. I. Radiative cooling model. The Astrophysical Journal 794, 91. Dullemond, C. P., Harsono, D., Stammler, S. M., and Johansen, A. (2016). Forming Chon- drules in Impact Splashes II Volatile Retention. The Astrophysical Journal, 832, article id. 91. Faure, F., Tissandier, L., Libourel, G., Romain, M., and Welsch, B. (2012). Origin of glass inclusions hosted in magnesian porphyritic olivine chondrules: Deciphering planetesimal compositions. Earth and Planetary Science Letters 319,1–8. Faure, F., Tissandier, L., Florentin, L., and Devineau, K. (2017). A magmatic origin for silica- rich glass inclusions hosted in porphyritic magnesian olivines in chondrules: An experi- mental study. Geochimica et Cosmochimica Acta 204,19–31. Florentin, L., Faure, F., Deloule, E., et al. (2017). Origin of Na in glass inclusions hosted in olivine from Allende CV3 and Jbilet Winselwan CM2: Implications for chondrule forma- tion. Earth and Planetary Science Letters 474, 160–171. Gerber, S., Burkhardt, C., Budde, G., Metzler, K., and Kleine, T. (2017). Mixing and transport of dust in the early solar nebula as inferred from titanium isotope variations among chondrules. The Astrophysical Journal Letters, 841, L17. Grossman, J. N. (1988). Origin of chondrules. In J. F. Kerridge and M. S. Matthews (Eds.), Meteorites and the early Solar System, 680–696. Tucson, AZ: University of Arizona Press. Grossman, L., Fedkin, A. V., and Simon, S. B. (2012). Formation of the first oxidized iron in the solar system. Meteoritics and Planetary Science 47, 2160–2169. Hevey, P. J., and Sanders, I. S. (2006). A model for planetesimal meltdown by 26Al and its implications for meteorite parent bodies. Meteoritics and Planetary Science 41, 95–106. Hewins, R. H., Zanda, B., and Bendersky, C. (2012). Evaporation and recondensation of sodium in Semarkona Type II chondrules. Geochimica et Cosmochimica Acta 78,1–17. Hutcheon, I. D., and Hutchison, R. (1989). Evidence from the Semarkona ordinary chondrite for 26Al heating of small planets. Nature 337, 238–241. Johnson, B. C., Minton, D. A., Melosh, H. J., and Zuber, M. T. (2015). Impact jetting as the origin of chondrules. Nature 517, 339–341. Kennedy, A. K., Hutchison, R., Hutcheon, I. D., and Agrell, S. O. (1992). A unique high Mn/Fe microgabbro in the Parnallee (LL3) ordinary chondrite: Nebular mixture or planetary differentiate from a previously unrecognized planetary body? Earth and Planetary Science Letters 113, 191–205. Kiefer, W. S., and Mittlefehldt, D. W. (2017). Differentiation of asteroid 4 Vesta: Core formation by iron rain in a silicate magma ocean. Lunar and Planetary Science Conference XLVIII, abstract # 1798. The Dirty Impact Plume Model 373

Kita, N. T., Nagahara, H., Tachibana, S., et al. (2010). High precision SIMS oxygen three isotope study of chondrules in LL3 chondrites: Role of ambient gas during chondrule formation. Geochimica et Cosmochimica Acta 74, 6610–6635. Kita, N. T., Tenner, T. J., Defouilloy, C., et al. (2015). Oxygen isotope systematics of chondrules in R3 clasts: A genetic link to ordinary chondrites. Lunar and Planetary Science Conference XLVI, abstract #2053. Kleine, T., Mezger, K., Palme, H., Scherer, E., and Münker, C. (2005). Early core formation in asteroids and late accretion of chondrite parent bodies: Evidence from 182Hf-182W in CAIs, metal-rich chondrites, and iron meteorites. Geochimica et Cosmochimica Acta 69, 5805–5818. Kleine, T., and Wadhwa, M. (2017). Chronology of Planetesimal Differentiation. In L. T. Elkins-Tanton and B. P. Weiss (Eds.), Planetesimals: Early differentiation and Conse- quences for Planets, 224–245. Cambridge, UK: Cambridge University Press. Kruijer, T. S., Burkhardt, C., Budde, G., and Kleine, T. (2017). Age of Jupiter inferred from the distinct genetics and formation times of meteorites. Proceedings of the National Academy of Sciences 114, 6712–6716. Larsen, K. K., Trinquier A., Paton C., et al. (2011). Evidence for magnesium isotope heterogeneity in the solar protoplanetary disk. The Astrophysical Journal Letters 735, L37–L43. LaTourrette, T., and Wasserburg, G. J. (1998). Mg diffusion in anorthite: Implications for the formation of early solar system planetesimals. Earth and Planetary Science Letters 158, 91–108. Lichtenberg, T., Golabek, G. J., Gerya, T. V., and Meyer, M. R. (2016). The effects of short- lived radionuclides and porosity on the early thermo-mechanical evolution of planet- esimals. Icarus 274, 350–365. Lichtenberg, T., Golabek, G. J., Dullemond, C. P., Schönbãchler, M., Gerya, T. V., and Meyer, M. B. (2018). Impact splash chondrule formation during planetesimal recycling. Icarus, 302,27–43. Lugmair, G. W., and Shukolyukov, A. (2001). Early solar system events and timescales. Meteoritics and Planetary Science 36, 1017–1026. Metzler, K. (2012). Ultra-rapid chondrite formation by hot chondrule accretion? Evidence from unequilibrated ordinary chondrites. Meteoritics and Planetary Science 47, 2193–2217. Nagahara, H. (1981). Evidence for secondary origin of chondrules. Nature 292, 135–136. Nagashima, K., Krot, A. N., and Huss, G. R. (2015). Oxygen-isotope compositions of chon- drule phenocrysts and matrix grains in Kakangari K-grouplet chondrite: Implication to a chondrule-matrix genetic relationship. Geochimica et Cosmochimica Acta 151,49–67. Niemeyer, S. (1985). Systematics of Ti isotopes in carbonaceous chondrites. Geophysical Research Letters, 12, 733–736. Olsen, M. B., Wielandt, D., Schiller, M., Van Kooten, E. M. M. E., and Bizzarro, M. (2016). Magnesium and 54Cr isotope compositions of carbonaceous chondrite chondrules – Insights into early disk processes. Geochimica et Cosmochimica Acta 191, 118–138. Palme, H., Hezel, D. C., and Ebel, D. S. (2015). The origin of chondrules: Constraints from matrix composition and matrix-chondrule complementarity. Earth and Planetary Science Letters 411,11–19. Rambaldi, E. R. (1981). Relict grains in chondrules. Nature 293:558–561. Schiller, M., Connelly, J. N., Aslaug, C. G., Mikouchi, T., and Bizzarro, M. (2015). Early accretion of protoplanets inferred from a reduced inner solar system 26Al inventory. Earth and Planetary Science Letters 420,45–54. Sanders, I. S. (1996). A chondrule-forming scenario involving molten planetesimals. In R. H. Hewins, R. H. Jones, and E. R. D. Scott (Eds.), Chondrules and the Protoplanetary Disk, 327–334. Cambridge, UK: Cambridge University Press. 374 Ian S. Sanders and Edward R. D. Scott

Sanders, I. S., and Scott, E. R. D. (2012). The origin of chondrules and chondrites: Debris from low-velocity impacts between molten planetesimals? Meteoritics and Planetary Sciences 47, 2170–2192. Sanders, I. S., and Taylor, G. J. (2005). Implications of 26Al in nebular dust: Formation of chondrules by disruption of molten planetesimals. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk. Astronomical Society of the Pacific Conference Series 341, 821–838. San Francisco, CA: Astronomical Society of the Pacific. Schrader, D. L., Nagashima, K., Fu, R. R., Davidson, J., and Ogliore, R. C. (2017). Evidence for chondrule migration from dusty olivine chondrules. Lunar and Planetary Science Confer- ence XLVIII, abstract# 1271. Soulié, C., Libourel, G., and Tissandier, L. (2017). Olivine dissolution in molten silicates: An experimental study with application to chondrule formation. Meteoritics and Planetary Sciences 52, 225–250. Tachibana, S., Nagahara, H., Mostefaoui, S., and Kita, N. T. (2003). Correlation between relative ages inferred from 26Al and bulk compositions of ferromagnesian chondrules in least equilibrated ordinary chondrites. Meteoritics and Planetary Science 38, 939–962. Taylor, G. J., Scott, E. R. D., and Keil, K. (1983). Cosmic setting for chondrule formation. In E. A. King (Ed.), Chondrules and their origins, 262–278. Houston, TX: Lunar and Planetary Institute. Tenner, T. J., Nakashima, D. N., Ushikubo, T., Kita, N. T., and Weisberg, M. K. (2015). Oxygen isotope ratios of FeO-poor chondrules in CR3 chondrites: Influence of dust enrichment and H2O during chondrule formation. Geochimica et Cosmochimica Acta 148, 228–250. Trinquier, A., Birck, J. -L., and Allègre, C. J. (2007). Widespread 54Cr heterogeneity in the inner Solar System. The Astrophysical Journal 655, 1179–1185. Villeneuve, J., Chaussidon, M., and Libourel, G. (2012). Absence de relation entre les âges 26Al de cristallisation des chondres et leurs compositions minéralogiques et chimiques. Comptes Rendus Geoscience 344, 423–431. Villeneuve, J., Libourel, G., and Soulié, C. (2015). Relationships between type I and type II chondrules: Implications on chondrule formation processes. Geochimica et Cosmochimica Acta 160, 277–305. Wänke, H., Dreibus, G., Jagoutz, E., Palme, H., and Rammensee, W. (1981). Chemistry of the Earth and the significance of primary and secondary objects for the formation of planets and meteorite parent bodies (abstract). Lunar and Planetary Science Conference XII, 1139–1141. Wänke, H., Dreibus, G., and Jagoutz, E. (1984). Mantle chemistry and accretion history of the Earth. In A. Kroner, G. N. Hanson, and A. M. Goodwin (Eds.), Archaean Geochemistry, 1–24. Berlin, Germany: Springer Verlag. Warren, P. H. (2011). Stable-isotopic anomalies and the accretionary assemblage of the Earth and Mars: A subordinate role for carbonaceous chondrites. Earth and Planetary Science Letters 311,93–100 Wilson, L., and Keil, K. (2012). Volcanic activity on differentiated asteroids: A review and analysis. Chemie der Erde 72, 289–322. Zook, H. A. (1980). A new impact model for the generation of ordinary chondrites (abstract). Meteoritics 15, 390–391. Zook, H. A. (1981). On a new model for the generation of chondrules (abstract). Lunar and Planetary Science Conference XII, 1242–1244. 15 Formation of Chondrules by Shock Waves

melissa a. morris and aaron c. boley

Abstract

Nebular shock heating is one of the most fully developed and rigorous models for chondrule formation, and is also the most consistent with the meteoritic record. In this review, we compare the results of current shock modeling to the wealth of meteoritic observations, to highlight where there is agreement and where there is potential failure of the models. The discussion is focused on gravitational disk instability-driven, large-scale shocks and on local bow shocks, with attention to the astrophysical setting for both. We suggest that more than one shock driver may be physically motivated and necessary to explain the variety of chondrules.

15.1 Introduction

Understanding the “chondrule-formation mechanism” has been an elusive task for over a century (Farrington, 1915). Proposed formation mechanisms include lightning, interactions with the young sun, planetesimal impacts, and nebular shocks (Shu et al., 1996; Desch and Cuzzi, 2000; Morris and Desch, 2010; Sanders and Scott, 2012). Among these, the heating of chondrule precursors in nebular shocks is one of the most fully developed and rigorous models for chondrule formation and is the most consistent with the meteoritic record (Morris et al., 2016), although significant questions certainly remain. There are several possible mechanisms for driving shocks in the solar nebula, including gravitational disk instabilities (Boss and Durisen, 2005; Boley and Durisen, 2008), X-ray flares or accretion shocks (e.g., Nakamoto et al., 2005), and bow shocks around planetesimals or protoplanets on eccentric orbits (Hood et al., 2009; Morris et al., 2012). X-ray flares and accretion shocks have largely been eliminated as possible drivers, because these shocks would take place at the surface of the disk, and chondrule precursors are expected to have settled to the midplane. Very little millimeter-sized material is expected at the surface of a disk, so if chondrules were to form there, they should not be as abundant as seen in the meteoritic record, and would show evidence for having formed in low density environments (e.g., elemental/ isotopic evidence for evaporation). Therefore, we concentrate our attention on shocks driven by gravitational instabilities (GIs) and bow shocks. Spiral asymmetries arising from disk instabilities can spontaneously form in massive proto- planetary disks whenever the destabilizing effects of self-gravity overwhelm the stabilizing effects of disk shear and sound waves (Durisen et al., 2007). This occurs roughly when the magnitude of the vertical component of the disk’s self-gravity, measured at the disk scale

375 376 Melissa A. Morris and Aaron C. Boley height, is comparable to or greater than the central star’s contribution (Boley and Durisen, 2008). These types of instabilities may have occurred in the early solar nebula (Gammie, 1996), driving strong spiral shocks. Recent Atacama Large Millimeter Array (ALMA) observations lend weight to this idea, in which a spiral structure that is morphologically similar to that expected for disk instabilities is clearly seen in the young system Elias 2–27 (Smith, 2016). Bow shocks can form if planetesimals and protoplanets are placed on eccentric orbits while embedded in a gaseous disk. At 2.5 AU, within the presumed chondrule-forming region, gas orbits the Sun at the circular Keplerian speed of vK ~ 20 km/s. A planetary body with eccentricity e will have phases during its orbit when speeds relative to the gas are ~evK.As an example, planetesimals in resonance with a proto-Jupiter would have had their eccentricities driven as high as e ~ 0.3–0.5 (Hood et al., 2009), inducing strong bow shocks (~8 km/s). Large-scale GI-driven shocks and small-scale bow shocks have both been shown to be consistent with the inferred range of thermal histories of porphyritic chondrules under certain conditions, with cooling rates at the low end and high end respectively (Morris and Desch, 2010; Mann et al., 2016; Morris et al., 2016). These two mechanisms would occur during different phases of disk evolution, allowing for the possibility of a range of formation times and conditions for chondrules. In this chapter, we discuss the wealth of meteoritic observations that constrain chondrule formation in general, and compare these to the current results found through modeling of chondrule formation in nebular shocks. In particular, we discuss the inferred thermal histories of chondrules (Hewins and Connolly, 1996) and their retention of volatiles (Alexander et al., 2008), including recent studies that question the long-accepted thermal constraints (Miura and Yamamoto, 2014; Villeneuve et al., 2015). We provide an overall assessment of current shock models for chondrule formation and discuss further work needed.

15.2 The Physics of Shocks

Shocks occur when a wave propagates faster than the local sound speed cs, resulting in an abrupt, sharp change in gas properties (e.g., density, pressure, and temperature). Strong shocks (Mach number, M 1), such as those considered in this chapter, occur when the gas abruptly changes from a supersonic to a subsonic flow, and the density across the shock front increases by at least a factor of four. Jump conditions are used to ensure that mass, momentum, and energy are conserved across the shock front discontinuity. In the case of pure hydrodynamic shocks in 1D flows, such as large-scale, GI-driven shocks, the Rankine-Hugoniot jump condi- tions are typically used:

ρ ¼ ρ ðÞ 1v1 2v2 Conservation of mass ; (15.1)

ρ 2 þ ¼ ρ 2 þ ðÞ 1v1 P1 2v2 P2 Conservation of momentum ; and (15.2)

1 γ P 1 γ P v2 þ 1 ¼ v2 þ 2 ðÞ 1 ðÞγ ρ 2 ðÞγ ρ Conservation of energy , (15.3) 2 1 1 2 1 2 Formation of Chondrules by Shock Waves 377 where ρ is the gas density, v is the velocity of the gas, P is the gas pressure, and γ is the adiabatic index for the gas. The subscripts 1 and 2 refer to the preshock and postshock regions, respectively.

In terms of the Mach number M = vs/cs, where vs is the shock speed and cs is the local sound speed, the jump conditions for gas density, pressure, and temperature can be expressed as follows:

ðÞγ þ 1 M2 ρ ¼ ρ ; (15.4) 2 1 ðÞγ 1 M2 þ 2

2γM2 ðÞγ 1 P ¼ P ; (15.5) 2 1 ðÞγ þ 1 ðÞγ 1 M2 þ 2 2γM2 ðÞγ 1 T2 ¼ : (15.6) ðÞγ þ 1 2M2

For a gas of solar composition (H, H2, and He)

n þ 2n þ 4n m v2 M2 ¼ H H2 He H 1 (15.7) þ þ 2 nH nH2 nHe kT1

5 þ 7 þ 5 nH nH2 nHe γ ¼ 2 2 2 , (15.8) 3 5 3 n þ n þ n 2 H 2 H2 2 He where n is the number of atoms or molecules for each species. In the case of strong shocks that are optically thick, a radiation front can propagate from the hot, postshock region into the preshock gas (Marshak wave). This situation can be well described by using the diffusion approximation to model the radiation field (Mihalas and Mihalas, 1984). The opacity of the material in the preshock region will determine how quickly the infrared radiation can diffuse, and therefore the distance the Marshak wave can penetrate. The distance that the Marshak wave can travel in the preshock region, in turn, determines the extent of heating that particles will experience before they encounter the hydrodynamic shock front. So-called “line photons”, emitted from molecular species such as H2O, play a role in transporting energy from the hot postshock gas to the preshock region as well (Morris et al., 2009; Morris and Desch, 2010). Strong nebular shocks can also lead to H2 dissociation. If the gas were to remain in thermodynamic equilibrium throughout the shock, then dissociation alone would keep the peak gas temperature in the postshock region between 2,000 K and 2,500 K. However, immediately behind the shock, the number density of H2 is out of equilibrium. As such, the peak gas temperature initially spikes toward that expected for a gas with a fixed adiabatic index. H2 quickly dissociates in the postshock environment, limiting the temperature spike and lowering 378 Melissa A. Morris and Aaron C. Boley the temperature toward the equilibrium value. All of these effects must be considered when determining the thermal histories of chondrules in shocks.

15.2.1 Heating of Particles in Shocks

When particles hit the shock front, they are “flash-heated” by friction as they are slowed to the postshock gas speed. Following this short episode of drag heating, the particles are then heated by thermal exchange between the hot, dense, postshock gas and through the absorption of infrared radiation emitted by surrounding particles. The dissociation of H2 will effectively buffer the gas temperature, which places limits on the peak temperature of solids, as well as the duration at peak temperature. Dissociation of H2 from nonequilibrium values causes solids that are partially thermally coupled to the gas to experience cooling rates >5,000 K/hr for less than a minute as the gas approaches a quasichemical equilibrium. The initial phases of cooling are dominated by this H2 dissociation, with line-cooling having minimal affect on the cooling rates of solids (Morris and Desch, 2010). Because particles are heated in shocks through interactions between the gas and surrounding particles, shock models of chondrule formation must account for the dynamics and energetics of particles and gas separately, and include interactions between the two. The transfer, absorption, and emission of radiation must be included in any rigorous shock model, as well as H2 dissociation and recombination and the kinetics/energetics of evaporation and recondensation of solids. Once fully developed, the predictions made by the shock model must then be tested against the meteoritic record.

15.3 Astrophysical Setting

The starting parameters for any nebular shock model are tied to assumptions made about the properties of the solar nebula. These assumptions are rooted in our understanding of the composition, temperature, and solids-to-gas ratio in the nebula prior to and during planet formation, which have a range of possibilities (e.g., Weidenschilling, 1977; Hayashi, 1981; Desch, 2007). The solids-to-gas ratio – at least in local, chondrule-forming regions – is a matter of particular debate. For example, several recent lines of evidence (e.g., Alexander et al., 2008) may indicate a much higher solids-to-gas ratio during chondrule formation than traditionally assumed. We discuss this further in the following summary of the meteoritical constraints on chondrule formation.

15.4 Meteoritic Constraints on Chondrule Formation

Although the detailed properties of chondrules are discussed in previous chapters, we provide a brief review of the constraints placed on chondrule formation by the meteoritic record. We begin with a discussion of the timing of chondrule formation, move on to review so-called “non- thermal” constraints, and end with a discussion of what has been considered the strongest constraint on formation mechanisms – the inferred chondrule thermal histories. Formation of Chondrules by Shock Waves 379 15.5 Timing of Chondrule Formation

Various isotopic systems (e.g., Al–Mg, Hf-W, U–Pb) are utilized to determine the timing of chondrule formation. Isotopic dating results indicate that chondrule formation took place over a wide period of time during the evolution of the solar nebula. Chondrules have been isotopically dated as early as contemporaneous with calcium-aluminum-rich inclusions (CAIs) to as late as ~ 5 Myr after CAIs (Amelin et al., 2002; Kita et al., 2005; Krot et al., 2005; Russell et al., 2006; Wadhwa et al., 2007; Connelly et al., 2008; Amelin et al., 2010; Connelly et al., 2012; Bollard, Connelly, and Bizzarro, 2014, Budde et al., 2016), although this wide range in ages is debated (Alexander and Ebel, 2012). It is clear, though, that the majority of porphyritic chondrules formed within 1.5–4 Myr after CAIs (Kurahashi et al., 2008; Villeneuve et al., 2009; Kita and Ushikubo, 2012; Schrader et al., 2017). The implied range in ages of chondrule formation is consistent with the lifespans of protoplanetary disks (3–6 Myr: Haisch et al., 2001; Williams and Cieza, 2011). Thus, chondrule formation may provide clues to the environment over a significant period of evolution of the solar nebula, from the formation of the earliest solids to the dissipation of our protoplanetary disk (Morris et al., 2015).

15.6 Nonthermal Constraints on Chondrule Formation

Chondrule formation seems to have been an ongoing process rather than a one-time event. This is inferred not only by the wide range in chondrules’ ages, but by the existence of fragments of relict chondrules found within later generations of chondrules (Jones et al., 2000; Connolly and Jones, 2016). Elevated gas partial pressures are required during chondrule formation in order to stabilize the silicate melt against evaporation (e.g., Ebel and Grossman, 2000; Connolly and Jones, 2016). Although dependent on the amount of dust enrichment in the chondrule-forming region, the pressures required are higher than believed to be plausible for ambient solar nebular conditions at a few AU. Such elevated gas pressures, however, could be achieved in the postshock region following passage of a shock front. The oxygen fugacity in the chondrule-forming region was likely more oxidizing than a solar composition gas (Krot et al., 2000), although it is thought to have been variable, either temporally or spatially. In order to retain volatiles and prevent isotopic fractionation in chondrules, Cuzzi and Alexander (2006) proposed that the size of the chondrule-forming region was 103 km and – that the density of solids must have exceeded ~10 m 3. Desch et al. (2012) argued for a similarly large, dense chondrule-forming region to explain the observed frequency of compound chon- drules. Several recent studies indicate chondrule formation took place in the presence of high vapor pressures of several moderately volatile species, such as Na, K, Si, Fe, and Li (e.g., Alexander et al., 2008; Alexander and Ebel, 2012; Grossman et al., 2012; Hewins et al., 2012; Fedkin and Grossman, 2013, 2016; Schrader et al., 2013; Schoelmerich et al., 2016). The most likely explanation for these high vapor pressures is a higher-than-average solids-to-gas ratio in the localized chondrule-forming region. In some cases, these results suggest orders-of-magni- tude higher concentrations of solids than the ~10 m–3 advocated by Cuzzi and Alexander (2006) and Desch et al. (2012). 380 Melissa A. Morris and Aaron C. Boley

As discussed in Chapter 4, some have interpreted chondrule-matrix complementarity as evidence that chondrules and matrix dust within a given chondrite formed in the same region of the solar nebula (Wood, 1985; Palme et al., 1993; Klerner and Palme, 1999; Hezel and Palme, 2008). Others, however, have argued that chondrule-matrix complementarity is due to redistri- bution on the parent body – at least for some chemical systems (Zanda et al., 2009), or that the roughly CI abundances of presolar grains, organic matter and volatiles in matrix suggests that most was not heated to anything close to peak chondrule temperatures (e.g., Alexander et al., 2001; Alexander, 2005). This disagreement has led to controversy regarding the veracity of complementarity as a constraint on chondrule formation. Nevertheless, recent work by Budde et al. (2016) on 183W complementarity seems to rule out parent-body processes because W is not readily mobilized in such a setting. It is interesting to note that planetary collisions were also dismissed as a chondrule-forming mechanism in their study because bulk meteorites (and presumably any colliding parent bodies) have no 183W anomalies, whereas chondrules have large (and complementary) 183W anomalies (Budde et al., 2016). By extension, there is not a clearly identified mechanism for causing 183W anomalies in chondrule and matrix precursors. Zanda et al. (Chapter 5) argue that the measurements of chondrules and matrix alone cannot be compared to the bulk meteorite, but must also include other components (e.g., CAIs, metal). Further work is required before complementarity can be considered to be a strong constraint on chondrule formation, as virtually all formation mechanisms can meet (or be eliminated by) this constraint, depending on interpretation.

15.7 Thermal Constraints on Chondrule Formation

Several previous chapters (particularly Chapter 1) have addressed in detail the meteoritical constraints on the thermal histories of chondrules. However, since the inferred thermal histories of chondrules are considered the first-order (or zeroth-order) constraint against which formation mechanisms are to be tested (Connolly and Jones, 2016), we present a brief summary here. Thermal histories of chondrules – the rate of precursor heating, peak temperatures achieved, duration of heating, and cooling rates – are determined mainly by chondrule chemistry and textures (Lofgren, 1982, 1989, 1996; Hewins and Connolly, 1996; Hewins, 1997; Connolly, Jones, and Hewins, 1998; Desch and Connolly, 2002; Connolly and Desch, 2004; Connolly et al., 2006; Lauretta et al., 2006). Chondrules with porphyritic textures, which make up approximately 84 percent of ordinary chondrite chondrules (Gooding and Keil, 1981), are inferred to have experienced peak tempera- tures in the range of 1,770–2,120 K for only seconds to minutes (Lofgren and Lanier, 1990; Radomsky and Hewins, 1990; Hewins and Connolly, 1996; Lofgren, 1996; Hewins, 1997; Connolly and Love, 1998; Jones et al., 2000; Connolly and Desch, 2004; Ciesla, 2005; Hewins et al., 2005; Connolly et al., 2006; Lauretta et al., 2006). Barred chondrules are thought to have experienced higher peak temperatures than porphyritic chondrules – over a range between 1,820 and 2,370 K (Hewins and Connolly, 1996). Reproduction of chondrule textures through heating of analog materials in furnace experi- ments suggest that porphyritic chondrules cooled at rates ranging from 10 to 1,000 K/hr Formation of Chondrules by Shock Waves 381

(Hewins et al., 2005), although this has recently been challenged (see discussion in next paragraph). For plagioclase-bearing CO chondrules, the cooling rates may have been as low as 1 K/hr (Wick and Jones, 2012). Chondrules with barred textures are inferred to have experienced more rapid cooling, at rates ranging from 300 to 3,000 K/hr (see review by Desch et al., 2012). The duration of heating, time at peak temperature, and starting temperature are determined mainly through chemical and isotopic studies. The presence of primary sulfides indicates that chondrule precursors were relatively cold before melting, forming in an environment with an ambient temperature <650 K (Rubin et al., 1999). Under conditions of free evaporation, the retention of volatile elements, such as Na and S, by chondrules has been interpreted by some to mean that they did not remain above the liquidus for more than minutes, and cooled extremely rapidly (≳104 K/hr) to their liquidus temperature (Yu and Hewins, 1998). The retention of S suggests that chondrule precursors did not experience heating between the starting tempera- ture of ~ 650–1,200 K for more than several minutes (Hewins and Connolly, 1996; Connolly and Love, 1998; Jones et al., 2000; Lauretta et al., 2001; Tachibana and Huss, 2005; Connolly et al., 2006). The observed lack of isotopic fractionation caused by free evaporation limits the duration chondrule precursors were between 1,300 and 1,600 K before melting (Tachibana et al., 2004; Tachibana and Huss, 2005) to minutes. The constraints on retention and fractionation of volatiles can be avoided, however, if chondrules formed in an environment with high partial pressures of volatiles that suppress evaporation. A solids-to-gas ratio that is much higher than that typically assumed ( ρd/ρg =5 10–3) as proposed by Alexander et al. (2008), for instance, could provide such an environment. Based on several recent cosmochemical investigations that suggest high vapor pressures of several moderately volatile species (e.g., Alexander et al., 2008; Alexander and Ebel, 2012; Grossman et al., 2012; Hewins et al., 2012; Fedkin and Grossman, 2013, 2016; Schrader et al., 2013; Schoelmerich et al., 2016), the preshock thermal histories as shown in Figure 15.1 may be unnecessary.

Figure 15.1 Chondrule thermal histories as inferred from experimental constraints. The shaded region in the postshock region indicates the range of inferred cooling rates (10–1,000 K/hr). 382 Melissa A. Morris and Aaron C. Boley

The long-accepted cooling rates of porphyritic chondrules (10–1,000 K/hr) have also been called into question. Villeneuve et al. (2015) argue that chondrule textures are not necessarily controlled by cooling rates and that more rapid cooling rates are possible. Miura and Yamamoto (2014) developed a formula for cooling rates, based on compositional zoning, that indicated Type II (FeO-rich) chondrules cooled at rates of 200–2,000 K/s. If these claims survive further tests, thermal histories of chondrules will need to be revisited.

15.8 Shock Mechanisms

We consider two main shock-inducing mechanisms to be feasible for chondrule formation: Shocks driven by large-scale asymmetries, such as gravitational disk instabilities, and bow shocks induced by planetary bodies on eccentric orbits.

15.8.1 Large-Scale Shocks

Strong, large-scale shocks in protoplanetary disks can be driven, for example, by gravitational disk instabilities, which manifest as spiral structures (Boss and Durisen, 2005; Boley and Durisen, 2008). Morris et al. (2016) investigated such shocks in detail using a 1-D steady-state –9 –3 flow with a shock of 8 km/s propagating through an ambient gas with density ρg =10 gcm (Desch, 2007) and temperature 300 K. A 1-D model is appropriate in this case due to the scale of the spiral shock. The model assumes equal proportions of micron-sized dust grains and particles of radius a =10μm, with a total mass fraction 1.25 10–3 relative to the gas, and a –3 population of chondrules of radii ac = 300 μm and a mass fraction (3.75 10 )C, where C is the concentration of chondrules, relative to the solar abundance of condensable elements at that temperature. A chondrule concentration C = 10 was used to include a slight enhancement over the assumed solar solids-to-gas ratio, which could be achieved solely through the settling of particles to the midplane (Weidenschilling, 1980) or turbulent concentration (Cuzzi et al., 2008). Chondrule precursors start at the temperature of the ambient gas, and as the shock approaches, are heated by infrared radiation emitted by solids behind the shock. During this stage, chondrules are heated nearly to their peak temperatures. At the shock front, the gas velocity instantaneously changes (within meters) but the chondrule precursors change velocity only after colliding with their own mass of gas. Thus, they slow to the postshock gas velocity only after propagating for an aerodynamic stopping time, tstop ~(ρsac)/(ρgcg)~1 minute. Here, –3 ρs 3gcm is the material density of the chondrules, ac = 300 μm are their radii, ρg 6 –9 –3 10 gcm is the post-shock gas density, and cg 3 km/s is the post-shock sound speed. For this short time period, chondrules are heated by gas-drag friction in addition to thermal exchange with the shocked gas and absorption of infrared radiation. The gas-drag friction heats chondrules from their preshock temperatures by a few hundred K, to their peak temperatures; but, by the end of the first tstop, the chondrules have cooled to nearly the equilibrium postshock gas temperature, experiencing cooling rates from their peak temperatures of ~104 K/hr. There- after, chondrules are in thermal equilibrium with the gas, and cool only as fast as they can leave the shocked region, several optical depths downstream from the shock front. Formation of Chondrules by Shock Waves 383

The 1-D shock code of Morris et al. (2016) treats the gas and solids as separate fluids. The composition of the gas is represented by four populations: atomic hydrogen (H), molecular hydrogen (H2), helium atoms (He), and molecules resulting from the evaporation of solids, represented by SiO. Solids are divided into submicron- to micron-sized dust grains, intermediate-sized particles of radius a =10μm, and larger, spherical chondrule precursors of radius ac = 300 μm. Dust grains are assumed to be dynamically and thermally coupled to the gas. An approximation to the temperature-dependent mean Planck opacity is used to calculate dust opacity, as described in Morris and Desch (2010). If the dust grains (represented by submicron forsterite grains, for simplicity) exceed 1,500 K, they are assumed to have evaporated completely. The intermediate- sized and chondrule-sized particles are unique in the treatment of their number density n, their velocity V, their temperature T, and their radii a, as well as material properties. All

fluids are initialized with speed Vs and temperature Tpre. The densities, velocities, and tempera- tures are integrated forward using a fourth-order Runge-Kutta integration, assuming a steady- state flow and applying the equations of continuity, a force equation including the drag force between gas and solids, and appropriate energy equations. The gas can absorb radiation energy (via the dust opacity), can be heated by interaction with intermediate-sized particles as well as chondrules and their precursors, and can lose or gain heat energy through dissoci- ations or recombinations of H2 and molecular line radiation of H2O. Intermediate-sized par- ticles, chondrules, and their precursors, can emit or absorb radiation, or be heated by thermal exchange with the gas or by frictional drag heating. Calculation of the radiation field follows the approach outlined in Mihalas (1978), assuming plane-parallel, temperature-stratified slab atmospheres. Physically, large-scale nebular shocks are not truly one dimensional, as radiation will be lost from the top and bottom of the disk (Morris and Desch, 2010; Stammler and Dullemond, 2014). Therefore, the model of Morris et al. (2016) does not properly account for vertical energy losses, so does not allow for the precise determination of the radiation field and postshock temperature. A 2-D model is needed for chondrule formation in large-scale shocks, incorporating all the relevant physics as in the 1-D model of Morris et al. (2016), in order to properly determine the effects of vertical energy loss.

15.8.1.1 Comparison to Meteoritic and Astrophysical Constraints The model of Morris et al. (2016) predicts chondrule cooling rates of 20–40 K/hr through most of the crystallization temperature range (1,600–1,400 K), for solids-to-gas ratio 10 solar. Cooling rates through 1,800–1,600 K are somewhat higher, but are still well within the range consistent with the meteoritic constraints on thermal histories of chondrules. The excess preheating of chondrule precursors predicted by the model of Morris and Desch (2010), which was inconsistent with isotopic constraints, has been eliminated. Chondrule temperatures throughout heating and cooling are in agreement with long-accepted thermal histories (Figure 15.2). Thus, the model meets the first-order (or zeroth-order) constraint on formation of chondrules. The model also predicts higher cooling rates in regions of high solids density (as did Desch and Connolly, 2002; Morris and Desch, 2010; Desch et al., 2012), consistent with the prevalence of barred textures among compound chondrules. 384 Melissa A. Morris and Aaron C. Boley

Figure 15.2 The data points indicate chondrule thermal histories as predicted by the large-scale shock model of Morris et al. (2016), as compared to those inferred from experimental constraints (solid curve). The shaded region in the postshock region indicates the range of inferred cooling rates (10–1,000 K/hr). Chondrule precursors start at an ambient temperature of 300 K, and remain at ~300 K until ~15 minutes before the shock front. Figure adapted from Morris et al. (2016)

The large-scale shock model of Morris et al. (2016) makes no predictions regarding the retention of moderately volatile species and lack of isotopic fractionation, as the large solid-to- gas ratios deemed necessary have not been included in the model. However, as noted in Section 15.7, the model predicts that high concentrations of solids increase cooling rates (although eventually the optical depth can become so high that no cooling occurs). In the case of a high concentration of large precursors that survive the passage of the shock, postshock cooling rates of chondrules might increase to the point that they no longer meet the meteoritic constraints, or in the extreme case, may not cool at all. It may even be that concentrations of chondrule precursors high enough to provide the necessary vapor pressure of volatiles would inhibit the shock altogether (Mihalas and Mihalas, 1984). These conditions need to be included in the large-scale shock model.

15.8.2 Feasibility of Large-Scale Shocks

Large-scale spiral shocks can be produced by perturbations from a nearby (or flyby) star, by a massive, embedded planet, or through instabilities within the disk itself. A stellar flyby would represent one (or potentially a few) events, so it is obviously not a favored mechanism. Perturbations from planets may only be strong enough to form chondrules within a narrow region around the planet (Nelson and Ruffert, 2005). Gravitational instabilities, on the other hand, are a mechanism that can persist during the early stages of disc evolution (~first Myr) and could be ubiquitous in the young solar nebula during that limited time. While the existence of such a massive disk stage has been hotly debated in the literature (see Durisen et al., 2007 and Helled et al., 2014), recent observations do provide evidence through imaging of the corres- ponding spiral structure (Smith, 2016). Formation of Chondrules by Shock Waves 385

With gravitational instabilities, most of the spiral shocks are expected to be too low in strength (typical Mach numbers M2 ~ 2) due to the spiral structure having shallow pitch angles (Boley and Durisen 2008). Nonetheless, occasional strong shocks are possible and do occur (Boss and Durisen, 2005).

15.8.3 Bow Shocks

The bow shock hypotheses for chondrules posit that planetesimals and/or protoplanets in the solar nebula were excited into very eccentric and/or inclined orbits during the planet formation process (e.g., Hood, 1998). The gas in planet-forming discs is expected to follow a Keplerian rotation curve with only minor deviations due to, for example, radial pressure gradients (e.g., Adachi et al., 1976; Weidenschilling, 1977), and as such, the gas flow can be approximated as being on a circular orbit at any given location. Planetesimals (here with sizes 10 km) and protoplanets (sizes ≳1,000 km) embedded in this gaseous disc will also be on Keplerian orbits that can become (at least temporarily) elliptical. Any planetesimal can thus experience a relative gas wind speed Vw that depends on the phase and characteristics of its orbit. For large enough eccentricities and/or inclinations, the relative wind speed can become hypersonic, which results in the formation of a hydrodynamic bow shock around the planetesimal. Gas drag effects on planetesimals, as well as disc torques for protoplanets, will eventually damp down elliptical orbits. Nonetheless, this process can take thousands of orbits (possibly much longer), depending again on the orbit itself, the size and mass of the planetesimal, and the regional disc conditions (see, e.g., Morris et al., 2012). Furthermore, for any given elliptical orbit, a planetesimal can experience a wide range of shock conditions. This is highlighted in Figure 15.3. Eccentricities and inclinations are chosen such that the shock speeds will be between 6 and 8 km/s for at least part of the orbit. At high enough inclinations, the planetesimal will enter a low-density region. To take this into account on the plots, we assume that the planet-forming disc has a fixed scale ratio H/R = 0.1, where H is the scaleheight of the disk and R is the distance from the central star. Whenever the planetesimal travels beyond this limit (either above or below), we set Vw to zero. An important point to consider in this section is that there is not a single bow shock speed for any given orbit. Moreover, all planetesimals will damp down over time due to gas damping (Morris et al., 2012), so a planetesimal on a very excited orbit will eventually go through a wide range of shock speeds. Figure 15.3 shows that eccentricities and/or inclinations need to be moderately high to achieve shock speeds ~8 km/s, the optimal shock speed to produce chondrules in 1-D shock calculations.

15.8.3.1 Anatomy of a Bow Shock Bow shocks are fundamentally different from large-scale shocks, and cannot be treated with the 1-D models discussed so far. For bow shocks, the planetesimal or protoplanet (the obstruction) causes a sudden diversion in the gas flow in the region surrounding the hypersonic object (Figure 15.4). Grains that encounter the bow shock can have a range of thermal histories depending on the grain’s impact radius b (i.e., the pre-encounter distance a grain has from the obstruction measured normal to the planetesimal’s trajectory). At b ~ 0, gas enters a narrow 386 Melissa A. Morris and Aaron C. Boley

Figure 15.3 Wind speeds for planetesimals in a gaseous disc for a range of conditions. All calculations assume that the gas is on circular, Keplerian orbits with a fixed scale height ratio H/R = 0.1, where R is the cylindrical distance from the star. Two different semimajor axes, a, are explored, as well as a range of eccentricities and inclinations. For inclined orbits, the argument of pericenter, ω, is also indicated in each subplot, as this affects phase of the orbit that passes through the midplane. Whenever the planetesimal exceeds the local scaleheight of the disc, we set the wind speed to zero to illustrate that it has entered a low- density region. With eccentricities between ~0.3, we can expect bodies to have regions of shock speeds that are comparable to 8 km/s. Formation of Chondrules by Shock Waves 387

Adiabatic, 7 km/s Wind 8 14000 7 12000 6 10000 5 8000 4 6000 3 Distance (km) 4000 2 Density/(Wind Density)

2000 1

0 0 -15000 -10000 -5000 0 5000 10000 15000 Distance (km)

Adiabatic, 7 km/s Wind

14000 2000 12000

10000 1500

8000

6000 1000 Distance (km)

4000 Gas Temperature (K)

2000 500

0 -15000 -10000 -5000 0 5000 10000 15000 Distance (km)

Figure 15.4 The adiabatic shock structure around a 3,000 km radius protoplanet that is roughly half the mass of Mars. The simulation, run using the methods of Boley et al. (2013) and Mann et al. (2016) assumes that the relative wind speed Vw is 7 km/s and that the gas is adiabatic (no cooling or heating sources, e.g., radiation). The equation of state for the gas uses a mixture of molecular hydrogen/hydrogen, helium, and metals. The rotational and vibrational modes for H2 are included for an ortho:para mixture of 3:1, and H2 dissociation is also included. The preshock wind is assumed to be at a temperature T = 300 K and to have a pre-shock gas density ρ =10–9 g/cc. The simulation has a resolution of 37.4 km. (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.)

postshock region. This region is the minimum standoff distance Δ for the bow shock (i.e., the distance between the planetesimal’s surface and the shock front). The physical barrier of the planetesimal forces the gas upward and around the planetesimal. Solids that flow into this region and survive the shock (i.e., do not vaporize) must couple with the flow or become accreted by the planetesimal. Thus, for any given Δ, there is a gas drag stopping distance Ds that roughly corresponds to a particle size for the local gas conditions. Particles with Ds > Δ will be accreted. In the case of the protoplanet in Figure 15.4, Δ ~ 400 km, consistent with general expectations ρ 1 for bow shock models in which Δ ρ R (Verigin et al., 2003), where R is the effective radius of e 2 the obstruction. Planetesimals will have very small Δ values as a result of this general scaling. The stopping distance for ~0.3-mm particles in this bow shock is comparable to ~200 km, 388 Melissa A. Morris and Aaron C. Boley suggesting that planetesimals and protoplanets greater than about 1,000 km in radius are necessary to avoid incoming chondrule-like solids from striking the planetesimal/protoplanet directly. The main high-temperature region of the bow shock occurs in the standoff region and extends to impact radii b ~ R. At larger impact radii, the shock quickly becomes oblique, which results in low postshock temperatures and densities. An additional region of high temperature can exist immediately behind the obstruction, but this gas is rarefied. Secondary shocks can also be present where the gas flow begins to curve downward on the trailing side of the obstruction (see Boley et al. (2013) for details). This creates the potential for secondary heating events for any solids that flow through these structures, although temperatures do not approach the peak temperatures seen in the initial shock front. Overall, the bow shock environment is drastically different from 1-D shock assumptions, and there is no guarantee that such 1-D calculations can accurately capture the shock’s effect on solids. To emphasize this point further, consider again the large protoplanet in Figure 15.4. Using the methods described in Boley et al. (2013), we track 0.3-mm particles through the adiabatic bow shock. Figure 15.5 shows the resulting trajectories and Figure 15.6 shows the corresponding temperature, density, and pressure conditions the particles experience. After about 12–20 minutes of simulation time, depending on the impact radius, the particles first encounter the bow shock. At low impact radius, the particles enter a postshock region with a sevenfold increase in gas density, a gas temperature of ~2,200 K, and a pressure just above 0.5 mbar (50 times the preshock pressure). Some of these particles are, nonetheless, accreted. As the impact radius increases, the shock decreases in strength, as expected from the bow shock morphology. Particles typically see a very rapid (seconds) rise in pressure, temperature, and density, followed by decreases that last about 20 minutes for the given bow shock. Smaller protoplanets/planetesimals will have an even faster return to lower temperature, pressure, and

Adiabatic, 7 km/s Wind 100 14000

12000 80

10000 bar 60 µ 8000

6000 40 Distance (km)

4000 Gas Pressure 20 2000

0 0 -15000 -10000 -5000 0 5000 10000 15000 Distance (km)

Figure 15.5 Similar to Figure 15.4, but for the gas pressure with particle trajectories superposed to illustrate the flow of solids through the bow shock. Solids penetrate into the postshock region before recoupling to the gas flow. Most solids are diverted around the protoplanet, but the ones with the smallest impact angles are accreted. The slight change in the trajectories just before the shock encounter is due to the gravity of the protoplanet. (A black-and-white version of this figure will appear in some formats. For the colour version, please refer to the plate section.) Formation of Chondrules by Shock Waves 389

2500

2000

1500

1000 Temperature (K) 500

0 0 102030405060708090 Time (minutes) 8

7

6

5

4

3

2 Density Enhancement 1

0 0 102030405060708090 Time (minutes) 1000 bar) µ 100 Pressure (

10 0 10 20 30 40 50 60 70 80 90 Time (minutes)

Figure 15.6 Gas temperature (top), density (middle), and pressure (bottom) conditions experienced by chondrule-sized particles as they pass through the bow shock. Curves that end abruptly were accreted by the protoplanet. The profiles correspond to the trajectories shown in Figure 15.5. A variety of conditions can be experienced by the dust depending on the impact radius. These profiles are for an adiabatic simulation, with the changing conditions due solely to the gas dynamics of the bow shock. density conditions. Temperatures tend to remain above the preshock values far downstream of the shock front, but all particles drop below gas temperatures of 1,500 K. All particles penetrate well into the postshock region before recoupling to the gas flow. Particles with b ~ 0 are accreted by the protoplanet. 390 Melissa A. Morris and Aaron C. Boley 15.8.4 Further Complexities

So far, we have only considered bow shocks in an adiabatic gas. As discussed in the 1-D shock model, chondrule precursors acquire their temperatures though a combination of processes, including thermal contact with the gas, frictional gas-drag heating, and importantly, radiation. Consider first an approximate gray dust opacity consistent with the approximate solar con- densable mass fraction of dust inside the water ice line (Lodders, 2003). Assuming that the opacity is dominated by μmto10μm-sized grains, the opacity will be approximately between κd ~1and 2 –1 10 cm g , depending on the actual abundance of small grains. Using the higher value for κd and the shock conditions found in Figure 15.4, the optical depth through the standoff distance is Δτ ~ 3, assuming that the grains are not vaporized. While this is not an optically thin environment, it is still not optically thick enough to stave off significant energy loss from the shock region. Suppose, instead, that small grains are not present and that the optical depth is dominated by 2 –1 chondrules of sizes s ~ 300 μm. In this case, the effective opacity is only κd ~ 0.04 cm g , assuming a solar dust-to-gas mass ratio (or chondrule-to-gas ratio) of 0.005. In this case, the optical depth through the standoff region is optically thin, and only approaches unity if the midplane is enhanced with chondrules by a factor of 100 or more. Regardless of whether the standoff region is optically thick, chondrules will radiate away energy efficiently as they flow past the obstacle, resulting in very fast cooling times (Boley et al., 2013; Mann et al., 2016). Only if the gas has a very high dust opacity from small grains (greater than 10 times the nominal dust opacities from, e.g., Pollack et al., 1994) do we expect the bow shock to converge to the adiabatic limit for the large protoplanet bow shocks (Mann et al., 2016). For bow shocks around planetesimals, the effective opacity must be even higher, as the size of the shock scales with the planetesimal’s size. This in turn scales with the optical depth through the shock. Mann et al. (2016) investigated thermal profiles for bow shocks around protoplanets (as in Figure 15.4) under adiabatic and radiative conditions. They further explored the consequences for different assumptions about the dust and chondrule opacities. As shown in Figure 15.7, many radiative conditions lead to very efficient cooling with cooling rates that are potentially much too fast to be consistent with chondrule furnace experiments. However, if the temperature, density, and pressure profiles approach those found for the adiabatic conditions (such as in an extremely dusty environment), then the protoplanet bow shocks could be consistent with chondrule cooling constraints. Finally, we note that large enough planetesimals and protoplanets are expected to have primitive atmospheres, which may lead to high mixing ratios of volatiles in the postshock regions (Morris et al., 2012). This could be advantageous for explaining the high partial pressures of species such as, for example, Na and K. However, Mann et al. (2016) showed that primitive atmospheres would be efficiently stripped from even protoplanets due to the ram pressure with the gas. There nonetheless remains the possibility that, for certain orbits, an atmosphere can be formed through accretion of nebular gas and some outgassing during low Vw phases, and then stripped during phases of high Vw.

15.8.5 Feasibility of Bow Shocks

As with planetesimal collision models, an immediate implication of the bow shock model requires that significant planet building predates, and is contemporaneous with, the formation of Formation of Chondrules by Shock Waves 391

Figure 15.7 The vertical bars show allowed chondrule cooling rates, for different chondrule textures, based on furnace experiments. The horizontal bars show cooling rates derived from protoplanet bow shock simulations for different assumptions about the radiative cooling. The adiabatic cases are consistent with many chondrule textures, but these require an extreme optically thick environment. The other bars labelled “Rad X, k=y” represent different assumptions for the dust and chondrule opacities (see Mann et al. (2016) for details). Radiative simulations for “normal” opacities show cooling rates that are much larger than the furnace experiment constraints. While one radiative simulation does show some potential overlap with the experiments, the cooling rate might only be attainable for unrealistic thermal coupling between the gas and the melts. Reproduced by permission of the AAS. (A black-and- white version of this figure will appear in some formats. For the colour version, please refer to the plate section.)

chondrules found in chondrites. Independent evidence suggests that this may be reasonable. For example, Hf-W dating of iron meteorites suggests that their parent bodies formed within 1.5 Myr of the bulk of CAI formation (Scherstén et al., 2006), and likely much earlier (0.1–0.3 Myr; Kruijer et al., 2014), which would predate the formation of most chondrules, at least based on Al–Mg relative age dating (Villeneuve et al., 2009). Furthermore, systems such as HL Tau (ALMA Partnership) suggest that planet formation may occur very early (<1 Myr) in the lifetime of at least some discs, which would be on timescales that are fast compared with at least the solar system’s meteoritic record. The bow shock model is critically dependent on planetesimals and/or protoplanets being placed on excited orbits. This requires that massive bodies or large potential fluctuations from spiral arms are already present at the time of chondrule formation. A proto-Jupiter would be capable of exciting such orbits (Weidenschilling et al., 1998), in which resonant interactions can increase the eccentricities of planetesimals and protoplanets to high values in 100,000 years (Hood and Weidenschilling, 2012). 392 Melissa A. Morris and Aaron C. Boley 15.8.6 Processed Mass

We can assess how much shock-processed material could be produced by a single object placed on a high-eccentricity orbit under the bow shock model, provided that such processing leads to chondrule formation. As discussed in Morris et al. (2012), planetesimals will have their eccentricity and inclination most effectively damped through aerodynamic drag, while protoplanets of about the mass of Mars and larger will have their orbital elements damped most effectively by disc torques. The actual damping timescale depends on the mass of the planetesimal/protoplanet, the orbital location in the disc, and the disc mass and structure. For a Mars-size embryo, damping timescales can range from about a few tens of thousands of years at 1 AU to about a million years at 3 AU in a minimum-mass solar nebula (Morris et al., 2012). In contrast, a 100 km planetesimal at 1AU in a gas with a density of 109 g/cc will have a damping time on the order of 2,000 years, scaling linearly with the size of the planetesimal. Detailed estimates for the mass produced in any single scattered embryo are given in Morris et al. (2012). For a rough estimate, consider a nebular model such that the surface density of – solids is σ =1gcm 2(5.2 AU/r)1.5, where r is the local disc radial distance. Assuming the dust mass is mainly in chondrule precursors and that the dust scale height is small due to settling, the excited planetesimal/protoplanet will pass through the dust disc twice per orbit. Thus, the total mass processed in the bow shock will be M ~2σπs2η, where η is the number of orbits during which the planetesimal creates chondrule-producing bow shocks. At about 1.5 AU with a gas midplane density of e4 10 10 g/cc, the damping timescale can be approximated as about 8 5; 000 yr for an se100 km planetesimal. As such, the planetesimal could produce e10 M of chondrules for a single scattering event. For a single scattering event of a planetesimal with 6 se1; 000 km, a mass of 10 M of dust could be processed. The bow shock model, therefore, relies on many planetesimals (but can potentially be a small fraction overall) being placed onto excited orbits throughout the lifetime of the solar nebula.

15.9 Conclusions

We find that models of large-scale shocks, such as those driven by GIs, are most consistent with the first-order (zeroth-order) constraint on cooling rates of chondrules, as well as most of the nonthermal constraints. Cooling rates predicted by models of large-scale shocks range from 10–40 K/hr for solids-to-gas ratios of 10 solar (at P ~10–4 to 10–3 atm) and are faster (~50–400 K/hr) for higher concentrations (up to 300 solar). Predictions are therefore consistent with most of the meteoritic constraints on thermal histories (see Figure 15.7). However, large-scale shocks have yet to be modeled with the high densities of solids thought to be necessary for volatile retention, and have not accounted for vertical energy losses. It is also important to note again that the GI-driven shocks are more likely early in the evolution of the solar nebula, becoming much less likely over time, so may not be a feasible formation mechanism for the youngest porphyritic chondrules. Bow shocks can occur in an environment that is conducive to volatile retention if an outgassed atmosphere is present, over a magma ocean, during chondrule heating and cooling. Such primitive atmospheres are likely to be stripped rapidly (see Mann et al., 2016), but may Formation of Chondrules by Shock Waves 393 persist for some phases of the protoplanet’s orbit. In the presence of dusty, high-mass atmos- pheres, cooling rates are predicted at the very upper end of (or above) the range of the inferred thermal histories of chondrules. Without such a dusty, high-mass atmostphere, cooling rates are predicted to greatly exceed those thought possible for porphyritic chondrules. Both GI-driven shocks and bow shocks could have occurred – in fact, are likely to have occurred – in the solar nebula. Not only can these types of shocks produce chondrules under the right set of circumstances, they also predict formation of solids that would not be recognized as chondrules. Shocks weaker than 8 km/s would predict peak temperatures below those inferred for chondrules, and would result in products of incomplete melting. Particles entering the periphery of a bow shock would experience effective shock speeds much lower than those hitting head-on, so would make similar predictions to those of weak shocks. In the case of bow shocks absent a dusty, high-mass atmosphere, models predict extremely rapid cooling and complicated thermal histories, including secondary heating. We encourage sample analyses to test these predictions, while acknowledging that although non–chondrule-producing shocks may have been much more frequent than chondrule-producing shocks, materials that would have preserved this record may be in low abundance in meteorites, due to, for example, size sorting during chondrite formation. We also ask “How constraining are the constraints?” How accurate are the ages of chon- drules found through radiometric dating using the various isotopic systems – which do not always agree between systems? Are the samples that are dated appropriate for such dating? How firm are the constraints on thermal histories? Many analyses indicating the necessity for high vapor pressures of moderately volatile species have not been duplicated. Do these constraints on high vapor pressures apply to all chondrules? Might there be a correlation with age or type? Does chondrule/matrix complementarity provide any constraints on chondrule formation? If so, can this condition be met by more than one proposed mechanism? The resolution of these questions is certainly needed in order to make significant progress in understanding the conditions of chondrule formation. Finally, perhaps we must consider that there was more than one formation mechanism for (even porphyritic) chondrules. Perhaps chondrule formation evolved temporally as the proto- planetary disk evolved, beginning with GI-driven shocks in the young, massive disk producing the oldest chondrules, bow shocks producing chondrules during planet migration, and at least one impact late in the disk, producing the components of the metal-rich chondrites (a similar conclusion was reached by Connelly and Bizzarro; see Chapter 11). Perhaps we should begin to examine chondrules in this context, determining which formation mechanism is most consistent with particular chondrules.

Acknowledgments

We thank Sara Russell, Sasha Krot, and Harold Connolly for the invitation to participate in this publication and the associated workshop. We also thank reviewers Conel Alexander and Alexander Hubbard for comments and suggestions that have greatly improved the chapter. MAM was supported by NASA Cosmochemistry Grant NNX14AN58G and NASA Emerging Worlds Grant NNX15AH62G. 394 Melissa A. Morris and Aaron C. Boley

References

Adachi, I., Hayashi, C., and Nakazawa, K. (1976). The gas drag effect on the elliptical motion of a solid body in the primordial solar nebula. Progress of Theoretical Physics, 56, 1756–1771. Alexander, C. M. O’D., Boss, A. P., and Carlson, R. W. (2001). The early evolution of the inner Solar System: a meteoritic perspective. Science, 293,64–69. Alexander, C. M. O’D. (2005). From supernovae to planets: the view from meteorites and interplanetary dust particles. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, Astronomical Society of the Pacific Conference Series 341, 972. (San Francisco, CA: Astronomical Society of the Pacific. Alexander, C. M. O.’D., Grossman, J. N., Ebel, D. S., and Ciesla, F. J. (2008). The formation conditions of chondrules and chondrites. Science, 320, 1617–1619. Alexander, C. M. O’D., and Ebel, D. S. (2012). Questions, questions: Can the contradictions between the petrologic, isotopic, thermodynamic, and astrophysical constraints on chon- drule formation be resolved? Meteoritics and Planetary Science, 47, 1157–1175. Amelin, Y., Krot, A. N., Hutcheon, I. D., and Ulyanov, A. A. (2002). Lead Isotopic Ages of Chondrules and Calcium-Aluminum-Rich Inclusions. Science, 297, 1678–1683. Amelin, Y., Kaltenbach, A., Iizuka, T., et al. (2010). U-Pb chronology of the Solar System’s oldest solids with variable 238U/235U. Earth and Planetary Science Letters, 300, 343–350. Boley, A. C., and Durisen, R. H. (2008). Gravitational instabilities, chondrule formation, and the FU Orionis phenomenon. The Astrophysical Journal, 685, 1193–1209. Boley, A. C., Morris, M. A., & Desch, S. J. (2013). High-temperature processing of solids through solar nebular bow shocks: 3D radiation hydrodynamics simulations with particles. The Astrophysical Journal, 776, 101–124. Bollard, J., Connelly, J. N., and Bizzarro, M. (2014). The Absolute Chronology of the Early Solar System Revisited. 77th Annual Meeting of the Meteoritical Society, 1800, 5234. Boss, A. P., and Durisen, R. H. (2005). Chondrule-forming shock fronts in the solar nebula: a possible unified scenario for planet and chondrite formation. The Astrophysical Journal, 621, L137-L140. Budde, G., Kleine, T., Kruijer, T. S., Burkhardt, C., and Metzler, K. (2016). Tungsten isotopic constraints on the age and origin of chondrules. Proceedings of the National Academy of Sciences, 113, 2886–2891. Ciesla,F.J.(2005).Chondrule-forming Processes – An Overview. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, Astronomical Society of the Pacific Conference Series, 341, 811–820. (San Francisco, CA: Astronomical Society of the Pacific). Connolly, H. C., Jr., and Love, S. G. (1998). The formation of chondrules: Petrologic tests of the shock wave model. Science, 280,62–67. Connolly, H. C., Jr., Jones, B. D., and Hewins, R. H. (1998). The flash melting of chondrules: An experimental investigation into the melting history and physical nature of chondrule precursors. Geochimica et Cosmochimica Acta, 62, 2725–2735. Connolly, H. C., Jr., and Desch, S. J. (2004). On the origin of the “kleine Kügelchen” called Chondrules. Chemie der Erde / Geochemistry, 64,95–125. Connolly, H. C., Jr., Desch, S. J., Ash, R. D., and Jones, R. H. (2006). Transient heating events in the protoplanetary nebula. In D. S. Lauretta and H. Y. McSween Jr. (Eds.), Meteorites and the Early Solar System II, 383–397. Tucson, AZ: University of Arizona Press. Connelly, J. N., Amelin, Y., Krot, A. N., and Bizzarro, M. (2008). Chronology of the Solar System’s Oldest Solids. The Astrophysical Journal Letters, 675, L121–L124. Connelly, J. N., Bizzarro, M., Krot, A. N., et al. (2012). The absolute chronology and thermal processing of solids in the solar protoplanetary disk. Science, 338, 651–655. Connolly, H. C., and Jones, R. H. (2016). Chondrules: The canonical and noncanonical views. Journal of Geophysical Research (Planets), 121, 885–1899. Formation of Chondrules by Shock Waves 395

Cuzzi, J. N., and Alexander, C. M. O’D. (2006). Chondrule formation in particle-rich nebular regions at least hundreds of kilometres across. Nature, 441, 483–485. Cuzzi, J. N., Hogan, R. C., and Shariff, K. (2008). Toward planetesimals: Dense chondrule clumps in the protoplanetary nebula. The Astrophysical Journal, 687, 1432–1447. Desch, S. J., and Connolly, H. C., Jr.(2002). A model of the thermal processing of particles in solar nebula shocks: Application to the cooling rates of chondrules. Meteoritics and Planetary Science, 37, 183–207. Desch, S. J., and Cuzzi, J. N. (2000). The generation of lightning in the solar nebula. Icarus, 143,87–105. Desch, S. J. (2007). Mass distribution and planet formation in the solar nebula. The Astrophys- ical Journal, 671, 878–893. Desch, S. J., Morris, M. A., Connolly, H. C., and Boss, A. P. (2012). The importance of experiments: Constraints on chondrule formation models. Meteoritics and Planetary Science, 47, 1139–1156. Durisen, R. H., Boss, A. P., Mayer, L., et al. (2007). Gravitational instabilities in gaseous protoplanetary disks and implications for giant planet formation. In B. Reipurth, D. Jewitt, and K. Keil (Eds.), Protostars and Planets V, 607–622. Tucson, AZ: University of Arizona Press. Ebel, D. S., and Grossman, L. (2000). Condensation in dust-enriched systems. Geochimica et Cosmochimica Acta, 64, 339–366. Farrington, O. C. (1915). Meteorites, their structure, composition, and terrestrial relations. Chicago, IL: Lakeside Press. Fedkin, A. V., and Grossman, L. (2013). Vapor saturation of sodium: Key to unlocking the origin of chondrules. Geochimica et Cosmochimica Acta, 112, 226–250. Fedkin, A. V., and Grossman, L. (2016). Effects of dust enrichment on oxygen fugacity of cosmic gases. Meteoritics and Planetary Science, 51, 843–850. Gammie, C. F. (1996). Linear Theory of magnetized, viscous, self-gravitating gas disks. The Astrophysical Journal, 462, 725–731. Gooding, J. L., and Keil, K. (1981). Relative abundances of chondrule primary textural types in ordinary chondrites and their bearing on conditions of chondrule formation. Meteoritics, 16,17–43. Grossman, L., Fedkin, A. V., and Simon, S. B. (2012). Formation of the first oxidized iron in the solar system. Meteoritics and Planetary Science, 47, 2160–2169. Haisch, K. E., Jr., Lada, E. A., and Lada, C. J. (2001). Disk frequencies and lifetimes in young clusters. The Astrophysical Journal Letters, 553, L153-L156. Hayashi, C. (1981). Structure of the solar nebula, growth and decay of magnetic fields and effects of magnetic and turbulent viscosities on the nebula. Progress of Theoretical Physics Supplement, 70,35–53. Helled, R., Bodenheimer, P., Podolak, M., et al. (2014). Giant planet formation, evolution, and internal structure. In H. Beuther, R. S. Klessen, C. P. Dullemond, and T. Henning (Eds.), Protostars and Planets VI, 643–665. Tucson, AZ: University of Arizona Press. Hewins, R. H., and Connolly, H. C., Jr. (1996). Peak temperatures of flash-melted chondrules. In R. H. Hewins, E. Scott, and R. Jones (Eds.), Chondrules and the Protoplanetary Disk, 197–204. Cambridge, UK: Cambridge University Press. Hewins, R. H. (1997). Chondrules. Annual Review of Earth and Planetary Sciences, 25,61–83. Hewins, R. H., Connolly, H. C., Lofgren, G. E., Jr., and Libourel, G. (2005). Experimental Constraints on Chondrule Formation. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, Astronomical Society of the Pacific Conference Series, 341, 286–316. San Francisco, CA: Astronomical Society of the Pacific. Hewins, R. H., Zanda, B., and Bendersky, C. (2012). Evaporation and recondensation of sodium in Semarkona Type II chondrules. Geochimica et Cosmochimica Acta, 78,1–17. 396 Melissa A. Morris and Aaron C. Boley

Hezel, D. C., Palme, H., Brenker, F. E., and Nasdala, L. (2003). Evidence for fractional condensation and reprocessing at high temperatures in CH chondrites. Meteoritics and Planetary Science, 38, 1199. Hezel, D. C., and Palme, H. (2008). Constraints for chondrule formation from Ca-Al distribu- tion in carbonaceous chondrites. Earth and Planetary Science Letters, 265, 716–725. Hood, L. L. (1998). Thermal processing of chondrule and CAI precursors in planetesimal bow shocks. Meteoritics and Planetary Science, 33,97–107. Hood, L. L., Ciesla, F. J., Artemieva, N. A., Marzari, F., and Weidenschilling, S. J. (2009). Nebular shock waves generated by planetesimals passing through Jovian resonances: Possible sites for chondrule formation. Meteoritics and Planetary Science, 44, 327–342. Hood, L. L., and Weidenschilling, S. J. (2012). The planetesimal bow shock model for chondrule formation: A more quantitative assessment of the standard (fixed Jupiter) case. Meteoritics and Planetary Science, 47, 1715–1727. Jones, R. H., Lee, T., Connolly, H. C., Jr., Love, S. G., and Shang, H. (2000). Formation of chondrules and CAIs: Theory vs. observation. In V. Mannings, A. P. Boss, and S. S. Russell (Eds.), Protostars and Planets IV, 927–962. Tucson, AZ: University of Arizona Press. Kita, N. T., Huss, G. R., Tachibana, S., et al. (2005). constraints on the origin of chondrules and CAIs from short-lived and long-lived radionuclides. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, Astronomical Society of the Pacific Conference Series, 341, 558–587. San Francisco, CA: Astronomical Society of the Pacific. Kita, N. T., and Ushikubo, T. (2012). Evolution of protoplanetary disk inferred from 26Al chronology of individual chondrules. Meteoritics and Planetary Science, 47, 1108–1119. Klerner, S., and Palme, H. (1999). Origin of Chondrules and Matrix in Carbonaceous Chon- drites. Lunar and Planetary Science Conference, 30, 1272. Krot, A. N., Fegley, B., Jr., Lodders, K., and Palme, H. (2000). Meteoritical and astrophysical constraints on the oxidation state of the solar nebula. In V. Mannings, A. P. Boss, and S. S. Russell (Eds.), Protostars and Planets IV, 1019–1054. Tucson, AZ: University of Arizona Press. Krot, A. N., Amelin, Y., Cassen, P., and Meibom, A. (2005). Young chondrules in CB chondrites from a giant impact in the early Solar System. Nature, 436, 989–992. Kruijer, T. S., Kleine, T., Fischer-Gödde, M., Burkhardt, C., and Wieler, R. (2014). Nucleosyn- thetic W isotope anomalies and the Hf-W chronometry of Ca-Al-rich inclusions. Earth and Planetary Science Letters, 403, 317–327. Kurahashi, E., Kita, N. T., Nagahara, H., and Morishita, Y. (2008). 26Al-26Mg systematics of chondrules in a primitive CO chondrite. Geochimica et Cosmochimica Acta, 72, 3865. Lauretta, D. S., Buseck, P. R., and Zega, T. J. (2001). Opaque minerals in the matrix of the Bishunpur (LL3.1) chondrite: Constraints on the chondrule formation environment. Geo- chimica et Cosmochimica Acta, 65, 1337–1353. Lauretta, D. S., Nagahara, H., and Alexander, C. M. O’D. (2006). Petrology and origin of ferromagnesian silicate chondrules. In D. S. Lauretta and H. Y. McSween Jr. (Eds.), Meteorites and the Early Solar System II, 431–459. Tucson, AZ: University of Arizona Press. Lodders, K. (2003). Solar system abundances and condensation temperatures of the elements. The Astrophysical Journal, 591, 1220–1247. Lofgren, G. E. (1982). The importance of heterogeneous nucleation for the formation of microporphyritic chondrules. Chrondrules and their Origins, 493, 41. Lofgren, G. (1989). Dynamic crystallization of chondrule melts of porphyritic olivine compos- ition: Textures experimental and natural. Geochimica et Cosmochimica Acta, 53, 461–470. Lofgren, G., and Lanier, A. B. (1990). Dynamic crystallization study of barred olivine chon- drules. Geochimica et Cosmochimica Acta, 54, 3537–3551. Formation of Chondrules by Shock Waves 397

Lofgren, G. E. (1996). A dynamic crystallization model for chondrule melts. In R. H. Hewins, E. Scott, and R. Jones (Eds.), Chondrules and the Protoplanetary Disk, 187–196. Cam- bridge, UK: Cambridge University Press. Mann, C. R., Boley, A. C., and Morris, M. A. (2016). Planetary embryo bow shocks as a mechanism for chondrule formation. The Astrophysical Journal, 818, 103–123. Mihalas, D. (1978). Stellar atmospheres, 2nd edition. San Francisco, CA: W. H. Freeman and Co. Mihalas, D., and Mihalas, B. W. (1984). Foundations of Radiation Hydrodynamics. New York, NY: Oxford University Press. Miura, H., and Yamamoto, T. (2014). A new estimate of the chondrule cooling rate deduced from an analysis of compositional zoning of relict olivine. The Astronomical Journal, 147, 54–63. Morris, M. A., Boley, A. C., Desch, S. J., and Athanassiadou, T. (2012). Chondrule formation in bow shocks around eccentric planetary embryos. The Astrophysical Journal, 752,27–44. Morris, M. A., and Desch, S. J. (2010). Thermal histories of chondrules in solar nebula shocks. The Astrophysical Journal, 722, 1474–1494. Morris, M. A., Desch, S. J., and Ciesla, F. J. (2009). Cooling of dense gas by H2O line emission and an assessment of its effects in chondrule-forming shocks. The Astrophysical Journal, 691, 320–331. Morris, M. A., Garvie, L. A. J., and Knauth, L. P. (2015). New insight into the solar system’s transition disk phase provided by the metal-rich carbonaceous chondrite isheyevo. The Astrophysical Journal, 801, L22–L27. Morris, M. A., Weidenschilling, S. J., and Desch, S. J. (2016). The effect of multiple particle sizes on cooling rates of chondrules produced in large-scale shocks in the solar nebula. Meteoritics and Planetary Science, 51, 870–883. Nakamoto, T., Hayashi, M. R., Kita, N. T., and Tachibana, S. (2005). Chondrule-forming shock waves in the solar nebula by x-ray flares. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk, Astronomical Society of the Pacific Conference Series, 341, 883–892. San Francisco, CA: Astronomical Society of the Pacific. Nelson, A. F., and Ruffert, M. (2005). A proposed origin for chondrule-forming shocks in the solar nebula. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk. Astronomical Society of the Pacific Conference Series, 341, 903–912. San Francisco, CA: Astronomical Society of the Pacific. Palme, H., Spettel, B., and Ikeda, Y. (1993). Origin of chondrules and matrix in carbonaceous chondrites. Meteoritics, 28, 417. Pollack, J. B., Hollenbach, D., Beckwith, S., et al. (1994). Composition and radiative properties of grains in molecular clouds and accretion disks. The Astrophysical Journal, 421, 615–639. Radomsky, P. M., and Hewins, R. H. (1990). Formation conditions of pyroxene-olivine and magnesian olivine chondrules. Geochimica et Cosmochimica Acta, 54, 3475–3490. Raymond, S. N., Kokubo, E., Morbidelli, A., Morishima, R., and Walsh, K. J. (2014). Terrestrial planet formation at home and abroad. In H. Beuther, R. S. Klessen, C. P. Dullemond, and T. Henning (Eds.), Protostars and Planets VI, 595–618. Tucson, AZ: University of Arizona Press. Rubin, A. E., Sailer, A. L., and Wasson, J. T. (1999). Troilite in the chondrules of type-3 ordinary chondrites: Implications for chondrule formation. Geochimica et Cosmochimica Acta, 63, 2281–2298. Rubin, A. E., Kallemeyn, G. W., Wasson, J. T., et al. (2003). Troilite in the chondrules of type-3 ordinary chondrites: Implications for chondrule formation. Geochimica et Cosmochimica Acta, 67, 3283. Russell, S. S., Hartmann, L., Cuzzi, J., et al. (2006). Timescales of the solar protoplanetary disk. In D. S. Lauretta and H. Y. McSween Jr. (Eds.), Meteorites and the Early Solar System II, 233–251. Tucson, AZ: University of Arizona Press. 398 Melissa A. Morris and Aaron C. Boley

Sanders, I. S., and Scott, E. R. D. (2012). The origin of chondrules and chondrites: Debris from low-velocity impacts between molten planetesimals? Meteoritics and Planetary Science, 47, 2170–2192. Scherstén, A., Elliott, T., Hawkesworth, C., Russell, S., and Masarik, J. (2006). Hf W evidence for rapid differentiation of iron meteorite parent bodies. Earth and Planetary Science Letters, 241, 530–542. Schrader, D. L., Connolly, H. C., Lauretta, D. S., et al. (2013). The formation and alteration of the Renazzo-like carbonaceous chondrites II: Linking O-isotope composition and oxida- tion state of chondrule olivine. Geochimica et Cosmochimica Acta, 101, 302–327. Schrader, D. L., Nagashima, K., Krot, A. N., et al. (2017). Distribution of 26Al in the CR chondrite chondrule-forming region of the protoplanetary disk. Geochimica et Cosmochi- mica Acta, 201, 275–302. Schoelmerich, M. O., Seitz, H.-M., and Klimm, K. (2016). Evaporational loss of lithium during high temperature experiments: Implications for chondrule formation. Lunar and Planetary Science Conference XLVII, abstract #1461. Smith, K. T. (2016). Spiral arms in a disk around a young star. Science, 353, 1509–1511. Stammler, S. M., and Dullemond, C. P. (2014). A critical analysis of shock models for chondrule formation. Icarus, 242,1–10. Shu, F. H., Shang, H., and Lee, T. (1996). Toward an astrophysical theory of chondrites. Science, 271, 1545–1552. Tachibana, S., Huss, G. R., Miura, H., and Nakamoto, T. (2004). Evaporation and accompany- ing isotopic fractionation of sulfur from Fe-S melt during shock wave heating. Lunar and Planetary Science Conference, 35, 1549. Tachibana, S., and Huss, G. R. (2005). Sulfur isotope composition of putative primary troilite in chondrules from Bishunpur and Semarkona. Geochimica et Cosmochimica Acta, 69, 3075–3097. Verigin, M., Slavin, J., Szabo, A., et al. (2003). Planetary bow shocks: Gasdynamic analytic approach. Journal of Geophysical Research, 108, 1323. Villeneuve, J., Chaussidon, M., and Libourel, G. (2009). Homogeneous distribution of 26Al in the solar system from the Mg isotopic composition of chondrules. Science, 325, 985. Villeneuve, J., Libourel, G., and Soulié, C. (2015). Relationships between type I and type II chondrules: Implications on chondrule formation processes. Geochimica et Cosmochimica Acta, 160, 277–305. Wadhwa, M., Amelin, Y., Davis, A. M., et al. (2007). From dust to planetesimals: implications for the solar protoplanetary disk from short-lived radionuclides. In V. B. Reipurth, D. Jewitt, and K. Keil (Eds.), Protostars and Planets V, 835–848. Tucson, AZ: University of Arizona Press. Walsh, K. J., Morbidelli, A., Raymond, S. N., O’Brien, D. P., and Mandell, A. M. (2011). A low mass for Mars from Jupiter’s early gas-driven migration. Nature, 475, 206–209. Weidenschilling, S. J. (1977). The distribution of mass in the planetary system and solar nebula. Astrophysics and Space Science, 51, 153–158. Weidenschilling, S. J. (1980). Dust to planetesimals: Settling and coagulation in the solar nebula. Icarus, 44, 172–189. Weidenschilling, S. J., Marzari, F., and Hood, L. L. (1998). The Origin of Chondrules at Jovian Resonances. Science, 279, 681–684. Wick, M. J., and Jones, R. H. (2012). Formation conditions of plagioclase-bearing type I chondrules in CO chondrites: A study of natural samples and experimental analogs. Geochimica et Cosmochimica Acta, 98, 140–159. Williams, J. P., and Cieza, L. A. (2011). Protoplanetary disks and their evolution. Annual Review of Astronomy and Astrophysics, 49,67–117. Formation of Chondrules by Shock Waves 399

Wood, J. A. (1985). Meteoritic constraints on processes in the solar nebula. In D. C. Black and M. S. Mathews, Protostars and Planets II, 687–702. Tucson, AZ: University of Arizona Press. Yu, Y., and Hewins, R. H. (1998). Transient heating and chondrite formation: Evidence from sodium loss in flash heating simulation experiments. Geochimica et Cosmochimica Acta, 62, 159–172. Zanda, B., Le Guillou, C., and Hewins, R. H. (2009). The relationship between chondrules and matrix in chondrites. Meteoritics and Planetary Science Supplement, 72, 5280. 16 Evaluating Non-Shock, Non-Collisional Models for Chondrule Formation

alexander hubbard and denton s. ebel

Abstract

Chondrule formation and meteorite parent body assembly link the historically geochemically- and petrologically-oriented field of meteoritics to the observationally- and theoretically-oriented field of astrophysics. Laboratory measurements’ high precisions and resolutions constrain planet formation on scales and in parameter spaces inaccessible to even the most powerful telescopes. The dynamic and cosmochemical canvas confronting theoretical and numerical studies of proto- planetary disks and planet formation is too daunting to face without meteoritic signposts. Conversely, with only ancient solid evidence in hand, meteoriticists need astrophysical observa- tions and protoplanetary disk models to give context to their witch’s brew of chemical, isotopic, and petrologic constraints. The ever-increasing wealth of protoplanetary disk observations, along with major advances such as the (re)discovery of the Magneto-Rotational Instability (or MRI), numerical simulations of disk dynamics, and laboratory investigations of dust coagulation have expanded our understanding of protoplanetary disks, leaving us well positioned to review proposed chondrule formation scenarios both from the perspective of astrophysical plausibility (i.e., could they occur and produce melted grains at a meaningful rate), and from a cosmochemical and petrologic perspective (i.e., would the melted grains look like chondrules and be incorporated into chondrite-like planetesimals). In this chapter, we evaluate several chondrule formation scenarios for our own solar system, but one inherent truth in astrophysics is that the universe is large enough for almost any conceivable process to find a home. Potential chondrule-formation mechanisms that fail to make the cut here may yet be important elsewhere.

16.1 Introduction

Chondrules and chondrites pose some of most perplexing meteoritic challenges to the study of the solar nebula specifically, and planet formation more generally. Testifying simultaneously to temperatures both above 1,750 K and below 650 K (Huss and Lewis, 1994; Hewins, 1997; Rubin et al., 1999; Schrader et al., 2016), they erase any hope that the solar nebula from which our solar system emerged was simple; and they demonstrate that the widespread, effectively laminar α-disk conceptualization of protoplanetary disks (Shakura and Sunyaev, 1973) is at best a useful shorthand for a family of active, evolving objects. Chondrules and chondrites contain clues not only about our own solar system’s formation, but also presumably about the formation of extrasolar planetary systems in general. Indeed, they

400 Evaluating Non-Shock, Non-Collisional Models for Chondrule Formation 401 provide nearly the only hard constraints on the small-scale dynamics of protoplanetary disks relevant to the growth of solids: dust coagulation, turbulent transport, and planetesimal assem- bly. Even the Atacama Large Millimeter Array (or ALMA) can at best observe protoplanetary disks at circa 1 astronomical unit (AU) resolution, tens to hundreds of times the dynamical scales associated with disk evolution and rocky planetesimal formation. Those clues testify to complicated and varied formation conditions, requiring a framework of disk models to put into context. In this chapter, we review several chondrule formation mechanisms from both an astrophysical dynamics perspective and from a cosmochemical and petrologic one to help strengthen the bridge between these two approaches to studying the formation of planets in the solar system. This is hardly the first endeavor to review chondrule formation processes incorporating both astrophysical and meteoritic constraints, and we owe much to previous work (e.g., Boss, 1996; Desch et al., 2012). The combination of immense chondrule and chondrite heterogeneity (Jones, 2012) with the uncertainties that plague chemical and dynamic protoplanetary disk modeling makes it difficult to determine how to weigh the different cosmochemical constraints. It is also not clear whether a single mechanism formed most chondrules, or whether multiple formation processes contrib- uted significantly to the observed chondrule population. If the diversity of chondrule reservoirs is due to differences between formation mechanisms, then those mechanisms need not indi- vidually match all cosmochemical constraints. On the other hand, a single chondrule formation mechanism scenario needed to have threaded many needles.

16.1.1 Chondrule Formation Mechanisms

We leave planetesimal collisions and shocks to their own chapters (Sanders and Scott, Chapter 14; Morris and Boley, Chapter 15). After elaborating on the constraints we will consider in Sections 16.2 and 16.3, we will examine the following mechanisms for chondrule formations in no particular order:

Innermost Disk (Section 16.4): Models such as the X-wind model (Shu et al., 1996) posit that chondrules formed in or above the innermost disk, where the temperature was naturally suffi- ciently elevated to melt chondrules. These models require a mechanism to have transported the newly formed chondrules to colder regions where they could mix with the ambient dust which became matrix material. Mechanisms to transport chondrules include disk winds and photophor- esis (Shu et al., 1994; Wurm et al., 2013; McNally and Hubbard, 2015). Lightning (Section 16.5): Electrical discharges similar to terrestrial lightning in the solar nebula would naturally generate regions of intense heating while leaving nearby regions of the disk relatively unscathed (Gibbard et al., 1997; Pilipp et al., 1998; Desch and Cuzzi, 2000). These regions are small, so lightning can easily satisfy chondrule–matrix comple- mentarity constraints (Hezel et al., Chapter 4) and the need to preserve cold-matrix material (Pilipp et al., 1992). Interestingly, recent work has shown that if lightning occurs over a sufficiently broad swath of an extrasolar protoplanetary disk, the associated electric fields would have observable consequences (Muranushi et al., 2015). Radiant Heating (Section 16.6): Herbst and Greenwood (2016) recently suggested that planetesimals with surface magma oceans radiantly heated free-floating dust passing close 402 Alexander Hubbard and Denton S. Ebel

to their surfaces. Such planetesimals would be outgassing, with consequences for the resulting chondrule compositions, and chondrule precursors were likely sufficiently tightly coupled to the gas flowing around the planetesimals that their aerodynamics determined their trajectory (and hence heating and cooling rates, as well as the amount of outgassed material encountered). This model predicts that chondrule sizes should correlate with their thermal histories and compositions. Localized Turbulent Dissipation (Section 16.7): Protoplanetary disks are broadly accepted to be turbulent in some fashion (Shakura and Sunyaev, 1973), although the specifics of that statement remains the focus of intense research. It seems likely that some variant of the MRI (re)discovered by Balbus and Hawley (1991) is the primary player. Turbulence naturally dissipates its energy in a localized fashion, and if that dissipation is sufficiently localized, it would be able to drive chondrule-forming events. Purely hydrodynamic turbulence is unlikely to be sufficiently super-sonic as to reach chondrule melting temperatures and large-scale shocks are covered elsewhere, so this leaves the dissipation of magnetized turbulence (Joung et al., 2004; McNally et al., 2013, 2014).

16.1.2 Background Disk and Chondrule Models

We assume a Hayashi Minimum Mass Solar Nebula (MMSN; Hayashi, 1981), while acknow- ledging that other MMSN models exist (e.g. Desch, 2007): gas surface densities and tempera- –2 tures of Σg ’ 430 g cm and T ’ 180 K at R = 2.5 AU, which imply midplane gas densities of –10 –3 –3 ρg =10 gcm . We adopt a = 0.25 mm and ρ.=3gcm as characteristic chondrule radii and densities (Friedrich et al., 2015), and use the silicate-to-gas mass ratio ε = 0.005 from Lodders

(2003). The gas mean molecular mass will be approximated as mg = 2.3 amu. Overbars represent normalizing to the above values.

16.2 Meteoritic Constraints

We have selected several primary meteoritic constraints against which we will evaluate chon- drule formation mechanisms: cooling rates of newly formed chondrules, full and partial preservation of isotopically unfractionated moderately volatile potassium and sodium, and the elemental and isotopic complementarity between chondrules and matrix within chondrites. We supplement these primary meteoritical constraints with several secondary ones. One significant thermal history consideration is that significant, but not yet certain, fractions of matrix material remained relatively cold, below ~650 K (Huss and Lewis, 1994; Rubin et al., 1999). We will refer to this component as cold-matrix. Chondrules and matrix, while complementary, none- thless have different compositions (Budde et al., 2016a, 2016b; Ebel et al., 2016). Some set of mechanisms had to have allowed dust families with distinct compositions, and chondrule formation had to have interacted with those families differently, leading to chondrule-destined and matrix-destined grains with (for example) distinct nucleosynthetic W and Mo patterns. We will further discuss how efficiently mechanisms produce chondrules: melting chondrules by heating gas is energetically expensive. While finding an energy reservoir sufficient to produce the chondrules of the modern asteroid belt is straightforward, finding an energy Evaluating Non-Shock, Non-Collisional Models for Chondrule Formation 403 reservoir sufficient for chondrule formation to have been a general process in the solar nebula is significantly less so (King and Pringle, 2010), requiring tapping a reservoir comparable to the gravitational potential energy of the disk itself. While the chondrule-or-no history of the planets in our own solar system has long since been erased, this question is key to understanding if we can generalize meteoritic results to extrasolar planet formation. One key constraint we will not invoke is the required concentration of dust and gas. Cosmochemistry results imply that chondrule-forming events needed to have occurred in regions hugely (by factors of hundreds; Wood, 1963; Ebel, 2006) enriched in condensables above the expected solar ratio (Ebel and Grossman, 2000; Tenner et al., 2015). However, in this chapter, we are attempting to merge the results of cosmochemistry and dynamics. Dust dynamics in turbulent gas flows is reasonably well understood, and dust dynamics models struggle to enrich chondrule-precursor-sized objects, even in the aftermath of a strong shock, by more than a few tens, and only a factor of a few in the absence of a shock (Hubbard et al., 2018, in press). We therefore note that the dust enrichment factors required by current cosmochemistry models are inconsistent with all nebular chondrule formation processes given current dust dynamics models, and leave untangling that troubling observation to future research.

16.2.1 Cooling Rates

We can estimate cooling rates for mechanisms that instantaneously heat the gas, which then collisionally heats the chondrules, which in turn radiate away the energy. Most mechanisms actually inject their energy over a more substantial timescale, and thus these estimates should be thought of as upper bounds to the cooling rates. In particular mechanisms that tap the gravita- tional potential energy of the disk, such as turbulent dissipation or gravitational shocks, should evolve on orbital timescales.

16.2.1.1 Cooling Constraints The rate at which chondrules cooled has been constrained by comparing their textures and chemical zoning to those of chondrule analogs heated and cooled in the laboratory. Desch and Connolly (2002) assembled cooling rate estimates from the literature, finding that the bulk of – – chondrules cooled at rates between 5 K h 1 and 3,000 K h 1 (Hewins et al., 2005), although recent work (e.g., Humayun, 2012) has suggested that at least some chondrules may have cooled slower. As we are primarily interested in explaining the bulk of the chondrule population, we will require that chondrule formation mechanisms manage to overlap the 5–3,000 K h–1 range.

16.2.1.2 Optically Thin Cooling It is instructive to examine the case of optically thin cooling, although it is unlikely to apply to chondrules. We can derive a lower bound on the cooling rate by assuming that chondrules drive all cooling, and that chondrules and gas are thermally coupled. A chondrule at temperature Tc radiates with luminosity:

¼ π 2σ 4: Lc 4 a SBTc (16.1) 404 Alexander Hubbard and Denton S. Ebel

At a chondrule-to-gas mass ratio εc, the heat capacity CV of the corresponding gas (neglecting H2 disassociation energy) is:

3 3 4πa ρ: kB 10πa ρ: kB CV e ffi (16.2) 3εcðÞγ 1 mg 3εc mg where the disk adiabatic index γ ffi 1.4. Thus, the cooling rate is: ε 4 _ ¼ Lc ¼ σ 4 3 c mg ffi  3ε 1ρ:1 Tc 1 T SBTc 4 10 ca Kh (16.3) CV 2:5aρ: kB 1; 750 K In Equation (16.3), we invoked relatively large sub-millimeter sized chondrules typical for LL chondrites, so to get the cooling rate down to 3,000 K h–1 we need low chondrule-to-gas mass ratios, or to continue pumping energy into the system to partially balance cooling. Good thermal coupling between gas and chondrules, which we assumed, is, however, inconsistent with this picture. Balancing chondrule radiation with heating from collisions with gas molecules implies: π γ þ ρ 2 1 gkB Lc ¼ Lcg ¼ a θ Tg Tc vth, (16.4) 2 γ 1 mg where θ ~ 1.7 is a correction coefficient (Jacquet and Thompson, 2014) and vth is the gas thermal speed: sffiffiffiffiffiffiffiffiffiffiffiffi 8kBTg vth ¼ : (16.5) πmg

4 For Tc ~ 1,750, Equation (16.4) implies a gas temperature close to 10 K, by which point the gas opacity can easily surpass that of the chondrules, and radiative heating of the chondrules by gas cannot be neglected. If the gas–chondrule combined system were optically thin, Equation (16.3) would underestimate the chondrule cooling rate because chondrule radiative cooling would completely outpace heating from the gas unless the gas were prohibi- tively hot. If, instead, the chondrule-forming region were optically thin (counting only the chondrules) but optically thick (including the hot gas), then we would find ourselves in the optically thick case.

16.2.1.3 Optically Thick Cooling The opacity of nebular gas-plus-dust mixtures is expected to depend strongly on the tempera- ture, but it is also instructive to perform basic cooling rate calculations assuming a constant opacity χ. Writing ℓ for the chondrule-forming region’s shortest length from center to surface (a thin sheet’s shortest length is its half-thickness), the optical depth is: τ ¼ ρ χℓ c g , (16.6) and the temperature as a function of optical depth from the surface is:

= τ þ 2=3 1 4 TðÞ¼τ Tc, (16.7) τc þ 2=3 Evaluating Non-Shock, Non-Collisional Models for Chondrule Formation 405 where Tc is the temperature at the center. The region cools radiatively with a luminosity-per-unit surface area of:

4 4 4 L ¼ σSBT ðÞffi2=3 σSBTc (16.8) 3τc

Except at the edges, the temperature is only a weak function of τ, so we can approximate T ~ Tc throughout the bulk of the chondrule-forming region, leading to a cooling rate of: 4χm σ T 4 χ T 4 K T_ ffi g SB c ffi 3 Â 104τ 2 : (16.9) c : τ2 c 2 1 ; 7 5 c kB 1cm g 1 750 K h

An optical depth τc ~ 10 corresponds nicely to measured chondrule cooling rates. Optical depths well below τc = 10 would cool too rapidly, and depths above 10 would cool very slowly unless they have opacities higher than generally expected for high temperature nebular material. Given 9 2 –1 R = 2.5 AU, τc ffi 10 implies ℓ ffi 5 Â 10 cm for χ =20cm g . The temperature dependence of the opacity is a significant concern. At R = 2.5 AU, a disk 2 –1 with opacity χ =20cm g would have a maximal optical depth of τc ffi 4000  1. At chondrule-forming temperatures though, dust sublimation and processing drop those numbers to –2 2 –1 χ ~10 cm g , and τc ~ 2 (very approximate, Semenov et al., 2003). Not merely chondrule- forming regions, but even the full disk could become optically thin if hot. Models for chondrule formation struggle with the need to track temperature- (and time-) dependent opacities, an aspect which has not yet been fully explored. This suggests a way to reconcile the long lifetime of gas dynamics with the rapid cooling of chondrules: chondrule-forming regions likely heated while optically thick, but became marginally optically thick as the solids evaporated or compactified, allowing the region to cool on a radiative, rather than dynamic, time. Note that for τc ffi 10 and Tc = 1,750 K, the optical surface of the chondrule-forming region will be at T ffi 890 K, easily hot enough to thermally process cold-matrix material. Accordingly, the observed cold-matrix would need to have been mixed in well after the chondrule formation event.

16.2.2 Sodium and Potassium

Chondrules’ moderately volatile alkali element abundances are difficult to explain. Chondrules vary from approximately undepleted in those elements to significantly, but not totally, depleted (Hewins, 1991). This despite the speed with which Na and K are expected to evaporate from chondrule melts under nebular conditions: very strong depletion within 30 min (Yu et al., 2003). Estimating a recondensation temperature of T ~ 1,000 K (Lodders, 2003), those 30 minutes correspond to a cooling time faster than 1,500 K h–1, near the maximally allowed rate. Na and K evaporation/recondensation sequences are complicated by the lack of any measured potas- sium isotopic fractionation, which means that the potassium cloud surrounding the chondrule melts must have been large enough to avoid Rayleigh fractionation during the cooling time: Cuzzi and Alexander (2006) estimated a minimum size of about 1,000 km. This length is significantly smaller than the ℓ ffi 5 Â 109 cm estimate from cooling rate considerations (Section 16.2.1.3). Thus, chondrule-forming regions’ being large enough to prevent Rayleigh fractionat- ing through the diffusion of gaseous evaporated potassium is not unexpected, but it does raise the question of how depletion occurred in the first place. 406 Alexander Hubbard and Denton S. Ebel

Maintaining Na and K in the undepleted chondrules required either a sufficient density of chondrules, or a sufficiently alkali pre-enriched gas to prevent evaporation, or complete recondensation. Recondensation is complicated by the lack of sodium zoning within chondrules (Alexander et al., 2008), which is inconsistent with most evaporation/recondensation models. Preventing evaporation by pre-enrichment requires temperature-dependent pre-enrichment, and so the pre-enrichment seen by the cooling chondrules has to have varied as they cooled through the alkali recondensation temperatures, implausible given the number of chondrule reservoirs (Jones 2012). Sufficiently preventing evaporation to avoid zoning by concentrating dust would imply optical depths in the thousands or more over the more than 1,000-km radius required to avoid Rayleigh fractionation (Alexander et al., 2008). As a consequence (Section 16.2.1.3), the cooling timescale would be prohibitively long.1 Thus, we are left with the need to evaporate and recondense the alkalis– sometimes (but not always) depleting the chondrules – without Rayleigh fractionating K. As shown in Hubbard (2016b), maintaining unzoned sodium is not the only difficulty: depleting chondrules of their alkalis is also difficult, and requires depleted chondrules to have separated from the evaporated Na and K clouds. Further, the lack of Rayleigh fractionation means that cooling depletion through recondensation temperatures had to be slow enough to allow continuous isotopic equilibration. These complications are sufficient that no model we will consider can be safely assumed to match the laboratory measurements.

16.2.3 Complementarity

Complementarity is the set of observations that chondrules and matrix within a given chondrite have different abundances and yet chondrite bulk abundances approximate the solar system average (i.e., CI composition, Hezel and Palme, 2010). Further, chondrules and matrix samples vary from chondrite class to chondrite class. If one were to disaggregate a chondrite of one class into chondrules and matrix and then use those components to construct a new bulk with the chondrule-to-matrix ratio of a different chondrite class, that bulk would have a non-CI compos- ition (Hezel and Palme, 2010). This existence of an apparent reference bulk CI composition is a significant constraint: given the large set of reservoirs associated with chondrule formation (Jones, 2012), it is implausible that chondrite assembly would have preserved bulk CI compos- ition unless that matrix material were correlated with (or co-genetic with) chondrule material. Complementarity has been observed in Mg/Si ratios (Hezel and Palme, 2010) and separately in Mo and W isotope ratios (Budde et al., 2016a, 2016b). However, not all elemental or isotopic systems are complementary: e.g. chondrites are depleted in the moderately volatile Na and K, and complementarity remains controversial (Zanda et al., 2017). Chondrule formation involves temperatures hot enough to evaporate solids (certainly including Mg and Si), and parent body processing provides alternate routes to redistributing elements. However, the more recent highly refractory, mass-independent isotopic results of Budde et al. (2016a, 2016b) argue persuasively that at least some systems are

1 The cooling time estimate in Alexander et al. (2008) is in error: it assumes that the center and surface of the cloud were at the same temperature, overestimating the cooling rate by several orders of magnitude. Evaluating Non-Shock, Non-Collisional Models for Chondrule Formation 407 complementary in a meaningful manner. Radial transport compromises complementarity on relatively modest timescales (Goldberg et al., 2015): if drift were rapid, complementarity would have been rapidly lost. If, instead, turbulent mixing dominated, it would have mixed chondrite reservoirs on a mixing time, leading to a homogenous chondrite population, contrary to Jones (2012). This implies that chondrite assembly must have occurred relatively soon after chondrule formation, before complementarity was lost or chondrule reservoirs mixed. Complementarity also gives teeth to the observation that the Tc ≳ 1; 750 K chondrule formation temperature floor (Hewins, 1997) is much higher than the Tm < 650 K cold-matrix temperature ceiling (Huss and Lewis, 1994; Rubin et al., 1999): cold-matrix–preserving regions must have been spatially and temporally close to hot chondrule-forming regions. Long spatial scales would imply mixing, which would destroy complementarity, while long temporal time scales would result either in the loss of complementarity or in the homogenization of the chondrule reservoirs.

16.2.3.1 Elemental and Isotopic Partitioning A further intriguing aspect of complementarity is that, while chondrules and matrix compos- itions may be correlated, they nonetheless differ. This difference occurs not only in bulk composition such as the Mg/Fe ratio (Ebel et al., 2016), but also in isotope ratios of nucleosyn- thetic, presolar origin (Budde et al., 2016a, 2016b), and places significant constraints on dust coagulation and chondrule formation mechanisms. Dust coagulation theory predicts that coagu- lation should not discriminate based on the composition of presolar collisional partners, which would drive toward homogeneity, contradicting observations. Even two families of grains, at least one chondrule precursor-sized, managing to maintain distinct compositions in the face of coagulation while existing close enough together to be co- genetic, are insufficient to explain the differing compositions of matrix and chondrules. The two families also need to have interacted with chondrule formation processes at different rates, with one family comparatively chondrule-destined, and the other relatively matrix-destined, lest chondrule and matrix compositions overlap. The two families of dust grain need to have been close enough to be neighbors, while simultaneously being sufficiently isolated from each other to avoid overlap of their compositions.

16.2.4 Energetic Constraints

In this chapter, outside the X-wind model, we consider chondrule formation processes that heat both the ambient gas and free-floating solids. The energy required to melt solid chondritic material is approximately (Wasson, 1996):

10 1 Emelt ffi 2 Â 10 erg g (16.10) while the energy required to heat disk gas a temperature ΔT is (neglecting H2 disassociation that occurs as the temperature nears 2,000 K): Δ Δ kg T 10 T 1 Ereq ffi ffi 9 Â 10 erg g (16.11) ðÞγ 1 mg 1; 000 K 408 Alexander Hubbard and Denton S. Ebel

Even if solids were concentrated to ε ffi 1, most of the energy associated with chondrule formation through melting free-floating dust must have gone into heating the ambient gas (King and Pringle, 2010). Transporting gas from infinity to R ~ 2.5 AU taps the disk’s gravitational potential energy, releasing: GM R E ffi e ffi 1:8 Â 1012 erg g1 ffi 20 E (16.12) grav 2R 2:5 AU req

Note that Egrav is greater than Ereq, but not by many orders of magnitude, and few processes could tap the entirety of Egrav for the purposes of chondrule melting. Chondrule heating processes that act through the medium of ambient gas require an energy reservoir comparable to the gravitational potential energy of the disk to process significant fractions of the solids.

Even then, the heating process has to have been reasonably efficient. Any H2 disassociation would increase the heating required, further exacerbating the problem. Processes which act through the medium of ambient gas with significantly smaller energy reservoirs, or which are inefficient, could still have generated the chondrules extant in the asteroid belt. However, those chondrules would only be evidence of a rare and unusual phenomenon rather than recording general disk processes and dynamics.

16.3 Dynamic Constraints

Thanks to the past few decades of research, we are in an excellent position to constrain the dynamic environment in which chondrules formed. This includes considerations of dust aero- dynamics, growth, and transport; and considerations of planetesimal formation.

16.3.1 Dust Aerodynamics

Dust grains interact with gas through drag:

v u Fd ∂tv ¼ þ g þ (16.13) τd md where md, v, and τ are the mass, velocity and frictional stopping time of a dust grain, and u the gas velocity at the dust grain’s position. Gravity (and any other forces that act equally on the gas and dust) is captured by g, while Fd captures forces that only act on the dust. The gas’s equation of motion is:

∂ þ ∙r ¼1 r þ tu u u ρ p g (16.14) g where ρg and p are the local gas density and pressure, and we have neglected any other forces that act solely on the gas. Chondrules (with radii a ≲ 0.25 mm) and chondrule precursors are generally expected to be significantly smaller than the local gas mean free path λ of centimeters to decimeters, and thus are in the Epstein drag regime, with: Evaluating Non-Shock, Non-Collisional Models for Chondrule Formation 409 rffiffiffi π ρ: τ ¼ τ ¼ a d E ρ (16.15) 8 gvth At constant mass, the higher a grain’s porosity, the lower its drag time. Very dense inner-most disk gas (R  1 AU) and very high porosities can both challenge the assumption that a < λ, landing one instead in the Stokes’ drag regime.

To couple Equation (16.15) to disk dynamics, the stopping time τd is generally normalized to the local Keplerian frequency Ω,defining the Stokes number:

St  τdΩ: (16.16) St  1 dust is tightly coupled to gas motions with characteristic time scales of an orbit, including the largest scale turbulent motions. At the midplane of a vertically isothermal disk, Equation (16.15) becomes:

π aρ• St0 ¼ : (16.17) 2 Σg

–4 At R = 2.5 AU our canonical chondrule would have St0 < 3 Â 10 , and be tightly coupled to all but the briefest gas motions.

We can combine Equations (16.13) and (16.14) under the assumption that τd is short compared to gas dynamic timescales, writing: τ w  v u ffi d rp þ terms in τ2 and higher: ρ d (16.18) g As a consequence, dust grains drift toward gas pressure maxima. This explains radial inspiral (the inner disk is higher pressure than the outer disk), dust settling (the midplane is higher pressure than surface layers) and dust trapping in high pressure anticyclonic vorticies. Different dust families with different St would drift at different speeds, so Equation (16.18) allows for aerodynamic sorting of dust grains. However, chondrules and their precursors are small enough that their relative drift was slow and significant segregation would have been difficult to achieve (Hubbard, 2016c).

16.3.2 Dust Coagulation

Dust coagulation theory envisions protoplanetary dust grains as aggregates of individual sub- micron monomers. These aggregates collide, due to thermal motion, turbulent stirring, or relative drift. When the collisions are slow enough, the grains stick together, forming a new aggregate. Other possible outcomes include various combinations of bouncing, compaction, fragmentation, erosion, or mass transfer from one aggregate to the other (Güttler et al., 2010). Significantly, dust collision speeds increase with St outside of the Brownian motion regime, so large grains collide faster than small ones. This naturally sets the maximum size of dust grains in a disk. Two problematic aspects of dust coagulation theory are particularly salient for considerations of chondrules and chondrites. Firstly, dust grains are expected to be arbitrarily sticky at the individual monomer and tiny aggregate stage: the collision speeds and kinetic energies are 410 Alexander Hubbard and Denton S. Ebel small, while the interaction surface areas and energies are relatively large. This makes the observed partitioning between chondrules and matrix of presolar grains with different nucleo- synthetic signatures (Budde et al., 2016a) puzzling. A mechanism is needed to maintain interaggregate heterogeneity via the dust aggregation process (Hubbard, 2016a) or to generate it via the chondrule formation process itself. The chondrule formation process could only generate chondrule/matrix compositional vari- ation if the chondrule-forming region were small enough (and the cooling process long enough) for the newly formed chondrules to exit the cloud of evaporated species, while also being large enough to prevent Rayleigh fractionation (Cuzzi and Alexander, 2006). While Hubbard (2016b) puts forth a model for potassium depletion through the chondrule formation mechanism which includes a long baking period for isotopic equilibration, that scenario cannot drive metal/silicate partitioning between matrix and chondrules: Fe would have completed recondensing long before the alkalis even began. Evaporative models also cannot explain the mass-independent isotopic partitioning seen by Budde et al. (2016a). A second problematic aspect of current dust growth theory is that bare silicates aggregates are extremely fragile (Guttler et al., 2010), and are predicted to be unable to grow (at least in number) to the size of at least the precursors of larger (e.g., CV or ordinary chondrite) chondrules (Jacquet, 2014). Growth of such grains should stall before the collision speed reaches about 1 cm s–1. Significantly, the literature focuses on growing icy dust grains, but the meteoritic record includes dry chondrites (Ebel et al., 2016), spotlighting a major hole in existing theory. Possible solutions to this problem under consideration include organic rims that act as adhesives (Hill and Nuth, 2000; Flynn et al., 2013), extensions from spherical monomers to more complicated geometries, and extensions from pure SiO2 monomers to softer and more adhesive compositions. This constraint is all the more pressing when we consider parent body formation.

16.3.3 Turbulent Transport

Protoplanetary disks are expected to be turbulent, with the most likely candidate some form of the MRI (Balbus and Hawley, 1991). As noted by Gammie (1996), protoplanetary disks are unlikely to be sufficiently ionized to support global MRI. Thermal ionization can support the MRI, but only for temperatures above about 1,000 K, notably above the background temperature limits posed by troilite or presolar grains (Huss and Lewis, 1994; Hewins, 1997; Rubin et al., 1999; Schrader et al., 2016). Over much of the orbital range of interest for terrestrial planet formation or meteoritics, magnetized turbulence is likely restricted to disk surface layers where nonthermal ionization can dominate. Recent work on magnetized turbulence has shown a large array of possible behaviors depending on the details of the nonideal processes involved (i.e., Ohmic, Hall, and ambipolar resistivities). Nonmagnetized instabilities have recently been uncovered, but are relatively weak (Nelson et al., 2013; Klahr and Hubbard, 2014). Turbulence drives turbulent diffusivities, often parameterized through a Shakura and Sunyaev (1973)-style α:

D ¼ αvthH: (16.19) From this we can define the radial and vertical mixing time scales: Evaluating Non-Shock, Non-Collisional Models for Chondrule Formation 411

1 R2 t ¼ R2=D ¼ Ω1, (16.20) R α H2

¼ 2= ¼ 1 : tH H D αΩ (16.20) It should be noted that turbulence likely varies as a function of altitude and orbital position, and is likely anisotropic, so the effective α values in Equations (16.20) and (16.21) could differ by several factors. Estimates for α range from 10–1 (strong transsonic turbulence) to 10–4 for weak midplane turbulence driven by active surface layers (Oishi and Mac Low, 2009). Turbulent diffusion plays into complementarity by limiting the possible delay between chondrule formation and chondrite assembly. If the delay were too long, either radial drift would have disrupted complementarity or turbulent mixing would have erased the difference between distinct chondrite reservoirs (Jones, 2012; Goldberg et al., 2015).

16.4 Innermost Disk Models 16.4.1 X-Wind

The X-wind model of Shu et al. (1996), heavily critiqued in Desch et al. (2010), proposes that dust grains were entrained by a magnetocentrifugal wind launched from the “X point” where the protosolar magnetosphere truncated the solar nebula. As the grains rose above the disk, they were heated to chondrule-forming temperatures by direct illumination from the protosun. This would have occurred only at only a few percent of an AU, where the background disk temperature was well above the maximum temperature for cold-matrix preservation. The wind is thus needed not only to loft the grains above the disk, but also to throw them out to a cool region where they could mix with matrix material. The existence of outflows from protoplanetary disks is not in question, having been verified by direct observations (Brown et al., 2013), and being one result of numerical simulations (Bai and Stone, 2013). Similarly, the existence of regions where direct illumination would heat solids to chondrule-melting temperatures is also certain. However, the innermost disk defies detailed numerical modeling, so the details of the wind speed or the fraction of solid material entrained are poorly constrained but quite significant (Hu, 2010). Further, it is unclear why and where the solids would decouple from the outflow and fall back to the disk. The X-wind model predicts that the innermost disk’s atmosphere might have seen sufficient high-energy particles to create meaningful quantities of some short-lived radionuclides (Gounelle et al., 2001), so solids processed through an X-wind type scenario might be identifiable.

16.4.2 Disk Interior

Another Innermost Disk location is the interior of a highly optically thick disk dissipating its accretion energy at the midplane, as per a constant-α disk model. We can write the local effective (surface) disk temperature for such a disk as a function of the mass accretion rate, 412 Alexander Hubbard and Denton S. Ebel and then calculate the midplane temperature as a function of the disk optical depth τ for an α-disk. From Frank et al. (2002), we have:

m_ ffi 3παcsΣgH (16.22)

 _ 1=4 ffi 3GMm Teff 3 (16.23) 8πR σSB

1=4 Tmid ffi ðÞ3τ=4 Teff (16.24)

where m_ is the mass flow rate through the disk, Teff and Tmid are the effective surface and midplane temperatures, and τ is the optical depth from the disk surface to the midplane. For large enough τ, Tmid could reach chondrule-forming temperatures even relatively far out in the disk. As we can see from Equation (16.21) though, vertical mixing is quick. Even in a scenario where the surface temperature was low enough to preserve cold-matrix material, mixing would have rapidly sunk that material deep enough to be thermally processed and chondrules would need to be radially transported outwards to mix with matrix in cooler regions. Further, α-disks can have midplanes hot enough to be in tension with cold-matrix preservation out to the asteroid belt. Taking MMSN-scale surface densities with κ ffi 20 cm2 g–1, the midplane temperature in an α-disk would be above the cold-matrix preservation temperature at R ¼ 1 and 2:5 AU for α values of 2 Â 104 and 102, respectively. In this scenario, outward transport mechanisms include meridional circulation, turbulent diffusion, and photophoresis, be it optically thick or thin. Meridional circulation in disks has long been hypothesized, and is a natural prediction for α-disks (Kley and Lin, 1992). It remains numerically controversial, but that should not be weighed too heavily: meridional circulation is slow, and easily overwhelmed by radial boundary conditions. That same slowness, however, means that it is unlikely to play a major role in transporting solids (Jacquet, 2013). Photophor- esis acts when the directed radiation field sets up a temperature gradient within dust grains, and can act in both the optically thin (Wurm et al., 2013) and optically thick (McNally and Hubbard, 2015) regimes. Gas molecules recoil more vigorously off the hotter side of the dust grains than off their colder side, generating a net force pushing the grains from hot regions to colder ones. Turbulent diffusion naturally transports solids in all directions. However, it competes with the inwards drift of the accretion flow and dust inspiral. Jacquet et al. (2012) derived the surface density of inspiraling dust grains in an α-disk. The steady state condition is: α Σd ðÞur þ wr Σd csHΣg∂r ¼ 0, (16.25) Scr Σg where Scr is the radial Schmidt number, of order unity, we recall that w = v – u and 3 H u ffi α c (16.26) r 2 R s Evaluating Non-Shock, Non-Collisional Models for Chondrule Formation 413 is the gas accretion velocity. Recall that turbulent diffusion doesn’t try to equalize dust density, p q but rather dust concentration. Writing Σg / R and T / R , we have: Σ ðÞ =α Scr d / exp cSt0 3=2 , (16.27) Σg R where c ¼ ðÞ2p þ q þ 3 =ðÞ3 2q e3 4. To mix with cold-matrix material, the chondrules need to be transported more than 10 times farther out to regions at least a factor of 3 cooler. That drops the chondrule-to-gas mass ratio by at least a factor of 30, problematically low for planetesimal formation purposes, but significant enough to contaminate the local dust popula- tion as long as the exponential term does not dominate.

16.4.3 Constraints

Thermal processing of solids in the innermost regions of disks and the existence of outflows are both uncontroversial. Photophoresis depends on the disk opacity and on the internal thermal conductivity of the grains, which are less certain, and meridional circulation remains unverified. While it is likely that some innermost disk material was transported outwards, it seems unlikely that it occurred in quantity and, as we will show, these processes did not generate meaningful numbers of chondrule-like objects.

16.4.3.1 Cooling Rates and Complementarity Blackbody temperatures and disk temperatures vary with orbital position approximately as 1=2 T / R . A body moving outwards with speed vr cools at a rate: ∂ _ T vr H T vr T T ¼ vr ¼π ¼π , (16.28) ∂R cs R Orb vKep Orb where vKep ¼ RΩ is the orbital velocity and Orb the local orbital period. For T ¼ 1; 750 K at R ¼ 0:05 AU, Equation (16.28) becomes v v H T_ ffi59 r Kh1 ¼59 r Kh1 (16.29) vKep vth R Protoplanetary disks are thin, with H/R  1, so significantly supersonic, approximately orbital radial velocities are needed to match the lower end chondrule cooling rates. Only winds or jets, launched at significant fractions of the local orbital speed, can drive sufficient cooling to explain a significant number of chondrules. Desch et al. (2012) estimates a cooling rate of 6 K h 1 for the X-wind model, which only flirts with the slowest chondrule cooling rates. Cooling therefore rules out all but X-wind–style Innermost Disk models for chondrule formation, and strongly limits the importance of X-wind–style models. All Innermost Disk models must transport chondrules outwards to regions with surviving cold-matrix. This makes fractionating elements and isotopes between chondrules and matrix trivial, but violates the constraint of complementarity. Complementarity therefore rules out all Innermost-Disk–style models for chondrule formation. 414 Alexander Hubbard and Denton S. Ebel

16.4.3.2 Sodium and Potassium In an X-wind model, with dust grains melted above the surface of the disk in very-low-density regions by direct illumination, all the moderately volatile elements would evaporate and be lost. In a midplane model with turbulent transport, the gas moves with the dust and we can imagine the evaporated potassium and sodium recondensing.

16.4.3.3 Energetics Innermost Disk models invoke the energy of the protosun or the gravitational potential energy of the disk very deep in the protosun’s gravitational potential, and therefore are energetically unconstrained. However, as material needs to fall into the potential well to power any outflows, only a small fraction of the material can be exported. In the case of a wind or jet, some of the entrained material could even be lost from the solar nebula entirely, further limiting the total amount of processed material available to be incorporated into planetesimals.

16.4.4 Innermost Disk Conclusions

Innermost-Disk–type models struggle matching laboratory constraints on cooling rates, and are unable to match the constraint of complementarity. Wind-driven versions are marginally capable of matching cooling rates, but are incapable of retaining any sodium or potassium. Embedded models could potentially retain the alkalis at the cost of prohibitively slow cooling rates. Thus, Innermost-Disk–type models are ruled out as a major chondrule formation mech- anism. Nonetheless, we expect some transport of solids processed in the innermost disk to cooler regions potentially associated with meteorite parent body formation. Wind-based models especially might be observable in extrasolar protoplanetary disks. Any experimental constraints would provide information about the inner edge of the solar nebula, making them extremely valuable to disk modelers. Further, molten grain collisions in the innermost disk have been proposed as a partial solution to the bouncing barrier for dust grain growth (Boley et al., 2014; Hubbard, 2017).

16.5 Lightning

Lightning-like electrical discharges in the solar nebula could have melted chondrules, and have excited significant interest (Gibbard et al., 1997; Pilipp et al., 1998; Desch and Cuzzi, 2000; Muranushi, 2010). These discharges occur when the electrical field exceeds a critical value, often estimated to be enough to accelerate electrons enough to ionize neutral gas molecules over one mean free path (Gibbard et al., 1997; Pilipp et al., 1998; Desch and Cuzzi, 2000). However, that is likely an overestimate because high mass-ratio inelastic collisions are very inefficient at thermalizing energy (Inutsuka and Sano, 2005). As long as collisions between electrons and neutrals are nearly elastic, electrons have more than one mean free path to accelerate, reducing the critical electric field by about an order of magnitude. Evaluating Non-Shock, Non-Collisional Models for Chondrule Formation 415 16.5.1 Charge Separation

Generating electrical discharges requires both charge separation and gas that is sufficiently neutral to prevent return currents from draining the electric field. Thus, we require relatively immobile charge carriers (i.e. dust grains). While generating families of positively and negatively charged dust is necessary for lightning, it is not sufficient. In addition, at least one of the positively or negatively charged families must also be strongly spatially concentrated to spatially concentrate charge. From Equation (16.18) we can see that that spatial concentration requires that the two dust families have different Stokes numbers (St). The requirement of two families of grains with strongly differing aerodynamics is problem- atic: numerical simulations which include bouncing generally do not see a large population of very small grains (Zsom et al., 2010). Small grains are stickier than large ones, while collision velocities depend on the larger grains (Güttler et al., 2010), so small grain–large grain collisions can lead to sticking even when large grain–large grain collisions lead to bouncing and stalled growth. Thus, the available pool of small grains is expected to be depleted.

16.5.1.1 Charging Grains Desch and Cuzzi (2000) proposed triboelectric charging to charge dust grains, which works when two materials with different contact potentials collide. As they pointed out, iron and silicates have a significant contact potential difference of about ΔΦcp ffi 2V.Thatis intriguing because matrix and chondrules differ significantly and consistently across chondrite classes in their iron-to-silicates ratio (Grossman and Wasson, 1982). While the potential origin of such a dichotomy is not understood, it is reasonable to hypothesize two families of dust grains in the solar nebula, one iron-rich and the other silicate-rich, and even potentially that they would have had differing aerodynamics (Hubbard 2016c). Note some metal-silicate fractionation mechanisms (such as metal expulsion from molten grains) rely on chondrule formation scale events and so pose a chicken-and-egg problem for lightning (Uesugi et al., 2008). A second mechanism collides grains with internally separated charges, such as occurs with water ice (Muranushi, 2010). Here, the grain surfaces have net charges which depend on their history. When two grains collide, they share their surface (but not interior) charges, transferring charge. As long as the electrical properties of the grains correlate with their aerodynamics, this would also lead to charge separation. Regardless of the mechanism, it is easily plausible that, if a broad dust-size spectrum exists, collisions between small and large grains would lead to charge transfer between the grains. However, evaluating the degree of charge transfer between grains requires accurate models of both the dust-size spectrum and surface characteristics.

16.5.1.2 Grain Concentration Concentrating grains of one type requires moving them relative to other grains, implying also that they move differently relative to the gas. As can be seen through Equation (16.18), grains concentration mechanisms such as settling and preferential concentration act through pressure gradients. The pressure gradients associated with orbital position or zonal flows are significantly shallower than vertical pressure gradients, and we neglect them here. One interesting aspect of 416 Alexander Hubbard and Denton S. Ebel lightning is that it explicitly assumes two different families of dust grains which are concen- trated with respect to each other in a way correlated with chondrule formation, offering an explanation for matrix/chondrule partitioning. Settling is an attractive means to generate charge separation because vertical pressure gradients in disks are reliably strong, and because it implies that any lightning would occur well above the midplane, allowing the midplane to remain a cold region for cold-matrix preservation and chondrite assembly. Large grains are more settled than smaller grains, so we expect a layer of fine dust above the settled large dust column. However, for large dust to settle significantly we require α < St0 at the midplane (Takeuchi and Lin, 2002). That is possible as long as the disk has turbulent surface layers and a quiescent midplane, but is otherwise unlikely given the small size of chondrules. Thus, vertical charge separation would be expected to occur high up in the disk, with the threat of significant nonthermal ionization equilibrating charge before a critical electric field could be built up. A second dust concentration mechanism noted by Desch and Cuzzi (2000) is preferential concentration (Maxey, 1987; Fessler et al., 1994; Cuzzi et al., 2010). We can imagine turbulence as a collection of low-pressure vortices maintained against centrifugal forces by outside pressure. Thus, for vortices with radius l (and velocity and time scales ul and tl), we have:

u2 1 r l : ρ pl e (16.30) g l

We can use Equation (16.18) to estimate that on a turbulent timescale the pressure gradient moves dust with stopping time τd  tl a length r τ ℓ ffi  ffi τ pl ffi d : tl w dtl ρ l (16.31) g tl The pressure gradient that supports vortices of length l is sufficient to move dust grains with stopping time τdetl out of the vortices and concentrate them in the high-pressure regions between vortices. This can drive extreme, almost arbitrary dust concentrations as long as the dust grains have sufficiently similar aerodynamics. The assumptions built into models of preferential concentration break down once the dust concentration becomes large enough to backreact on the gas. Further, small differences in aerodynamics rapidly decrease peak concentrations because the dust grains couple to different scale turbulent eddies (Hubbard, 2013): the peak dust concentration depends on the dust size spectrums. Dust coagulation theory predicts that dust mass will pile up in grains with a narrow size distribution, just large enough to mostly bounce, so preferential concentration is a plausible but tricky mechanism for concentrating dust that depends sensitively on the uncertain details of dust coagulation. Preferential concentration is not vertically restricted in the same way that relative settling is, meaning that to separate the chondrule-forming region from the chondrite assembly/cold-matrix preservation region we need turbulent activity to vary with altitude. That is reasonable for surface layer magnetized turbu- lence, but magnetized turbulence assumes high ionization fractions, counteracting charge separation. Evaluating Non-Shock, Non-Collisional Models for Chondrule Formation 417

16.5.1.3 Maintaining Charge Separation The background ionization state of the disk must be low enough to keep free electrons and ions in the disk from discharging the grains. Pilipp et al. (1998) examined charge separation via settling counteracted by ionization from the decay of 40K, while Desch and Cuzzi (2000) examined triboelectric charging and preferential concentration counteracted by the decay of 26Al. Both found that charge separation could not be maintained, except in the case of tribo- electric charging of preferential concentrated dust occurring after 26Al had mostly decayed. Large concentrations of dust can soak up charges, suggesting an alternate route to low ioniza- tion: Muranushi (2010) found that two or more orders of magnitude dust concentration could permit charge separation. Even more recently, it has been pointed out that not only is the relationship between current and electric field (Ohm’s law) not linear in protoplanetary disks, it is not even reliably single-valued (Okuzumi and Inutsuka, 2015). Determining whether a mechanism of charge separation can generate sufficient electrical fields to drive chondrule-forming discharges is a difficult task which depends sensitively on the nature of the dust, its spatial and size distributions, and the amount of charge transfer on collision. It also depends on the background disk chemistry and ionization state in a nonlinear fashion. However, it is clear that sufficient charge separation is not expected to be possible in the abundantly ionized surface layers. Even in the shielded midplane, large dust concentrations are needed both to increase the rate of charging dust–dust collisions and to soak up free electrons and ions.

16.5.2 Constraints

16.5.2.1 Cooling Rates, and Sodium and Potassium A major challenge in invoking lightning for chondrule formation is the small diameter of lightning bolts. The initial channel is only expected to be perhaps 103 electron mean-free-paths wide, or about 10 m. This places lightning solidly in the optically thin regime, with all the attendant cooling rate issues of Section 16.2.1.2. This small size means that photons are not the only free-streaming players: potassium atoms at 1; 500 K have thermal speeds about 5 Â 104 cm s 1, and a thermal diffusion coefficient of about:

6 2 1 DK ffi 10 cm s : (16.32)

The diffusion time of potassium out of the channel is only e1 s. Worse, even assuming that the – – gas density was an elevated 10 9 gcm 3, it would take a potassium atom over a minute to collide with a dust grain, making recondensation that much less likely. Thus, any sodium or potassium depletion would be accompanied by Rayleigh fractionation: the moderately volatile elements testify strongly against lightning as a chondrule formation mechanism. Even the extremely quick cooling time associated with lightning cannot trap sodium over multiple heating events (Jacquet et al., 2015).

16.5.2.2 Complementarity Because lightning bolts are low-volume, lightning can easily satisfy the basic constraint of complementarity: most of the disk volume would avoid heating, and thus cold-matrix could be 418 Alexander Hubbard and Denton S. Ebel preserved even very near where chondrules were melted. Further, as lightning requires shielding from X-ray and cosmic ray ionization, it must occur near the midplane. A midplane-located chondrule formation mechanism should lead to all types of solids being thermally processed, and so raises the question of why we don’t see planetesimal-precusor–sized chondrules (Hub- bard and Ebel, 2015). On the other hand, the requirement for spatially separated distinct dust families could drive matrix/chondrule partitioning.

16.5.2.3 Energetics Lightning releases energy stored in the electric field, put there by separating oppositely charged grains. Gas carries most of the kinetic energy associated with turbulence, and turbulent lofting, a diffusive process, is an extremely inefficient way to convert the disk’s gravitational potential energy into electrical energy. Similarly, as the overall dust-to-gas mass ratio is low, preferential concentration can at most tap a small part of the turbulent energy. However, lightning is only expected in regions where the dust is highly concentrated, so that the lightning bolts occur in regions where the solid-to-gas mass ratio ε is well above unity, increasing the number of grains affected by each bolt. Lightning might be able to satisfy energetics requirements given extreme preferential concentration or extreme dust settling.

16.5.3 Lightning Conclusions

As long as dust growth allows for a significant range of dust sizes, and we note that this is in tension with current dust growth theory, the charging of dust grains and at least some charge separation seems likely. Generating sufficient charge separation in the face of background electrical current, however, seems unlikely unless the dust density can be increased by about two orders of magnitude. Even if lightning is possible, it is unlikely that enough energy can be put into the electric fields to process significant fractions of the solids. Further, the cooling time constraints and the existence of non-Rayleigh fractionated, moderately depleted K in chondrules rule out lightning as a chondrule formation mechanism. Intriguingly however, recent work has shown that lightning in extrasolar protoplanetary disks should have observation signatures (Muranushi et al., 2015).

16.6 Radiant Heating

Herbst and Greenwood (2016) suggested that differentiated planetesimals with surface magma oceans (Greenwood et al., 2005; Scherstén et al., 2006) could radiantly heat and melt nebular solids as they flowed past. A recent addition to the chondrule formation mechanism zoo, this model has not yet been examined in detail, so we restrict outselves very simple statements indicating a need for greater investigations.

16.6.1 Feasibility

In the optically thin limit, a chondrule precursor passing close to a surface at temperature Ts will 1=4 heat to a temperature 2 Ts, so a chondrule melting temperature of Tc ¼ 1; 750 K implies Ts ≳ 2,080 K. While it is clear that planetesimals differentiated early in the solar nebula (Wadhwa Evaluating Non-Shock, Non-Collisional Models for Chondrule Formation 419 et al., 2006), the existence of sufficiently hot surface magma oceans is less clear, making it hard to evaluate the feasibility of radiant heating. Optically thick volumes would lower the constraint to Ts ffi Tc. If we assume a r ¼ 100 km molten planetesimal moving at the headwind speed of ve50 m s 1 through the disk, it will take the dust about t e1=2 h to drift past the planetesimal, much less than the 1; 750 K 180 K e50 h (16.33) Lcg=Cv time for the radiatively heated chondrules to collisionally heat the initially cool ambient gas. In Equation (16.33), we have used Equations 16.2 and 16.4. Unusually, radiant heating assumes chondrules significantly warmer than the gas. Assuming that degree of gas heating, we can estimate the volume of solids heated radiantly by balancing the energy radiated by the planetesimal with the energy required (Equation 16.11) to heat the gas flowing past it in a cylinder of radius ℓ:

π 2σ 4 ¼ πℓ2 ρ 4 r SBTs v gEreq, (16.34) ρ ¼ 10 3 which for g 10 gcm gives ℓ ffi 7 Â 105ðÞT =K 2 ffi 200: (16.35) r s

Planetesimals with Ts e1; 750 K and r > 25 km can heat ℓ ¼ 5; 000 km radius, optically thick regions before they flow past. Thus, while much work remains to be done, sufficiently hot planetesimals would have heated nebulas solids to chondrule-forming temperatures across optically thick regions. A second problem is the trajectory of the grains. Chondrules with a ¼ 0:025 cm had a stopping time τd e4; 000 s. That time would drop to perhaps τd e800 s for their porous precursors or for chondrules associated with CM or CO chondrites, and finally to perhaps τd e200 s for the porous precursors of CM or CO chondrules. If a planetesimal of radius r moves through disk gas at a speed v, then solid material with a stopping length vτd  r will have its trajectory altered by the gas flow around the planetesimal. For v ¼ 50 m s 1 the above stopping times imply planetesimal radii of 200, 40 and 10 km, respectively. Thus, the simple geometry studied in Herbst and Greenwood (2016) does not apply, and more elaborate models are needed to quantify the trajectories and hence thermal histories of solids (Ormel and Klahr, 2010). Dust grains of different sizes will evolve differently in radiant heating models because their aerodynamics are important, implying chondrule size/thermal history correlations.

16.6.2 Constraints and Conclusions

Herbst and Greenwood (2016) calculate intriguing heating and cooling rates, but as we have developed in this chapter, significantly more elaborate models are needed. If ℓ  r as suggested by Equation (16.35), then Herbst and Greenwood (2016) likely overestimate the cooling rates by about two orders of magnitude, making it difficult to match laboratory constraints. While the 420 Alexander Hubbard and Denton S. Ebel heated region is likely large enough to avoid Rayleigh fractionation of evaporated potassium through diffusion into the unheated nebular gas, the planetesimals are large enough to aerody- namically sort dust grains. Any volatile-rich outgassing from the assumed magma oceans would lead to size-dependent volatile enrichment and likely fractionation. Radiant heating mechanisms only effect regions close to the molten planetesimals, so complementarity is straightforward as long as the planetesimals are sufficiently rare. If radiant heating aerodynamically sorts the grains it heats sufficiently strongly, it might provide a way to partition elements between chondrules and matrix. Finally, the rate of chondrule production would be limited by the rate of molten planetesimal production, and it is unclear if that production occurred. In sum, the proposal is intriguing, but underdeveloped.

16.7 Localized Turbulent Dissipation

The energetic considerations of Section 16.2.4 lead one naturally toward mechanisms which tap the accretion energy of the solar nebula. Accretion is generally accepted to proceed through turbulence, be it gravitational, hydrodynamical, or magnetically mediated. Winds have emerged as a potential alternative (Bai and Stone, 2013) which would be hostile to chondrule formation by spreading out energy dissipation.

16.7.1 Feasibility: Hydrodynamical versus Magnetic Dissipation

In the absence of thermal physics, accretion disks are expected to be hydrodynamically stable (Balbus and Hawley, 1998), which made rediscovering the magnetically mediated MRI par- ticularly exciting (Balbus and Hawley, 1991). However, only a few years later Gammie (1996) showed that the midplanes of protoplanetary disks are unlikely to be sufficiently ionized to allow the MRI to act. We now have a picture where X-ray, extreme and far ultraviolet, and cosmic ray ionized surface layers support the MRI, but the nearly neutral midplanes remain quiescent (Gammie, 1996; Gressel et al., 2015). Protoplanetary disks are, however, sufficiently inviscid that surface turbulence can propagate meaningfully into the midplane (Oishi and Mac Low, 2009). More recent work has shown that disks can actually become hydrodynamically unstable or even overstable (Nelson et al., 2013; Klahr and Hubbard, 2014) in the presence of thermal diffusivities. In these cases, the energy source is whatever sets the background temperature gradient (likely irradiation from the central protostar). Hydrodynamical turbulence dissipates its energy either in a localized high-intensity fashion through shocks or more broadly through viscous dissipation. Current models for hydrodynamical, and indeed non- GI, turbulence in general do not see sufficiently high Mach number shocks to drive chondrule formation. If individual turbulent events do not raise the temperature to chondrule melting levels, we can use a turbulence-as-α prescription to sum over events. As long as the turbulent layers are optically thick toward the surface of the disk, turbulent dissipation could raise the temperature to chondrule melting levels at the cost of destroying local cold-matrix, and very slow cooling. This is equivalent to an Innermost Disk model. Evaluating Non-Shock, Non-Collisional Models for Chondrule Formation 421 16.7.2 Magnetic Dissipation

In protoplanetary disks, Ohmic and ambipolar resistivies dominate over fluid viscosities, and turbulent energy dissipation occurs through the magnetic field (Brandenburg, 2009). This takes the form of quasi–two-dimensional structures known as current sheets, which have been proposed to lead to chondrule formation (Joung et al., 2004). Current research on magnetized turbulence in protoplanetary disks has largely focused on the effect of the dissipative terms on the magnetic field and accretion stresses, leaving the thermal implications relatively unexplored. However, long-lived dissipation regions which would have significant effects on the tempera- ture have been found (McNally et al., 2014). Further, the sharp temperature dependence of the ionization fraction (and hence the nonideal MHD terms) leads to at least one instability, which can focus the energy dissipation into a very small region (Hubbard et al., 2012; McNally et al., 2013). The combination of cold-matrix preservation and complementarity means that at least most of the disk must be well below thermal ionization temperatures of about 1; 000 K. This restricts magnetic dissipation to radiatively ionized surface layers. It also strongly complicates numerical investigations by forcing strong spatial variation in the dissipation coefficients.

16.7.3 Constraints

16.7.3.1 Cooling Rates As noted in Sections 16.2.1.2 and 16.2.1.3, chondrule heating is only expected to work in optically thick regions, which also fits with the need to avoid Rayleigh fractionation of potassium isotopes. Turbulent dissipation should heat the disk on dynamic timescales. However, if the magnetic dissipation is limited to occurring at altitude, the optical depth to the disk surface would be modest, and evaporating and liquefying the dust would drop the opacity significantly. A delay between the heating and the evolution of the opacity would allow the dissipation region to heat, become transparent, and cool regardless of the continued energy input. This reinforces the need for a better understanding of the opacity’s (time-dependent) temperature dependence.

16.7.3.2 Complementarity and the Alkalis Turbulent dissipation is localized, allowing regions of lower temperature to coexist near dissipation regions. However, thermal ionization only supports the MRI in regions above about 1; 000 K, easily sufficient to destroy the cold-matrix. As noted in Section 16.7.2, this restricts dissipation to nonthermally ionized surface layers, protecting cold-matrix and allowing newly formed chondrules to settle to the midplane, where they could form aggregates and trigger the streaming instability. If the turbulent dissipation is restricted to sufficiently high altitudes, differential settling could separate families of grains with different elemental and isotopic compositions and allow alkali depletion (Hubbard, 2016b, 2016c).

16.7.3.3 Energetics Magnetized turbulence can tap the entire gravitational potential energy of the disk, and hence can process significant amounts of solids, so long as radiative losses are not too high. That is 422 Alexander Hubbard and Denton S. Ebel difficult to reconcile with the need for the chondrule-forming region to tend toward being optically thin to match cooling rates.

16.7.4 Grain Magnetization

One laboratory diagnostic associated with magnetized disks is the fossil magnetization of meteoritical solids, which can be used to estimate the magnetic field seen by chondritic inclusions as they cooled through their Curie temperature (Fu et al., 2014). The Curie tempera- ture of iron metal is about 1; 000 K, well above the cold-matrix preservation temperature, well below chondrule formation temperature, comparable to the alkali recondensation temperature, and near the thermal ionization temperature. One difficulty associated with this approach is that Ohmic resistivity concentrates dissipation in regions of field reversal, and so chondrules melted through Ohmic dissipation should not sample the full magnetic field of the disk. Ambipolar resistivity, however, concentrates the dissipation in regions of maximal Lorentz forces, and will sample a significant fraction of the field.

16.7.5 Turbulent Dissipation Conclusions

The dissipation of magnetized turbulence may be able to match many of the requirements, but numerical simulations have not yet focused on chondrule formation. The restriction to altitude fits well with the requirements of complementarity, cold-matrix preservation, and a cold planetesimal formation region. The restriction to altitude further assists in the partitioning of distinct chondrule-destined and matrix-destined families, and turbulence can tap the gravita- tional potential energy of the disk. However, it is not yet certain the turbulence can sufficiently concentrate its energy dissipation. Further, to match cooling rates, turbulent dissipation requires a delicate transition from optically thick toward optically thin as the region heats.

16.8 Conclusions

We have examined several proposed chondrule formation mechanisms for both cosmochemical and dynamic plausibility, with our conclusions summarized in Table 16.1, where ? implies “plausible but problematic,” and “extremely unlikely but not strictly ruled out”.Notethatthe Energetics column encompasses the bulk production of chondrules, and thus the difficulty the

Table 16.1 Chondrule Formation Mechanisms in the Solar Nebula

Complementarity Na and K

Co- Retention/ Mechanism Feasibility Cooling genetic Partitioning Depletion Fractionation Energetics Conclusions

Innermost Wind ✓ ✗ ✓ ✗✗? ✗ Disk Embedded ✓ ✗✗ ✓ ✓ ? ✗ Lightning ✗ ✓ ? ✗ ✗ Radiant Heating ✓ ?? Turbulent Dissipation ? ? ✓✓ ? ✓✓? Evaluating Non-Shock, Non-Collisional Models for Chondrule Formation 423

Table 16.2 Heating Mechanisms in Extrasolar Protoplanetary Disks

Mechanism Feasibility Energetics Observability Significance

Innermost Disk Wind ✓ ? ✓✓ Embedded ✓ ? ✗✗ Lightning ✓ Radiant Heating ✗✗ Turbulent Dissipation ? ✓ ??

Innermost Disk models have in radially transporting chondrules is registered there. Usefully, we find that most models, at least in their current state, can be ruled out for chondrule production. Innermost Disk models, while they should have operated at some level, are ruled out as more than trace sources of chondrules by their cooling timescales, alkali retention, and complementarity. Lightning-based models are ruled out by cooling timescales and potassium Rayleigh fractionation. Further, it is not clear that lightning is feasible in the first place and even if it was, the production rate is constrained to be quite low. Radiant heating is not well studied as of yet. While simple estimates show that planetesimals with molten surfaces would have an interesting effect on the solids, it seems likely that such a model would make predictions about aerodynamic sorting which are not seen in the meteoritic record. Turbulent dissipation remains as a strong subject for future work studying the actual thermal history of particles in disks with accurately modeled thermo- dynamics and radiative transport with temperature- and time-dependent opacities. Historical work has placed a strong emphasis (for good reason) on matching chondrule thermal histories. Recent studies suggest that matrix and chondrules are co-genetic, but also require isotopic partitioning between chondrule-destined and matrix destined grains. That combines to form a very potent and exciting constraint on chondrule formation mechanisms. Mechanisms that separate chondrule formation from chondrite assembly find partitioning easy to satisfy. Mechanisms that heated small isolated portions of the solar nebula allowing local chondrite assembly naturally allow matrix and chondrule to be co-genetic, but must have distinguished between chondrule-destined and matrix-destined grains. One useful aspect of studying chondrule formation is that chondrules require the solar nebula not to have been boring. The mere existence of chondrules has inspired and justified studies which, while not always relevant to the solar nebula, should find a home in at least some extrasolar protoplanetary disks, as summarized in Table 16.2 where we strongly emphasize observability over production rates (i.e., energetics). It is our belief and hope that advances in meteoritics continue to inspire research which remains relevant to the larger study of planet formation across our galaxy, even if it ends up failing when confronted with the extremely detailed measurements possible in terrestrial laboratories.

Acknowledgments

The authors thank the anonymous reviewers for constructive and corrective reviews. The assist- ance, support, and patience offered by Alexander Krot and Sara Russell was instrumental in the production of this manuscript. This work was supported by NASA OSS grant NNX14AJ56G. 424 Alexander Hubbard and Denton S. Ebel

References

Alexander, C. M. O’D., Grossman, J. N., Ebel, D. S., and Ciesla, F. J. (2008). The formation conditions of chondrules and chondrites. Science, 320, 1617. Bai, X. -N., and Stone, J. M. (2013). Wind-driven accretion in protoplanetary disks. I. Suppression of the magnetorotational instability and launching of the magnetocentrifugal wind. Astrophys. J., 769, 76. Balbus, S. A., and Hawley, J. F. (1991). A powerful local shear instability in weakly magnetized disks. I - Linear analysis. II - Nonlinear evolution, Astrophys. J., 376, 214. Balbus, S. A., and Hawley, J. F. (1998). Instability, turbulence, and enhanced transport in accretion disks, Rev. Mod. Phys., 70,1. Boley, A. C., Morris, M. A., and Ford, E. B. (2014). Overcoming the meter barrier and the formation of systems with tightly packed inner planets (STIPs), Astrophys. J. Lett., 792, L27. Boss, A. P. (1996). A concise guide to chondrule formation models. In R. H. Hewins, R. H. Jones, and E. R. D. Scott (Eds.), Chondrules and the Protoplanetary Disk, 257–263. Cambridge, UK: Cambridge University Press. Brandenburg, A. (2009). Large-scale dynamos at low magnetic prandtl numbers. Astrophys. J., 697, 1206. Brown, J. M., Pontoppidan, K. M., van Dishoeck, E. F., et al. (2013). VLT-CRIRES survey of rovibrational CO emission from protoplanetary disks, Astrophys. J., 770, 94. Budde, G., Burkhardt, C., Brennecka, G. A., et al. (2016a). Molybdenum isotopic evidence for the origin of chondrules and a distinct genetic heritage of carbonaceous and non- carbonaceous meteorites, Earth Planet. Sci. Lett., 454, 293. Budde, G., Kleine, T., Kruijer, T. S., Burkhardt, C., and Metzler, K. (2016b). Tungsten isotopic constraints on the age and origin of chondrules, Proc. Natl. Acad. Sci., 113, 2886. Cuzzi, J. N., and Alexander, C. M. O’D. (2006). Chondrule formation in particle-rich nebular regions at least hundreds of kilometres across, Nature, 441, 483. Cuzzi, J. N., Hogan, R. C., and Bottke, W. F. (2010). Towards initial mass functions for asteroids and Kuiper Belt Objects, Icarus, 208, 518. Desch, S. J. (2007). Mass distribution and planet formation in the solar nebula, Astrophys. J., 671, 878. Desch, S. J., and Connolly, H. C. Jr. (2002). A model of the thermal processing of particles in solar nebula shocks: Application to the cooling rates of chondrules, Meteorit. Planet. Sci., 37, 183. Desch, S. J., and Cuzzi, J. N. (2000). The generation of lightning in the solar nebula. Icarus, 143, 87. Desch, S. J., Morris, M. A., Connolly, H. C., and Boss, A. P. (2012). The importance of experiments: Constraints on chondrule formation models, Meteorit. Planet. Sci., 47, 1139. Desch, S. J., Morris, M. A., Connolly, H. C. Jr., and Boss, A. P. (2010). A critical examination of the X-wind model for chondrule and calcium-rich, aluminum-rich inclusion formation and radionuclide production. Astrophys. J., 725, 692. Ebel, D. S. (2006). Condensation of Rocky Material in Astrophysical Environments. In D. S. Lauretta and H. Y. McSween (Eds.), Meteorites and the Early Solar System II, 253–277. Tucson, AZ: University of Arizona Press. Ebel, D. S., Brunner, C., Konrad, K., et al. (2016). Abundance, major element composition and size of components and matrix in CV, CO and Acfer 094 chondrites. Geochim. Cosmo- chim. Acta, 172, 322. Ebel, D. S., and Grossman, L. (2000). Condensation in dust-enriched systems. Geochim. Cosmochim. Acta, 64, 339. Fessler, J. R., Kulick, J. D., and Eaton, J. K. (1994). Preferential concentration of heavy particles in a turbulent channel flow. Phys. Fluids, 6, 3742. Evaluating Non-Shock, Non-Collisional Models for Chondrule Formation 425

Flynn, G. J., Wirick, S., and Keller, L. P. (2013). Organic grain coatings in primitive interplan- etary dust particles: Implications for grain sticking in the Solar Nebula. Earth, Planets Space, 65, 1159. Frank, J., King, A., and Raine, D. J. (2002). Accretion Power in Astrophysics: Third Edition, Cambridge UK: Cambridge University Press. Friedrich, J. M., Weisberg, M. K., Ebel, D. S., et al. (2015). Chondrule size and related physical properties: A compilation and evaluation of current data across all meteorite groups. Chem Erde, 75, 419. Fu, R. R., Lima, E. A., and Weiss, B. P. (2014). No nebular magnetization in the Allende CV carbonaceous chondrite. Earth Planet. Sci. Lett., 404, 54. Gammie, C. F. (1996). Layered accretion in T Tauri disks. Astrophys. J., 457, 355. Gibbard, S. G., Levy, E. H., and Morfill, G. E. (1997). On the possibility of lightning in the protosolar nebula. Icarus, 130, 517. Goldberg, A. Z., Owen, J. E., and Jacquet, E. (2015) Chondrule transport in protoplanetary discs. Mon. Notices Royal Astron. Soc., 452, 4054. Gounelle, M., Shu, F. H., Shang, H., et al. (2001). Extinct radioactivities and protosolar cosmic rays: self-shielding and light elements. Astrophys. J., 548, 1051. Greenwood, R. C., Franchi, I. A., Jambon, A., and Buchanan, P. C. (2005). Widespread magma oceans on asteroidal bodies in the early Solar System. Nature, 435, 916. Gressel, O., Turner, N. J., Nelson, R. P., and McNally, C. P. (2015). Global simulations of protoplanetary disks with ohmic resistivity and ambipolar diffusion. Astrophys. J., 801,84. Grossman, J. N., and Wasson, J. T. (1982). Evidence for primitive nebular components in chondrules from the Chainpur chondrite. Geochim. Cosmochim. Acta, 46, 1081. Güttler, C., Blum, J., Zsom, A., Ormel, C. W., and Dullemond, C. P. (2010). The outcome of protoplanetary dust growth: Pebbles, boulders, or planetesimals?. I. Mapping the zoo of laboratory collision experiments. Astron. Astrophys, 513, A56. Hayashi, C. (1981). Structure of the solar nebula, growth and decay of magnetic fields and effects of magnetic and turbulent viscosities on the nebula. Progr. Theoret. Phys. Suppl., 70, 35. Herbst, W., and Greenwood, J. P. (2016). A new mechanism for chondrule formation: Radiative heating by hot planetesimals. Icarus, 267, 364. Hewins, R. H. (1991). Retention of sodium during chondrule melting, Geochim. Cosmochim. Acta, 55, 935. Hewins, R. H. (1997). Chondrules. Ann. Rev. Earth Planet. Sci., 25, 61. Hewins, R. H., Connolly, H. C. Jr., Lofgren, G. E., and Libourel, G. (2005). Experimental constraints on chondrule formation. In A. N. Krot, E. R. D. Scott, and B. Reipurth (Eds.), Chondrites and the Protoplanetary Disk. ASP Conf. Ser., 341, 286. San Francisco, CA: Astronomical Society of the Pacific. Hezel, D. C., and Palme, H. (2010). The chemical relationship between chondrules and matrix and the chondrule matrix complementarity. Earth Planet. Sci. Lett., 294, 85. Hill, H. G. M., and Nuth, J. A. III, (2000). Nebular hydrocarbon synthesis in the laboratory: the catalytic potential of synthetic silicate dust. Meteorit. Planet. Sci. Supp., 35, A73. Hu, R. (2010). Transport of the first rocks of the solar system by X-winds. Astrophys. J., 725, 1421. Hubbard, A. (2013). Turbulence-induced collision velocities and rates between different sized dust grains. Mon. Notices Royal Astron. Soc., 432, 1274. Hubbard, A. (2016a), Ferromagnetism and particle collisions: Applications to protoplanetary disks and the meteoritical record. Astrophys. J., 826, 152. Hubbard, A. (2016b). Generating potassium abundance variations in the Solar Nebula. Mon. Notices Royal Astron. Soc., 460, 1163. Hubbard, A. (2016c). Partitioning tungsten between matrix precursors and chondrule precursors through relative settling. Astrophys. J., 826, 151. Hubbard, A. (2017). Making terrestrial planets: high temperatures, FU Orionis outbursts, earth, and planetary system architectures. Astrophys. J. Lett., 840, L5. 426 Alexander Hubbard and Denton S. Ebel

Hubbard, A., and Ebel, D. S. (2015). Semarkona: Lessons for chondrule and chondrite forma- tion. Icarus, 245, 32. Hubbard, A., Mac Low, M.-M., and Ebel, D. S. (2018). Dust concentration and chondrule formation, Meteoritics and Planetary Sciences, ArXiv e-prints arXiv:1803.10047. Hubbard, A., McNally, C. P., and Mac Low, M.-M. (2012). Short circuits in thermally ionized plasmas: a mechanism for intermittent heating of protoplanetary disks. Astrophys. J., 761,58. Humayun, M. (2012). Chondrule cooling rates inferred from diffusive profiles in metal lumps from the Acfer 097 CR2 chondrite. Meteorit. Planet. Sci., 47, 1191. Huss, G. R., and Lewis, R. S. (1994). Noble gases in presolar diamonds II: Component abundances reflect thermal processing. Meteoritics, 29, 811. Inutsuka, S. -i., and Sano, T. (2005). Self-sustained ionization and vanishing dead zones in protoplanetary disks. Astrophys. J. Lett., 628, L155. Jacquet, E. (2013). On vertical variations of gas flow in protoplanetary disks and their impact on the transport of solids. Astron. Astrophys., 551, A75. Jacquet, E. (2014). The quasi-universality of chondrule size as a constraint for chondrule formation models. Icarus, 232, 176. Jacquet, E., Alard, O., and Gounelle, M. (2015). Trace element geochemistry of ordinary chondrite chondrules: The type I/type II chondrule dichotomy. Geochim. Cosmochim. Acta, 155, 47. Jacquet, E., Gounelle, M., and Fromang, S. (2012). On the aerodynamic redistribution of chondrite components in protoplanetary disks. Icarus, 220, 162. Jacquet, E., and Thompson, C. (2014). Chondrule destruction in nebular shocks. Astrophys. J., 797, 30. Jones, R. H. (2012). Petrographic constraints on the diversity of chondrule reservoirs in the protoplanetary disk. Meteorit. Planet. Sci, 47, 1176. Joung, M. K. R., Mac Low, M. -M., and Ebel, D. S. (2004). Chondrule formation and protoplanetary disk heating by current sheets in nonideal magnetohydrodynamic turbu- lence. Astrophys. J., 606, 532. King, A. R., and Pringle, J. E. (2010). The accretion disc dynamo in the solar nebula, Mon. Notices Royal Astron. Soc.. 404, 1903. Klahr, H., and Hubbard, A. (2014). Convective overstability in radially stratified accretion disks under thermal relaxation. Astrophys. J., 788, 21. Kley, W., and Lin, D. N. C. (1992). Two-dimensional viscous accretion disk models. I – On meridional circulations in radiative regions. Astrophys. J., 397, 600. Lodders, K. (2003). Solar system abundances and condensation temperatures of the elements. Astrophys. J., 591, 1220. Maxey, M. R. (1987). The gravitational settling of aerosol particles in homogeneous turbulence and random flow fields. J. Fluid Mech., 174, 441. McNally, C. P., and Hubbard, A. (2015). Photophoresis in a dilute, optically thick medium and dust motion in protoplanetary disks. Astrophys. J., 814, 37. McNally, C. P., Hubbard, A., Mac Low, M. -M., Ebel, D. S., and D’Alessio, P. (2013). Mineral processing by short circuits in protoplanetary disks. Astrophys. J. Lett., 767, L2. McNally, C. P., Hubbard, A., Yang, C. -C., and Mac Low, M.-M. (2014). Temperature fluctuations driven by magnetorotational instability in protoplanetary disks. Astrophys. J., 791,62. Muranushi, T. (2010). Dust-dust collisional charging and lightning in protoplanetary discs. Mon. Notices Royal Astron. Soc., 401, 2641. Muranushi, T., Akiyama, E., Inutsuka, S. -i., Nomura, H., and Okuzumi, S. (2015). Develop- ment of a method for the observation of lightning in protoplanetary disks using ion lines. Astrophys. J., 815, 84. Nelson, R. P., Gressel, O., and Umurhan, O. M. (2013). Linear and non-linear evolution of the vertical shear instability in accretion discs. Mon. Notices Royal Astron. Soc., 435, 2610. Oishi, J. S., and Mac Low, M. -M. (2009). On hydrodynamic motions in dead zones. Astrophys. J., 704, 1239. Evaluating Non-Shock, Non-Collisional Models for Chondrule Formation 427

Okuzumi, S., and Inutsuka, S. -i. (2015). The nonlinear Ohm’s Law: Plasma heating by strong electric fields and its effects on the ionization balance in protoplanetary disks. Astrophys. J., 800, 47. Ormel, C. W., and Klahr, H. H. (2010). The effect of gas drag on the growth of protoplanets. Analytical expressions for the accretion of small bodies in laminar disks. Astron. Astro- phys., 520, A43. Pilipp, W., Hartquist, T. W., and Morfill, G. E. (1992). Large electric fields in acoustic waves and the stimulation of lightning discharges. Astrophys. J., 387, 364. Pilipp, W., Hartquist, T. W., Morfill, G. E., and Levy, E. H. (1998). Chondrule formation by lightning in the Protosolar Nebula? Astron. Astrophys., 331, 121. Rubin, A. E., Sailer, A. L., and Wasson, J. T. (1999). Troilite in the chondrules of type-3 ordinary chondrites: Implications for chondrule formation. Geochim. Cosmochim. Acta, 63, 2281. Scherstén, A., Elliott, T., Hawkesworth, C., Russell, S., and Masarik, J. (2006). Hf W evidence for rapid differentiation of iron meteorite parent bodies. Earth Planet. Sci. Lett., 241, 530, Schrader, D. L., Fu, R. R., and Desch, S. J. (2016). Evaluating chondrule formation models and the protoplanetary disk background temperature with low-temperature, sub-silicate solidus chondrule cooling rates. Lunar and Planetary Science Conference XLVII, 1180. Semenov, D., Henning, T., Helling, C., Ilgner, M., and Sedlmayr, E. (2003). Rosseland and Planck mean opacities for protoplanetary discs. Astron. Astrophys., 410, 611. Shakura, N. I., and Sunyaev, R. A. (1973). Black holes in binary systems. Observational appearance. Astron. Astrophys., 24, 337. Shu, F., Najita, J., Ostriker, E., et al. (1994). Magnetocentrifugally driven flows from young stars and disks. 1: A generalized model. Astrophys. J., 429, 781. Shu, F. H., Shang, H., and Lee, T. (1996). Toward an astrophysical theory of chondrites. Science, 271, 1545. Takeuchi, T., and Lin, D. N. C. (2002). Radial flow of dust particles in accretion disks. Astrophys. J., 581, 1344. Tenner, T. J., Nakashima, D., Ushikubo, T., Kita, N. T., and Weisberg, M. K. (2015). Oxygen isotope ratios of FeO-poor chondrules in CR3 chondrites: Influence of dust enrichment and H2O during chondrule formation. Geochim. Cosmochim. Acta, 148, 228. Uesugi, M., Sekiya, M., and Nakamura, T. (2008). Kinetic stability of a melted iron globule during chondrule formation. I. Non-rotating model. Meteorit. Planet. Sci, 43, 717. Wadhwa, M., Srinivasan, G., and Carlson, R. W. (2006). Timescales of planetesimal differenti- ation in the early solar system, In D. S. Lauretta and H. Y. McSween (Eds.), Meteorites and the Early Solar System II, 715–731. Tucson, AZ: University of Arizona Press. Wasson, J. T. (1996). Chondrule formation: Energetics and length scales. In R. Hewins, R. Jones, and E. Scott (Eds.), Chondrules and the Protoplanetary Disk,45–54. Cambridge, UK: Cambridge University Press. Wood, J. A. (1963). On the origin of chondrules and chondrites. Icarus, 2, 152. Wurm, G., Trieloff, M., and Rauer, H. (2013). Photophoretic separation of metals and silicates: The formation of mercury-like planets and metal depletion in chondrites. Astrophys. J., 769, 78. Yu, Y., Hewins, R. H., Alexander, C. M. O’D., and Wang, J. (2003). Experimental study of evaporation and isotopic mass fractionation of potassium in silicate melts. Geochim. Cosmochim. Acta, 67, 773. Zanda, B., Zanetta, P. -M., Leroux, H., et al. (2017). The chondritic assemblage. LPI Contribu- tions, 1963, 2035. Zsom, A., Ormel, C. W., Güttler, C., Blum, J., and Dullemond, C. P. (2010). The outcome of protoplanetary dust growth: Pebbles, boulders, or planetesimals? II. Introducing the boun- cing barrier. Astron. Astrophys., 513, A57. 17 Summary of Key Outcomes

harold c. connolly jr., alexander n. krot, and sara s. russell

Abstract Here we review the conclusions from the book chapters and outline some ideas for directions of future research. We discuss what we know and do not know about chondrule precursors, the chronology of chondrule formation, physical (temperature, dust/gas ratio, total pressure) and chemical conditions (fO2, partial pressure of Na, S, SiO, ...) in chondrule-forming regions during chondrule formation, chondrule thermal histories, and chondrule formation models.

17.1 Introduction This book is focused on the earliest stages of planet formation, relating chondrules and their formation to other chondritic components (refractory inclusions and matrix), chondrites as a whole, and the astrophysical context of processes that may have shaped them. Researchers who study the formation of chondrites and their components often have a propensity for detailed and abundant data, which is actually critical to constraining the origins and evolution of chondritic components and chondrites in general. There is a clear need for the combined approach of using sample analysis, experimental studies, and numerical modeling to constrain the origins of chondrules and chondrites. In this chapter, we review the key scientific outcomes of each of the three major themes of the book – (1) Chondrule precursors: relationship between chon- drules, matrix, and refractory inclusions; (2) Chronology of chondrule formation, and (3) Mechanisms of chondrule formation – and discuss pathways forward for future research with an emphasis on understanding chondrites as a system that, through this system approach, can better constrain chondrule formation. At the end of each of the following sections, we present the major findings of each chapter relevant to the section.

17.2 Chondrule Precursors: Relationship Between Chondrules, Matrix, and Refractory Inclusions A hypothesis that there may be more than one chondrule formation mechanism (e.g., Krot et al., 2005), is becoming more openly embraced and discussed in detail than it has been in the past. If indeed chondrules were produced by more than one mechanism, the community needs to define which characteristics of chondrules and/or chondrites would help to constrain or define these mechanisms.

428 Summary of Key Outcomes 429

Another major issue identified as something requiring more research is defining if there is a relationship between the formation mechanism and the environment in which chondrules and other chondrite components formed within. As was discussed by Connolly and Desch (2004), chondrule melting mechanism may have nothing to do with the environment in which chon- drules formed. For example, the melting mechanism may either directly or indirectly control total pressure or the partial pressure of elements, but then again it may have nothing to do with such factors. It is important for the modeling of melting mechanism to define what characteris- tics of chondrites could have been directly formed through the chondrule (and igneous CAIs) melting process and which are not related. Considerable progress is being made on defining potential petrogenetic relationships between the major chondritic components – chondrules, refractory inclusions (CAIs and AOAs), and matrix. However, no unique solution or consensus was brought forward as to the origin of these components or a definitive understanding of their relationships with each other (e.g., formation conditions, place or time of formation, and so on). The relationship (if any) of organics, presolar grains, and ferromagnesian amorphous silicates within chondrites to other chondritic components and potentially to the formation of chondrules (or that of igneous CAIs) needs to be explored in more detail. Future work needs to emphasize a systematic approach to defining the relationship (if any) of all the components in chondrites to better constrain formation mechanisms and environmental conditions during formation and processing all the way to the accretion stage and geologically active period of parent bodies. Another important issue, for which no answer is yet provided, is the potential relationship of chondrite formation (and/or their components) to the formation of proto-gas giant planets, if there is a relationship. The Grand Tack hypothesis of Walsh et al. (2011) states that carbon- aceous chondrites could have formed in the outer solar system, outside Jupiter’s orbit. If so, was the inferred difference in the accretion regions of carbonaceous and noncarbonaceous chon- drites due to the formation of the proto-gas giant planets? Were chondrule formation and thermal processing of dust in general limited to the inner Solar System or were they ubiquitous throughout the protoplanetary disk? Chapter 2 (Krot et al.) summarized recent data on the mineralogy, petrography, oxygen- isotope compositions, and trace element abundances of precursors of chondrules and igneous Ca,Al-rich inclusions (CAIs). It is inferred that porphyritic chondrules, the dominant textural type of chondrules in most chondrite groups, largely formed by incomplete melting of isotopic- ally diverse solid precursors, including refractory inclusions (CAIs and amoeboid olivine aggregates (AOAs)), fragments of chondrules from earlier generations, and fine-grained matrix-like material during highly-localized transient heating events in dust-rich disk regions characterized by 16O-poor average compositions of dust (Δ17O~‒5‰ to +3‰). These observations preclude formation of the majority of porphyritic chondrules by splashing of differentiated planetesimals; instead, they are consistent with melting of dustballs during localized transient heating events, such as bow shocks and magnetized turbulence in the protoplanetary disk, and (possibly) by collisions between chondritic planetesimals. Like porphyritic chondrules, igneous CAIs formed by incomplete melting of isotopically diverse solid precursors during localized transient heating events. These precursors, however, consisted exclusively of refractory inclusions, and the melting occurred in an 16O-rich gas (Δ17O~‒24‰) 430 Harold C. Connolly JR., Alexander N. Krot, and Sara S. Russell of approximately solar composition, most likely near the protosun. There is no evidence that chondrules were among the precursors of igneous CAIs, which is consistent with an age gap between CAIs and chondrules. In contrast to typical (non–metal-rich) chondrites, the CB metal-rich carbonaceous chon- drites contain exclusively magnesian nonporphyritic chondrules formed ~ 5 Ma after CV CAI during a single-stage event, most likely in an impact-generated gas–melt plume. Bulk chemical compositions of CB chondrules and equilibrium thermodynamic calculations suggest that at least one of the colliding bodies was differentiated. The CH metal-rich carbonaceous chondrites represent a mixture of nonporphyritic chondrules and metal grains formed in the CB impact plume and typical chondritic materials (magnesian, ferroan, and Al-rich porphyritic chondrules, uniformly 16O-rich refractory inclusions, and chondritic lithic clasts) that appear to have largely predated the impact plume event. It is concluded that there are multiple mechanisms of transient heating events that operated in the protoplanetary disk during its entire lifetime and resulted in formation of chondrules and igneous CAIs. Identification of mineralogical, chemical, and isotopic characteristics of chondrules formed by a specific mechanism remains challenging. Chapter 4 (Hezel et al.) reviewed chemical and isotopic compositions of chondrules and matrices in various chondrite groups, as possible indicators of their genetic relationship. These show that chondrules and matrix in a chondrite group are chemically and isotopically comple- mentary. In other words, both have nonsolar (non-CI) compositions but when combined together are close to solar (CI) composition in many elemental and some isotopic ratios. Aqueous and metasomatic alteration processes on the chondrite parent bodies appear to have redistributed some elements and isotopes, but apparently are not responsible for most of the observed complementarities. These characteristics are interpreted as evidence for a close genetic relationship between chondrules and matrix in a chondrite group: (i) both components origin- ated from a common parental reservoir; (ii) matrix was one of the chondrule precursors; and (iii) significant, but still poorly defined, fraction of matrix was thermally processed (vaporized and recondensed) during melting of chondrule precursors. The inferred chondrule-matrix comple- mentarity supports the category of models in which chondrules and matrix form in the same location in the protoplanetary disk, and excludes models that require different locations for chondrule and matrix formation and later mixing of these two components. Chapter 5 (Zanda et al.) reviewed chondrites as a system to establish if a relationship exists between chondrules and matrix, as did Hezel et al. (Chapter 4), but came to a different conclusion concerning the nature of the relationship between the major structural components of chondrites. Through a review of the literature and new research, Zanda et al. suggest that chondrules and matrix are not genetically related, and thus did not form from a common reservoir within the protoplanetary disk. Instead, Zanda et al. see the overall relationship of the moderately volatile element abundances of bulk chondrites to solar composition as a result of exchange within a geologically active parent body between matrix and chondrules. It is proposed that bulk chondrite compositions are best reproduced by admixtures of generic chondrule compositions with CI- composition matrix. The findings of Zanda et al. suggest that chondrules and matrix could have formed at different times and locations within the protoplanetary disk, and thus long-range relative transport of these components cannot be ruled out by existing geochemical constraints. Chapter 7 (Jacquet et al.) showed that enstatite chondrites are similar to other chondrite groups in their broad texture, but are mineralogically quite distinct. While chondrules in all Summary of Key Outcomes 431 other chondrites are dominated by ferromagnesian olivine- and low-Ca pyroxene-bearing varieties, those in enstatite chondrites are unusual in having nearly pure enstatite (MgSiO3)as the main mineral. Chondrules from enstatite chondrites are remarkably reduced: olivine and pyroxene phenocrysts are nearly iron-free, and opaque nodules contain Ca-, Mn-, and Mg- bearing sulfides and Si-rich metal. Earlier work suggested that the mineralogy of enstatite chondrites resulted from condensation under highly reducing conditions in a region with a supersolar C/O ratio. Rather than high C/O ratios, Jacquet et al. prefer a model in which typical chondrule precursors (ferromagnesian silicates + Fe,Ni-metal + sulfides) have been melted in an environment rich in sulfur and poor in oxygen during the chondrule-forming process. The exact mechanism of producing the required enrichment in sulfur and depletion in oxygen remains unclear. However, irrespective of how enstatite chondrites formed, they are likely to be inner solar system products, based on their isotopic similarities to Earth. Chapter 8 (Tenner et al.) reviewed recent high-precision in situ oxygen–isotope measurements of mineral phases in individual chondrules from different chondrite groups using secondary ion mass-spectrometry. It is found that porphyritic chondrules from the least metamorphosed carbon- aceous chondrites (petrologic type <3.0) exhibit oxygen–isotope homogeneity among olivine, pyroxene, and plagioclase phenocrysts and mesostasis, and on a three-isotope oxygen diagram plot along slope-1 line; most relict grains are in isotopic disequilibrium with the chondrule phenocrysts. As chondrule melts experienced extensive gas–melt interaction during crystallization, these obser- vations indicate that: (1) porphyritic chondrules formed by incomplete melting of isotopically heterogeneous solid precursors dominated by 16O-poor dust; (2) oxygen–isotope ratios of chon- drules remained nearly constant as minerals crystallized from the host chondrule melt, suggesting that oxygen–isotope compositions of a given chondrule melt and its surrounding gas were similar; and (3) gas–melt oxygen–isotope exchange effectively erased evidence for kinetic mass-dependent fractionation. There are systematic differences in Δ17O values of chondrules from ordinary, enstatite, and carbonaceous chondrites. These observations are generally consistent with formation of chondrules in different dust-rich regions (dust/gas ratio 100‒1,000 Â solar) of the protoplanetary disk with limited mixing between them after chondrule formation. The nature of nonsolar, 16O-depleted composition of the disk dust remains poorly understood.

17.3 Chronology of Chondrule Formation In the first section, Krot et al. (Chapter 2) report petrographic and oxygen isotopic observations that tell us that refractory inclusions (CAIs + AOAs) were present in the regions where and when chondrules were formed, but no chondrules appear to have been present in the regions where and when refractory inclusions were formed. This is an important geological fingerprint, the importance of which cannot be ignored. The lack of chondrules in the region where and when refractory inclusions formed is in agreement with the short-lived 26Al–26Mg and 182Hf–182W isotope chronologies of chondrules and CAIs reported by Nagashima et al. (Chapter 9) and Kleine et al. (Chapter 10), who concluded that CAIs are older than the vast majority of chondrules by at least 1 Myr. However, the combined U-corrected Pb–Pb absolute and Al–Mg relative chronologies of still relatively small number of individual chondrules and CAIs (Connelly and Bizzarro (Chapter 11)) suggest that a high proportion of chondrules in a chondrite group has the same age as CAIs, yet other chondrules are younger (an important 432 Harold C. Connolly JR., Alexander N. Krot, and Sara S. Russell exception are CB chondrites containing young chondrules of a single generation, e.g., Bollard, Connelly, and Bizzarro, 2015). These apparent contradictory findings make a clear statement as to the state of our ability to have a consensus between the physical characteristics of chondrites to that of their isotope geochemistry: None exist and there appears to be no unique answer on the age differences between CAIs and chondrules. If some chondrules are older than others, we have to ask the question: How were they stored before mixing with younger chondrules to form the chondrites we observe today? This same question applies to refractory inclusions as well. Chapter 9 (Nagashima et al.) reviewed data for Al–Mg system in chondrules. The decay of the extinct nuclide 26Al to 26Mg can potentially be used as a chronometer for chondrule formation: however, there are two major issues to its use. Firstly, this isotope system is easily disturbed by asteroidal alteration, and so studies focusing on the least altered (type 3.0‒3.1) chondrites are most reliable. Secondly, the chronometer relies on an assumption of a homogen- ous distribution of 26Al/27Al in the early solar system, an assumption the authors point out requires further research. Different chondrite groups contain chondrules with apparently distin- guishable initial 26Al/27Al, with UOCs (unequilibrated ordinary chondrites)  COs  Acfer 094  CVs > CRs  CHs. The data imply that chondrules started to form around 1 Myr after CAIs and the chondrule-forming process continued for several Myr. Chapter 10 (Kleine et al.), in contrast to the conclusion of chondrites sampling chondrules of many separate ages and most ages are indistinguishable from those of CAIs (Connelly and Bizzarro, Chapter 11), showed through research on the Hf–W isotopic system that most chondrules from a single chondrite group may have formed around the same time, 2‒4 Myr after CAIs. For example, CV chondrules formed at around 2.2 Æ 0.8 Ma after CAIs, and CR chondrules formed distinctly later at 3.6 Æ 0.6 Ma after CAIs. These ages are compatible with results from 26Al–26Mg dating. Kleine et al. also reviewed data demonstrating that chondrules and matrix have comple- mentary nucleosynthetic tungsten isotope anomalies: chondrules tend to have an enrichment in ε183W (i.e., s-process depleted), and matrix has a depletion in this isotope (i.e. is s-process enriched), compared to bulk planets. This is probably because of a heterogeneous distribution of a presolar metallic component between chondrules and matrix, which the authors suggest occurred as a result of metal-silicate fractionation during chondrule formation. Chapter 11 (Connelly and Bizzarro) reviewed U-corrected Pb–Pb isotope data from CV, CR, and L chondrites indicating that chondrule formation started at around the same time as CAIs (~4,567 Ma) and continued for around 4 Myr. They suggested that an initial, highly active period for chondrule formation around 1 Myr after CAIs was followed by an extended period of chondrule recycling. Chondrites from a single group sampled chondrules with ages varying over several Myr. Chondrules from the CB meteorite Gujba all formed 4,564 Myr ago in the absence of protoplanetary disk material, indicating these unusual chondrules may have formed in one single event and the disk had dissipated by this time.

17.4 Mechanisms of Chondrule Formation The community does not appear any closer to providing a unique solution for a mechanism or mechanisms that could have produced chondrules, even though establishing the thermal Summary of Key Outcomes 433 histories chondrule experienced has made considerable progress. However, this is not discour- aging because it is clear that a kind of new genre of mechanisms is beginning to be taking much more seriously than a decade ago. Traditionally, within cosmochemistry, the kinds of mechan- isms that could have produced chondrules are divided into two broad categories: ‘nebular’ and ‘parent body’. A nebular mechanism means that chondrules (and by extension of the same reasoning, igneous CAIs) formed as free-floating wanderers before the accretion of chondrites. A parent body mechanism means that chondrules (and again by extension of the same reasoning, CAIs) were formed by the interaction (e.g., collisions, and so on) of planetesimals. The important implication of mechanisms that fall into each category is that if chondrules formed by the interaction of parent bodies, then they (and we cannot entirely eliminate igneous CAI formation by such mechanisms) are the by-product of planetesimal formation. However, if chondrules formed before accretion, then their formation and existence (and that of igneous CAIs) put them on the direct path to making parent bodies and, ultimately, planets. The latter seems to provide more significance to chondrules. However, the phrase ‘by-product of planet- esimal formation’ and ultimately planet formation is misleading. If chondrules formed by the interaction of rocky bodies, then they are the evidence of the dynamic evolution of these bodies, and thus of planet formation. Being the evidence of planet formation maintains the important scientific status of these igneous rocks in understanding the earliest time of solar system formation. A change founded in science and in philosophical thinking of the view and construct of the importance of chondrules being produced – either as pre-accretionary objects or as the recorders of the dynamical evolution of the earliest rocky bodies in the solar system – is important to embrace. Such a change in thinking provides key support for additional modeling efforts to define chondrule thermal histories through a variety of hypotheses found in post-accretion mechanisms, either known to have occurred (e.g., collisions) or hypothesized to have occurred (e.g., jetting). It is the responsibility of science to continue to explore all hypotheses for the formation of chondrules and other chondrite components, and that requires continued analysis of samples. Chapter 3 (Jones et al.) reviewed what is known about the thermal histories of chondrules. These have been well-constrained using experimentally produced analogues that can replicate the texture and chemistry of chondrules. They showed that porphyritic chondrules were heated to temperatures close to, but lower than, the liquidus. Nonporphyritic chondrules (such as barred and radial textures) formed at temperatures slightly exceeding the liquidus. The survival of relict grains in porphyritic chondrules requires cooling rates of up to thousands of degrees C/hr for at least some chondrules around the peak temperature. However, other constraints such as the presence of chondrule plagioclase suggest that cooling rates may have decreased (to less than 50 oC/hr) as the chondrules cooled. Overall, chondrules show variable cooling rates, even within single chondrites. These experimental data, being produced on an ongoing basis, provide some of the strongest constraints on chondrule-forming models. Chapter 6 (Ebel et al.) examined the volatile abundances and stable isotope compositions of chondrules. They reviewed evidence demonstrating that there was little evaporation occurring during chondrule formation, even for volatile elements such as sodium. In addition, the degree of isotopic fractionation in chondrules is very low in all elements measured, including Mg, Fe, Zn, K, and so on, pointing to suppression of open-system evaporative behavior. These 434 Harold C. Connolly JR., Alexander N. Krot, and Sara S. Russell observations imply that chondrules formed in an environment with a very high dust density, so that the chondrules and gas achieved equilibrium with each other. However, petrographic evidence for recondensation of silica to form pyroxene mantles in chondrule is at odds with this model, since silica is relatively refractory. Chapter 12 (Fu et al.) reviewed paleomagnetic studies of primary natural magnetization of isolated chondrules from ordinary and carbonaceous chondrites, which provide the potential basis for understanding the strength of nebular magnetic fields and the environment in chondrule-forming regions, and for constraining possible mechanisms of chondrule formation, including x-wind, magnetic reconnection flare, short circuit instability, shock wave, and impact jetting. Recent results are promising but still very limited. More detailed modeling and studies of primary natural magnetization in chondrules of different ages from meteorites of minimal secondary alteration from different chondrite groups can potentially constrain temporal and spatial coverage of the paleomagnetic records. Chapter 13 (Johnson et al.) explored the impact jetting model of chondrule formation, suggesting that chondrules formed by impacts between growing Moon- to Mars-sized planetary embryos. The predictions of this model are generally consistent with meteorite observations (chondrule sizes, abundance, thermal history, retention of volatile elements, lack of mass- dependent fractionation effects, chondrule-matrix complementarity), however, more detailed modeling of the jetting process is needed to evaluate this model properly. Chapter 14 (Sanders and Scott) summarized arguments in favor and against a splashing model of chondrule formation. According to this model, chondrules formed in the impact plumes generated by low-velocity collisions between molten planetesimals covered by a thin chondritic crust. The impact plumes were “dirty” (i.e., contaminated with earlier-formed dust from a variety of sources). The proposed concept of dirty plumes may explain the common presence of relict grains in porphyritic chondrules, and the observed W and Mo isotopic complementarity between the CV chondrules and matrix. Chapter 15 (Morris and Boley) reviewed different shock wave models of transient heating events in the protoplanetary disk, including (i) large-scale shocks due to disk gravitational instability and (ii) bow shocks. (i) Large-scale shocks appear to be consistent with the inferred cooling rates of chondrules and igneous CAIs, and could explain the nature of the earliest transient heating events near the protosun, the likely formation region of igneous CAIs. These models, however, have not yet been developed for disk regions with high dust-to-gas ratios invoked to explain the retention of volatiles in chondrule melts and the lack of mass-dependent fractionation effects in common and volatile elements (Ebel et al., Chapter 6). In addition, they may not be consistent with a prolonged duration of chondrule formation (Chapter 11; Chapter 9; Chapter 10). (ii) Chondrule-forming bow shocks could satisfy the chronology of chondrule- forming events, but require planetary embryos with magma ocean and outgassed atmospheres. This mechanism appears to produce cooling rates which are at the very upper end of the inferred cooling rates for chondrules (Chapter 3). Both models may be tested by future petrologic observations. It is concluded that: (1) both mechanisms could have resulted in transient heating events in the disk at different stages of its evolution capable of melting precursors of chondrules and igneous CAIs; (2) there could be multiple mechanisms of chondrule formation; and (3) determining the characteristics of chondrules formed by a specific mechanism becomes very important. Summary of Key Outcomes 435

Chapter 16 (Hubbard and Ebel) discussed several transient heating mechanisms, including lightning, radiant heating, and turbulent dissipation (note that planetesimal collision and shock mechanisms are discussed in Chapters 13–15); these mechanisms may operate in protoplanetary disks and might have been responsible for chondrule formation in the solar protoplanetary disk. These mechanisms are evaluated using (1) selected meteoritic constraints: (i) cooling rates of chondrule melts; (ii) full or partial preservation of isotopically unfractionated moderately volatile potassium and sodium in chondrules; and (iii) elemental and isotopic complementarity between chondrules and matrix within chondrites, and (2) the dynamic environment in which chondrules formed: (i) dust aerodynamics; (ii) dust coagulation; and (iii) turbulent transport. Hubbard and Ebel find that most mechanisms discussed in the chapter can be ruled out for chondrule formation, with only turbulent dissipation remaining as a possibility, subject to future work studying the actual thermal history of particles in disks with accurately modeled thermodynamics and radiative transport with temperature- and time-dependent opacities.

17.5 Conclusions While considerable progress has been made in the field since Chondrules and the Protoplane- tary Disk (Hewins et al., 1996) was published, there are still many major puzzles in our understanding of the formation of chondritic meteorites and their major components – chon- drules, matrices, and refractory inclusions. Much progress in science and the philosophy behind the science of chondrule formation has occurred over the last two decades. The modeling of formation mechanisms that involve collisions of planetary bodies has taken hold in a quantita- tive fashion (Asphaug et al., 2011; Johnson et al., 2015) and provided clues into potential mechanisms for forming chondrules. The first 3-D modeling of the shock wave mechanism has been reported (Boley et al., 2013). In addition, the science community has become open to the potential that chondrules may have formed by more than one mechanism—an idea that was not fashionable in any way 20 years ago. Continuing to define the physical, chemical, and chronological characteristics of chondrules, some of which may have formed by different mechanisms, becomes very important to constrain which chondrules may have formed by what mechanisms. Our level of constraining chondrites as a system has increased significantly. Petrology and geochemistry have played a key role in offering new potential constraints on the age of structural components of chondrites and of possible relationships of these components, and how those petrological and geochemical relations were established. It is left to be determined if chondrules or chondrites have any cosmic significance outside the solar system. If chondrule formation can be definitively linked to the dynamic evolution of early planetary bodies through collisions, thus making chondrules the natural products of accretion and planet formation, then with confidence they can be argued to be of cosmic significance in any solar system that contains rocky planets – something that was not even discussed in writing 20 years ago. Just the idea that chondrules and chondrites could have cosmic significance should be more food for increased attention to them and their origins. Finally, at the writing of this book the world was awaiting the arrival of Hayabusa2 and OSIRIS-REx at their respect target asteroids with the returned samples arriving on Earth in 436 Harold C. Connolly JR., Alexander N. Krot, and Sara S. Russell

2020 and 2023. Ultimately, these samples from the Hayabusa2 and OSIRIS-REx missions will provide exciting new insights into the chondrite-asteroid connection, the dynamic evolution of the protoplanetary disk, and the centuries-old problems of chondrule and chondrite formation.

References Asphaug, E., Jutzi, M., and Movshovitz, M. (2011). Chondrule formation during planetesimal accretion. Earth and Planetary Science Letters, 308, 369–379. Boley, A. C., Morris, M. A., and Desch, S. J. (2013). High-temperature processing of solids through solar nebular bow shocks: 3D radiation hydrodynamics simulations with particles. Astrophysical Journal, 776, 101‒124. Bollard, J., Connelly, J., and Bizzarro, M. (2015). Pb-Pb dating of individual chondrules from the CBa chondrite Gujba: Assessment of the impact plume formation model. Meteoritics and Planetary Science, 50, 1197‒1216. Connolly, H. C. Jr., and Desch, S. J. (2004). On the origin of the “Kleine Kügelchen” Called chondrules. Chemie der Erde-Geochemistry, 64,95‒125. Hewins R., Jones, R. H., and Scott, E. R. D. (Eds.). (1996). Chondrules and the Protoplanetary Disk. Cambridge, UK: Cambridge University Press. Johnson, B. C., Minton, D. A., Melosh, H. J., and Zuber, M. T. (2015). Impact jetting as the origin of chondrules. Nature, 517, 339–341. Kargel, J. S., Fegley, B. Jr., and Schaefer, L. (2003). Ceramic volcanism on refractory worlds: The cases of Io and chondrite CAIs. Lunar and Planetary Science Conference XXXIV, abstract #1964. Krot, A. N., Amelin, Y., Cassen, P., and Meibom, A. (2005). Young chondrules in CB chondrites from a giant impact in the early Solar System. Nature, 436, 989–92. Walsh, K. J., Morbidelli, A., Raymond, S. N., O’Brian, P., and Mandell, A. M. (2011). A low mass for Mars from Jupiter’s gas-driven migration. Nature, 475, 206–209. Index

achondrites, 2, 203, 251, 259–60, 288, 305, 363 Delta notation, 197 aerodynamic drag, 392 diopside, 16–17, 36, 38, 42–4, 50, 191, 243 albite, 19, 22, 242 dust aerodynamics, 408 ALMA, 376, 391, 401 dust coagulation, 409 Al-rich chondrules, 4 dusty olivine, 23, 70, 201–2, 332–5, 338, 374 amoeboid olivine aggregates, 5, 7, 11, 51, 70, 208, 258, 367, 429 Edward Howard, 1 amorphous, 75, 100–1, 119, 179, 429 eucrite, 305, 354, 359 angrites, 259–62, 265, 268, 271, 274, 288, 290, 305, 308, 319–20, 366 fayalite, 70, 212, 229, 232, 237–9, 270, 275, aqueous alteration, 3, 94, 105, 115, 117, 123, 158–9, 297–8 187, 199, 203, 213, 237, 252–3, 291, 293–4, 326, Fe,Ni metal, 1, 12, 17, 26, 29–30, 34, 37, 44, 61, 160, 331–2 225–6, 344, 431 atmospheres, 371, 383, 390, 392, 397, 434 feldspar, 4, 100, 253 fireballs, 5 basalts, 39, 41, 66, 76, 89 brachinite, 188 grain concentration, 415 breccia, 30, 114, 143, 179, 193, 252 Grand Tack, 233, 350, 429 granoblastic, 28–9 Ca,Al-rich inclusions, 5–6, 11–12, 14, 16–19, 21, 28, graphite, 178–9, 184 31, 35–7, 39, 42, 44–7, 49, 51–2, 92–3, 95, 100, Gustav Rose, 1 112–13, 117, 145, 148, 158, 161, 184, 190–1, 197, 208, 210, 225, 235, 237, 239, 241–2, 247–9, 251, helium, 383, 387 254–6, 258–65, 267, 269–71, 273–4, 279, 281, Henry Clifton Sorby, 2 288–90, 293–4, 296–8, 300–2, 305, 310, 313, 315, hydrogen, 228, 383, 387 317, 321, 324, 326, 328, 336, 344, 355–6, 362, 366–7, 373, 379–80, 396, 429–34, 436 ice, 215 cathodoluminescence, 23, 180, 208, 215 igneous rims, 23, 30, 219 CCAM line, 14, 181–2, 184, 197, 199, 203, 211, impact jetting model, 335, 343, 347, 352–3, 356, 434 214–15, 217, 368 interstellar material, 101, 112 chromite-rich chondrules, 28, 30 iron meteorites, 118–19, 171–2, 271, 277, 298, 344, chromium isotopes, 113, 119–20, 248, 259, 273, 285, 354, 361–2, 365, 371, 373, 391 288, 299, 310, 316–17, 322, 373–4 isochron, 247, 249, 256, 258, 279, 286, 297, 300, CO self-shielding, 197 303–4, 306, 308–11, 313, 319 comets, 5 Cosmogenic effects, 277 Jupiter, 12, 233, 350, 361, 365, 376 craters, 345 cristobalite, 176, 178 kamacite, 31, 178, 184, 192, 325, 327, 330 Knudsen cell, 34 dead zone, 49, 328 decay constant, 248, 302 lightning, 395, 401, 414, 418, 422–6

437 438 Index magnetite, 61, 101, 212–13, 232, 325 Parnallee, 362, 371 Magneto-Rotational Instability, 400, 402, 410, QUE 93351, 177 420–1 QUE 97008, 309 main belt, 188, 349–50 Renazzo, 54, 73, 84, 88, 102, 116, 118–19, 121, Mars, 2, 281, 318, 345, 356 126, 129, 133, 135–6, 138–40, 142, 148, 167–8, Marshak wave, 377 170, 243–5, 252, 335, 338, 398 MC-ICPMS, 174, 245, 248–9, 255, 257 Sahara 97159, 305 melilite, 5, 16–17, 19–21, 36, 38, 42–5, 52 Semarkona, 26, 49, 66, 68, 77, 83, 85–8, 98, 109, Mercury (Planet), 188 114–15, 148, 157, 159–61, 166, 168–71, 173, metamorphism, 3, 63, 65, 68, 72, 78–82, 85, 148, 167, 213–14, 216, 236, 239, 249, 253–4, 258–9, 195, 237, 265–6, 268, 279, 292, 294, 296–7, 299, 271, 298, 309, 332–7, 358, 366, 372, 395, 318, 402, 419, 423, 434–5 398, 426 Meteorites, named Vigarano, 28, 37, 49, 75–6, 99, 114, 124, 128, 149, Acfer 094, 31, 48, 51, 53, 55, 70, 86, 89, 100–1, 214, 220, 228, 248, 251, 258, 264 109, 116, 143, 147, 150, 181–2, 194, 198–200, microchondrules, 32, 109 202–3, 205, 208, 211, 213, 220, 224, 227, 233–6, minimum mass solar nebula, 326, 349, 402 242–5, 253–5, 267, 297, 321, 424, 432 ALH 81189, 177 nucleosynthetic isotope variations, 144–5, 186, 278, ALH 90411, 158 310 Allende, 36, 47, 49, 52–5, 74–5, 87–8, 94–5, 98–9, 105, 107–8, 115, 118, 120, 124, 126–7, 132–3, oldhamite, 175, 177–80, 184–6, 188–9, 144–5, 148–9, 153, 158, 161, 167–8, 170–1, 191–2 173–4, 182, 194, 198, 203, 205, 210–11, 220, organics, 100, 111, 154, 187, 231–2, 380 228–30, 236–8, 240, 242–3, 245, 251, 253, 256, 258, 264–6, 269–72, 274, 279–81, 286, 288–9, partition coefficients, 68 291, 296, 305–7, 310, 312, 320, 330 –1, 333–4, pebble accretion, 300–1, 318, 320, 322, 337–9, 354, 361, 367–72, 425 350 Bishunpur, 88, 98, 109, 114, 149, 167, 173, 213–14, perovskite, 5, 16–17, 20, 43, 113 236, 309, 336, 396, 398 phosphate, 30 Bjurböle, 98, 331 Planck opacity, 383 Chainpur, 98, 149, 161, 331, 425 potassium, 158, 170 Clovis, 158 presolar grains, 101, 108, 111, 113, 123, 144, 154, Efremovka, 49, 126, 128, 131, 133, 135–6, 139–40, 281–2, 284, 286, 380, 410, 429 148, 220, 228, 236, 251, 273, 310 primitive chondrule mineral line, 203 GRA 95208, 158 GRO 95551, 7, 15, 55, 206, 209, 218, 245 quartz, 176 Gujba, 39, 41, 47, 53, 300, 305, 307–8, 315, 320, 357, 432, 436 rare earth elements, 39, 68, 185 Isheyevo, 13, 16, 44, 48–9, 51, 53, 225, 239, 241, Rayleigh fractionation, 156, 159, 214, 406, 410, 417, 272, 322, 397 423 Ivuna, 123, 147, 305 refractory elements, 35, 39, 107, 124, 142–3, 152, Kaba, 128, 148, 237, 250–1, 253, 255, 264, 267, 180, 354 270–1, 273, 288, 290, 298 refractory metal nuggets, 101 Kainsaz, 126, 128, 132–3, 135–6, 138–9 relict grains, 4, 15–16, 24, 27, 30, 37, 53, 57–9, LEW 86134, 309 66, 70, 79–80, 89, 113, 173, 183, 202, MAC 88136, 177 205, 224, 235, 361–2, 364, 367, 371, Murchison, 98 431, 433–4 NWA 2976, 305 Robert Hutchison, 371 NWA 3118, 36, 38, 310 NWA 5492, 7, 15, 55, 206, 209, 218, 245 shock, 3 NWA 5697, 265, 305, 307–9 sodium, 19, 21, 49, 52, 82, 84, 116, 120, 164, 168–9, NWA 5717, 109 190, 334, 352, 357, 366, 372, 395, 399, 402, 406, NWA 6043, 307 414, 417, 425, 433, 435 NWA 7655, 307–8 Space missions NWA 801, 158 Hayabusa, 5, 7 Orgueil, 5–6, 124–7, 130, 134–5, 139–41, 149, 158, Hayabusa2, 435 231, 237 OSIRIS-REx, 435 Index 439

SQUID microscope, 336 type II chondrules, 4, 21, 23, 27, 30–1, 35, 55, 59, 61, Standard Mean Ocean Water, 16 66–71, 79, 88–9, 104, 109, 173, 178, 202, 207, storage, 312, 317 215, 217, 220, 223, 225, 227, 232–3, 238, 249, Stuart Agrell, 371 251, 359, 374, 398 sulfides, 1, 24, 26, 57, 61, 79, 82, 88, 116, 141, 147, 157–8, 171, 175, 177, 179, 183–6, 188–90, ureilite, 259–60 192–4, 207, 231, 252 ureilites, 363 textures of chondrules Vesta, 119, 354, 358–9, 372 barred olivine, 4 volatile elements, 4, 82, 92, 104, 107, 119, 122–3, 135, cryptocrystalline, 4, 12, 14–15, 37, 59–60, 199, 206, 142, 149–54, 157, 167, 171–3, 180, 255, 366, 209, 225–6, 236, 343 381, 414, 417, 433–4 glassy, 4, 343 porphyritic, 4, 15, 45, 57, 59, 62–3, 65–6, weathering, 3 225 radial pyroxene, 4, 14–15, 86, 218 x-wind model, 142, 283, 285, 295, 328, 330, 334, 401, titanium isotopes, 241, 365 407, 411, 413–14, 434 tridymite, 115, 176, 179 type I chondrules, 4, 23, 27, 31, 59, 61, 63, 65, 68–70, Young and Russell line, 197, 199, 203–4, 215, 217 73, 80, 82, 84, 90, 109, 163, 165, 180, 183, 205, 214, 220, 223, 225, 227, 233, 235, 245, 249, 398 zinc, 157

Figure 2.1 (a‒c, e) Combined x-ray elemental maps in Mg (red), Ca (green) and Al (blue) and x-ray elemental maps in Ni (d, f) of PCA 91082 (CR2), Tieschitz (H/L3), Hammadah al Hamra (HH) 237 (CB), and Isheyevo (CH/CB) chondrites. Typical chondrites consist of three major components: (1) chondrules + Fe,Ni-metal (Fe,Ni), (2) refractory inclusions (CAIs and AOAs), and (3) interstitial matrix (mx). In contrast, the metal-rich CB, CH/CB and CH carbonaceous chondrites lack matrix and contain abundant Fe,Ni-metal. Figure 2.5 Δ17O of CAIs incompletely melted during formation of CH porphyritic chondrules. These include (i) CAIs surrounded by chondrule-like ferromagnesian silicate igneous rims, (ii) relict polymineralic CAIs in chondrules, and (iii) objects intermediate between Type C-like CAIs and plagioclase-rich chondrules with clusters of relict spinel grains. Data from Krot et al. (2016a).

Figure 2.6 Oxygen–isotope compositions of porphyritic chondrules and CAIs surrounded by primordial Wark-Lovering rims (unaffected by chondrule melting events) from the CH and CB metal-rich carbonaceous chondrites. Data from Krot et al. (2010, 2017b). Figure 2.16 (a) Combined x-ray elemental map in Mg (red), Ca (green), and Al (blue) of the Kakangari (K) chondrite. Region outlined in (a) is shown in detail in (d). (b, c) Oxygen–isotope compositions of olivine and low-Ca pyroxene in chondrules and matrix from Kakangari. There are two isotopically distinct populations of matrix grains, 16O-rich and 16O-poor. Chondrules are dominated by 16O-poor silicates, but contain relict 16O-rich olivines isotopically similar to those in the matrix. Figure 2.16 (cont.) (d) Combined x-ray elemental map in Mg (red), Si (green) and Fe (blue), (e) BSE image, and (f, g) secondary ion images in “f” 27Al‒ and (g) δ18O of a porphyritic olivine-pyroxene chondrule and its coarse-grained igneous rim composed of magnesian olivine, low-Ca and high-Ca pyroxenes, and mesostasis. The rim contains abundant relict 16O-rich olivine grains (some of them are indicated by red arrows) overgrown by 16O-poor olivines. cpx = high-Ca pyroxene; mes = mesostasis; met = Fe,Ni-metal; ol = olivine; px = low-Ca pyroxene. Data from Nagashima et al. (2015b). Figure 2.17 Combined x-ray elemental maps in Mg (red), Ca (green), and Al (blue) of the complex CV CAIs (a) SJ101 from Allende and (b) 3N from NWA 3118. Region outlined in (b) is shown in BSE in (c). (a) CAI SJ101 consists of the pyroxene-melilite-anorthite-spinel and forsterite-pyroxene lithologies surrounded by a forsterite-free pyroxene-anorthite-spinel mantle. The CAI formed by melting of a Type B-like CAI surrounded by forsterite-rich accretionary rim (after Bullock et al., 2012). (b, c) CAI 3N encloses of about 25 relict igneous CAIs of different types – Type B, Type C, compact Type A, and ultrarefractory (UR). It formed by melting of a dustball composed of multiple CAIs and forsterite-rich dust (after Ivanova et al., 2015). (d) Oxygen–isotope compositions of individual minerals in the relict ultrarefractory CAI 3N-24 and the host forsterite-bearing CAI 3N. The relict CAI is 16O-depleted relative to the host inclusion. Data from Ivanova et al. (2012). px = Al,Ti-diopside; an = anorthite; fo = forsterite; mel = melilite; sp = spinel; Zr,Sc-px = Zr-Sc-rich pyroxene; Zr,Sc,Y-ox = Zr-Sc-Y-rich oxides. Figure 2.18 Combined x-ray elemental maps in Mg (red), Ca (green), and Al (blue) overlaid on BSE images of a coarse-grained igneous CAI composed of Type A, B, and C portions from Vigarano (CV). The Type C mantle contains abundant chondrule fragments (indicated by yellow arrows in (c)). From MacPherson et al. (2012).

Figure 3.9 Intergrowth of clinoenstatite (CEn) and orthoenstatite (OEn) in porphyritic chondrule 42 from Vigarano CV3 chondrite. a) BSE image showing the fracturing network in low-Ca pyroxene (arrows). b) EBSD image (Euler angles) showing polysynthetic twins in the clinoenstatite. Blue and green colours show CEn twin orientations. c) EBSD map of the entire chondrule. Clinoenstatite: yellow; orthoenstatite: gray; olivine: green; Fe-Ni metal: red. Figure 2.19 (a) Combined x-ray elemental map in Ti (red), Ca (green), and Al (blue), and (b, c) BSE images of a plastically deformed coarse-grained igneous Compact Type A CAI from the CV chondrite NWA 3118. Regions outlined in (a) are shown in detail in (b) and (c). Hibonite-rich mantle is thicker on the convex side of the CAI than on the concave side, suggesting asymmetric heating of the convex and concave sides. cpx = Al,Ti-diopside; hib = hibonite; mel = melilite; sp = spinel. Figure 2.22 (a, c) Combined x-ray elemental maps in Mg (red), Ca (green), and Al (blue), (b) x-ray elemental map in Ti, and (d‒f ) BSE images of the uniformly 16O-depleted igneous CAIs surrounded by igneous rims composed of Al-diopside and forsterite. Regions outlined in (c) and (e) are shown in detail in (d) and (f ), respectively. Numbers correspond to CAI numbers listed in Figure 2.21. cpx = high-Ca pyroxene; fo = forsterite; grs = grossite; hib = hibonite; mel = melilite; sp = spinel. Figure 2.23 (a) Oxygen–isotope compositions of magnesian nonporphyritic (SO and CC) chondrules and 16O-depleted (relative to typical CAIs carbonaceous chondrites having Δ17Oof~‒24‰) igneous CAIs surrounded by igneous rims from the CH and CB carbonaceous chondrites. Most chondrules have similar Δ17Oof~‒2.5‰; the CAIs are 16O-enriched relative to the chondrules. Individual CAIs and their igneous rims have similar O-isotope compositions; there are, however, inter-CAI compositional differences. It is suggested that magnesian CC and SO chondrules in CB and CH chondrites formed in a gas–melt plume generated by a collision between planetesimals, while the rimmed CAIs resulted from complete melting of preexisting CAIs followed by gas–melt interaction in the plume. The terrestrial fractionation line (TF) is shown for reference. Numbers correspond to numbers of individual inclusions; some of them are shown in Figure 2.20. (b) Oxygen–isotope compositions of magnesian porphyritic chondrules 1573–4-9 and 1–3-3 with coarse relict 16O-depleted spinels, and uniformly 16O-depleted spinel-rich igneous CAIs 1–3-4 and 1573–1-1 surrounded by igneous rims. These chondrules and CAIs are shown in Figure 2.22. Oxygen- isotope compositions of the relict spinel grains are similar to those of the uniformly 16O-depleted igneous CAIs in CB and CH chondrites. cpx = high-Ca pyroxene; fo = forsterite; grs = grossite; hib = hibonite; mel = melilite; ol = ferromagnesian olivine; pl = plagioclase; pv = perovskite; px = low-Ca pyroxene. Data from Krot et al. (2010, 2017a, 2017b); Krot, Nagashima, and Petaev (2012).

2.0

1.8 Fe 1.6

1.4 Ca, Al Si 1.2

1.0

0.8

0.6 CI- and Mg-normalized Zn 0.4

0.2

0.0 SunCICMCVCOCKCRCHH L LLEHELR

Figure 4.2 Differences and similarities among the solar photosphere (Sun) and bulk chondrite element compositions. Data taken from: Wasson and Kallemeyn (1988); Wolf and Palme (2001). Red: Ca; Green: Al; Blue: Si; Red: Fe. Figure 4.8 Proposed sequence of events and processes in the protoplanetary disk. min: minutes; h: hours; comp. est. by e.g. incorp. of: complementarity established by e.g., incorporation of.

Figure 5.1 Chondrite and Orgueil normalized concentrations of major and minor elements in the main carbonaceous chondrite groups. The right panel lists 50% equilibrium condensation temperature from Lodders (2003). Data from Wolf and Palme (2001) and Wasson and Kallemeyn (1988). After Palme (2001). Figure 4.3 Bulk chondrule, matrix, and chondrite element compositions of four different carbonaceous chondrites. Chondrules typically have superchondritic refractory and subchondritic volatile element concentrations. The opposite is observed in matrix. Figure taken from Hezel et al. (2017a). Figure 4.5 Examples of element chondrule-matrix complementarities. Data taken from: a) Hezel et al., 2008; c) Palme et al., 2015; d) Becker et al., 2015; Budde et al., 2016a. Figure b) taken from Ebel et al., 2016. Figure 7.1 Mg-Ca-Al (red-green-blue) composite element map of the (a) Allan Hills (ALH) 81189 (,3) EH3 chondrite and (b) Queen Alexandra Range (QUE) 93351 (,6) EL3 chondrites. BSE images of porphyritic chondrules in ALH 81189 (c) and MAC 88136 (EL3, d). En = enstatite, Ol = olivine, Ca-pyx = Ca-pyroxene, Mes = albitic mesostasis, Troil = troilite, Daub = daubréelite.

Figure 7.3 Representative rare earth element patterns of enstatite chondrite chondrule phases. (See detailed caption with data sources in main text). Figure 7.4 (a) Oxygen 3-isotope diagram showing SIMS data for chondrules in E3 (EH3 and the anomalous EC Lewis Cliff 87223) chondrites compared to other chondrites. Also shown are the terrestrial fractionation (TF), carbonaceous chondrite anhydrous mineral (CCAM) and enstatite chondrite mixing (ECM) lines. (See detailed caption with data sources in main text).

Figure 12.4 (a,b) Reflected light optical and QDM magnetic field maps showing a barred olivine chondrule in the Allende meteorite. (c,d) Transmitted light optical and QDM magnetic field maps of an isolated dusty olivine-bearing chondrule from the Semarkona meteorite. Adiabatic, 7 km/s Wind 8 14000 7 12000 6 10000 5 8000 4 6000 3 Distance (km) 4000 2 Density/(Wind Density)

2000 1

0 0 -15000 -10000 -5000 0 5000 10000 15000 Distance (km)

Adiabatic, 7 km/s Wind

14000 2000 12000

10000 1500

8000

6000 1000 Distance (km)

4000 Gas Temperature (K)

2000 500

0 -15000 -10000 -5000 0 5000 10000 15000 Distance (km)

Figure 15.4 The adiabatic shock structure around a 3000 km radius protoplanet that is roughly half the mass of Mars. The simulation, run using the methods of Boley et al. (2013) and Mann et al. (2016) assumes that the relative wind speed Vw is 7 km/s and that the gas is adiabatic (no cooling or heating sources, e.g., radiation). The equation of state for the gas uses a mixture of molecular hydrogen/hydrogen, helium, and metals. The rotational and vibrational modes for H2 are included for an ortho:para mixture of 3:1, and H2 dissociation is also included. The preshock wind is assumed to be at a temperature T = 300 K and to have a pre-shock gas density ρ =10–9 g/cc. The simulation has a resolution of 37.4 km. Adiabatic, 7 km/s Wind 100 14000

12000 80

10000 bar 60 µ 8000

6000 40 Distance (km)

4000 Gas Pressure 20 2000

0 0 -15000 -10000 -5000 0 5000 10000 15000 Distance (km)

Figure 15.5 Similar to Figure 15.4, but for the gas pressure with particle trajectories superposed to illustrate the flow of solids through the bow shock. Solids penetrate into the postshock region before recoupling to the gas flow. Most solids are diverted around the protoplanet, but the ones with the smallest impact angles are accreted. The slight change in the trajectories just before the shock encounter is due to the gravity of the protoplanet.

Figure 15.7 The vertical bars show allowed chondrule cooling rates, for different chondrule textures, based on furnace experiments. The horizontal bars show cooling rates derived from protoplanet bow shock simulations for different assumptions about the radiative cooling. The adiabatic cases are consistent with many chondrule textures, but these require an extreme optically thick environment. The other bars labelled “Rad X, k=y” represent different assumptions for the dust and chondrule opacities (see Mann et al. (2016) for details). Radiative simulations for “normal” opacities show cooling rates that are much larger than the furnace experiment constraints. While one radiative simulation does show some potential overlap with the experiments, the cooling rate might only be attainable for unrealistic thermal coupling between the gas and the melts.