Advances in Isotope Geochemistry

Paul B. Tomascak Tomáš Magna Ralf Dohmen Advances in Lithium Isotope Geochemistry Advances in Isotope Geochemistry

Series editor Jochen Hoefs, Göttingen, Germany More information about this series at http://www.springer.com/series/8152 Paul B. Tomascak • Tomáš Magna Ralf Dohmen

Advances in Lithium Isotope Geochemistry

123 Paul B. Tomascak Ralf Dohmen State University of New York at Oswego Institute of Geology, Mineralogy Oswego, NY and Geophysics USA Ruhr-University Bochum Bochum, Nordrhein-Westfalen Tomáš Magna Germany Czech Geological Survey Prague Czech Republic

ISSN 2364-5105 ISSN 2364-5113 (electronic) Advances in Isotope Geochemistry ISBN 978-3-319-01429-6 ISBN 978-3-319-01430-2 (eBook) DOI 10.1007/978-3-319-01430-2

Library of Congress Control Number: 2015953806

Springer Cham Heidelberg New York Dordrecht London © Springer International Publishing Switzerland 2016 This work is subject to copyright. All rights are reserved by the Publisher, whether the whole or part of the material is concerned, specifically the rights of translation, reprinting, reuse of illustrations, recitation, broadcasting, reproduction on microfilms or in any other physical way, and transmission or information storage and retrieval, electronic adaptation, computer software, or by similar or dissimilar methodology now known or hereafter developed. The use of general descriptive names, registered names, trademarks, service marks, etc. in this publication does not imply, even in the absence of a specific statement, that such names are exempt from the relevant protective laws and regulations and therefore free for general use. The publisher, the authors and the editors are safe to assume that the advice and information in this book are believed to be true and accurate at the date of publication. Neither the publisher nor the authors or the editors give a warranty, express or implied, with respect to the material contained herein or for any errors or omissions that may have been made.

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Springer International Publishing AG Switzerland is part of Springer Science+Business Media (www.springer.com) Preface

In this book we decided to attach the permil sign (‰) to all Li isotopic quantities. One way of viewing stable isotopes denoted by δ is that the arithmetic sets the results as being part-per-thousand quantities, so to place the ‰ on a value is redundant. However, this implies a certain familiarity from the reader. Our decision regarding the ‰ in this volume was guided by the potential that the audience may include those not so steeped in the thinking of stable isotopes. This calls to mind a historical note regarding Li isotopes. Readers of the early literature on the subject (beginning with Chan in 1987) will find papers that use δ6Li. Prior to 2000, using the now-accepted δ7Li notation was viewed as an unwanted usurpation by at least one prominent geochemist. Nevertheless, being clear is important, and although δ7Li was not the first notation employed, it follows stable isotope convention. We find that students have a hard enough time understanding isotope geochemistry, so to oppose the notation used in virtually all systems (positive values are isotopically heavier than negative values) invites con- fusion. Hence, our use of the ‰ is a further step to make this compilation clear for all.

v Acknowledgements

Once upon a time Jochen Hoefs approached one of us (P.T.) about putting together a book for this Springer Advances series. Thus began an unex- pectedly long sojourn, during which the use of Li isotopes has come of age. Jochen had faith enough in us to persevere (keeping in mind the invitation to write this came in 2008!). For this opportunity and for his overall encour- agement we are extremely thankful. Two other names that need to be mentioned prominently are those of Lui-Heung Chan and Jan Košler, who passed away in 2007 and 2014, respectively. Anyone who works with Li isotopes owes Lui a debt of grat- itude. Her passing saddened all with whom she had come in contact and was a great loss for the geochemical community as a whole. Jan was the M.Sc. supervisor for T.M., who brought him the idea of developing Li analytical chemistry for mantle rocks and who, with time, became a close friend. Jan’s scientific contributions, especially to the ICP-MS community, are diverse and lasting. His presence as a colleague and friend will be badly missed. Throughout the construction of this work we have asked for input of various kinds from a number of individuals, who have graciously helped. To this end, we would like to thank Natalie Caciagli-Warman, Rick Carlson, Alain Coc, Brian Fields, C. Eric Hellquist, Nick Prantzos, John Rakovan, Gary Steigman, Elisabeth Vangioni, Aurelie Verney-Carron and Nathalie Vigier for their assistance. We have received a high level of support and almost unreasonable posi- tivity from the editorial staff at Springer. How they had the patience to put up with our constant delays speaks of their desire to see the project through to completion. In particular we thank Janet Sterritt and Annett Buettner. R.D.: I would like to thank Joana Polednia for help with compiling the experimental data. T.M.: I wish to acknowledge my family in the first place, as their support is the most prominent and most important everlasting key for my human being. Numerous colleagues with whom I have had multiple disputations and discussions over the years deserve thanks for their inspiration. While not possible to list all of these, here are some of the prominent ones: Lukáš Ackerman, James Day, Alex Deutsch, Vojtěch Erban, Tim Grove, Nikolaus Gussone, Alex Halliday, Darrell Harrison, Vojtěch Janoušek, Timm John, Klaus Mezger, Fred Moynier, Clive Neal, Felix Oberli, Andreas Pack, Franck Poitrasson, Vladislav Rapprich, Erik Scherer, Michael Seitz, Zach

vii viii Acknowledgements

Sharp, Tony Simonetti, Andreas Stracke, Larry Taylor, Dewashish Upad- hyay, Uwe Wiechert, Rainer Wieler and Karel Žák. I am grateful to the Czech Science Foundation for partial support through projects P210/12/1990, 13-22351S and 15-08583S. P.T.: I thank my university for supporting my ability to continue to conduct research and to the many students with whom I have had the pleasure to interact over the years. The goal of bringing geochemistry from the hazy and somewhat scary distance to clear focus continues to be a beloved challenge, which is made worthwhile by passionate students. Contents

1 Introduction...... 1 References ...... 4 2 Methodology of Lithium Analytical Chemistry and Isotopic Measurements...... 5 2.1 Historical Perspective ...... 5 2.2 Lithium Isolation by Cation Exchange Chromatography ...... 6 2.3 Methodology of Li Isotopic Measurements ...... 8 2.3.1 Mass Spectrometry-Based Methods of Widest Use ...... 9 2.3.2 “Other” Analytical Methods ...... 13 References ...... 14 3 Cosmochemistry of Lithium ...... 19 3.1 Cosmology of Lithium: Big Bang Nucleosynthesis, Evolution of the Universe, and Stellar Processes ...... 19 3.1.1 The Earliest History and Further Evolution of Lithium ...... 19 3.1.2 The Vicinity of the Solar System ...... 23 3.1.3 Searching for Exoplanets with Lithium ...... 25 3.2 Cosmochemistry of Lithium in Meteorites ...... 26 3.2.1 Lithium in Bulk Chondritic Meteorites ...... 27 3.2.2 Lithium Distribution and Isotopic Systematics in Constituents of Chondritic Meteorites ...... 30 3.2.3 Lithium in Other Meteorites ...... 31 3.3 Cosmochemistry of Lithium in Planetary Bodies ...... 31 3.3.1 Lithium in Earth’s Moon ...... 32 3.3.2 Lithium in Mars ...... 35 3.3.3 Lithium in Vesta ...... 38 References ...... 40 4 Li Partitioning, Diffusion and Associated Isotopic Fractionation: Theoretical and Experimental Insights ..... 47 4.1 Introduction...... 47 4.2 Equilibrium Partitioning Behavior of Li ...... 49 4.2.1 Thermodynamic Background ...... 49 4.2.2 Lattice Strain Model for Element Partitioning . . . . 49 4.2.3 Mineral-Melt Partitioning ...... 52

ix x Contents

4.2.4 Mineral-Fluid Partitioning and Associated Stable Isotopic Fractionation ...... 67 4.2.5 Inter-mineral Partitioning ...... 76 4.3 Diffusion of Li and Associated Effects on δ7Li...... 80 4.3.1 Theoretical Background ...... 80 4.3.2 Experimental Diffusion Coefficients ...... 83 4.3.3 Characteristic Time Scales of Li Diffusion in Silicates (Ol, Cpx, Pl)...... 91 4.3.4 Two Diffusion Mechanisms of Li in Silicates . . . . 93 4.3.5 Bulk Diffusion of Li Through Rocks and Characteristic Length Scales ...... 97 4.4 Fluid-Rock Interaction ...... 98 4.4.1 Weathering of Igneous Rocks (Predominantly Basaltic Compositions) ...... 99 4.4.2 Weathering of Olivine (Serpentinization)...... 109 4.4.3 Alteration of Sediments ...... 110 4.4.4 Synopsis ...... 111 References ...... 112 5 Lithium in the Deep Earth: Mantle and Crustal Systems . . . 119 5.1 Lithium in the Mantle...... 119 5.1.1 Mantle Xenoliths and Related Samples ...... 119 5.1.2 Mantle-Derived Magmas...... 128 5.2 Lithium in the Continental Crust ...... 141 5.2.1 Products of Crustal Melting...... 141 5.2.2 Crustal Metamorphism and Fluid Flow ...... 146 References ...... 149 6 The Surficial Realm: Low Temperature Geochemistry of Lithium ...... 157 6.1 Meteoric Water ...... 157 6.2 Groundwater and Pore Waters ...... 158 6.2.1 Shallow Groundwaters ...... 158 6.2.2 Deep Groundwaters ...... 159 6.2.3 Pore Waters ...... 161 6.3 Continental Weathering...... 162 6.3.1 Soils and Sediments ...... 163 6.3.2 Sedimentary Rocks ...... 167 6.3.3 Lakes and Rivers: Dissolved Loads ...... 168 6.4 Hydrothermal Systems ...... 171 6.4.1 Rocks in Sea Floor Hydrothermal Systems ...... 171 6.4.2 Fluids in Sea Floor Hydrothermal Systems ...... 173 6.4.3 Continental Geothermal Systems ...... 174 6.5 Marine Geochemistry of Li ...... 176 6.5.1 The Marine Li Cycle ...... 176 6.5.2 Secular Variation of Li in Seawater ...... 177 Contents xi

6.6 Biological and Anthropogenic Li Isotopic Fractionation ...... 181 6.6.1 Lithium in Plants and Animals ...... 181 6.6.2 Tracer Applications in Water Treatment and Agriculture ...... 183 6.6.3 Nuclear Forensics ...... 184 References ...... 184 7 Summation: What Have We Learned and Where Can We Go? ...... 191 7.1 Retrospective...... 191 7.2 Analytical Prospects ...... 194 7.3 Final Words ...... 194 References ...... 194 Introduction 1

Lithium (Li) was discovered by J. August industry, organic and polymer chemistry, the Arfvedson in 1817 (Berzelius 1817). The newly nuclear industry, as well as in the production of isolated alkaline compound in the Swedish grease and lubricants, and as a desiccant (halo- petalite sample was given the Greek name for gens, hydroxide, peroxide, perchlorate) in air stone: λίθος (lithos); the metal was dubbed cleaning devices for submarines, spacecraft, etc. lithium. In the following year Sir Humphry Davy Lithium ion batteries constitute a major and described the isolation of metallic Li using growing industrial channel for the element, with electrolysis (Anonymous 1818), although this the proliferation of mobile electronic devices and development is commonly considered contem- the diversification of electric-powered vehicles. poraneous with the work of Brande (Weeks These demands for Li are reflected in a dramatic 1956). Using a more modern electrolytic process, increase in Li production during the last two Bunsen and Matthiessen succeeded in isolating decades (Fig. 1.1). Earlier commercial exploita- multi-gram quantities of Li metal (Bunsen 1855). tion of Li came from primary magmatic sources, Commercial production of Li started only after such as large pegmatite bodies with Li-bearing World War I. A more massive production com- minerals (amblygonite, eucryptite, lepidolite, menced after World War II with the findings of petalite, spodumene), and less commonly from potential Li utility in the nuclear industry. evaporites. More recently exploitation of brine Lithium as an element has a number of pools (salars) have become by far the dominant peculiar properties. For example, it is the lightest source of the element, with principal resources of all solids of the periodic table of elements with currently in Australia, China, and South Amer- a density of just 0.543 g cm−3, meaning that it ica. Brine resources appear to be the stable pro- floats on the water. It also has the highest specific ducers of Li also in the near future (Ebensperger heat capacity of all solids (5.38 kJ kg−1 K−1). Its et al. 2005; Kesler et al. 2012) although alter- melting and boiling temperatures (180.5 and native sources of Li as a side product of other 1342 °C, respectively) are the highest of alkali economic activities are also being explored (Qin metals of group 1A of the periodic table and, et al. 2015). Future demands appear likely to similarly, its first ionization energy (5.39 eV) is continue to grow, and the recycling of the ele- also higher compared with other elements of this ment may become an integral part of the Li group. business (Mohr et al. 2012). Because of some of these properties, the Just over a decade has passed since the pub- utility of Li has found its peak applications in the lication of a broadly inclusive summary of Li ceramic industry, electronics and electrical isotopic research in the world (Tomascak 2004). industry, metallurgy, medicine, the optical Despite the short time, the use of this isotopic

P.B. Tomascak et al., Advances in Lithium Isotope Geochemistry, 1 Advances in Isotope Geochemistry, DOI: 10.1007/978-3-319-01430-2_1 © Springer International Publishing Switzerland 2016 2 1 Introduction

The title of this work accurately defines our 600 focus, which is primarily the compilation and interpretation of the isotopic results in geo- and

400 cosmochemical studies. Lithium as an element has its own important role in studies of these kinds. Although we touch on this role in places 200 (necessarily in the discussion of laboratory

tonne Li /year (× 1000) experiments, Chap. 4), the reader is referred to the literature for additional detail on Li—the 1940 1980 2020 element—as used in the Earth sciences. Addi- Year tionally, Li is of great importance in many other fi Fig. 1.1 Worldwide production of lithium. A dramatic elds, ranging from ecology, biomedicine, and increase after World War II was caused by the demand for pharmacology, to materials, energy, and nuclear Li in the nuclear industry, whereas more recent increase science and engineering. The reader is again reflects rapid diversification of commercial applications. directed to the primary literature of these areas Data from Jaskula (2015) for information. This volume is broken down along conven- system in the investigation of geo- and cosmo- tional lines. We have endeavored to the greatest chemical questions has truly exploded, hinging extent to repeat as little as possible from previous in part on the arrival of analytical technology at compilations, attempting to integrate and high- the close of the millennium (Fig. 1.2). Prior to light newer findings. We begin in Chap. 2 with a 2004, approximately 50 publications with Li summary of the analytical techniques used, past isotope data existed in the accessible and present, to isolate and measure Li isotopic peer-reviewed literature, of which almost 30 % ratios. Unlike many other fields in analytical were purely analytical studies (i.e., not specifi- geochemistry, the community has not converged cally addressing research questions, but rather on a single “standard operating procedure” when reporting laboratory methods). Articles with it comes to either the chemical separation and dates of 2004 and later currently number over purification of the element or to the quantifica- 220. Technically-based studies have not disap- tion of its isotopes. There remains a variety of peared in the era of advanced analytical capa- methods still in wide use. bilities (about 6 % of the post-2003 studies are The summary of Li isotope cosmochemistry principally technique articles), but clearly many in Chap. 3 sets the stage for the deeper focus of of the earlier analytical issues have been over- Li in our planet. A general outlay of Li in the come, facilitating the current research boom. cosmos is established, giving the major back- In response to this explosion, we have ground necessary to appreciate Li in the terres- endeavored to accumulate the trove of new trial neighborhood. This chapter attempts to cast findings into an updated summary of the state of the ultimate origins of the isotopes of Li in terms Li isotopic science. Although this engenders a that are understandable for a non-astrophysicist degree of repetition of results and findings from audience. existing summary works (e.g., Burton and Vigier Chapter 4 takes as its focus the synthetic 2011; Chan 2004; Elliott et al. 2004; Tang et al. studies that are designed to help understand Li 2007; Tomascak 2004), the focus of this com- behavior in natural settings. In order to use Li pilation is to distill the more modern studies and isotopes in a quantitative way in natural systems, to highlight major developments to date. Given constraints must exist from experiments, and as that the history of the development of Li isotopic we probe deeper into natural samples the demand geochemistry is covered in detail in other works, for accurate and meaningful laboratory con- our goal here is to concentrate on modern straints intensifies. The diversification of experi- research and the directions it currently takes. mental approaches to understanding Li isotopes 1 Introduction 3

Fig. 1.2 Lithium isotopic composition of various extra-terrestrial and terrestrial reservoirs. The details of individual reservoirs and δ7Li variations are discussed throughout the volume. For marine sediments, samples with obvious metamorphic effects are excluded (see Chap. 6). Soil samples include loess and similar detritus. For river waters and continental groundwaters and formation waters, only samples without discernible anthropogenic influence were included. Also, for river waters, those affected by hydrothermal input were ommitted (see Chap. 6)

-30 -20 -10 0 10 20 30 40 50 60 δ7Li (‰) has been great in recent years, particularly with rocks. As the mantle is the primary reservoir of the acceptance that diffusion produces, in many Li in the Earth, it is critical to understand its bulk settings, effects that must be considered. This Li isotopic composition in order to guide solar chapter summarizes existing results and draws system-scale interpretations. The bulk composi- attention to where conclusions drawn from lab- tion of the mantle and the transit of Li between oratory and theoretical studies appear to most mantle and crustal reservoirs are major topics significantly affect the interpretation of data from addressed in this chapter. natural systems. The fast pace of change in the In Chap. 6 we consider the many studies that field as a whole dictated by experimental studies deal with the domain in which the bulk of Li is highlighted. isotope fractionation takes place on Earth: The solid Earth is the target of Chap. 5, with near-surface processes. Since stable isotopic emphasis on the interpretation of data from fractionation predominates in conditions of low mantle and crustal igneous and metamorphic temperature, the bulk of the Li isotopic “action” 4 1 Introduction in the Earth system can be said to exist within the Chan LH (2004) Mass spectrometric techniques for the narrow fringe where rock, water, and atmosphere determination of lithium isotopic composition in geo- logical material. In: de Groot PA (ed) Handbook of closely coexist. Applications of Li isotopes in stable analytical techniques, vol 1, issue 6, pp 122–141 this realm are manifold and extend to studies of Ebensperger A, Maxwell P, Moscoso C (2005) The the Earth’s evolving climate system, as well as lithium industry: its recent evolution and future – the potential for use as a tracer of environmental prospects. Resource Policy 30:218 231 Elliott T, Jeffcoate A, Bouman C (2004) The terrestrial Li contaminants. isotope cycle: light-weight constraints on mantle The volume concludes in Chap. 7 with a convection. Earth Planet Sci Lett 220:231–245 distillation of major findings with indications of Jaskula BW (2015) Lithium. USGS Minerals Yearbook, fertile areas for continued discovery. This section 2013, pp 11 Kesler SE, Gruber PW, Medina PA, Keoleian GA, partly reprises an earlier set of recommendations Everson MP, Wallington TJ (2012) Global lithium for future study (Tomascak 2004), in an attempt resources: relative importance of pegmatite, brine and to gauge what the geochemical community has other deposits. Ore Geol Rev 48:55–69 learned from this stable isotopic system. Mohr SH, Mudd GM, Giurco D (2012) Lithium resources and production: critical assessment and global projec- tions. Minerals 2:45–64 Qin S, Zhao C, Li Y, Zhang Y (2015) Review of coal as a References promising source of lithium. Int J Oil Gas Coal Technol 9:215–229 Tang YJ, Zhang HF, Ying JF (2007) Review of the Anonymous (1818) An account of the new alkali lately lithium isotope system as a geochemical tracer. Int discovered in Sweden. Q J Sci Arts 5:337–340 Geol Rev 49:374–388 Berzelius JJ (1817) Ein neues mineralisches Alkali und Tomascak PB (2004) Developments in the understanding ein neues Metall. J Chem Phys 21:44–48 and application of lithium isotopes in the Earth and Bunsen R (1855) Darstellung des Lithiums. Justus Liebig planetary sciences. In: Johnson CM, Beard BL, Annalen der Chemie 94:107–111 Albarède F (eds) Geochemistry of non-traditional Burton KW, Vigier N (2011) Lithium isotopes as tracers stable isotopes. Rev Mineral Geochem 55:153–195 in marine and terrestrial environments. In: Baskaran M Weeks ME (1956) Discovery of the elements. Easton, (ed) Handbook of environmental isotope geochem- Pennsylvania. J Chem Educ 1956, pp 910 istry; Adv Isotope Geochem 4:41–59 Methodology of Lithium Analytical Chemistry and Isotopic 2 Measurements

2.1 Historical Perspective separation and precise isotopic measurements with *2.5 ‰ external uncertainty. Although earlier measurements which used mass spec- Lithium was discovered as an elemental species in trometry were performed on a range of terrestrial 1817 by J.A. Arfvedson and isolated as a metal one and extraterrestrial samples (e.g., Balsiger et al. year later by Sir H. Davy. The existence of two 1968; Brown et al. 1977; Eugster and Bernas naturally occurring isotopes of lithium, 6Li and 1971; Krankowsky and Müller 1967; Michiels 7Li, was unambiguously proven by A.J. Dempster and De Bièvre 1983), data reported in these (1921) and the atomic weight of Li and the indi- studies were usually expressed in absolute vidual isotope masses were first determined with 7Li/6Li ratios with errors exceeding several per reasonable accuracy and precision by F.W. Aston mil. This severely hampered data comparison (1932) (few available data are referenced therein) and the overall utility of Li isotopes in at 6.928 ± 0.008, 6.012, and 7.012 atomic mass geochemistry. units, respectively (atomic mass unit, amu,is − A major problem of early reports on Li iso- defined as 1.660538921 ± 73 × 10 27 kg and cor- topic compositions was the lack of a widely responds to 1/12 of mass of unbound neutral 12Cin distributed reference material that would allow nuclear and electronic ground state; modern phy- for reliable inter-laboratory comparison of mea- sics uses the term “unified atomic mass” abbrevi- sured isotopic ratios. In 1973, Flesch and ated as ‘u’ or dalton). In the decades that followed co-workers isolated Li from “virgin” this, however, only a few attempts to refine the spodumene-rich ores from the granitic pegmatite absolute atomic weight as well as its uncertainty at Kings Mountain, North Carolina, USA, and (see summary in Svec and Anderson 1965) were determined its isotopic composition at implemented. The actual atomic weights of 6Li 7Li/6Li = 12.0192 ± 0.0002 (Flesch et al. 1973). and 7Li are 6.0151223 ± 5 and 7.0160040 ± 5 u, The isolated Li was made available by the respectively. The actual atomic weight of Li, cur- National Bureau of Standards (now NIST) as rently accepted by IUPAC, is 6.941 ± 0.002 L-SVEC (since reassigned as SRM 8545). The (Wieser 2006) and it may well be that this level of availability of an external standard permitted precision will not improve significantly in the near calibration of measured 7Li/6Li and presentation future due to natural variations in 7Li/6Li exceed- of data relative to L-SVEC in permil deviations. ing *80 ‰ (Tomascak 2004, this volume). Chan (1987) and several papers that followed The natural variations of 7Li/6Li were not employed the normalized 6Li/7Li (δ6Li). Since measured extensively or with clearly established that time, though, the use of the δ7Li value has reproducibility until late 1980s (Chan 1987). prevailed, as it is consistent with the other major This article reported a method for chemical

P.B. Tomascak et al., Advances in Lithium Isotope Geochemistry, 5 Advances in Isotope Geochemistry, DOI: 10.1007/978-3-319-01430-2_2 © Springer International Publishing Switzerland 2016 6 2 Methodology of Lithium Analytical Chemistry and Isotopic Measurements stable isotopic systems (i.e., positive values exposure to such reagents may result in erratic reflect enrichment in the heavy isotope). This 7Li/6Li (Košler and Magna 2014). Vigilance in nominal δ7Li = 0 ‰ mimics the Li isotopic areas of sample preparation is just as critical as in composition of global magmatic continental crust the chemistry laboratory (as discussed below) for (Bryant et al. 2004; Magna et al. 2010; Teng the production of verifiable Li isotopic data. For et al. 2004, 2008, 2009). Considering that the example, in many rock processing facilities the δ7Li values in mineral phases from worldwide preparation of samples for XRF or other granitic pegmatites vary greatly (e.g., Magna flux-fusion methods has taken place. Traces of et al. 2013), the choice of this particular spo- residual dust from Li-borate flux, even years dumene by Flesch et al. was quite serendipitous. later, threaten samples being processed for min- The L-SVEC reference standard has proven, eral separates or rock powders in such spaces. perhaps fortuitously, robust, with proven Li iso- topic homogeneity within better than ±0.1 ‰ (2σ) for aliquots allocated to different laboratories 2.2 Lithium Isolation by Cation (Magna et al. 2004). This is noteworthy in the Exchange Chromatography view of homogeneity problems encountered for reference materials used in several other stable Separation of Li from other elements in natural isotopic systems (e.g., NIST SRM 980 for Mg; materials is challenging owing to M/Li ratios Galy et al. 2003), Si isotopic discrepancies for commonly in excess of 104 (e.g., Na/Li in sea- IRMM-018 versus NBS-28 (see Reynolds et al. water), as well as the similarity of the ion 2006 and discussion therein), or different values exchange partitioning of Li to some major in newly marketed reference materials compared cations. Pioneering work on partitioning of Li and to older, exhausted resources, such as for Ca (note other elements between ion exchangers and var- differences in δ44/40Ca between SRM 915a and ious types of elution media stressed the efficacy of 915b; Heuser and Eisenhauer 2008). Recently, mineral acids such as HCl and HNO3 in separat- however, the L-SVEC reference material has run ing Li (Strelow et al. 1974; Šulcek et al. 1965; out and was replaced by the new IRMM-016 Šulcek and Rubeška 1969), in particular when 7 artificial material with δ LiL-SVEC = –0.2 to mixed with organic solvents (e.g., methanol and +0.3 ‰ (Aulbach and Rudnick 2009; Aulbach ethanol, although other organic elution media et al. 2008; Caciagli et al. 2011; Huang et al. may be used for specific purposes; Kim 2001). 2010; Jeffcoate et al. 2004; Kasemann et al. 2005; All major subsequent analytical developments Millot et al. 2004; Penniston-Dorland et al. 2010; adopted this approach with different combina- Qi et al. 1997b; Simons et al. 2010; Teng et al. tions of these mineral acids and simple alcohols. 2006; Zack et al. 2003). This can thus be con- A major consideration when ion exchange is sidered identical within external analytical applied to Li isolation is the large isotopic frac- uncertainty of Li isotopic measurements and no tionation introduced during chromatographic further recalculation is required. Care should be separation as a consequence of greater affinity of exercised, however, to juxtapose L-SVEC and 6Li to stationary phase (Taylor and Urey 1938). IRMM-016 when significantly different δ7Li As a result of incomplete Li recovery, unintended values emerge (e.g.,<–0.5 ‰; Liu et al. 2010; isotopic fractionation can be caused. For example, Marks et al. 2007) that are slightly beyond the Moriguti and Nakamura (1998) and Košler et al. level of current analytical uncertainties. (2001) have shown that Li isotopic fractionation Yet another problem for Li isotopic analyses may reach several tens of percent with 7Li eluting is represented by laboratory-induced contamina- first and 6Li tailing (Fig. 2.1) as a result of equi- tion that may seriously compromise the intrinsic librium fractionation of Li between the solid 7Li/6Li in natural samples as these ratios often phase and solvent (Schauble 2004); thus, show large variations in commercially available *100 % yields are indeed essential. This differs Li materials (Qi et al. 1997a) and even minimal from elements with more than two isotopes where 2.2 Lithium Isolation by Cation Exchange Chromatography 7

14000 1998; Nishio and Nakai 2002; Rosner et al. 2007; Tomascak et al. 1999). In particular, the removal 10500 50 of Na, as the nearest major element to be eluted δ

7 after Li, has received attention; avoiding or Li (‰) 7000 0 eliminating Na tailing in the Li fraction during chromatographic separation is the goal here. The Counts/s extent to which contaminant elements may com- 3500 -50 promise successful isotopic analysis depends largely on the measurement technique. Whereas 0 02 46 8 10 12 there is some degree of freedom in matrix con- eluent volume (ml) centrations for plasma-based techniques (Huang et al. 2010; Jeffcoate et al. 2004;Košler et al. Fig. 2.1 Example of Li isotopic fractionation during ion 2001; Magna et al. 2004; Rosner et al. 2007; exchange chromatography (after Košler et al. 2001). A progressive shift from isotopically heavy to light Tomascak et al. 1999), little tolerance is apparent fractions during ongoing chromatographic isolation of Li, with thermal ionization methods (Hoefs and 7 with a δ Li range of *100 ‰ during elution. A failure to Sywall 1997; James and Palmer 2000; Moriguti complete recovery of Li would result in the collection of a and Nakamura 1998; Xiao and Beary 1989). sample with an erroneous isotopic composition Early attempts to measure Li isotopic com- positions in natural samples required several internal corrections for incomplete recovery may hundreds of nanograms Li (e.g., Chan 1987; be applied. Unfortunately for the laboratory geo- Chan and Edmond 1988; You and Chan 1996) chemist, achieving quantitative recovery is not but these requirements were alleviated with simply a matter of calibrating ion exchange col- technical improvements; high-sensitivity umns. Position of the elution peak for Li from ICP-based mass spectrometers now allow accu- exchange columns is affected by the bulk chemi- rate analyses with less than 5 ng Li. The need to cal composition of the sample and the ion load to prepare such large samples, especially for rela- the column. Thus, separate elution recipes may be tively low-Li materials like mantle rocks, necessary for different sample types (Fig. 2.2). required large-volume columns packed with The reality of this issue was driven home by Chan ion-exchange resins and large amounts of col- 7 et al. (1999), where biased δ Li results resulted in lection media (sometimes several hundred milli- the publication of a correction (Chan et al. 2002). liters). This made early techniques rather time Analyses of rocks yielding distinct (and geologi- consuming and prone to elevated procedural cally reasonable) isotopic compositions were blanks. With diminishing sample size came the shown, after chemical separation issues were capacity to analyze materials with very low Li solved, to be isotopically homogeneous and with abundance and/or samples of very small size, no anomalous samples. critical for high spatial resolution as well as when The sensitivity of different instrumental limited amounts of sample were available. methods to analyte solutions that lack complete Exchange media with Li-specific characteris- purification is variable, but, for plasma-based tics are not available, unlike for certain elements methods, it is clear that significant amounts of or elemental groups (e.g., Dow AMBERLITETM contaminant elements may compromise the for B, Sr. Spec for Sr and Pb, and Eichrom TRU accuracy and precision of Li isotopic determina- Resin for actinides). The use of cation exchange tions. Hence, the quality of ion exchange sepa- resins such as AG50W-X8 appears to provide ration is an important issue in sample preparation. high efficiency in separating Li from other ele- The effects of imperfect separation of Li from ments, but optimization of the geometry of other elements have been demonstrated by many chromatographic columns also plays a role in studies (e.g., Huang et al. 2010; Jeffcoate et al. quality of element separation, blank suppression, 2004; Magna et al. 2004; Moriguti and Nakamura and time cost. 8 2 Methodology of Lithium Analytical Chemistry and Isotopic Measurements

(a)2.0 (b) 300 Li Na

1.5 200

1.0

100 0.5 Relative abundance Relative abundance

0.0 0 100 200 300 400 500 20 30 40 50 60 70 80 eluent volume (ml) eluent volume (ml)

Fig. 2.2 Examples of the difference in cation exchange column calibration may lead to incomplete collection of elution peaks for Li resulting from various matrices (e.g., the Li fraction (Chan et al. 1999, 2002) peridotites and other silicate rocks, seawater). Improper

Mixtures of mineral acids with methanol or 1995a, 1998), HCl–ethanol (Jeffcoate et al. 2004 ethanol provide better separation of Li from other in second step), pure HNO3 (Hoefs and Sywall elements than strictly inorganic eluants. The 1997 and Magna et al. 2006 in second step) and increasing concentration of alcohols at a given HNO3–methanol (Huang et al. 2010; Choi et al. molarity of the respective mineral acid increases 2010;Košler et al. 2001; Magna et al. 2004; the separation factor between Li and Na (as the Nishio and Nakai 2002; Seitz et al. 2004; next element to elute), such that 80 % by volume Tomascak et al. 1999). These studies encompass a methanol solution roughly triples the separation wide variety in resin mesh size (100–200 vs. 200– over 60 % methanol solution. This is 400), cross-linkage (X8 vs. X12; see Schönbäch- counter-balanced by decreasing separation factor ler and Fehr 2014 for a more detailed review of when the molarity of the respective acid increa- issues related to ion exchange chromatography), ses. Therefore, a near-ideal formula entails and column size/volume, but they all seem to low-molarity acid mixed in high percentage provide a reasonable means of Li separation. organic solvent. There is apparently little differ- Inasmuch as no single method presents a clear ence between HCl and HNO3, the former pro- best-case separation (i.e., one that minimizes viding slightly larger separation factor between reagents and time whilst maximizing ease of use Li and Na and the latter quantitatively eliminat- and flexibility of sample type), there appears to be ing some elements such as Fe, Zn and Cd, which room for further analytical improvements. Per- may appear in HCl-based elution schemes haps in the near future a selective complexing (Strelow et al. 1974). One practical downside of agent or the use of other solvents (e.g., acetone) the acid–organic mixture is the formation of will further improve the process of Li separation small bubbles in the resin. This phenomenon for isotopic measurement. restricts both choice of the internal diameter of a column (which cannot be too narrow) and the column material (e.g., non-wetting material like 2.3 Methodology of Li Isotopic PFA Teflon versus quartz glass). Measurements A glut of methods have been successfully employed in geochemistry and cosmochemistry: At present, a plethora of technical possibilities pure HCl (Chan 1987; James and Palmer 2000; exists for measurements of Li isotopic composi- Misra and Froehlich 2009; Moriguti and Naka- tions, but superior performance with respect to mura 1998; Oi et al. 1997; Sahoo and Masuda the accuracy and precision of isotopic data 2.3 Methodology of Li Isotopic Measurements 9 collection is represented by plasma-based tech- remains the only reliable method for determina- niques (Q-ICPMS, SF-ICPMS, MC-ICPMS), tion of absolute Li isotopic abundances, as the thermal ionization (TIMS), and secondary ion- plasma-based instruments always return ization (SIMS). Although this instrumentation mass-biased results and the instrumental frac- and its technical aspects are rather complex and tionation is too great and variable to effectively expensive, these methods provide the data output correct back to absolute ratios. As such TIMS has required for geochemistry and cosmochemistry. represented a cornerstone for resolving even Nevertheless, other methods have also been small differences in δ7Li considering ±1 ‰ errors developed for specific purposes and we discuss that were routinely obtained (Chan 1987). It is them here briefly. Also, it is important to note the unfortunate that, despite the generation of the prevalence of the δ6Li notation, where negative L-SVEC standard in the early 1970s, TIMS values represent isotopically heavier values, in studies throughout the 1970s and 1980s failed to most of the earlier publications. Following the measure this material, making it impossible to Goldschmidt 2002 geochemical conference a recalculate data to modern systematics. more logical δ7Li notation has uniformly been Several strategies have been developed that + accepted (Coplen et al. 2002). The conversion employed polyatomic species, such as Li2BO2 can simply be performed with the following (Bickle et al. 2000; Datta et al. 1992; Chan 1987; + equation: Sahoo and Masuda 1995b), LiNaBO2 (Chan + et al. 1992), Li2F (Green et al. 1988), or metal d6Li + d7Li ¼À Li ion (Ahmed et al. 2002; Jabeen et al. 2003; d6Li James and Palmer 2000; Michiels and De Bièvre 1 þ 1000 1983; Moriguti and Nakamura 1998; Sahoo and Masuda 1995a, 1998; Xiao and Beary 1989; You This conversion must be performed for δ6Li and Chan 1996). These approaches combined values outside the range *–10 to *+10 ‰. different loading procedures, utilizing Li B O , Within this range the sign-changed difference 2 4 7 Li PO , LiCl + H BO , LiOH, LiI, LiF, LiCl (but between δ6Li and δ7Li values results in shifts 3 4 3 3 other possibilities were also explored, see Xiao <0.1 ‰ whereas for δ6Li values in the range of and Beary 1989), with various filament config- seawater (–31.0 ‰; Millot et al. 2004) the dif- urations: single Re, Ta–Re or double Re, or triple ference is *1 ‰. It increases to *+10 ‰ at Re filament assembly. δ7Li > 100 ‰ and to *40 ‰ at δ7Li > 200 ‰, The developments by Chan et al. in mea- found for anthropogenically fractionated samples surement by TIMS, making use of the L-SVEC (Millot et al. 2010;Négrel et al. 2010). standard, broke the logjam and ushered in a new era of Li isotopic geochemistry. Studies using this technique introduced several first-order 2.3.1 Mass Spectrometry-Based observations on the geochemical cycle of Li, Methods of Widest Use making the first steps into this terra incognita (e.g., Chan et al. 1992, 1993, 1994; You et al. 2.3.1.1 TIMS 1995). However, with the advent and rapid Prior to the diversification of ICP and SIMS development of MC-ICPMS instrumentation instruments, thermal ionization was the primary (e.g., Halliday et al. 1995, 1998), TIMS tech- method of choice in yielding high-precision Li niques quickly became a somewhat obsolete isotopic compositions for several decades. approach for determination of Li isotopic com- Indeed, the first absolute Li isotopic composi- positions in natural samples considering the more tions were obtained by TIMS (Balsiger et al. time consuming mass spectrometry and lesser 1968; Svec and Anderson 1965, 1966) or its opportunity to monitor shifts in instrumental modifications (Brown et al. 1977) and this mass bias. Despite these drawbacks, TIMS 10 2 Methodology of Lithium Analytical Chemistry and Isotopic Measurements remains the only method capable of directly isotopes (Tomascak et al. 1999). The possibility measuring the absolute Li isotopic abundances of monitoring the varying instrumental bias by and as such, it will continue to have importance alternating samples and reference materials for precise validation of Li isotopic abundances (L-SVEC; Fig. 2.3) on short time scales (plus the in reference materials. capability of monitoring abrupt changes in instrumental bias) promoted a massive increase 2.3.1.2 ICPMS in number of available data for Li isotopic Despite the explosion in the use of plasma-based compositions in all manner of geological set- technologies in early 1980s, these methods did tings, as well as some of the first credible data on not find broad application to Li isotopic mea- non-terrestrial materials (cf. Tomascak 2004 and surements until much later. The early generation later development discussed in several sections of single-collector quadrupole ICP instruments of this volume). The major advantages of this was not well suited for isotopic determinations in technique lie in rapid throughput of solute sam- geological samples but the few attempts returned ples, possibility to monitor instrumental bias data that were fully in the context of contempo- during the analytical session, and low consump- raneous investigations (Grégoire et al. 1996; tion of material that may, under suitable cir- Vanhoe et al. 1991). Despite larger inaccuracy of cumstances (sample/noise ratio), utilize less than Li isotopic measurements by this technique, it one nanogram of Li with sub-permil external was applied to, for example, serious social– reproducibility. Furthermore, MC-ICPMS medical questions (Sun et al. 1987). appears less susceptible to undesired shifts in Košler et al. (2001) presented an improved measured 7Li/6Li ratios resulting from the pres- quadrupole ICPMS protocol for determining Li ence of matrix elements due to imperfect chem- isotopic compositions in natural samples (for- ical separation of Li (e.g., Bryant et al. 2003; aminiferal species) with external reproducibility Jeffcoate et al. 2004; Magna et al. 2004; Nishio approaching the ±2 ‰ (2σ) barrier which would and Nakai 2002; Rosner et al. 2007; Tomascak be adequate for many geological applications. et al. 1999). Despite the strong fractionation of Li Further optimization has improved these statis- isotopic ratios with MC-ICPMS compared to tical parameters (Carignan et al. 2007; Misra and TIMS, the method of measuring samples brack- Froehlich 2009). The higher sensitivity eted by standards has effectively circumnavi- sector-field (single-collecting) ICPMS does not gated this problem. appear to represent significant improvement in The issue of accuracy and inter-laboratory accuracy and long-term precision of Li isotopic comparison is only secured by the availability of measurements compared to quadrupole instru- reference rock materials. At present, several of ments (Janoušek et al. 2010; Magna et al. 2010). these are utilized worldwide (e.g., the basalts Nonetheless, Li isotopic measurement by BHVO-2, JB-2, BCR-2, and IRMM BCR-403 or Q-ICPMS is not a “routine” practice, and as such IAPSO seawater) and many others have been the diversification of MC-ICPMS instruments characterized for their δ7Li through multiple leaves ICPMS a less used alternative today. independent measurements. Considering this However, for studies for which precision of wide range of materials of distinctive chemical ±1.0 ‰ is sufficient, it is a viable method that is compositions, it may be a prerequisite to certain to continue to be applied. “matrix-match” complete analytical procedures with corresponding reference materials instead of 2.3.1.3 MC-ICPMS relying solely on one reference material. With the advent of MC-ICPMS instrumentation Technical approaches to the measurement by and subsequent rapid developments in accurate, MC-ICPMS are diverse. Overall, solutes are high-precision stable isotopic measurements aspired into the plasma via low-flow nebulizers, (Halliday et al. 1995, 1998), it was only a matter nominally 30–100 μL min−1, where the plasma of time before the tool was brought to bear on Li may be “hot” (*1200–1400 W) or “cool” 2.3 Methodology of Li Isotopic Measurements 11

(a) 14.86 (b) 13.52

14.82 13.50 Li Li 14.78 6 6 13.48 Li/ Li/ 7 7 14.74 13.46

14.70 13.44 (c) 14.84

14.82 Li 6 14.80 Li/ 7

14.78

14.76

Fig. 2.3 Example of analytical sequences with standard– alternatives should yield identical results. Stability of the sample bracketing, as commonly used with MC-ICP-MS instrumental bias is among the key factors for accurate measurement (a). Calculation of the final δ7Li may be determination of Li isotopic composition, here exempli- performed in two ways: (i) relative to the average 7Li/6Li fied by a maximum 7Li/6Li variability of *0.4 ‰ over of two bracketing reference solution measurements irre- >5 h. In b and c, example of noisy and stable measure- spective of the actual measured 7Li/6Li (*14.80 for the ment is apparent by a *10-fold difference in internal depicted analytical sequence), or (ii) recalculating the errors of the measurement of the same solution over two measured values to the natural 7Li/6Li of *12.1. Both different analytical sessions, performed within 1 week

(*700–800 W; Bryant et al. 2003; Choi et al. (Magna et al. 2004) despite lower mass resolving 2013); aspiration is often aided by desolvation, power required to delimit these individual peaks. resulting in more homogeneous droplet size just When Tomascak et al. (1999) introduced prior to dispersion in plasma. During an analyt- MC-ICPMS measurement protocols for Li iso- ical session, a solution is introduced after the topic measurements with the first generation background signal level is reached that may instruments, a reproducibility of ±1.1 ‰ (2σ), require significant washout time, but alternative basically equivalent to TIMS analysis, was switch between sample and standard solution achieved. Currently, the long-term reproducibil- without rinsing was also explored (Rosner et al. ity of MC-ICPMS measurements using more 2007). This latter approach requires a complete modern instruments appears to be at the ±*0.3– signal intensity match to better than ±5%in 0.4 ‰ (2σ) level (Jeffcoate et al. 2004; Magna order to avoid instrumental bias introduced by et al. 2004; Millot et al. 2004). Although improperly balanced signals of bracketing stan- short-term reproducibility may surpass this level, dard and unknown sample (Huang et al. 2010; it is not apparent in the current literature that any Magna et al. 2004; Rosner et al. 2007). The group is able to do demonstrably better for true, isobaric interferences from 12C++, 14N++ and long-term conditions. 6LiH+ appear to pose a minor problem and were At present the literature on Li isotopes is indeed not detected at mass resolution of *1400 dominated by data generated with solution 12 2 Methodology of Lithium Analytical Chemistry and Isotopic Measurements

MC-ICPMS. Recently, le Roux (2010) presented measurement of Li isotopes emerged (e.g., Barrat a laser ablation (LA) MC-ICPMS method. In this et al. 2005; Beck et al. 2004; Decitre et al. 2002; study, they performed in situ measurements of Li Chaussidon and Robert 1999; Kobayashi et al. isotopic composition in reference glasses whose 2004; Richter et al. 2003). Li isotopes had been characterized by means of Whereas SIMS requires precise matrix solution MC-ICPMS. It appears that combination matching for any elemental and/or isotopic of LA with MC-ICPMS may result in rapid determinations, little effort has been exerted for determination of Li isotopic compositions on the characterization of suitable geological reference sub-0.1 mm scale if followed by technical materials in terms of Li isotopic compositions improvements in detection limits and sensitivity. (Jochum et al. 2006; Kasemann et al. 2005). This For example, minerals with 10–20 ppm Li gen- appears a persistent problem and may require erated *0.1 V signals on mass 7Li against 6–8V further experimental work. Such inferences are for 10 ppb Li solutions routinely obtained with stressed by apparent matrix-induced fractionation solution MC-ICPMS. Also, diligent matrix of Li isotopes (Bell et al. 2009) which may matching appears to be essential for laser sam- indeed impart an additional uncertainty to in situ pling, which demands a suite of homogeneous measurements unless these matrix effects are not and well characterized glass reference standards. disentangled correctly. It would follow that an The recent report of Xu et al. (2013) showed the instrumental mass fractionation factor may be spatial capabilities coupled with high sample estimated from false data and that not only must throughput weighed against still rather high a mineral phase be matched, but also its major analytical errors (mostly >1.8 ‰,2σ) and poor element composition should be more or less reproducibility for a large selection of reference identical with unknown samples. These matrix glass standards. It shows that this methodology effects may partly be responsible for a larger yet awaits further detailed analytical efforts uncertainty linked to SIMS, that in most appli- before it may become a fast and lower-cost cations is limited to *±2 ‰ (2σ). It is expected analytical alternative. that this can be significantly reduced in carefully controlled experiments (Marks et al. 2008). 2.3.1.4 SIMS Marks et al. (2008) provided a detailed Although it may seem that implementation of investigation and SIMS–MC-ICPMS laser ablation into MC-ICPMS could represent a cross-calibration of pegmatitic Na-rich arfved- new and straightforward in situ means of Li sonite and aegirine in order to determine the isotopic determination, SIMS technology is far matrix effects caused by SIMS. These authors ahead in terms of spatial resolution, sensitivity, found a significant difference in resulting δ7Li of and overall technical abilities in investigations of the two methods at the level of 4–5 ‰ (see also Li isotopic compositions. The first real attempts Kasemann et al. 2005), and differences on the to measure Li isotopic compositions with a sec- order of nearly 10 ‰ for NIST 610, 612 and 614 ondary or sputtering ion source were performed glasses. Similarly, Bell et al. (2009) applied a on meteorites (Gradsztajn et al. 1967; Poschen- differential matrix correction factor deduced from rieder et al. 1965) and lunar rocks (Eugster and experimentally verified fractionation of Li iso- Bernas 1971), although the precision and accu- topes in response to major chemical composition racy were apparently too low for any valuable of olivines (Mg#). Recently, Su et al. (2015) scientific statements apart from the 7Li/6Li ratios have provided a detailed SIMS, ICPMS and identical to those found for terrestrial rocks MC-ICPMS comparison of multiple grains of within large uncertainties. Nevertheless, further olivine, clinopyroxene and orthopyroxene from measurements of Li isotopic compositions in ultramafic rocks collected in China, with the aim geological samples were mostly obtained by of establishing well-characterized mineral refer- TIMS and later MC-ICPMS; only relatively ence materials for in situ techniques. While recently have studies utilizing SIMS laborious by nature, these analyses showed that 2.3 Methodology of Li Isotopic Measurements 13

(i) SIMS results for olivine may strongly depend be disputed (Kushita 1986). Chapman and Dale on Mg# [molar MgO/(MgO + FeOt) × 100] (1976) reported on an AAS methodology for fast which would demand careful daily calibration determination of 7Li/6Li in a range of samples over a range of Mg# values, and (ii) significant which Chapman et al. (1980) suggested to δ7Li variations of several ‰ persisted between approach analytical precision obtained by clas- different laboratories and even between different sical TIMS techniques. However, the data of sessions conducted in a single laboratory. This Kushita (1986) show that this expectation was may, in part, be linked with different instrumental largely over-estimated. optimization parameters because the reported Li In later progress, many problems were over- contents in mineral phases were shown to be come (Wizemann and Niemax 2000), but the reproducible. precision of Li isotopic determinations reached The homogeneity in Li contents is also an with AAS remains an issue. Moreover, results important factor to consider, as has been under- obtained for many natural samples (Meier 1982) scored by inter-laboratory consistency of Li that are likely to have non-extreme isotopic abundances in a natural quartz specimen (Audé- compositions (e.g., natural waters) show extre- tat et al. 2015). Therefore, combined analytical mely light compositions. Therefore, the reliabil- tools may become a truly essential strategy for ity of AAS (and also atomic emission further SIMS studies of, for example, meteoritic spectrometry; ur Rehman et al. 2009)or constituents such as chondrules and calcium– laser-induced breakdown/atomic fluorescence aluminum-rich inclusions, considering both their (Smith et al. 1998) may be adequate for nuclear small sizes and the important message they may research, but geochemical applications require carry since their formation in the earliest era of much more precise data, largely beyond capa- the Solar system. In fact, Li partitioning within bilities of these techniques. single crystals of a mineral species may depend on major element composition, such as shown 2.3.2.2 Nuclear- and Charged- for plagioclase (Coogan 2011). Particle-Based Techniques The general use of neutron activation analysis (NAA) in 7Li/6Li determination has been 2.3.2 “Other” Analytical Methods explored, with primary application to the nuclear industry. In principle, the abundance of one of 2.3.2.1 Atomic Absorption Spectrometry the Li isotopes is determined precisely and total (AAS) Li abundance is then measured by an indepen- This technique has been successfully applied to dent method; the final 7Li/6Li can then be cal- rapid determination of Li isotopic compositions culated (Wiernik and Amiel 1970). Wölfle and in the nuclear industry. Lithium-6 has large cross Neubert (1977) introduced an activation analysis section for thermal neutrons (*942 barns) and method for simultaneous determination of 6Li the knowledge of 6Li level is important for its and 7Li, and applied it to aqueous solutions, but burnup rate and for estimation of the production the approach is less effective for natural samples. rate of tritium through the 6Li(n,α)3H reaction. On the other hand, Rajan et al. (1980) applied The possible usage of AAS for determining Li nuclear techniques to 7Li/6Li determination of isotopic ratios was first explored by Wheat several stony meteorites, showing largely (1971) who employed a mono-isotope hollow homogeneous Li isotopic compositions within cathode lamp. For AAS, it is essential to pre- ±10 ‰ of terrestrial value (given by analysis of cisely know the Li content in unknown samples; spodumene) which is broadly compatible with Li isotopic ratios are calculated from measured recent high-precision MC-ICPMS investigations absorbance of 6Li (Meier 1982) but the linearity (e.g., Seitz et al. 2007; see Chap. 3) but any of these absorbance determinations, which is inter-sample relations cannot be accounted for by prerequisite for exact data acquisition, appears to this method. It may be that NAA techniques 14 2 Methodology of Lithium Analytical Chemistry and Isotopic Measurements

(Chao and Tseng 1995; Itoh et al. 1993) and Aulbach S, Rudnick RL, McDonough WF (2008) other more or less related methods, such as Li-Sr-Nd isotope signature of the plume and cratonic lithospheric mantle beneath the margin of the rifted proton induced gamma-ray emission (Trompetter Tanzanian craton (Labait). Contrib Mineral Petrol et al. 1999) or nuclear track detection (Kjellman 155:79–92 et al. 1985), remain permissible methods for fast Balsiger H, Geiss J, Groegler N, Wyttenbach A (1968) Distribution and isotopic abundance of lithium in determination of Li contents in geological and – ä ä stone meteorites. Earth Planet Sci Lett 5:17 22 other samples (see R is nen 1992 for review), Barrat JA, Chaussidon M, Bohn M, Gillet P, Göpel C, but they are less useful in determining Li isotopic Lesourd M (2005) Lithium behavior during cooling of compositions and their subtle variations at pre- a dry basalt: an ion-microprobe study of the lunar cisions of less than several permil, seriously limit meteorite Northwest Africa 479 (NWA 479). Geochim Cosmochim Acta 69:5597–5609 their utility in the examination of a large range of Beck P, Barrat JA, Chaussidon M, Gillet P, Bohn M geochemical processes. (2004) Li isotopic variations in single pyroxenes from the Northwest Africa 480 shergottite (NWA 480): a 2.3.2.3 Resonance Ionization and Other record of degassing of Martian magmas? Geochim Cosmochim Acta 68:2925–2933 Techniques Bell DR, Hervig RL, Buseck PR, Aulbach S (2009) Of the less-commonly applied methods to mea- Lithium isotope analysis of olivine by SIMS: calibra- sure Li isotopes, resonance ionization mass tion of a matrix and application to magmatic phe- – spectrometry (RIMS), could offer an alternative to nocrysts. Chem Geol 258:5 16 Bickle MJ, Chapman HJ, You C-F (2000) Measurement destructive methods (Suryanarayana et al. 1998) of lithium isotopic ratios as lithium tetraborate ions. due to its extremely low detection limits (on the Int J Mass Spectrom 202:273–282 order of several tens femtograms of Li). The Brown HL, Biltz C, Anbar M (1977) A precision isotope 6 7 possible application of this technique to cosmo- ratio mass spectrometer for the analysis of Li/ Li. Int J Mass Spectrom Ion Phys 25:167–181 chemical studies (Knight et al. 2007; Levine et al. Bryant CJ, McCulloch MT, Bennett VC (2003) Impact of 2009) could provide answers to some funda- matrix effects on the accurate measurement of Li mental questions of the origin of the Solar system, isotope ratios by inductively coupled plasma mass “ ” presuming continued technical advancement. spectrometry (MC-ICP-MS) under cold plasma conditions. J Anal Atom Spectrom 18:734–737 Laser spectroscopy was developed for studies of Bryant CJ, Chappell BW, Bennett VC, McCulloch MT nuclear properties of different Li isotopes (Her- (2004) Lithium isotopic compositions of the New genröder et al. 1993;Nörtershäuser et al. 2011) England Batholith: correlations with inferred source — but seems to be less applicable to geochemistry. rock compositions. Trans R Soc Edinburgh Earth Sci 95:199–214 Caciagli N, Brenan JM, McDonough WF, Phinney D (2011) Mineral-fluid partitioning of lithium and References implications for slab-mantle interaction. Chem Geol 280:384–398 Carignan J, Vigier N, Millot R (2007) Three secondary Ahmed S, Jabeen N, ur Rehman E (2002) Determination reference materials for lithium isotope measurements: of lithium isotopic composition by thermal ionization Li7-N, Li6-N and LiCl-N solutions. Geostandards mass spectrometry. Anal Chem 74:4133–4135 Geoanal Res 31:7–12 Aston FW (1932) The isotopic constitution and atomic Chan LH (1987) Lithium isotope analysis by thermal weights of caesium, strontium, lithium, rubidium, ionization mass spectrometry of lithium tetraborate. barium, scandium and thallium. Proc R Soc Lond A Anal Chem 59:2662–2665 134:571–578 Chan LH, Edmond JM (1988) Variation of lithium Audétat A, Garbe-Schönberg D, Kronz A, Pettke T, isotope composition in the marine environment: a Rusk B, Donovan JJ, Lowers HA (2015) Characteri- preliminary report. Geochim Cosmochim Acta sation of a natural quartz crystal as a reference material 52:1711–1717 for microanalytical determination of Ti, Al, Li, Fe, Mn, Chan LH, Edmond JM, Thompson G, Gillis K (1992) Ga and Ge. Geostand Geoanal Res 39:171–184 Lithium isotopic composition of submarine basalts: Aulbach S, Rudnick RL (2009) Origins of implication for the lithium cycle in the oceans. Earth non-equilibrium lithium isotopic fractionation in Planet Sci Lett 108:151–160 xenolithic peridotite minerals: examples from Tanza- Chan LH, Edmond JM, Thompson G (1993) A lithium nia. Chem Geol 258:17–27 isotope study of hot springs and metabasalts from References 15

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3.1 Cosmology of Lithium: Big synthesized during the Big Bang Nucleosynthe- sis (BBN), the others being 2H(≡D; deuterium), Bang Nucleosynthesis, 3 4 3 Evolution of the Universe, He, He, and a few radioactive nuclei like H ≡ 3 and Stellar Processes ( T; tritium, decaying to He). A number of other very short-lived nuclei (e.g., 11C) that either decayed or participated in subsequent fusion The purpose of this chapter is to inform the reactions (Fig. 3.1) and perhaps paltry amounts reader in a simplified way about the importance of stable 9Be (9Be/H * 10−18) were also pro- of Li in astrophysics and stellar physics and its duced by the end of the BBN process. impact (i) on extreme processes and events tak- Lithium-7 is an important child of BBN and ing place during Big Bang Nucleosynthesis can be produced by two chains of reactions: (i) T (hereafter referred to as BBN) and shortly after (α,γ)7Li, and (ii) 3He(α,γ)7Be(n,p)7Li. The the BBN ceased, and (ii) after more regular 7Be → 7Li decay via electron capture appears to processes of forming new matter of the Universe be responsible for c. 90 % of the 7Li formed were launched. For a more interested reader the during BBN; the remaining *10 % was synthe- literature is rich in studies dealing with the his- sized directly as primordial 7Li. The 7Be half-life tory of light nuclei during the evolution of our of c. 53.1 days was measured by Jaeger et al. Galaxy (e.g., Prantzos 2012) as well as the per- (1996) while the half-life of 53.22 days is given by spective on evolution of Li during cosmic history Audi et al. (2003). Despite its consumption much (see reviews by Reeves 1994; Iocco et al. 2009, 7Be remained alive until BBN ceased (Fig. 3.2). for example, or recent proceedings edited by Conversely, 7Li is mostly destroyed by the Charbonnel et al. 2009). reaction 7Li(p,α)4He. Overall, the T → 7Li reaction is the most important, but at high bary- 3.1.1 The Earliest History and Further onic densities (roughly speaking a ratio of pro- Evolution of Lithium tons and neutrons to photons) the 7Li → 4He reaction appears to reach approximately similar With the advent of and further improvements in leverage, balancing the contribution from astronomical and astrophysical observations, it T → 7Li reaction, and under such conditions Li became clear that Li belongs to the few key is mainly produced by the decay of 7Be. elements that may provide critical information on The BBN contribution to cosmic Li abun- the earliest history of cosmic matter as well as dances can be observed from a so called “Spite place stringent constraints on stellar evolution. plateau” (Fig. 3.3) in the Li/H versus [Fe/H] plot This stems from the fact that Li, more precisely of old halo stars. The ratio [Fe/H] is called 7Li, is among the very few stable species metallicity, the abundance of all elements in stars

P.B. Tomascak et al., Advances in Lithium Isotope Geochemistry, 19 Advances in Isotope Geochemistry, DOI: 10.1007/978-3-319-01430-2_3 © Springer International Publishing Switzerland 2016 20 3 Cosmochemistry of Lithium

Fig. 3.1 Big Bang 11 12 Nucleosynthesis C C β + (BBN) nuclear reactions up β − to CNO (modified from γ p, d,n Coc et al. 2012). Deuterium γ α, 10 11 12 formation represents a key B B d,p B event for the subsequent α p, γ chain of nuclear reactions γ d, T,γ p, 7 9 Be Be ,n T,p α n,p ,γ αα p, γ d, γ T,n α, 6 7 8 Li α Li γ Li p, n, α p, γ d, α d,n 3 d,p 4 γ He He α, n,p γ p, d,n d,n n,γ 1H 2H d,p 3H β − β +

n heavier than H and He. It is defined as follows: metal-poor ([Fe/H] ≤ –1.3) Population II stars

[Fe/H] = log10 (NFe/NH)star – log10(NFe/NH)Sun (Spite and Spite 1982). The revised calculations, (N = number of atoms per unit of volume; Fe is underscored by advancements both in the exper- taken as an element that is easy to measure in the imentally determined cross-section of the 3He(α, visible spectrum and, more importantly, an ele- γ)7Be reaction (important for production of 7Li) ment fully synthesized in stars and not during and tighter constraints placed on the cosmic bar- BBN). Ancient metal-poor halo stars have con- yon density using Wilkinson Microwave Aniso- stant 7Li abundances for a large range of metal- tropy Probe (WMAP) measurements, have licities whereas abundances of those nuclei that delimited the primordial 7Li/H at *5.2 × 10−10 are produced during the galactic epoch (such as (Cyburt et al. 2008). Be and B) increase almost constantly (Spite and Yet, a so-called “cosmic lithium problem” still Spite 1982). Earlier calculations implied that for exists: the predicted Li abundance is a factor of 2– the standard BBN theory, primordial 7Li/H 4 higher than the observed (Asplund et al. 2006; abundance is between 5 × 10−10 and 2 × 10−9 Coc 2013; Fields 2011). This appears to impart and that over the history of the Universe 7Li unsolved problems to the existing models of abundance has steadily grown and decreased particle physics. Fields (2011) offers several (Delbourgo-Salvador et al. 1985; Steigman solutions to this “cosmic lithium problem” with 1993). Predictive theoretical models for the pri- respect to astrophysical observations, cosmolog- mordial BBN 7Li abundance suggest a very nar- ical models of BBN predictions and more com- row range between 8 × 10−11 < 7Li/H < 2 × 10−10 plex solutions beyond standard models of particle (Steigman and Walker 1992). This range is con- physics. The recent observations of up to sistent with the estimates of Audouze et al. (1983) several-times larger 7Li depletions (lower than and Reeves (1994)at7Li/H between 1 and would be predicted solely by BBN) in very old 3 × 10−10, with more elaborate calculations of Population II dwarfs with [Fe/H] <–4 (e.g., Caf- Ryan et al. (2000)at*1.2 × 10−10, and with fau et al. 2011; Sbordone et al. 2010) combined empirical observations for warm (T ≥ *5500 K) with a significant scatter of the Spite plateau at 3.1 Cosmology of Lithium: Big Bang Nucleosynthesis … 21

Time (s) 1 2 3 4 10 10 10 10 -9 WMAP + BBN p n 4He 10 -2 Li -11 10 -4 2H (D) B 3He -6 10 3

H (T) log [n(X)/n(H)] -13 Be n 10 -8 7Be

Mass fraction -3 -2 -1 0 10 -10 7Li Metallicity [Fe/H] 10 -12 6Li Fig. 3.3 Schematic plot of abundance of light elements (Li, Be, B) as a function of stellar metallicity. The 3.0 1.0 0.3 0.1 so-called Spite plateau is apparent for Li whereas Temperature (109 K) abundance of Be and B grows constantly over a range of metallicities. The difference between modern predic- Fig. 3.2 Production of light nuclei during BBN (modified tions (dashed line) and long-term observations of Li from Coc et al. 2012). The abundance of other stable nuclei abundance in stars (dark grey solid line) is depicted by a (e.g., 9Be, CNO) was several orders of magnitude lower stippled area such low and even lower metallicities (so-called the destruction of 7Li in stars (in particular “meltdown”) may imply not-yet completely important at high metallicities), but was con- understood astrophysical processes causing the stantly replaced by production from other sour- dispersion of Li abundance in these otherwise ces. For low-metallicity stars, 7Li started to be very similar, metal-deficient old stars rather than continuously produced during stellar nucleosyn- invoking new cosmological theories. These latter thesis and through spallation reactions on C, N, observations may also imply high temperatures and O nuclei as well as α + α fusion reactions of (>2 × 106 K) in the prior history of stellar material Galactic Cosmic Rays (GCR) interacting with the with extreme Li depletions on the order of up to interstellar medium (ISM). In this respect, it is 105 or, alternatively, distinct stellar processes that noteworthy that stellar Li–Be–B depletion is not could once have led to efficient depletion in Li. paralleled by lowered GCR Li–Be–B abundances Whereas recent findings of Howk et al. (2012) (Fig. 3.4). This clearly supports GCR as an are in broad agreement with combined important input of these light nuclei. It is also WMAP + BBN predictions, their observations speculated that a less significant proportion of 7Li also tell us very little about primordial lithium. may perhaps be formed in core-collapse Type II The reason is that Howk et al. (2012) observed supernovae by neutrino interactions with pre- lithium in a gas cloud system whose metallicity is dominantly 4He and 12C (Woosley et al. 1990)in intermediate between primordial (i.e., very low AGB stars (e.g., Romano et al. 2001) and by metallicity of the Spite plateau stars) and Solar. It 7Be → 7Li reaction in novae, the latter type of has been known that in its post-BBN evolution, stars being among significant Li suppliers the lithium abundance increased from a low pri- (Romano et al. 2001; Tajitsu et al. 2015). mordial value to the value inferred from the Solar Contrary to the clear variability in 7Li pro- System or from the interstellar medium of the duction owing to three different mechanisms Galaxy. The results of Howk et al. (2012) are (BBN, GCR, stellar nucleosynthesis), 6Li is expectedly intermediate between the Spite plateau nearly solely produced by spallative interactions and the Solar abundance but they lack a model for of GCR with ISM. A very subordinate proportion evolution of the primordial abundance of Li. of 6Li was probably synthesized during BBN. During the subsequent evolution of the Uni- The BBN production of 6Li is dominated by the verse, primordial 7Li decreased due to astration, D(α,γ)6Li reaction, which is largely ineffective, 22 3 Cosmochemistry of Lithium

H Moreover, it may be plausible that a large frac- 10 6 He tion of Li is bound in molecules like LiH, or is GCR abundances 10 4 depleted from the gas into microscopic dust O C Ne particles known to commonly exist in ISM. The Mg 2 Si 10 S Fe extent of this depletion is largely unknown, Ar Ni Cr however, but may play important role. 0 10 K Ti F Zn Also interesting are the production rates of V -2 Ge individual Li isotopes for GCR processes. 10 Sc Kr Abundance (Si=100) B Sr Y Be Audouze et al. (1983) calculated that the maxi- -4 7 6 10 Li mum production Li/ LiGCR cannot exceed 6–7 7 6 and Li/ LiGCR production ratios as low as *1 Fig. 3.4 Solar abundance of elements (in order of were also reported (e.g., Garcia-Munoz et al. increasing Z, up to 39) normalized to Si, relative to 1975; Meneguzzi et al. 1971; Webber et al. dark grey line Galactic Cosmic Ray abundance ( ). 2002). This would mean that either an additional Notably, light nuclei (Li, Be, B) are depleted in the 7 fi Sun, attesting to significant production of these nuclei in source of Li, or a mechanism to ef ciently GCR reactions and consumption in the Sun destroy 6Li, or a combination of both, are nee- ded, during Galactic evolution. In fact, lifetime in producing only vanishingly small amounts of 6Li the stellar environment is *100 × shorter for 6Li (Fig. 3.1). Thus, the primordial 6Li/H was vari- compared to 7Li given the greater fragility of the − ously estimated at *10 14 levels (i.e., former nuclide at high temperatures, and it 7Li/6Li * 104; Hammache et al. 2010; Nollett appears that in massive stars initial 7Li was not et al. 1997; Vangioni-Flam et al. 2000, 1999). destroyed (Audouze et al. 1983). Alternatively, it Most 6Li is produced by much more efficient was destroyed but was subsequently compen- spallation-induced energetic collisions of GCR sated for by synthesis of 7Li in the neutrino through p, α + He and CNO → 6Li, 7Li, and fast process. Several measurements of Li abundance nuclei (α, C, O) fragmenting on H and He in the in ISM regions yielded Li/H of *0.6 to ISM (Vangioni-Flam et al. 1999). We refer to a 4 × 10−10 (Knauth et al. 2003) although Li detailed review by Prantzos (2012) for the his- abundance may vary between distinct ISM tory and physical processes of Li evolution. regions. The measurements of ISM appear to find It is generally accepted that the combined Li isotopic ratios that are somewhat lower than reactions of energetic cosmic ray C–N–O nuclei, meteoritic 7Li/6Li (e.g., Ferlet and Dennefeld and cosmic ray H–He interacting with ambient 1984). Meyer et al. (1993) reported 7Li/6Li val- interstellar C, N and O species produce ues for two ISM regions between 5.5 and 6.8 low-energy Li, Be and B nuclei. Indeed, it has whereas Lemoine et al. (1993) found ISM ratios been suggested by some that the latter process, which were very close to meteoritic (*12). which is an important part of the galactic These variations were further underscored by chemical evolution, may dominate the overall later data from Knauth et al. (2000), who reported production of modern Li in the Galaxy. This is on observations of 6Li and 7Li abundance in sev- depicted in Fig. 3.4 where the six orders differ- eral interstellar clouds, with 7Li/6Li ratios varying ence between Solar abundances and those between 1.7 and 10.6 (i.e., broadly meteoritic vs. determined for GCR manifests an important highly non-Solar). The decreased 7Li/6Li is sug- contribution by GCR reactions to the total Li gested to derive from recent massive interactions budget (and also budgets of Be and B). We stress with energetic particles that may result in 7Li/6Li here that the ISM abundance of Li is largely ratios as low as *1 through cosmic ray spallation unknown or, at least, less well constrained due to (see above). Knauth et al. (2000) also estimated technical difficulties, such as extreme proximity the interstellar Li/H abundance at *1×10−9 and of spectroscopic resonance lines accessible for Li speculated that low 7Li/6Li *2 may be intrinsic to abundance measurements (Steigman 1993). the interstellar gas and that contamination with 3.1 Cosmology of Lithium: Big Bang Nucleosynthesis … 23 high-7Li/6Li material would be warranted. The cosmology. Ahrens et al. (1992) discussed that Li weighted mean 7Li/6Li of 7.6 ± 2.3 (1σ) in several depletion of the Sun may have been completed ISM regions was more recently given by Kawa- during its pre-main sequence evolution, but this nomoto et al. (2009). Such Li isotopic variability depends on the abundances of He and other of the ISM was corroborated by Knauth et al. metals, on the mixing length, and the thickness of (2003) who investigated 7Li/6Li ratios of ISM in an extra mixing layer. Some *200,000 km for the vicinity of active star-forming regions. These the convective zone in the Sun, responsible for Li authors confirmed low 7Li/6Li ratios in several depletion, was calculated. However, if Li deple- ISM regions but they also provided additional tion by a factor of *150× is considered (roughly support for 7Li/6Li ratios that are approximately the extent of depletion found in the Sun relative chondritic (c. 12.5). to meteoritic matter; Anders and Grevesse 1989), then parallel problems arise with too little Be depletion compared to the observations. Here, we 3.1.2 The Vicinity of the Solar System stress that beside Li, only a few other elements show depletion in the Sun relative to chondritic The 7Li/6Li ratios of the local ISM appear to be meteorites by the factor of *100 (Te, I, Cs and generally in the Solar range, based on observations Hg; see a recent compilation in Palme et al. of other elements (Be, C, O) whose abundances 2014). This is important to note because Solar are consistent with an input from environment elemental abundances in fact are similar to those where some Li is also synthesized, such as in the meteorites; from the group of light ele- asymptotic giant branch stars, Type II supernovae ments Li–Be–B, Li is the most depleted in the or GCR. Webber et al. (2002) have used data Sun (Asplund et al. 2009). collected with the Voyager 1 and 2 spacecraft and included periods where levels of the Solar modu- lation of energy spectra of various nuclei (6Li, 7Li, 7 9 10 11 Be, Be, B and B) were at minima (1986– 3.5 1988, 1996–2000). From these measurements, interstellar production of 7Li/6Li was estimated to – be about 1.4 1.5, consistent with most models of log [Li/H] +12 3.0 spallogenic 7Li/6Li (see Vangioni-Flam et al. 1999, for a discussion). Consistently, for popula- 7 6 tion I stars such as the Sun, a spallogenic Li/ Li of 0.20 1.38 was calculated (Steigman and Walker 1992). 0.15

An additional free parameter requiring consider- Li 7

ation is the possible galactocentric variation in Li/ 6 0.10 7Li/6Li. Kawanomoto et al. (2009) suggested that such variation appears to be nonexistent for Li 0.05 isotopes whereas it has been clearly demonstrated for D/H and T/H. On the other hand, advanced 4 8 12 modelling of Prantzos (2012) shows that there Radius (kpc) may be a gradient in our Galaxy both in Li abun- 7 6 Fig. 3.5 Model calculations of Li abundance (a) and dance and Li/ Li which would be consistent with 6 7 an increased metallicity toward the inner disc of Li/ Li ratio (b) as functions of radial distance from the center of the Milky Way disc (after Prantzos 2012). the Milky Way (Fig. 3.5). These theoretical Dashed curve = 4.5 Ga, solid curve = present. Thick predictions remain to be yet evaluated. segments were used to calculate the slope in a. More exact evaluation of Li isotopic ratios is hindered by uncertain- Reproduction of the depletion in the Solar Li 7 6 7 abundance relative to meteoritic Li contents is an ties in Li yields of low-mass sources, Li/ Li ratios of ISM, and possible reactions between ISM and GCR important aspect in stellar evolution models and (Prantzos 2012) 24 3 Cosmochemistry of Lithium

The major outcome of the calculations of where the freshly synthesized 7Li is not con- Ahrens et al. (1992) is that any generalization for sumed in nuclear reactions; it is called the stars other than Sun should be avoided, or, at Cameron–Fowler transport mechanism least, carefully examined. Indeed, some stars of a (Cameron and Fowler 1971). Capture of a Sun-like mass still maintain Li abundances more sub-stellar body would enhance abundances of or less in the range of the original starting values. other refractory elements or would have other It follows that the problem lies in reducing or consequences, none of which, however, are avoiding Li burning in the pre-main sequence observed (Koch et al. 2011; Kumar et al. 2011). evolution rather than proposing mechanisms of The photosphere of the Sun (and by inference Li burning in the main sequence. As a conse- the Sun as a whole) is depleted in Li by two quence, the Sun may not be the proper choice for orders of magnitude compared with meteoritic calibrating the extent of Li depletion in stellar abundance (*150×; Anders and Grevesse 1989; regions because in this range of solar masses the Asplund et al. 2009), yet it carries distinctively dependence of Li depletion on the depth of high 7Li/6Li > 33 as measured in sunspots mixing is scaled up. This was also noted by (Ritzenhoff et al. 1997). Technically, sunspots Steigman (1993) who summarized that stellar are cooler surface regions and their Li spectra production of 7Li was very subordinate dur- should thus be better resolved; indeed, evidence ing *4.6 Gyr existence of the Solar System. is provided for small but resolvable amounts of They suggested that, perhaps, the Sun is not 6Li (Ritzenhoff et al. 1997). The high 7Li/6Li representative of the local ISM at *4.6 Ga so ratios in the Sun probably are the result of dif- that Li–Be–B abundances may in fact be over- ferential destruction rates of 6Li and 7Li at abundant. Moreover, there is a difference 2 × 106 and *2.5 × 106 K, respectively between the primordial Li abundance of the (Delbourgo-Salvador et al. 1985). The isotopi- Galaxy and the Solar System; Audouze et al. cally heavy values found for sunspots were also (1983) estimated the primordial Galactic Li obtained in physical samples of implanted solar abundance at Li/H of 11.8 ± 3.2 × 10−11 whereas wind particles in lunar soils. Chaussidon and the Solar System Li/H starting value was esti- Robert (1999) provided an estimate for mated to be as high as *2×10−9, requiring 7Li/6Li > 31 in the solar wind through ion enrichments in Li during the Galactic history. microprobe shallow-depth in situ profiling of In contrast to general Li depletion in the Sun, exposed lunar soil grains, resulting in a reason- some other classes of stars may carry far higher able match between astrophysical predictions and Li loads. Values up to 400× higher than pri- cosmochemical approaches. The analyses also mordial 7Li (Scalo 1976) are quite often observed showed that both the solar wind and spallogenic in giants of various classes (e.g., Brown et al. Li penetrate the very surface of lunar materials 1989; Carlberg et al. 2010; Koch et al. 2011; no deeper than a few tenths of a micrometer. The Monaco et al. 2011; Smith et al. 1995; Waller- observation of heavy Li in the outer shells of stein and Sneden 1982), although the nature of lunar soil particles was confirmed later by Magna these enrichments is not yet completely clear. et al. (2006) although not yielding such high The 7Li/6Li ratios in these massive stars do not 7Li/6Li ratios given the different methodology differ from Li-poor regular stellar bodies used to investigate lunar samples. (Andersen et al. 1984). It may be that 3He cir- It is critical to note that, through time, stellar culating from stellar atmospheres into deeper 7Li/6Li ratios might have undergone significant zones transforms into 7Be which decays to 7Li modifications due to stellar processes, as dis- through (α,γ) reaction (Monaco et al. 2011). An cussed above, while meteoritic Li isotopic ratios important factor in this reaction (see also appear to be more or less constant over the life- Sect. 3.1.1) is the need for a hot stellar envi- time of our Solar System. However, the 7Li/6Li at ronment (>107 K) to produce 7Be and its fast the Solar System start is unknown and alternative transport to cooler subsurface layers (<3 × 106 K) 7Li/6Li evolution of the Sun can be sketched 3.1 Cosmology of Lithium: Big Bang Nucleosynthesis … 25

(Fig. 3.6). We cannot therefore postulate unam- The uncertainty in deconvolving the history biguously that the Sun represents the interstellar and evolution of 7Li/6Li of the Sun may, to some gas composition at the time of formation extent, be linked with variations in 7Li/6Li ratios of at *4.6 Ga. This may imply a process-oriented other stars. It is generally difficult to obtain 7Li/6Li dichotomy in Li isotopic evolution of gaseous and ratios for stars (Piau 2008) due to technical com- solid materials of the Solar System. Of course, for plexities in resolving the spectra of 6Li and 7Li. As the latter we do have physical samples in the form a consequence, the measured 7Li/6Li ratios are of primitive and evolved meteorites such as usually very high and burdened by large analytical chondrites and meteorites from Mars and as- errors. In fact, lower 7Li/6Li limits at *20 are teroidal bodies such as Vesta, as well as samples typical (e.g., GarcíaPérez et al. 2009; Hobbs and from the Moon, collectively spanning time from Thorburn 1994; Hobbs et al. 1999; Maurice et al. the beginning of the Solar System until present. In 1984; Nissen et al. 1999). In this respect, both the general, although the Li isotopic compositions of low-mass halo stars and metal-poor stars appar- Solar System solids and the Sun differ signifi- ently play the determining role in models of evo- cantly today, it may well be that the bulk Solar lution of Li and its isotopes because their lifetimes System started with a more uniform and perhaps are comparable to those of the Galaxy itself and chondritic 7Li/6Li *12, or perhaps even a lower because precise measurements of individual Li value reflecting low 7Li/6Li ratios typical of ISM, abundances in such stars may be critical in con- and evolved to the Sun’s present-day straining the Galactic history and evolution. Lind 7Li/6Li *33 (see above) as a consequence of et al. (2013) found very high 7Li/6Li ratios in stellar processes. Over the same time the com- several metal-poor stars, showing that these were position would have remained essentially con- indeed devoid of 6Li, attesting to standard models stant for solid residua of the Solar proto-nebula of BBN evolution. High 7Li/6Li ratios, practically from which planets accreted because the accretion at the edge of current instrumental analytical process, including multiple impacts, runaway capabilities, appear to be the common feature of collisions and oligarchic growth of planets, did stars. For example, Ghezzi et al. (2009) measured not impose conditions necessary for major chan- 7Li/6Li ratios in several stars and despite many ges in the ratio of Li isotopes. technical barriers and the need for many correc- tions, they found very high 7Li/6Li ratios (≥125) within analytical errors for all studied cases. Li

6 ? Sun Li/ 7 3.1.3 Searching for Exoplanets

present with Lithium

Moon Mars The stellar Li abundance could eventually be ISM ? Vesta (and other asteroids?) Earth chondrites useful in searching for the existence of stellar systems with orbiting planets, as spectral mea- Solar System time start surements of Li abundance are, in some cases, relatively easy. Theoretically, the existence of an Fig. 3.6 Schematic evolution of 7Li/6Li in our Solar orbiting planetary system may affect the stellar System. Solid materials (Earth, Mars, Moon, asteroids, angular momentum as well as surface convective meteorites) all have a constant Li isotopic ratio at *12 whereas Solar 7Li/6Li ratio has apparently evolved toward mixing because rotational forces produced by the higher values as a consequence of preferential destruction movement of large satellites could impart sig- of 6Li over 7Li. It is unclear, however, what 7Li/6Li ratio nificant additional energy to mixing of stellar our Sun started with. Also, local ISM prior to formation of layers. By inference, this would affect the avail- Solar System might have had a range of 7Li/6Li ratios and the present-day 7Li/6Li value may not necessarilly reflect ability of Li for burning. However, alternative that of a new-born Solar nebula explanations may seek Li depletion to be 26 3 Cosmochemistry of Lithium inherited from the pre-main sequence stage of toward the star. However, one must keep in mind stellar evolution. that (i) some of the former stellar systems may In order to establish Li as a practical proxy for host planets not yet detected by our analytical stellar planet-harboring systems, Israelian et al. tools and, (ii) depletion in Li of stars may be (2009) reported on a pronounced Li depletion in related to massive Jupiters which may impart stars with proven orbiting planets. These authors more gravitational influence on stellar mixing. suggested that stars more massive than the Sun Collectively, some limitations resulting from do not suffer from comparable Li depletion due both the model parameters and data evaluation to convective layers that are too shallow to call for further analyses and a better under- transport Li into greater depths where Li-burning standing of stellar environments. is triggered. It may appear that Li depletion is peculiar to Sun-like stars within a very restricted temperature range (±80 K) and wherein neither 3.2 Cosmochemistry of Lithium age nor metallicity are the major determining in Meteorites parameters of Li depletion (Israelian et al. 2009). The results of Israelian et al. (2009) were later Lithium is rendered an important element for critically reviewed by Baumann et al. (2010) who stellar nucleosynthesis (see previous section), concluded that planet-harboring stars do not differ and holds promise as a petrogenetically mean- in their Li abundances from stars without ingful parameter of planetary magmatic differ- anchored planetary systems and that these fea- entiation, crustal evolution and tures may be coupled with different rotation alteration/metasomatism. These assumptions profiles of stars (Charbonnel and Talon 2005). partly follow from its moderately volatile Lithium depletion may be pronounced in stars behavior with a 50 % condensation temperature with thicker convective zones and in stars with a from Solar System gas of 1142 K (Lodders 2003) higher degree of differential rotation between the and its incorporation into forsterite and enstatite core and envelope so that Li is entrained in dee- where it substitutes for Mg. Despite its broad use per, hotter regions. More generally, Baumann in understanding magmatic and post-magmatic et al. (2010) provided evidence against stellar processes, only a handful of detailed studies of planet-hosts possessing surface Li abundance elemental and isotopic distribution of Li in different to those which would be intrinsic to meteorites, considered the building blocks of solitary stars. Also, Ramírez et al. (2012) were not terrestrial planets, has emerged. able to statistically distinguish between In 1960s and early 1970s, it became clear that planet-hosting and planet-free stellar systems (i) chondritic meteorites of diverse classes (car- despite multiple constraints on surface tempera- bonaceous, enstatite, ordinary) are uniformly low ture, mass and metallicity. They concluded that in Li abundance, generally in the range between some of the foreseen differences in Li abundance 1 and 3 ppm (e.g., Balsiger et al. 1968; Dews may stem from subtle disparity in these key stellar 1966; Krankowsky and Müller 1964, 1967; parameters between the two compared groups. Nichiporuk and Moore 1970, 1974; Tera et al. Recently, Figueira et al. (2014) employed 1970), (ii) Li resides primarily in the silicate more stringent statistical tools for larger data set portion of meteorites, (iii) chondrules are gener- and found that differences in stellar surface Li ally depleted in Li relative to the bulk sample abundance exist at statistically significant levels (Krankowsky and Müller 1967; Shima and (>3σ) for planet-devoid and planet-harboring Honda 1966), and (iv) some achondrites may stars whereas a difference in Li abundances carry significant quantities of Li (up to 15 ppm), at *7σ level was found for a selected population for example shergottites and nakhlites coming of stars analyzed by Delgado Mena et al. (2015) from Mars (see Sect. 3.3.2) and eucrites coming and ascribed by the authors to the presence of from asteroid 4-Vesta (see Sect. 3.3.3). Most of massive hot Jupiter-size planets migrating closer these preliminary conclusions were later 3.2 Cosmochemistry of Lithium in Meteorites 27 confirmed by analyses of a larger number of (2006) who postulated that corrections for cos- specimens, but the advent and wide application mogenic contribution performed by Chaussidon of modern state-of the-art analytical instruments et al. (2006) were overestimated and that there is made it clear that Li is distributed heteroge- no real isochron through the data set that would neously on a small scale and that it behaves in a imply the presence of former live 7Be through a complex fashion (e.g., Hanon et al. 1999; correlation between 7Li/6Li and 9Be/6Li. Careful Chaussidon et al. 2006; Maruyama et al. 2009; re-examination of the data of Chaussidon et al. McDonough et al. 2006; Murty et al. 1983; Seitz (2006) by Leya (2011) partly re-assessed the et al. 2007, 2012; Sephton et al. 2013). validity of such an isochron although care must be exercised in proper handling of cosmogenic effects as well as statistical evaluation, combined 3.2.1 Lithium in Bulk Chondritic with cautious petrography. Meteorites The non-trivial task of obtaining information from presolar grains has been evaluated by A few early studies of meteorites produced a Fujiya et al. (2011) who, using the NanoSIMS significant load of Li isotopic data (Balsiger et al. technique, investigated Li (and B) isotopic 1968; Dews 1966; Gradsztajn et al. 1967; Kran- compositions in 12 SiC X grains, thought to kowsky and Müller 1964, 1967; Poschenrieder originate in Type II supernovae. However, all of et al. 1965; Rajan et al. 1980; Shima and Honda the grains analyzed by Fujiya et al. (2011) carried 1966; Yanagita and Gensho 1977) but given the Solar System 7Li/6Li * 12 within analytical contemporary precision of *±10 ‰ and con- errors which the authors ascribed to meteoritic or sidering the lack of a recognized international laboratory contamination. Considering consider- reference material, the only significant conclusion ably higher Li concentrations in SiC grains of reached from those investigations was that the Li Fujiya et al. (2011) compared with those mea- isotopic signature in meteorites overlaps, within a sured by Gyngard et al. (2009), any contribution broad uncertainty; hence it was inferred that the by GCR spallation-produced Li has been con- chondritic (and by extension the bulk Earth) strained to < 4 % but will probably be at sig- 7Li/6Li ratio was *12 (Anders and Grevesse nificantly lower levels. 1989). These early results also placed a bench- Only in 2003, McDonough et al. (2003) pre- mark for the astrophysical observations that some sented the first quasi-systematic investigation of interstellar regions of the Universe have 7Li/6Li Li isotopic compositions in several chondritic ratios far below the chondritic value, some as low meteorites of different petrologic types. These as *2 (e.g., Knauth et al. 2003, 2000). results implied a shift in δ7Li depending on the Further support for varying 7Li/6Li ratios degree of alteration and/or thermal metamor- across the Galaxy, or at least across the Solar phism such that aqueous alteration introduced System, has since come from in situ analyses of higher δ7Li whereas low δ7Li values were asso- presolar SiC grains thought to represent the ciated with higher degrees of thermal metamor- oldest and chemically and physically extremely phism. Moreover, the preliminary data of resistant material of the Solar System available McDonough et al. (2003) invoked δ7Li *0 ‰ (Gyngard et al. 2009; Lyon et al. 2007). These for the Solar System with profound implications studies have shown that some sub-chondritic for the inner Solar System planets. However, 7Li/6Li ratios (down to *9) may be related to more recent studies (McDonough et al. 2006; contribution from GCR spallation reactions, Seitz et al. 2007) did not lend support to these yielding 6Li excess. Chaussidon et al. (2006), preliminary conclusions and have inferred a though, presented potential evidence of the for- Solar System δ7Li of *+3–4 ‰, consistent with mer existence of live 7Be during the formation of a large body of other evidence from terrestrial, chondrules in the Allende meteorite. These latter lunar, martian and asteroidal rocks (see the fol- results were disputed by Desch and Ouellette lowing sections in this chapter). 28 3 Cosmochemistry of Lithium

One of major obstacles in analyzing meteorites δ7Li = 3.9 ± 0.8 ‰, 2SD; n = 5). Collectively, is their petrologic diversity, record of most chondritic meteorites span δ7Li range lar- post-solidification processes that could have dis- gely between +1 and +5 ‰, and only carbona- turbed the original Li distribution, and ceous chondrites of the Renazzo (CR) type (Fehr intra-sample variability. The latter has clearly been et al. 2009), some enstatite chondrites, and a identified in somewhat contrasting results for dark single aubrite (enstatite achondrite) show con- inclusions in the Allende chondritic meteorite sistently light Li isotopic signatures (Pogge von (Seitz et al. 2012; Sephton et al. 2006). Further- Strandmann et al. 2011) that may provide further more, the remobilization of Li isotopes during insights into the distribution of chondrite parent thermal metamorphism or aqueous processes bodies within the early Solar System. From the Li could potentially mask the original Li isotopic perspective it is difficult to assign enstatite signature of meteorites. However, in such cases it chondrites a major role as building material for may provide important information on time scales the Earth and other large planetary bodies in the and temperatures of these secondary events. Solar System (see also Magna et al. 2015) The fact that Li isotopic compositions of because their major element composition is in different meteorite classes do not simply stark contrast with that of the Earth (see discus- reflect their individual and Li-specific sion in Palme and O’Neill 2014). On the contrary, post-crystallization histories, often involving many other geochemical aspects (in particular thermal processing, reheating and/or fluid perco- isotopic compositions of many refractory ele- lation, is mirrored in roughly consistent results for ments such as Ti, Cr, Ni, Ba, etc.) are not at odds several classes of chondritic meteorites (Pogge with a view of enstatite chondrites as major von Strandmann et al. 2011; Seitz et al. 2007). In building blocks of terrestrial planets (see also fact, all classes of chondrites, (i.e., carbonaceous, opposing views of Fitoussi and Bourdon 2012 vs. ordinary and enstatite) show a similar extent Javoy et al. 2010, and discussion therein). of δ7Li variation, paralleled by only slightly dif- Although it seems that the Earth and Mars were ferent mean values (Fig. 3.7). The carbonaceous formed from a complex mixture of primitive chondrites include Allende, Murchison, Orgueil, materials (e.g., Warren 2011), a significant effort Karoonda, Ornans, Efremovka, Vigarano, Tagish should be made toward a better understanding of Lake and several other specimens and have planetary-forming processes; if the behavior of Li δ7Li = +3.2 ± 1.7 ‰ (2σ, n = 35); among ordinary is better understood it could aid in this effort. chondrites, Bruderheim, Bjurböle, Semarkona, Lithium data for other meteorite classes are Tuxtuac, Parnallee, Chainpur, Saint-Séverin For- scant. Olivine from the Admire pallasite showed est Vale and Kernouvé were also analyzed, the a rather Li-poor nature (0.8 ppm) and a δ7Li mean δ7Li = +2.4 ± 1.6 ‰ (2σ, n = 34); Quingz- value (+3.4 ‰) fully in the range of chondritic hen, Indarch, Abee, Hvittis or Neuschwanstein are meteorites (Seitz et al. 2007). Further analyses, among the analyzed enstitate chondrites and the conducted using SIMS instrumentation, have group’s δ7Li = +1.9 ± 1.5 ‰ (2σ, n = 10). confirmed this value although with some scatter A comparison of larger data sets collected for and with a few pallasites showing lighter Li individual samples seems to indicate a isotopic compositions (δ7Li < + 0.6 ‰; Bell et al. sample-scale heterogeneity. For example, avail- 2008). Preliminary data for olivine and pigeonite able data for carbonaceous chondrite Orgueil extracted from ureilites show somewhat higher yield a mean δ7Li = +3.6 ± 1.4 ‰ (2σ,n=5) (and similar) δ7Li values for both co-existing and a similar extent of variability has been minerals (+6.3 and +6.7 ‰, respectively), found for the ordinary chondrite Parnallee although pigeonite displayed significantly larger (δ7Li = +2.9 ± 1.7 ‰,2σ; n = 3). Slightly scatter of the Li isotopic compositions (Single- less heterogeneous results were found for tary et al. 2008). Whether or not these signatures other carbonaceous chondrites, Allende reflect kinetic processes on the respective parent (δ7Li = 2.6 ± 1.0 ‰, 2SD; n = 5) and Murchison bodies remains to be tested. However, only a 3.2 Cosmochemistry of Lithium in Meteorites 29 ENSTATITE CHONDRITES ORDINARY CHONDRITES CHONDRITES CARBONACEOUS

-4 -2 0 2 4 6 δ7Li (‰)

Fig. 3.7 Lithium isotopic composition of different Indarch). This indicates sample-scale heterogeneity; con- groups of chondritic meteorites. Several features are sequently, care should be exercised when analyzing apparent. First, within rather large uncertainty, all major chondritic materials. It also appears that the data of chondrite groups (carbonaceous, ordinary, enstatite) have McDonough et al. (2003) may suffer from yet unac- broadly similar mean δ7Li, although the latter group has counted analytical artifacts because other analyses of on average slightly lighter lower mean δ7Li. Second, Karoonda, Felix, Efremovka, and Chainpur are consistent some samples show particular variation in δ7Li (e.g., with the chondrite group means. Data sources: James and Orgueil, Murchison, Karoonda, Allende, Parnallee) Palmer 2000; Seitz et al. 2007, 2012; Pogge von whereas published data for some others are nearly Strandmann et al. 2011; Magna et al. 2006; Sephton identical (e.g., Ornans, Dhurmsala, Bremervörde, et al. 2004, 2006, 2013; McDonough et al. 2003 more detailed survey of a larger population of paralleled by considerations of intra-sample chondrites (and primitive achondrites, such as heterogeneities, could provide further clues to aubrites, angrites, brachinites, and ureilites), the history of Li in the early Solar System. 30 3 Cosmochemistry of Lithium

3.2.2 Lithium Distribution They suggested that this variation reflected and Isotopic Systematics mixing of three different nucleosynthetic sources in Constituents (BBN, ISM, and dense molecular clouds); at of Chondritic Meteorites least two different sources (high-energy spalla- tion reactions and a common silicate reservoir) The early work of Dews (1966) showed that were also invoked for chondrules from the chondrules are depleted in Li (< 0.27 ppm) rel- Mokoia meteorite with δ7Li values as low as ative to bulk chondrites and that there are some −54 ‰ (Robert and Chaussidon 2003). Different variations in Li abundance in chondrules from distribution of Li and δ7Li was reported for individual samples. These findings were later Allende by Bell et al. (2008) who found gener- reproduced for other meteorites, though with less ally low Li concentrations (< 1 ppm) for Mg-rich pronounced Li depletion in chondrules relative to olivine (with forsterite contents > 90 mol%) and matrix (Nichiporuk and Moore 1970). On the between *0.5 and *4.5 ppm Li was found in contrary, Krankowsky and Müller (1967), Bal- olivine with a larger fayalite component siger et al. (1968) and Murty et al. (1983) (>10 mol%). The Li isotopic variability in for- reported analyses of chondrules and bulk speci- steritic olivine (*0to*30 ‰) is contrasted by mens with somewhat contrasting results, even rather uniform δ7Li = +3 ± 2 ‰ in Fe-rich oli- with up to two-fold Li enrichments in chondrules vine. Bell et al. (2008) discussed these findings relative to bulk specimens. More recently, Hanon as reflecting preferential loss of 6Li by evapora- et al. (1999) used SIMS to determine Li (and Be– tion at high temperatures. B) elemental distribution in several meteorites, In situ measurements of Li abundance and including Allende and Semarkona. Their results isotopic composition in several chondrules from indicated a rather narrow range of Li concentra- the Allende meteorite by Maruyama et al. (2009) tions in chondrites (*0.2–2.6 ppm), paralleled revealed Li-poor character of olivine in three by a lesser elemental variability of Li within chondrules (generally below 0.3 ppm) whereas individual chondrules compared with B. In con- olivine in three other chondrules had mean Li trast to preferential incorporation of Be and B abundance from 0.6 to 0.8 ppm. Elevated Li into mesostasis, Li partitioning into crystals contents were found in low-Ca pyroxene of two versus mesostasis between different chondrites is chondrules (>3.6 ppm), matrix (*2 ppm), and in not uniform. The coupled Li–Be–B elemental Na-rich mesostasis (>1 ppm). Parallel measure- systematics also indicated immobile character of ments of δ7Li in selected olivines revealed a Li and Be during metamorphism of the parent significant variation between −32 and +21 ‰ body of ordinary chondrites (Hanon et al. 1999). without any clear systematics. Because most The Li–Be–B systematics in Semarkona chon- mesostasis in Allende is rather low in Li, Mar- drules were independently also investigated by uyama et al. (2009) invoked leaching of Li by Chaussidon and Robert (1998) who found much aqueous fluids. The Na-rich phases located at larger variations in Be–B contents relative to Li chondrule rims appear to have captured some Li (from *0.5 to *3.7 ppm Li vs. two and three although the process of Li mobilization and orders of magnitude variations for Be and B, re-incorporation may have operated at different respectively). Moreover, no uniform partitioning times. Recently, Seitz et al. (2012) investigated between mesostasis and pyroxene was observed Li elemental and isotopic systematics in *90 (similar to findings of Hanon et al. 1999), chondrules by means of MC-ICPMS technique. implying the lack of magmatic equilibrium due Again, three observations are apparent: (i) no to rapid cooling of chondrules. systematic relationship of Li contents and pet- Chaussidon and Robert (1998) also reported rogenetic parameters (e.g., magnesium number δ7Li values for multiple measurements of three [Mg#], defined as (molar MgO/(molar chondrules from Semarkona with a mean value MgO + molar FeOtot)), (ii) overall low Li con- for the majority of data points at +10 ± 10 ‰. tents in chondrules relative to bulk chondrites, 3.2 Cosmochemistry of Lithium in Meteorites 31 and (iii) a large range in δ7Li values among the acting on the meteorite parent bodies at different, chondrules without any peculiar systematics and albeit low, temperatures (Sephton et al. 2013). relation to the degree of alteration/ metamorphism, degree of melting or major ele- ment chemistry. From these observations, Seitz 3.2.3 Lithium in Other Meteorites et al. (2012) implied compositional hetero- geneities in chondrule forming regions in the Beside chondrites and achondrites, a Solar nebula. Collectively, the data for individual well-populated clan of other meteorite types, components of chondrites show a large scatter in pallasites, mesosiderites and iron meteorites, both Li contents and isotopic compositions, exists (Krot et al. 2014). A significant proportion indicating non-uniform behavior of Li during of these meteorites is formed by Fe–Ni metal in chondrule formation and subsequent processes which Li will likely be highly incompatible. This after the chondrules were embedded into bulk was proven by comparing Li contents in silicate chondritic matter. There is, somewhat surpris- and troilite from Odessa IA iron meteorite ingly, limited evidence for post-magmatic with *1–2 ppm Li in silicate fraction ver- re-distribution of Li in chondrites through aque- sus *40 ppb in the metal fraction (Krankowsky ous alteration or thermal metamorphism whereby and Müller 1967). Murty et al. (1982) investigated the former process could be efficiently tracked iron meteorites from several classes and found with Li (see Chap. 6). This can be linked to the that Li is distributed heterogeneously in these assumption of Seitz et al. (2012) and Sephton meteorites. Moreover, most Li appeared to be et al. (2013) that the whole-rock Li systematics concentrated in a non-magnetic fraction whereas appear to be a mixture of several distinct com- the magnetic fraction retains a subordinate pro- ponents with largely different Li characteristics. portion of Li. Despite very low Li contents (low Important information on post-magmatic his- tens of ppb Li, as determined from the measured tories of chondrites may also be gathered by 6Li contents by neutron activation analysis), a applying sequential leaching procedures which resolved difference in Li contents between IA seek to selectively dissolve individual compo- (e.g., Canyon Diablo, Odessa, Toluca) and IIAB nents. For example, acetic acid usually removes a (e.g., Coahuila, Sikhote Alin) versus IIIA (e.g., non-silicate fraction out of the bulk material. By Cape York, Henbury) iron meteorites was found this technique, Sephton et al. (2004) isolated (>*20 ppb in IA–IIAB vs. *5 ppb in IIIA). carbonates, and Fe oxides and oxyhydroxides However, no further investigations were made from the Murchison meteorite. They found sub- that considered these meteorite classes. Given the stantial δ7Li variations between the individual low abundances of Li in metal, consistent with its components, i.e., carbonate-rich, phyllosilicate- lithophile character, the possible implications for rich matrix and chondrules that simply constitute histories of the respective parent bodies of iron the bulk mass of the meteorite. Carbonates in their meteorites may be rather limited. study carried a heavy Li isotopic signature (>+12.6 ‰), which is largely consistent with terrestrial systematics of Li in low-temperature 3.3 Cosmochemistry of Lithium carbonates. A supplementary study by Sephton in Planetary Bodies et al. (2013) found that Fe mineralogy (i.e., the type of the Fe oxide or oxyhydroxide) exerts the Comparatively little has so far been done in order major control on abundance of Li in non-silicate to improve our understanding of the behavior of (i.e., acetic acid-leachable) phases. Also, acetic Li and its isotopes in non-terrestrial materials acid-leachable phases carried high δ7Li, similar to despite at least three decades of successful mea- the findings of Sephton et al. (2004). These surements of Li isotopic compositions and despite observations may have implications for post- the importance of Li in stellar processes. magmatic processes, such as aqueous alteration Only with more recent developments and 32 3 Cosmochemistry of Lithium advancements in analytical techniques and Early lunar Li isotopic research was focused instrumentation (see Chap. 2), several studies on the detection in regolith of a solar wind have presented precise and accurate measure- component (Eugster and Bernas 1971) although ments of Li isotopic compositions in lunar rocks measurements were compromised by large ana- as well as in martian, Vestan (howardite–eucrite– lytical errors. Solar wind was theoretically diogenite [HED] clan) and other meteorites, with modeled with high 7Li/6Li ratios resulting from some implications for the histories and evolution preferential destruction of 6Li at the bottom of of their respective parent bodies. Whereas the convective zone in the Sun (Bernas et al. numerous samples of these bodies exist for direct 1967). This overall depletion of Li as well as laboratory investigations, recent efforts have also high 7Li/6Li in the Sun was later confirmed the- been made to detect Li abundance in Mercury, the oretically (Delbourgo-Salvador et al. 1985), by innermost planet of the Solar System, for which direct astrophysical observations (e.g., Ritzenhoff no physical samples are at present available. et al. 1997). The early measurements of Eugster The preliminary report by Irving et al. (2013) and Bernas (1971), however, were unsuccessful on a recent find of North West Africa in determining 7Li excess in two lunar breccias (NWA) 7325 requires further investigations but (10046 and 10074) and a coarse-grained soil Li-depleted character of this meteorite (0.53 ppm) (10084), potentially due to long exposure of is unique in the context of its provenance from a surface. Empirical support for heavy Li enrich- differentiated planet. Whereas there may be little ment in solar wind arrived more recently, direct link between the surface Li and atmo- through high-resolution spatial analysis of indi- spheric Li abundance, Sprague et al. (1996) were vidual lunar soil particles (Chaussidon and not able to detect Li in the atmosphere of Mer- Robert 1999). Estimates from the analysis by cury. This would be consistent with the extremely Chaussidon and Robert (1999) of Apollo soils depleted nature of NWA 7325, but further 10060, 79035, 79221 and 79261 suggested that investigations are required to answer ticklish solar wind had very high 7Li/6Li of *31, questions concerning this meteorite. whereas the bulk Earth value is *12. The data of Eugster and Bernas (1971) showed 7Li/6Li ratios similar to those found 3.3.1 Lithium in Earth’s Moon previously for terrestrial materials. Also, Garner et al. (1975) measured 7Li/6Li ratios in several Since the return of Apollo samples, a number of lunar soils and rocks; these data too were indis- contributions to the annual Lunar and Planetary tinguishable from the terrestrial reference mate- Science Conference contained Li concentrations rial. These first measurements, despite the lack of for many lunar samples. Some significant obser- internationally recognized and accepted stan- vations arise from these studies. For example, dardization material, and being far from the Dreibus et al. (1976) estimated 1.6 ppm Li for current levels of precision and accuracy, pro- bulk Moon and implied from Li–Zr elemental vided the first glimpse at the Li isotopic com- correlations that the Moon inherited *40 % of its position of the Moon. The lack of distinction mass from high-temperature condensates and the between Li isotopes in lunar and terrestrial rest of its matter came from Mg-silicates. How- materials was only reinforced three decades later ever, there were some discrepancies between by Magna et al. (2006) who analyzed a suite of Li/Zr, which scattered significantly, and Li/Be lunar rocks and soils with different solar wind (considered constant at *4.6 for the Moon, as exposure ages. The results for soils did not evidenced from maria basalts and KREEP-rich deviate drastically from values reported for the and KREEP-poor samples; KREEP = potassium– bulk silicate Moon and, by inference, from the rare earth elements–phosphorus-enriched), which Earth. The contribution of the solar wind to bulk may point to dissimilar behavior of these ele- Li in lunar soils was detected for different size ments despite their refractory behavior. fractions of the lunar soil 67601, with slight 7Li 3.3 Cosmochemistry of Lithium in Planetary Bodies 33 excess in the finest fractions (<42 μm), most +2 ‰), and appears to reflect long-term secular likely formed by degradation of surface shells of evolution of the crust–mantle system. For the regolith particles through long-term solar wind Moon, chemically evolved lithologies are quite abrasion. Nevertheless, translated into bulk lunar rare in the Apollo collection, and their origin geology, there is no detectable shift in δ7Li for through silicate liquid immiscibility (e.g., Neal magmatic rocks (Magna et al. 2006; Seitz et al. and Taylor 1989) is likely to be dissimilar to the 2006), making them suitable candidates in which formation of igneous rocks of the terrestrial to study processes of extreme magmatic differ- continental crust (Taylor and McLennan 2009). entiation in the lunar magma ocean (Snyder et al. Whereas many parameters suggest a comple- 1992) under nominally dry conditions, although mentary relationship of mare basalts and anor- the genuine water content of the Moon is con- thosites (Taylor and Jakeš 1974), low Li tested (e.g., Hauri et al. 2011; Saal et al. 2008; abundance in lunar highland rocks does not find Sharp et al. 2010). This may be important by its equivalent on Earth. Instead, low Li contents considering a possible role of chemistry of min- in pristine ferroan anorthosites, paralleled by eral phases that build Li into their structures and elevated δ7Li (up to +9 ‰; Magna and Neal that may thus reflect changes in conditions dur- 2011; Magna et al. 2006), may imply kinetic ing magmatic fractionation and/or thermal effects in a plagioclase-rich layer, although rapid metamorphism (Chaklader et al. 2006). preferential diffusion of 6Li into newly formed Steele et al. (1980) measured Li contents in plagioclase (Coogan 2011; Coogan et al. 2005) plagioclase from a wide spectrum of lunar rocks, would likely drive 7Li/6Li instead towards low including anorthosites, troctolites, norites, brec- values. Therefore, the heavy Li isotopic signature cias, and mare basalts. They found surprisingly of the lunar ferroan crust remains somewhat large variation between 1.2 and 21 ppm Li, with enigmatic and needs to be cautiously modeled. an extraordinary plagioclase measurement from a Shearer et al. (1994) measured the Li, Be and B KREEP (K–REE–P-rich) basalt clast from sam- contents in picritic lunar volcanic glasses and ple 15344 having 54 ppm Li. The latter con- found that Li and other light lithophile element centration reinforces the incompatible contents vary greatly among different sampling element-enriched nature of KREEP (Warren sites (Apollo 11, 12, 14, 15 and 17, respectively). and Wasson 1979). Steele et al. (1980) also These authors suggested that effusive volcanic found a positive correlation between Li and Na, glasses cooled quickly under near-vacuum con- although mafic mineral phases host significant ditions and that they are genetically unrelated to amounts of Li and plagioclase is the principal mare basalts; volcanic glasses appear to more carrier for Na. This may mean that buoyant closely reflect the intrinsic Li content of the Moon, anorthositic lunar crust was derived from a estimated at *0.8 ppm (Taylor 1982), although roughly common reservoir whereas high but other estimates have been made (see above). The variable Li contents in plagioclase from Mg-rich somewhat heterogeneous nature of these glasses, lithologies either reflect several isolated and however, is underscored by significant differences localized sources, or were actively communicat- in Li abundance and isotopic compositions yiel- ing with those regions of residual lunar magma ded for different aliquots of green glass 15426 and ocean that were not yet solidified. orange glass 74220 (Magna et al., 2006; Seitz Very low Li contents (Magna and Neal 2011; et al. 2006). The nature of this discrepancy Magna et al., 2006) in very ancient anorthosites remains unresolved at present, although these (>4.3 Ga; e.g., Borg et al. 2011; Carlson and materials may represent a closer window into the Lugmair 1988; Norman and Taylor 1992) appear deep lunar mantle (Longhi 1992). to be an intrinsic feature of anorthositic portion Only rather recently, Magna et al. (2006) and of the lunar crust. Terrestrial upper continental Seitz et al. (2006) published the first crust carries high Li abundance high-precision Li isotopic data for bulk lunar (mean *35 ppm) and a diagnostic δ7Li (0 to samples (Fig. 3.8). Both studies showed 34 3 Cosmochemistry of Lithium

FANs 16 Mg-suite glass

12

8 Li (‰) 7 δ high-Ti

4 KREEP-rich

low-Ti

0 0 5 10 15 20 25 30 Li (ppm)

Fig. 3.8 Lithium content versus δ7Li in lunar samples found for ferroan anorthosites (FANs), representative of a (Day et al. 2008; Magna and Neal 2011; Magna et al. significant portion of the lunar highlands. A rather limited 2006; 2009; Seitz et al. 2006). An apparent difference variability has been observed for plutonic Mg-suite rocks between low-Ti and high-Ti basalts, as well as a more (dunites, norites, troctolites). A trend toward Li-enriched detailed assessment of Li systematics in volcanic glasses character of the KREEP reservoir is obvious and is (Shearer et al. 1994), appear to rule out a single source for consistent with modelled Li abundance of 56 ppm these major reservoirs. Notable depletion in Li has been (Warren and Wasson 1979) independently that Li isotopic composition of the isotopically lighter than the latter, and that silicate lunar mantle is nearly identical to that of low-Ti mare basalts may better reflect the silicate the Earth, with consequences for the Earth–Moon mantle of the Moon. This was also invoked from formation in a giant impact (Canup 2004) and Hf isotopic compositions (Beard et al. 1998) and subsequent turbulent exchange of hot ionized gas the distinctly greater proportion of low-Ti mare after the collision (Pahlevan and Stevenson basalts at the lunar surface (Giguere et al. 2000). 2007). Importantly, both studies have also shown A note should be made that the bimodality of that lunar rocks may be quite heterogeneous on a lunar mare basalt chemistry based on Ti contents specimen scale as is obvious from several sam- is somewhat ad hoc. More recent remote sensing ples that were measured independently (such as data reveal a rather continuous range in TiO2 KREEPy impact melt 14310, orange glass contents, suggesting that the bimodality is a 74220, green glass 15426, quartz-normative sampling bias from the locations of sample return mare basalt 15475, and olivine-normative mare sites (Giguere et al. 2000). The Li isotopic basalt 15555) and that, therefore, the δ7Li of the dichotomy related to the differences in mantle bulk silicate Moon must be interpreted with sources was later confirmed with oxygen, iron caution. This is further stressed by the fact that and magnesium isotopic differences between there are no direct mantle rocks in the Apollo these distinctive Ti varieties (Liu et al. 2010; collection and that the chemical and isotopic Sedaghatpour et al. 2013; Spicuzza et al. 2007). fingerprints of the deep lunar mantle may thus However, anti-correlated Li and O isotopic only be inferred indirectly. Indeed, Magna et al. compositions may reflect different processes (2006) and Day et al. (2008) have shown that and/or crystallographic control that individual there are statistical differences between the Li mineral phases may exert on the respective bulk isotopic compositions of low-Ti and high-Ti isotopic signatures. For example, a heavy Fe mare basalts, such that the former are in general isotopic signature in high-Ti mare basalts (Liu 3.3 Cosmochemistry of Lithium in Planetary Bodies 35 et al. 2010; Poitrasson et al. 2004) appears to broadly similar (or lighter) Li isotopic signature reflect a control by ilmenite (Craddock et al. compared to high-Ti mare basalts may imply (i) a 2010). Although ilmenite segregation would not role of water that exsolved only after crystal- affect Li due to incompatibility, it could imply a lization of high-Ti mare basalts, or (ii) a change generic feature of lunar magmas to become a in conditions of KREEP crystallization. useful proxy for studying crystal fractionation Collectively, the lunar crystalline rocks have a processes in settings where water can be ignored rather restricted range in δ7Li that appears to (unlike in most terrestrial igneous systems). reflect differences in crystallization assemblages A more comprehensive study of Li isotopes in of the two major sampled groups of mare basalts. basalts from different Apollo sites was performed In order to estimate δ7Li of the bulk silicate Moon, by Magna et al. (2009), who reinforced the Li we tentatively consider low-Ti mare basalts more isotopic dichotomy between isotopically light reliable because their mantle sources were derived low-Ti and heavy high-Ti mare basalts (Fig. 3.8) earlier in the sequence than those of high-Ti mare resulting from distinct periods of crystallization basalts (Snyder et al. 1997, 1992; Taylor and of their respective cumulate sources from the Jakeš 1974) and because they form *90 % of residual lunar magma ocean, modeled earlier exposed mare basalts (Giguere et al. 2000). The (Day et al. 2008). Magna et al. (2009) also estimated δ7Li of the bulk silicate Moon is * confirmed variability of Li isotopic compositions + 4.0 ‰ which is slightly heavier than that of the within the respective low-Ti and high-Ti mare Earth’s upper mantle but within uncertainty. It basalt groups, in many cases on the order thus appears that the exotic δ7Li * +15 ‰ of *3 ‰. An important observation of the inferred for the source of low-Ti basaltic meteorite modern lunar Li data set is that KREEP-rich NWA 479 (Barrat et al. 2005) is less realistic and rocks such as 15386, 14310, and high-potassium strong kinetic modifications seem likely. Apollo 11 mare basalts with a (K)REEP signa- ture (10017, 10057, 10069) (Jerde et al. 1994) appear to carry a uniform Li isotopic signature 3.3.2 Lithium in Mars paralleled by elevated Li abundance relative to other Apollo 11 basalts. The high Li abundance Contrary to lunar rocks for which we have spatial in KREEP is a direct consequence of its forma- control through manned collection by the astro- tion during the last stages of the lunar magma nauts of six successful landing missions and three solidification (>99 %; Snyder et al. 1992; Warren robotic sample return missions, it is not easy to and Wasson 1979) combined with the moder- assign a specific location on the surface of Mars ately incompatible character of Li in magmatic for the >150 currently known martian specimens, process. The key aspect of this observation is that including pairings. This poses problems in putting the mean δ7Li of KREEP is slightly lower than the petrologic and geochemical information into a that of high-Ti mare basalts although their rela- broader context of martian magmatic history. To tive chronology is basically the opposite. The date in terrestrial collections there are basaltic, δ7Li homogeneity observed for KREEP-rich olivine-phyric, olivine–orthopyroxene-phyric, lithologies is important in the view of differ- lherzolitic, diabasic and gabbroic shergottites, ences in magmatic ages for Apollo 11 and Apollo cumulate clinopyroxenites (nakhlites), dunites 14 and 15 rocks, respectively (Shearer et al. (chassignites), collectively called Shergotty– 2006, and references therein), suggesting a rela- Nakhla–Chassigny (SNC) meteorites, orthopy- tively uniform process of formation of chemical roxenite Allan Hills (ALH) 84001, and polymict signatures of this spatially and temporally dis- breccia Northwest Africa (NWA) 7034; some of tinctive reservoir. By considering a Rayleigh the specimens are paired whilst some are unique. process for residual lunar magma ocean (Day Apart from petrographic distinctions, shergottites et al. 2008), KREEP should have distinctly are also classified into depleted, intermediate and higher δ7Li than high-Ti mare basalts. Yet, a enriched shergottites, respectively, according to 36 3 Cosmochemistry of Lithium their incompatible trace element patterns. In this studies focused on their light lithophile element respect, it is important to note that Li isotopes are geochemistry (LLE: Li, Be and B). From detailed not fractionated during impact events (Magna in situ analysis of LLE abundances in pyroxene et al. 2011) which are thought to be the major from the Shergotty meteorite fall, McSween et al. process by which martian meteorites are exca- (2001) inferred that there was significant water vated from the surface. We stress that martian present at the time of crystallization of the min- meteorites did not undergo melting but are heavily eral core at depth and that degassing operated en shocked whereas terestrial impacts produced large route to the surface. Lentz et al. (2001), Beck amounts of melts (tektites), which resulted from et al. (2004) and Herd et al. (2005) extended instantaneous melting of surface materials. The these observations of degassing- and original information during these latter events can water-control on Li elemental and isotopic sys- thus be lost through homogenization of melted tematics in Mars to nakhlite meteorites. This idea sediments and soils, as reported for Li isotopes was, however, soon abandoned (Beck et al. 2006; (Magna et al. 2014b). Therefore, the situation with Herd et al. 2004) through elaboration of the martian meteorites is fortuitous for further petro- concept of diffusive kinetic modifications of Li genetic and geochemical studies despite highly (Coogan et al. 2005; Jeffcoate et al. 2007; shocked nature of these precious samples. Lundstrom et al. 2005; Parkinson et al. 2007). Since it was recognized that some meteorites The bulk behavior of Li during planetary very probably come from Mars (Becker and crystallization processes remained less con- Pepin 1984; Bogard and Johnson 1983), several strained (Herd et al. 2004) and the implications

8.0 MORB

6.0

4.0 Li (‰) 7 δ 2.0

0.0

-2.0 02468101214 Li (ppm)

Fig. 3.9 Lithium content versus δ7Li in Martian mete- exception of Shergotty, Zagami, EETA 79001A and ALH orites. A rather restricted variation in Li abundance and 84001. Some samples appear to have experienced isotopic composition is observed for intermediate sher- terrestrial weathering (e.g., the low δ7Li of NWA 4925, gotittes compared to the larger range of δ7Li found for discussed in more detail by Magna et al. 2015). Another enriched and depleted shergottites. Nakhlites (clinopy- interesting feature is Li-depleted character of most roxenites) carry higher Li contents whereas orthopyrox- Martian meteorites when compared to terrestrial MORB enite ALH 84001 has low Li content and δ7Li. The (see Sect. 5.2.1). This may imply a distinct magmatic polymict breccia NWA 7034 seems to provide first-order evolution of the Martian mantle because this Li depletion constraints on Li systematics of the Martian surface comprises all basaltic rocks (primarily shergottites) and because its major element composition closely resembles extends over the range of magmatic ages (from about 600 some regions of Mars (Agee et al. 2013; McSween et al. to roughly 150 Ma), recorded for basaltic products of 2009). Multiple data obtained for different aliquots of the Martian magmatism. Data sources: Magna et al. (2006, same samples are not distinguished but they show rather 2015), Seitz et al. (2006), Reynolds et al. (2006) and consistent δ7Li results to within ±1.2 ‰ (2σ), with the Filiberto et al. (2012) 3.3 Cosmochemistry of Lithium in Planetary Bodies 37 for Li isotopic compositions were largely unex- SNC meteorites are less representative of the plored. Magna et al. (2006), Reynolds et al. martian surface (McSween et al. 2009), although (2006) and Seitz et al. (2006) published the first petrologically and chemically distinguished Li isotopic data for several martian meteorites lithologies are being increasingly discovered (Fig. 3.9) and found that Mars has a similar range (e.g., Stolper et al. 2013). This may also have in δ7Li as both the Earth and the Moon, implying profound consequences on bulk planetary Li dis- that (i) planetary size and timing of bulk planet tribution as the Earth’s continental crust carries a solidification have little effect on the final Li significant proportion of Li (Teng et al. 2004) that isotopic composition of the respective silicate accumulated during long-term cycles of crustal mantles, and (ii) all three major planetary bodies growth, involving multiple episodes of subduc- appear to have broadly similar Li isotopic com- tion, re-melting, recycling etc. It is clear that none position close to chondritic values with only of the SNC meteorites that were delivered to the limited variation (Magna et al. 2015). The δ7Li of Earth via a series of independent impacts on the the bulk silicate Mars (BSM) has been estimated martian surface between *0.7 and *20 Ma to be similar to that of the bulk silicate Earth (Nyquist et al. 2001) represents a true heritage of (Magna et al. 2006; Seitz et al. 2006) from a the martian crust. This has become evident from rather limited number of meteorites analyzed. Li analyses of a larger number of SNC meteorites More recently, Magna et al. (2015) have esti- (Magna et al. 2015; Seitz et al. 2006) that do not 7 mated δ LiBSM = +4.2 ± 0.9 ‰ (2σ) by using a deviate significantly from the Li isotopic range larger meteorite population and considering only found for chemically similar terrestrial (i) falls and cold desert meteorites because these mantle-derived rocks (cf. terrestrial continental meteorites are less prone to alteration (Fehr et al. crust with different systematics, see Chap. 5). 2009) as well as wall-rock contamination, Although it is not necessary to expect the (ii) insignificant fractionation of δ7Li during martian crustal systematics to be similar to those magmatic process (Tomascak et al. 1999), and of the Earth (compare, for example, the vastly (iii) meteorites with high MgO contents (>15 wt different and reversed Li systematics of lunar %) because these rocks are thought to most clo- anorthositic crust; see Sect. 3.3.1 and Fig. 3.8), sely approach primitive melt compositions the silicate character of bulk Mars may be (Nyquist et al. 2009). It is interesting to note that indicative of a Li signature broadly similar to that this estimate involves enriched, intermediate as of the Earth. This is underscored by a larger suite well as depleted shergottites, meaning that of martian meteorites analyzed for Li (Magna whatever process was responsible for the et al. 2015). An appealing feature is a low Li incompatible element patterns in these high-MgO content of some of the primitive martian mete- rocks, it apparently did not impose any effect to orites, often < 2 ppm (Fig. 3.9). These low Li δ7Li. Pre-terrestrial as well as terrestrial weath- contents are largely missing in pristine ering may become an issue under particular mantle-derivative melts on Earth, such as conditions but most martian meteorites appear to mid-ocean ridge or oceanic island basalts, gen- be devoid of these effects (Magna et al. 2015). erally with >3 ppm (see Chap. 5). Contrary to lunar stratigraphy with abundant More recently, two peculiar martian meteorites early crystallized crust (Shearer et al. 2006) and were found that may bear some important infor- the important role of crustal evolution for global mation on the chemistry of the surface of Mars. terrestrial geological history (Taylor and The Tissint fall is a depleted shergottite akin to McLennan 1985), no clear evidence for a similar other depleted shergottites in its chemical com- major reservoir with vastly different chemistry position but it contains peculiar black glass relative to the mantle has so far been identified on enclosed in the main mass that is thought to carry Mars. Orbital remote sensing as well as rover a chemical signature of martian surface and in situ data have shown that surface lithologies of atmosphere (Chennaoui Aoudjehane et al. 2012); Mars may have basalt-andesitic affinity and that we note here the recent disagreement with such a 38 3 Cosmochemistry of Lithium view (Barrat et al. 2014). The chemical compo- chassignites (dunites) are co-magmatic sition of *4.4 Ga (Nyquist et al. 2013) polymict (McCubbin et al. 2013) remains to be further breccia NWA 7034 fits with in situ and remote investigated but available Li data are consistent sensing data (Agee et al. 2013; Humayun et al. with that interpretation. Saline brines have also 2013), thus providing evidence for a component been advocated for ancient ALH 84001 cumulate of Mars, temporally, chemically and petrologi- orthopyroxenite (Lapen et al. 2010), once cally distinct from the SNC meteorites. In par- thought to carry signatures of early pre-terrestrial ticular, the Li isotopic composition of NWA 7034 life (McKay et al. 1996). The light Li isotopic (−0.2 ‰; Magna et al. 2015) points toward Li signature in this particular meteorite (Magna elemental and isotopic distinction between the et al. 2015; Seitz et al. 2006) is consistent with martian mantle and crust (Fig. 3.9), similar to that crustal contamination or the imprint of fluid observed for the Earth (Teng et al. 2004) despite percolation. Moreover, local brines have also the smaller magnitude of difference in Li abun- been inferred for NWA 7034 (Agee et al. 2013) dance. Here we note moderate Li contents (*10 and its low δ7Li may confirm large-scale fluid ppm) in martian surface lithologies, obtained with activity during the earlier, “wetter” history of the Curiosity mission (e.g., Ollila et al. 2014). Mars, considering abundant water dissolved in It is viable that part of the δ7Li variation martian magmas (McCubbin et al. 2012; recorded in martian meteorites may hint at McSween et al. 2001). A recent theoretical metasomatic activity during or shortly after the treatment by Fairén et al. (2015) modelled the emplacement of lavas at very shallow depths. evolution of Li elemental and isotopic signature Such a view is consistent with tight positive during weathering of a basalt on the martian correlation between Li and Cl abundances in the surface because the early atmosphere of Mars martian suite, paralleled by broadly positive was different from the current one and also from correlation of δ7Li with Li/Cl ratios. Curiously, the Earth’s atmosphere in terms of chemical and these observations were also made for geother- physical parameters. The formation of various mal fluids from the Taupo Volcanic Zone, New low-temperature clay minerals could impart dis- Zealand, and Yellowstone, USA (Bernal et al. tinctive Li systematics to these weathering 2014; Millot et al. 2012; Sturchio and Chan products in the presence of water. Whether this 2003), implying a low-tempered origin of Cl-rich computational modelling accurately simulates fluid percolation in Mars. This suggests that true martian surface conditions remains to be materials relatively enriched in isotopically light validated with future in situ as well as sample Li were available at near-surface levels, such as return missions. found for those nakhlites (Magna et al. 2015), which have shallower estimated depths of crys- tallization (Day et al. 2006). These shallow- 3.3.3 Lithium in Vesta seated clinopyroxenites carry a greater propor- tion of mesostasis with secondary minerals Analyses of howardites–eucrites–diogenites compared to their deeper counterparts having (HED meteorites) that most likely come from δ7Li signatures more closely resembling that of asteroid 4-Vesta are less common than those of bulk Mars. Deeper nakhlites have slightly higher lunar and martian rocks. With a diameter δ7Li values whereas both known dunites have of *530 km, Vesta is the second largest asteroid δ7Li within the range of mantle olivine from in the Solar System after Ceres. Its ancient age Earth. One potential means of producing this of *4565 Myr was indicated independently by relationship is settling of olivine and clinopy- extant as well as extinct chronometers such as roxene from a magma chamber, perhaps under Sm–Nd, Lu–Hf, Hf–W, and Mn–Cr equilibrium conditions (see Chap. 4 for (e.g., Blichert-Toft et al. 2002; Day et al. 2012; inter-mineral Li isotopic fractionation). Whether Kleine et al. 2004). The HED meteorites differ in or not clinopyroxenites and cumulate chemistry and petrography from other 3.3 Cosmochemistry of Lithium in Planetary Bodies 39 achondrites and it is apparent that they originated 2014a; Magna et al. 2006) appears to have been from a moderately small body that experienced modified by a currently unidentified process, as igneous processes, formation of a small metallic also supported by its aberrant oxygen isotopic core, and crust–mantle segregation, followed by composition (Wiechert et al. 2004). Whether or thermal metamorphism, the latter likely caused not this signature results from low temperature by large impacts. Basaltic eucrites that are post-magmatic processes is to be further inves- thought to originate in the shallow crust of Vesta tigated but it could be a hint to secondary pro- are significantly enriched in Li (gener- cesses acting on an asteroidal body. Second, ally >7 ppm and up to 15 ppm; Magna et al. Stannern and Bouvante, representative of possi- 2014a; Magna et al. 2006; Murty et al. 1983; ble mixing/contamination of main-trend eucrites Pogge von Strandmann et al. 2011; Seitz et al. with crustal melts enriched in incompatible ele- 2007; Tera et al. 1970), yet their Li isotopic ments (the so-called Stannern trend), have δ7Li signature is largely invariable with δ7Li values values that are irresolvable from other eucrites, reported between +2.2 and +4.9 ‰ (Magna et al. but appear to carry slightly elevated Li contents 2014a; Magna et al. 2006; Pogge von Strand- (Magna et al. 2014a; Magna et al. 2006; Pogge mann et al. 2011; Seitz et al. 2007), irrespective von Strandmann et al. 2011; Tera et al. 1970). of major element chemistry. This is contrary to heavy δ57Fe values measured Several observations can be made for eucrites for Stannern-trend eucrites (Wang et al. 2012) despite the rather limited Li isotopic range in which were probably caused by late-stage ilme- HED clan meteorites (Fig. 3.10). First, the iso- nite accumulation because these particular topically heavy eucrite Cachari (Magna et al. eucrites experienced re-melting of a former asteroidal crust with significant amounts of 7.0 ilmenite (Barrat et al. 2007). Third, the Li iso- topic signature of the eucrite Pasamonte (Magna 6.0 et al. 2006) is similar to other eucrites although 5.0 its distinctive oxygen isotopic composition

4.0 (Greenwood et al. 2005; Scott et al. 2009) may indicate its origin from a different parent body

Li (‰) 3.0 7 than the other eucrites (cf. discussion in Barrat δ 2.0 et al. 2011, on this issue). If correct then Li would provide evidence for multiple asteroidal 1.0 bodies in the early Solar System history that 0.0 experienced melting and magmatic evolution at 048121620 asteroidal/planetesimal levels leading to the Li (ppm) development of basaltic crusts with largely uni- Fig. 3.10 Lithium content versus δ7Li in meteorites from form Li contents and isotopic compositions. asteroid Vesta. A range in Li abundance for all major Fourth, lower Li abundances found in cumulate suites (howardites, eucrites, diogenites) is not paralleled eucrites (i.e. those which inherited a large portion by significant variation in δ7Li. This limited range of δ7Li values is uncommon to all other major planetary bodies of crystals from other sources, such as Serra de for which multiple data exist (Earth, Moon, Mars) and Magé and EET 87548) are not reflected in dis- may hint to short-term magmatism and limited tinctive Li isotopic composition. post-magmatic disturbance. The incompatible-element- Diogenites are igneous plutonic enriched Stannern trend eucrites do not plot significantly off the major Li trend found for other eucrites. Dashed orthopyroxene-rich deep-crustal cumulates (Bar- line connects bulk samples and orthopyroxene from the rat et al. 2010, 2008). Their Li systematics has Johnstown and Bilanga meteorites. La Paz Icefield remained largely unknown with the exception of (LAP) 03569 plots below the depicted range two values given in Seitz et al. (2007) that (δ7Li = −4.25 ‰). Data sources: Magna et al. (2006, 2014a), Seitz et al. 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fl 4.1 Introduction temperatures but requires a uid or melt that is far from equilibrium with the given mineral, as is typically the case for metasomatic processes or Laboratory experiments on the partitioning of open systems in general. The replacement of elements and isotopes between various phases minerals and other alteration features can be often (minerals, fluids, and melts) at thermodynamic identified by petrography and textural observa- equilibrium are fundamental to the interpretation tions. The occurrence of mechanism (ii) may also of geochemical data. In addition to the equilib- leave textural evidence. However, whether and to rium fractionation of stable isotopes, their com- what extent the chemistry of the relevant mineral position in a mineral depends on the partitioning is changed during dynamic recrystallization is a behavior of the element for simple mass balance topic that has not yet been studied in great detail. reasons. However, the signature of Li concen- It is commonly assumed that the presence of trations and isotopes stored in minerals also crit- fluids facilitates re-crystallization, which, as a ically depends on the mechanism of the exchange secondary process, could also affect trace element with the given environment (fluids, melts or and isotopic contents of the respective mineral. adjacent minerals). What we measure with our Mechanism (iii) does not affect the shape of the analytical tools is related to the closure of the mineral grain, but can only be identified by in situ mineral with respect to exchange of Li with the measurements of the spatial distribution of environment (Dodson 1973), which is strongly elements/isotopes. Zoning of Li preserved within sensitive to the active exchange mechanisms. the mineral may be generated by diffusion, but in After a mineral has formed by growth from a order to exclude other reasons for intra-grain fluid, a melt, or a gas, there are three end-member variations (e.g., growth zoning), in situ mea- scenarios that can be generally considered to surements and simulations of the isotopic zoning affect Li (or any other element) within the min- are necessary. The combined information on Li eral: (i) Recrystallization driven by a chemical concentration and isotopic abundance is actually gradient—mineral replacement by dissolution a very powerful tool to distinguish between dif- re-precipitation processes (e.g., Putnis and John ferent origins of chemical zoning, as illustrated by 2010); (ii) Recrystallization driven by interfacial the example of Mg and Fe isotopes (Oeser et al. energies or strain—normal grain growth/ 2015). deformation; (iii) Intra-crystalline diffusion dri- Unlike most other elements, Li appears to be a ven by a chemical or isotopic gradient. Mecha- special case since its mobility within mineral nism (i) provides element exchange even at low

P.B. Tomascak et al., Advances in Lithium Isotope Geochemistry, 47 Advances in Isotope Geochemistry, DOI: 10.1007/978-3-319-01430-2_4 © Springer International Publishing Switzerland 2016 48 4 Li Partitioning, Diffusion and Associated Isotopic … structures is very high, even affecting Li in the radius of Mg2+ at 0.89 and 0.57 Å, respectively. core of mm-sized crystals at moderate tempera- Therefore, Li tends to substitute for Mg in min- tures (e.g., plagioclase; Giletti and Shanahan erals but charge balancing requires a coupled 1997). Therefore, solid-state diffusion related to substitution with, for example, a trivalent cation secondary processes has a strong potential to for Mg. For example, in forsterite a coupled affect the Li chemistry of minerals and bulk substitution with Sc3+ was evidenced (Grant and rocks, unrelated to their formation. On the other Wood 2010) and in Fe-bearing olivine, Fe3+ hand, any zoning of Li preserved in minerals also could be of importance. Alternatively, Li+ can provides a way to determine the time scales of substitute for Na+ if the latter is present as a major the related geologic process by a diffusion model cation but due to the relatively large difference in (“diffusion chronometry”; Chakraborty 2008; the ionic radius (1.18 Å for octahedral coordina- John et al. 2012). tion; Shannon 1976) this substitution is very For these reasons, we compile and discuss in limited. Minerals that contain Li as a major ele- this chapter the measured partition coefficients of ment in nature are essentially limited to pegmatite Li and Li isotopes between various phases occurrences, like petalite, lepidolite, zinnwaldite (mineral–melt, mineral–fluid, and mineral–min- (a Li-rich trioctahedral mica), spodumene, eral), and the measured diffusion coefficients of amblygonite, lithiophilite, and lithiophosphate. Li in minerals, fluids, and melts. Furthermore, we Unlike the stable isotopes of most other elements, present and discuss consequences of diffusive in many minerals the light isotope (6Li) partitions fractionation of Li isotopes. Based on the infor- preferentially into the solid compared to the fluid mation compiled for partitioning and diffusion (e.g., Wunder et al. 2006; Kowalski and Jahn behavior of Li, we finally discuss experiments on 2011). This behavior is consistent with natural fluid–rock interaction with special emphasis on observations that clearly document a positive flux the behavior of Li. of δ7Li into the seawater because of alteration of Before going into the details of the experi- continental crust, oceanic crust and terrigenous mental work dedicated to Li we give a short sediments (see Chap. 6). overview of the general geochemical behavior of The diffusion behavior of Li, to the first order, Li to provide a framework for the detailed dis- is related to its ionic charge and size. The com- cussions. The alkali element Li is exclusively bination of a low charge (1+) and a relatively present as a monovalent cation in minerals (e.g., small ionic radius compared to other monovalent Wenger and Armbruster 1991) and melts, with cations is responsible for the relatively high chemical bonding dominated by undirected mobility of Li in solids and melts compared to electrostatic interactions (ionic bonding). In other cations (see for example Figs. 5, 6, and 7 in aqueous solutions Li forms a hydrated complex, Brady and Cherniak 2010). The large mobility of + [Li(H2O)n] , and coordinated approximately by 4 Li in minerals makes the preservation of primary oxygens, but the exact coordination changes with isotopic signatures difficult (as will be discussed the temperature and pressure (Jahn and Wunder below based on the experimental data). However, 2009; Kowalski and Jahn 2011). Lithium is this may support the potential for Li to be used as generally considered a fluid mobile element, a geo-chronometer for low- and medium- especially at hydrothermal conditions (e.g., Chan temperature processes or for very short-term et al. 1994; Brenan et al. 1998a; Seyfried et al. processes (e.g., Beck et al. 2006; Jeffcoate et al. 1991; see also Chap. 6), but when clay minerals 2007; John et al. 2012; Charlier et al. 2012). In like smectite are formed they can capture Li at these studies diffusion of Li in minerals has been lower temperatures (Sect. 4.2.4). The ionic radius identified by associated isotopic fractionation of Li+ for octahedral and tetrahedral coordination effects producing large anomalies for δ7Li of within crystals is 0.92 and 0.59 Å, respectively around ±10–20 ‰, which will be illustrated in (Shannon 1976), which is very close to the ionic Sect. 4.3. 4.2 Equilibrium Partitioning Behavior of Li 49

4.2 Equilibrium Partitioning 4.2.2 Lattice Strain Model Behavior of Li for Element Partitioning

4.2.1 Thermodynamic Background It has been shown by numerous studies that mineral–melt partition coefficients of trace ele- The Nernst partition coefficient is defined as the ments strongly correlate with their misfit to the ratio of the mass fractions of the element i (or respective lattice site in the mineral (e.g., Blundy any suitable respective chemical component) in and Wood 1991; Beattie 1994). This dependence α a β b on the ion size can be illustrated for example by phase , Ci , and phase , Ci : a “Onuma” diagrams (Onuma et al. 1968) where a=b C ¼ i ð : : Þ the logarithm of K ð Þ for an isovalent series of Kp ðiÞ b 4 2 1 p i Ci cations is plotted against their ionic radius. This shows a parabolic shape, where the position of Here we use K ð Þ instead of the more com- p i the maximum is characteristic of the size of the monly used symbol D because the latter is also i respective lattice site (Fig. 4.1). This systematic used for the diffusion coefficient. It reminds us of relationship can be quantified by the lattice strain the relation to the equilibrium constant, K, of the model (Blundy and Wood 1994), which is based corresponding partitioning reaction: on an expression of Brice (1975) for the iðbÞ¼iðaÞ; ð4:2:2Þ mechanical energy around a homovalent substi- tution in an elastically isotropic medium. which is given, in general, as follows: According to their model, the partition coefficient a a a of a misfit cation entering crystal lattice site M can a C c KðP; TÞ¼ i ¼ i i ; ð4:2:3Þ be approximated by the following expression: b b cb ai Ci i a=Liq ¼ n þ Á Kpðin þ Þ Kpð0Þ exp a b 8 9 rn þ 2 3 where a and a are the activities of the compo- > þ 0ðMÞ þ 1 þ > i i >À4p n þ À n þ þ À n > < EM NA 2 rin r0ðMÞ 3 rin r0ðMÞ = nent i in phase α and phase β, respectively, which ; > > depend in general on the chosen reference state :> RT ;> (fixed chemical composition of phase α and phase ð4:2:5Þ β at any P and T, see for example Ganguly 2008), a b fi c c þ as do the respective activity coef cients i and i . n where NA is Avogadro’s number, K ð0Þ is the For a trace element Henry’s law is typically p partition coefficient for an ion with charge n +, assumed, which means that the latter activity rn þ its respective ionic radius, and En þ the coefficients are constant but only for a fixed major 0ðMÞ M element composition. We can thus conclude from effective Young’s modulus of the site. The fun- Eq. 4.2.3 that for trace elements the Nernst par- damental basis of the model is that (i) the ionic tition coefficient is independent of their concen- radius of the fictive ion 0 is ideal in the sense that trations and in general is a function of P, T and it enters site M without inducing any elastic strain i 0 composition of the solvents, Xa and Xb: and (ii) element and can be replaced within the b melt without any resulting energy change. a=b c ðXbÞ K ¼ KðP; TÞ i : ð4:2:4Þ Although the additional assumption of an elasti- p ðiÞ cað Þ i Xa cally isotropic medium is not fulfilled in basically all relevant mineral sites, this model was suc- However, we will later demonstrate an exam- cessfully used to understand the systematics of ple where the partition coefficient of Li is also numerous mineral–melt partition coefficients dependent on the concentration of another trace (e.g., Blundy and Wood 2003a, b). The success of element (Grant and Wood 2010), which is related this model may be related to the fact that the to a coupled incorporation of two trace elements. anisotropic elastic behavior of the crystal site is 50 4 Li Partitioning, Diffusion and Associated Isotopic …

(a) (b)

n þ n þ n þ Fig. 4.1 Onuma diagrams of experimentally determined optimization of K ð0Þ, EM , and r0ð Þ. The effective ionic fi p M partition coef cients of a 1+, 2+, and 3+ cations between radii were taken from Shannon (1976) for eightfold clinopyroxene and melt from experiment DC22 of Blundy coordination. Note the general trend that for monovalent and Dalton (2000) and b 1+ cations between amphibole cations the site is more tolerant to misfit ions compared to (run# 1950, Adam and Green 2006), plagioclase (run# highly charged cations (2+ and 3+ here), reflected in 93-8a, Dohmen and Blundy 2014), clinopyroxene (same tighter parabola for the latter and hence larger effective as in a), orthopyroxene (run# 1949, Adam and Green Young’s moduli. The ideal cation radius (here the position Dashed lines solid 2006) and melt, respectively. in (a) and of the maximum of the parabola) varies for different lines in (b) are regressions of the data for each homova- crystallographic structures and sites lent series using the lattice strain model (Eq. 4.2.5) and

similar for ions of slightly different sizes such that m þ ¼ m þ þ Kpð0Þ Kex it can be integrated into a unique En . Note that 8 9 M > À12DG > fi > XM þ XM þ XM exp elec > Eq. 4.2.5 does not imply that partition coef cients < m þ ðmÀ1Þþ ðm þ 1Þþ RT = Á  of, for example, Li are insensitive to the melt > À22D > :> M M Gelec ;> composition (e.g., Blundy et al. 1995; Wood and þ Xð À2Þþ þ Xð þ 2Þþ exp þ ÁÁÁ m m RT Blundy 1997). The effect of melt composition is ð : : Þ n þ 4 2 6 integrated into Kpð0Þ, but for cations with the same charge, the respective dependence on melt com- m þ fi where Kex is a site- and mineral-speci c con- position is assumed to be the same in Eq. 4.2.5. M stant, and Xj þ is the fraction of M sites For limitations of this approximation and corre- charge-balanced by a j+ ion. This fraction sponding uncertainties see, for example, Schmidt therefore considers the contribution to the parti- et al. (2006) or Miller et al. (2006). m+ tion coefficient of an ion j by adding the var- The lattice strain model was extended to ious different combinations of local atomic heterovalent substitutions of an ion with charge configurations around the M site to m+ (Wood and Blundy 2001), by considering the charge-balance 1+, 2+, 3+,… cations. Heterova- electrostatic work required to replace the fictive lent substitutions with atomic configurations of ideal cation with charge n+byafictive ideal ion local charge balance are more stable but with charge m+ in the crystal and the melt. The depending on the magnitude of DGelec and the difference of this electrostatic work was defined probability of the different configurations (rep- D fi as Gelec and the partition coef cient for the M resented here by the magnitude of X þ ), other fictive ideal ion with charge m+, Km þ (compare j pð0Þ atomic configurations can become important with Eq. 4.2.5), is given by: 4.2 Equilibrium Partitioning Behavior of Li 51

(e.g., for REE partitioning between clinopyrox- Na ðaÞþLi ðLiquid) ¼ Li ðaÞþNa ðLiquid) ene and melt see Wood and Blundy 2001). ð4:2:7Þ For example, for Li partitioning into clinopy- roxene with the structural formula (M2M1) where, from the corresponding mass action law, [T O ], the following combinations could be 2 6 the ratio of the Li and Na partition coefficient is considered: (LiAl)[Si2O6], (LiTi)[AlSiO6], − − given by: (CaLi) [Si2O6], (LiAl) [AlSiO6], where the latter a= two have an effective negative charge. The Liq a Liq K ð Þ c ðXaÞ c ðX Þ probability of these various local configurations is p Li ¼ ð ; Þ Na Li Liq a= K4:2:7 P T a Liq cLiqðX Þ c ðXaÞ calculated by the assumption of random mixing KpðNaÞ Na Liq Li of the cations on the respective lattice sites, e.g. ð : : Þ ÀÁ2 4 2 8 ½ 1 þ Š¼2 M1 Á T þ 2 M1Á the molar fraction of X XAl XSi XTi T Á T ½ 2 þ Š¼ð1 À M1 À M1 À M1 Þ XSi XAl and X XAl XCr XFe3 þ According to the lattice strain model and ÀÁ2 ÀÁ2 ÀÁ2 Á T þ 2ð M1 þ M1 þ M1 ÞÁ T þ M1 Á T assuming that Li and Na occupy the same crys- XSi XAl XCr XFe3 þ XSi XTi XAl a=Liq= a=Liq (Wood and Blundy 2001). Because Mg occupies tallographic site, this ratio KpðLiÞ KpðNaÞ is also the smaller M1 site in clinopyroxene, replacement given, as follows: of Mg by Li would involve the formation of a − a=Liq charged species, (CaLi) [Si2O6], since no triva- Kp ðLiÞ lent cation present as a minor or major element a=Liq K ð Þ (Al3+ or Cr3+) would occupy the larger M2 site. p Na 8 9 r1 þ ÀÁÀÁ >À4p 1 þ 0ðMÞ 2 À 2 À 1 3 À 3 > An alternative candidate would be REE, but < EM NA 2 rNa rLi 3 rNa rLi = ¼ exp ; unless Li and REE are strongly associated as a > > :> RT ;> defect complex in the structure (e.g., as indicated for garnet in Cahalan et al. 2014), the probability ð4:2:9Þ of the species (REELi)[Si2O6] would be very 2 M a=Liq small because XREE is small. Therefore, although which provides a formula to calculate Kp ðLiÞ if the smaller M1 site should be more suitable for a=Liq Li, the larger probability of charge balance with Kp ðNaÞ and the lattice strain parameters for min- Al3+ on the M1 site, makes the M2 site more eral α are known from experiments. The right favorable for Li in clinopyroxene. This preference hand sides of Eqs. 4.2.8 and 4.2.9 should be the of Li for the M2 site is reflected in Fig. 4.1a where same if the lattice strain model is correct. Since Li fits on the Onuma parabola defined by the the lattice strain model ignores contributions of larger monovalent cations. The respective size of the melt chemistry on homovalent exchanges, we 1 þ ¼ 1:139 Å can immediately conclude that for the validity of the ideal cation, r0ðMÞ , is characteristic cLiqð Þ * cLiqð Þ for the larger M2 site in clinopyroxene. this model Li XLiq Na XLiq . Equation 4.2.9 The lattice strain model was also used to infer can be also rewritten in the following form: from mineral-melt partition coefficients of proxy 0 1 a= fi K Liq cations respective mineral-melt partition coef - ln@ p ðLiÞ A RT a=Liq cients of any other cation (e.g., Blundy and K ð Þ p Na ! Wood 2003a; Gallagher and Elliott 2009; Doh- 1 þ r0ð Þ ÀÁ1 ÀÁ men and Blundy 2014). For this purpose it is ¼À4pE1 þ N M r2 À r2 À r3 À r3 ; M A 2 Na Li 3 Na Li useful to consider homovalent exchange reac- tions, for example the exchange of Li with Na ð4:2:10Þ between mineral α and a liquid (melt or aqueous fluid): 52 4 Li Partitioning, Diffusion and Associated Isotopic …

60 example, the variability of plagioclase melt par- fi ) tition coef cients with anorthite content and l Ol

o

m 40 temperature can be explained by the decreasing

/

J

k size of the metal site and the corresponding

(

)

a

l 1 þ

N

/ ( 20

α p Opx decrease of r0ðMÞ with increasing anorthite com-

K

/

)

i

l

L

/

( ponent and decreasing temperature (Blundy and

α p 0 K Crd Cpx Wood 1994; Dohmen and Blundy 2014). An ln 1 þ increase of r0ð Þ with temperature could explain RT Amp M -20 Pl negative trends in Fig. 4.2 since increasing the site 600 1000 1400 1800 2200 would make it a better fit for the larger Na+ ion. Temperature (°C) To summarize, in the framework of the lattice Fig. 4.2 Experimentally observed values of RTlnðKpðÞLi Þ strain model, the mineral structure and detailed KpðÞNa versus temperature for selected minerals. Solid symbols mineral chemistry have two main impacts for mineral-melt partition coefficients, open symbols trace element partitioning: (i) the ideal ionic mineral-fluid partition coefficients. The scatter of data radius and the effective Young’s modulus (e.g., around the general trend for minerals like plagioclase or Fig. 4.1b), and (ii) depending on the major, minor clinopyroxene could be related to the respective wide variation in mineral chemistry (see also the detailed and even trace element chemistry, different local discussion following for each mineral) atomic configurations are possible to allow charge-balanced as well as charged incorporation To illustrate the validity of Eq. 4.2.10,we of a trace element. The melt chemistry also plays plotted the left hand side for various rock-forming a role in trace element partitioning, but under minerals (actually mineral groups, except cor- thermodynamic equilibrium, a given mineral fl dierite), as calculated from experimentally deter- chemistry is coupled to the melt/ uid chemistry. fl mined partition coefficients between minerals and Although in the general case the melt/ uid com- fi melts/fluids, against temperature (Fig. 4.2). position is not uniquely de ned for a given min- Although for some mineral groups the mineral eral composition, thermodynamic equilibrium fl composition varies strongly, each mineral group poses additional constraints on the melt/ uid defines a separate range in this plot. Furthermore, chemistry, in particular the chemical potential of apparent trends as a function of temperature can the chemical components. Thus the total degree fl be recognized irrespective of the liquid compo- of freedom for the melt/ uid composition is sition (fluid versus melt). This observation sup- reduced when a certain mineral composition is ports the prediction given by the lattice strain stable. For the latter reason, we discuss the fi model because the strain induced by the exchange experimental partition coef cients mostly in in fluid or melt should be very small compared to respect to their variation with mineral chemistry. that induced to the lattice. The numbers for RTlnðKpðÞLi Þ follow roughly the sequence KpðÞNa 4.2.3 Mineral-Melt Partitioning Ol > Opx > Cpx > Pl > Amp, which can be explained in terms of the lattice strain model by Min=Melt K was measured for the following min- the respective increase in the optimum ionic pðLiÞ radius of the respective site for monovalent erals or mineral groups (Table 4.1): amphibole cations (compare also with Fig. 4.1b). According (Amp), apatite (Ap), clinopyroxene (Cpx), cor- dierite (Crd), garnet (Grt), leucite (Lct), melilite to Eq. 4.2.10, RTlnðKpðÞLi Þ should be constant, if KpðÞNa (Mll), olivine (Ol), orthopyroxene (Opx), per- 1 þ 1 þ both r0ðMÞ and EM are constant, but both the ovskite (Prv), biotite (Bt), muscovite (Ms), pla- detailed mineral chemistry and the temperature gioclase (Pl), and spinel (Sp) (mineral have their own impact on these quantities. For abbreviations as recommended by Whitney and 4.2 Equilibrium Partitioning Behavior of Li 53

Table 4.1 Published Reference Mineral/mineral group P (MPa) T (°C) Kp(Li) partitioning coefficients of Li between minerals and Adam and Green (2006) Amp 1000 1025 0.09 melts Adam and Green (2006) Amp 2000 1050 0.098 Brenan et al. (1998a) Amp 1500 1000 0.19 Brenan et al. (1998b) Amp 1500 1000 0.13 LaTourette et al. (1995) Amp 1500 1092 0.124 Klemme and Dalpe (2003) Ap 1000 1250 0.039 Klemme and Dalpe (2003) Ap 1000 1250 0.084 Klemme and Dalpe (2003) Ap 1000 1250 0.027

Corgne et al. (2005) CaSiO3 Prv 21,000 20 0.13

Corgne et al. (2005) CaSiO3 Prv 20,000 2100 0.15

Corgne et al. (2005) CaSiO3 Prv 25,000 2300 0.05

Taura et al. (2001) CaSiO3 Prv 25,000 2350 0.17

Corgne et al. (2005) CaSiO3 Prv 25,000 2300 0.16

Corgne and Wood (2005) CaTiO3 Prv 0.1 1410 0.05

Corgne and Wood (2005) CaTiO3 Prv 0.1 1410 0.05

Corgne and Wood (2005) CaTiO3 Prv 0.1 1410 0.052

Corgne and Wood (2005) CaTiO3 Prv 3000 1800 –

Corgne and Wood (2005) CaTiO3 Prv 3000 1800 0.13

Corgne and Wood (2005) CaTiO3 Prv 0.1 1520 0.057

Corgne and Wood (2005) CaTiO3 Prv 3000 1750 –

Corgne and Wood (2005) CaTiO3 Prv 3000 1600 –

Corgne and Wood (2005) CaTiO3 Prv 3000 1600 0.05

Corgne and Wood (2005) CaTiO3 Prv 0.1 1600 0.06

Corgne and Wood (2005) CaTiO3 Prv 0.1 1600 0.11

Corgne and Wood (2005) CaTiO3 Prv 3000 1620 0.043

Corgne and Wood (2005) CaTiO3 Prv 3000 1620 0.17

Corgne and Wood (2005) CaTiO3 Prv 3000 1660 0.025

Corgne and Wood (2005) CaTiO3 Prv 3000 1660 0.3

Corgne and Wood (2005) CaTiO3 Prv 0.1 1470 0.05

Corgne and Wood (2005) CaTiO3 Prv 0.1 1470 0.057

Corgne and Wood (2005) CaTiO3 Prv 0.1 1400 0.045

Corgne and Wood (2005) CaTiO3 Prv 0.1 1400 0.069

Corgne and Wood (2005) CaTiO3 Prv 0.1 1400 0.023 Adam and Green (2006) Cpx 2500 1160 0.23 Adam and Green (2006) Cpx 2700 1160 0.26 Adam and Green (2006) Cpx 2000 1050 0.17 Adam and Green (2006) Cpx 2500 1180 0.27 Adam and Green (2006) Cpx 2000 1100 0.16 Adam and Green (2006) Cpx 2500 1100 0.21 Adam and Green (2006) Cpx 1000 1075 0.14 Adam and Green (2006) Cpx 3000 1170 0.35 (continued) 54 4 Li Partitioning, Diffusion and Associated Isotopic …

Table 4.1 (continued) Reference Mineral/mineral group P (MPa) T (°C) Kp(Li) Bennett et al. (2004) Cpx 3000 1330 1.052 Bennett et al. (2004) Cpx 3000 1298 0.0224 Bennett et al. (2004) Cpx 3000 1330 1.03 Blundy and Dalton (2000) Cpx 3000 1375 0.2 Blundy and Dalton (2000) Cpx 3000 1375 0.095 Blundy and Dalton (2000) Cpx 3000 1375 0.09 Blundy and Dalton (2000) Cpx 3000 1635 0.074 Blundy and Dalton (2000) Cpx 800 1375 0.043 Blundy and Dalton (2000) Cpx 3000 1640 0.081 Blundy et al. (1998) Cpx 1500 1255 0.254 Brenan et al. (1998a) Cpx 1000 1320 0.27 Brenan et al. (1998a, b) Cpx 0.1 1275 0.14 Brenan et al. (1998a, b) Cpx 0.1 1275 0.2 Elkins et al. (2008) Cpx 2500 1420 0.22 Gaetani et al. (2003) Cpx 1200 1315 0.34 Gaetani et al. (2003) Cpx 1200 1185 0.26 Hart and Dunn (1993) Cpx 3000 1380 0.59 Hill et al. (2000) Cpx 0.1 1218 0.027 Hill et al. (2000) Cpx 0.1 1218 0.027 Hill et al. (2000) Cpx 0.1 1218 0.032 Hill et al. (2000) Cpx 0.1 1218 0.207 Huang et al. (2006) Cpx 1100 1150 0.25 Huang et al. (2006) Cpx 1100 1155 0.16 Klemme et al. (2002) Cpx 3000 1400 0.75 Kuzyura et al. (2010) Cpx 7000 1265 0.52 Kuzyura et al. (2010) Cpx 7000 1265 0.43 McDade et al. (2003) Cpx 1500 1325 0.227 Schmidt et al. (1999) Cpx 1500 1040 0.173 Wood and Trigila (2001) Cpx 0.1 1140 0.19 Wood and Trigila (2001) Cpx 200 1070 0.13 Wood and Trigila (2001) Cpx 200 1042 0.15 Wood and Trigila (2001) Cpx 200 1042 0.15 Evensen and London (2003) Crd 200 850 0.12 Evensen and London (2003) Crd 200 800 0.13 Evensen and London (2003) Crd 200 700 0.4 Evensen and London (2003) Crd 200 750 0.34 Evensen and London (2003) Crd 200 800 0.21 Evensen and London (2003) Crd 200 700 0.44 Evensen and London (2003) Crd 200 800 0.26 Evensen and London (2003) Crd 200 700 0.33 Evensen and London (2003) Crd 200 750 0.3 (continued) 4.2 Equilibrium Partitioning Behavior of Li 55

Table 4.1 (continued) Reference Mineral/mineral group P (MPa) T (°C) Kp(Li) Evensen and London (2003) Crd 200 800 0.31 Adam and Green (2006) Grt 3500 1190 0.0410 Adam and Green (2006) Grt 2500 180 0.0350 Bennet et al. (2004) Grt 3000 1330 0.0610 Bennet et al. (2004) Grt 3000 1402 0.1400 Gaetani et al. (2003) Grt 1600 1230 0.0200 Klemme et al. (2002) Grt 3000 1400 0.1200 Kuzyura et al. (2010) Grt 7000 1265 0.1300 Kuzyura et al. (2010) Grt 7000 1265 0.0600 Van Westrenen et al. (1999) Grt 3000 1560 0.0180 Van Westrenen et al. (1999) Grt 3000 1565 0.0052 Van Westrenen et al. (1999) Grt 3000 1545 0.4800 Van Westrenen et al. (1999) Grt 3000 1530 0.0830 Van Westrenen et al. (1999) Grt 3000 1530 0.3800 Van Westrenen et al. (2000) Grt 2900 1540 0.0220 Van Westrenen et al. (2000) Grt 3000 1538 0.0084 Yurimoto and Ohtani (1992) Grt 20,000 2000 0.5500 Yurimoto and Ohtani (1992) Grt 16,000 1900 0.6800 Yurimoto and Ohtani (1992) Grt 16,000 1900 0.3800 Wood and Trigila (2001) Lct 200 1070 0.032 Lundstrom et al. (2006) Mll 0.1 1213 0.48 Lundstrom et al. (2006) Mll 0.1 1200 0.5 Lundstrom et al. (2006) Mll 0.1 1199 0.62

Corgne et al. (2005) MgSiO3 Prv 25,000 2300 0.02

Corgne et al. (2005) MgSiO3 Prv 25,000 2300 0.06

Corgne et al. (2005) MgSiO3 Prv 25,000 2300 0.17

Corgne et al. (2005) MgSiO3 Prv 25,000 2300 0.06

Corgne et al. (2005) MgSiO3 Prv 25,000 2300 0.16

Corgne et al. (2005) MgSiO3 Prv 25,000 2350 0.07

Liebske et al. (2005) MgSiO3 Prv 26,000 2300 0.055

Liebske et al. (2005) MgSiO3 Prv 26,000 2200 0.08

Liebske et al. (2005) MgSiO3 Prv 26,000 2300 0.04

Taura et al. (2001) MgSiO3 Prv 25,000 2300 1.2 Adam and Green (2006) Ol 2000 1100 0.29 Adam and Green (2006) Ol 1000 1075 0.42 Blundy and Dalton (2000) Ol 3000 1375 0.127 Blundy and Dalton (2000) Ol 3000 1375 0.126 Brenan et al. (1998a, b)Ol ––0.15 Brenan et al. (1998a, b) Ol 0.1 1349 0.15 Brenan et al. (1998a, b) Ol 0.1 1349 0.13 Brenan et al. (1998a, b) Ol 0.1 1349 0.15 (continued) 56 4 Li Partitioning, Diffusion and Associated Isotopic …

Table 4.1 (continued) Reference Mineral/mineral group P (MPa) T (°C) Kp(Li) Brenan et al. (1998a, b) Ol 1000 1320 0.35 Brenan et al. (1998a, b) Ol 0.1 1349 0.15 Grant and Wood (2010) Ol 0.1 1390 0.06 Grant and Wood (2010) Ol 0.1 1390 0.13 Grant and Wood (2010) Ol 0.1 1390 0.12 Grant and Wood (2010) Ol 0.1 1390 0.64 Grant and Wood (2010) Ol 0.1 1390 0.16 Grant and Wood (2010) Ol 0.1 1390 0.14 Grant and Wood (2010) Ol 0.1 1390 0.11 Grant and Wood (2010) Ol 0.1 1390 0.18 Grant and Wood (2010) Ol 0.1 1390 0.19 Grant and Wood (2010) Ol 0.1 1390 0.01 Grant and Wood (2010) Ol 0.1 1390 0.02 Grant and Wood (2010) Ol 0.1 1390 0.02 Grant and Wood (2010) Ol 0.1 1390 0.02 Grant and Wood (2010) Ol 0.1 1390 0.01 Grant and Wood (2010) Ol 0.1 1390 0.02 Grant and Wood (2010) Ol 0.1 1390 0.02 Grant and Wood (2010) Ol 0.1 1390 – Grant and Wood (2010) Ol 0.1 1390 – Grant and Wood (2010) Ol 0.1 1390 0.13 Grant and Wood (2010) Ol 0.1 1390 0.18 Grant and Wood (2010) Ol 0.1 1390 0.13 Grant and Wood (2010) Ol 0.1 1390 0.01 Grant and Wood (2010) Ol 0.1 1390 0.02 Grant and Wood (2010) Ol 0.1 1390 0.17 Grant and Wood (2010) Ol 0.1 1390 0.53 Grant and Wood (2010) Ol 0.1 1390 0.39 Grant and Wood (2010) Ol 0.1 1390 0.68 Grant and Wood (2010) Ol 0.1 1390 0.78 Grant and Wood (2010) Ol 0.1 1390 0.03 Grant and Wood (2010) Ol 0.1 1390 0.1 Grant and Wood (2010) Ol 0.1 1390 0.25 Grant and Wood (2010) Ol 0.1 1390 0.67 Grant and Wood (2010) Ol 0.1 1390 1.04 Grant and Wood (2010) Ol 0.1 1390 0.02 Grant and Wood (2010) Ol 0.1 1390 0.02 Grant and Wood (2010) Ol 0.1 1390 0.05 Grant and Wood (2010) Ol 0.1 1390 0.13 Grant and Wood (2010) Ol 0.1 1390 0.2 Grant and Wood (2010) Ol 0.1 1390 0.25 (continued) 4.2 Equilibrium Partitioning Behavior of Li 57

Table 4.1 (continued) Reference Mineral/mineral group P (MPa) T (°C) Kp(Li) Grant and Wood (2010) Ol 0.1 1390 0.36 Grant and Wood (2010) Ol 0.1 1390 0.5 Grant and Wood (2010) Ol 0.1 1390 0.82 Grant and Wood (2010) Ol 0.1 1390 0.13 Grant and Wood (2010) Ol 0.1 1390 0.24 Grant and Wood (2010) Ol 0.1 1390 0.19 Grant and Wood (2010) Ol 0.1 1390 0.15 Grant and Wood (2010) Ol 0.1 1390 0.1 Grant and Wood (2010) Ol 0.1 1390 0.23 Grant and Wood (2010) Ol 0.1 1390 0.07 Grant and Wood (2010) Ol 0.1 1390 0.08 Grant and Wood (2010) Ol 0.1 1390 0.08 Grant and Wood (2010) Ol 0.1 1390 0.09 Grant and Wood (2010) Ol 0.1 1390 0.11 Grant and Wood (2010) Ol 0.1 1390 0.07 Grant and Wood (2010) Ol 0.1 1390 0.1 McDade et al. 2003 Ol 1500 1325 0.289 Spandler and O’Neill (2010) Ol 0.1 1300 0.29 Taura et al. (1998) Ol 7000 1900 0.417 Taura et al. (1998) Ol 5000 1900 0.4 Taura et al. (1998) Ol 5000 1800 0.48 Taura et al. (1998) Ol 5000 1900 0.54 Taura et al. (1998) Ol 3000 1600 0.33 Taura et al. (1998) Ol 5000 1700 0.47 Taura et al. (1998) Ol 7000 1800 0.49 Taura et al. (1998) Ol 3000 1800 0.21 Taura et al. (1998) Ol 9700 2000 0.55 Taura et al. (1998) Ol 9700 1800 0.56 Taura et al. (1998) Ol 14,400 1900 0.45 Zanetti et al. (2004) Ol 1400 1055 0.22 Adam and Green (2006) Opx 2700 1160 0.28 Brenan et al. (1998a, b) Opx 1000 1320 0.2 Frei et al. (2009) Opx 1100 1230 0.185 Frei et al. (2009) Opx 2300 1440 0.2 Frei et al. (2009) Opx 2700 1510 0.175 Frei et al. (2009) Opx 3200 1535 0.184 Frei et al. (2009) Opx 2000 1330 0.2200 McDade et al. 2003 Opx 1500 1325 0.166 van Kan Parker et al. (2010) Opx 0.1 1390 0.22 van Kan Parker et al. (2010) Opx 0.1 1350 0.208 van Kan Parker et al. (2010) Opx 0.1 1326 0.13 (continued) 58 4 Li Partitioning, Diffusion and Associated Isotopic …

Table 4.1 (continued) Reference Mineral/mineral group P (MPa) T (°C) Kp(Li) van Kan Parker et al. (2011) Opx 1100 1430 0.21 van Kan Parker et al. (2011) Opx 2100 1550 0.22 van Kan Parker et al. (2011) Opx 2800 1600 0.26 Van Westrenen et al. (2000) Opx 3000 1538 0.48 Adam and Green (2006) Phl 2700 1160 0.21 Adam and Green (2006) Phl 2000 1050 0.19 Adam and Green (2006) Phl 2500 1100 0.21 LaTourette et al. (1995) Phl 2000 1109 0.064 Schmidt et al. (1999) Phl 1500 1140 0.213 Schmidt et al. (1999) Phl 1500 1040 0.46 Schmidt et al. (1999) Phl 1500 1140 0.169 Aigner-Torres et al. (2007) Pl 0.1 1200 0.2944 Aigner-Torres et al. (2007) Pl 0.1 1180 0.5754 Aigner-Torres et al. (2007) Pl 0.1 1180 0.6784 Bindeman and Davis (2000) Pl 0.1 1257 0.2880 Bindeman and Davis (2000) Pl 0.1 1257 0.2710 Bindeman and Davis (2000) Pl 0.1 1257 0.3050 Bindeman and Davis (2000) Pl 0.1 1297 0.2960 Bindeman and Davis (2000) Pl 0.1 1299 0.2250 Bindeman and Davis (2000) Pl 0.1 1299 0.2470 Bindeman and Davis (2000) Pl 0.1 1299 0.2520 Bindeman and Davis (2000) Pl 0.1 1153 0.2450 Bindeman and Davis (2000) Pl 0.1 1187 0.3170 Bindeman and Davis (2000) Pl 0.1 1187 0.2680 Bindeman and Davis (2000) Pl 0.1 1187 0.2540 Bindeman and Davis (2000) Pl 0.1 1187 0.2720 Bindeman et al. (1998) Pl 0.1 1190 0.2530 Bindeman et al. (1998) Pl 0.1 1190 0.2750 Bindeman et al. (1998) Pl 0.1 1297 0.2430 Bindeman et al. (1998) Pl 0.1 1297 0.2570 Bindeman et al. (1998) Pl 0.1 1250 0.2180 Bindeman et al. (1998) Pl 0.1 1250 0.2470 Bindeman et al. (1998) Pl 0.1 1153 0.2100 Bindeman et al. (1998) Pl 0.1 1153 0.2430 Bindeman et al. (1998) Pl 0.1 1153 0.2260 Bindeman et al. (1998) Pl 0.1 1297 0.1600 Bindeman et al. (1998) Pl 0.1 1297 0.1380 Blundy (1997) Pl 0.1 1208 0.130 Blundy (1997) Pl 0.1 1212 0.280 Blundy et al. (1998) Pl 1500 1255 0.1510 Caciagli et al. (2011) Pl 1000 1000 0.02 (continued) 4.2 Equilibrium Partitioning Behavior of Li 59

Table 4.1 (continued) Reference Mineral/mineral group P (MPa) T (°C) Kp(Li) Caciagli et al. (2011) Pl 1000 900 0.017 Dohmen and Blundy (2014) Pl 0.1 1251 0.3614 Dohmen and Blundy (2014) Pl 0.1 1086 0.1483 Dohmen and Blundy (2014) Pl 0.1 1251 0.1331 Dohmen and Blundy (2014) Pl 0.1 1251 0.1427 Dohmen and Blundy (2014) Pl 0.1 1251 0.0978 Dohmen and Blundy (2014) Pl 0.1 1251 0.1247 Dohmen and Blundy (2014) Pl 0.1 1290 0.1422 Dohmen and Blundy (2014) Pl 0.1 1290 0.1400 Dohmen and Blundy (2014) Pl 0.1 1290 0.1325 Dohmen and Blundy (2014) Pl 0.1 1290 0.1404 Dohmen and Blundy (2014) Pl 0.1 1226 0.1264 Dohmen and Blundy (2014) Pl 0.1 1184 0.1235 Dohmen and Blundy (2014) Pl 0.1 1184 0.1541 Dohmen and Blundy (2014) Pl 0.1 1108 0.1275 Dohmen and Blundy (2014) Pl 0.1 1108 0.1210 Dohmen and Blundy (2014) Pl 0.1 1282 0.0892 Dohmen and Blundy (2014) Pl 0.1 1286 0.1413 Tepley et al. (2010) Pl 0.1 1264 0.2970 Tepley et al. (2010) Pl 0.1 1265 0.1810 Elkins et al. (2008) Spl 2500 1420 0.13 Lundstrom et al. (2006) Spl 0.1 1213 0.1 Lundstrom et al. (2006) Spl 0.1 1225 0.184 Icenhower and London (1995) Bt 200 650 0.8 Icenhower and London (1995) Bt 200 750 1.01 Icenhower and London (1995) Bt 200 660 1.67 Icenhower and London (1995) Bt 200 650 1.64 Icenhower and London (1995) Ms 200 650 0.82

Min=Melt Min=Melt Evans 2010). KpðLiÞ was mostly measured for on their composition, KpðLiÞ varies over a silicate melts and in some cases for carbonate large range (0.01–1.0) and it is thus not appro- melts (e.g., Blundy and Dalton 2000). For the priate for geochemical modelling to assign a investigated range of experimental conditions Min=Melt single typical value of KpðLiÞ for a given Min=Melt and minerals, KpðLiÞ cover a range from 0.005 mineral. Simplified statements like ‘Li is more to *1.7. Therefore, during partial melting Li compatible in orthopyroxene than in clinopy- could behave as an incompatible, moderately roxene’ appear to be incorrect in the general case. incompatible, neutral, or even compatible ele- Min=Melt Most values for KpðLiÞ , however, consider ment depending on the mineral assemblage and only common rock-forming minerals of the crust detailed mineral chemistry. Even for the same and the upper mantle with respective typical mineral, e.g., clinopyroxene or garnet, depending compositions (excluding synthetic compositions 60 4 Li Partitioning, Diffusion and Associated Isotopic … such as pure forsterite), and are confined to a owing mainly to the small crystal sizes obtained. much smaller range (0.1–0.7). Lithium may be a special case compared to other After a short overview of the experimental elements, since it tends to diffuse very quickly Min=Melt within most minerals. At the relatively high methods, we discuss the data for KpðLiÞ of each temperatures before the final quench, Li should mineral individually based on crystal chemistry. re-equilibrate very quickly if any zoning was Here, we often refer to the lattice strain model to formed during the growth (see the characteristic discuss the variability of the experimental data diffusion time scales of Li discussed below), and with respect to the mineral chemistry and possible therefore reflect the equilibrium partitioning at the incorporation mechanisms of Li. final temperature. Experimental methods. In basically all stud- In most of the papers that are summarized ies measuring mineral–melt partition coefficients here, the authors focused on the discussion of a mineral/rock powder or synthetic glass powder partition coefficients for other alkali elements was heated in a chemically inert container (e.g., a (Rb, Cs) or trivalent cations, in particular REE. noble metal capsule, Au, Pt, etc.) above the Min=Melt solidus or even liquidus temperature at a given Studies that also discuss KpðLiÞ in more detail pressure for the given bulk composition and then are those by Brenan et al. (1998a, b), Taura et al. cooled slowly (typically with rates on the order of (1998), and Grant and Wood (2010). The data for a few °C/hour) to a final temperature below the Min=Melt KpðLiÞ are discussed separately for the differ- liquidus. In some experiments the starting mate- ent mineral groups; abbreviations for end mem- rial was directly heated to the final temperature. bers of mineral groups are italicized. After the sample reached the final temperature it = Plagioclase (Pl) KPl Melt Ab was immediately quenched or annealed for some . pðLiÞ along the albite ( , — An time (a couple of hours to days) to completely NaAlSi3O8) anorthite ( , CaAl2Si2O8) join equilibrate the sample, e.g., by diffusion and was measured for An contents between 9 and recrystallization. The slow cooling rate is largely 100 mol% (Blundy 1997; Blundy et al. 1998; chosen to ensure the growth of large crystals, at Bindeman et al. 1998; Bindeman and Davis least large enough to allow in situ measurement of 2000; Aigner-Torres et al. 2007; Tepley et al. the mineral composition. All of the Li partition 2010; Dohmen and Blundy 2014) at atmospheric coefficients reported here were measured using pressures (except Blundy 1997 at 1.5 GPa) and either SIMS or LA-ICP-MS; the major element temperatures between 1090 and 1300 °C. composition of the respective minerals and the Pl=Melt – Observed values for KpðLiÞ are between 0.09 quenched melt was typically measured using 0.6, which is a relatively small range compared electron microprobe (EPMA). It is a fundamental to, for example, garnet, olivine, and clinopyrox- issue of these types of experiments that the tem- ene (see below). Based on the work of Blundy perature of equilibration between the mineral and and Wood (1991) plagioclase–melt partition fi the melt is not well de ned since the mineral coefficients are presented by Bindeman et al. grows during cooling and growth zoning of trace (1998) in the following form: and major elements may result. This means that Pl=Melt ln ¼ þ þ þ Á ð : : Þ the core composition of the mineral was not RT Kpðin þ Þ Ain Bin XAn 4 2 11 necessarily in thermodynamic equilibrium with fi the melt at the nal temperature if the time at this where XAn is the molar fraction of anorthite in fi fi nal temperature was not suf cient to plagioclase, and ALi = –12.1 ± 1.0 kJ/mol and re-equilibrate the core with the melt by solid-state BLi = –6.9 ± 1.9 kJ/mol according to Bindeman diffusion. Rarely, compositional profiles across et al. (1998). However, the relation in Eq. 4.2.11 the minerals and quenched melts were measured does not fully cover the explicit dependence on in these experimental studies to identify any the An content but also reflects that the An con- chemical zoning (e.g., LaTourette et al. 1995), tent is an implicit function of T as discussed in 4.2 Equilibrium Partitioning Behavior of Li 61

Bindeman et al. (1998) and Dohmen and Blundy in general on the stable point group related to the (2014). Dohmen and Blundy (2014) fitted the Al–Si ordering state, C1 or I1, at the given T and An lattice strain parameters as a function of the XAn and the interaction parameters of the anor- content from experiments performed in the thite–albite solid solution, WAnÀAb. The optimized pseudo-ternary system diopside–albite–anorthite values for the diopside–albite–anorthite system at 1 þ ° ¼À4:71 ¼À18:4 and obtained a linear dependence of r0ðMÞ and 1200 C are BLi kJ/mol, ALi – ¼ 4:0 E1 þ on the anorthite content, X (see also and 16.56 kJ/mol, CLi and 0.0 kJ/mol for M An 1 1 fi Blundy and Wood 1994, 2003a, b), as follows: the C and the I stability eld, respectively. The parameters BLi and CLi reflect how the chemical 1 þ ½ ˚ Š¼ðÞÀð1:2370 Æ 0:0020 0:0171 potential of the Li-bearing component (LiAl- r0ðMÞ A Si3O8) in plagioclase changes with XAn. Recent Æ 0:0032ÞÁXAn experiments of Coogan (2011) in which plagio- ð4:2:12aÞ clase with variable XAn re-equilibrates Li with each other by solid state diffusion indicate a much 1 þ stronger dependence of this chemical potential on E ½Š¼GPa ðÞþð49:5 Æ 0:5 17:16 Æ 0:88Þ M X , equivalent to B À70 kJ/mol. More Á X An Li An experiments are necessary to resolve this dis- ð : : Þ 4 2 12b crepancy (see also Sect. 4.2.5). With the set of parameters given by Dohmen and Blundy (2014), With Eqs. 4.2.9 and 4.2.12a, b, respectively, Pl=Melt = the net effect of crystal chemistry for K is KPl Melt can be predicted for a given An content if pðLiÞ pðLiÞ relatively small and not drastically different from Pl=Melt KpðNaÞ is known from other sources (see the the empirical relationship of Bindeman et al. Pl=Melt (1998). possible strategies to calculate K ð Þ in Doh- p Na Olivine (Ol). With the structural formula men and Blundy 2014). A more general equation Pl=Melt M1M2[TO4], where both the M1 and M2 site are than Eq. 4.2.12a, b to calculate KpðLiÞ at sixfold coordinated, olivine usually has low trace atmospheric pressures was also given by Doh- element concentrations (in particular REE) men and Blundy (2014) who developed a pre- compared to pyroxene or garnet. The major ele- dictive thermodynamic model for 1+, 2+, and 3+ ment chemistry follows the binary join forsterite cations occupying the A site in plagioclase. For (Fo,Mg2SiO4)—fayalite (Fa,Fe2SiO4). Even the 1+ cations the model is based on the fusion though the mineral chemistry of olivine appears reaction of the Ab component combined with the Ol=Melt to be simple, a huge variability of K ð Þ was lattice strain model (Eq. 4.2.9), which allows p Li found, 0.01–1.0, which is similar to clinopyrox- estimation of the effect of T and compositional Ol=Melt\0:13 parameters according to the following equation: ene or garnet. KpðLiÞ was exclusively  found in the very detailed study of Grant and ln H Á Pl=Melt ¼ 29725 À 1:07928 Á Wood (2010) where they explored the effect of RT KpðLiÞ T Ol=Melt 2 trace element chemistry on KpðLiÞ . All of their À 0:01499 Á T þ ALi experiments were performed with almost pure þ Á þ Á ðÞ2 BLi XAn CLi XAn forsterite at atmospheric pressures and the same ð4:2:13Þ temperature of 1390 °C, thus isolating composi- = tional effects of trace elements on KOl Melt from where θ is a factor that considers the melt pðLiÞ chemistry, which can be assumed to be a constant other factors. Most importantly they found a ’ for a limited compositional range of the melt. The doping effect and deviation from Henry s law for a system with almost pure forsterite containing values for ALi, BLi,andCLi, respectively, depend 62 4 Li Partitioning, Diffusion and Associated Isotopic … defined doping levels of Sc or Ga (Fig. 2 in Grant available to be incorporated into the M site and Ol=Melt charge balance Li+ on the same site. Other factors and Wood 2010). Mostly, KpðLiÞ appears to like pressure and temperature as well as fayalite depend on the concentration of Sc3+ in the melt content seem to play a subordinate role for due to the formation of an associated defect = 3+ + Ol Melt containing Sc and Li at concentration levels KpðLiÞ . For example, Taura et al. (1998) above *500 μg/g Sc in the melt (the affinity of specifically investigated the effect of pressure on Li and Sc in the olivine structure is also reflected olivine–melt partition coefficients (experiments at by the synthesis of a pure LiScSiO4 endmember; pressures between 3 and 14 GPa) and found on Ito 1977). At lower concentrations it is argued Ol=Melt average a small increase of KpðLiÞ (factor 2 that two Li atoms substitute for one Mg atom between 3 and 14 GPa) with increasing pressure. where one Li atom goes to an interstitial site. The Clinopyroxene (Cpx). This has the general formation of Li interstitials in olivine is also structural formula (M1, M2) [T O ] where sev- indicated by diffusion experiments (Dohmen 2 6 eral combinations of occupations of the M sites et al. 2010) as discussed in detail below. Similar = KCpx Melt to Sc, incorporation of Ga into olivine is also and T sites are possible. pðLiÞ was measured coupled to the presence of Li, and association of at pressures between 0.1 MPa and 7 GPa, tem- Li and Ga at relatively high concentrations of peratures between 1040 and 1640 °C, and a great around 500 μg/g is suggested (Grant and Wood diversity of Cpx compositions as illustrated in 2010). This behavior highlights again the Fig. 4.3a, b. The strong variation of the Cpx importance of a charge balance on mineral–melt chemistry is closely related to the melt chemistry Ol=Melt in these experimental studies. For example, the partitioning and shows that K ð Þ tends to p Li SiO content varies between 4.8 and 26.2 wt% increase with the increasing concentration of 2 (Blundy and Dalton 2000; Kuzyura et al. 2010) trivalent cations in the melt that are capable of in carbonate melts and between 39 and 64 wt% association with Li within the olivine structure. (Adam and Green 2006; Huang et al. 2006)in In natural olivine Fe3+ is potentially important silicate melts. In accordance with compositional to charge balance Li+ in olivine. The concentra- 3+ variability of Cpx, experimentally measured tion of Fe in olivine is strongly dependent on Cpx=Melt values of K ð Þ cover a range from 0.02 to the fayalite content, fO2 and T, with concentra- p Li tions in the range of *0.1–10000 μg/g (roughly 1.05. For most experiments the compositions of estimated from point defect calculations in Doh- Cpx can be mainly described by the end mem- men and Chakraborty 2007). Other experimental bers Di+Hd, En+Fs, Jd, and CaTs (Fig. 4.3) and Ol=Melt even for these Cpx with minor contents of the data for K ð Þ (Brenan et al. 1998a, b; Taura p Li more exotic components Esk, Acm, Ca-Acm, and et al. 1998; Blundy and Dalton 2000; McDade CaTi-Ts (see Fig. 4.3 caption for definition of the et al. 2003; Zanetti et al. 2004; Adam and Green components and chemical formulas) the variation 2006; Spandler and O’Neill 2010) measured over = KCpx Melt a wide pressure and temperature range (0.1 MPa– of pðLiÞ is between 0.04 and 1.05. According 14.4 GPa and 1055–1900 °C, respectively) and to the model of Wood and Blundy (2001), with Fo contents ranging from 78 to 99.9 mol% Cpx=Melt KpðLiÞ should be strongly dependent on the vary much less than those of Grant and Wood availability of sites that allow local charge bal- – (2010), between 0.13 0.56. This smaller varia- ance. For example, when the Jd component is Ol=Melt tion of KpðLiÞ compared to that of Grant and present in Cpx it would provide a Wood (2010) seems surprising because of the charge-balanced site for other 1+ cations. We can very different experimental conditions in these further illustrate this feature of Li partitioning by other studies. This observation may also indicate Cpx=Melt Jd plotting KpðLiÞ against the component. For that in these experiments enough trivalent ele- most experiments we can observe a general trend ments (e.g., Fe3+,Sc3+,Ga3+) in the melt were 4.2 Equilibrium Partitioning Behavior of Li 63

(a) 0.00 1.0 (b) 1

0.8 0.25 X Adam & Green (2006) Bennet et al. (2004) En Hd + X Blundy & Dalton (2000)

l Blundy et al. (1998)

/ + X Brenan et al. (1998) Fs

cpx 0.1 Elkins et al. (2008) X 0.6 Di p(Li) Gaetani et al. (2003) K Hart & Dunn (1993) 0.50 Hill et al. (2000) Huang et al. (2006) Klemme et al. (2002) 0.4 Kuzyura et al. (2010) McDade et al. (2003) Schmidt et al. (1999) Wood & Trigila (2001) 0.75 0.01 0.0 0.2 0.4 0.6 0.0 0.1 0.2 0.3

XCaTs XJd

Fig. 4.3 Illustration of the compositional variation of Fe2Si2O6), Ca-Tschermak (CaTs, CaAl2SiO6), Cpx=Melt TiTs Esk Cpx for which K was measured and the corre- CaTi-Tschermak ( , CaTiAl2O6), Ca-eskolaite ( , pðLiÞ Acm 3+ Cpx=Melt Ca0.5[]0.5AlSiO), and acmite ( , NaFe Si2O6). For sponding variation of KpðLiÞ . a Ternary plot of the the Cpx of Wood and Trigila (2000) instead of Acm a 3+ 3+ molar fraction of the components jadeite + hedenbergite, CaFe -Tschermak component (CaFe AlSiO6) was enstatite + ferrosilite, and Ca-Tschermak component in needed to reproduce their respective cpx stoichiometry. = Cpx, normalízed to 1. b Variation of KCpx Melt with the For the calculation of the normative composition it was pðLiÞ 2+ 2+ molar fraction of jadeite in Cpx. The oxide compositions assumed that the fraction Fe /(Fe + Mg) is the same on of the respective cpx were recalculated to a normative the M1 and M2 site. The same type of symbols was used composition considering the 9 end-members: diopside in (a)asin(b) for the data sources, which are given in the respective symbol legend. A rough trend of increasing (Di, CaMgSi2O6), hedenbergite (Hd, CaFeSi2O6), jadeite Cpx=Melt X (Jd, NaAlSi2O6), enstatite (En,Mg2Si2O6), ferrosilite (Fs, KpðLiÞ with Jd can be observed

Cpx=Melt Jd Cpx=Melt that KpðLiÞ increases with the content. Hill et al. 2000), but for KpðLiÞ there is not a Exceptions are the experiments of Hill et al. similar obvious trend visible (cf. Fig. 4.3 with (2000) where the Cpx contains the highest data of Hill et al. 2000). In addition, deviations amount of CaTs (>0.25), one experiment of Cpx=Melt from the KpðLiÞ versus XJd trend seems to be Bennett et al. (2004) in which Cpx contains the unrelated to the melt composition, since for highest amount of the Esk (0.37), and experi- example data for carbonate melts and silicate ments of Klemme et al. (2002) and Kuzyura et al. melts follow the same trend. Blundy and Dalton (2010) in which Cpx contains a considerable (2000) also concluded in their comparative study amount of Esk, in the range of 5–31 mol%. of KCpx=Melt for silicate and carbonate melts in the It should be considered that a variable major p – – – element composition can also affect the optimum diopside albite and diopside albite dolomite ionic radius for a given site, which may explain system that for alkali elements the difference is = small (<3) and similar to the partitioning the deviation for the general KCpx Melt versus X pðLiÞ Jd behavior between immiscible silicate and car- trend for Cpx with high contents of components bonate melts (e.g., Veksler et al. 1998). Esk CaTs Jd like and . In addition to , as illustrated Garnet (Grt). This has the general structural here for Li, a number of studies demonstrated formula X3Y2Z3O12 where X, Y, and Z are Cpx=Melt that for many elements Kp increases with coordinated by 8, 6, and 4 oxygen atoms, an increasing amount of CaTs (or basically the respectively. The major element composition of amount of tetrahedrally coordinated Al; e.g., Grt can be mostly described by the four end-

Lundstrom et al. 1998; Blundy and Dalton 2000; members, pyrope (Prp,Mg3Al2Si3O12), 64 4 Li Partitioning, Diffusion and Associated Isotopic … almandine (Alm,Fe3Al2Si3O12), grossular (Grs, garnets with Maj contents of 56–59 mol%. It Ca3Al2Si3O12), and spessartine (Sps,Mn3Al2- appears that, in general, Li becomes more com- Si3O12), but at high pressures the majorite com- patible in Grt with increasing Grs or Maj content. ponent (Maj,Mg3MgSiSi3O12) becomes The effect of Maj on the incorporation of Li into Grt=Melt garnet was also studied in Hanrahan et al. (2009) relevant. KpðLiÞ was measured for the follow- and will be discussed in the Sect. 4.2.5 on ing garnet compositions: along the Prp–Grs join inter-mineral partitioning of Li. with Grs contents varying from 16 to 90 mol% Orthopyroxene (Opx). This can be described (van Westrenen et al. 1999; Bennett et al. 2004; by the same structural formula as Cpx, (M1, M2) Klemme et al. 2002), with some additional Alm Opx=Melt component varying between 9 and 33 mol% (van [T2O6]. KpðLiÞ was measured by Adam and Westrenen et al. 2000; Gaetani et al. 2003; Adam Green (2006), Brenan et al. (1998a, b), Frei et al. and Green 2006; Kuzyura et al. 2010), and (2009), McDade et al. (2003), van Kan Parker experiments at very high pressures between 16 et al. (2010, 2011), and van Westrenen et al. and 20 GPa with garnet containing a significant (2000) at pressures between 0.1 MPa and 3.2 proportion of Maj component with sixfold GPa, temperatures between 1160 and 1600 °C, coordinated Mg and Si on the Y site (Bennett and diverse major element compositions. The et al. 2004; Kuzyura et al. 2010; Yurimoto and Opx stoichiometry can be mainly expressed by Grt=Melt the following endmember components: En, Fs, Ohtani 1992). The overall variation of KpðLiÞ is Di, Hd and Mg-Tschermak (MgTs, MgAl SiO ). between 0.005–0.68. Even for experiments along 2 6 The compositional variables X , X , and the Prp–Grs join at the same pressure (3 GPa) Di MgTs Mg# ¼ Mg=ðFe þ MgÞ, vary between 0.03– and a moderate temperature range (1400–1550 ° = 0.23, 0–0.2, and 0.0–0.17, respectively. Despite C), KGrt l varies between 0.005 and 0.48, indi- pðLiÞ this variable composition and the very different = cating a strong compositional dependence of P–T KOpx Melt = conditions, pðLiÞ is remarkably similar KGrt Melt. The crystal chemistry of garnet domi- pðLiÞ for the 15 published experiments, varying nates the partitioning behaviour of many trace between 0.13 and 0.48 with an average value of elements. For example, the partition coefficients Opx=Melt 0.22 ± 0.08. There is no clear trend of K ð Þ of trivalent cations can be reproduced by a sys- p Li with any of the compositional variables, but the tematic variation of the lattice strain parameters highest value of 0.48 was measured by Van (r3 þ , E3 þ , and K3 þ ) with the content of the 0ðXÞ X pð0Þ Westrenen et al. (2000) for Opx with the highest four end-members, Prp, Alm, Grs, and Sps (van XDi (0.23) and XMgTs (0.08). Due to the high MgTs Westrenen et al. 1999, 2000). Unfortunately, content in this Opx composition, more charge there are no data available for an isovalent ele- balanced sites could be provided for Li in the ment series of 1+ cations to perform a similar form of LiAlSi2O6 thereby enhancing the com- calculation and quantify the effect of composition patibility of Li in Opx. Grt=Melt Maj for KpðLiÞ . The presence of the component Perovskite (Prv). This has the structural for- could have a twofold effect: (i) a change of the mula ABO3 where A and B are 12-fold and six- lattice strain parameters (e.g., as calibrated for Prv=Melt fold coordinated sites, respectively. KpðLiÞ was Prp, Alm, Grs, Sps); (ii) due to the presence of Si measured for three different perovskites: CaSiO3 on the Y site, the number of local atomic con- (CaPrv) at P between 21–25 GPa and T between fi 1+ gurations to charge balance Li on the X-site is 2100–2350 °C (Taura et al. 2001; Corgne et al. increasing by the incorporation of Li as (Mg2Li) 2005); CaTiO3 (CaTiPrv) for P between 0.1 MPa (AlSi)Si3O12. The latter argument seems to be and 3 GPa, T between 1500–1800 °C (Corgne Grt=Melt consistent with the high values for KpðLiÞ and Wood 2005), and MgSiO3 (MgPrv) for (0.44–0.68) of Yurimoto and Ohtani (1992) for P between 25–26 GPa and T = 2300–2350 °C 4.2 Equilibrium Partitioning Behavior of Li 65

(Taura et al. 2001; Corgne et al. 2005; Liebske content of around 1 wt% in the perovskites of MgPrv=Melt Taura et al. (2001), and we can conclude that at et al. 2005). KpðLiÞ measured by Corgne identical conditions Li is more compatible in et al. (2005) and Liebske et al. (2005) are very CaTiPrv=Melt similar and within the range 0.05–0.08 (6 exper- MgPrv compared to CaPrv. KpðLiÞ mea- iments, where for Corgne et al. 2005 we consider sured by Corgne and Wood (2005) was found to only three values measured by SIMS; in their CaPrv=Melt be very similar to KpðLiÞ and varies for 20 Prv=Melt study KpðLiÞ measured by LA-ICP-MS was experiments from 0.025 to 0.3. This variation is systematically biased to higher values compared not correlated with the strongly different P and to SIMS data for unknown reasons). Interestingly, T conditions of these experiments (0.1 MPa, c. CaPrv=Melt 1400–1600 °C vs. 3 GPa, 1600–1800 °C). For the two values of KpðLiÞ measured by Corgne trivalent and tetravalent cations, Corgne and et al. (2005) are 0.05 and 0.13, very similar to CaTiPrv=Melt MgPrv. This seems surprising as the Mg site in Wood (2005) demonstrated that for Kp MgPrv (at the center of a distorted dodecahedron) the most relevant variable is the CaO content in should be apparently a better fit for Li than the Ca the melt. This dependence is the result of the site in CaPrv (at the center of a regular dodeca- substitution of these elements into the A site by a hedron), when inferred from the ionic radii of Mg coupled substitution that involves vacancies on and Ca (see also Chap. 5). In fact, fitting of the the A site. For Li this kind of mechanism is not partition coefficients of 1+ cations for MgPrv and possible but, as mentioned above, Al on the A CaPrv by the lattice strain model show that the site and by inference Al2O3 content in the melt optimum ionic radius for the A site is much could potentially be a more relevant variable. smaller in MgPrv compared to CaPrv, c. 1.1 Å Mica. These can be generally described by the versus c. 1.3 Å, respectively (Corgne et al. 2005). structural formula X2Y4–6[Z8O20](OH, F)4. fi ’ 1 þ Mica–melt partition experiments of Li were The tted effective Young s modulus, EA is 1 þ measured for biotite (Bt) and muscovite (Ms) by rather similar for the two, but K is significantly pð0Þ Icenhower and London (1995), and for phlogo- larger for CaPrv compared to MgPrv, 0.07 versus pite (Phl) by LaTourette et al. (1995), Schmidt Prv=Melt 0.9, thus compensating as a net result for KpðLiÞ et al. (1999) and Adam and Green (2006). For the misfit of Li for CaPrv compared to MgPrv. Phl, Bt, and Ms the X site is mainly occupied by 1 þ K and minor amounts of Na. The Y site is This higher value of Kpð0Þ could be related to the melt chemistry stabilizing CaPrv or the higher occupied mainly by Al in Ms but Fe, Mg, Al, and availability of sites that charge balance 1+ Ti in Bt and Phl. The main difference between Bt and Phl is the higher degree of the Tschermak cations. For example, the presence of Al on the A 3+ 2+ 4+ site could be relevant to form structural units like substitution of 2 Al by Mg +Si ions in Bt. Phl=Melt P – Li0.5Al0.5CaO3, but in the study of Corgne et al. KpðLiÞ was measured at = 1.5 7 GPa, (2005) the Al content in CaPrv is not generally T = 1040–1150 °C, and varies between 0.06 and larger than in MgPrv. 0.46, which is negatively correlated with the Al In Taura et al. (2001), a single experiment was content of the Z site. This correlation can be published where MgPrv and CaPrv coexisted and explained by considering the coupled substitu- MgPrv=Melt ¼ 1:2 CaPrv=Melt ¼ 0:17 tion of 2 Al3+ with Si4+ + 2Li+ ions, similar to KpðLiÞ and KpðLiÞ were measured simultaneously, and thus effects of P, derivation of the Li-mineral lepidolite (K2(Li, T, and melt chemistry could be excluded to Al)5–6[Si6–7Al2–1O20](OH, F) from pure Ms Bt=Melt explain this difference. This significantly larger K2Al4[Si6Al2O20](OH, F). KpðLiÞ was mea- MgPrv=Melt sured at P = 0.2 GPa, T = 650–750 °C, and varies value of KpðLiÞ compared to those given by Corgne et al. (2005) and Liebske et al. (2005) between 0.8 and 1.64; hence Li is generally could be explained by the relatively large Al much more compatible in Bt than in Phl. In Bt 66 4 Li Partitioning, Diffusion and Associated Isotopic … the Al content on the Y site is much higher, thus Fe2+ (as well as Mn and Li) the B site, generally allowing the potential substitution of 2Mg2+ ions smaller cations than Ca, like Fe2+,Mg2+,Fe3+, by coupled Al3+ +Li+, all on the Y site. and Al3+ (as well as Mn and Li) the C site, Si and Ms=Melt ¼ 0:8 Amp=Melt KpðLiÞ was measured in a single experi- Al the T site. KpðLiÞ was measured mostly for ment at P = 0.2 GPa, T = 650 °C. Onuma dia- Na–Ca amphiboles at P =1–2 GPa, T = 1000– Phl=Melt ° – grams of Kp for the alkali elements show 1100 C, and varies a little (0.09 0.19; Adam = and Green 2006; Brenan et al. 1998a, b; that KPhl Melt does not fit into the parabola for pðLiÞ LaTourette et al. 1995). Naturally occurring larger 1+ cations like Na, K, Rb, and Cs (Fig. 6 compositions of Amp are broadly variable and in Schmidt et al. 1999; Fig. 8 in LaTourette et al. hence partition coefficients may vary strongly 1995), which occupy the X site. The optimum between different amphiboles. The occupation of Å ionic radius for this site (c. 1.7 ) is much larger various cations on the different sites in Amp is than the ionic radius of Li and hence it is strongly well illustrated by Onuma diagrams (e.g., Fig. 7 incompatible for the X site. However, as dis- in Brenan et al. 1998a, b; Fig. 3 in Adam and cussed before, Li can be substituted into the Green 2006). Similar to micas, alkali elements much smaller Y site, and therefore, for micas prefer the large A site in amphiboles, but, Min=Melt KpðNaÞ together with Eq. 4.2.10 would be not a apparently, Li tends to substitute into the smaller Min=Melt B site. This may be the reason for the poorer fit good proxy to predict K ð Þ . p Li of the monovalent series in Fig. 4.1b compared to Cordierite (Crd). It has the structural formula Cpx or Pl. More details related to the crystal (Mg, Fe) [Si Al O ]·nH O. Evensen and Lon- 2 5 4 18 2 chemistry of Li-bearing amphiboles can be found Crd=Melt P don (2003) measured KpðLiÞ at = 200 MPa in Hawthorne et al. (1992, 1993, 1994, 1996a, T – ° Crd=Melt 1996b), Oberti et al. (2003), and Ottolini et al. and = 700 850 C. Values for KpðLiÞ are in – X (2009). the range of 0.13 0.44 where Mg of the Crd ApÀMelt varies between 0.69 and 0.85. Evensen and Apatite (Ap). In three experiments KpðLiÞ Crd=Melt was in the range of 0.027–0.084 (LA-ICP-MS), London (2003) parameterized K ð Þ as a p Li measured by Klemme and Dalpe (2003) for function of temperature (in °C), as follows, carbonatite melts at 1250 °C and 1 GPa. Crd=Melt ¼ 1:61À1:733 Â 10À3 Á KpðLiÞ T. Leucite (Lct). In a single experiment per- Several studies suggest that Li enters the cor- formed by Wood and Trigila (2001) at 200 MPa dierite structure by substitution for octahedral and 1070 °C, leucite was formed and cations like Mg and Fe2+ (Armbruster and Irou- LctÀMelt KpðLiÞ = 0.032 was obtained. schek 1983; Kirchner et al. 1984; Gordillo et al. Melilite (Mll) and Spinel (Spl). A handful of + 1985). To charge balance Li the following sub- mineral–melt partitioning experiments were per- stitution was suggested by Evensen and London formed by Lundstrom et al. (2006), motivated by 2+ → + + + (2003): Mg Li +Na , where Na occupies studying the trace element distribution in Ca– non-regular sites within the channels formed by Al-rich inclusions (CAIs) in chondrites. They six-membered rings of (Si, Al)O4 tetrahedra in measured melilite–melt and spinel–melt partition Crd, like other large alkali ions. Therefore, similar coefficients at atmospheric pressures and the final Min=Melt T ° to mica, KpðNaÞ cannot be used as a proxy for between 1200 and 1225 C. Melilite consisted Min=Melt of a solid solution between åkermanite (Ak, K ð Þ together with Eq. 4.2.10 because Na and p Li Ca MgSi O ) and gehlenite (Gh,CaAl SiO ) Li do not enter the same site. 2 2 7 2 2 7 with Ak contents of 38–44 mol% and reported Amphibole (Amp). These have the general Mll=Melt K ð Þ are 0.48–0.62. Two measurements of structural formula A0–1B2C5[T8O22](OH, F)2, p Li where Na and K occupy the A site, Na, Ca, Mg, Spl=Melt KpðLiÞ for pure MgAl2O4 spinel were 4.2 Equilibrium Partitioning Behavior of Li 67 performed by Lundstrom et al. (2006), 0.1 and only depends on the starting materials but also 0.184, respectively, and in one experiment of (mainly) on the temperature as a fundamental

Elkins et al. (2008) for (Fe, Mg)Al2O4 spinel parameter to control the efficiency of the different Spl=Melt ¼ 0:1 ° kinetic processes. KpðLiÞ at 1420 C and 2.5 GPa Aside from the problem of reaching equilib- (Table 4.1). rium an additional complication is the analysis of the fluid composition. In several studies the fl 4.2.4 Mineral-Fluid Partitioning composition of the uid was not directly mea- and Associated Stable sured after annealing but inferred from mass Isotopic Fractionation balance using the modal abundance and chem- istry of the minerals produced and the initial fluid composition (e.g., Brenan et al. 1998b; Martin Experimental methods. In Table 4.2 the et al. 2011; Caciagli et al. 2011; Decarreau et al. experimental studies where Li partitioning 2012). With a large water/mineral ratio, the fluid between a mineral and an aqueous fluid, Min=Fluid 7 7 7 composition for incompatible elements typically K , and Δ LiMin/Fluid = δ LiMin − δ LiFluid pðLiÞ underwent no significant change. Lithium iso- were measured are listed, including details on the topic data were mostly collected from bulk experimental conditions as well as the analytical analysis of the product minerals using various and experimental method used. These experi- chemical procedures (Teng et al. 2004; Wunder ments can be broadly categorized based on the et al. 2006, 2007, Vigier et al. 2008). A funda- starting material used: (i) gels or mixtures of mental issue of experiments with fluids is that solid phases (oxides, hydroxides, etc.) of which some solid material may precipitate from the the stable phase, the mineral of interest, is sup- fluid during the quenching of the sample. The posed to form directly at the given experimental quench products are difficult to identify and condition; (ii) mineral of interest and a reactive separate completely from the mineral of interest, fluid that are supposed to form the stable mineral and therefore could affect the measured partition composition by complete replacement of the coefficient. Furthermore, the composition of initial crystals; (iii) mineral of interest and a less remaining fluid is not representative of the P, reactive fluid, where trace element equilibrium is T conditions of the experiment. To overcome supposed to be reached by exchange with the these issues, a more sophisticated approach was fluid and solid-state diffusion. Ideally, to ensure applied by Kessel et al. (2005) to measure the that thermodynamic equilibrium has been partition coefficients of trace elements between attained reversal experiments are required. In a clinopyroxene, garnet, aqueous fluid, hydrous reversal experiment the equilibrium distribution melt and supercritical liquid. They performed the of the two Li isotopes is achieved from opposite partitioning experiments in a ‘rocking’ directions in terms of the starting compositions of multi-anvil apparatus to homogenize the liquids mineral and fluid. In only a very few of the (Schmidt and Ulmer 2004). To measure the liq- experimental studies presented below a reversal uid composition after the experiment the capsule approach was applied (e.g., a part of the experi- and hence the liquid was frozen and directly ments in Caciagli et al. 2011). Attainment of analyzed in a special setup for LA-ICP-MS (the thermodynamic equilibrium becomes usually diamond-trap method; Kessel et al. 2004). more problematic at lower temperatures. The Min=Fluid Experimental data for K ð Þ were mea- type of starting material (i), (ii) or (iii) can be also p Li sured for amphibole, carbonate, epidote (zoisite), considered as three end-member mechanisms to olivine, pyroxene, quartz, serpentine and stauro- reach equilibrium since in any real experiment a Min=Fluid combination of these three end-member mecha- lite. For these minerals, KpðLiÞ covers a large nisms occurs but to different degrees, which not – Min=Fluid range of 0.0004 152. High KpðLiÞ > 1 were Table 4.2 Published partitioning coefficients of Li between minerals and fluids including experimental and analytical Isotopic Associated and Diffusion Partitioning, Li methods 4 68

Reference Mineral/mineral # data Kp(Li) Δ7Li P T(°C) w/r Exp. setup Solid Fluid Experiment group points (MPa) type (see Starting material Anal. method Type of Starting material Anal. method text) analysis

Brenan et al. Cpx 4 0.09–0.25 – 2000 900 50 PC Natural SIMS In-situ Water or one MB (ii) (1998a, b) diospide + silica + albite experiment with 0.5 M NaCl

Brenan et al. Grt 1 0.0083 – 2000 900 PC Pyrope almandine garnet SIMS In-situ Pure H2O MB (ii) (1998a, b)

Caciagli Cpx 7 0.07–0.25 – 1000, 800–1100 4–6 PC, DD + SiO2 ± SCO SIMS In-situ H2O+Li2CO3 MB (ii) et al. (2011) 200 CSHV

Caciagli Cpx 6 – 0to 1000 900–1000 4–6 PC DD + SiO2 ± SCO MC-ICP-MS Bulk H2O+Li2CO3 MB (ii) rev. et al. (2011) −3

Caciagli Ol 6 0.17–1.34 – 1000 800–1100 4–6 PC SCO + SiO2 ± SCO ± Al2O3 SIMS In-situ H2O+Li2CO3 MB (ii) et al. (2011)

Caciagli Pl 9 0.09–0.3 – 1000 800–1100 4–6 PC CBB ± SiO2 ± Al2O3 ± Ab SIMS In-situ H2O+Li2CO3 MB (ii) et al. (2011)

Caciagli Opx 1 0.02 – 1000 1000 4–6 PC DD + SCO SIMS In-situ H2O+Li2CO3 MB (ii) et al. (2011)

Decarreau Sme 18 0.67–22 – sat. 75–150 150 TB, C Si4Mg3O11;nH2O AAS Bulk H2O AAS (ii) et al. (2012) (gel) + LiCl Marriott Cal 7 – −7.6 atm 5–30 40 TP Cal seed MC-ICP-MS Bulk Syn. Seawater NA (i) et al. (2004a) to −10.1 Marriott Cal 5 – −1.9 atm 25 100 TP Cal seed MC-ICP-MS Bulk Syn. Seawater NA (i) et al. to (2004b) −3.1 Marriott Arg 5 – −10.9 atm 25 100 TP Cal seed MC-ICP-MS Bulk Syn. Seawater NA (i) et al. to −12 (2004b) Williams Sme-Ilt 9 0.1 −0.2 100 300 1 CSHV SWy-1 sme SIMS Bulk DDI + LiOH MC-ICP-MS (ii) and Hervig to (2005) −11.6

Kessel et al. Cpx 2 2 – 6 800–1000 0.2 MA Basalt powder LA-ICP-MS In-situ H2O LA-ICP-MS (ii) (2005)

Kessel et al. Grt 7 0.13–0.52 – 4,6 700–1200 0.2 MA Basalt powder LA-ICP-MS In-situ H2O LA-ICP-MS (ii) (2005)

Berger et al. Chl 12 0.1–2 – sat. 50–260 2– PPV Silicate glass Flame Bulk H2O Flame (ii) (1988) 50 spectrometry spectrometry (continued) … Table 4.2 (continued) 69 Li of Behavior Partitioning Equilibrium 4.2

Reference Mineral/mineral # data Kp(Li) Δ7Li P T(°C) w/r Exp. setup Solid Fluid Experiment group points (MPa) type (see Starting material Anal. method Type of Starting material Anal. method text) analysis

Berger et al. Sme 14 1.6–15 – sat. 50–260 2– PPV Silicate glass Flame Bulk H2O Flame (ii) (1988) 50 spectrometry spectrometry

Berger et al. Zeo 8 3–33 – sat. 50–260 2– PPV Silicate glass Flame Bulk H2O Flame (ii) (1988) 50 spectrometry spectrometry

Fabbrizio Ol 5 0.07–0.35 – 2000 900–1300 0.2 PC Peridotite/MSH LA-ICP-MS In-situ H2O LA-ICP-MS (ii) et al. (2013)

Fabbrizio Opx 3 0.18–0.48 – 2000 900–1100 0.2 PC Peridotite/MSH LA-ICP-MS In-situ H2O LA-ICP-MS (ii) et al. (2013)

Fabbrizio Cpx 2 0.25–0.72 – 2000 900–1000 0.2 PC Peridotite/MSH LA-ICP-MS In-situ H2O LA-ICP-MS (ii) et al. (2013)

Fabbrizio Amp 1 0.82 – 2000 900 0.2 PC Peridotite/MSH LA-ICP-MS In-situ H2O LA-ICP-MS (ii) et al. (2013)

Lynton et al. Qtz 3 – +8 to 50– 500 0.5 CSHV Nat. Msc, Qtz, Kfs, Phl TIMS Bulk H2O + 1MCl TIMS (iii) (2005) +12 100

Lynton et al. Msc 3 – +18 to 50– 500 0.5 CSHV Nat. Msc, Qtz, Kfs, Phl TIMS Bulk H2O + 1MCl TIMS (iii) (2005) +20 100 Martin et al. Zo 2 0.003, – 3000 700–850 0.5 PC Lws-gel, Qtz LA-ICP-MS In-situ MilliQ MB (ii) (2011) 0.004 Martin et al. Lws 2 0.0004, – 3000, 650–700 0.5 PC Lws-gel, Qtz LA-ICP-MS In-situ MilliQ MB (ii) (2011) 0.01 3500

Vigier et al. Sme 14 – −1.6 sat. 25–250 60 TB Si4Mg3O11;nH2O AAS, Li Seawater, DDI AAS, (ii) (2008) to (gel) + LiCl QUAD-ICP-MS, separates QUAD-ICP-MS, −10.9 MC-ICPMS MC-ICPMS

Wunder Atg 1 65 −1.1 4000 500 0.1 PC Syn. powder MC-ICP-MS Bulk H2O MC-ICP-MS (ii) et al. (2010)

Wunder Ctl 2 16, 34 +0.6, 400 400 0.1 CSHV Syn. powder MC-ICP-MS Bulk H2O MC-ICP-MS (ii) et al. (2010) +0.9

Wunder Lz 2 83, 152 −2.8, 200, 200–300 0.1 PC Syn. powder MC-ICP-MS Bulk H2O MC-ICP-MS (ii) et al. (2010) −4.1 400

Wunder Spd 12 – −0.7 2000 500–900 0.1– PC Syn. powder MC-ICP-MS Bulk H2O + LiCl, LiOH MC-ICP-MS (ii) et al. (2006) to 0.2 −3.8

Wunder St 6 – +1.1 3500 670–880 0.2 PC Syn. powder MC-ICP-MS Bulk H2O + LiCl, LiOH MC-ICP-MS (ii) et al. (2007) to +1.6 (continued) Table 4.2 (continued) Isotopic Associated and Diffusion Partitioning, Li 4 70

Reference Mineral/mineral # data Kp(Li) Δ7Li P T(°C) w/r Exp. setup Solid Fluid Experiment group points (MPa) type (see Starting material Anal. method Type of Starting material Anal. method text) analysis

Wunder Lpd 16 – −1.5 2000 350–500 0.1 PC Syn. powder MC-ICP-MS Bulk H2O + LiCl, LiOH MC-ICP-MS (ii) et al. (2007) to −5.3

Wunder Spd 3 – −1.9 1000– 500–625 0.2 PC Syn. powder MC-ICP-MS Bulk H2O + LiCl, LiOH MC-ICP-MS (ii) et al. (2011) to 8000 −3.5

Wunder Amp 1 – −1.7 2000 700 0.2 PC Syn. powder MC-ICP-MS Bulk H2O + LiCl, LiOH MC-ICP-MS (ii) et al. (2011) CBB Crystal Bay bytownite, DD Dekalb diopside, DDI Distilled deionized water, SCO San Carlos olivine, C sealed container, CSHV cold seal hydrothermal vessel, MA multi-anvil press, PC piston cylinder press, PPV Prolabo pressure vessel, TBteflon bomb, TP setup of Tesoriero and Pankow (1996), sat. saturated water vapor pressure, AAS Atomic absorption spectrometry, LA-ICP-MS Laser ablation inductively coupled plasma mass spectrometry, MB Mass balance, NA Not analyzed, SIMS Secondary ion mass spectrometry, TIMS Thermal ionization mass spectrometry … 4.2 Equilibrium Partitioning Behavior of Li 71 measured exclusively for hydrous minerals, like Kessel et al. (2005) was that Li and most trace smectite, serpentine and zeolite, which are typi- elements (except heavy REE, Y and Sc) become cal products of aqueous alteration. For pyroxene, incompatible in the bulk solid material when a Min=Fluid supercritical liquid is stabilized at 6 GPa and amphibole and olivine, KpðLiÞ is rather similar = T ≥ 800 °C. For Li this result is partly related to KMin Melt to pðLiÞ when applied to a similar major the decreasing modal abundance of Cpx in their element chemistry. Below, the data are discussed experiments at these conditions. The modal in more detail separately for each mineral or abundance of Cpx relative to Grt is of particular = mineral group. importance for Li, since KCpx Liquid [ 1 at any of = pðLiÞ Clinopyroxene (Cpx) KCpx Fluid ¼ . pðLiÞ experimental conditions applied by Kessel et al. 0:07À2:02 was measured by Brenan et al. Cpx=Supercritical (2005). In fact, K ð Þ = 1.98–2.02 at 6 (1998a, b), Kessel et al. (2005), Caciagli et al. p Li GPa and T = 800 and 1000 °C, respectively, (2011), and Fabbrizio et al. (2013) for T = 800– Cpx=Melt = ° ° – Cpx Melt whereas KpðLiÞ = 1.17 at 4 GPa and 1000 Cat 1100 C and P = 1 6 GPa. Similar to KpðLiÞ Cpx=Fluid conditions where the silicate melt is stable. These the large variability of KpðLiÞ is most likely Cpx=Liquid high values of K ð Þ at 6 GPa are higher related to the variability in major element p Li Cpx=Melt chemistry as will be discussed below. than most of KpðLiÞ measured in the pressure Cpx=Fluid ¼ range 3–7 GPa (Sect. 4.2.3). However, as illus- Brenan et al. (1998b) measured KpðLiÞ Cpx=Melt 0:09À0:25 at P = 2 GPa and T = 900 °C where trated in Fig. 4.3, KpðLiÞ is strongly correlated Cpx consisted mainly of the Di component. to the Jd content. The mineral composition of the A slight zoning of Al in Cpx was observed, Cpx and Grt was not given by Kessel et al. which was interpreted as growth zoning due to (2005), but as the starting composition was an the large compatibility of Al. They found a average MORB composition with 3.2 wt% strong correlation of the Li content with Al, Na2O, and Grt and Cpx were the only major which was associated with a dependence of minerals formed, it is obvious that Cpx contained Cpx=Fluid a significant portion of Jd. The corresponding KpðLiÞ on the Al content in Cpx. This obser- increase of Al on the M site makes the formation vation indicates a coupled substitution for Li to of LiAlSi2O6 more likely. form the neutral component LiAlSi2O6 in Cpx as = Fabbrizio et al. (2013) performed experiments KCpx Melt was also inferred in Sect. 4.2.3 from pðLiÞ .In with and without the diamond trap method to = one experiment 0.5 molal NaCl was added to the measure KMin Fluid ¼ 0:72 at 2 GPa and T = 900 = pðLiÞ fl KCpx Fluid uid, but no effect was apparent for pðLiÞ . °C. The starting material consisted of a peridotite Cpx=Fluid ¼ or a synthetic Mg(OH) –SiO mixture both Caciagli et al. (2011) measured KpðLiÞ 2 2 0:07À0:3 at T = 800–1100 °C and P = 1 GPa, doped with TiO2, NaCl, and/or CaF2. The main where similar to Brenan et al. (1998a, b) the Cpx aim of the study of Fabbrizio et al. (2013) was to consisted of 95–100 % Di. investigate the partitioning of halogens into In Kessel et al. (2005) only bulk rock–liquid upper mantle minerals; in general, these authors partition coefficients but no mineral–liquid found strong fractionations between Amp, Ol, fi Min=Fluid coef cients were reported. We have calculated Cpx, and Opx for trace elements but KpðLiÞ Cpx=Liquid Grt=Liquid K ð Þ (as well as K ð Þ , see below) from was relatively similar for all these phases. For p Li p Li ° the information given in the electronic appendix example, in their experiment at 900 C where all Min=Fluid of their paper. The experiments were performed these phases coexisted, KpðLiÞ varied only at P = 4 and 6 GPa, T = 700–1200 °C, which Ol=Fluid Amp=Fluid between 0.35 (K ð Þ ) to 0.82 (K ð Þ ) (see produced an aqueous fluid, a hydrous silicate p Li p Li also Sect. 4.2.5). melt or a supercritical liquid. A main result of 72 4 Li Partitioning, Diffusion and Associated Isotopic …

= Garnet (Grt). KGrt Fluid ¼ 0:0083 was mea- 100 pðÞLi Decarreau et al. (2012) sured by Brenan et al. (1998b)atT = 900 °C and 150°C 90°C Grt=Fluid 75°C P = 2 GPa and K ðÞ ¼ 0:12À0:52 by Kessel p Li 10 et al. (2005)atT = 700–900 °C and P = 4 GPa.

/ Grt=Melt

Sme

Maj p(Li) As discussed before for KpðÞLi , the com-

K ponent may be of relevance here, which gener- 1 Berger et al. (2012) ally increases with increasing P (see also 260°C 150°C Hanrahan et al. 2009 and discussion below in 50°C Sect. 4.2.5). However, Kessel et al. (2005) report 0.1 no major element data and Grt composition is 0.1 110100 μ thus largely unknown. Lisol ( g/g) = Olivine (Ol). KOl Fluid ¼ 0:14À1:34 was Sme=Fluid pðÞLi Fig. 4.4 Measured KpðLiÞ plotted against Li in measured by Caciagli et al. (2011) and Fabbrizio solution for different temperatures as indicated in the fi dashed lines et al. (2013). In both studies very Fo rich olivine gure. The are linear regressions of data from Decarreau et al. (2012) for each temperature was produced with Mg# = 97–100 and both data = fi − KSme Fluid Ol=Fluid assuming a xed slope of 1/2. Note that pðLiÞ of sets show a consistent change of KpðÞLi with Berger et al. (1988) also depends on the Li content of the T (at P = 1 and 2 GPa, respectively). Whether solution but follow a different and less systematic trend this trend relates to the dependence on T related for each temperature to increasing solute content in the fluid with T (see Fig. 4 in Fabbrizio et al. 2013 and related Smectites (Sme). Decarreau et al. (2012) and discussion) or to the detailed olivine chemistry is Sme=Fluid Berger et al. (1988) measured K = 0.67– unclear. Caciagli et al. (2011) parameterized this pðLiÞ dependence on T, as follows: 22 in the temperature range between 50 and = 260 °C (Fig. 4.4). In both studies a systematic Ol Fluid ¼À6:0ðÞþÆ2 6:5ðÆ2ÞÂ1000= ln KpðLiÞ T. Sme=Fluid = change of KpðLiÞ with the Li concentration in Plagioclase (Pl). KPl Fluid ¼ 0:09À0:30 for pðLiÞ the solution was observed (Fig. 4.4), but Berger An rich plagioclase (91–99 % An) was measured et al. (1988) reported only average values at a by Caciagli et al. (2011)atT=800–1000 °C and given T. Decarreau et al. (2012) measured An P = 1 GPa. A very strong apparent effect of Sme=Fluid – KpðLiÞ for hectorite, a Mg Li smectite with was observed that can be parameterized by the exch the formula (Mg3−xLix)5[Si4O10](OH)2Mx following relationship in kJ/mol (cf. with exch Eq. 4.2.11): (Mx stands for exchangeable cations on the interlayer site, mainly Na), by synthesis from a Ol=Fluid gel at 75–150 °C and equilibrium water pressure. RTln K ¼ 162À188 Á XAn ð4:2:14Þ pðÞLi The Li concentration in hectorite increased with time until a stable concentration was reached, This dependence cannot reflect the explicit An which was interpreted to represent thermody- effect of according to Dohmen and Blundy Sme=Fluid (2014). Even when compared to the study of namic equilibrium. KpðLiÞ decreased system- Coogan (2011) this effect is much greater (−188 atically with increasing Li content in the solution, vs. −70 kJ/mol, see also discussion in which was explained by the heterovalent Sect. 4.2.5). The variation of the chemical exchange of Mg2+ with Li+ charge balanced by potential of the Li component in plagioclase is Na+ on the exchangeable interlayer site. Based controversial and the presence of some amounts on the mass action law of this exchange reaction, of interstitial Li (as indicated for other silicates, Decarreau et al. (2012) derived the following Sect. 4.3) may be an explanation. relation: 4.2 Equilibrium Partitioning Behavior of Li 73  ÀÁ2 hence these values should be treated with cau- CSme  CSmeCFluid=CFluid Á CFluid Li Mg Na Mg Li tion. However, it seems obvious that Li is very ð4:2:15Þ incompatible in epidotes and structurally similar minerals like lawsonite. The only available site After dividing the above relation by the for Li is the large Ca site, which needs charge Fluid fi square of CLi and using the de nition of balance by one 4+ trace cation on the Al site, or Sme=Fluid one 5+ trace cation on the Si site. K ð Þ according to Eq. 4.2.1, we obtain a p Li Serpentine (Srp). Wunder et al. (2010) similar relationship for the Li partition measured mineral fluid partitioning for three coefficient: different serpentines, chrysotile (Ctl) (at 0.4 GPa, rffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi ° ° 1 400 C), antigorite (Atg) (at 4 GPa, 500 C) and Sme=Fluid  Sme Fluid= Fluid Á qffiffiffiffiffiffiffiffiffiffiffi – ° KpðLiÞ CMg CNa CMg lizardite (Lz) (at 0.2 GPa, 200 300 C) and Fluid CLi found that Li is generally very compatible in ð4:2:16Þ serpentines with Ctl=Fluid ¼ 16À34\ Atg=Fluid ¼ 65\ Lz=Fluid In the experiments of Decarreau et al. (2012) KpðLiÞ KpðLiÞ KpðLiÞ the factor on the left side of equation was rela- ¼ 83À152 tively constant, which can thus explain the con- − stant slope of 1/2 in Fig. 4.4. Carbonate (Cb). In study of Mariott et al. Berger et al. (1988) reported an average value (2004a, b) calcite and aragonite were precipitated – ° of 2.4 1.9 between 50 and 260 C, respectively. inorganically from continuous influx of Consistent with the study of Decarreau et al. carbonate-saturated fluid that was doped with Li Sme=Fluid T – ° (2012), KpðLiÞ increases with decreasing Li at =530 C. They determined the exchange content of the solution (Fig. 4.4). The discrep- partition coefficient (or distribution coefficient) of ancy between the data of Decarreau et al. (2012) Cb=Fluid ¼ Cal=Fluid= Cal=Fluid Li with Ca, KpðLiÀCaÞ KpðLiÞ KpðCaÞ but and Berger et al. (1988) can be explained = KCal Fluid values were not reported separately. according to Eq. 4.2.16 by variable major ele- pðLiÞ ment chemistry of smectite and solution with Li Since the Ca content in the solution and in the content as given in Table III in Berger et al. carbonates were effectively constant in these Cb=Fluid (1988). experiments KpðLiÀCaÞ also gives an indirect Based on Fig. 4.4 we can conclude that for Cal=Fluid inference on the behavior of K : In Mar- typical seawater compositions or hydrothermal pðLiÞ Cb=Fluid vent fluids with Li concentrations in the range of riott et al. (2004a) KpðLiÀCaÞ was shown to depend Sme=Fluid T 0.1–1.0 mg/L, K ð Þ  10; therefore, smec- exponentially on , decreasing by a factor of p Li T ° tite can represent a major sink for Li during about three with increasing from 5 to 30 C, formation by aqueous alteration (see also and therefore the Li/Ca ratio in carbonate was Sect. 4.4). suggested as a thermometer. However, the Li/Ca Amp=Fluid ratio is also sensitive to the salinity. Marriot et al. Amphibole (Amp). K ð Þ ¼ 0:82 was = p Li (2004b) revealed that KCb Fluid for calcite measured at T = 900 °C and P = 2 GPa in a single pðLiÀCaÞ experiment by Fabbrizio et al. (2013) for a increases by a factor of four as salinity increases ‰ NaCa-amphibole. from 10 to 50 . Inorganic aragonite demon- fi Cb=Fluid Lawsonite (Lws) and Zoisite (Zo). Martin strates no signi cant change in KpðLiÀCaÞ over the Zo=Fluid  0:003 same salinity range (Marriott et al. 2004b). The et al. (2011) measured KpðLiÞ and = different behaviour for aragonite and calcite was KLws Fluid  0:0004 À 0:01 with standard devia- pðLiÞ explained by the different incorporation tions that are similar to the measured values and 74 4 Li Partitioning, Diffusion and Associated Isotopic … mechanism of Li. Substitution of Li+ in the Ca2+ fractionation of stable isotopes between two site in aragonite makes the Li/Ca of phases. For example, the lower coordination of growth-solution the key variable, while intersti- Li in fluid (approximately fourfold) than in many tial incorporation of Li+ in calcite means that the crystal structures (mostly 6–8 fold) is responsible Li concentration of growth-solution is more for the preference of light Li in most minerals important. Based on the comparison of the compared to the fluid (e.g., Wunder et al. 2006). experimental results and Li/Ca ratios of bio- Theoretical calculations of Jahn and Wunder genically produced carbonates, Marriott et al. (2009) have also shown that at high fluid densi- (2004b) argue that there is little biological con- ties the fractionation behavior may be reversed trol during Li incorporation into calcium and 6Li preferentially partitions into the fluid carbonates. compared to minerals where Li is fourfold Zeolite (Zeo) and Chlorite (Chl). Berger coordinated (e.g., staurolite). If fluid densities Zeo=Fluid ¼ 3:5À33 increase above 1.2 g/cm3, the proportion of et al. (1988) measured KpðLiÞ and = fivefold and sixfold coordinated Li is larger than KChl Fluid ¼ 0:35À1:7 T – ° pðLiÞ at =50260 C. For those of threefold and fourfold, and hence the Min=Fluid average coordination in the fluid is greater than both minerals KpðLiÞ show a pronounced increase with decreasing T, which is opposite to 4.5. In a subsequent theoretical study Wunder the trend observed usually for partition coeffi- et al. (2011) (e.g., see their Fig. 4.3) have shown cients. Berger et al. (1988) argued that these that coordination is a good proxy for the mag- – partition coefficients do not reflect true equilib- nitude of fractionation but the mean Li O bond = distance is an even better parameter to estimate rium partition coefficients (also KSme Fluid, see pðLiÞ mineral–fluid fractionation of Li isotopes (and above) at these conditions but rather effective potentially other isotopes). Equilibrium fraction- fi partition coef cients that would be relevant for ation of Li isotopes was calculated for staurolite, alteration processes at low temperatures. mica, and spodumene by Kowalski and Jahn Theoretical constraints for equilibrium (2011) using an ab initio approach. The predicted 7 fractionation of Li isotopes. Equilibrium frac- values for Δ LiMin/Fluid agreed with the experi- tionation of stable isotopes is a mental values to within ±1.0 ‰ where, in gen- quantum-mechanical effect and is related to the eral, the theoretical values were higher. The dependence of the vibrational modes of atoms in theoretical values were corrected by considering a phase on their mass (Urey 1947; Schauble the thermal expansion of the crystals and com- 2004). In general a harmonic oscillator serves as pression of the fluid, which lowered the theo- 7 a model to explain this dependence for atoms in retical values for Δ LiMin/Fluid and lead to an even crystals, which predicts for most cases that heavy better agreement with the experimental data. ions concentrate in solids compared to Fractionation of Li isotopes due to parti- fl 7 gas/ uids/melts due to the stronger bonding of tioning between minerals and fluid, Δ LiMin/Fluid, elements within the crystalline structure. The was measured for amphibole, carbonate, a isotopic fractionation between two phases and clinopyroxene, mica, quartz, serpentine, stauro- b fi Δ7 δ7 − δ7 de ned for Li as Liα/β = Liα Liβ is in lite and clay minerals (illite/smectite); most of general strongly temperature dependent. Based these data are shown in Fig. 4.5. In general, each on thermodynamic constraints, Bottinga and data set follows the theoretically expected linear 7 2 Javoy (1973) have shown that for a limited trend when Δ LiMin/Fluid is plotted against 1/T , temperature interval the temperature dependence and the slopes appear almost parallel. Within the 7 2 can be expressed as Δ Liα/β = A þ B=T , where A experimental T range each data set could be also and B are material specific constants reflecting fitted assuming a reciprocal T dependence as the magnitude of isotopic fractionation. illustrated in Fig. 4.5 for the data from spo- The coordination as well as the bond distance dumene by Wunder et al. (2011). Exceptions are of an element have been used to infer the the scarce data for T < 100 °C of smectite and 4.2 Equilibrium Partitioning Behavior of Li 75

7 2 Fig. 4.5 Δ LiMin/Fluid versus 1/T for various minerals. T. within the experimental T range both regressions can Note that for T < 100 °C the observed isotope fraction- reproduce the data within the error. References: stauro- ation seems to be relatively constant for smectite and lite + Li-mica, Wunder et al. (2007), spodumene, Wunder calcite, inconsistent with the strong temperature effects et al. (2006, 2011), diopside, Caciagli et al. (2011), observed at higher temperatures following the theoreti- amphibole, Wunder et al. (2011), smectite, Vigier et al. cally expected relationship. The solid line is a linear (2008), calcite, Marriott et al. (2004a), regression of the data for spodumene, the dashed line is a chrysotile + lizardite + antigorite, Wunder et al. (2010) 7 regression assuming Δ LiMin/Fluid is proportional to 1/ calcite, which do not show any significant experiments of Marriott et al. (2004b) compared dependence on T. This feature might indicate to those of Marriott et al. (2004a). Thus the faster disequilibrium in the respective experiments, precipitation rate (Marriott et al. 2004b) may be because of slower reaction rates at these low responsible for the smaller fractionation temperatures. Marriott et al. (2004a) discussed observed. If the growth rate is very fast, trace possible kinetic effects on δ7Li for the precipi- elements and their isotopic concentrations cannot tation of calcite. Coralline aragonite also does not re-equilibrate with the fluid during growth but show a strong effect of T on δ7Li. If it is related are “randomly” incorporated into the mineral to kinetic effects, the observed δ7Li in carbonates structure. should be a function of the precipitation rate and In most minerals 6Li is preferred compared to hence oversaturation. the fluid. Exceptions are staurolite and chrysotile. In a second set of experiments, Marriott et al. The Li–O bond distance, roughly characterized (2004b) precipitated calcite and aragonite at 25 ° by the coordination of Li, in the respective C with δ7Li of around −3 and −12 ‰, respec- crystallographic site relative to those in the fluid tively, showing that 6Li prefers aragonite com- is the critical parameter which controls the frac- pared to calcite. The growth rates of calcites in tionation of the Li isotopes (Wunder et al. 2011). experiments of Marriott et al. (2004a, b) were not Therefore the preference of 7Li over 6Li in directly given but after the similar time (about staurolite compared to the fluid was explained by 5 h) twice as much calcite was grown in the the fourfold coordination in Wunder et al. (2007) 76 4 Li Partitioning, Diffusion and Associated Isotopic … and this behavior was confirmed by theoretical surfaces to the bulk δ7Li was investigated by calculations of Kowalski and Jahn (2011). For Williams and Hervig (2005), who measured 7 chrysotile the positive Δ LiCtl/Fluid was explained re-partitioning of Li isotopes during reaction by the presence of Li in nano-tubes in their from smectite to illite. It was speculated that the crystallographic structure. For the different types isotopic signature of Li in the exchangeable of serpentine Wunder et al. (2010) observed the interlayer site would be similar to the isotopically 7 7 sequence Δ LiCtl/Fluid = 0.6–0.9 ‰ > Δ LiAtg/- heavy waters, whereas the octahedrally coordi- 7 Fluid = −1.1 ‰ > Δ LiLz/Fluid = −2.8 to −4.1 ‰. nated Li is light. Partitioning experiments of Additional data not shown in Fig. 4.5 are Decarreau et al. (2012) do not indicate a signif- those of Lynton et al. (2005) who observed icant amount of Li present in the interlayer site. 7 7 Δ LiQtz/Fluid =8–12 ‰ and Δ LiMs/Fluid =18– In addition, according to Pistiner and Henderson 20 ‰ at 500 °C and 50–100 MPa. The enrich- (2003) surface sorption does not induce isotopic ment of 7Li relative to fluid in Qtz and Ms is in fractionation between smectite and fluid. In the contradiction to Wunder et al. (2007) and what study of Williams and Hervig (2005) Li content would be expected from theoretical constraints and δ7Li increased with decreasing average grain based on the Li–O bond distance (Wunder et al. size, which most likely shows the effect of 2011; Kowalski and Jahn 2011). Lynton et al. adsorbed Li, since with decreasing grain size the (2005) used a different experimental strategy: contribution of surficially adsorbed elements to natural minerals were exposed to a Li-rich fluid the bulk increases. and did not equilibrate with the fluid by reaction and re-crystallization but by solid-state diffusion of Li into the Qtz and Ms. The 6Li and total Li 4.2.5 Inter-mineral Partitioning content in the solid changed, with an overall slight increase of total Li and decrease of δ7Li Inter-mineral partitioning of Li may be relevant after run durations of 15 or 30 days at 500 °C. for closed system processes where intensive Whether thermodynamic equilibrium was parameters like P and T change during cooling reached in this study is unclear and Wunder et al. and decompression of the rocks. For example, it (2007) speculated that diffusive fractionation has been suggested that Li in mantle xenoliths contributed to the observed fractionation. How- was redistributed between Ol and Cpx during ever, as is illustrated in Sect. 4.3, diffusion of Li cooling such that Li leaves Ol and enters Cpx, into a mineral should produce a negative anom- leading to significant diffusive isotopic fraction- aly for δ7Li since the 6Li diffuses slightly faster ation (Jeffcoate et al. 2007). The temperature than 7Li. Therefore the origin of the positive dependence of Li partitioning between minerals 7 values for Δ LiMs/Fluid remains unclear. For and the resulting diffusive exchange during quartz it should also be considered that Li cooling can be used to constrain the respective occupies interstitial positions in channels parallel rate. For example, Coogan et al. (2005) used the to the c axis, and its presence has been shown to re-distribution of Li between Pl and Cpx during correlate with that of Al, suggesting a coupled cooling to infer the hydrothermal cooling of new substitution of Li + Al3+ with Si4+ (Dennen upper oceanic crust. So far there are only three 1966; Perny et al. 1992). studies that examined the partitioning of Li One data point measured at 300 °C plots off between mineral phases (Coogan et al. 2005; the general trend for other micas. Wunder et al. Hanrahan et al. 2009; Yakob et al. 2012). Addi- (2007) speculated about disequilibrium at these tional inter-mineral partitioning data for Li can be low T conditions and surface energy effects extracted from multiply saturated mineral–melt or related to the small grain size of the mica. Bulk mineral–fluid partitioning experiments assuming analytical data are clearly more affected by bulk thermodynamic equilibrium. Inter-mineral adsorbed Li if the surface area/volume ratio is partitioning also provides clues to the relative large. The contribution of adsorbed Li on mineral compatibility of Li in minerals. In Fig. 4.6aan 4.2 Equilibrium Partitioning Behavior of Li 77

Ol=Cpx Pl=Cpx piston cylinder apparatus and a pressure cell, overview of measured KpðLiÞ , KpðLiÞ , and = which was assumed to buffer the fO2 close to KGrt Cpx T pðLiÞ as a function of at various pressures is NNO. The starting mix consisted of natural Ol, shown. Since Cpx was the most common mineral Cpx, Kfs, Qtz, and a Li-solution. As run product present in the partitioning experiments, we can only Cpx and Ol were found together with use it as a reference mineral here, and these par- quench products from the fluid. The final Li titioning coefficients provide direct information concentrations in Ol and Cpx were in the range on the relative compatibility of Li in the most of 7–30 μg/g and most importantly it was found relevant minerals of the upper mantle and the Ol=Cpx that K ð Þ was effectively constant at 2.0 ± 0.1 oceanic crust. The disadvantage of using Cpx is p Li T Ol=Cpx that its major element chemistry and hence the and therefore independent of . Most KpðLiÞ compatibility of Li can be quite variable, as dis- values calculated from multiply saturated min- cussed before (see Sect. 4.2.3 and Fig. 4.6). Since eral–fluid (Caciagli et al. 2011; Fabbrizio et al. Opx=Cpx 2013) or mineral–melt partitioning experiments the present data for KpðLiÞ are close to 1 (Bre- (Brenan et al. 1998a; Blundy and Dalton 2000; nan et al. 1998a: 0.74 ± 0.21; Fabbrizio et al. Ottolini et al. 2009) carried out at different 2013: 0.72 ± 0.55, 0.66 ± 0.27; Adam and Green pressures and higher temperatures are in general 2006: 1.08 ± 0.24), Li seems to not partition Ol=Cpx significantly between two pyroxenes with a typ- agreement with this result. KpðLiÞ measured at ical mantle chemistry. Therefore, at least for the higher temperatures by Brenan et al. (1998a) and experimental conditions investigated it can be Blundy and Dalton (2000) are slightly lower at generally concluded that Ol is the most compat- about 1.3. The recent data of Fabbrizio et al. ible phase for Li among major mantle minerals, Ol=Cpx (2013) and Caciagli et al. (2011) for KpðLiÞ are which is also consistent with observations in significantly different, varying between 0.5 and mantle xenoliths (see Chap. 5). Considering the 10. Considering the discussion above on the modal abundance of olivine in mantle peridotites = KCpx Melt it is obvious that the Li budget of the upper large variability of pðLiÞ , this minor variation mantle is dominated by olivine. For basalts and within a factor of 2–5 can probably be assigned gabbros, Cpx is more relevant than Pl but the to the variable chemistry of Cpx in the different presence of Ol may also contribute to their bulk studies. In addition, for incorporation of Li into f Li inventory. Lithium appears to become more olivine, O2 or doping with Y2O3 seem to be compatible with pressure in Cpx and Grt, which is Ol=Melt relevant. As discussed before for KpðLiÞ and related to the respective mineral chemistry (see particularly trivalent cations like Fe3+ or Sc3+ below). This could have consequences for the (Grant and Wood 2010) are important in olivine compatibility of Li in eclogite. to charge balance Li. Since the variation in the Next, we discuss the data shown in Fig. 4.6 mineral chemistry of the re-crystallized Ol and separately for the three different mineral pairs, Cpx in the experiments of Yakob et al. (2012)is with the emphasis on the role of the thermody- relatively small (Cpx with XDi ¼ 0.9–0.99 and namic variables on the inter-mineral partitioning small amounts of En and Mg-Ts) and does not P data. It should be noted that in addition to and depend systematically on the T of the experi- T additional thermodynamic variables may be ment, their study reflects the intrinsic dependence relevant for inter-mineral partitioning depending = of the equilibrium constant for KOl Cpx on on the incorporation mechanism of the respective pðLiÞ T element as shown for Mg partitioning between Pl . Therefore, we would not expect re-distribution Pl=Cpx of Li between Cpx and Ol during cooling and and Cpx (Faak et al. 2013, where K pðMgÞ agree with the conclusion by Yakob et al. (2012) depends on the activity of silica) that “partitioning of Li between olivine and Ol-Cpx. Yakob et al. (2012) measured diopside is not temperature dependent over a Ol=Cpx – ° KpðLiÞ at 1.5 GPa between 700 1100 C using a 78 4 Li Partitioning, Diffusion and Associated Isotopic …

Fig. 4.6 Experimental (a) data for inter-mineral partitioning data of Li between Ol and Cpx, Pl and Cpx, Grt and Cpx as a function of temperature and pressure. a Black solid line (Yakob et al. 2012) and open squares represent Ol=Cpx Orange solid lines KpðLiÞ ; (parameterization of Hanrahan et al. 2009 for three pressures as given next to the lines) and open triangles represent Grt=Cpx Violet solid line KpðLiÞ ; represents the parameterization of Coogan Pl=Cpx et al. (2005) for KpðLiÞ . The color of the symbols indicates the source of the data as given in the legend panel in the figure together (b) with the experimental pressures. b Individual data of Coogan et al. (2005) for Pl=Cpx log KpðLiÞ and their corrected data, considering Pl=Cpx that KpðLiÞ depends on the anorthite content of Pl (for details of the calculations see text). The solid line is the same as shown in panel (a)

range of temperatures relevant to mantle and from mantle conditions as argued for example by igneous processes.” However, for a significantly Seitz and Woodland (2000). different Cpx chemistry this conclusion may not Pl-Cpx. Coogan et al. (2005) measured Ol=Cpx Pl=Cpx – ° be strictly valid. Furthermore, KpðLiÞ is sensitive KpðLiÞ at 0.1 MPa and T = 800 1200 C and to detailed mineral chemistry and to infer if Li of found a significant temperature dependence, Ol and Cpx in mantle samples reflect equilibrium which was parameterized as follows (Fig. 4.4b): at mantle conditions, this should be taken into Pl=Cpx ¼ 7 þ À11534 ln KpðÞLi T . account. Consequently, a deviation of the The starting material consisted of hand-picked Ol=Cpx apparent KpðLiÞ from about 2 in mantle samples and pulverized Pl and Cpx from mid-ocean ridge may not necessarily indicate a re-setting of Li gabbro and basalt. The composition of Cpx was 4.2 Equilibrium Partitioning Behavior of Li 79 constant, given as XJd ¼ 0:03, material together with variable doping of Li3PO4. XDi þ XHd ¼ 0:775, XEn þ XFs ¼ 0:146, Experiments were performed in a multi-anvil XCaTs ¼ 0:029, XTiTs ¼ 0:016, and XAn of Pl device at P =4–13 GPa and T = 1100–1400 °C. varied between 0.72–0.91. Even for this limited Grt=Cpx A strong variation of KpðLiÞ was found with range in An, a significant range in measured changing T and P, but in addition the mineral Pl=Cpx KpðLiÞ at isothermal conditions can be expected chemistry of both Grt and Cpx changed sys- according to Coogan (2011), where partitioning tematically. The composition of Grt and Cpx can of Li between Pl with different An (0.6–0.9) was be described here by the following endmembers: experimentally investigated. He predicted that Grs, Pyr, Maj and Di, En, CaTs, respectively. ÀÁ Pl=b ¼ ; ; À 6:6 Á Grt=Cpx P ln KpðLiÞ fPT Xb XAn. If this rela- Measured KpðLiÞ < 1 but with increasing and tion is correct, this dependence needs to be also decreasing T, Li becomes increasingly compati- Pl=Cpx ble in Grt compared to Cpx in the MASH system considered for ln KpðLiÞ to obtain the explicit = (Fig. 4.4a), as given by the parameterization of KPl Cpx T dependence of pðLiÞ on . In Fig. 4.4b, the data Hanrahan et al. (2009): of Coogan et al. (2005) are normalized to XAn ¼ 0:8 using this strong dependence on An, but the Grt=Cpx ¼ 0:2351 Á À 0:00253 Á ln KpðÞLi P T resulting data are far less systematic compared to ð4:2:17Þ the uncorrected data. Alternatively, the thermo- dynamic model of Dohmen and Blundy (2014) Grt=Cpx Hanrahan et al. (2009) suggested K could be used to constrain the effect of An on pðLiÞ could be used as a geobarometer but emphasized Pl=Cpx KpðLiÞ , which predicts that that this calibration is only strictly valid for the MASH system and specifically the role of Na ÀÁ Pl=b kJ ln K ¼ fP; T; Xb À 4:71 =ðRTÞ needs to be explored in more detail. The pðLiÞ mol kJ increasing compatibility of Li in garnet at high Á þ 4:61 =ð ÞÁ 2 P X XAn mol RT XAn is directly related to Maj in Grt, which increases in the MASH system continuously and from about 0 at 4 GPa to about 0.09 at 13 GPa. Four different substitutions of Li into Grt were ÀÁ kJ Pl=b ¼ ; ; À 4:71 =ð Þ suggested by Hanrahan et al. (2009): 1. (Mg2Li) ln KpðLiÞ fPT Xb RT mol Al2(PSi2)O12; 2. (Mg3)(LiSi)(PSi2)O12;3. kJ 2 Á þ 0:61 =ð ÞÁ (Mg2Li)(AlSi)Si3O12;4.(Li2Mg)(Si2)Si3O12, XAn RT XAn mol where the substitution 1 could be more relevant at low pressures whereas substitutions 2–4 for the C1 and the I1 stability field, respectively. require the presence of the majorite component to Using this calibration the normalization of have Si on the Y site in Grt. Correspondingly, in Pl=Cpx ¼ 0:8 KpðLiÞ to XAn results in a marginal Cpx the CaTs component decreases with change. We can therefore conclude that the increasing P, which reduces the probability to experimental data of Coogan (2011) are incon- incorporate Li as LiAl2SiO6 in Cpx. Pl=Cpx The presence of Na could have a strong effect sistent with the data set for KpðLiÞ of Coogan on this relationship since it would stabilize the et al. (2005), which indicates a minor effect of An = jadeite component in Cpx, which according to on KPl Cpx, more consistent with the prediction of pðLiÞ Fig. 4.3b makes Cpx more favorable for Li. Dohmen and Blundy (2014). Other experimental data shown in Fig. 4.6a for = Grt-Cpx. In study of Hanrahan et al. (2009), KGrt Cpx = pðLiÞ are roughly consistent with the predic- KGrt Cpx was measured in the synthetic MASH pðLiÞ tions by Eq. 4.2.17. With the exception of the system. Synthetic glass was used as starting results of Bennett et al. (2004), Na was present in 80 4 Li Partitioning, Diffusion and Associated Isotopic … these other experimental studies and Cpx con- a one-dimensional isotropic diffusion medium, tained a significant portion of Jd (Fig. 4.3a). In this net flux can be described by Fick’s 1st law: addition, Grt had in general a much lower Grs @ component (*15 mol% vs. *30–40 mol% for j Cj J ¼ÀDj Á ð4:3:1Þ most experiments in Hanrahan et al. 2009), x @x which could affect the ideal cationic radius for 2+ which states that the flux of an element j along the octahedral site due to the much larger Ca x j (1.12 Å) compared to Mg2+ (0.9 Å). So the dis- dimension , Jx (units: mass or amount per unit crepancy between these data and Hanrahan et al. surface area per unit time), is directly proportional (2009) are most likely related to the different to the concentration gradient along this direction mineral chemistry. and the proportionality constant is called the dif- fusion coefficient, Dj. In the general case the dif- fusion coefficient is a 2nd rank tensor, symmetry 4.3 Diffusion of Li and Associated of which is constrained by the symmetry of the Effects on δ7Li diffusion medium (for more details, see for example Nye 1957). Hence, fluids, melts, and 4.3.1 Theoretical Background cubic crystals all represent isotropic diffusion media characterized by a single diffusion coeffi- Here we give only a very short introduction into cient. For other than cubic crystals, in principle various aspects of diffusion that are relevant to Li the diffusion anisotropy has to be accounted for. diffusion and associated fractionation of Li iso- However, to date all measurements of diffusion fi topes in minerals, fluids, and melts. A fundamen- coef cients in minerals show that the anisotropy is fi tal introduction into the mathematical treatment of often below the accuracy of the diffusion coef - diffusion can be found in Crank (1975). There are cients and almost never larger than a factor of 10 also many textbooks about diffusion written (Brady and Cherniak 2010). A notable exception mainly by materials scientists and physicists is rutile (e.g., Sasaki et al. 1985). It should be focusing on diffusion in metals, metal alloys and noted that diffusion anisotropy in a 3D grain can fi simple oxides (e.g., Manning 1968; Flynn 1972; produce chemical and isotopic pro les along Philibert 1991; Bokshtein et al. 1985; Glicksman certain cross sections that can be grossly misin- 2000). A recent monograph on diffusion in min- terpreted (e.g., Costa and Chakraborty 2004). erals and melts comprises a basic introduction into the fundamentals of diffusion and a comprehen- 4.3.1.2 The Isotopic Effect fi I sive data set for diffusion in minerals and melts The diffusion coef cient of an isotope of Li in a fl (Zhang and Cherniak 2010). given uid or solid phase is directly proportional to the jump frequency, CI, which is thermally acti- 4.3.1.1 Diffusion Coefficients vated. From dynamic theory it can be derived that Diffusion of atoms in a phase is a random process the ratio of the jump frequency of any two isotopes related to thermally activated atomic jumps. In of an element is related to the square root of their crystals at thermodynamic equilibrium certain atomic masses, as follows (e.g., Flynn 1972): fi () point defects are present at de ned concentra- 0:5 C7 6 tions that allow the movement of atoms within Li ¼ 1 þ DK Á À1 ; ð4:3:2Þ C6 7 the structure. Vacancies or interstitials are the Li most common point defects that take an active part in the diffusion of atoms. Diffusion is a where the factor ΔK is necessarily ≤1 and char- process that always occurs, atoms are moving acteristic for the active diffusion mechanism relative to other atoms, but when a concentration (e.g., Philibert 1991). For example, the factor gradient of isotopes or elements exists within a ΔK ≈ 1 for a very light isotope diffusing as an phase a net flux is created. In the simplest case of interstitial in a heavy matrix. This case defines 4.3 Diffusion of Li and Associated Effects on δ7Li 81 the maximum possible difference between the relation given in Eq. 4.3.4 was also experimen- jump frequency of two isotopes but atomic jumps tally confirmed (e.g., for Li diffusion in Si; Pell in any phase are correlated to the movement of 1960). High temperature fractionation of isotopes other atoms due to geometrical constraints and by diffusion in Earth materials has long been electrochemical coupling (e.g., Schoen 1958; see ignored, but experimental studies have docu- also the full formalism for diffusion in melts in mented this effect for diffusion of Li, Mg, Ca, Watkins et al. 2014). Consider that Li as a trace and Fe in aqueous solutions, melts, and silicates element occupies a regular lattice site in a crystal, (Richter et al. 1999, 2003, 2006, 2008, 2009, which is otherwise filled by Mg (such as in for- 2014; Chopra et al. 2012). In these experimental sterite), and that diffusion occurs by the vacancy studies the fractionation of Li isotopes (and in mechanism. A jump of a Li isotope into an analogous ways isotopes of other elements) is adjacent vacant site would leave a vacancy described by a similar equation like Eq. 4.3.4: behind and Li has moved by a distance equal to à b the distance between the two sites. However, D ð7Þ 6 Li ¼ ; ð4:3:5Þ there is also the possibility that an adjacent Mg à 7 DLið6Þ jumps into this vacant site. In the extreme case if the jump probability of the Mg is negligible, the where β is an empirical parameter that is fitted to Li isotope cannot move forward but could jump best describe the observed fractionation effects. back and forth between two adjacent sites. Experimentally determined values of β for Li in Therefore, for a net flux of Li atoms, jumps of different diffusion media are reported in Mg atoms are necessary. As a consequence, the Table 4.3. Equation 4.3.5 is not equivalent to diffusion coefficients of the two Li isotopes are Eq. 4.3.3 but it can be numerically shown that the related to each other, as follows: dependence on the mass ratio is roughly the same () when we set b ¼ f DK=2. 0:5 D7 6 Li ¼ 1 þ f DK À1 ; ð4:3:3Þ 6 7 D Li 4.3.1.3 Diffusive Coupling Because of the correlation of individual atomic where f is the correlation factor that depends in jumps, different diffusion coefficients have been general on the relative jump frequencies and the defined depending on the concentration and total number of the atoms that are involved in the chemical gradients present, for example, self active diffusion mechanism for the element of diffusion, tracer diffusion, chemical diffusion, interest. As a consequence, f is principally inter-diffusion, multicomponent diffusion (for dependent on the crystallographic arrangement of definitions of the different types of diffusion atoms and the diffusion direction and is one coefficients see Flynn 1972; Chakraborty 1995; factor that contributes to the diffusion anisotropy. Ganguly 2002; Costa et al. 2008). As explained For interstitial diffusion of a light element in a above for the case of Li diffusion in an Mg heavy matrix f = 1, and hence for this special sub-lattice, in principle the net transfer of ele- case the simple isotope effect is obtained: ments by diffusion in crystals, fluids, and melts, cannot be treated as an isolated process, inde- 0:5 D7 6 pendent from other elements. Charge balance for Li ¼ : ð4:3:4Þ 6 7 ionic diffusion and site balance for crystals are D Li important constraints that lead to the coupling of The isotopic effect as given by Eq. 4.3.3 has the diffusive fluxes of various elements within a long been recognized in the material science and diffusion medium (e.g., Lasaga 1979). Therefore, physics literature and was experimentally veri- in the more general case of a concentration gra- fied mainly for diffusion in various metals and dient of an element, Eq. 4.3.1 is a simplification, semiconductors (e.g., Johnson 1967). The because it treats the diffusive flux of element j as 82 4 Li Partitioning, Diffusion and Associated Isotopic …

Table 4.3 Experimentally measured values of β for Li isotopes in different diffusion media References Phase T (°C) β D7Li/D6Li Richter et al. (2006) Water 0.015 0.99772 ± 0.00026 Richter et al. (2003) Rhyolite/basalt 0.215 0.9674 Richter et al. (2014) Clinopyroxene 900 0.27 0.9592 Dohmen et al. (2010) Olivine 800–1000 0.3 0.95 ± 0.05 Johnsons and Kruse (1966) Rutile 450 0.5 0.93 ± 0.03 Pell (1960) Si 800 0.5 0.93 ± 0.02 independent of the concentration gradient of elements in an otherwise effectively homoge- other elements. However, in most cases of neous ionic crystal can be described by Fick’s 1st interest Li can be considered as a trace element, law (Eq. 4.3.1) with a constant diffusion coeffi- fi à which apparently simpli es how to handle Li cient Dj , which is basically equivalent to the diffusion processes in minerals. If for example Li tracer diffusion coefficient defined in Eq. 4.3.1. partially substitutes for Na in a Na end-member However, it should be noted that the critical mineral (e.g., albite), the net diffusive flux of Li concentration to justify this simplified treatment in response to a concentration gradient within might be extremely small depending on the this mineral is an inter-diffusion process where mineral. As such, for the natural concentration Na and Li have to exchange on the respective range of Li in minerals, coupling of diffusive alkali site. Inter-diffusion of ions in a binary fluxes with other elements may be relevant, in system can be described by a single diffusion particular in case of a heterovalent substitution of coefficient, as follows (e.g., Lasaga 1979): Li into a mineral (for example when Li substi- tutes Mg, see Sect. 4.2) à Á à @ c ¼ DLi DNa 1 þ ln Li ; DLiÀNa Á à þ Á à @ XLi DLi XNa DNa XNa 4.3.1.4 Diffusion, Diffusive Fractionation ð4:3:6Þ and Time The concentration distribution with time is gov- à where Xj and Dj are the mole fraction of element erned by the diffusion equation, a partial differ- j on the alkali site and the tracer diffusion coef- ential equation (or in the case of diffusion ficient of element j, respectively (note that the coupling, a set of coupled partial differential term tracer is not related to the term trace element equations) that can be derived by applying the here but to the isotopic tracer). For the limiting mass balance on an infinitesimal volume unit and fl case when XLi ! 0 and hence XNa ! 1, combining it with the appropriate diffusive ux Eq. 4.3.6 further simplifies: (i) For low concen- equation (Crank 1975; Costa et al. 2008). For the trations of Li where Henry’s law is valid the flux given in Eq. 4.3.1 for a one-dimensional activity coefficient of Li is constant in this binary system we obtain Fick’s 2nd law: system and hence the thermodynamic factor on @ @ j @2 the right hand side in Eq. 4.3.6 becomes 1. Cj ¼À Jx ¼ Á Cj ; ð : : Þ Ã = à @ @ Dj @ 2 4 3 7 (ii) Depending on the ratio of DLi DNa a critical t x x concentration of Li (XLi) exists, where from ffi à where it is assumed that the diffusion coefficient is Eq. 4.3.6 we can show that DLiÀNa DLi (for numerical examples see Costa et al. 2008) and in constant, as is the case for a limiting concentration some analogy to Henry’s law, the diffusion range of a trace element (see discussion before, coefficient of Li is independent of the concen- Eq. 4.3.6). Solutions of the diffusion equation are tration. Based on these two approximations, it given for a number of different geometries, initial has been generalized that diffusion of trace and boundary conditions in Crank (1975)or 4.3 Diffusion of Li and Associated Effects on δ7Li 83

Carslaw and Jaeger (1959). The latter book pre- (Coogan 2011; Charlier et al. 2012), in diopside sents solutions to thermal conductivity problems (Coogan et al. 2005; Beck et al. 2006), and across which can be transferred easily to diffusion the bulk rock (Teng et al. 2006; Halama et al. problems as explained in Crank (1975). In gen- 2009; John et al. 2012; Lai et al. 2015 ) have been eral, any observed concentration distribution used to infer cooling and ascent rates as well as created by diffusion (for example in a mineral) can time scales of fluid release in subduction zones. be fitted by solving diffusion equations of the type given in Eq. 4.3.7, if appropriate initial and boundary conditions are chosen and the diffusion 4.3.2 Experimental Diffusion coefficient is known. The latter quantity is a strong Coefficients function of temperature and hence additional constraints for the temperature or thermal history Any type of diffusion coefficient as a material are required to obtain information on the time constant depends on a number of intrinsic ther- scale of the related geologic process, or, for modynamic parameters, like T, P, fO2, and fH2O, example, the cooling rate of the system (example but among these, T is clearly the most relevant shown in Fig. 4.7). Diffusive fractionation of Li variable for minerals and melts of a given major isotopes can be modeled by solving Eq. 4.3.7 for element composition since the atomic jumps are both Li isotopes separately (e.g., Richter et al. thermally activated (see for example Chakraborty 2003), assuming any value of β < 0.5 describing 2008). Experimental diffusion coefficients are their relative diffusivities according to Eq. 4.3.6. usually presented by an Arrhenius law, as Concentration gradients of Li in plagioclase follows:

Fig. 4.7 Measured and MIL 03346 Model modeled diffusion profiles of Li and δ7Li in augite crystals from the Martian 10 meteorite MIL03346 (Beck 4 K/h et al. 2006). On the right side various profiles for different cooling rates are 10 K/h shown. The dashed lines on 5 fi Li (ppm) the left are the best ts from 50 400 K/h which a cooling rate of K/h 20 K/h 50 °C/h from an initial temperature of 1000 °C t = 0 was derived. Note the 0 strong variation of δ7Li within 20 ‰ produced by Li diffusion requiring a β of 10 about 0.15. Reprinted with t = 0, Vcool = 1 K/h permission from Elsevier

0 4 K/h Li (‰) 7 20 K/h δ -10 400 K/h 50 K/h

-20 10 K/h

400 2000 200 2000 200 400 Distance to core (μm) 84 4 Li Partitioning, Diffusion and Associated Isotopic …

¼ ð O ; H O, ; ...Þ Dj D f 2 f 2 Xi E þ PDV Á exp À ; ð4:3:8Þ RT where E is the activation energy at zero pressure (or effectively at atmospheric pressure), ΔV is the activation volume, and D° is the pre-exponent, which may incorporate dependencies in other intrinsic thermodynamic variables like fO2. The point defect chemistry of a mineral is in general a function of the intrinsic thermodynamic variables and hence are the diffusion coefficients. For example, fO2 has a control on the concentration of vacancies on the metal site in olivine or spinel (e.g., Dohmen and Chakraborty 2007), whereas fH2O affects the structure of silicate melts and hence their viscosity and diffusion properties (e.g., Zhang and Ni 2010). Fig. 4.8 Range of diffusion coefficients for minerals, Experimental data for Li diffusion in minerals melts and fluids, as adapted from Watson and Baxter fi and melts are illustrated in Fig. 4.8 and details of (2007) in comparison to measured Li diffusion coef - cients shown in an Arrhenius plot (logarithm of diffusion experimental conditions are given in Table 4.4. coefficient vs. reciprocal temperature, compare with In general for a given element, diffusion in fluids Eq. 4.3.8): (1)–(5) silicate glasses (see Fig. 4.12 for more is faster than in melts, which is again faster than details; Rt: Johnson 1964; Qz: Verhoogen 1952; Ol: in minerals, and correspondingly, activation Dohmen et al. 2010; Pl: Giletti and Shanahan 1997; Cpx: Coogan et al. 2005; Zrn: Cherniak and Watson 2010. The energies increase in the same order (Watson and grey areas indicate the range of diffusion coefficients of Baxter 2007; Brady and Cherniak 2010). The other elements in minerals and melts as given in Watson D same general trend can be also found for Li with and Baxter (2007). Obviously, Li for minerals and melts a few exceptions (Fig. 4.8). Diffusion of Li par- is in general larger than for other elements, except H allel to the c-axis in rutile seems to be extremely fast (Johnson 1964), similarly to or even faster than diffusion in aqueous fluids. Diffusion of Li homogenous. Before taking these experimental in quartz (Verhoogen 1952) is as fast as in a observations as fact, it is worth investigating the rhyolite glass (Jambon and Semet 1978). experimental details and discussing the possible Extrapolation of Li diffusion in plagioclase incorporation and diffusion mechanisms of Li in (Giletti and Shanahan 1997), clinopyroxene the various minerals. (Coogan et al. 2005) and olivine (Li interstitials; Rutile (Rt). Diffusion of Li in rutile was Dohmen et al. 2010) to experimental data for investigated by an unusual experimental setup basaltic melts around 1200–1300 °C shows that where rutile was first doped with Li under vac- they are comparably fast. Therefore, Li appears uum at 450 °C for only 10–20 h and then the to be special in the sense that the difference in its diffusion annealing was carried out in various diffusivity between some crystals and melts is baths (distilled water, KNO3–NaNO2, and much lower compared to other elements. This KNO2) to leach Li out of rutile at the surface and specific behavior would have consequences for maintain approximately a zero concentration at diffusion models considering the Li exchange the surface as the boundary condition of the between melt and crystals since in this case the experiment (Johnson 1964). The crucial point melt is typically treated as an infinite reservoir, in here is that the concentration of Li in rutile was which diffusion is sufficiently fast that the melt is not measured directly but inferred indirectly by 4.3 Diffusion of Li and Associated Effects on δ7Li 85 optical absorption measurements using plates that were in contact with LiCl, NaCl, or monochromatic light. Johnson (1964) assumed KCl, on only one side. The diffusion coefficient that the optical absorption of the rutile crystal is of positive ions was calculated from the electrical directly proportional to the Li content since Li conductivity σ using the relation valid for interstitials acts as a donor for electrons in rutile monovalent ions, D = σkT/(e2c), where e is the and should increase the optical absorption. charge of an electron, k is the Boltzmann’s Whether the optical absorption is directly pro- constant, and c the concentration of the diffusing portional to the Li concentration has not been ion. The concentration needs to be known to tested but the solutions have been analyzed after calculate D, but Verhoogen (1952) used a con- an apparently complete removal of Li from the stant concentration, which is only indirectly rutile to calibrate the optical absorption to the estimated from conductivity measurements apparent Li content. There are two major issues without a Li source, arguing that the concentra- here: (i) It is not clear whether the rutile was tions of the ions is equal to some maximum completely homogenized with Li in the number of point defects in quartz. It is obvious pre-anneal under vacuum, since it is only the that the derived diffusion coefficients of Li rely optical absorption which was measured. The on many assumptions there. In addition, it has to optical absorption of rutile changes significantly be assumed that the charge transport occurs in by just heating in vacuum even without a Li these experiments through the bulk of the crystal. source present. (ii) During the diffusion anneal in However, Verhoogen (1952) used natural crys- the baths the redox conditions are probably dif- tals that may contain a number of linear and ferent, leading also to changes in the electron planar defects that act as short circuit paths. Yet structure of rutile that may be completely unre- again, the concentration distributions within the lated to Li. Therefore, the data of Johnson (1964) quartz crystals were not directly measured and should be treated with great caution. There are the author himself states that “These values are two additional indirect arguments that suggest not believed to represent more than the order of their data are wrong. If their data were correct it magnitude of the diffusion coefficients; numeri- would imply that (i) the diffusion of H in rutile is cal values may be in error by a factor of 2 or 3”. several orders of magnitude slower than that of Therefore, we cannot rely on these data as Li (see comparison in Van Orman and Crispin accurate diffusion coefficients. 2010); (ii) diffusion of H or Li in water would be An independent constraint for Li diffusion in smaller than that of Li in rutile. Both implications quartz comes from the observation of Charlier seem to be very unlikely from all existing et al. (2012) who compared Li concentration empirical evidence. In a follow-up study by profiles in quartz and plagioclase from the same Johnson and Krouse (1966), applying a similar pumice (P1961, Oruanui eruption, New Zealand) experimental setup, the diffusive fractionation of and concluded from diffusion modeling that “If Li isotopes was investigated. They inferred a β the feldspar and quartz crystals in P1961 are value of around 0.5, consistent with theoretical assumed to have undergone an identical thermal predictions for interstitial diffusion. However, for and decompression history, then the the same reasons as for the work of Johnson quartz-hosted Li concentration gradients imply a (1964) these data should be treated with caution. diffusion coefficient that is similar to, or up to New experimental work in which the concen- about 4 times faster than, that in the feldspar”.In tration of Li and Li isotopes are spatially resolved another experimental study using a combination remains to be undertaken. of measurements by electrical conductivity and Quartz (Qz). To measure Li (and Na + K) SIMS, two mechanisms of Li diffusion were diffusion in quartz Verhoogen (1952) used inferred (Sartbaeva et al. 2005; no diffusion another experimental setup in which diffusion of coefficients reported): fast diffusion of Li inter- ions is indirectly measured by imposing an stitials along crystallographic channels perpen- electrical gradient across 2–3 mm thick quartz dicular to the c-axis and a slower diffusion 86 4 Li Partitioning, Diffusion and Associated Isotopic … process related to the coupled movement of Li+ the powder source technique was applied (de- and Al3+. tailed discussion of this technique in Watson and Plagioclase (Pl). In the experimental study of Dohmen 2010). A pre-annealed powder mix of 6 6 Giletti and Shanahan (1997), diffusion of Li as an Li-enriched olivine and Li2SiO3 acted as the isotopic tracer was measured in natural albite diffusion source for polished, crystallographi-

(Ab98.5Or0.8An0.6) and anorthite (An95.6) crystals. cally oriented San Carlos olivine single crystals They used films doped with 6Li as a diffusion (hereafter called as Type I experiment). Diffusion source and the diffusion couples were placed in experiments were carried out in gas mixing fur- welded Au or Pt capsules and annealed at atmo- naces at atmospheric pressures under controlled spheric pressures. The concentration distribution oxygen fugacity, which in all experiments was of 6Li after the diffusion anneal was analyzed by kept close to the wüstite–magnetite buffer (WM). SIMS either in depth profiling or step scanning The resulting diffusion profiles parallel to the mode, where Giletti and Shanahan (1997) avoided c-axis were analyzed with SIMS in the step several of the potential pitfalls in profile mea- scanning mode. In addition to the samples from surements of Li with SIMS. It is unknown whe- the powder setup, they also measured Li diffu- ther in their experiments a concentration gradient sion profiles in olivine crystals that were for the total Li was imposed and if they fitted the pre-annealed in a Li contaminated furnace normalized concentration of 6Li/7Li, 6Li/27Al, or (hereafter referred to as Type II experiment) and simply the count rate of 6Li. This would make a in two single crystal diffusion couples consisting difference since we can consider two scenarios. of synthetic forsterite and San Carlos olivine The first one would be just an isotopic exchange from the experimental study of Petry et al. (2004) of 6Li with 7Li, originally present in the natural (hereafter called as Type III experiment). crystal. In this case 6Li could diffuse as an inter- Overall they observed a diversity in profile stitial and exchange with 7Li on either interstitial shapes (Fig. 4.9) and diffusion rates that can be sites or alkali sites in plagioclase but the Li con- consistently understood and simulated with a centration stays homogeneous as it should be for two-species model considering Lii an ideal tracer diffusion experiment. The second (Li-interstitials) and LiMe (Li on regular metal scenario would be that the diffusion source sites). Although the model contains more than imposes also a chemical gradient for Li and the one fitting parameter it was shown by the profile total Li concentration increases, which involves shape that most of them are independently con- the exchange with for example Na atoms on the strained. For all the different cases observed, a alkali site (see discussion above related to consistent set of diffusion coefficients for Lii was Eq. 4.3.6) and the diffusion rate could differ from determined by fitting profiles, and a single that of the tracer diffusion rate. In any case, Giletti Arrhenius relationship can be used to calculate and Shanahan (1997) found very similar Li dif- them (Table 4.4, Fig. 4.10a). The diffusion fusion rates for albite and anorthite, which are coefficient of LiMe is much lower than that of Lii very fast compared to diffusion of other cations in (see Fig. 4.10a) and varies strongly depending to plagioclase, or for Li diffusion in clinopyroxene or first order on the Li concentration at the rim. For olivine (Fig. 4.8). Thus, at temperatures of *770 different experimental setups (Type I, II, +III) a °C, according to this data set, Li in plagioclase different rim composition was imposed, varying would record only the final seconds to minutes of from hundreds to a few μg/g of Li. For a given a volcanic eruption (Charlier et al. 2012). setup, and hence roughly constant rim composi- Olivine (Ol). Two studies (Dohmen et al. tion, a single Arrhenius relationship was found, 2010; Spandler and O’Neill 2010) investigated indicating a compositional dependence of this Li diffusion in olivine with various different diffusion coefficient. Such a compositional experimental setups imposing different concen- dependence can be related to the coupling of the tration levels of Li in olivine and corresponding diffusive flux of Li with the diffusivity of other boundary conditions. In Dohmen et al. (2010), ions (e.g., Fe and Mg), but the origin is yet 4.3 Diffusion of Li and Associated Effects on δ7Li 87

100 Since it was shown by Dohmen et al. (2010) that the net diffusion rate of Li is strongly dependent on the detailed boundary conditions at the rim and decrease with decreasing Li concen- trations imposed on the rim by the experimental ) setup, it could be expected that for diffusion of Li

m

p 10

p out of olivine and a Li concentration at the rim of

(

i μ D L around 0.6 g/g, Li becomes smaller compared to the simple Arrhenius relation given by Doh-

men et al. (2010). In fact, DLi for the M site // c-axis from the Type III experiment at 1200 °C (Fig. 4.8), with a similar concentration range of Li 1 ’ 0 100 200 300 as in Spandler and O Neill (2010), is consistent Distance from surface (μm) with the latter study, considering also the appar- ent activation energy of Dohmen et al. (2010). Fig. 4.9 Measured Li profiles in Ol from two experi- Clinopyroxene (Cpx). Two studies by Coo- ° ments of Dohmen et al. (2010) performed at 1000 C gan et al. (2005) and Richter et al. (2014) were showing an unusual step like profile shape (compare with Fig. 4.7). Solid lines are simulated fits with a diffusion dedicated to the measurement of Li diffusion in model considering two Li species, Lii and LiMe clinopyroxene, both using the powder-source technique but that of Coogan et al. (2005) addi- tionally used a 6Li-enriched powder. In the unclear. Because of this complexity, Dohmen experiments of Coogan et al. (2005), diopside et al. (2010) recommended a preliminary Arrhe- single crystals from Madagascar, with stoi- fi nius relationship for the diffusion coef cient of chiometry of (Fe0.07Mg0.88Al0.05Na0.016Mn0.002) LiMe that applies for typical natural Li concen- Ca0.99(Si1.94Al0.06)O6, were annealed in a gas trations in olivine (Table 4.3, Fig. 4.9) and has an mixing furnace at controlled and variable fO2. arbitrarily estimated uncertainty of ±1 log unit. Concentration profiles of 6Li/30Si were measured A single data point of DLi for each crystallo- by SIMS in the step scanning mode and “nor- graphic axis was measured by Spandler and mal” diffusion profile shapes were observed. O’Neill (2010). A San Carlos olivine single Four experiments, each at a different temperature crystal with a drilled hole was filled with a melt in the range between 800 and 1100 °C, were that was in chemical equilibrium with the crystal performed and two different fO2 were chosen. for the major elements at the T = 1300 °C. The The four data points can be reasonably well fitted melt was doped with various trace elements that by one Arrhenius equation (see Table 4.4)andLi diffused into the olivine, but some elements, for diffusion rates are significantly faster compared example Li, were doped with a low concentration to LiMe in olivine but slightly slower than Lii in (1.8 μg/g) and diffused out of the olivine. The olivine (Fig. 4.10b). diffusion anneal was carried out in a gas mixing Since Richter et al. (2014) used a natural furnace at an fO2 equal to FMQ–1 and the main isotopic composition they were able to measure observation of this study was that most of the 19 also the diffusive fractionation effect (Fig. 4.11) elements investigated, including REE and Li, by SIMS. Their diffusion anneals were carried diffused at similar rates as Fe and Mg, whose out at a single T (*900 °C) in sealed silica – rates were consistent with recent studies on Fe ampules containing a solid-state fO2 buffer. Two Mg diffusion in San Carlos olivine (Dohmen and types of single crystals were used, Templeton Chakraborty 2007; Petry et al. 2004). A small augite and Dekalb diopside with composition D – diffusion anisotropy for Li (log Li = 15.27 // (Fe0.11Mg0.86Al0.01Na0.04)Ca0.97(Si1.97Al0.03)O6 – [010] and 14.82 //[001]) was found, but none and (Fe0.02Mg0.97Al0.01Na0.03)Ca0.98(Si1.99Al0.01) for the REE, for example. O6, respectively, as recalculated from Table 4.2 84L attoig ifso n soitdIsotopic Associated and Diffusion Partitioning, Li 4 88

Table 4.4 Published diffusion coefficients of Li in minerals (experimental conditions and Arrhenius parameters) 2 2 References Mineral/mineral Cryst. orient. T (°C) P log fO2/fO2 E (kJ/mol) log D° (m /s) log D (m /s) Notes group (MPa) buffer Giletti and Shanahan Ab ⊥to (010), (001), 200–800 0.1 Air 146 ± 14 −3.8 ± 1.0 –– (1997) (100) Giletti and Shanahan An ⊥to (010) 500–850 0.1 Air 151 ± 62 −3.6 ± 3.6 –– (1997) Coogan et al. (2005) Cpx II to [010] 800–1100 0.1 −13, −17 258 −1.54 –– Richter et al. (2014) Cpx – 900 0.1 NNO –– −7.92 to Effective diffusion −9.33 coefficient Richter et al. (2014) Cpx – 900 0.1 WM –– −11.19 Effective diffusion coefficient Richter et al. (2014) Cpx – 900 0.1 IW –– −7.43, Effective diffusion −7.48 coefficient Verhoogen (1952) Qtz II to c 300–500 0.1 Air 86 −6.16 –– Johnson (1964)RtIItoc80–360 –– 31.8 ± 0.3 −4.53 ± 0.04 –– Cherniak and Watson Zrn ⊥to c 700–1150 0.1 – 275 ± 11 −6.14 ± 0.48 –– (2010) Dohmen et al. (2010) Ol II to [001] 800–1200 0.1 WM 214.2 ± 6.2 −2.32 ± 0.27 – Li interstitial Dohmen et al. (2010) Ol II to [001] 800–1200 0.1 WM 246 −5.9 ± 1.0 – Li Me for Ol with Li of 1–10 μg/g Spandler and O’Neill Ol II to [100] 1300 0.1 −8.3 –– −14.99 – (2010) Spandler and O’Neill Ol II to [010] 1300 0.1 −8.3 –– −15.27 – (2010) Spandler and O’Neill Ol II to [001] 1300 0.1 −8.3 –– −14.82 – (2010) NNO Ni-NiO WM Wuestite magnetite IW Iron-wuestite … 4.3 Diffusion of Li and Associated Effects on δ7Li 89

(a) rates of Li depending on the type of diffusion reservoir used and the fO2 of the experiment and the crystal used. Thus, Richter et al. (2014) adopted the two-species model of Dohmen et al. (2010) to quantitatively reproduce their obser- vations, including the different magnitude of δ7Li anomalies, varying from about 0 to –35 ‰. They obtained this variable behavior by the model assuming different initial vacancy concentrations

and different diffusion coefficients of Lii.Asa consequence they determined diffusion coeffi-

cients of Lii which varied over four orders of magnitude depending on the fO2,defined by Ni– NiO (NNO), wüstite–magnetite (WM), or iron– wüstite (IW) buffer (Fig. 4.10b). The special (b) diffusion behavior of Li in diopside and olivine is discussed in more detail in Sect. 4.3.3. Zircon (Zrn). In the experimental setup of Cherniak and Watson (2010) the powder-source technique was used with natural spodumene powder (experiments at atmospheric pressure),

Li4SiO4 mixed with REE-oxides and ZrSiO4 powder (one experiment), or spodumene with quartz (42 wt%) and zircon (17 %) (hydrothermal experiments). For the experiments at atmospheric

Fig. 4.10 Experimentally measured DLi for Ol and Cpx pressure, the diffusion couples were sealed in (diopside) shown as a function of a T for various silica glass capsules and the hydrothermal f f ° experimental setups and O2 and b O2 at 900 C: experiments were carried out in a piston cylinder = D of Lii //[001] in Ol; = D of LiMe //[001] in Ol (Dohmen et al. 2010). Symbol fill denotes the experimental apparatus where the fluid was produced by add- setup: = Type I with San Carlos olivine; = Type I ing oxalic acid to the powder mixture, which with Pakistan Ol; = Type IIb with San Carlos Ol; should buffer H2O:CO2 in the fluid at 1:1. Depth = Type IIa experiment with SC and Pakistan Ol; profiles parallel and normal to the c-axis of the = Type III with forsterite; = Spandler and O’Neill (2010) with SC olivine parallel to the three main crystal- oriented and polished single crystals were mea- lographic axes; triangles = effective D of Li in Cpx: =in sured by nuclear reaction resonance analysis Madagascar diopside (Coogan et al. 2005); = Richter (NRA) using the reaction 7Li(p, γ)8Be. The f et al. (2014) at different O2 with Templeton augite and maximum Li concentrations that they obtained Dekalb diopside. For the details of Type I, II, and III experimental setups see Dohmen et al. (2010). The solid after the experiment were in the range of 100– lines are Arrhenius relationships given by Coogan et al. 600 μg/g. According to this study Li diffusion in (2005) for Li in diopside and Dohmen et al. (2010) for Lii zircon is significantly slower than in all minerals blue area and for Li at natural concentration levels. The studied so far (Fig. 4.8), and probably defines a indicates a rough estimate of the uncertainty given by Dohmen et al. (2010) for this relationship, since the exact lower limit of Li diffusion in minerals. This dependence of Li on its concentration has not been carefully observation is consistent with the trend for other calibrated. Note that the data points of Spandler and O’Neill diffusion data for zircon, which are generally (2010) would be consistent with this relationship much lower than for other minerals (e.g., for REE and Pb; Cherniak 2010). in Richter et al. (2014). Similar to Dohmen et al. Melts. Li tracer diffusion in melts or glasses (2010) for olivine, they found various profile was determined for the following compositions shapes (e.g., Fig. 4.11) and effective diffusion (Fig. 4.12): albitic glass, orthoclase glass, and 90 4 Li Partitioning, Diffusion and Associated Isotopic …

10

2.0 0

n

-10

Li (‰) -20 7 δ

Normalized Li concentratio 1.0 -30

-40 0 0.2 0.4 0.6 0.8 1.0 0 0.2 0.4 0.6 0.8 1.0 Normalized distance Normalized distance

Fig. 4.11 From Richter et al. (2014), showing an unusual step like profile shape: Li profile in Cpx (Templeton augite) at 900 °C on left, δ7Li profile from the same sample. Reprinted with permission from Elsevier rhyolitic glass (Jambon and Semet 1978); jadeite is dissolved into the melt. Various physical glass (Roselieb et al. 1998); andesitic and dacitic models have been used to predict diffusion melt (Cunningham et al. 1983); basaltic melt coefficients from the viscosity of melts (e.g., (Lowry et al. 1981; Lundstrom 2003). In addi- Einstein 1905; Eyring 1936). The empirical tion, a number of studies exist where Li model of Mungall (2002) relates tracer self-diffusion was investigated in various Li sil- icate glasses (e.g., Welsch et al. 2012 and refer- ences therein), but here we summarize only data T (°C) for naturally relevant compositions. 1200 800600 400 The glasses and melts were mostly completely -7 dry, only the rhyolite contained some water (0.52 -8 wt% H O). In most experiments a thin film (6) (7) 2 -9 (2) enriched with Li was used, except for Lundstrom (1) (3) ) (8) s (4)

/ -10 (2003) where a basaltic melt was in contact with 2

(m a partially molten peridotite. In every study Li i

L -11 profiles were measured using SIMS after the D

g

o -12 (5) diffusion experiment. As for minerals, Li diffu- l sion in melts is enhanced compared to most other -13 cations due to its low ionic charge and size (see review of Lesher 2010). For example, in Li sili- -14 cate glasses the ionic conductivity is dominated -15 by Li diffusion (e.g., Welsch et al. 2012). How- 6 810121416 18 ever, at a given temperature the range of pub- 104/T (K) lished Li diffusion coefficients is very limited Fig. 4.12 Arrhenius plot of experimentally measured DLi when compared to the variability of Li diffusion in various melts and glasses: (1) basaltic melt (Lowry in minerals. Extrapolated data measured for et al. 1981); (2) albitic glass (Jambon and Semet 1978); albitic glass and those measured in basaltic to (3) jadeite glass (Roselieb et al. 1998); (4) rhyolitic glass dacitic melt are very similar (Fig. 4.12). Water is (Jambon and Semet 1978); (5) glass with 39 mol% albite and 61 mol% orthoclase component (Jambon and Semet known to affect diffusion in melts up to orders of 1978); (6) andesitic melt (Cunningham et al. 1983); magnitude (e.g., Behrens and Zhang 2001) (7) dacitic melt (Cunningham et al. 1983); (8) basaltic because it depolymerizes the melt and lowers the melt (Lundstrom 2003). The dashed lines are predictions D viscosity. Therefore, Li diffusion could be of Li for the rhyolitic melt and the basaltic melt as calculated from the empirical model of Mungall (2002). enhanced compared to the experimental studies if The arrow points to predictions considering 5 wt% H2O a significant amount of water (several wt% H2O) in the respective melts 4.3 Diffusion of Li and Associated Effects on δ7Li 91 diffusivities to temperature, viscosity, melt com- DLi and added 5 wt% H2O to these compositions, position, and cation size. He provides an equa- which produces a further shift to higher diffu- tion for alkali elements which is not sensitive to sivities by another half order of magnitude. viscosity at all, unlike high field strength or intermediate field strength elements, as follows: hi4.3.3 Characteristic Time Scales of Li logð Þ¼À3:02 þ 18:5 Áð À 1:03Þ2 þ 2:88 Á = DLi rLi M O Diffusion in Silicates (Ol, h 2 À 1=T Á 538 Á Al=ðNa þ K þ HÞþ11910 ÁðrLi À 1:03Þ Cpx, Pl) þ 3029 Š As discussed above, Li diffusion in minerals is The ratios M/O and Al/(Na + K + H) are the very fast compared to other elements with the molar ratio of non-bridging cations to oxygen major geochemical consequence that Li concen- 7 anions and molar ratio of the given elements, trations and δ Li can be significantly affected by respectively, and rLi is the ionic radius of Li in diffusion at even moderate temperatures within the melt. The empirical equation is inferred from reasonable time scales (e.g., in mantle xenoliths). 7 a diffusion model for alkalis (A) in melts where For δ Li we have to consider two effects: (i) the pairs of alkalis break a T–O–T bridge into a isotopic re-equilibration with a new chemical non-bridging T–O–AA–O–T group (Mungall environment, and (ii) diffusive fractionation. We 2002). Diffusion of alkalis can be therefore seen made attempts to estimate these characteristic as coupled diffusion of A2O. If the alkali con- time scales applying a simple model geometry. In centration is very low the formation of pairs Fig. 4.13a we illustrate by a normalized con- becomes unlikely and the diffusivity of alkalis centration distance plot how a spherical grain drops (Mungall 2002). In Fig. 4.12 we have also with diffusion isotropy (cubic crystal) homoge- plotted the predictions from Eq. 4.3.9 for the nizes with time (initial homogeneous concentra- C fi C experimental melt compositions of a rhyolitic tion 0, xed rim concentration, rim). The and basaltic melt. The predicted DLi is for these analytical solution of this diffusion problem is two cases about 0.5 log units too high. We also given in Crank (1975, p. 91, Eq. 6.18). From applied Eq. 4.3.9 to estimate the effect of H2Oon Fig. 4.13a we can infer directly that the time

Fig. 4.13 Normalized concentration distance plots for were performed for a value of β = 0.25. Numbers on the diffusion in a sphere and an infinite plane after various curves denote values of Dt/a2 for the sphere, where a is different times: Li flux into a sphere (left); δ7Li corre- the radius sponding to a ratio of Crim/C0 =2(right). All calculations 92 4 Li Partitioning, Diffusion and Associated Isotopic …

Fig. 4.14 Characteristic time scales of Li diffusion versus temperature for olivine, diopside, and plagioclase. See text for details

required to effectively homogenize the spherical depends also on the concentration gradient pre- grain is about 0.5 a2=D and to affect the com- sent (rim composition/initial composition). It is position of the region next to the rim (10 % of the also a rough estimate for the minimum time 7 radius) a time of about 0.001 a2=D has to elapse. needed to affect the bulk δ Li by diffusive Hence, assuming a grain size of for example fractionation. 500 μm, a minimum time of about 2.5 e−12 m2/ Using these two relationships we illustrate in D is needed to obtain a zoning within the 50 μm Fig. 4.14 how the characteristic time scales next to the rim that could be detected by change with T for olivine, diopside and plagio- LA-ICP-MS and SIMS, considering a spatial clase using the published Li diffusion coefficients resolution of about 10–20 μm. These simple of Dohmen et al. (2010) for LiMe, Coogan et al. relations will be used below to estimate the (2005), and Giletti and Shanahan (1997), characteristic time scales of Li diffusion in vari- respectively. The simplified model used here ous minerals as a function of T.Wedefine two does not consider the complexities of diffusion in characteristic time scales for a given spherical diopside and olivine (see below), but still gives a crystal with radius a, at a given temperature: a general sense of the relative time scales and their 2 minimum time scale, tmin = 0.001 · a /D(T), implications. For the calculations we have used a which would be the time necessary to affect the radius of 100 μm and to illustrate the effect of bulk or average composition of the crystal by grain size we have used also 1000 and 10 μm for 2 diffusion; a maximum time scale, tmax = 0.5 a /D olivine. (T), which would be the time necessary to In Fig. 4.14 the two characteristic time scales effectively homogenize the crystal. Note that the define the upper and lower boundary of a minimum time scale is only a rough estimate temperature-time window in which the mineral is since if the bulk composition is affected it open to diffusional exchange with the 4.3 Diffusion of Li and Associated Effects on δ7Li 93 environment, but does not equilibrate com- 4.3.4 Two Diffusion Mechanisms of Li pletely. Below this field there is no significant in Silicates effect of Li diffusion and above this field the mineral is re-equilibrated. A first obvious result For three types of silicates, quartz (Sartbaeva of this calculation is that at T > 1200 °C, even in et al. 2005), olivine (Dohmen et al. 2010), and larger crystals, Li is re-equilibrated within less clinopyroxene (Richter et al. 2014), experimental than a couple of years. However, at more mod- evidence was found that Li diffusion occurs via erate temperatures (400–700 °C) modification of two different mechanisms, as interstitial and as Li contents in olivine and diopside requires species occupying regular metal sites. The longer time scales. Plagioclase is a rather extreme incorporation of Li on interstitial as well as metal case since for Li it reacts very quickly to any sites in forsterite was also confirmed by quantum change in the chemical environment. mechanical methods and significant fractionation In the most simplified form we also simulated of Li isotopes between these two sites was pre- diffusive fractionation of Li isotopes by solving dicted (Zhang and Wright 2012). Such a behav- Eq. 4.3.7 for 6Li and 7Li separately (e.g., Gal- ior has not been recognized for any other lagher and Elliott 2009) assuming any value of element, apart from H diffusion in melts β < 0.5 describing their relative diffusivities (Wasserburg 1988; Zhang et al. 1991) or olivine according to Eq. 4.3.5. For the same case as (Kohlstedt and Mackwell 1998). It seems to be a above (e.g., initially homogeneous), we plotted general phenomenon that ions with small charge the evolution of δ7Li with time (Fig. 4.13b). Two and small radius can penetrate into ionic crystals very general features of isotopic fractionation interstitially. The presence of two diffusion patterns originated by diffusion can be recog- mechanisms can produce very unusual profile nized here: (i) Due to the larger mobility of 6Li shapes for Li but also produces isotopic frac- compared to 7Li we observe a region within the tionation effects that are less straightforward to crystal with a pronounced negative anomaly (or interpret compared to the simple case illustrated positive anomaly, respectively) for an influx (or in Fig. 4.13. This is worth exploring here in more flux out) of Li. (ii) The region of this anomaly detail. Apparently oxygen fugacity is a critical extends roughly over the region with the chem- parameter in that respect. The step-like diffusion ical zoning where the minimum (or maximum, profiles for Li in olivine (Fig. 4.9) and orthopy- respectively) of δ7Li coincides with the point of roxene (Fig. 4.11) with a spiky δ7Li anomaly is maximum curvature for the Li profile, @2C=@x2. characteristic for two diffusion mechanisms. The magnitude of δ7Li at this minimum is somehow defined by the ratio of the diffusion 4.3.4.1 Two Species Diffusion Model coefficients for 6Li and 7Li as well as by the The step profiles were successfully modeled by initial concentration gradient (see also Parkinson Dohmen et al. (2010) considering two diffusing et al. 2007), given by the difference between Crim Li species in olivine and their model was adapted and C0. The time scale over which the diffusive by Richter et al. (2014) for clinopyroxene (solid fractionation effect occurs is coupled to the lines in Figs. 4.9 and 4.11). Dohmen et al. (2010) chemical diffusion process and hence the time to assumed two different sites that Li can occupy: Li homogenize the Li isotopes is roughly equivalent on the metal site, LiMe, and Li on the interstitial to the time for chemical homogenization, as site, Lii. They further assumed that Lii is more 2 given here for the spherical model by t ¼ a =D. mobile than LiMe and that Li prefers the metal Note that in particular for Li this simplified site, hinging mostly on more favorable energetic model can be very misleading for the interpre- stability. Since olivine contains a certain number 7 tation of naturally occurring δ Li profiles, as will of vacancies on the metal site, VMe (number of be discussed in the next section. which is relatively well known; Nakamura and 94 4 Li Partitioning, Diffusion and Associated Isotopic …

Schmalzried 1983; Dohmen and Chakaraborty only on a single Li diffusion coefficient, but is

2007), Lii in a site adjacent to VMe can jump into very sensitive to the initial vacancy and Li con- this vacancy and thereby form a LiMe, as can be centrations, the equilibrium constant described by the following homogeneous reac- K (Eq. 4.3.11), and the new vacancy and Li tion between these three point defects: concentrations imposed at the rim, as defined by the chemical environment (the initial and V þ Li ¼ Li : ð : : Þ Me i Me 4 3 10 boundary conditions of the diffusion problem). To obtain the unique behavior that Li penetrates ’ fi Assuming Henry s law and de ning as the into the crystal in the form of a front with an j reference state for the point defect one point almost constant concentration and a sharp drop to defect per formula unit of the mineral (M2SiO4 or the initial concentration behind it, the set of × 6 CaMSi2O6) 10 , ppm atomic formula unit parameters has the following two constraints for (hereafter referred to as ppm afu), the corre- the diffusive fluxes and the equilibrium constant: sponding mass action law to Reaction 4.3.10 is  K ≫ fi JLii JLiMe and 1. The rst constraint obtained:  ; implies that DLii DLiMe DVMe but also that the ½ŠLi boundary conditions (rim compositions) impose ð ; Þ¼ Me ; ð : : Þ certain concentration gradients of the three point K P T ½ŠÁV ½ŠLi 4 3 11 Me i defects (Dohmen et al. 2010). Therefore, the overall diffusion behavior of Li in clinopyroxene where [j] denotes the concentration of species j in and olivine can drastically change depending on ppm afu. Because of this reaction the diffusion the chemical environment, in particular the equation (Eq. 4.3.4) has to be extended by a oxygen fugacity and the Li source. reaction term since each of the three relevant For brevity, we will not explain these features species, V ,Li, and Li , can be formed or Me i Me in detail, but the more interested reader can fol- destroyed over time, which is controlled by the low the discussions in Dohmen et al. (2010) and equilibrium constant of reaction 4.3.10. Thus, the Richter et al. (2014). Here, our aim is to illustrate governing equations for simulating the observed various consequences of these two mechanisms Li diffusion profiles in olivine and clinopyroxene for the rate of diffusive equilibration, diffusive are a coupled set of diffusion equations with fractionation and for the profile shapes that could j ={V ,Li, and Li }: Me i Me be observed in natural crystals. In Fig. 4.15 we 2 show strongly variable profile shapes and frac- @C @ C j ¼ D Á j þ R ; ð4:3:12Þ tionation of δ7Li by varying just the vacancy @t j @x2 j concentrations at the rim (see also Dohmen et al. where Rj is the rate of production/destruction of 2010 or Richter et al. 2014 for other examples). diffusing species j. The coupled sets of diffusion In all cases we used the following set of ¼ ¼ =100 equations were solved numerically using the parameters, DLiMe DVMe DLii and K method of finite differences with the assumption = 100, consistent with the experimental that at each grid point, local equilibrium of observations of Dohmen et al. (2010) and Richter reaction 4.3.10 was attained within each time et al. (2014), initially homogenous vacancy step of the calculation (Dohmen et al. 2010; concentrations of 50 ppm afu and Li of 21 ppm * μ Richter et al. 2014). afu (equivalent to 1 g/g in olivine and *0.7 μg/g in clinopyroxene), and a fixed con- 4.3.4.2 Modeling Results—Sensitivity centration for Li at the rim of 84 ppm afu (equivalent to *4 μg/g in olivine and *2.8 μg/g to fO2 Due to the complexity of Li diffusion in various in clinopyroxene). fi silicates the observed zoning of Li and its iso- Since we xed the equilibrium constant and topes can be strongly variable and depends not the total Li content at the rim, we can change the 4.3 Diffusion of Li and Associated Effects on δ7Li 95

(a) (b)

(c) (d)

(e) (f)

Fig. 4.15 Simulated profiles using the two species the fO2. a and b:[VMe] = 0.05 ppm afu; c and d: diffusion model (Dohmen et al. 2010) for various different [VMe] = 10 ppm afu; e and f:[VMe] = 0.01 ppm afu. The vacancy concentrations at the rim, showing that Li numbers on the curves denote Dt/l2 where l is the half diffusion and associated isotopic fractionation strongly diameter of the plane sheet depends on the chemical environment as defined by, e.g., 96 4 Li Partitioning, Diffusion and Associated Isotopic … relative concentrations of Lii and LiMe by just parameters (the same β value) but different varying the rim concentration of the vacancies. boundary or initial conditions. The model is For example in case (i), for a value of relatively simple considering the broad range of

[VMe] = 0.05 ppm afu at the rim, step-like profile different observations it can reproduce. shapes are obtained (Fig. 4.15a) and corre- Based on the simulations with the two species sponding isotopic fractionation patterns are cre- model, Dohmen et al. (2010) concluded that Lii ated (Fig. 4.15b) that are equivalent to those only becomes relevant if Li concentrations were found by Richter et al. (2014). For times where present in olivine that exceeded typical Li natural the Li concentration is effectively equilibrated concentrations (>10 μg/g). However, the simula- (Dt/l2 = 2.5 or 5), there is still a significant iso- tions presented in Fig. 4.15 with this model sug- topic anomaly present (down to –30 or –10 ‰, gest that even for a reasonable range of Li and respectively). In case (ii), when we increase vacancy concentrations, quite variable net diffu-

[VMe] at the rim to a value of 10 ppm afu, we sion rates and isotopic fractionations can be obtain “normal” profile shapes for both [Li] and expected that are related to the presence of Lii. δ7Li (compare Fig. 4.15c–d with Fig. 4.7), and A critical thermodynamic parameter appears to be the isotopic anomalies are smaller by a factor of the fO2, which strongly affects the concentrations 2–3 than in the former case. However, again, the of vacancies on the metal site (Nakamura and isotopic anomaly is still present after a critical Schmalzried 1983; Dohmen and Chakraborty time where with the common analytical methods 2007). The profile shapes of Li and δ7Li measured a concentration gradient would be not resolvable. in natural olivine and clinopyroxene do not The most interesting and counterintuitive case indicate that the interstitial diffusion mechanism

(iii) occurs when we decrease [VMe] to a value of was relevant in these cases. Eventually, more 0.01 (Fig. 4.15e–f). Here, the shape of the Li experimental data and detailed spatially-resolved concentration profile develops in a very complex measurements in natural samples are required to way, but the δ7Li profile looks very similar to the resolve this issue. From mantle xenoliths, relative case shown in Fig. 4.15b. The concentration of diffusion rates of Li in olivine and clinopyroxene Li decreases towards the rim, but the time evo- were inferred. According to these data clinopy- lution of the profile records Li diffusion into the roxene is faster than olivine by a factor of about grain. This is also evident from the isotopic 10, which is much smaller than the factor *1000 fractionation, since the grain becomes lighter predicted by the relations of Coogan et al. (2005) within the diffusion zone, showing the case of an and Dohmen et al. (2010). This apparent dis- uphill diffusion process. Contrary to the previous crepancy could be explained by the two species cases, after a certain time the isotopic ratio is mechanism and their sensitivity to fO2 as dis- homogeneous whereas there is still a significant cussed above. concentration gradient present for Li. Because of the complex diffusion behavior of The three different cases illustrated above Li one should be very cautious with the inter- have also been directly observed in experiments pretation of Li isotopic data. The numerical and with olivine and clinopyroxene: case (i) and experimental examples discussed show that some (ii) in Dohmen et al. (2010) and Richter et al. of the common assumptions for the interpretation (2014), where the different behavior was of isotopic data have to be abandoned: observed depending on the redox conditions or type of Li reservoir used in the experiments; case (i) Homogeneous Li concentration + nega- (iii) has been observed in an unpublished study tive or positive δ7Li anomaly does not on Li diffusion in clinopyroxene (Liu, Marschall, imply that δ7Li has been not affected by Gaetani, pers. communication to RD, 2014). solid-state diffusion. Note that in this case Most importantly, Richter et al. (2014) were able the δ7Li gradient can be relatively small and to reproduce Li and δ7Li profiles for case (i) and not well resolved with common analytical (ii) simultaneously with the same set of methods. 4.3 Diffusion of Li and Associated Effects on δ7Li 97

(ii) A negative Li concentration gradient interphase boundaries, which are fast diffusion towards the rim does not necessarily imply paths. The weighting includes the volume frac- Li loss from the mineral. Here, the δ7Li tion of the diffusion media and the detailed tex- profile can help if Li diffused into or out of ture (e.g., grain size, crystallographic preferred the mineral. orientation, etc.; Hashin and Shtrikman 1963; (iii) Finally, Li is a special case because of the Brady 1983; Dohmen and Milke 2010). How- two-species mechanism where both cases ever, if the modal composition of a rock is may occur: faster equilibration of the iso- dominated by one mineral (e.g., by plagioclase in topes or faster equilibration of the chemical basalt and gabbro or olivine in a peridotite composition. xenolith) we can use the respective D and aver- age grain size to calculate a first approximation Independent of these complexities, if the for the bulk diffusivity, as follows (modified Hart mineral stayed long enough at some critical equation from Mortlock 1960): temperature the memory of the crystal in terms of  7 d Li concentration and δ Li will be lost and the s Dgb Dbulk ¼ Dl Á 1 þ b Á À 1 ; ð4:3:13Þ crystal becomes homogeneous for both quanti- 2a Dl ties. The time to reach this point depends for the two-species model also on the boundary condi- where Dl is the diffusion coefficient of the min- tions (as indicated by the numbers on the curves eral, Dgb is the grain boundary diffusion coeffi- in Fig. 4.15) and is significantly different from a cient, s is the segregation coefficient, which is simple diffusion problem (Fig. 4.13). equivalent to the mineral–grain boundary parti- tion coefficient, δ is the width of the grain boundary (typically on the order of nanometers), 4.3.5 Bulk Diffusion of Li Through and b is a geometric factor that considers the Rocks and Characteristic detailed texture, which is on the order of 2–3in Length Scales most cases. The grain boundary diffusion coef- ficient is usually several orders of magnitude Measurements of Li isotopes across traverses in larger than Dl but, due to the small grain country rocks indicate that Li diffusion through boundary width, the overall contribution of grain bulk rocks affects the compositions of rocks on boundaries to the bulk diffusivity is very small if the outcrop scale (e.g., Teng et al. 2006). Diffu- there is not a strong tendency of the element to sion in a multi-phase polycrystalline rock is in segregate into the grain boundary (e.g., as pre- principle a complex process, but it can be dicted for strongly incompatible elements; Doh- mathematically shown that after some critical men and Milke 2010). We use Eq. 4.3.13 to time diffusion through such systems can be calculate the bulk diffusivity through an olivine characterized by a single bulk diffusion coeffi- and a plagioclase dominated rock, each with an cient (e.g., Hashin and Shtrikman 1963; Kaur average grain radius of 100 μm, and to estimate et al. 1995). Here one speaks of the Type A the spatial scale on which a bulk rock can be diffusion regime (e.g., Dohmen and Milke 2010) affected by diffusion (Fig. 4.16). This length and this critical time is roughly defined by the scale is calculated from the simple relation, d =4 0.5 time to equilibrate an individual grain by diffu- (Dbulkt) . These simplified calculations show sion. For an average grain radius ‘a’, this time is that, for plagioclase-dominated rocks and T ° equivalent to tmax as defined above, but using the >800 C, Li isotopes could be affected over diffusion coefficient for the mineral with the km scales. At even higher temperatures around lowest D. The bulk diffusion coefficient is a 1000 °C Li isotopic compositions of weighted average of the diffusion coefficients of olivine-dominated rocks could be affected by the various minerals and those for the grain and diffusion at outcrop scales. 98 4 Li Partitioning, Diffusion and Associated Isotopic …

dependent on the reaction progress. There are several possible experimental setups but the most typical approach of alteration experiments is to mix the solid material as a powder with water and anneal the sample inside a closed container at the desired temperature and pressure for vari- able times. The flexible cell hydrothermal solu- tion equipment (Seyfried et al. 1979, 1987)ora modified Prolabo hydrothermal apparatus (Ber- ger et al. 1987) allows periodic sampling of the fluid to monitor the reaction progress without quenching the sample. The fluid is usually fil- tered before analysis. Containers of the sample material are mostly Teflon bottles, Teflon bombs or Au capsules in a hydrothermal bomb. Critical experimental parameters for the reaction path and Fig. 4.16 Characteristic length scales of Li diffusion reaction progress in addition to the solid starting through bulk rocks dominated by either olivine or P T fl plagioclase with an average grain size of 100 μm material, and , are the uid/rock ratio, the solid surface to fluid volume ratio, and the starting composition of the fluid since it controls 4.4 Fluid-Rock Interaction the stable (or metastable) mineral assemblage as well as the timing of the formation of secondary To understand the effects of fluid–rock interac- minerals. For example, in a closed system at tion on Li and its isotopes a number of experi- relatively low fluid/rock ratio the solution is more ments at low temperatures (25–375 °C) were quickly saturated with secondary minerals due to performed with different starting materials. In dissolution (congruent or incongruent) of the most of these studies interaction between basalt solid material. Alternatively, mixed flow-through and seawater was investigated (e.g., Seyfried reactors are used (Wimpenny et al. 2010)to et al. 1984; Berger et al. 1987, 1988; Seyfried inhibit saturation of secondary minerals and et al. 1998; James et al. 2003; Millot et al. 2010; investigate primarily mineral dissolution pro- Wimpenny et al. 2010; Verney-Carron et al. cesses. Yet another experimental approach was 2011; Brant et al. 2012) in order to most faith- applied by Pistiner and Henderson (2003)to fully reproduce conditions during investigate surface sorption of Li on secondary low-temperature seafloor weathering and alter- minerals (smectite, gibbsite, ferrihydrite) and ation by hot vent fluids. In addition, weathering associated stable isotopic fractionation: the of terrigenous sediments (Chan et al. 1994; You powder of each mineral was immersed in NIST and Gieskes 2001; James et al. 2003), granite L-SVEC solution at room temperature for vari- (Pistiner and Henderson 2003), and olivine ous lengths of time. A complete list of experi- (Berger et al. 1988; Wimpenny et al. 2010) were mental studies with their specific conditions is studied experimentally to investigate associated given in Table 4.5. isotopic fractionation of Li. The experimental results are summarized in The alteration experiments are in some sense more detail below for the different starting not fundamentally different from the experiments materials and experimental approaches. The solid to measure equilibrium partitioning between reaction products are rarely investigated in detail, minerals and fluid discussed above, but since mostly using X-ray powder diffraction (XRD) or, they did not reach thermodynamic equilibrium, sometimes, transmission electron microscopy the measured isotopic composition is strongly (TEM). 4.4 Fluid-Rock Interaction 99

4.4.1 Weathering of Igneous Rocks fluid do not change considerably, which is (Predominantly Basaltic interpreted as an attainment of a steady state for Compositions) the reaction system. Berger et al. (1987, 1988) documented the leaching process in glass in more In a series of papers the behavior of Li during detail by experiments with basaltic glass slabs weathering of natural and synthetic basalt at and powder as starting materials for different temperatures between 25 and 350 °C was extent of compaction. The chemical composition investigated (Seyfried et al. 1984, 1998; Berger was studied across the sections of the altered et al. 1987, 1988; James et al. 2003; Millot et al. samples. The altered rim of the slab had chemical 2010; Wimpenny et al. 2010; Verney-Carron concentration gradients for major elements and et al. 2011; Brant et al. 2012). The modal com- also Cs (Li and Rb were not measured) where position of the basalt, especially the amount of cations were removed relative to SiO2. The glass, is assumed to have an effect on the reaction concentration profiles of major elements and H products and reaction progress (Brant et al. 2012) indicated that reaction progress was controlled by because different solid materials react in different diffusion through the bulk glass (solid state dif- ways with water and presumably have different fusion) in the early stage accompanied later by reaction rates and fractionation effects with diffusion of aqueous species in a porous gel respect to Li. Most studies used basaltic glass layer, which recrystallized continuously. The (synthetic or natural) as starting material (see amorphous silicate gel formed was interpreted to Table 4.5). be a precursor for crystalline clays in a later Pioneering experimental studies (e.g., Hajash stage. Based on these observations Berger et al. 1975; Seyfried and Bischoff 1979; see also a (1987) proposed a four-stage reaction model (see review by Seyfried 1987) have shown that during also Fig. 1 of Berger et al. 1994). weathering of basalt Mg is removed from sea- It is important to emphasize that the alteration water, and that Ca, Na, and K are leached from of basalt can be conceptually seen as a combi- basalt, which is related to the formation of nation of two fundamental mechanisms: palagonite (the amorphous precursor material of (i) leaching of the glass (or congruent or incon- smectite), chlorite, and clay minerals (mostly gruent dissolution), and (ii) secondary mineral smectite). At temperatures ≤ 70 °C smectite is formation (precipitation). Each of these two probably formed but is difficult to detect using mechanisms plays its role in partitioning of Li XRD because of its small amount and poor and its isotopes between the solid and fluid crystallinity, as also indicated from TEM inves- phase. Leaching of the glass produces a flux of Li tigations (e.g., Millot et al. 2010). The water to into the fluid, but during formation of a sec- rock ratio basically controls the amount of sec- ondary mineral Li is repartitioned between fluid ondary mineral formation. The amount of Mg in and solid; this depends on its partition coefficient the reacting water is also critical for the alteration and the initial Li content in the fluid and the process. For example, in the absence of dissolved solid. The distribution of Li may not reflect Mg, epidote and plagioclase are the dominant equilibrium partitioning between the bulk fluid secondary phases instead of smectite or chlorite and mineral, because it depends on the precipi- (Seyfried and Mottl 1982). tation rate of the newly formed solids relative to With increasing temperature, alteration of the homogenization rate of the fluid by diffusion basalt becomes more efficient, resulting generally and convection. Indeed, relatively slow transport in higher degree of leaching of Li (and Cs, Rb in the fluid could also be responsible for the and B) out of the basaltic glass As a conse- nucleation of metastable secondary phases (see quence, these elements become progressively Ruiz-Agudo et al. 2014 and references therein). enriched in the fluid (e.g., James et al. 2003; In that case an apparently higher partition coef- Millot et al. 2010). The experimental runs are ficient for Li between phases needs to be usually quenched when the concentrations in the considered. Table 4.5 List of experimental studies with their specific conditions Isotopic Associated and Diffusion Partitioning, Li 4 100

Reference Starting materials Experimental conditions Product analysis Comments

Solid Grain Li Fluid Li S/V (1/cm) w/r pH T (° P Total Container Exp. setup Solid Secondary d7Li size (mg/g) (ng/g) surface C) (MPa) run material minerals measured (μm) area/volume duration identified of solution (h)

Brant et al. 6Li doped MORB glass 63– 20 IAPSO std 8.90 9.2–9.4 5.7a 125 0.23 1986 Ti Hydrothermal n.a. Fluid Larger depletion of Mg, (2012) powder 250 seawater reaction vessel SO4

Li doped MORB powder,6 63– 20 IAPSO std 8.90 9.2–9.4 5.6a 125 0.23 1986 Ti Hydrothermal n.a. Fluid Similar texture but partly crystallized (200 ° 250 seawater reaction vessel smaller plag grains at C/min) rapidly cooled MORB

6Li doped MORB powder, 63– 20 IAPSO std 8.90 9.2–9.4 5.5a 125 0.23 1986 Ti Hydrothermal n.a. Fluid partly crystallized (30 ° 250 seawater reaction vessel C/min)

Verney-Carron Li enriched synthetic 40– 8950 Pure water 0 0.7b 1200 8.3 90 sat. 0.5–96 Savillex(R) n.a. Fluid et al. (2011) basaltic glass powder 100 PTFE reactor

40– 8950 Pure water 0 0.7b 1200 8.3 50 sat. 6–48 Savillex(R) n.a. Fluid 100 PTFE reactor

40– 8950 Pure water 0 7.4, 0.7b 120, 8.3 90 sat. 24–1512 Savillex(R) n.a. Fluid 100 1200 PTFE reactor

40– 8950 0.001 M 0 0.7b 1200 3 90 sat. 2–816 Savillex(R) n.a. Fluid 100 HCl PTFE reactor

40– 8950 0.0001 M 0 0.7b 1200 10 90 sat. 2–816 Savillex(R) n.a. Fluid 100 NaOH PTFE reactor

Seyfried et al. Tholeitic basaltic glass <62 7 Seawater 190 10 5–7 150 50 3564 Au Dickson XRD Smectite, n.a. (1984) powder hydrothermal anhydrite apparatus

Fresh diabase powder <62 5 Na–K–Ca– 0 1, 2, 5 n.a. 375 40 1541– Ti Dickson XRD Smectite, n.a. Cl-fluid 2041 hydrothermal chlorite, apparatus actinolite, albite

Seyfried et al. Weathered basalt powder n.a. 8.2 Na–K–Ca– 025– 350 50 893 Au Flexible cell XRD Albite, Fluid Initially 7Li more (1998) (cryst.) Cl-fluid 5.5 hydrothermal Ca-zeolite, efficiently leached than chlorite, 6Li smectite

Berger et al. Tholeitic basaltic glass 135– 8.2 Synthetic 04c 25 4a 200 sat. 23 Pt Prolab XRD, Smectite, n.a. H depth profiles NRA (1987) powder 360 seawater hydrothermal EMPA, chlorite (lowered apparatus STEM Mg, Ca, and – c a fi 135 8.2sulfate 0 4 25 3.5 260 sat. 22 Pt Prolab XRD, Smectite n.a. H depth pro les NRA 360 conc.) hydrothermal EMPA, apparatus STEM

135– 8.2 0 4c 25 3– 320 sat. 20 Pt Prolab XRD, Smectite n.a. H depth profiles NRA 360 4a hydrothermal EMPA, apparatus STEM (continued) … Table 4.5 (continued) Interaction Fluid-Rock 4.4

Reference Starting materials Experimental conditions Product analysis Comments

Solid Grain Li Fluid Li S/V (1/cm) w/r pH T (° P Total Container Exp. setup Solid Secondary d7Li size (mg/g) (ng/g) surface C) (MPa) run material minerals measured (μm) area/volume duration identified of solution (h)

Berger et al. Li–Sc forsterite (syn.) 3–10 3250 Synthetic n.a. 5 – 4 260 sat. 17 Pt Prolab n.a. (1988) powder seawater hydrothermal (lowered apparatus Mg and sulfate conc.)

− Berger et al. Basaltic glass slab – 33,000 10 4 M HCl 0 1.50E−02 – 4 200 sat. 21 Pt Prolab XRD, Hydrated n.a. (1988) solution hydrothermal EMPA silicate gel apparatus − Berger et al. Basaltic glass slab – 33,000 10 4 M HCl 0 6.00E−04 – 4 300 sat. 3.5 Pt Prolab XRD, Hydrated n.a. (1988) solution hydrothermal EMPA silicate gel apparatus

James et al. Terrigeneous sediments – 40.8 Synthetic 10 – 4 4.3– 22– 40 742 Au Flexible cell n.a. – Fluid Stepwise heating (2003) (fine grained turbidites and seawater 7.2 350 hydrothermal hemipelagic muds) (lowered multi-phase (quartz, Mg and feldspars, lithic fragments sulfate (volcanic + mica) conc.)

James et al. Basaltic powder (hyapilitic – 4.3 Synthetic 21 – 4 5.3– 22– 40 737 Au Flexible cell n.a. – Fluid Stepwise heating (2003) texture, plag + cpx, 3 % ol, seawater 8.6 350 hydrothermal 3 % glass, 10 % opaque (lowered minerals) Mg and sulfate conc.)

Millot et al. JB-2 tholeitic basalt – 7.8 Seawater 100 – 10 n.a. 25 sat. 245 PFA Oven TEM, Unknown Fluid + solid (2010) powder IRMM beaker FTIR, BCR-403 XRD diluted to – – 7.860 % with 100 10 n.a. 75 sat. 245 PFA Oven TEM, Unknown Fluid + solid beaker FTIR, H2O XRD

– 7.8 100 – 10 n.a. 200 sat. 7 PTFE Oven TEM, Smectite Fluid + solid bomb FTIR, XRD

– 7.8 100 – 10 n.a. 200 sat. 12 Au Internally TEM, Smectite Fluid + solid heated FTIR, pressure XRD vessel

– 7.8 100 – 10 n.a. 250 sat. 7 Au Internally TEM, Smectite Fluid + solid heated FTIR, pressure XRD vessel (continued) 101 Table 4.5 (continued) Isotopic Associated and Diffusion Partitioning, Li 4 102

Reference Starting materials Experimental conditions Product analysis Comments

Solid Grain Li Fluid Li S/V (1/cm) w/r pH T (° P Total Container Exp. setup Solid Secondary d7Li size (mg/g) (ng/g) surface C) (MPa) run material minerals measured (μm) area/volume duration identified of solution (h)

Wimpenny Syn. basaltic glass (similar 40– 40– Synthetic, 0 BET Open 2–4 25, atm. 70 Ti Mixed ICP-MS, Unknown Fluid + solid Dissolution experiment No et al. (2010) to MORB) 100 100 HCl, ionic 23,000 cm2/ system 55 through-flow XRD fractionation strength g 1.5– reactors detected 0.01 mol/kg 4 g/min

San Carlos olivine 40– 40– Synthetic, 0 BET Open 2–4 25, atm. 80 Ti Mixed ICP-MS, Fluid + solid Dissolution experiment No 120 120 HCl, ionic 800 cm2/g system 55 through-flow XRD fractionation strength 1.5– reactors detected 0.01 mol/kg 4 g/min

San Carlos olivine 40– 40– Synthetic, 0 BET Open 10– 25, atm. 200 Ti Mixed ICP-MS, Chrysotile Fluid + solid Dissolution-precipitation Fluid 7Li 120 120 HCl, ionic 800 cm2/g system 11 75 through-flow XRD experiment enriched strength 1.5– reactors potentially due 0.01 mol/kg 4 g/min to preferential enrichment of 6Li into chrysotile

Pistiner and Basalt powder pellet – n.a. 0.8 M 0 – 33, 50 n.a. 22, atm 168 – n.a. Fluid Henderson (cryst.): Pl, Ol, Pyx HNCO3 45 (2003) microphenocrysts solution

Basalt powder pellet – n.a. 0.8 M 0 – 33, 50 n.a. 22, atm. 168 – n.a. Fluid (cryst.): large Ol HNCO3 45 phenocrysts, medium solution grained Pl, Pyx

Granite powder pellet, Qtz, – n.a. 0.8 M 0 – 33, 50 n.a. 22, atm. 168 – n.a. Fluid Ap, Pl, oxides, Chl, Ep, HNCO3 45 Cal solution

Chan et al. Shikoku Basin – 74 Synthetic 0 – 3 n.a. 25– 80 25.4 Au Flexible cell n.a. Fluid (1994) hemipelagic mudstone NaCl– 350 hydrothermal composed of Qtz, Pl, CaCl – 2 – – carbonate, and clay 74solution, 0 3 n.a. 25 80 22.5 Au Flexible cell n.a. Fluid Heating and cooling minerals near 350 hydrothermal stage seawater salinity

You and Sediment from the – 74 Synthetic 0 – 35– 25– 80 – Au Dickson type n.a. Fluid 6Li doped fluid Stepwise Gieskes (2001) decollemnet zone of ODP NaCl– 7.3 350 rocking heating and site 808, Nankai Trough, CaCl2 autoclave cooling Japan (Qtz, Pl, carbonate, solution, and clay minerals) near seawater salinity apH measured at 25 °C bBET surface area n.a. not analyzed cGeometric surface area … 4.4 Fluid-Rock Interaction 103

δ7 6Lit The effect of rock alteration on Li of the we obtain an equation for sol as a function of fl t uid and the remaining rock is in principle the the weight fraction of the reacted basalt, falt: result of mass balance, equilibrium fractionation  and, potentially, kinetic fractionation effects. In 6Li0 Á t þ 6Li0 Á w=r b falt sol 1 þ w=r the absence of isotopic fractionation and in the 6Lit ¼ ð4:4:2Þ sol w=r þ 6 Á t case of a mineral–fluid partition coefficient equal 1 þ w=r Kp falt to unity, δ7Li in the fluid can be described by a 7 simple mixing relationship between the fresh For Li we obtain an analogous equation rock and the fluid. In the absence of kinetic where we introduce in addition the isotopic δ7 7Lit 7Lit fractionation effects, the evolution of Li and Li a ¼ alt = sol fractionation factor, 6Lit 6Lit : in the fluid and the alteration product can be alt sol modeled from mass balance. During the reaction  7Li0 Á t þ 7Li0 Á w=r progress the solid can be separated into reacted b falt sol 1 þ w=r 7Lit ¼ ð4:4:3Þ and residual unaltered basalt and we obtain the sol w=r 6 t 1 þ = þ a Á K Á f following relationship between the initial and the w r p alt final Li concentrations from mass balance in a The amount of reacted basalt, f t , is controlled closed system: alt by the reaction kinetics and, depending on the 6 0 0 6 0 0 6 6 rate controlling mechanism, is a certain function Li Á f þ Li Á f ¼ Lit Á f t þ Lit b b sol sol b b sol of time. This simple model ignores the possibility Á f t þ 6Lit Á f t sol alt alt that the altered part of the basalt can be hetero- ð4:4:1Þ geneous, composed for example of different layers, as observed by Berger et al. (1987). 6Li0 6Li0 0 0 where b, sol, and fb , fsol are the concen- However, even in that case an effective partition 6 trations of Li in the starting basalt and the coefficient and isotopic fractionation factor can starting solution and their weight fractions, be defined, but it does not reflect thermodynamic 6Lit 6Lit 6Lit t t respectively, and b, sol, alt, and fb, fsol, equilibrium anymore and these parameters may t 6 t falt are the concentrations of Li in the unreacted thus be a function of time. The evolution of falt part of the basalt, the remaining solution, and the with t can be described, for example, by a rate reacted part of the basalt at time t, and their law, which is specific for particular operating weight fractions, respectively. The weight frac- reaction mechanisms and their rate determining 0 0 tions fb and fsol can be expressed as a function of step. Such a rate law containing various relevant w r 0 ¼ 1 parameters (e.g., fluid flux, grain size, tempera- the water to rock ratio, / : fb 1 þ w=r and = ture) would be necessary to extrapolate the f 0 ¼ w r . Assuming that the weight fraction sol 1 þ w=r experimental results for different scenarios. From fl fi of the uid stays constant (which is justi ed for basic kinetic principles it cannot be expected that large water/rock ratios in which the mass transfer the results from powder experiments can be fl of H2O out of the uid/water into hydrous min- applied in a straightforward way to natural rocks erals like smectite or chlorite can be ignored), the (see, for example, Rubie and Thompson 1985). concentration of Li in the unreacted part of basalt Using this simple model we can illustrate the stays constant, and bulk equilibration between effect of various system parameters on the reac- fl the alteration product and the remaining uid, tive flux of Li and δ7Li consistent with mass 6Lit 6 alt fi K ¼ t balance and equilibrium partitioning. In Fig. 4.17 described by a partition coef cient, p 6Li , sol we simulate the evolution of [Li] and δ7Li in the 104 4 Li Partitioning, Diffusion and Associated Isotopic …

(a) (b) 10000 35

Kp = 1, w/r = 1 30

1000 25

) Kp = 1, w/r = 10

L / 20

g (‰)

m Kp = 1 or 10,w/r = 100

( sol

l 15

o Li

s

i 100 7 δ

L Kp = 100,w/r = 100 10 5

10 0 0 0.20.4 0.6 0.8 1 0 0.20.4 0.6 0.8 1 Fraction of basalt altered Fraction of basalt altered (c) (d) 10 18 16 8 14

) 12

L

/ 6

g 10 (‰)

m

( WR 8 R 4 Li

W

i 7 δ

L 6 2 4 2 0 0 0 0.20.4 0.6 0.8 1 0 0.20.4 0.6 0.8 1 Fraction of basalt altered Fraction of basalt altered Fig. 4.17 Illustration of the effect of the partition line in panel (a). The format of corresponding lines in coefficient, Kp, and the water to rock ratio, w/r, on the panel (b), (c), and (d) is the same as in panel (a). In all reaction path for Li and δ7Li in the solution and the bulk cases an isotopic fractionation factor of α = 0.995 and the rock (altered + fresh basalt) as a function of the fraction of same initial Li and δ7Li of fluid and basalt were chosen: altered rock from the bulk rock. This simulation assumes 100 μg/L and 7.8 μg/g, and δ7Li = +31 ‰ and bulk equilibration between the altered rock and the δ7Li = +4.8 ‰ solution. The values for Kp and w/r are indicated at each

fluid and the bulk rock with reaction progress respect to Li that, during reaction and after assuming different values for Kp and w/r and re-partitioning, Li is still enriched in the fluid. constant isotopic fractionation factor α = 0.995. The pronounced loss of Li into the fluid with This latter value applies to smectite–fluid parti- increasing temperature observed in experiments tioning at a temperature of *175 °C (Vigier could be related to the enhanced reaction pro- et al. 2008). According to this simple model four gress due to thermally activated kinetic processes factors control the terminal enrichment of Li in (i.e., diffusion). For example, the thickness of the the fluid: the partition coefficient Kp, the reaction altered rim increases with time and temperature, t progress, characterized here by falt, the initial which could be explained by the strong temper- concentrations of Li in the fluid and the fresh ature dependence of the diffusion processes. basalt. We can conclude that even though Li is However, from a general point of view Li is not compatible in smectite, the starting fluidinthe strongly fluid mobile since the alteration prod- experiments, similar to seawater, is so dilute with ucts retain most of the Li as inferred from 4.4 Fluid-Rock Interaction 105

(a) 1995). Here the mass balance is applied to each SEAWATER infinitesimal reaction step and after integration 30 0.02 one obtains an equation for the concentration of 0.05 6Li in solution: 0.1 20  6Lit ¼ 6Li0 À6 Li0=K6 0.3 ALTERATION sol solb p Li (‰) 0.5 0.05 PHASE 7 0.7 6 r t 6 0 6 δ 10 Á exp ÀK Á 1 þ Á f þ Li =K 0.1 p w alt b p BASALT Millot et al. (2010) ð4:4:4Þ 0.3 75°C 0 0.7 0.5 250°C 0.7 For 7Li we obtain an analogous equation: 0 0.002 0.004 0.006 0.008 0.010  1/Li 7 7 0 7 0 6 (b) Lit ¼ Li À Li =ða Á K Þ sol sol b p SEAWATER r 30 Á Àða Á 6ÞÁ 1 þ Á t exp Kp falt ALTERATION 0.02 w PHASE þ 7Li0=ða Á 6Þ 0.05 b Kp 20 0.1 ð4:4:5Þ Li (‰) 7

δ 10 0.7 0.3 0.5 where the same nomenclature as before is used. 1.0 The same kind of simulations as in Fig. 4.17 BASALT Millot et al. (2010) 75°C fi 0 250°C using Eqs. 4.4.4 and 4.4.5 show no signi cant difference between the Rayleigh model and the 0 0.002 0.004 0.006 0.008 0.010 bulk equilibrium model. Only for high 1/Li water/rock ratios (>10) marginal differences are Fig. 4.18 Modeled lithium isotopic compositions of the observed when the reaction has proceeded to fluid, alteration phase (smectite), and bulk rock (mixture more than 50 %. of altered + fresh basalt) with water to rock ratio of 10 for The potential effect of temperature was sim- different stages of the alteration process at a 75 °C with ulated using Eqs. 4.4.4 and 4.4.5 where we have Kp =1,α = 0.99, and b 250 °C with Kp = 18, α = 0.998, according to the experimental partitioning data of Decar- used the experimental calibrations for the equi- reau et al. (2012) and Vigier et al. (2008). Open librium partition coefficient for smectite–fluid of diamonds = fluid; open circles = bulk solid; open Decarreau et al. (2012) and the corresponding triangles = alteration phase; dashed lines = fluid (blue) or bulk solid (black) composition for a partition coeffi- isotopic fractionation factor of Vigier et al. 7 cient and isotopic fractionation factor α =1;open (2008). The reaction path for δ Li of the fluid, hexagons = data of Millot et al. (2010) from alteration alteration phase (smectite), and bulk rock (mix- experiments. Numbers next to the data points indicate the ture of altered + fresh basalt) is illustrated in fraction of basalt altered. Note that the reaction path is strongly dependent on the partition coefficient and Fig. 4.18. The isotopic fractionation factor and isotopic fractionation factor, as also illustrated in Fig. 4.17 the partition coefficient are important variables driving the evolution of δ7Li in the fluid and the mineral–fluid partition coefficients of typical rock apart from the simple mixing line. In all the alteration products (see Sect. 4.2.4). simulated cases δ7Li of the fluid is very sensitive The above model assumes bulk equilibration to basalt alteration and decreases with increasing of the alteration product and the fluid, which, reaction progress. The bulk solid, though, hardly especially at low temperature, may be not ful- changes unless, for example, at least 50 % of the filled. Alternatively, a Rayleigh fractionation basalt has been altered for the case simulated in model can be applied assuming successive for- Fig. 4.18a. This drastically different sensitivity to mation of alteration products (Magenheim et al. alteration of fluid compared to rock is related to 106 4 Li Partitioning, Diffusion and Associated Isotopic … the about 100 times larger concentration of Li in compared to 7Li in the fluid, compensating the basalt compared to seawater, the moderate slightly for the equilibrium fractionation effect. water to rock ratio, and the small contribution of Apparent smectite–fluid partition coefficients for t fl the altered part of the rock, falt, to the bulk rock Li calculated from Li contents in uid and composition for low extents of alteration. For smectite found in the alteration experiments of higher water to rock ratios and large partition Berger et al. (1987) are on the order of 10 and coefficients, a smaller degree of alteration is 800 at 320 and 200 °C, respectively, which is, at required to affect δ7Li in the rock, as can be also least for 320 °C, similar to the equilibrium par- seen from Fig. 4.17. The values for δ7Li deviate tition coefficients measured by Decarreau et al. from the simple mixing line depending on Kp, α, (2012) for Mg-rich smectite. However, the sig- and reaction progress characterized here by the, nificantly larger apparent partition coefficient t ° falt, the fraction of rock altered. Interestingly, for observed at 200 C further points to disequilib- 7 Kp = 1 and α = 0.99, δ Li of the altered basalt can rium partitioning with decreasing temperature. become lower as a consequence of high/low Another observation of Millot et al. (2010) water to rock ratio and the preference of clay was that δ7Li in the fluid ultimately increases minerals to incorporate light Li (Fig. 4.18a). For with decreasing temperature (compare also with higher partition coefficients the altered rock Fig. 4 in Millot et al. 2010). These authors typically becomes heavier. Thus, the effect of argued that a steady state is reached where dis- temperature according to the experimental parti- solution processes (or leaching) and tioning data is two-fold: equilibrium isotopic re-precipitation of smectite are balanced. They fractionation tends to increase with decreasing obtained an apparent trend of Δ7Li with smec- temperature, whereas the behavior of Kp is the tite–fluid pair as a function of T, which was used opposite. In summary, this results in very dif- as a thermometer for hydrothermal alteration of ferent possible scenarios for δ7Li of the altered basalt. However, based on the above model the rock dependent on temperature and the water to variation of this “steady state” δ7Li with tem- rock ratio. perature could be simply explained by a smaller Data for fluid compositions from the alteration amount of basalt transformed into clay minerals experiments of Millot et al. (2010) are broadly due to more sluggish kinetics of the transforma- consistent with the above modeling results and tion process. Based on our calculations we would plot close to the simple mixing line. However, argue that this trend is specific for the experi- the data measured at 75 °C seem to be more mental conditions (grain size and water to rock consistent with the predictions for 250 °C and ratio) given in Millot et al. (2010) and cannot be vice versa. These slight discrepancies could be applied in general. We further conclude that the explained by the uncertainty of the experimental “steady state” isotopic composition of the fluid in data, but kinetic effects potentially may also play Millot et al. (2010) does not reflect equilibrium a role here, because these may produce an fractionation between solids and fluid and δ7Li of effective partition coefficient and isotopic frac- the fluid would likely evolve further towards tionation factor different from those applicable at equilibrium if the experiments would be exten- thermodynamic equilibrium. For example, as ded for sufficiently long times. Interestingly, the 7 discussed above, at low temperatures diffusion in T dependence of the apparent Δ Lismectite–fluid of the fluid may not be efficient enough to homog- Millot et al. (2010) is similar to the equilibrium enize during secondary mineral precipitation and fractionation measured by Vigier et al. (2008) for apparently higher or smaller partition coefficients T >90°C, but the total extent of Li isotopic are obtained. At higher temperatures the min- fractionation is stronger in the former case, eral–fluid partitioning may be closer to equilib- which is probably related to the dominant con- rium but Li diffusion out of the glass becomes tribution of the unaltered basalt to δ7Li in the more efficient and enhances the levels of 6Li bulk sample (Fig. 4.19). 4.4 Fluid-Rock Interaction 107

Temperature (°C) concentration profiles in the altered glass and 300 250200 150 100 50 25 4 adjacent plagioclase (as in Berger et al. 1987, 1988). 0 The major element composition of differenti-

-4 ated glasses may also affect the kinetics of the leaching process. If chemical diffusion is -8 i responsible for the enhanced reaction rate it

L

7 Δ should also produce associated isotopic frac- -12 tionation effects. Since 6Li diffuses slightly faster Vigier et al. (2008) 7 -16 (see Sect. 4.3) than Li out of plagioclase, the Millot et al. (2010) plagioclase and, by inference, the bulk rock -20 4681012should become isotopically heavier. This effect 1/T2 was modeled by Brant et al. (2012) assuming a Fig. 4.19 Comparison of equilibrium fractionation of Li plane-sheet geometry. They concluded that in the fl isotopes between smectite and uid of Vigier et al. (2008) ° δ7 with apparent Δ7Li between bulk rock and fluid from temperature range between 350 and 450 C Li alteration experiments by Millot et al. (2010) of basalt could increase by up to 8.7 ‰ within a few hundred years, but for significantly longer times the Li content in plagioclase becomes very As mentioned above, kinetic fractionation small and homogenous. Therefore, this fraction- effects, related to diffusion in the fluid, gel, glass, ation effect would eventually be erased and δ7Li plagioclase, etc., or due to surface exchange of the basalt would become dominated by reactions, may have an additional impact on δ7Li clinoproxene (see also Fig. 11c in Brant et al. in the produced phases. Two experimental stud- 2012). ies were dedicated to potential effects of Li dif- In the experimental study of Verney-Carron fusion in solid materials (Verney-Carron et al. et al. (2011), leaching of Li out of synthetic 2011; Brant et al. 2012). Brant et al. (2012) basaltic glass was experimentally investigated performed experiments at 125 °C with water to and discussed in terms of associated kinetic rock ratios of about 9 using crystalline and fractionation effects. In these experiments syn- amorphous material and found surprisingly that thetic glass powder doped with 1 % Li2O was Li is extracted more efficiently from the crys- mixed with Li-free pure water at very high talline material. This observation seems to be in water/rock ratios (about 1000) and annealed at 50 contradiction with that of Seyfried et al. (1984), and 90 °C for 15 min to 63 days. The main who found that diabase is less reactive compared observations of Verney-Carron et al. (2011) were to basaltic glass at 375 °C and slightly lower twofold: (i) Li was dissolved at a higher rate water to rock ratios. The efficient leaching of Li compared to Si, and (ii) δ7Li of the reacted water out of partly crystallized basalt in Brant et al. was lower by up to 5 ‰ compared to the syn- (2012) was explained by diffusion of Li out of thetic basalt at the beginning of the experiment plagioclase although approximate calculations of and with increasing time approached the value of the diffusive fluxes show that two orders of the fresh basaltic glass. Both observations are magnitude higher diffusion rates than those consistent with observations made by Wimpenny experimentally observed by Giletti and Shanahan et al. (2010) in their dissolution experiments and (1997) would be required. Importantly, Li diffu- are indicative of diffusive loss of Li, whereby 6Li sion is supposed to be significantly faster in sil- was leached more efficiently than 7Li. This pro- icate glasses than in plagioclase at 125 °C cess was modeled by Verney-Carron et al. (2011) (Fig. 4.8) and, as a consequence, leaching of Li as a combination of diffusion of Li in the glass out of basaltic glass should be more efficient than and dissolution of the leached layer. The latter leaching out of plagioclase. This problem cannot process imparts a counter effect to δ7Li compared be resolved without determining the Li to the pure diffusion process, since the leached 108 4 Li Partitioning, Diffusion and Associated Isotopic … layer is relatively enriched in 7Li because of the was not analyzed after the experiment was ter- former diffusion process scavenging 6Li. With minated. Nevertheless, with increasing tempera- this combined diffusion–dissolution model ture it can be expected that alteration proceeds Verney-Carron et al. (2011) were able to repro- more efficiently and the effective partition coef- duce the measured evolution of δ7Li and ficient between the fluid and the alteration phases obtained from these fittings the relative diffusion must correspondingly decrease to support the coefficients for both Li isotopes in the basaltic enrichment of Li in the fluid. Such a behavior glass, which apparently differed by only 0.02 to could be explained by metastable formation of 0.1 % at the experimental conditions. They also secondary minerals at lower temperatures, as also concluded from these experiments that dissolu- indicated by the apparent partition coefficient tion rates became more important than diffusion found in Berger et al. (1987). at higher temperatures and eventually diffusion in Since weathered basalt (at T < 150 °C) retains basaltic glass imparted a minor effect to δ7Li of Li in smectite (and Cs preferentially over Rb) and natural systems. the basaltic glass is already leached with respect In the experimental study of James et al. to Li (and Cs, Rb), it was speculated by Seyfried (2003) mostly crystalline basaltic powder was et al. (1998), based on observed Cs/Rb, that altered at water/rock ratio of 4 stepwise from 51 continuing alteration of the basalt at higher to 350 °C. The starting solution was prepared temperatures neither released Li into the fluid nor such that it resembles fluids estimated to, affected its isotopic composition. Therefore, “recharge sediment and basalt-dominated Seyfried et al. (1998) investigated the alteration hydrothermal systems at mid-ocean ridges.” of weathered basalt at hydrothermal conditions This solution lacked Mg and sulfate and as a (350 °C, 500 bars, fluid/rock ratio = 2). The consequence the formation of smectite or chlorite basalt was considered to be altered based on trace were potentially inhibited unless Mg would be element contents (e.g., high Cs/Rb ratio) leached from the basalt as the first major cation. although X-ray diffraction analysis of the starting James et al. (2003) observed that after a moderate material did not identify alteration phases like decrease of Li content in the fluid at T = 150 °C smectite but only crystalline phases typical of from a starting value of 381 μg/L, its concentra- unaltered basalt, mainly plagioclase and tion increased continuously up to 790 μg/L at 350 clinopyroxene. In analogy to fresh basalt, a sig- °C. During stepwise cooling the Li concentration nificant amount of Li was leached from the of the fluid remained approximately constant. weathered basalt and Li isotopes were affected in These observations are generally consistent with a complex manner, with increasing δ7Li in the the experimental study of Seyfried et al. (1984). solution in the early stage and a subsequent In parallel to Li abundance, δ7Li first increased decrease in a later stage of the experiment (about during heating from room temperature to 150°C 1 week). This observation is consistent with that by about 7 ‰ from a starting value of −5.5 ‰ of James et al. (2003), where step-wise heating (the starting basalt had δ7Li = +4.6 ‰), and from 150 °C to higher temperatures also led to decreased from +1.4 to −0.6 ‰ until the peak the enrichment of Li into the fluid. temperature of 350 °C was reached. The initial Two studies were specifically designed to loss of Li from the fluid indicated uptake of Li by measure dissolution rates of Li from basaltic secondary mineral formation. According to the glass (Pistiner and Henderson 2003; Wimpenny mass balance model above, for the experimental et al. 2003). Pistiner and Henderson (2003) used conditions of James et al. (2003) an effective two experimental approaches: (i) Powdered partition coefficient of greater than about 100 is crystalline basalt or granite were altered at required to create this loss assuming a very low pH = 0.1 at 25 and 45 °C at water to rock ratios degree of alteration (few %). The secondary of 33 and 50 for 1 week. (ii) Surface exchange minerals formed during the experiments of James experiments at 22 °C where mineral separates of et al. (2003) are unknown since the solid product smectite, gibbsite and ferrihydrite were stirred in 4.4 Fluid-Rock Interaction 109 either pure water (for 24 h) or a solution (NIST Li isotopic data from these kinds of experiments. L-SVEC) with 0.5 mg/L Li and pH = 7 for Clearly, more carefully designed experiments are various times. The general idea of these two needed to understand and quantify the effects of setups was to separate the effect of the two granite alteration on δ7Li. competing mechanisms of rock alteration, dis- Variable fractionation effects for Li isotopes solution and re-precipitation (note that transport were found in the surface exchange experiments: through the reacted layer is an additional relevant strong for gibbsite (up to 12 ‰), weak for fer- mechanism, as indicated by Berger et al. 1987, ryhydrite (2 ‰), and absent for smectite. This 1988). The acidic composition in the alteration fractionation during physisorption must be experiment was chosen to prevent the precipita- clearly distinguished from equilibrium fraction- tion of new phases, although the water to rock ation between the bulk mineral and fluid as for ratio was similar to other alteration experiments example investigated by Vigier et al. (2008) for (e.g., Berger et al. 1987). Neither the Li con- smectites. Thus, these experiments cannot centrations in the reaction products were reported reproduce the behavior of Li isotopes during nor secondary minerals identified, but δ7Li in the precipitation of alteration phases. product solutions was determined. The partial Compared to the study of Pistiner and Hen- dissolution of basalt produced δ7Li in the fluid derson (2003), the experimental design of Wim- similar to the starting basalt, but the fluid penny et al. (2010) allowed a clearer separation of obtained from granite dissolution was lighter by the effects of dissolution from secondary mineral about 8 ‰ compared to the starting composition formation. In these experiments basaltic glass of the solid reactant. The result for basalt is in powder was placed in flow-through reactors with apparent contradiction to Verney-Carron et al. flow rates between 1 and 8 g/min at 25–55 °C and (2011), where at relatively similar experimental low pH (2–4). Dissolution was incongruent with conditions the fluid became typically lighter. In higher loss of Li, Mg, and Al compared to ele- the latter study it has also been shown that with ments like Na or Ca, which were almost increasing time δ7Li in the fluid approached that unchanged. Dissolution rates of the basalt were of the fresh basalt and this occurred faster for calculated based on the dissolved Si concentra- more acidic compositions. Thus, for the very tion. The concentrations of Li varied depending acidic conditions in Pistiner and Henderson on pH between 0.03 μg/L and 0.5 μg/L (higher (2003) it is reasonable to expect that the basalt concentrations for lower pH corresponding to dissolves more congruently and kinetic fraction- higher dissolution rates). No change of δ7Li ation effects are irrelevant in the leaching of Li (+4.7 ‰) was detected for pH of 4, but for lower from basalt. In addition, in the study of Pistiner pH a slight increase by up to 1.5 ‰ was observed. and Henderson (2003) mostly crystalline basalt The δ7Li in the fluid was slightly lower for all pH was dissolved and it is unclear if leaching of Li values and varied between +2 and +3 ‰. The out of crystalline phases occurred as suggested enrichment of 6Li in the fluid could be explained for plagioclase in Brant et al. (2012). by its faster diffusion rate in the leached layer of To explain the heavier isotopic composition in the basaltic glass compared to 7Li, as also the fluid from granite dissolution, Pistiner and demonstrated by Verney-Carron et al. (2011). Henderson (2003) suggested that pre-existing secondary minerals in the starting material (chlorite, epidote, and calcite) are preferentially 4.4.2 Weathering of Olivine dissolved from the granite. For example chlorite (Serpentinization) has a high Li concentration and, probably, carries a light Li isotopic signature, according to its In addition to the dissolution of basalt, Wim- equilibrium fractionation behavior (see Fig. 4.5). penny et al. (2010) investigated the weathering of These early endeavors preceded a fuller realiza- San Carlos olivine by two different approaches in tion of the potential complexities in interpreting the flow-through reactor to separate the effects of 110 4 Li Partitioning, Diffusion and Associated Isotopic … dissolution and precipitation. Olivine dissolution composed of quartz, plagioclase, carbonate, and experiments used similar experimental condi- clay minerals (illite, smectite, kaolinite, and tions (low pH) as for the dissolution of basalt. To chlorite). The mudstone was reacted with a allow re-precipitation of secondary minerals, a synthetic NaCl–CaCl2 solution devoid of Mg and set of experiments was performed at high pH near seawater salinity at temperatures from 25 to values (10–11). After the dissolution experiment 350 °C, P = 80 MPa, and a starting fluid/rock δ7Li in the olivine was about +2 ‰, on average ratio of 3. Two sets of experiments were per- heavier compared to the starting value of +1.5 ‰ formed with stepwise heating and cooling stages but similar within broader uncertainties. aimed to separate the effect of heating and During the precipitation experiments at high cooling on the Li exchange process. Similar to pH, chrysotile formed as indicated by XRD the weathering of basalt, Li was progressively measurements of the solid products. Related to lost to the fluid with increasing time and tem- the successive formation of secondary minerals, perature. During cooling the concentration of Li δ7Li in the fluid increased with time from about remained stable, indicating irreversible exchange +2 ‰ up to +12 ‰. According to the experi- of Li between the sediment and the fluid. The mental study of Wunder et al. (2010) the ser- final Li content of the solid was smaller in this pentine phases antigorite and lizardite type of experiment compared to the heating preferentially incorporate 6Li but chrysotile experiment, consistent with higher Li concen- slightly prefers 7Li because of the lower coordi- trations in the final fluid. The concentration of Li nation of Li incorporated in nano pores (see in the fluid was for both types of experiments discussion in Sect. 4.2 and Fig. 4.5). Apparently, significantly higher than in experiments of basalt in the open system and at the low temperature alteration (on the order of 30,000 μg/L vs. typi- conditions (25 and 75 °C) of Wimpenny et al. cally 100–1,000 μg/L), which was partly related (2010), chrysotile shows a different behavior to the low water/rock ratio but also to the higher compared to the closed system and 400 °Cas initial Li content of the sediment compared to used in Wunder et al. (2010). fresh basalt (75 μg/g vs. typically 5–8 μg/g). Berger et al. (1988) altered polycrystalline In the heating experiment the Li isotopic synthetic Li–Sc-forsterite in seawater devoid of composition of the first analyzed fluid after ca. 2+ 2− Mg and SO4 at relatively higher temperatures 14 days at 97 °C was heavier by 4 ‰ than that of (150–260 °C). A strong incongruent dissolution the initial sediment and approached the sediment was observed at these conditions where the composition with increasing T and time, corre- cations Mg2+ and Li+ were dissolved preferen- sponding to enhanced reaction progress. Inter- tially compared to silica. Based on equilibrium estingly, in the cooling experiment the fluid calculations Berger et al. (1988) argued that no started with slightly isotopically heavier Li, and secondary phases were formed and hence they became lighter at 350 °C, similar to the starting concluded that Li and Mg were leached from solid, while during the cooling it became heavier forsterite by diffusive processes. This statement again by up to 5 ‰. These observations indicate remains speculative in the absence of analysis of that while the net transfer of Li into the fluid the reacted forsterite. Congruent dissolution of effectively ceased during the longer cooling stage, forsterite with re-precipitation of silica gel would Li isotopes still exchanged between the fluid and be an alternative hypothesis. the solids and 6Li became enriched in the solid with decreasing T, consistent with equilibrium fractionation data for clay minerals (Fig. 4.5). 4.4.3 Alteration of Sediments In similar types of experiments at low water/rock ratios, hemipelagic sediments (You Chan et al. (1994) studied the Li isotopic geo- et al. 1996), diatomaceous ooze (Thornton and chemistry of marine sediments through alteration Seyfried 1987), and unaltered marine sediment experiments with hemipelagic mudstone (Von Damm 1991) were altered and Li released 4.4 Fluid-Rock Interaction 111 into the fluid at similar, but slightly lower con- experimental studies the solid reaction products centrations than in experiments of Chan et al. were not investigated and secondary minerals not (1994). These differences can be explained by the identified. Based on the evolution of the fluid smaller initial Li contents of the solid material. In chemistry You and Gieskes (2001) speculated two experiments, You and Gieskes (2001) altered about the reactions occurring with different hemi-pelagic sediment with similar mineral components of sediments, decomposition of assemblage as in Chan et al. (1994). The fluid organic matter, formation of secondary minerals, composition was also similar to that used by recrystallization of plagioclase, dissolution of Chan et al. (1994), but in one experiment the silica, and retrograde alteration. Based on the starting fluid was heavily enriched with 6Li. As evidence from basalt alteration experiments, the in Chan et al. (1994), the reaction system was most relevant process for Li should be repre- heated stepwise to 350 °C and cooled back to sented by the formation of secondary minerals, in room temperature. The observations on the particular clay minerals. In addition, exchange behavior of Li in the undoped experiment were with or recrystallization of the initially present exactly the same as in Chan et al. (1994). In case clay minerals should be important since they of the Li-enriched fluid, during the heating stage should control the Li budget of the bulk rock. up to 100 °C no change in the Li concentration The isotopic exchange observed during the was apparent, followed by a depletion of Li and cooling experiment of Chan et al. (1994) indi- apparent uptake of Li into secondary minerals up cates that Li exchange between clay minerals and to 250 °C. During heating to the final tempera- fluid is a very efficient process even at moderate ture Li was released again to the fluid and the temperatures. concentration remained stable during cooling. During heating the Li isotopic compositions approached those of the sediments and also 4.4.4 Synopsis remained constant during cooling. A terrigenous sediment, a mixture of Major elements in bulk rocks and fluids control fine-grained turbidites and hemipelagic muds the type of secondary minerals formed during (29 % quartz, 24 % feldspars, 34 % lithic frag- aqueous alteration. The behavior of Li is mainly ments), was reacted with synthetic fluid (devoid controlled by the formation of secondary miner- of sulfate ions) by James et al. (2003)at als. However, kinetic fractionation associated fluid/solid ratio of 4 and by applying stepwise with leaching processes out of the rock through heating up to 350 °C. Again, Li was progressively diffusion from “fast” diffusion media (e.g., glass released to the fluid with increasing time and or plagioclase) needs to be considered as well. temperature. Consistent with other alteration The contribution of leaching appears to be only experiments, δ7Li of the fluid increased (from important in the early stages of the reaction and −5.5 to +6.6 ‰) in the early stages of the at lower temperatures. The dissolution process by experiment at 50 °C and finally yielded a value itself (transition of Li from the surface into the 2 ‰ lower than that of the unaltered sediment solution) seems to produce no significant isotopic (+1.3 ‰) at the final temperature. This result fractionation for Li. However, especially at lower would be consistent with the observations that at temperatures, non-equilibrium partitioning lower temperature 6Li was taken up preferentially between secondary minerals and fluid is relevant, by secondary minerals, but with increasing T this which may lead to large apparent mineral–fluid preference decreases, as documented above. partition coefficients and, from mass balance, Although the general picture for the behavior significant effects on δ7Li of the bulk rock. In any of Li during the alteration of sediments of vari- case, even for large water/rock ratios, a large ous types and mineralogical composition seems fraction of the solid needs to be altered to change to be consistent, it is unclear by exactly which its δ7Li, whereas the fluid is universally very process Li exchange is controlled. In all of these sensitive to alteration. Depending on the different 112 4 Li Partitioning, Diffusion and Associated Isotopic … contributions of leaching, bulk dissolution and water: experimental investigations and modelization precipitation to the overall alteration process and between 50 °C and 300 °C. Chem. Geol. 71:297–312 Berger G, Claparols C, Guy C, Daux V (1994) Dissolu- the types of secondary minerals formed, very tion rate of a basalt glass in silica-rich solutions: different scenarios are possible for the evolution implications for long-term alteration. Geochim Cos- of δ7Li in the fluid and the altered rock. At the mochim Acta 58:4875–4886 moment no kinetic model exists that fully Bindeman I, Davis A (2000) Trace element partitioning between plagioclase and melt: investigation of dopant describes, for example, the alteration of basalt as influence on partition behaviour. Geochim Cos- a function of various relevant input parameters mochim Acta 64:2863–2878 (fluid flux, grain size, input fluid chemistry). 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5.1 Lithium in the Mantle peridotite xenoliths were fairly sparse and thus detailed estimates of Li systematics in Earth’s mantle were quite uncertain. This data deficit was It has been clear for some time that Earth’supper mainly due to analytical challenges presented by mantle is Li-poor relative to the bulk continental high-Mg, low-Li materials, although more ques- crust (<2 ppm compared to *18 ppm; Jagoutz et al. tions have arisen since that barrier has been bro- 1979; Ryan and Langmuir 1987;Tengetal.2008). ken. Ryan and Langmuir (1987) provided It has become similarly manifest that isotopic con- preliminary data for Li contents in coexisting trasts exist between surface and deep Earth reser- mineral phases of several mantle nodules that voirs (e.g., altered ocean crust with heavy isotopic showed a dominant control of olivine on Li dis- signatures compared to mid-ocean ridge basalts; tribution in spinel peridotites, although it may Chan and Edmond 1988). One of the great hopes for differ for garnet-bearing peridotites. This was Li isotope studies in the deep Earth has been that the later validated by Seitz and Woodland (2000) and system could unlock some of the questions related Eggins et al. (1998) for a larger suite of mantle to mantle evolution and quantify aspects of xenoliths, but these authors also showed domi- crust-mantle cycling (e.g., Elliott et al. 2004). nant control exerted by clinopyroxene in many Although this hope may ultimately be realized, a lithologies, such as pyroxenites. This finding in number of roadblocks have slowed progress toward itself may imply differences between pristine and it considerably. There are two major ways that can modified lithologies governed by post-magmatic be paved in constraining Li in the mantle. Either processes. At equilibrium, olivine appears to be pristine peridotites can be used or melt derivatives the major host of Li (Ottolini et al. 2004) followed of the mantle can used as proxies, considering the by near-equal distribution between clinopyroxene lack of significant Li isotopic fractionation during and orthopyroxene (Dolivine/melt *0.3–0.45 > mantle differentiation (Tomascak et al. 1999; Chan Dcpx/melt ≈ Dopx/melt *0.15–0.3; Brenan et al. and Frey 2003). Both approaches have positive and 1998a, b; Ottolini et al. 2009; Chap. 4); both negative aspects, however. spinel and garnet appear to generally contribute little to the total Li budget of mantle xenoliths and 5.1.1 Mantle Xenoliths and Related hence to the mantle itself (Seitz and Woodland Samples 2000; Woodland et al. 2002; Ottolini et al. 2004), although in opportune settings garnet may carry fi The review by Tomascak (2004) noted that the signi cant loads of Li (Barry et al. 2015). then-available data for Li abundance and isotopic Experimentally determined partitioning of Li composition in pristine, unmetasomatized suggests some pressure-dependence, at least for clinopyroxene (McDade et al. 2003), but more

P.B. Tomascak et al., Advances in Lithium Isotope Geochemistry, 119 Advances in Isotope Geochemistry, DOI: 10.1007/978-3-319-01430-2_5 © Springer International Publishing Switzerland 2016 120 5 Lithium in the Deep Earth: Mantle and Crustal Systems focused examination would certainly provide variations in δ7Li of bulk peridotites beyond clearer constraints. The implications of these analytical uncertainties. The interpretations of observations are twofold: (i) one can broadly Bouman et al. (2000) were tied to distinct ages of estimate the Li abundance of the planet in bulk different peridotites and temporal evolution of with a rough understanding of the modal pro- the mantle with respect to Li isotopic composi- portions and Li contents of only a few minerals tions. One additional peridotite analysis pub- (Foley et al. 2013), although an incomplete lished shortly afterwards (Chan et al. 2002b) was understanding of lower mantle Li has an uncer- consistent with the δ7Li * +5 ‰ for the Earth’s tain impact on such estimates; (ii) the absence of a upper mantle. It soon was realized that secondary unique mineral that would, at mantle conditions, processes such as serpentinization of oceanic represent the bulk of Li (such as garnet does for peridotites (Decitre et al. 2002) or metasomatism lanthanoids, for example) may sometimes hamper through melt extraction and overprinting by a more detailed assessment of a particular process hydrous fluids (Brooker et al. 2004) may signif- such as fractional crystallization of a specific icantly alter intrinsic δ7Li signatures. Moreover, mineral phase or more general conditions leading kinetic processes through sub-solidus diffusive to formation of such phases. fractionation were soon revealed to be capable of The Li elemental and isotopic variability in introducing further complexities to the original mantle assemblages was confirmed by pilot Li isotopic signatures (Lundstrom et al. 2005). studies of mantle xenoliths (Bouman et al. 2000; Because Li substitutes for Mg in major mantle Brooker et al. 2000) which clearly showed minerals such as olivine and pyroxene, it appears

Fig. 5.1 Lithium concentration and δ7Li of bulk mantle and Pogge von Strandmann et al. (2011), Zabargad from nodules, including peridotites (predominantly) with some Brooker et al. (2004), S. Korea from Kil (2010), San pyroxenites. Compilation includes data from both pristine Carlos (New Mexico) from Seitz et al. (2004), Russia samples and those affected by secondary processes, as (various localities) from Halama et al. (2009), Pogge von described in the text. Data sources: Bohemian Massif Strandmann et al. (2011), and Ionov and Seitz (2008) and from Ackerman et al. (2013), Gakkel Ridge from Gao Horoman from Lai et al. (2015) et al. (2011), Mongolia from Magna et al. (2006a, 2008) 5.1 Lithium in the Mantle 121 that a non-negligible fraction of Li could be hosted from the mantle, brought to the surface by in the lower mantle where high-pressure poly- erupted host lavas. Rapid transport from depths morphs of olivine and pyroxene, as well as garnet between *30 and *200 km often inhibits and other Mg-dominated silicates and oxides (sili- incipient low-degree partial melting which could cate perovskite now known as bridgmanite, and modify incompatible element contents. Criti- magnesiowüstite), constitute the majority. To date cally, large mantle xenoliths (commonly ≥0.5 m only the studies of Kaminsky et al. (2001) and Seitz in diameter) may successfully preserve mantle et al. (2003) have reported Li abundances in these signatures despite their fragility, whilst small uncommon minerals. It was clear, however, that peridotite xenoliths (<0.1 m in diameter) fre- there are significant differences among co-existing quently are affected by secondary processes. Ca-perovskite, Mg-perovskite and magne- Alternatively, massive alpine-type orogenic siowüstite (a.k.a. ferropericlase), ranging from peridotites can be used because they, at least in sub-ppm levels in the former species to *50 ppm in part, may have escaped interactions with melts. the latter. Whether or not this enrichment relates to a Nonetheless, pristine examples of such peri- particular Li-enriched source or hints at a yet dotites are rare (e.g., Horoman peridotite, Japan; unaccounted Li refertilization process remains to be Lai et al. 2015). answered by future Li isotopic analyses. A prelim- Magna et al. (2006a) and Jeffcoate et al. inary report on diamond-hosted ferropericlase (2007) have inferred that equilibrium fractiona- inclusions by Seitz et al. (2006)gavehighlyvari- tion translates into similar δ7Li values in able δ7Li (*−4to*+10 ‰). This may imply co-existing olivine and orthopyroxene whereas entrainment of sedimentary and eclogitic litholo- clinopyroxene commonly carries a heavier Li gies to the upper/lower mantle boundary through isotopic signature. A number of studies supple- subduction. In any case, the Earth’s lower mantle mented the early data set for pristine spinel may appear distinctly enriched in Li which could lherzolites (Brooker et al. 2004; Seitz et al. 2004; reflect pressure- (but not temperature-) dependent Magna et al. 2006a, 2008; Jeffcoate et al. 2007; behavior of Li coupled to complex substitutions Pogge von Strandmann et al. 2011) and collec- involving Li, Si, Mg and Al (Seitz et al. 2003; Yang tively have illustrated the Li elemental and iso- et al. 2009) and, presumably, Fe3+. At least for topic homogeneity of fertile upper mantle with a lower region of the upper mantle and the transition narrow range of δ7Li values between *+2.5 and zone, Li solubility increases in parallel with ele- *+4.9 ‰ (Fig. 5.1). More detailed investiga- vated majoritic garnet component (Seitz et al. tions of bulk peridotite compositions (Magna 2003). But truly, our understanding of et al. 2008; Pogge von Strandmann et al. 2011) high-pressure processes thatmay affect Li solubility constrained further the δ7Li value of the Earth’s and/or mobility in lower mantle mineral phases pristine unmetasomatized upper mantle at * requires dedicated studies. The extent of such +3.5 ‰ with limited variation of *±1.0 ‰ at 2σ explorations is at present hampered by the small-s- confidence. This is very similar to the range in cale nature of available samples relative to capabil- δ7Li values found by Gao et al. (2011) for ities of current analytical techniques. abyssal mantle xenoliths dredged from the Gak- kel Ridge. Combined with the earlier results for 5.1.1.1 Pristine Peridotites xenoliths collected in continental areas, these and the Primitive Upper observations are important for further unraveling Mantle Li Budget the secondary post-solidification effects in vari- The question as to the true Li isotopic signature ous tectonic settings in the mantle as well as for of pristine unmodified mantle, considering the comparison with respect to Li isotopic system- potential for secondary processes to modify Li atics among the terrestrial planetary bodies. isotopic systematics, has persisted. Although The alternatives to xenoliths are orogenic rather sparsely distributed across the globe, peridotite massif exposures and oceanic abyssal peridotite xenoliths are the best direct samples peridotites. One apparent problem with these 122 5 Lithium in the Deep Earth: Mantle and Crustal Systems exposures is the quick reaction of olivine with composition of the Earth’s upper mantle irre- water at elevated temperatures which gives rise spective of the nature of the peridotite samples to serpentinites and similar lithologies, and (massive or xenolith). Hence it appears that some which may partly or completely disguise the peridotite xenoliths faithfully record mantle Li, original Li isotopic signatures (e.g., Decitre et al. which is isotopically rather homogeneous. 2002; Benton et al. 2004). Abyssal peridotites may record ancient mantle melt depletion events 5.1.1.2 Metasomatized Peridotites (Stracke et al. 2011) that could influence the Li and Post-magmatic systematics. However, low-degree melting Modifications appears to impart limited variation to Li isotopes Soon after introducing the δ7Li of the Earth’s (Tomascak et al. 1999). Jeffcoate et al. (2007) mantle through analyses of pristine mantle peri- estimated that equilibrium fractionation would dotite xenoliths, it also became clear that much of generate magmas with <0.5 ‰ difference from the mantle may be isotopically modified as a con- their sources. Refractory harzburgitic residues sequence of secondary processes. Thus, isotopic and dominant silicate phases (olivine, orthopy- variations may not uniquely reflect intrinsic roxene) analyzed by Ionov and Seitz (2008) had heterogeneity. Brooker et al. (2004)attributedLi δ7Li that were not drastically different from each isotopic variations in modally metsomatized spinel other, providing evidence for insignificant mod- peridotites to overprinting by fluids or melts enri- ification to the original Li isotopic signature after ched in isotopically heavy Li, prospectively ulti- melt depletion. Unmetasomatized samples of the mately derived from subducted crust. The Horoman peridotite massif also show a rather comprehensive study of Seitz et al. (2004)docu- constrained variation of δ7Li values at mented post-magmatic shifts in δ7Li of coexisting 3.8 ± 1.4 ‰ (2σ) (Lai et al. 2015), further minerals which most notably affected clinopyrox- attesting to a globally similar Li isotopic ene (Fig. 5.2). This finding was particularly

Fig. 5.2 Lithium concentration and δ7Li of minerals Rudnick and Ionov (2007), Tang et al. (2007, 2011), from mantle nodules. Data sources: Magna et al. (2006a, Aulbach et al. (2008), Aulbach and Rudnick (2009), 2008), Halama et al. (2009), Jeffcoate et al. (2007), Ionov Wagner and Deloule (2007) and Xiao et al. (2015) and Seitz (2008), Seitz et al. (2004), Nishio et al. (2004), 5.1 Lithium in the Mantle 123 significant in the face of apparently undisturbed variations in δ7Li observed in studies of peridotitic bulk Li abundance systematics of the Mg-silicates. samples (as well as other rocks) and their indi- It was at this point that the hidden complexity vidual minerals (see Chap. 4). within the world of mantle Li isotopes became In one of the first detailed studies of diffusive apparent: samples which do not show obvious re-mobilization of Li, Lundstrom et al. (2005) evidence of metasomatism or secondary processes investigated a dunite–harzburgite–spinel lherzolite are not necessarily safe targets for study. Indeed, the transect from an utramafic body of the Trinity study of Nishio et al. (2004)revealedsignificant peridotite in northern California. These authors shifts in δ7Li values in Li-enriched clinopyroxene found significant variations in δ7Li of clinopyrox- from a variety of mantle xenoliths, initially ascribed enealongaprofile and concluded that solely melt to an exotic Li isotopic composition of the EM-1 extraction cannot account for large Li isotopic mantle component with the anticipated troughs observed at dunite/harzburgite interface δ7Li < −17 ‰. Rudnick and Ionov (2007)found and that diffusion must be considered a potentially that these low values may instead result fromkinetic important driving force for Li isotopic variations. effects caused by the diffusive ingress of isotopi- The numerical model of Lundstrom et al. (2005) cally light Li into clinopyroxene from recent showed that 6Li would preferentially diffuse out of (syn-eruptive) percolation of fluids. Nevertheless, melt conduits into wall-rock peridotites. the idea of an isotopically light component resident Several later papers focused in greater detail in the mantle over long time scales persists in some on the topic of kinetic redistribution of Li iso- data compilations (Tang et al. 2010). topes in mantle lithologies. Jeffcoate et al. (2007) It has been verified that Li isotopes diffuse very provided evidence for significant intra-mineral rapidly compared to most species (with the δ7Li variations in coexisting olivine, orthopy- exception of water and gases), paralleled by dif- roxene and clinopyroxene from a cryptically ferential isotopic diffusion rates: 6Li diffuses metasomatized xenoliths from San Carlos. The somewhat faster than 7Li (see Chap. 4). Further- in situ data showed dramatic δ7Li variations in more, and germane to studies such as Nishio et al. pyroxenes (−43 to −5 ‰ in orthopyroxene, −37 (2004), diffusion in clinopyroxene is faster than in to −3.5 ‰ in clinopyroxene) whilst δ7Li varied olivine at mantle conditions (Jeffcoate et al. 2007; between −6 and +8 ‰ in olivine. The Li Parkinson et al. 2007; Rudnick and Ionov 2007), abundances of pyroxenes varied in parallel with leading to sizeable isotopic differences in the final Li isotopic compositions, but for olivine this crystallization and/or post-solidification products variation was inconsistent. It should be noted that from an originally homogeneous reservoir. In fact, MC-ICPMS measurements of Magna et al. geochemically-relevant Li isotopic diffusion was (2006a) yielded δ7Li of +3.6 and +3.3 ‰ in bulk generated synthetically with the formation of aliquots of San Carlos olivine and orthopyrox- smooth Li concentration and isotope profiles ene, respectively, leaving the variations to be between molten basalt and rhyolite with vastly only resolved at microscales and mostly pre- different starting compositions (Richter et al. served at the rims of crystals. 2003). Even an ephemeral thermal event (*6 min Significant Li enrichment associated with in the RB-5 experiment of Richter et al. 2003) led disequilibrium partitioning has also been to a measurable re-distribution of Li, although observed for some spinel lherzolites from Boeun, other elements showed an abrupt concentration South Korea (Kil 2010), and explained to be a offset at basalt/rhyolite interface. This confirms consequence of diffusive fractionation rather than high rates of Li diffusion in silicate melts com- surface contamination or metasomatic process in pared with major cations such as Mg, Fe and alkali the mantle. Gao et al. (2011) analyzed a suite of elements which had been noted earlier (Jambon mantle xenoliths (spinel lherzolites, plagioclase and Semet 1978; Lowry et al. 1981; Cunningham lherzolites and a spinel harzburgite) from the et al. 1983). These observations must thus be Gakkel Ridge in the Arctic Ocean. Disequilib- taken into account when discussing large rium distribution of Li among clinopyroxene, 124 5 Lithium in the Deep Earth: Mantle and Crustal Systems orthopyroxene and olivine (in order of decreas- xenoliths (Ionov and Seitz 2008) has been found ing Li abundances) was found, paralleled by for olivine–clinopyroxene (and orthopyroxene in decreasing δ7Li from olivine to orthopyroxene to few cases) pairs in mantle xenoliths from far-east clinopyroxene. Interestingly, bulk rock δ7Li Russia (Rudnick and Ionov 2007). There, Li values are still in the range of equilibrated mantle isotopic compositions varied greatly not only in xenoliths which may be a consequence of limited pyroxenes (more vulnerable to post-magmatic modification of δ7Li in orthopyroxene combined effects) but also in the coexisting olivine (δ7Li with the heavy Li signature (and low Li contents) from −1.7 to +11.9 ‰), that in case of rapid in olivine counter-balanced by isotopically light cooling preserves the intrinsic isotopic signature. Li in clinopyroxene with high Li contents (up to Thick bodies of alkali basalts appeared to prevent 5 ppm). The unmodified bulk values simply quick cooling, which translated to significant imply redistribution of Li in a closed system, post-eruption kinetic redistribution of Li. Hence, which only affects Li and δ7Li on the mineral cooling rate should figure prominently in the scale but not hand specimen scale. Data of Gao interpretation of Li isotopic data from mantle et al. (2011) also imply temperature dependence xenoliths, although other parameters may be of Li isotopic fractionation between olivine and taken into account such as fertility of the xeno- 7 7 7 clinopyroxene (Δ Liol–cpx = δ Liol – δ Licpx) with liths and grain size (Rudnick and Ionov 2007). 7 significant increase in Δ Liol–cpx values with As explained in detail in Chap. 4, diffusion of Li decreasing temperature (i.e., during cooling). in olivine and clinopyroxene is complex, which Nonetheless these results may preferentially cannot be simply described by a single diffusion apply to closed-system evolution. According to rate. So there might be conditions in which the Gao et al. (2011), this would translate to large fast diffusion mechanism of Li in olivine is δ7Li difference between these co-existing min- activated, which would imply that δ7Li in both erals and possible preservation of negative δ7Li olivine and clinopyroxene could be affected in values in mantle clinopyroxene for long enough similar timescales (e.g., Dohmen et al. 2010). time to be sampled after surface emplacement. However, that is still speculation. Ionov and Seitz (2008) emphasized the role of These reports may lead to the question: are cooling rate of the final eruptive sequence car- light Li isotopic signatures an indicator of diffu- rying xenoliths as a key factor in the generation sive redistribution, or can isotopically light mantle of kinetic Li isotopic effects. They analyzed oli- rocks be generated by primary processes? Of vine and orthopyroxene in a suite of harzburgitic course, kinetic effects may ultimately generate the xenoliths from the Avacha volcano, Kamchatka feedstock for light-Li melts through processes arc, and from the Vitim volcanic field, Siberia. occurring in the sources of such rocks (e.g., sub- They interpreted that xenoliths in pyroclastic duction zone processes), but it is important to be eruptions, which experienced near-instantaneous able to distinguish isotopic signatures that derive cooling, had limited potential to develop Li iso- from sources rather than processes occurring topic disequilibria. On the contrary, xenoliths during transport and emplacement. Halama et al. brought to surface in lava flows experienced (2009) analyzed refractory xenoliths from Ava- protracted cooling times, which translated to cha, far-east Russia, including veined pyroxenites significant δ7Li variations at the mineral scale. representative of the modally metasomatized The difference was particularly apparent for mantle. Perhaps contrary to expectations, they orthopyroxene with δ7Li of +1.8 to +4.0 ‰ in found mantle-like homogeneity of δ7Li in both xenoliths of pyroclastic lavas versus −16.7 ‰ in olivine and orthopyroxene (+3.0 to +5.6 ‰), a single basaltic lava hosted harzburgite; these paralleled by similar δ7Li in veins and their min- differences were paralleled by elevated Li con- eral constituents (+3.0 to +5.1 ‰). To explain the tent in the latter sample (Ionov and Seitz 2008). lack of fractionated isotopic signatures in rocks Similar Li isotopic difference as that observed that had undergone metasomatism by melts car- for olivine–orthopyroxene pairs in lava-hosted rying a strong subduction signature in their trace 5.1 Lithium in the Mantle 125 element geochemistry (e.g., Ba/Th, Pb/Ce), they peridotites with those showing overprints by invoked diffusive Li equilibration between cryptic as well as modal metasomatism. Signifi- cm-scale ascending melt channels and pervasive cant intra-grain variations in δ7Li were observed pristine mantle. This exchange was interpreted as for amphibole-bearing xenoliths whilst only happening on geologically short time scales modest variations were found for amphibole-poor (perhaps as rapidly as 100 years at subarc mantle xenoliths. By excluding late-stage post-eruption temperatures), indicating that Li isotopic differ- hydrous alteration, the variations in olivine, ences between slab-derived, isotopically distinct orthopyroxene, clinopyroxene and amphibole can fluids or melts and ambient mantle are unlikely to be associated with diffusive fractionation during survive if mass transfer takes place at this spatial silicate melt percolation. The hydrous nature of scale (i.e., small veins with low melt/rock ratios). such melt enables precipitation of amphibole This point will be revisited in the following sec- which is otherwise unstable at mantle conditions. tion on mantle-derived melts. Wagner and Deloule (2007) further speculated that While diffusive effects certainly are important olivine and, to a lesser extent, orthopyroxene were in considering δ7Li variations in many mantle less susceptible to metasomatic modifications as samples, Su et al. (2012) argued that extremely these minerals record Li isotopic equilibrium; on light Li in olivine from central China (down to the contrary, clinopyroxene and amphibole mostly −42 ‰, measured by SIMS) and can hardly be showed larger variations in δ7Li. discussed in the framework of diffusive isotope Silicate melt has also been invoked as a meta- fractionation. Beside the possible existence of somatic agent for a suite of spinel peridotites, spinel low-δ7Li lithologies in the mantle, Su et al. (2012) websterites and pyroxenites from Harrad Uwayrid, offer an alternative explanation, seeking metaso- Saudi Arabia (Kaliwoda et al. 2008). Most of their matism by natrocarbonatite melts as a culprit for in situ measured data across olivine, orthopyrox- the ‘reversed’ metasomatic δ7Li systematics of ene, clinopyroxene and phlogopite grains showed co-existing olivine and clinopyroxene. They sug- convex Li elemental profiles, but often with gest that in the presence of high Na concentrations, homogeneous cores. In contrast, abundances of B 7Li tends to enter pyroxenes more easily than 6Li, and Be were very homogeneous across the whole which is left behind in the melts incorporated into grains analyzed. The Li isotopic profiles in most olivine. However, before solid conclusions can be analyzed grains do not show a typical diffusive made a thorough experimental investigation must pattern, nor do the δ7Li values depart significantly assess this intriguing possibility. Nevertheless, from mantle-like values (+1.1 to +9.7 ‰)with carbonatitic metasomatism is quite a common some uncertainty. For one particular sample Kali- phenomenon which may place peculiar constraints woda et al. (2008)revealedsignificant differences on Li partitioning in mantle lithologies (Woodland in abundances of Li, B and Be as well as δ7Li for et al. 2004). The very Li-rich, zoned olivine (tens two adjacent clinopyroxenes, a phenocryst and a of ppm) analyzed by Su et al. (2012) may, how- spongy specimen, with a clearly disturbed ele- ever, have been affected by diffusion. It was pre- mental and isotopic distribution in the latter. Whe- dicted by Dohmen et al. (2010) that the fast ther a hydrous fluid or hydrous silicate melt could diffusion mechanism becomes activated if large Li be responsible for the observed light element sys- contents are incorporated into olivine. tematics remains to be validated. Carbonatitic metasomatism may be quite dif- Another source of metasomatic agents has been ferent from silicate metasomatic events. Wagner invoked for mantle xenoliths (garnet-bearing and Deloule (2007) have used in situ analyses (by versus garnet-free peridotites, and spinel peri- SIMS) of Li isotopic compositions of co-existing dotites) from the Tanzanian part of the East Afri- minerals from hydrous as well as anhydrous can Rift system (Aulbach et al. 2008). Olivine peridotites from the Massif Central, France, in samples in their study showed subtle lithologi- order to constrain the effects of variable extent of cally correlated differences in Li contents and metasomatism and to compare unmetasomatized isotopic compositions. Although δ7Li of 126 5 Lithium in the Deep Earth: Mantle and Crustal Systems clinopyroxene samples varied, most were in iso- mineral grain cores. If correct, a notable inference topic equilibrium with coexisting olivine. of these examinations is that the xenoliths record a One-dimensional numerical modeling led Aul- low-δ7Li source located in the mantle, plausibly bach et al. (2008) to infer that Li isotopic dise- formed by melting of subducted lithologies such quilibrium, recognized by isotopically light Li in as eclogites and perhaps lower crustal granulites pyroxene, may be attained en route to the surface (Su et al. 2012). Nevertheless, the observed zoning through exchange with host melilitites. This pro- could be explained exclusively by diffusion; care cess would only minimally affect Li contents in must be exercised to avoid the potential masking the same crystals, which were slightly higher than effects introduced by grain sectioning. those typical of normal mantle. Based on major Contrasting results for several distinct litholo- and trace element constraints coupled with gies (normal lherzolites, clinopyroxene-rich lher- radiogenic isotope considerations, Aulbach et al. zolites, wehrlites) were obtained by Xiao et al. (2008) invoked a plume-derived silicate melt that (2015) using in situ SIMS technique. The data might have interacted with peridotites from Labait were screened for instrumental matrix-induced area. On the other hand, some other peridotites isotopic fractionation following the approach of from Tanzania appear to have escaped significant Bell et al. (2009). Wehrlites showed mantle-like metasomatic overprint despite interaction with δ7Li and slightly elevated Li contents at *3 ppm. silicate and/or carbonatite melts (Aulbach and Combined with previous petrological studies, Rudnick 2009). Modeling implies that low δ7Li in Xiao et al. (2015) implied metasomatic overprint clinopyroxene of some xenoliths coupled with of peridotites by melts derived from the astheno- apparent elemental equilibrium between olivine sphere, although the Li isotopic signature of the and clinopyroxene must be a pre-eruption feature asthenosphere remains largely unknown. Iso- taking place during transport to the surface in host topically heavy clinopyroxene-rich lherzolites magmas carrying high Li abundances with δ7Li may be representative of precipitation of residual values more or less in the mantle range. melts after previous loss of 6Li due to diffusion. A number of additional studies have revealed Contrasting results in two geochemically Li systematics that do not conform to those pre- diverse peridotite xenoliths from the North China dicted for equilibrium Li isotopic fractionation Craton were obtained by Zhang et al. (2010) for Li between major mantle minerals (Nishio et al. elemental and isotopic profiles across co-existing 2004; Tang et al. 2007, 2011, 2014a, b). Tang et al. mineral grains (olivine, orthopyroxene, clinopy- (2007, 2011) measured Li abundances and iso- roxene). Whereas the low-Mg# peridotite shows topic compositions in a suite of spinel lherzolites elemental and isotopic equilibrium with normal from the North China Craton. In all cases, Li distribution of Li and Li isotopes, the high-Mg# abundances were elevated in pyroxenes compared harzburgites clearly carry signatures of with olivine whereas δ7Li of coexisting pyroxenes post-magmatic modification, reflected in highly were shifted toward negative values, consistent variable Li abundances and disequilibrium δ7Li with kinetic ingress of 6Li from values among major minerals. This latter finding subduction-derived fluids. A later phase of inter- implicates sizeable modifications of refractory action of these modified mantle xenoliths with lithospheric mantle through interaction with fer- melts derived from the asthenosphere was sug- tile melts of subducted lower crust within the gested from elevated δ7Li values at rims of olivine North China Craton. A combination of diffusive and pyroxenes. This two-stage model was subse- Li isotopic fractionation during melt/fluid– quently applied to other North China Craton peridotite interaction and cooling-induced mantle xenoliths (e.g., Tang et al. 2011, 2012). sub-solidus redistribution of Li in xenoliths Tang et al. (2014a, b), using in situ SIMS deter- entrapped in host magmas en route to the surface minations, argued that although the xenoliths did has been invoked for mantle xenoliths from the record effects of diffusion, this was not the major Hannuoba, Fanshi and Hebi areas in the North control on isotopic compositions, at least in case of China Craton (Tang et al. 2011). Lithium 5.1 Lithium in the Mantle 127 elemental and isotopic disequilibria found in most requiring two temporally-distinct episodes of investigated samples are paralleled by Sr–Nd reaction between mantle and fluid/melts. The isotopic disequilibria between coexisting early metasomatic agent had low δ7Li; more clinopyroxene and orthopyroxene as a conse- recent diffusive re-equilibration was in the pres- quence of recent interaction between peridotites ence of isotopically more “normal” melt. These and asthenosphere-derived melts. authors suggested that the ultimate source of the Xu et al. (2013) applied an in situ approach isotopically light Li may be dehydrated residues using LA-MC-ICPMS to collect Li abundance of a subducted slab, although contradictory and isotopic data in a suite of spinel peridotites interpretations of the isotopic compositions of from Kuandian Basin in the North China Craton. eclogites complicate this explanation. Despite large analytical errors compared with Tang et al. (2014b) analyzed olivine cores and conventional MC-ICPMS (see Chap. 2), clearly rims from a suite of high-Mg harzburgites from resolved variations were found in δ7Li for oli- the Hebi area in the North China Craton. A re- vine, orthopyroxene, clinopyroxene and amphi- markable aspect of their dataset was the rather bole (where present). Coupled incompatible normal Li content in most olivines, in the range element indicators and Li systematics implied of olivine from equilibrated peridotites; more- metasomatic overprint of the Kuandian xenoliths over, no significant difference in Li contents and by silicate melts. In particular, the in situ char- δ7Li was found by Tang et al. (2014b) for olivine acter of the study allowed for detailed mapping cores (0.9–2.2 ppm Li; −27.2 to +2.4 ‰ (n = 11), of Li elemental profiles across different mineral two values at +1.2 and +21.5 ‰) and rims (0.9– grains. In almost all cases significant enrichment 3.6 ppm Li; −38.0 to −1.5 ‰ (n = 12), one value of Li in rims relative to cores, paralleled by at +19.5 ‰ related to the isotopically heavy highly negative δ7Li in rims of both clinopy- core). Thus, the observed δ7Li range, paralleled roxene and orthopyroxene (down to −56 ‰), by the lack of a developed diffusive isotopic hint to recent diffusive fractionation of Li profile, suggested participation of a component induced by percolation of metasomatic liquids with unusually low δ7Li, again pointing to sub- shortly before eruption and during cooling. In an ducted residues of oceanic crust. It thus appears attempt to unveil the source of these metasomatic that such processes may operate under appro- agents, in situ determination of Sr isotopes was priate conditions and that Li isotopic hetero- employed, combined with incompatible trace geneities in the mantle may survive long and be elements parameters (e.g., La/Nb, U/Nb). These more resistant to attenuation as a consequence of appeared to exclude recycled continental crustal diffusion (cf. Halama et al. 2008). sources and instead invoked a possible relation to These general inferences from the North subducted Pacific oceanic crust (Xu et al. 2013), China Craton may also apply to other areas. similar to findings of Tang et al. (2012) for Ackerman et al. (2013) measured Li contents and another area in the North China Craton. isotopic compositions in a suite of The study of Tang et al. (2012) showed ‘re- basanite-hosted xenoliths from NE Bavaria, versed’ sense of Li isotopic fractionation in Germany, and revealed some of the lowest δ7Li metasomatized mantle xenoliths, with preva- measured to date for bulk rocks (down to lently light signatures in olivine relative to −9.7 ‰). With the exception of two samples co-existing clinopyroxene and orthopyroxene. with δ7Li > +2.3 ‰, none of the peridotites The data were inconsistent with kinetic frac- showed mantle-like δ7Li. These compositions tionation having produced the entire range in Li were paralleled by extremely high Li/Yb ratios elemental and isotopic compositions, because of up to *140, resulting from combined Li ingress normal-mantle abundances in all minerals, as and Yb depletion by melting. This was inter- well as clinopyroxene being isotopically heavier preted to reflect high degrees of melting and than olivine despite faster Li diffusion in the concomitant kinetic isotopic fractionation. former. They favored an evolutionary model However, the ingress of isotopically light Li 128 5 Lithium in the Deep Earth: Mantle and Crustal Systems might have instead originated in subducted lithosphere, fluid-modified mantle wedge and dis- oceanic crust. No clear distinction could be made tinct continental components (terrigenous sedi- for Li systematics in terms of metasomatic agents ments and subcontinental lithospheric mantle). The (i.e., silicate or carbonate-rich), but not all idea behind a low-δ7Li recycled component xenoliths with low δ7Li were associated with accommodated a previously introduced model by unusually high Li abundance in clinopyroxene, Zack et al. (2003) whose analyses of omphacite which often carries low δ7Li values (as detailed (clinopyroxene) as a prime constituent of eclogitic above). A similar distribution of Li elemental and lithologies, together with low-Li garnet (e.g., isotopic compositions has been observed for the Woodland et al. 2002), showed highly negative vertically stratified mantle beneath the Kozákov δ7Li values as low as −11 ‰,thatwerethoughtto volcano, Bohemian Massif, close to a major represent dehydrated subducted residue after tectonic discontinuity. All investigated samples aqueous solutions with isotopically heavy Li had appear to carry a disturbed δ7Li signature (−5.8 been expelled off the slabs entering subduction to −0.8 ‰). This was interpreted to reflect the zones. This model of introducing light Li into a combined effects of mantle metasomatism and deeper mantle was critically reviewed (Marschall melt depletion (Ackerman et al. 2013), similar to et al. 2007b) by involving kinetic diffusive modi- the interpretation of samples from the NE fications and stating that dehydration alone would Bavaria suite (Medaris et al. 2015). be capable of decreasing the resulting δ7Li values in Taken as a whole, these observations provide residual slab by no more than c. 3 ‰. However, the evidence that significant volumes of the sub- progressive Li isotope fractionation coupled with a continental lithospheric mantle beneath large large proportion ofLi retained even after high-grade continental areas may retain a record of Li iso- moderate-temperature metamorphism applicable to topic modification, perhaps due to melt extrac- subduction zones (Marschall et al. 2006) may still tion associated with reactions between a remain an important alternative for explaining melt-depleted mantle column and fertile melts systematically decreasing δ7Li away from the en route to the surface. The evidence for pro- trench (e.g., Moriguti and Nakamura 1998;Magna cesses other than solely kinetic fractionation, et al. 2006b; Agostini et al. 2008). It is thus per- apparently strong in many cases, has generally missible that fractionated Li transferred further into not been backed up by diffusion modeling. It is the subduction factory may appear in distant time– our opinion that diffusive fractionation in space relations at the Earth’s surface through igneous systems has not been given appropriate MORB and/or OIB, but unraveling the ultimate consideration in many cases. Detailed in situ Li sources and histories of reservoirs carrying this Li and δ7Li profile measurements (performed on isotopic distinction still requires much work. multiple grains), accompanied by diffusion Rapidity of diffusive exchange in newly modeling, are required to clearly show that the formed melts has been documented by Magna observed δ7Li signatures can not be produced by et al. (2008) who analyzed three melt pockets diffusion. co-existing with either clinopyroxene or olivine. Notably, these newly formed melt pockets were isotopically lighter by 2–9 ‰ than the 5.1.2 Mantle-Derived Magmas co-existing minerals. That the two olivine-associated melt pockets had heavier δ7Li Deviations from a global mean δ7Li (*+3 to +5 ‰) than that associated with clinopyroxene suggests in products of mantle melting could provide clues that kinetic effects are faster into/out of on processes of recycling of individual components clinopyroxene than olivine, consistent with more of subduction-related materials. These include focused experimental studies (see Chap. 4). Jef- marine sediments, isotopically stratified oceanic fcoate et al. (2007) modeled peridotite mineral 5.1 Lithium in the Mantle 129 data using observations of Lundstrom et al. meteoric water because Iceland basaltic lavas are (2005) to indicate that diffusive loss of Li to wall much younger than HIMU lavas (in fact, most of rocks of ascending melts is unlikely to have a them are relatively recent at < *10 ka) and a >1 ‰ effect on systems where conduits trans- large part of these high-MgO (>7 wt% MgO) porting melts are spaced >50 m apart. Interest- basaltic to picritic lavas erupted subaerially, thus ingly, a glassy sample of Hawaiian basalt escaping secondary contamination by melt water analyzed by Jeffcoate et al. (2007) showed no Li from ice sheets. isotopic zoning, which attests to the key role of Finally, one potential shortcoming of using cooling rate in developing diffusive Li profiles. igneous derivatives to plumb Li isotopes in the Generating isotopically heavy Li that could mantle is the susceptibility of volcanic rocks to then be re-processed in and beyond the subduc- weathering. Although efforts are consistently tion zone appears to be a more complex issue. made to avoid “visibly altered” samples and parts The isotopically heavy Li mainly enters the of sample, it seems likely that some of the volcanic ocean floor rocks through exchange with iso- rocks from which Li isotopic information has been topically heavy seawater so that Li becomes extracted have been subject to subtle weathering, enriched and isotopically heavier in the ocean which could yield rocks with artificially lighter or floor compared to deeper ocean crustal levels heavier isotopic systematics. One early example (Chan et al. 2002a). Decitre et al. (2002) and of potentially unreliable isotopic measurements in Benton et al. (2004) showed that serpentinites what was described as unaltered MORB samples become isotopically quite heterogeneous with a (Chan et al. 1992) had δ7Li (+6.3 to +6.8 ‰) slight bias towards elevated δ7Li although mod- outside the subsequently verified global range, as ifications to Li budget of serpentinites only occur well as anomalously high bulk K2O content at high water/rock ratios (Vils et al. 2009) and (0.16–0.22 wt%). Although foundational in our serpentinization itself appears to provide subor- understanding of Li in the mantle, it is interesting dinate contribution of Li to the global Li budget to note that these two samples represented one in oceanic lithosphere (Vils et al. 2008). Inde- third of the data set for the MORB samples in that pendently of each other, Chan et al. (2009) and study (Chan et al. 1992). Vlastélic et al. (2009) have inferred that elevated δ7Li values (>+6 ‰) in some HIMU lavas are a 5.1.2.1 Mid-Ocean Ridge Rocks result of incorporation of altered oceanic crust Mid-ocean ridge basalts (MORB) are considered into the mantle source of HIMU basalts through to reflect melt-derivation from depleted mantle. recycling, an idea largely consistent with There appears to be a range in elemental and incompatible trace elements and radiogenic iso- radiogenic isotopic systematics reflecting com- topic constraints (e.g., Hofmann 1997; Stracke positional heterogeneity of the depleted mantle et al. 2005), and that these heterogeneities can that includes incompatible-element-depleted and indeed survive in the mantle for a period of variably enriched components (e.g., Salters and >1 Gyr without apparent attenuation (cf. Halama Stracke 2004). Thus the correlation between et al. 2008). Importantly, geographically distinct MORB and peridotitic mantle may be a some- occurrences of mantle plume-related volcanic what complex one and not all MORB may reflect rocks with high δ7Li values (up to +9 ‰ in rocks the isotopic composition of the majority of the with the highest MgO contents) were reported mantle. Nevertheless, considering minimal frac- from Iceland (Hansen et al. 2011; Magna et al. tionation of Li isotopes during low-degree melt- 2011) that cannot easily be assigned to a specific ing (Tomascak et al. 1999), they may still recycled material with HIMU affinity indicating possess important information about the mantle that heavy Li isotopic signature may be more that lies beneath the ridges. ubiquitous than it would appear from genetic Despite widespread availability of MORB connection to solely HIMU-type ‘habitats’.It samples from many distinct locations in the also cannot result from simple contamination by Atlantic, Pacific and Indian Oceans, limited Li 130 5 Lithium in the Deep Earth: Mantle and Crustal Systems

Fig. 5.3 Lithium concentration and δ7Li of MORB and and Nakamura (1998), Nishio et al. (2007), Chan et al. altered MORB. Data for altered MORB samples come (1992), Tomascak et al. (2008), East Pacific Rise from from the Mid-Atlantic Ridge (Chan et al. 1992) and East Chan et al. (1992), Elliott et al. (2006), Tomascak et al. Pacific Rise (Chan et al. 2002a, b; Brant et al. 2012). (2008), Indian Ridge from Nishio et al. (2007) and Sources of other data: Mid-Atlantic Ridge from Moriguti Tomascak et al. (2008) isotopic data were available until recently (Elliott 2006; Nishio et al. 2007; Tomascak et al. 2008) et al. 2006; Nishio et al. 2007; Tomascak et al. which may reflect moderately incompatible 2008; Fig. 5.3). Although these studies served to behavior of Li during melting of the mantle strengthen the earlier data set, the more recent column, but consistent δ7Li values reflect the data ultimately did not change the fundamental inability of solely low-degree mantle melting to characterization of Li in the MORB mantle rel- modify Li isotopes. Part of the δ7Li variation ative to the preliminary study of Chan et al. seen in MORB may reflect incorporation of other (1992; δ7Li = +3.4 to +4.7 ‰). This overlaps the lithologies and/or processes, most notably range in δ7Li estimated independently from subduction-modified mantle enriched by fluids peridotites. from the dehydrating slab (Elliott et al. 2006; The expanded MORB data set did provide Tomascak et al. 2008). However, the evidence some deeper insight, however. Tomascak et al. that aberrations in δ7Li of MORB stem exclu- (2008) provided evidence for a lack of provin- sively from recycled components resident in the ciality in distribution of Li isotopes in unaltered upper mantle falls far short of conclusive. For MORB from three distinct ridges. Additionally, example, based on correlations between δ7Li and no distinction could be made between halogen concentrations, Tomascak et al. (2008) normal-type MORB (more or less representative favored a shallow magma chamber contamina- of depleted mantle) and enriched MORB vari- tion model to explain some East Pacific eties, irrespective of indices of incompatible Rise MORB data. element enrichment/depletion (for example,

K2O/TiO2). In any case, *3–10-fold enrich- 5.1.2.2 Intraplate Volcanic Rocks ments in Li (*4to*13 ppm) relative to the The other possibility to constrain mantle com- precursor mantle could be observed (Elliott et al. position of Li is linked to numerous occurrences 5.1 Lithium in the Mantle 131 of hotspots. Although basalts from these loca- Iceland pointed to limited Li enrichments and tions derive from mantle melting, their charac- showed, with only few exceptions, concentra- terization is complicated by chemical signatures tions <5 ppm (Ryan and Langmuir 1987). In which may carry evidence of recycling of distinct particular locations, mostly confined to Polyne- parts of the crust and/or marine sediments sia, Li abundances up to *10 ppm have been (Willbold and Stracke 2006). Thus, intraplate reported (Dostal et al. 1996). No studies of Li volcanic rocks may not be expected to yield isotopes in OIB were published until 1999, constraints on Li in the average mantle, but they despite the global importance of such large-scale may instead be valuable tracers of the persistence mantle features. The results for a suite of Kilauea of subducted components. For example, the Iki lava lake samples, ranging from high MgO interpretation of subduction zone rocks by Mar- rocks to differentiated liquids with high SiO2 and 7 schall et al. (2007b) is that isotopically heavy Li K2O, showed δ Li values (+3.0 to +4.8 ‰) fully could later appear in OIB (ocean island basalts) in the range of pristine MORB (Tomascak et al. through entrainment of 7Li-enriched mantle 1999). Although no Li abundances were reported wedge materials. This is consistent with the in that study, later results (Chan and Frey 2003) interpretation of Li isotopes from subduction indicated that most Hawaiian basalts contained zone magmatic rocks (e.g., Tomascak et al. 3–5 ppm Li. These low values were supple- 2000). mented by later studies of Li systematics in The first systematic Li abundance data avail- various OIB locations (Fig. 5.4; Pistiner and able from various OIB such as Hawaii and Henderson 2003; Ryan and Kyle 2004; Jeffcoate

Fig. 5.4 Lithium concentration and δ7Li in volcanic (2012), Schuessler et al. (2009), Azores from Krienitz rocks from oceanic islands. These are primarily basaltic in et al. (2012), Genske et al. (2014), South Pacific (includes composition, but not exclusively (e.g., some Iceland rocks data of 14 different localities) from Vlastelic et al. (2009), with >50 wt% SiO2). Hawaiian rocks with indication of Nishio et al. (2005), Chan et al. (2009), Krienitz et al. surficial alteration have been excluded. Data sources: (2012), “other” (includes data from islands in the Atlantic Hawaii from Chan and Frey (2003), Kobayashi et al. and northern Pacific) from Ryan and Kyle (2004), (2004), Nishio et al. (2007), Iceland from Ryan and Kyle Krienitz et al. (2012) and Magna et al. (2011) (2004), Magna et al. (2011), Pogge von Strandmann et al. 132 5 Lithium in the Deep Earth: Mantle and Crustal Systems et al. 2007; Schuessler et al. 2009; Magna et al. lavas from Iceland (δ7Li up to +9 ‰; Hansen 2011). A bias towards higher Li contents, how- et al. 2011; Magna et al. 2011). This signature ever, continued to be evident in classical HIMU could not be explained by contamination by (high μ; μ ≡ 238U/204Pb; Zindler and Hart 1986) meteoric water and wall rock, and may instead occurrences in Polynesia (Nishio et al. 2005; appeal to recycling of altered oceanic crust at Chan et al. 2009; Vlastélic et al. 2009; Krienitz large scales, or parts of previously metasoma- et al. 2012). tized subarc mantle (Elliott et al. 2006; Marschall Considering the lithological diversity among et al. 2007b). Genske et al. (2014) found that the the inputs to subduction assemblages that con- majority of Azores lavas have MORB-like Li tribute to the final chemical signature of isotopes. Some samples in their study, however, subduction-modified mantle, and considering the range up to δ7Li of +8.8 ‰; the majority of these large variability in both the Li contents and iso- isotopically heavy samples were attributed to topic composition of these components (Bouman shallow-level crustal contamination, although et al. 2004; Chan et al. 2006), it is intriguing that when combined with radiogenic isotope data a most OIB worldwide carry a homogeneous Li recycled source component in the mantle could isotopic signature (Fig. 5.4). In particular, unal- not be ruled out. Kobayashi et al. (2004) implied tered Hawaiian basalts have particularly well a recycled component in the source of Hawaiian constrained δ7Li values, roughly between +2.5 basaltic lavas from in situ (SIMS) Li isotopic and +5.7 ‰ (Tomascak et al. 1999; Chan and analysis of glass inclusions enclosed in olivine Frey 2003; Kobayashi et al. 2004) with some phenocrysts. Their results showed a considerable subtle variations among individual volcanic spread in δ7Li values from −10.2 to +8.4 ‰, but systems (Mauna Kea, Mauna Loa, Koolau, with low to negative δ7Li prevailing among the Kilauea Iki). Interestingly, the Hawaiian certified data and broadly correlating with δ11B values. reference basalt BHVO-2 has been shown to These contrasting results indicate distinctly dif- have δ7Li in the range of the unmetasomatized ferent compositions of recycled reservoirs that peridotites (see Sect. 5.1.1.1). Most other OIB can only be more unambiguously distinguished appear to mimic the mantle δ7Li range, including with appropriate samples. Our understanding of lavas with (i) HIMU affinity (Cook–Austral behavior of Li is not yet sufficient to fully con- Islands, St. Helena; Ryan and Kyle 2004; Nishio strain the fate of subducted materials in the et al. 2005; Chan et al. 2009; Vlastélic et al. mantle. 2009; Krienitz et al. 2012), (ii) EM-1 affinity Collectively, the available analyses suggest the (Pitcairn; Krienitz et al. 2012), (iii) EM-2 affinity worldwide Li isotopic homogeneity of OIB (Society Islands, Réunion; Ryan and Kyle 2004; occurrences and indicate that the incorporation of Nishio et al. 2005; Vlastélic et al. 2009; Krienitz materials with distinct chemistry, such as sedi- et al. 2012), and (iv) mixed affinities (Easter ments (marine, terrigenous) and/or altered sea- Island; Krienitz et al. 2012). Also, some hot spot floor rocks, is not easily projected into the final localities associated with assumed, though not δ7Li signature. Unless these additives carry yet geophysically confirmed, mantle plume such peculiar Li systematics that can be sustained as Jan Mayen (Debaille et al. 2009) appear to through subduction processing and long-term have Li systematics (Magna et al. 2011) indis- residence in the Earth’s mantle without much tinguishable from those reported for typical modification of their original Li inventory, the plume-related OIB occurrences. isotopic signal appears most commonly to be Some OIB, though, carry distinctively heavy obliterated. This suggests that mixing in the Li isotopic signature such as HIMU-type basalts mantle is largely an effective process with very with δ7Li > +6 ‰ (Ryan and Kyle 2004; Nishio few localities showing resolved departures et al. 2005; Chan et al. 2009; Vlastélic et al. towards non-mantle δ7Li preserved through inef- 2009). However, an isotopically heavy Li sig- ficient mixing of large-scale heterogeneities nature was also found in pristine non-HIMU introduced into the deeper mantle. We cannot 5.1 Lithium in the Mantle 133 exclude the impact of changes in physics of the elevated Li contents (*9–12 ppm) relative to mantle throughout Earth's history, considering, MORB and OIB, but with somewhat variable for example, that the HIMU signature is known to δ7Li values between +2 and +5 ‰, mostly be of ancient origin (Hofmann and White 1982; hinging on resolvedly higher δ7Li of BCR-2G Zindler and Hart 1986; Sobolev et al. 2007). (+4 to +5 ‰). Liu et al. (2013) measured Li Compared with more or less localized OIB isotopes in Columbia River basalts at two sites, localities, continental basaltic volcanism includes with average δ7Li of +0.5 ± 1.0 ‰ and +1.7 ± large igneous provinces, some of the most 2.2 ‰. These values are in line with the low voluminous effusive igneous products on Earth. values measured in BCR-1 and BCR-2, and point These may exceed 1 × 106 km3 total volume, to sources with a component isotopically lighter cover areas often >1 × 106 km2, and vary in age than N-MORB. One possibility would be repre- from Mesoproterozoic to Neogene. Given such sented by sources that assimilated continental volumes present at the Earth’s surface it is sur- crust, consistent with the interpretation of Carl- prising the Li systematics of these occurrences son et al. (1981), based on radiogenic isotopic are highly under-represented in the literature systematics. database. One of the commonly used basalt cer- The somewhat Li-enriched nature of alkali tified reference rocks, BCR, derives from the basalts in continental settings (Fig. 5.5) was Columbia River large igneous province. Three already observed by Ryan and Langmuir (1987) different aliquots of the rock have been issued for a few occurrences and later in more detail by (BCR-1, BCR-2 and BCR-2G), all showing Mengel and Hoefs (1990), who noted that

Fig. 5.5 Lithium concentration and δ7Li in continental et al. (2011), Massif Central from Hamelin et al. (2009), igneous rocks, primarily ultramafic to intermediate com- central Europe from Abdelfadil et al. (2014), Ackerman position volcanic rocks, many associated with extensional et al. (2015), carbonatites from Halama et al. (2007, tectonic settings, but also including carbonatites and 2008), “other” (includes data from silicate rocks from selected mafic plutonic rocks. Data sources: Antarctica eight globally dispersed regions) from Ackerman et al. from Ryan and Kyle (2004), Columbia River flood basalts (2013), Magna et al. (2008), Halama et al. (2008, 2009), from Liu et al. (2013), Mediterranean (includes data from Marks et al. (2007), Rudnick and Ionov (2007), Kil Sardinia and Stromboli) from Schiavi et al. (2010), Kim (2010) and Tomascak and Magna (2014) 134 5 Lithium in the Deep Earth: Mantle and Crustal Systems contamination of alkaline magmas by continental Cenozoic basalt from Mongolia had an isotopi- crust may have operated for Cenozoic alkaline cally heavy signature (δ7Li = +8.6 ‰; Magna et al. 18 basalts from Germany based on Li–SiO2–δ O 2008) but the nature of this distinctive value variations. Contamination of this type would remains to be explored. mask the indigenous Li inherited from alkali Several suites of Antarctic intraplate alkaline basalt sources due to the high Li content of the lavas were examined by Ryan and Kyle (2004), continental crust (Teng et al. 2004). However, who noted moderately incompatible behavior of the Li contents in such assemblages may be quite Li in magmatic processes and distinctive Li/Yb in low. This is apparent from the study of Weyer intra-plate rocks versus MORB, tentatively and Seitz (2012) who analyzed matrix and oli- ascribed to different mantle sources. Chemically vine of several basaltic rocks from alkaline less evolved Antarctic intraplate lavas (e.g., from complexes in Germany (Vogelsberg, Hegau, the Crary Mountains with MgO > 6 wt%, but also Westerwald), related to pan-European Cenozoic other occurrences with MgO contents 3–5 wt%) volcanism (e.g., Wilson and Downes 1991). confirm the largely homogeneous Li isotopic These data showed Li abundances to be at the signature of source mantle of these alkaline rocks lower end of alkaline volcanic rocks (6.2– (δ7Li = +3.3 to +5.0 ‰). Together, these samples 7.9 ppm), paralleled by rather homogeneous δ7Li were considered least affected by crustal con- values between +2.5 and +4.7 ‰. In all instances tamination and open system processes. The the olivine had lower Li concentration than homogeneity of Li isotopes in intraplate settings matrix but highly variable δ7Li (from −10.5 to has further been validated by Halama et al. (2009) +6.5 ‰), attesting to significant exchange for several basaltic andesites from the Avachinsky between olivine and melt by diffusion, as pre- volcano in Kamchatka that mimic Li elemental dicted by Gallagher and Elliott (2009). In par- and isotopic systematics of the Antarctic rocks. ticular, the Li-rich olivine from Vogelsberg These suites that show only limited Li iso- nephelinite carried the lightest Li isotopic sig- topic variability contrast to a certain extent with nature but Weyer and Seitz (2012) estimated isotopically heavy Li found in alkaline basaltic *50 % of this olivine to be xenocrystic in origin, lavas from the Massif Central in France (Hamelin supported by unusually high Ni contents and et al. 2009). The δ7Li values up to +7.4 ‰ were Mg# significantly above the mantle value. The discussed in the context of a HIMU component ensemble data implied more pronounced effects in the source of Cenozoic basalts across Europe of disequilibrium on lavas containing xenocrysts that has also emerged from the Sr–Nd–Pb iso- rather than phenocrysts. topic perspective (e.g., Blusztajn and Hart 1989; The preliminary report of Magna and Rapprich Wilson and Downes 1991). On the other hand, a (2012) for a suite of chemically pristine alkaline light Li isotopic signature (δ7Li as low as lavas (melteigites, ijolites, essexites) from the +0.4 ‰) found in chemically more evolved Bohemian Massif revealed low to moderate (*4– lithologies of the Massif Central, such as ben- 16 ppm) Li abundances. Samples from the geo- moreites and trachytes, provides evidence for the chemically complex Doupovské hory Volcanic assimilation of lower continental crust compo- Complex, Czech Republic, yielded δ7Li values nents, as gauged through metasedimentary that varied from mantle-like towards both xenoliths. These inferences are also in accord 7Li-depleted and 7Li-enriched compositions (−1 with variations in Li concentration from 6 to to +11 ‰), whilst pristine high-MgO (>12 wt%) 20 ppm, increasing consistently with the degree basanites from another suite showed little varia- of fractional crystallization (Hamelin et al. 2009). tion in δ7Li, all within the mantle range. Similar Li Contrary to findings of Hamelin et al. (2009), elemental and isotopic homogeneity has been Ackerman et al. (2015) did not observe a similar found for three host basanites from NE Bavaria, unidirectional evolution of Li abundance and Germany, all with *6 ppm Li (Ackerman et al. isotopic composition in a suite of trachytes, tra- 2013). On the other hand, a xenolith-bearing chyandesites and phonolites from the Bohemian 5.1 Lithium in the Mantle 135

Massif. This may, in part, be due to the fact that olivine-hosted melt inclusions from the Strom- these samples were not co-genetic (cf. a boli volcano pumices (Schiavi et al. 2012) when closed-system suite analyzed by Hamelin et al. compared with their host lithologies (Schiavi 2009), although modelling of fractional crystal- et al. 2010). Modeling suggested that this lization showed clear overall evolution trends. prevalently light Li (and B) isotopic signature Large extents of fractional crystallization of oli- may have originated through metasomatic over- vine + clinopyroxene + amphibole + plagioclase printing by high-pressure fluids released from appear to have been detected in elevated Li dehydrated metasediments and metabasites. This contents of some phonolites (up to 75 ppm) notion may be reasonable considering fraction- whereas open-system behavior is evidenced by ated Li signature in fluids derived from a sub- (i) a peculiar dichotomy in Li–Cs elemental ducted slab (John et al. 2012). systematics, and (ii) a large variation of δ7Li The complex behavior of Li in alkaline com- between −5.3 and +8.4 ‰ without any particular plexes has been illustrated by Marks et al. systematic behavior (Ackerman et al. 2015). (2007), who analyzed amphibole and clinopy- Interestingly, isotopically light mantle roxene from feldspathoid-bearing rocks (ag- domains beneath some parts of Europe have also paites: lujavrites, kakortokites, naujaites etc.) of been invoked for tholeiites collected from the the Ilímaussaq plutonic complex, South Green- vicinity of Etna volcano on Sicily (Gurenko and land. The particularly Li-enriched character of Schmincke 2002), based on coupled Li–B iso- sodic amphiboles (up to *3100 ppm Li, broadly topic compositions in glass inclusions and their correlating with increasing degree of differentia- host glasses (δ7Li from −3.3 to +1.2 ‰). How- tion) contrasted strongly with low Li contents in ever, Gurenko and Schmincke (2002) implied Ca-amphibole (7–68 ppm) and consistently low former subduction as an ultimate source of light Li contents in co-existing clinopyroxene (15– Li rather than assimilation of continental crustal 84 ppm). Moreover, Li correlated positively with rocks altered under hydrothermal conditions. Fe3+/ΣFe in Na-amphibole, indicative of a pos- Abdelfadil et al. (2014) reached the same con- sible proxy for redox conditions during magma clusion from the isotopic compositions of Pale- crystallization (Marks et al. 2007). Homogeneity ozoic calc-alkaline lamprophyres in the northern in δ7Li values of amphiboles from the inner part Bohemian Massif (δ7Li from −5.5 to +1.9 ‰). In of the Ilímaussaq complex (*+2 ± 2 ‰) were their study, correlations between Li concentra- thought to reflect mantle-derived origin, a con- tion, δ7Li, radiogenic isotopes and other trace clusion supported by coupled Nd–O isotopic elements were interpreted to rule out diffusional systematics which largely precluded significant control and instead argued for a mantle compo- crustal contamination. Open system nent derived ultimately from subducted compo- post-magmatic kinetic modifications resulting nents, principally originating in sediments. from the ingress/circulation of 7Li-enriched fluids These findings perhaps explain predominantly were advocated for outer lithologies of the light Li isotopic composition in basalts from complex (Marks et al. 2007). Sardinia (δ7Li from +1.5 to +3.6 ‰, with a single Collectively, Li isotope systematics in basalt with a distinct REE pattern at +8.0 ‰; intra-plate alkaline volcanic rocks may in favor- Kim et al. 2011). These authors implied that the able situations disclose important processes of light Li and EM-1-like incompatible trace ele- tapping heterogeneous mantle components as ment signature could be reconciled with the well as assimilation of evolved lithologies asso- absence of altered oceanic crust. They suggested ciated with fractional crystallization although that diffusion may have caused 7Li depletion in from a global perspective, the available Li the initial mantle source of the Sardinian basalts abundances and isotopic compositions appear to although at these scales complete diffusive be broadly consistent for distinct locations. Yet, modification may not be easily achieved. Pre- the number of studies that consider tectonic set- dominantly light Li was also observed in tings with thick continental crust penetrated by 136 5 Lithium in the Deep Earth: Mantle and Crustal Systems alkaline liquids derived from asthenospheric variation in the original carbonatite source has depths is so far very limited. Further studies of been homogenized with time as a consequence of intra-plate plutonic and volcanic sequences car- the high mobility of Li. rying chemical signatures of distinctive sec- A more detailed assessment of Li elemental ondary processes await more in-depth and isotopic distribution in carbonatites from exploration. southwestern China and their relation to coex- Carbonatites have also been used to unravel isting syenites was explored by Tian et al. the Li isotopic evolution of the Earth’s upper (2015). Bulk carbonatites had between 1 and mantle through time (Halama et al. 2007, 2008). 120 ppm Li; arfvedsonite and biotite had Carbonatites have long been considered potential >360 ppm Li, with apparent partition coefficient arfv–WR probes of the sub-continental mantle. Taking into Kd * 8–12. An even larger difference of account their rapid ascent and extremely low *22 was found for a single analyzed biotite– viscosities, carbonatite magmas are less suscep- whole rock pair. The large range in Li contents of tible to chemical modification during transport carbonatites contrasted with rather small abun- through the crust and can therefore preserve dance variation in associated syenites (4– pristine mantle signatures better than other 39 ppm). Notable variation in δ7Li for the whole mantle-derived magmas (Bell 1989, 1998). Apart suite (−4.5 to +10.8 ‰) was found, although the from massive enrichments in large ion lithophile majority of data are between *0 and *+4 ‰. elements (LILE; K, Rb, Cs, Sr, Ba), By excluding kinetic effects by diffusion, Tian high-field-strength elements (HFSE; Zr, Hf, Nb, et al. (2015) advocated for isotopically hetero- Ta, Th, U) and the light rare earth elements geneous sub-continental lithospheric mantle (LREE; La–Sm) (e.g., Bizimis et al. 2003; modified by subduction-derived fluids that bore Chakhmouradian 2006), available Li data for the imprint of marine sediments and subducted carbonatites show variable depletions and oceanic crust. Isotopically anomalous enrichments from sub-ppm levels in calcic and sub-continental lithosphere has been invoked silicic carbonatites to several hundred ppm Li in also for other regions of China (see Sect. 5.1.1) sodic carbonatites (up to *300 ppm; Halama based on detailed analyses of minerals from et al. 2007, 2008). It appears, too, that these mantle xenoliths. values reflect the mantle source of carbonatites These findings may collectively suggest that rather than kinetic processes and/or contamina- Earth’s upper mantle may, in places, be far from tion (Cooper et al. 1995; Halama et al. 2008). a homogeneous reservoir and that under partic- The geochemical signatures of carbonatites have ular conditions, regions in the mantle may exist been explained by the involvement of enriched that carry significantly different load of Li. mantle, lithospheric mantle and/or HIMU mantle Whether such ‘oddities’ are rather rare and rela- (e.g., Bell and Simonetti 1996; Bell 1998). ted to specific petrogeny and what consequences Considering the Li isotopic disparity between for mantle dynamics these unusual magmas may some of these “global” mantle end members have, remains to be investigated in the future. (Chan et al. 2009; Vlastélic et al. 2009; Krienitz et al. 2012), differences in δ7Li could eventually 5.1.2.3 Arc and Back-Arc Rocks distinguish between individual contributions Volcanic arcs are among the dominant morpho- from these geochemically distinctive reservoirs logic features shaping the Earth’s surface and (Zindler and Hart 1986). Yet, carbonatites ana- volcanism at convergent margins gives rise to lyzed by Halama et al. (2007, 2008) fall into a large batches of magma with peculiar elemental narrow range of δ7Li (+4.1, ± 2.6 ‰,2σ); this budgets which contribute significantly to the isotopic homogeneity extends from recent back construction of the continental crust (Gill 1981). to *2.7 Ga. This finding may suggest that either Because these massive collisional arcs include an the mantle source of carbonatites is homoge- extraordinary lithological diversity from marine neous across the time and space, or that any δ7Li and continental sediments through magmatic 5.1 Lithium in the Mantle 137 rocks to metamorphic sections, large variability conditions related to the loss of fluid-mobile ele- in Li elemental distribution and isotopic com- ments such as Li, K, Rb, Cs and B and implied position of terminal products of arc magmatism significant Li isotopic fractionation during alter- may be expected. The variability of sedimentary ation of seafloor basalts (weathered basalt with input in Li contents from sub-ppm to a few starting δ7Li = +7.4 ‰) as an important input to hundred ppm, and in δ7Li at the >20 ‰ scale the subduction factory. The rapid development of (Chan et al. 1994; Chan and Kastner 2000; fluid containing isotopically heavy Li (δ7Li up to Bouman et al. 2004; Chan et al. 2006; Chan and +14 ‰), paralleled by high Cs/Rb, suggested Hein 2007) contrasts with largely homogeneous preferential decomposition of low-temperature Li contents and isotopic compositions of alteration phases such as clay minerals. The role non-sedimentary portion of the oceanic litho- of these secondary phases in cycling of Li during sphere, although important vertical Li variability seafloor weathering and hydrothermal circulation of oceanic crust exists (Chan et al. 2002a; Gao was further constrained by James et al. (2003) who et al. 2012). implied that fluid/solid ratios may bear important The importance of sedimentary input to the Li control on Li isotopic fractionation as well as the budget of arc volcanic rocks when combined partitioning of Li into fluid or solid phases at with other trace element indicators such as Li/Y elevated temperatures. has recently been evaluated by Plank (2014), The expulsion of fluids coupled with com- noting that Li enrichment from the sedimentary paction of mainly the loose sedimentary cover in column may become significant in some loca- incipient stages of subduction is followed by tions (e.g., Cascades, eastern Sunda, Aleutians) prograde metamorphic dehydration reactions while it is less important for other arcs (e.g., ultimately leading to ‘dry’ eclogites, representing Tonga, Mariana, Central America). Apart from dehydrated residues of subducted slabs. This sedimentary input (and altered oceanic crust in paradigm originally led Zack et al. (2003)to bulk), fluids operating in subduction zones may propose a model of Li isotopic fractionation bear collateral information on their sources and through loss of aqueous fluids during subduction. processes that take place across the oceanic They found large Li isotopic variations in lithosphere as a consequence of progressive omphacite (a complex Ca–Na–Mg clinopyrox- dehydration and preferential loss of isotopically ene) from eclogites (δ7Li between −11 and heavy Li into the fluid (aqueous) phase for which +5.3 ‰, with negative δ7Li measured in 9 out of 7Li has greater affinity than 6Li (Yamaji et al. 11 whole-rock samples) and interpreted these 2001; Wunder et al. 2006). Ground truth for this findings as a result of preferential loss of 7Li into relationship is inferred from the isotopically the aqueous phase, leaving the residual solid heavy muds exuded by serpentinite seamounts isotopically light Li. However, it is not yet clear (Benton et al. 2004). whether Li should be fractionated during sub- A prominent feature of Li is its dualistic duction at all. Employing more recent experi- behavior as (i) a moderately incompatible element mental constraints, Marschall et al. (2007b) in melting process (Brenan et al. 1998a) and (ii) a examined a large set of worldwide orogenic fluid-mobile element in aqueous transfer (Berger ecogites and provided an alternative interpreta- et al. 1988; Brenan et al. 1998b). Therefore, the tion of their Li isotopic signatures. They esti- changes in its distribution during dehydration can mated a maximum δ7Li decrease of *3 ‰ in track the ‘wet’ history of the respective lithology eclogites as a consequence of metamorphic and provide constraints on the nature of hydrously dehydration reactions and stressed that much of altered crust in early stages of subduction and the δ7Li variation towards lower values observed beyond (Marschall et al. 2007a). An early exper- in eclogites may be accommodated through imental work of Seyfried et al. (1998) revealed kinetic re-distribution and addition of Li from large variations in δ7Li for hydrothermal ambient country rocks. 138 5 Lithium in the Deep Earth: Mantle and Crustal Systems

However, not all δ7Li variations in Sect. 5.2.2) incorporated high-MgO lavas, less subduction-related lithologies can be explained prone to wall-rock contamination, in their Cen- by the mobility of Li isotopes via diffusion (e.g., tral American arc study, the samples had δ7Li of Agostini et al. 2008), and it is becoming roughly between +4 and +6 ‰. With a data set increasingly clear that Li may be retained in that included three intra-oceanic arcs and sam- subducted residues which resist attenuation of ples ranging from 43.9 to 64.7 wt% SiO2, the original Li isotopic signature through diffu- Tomascak et al. (2002) illustrated a general lack sion over extended time scales (e.g., Marschall of correspondence between Li isotopes and et al. 2006, 2007a; Chan et al. 2009; Vlastélic indicators of slab fluid involvement (B/Be, et al. 2009; Tang et al. 2012). Moreover, physical Ba/La, Li/Y). constraints such as the angle and rate of sub- The results of these earlier investigations pro- duction, thickness of oceanic crust and sedi- duced a pessimistic picture of the viability of arcs mentary cover, temperature, heat flow, age, for elucidating the pathways of Li cycling. More degree of alteration, etc. all play roles in deter- recent studies, however, produce a more complex mination of the fate of Li in the subduction (and perhaps somewhat more optimistic) view (see factory. Tomascak 2015). Primitive high-MgO tholeiites The pilot studies of Li systematics at con- and basaltic andesites (>6 wt% MgO) from two vergent margins were mixed on whether impor- different volcanic systems in the southern Cas- tant across-arc variations occur in volcanic lavas. cades oriented perpendicular to the volcanic front Moriguti and Nakamura (1998) noted strong were formed by near-anhydrous melting of spinel correlation of δ7Li and slab depth in the Izu– peridotite (Grove et al. 2002) and yet showed Bonin arc, concluding that the samples reflected decreasing δ7Li with increasing depth of the mixtures of an isotopically light mantle compo- Wadati–Benioff zone (i.e., increasing distance nent with an isotopically heavy slab fluid com- from the arc trench; Magna et al. 2006b). By ruling ponent. Their initial interpretation would require out contamination by continental crustal materials, an uncontaminated mantle with δ7Li * 0 ‰, addition of slab-derived Li with apparently frac- which is inconsistent with our understanding of tionated signature was suggested. A report by the pristine mantle. Of note, the two isotopically Leeman et al. (2004) for a suite of tholeiites, lightest lavas were rhyolitic in composition. If, basaltic andesites and nephelinites from the therefore, these samples reflect some degree of northern section of the Cascades across southern crustal contamination, the implied mantle Washington (some 500 km north of Mt. Shasta end-member shifts to δ7Li * +2 ‰. region studied by Magna et al. 2006b) revealed no Tomascak et al. (2000) also examined litho- apparent across-arc variations in δ7Li. Indeed, the logically heterogeneous arc volcanic rocks (all study of Leeman et al. (2004) traced the weakened with SiO2 > 50 wt%) and reached different signal of material from incoming slab to vanishing conclusions on the nature of Li delivery to the amounts in the back-arc region by using trace sources of arc magmatism. The lack of correla- element parameters (e.g., Li/Y, B/Nb, Ba/Nb) tion between δ7Li and typical indicators such as combined with isotopic compositions of B and B/Be among the Panama arc samples was inter- radiogenic lithophile elements (Sr–Nd–Pb). preted to reflect in general a sequestration of It is interesting to note that similar discrep- slab-derived Li in the subarc mantle. That some ancy between Li isotopic systematics in two samples showed elevated δ7Li (up to +11.2 ‰) distinct sections of a single arc system (or spa- despite low B/Be was interpreted to reflect the tially closely related arcs) was also found for the influence of tectonics and the thermal state of the northeastern Japan arc which is nearly perpen- arc, with slab-derived 7Li-enriched mantle being dicular to the Izu arc. Compared with tapped after it had lost most of its slab B. well-defined negative correlation between δ7Li Whereas Chan et al. (2002b) (a re-reporting of and distance from the trench found for the Izu arc corrected data from Chan et al. 1999; see (Moriguti and Nakamura 1998). Moriguti et al. 5.1 Lithium in the Mantle 139

(2004) found no across-arc variations in δ7Li. somewhat contrasting results. Gurenko et al. Thus, more generally, it appears that there may (2005) analyzed glass inclusions hosted from a be substantial differences in the products of arc rhyolite from the island of Dominica; they volcanism in terms of their Li inventory. These obtained generally lower δ7Li values for inclu- differences may, in part, stem from distinctively sions entrapped in orthopyroxene (+4.2 to different sedimentary inputs that could vary +11.6 ‰) relative to those in plagioclase (+7.4 to greatly not only among different arc systems but +14.7 ‰), paralleled by variable but overall high also along a single arc (Bouman et al. 2004; Li contents (20–60 ppm). Hydrothermally altered Chan et al. 2006). The Li budget of the complete volcanic sequences of the subducted oceanic sedimentary section entering individual arc seg- crust were advocated to contribute significantly ments should thus be ascertained in order to to 7Li-enriched character, also supported by high better understand (or indeed to more accurately δ11B values. The lower δ7Li values in quantify) the Li cycle in arcs. No better charac- plagioclase-hosted inclusions were interpreted to terization of a sedimentary component in the reflect post-solidification kinetic effects that did source of arc magmas has been made than the not affect inclusions in orthopyroxene (Gurenko study of Tang et al. (2014a). Their results show et al. 2005). Somewhat different results were conclusively that isotopically light sea floor presented for St. Vincent Island, c. 300 km south sediment transferred its Li payload to the subarc of Dominica (Bouvier et al. 2008). mantle, with approximately 4 % sediment com- Olivine-hosted glass inclusions from St. Vincent ponent in the lightest lavas, as deduced from had consistently lower Li abundances (2.6– relationships between δ7Li and Y/Li. Similar 9.2 ppm) compared to those examined by Gur- element-isotope coupling was demonstrated by enko et al. (2005), and δ7Li appeared to vary the earlier findings of Clift et al. (2005) from greatly from MORB-like values (roughly Costa Rican tephra. between +2 and +6 ‰) towards significantly Agostini et al. (2008) analyzed Li–B isotopes lighter ones (as low as −10 ‰; Bouvier et al. in a suite of alkali-rich basaltic lavas through 2008). These latter authors advocated for fluids calc-alkaline to ultra-potassic lithologies from the from dehydrated altered oceanic crust, without convergent margin of western Anatolia (Turkey) appeal to Li-rich sedimentary lithologies. To that revealed a significant range in δ7Li values account for the abundances of Cl and S observed (>15 ‰), paralleled by moderate to high Li in some melt inclusions, seawater-like fluids abundances (7–53 ppm). The model developed to were invoked in the mantle source of such melts. explain these results involved the interplay of The lack of a distinguishable Li elemental and/or tectonics on the variable role of slab-derived isotopic signature from such a source was con- fluids. The authors ruled out kinetic modification sidered a reasonable consequence of the low Li of the Li isotopic signature in these rocks based content of seawater; at very low mass fraction, on other geochemical parameters (δ11B, B/Nb, such a component would be incapable of per- B/Be, etc.), as well as a general skepticism of the turbing Li systematics. Collectively, it appears scale of such effects. They also sidestepped the that understanding the exact process of transfer conclusion of Marschall et al. (2007b), that of fluids with peculiar Li systematics and trans- subducting slabs will not develop very isotopi- port efficiency from the subducted slab into and cally light signatures by a distillation process, by through the overlying mantle wedge still requires implying that the cessation of subduction would experimental work as well as field studies. more effectively dehydrate the slab. In this case, A rapid rise of fluids produced in the dehydration protracted dehydration would produce slab fluids of a subducting slab, during discrete events, has with δ7Li < MORB (Agostini et al. 2008). been invoked by John et al. (2012). An estimate Similar conclusions have been reached by of *200 years for fluid transport through the analysis of melt inclusions in arc lavas, although system was derived from modeling diffusive Li two studies of the Lesser Antilles arc returned elemental and isotopic data from veined 140 5 Lithium in the Deep Earth: Mantle and Crustal Systems subduction zone rocks in China, taking into were analyzed for their Li contents and isotopic account changes in Li distribution and porosity. compositions by Košler et al. (2009). These Dehydrated seawater-altered oceanic lithosphere authors did not find any significant variation was inferred as a source of these fluids although beyond the values reported elsewhere (Li = 4.6– the exact origin (volcanic section of the oceanic 7.3 ppm; δ7Li = +3.3 to +4.9 ‰). Two outliers crust or lithospheric mantle) could not be with somewhat more variable δ7Li (+1.5 and unequivocally distinguished (John et al. 2012). +6.8 ‰) were explained through the incorpora- A limited communication between the resid- tion of either a sedimentary component or ual dehydrated slab and the sources of back-arc hydrothermally altered oceanic crust, as also lavas appears to be a distinctive signature of Li inferred from trace element systematics, but it isotopic system when compared to volcanic could also reflect local inhomogeneity at rocks from the arc front. Whilst Li systematics intra-sample scale. Janoušek et al. (2010) inves- have been explored in a number of studies of tigated the Li geochemistry and isotopic sys- convergent margins, little effort has been made in tematics of back-arc lavas from Nicaragua and characterizing the fate of Li in back-arc regions adjacent Caribbean Sea occurrences. The homo- where slab-derived liquids may already impart geneous δ7Li values of these samples (+3.8 to insignificant control on the bulk chemistry of +5.2 ‰) underscore the lack of mass transfer back-arc magmas. The first reported data from from frontal arc systems into the source region of back-arc region of the Lau basin (Chan et al. back-arc lavas. A localized sub-group of samples 2002b) centered at δ7Li * +4.5 ‰ and modest shows a slightly lighter average Li isotopic sig- Li contents (*5 ppm) that are at the lower end of nature (δ7Li * +2 ‰), paralleled by resolved what has been reported for subduction-related differences in Sr–Nd–Pb isotopic compositions magmas. Further analyses of back-arc lithologies compared with the rest of the suite. This may (OIB-like basalts, basaltic andesites, high-K imply incorporation of a seafloor sediment calc-alkaline basalts) were performed by Lee- component into the source of some Nicaraguan man et al. (2004) for the northern Cascades but back-arc lavas. Walker et al. (2009) noted simi- neither Li contents (5.6–11.1 ppm) nor δ7Li larly light Li isotopic compositions values (+2.9 to +4.9 ‰) implied distinctive (δ7Li < +3 ‰) in a subset of behind-arc lavas variation in Li isotopes that could be caused by from southeastern Guatemala and western Sal- large-scale influx of subduction-related compo- vador whilst volcanic front lavas from the same nents. The MORB-like δ7Li values were, in part, area generally had MORB-like δ7Li. Collec- paralleled by δ11B values (<−8 ‰) more typical tively, these studies have shown uniformly that of common OIB (Leeman et al. 2004). mass transfer between the subduction region and Rift-related Anatolian alkalic lavas (6–9 wt% the back-arc environment apparently is subordi- MgO) (Agostini et al. 2008) showed moderate Li nate and does not result in volumetrically sig- enrichment (*11 ppm) and broadly MORB-like nificant source components with isotopically δ7Li (*+3 ‰ save for one sample at −0.4 ‰). distinguishable signature in back-arc lavas. The preliminary results of Brens et al. (2013) for Whereas there is an increasing number of an across-arc profile in the Tonga–Kermadec arc studies that show discernible Li isotopic variations show large δ7Li scatter in arc lavas as a conse- in arcs, the number of locations with δ7Li essen- quence of possible fluid transfer from sediments tially equivalent to MORB is significant. This (consistent with trace elements constraints) while observation begs the question: why does such the Fonualei back-arc lavas appear to have δ7Li limited variation occurs so commonly? If slab Li is indistinguishable from the MORB range. transferred into the sources of arc magmas, either Two studies to date have focused exclusively its net isotopic signature is sufficiently similar to on arc lavas related to processes behind the MORB that it leaves no isotopic trace or there is an volcanic front. Back-arc Cenozoic alkaline lavas isotopic fractionation in the subduction process from the James Ross Island (Antarctic Peninsula) for which we are inadequately accounting. The 5.1 Lithium in the Mantle 141 following scenarios summarize the most probable (3) Isotopic disequilibrium from kinetic pro- resolutions to arcs where no Li isotopic “slab cesses during slab fluid release and/or signature” is exhibited. transport. The model proposed by Mar- schall et al. (2007b) suggests minimal iso- (1) The Li isotopic signal of subducted material topic fractionation into fluids from high is attenuated before reaching the region of pressure metamorphism of subducted sedi- melting in the subarc mantle. As proposed ments. This implies that fluids separated originally by Tomascak et al. (2000, 2002), from metamorphosed MORB would lack a this model calls for the fluid mobility of slab distinctive isotopic signature. Similarly, Li to be counteracted by its moderate isotopically light fluids en route to the compatibility in mantle silicate minerals. region of melt generation in the subarc Evidence in favor of this may be manifest in mantle might preferentially lose an isotopi- certain xenoliths from the subcontinental cally light component through diffusion, lithosphere (e.g., Tang et al. 2007), whereas potentially wiping out or at least reducing Plank (2014) argued against this mechanism the slab signature. on the basis of elevated Li/Y in all arc (4) Contamination and isotopic re-equilibra- magmas. tion during melt transport to and storage (2) Equilibrium slab inputs may not be iso- within the crust. Walker et al. (2009) topically distinct from MORB. Based implied that much of the δ7Li variation mainly on elemental systematics, Plank within part of the Central American arc (2014) argued that a slab Li signature is system was likely related to kinetic over- reasonably well preserved in most arc set- printing of original Li isotopic systematics. tings. The original studies on the topic Similarly, Cabato et al. (2013) indicated predicted isotopically heavy slab fluids, that variations in Li–Be–B concentrations, dominated by altered oceanic crust (Mor- coupled with Li isotopes, in plagioclase iguti and Nakamura 1998; Tomascak et al. phenocrysts from Santorini, are best 2000; Chan et al. 2002b). Owing signifi- explained by late-stage processes (e.g., cantly to the paucity of early data on marine magma chamber decompression). Detecting sediments, as well as the preliminary indi- Li isotopic influences from this group of cations that such materials were isotopically processes is facilitated by in situ analysis of heavy compared to MORB (Zhang et al. crystals, where diffusion profiles can be 1998), the Li perspective on the role of assessed relative to the host magma. sediments in arc magma petrogenesis was initially hampered. With an improved esti- mate of Li in global marine sediments (mean δ7Li * +4: Chan et al. 2006; +2.4: 5.2 Lithium in the Continental Plank 2014), it is clear that, depending on Crust the incoming lithologic assemblage, the incorporation of sedimentary source com- 5.2.1 Products of Crustal Melting ponents in arcs may generate imperceptible isotopic effects (indeed it appears that a Broadly speaking, melting of the continental crust sediment signal traceable with Li isotopes produces silica-rich magmas, many of which form may be rare; Tang et al. 2014a). Addition- major granitic bodies. Attempts to use elemental ally, mixtures of isotopically heavy altered and isotopic geochemistry to unravel the sources oceanic crust components with isotopically and petrogenetic histories of such bodies have met light sediment components are likely to with mixed results (Taylor and McLennan 1985), yield fluids with broadly MORB-like Li although a paradigm that is widely used separates isotopes. granitic rocks into generally distinct groups (e.g., 142 5 Lithium in the Deep Earth: Mantle and Crustal Systems

Chappell and White 2001). The reader is directed separation of mafic minerals like biotite and to summaries such as that of Barbarin (1999)for amphibole was inadequate to produce measurable more detail on granite typology. Despite the sig- (±1 ‰) isotopic variation. This conclusion has nificance of granitic rocks to the architecture of been echoed by other studies of granitic suites the continental crust and their general enrichment (Teng et al. 2004, 2009; Magna et al. 2010). This in Li relative to the mantle, it has only been is good news, in that it verifies that granitic rocks recently that Li isotopic research has turned to this should faithfully inherit the Li isotopic composi- subject in earnest. Some of these studies have tions of their sources. focused on using granitic bodies as grand aver- Studies conducted in terranes with multiple agers of the bulk continental crust, whereas others granite types have yielded no consistent conclu- have attempted to understand the extent to which sion with regard to the interpretation of Li iso- Li isotopes may inform us about both source and topic compositions. Lithium isotopes in magmatic processes. Australian S-type samples were more homoge- Although before the era of well-quantified Li neous than those of I-type; δ7Li in the I-type isotope measurements, Plyusnin et al. (1979) granites examined overlapped completely the reported on Li isotopes in granitic rocks. Their S-type range, but extended to substantially higher results suggested that the more geochemically values (up to c. +8 ‰; Bryant et al. 2004). evolved samples of a suite of evolved granites A smaller data set of samples of southeastern had the lightest Li isotopic signatures. They Australian granites (Teng et al. 2004) reproduced suggested that a linear relationship between the the offset between isotopically lighter S-type and Li isotopic composition and the content of F (and heavier I-type suites, although the mean values of Li) in a suite of granitic rocks was best explained δ7Li for both were lower (Fig. 5.6a). This would by closed system fractionation. This interpreta- be largely consistent with the Li isotopic sys- tion appears out of step, however, with more tematics of the precursors of I and S-type gran- recent studies of high-F granitic suites (e.g., ites: the former granites would reflect Romer et al. 2014) and of granites in general. meta-igneous lithologies with a predominantly Nevertheless, the suggestion that Li isotopes positive δ7Li range whilst generally isotopically could be used to distinguish rocks that had expe- light meta-sedimentary parentage would be rienced metasomatic or other secondary phenom- sought in the latter granites. However, granitic ena from those which record purely magmatic rocks from the Western Carpathians (Magna histories (Plyusnin et al. 1979) remains viable. The et al. 2010) showed a contrary Li isotopic rela- calculations of Teng et al. (2006b) support this tionship, with much more variable signatures of relationship: minerals that form from late-stage S-type granites than I-type. It should be noted fluids or interact with sub-solidus fluids should be that the majority of these studies concentrated on discernible through the development of heavier dispersed granitic bodies, typically using a single isotopic compositions. Data from the peralkaline representative sample for an individual pluton. In Ilímaussaq complex in Greenland (Marks et al. cases where multiple samples were analyzed 2007) support the action of late-stage, externally from one intrusion, internal δ7Li variation of 3– derived fluids to account for isotopically heavy 4 ‰ was not uncommon (e.g., Qingshui, China: pluton margin samples. +0.05 to +4.6 ‰ for six samples; Teng et al. The first systematic study of granitic rocks in 2009). The extreme manifestation of this came the more modern analytical era came from Bryant from the detailed study of the layered Harney et al. (2004), with a data set from the New England Peak Granite intrusion, where Teng et al. (2006b) Batholith of southeastern Australia. They exam- found a total range in δ7Li of 9.7 ‰ for 25 ined both S-type (broadly sedimentary-derived) samples (Fig. 5.6b). Thus, the results of studies and I-type (broadly igneous-derived) sample of granitic terranes which use individual samples suites. Their results provided further evidence that to represent entire plutons may be misleading isotopic fractionation driven by fractional and should be interpreted with great caution. 5.2 Lithium in the Continental Crust 143

(a)

(b)

Fig. 5.6 Lithium concentration and δ7Li in rocks of from Teng et al. (2006), Variscan granites from Romer broadly granitic composition, ranging from true granites to et al. (2014), Greenland, including syenitic compositions, more intermediate and more alkaline compositions, and from Marks et al. (2007), Wimpenny et al. (2010), their volcanic and metamorphic equivalents. a Granitic composites are pooled samples from Teng et al. (2004), samples classified using the I, S, A nomenclature. Data “other” (includes data from a number of globally dispersed sources: Australia from Bryant et al. (2004), Teng et al. locations) from: Bottomley et al. (2003), Sturchio and Chan (2004), Western Carpathia from Magna et al. (2010), China (2003), Millot et al. (2007), Bryant et al. (2004), Teng et al. from Teng et al. (2009). b Other igneous rocks of broadly (2004), Magna et al. (2010), Kisakurek et al. (2004), granitic composition. Data sources: Harney Peak granites Rudnick et al. (2004) and Godfrey et al. (2013)

Bryant et al. (2004) found their data consistent compared to I-type plutons. The latter they took with the more isotopically homogenized sedi- to represent what were originally mantle signa- mentary sources of the S-type granite suites tures variably contaminated by isotopically light 144 5 Lithium in the Deep Earth: Mantle and Crustal Systems upper crust. Magna et al. (2010) underscored that isotope fractionation (e.g., Foustoukos et al. granites with substantial source input from the 2004), experimental data indicate measurable lower crust (which, they point out based on the mineral–fluid isotopic partitioning under geo- data of Teng et al. (2008), has been shown to logically relevant conditions (Lynton et al. 2005; exhibit, “rampant Li isotope variability”) should Wunder et al. 2006, 2007). Nevertheless, leuco- be expected to possess a degree of isotopic granitic rocks analyzed by Bryant et al. (2004) heterogeneity. Indeed, that granitic suites show showed no marked Li isotopic offset from the as limited isotopic variation as they do suggests ordinary granites they cross-cut. The absence of an important disconnect from the lower crustal extreme fractionation among evolved and min- xenolith data set. Magna et al. (2010) ultimately eralized granites of central Europe (Magna et al. concluded that there is no simple, universal 2010; Romer et al. 2014) further tempers interpretation of Li isotopes among S- and I-type expectations that fluid–melt processes will granite suites. invariably strongly affect Li isotope systematics. Teng et al. (2009) indicated that Li isotopes Teng et al. (2006b) suggested that the degree of appear unable to resolve the ongoing uncertainty fluid exsolution would need to be high in order to over the derivation of I-type and A-type granites produce significant late-stage Li isotopic frac- (classically considered “anorogenic” granites). In tionation in granitic systems. their study of A-type granitic rocks in northeastern The exsolution of fluids or low-temperature China, Li isotopes were not useful in constraining melts from magma is inferred in many models petrogenesis or otherwise improving conclusions attempting to explain the geochemistry of highly based on other geochemical data. Romer et al. fractionated igneous systems, including granitic (2014) reached similar conclusions based on the pegmatites (e.g., Sirbescu and Nabelek 2003), majority of central European granites in their although this is only one potential mechanism for study. Only Magna et al. (2010) were able to use Li Li isotopic fractionation in such systems. As isotopic data to significantly constrain interpreta- pegmatites are among the best examples of tions of the petrogenesis of A-type granites from extreme enrichment of Li in the crust, it is the Western Carpathians. The Li isotope–element somewhat paradoxical that they too have gone relationships allowed Magna et al. (2010) to rule largely isotopically overlooked prior to very out several potential petrogenetic models, includ- recently. Although there were early indications ing some which involve “standard” A-type granite that pegmatites may record significant Li isotopic origins (e.g., reactivated lower crust and fractional fractionation (Vocke et al. 1990; Tomascak et al. crystallization of primary mantle melts). In their 1995), only more recent studies show this clearly study, relatively high values of δ7Li (+1.1 to and begin to offer explanations. Whereas the two +6.7 ‰) were essential to their deductions. In pegmatitic rocks analyzed by Bottomley et al. other areas A-type granites show generally lower (2003) and Magna et al. (2010) were notably δ7Li, which, because of overlap with the isotopic isotopically heavy compared to ordinary granites composition of sedimentary and metasedimentary (δ7Li of +30.2 ‰ and +8.9 ‰, respectively), rocks, makes it effectively impossible to simple pegmatites in the Harney Peak Granite unequivocally rule out a variety of possible had a similar range in δ7Li as the granites with sources. which they coexist (Teng et al. 2006b). However, Are there conditions under which granitic whole rocks and minerals from the Tin Mountain rocks may not bear a true isotopic signal of their pegmatite are uniformly isotopically heavier than parentage? Bryant et al. (2004) posited that iso- all samples from the parental Harney Peak topic variability in a limited number of Aus- Granite. On average higher δ7Li in two peg- tralian leucogranite samples could have been matites relative to those in neighboring country introduced by isotopic fractionation driven by rocks were also revealed in study by Liu et al. fluid separation. Although it has been argued that (2010) and systematic decrease away from peg- vapor separation introduces no significant Li matites was explained by diffusion. 5.2 Lithium in the Continental Crust 145

Of the small number of Li isotopic studies on reasons. Gordienko et al. analyzed different pegmatites to date, the majority deal not with minerals in the different zones (amphibole, issues of granite–pegmatite source relationships, pyroxene, mica). It is clear that, due to the but with internal processes. There is still con- crystallographic influence on isotopic fractiona- siderable uncertainty on aspects of the internal tion, coexisting minerals in which Li inhabits evolution of large, zoned granitic pegmatites differently coordinated sites will possess distinct (London 2008, 2009); there has been some hope δ7Li despite crystallizing from the same medium that Li isotopes could shed light on this. Lithium (Teng et al. 2006b). The fact that δ7Li in spo- isotopic variations among pegmatite minerals dumene samples from the two localities exam- hinges on a variety of factors, many of which are ined by Gordienko et al. (2007) showed limited difficult to quantify with precision. The overrid- variation (0.2 ‰ difference for one pair, 2.4 ‰ ing control, however, appears to be the difference for the other) appears to highlight their miscon- between the bonding of Li in the crystallizing ception. Nonetheless, Barnes et al. (2012) show medium (melt or fluid) compared to the that of that the positive correlation between bulk peg- the crystallographic site in a given mineral matite F content and δ7Li, although weak, is (Wenger and Armbruster 1991; Wunder et al. consistent with inferences from melt structure 2007, 2011); heavy Li appears to partition (Henderson 2005). The role of F-complexing in a favorably into less coordinated sites with shorter silicate melt may have implications for the Li–O bonds. If Li in silicate melts is predomi- development of distinct isotopic signatures of nantly in tetrahedrally coordinated groups (Sol- highly evolved pegmatite populations. Reflecting tay and Henderson 2005), then minerals that back to earlier studies, although Plyusnin et al. incorporate Li in sites of higher coordination (1979) suggested a link between Li isotopic should be isotopically lighter than the melt from fractionation and F content, their prediction was which they crystallize (Teng et al. 2006b) in the opposite sense (i.e., that higher δ7Li would because of the difference in bond lengths result in systems with lower bulk F). involved. Generally positive correlation between δ7Li In addition to equilibrium effects, it has been and concentration of Rb (as well as Li) in bulk suggested that kinetic effects may also be pegmatitic rocks in and around the Tin Mountain important in the Li isotopic geochemistry of pegmatite in South Dakota was interpreted by granitic pegmatites. This follows from the like- Teng et al. (2006b)toreflect the effects of large lihood that rapid crystal growth is a typical fea- extents of equilibrium crystal fractionation in the ture of pegmatites (London 2009). Such a driving formation of pegmatite magmas. The same extent force was suggested by Maloney et al. (2008)to of fractionation is not apparent within evolving explain the 3–10 ‰ difference between elongate pegmatites, however. For example, minerals tourmaline samples and those exhibiting radial crystallized in more elementally fractionated growth. They interpreted the tourmaline data to mineral assemblages of the Tin Mountain peg- have been produced in a system where Li diffu- matite had δ7Li that overlapped values from the sion was outpaced by crystal growth. Alterna- same minerals (quartz, plagioclase, muscovite) in tively, it could point to crystallography-assisted less fractionated zones. The same lack of sig- fractionation along different axes, as has been nificant interzonal isotopic fractionation was seen reported for boron (Hinsberg and Marschall among tourmaline samples from southern Cali- 2007). fornia pegmatites (Maloney et al. 2008). Thus, Based on positive linear correlations between from the data accumulated to date, there is no mineral δ7Li and the K/Rb of coexisting miner- evidence that Li isotopes can aid in better als, Gordienko et al. (2007) inferred isotopic understanding the internal crystallization histo- fractionation in response to progressive peg- ries of granitic pegmatites. matite crystallization. Although this conclusion Teng et al. (2006b) also measured Li isotopes may be valid, it was certainly for the wrong in amalgamated fluid inclusions from quartz, 146 5 Lithium in the Deep Earth: Mantle and Crustal Systems concluding that the offset between fluids (+8.1 to isotopic behavior in minerals (Wunder et al. +13.4 ‰) and the pegmatite wall zone bulk rock 2011), the relationship had not previously been with the lowest δ7Li (+7.5 ‰) could reflect demonstrated empirically. Magna et al. equilibrium fluid–mineral fractionation. The (2013) showed that shorter bond lengths would pooling of magmatic and possibly much later favor heavy isotopes, regardless of the specific fluids in these analyses, however, creates signif- mineral or specific pegmatite, and consistent with icant uncertainty in their isotopic interpretation. experimental evidence (Wunder et al.2011). Masukawa et al. (2013) analyzed Li and Sr iso- Second, δ7Li in residual pegmatite liquids does topes in a set of ore-forming fluids from inclu- not evolve in a particular direction. Instead, δ7Li sions in vein quartz from the Takatori mine, of co-crystallizing phases may differ substantially Japan, in order to understand possible sources of (cf. muscovite and beryl as the first phases to tungsten ores. By selecting temporally disparate solidify consistently showing the lowest and samples, Masukawa et al. (2013) associated low highest δ7Li, respectively). Third, significant δ7Li together with high contents of metals (W, variations in δ7Li (up to 20 ‰) have been found Ni, Cu, etc.) in the earlier stages of hydrothermal for the same mineral species from different peg- mineralization as a result of sourcing an S-type matite bodies. Although the cause is unknown, it granitic magma. Later phases of mineralization appears to be associated with sourcing distinc- with progressively higher δ7Li were probably tively different precursors. However, the exact more consistent with exchange of Li between link between individual pegmatites and their metasediments with previous seafloor alteration possible source regions cannot be easily pro- and ore-bearing fluids, coupled with concomitant posed. Alternatively, it may hinge on the extent precipitation of Li-bearing phases. of melt/fluid–solid Li isotopic fractionation dur- Shabaga et al. (2010) used SIMS techniques ing the evolution of Earth’s continental crust to determine the Li isotopic composition in coupled to physical parameters (e.g., differences tourmaline from Brazil, Nigeria, and Mozam- in heat production and transfer, paleogeography). bique in an attempt to establish Li as a prove- The latter interpretation takes into account that nance tool for precious gemstones. These authors large pegmatite bodies are predominantly implied significant differences in coupled Archean and Proterozoic phenomena; average δ7Li–δ11B relationships that might help unravel post-Proterozoic pegmatite occurrences are sig- source regions for certain type of tourmaline. nificantly smaller (e.g., Tkachev 2011; Bradley However, a later study of Ludwig et al. (2011), and McCauley 2013; McCauley and Bradley conducted on the same samples with the same 2014). analytical approach, revealed distinctly different Li and B isotopic compositions; the three loca- tions could still be resolved with coupled Li–B 5.2.2 Crustal Metamorphism isotopic systematics, however. It follows that and Fluid Flow such analytically-based differences (isobaric molecular interferences) have the potential to Metamorphic effects are commonly so intimately seriously affect geological explanations. related to the igneous process that discriminating Magna et al. (2013) made a comprehensive the geochemical consequences of the two con- study of minerals from a variety of granitic dition sets is problematical. Attempts to use Li pegmatites representing varied crystallization isotopes to understand aspects of the metamor- ages from Archean to recent. Their preliminary phic histories of crustal rocks have had to deal results are three-fold. First, within individual with distinct potential mechanisms of isotopic pegmatite bodies, the Li isotopic composition of fractionation: those related to processes that can minerals is correlated strongly with Li-φ (O, OH, be quantified through equilibrium models, and F) bond length. Although it was known that bond kinetic models, including those that rely on the distance was a significant component of Li important difference in diffusion rates of 6Li and 5.2 Lithium in the Continental Crust 147

7Li. In certain instances, the magnitude of the Catalina Schist (Penniston-Dorland et al. 2012), isotopic effects brought on by diffusional pro- California, is indistinguishable from that of cesses appears to create a diagnostic signature, clastic sedimentary rocks from Teng et al. although modeling the effect does not universally (2004), this in spite of the Catalina rocks being reach a satisfactory conclusion. metamorphosed up to *750 °C. The relative Contact zones around plutonic bodies have homogeneity in Li isotopic composition across been popular targets for understanding metamor- metamorphic grade for the Catalina Schist, cou- phic processes, particularly cases where the host pled with only minor decrease in Li concentra- rocks are lithologically homogeneous and well tion, was interpreted to reflect predominance of defined (Nabelek et al. 1984; Ferry 1994), and so it relatively closed-system behavior is logical that these geological scenarios have been (Penniston-Dorland et al. 2012). Lithium employed in Li isotopic studies. Romer et al. released with reactions consuming lower-grade (2005) discussed the geochemistry of an intrusion chlorite was taken up, without isotopic effects, in of alkali syenite (and pegmatite) magma into the crystallization of higher-grade phengite. Xiao dominantly carbonate rocks in Mexico. The study et al. (2011) found restricted Li isotopic variation encompassed a wide array of elemental and iso- between eclogites and coexisting high pressure topic tools and showcased the meter- to veins, as well as between coexisting mantle millimeter-scale complexity that can be recorded peridotite and eclogite, although the data set was in what might appear a geologically straightfor- small. Thus, in abundant metamorphic settings, it ward setting; complexity, in this case, interpreted appears that no significant geochemical driving to be caused by multiple chemically distinct fluids. forces exist to produce substantive variations in The variations in Li isotopes in samples from this Li isotopes. Whether such isotopic effects are contact zone were attributed to inputs from produced relates vitally to the nature of syn- and decarbonation reactions as well as from isotopi- post-metamorphic fluid flow. cally distinct magmatic and metamorphic fluids. On the contrary, Romer and Meixner (2014) In contrast to that study, the intrusion of the interpreted increasing metamorphic temperatures granodioritic Onawa pluton in Maine (Teng et al. to have driven isotopic fractionation in Variscan 2007) introduced minimal Li isotopic variation metapelites from central Europe. They made despite causing kilometer-scale contact meta- clear that the bulk composition (and hence min- morphism up to the formation of migmatites. eralogy) of the protoliths was crucial in deter- Teng et al. (2007) saw no correlated Li isotopic mining the ultimate extent of isotopic change, variation with the extent of contact metamor- owing to the differences in Li availability in the phism (i.e., distance from contact), although original sediment. Sedimentary rocks enriched in contact metapelites experienced apparent Li loss Li may manifest minimal isotopic fractionation during dehydration. This was a somewhat ideal during progressive metamorphism as long as Li study area in that little retrogression took place, in the detrital minerals was primarily structurally which permitted unaltered prograde isotopic incorporated, as opposed to being present on signatures to be preserved. The metamorphic surfaces or other readily-exchangeable positions. dehydration in rocks hosting the Onawa pluton The data of Romer and Meixner (2014) demon- caused only 1–2 ‰ decrease in δ7Li of aureole strate that the process of metamorphism through rocks and was successfully modeled by Rayleigh eclogite facies may introduce only a <2 ‰ fractionation. The model suggested that very decrease in δ7Li, consistent with the model of high fluid/rock Li distribution coefficients cou- Marschall et al. (2007b). A similar magnitude of pled with modest isotopic fractionation factors fractionation was seen in the Onawa aureole, would be required to produce more significant although in other protoliths the total fractionation change in δ7Li in the absence of, for example, introduced was perhaps as great as −20 ‰. input of isotopically distinct magmatic fluid. The Because they contain high Li contents and mean δ7Li of metasedimentary rocks from the commonly exsolve fluids which cause 148 5 Lithium in the Deep Earth: Mantle and Crustal Systems exomorphic effects in their host rocks (e.g., been influenced to some extent by fluid-mediated Morgan and London 1987), many of the studies of exchange with wall rocks involving concomitant metamorphic fluid–rock interaction involve diffusion-driven isotopic fractionation. Assuming granitic pegmatites. Host metamorphic rocks of Li in the melt behaves like Li in an aqueous fluid, the Tin Mountain pegmatite in South Dakota exchange would favor 6Li partitioning into wall (Teng et al. 2006a) were invaded by Li-rich fluids rocks. Such a process would be expected to drive following the intrusion of the spodumene-bearing melt δ7Li to higher values, although no direct pegmatite, resulting in perturbations of their ele- evidence (i.e., analysis of surrounding rocks) was mental and isotopic systematics to a distance of at given to support this speculation. least 30 m. The possibility that these perturbations Light Li isotopic compositions are character- were the result of either mixing (between of fluid istic in eclogite facies rocks, largely regardless of and rock Li) or Rayleigh-type fractionation could protolith (Zack et al. 2003; Marschall et al. be ruled out, as both would require the host rocks 2007b; Xiao et al. 2011; Romer and Meixner to have geochemically unreasonable 2014), with a range from −21.5 to +8.2 ‰ (the δ7Li < −20 ‰. Teng et al. (2006a) instead argued majority falling between −9.7 and +6.9 ‰; that the data were more consistent with Fig. 5.7). Zack et al. (2003) explained the iso- diffusion-driven isotopic fractionation. The hori- topic variations in eclogite facies rocks from the zontal magnitude of the resultant diffusional halo Alps as reflecting a Rayleigh process that took depended on the nature of the host rock, but both place during progressive slab dehydration; a mica schist and amphibolite achieved similar similar conclusion was reached by Bouman et al. “peak” δ7Li (*−20 ‰). (2004) with their modeling of subduction zone A similar conclusion was reached by Liu et al. processes. This interpretation, however, was (2010) in a study of the rocks hosting a series of questioned by Marschall et al. (2007b). Mar- much smaller Li-rich pegmatite bodies in Wis- schall et al. (2007b) modeled Li fractionation consin (2–3 m wide dikes compared to the during eclogite facies metamorphism using *30 m thick Tin Mountain pegmatite). In both experimental data published after the work of the case of Teng et al. (2006a) and Liu et al. Zack et al. (2003). Their results indicated that at (2010), the interpreted process involved Li-rich relevant temperatures and pressure, slab dehy- fluids reacting with surrounding rocks, with the dration was unlikely to involve Li isotopic frac- advance of the faster-diffusing 6Li being recor- tionation of more than *3 ‰. Hence the ded in metasomatized schists and amphibolites as explanation of very isotopically light data from the very light isotopic compositions. Liu et al. most eclogites required other mechanisms. (2010) pointed out that the extent of Li transport Marschall et al. (2007b) proposed two factors into country rocks for the Wisconsin pegmatites to influence Li contents and Li isotopes in sub- was similar to that at Tin Mountain, despite the duction zone metamorphic complexes. First, they difference in scale of the pegmatites. Thus mul- called for significant flux of Li derived from tiple factors must govern the magnitude of Li subducted sediments or sedimentary rocks. These export from Li-rich magmas/fluids into host materials both carry on average higher Li con- rocks. Nonetheless, the results of both Teng et al. centrations than coexisting mafic rocks and have (2006a) and Liu et al. (2010) were inconsistent lower δ7Li (see Sect. 6.3). A prominent source of with volume diffusion, further implicating a enriched, isotopically light sedimentary-source fluid-mediated process. Li was called for by both Penniston-Dorland Maloney et al. (2008) entertained the possi- et al. (2010, 2012) and Simons et al. (2010)in bility that high δ7Li values in minerals from their studies of high pressure rocks from other southern California pegmatites (and hence high subduction zones. δ7Li in pegmatite magmas) may result from pre- A simple fluid exchange model is inadequate or syn-emplacement open system behavior. They to account for the nearly 20 % of eclogite facies suggested that Li in pegmatite magmas may have data with very low δ7Li (< c. −5 ‰). Neither do 5.2 Lithium in the Continental Crust 149

Fig. 5.7 Lithium concentration and δ7Li in high pressure (2007b), Xiao et al. (2011), Western Europe (multiple metamorphic rocks (primarily eclogite facies metamor- localities) from Romer and Meixner (2014), Marschall phosed sedimentary and mafic igneous rocks, with some et al. (2007b), Zack et al. (2003), Franciscan from blueschists). Data sources: Syros from Marschall et al. Penniston-Dorland et al. (2010), Simons et al. (2010) and (2007b), China (multiple localities) from Marschall et al. Guatemala from Simons et al. (2010) geochemically reasonable sources exist to gen- erate fluids sufficiently isotopically light, nor do References known Li isotopic fractionation factors permit isotopically heavier fluids developing such Abdelfadil KM, Romer RL, Glodny J (2014) Mantle characteristics. Although Marschall et al. (2007b) wedge metasomatism revealed by Li isotopes in – – did not quantitatively demonstrate the feasibility orogenic lamprophyres. Lithos 196 197:14 26 Ackerman L, Špaček P, Magna T, Ulrych J, Svojtka M, of diffusive fractionation of Li isotopes in the Hegner E, Balogh K (2013) Alkaline and eclogitic rocks, it seems inescapable that kinetic carbonate-rich melt metasomatism and melting of fractionation driven by diffusive influx of Li is subcontinental Lithospheric mantle: evidence from required to explain such extremely light isotopic mantle Xenoliths, NE Bavaria, Bohemian Massif. J Petrol 54:2597–2633 compositions in rocks with elevated Li contents. Ackerman L, Ulrych J, Řanda Z, Erban V, Hegner E, Such driving forces had not been brought into the Magna T, Balogh K, Frána J, Lang M, NovákJK mainstream when the study of Zack et al. (2003) (2015) Geochemical characteristics and petrogenesis Č é was published. Similarly, Penniston-Dorland of phonolites and trachytic rocks from the esk Středohoří Volcanic Complex, the Ohře Rift, Bohe- et al. (2012) resorted to diffusive fractionation mian Massif. Lithos 224–225:256–271 7 to explain the low δ Li (< c. −4 ‰) of a minority Agostini S, Ryan JG, Tonarini S, Innocenti F (2008) of Catalina Schist samples. In conclusion, it is Drying and dying of a subducted slab: coupled Li and clear that diffusion-driven isotopic fractionation B isotope variations in Western Anatolia Cenozoic Volcanism. Earth Planet Sci Lett 272:139–147 is an important consideration in understanding Li Aulbach S, Rudnick RL, McDonough WF, (2008) Li–Sr– dynamics in many metamorphic terranes, Nd Isotope Signatures of the Plume and Cratonic although the quantification of this process has not Lithospheric Mantle Beneath the Margin of the Rifted yet been rigorously evaluated in most settings. Tanzanian Craton (Labait). Contrib. Mineral. Petrol. 155:79–92 150 5 Lithium in the Deep Earth: Mantle and Crustal Systems

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In this section we consider the variations shown compartments: hydrothermal (by which we intend by Li isotopes in materials related to near-surface to describe shallow crustal systems of temperature (classically referred to as “low temperature”) ≥ *200 °C) and, essentially, everything systems, ranging from sediments and sedimentary sub-hydrothermal. We begin with the latter, mov- rocks to natural waters and possible anthro- ing from the ultimate feedstock for surficial waters pogenic impacts. The combination of modest and then move on to other reservoirs. abundance in many minerals and fluid solubility make Li an element of interest in many surficial environments. The strong fractionation of Li 6.1 Meteoric Water isotopes has been exploited substantially in the examination of a variety of near-Earth-surface The Li isotopic composition of atmospheric pre- processes. In addition to their intrinsic value, cipitation has been only rarely reported. To some many of these studies are critical toward the extent this may reflect the low Li concentration in understanding of the cycling of Li between shal- average rainfall (generally <3 ppb; Millot et al. low and deep Earth reservoirs, as, relative to the 2010a), which made analysis more challenging latter, the bulk of the mass fractionation goes on prior to the development of modern analytical in the former. Many studies aim also to develop methods. Single samples of rainfall from Hawaii Li isotopic tools for better quantifying processes (Pistiner and Henderson 2003; δ7Li = +14.3 ‰ in in the shallow Earth, such as geothermometry of filtered aliquot) and Sao Miguel (Pogge von hydrothermal systems. Strandmann et al. 2010; δ7Li = +32.8 ‰) were The division of major classes of processes analyzed, the latter noting the congruence with affecting isotopic distributions inevitably becomes seawater. Although Pistiner and Henderson (2003) somewhat arbitrary, given that the Earth system did not interpret the absolute isotopic composition presents few situations in which a sole process can of the sample they reported, they made the obser- be identified. Thus, although we acquiesce that vationthatfiltration oftheoriginalsampleremoved elemental and isotopic fractionation at hydrother- an isotopically light atmospheric dust component mal temperatures is distinct from that occurring at (δ7Li = +10.2 ‰ in unfiltered aliquot). As will be conditions of shallow burial, diagenesis or surficial seen below, the same relationship is manifest in weathering, many (most?) geological settings riverwaterdissolved/suspendedloadcomparisons. contain superimposed aspects of some or all of To date only one systematic study has con- these pathways. For the purpose of this summary, sidered Li isotopes in rainfall (Millot et al. we have chosen to divide conditions into two 2010a), which produced several interesting and

P.B. Tomascak et al., Advances in Lithium Isotope Geochemistry, 157 Advances in Isotope Geochemistry, DOI: 10.1007/978-3-319-01430-2_6 © Springer International Publishing Switzerland 2016 158 6 The Surficial Realm: Low Temperature Geochemistry of Lithium unexpected outcomes. Analyzing filtered sam- well with generally low Li contents in airborne ples from different physiographic regions in particles, although speciation of Li in atmo- France, Millot et al. (2010a) found that even in spheric deposits may, to a certain extent, depend sea-coast sampling locations marine sources did on climatic conditions (Siggaard-Andersen et al. not dominate the Li budget. The largest contri- 2007). These cumulative findings imply a rather bution of seawater aerosol Li to samples only high reactivity of Li particles in the air and their approached 50 %, and in one of two sites within effective removal from the atmosphere. 100 km of the Atlantic Ocean the proportion was at most *20 %. Lithium in rainfall from sites >300 km distant from the ocean contained at 6.2 Groundwater and Pore Waters most 10 % sea-salt-derived Li. In samples from Clermont-Ferrand in particular it was clear that 6.2.1 Shallow Groundwaters anthropogenically-influenced sources were sig- nificant, with δ7Li as high as +95 ‰ (see Most groundwaters within the upper *100 m of Sect. 6.5.2 for more detail on this process). the Earth’s surface are characterized by low to Witherow et al. (2010) reported δ7Li of sur- modest dissolved solid contents (generally face snow from three Antarctic glaciers <1000 ppm) in all but very arid regions (δ7Li = +0.8 to +2.9 ‰), and Pogge von (Fig. 6.1). Relatively short storage times limit the Strandmann et al. (2006) analyzed Li in one role of mineral–fluid reactions in shaping the water sample directly derived from melting of geochemistry of these waters compared to their glacial ice in Iceland (δ7Li = +33.3 ‰). All of deep brethren. these waters have very low Li concentrations In arid areas the lack of abundant recharge (<1 ppb). Whereas the Iceland sample might coupled with evaporation may produce ground- have seawater aerosols as a major source of Li, waters that are, in some cases, as saline as for- clearly the Antarctic samples do not. The low Li mation waters. For example, none of the Great abundance in modern rain water corresponds Artesian Basin groundwaters sampled by

Fig. 6.1 Lithium concentration and δ7Li of shallow groundwaters. Data sources: France from Negrel et al. (2012), Andes from Godfrey et al. (2013), Great Artesian Basin (GAB) from Meredith et al. (2013), “other” (includes data from four separate localities) from Hogan and Blum (2003), Kloppmann et al. (2009), Rad et al. (2013), Tomascak et al. (2003) 6.2 Groundwater and Pore Waters 159

Meredith et al. (2013) was ancient (all <10 ka), composition, but tended to be characterized by but the total dissolved solid content of most δ7Li ≥ +30 ‰. In the absence of marine sources, samples was >30,000 ppm. The deepest of these such heavy Li signatures may derive from sorp- waters was withdrawn from only 126 m. tion reactions or through the formation of low Water-sediment or water-rock interactions temperature minerals in the shallow subsurface; have the capacity to complicate interpretation of without more detail on the mineralogy of aqui- Li isotope data from shallow groundwaters. In a fers, neither interpretation can be invalidated. feasibility study for the use of intrinsic stable Whereas most of the focus on Li isotopic isotopic tracers in groundwaters, Kloppmann fractionation during mineral crystallization at low et al. (2009) interpreted that Li isotopic signa- temperatures centers on silicates, Huh et al. tures showed cryptic effects of interaction (1998) suggested that the crystallization of car- between injection water and sediments. With bonates or halides could drive Li isotopic frac- prolongation of water-sediment interaction, Li tionation. The laboratory results of Marriott et al. was removed from groundwaters, and isotopic (2004a) appear to support this in terms of arag- signatures shifted to lighter values by c. 15 ‰. onite, and to a lesser extent calcite. One attempt Similar cryptic isotopic effects were seen by to quantify fractionation effects during halite Meredith et al. (2013). The waters in the latter crystallization from brine showed only minimal study appeared, based solely on elemental data, isotopic effect (Godfrey et al. 2013). It is, how- to conform to two-component mixing between ever, correct to say that carbonates themselves relatively fresh (shallow) and saline (deep) are in most cases depleted in Li relative to sili- reservoirs. Lithium isotopes invalidated such a cates (e.g., Hoefs and Sywall 1997; Chan et al. simple interpretation, suggesting the retention of 2006) and thus, their contribution to the total isotopically distinct signatures by components of mass balance of Li at low temperatures appears the hydrologic system separated both by depth to be minor in the majority of geological and horizontal distance. This could be at least environments. partially accounted for by variable mineralogy of In an area where rigorous monitoring of aquifer sediments, but mapping this out would groundwater quality is in place, Négrel et al. require a more involved campaign of sampling of (2012) were able to promulgate a much more both waters and solids. The results of Meredith complete assessment of Li isotopic variations in a et al. (2013) were broadly confirmed by a study large aquifer system. The groundwaters of of deeply issued waters (>427 m) from the Great southwestern France showed the effects of mix- Artesian Basin (Pogge von Strandmann et al. ing between rainwaters of variable δ7Li and 2014). The δ7Li of these samples ranged between waters that carry an imprint of interaction with +9 and +16 ‰ with the highest Li contents in sediments and/or bedrock (all isotopically lighter low-δ7Li waters. The ratio of than rainwaters) at elevated temperature. weathering/dissolution of primary minerals and precipitation of secondary minerals controlled Li abundance whereas δ7Li was largely controlled 6.2.2 Deep Groundwaters by the extent to which secondary minerals (e.g., kaolinite, gibbsite, smectite, illite, chlorite, etc.) Many groundwater systems are completely tran- fractionate Li isotopes. An important considera- sitional with hydrothermal systems, which were tion in the study of Pogge von Strandmann et al. documented in Sect. 6.1. We include in this (2014) was that these large aquifer systems may section settings that range down-temperature be an important input of isotopically light Li to from “passive” geothermal waters (generally seawater. ≤100 °C). Formation waters, basinal brines and Ambient temperature groundwaters in the oilfield brines differ from ordinary groundwaters western U.S. Great Basin analyzed by Tomascak in that their histories of water-rock interaction are et al. (2003) were not uniform in Li isotopic more extensive. They show a range in Li isotopic 160 6 The Surficial Realm: Low Temperature Geochemistry of Lithium character, reflecting, in general, variability in seawater Li isotopic signature in formation starting water isotopic composition, lithology, waters that derive Li from rocks deposited during and the extent of water-mineral interaction. that era. This has a direct impact on the inter- Inasmuch as the Li isotopic composition of pretation of sources of groundwaters such as modern seawater is distinctively heavy, the sys- those of the Canadian Shield (Bottomley et al. tem has a significant role in assessing the con- 1999, 2003). In this study, data from deep waters tribution of marine sources to basinal waters were interpreted to possess significant Devonian (Fig. 6.2). In many cases it is clear that seawater seawater components. was the primary source of formation waters Using Li isotopes to constrain the tempera- (Chan et al. 2002b), although whether waters tures of deep basin fluids has received substantial reached concentrations 24× that of seawater attention. Considering the frequent lack of con- through evaporation or freezing could not be vergence among geothermal temperatures resolved (Bottomley et al. 1999). acquired through other geochemical techniques It should be noted that although seawater Li (e.g., Helgeson et al. 1978; Fournier and Potter has been isotopically consistent over the last c. 1979), if successful this would indeed be a major 6 Myr (Misra and Froelich 2012), shifts to advancement. compositions as light as δ7Li = +5 ‰ appear to Millot and Négrel (2007) took a novel have occurred during the last 100 Myr (Pogge approach with French geothermal waters of von Strandman et al. 2013). This tempers the well-understood temperature to suggest that Li utility of Li as a direct seawater fingerprint. isotopes could be used as a geothermometer. By These short-term variations may be compromised their results, for waters in reservoirs with T < c. by conservative behavior of Li in seawater, with 250 °C which were least affected by secondary the mean residence time of c. 1–1.5 Myr processes, the system appeared to yield viable (Stoffyn-Egli and Mackenzie 1984; Huh et al. results, with a precision of ±50 °C. Although this 1998), allowing for long-term processes to be precision is poorer than many of the conventional accounted for whilst temporally restricted events methods (e.g., silica solubility; Fournier and may remain unresolved. For example, Devonian Rowe 1966), Li isotopic thermometry may apply seawater Li remains isotopically uncharacterized; to a wide array of thermal waters where the thus we cannot unambiguously determine a applicability of some other methods is limited by

Fig. 6.2 Lithium concentration and δ7Li of deep groundwaters and basinal brines. Data sources: France from Millot et al. (2007, 2011), Millot and Négrel (2007), Italy (serpentine springs) from Boschetti et al. (2013), Israel from Chan et al. (2002b), Canada from Bottomley et al. (1999, 2003) 6.2 Groundwater and Pore Waters 161 major ion composition. The same approach had basal sediments in the Gulf of Cadiz reached been outlined previously for sea floor systems only slightly over 200 °C; Scholz et al. 2009). (e.g., Chan et al. 1994), but in these cases there Lithium isotopes have contributed to better appear more prominent pathways to obliterate understanding the geochemical evolution of meaningful thermometric potential of many these waters owing to the divergent isotopic thermal waters (e.g., Decitre et al. 2004). Millot fingerprints of the processes involved. Chan and et al. (2011) achieved very geologically reason- Kastner (2000), expanding on earlier work by able temperature estimates from Paris basin Zhang et al. (1998), proposed two isotopically waters from two localities, in the range 45–95 °C, distinct drivers for Li isotopes in pore water. Low with precisions of ±10–20 °C. Using this temperature alteration of volcanic material in sea approach to thermometry, however, does not floor sediments will generate clay minerals that always yield conclusive results, as demonstrated favor incorporation of light Li, driving pore by data from saline waters from the Appalachian water compositions to δ7Li > seawater (Fig. 6.3). Plateau and Gulf Coast of the U.S. (Macpherson Contrarily, reactions wherein clay minerals et al. 2015). Their samples did not fall on model exchange Li from crystallographic sites in favor + curves predicted from Li isotopic equilibrium of NH4 will liberate light Li and drive pore between water and average shales. water to isotopically lighter compositions (see To the extent they have been studied, springs also Vigier et al. 2008). The same conclusion issuing from terrestrial (that is, continental) ser- was reached for pore waters in the Escanaba pentinites have a peculiar geochemical signature, trough (James and Palmer 2000). This process is being highly alkaline with generally strongly especially important in sediments where bacterial elevated trace elements (B, Li) yet low dissolved processes yield high contents of dissolved solid contents. Apennine serpentinites do not ammonium. You et al. (1995) called on influx of show strong Li concentration variation, but do dilute, isotopically light fluid from deep within have a wide range in Li isotopes (Boschetti et al. the Nankai accretionary prism to explain varia- 2013). The Li isotopes strongly suggested that tion of pore fluids compositions in the vicinity of serpentinite springs represent variable degrees of the decollement. fractionation during crystallization of secondary Scholz et al. (2010) discovered, from com- minerals. Initial waters would thereby have had parison of compositions from a global pore water δ7Li < +20 ‰. data, a relationship between Li in pore waters and the nature of tectonics in the margin delivering sediment. Their analysis showed that passive 6.2.3 Pore Waters margin pore waters were enriched in Li relative to those at convergent margins. The alteration of Pore waters have been sampled predominantly in volcanogenic material, inherently poorer in Li the sediment piles of convergent margins, but than sediment derived from continental weath- also as subaerial and submarine seeps and mud ering, was shown to produce low temperature volcanoes. The compositions of these waters minerals that strongly retain Li without imparting reflect a continuum of processes. With seawater isotopic fractionation. Diagenetic alteration of as the principle starting fluid, these waters may carbonate-rich sediments was implicated by You evolve along distinct paths depending on the et al. (2003) to explain down-core variations in influences of mineral alteration and ion Li concentration and isotopic composition of exchange; their geochemistry may also be influ- pore waters from the eastern Pacific. enced by influx of deeper, hydrothermal waters. Mud volcano fluids from the Nankai accre- Even the deepest of these accumulations do not tionary prism (Nishio et al. 2015) had low δ7Li reach hydrothermal temperatures (e.g., even with (seawater-corrected values < +8 ‰), equivalent a thick accumulation and steep geotherm, the to those from hydrothermal samples from the 162 6 The Surficial Realm: Low Temperature Geochemistry of Lithium

Fig. 6.3 Lithium concentration and δ7Li of marine influence) from James et al. (1999), James and Palmer sediment pore waters. Data sources: Greenland (Irminger (2000), Gulf of Cadiz and Mediterranean Sea from Scholz Basin) from Zhang et al. (1998), Central America from et al. (2009, 2010), Gulf of Mexico, Black Sea from Chan and Kastner (2000), Scholz et al. (2010), Pacific Scholz et al. (2010) (Escanaba Trough with and without hydrothermal

Escanaba trough (James and Palmer 2000) and from hydrothermal alteration of sediments. some fluids in the Black Sea (Scholz et al. 2010). Although low-temperature fractionation may Nishio et al. (2015) suggested that their results drive processes within individual sediment piles, reflect the identifiable influx of geochemically when viewed collectively, the global pore water distinct fluids during times of seismic slip. One data set shows that an isotopically light, value reported from the subaerial mud volcanoes Li-enriched advected component is present in of Taiwan was slightly isotopically heavier than most systems (Fig. 6.3). The consensus of these these (c. +13 ‰; You et al. 2004). studies is that pore water processes have only Although a significant proportion of the ear- small influence on the marine budget of Li. lier pore water data show dilute compositions and heavy Li isotope signatures (i.e., δ7Li ≥ that of seawater), more recent data show that waters 6.3 Continental Weathering with Li concentrations greater than modern sea- water make up less than half of the ensemble. The quantitative understanding of continental There is a reasonable inverse correlation between weathering processes has become a major focus pore water δ7Li and concentration (R2 = 0.69; of geochemical studies, spurred by the potential Fig. 6.3). Scholz et al. (2010) applied a reactive for connecting weathering on the continents to transport model to quantitatively explain this, as marine geochemical records. Extraterrestrial did James et al. (1999), using a smaller data set. application of the system has recently been dis- The latter study concluded that pore waters from cussed (Fairén et al. 2015). The interest in these part of the Escanaba trough system derived their records is clear given the feedback processes high Li contents (some >100× seawater) through between chemical weathering reactions and upward advection of waters carrying Li liberated atmospheric greenhouse gas contents. Lithium 6.3 Continental Weathering 163 and its isotopic composition have begun to be fractionation during weathering in a simple nat- integrated into studies of this kind through both ural setting. Exposed transects from unweathered major world watersheds and smaller catchments to completely saprolitized rocks (granite and with greater controls on geological variables, as meta-diabase) permitted examination of Li well as intact soil sections. through the progressive alteration of the original rocks. In both saprolites there was a general correlation between the degree of weathering 6.3.1 Soils and Sediments (lower bulk rock density and higher chemical index of alteration) and the decrease in δ7Li, with A limited number of studies have considered Li overall δ7Li from −9to−16 ‰. Secondary isotopes in the realm of soils (Fig. 6.4). Pistiner mineral formation in the granite could plausibly and Henderson (2003) measured Li in weathered be described by a Rayleigh-type process, in volcanic materials from Iceland and Hawaii, but which Li is lost during progressive fractionation. did not find compelling isotopic fractionation. The diabase transect only partially followed this Since that study it has become apparent that the scheme; post-saprolitization overprinting from secondary mineralogy is a master variable in the interaction with groundwater appears to have quantification of isotopic fractionation during complicated the Li elemental and isotopic sys- weathering (Wimpenny et al. 2010a). Given the tematics of many of those samples. many factors that influence secondary mineral Considering the importance of weathering formation (source mineralogy, solution chem- basaltic rocks on global CO2 dynamics, exposed istry, etc.), variability in the products of chemical basalt terranes have been popular targets for Li weathering is to be expected. isotopic studies. Huh et al. (2004) performed a The first systematic study (Rudnick et al. targeted study of Hawaiian soils in an attempt to 2004) was an attempt to understand isotopic understand Li mass balance during weathering.

Fig. 6.4 Lithium isotopes relative to CIA (chemical Kısakürek et al. (2004), Hawaii from Huh et al. (2004), index of alteration) in soil samples. Data sources: Iceland South Carolina (SC) from Rudnick et al. (2004), Also from Pogge von Strandmann et al. (2012), Columbia included are loess paleosol samples (non-carbonate frac- River (CR) from Liu et al. (2013), Deccan traps from tion) from Tsai et al. (2014) 164 6 The Surficial Realm: Low Temperature Geochemistry of Lithium

They found that, like the saprolites studied by the land mass (continent versus localized island) Rudnick et al. (2004), under wet conditions (e.g., may impart a critical control on δ7Li of the mean annual precipitation >200 cm) weathered atmospheric deposition through differences in materials generally lose Li, and secondary min- climatic cycles. erals prefer the light isotope. Nonetheless, loss of In a similar study of a soil section developed Li is not always synonymous with shift in iso- from intensive weathering of flood basalt in topic composition. Ryu et al. (2014) also studied India, Kısakürek et al. (2004) saw somewhat tropical soils from Hawaii Islands and found that, different Li isotopic behavior. In the Deccan in young sequences (c. 3 ka), shallow soil sam- Traps they noted an overall decrease in Li con- ples that experienced as much as 30 % loss of Li centration with progress of weathering, coupled showed no isotopic deviation from pristine with an influx of isotopically heavy Li in samples source basalt. representing the paleo-water table. Lithium Isotopically heavier soil samples from Huh redistribution interpreted in the meta-diabase et al. (2004), particularly those from more arid samples of Rudnick et al. (2004) yielded higher sampling sites, were thought to reflect input of Li concentrations and δ7Li values. Kısakürek seawater-derived aerosol Li, with its character- et al. (2004) used mass balance principles to istically high δ7Li (c. +31 ‰). The same con- implicate an isotopically light atmospheric input clusion was reached by Pogge von Strandmann (δ7Li * +1 ‰) as a major Li source in these et al. (2012) in their analysis of a basaltic soil soils. In addition to bulk samples, Kısakürek section in Iceland. Soils least affected by aero- et al. (2004) leached samples with dilute acetic sols, comprising mainly glass and mixed acid; in general the leachates were Li depleted Al-silicate clays and oxyhydroxides, maintained and slightly isotopically heavier than the resi- δ7Li up to 5 ‰ lighter than pristine tephra. These dues. The implication of isotopically distinct Li results, however, are somewhat opposite to those imported from aerosol or dust sources in these of Millot et al. (2010a; Sect. 6.1) who implied a studies (Huh et al. 2004;Kısakürek et al. 2004; minor fraction of seawater-borne aerosols in rain Pogge von Strandmann et al. 2012) was not waters in France. This suggests that the size of shared by Ryu et al. (2014), whose mass balance

Fig. 6.5 Lithium concentration and δ7Li of marine sediments. Data sources: Aleutians, Tonga, Central America, South Sandwich, Guaymas and Greenland all from Chan et al. (2006), Cascadia (includes Escanaba Trough) from Chan et al. (2006), James et al. (1999), Also included are loess samples from Teng et al. (2004) and Aeolian dust from Liu et al. (2013) 6.3 Continental Weathering 165 model allowed <20 % total contribution from affect interpretations derived from studies of river atmospheric deposition, with a majority of the six sediment. Hawaiian sites yielding <10 %. The study of Dellinger et al. (2014) brought to Bedload and suspended loads in rivers help to light several interesting details in the considera- understand aspects of the soil data. Of the four tion of Li isotopes in river sediments, and, hence, published studies that include pairs of suspended in the study of continental weathering. They were and bedload sediment from the same rivers able to use isotope–element relations to define (Kısakűrek et al. 2005; Pogge von Strandmann contributions of different major rock types in the et al. 2006; Wimpenny et al. 2010b; Millot et al. catchments of tributaries, showing the variability 2010c), there is a reasonable positive correlation of Li isotopic yield from different fundamental between Li isotopes in the two (R2 = 0.81; n = 61 rock types (primarily igneous versus sedimen- pairs; Fig. 6.6). This relationship indicates that tary, with some contribution from metasedi- the average bedload is *2 ‰ heavier than the mentary rocks). Additionally, their data coexisting suspended sediment. Dellinger et al. suggested that the extent of recent chemical (2014) used a combination of new and existing weathering was not a primary driving force data to demonstrate that the sediments carried by modifying the δ7Li of suspended loads. This is three major world rivers showed strong correla- particularly so for particles derived from sedi- tion of δ7Li and Al/Si, a proxy for grain size. The mentary lithologies, where, logically, minerals sediment with lowest Al/Si corresponded to the had been through the sedimentary process before, highest δ7Li in bedload sands (up to +5.3 ‰) and achieving a degree of stability, and hence limit- the corresponding suspended material had high ing their capacity for Li isotopic fractionation. Al/Si and δ7Li as low as −3.6 ‰. This study The majority of river sediment studies were emphasized the potential sampling bias that can conducted in areas where bedrock geology was

Fig. 6.6 Lithium isotopes in river suspended loads (2010a, b); Iceland from Pogge von Strandmann et al. relative to bedloads. Dashed line is the regression of all (2006), Canada from Millot et al. (2010c), Sao Miguel data (R2 = 0.81), which essentially has a slope of 1. Data from Pogge von Strandmann et al. (2010), South America sources: Tibet/India from Kısakȕrek et al. (2005), from Dellinger et al. (2014) Dellinger et al. (2014), Greenland from Wimpenny et al. 166 6 The Surficial Realm: Low Temperature Geochemistry of Lithium largely invariant, and isotopic fractionation decreased chemical weathering and hence effects could be explicitly dissected. The pro- decreased removal of atmospheric CO2. duction of secondary minerals in both suspended Chan et al. (2006) compiled the most particles and bedload mineral grains involves authoritative data set on Li in marine sediments, fractionation of variable extents. Wimpenny et al. with 69 samples from seven globally-dispersed (2010b) measured offsets of 2 to 6 ‰ between sites. The samples represent the primary marine source basalts and river sediments, with the latter lithologies, with an emphasis on clays and muds uniformly isotopically lighter than gneissic bed- and comparatively few purely carbonates (four rock of western Greenland. Vigier et al. (2009) samples). The overall data set has a mean found negligible source rock/sediment fraction- δ7Li = +4.4 ± 3.6 ‰; without the strictly car- ation in Iceland. In the tropical (i.e., bonate samples the mean is +4.0 ± 2.6 ‰ non-glaciated) island of Sao Miguel, Pogge von (Fig. 6.5). Variations in the Li isotopic signatures Strandmann et al. (2010) observed river sedi- of samples from individual collection sites from ments to be up to 9 ‰ heavier than the volcanic this study are apparent. For example, nine sam- source bedrock. Whereas the specific secondary ples from off the Tonga arc record a mean mineralogy plays a role in the fractionation that δ7Li = +7.1 ± 2.4 ‰, whereas the same number takes place during weathering, the difference of samples collected from the Pacific seafloor off observed between these studies illustrates the of the Cascadia arc had a mean significance of the intensity of chemical weath- δ7Li = +2.6 ± 1.1 ‰. All of the sample suites ering on Li isotopic behavior. This will be show a consistent, if slight, decline in δ7Li with examined further in sections to follow. increasing Li concentration. Similar Li elemental In the headwaters of the Ganges River in Tibet, and isotopic systematics were observed by Li isotopes of sediments corresponded to the Bouman et al. (2004) for marine sediments from general variability in bedrock (Kısakűrek et al. East Sunda and the Lesser Antilles. Terrigenous 2005). Suspended loads and bedload in tributaries sediments with distinctly high Th/Sc and La/Sm draining dolomitic substrates had higher δ7Li than ratios also carry high loads of Li with isotopi- those draining silicate-dominant rocks. However, cally light signature (Bouman et al. 2004), con- the absence of data for rocks inhibits deeper sistent with derivation from mature continent by interpretation of the sediment data. weathering. A distinct approach to the use of terrigenous Complimentary to the data from marine and sediments to evaluate paleoclimate was that of fluvial sediments, Teng et al. (2004) analyzed Dosseto et al. (2015), in which material from bulk samples of loess from various localities. The terraces of rivers draining a modern mountain mean δ7Li for loess (+0.3 ‰), analytically range (the Himalayas) was examined. Given the indistinguishable from aeolian dust from Wash- problem of time lag between weathering and ington (Liu et al. 2013), is isotopically lighter deposition, Dosseto et al. (2015) used deposits in than both the suspended load (+1.3 ‰) and the river headwaters. The variation of Li isotopes bedload (+2.9 ‰) averages. Teng et al. argued in bulk sediments (δ7Li = −1.5 to +2.0 ‰) was that loess ought to be a reliable estimator for the consistent with but more restricted than in bulk Li isotopic composition of the upper crust. clay-size separates of the same samples Rather than integrating global sources, Tsai et al. (δ7Li = −4.3 to +2.7 ‰). The covariation of (2014) suggested that loess samples from China sediment δ7Li with Tibetan ice core δ18O over track regional sources in proximal Chinese the last c. 40 ka was interpreted to support the deserts. In their study, Tsai et al. (2014) analyzed predominance of runoff and physical erosion both detrital (δ7Li from −3.5 to +5.6 ‰) and (and hence tectonics) as drivers of chemical carbonate components (δ7Li from −4.1 to weathering, more so than changing temperature. +14.2 ‰), unlike the other studies that used only Thus accelerated tectonic uplift will promote bulk samples. 6.3 Continental Weathering 167

In the most unique use of Li and Li isotopes to only to the processes inherent in diagenesis and date in relation to weathering proxies, Ushikubo lithification, but in all of the subsequent fluid– et al. (2008) used detrital zircon grains in ancient mineral interactions that may occur, both in deep quartzites to investigate the early history of the burial and after exhumation. Earth’s continental crust. Although the main The data of Qiu et al. (2009) were interpreted argument from Ushikubo et al. (2008) for an to reflect a rather simple relationship between Li evolved continental crust on Earth at c. 4.3 Ga isotopes and the magnitude of weathering of the was the high Li contents in the zircon, the Li source lithologies. Importantly, their results isotopes from these grains were far more variable suggested, at least for the case of the Paleozoic of than would be expected if derived from primitive southern England, that low grade metamorphism sources. The large number of grains and spots does not affect Li isotopic systematics of sedi- analyzed turned up a significant range in δ7Li, ments (see Sect. 5.2.2 for more on metamor- from −20 to +15 ‰. This study contained phism). The results from Teng et al. (2004), in comparative data from younger zircons from their examination of globally distributed exam- igneous and high grade metamorphic rocks, tes- ples of sedimentary rocks, were slightly isotopi- tifying to the resiliency of Li to diffusion in zir- cally heavier (average δ7Li = −0.2 ‰ compared con and opening the pathway to future studies to −1.0 ‰ from Qui et al. 2009). More variable using Li in zircon as a monitor of magmatic and still are the isotopic compositions recorded in metamorphic processes in crustal terranes. sedimentary rocks from the Massif Central (Millot et al. 2007) and northern Canada (Millot et al. 2010c), with a total range of δ7Li from −7.0 6.3.2 Sedimentary Rocks to +9.8 ‰. These rocks had an average δ7Li of +1.8 ‰, although 1/3 of the rocks in these two Although sediments themselves have, in various studies were carbonates. forms, been extensively analyzed for Li isotopes, Given the variability in degree of weathering the product of lithification of these materials has of source rocks, coupled with the likelihood of garnered comparatively little study (Fig. 6.7). variable post-lithification processes, the observed Sedimentary rocks carry with them added layers range in Li isotopic composition is surely to be of potential stable isotope complexity owing not expected. Similar isotopic diversity superimposed

Fig. 6.7 Lithium concentration and δ7Li of (largely) unmetamorphosed sedimentary rocks. Data sources: Canadian Shield from Millot et al. (2010c), UK from Qiu et al. (2009), Massif Central from Millot et al. (2007), Schwarzburg from Romer et al. (2014b), “other” are data from globally dispersed shales and Chinese sedimentary rock composites from Teng et al. (2004) 168 6 The Surficial Realm: Low Temperature Geochemistry of Lithium on overall homogeneity was clear from the data of of the unweathered rocks (estimated global Romer et al. (2014b) from Paleozoic shales from flow-weighted average δ7Li = +23.5 ‰; Huh central Europe (δ7Li ranged from −6.1 to et al. 1998; Fig. 6.8). It is clear that for most +14.6 ‰, with an average of −0.1 ‰). Although major rivers the bulk of the transported Li (typ- the literature on sedimentary basins focuses sig- ically >80 %) is in suspended phases rather than nificantly on temperature, Romer et al. (2014b) the dissolved load (e.g., Dellinger et al. 2015). made the important observation that an equally if The factors that influence how this fractionation not more important parameter in the generation of develops are varied, and understanding how river Li isotopic signatures of clastic sedimentary rocks fluxes affect marine Li hinges, in part, on more is the timing of dehydration in the sequence of clear resolution of these factors. diagenetic events. The loss of Li from In examining rivers of the world, Huh et al. exchangeable sites as opposed to structural posi- (1998) found rather poor correlation between tions in the crystal lattice is a major factor in dissolved riverine Li and the major catchment determining the ultimate δ7Li of the sedimentary lithology. It was not until the more detailed study rock, and hence the evolution of Li in subaerial of Huh et al. (2001) on the Orinoco River system versus subaqueous environments will be that it became apparent that weathering out- significant. weighs lithology when assessing the isotopic Using laboratory experiments, Williams and composition of dissolved Li. This is good news Hervig (2005) evaluated exchange and for global paleoclimate research, in that it sug- dissolution/precipitation of smectite/illite relative gests that the sensitivity of Li isotopes to to Li and B isotopes at conditions relevant to the weathering intensity can be a useful proxy: rivers lithification of sedimentary materials (300 °C, that drain stable, intensely weathered terranes 100 MPa). The transition to illite was accom- (transport-limited) will contribute lighter Li than plished by smectite dissolution and illite those that drain actively eroding terranes (reac- neocrystallization, and was accompanied by Li tion- or weathering-limited). This assessment isotopic fractionation. They observed effects was further strengthened by results from Kısa- introduced by size distributions, with very fine kűrek et al. (2005) on a detailed study of the grain size resulting in slower approach to Ganges River system (Fig. 6.8a). apparent equilibrium. Their results suggested a Nevertheless, with more detailed study has fractionation factor at 300 °C that is *6 ‰ lar- come evidence for greater complexity. The study ger (i.e., illite is 6 ‰ lighter) than predicted from of Pogge von Strandmann et al. (2010) drew no the empirical data (Chan et al. 1992, 1994), but single connection between dissolved Li in rivers consistent with theoretical estimates (Yamaji draining young basaltic terranes and Li isotopes et al. 2001). (Fig. 6.8b). They indicated that the balancing role of pH, as a determinative factor in the formation of secondary minerals, dominates in some settings. In 6.3.3 Lakes and Rivers: Dissolved something of an all’s-well-that-ends-well conclu- Loads sion, Pogge von Strandmann et al. (2010) added that due to the low Li content of glacial rivers that The first publication containing Li isotopic data tend not to follow the Huh et al. (2001) paradigm, for river waters (Falkner et al. 1997) gave a hint these rivers are not apt to contribute significantly to at the potential for Li as a tool for understanding the global marine Li pool, and thus the potential weathering processes. Subsequently, many stud- paleoclimate proxy probably remains viable. The ies have geochemically taken apart catchments restricted area of catchments such as those in Ice- on various scales toward this ultimate goal. One land also limit their global impact compared to, for unavoidable finding from any of these studies is instance, the Ganges, Amazon and Mackenzie. that the Li released during weathering is iso- An extensive study of Li in the Amazon topically heavier than that present in the minerals system by Dellinger et al. (2015) yielded several 6.3 Continental Weathering 169

Fig. 6.8 Lithium (a) concentration and δ7Li of dissolved loads from rivers. a Rivers in continental basins. Data sources: Himalayas (including high Himalayan and Ganges-Brahmaputra) from Huh et al. (1998), Kısakȕrek et al. (2005), China from Huh et al. (1998), Wang et al. (2015), South America from Huh et al. (1998, 2001), Dellinger et al. (2015), North America, Russia from Huh et al. (1998), Millot et al. (2010c), Greenland from Wimpenny et al. (Wimpenny et al. 2010a, b), Antarctica from Witherow et al. (2010), France from Lemarchand et al. (2010), Négrel et al. (2010), Jordan from Chan et al. (1992). b Rivers in volcanic (mainly basaltic) (b) terranes (without hydrothermal impact). Data sources: Iceland from Pogge von Strandmann et al. (2006), Vigier et al. (2009), Sao Miguel from Pogge von Strandmann et al. (2010), Java, Reunion form Henchiri et al. (2014), Martinique from Rad et al. (2013)

important conclusions concerning the control of altitude floodplain processes. The relationship Li in the dissolved load ultimately delivered to they instead identified indicated that the lowest the oceans. They found no single relation δ7Li waters were from both tributaries that record between the δ7Li of dissolved Li and factors such low and high weathering intensities, and hence as Li flux or weathering rate, as the behavior of δ7Li could not be used independently to charac- high relief tributaries were moderated by lower terize weathering regime. Thus low δ7Li in 170 6 The Surficial Realm: Low Temperature Geochemistry of Lithium paleocean sediment records interpreted to reflect equally plausible interpretations of the geo- periods of enhanced continental erosion becomes chemistry of the waters (through ternary mixing a non-unique solution. or through a Rayleigh-type fractionation pro- Lithium isotopic variability in dissolved loads cess). These complications do not, however, of other river systems (Vigier et al. 2009; Millot appear to derail the potential usefulness of Li as a et al. 2010c; Liu et al. 2011; Wimpenny et al. weathering proxy. 2010b; Witherow et al. 2010) have largely Pogge von Strandmann and Henderson (2015) upheld the lack of correlation with catchment used combined Li isotopic and U activity ratios of lithology first described by Huh et al. (1998). river waters from New Zealand to demonstrate a Fractionation has been primarily linked to parti- potential of this geochemical tandem in address- cles: the nature and extent of precipitation of ing key questions in continental erosion. Rivers in secondary minerals and/or water/particle inter- areas with rapid uplift have low δ7Li whereas the action history. However, when examined closely, opposite is observed for regions with slow uplift the Li isotopic systematics of river systems rates. Pogge von Strandmann and Henderson commonly show intricacy that transcends (2015) implied that more fresh material with single-step interpretations. For example, Pogge primary mineralogy is exposed at high uplift rates von Strandmann et al. (2006) noted that the to silicate weathering, limiting the formation of aggregate Li isotopic composition of a river (at secondary mineral phases. Thus, tectonically its mouth) could not be reconstructed by a simple stabilized regions would allow for saturated wa- source dilution model, and that fractionation ters and a greater opportunity for the precipitation processes going on in the channel were essential of isotopically light minerals. Lithium isotopes components of the Li isotopic evolution of the may thus provide a novel proxy for estimating the water. Also, Yoon (2010) and Wang et al. (2015) efficiency of continental weathering in CO2 showed that some mineralogical control appears removal. Note, however, that the findings of to be present in catchments with significant Pogge von Strandmann and Henderson (2015) are exposed evaporite (halite, gypsum). contrary to what Huh et al. (2001) determined for The average sedimentary rock δ7Li of −0.1 ‰ rivers in the Orinoco drainage. is lower than the mean values for marine clastic In a unique approach to the assessment of sediments (+4.1 ‰), river bedloads (+2.9 ‰) and paleoclimate, Godfrey et al. (2013) used data river suspended loads (+1.3 ‰). As pointed out from bulk fluid inclusions in lacustrine halite previously, the Li isotopic compositions of sed- from an intensely arid salar in northwestern imentary rocks incorporate both the variability in Argentina. Based on laboratory experiments source material, with its particular weathering Godfrey et al. determined that Li isotopic frac- history, and the many processes that affect the tionation during halite crystallization was mini- sediment following deposition. As most of these mal. The data from two Pleistocene halite processes tend to favor extraction of an isotopi- samples were up to *10 ‰ heavier than modern cally heavy component, it is no surprise that waters in the salar. This difference was inter- these rocks have a lower mean δ7Li. preted to be the record of wetter conditions in the Studies that couple Li and other stable iso- region 60–90 ka. A different approach to under- topic systems have shown the capacity to clarify standing the sources of Li in evaporites from the some of the questions generated among the western U.S. Great Basin (Araoka et al. 2014) purely Li isotopic studies. For example, Tipper involved the comparison of Li from pure water et al. (2012) measured Mg isotopes in their and nitric acid leaches of playa sediments. The re-examination of waters from the MacKenzie small data set from this study did not provide a River system whose Li isotope systematics were conclusive view of Li in this environment, originally investigated by Millot et al. (2010c). although the authors suggested that generally The Mg data produced the opposite trend to what light Li isotopic compositions (δ7Li of leaches was expected from theory, leading to separate, between −1.0 and +9.3 ‰) were more consistent 6.3 Continental Weathering 171 with hydrothermal sources rather than with Li yet, they are isotopically essentially unperturbed derived from chemical weathering. (Brant et al. 2012). Among the natural systems Tomascak et al. (2003) had suggested the use that have been investigated, though, this result is of sedimentary carbonate in a similar capacity for rare. paleoclimatology. The equilibrium fractionation One of the original foci of Li isotopic research associated with carbonate crystallization has only was on the nature of Li in seafloor basalts (Chan been broadly studied at this point (Marriott et al. and Edmond 1988; Chan et al. 1992). Despite 2004a). The one aragonitic carbonate analyzed this extensive history of study, it is an area that by Godfrey et al. (2013) had δ7Li which cor- has proved slippery in arriving at a consensus rected to a value similar to that predicted for the (not only for Li!). Until recently, the prevailing Pleistocene lake, but also consistent with those of model invoked uptake of seawater Li into modern regional spring waters. neocrystallized hydrothermal minerals to explain a correlation between Li concentration and iso- topic composition (Chan et al. 1992, 2002a; 6.4 Hydrothermal Systems Fig. 5.3). The same trend is seen when sample age is compared to the Li concentration and Geothermal systems have been targeted in several isotopic ratio (Chan et al. 1993). This model Li isotopic studies, on the continents and in the implies equilibrium Li isotopic fractionation oceans. In hydrothermal systems the temperature during sea floor recrystallization with α * 0.981 of exchange is significant toward the specific at 2 °C, and α * 0.984 at 80 °C. Low temper- Li concentration and isotopic outcome, as is the ature (*80 °C) alteration of sea floor basalt has ratio of solid to fluid. Details have been estab- been shown to result in Li uptake by solids lished by many empirical and laboratory studies. (Seyfried et al. 1984; Chan et al. 1994). The majority of these have considered the most It is uniformly seen that the globally relevant case of interaction between recrystallization-induced fractionation of basalt seawater and basalt, although some studies have protoliths at low greenschist facies produces introduced other geological complications to try rocks that are isotopically lighter than the unal- to better understand Li in the shallow Earth. tered basalt. Rocks such as these, with Li con- centrations and δ7Li < MORB, are uncommon among the current catalog of seafloor drilled and 6.4.1 Rocks in Sea Floor dredged samples (Chan et al. 1992, 1994, Hydrothermal Systems 2002a), probably because subsequent, lower temperature reactions overprint the higher tem- If the discussion is limited to high temperature perature mineralogical and chemical composi- alteration (>300 °C), resultant altered sea floor tion. Nevertheless, the specific fractionation that rocks are typically leached of Li (Seyfried et al. takes place is highly dependent on local condi- 1984; Berger et al. 1988). Lithium does not show tions. For example, neocrystallized smectites at a strong tendency for incorporation in secondary T > 200 °C were shown to not fractionate Li minerals at these temperatures. In the absence of isotopes to a measurable degree, regardless of the crystallization of secondary minerals, the fluid composition (Vigier et al. 2008). altered rocks will not be isotopically fractionated, Isotopically heavy Li, derived from seawater, similar to what is observed during laboratory is incorporated in low temperature minerals, such experiments of congruent dissolution at low that altered sea floor basalts have δ7Li as high as temperatures (Pistiner and Henderson 2003). +21 ‰ (Chan et al. 2002a). Smectite separates This observation appears to be borne out in MOR from altered basalts are similarly isotopically basalts at Hess Deep (East Pacific Rise) which heavy (δ7Li = +15 to +19 ‰; Chan et al. 2002a). show minimal petrographic evidence for recrys- These were bulk mineral samples, so the content tallization, but which have been leached of Li; of structurally-incorporated Li relative to a 172 6 The Surficial Realm: Low Temperature Geochemistry of Lithium sorbed fraction is unknown. Pistiner and Hen- show bulk compositions that overlap this range derson (2003) gave a glimpse at the significance (Benton et al. 2004). of sorption relative to crystallographic incorpo- The isotopically light values are consistent ration in fractionation of Li isotopes. The anal- with bulk rock serpentinite data from Vils et al. ysis of purified individual alteration minerals has (2009), in samples drilled from the mid-Atlantic lagged behind the analysis of bulk rock samples, ridge. They found samples with δ7Li as low as suggesting that there may be more to learn here, −28.5 ‰. The data suggested, contrary to especially as the precision of spatially-resolved expectations, that Li in the serpentine was analytical methods improves. derived from oceanic basalts rather than primar- The model of Brant et al. (2012) was needed ily from seawater. Furthermore, the crystallog- to explain data from Hess Deep basalts that did raphy of the serpentine may exert a control on the not agree with the trend established by previous fluid/serpentinite Li isotopic behavior of the investigators (Fig. 5.3). It is interesting to note system. In a study of laboratory synthesis prod- that several of the ODP Hole 504B and 896A ucts, Wunder et al. (2010) suggested that rocks fall similarly off the trend. Additionally, the serpentine polymorph chrysotile, with its sheeted dike basalts at Hess Deep had δ7Li tubular growth, does not fractionate Li isotopes generally higher than unaltered MORB, as as do its siblings antigorite and lizardite. Exper- opposed to previous studies where deep MOR imental fluid-mineral reactions with the latter two dike rocks were mostly isotopically lighter than polymorphs show strong preference for 6Li in MORB (e.g., Chan et al. 2002a). solids, with 7Li concentrated in coexisting If their model is accurate, submarine lava aqueous fluids in the range 200–500 °C and flows had Li extracted by moderate-temperature crustal to shallow mantle pressures. The lack of fluids (50–100 °C), but the lack of recrystalliza- strong fluid/serpentine isotopic fractionation tion resulted in minimal Li isotopic fractionation. when chrysotile was involved implicated This is consistent with the results of Li-water complexes in nanotubes. Verney-Carron et al. (2015). By extrapolation to Contrarily, in laboratory experiments in which other MOR systems, Li is generally lost to dis- basalt glass and olivine crystals were weathered charging fluids in the hot, axial portion of the at nominally Earth surface conditions and with ridge and the gain of Li in MORB is almost controlled fluid composition, Wimpenny et al. entirely an off-axis phenomenon. This contrasts (2010a) generated secondary chrysotile which with the interpretation of Chan et al. (2002a), fractionated Li isotopes. The sense and magni- which viewed the axial region as a net sink for tude of the fractionation was consistent with marine Li, rather than a source. Isotopic frac- empirical results from natural serpentinites, with tionation of Li seen in sheeted dike samples at fluids reaching compositions *10 ‰ heavier Hess Deep was interpreted to reflect combina- than coexisting chrysotile. Collectively, the tions of crystallization of low temperature alter- results on serpentinites have implications for the ation minerals as well as, importantly, diffusion marine Li cycle (Sect. 6.4) as well as for Li in from plagioclase. subduction systems (Chap. 5). Hydrothermal alteration of the upper mantle Marine ferromanganese encrustations may be to produce serpentinites has been studied with Li used to assess hydrothermal processes, although isotopes. Decitre et al. (2002) observed widely Li isotopic systematics in these materials appear variable Li isotopic compositions in serpentinites complicated. You and Chan (1996) published the from dredges of the southwest Indian ridge (δ7Li first data from ferromanganese crusts in a paper from +2.9 to +14.2 ‰ in bulk samples). In situ devoted to analytical technique development. analyses of serpentinized minerals in the same A more comprehensive analysis of marine fer- samples yielded δ7Li as low as −19 ‰. Serpen- romanganese crusts (Chan and Hein 2007) found tine muds and clasts from the Mariana forearc that, although primary Mn-oxides appear to 6.4 Hydrothermal Systems 173 initially record an isotopic record of the seawater basalt (±sediments) as significant sources of Li in from which the minerals form, this signature is these fluids. The overall interpretations of Brant almost certain to be overprinted during subse- et al. (2012) are consistent with the isotopic quent re-equilibration. composition of hydrothermal vent fluids as well Although not a robust tracer of paleoseawater as the products of laboratory experiments. The Li, the data of Chan and Hein (2007) indicate that reactive transport model of Verney-Carron et al. theconcentration ofleachableLiinFe–Mncrustsis (2015) appears to reproduce the relative isotopic a strong indicator of process, with purely consistency of MOR hydrothermal fluids from hydrothermally-generated precipitates yielding systems with diverse Li contents. nearly two orders of magnitude more Li than Significantly in the interpretation of data from comparable hydrogenetic/diagenetic and weakly such systems, the separation of vapor and brine hydrothermal encrustations. Leaching of these has been examined. Several studies have indi- samples by hydrochloric acid yields isotopically cated that vapor separation does not introduce Li lighter Li (δ7Li as low as +8.9 ‰) than that isotopic fractionation (James et al. 1999; Fous- removed by acetic acid (which has seawater-like toukos et al. 2004; Millot et al. 2010b), although δ7Li * +31 ‰). Presumably, Li sorbed to it can introduce elemental variations that are Mn-oxides records seawater conditions, whereas being used in monitoring the evolution of vol- isotopic fractionation takes place when incorpo- canic systems (e.g., Berlo et al. 2004). Schiavi ration into structuralsites inMn- and Fe-oxidesand et al. (2010), however, interpreted data from oxyhydroxides occurs, consistent with laboratory Stromboli samples to require melt–fluid (vapor) studies (e.g., Pistiner and Henderson 2003). isotopic fractionation. Similarly, Vlastélic et al. (2011) concluded that Li isotopic variation in Piton de la Fournaise eruptive samples stems 6.4.2 Fluids in Sea Floor from vapor separation. Nevertheless, it is appar- Hydrothermal Systems ent that in eruptive settings such as MOR, Li is not appreciably lost in a vapor phase, and hence After relatively small reaction times, hydrother- not isotopically fractionated (as compiled by mal alteration releases fluids that are isotopically Brant et al. 2012). analogous to the rocks from which they derive Vent fluids in the Lau back-arc basin spread- much of their dissolved Li (Seyfried et al. 1998; ing show similar δ7Li to MOR vent fluids (+2.9 James et al. 2003). However, the light isotope to +7.0 ‰; Mottl et al. 2011), although with a may be appreciably enriched in minerals gener- lower mean δ7Li value (Fig. 6.9a). Temperatures ated at these conditions (Chan et al. 1994; Scholz of the Lau Basin fluids are generally in the range et al. 2009). At lower temperature hydrothermal 200–350 °C (after correction for bottom water conditions, release and uptake of Li from/into mixing). It is possible that earlier data for MOR basalts may operate concurrently. vent fluids (e.g., Chan et al. 1994) are on average Pristine marine vent fluids (those for which 1–2 ‰ too high as a consequence of the lack of seawater incorporation is corrected) typically minor bottom water correction. Mottl et al. show restricted Li isotopic variation, δ7Li in the (2011) interpreted Li isotopic variations along range +3 to +11 ‰ (Chan et al. 1993, 1994; strike in the Lau vent fluids to reflect increased Foustoukos et al. 2004) and have Li concentra- input of sediment-derived fluid, largely on the tions >10× that of seawater (c. 180 ppb; Morozov basis of geochemical variations of other ele- 1968). Lithium in these fluids thus cannot have ments. This appears plausible, given the slightly been derived entirely from either seawater higher average δ7Li of sediment from the Tonga (δ7Li = + 31 ‰) or from unaltered MORB trench compared to marine sediments elsewhere (δ7Li = +4 ‰), implicating seafloor alteration of (Bouman et al. 2004; Chan et al. 2006). 174 6 The Surficial Realm: Low Temperature Geochemistry of Lithium

Fig. 6.9 Lithium (a) concentration and δ7Li of hydrothermal fluids. Model binary mixtures are between seawater (SW) and a sea floor vent source and between average river water and a source like average continental crust (CONT). a Data sources: MOR (Mid-ocean ridge) vents from Chan et al. (1993, 1994), Foustoukos et al. (2004), BAB (back-arc basin) vents from Millot et al. (2011), Lake Baikal from Falkner et al. (1997), Yellowstone from Sturchio and Chan (2003), Japan (Ontake volcano) from Nishio et al. (2010), Lesser Antilles from Millot et al. (2010b), Rad et al. (2013), “other” (includes data from multiple (b) continental and ocean island localities) from Godfrey et al. (2013), Henchiri et al. (2014), Kısakűrek et al. (2005), Tomascak et al. (2003). b Hydrothermal fluids from springs and well at Taupo, New Zealand. Data sources: Bernal et al. (2014), Millot et al. (2012)

6.4.3 Continental Geothermal through which they flow may vary far from the Systems comparatively limited starting compositions present on the ocean floor. Hence continental Continental geothermal systems have received hydrothermal waters show a range in chemical comparatively less attention than their marine and isotopic compositions. analogs, but generally mimic the findings from Individually, though, continental and conti- the sea floor. On the continents, however, both nental margin geothermal systems show a ten- the composition of initial waters and the rocks dency toward equilibration with respect to Li 6.4 Hydrothermal Systems 175 isotopes (Fig. 6.9a). That is, Li in thermal waters Among 23 New Zealand geothermal well from individual localities or regions shows waters of variable temperature (205–320 °C) restricted isotopic heterogeneity. reported by Millot et al. (2012), there was virtu- Adefinitive study to use Li in exploring high ally no Li isotopic variation (δ7Li = 0.2 ± 0.5 ‰; temperature terrestrial water-rock interaction is Fig. 6.9b). Bernal et al. (2014) also reported data that of Sturchio and Chan (2003). They analyzed from the Taupo Volcanic Zone, both thermal 11 Yellowstone geothermal springs, finding springs and geothermal wells, whose δ7Li were average δ7Li of +3.7 ± 2 ‰. Altered rhyolites ostensibly identical to those of Millot et al. through which these waters flowed had generally (2012), with average δ7Li = − 0.6 ± 1.0 ‰ slightly lower δ7Li (average −0.4 ‰), which (Fig. 6.9b). The minimal isotopic variation was correlated strongly with Li concentration consistent with the source waters having equili- (R2 = 0.88). They concluded that the Li isotopic brated with regional volcanic and/or metasedi- budget could be explained primarily through mentary rocks (global average δ7Li = −0.1 ‰) progressive addition of Li to altered rocks with with no significant seawater components. concomitant isotopic fractionation, producing A diverse group of waters in an island arc setting isotopically heavier waters at higher water/rock was investigated by Millot et al. (2010b). These ratios. One point not clear from this study was included hydrothermal waters from continental the Li isotopic signature of regional meteoric wells as well as shallow sea floor vent fluids, dis- water; their calculations used a starting δ7Li of tributed between two islands of the Antilles arc. +1 ‰ for waters, which is not constrained by Watersrangedfromhigh(upto260°C)tomoderate measurements from the area, and perhaps not (as low as 90 °C) estimated reservoir temperatures. consistent with other studies. Despite substantial homogeneity of the Li isotopic Bullen and Kharaka (1992) presented pre- compositions of waters from some of the sites liminary data from several of the same Yellow- (overall δ7Li average = +6.5 ‰), the concentration stone localities, with all values some *12 ‰ data for the Bouillante springs fell along a mixing lower than those of Sturchio and Chan (2003). curve with seawater. Given the relatively consistent offset between the This study was particularly noteworthy in that two data sets, combined with the fact that the the authors used data from laboratory seawater– Bullen and Kharaka (1992) data are so different basalt interaction experiments to develop a pre- from any published since, it seems likely that the dictive model for the Li source rock isotopic Sturchio and Chan (2003) data are correct and composition. This required accurate estimation of that an analytical or other laboratory issue the reservoir temperature and isotopic composi- resulted in the apparent very isotopically light tion. The results indicated source rocks with light water values from Bullen and Kharaka. Li isotopic signatures (δ7Li from about −2to Two hydrothermal springs from an exclu- −7 ‰). Few rocks have been analyzed that sively volcanic setting in the Azores (Pogge von extend to such low values, among them sedi- Strandmann et al. 2010) were isotopically mentary rocks from Millot et al. (2007); their inconsistent, but the heavier value was from a marls averaged δ7Li = −6.3 ‰. It is also possible cool spring (33 °C) that likely reflects greater that the fractionation model, based on data from down-temperature fractionation (δ7Li = + 6.9 ‰, basalts, is not suitable for systems with sedi- compared to +1.0 ‰ for T > 100 °C). Thermal mentary host rocks. springs from multiple sites hosted by volcanic Interplay between hydrothermal activity on rocks in the western U.S. (Tomascak et al. 2003) active volcanic islands and the composition of had δ7Li ranging from +3 to +17.1 ‰. The associated river waters has been documented heavier compositions measured in this study (Rad et al. 2013; Henchiri et al. 2014). The study appear to reflect samples containing variable of three distinct volcanic islands indicated contamination of isotopically heavier lake water covariation between δ7Li and Li/Na, with iso- (δ7Li * +20 ‰). topically heavy river water generally having the 176 6 The Surficial Realm: Low Temperature Geochemistry of Lithium lowest elemental ratio. The river waters analyzed to implicate diffusion in controlling the isotopic by Henchiri et al. (2014) that were most affected composition of some experimental fluids. by hydrothermal sources had δ7Li c. 10 ‰ lower When marine and continental hydrothermal than waters that had negligible hydrothermal fluids are compared (Fig. 6.9), similar systemat- influence. Although the Rayleigh fractionation ics, as would be expected, are clear. The major model used by Henchiri et al. (2014) adequately difference is likely due to isotopically lighter host explained Li isotopic–elemental relations in most rocks in the continental crust. Mixing arrays data sets, it failed in the case of Iceland (Pogge given dilution of rock-derived Li by marine or von Strandmann et al. 2006; Vigier et al. 2009), meteoric Li yield plausible fits to the observed which shows broadly overlapping Li isotopic data (Fig. 6.9a). compositions for rivers with both strong and weak hydrothermal impacts. Nishio et al. (2010) have attempted to use Li 6.5 Marine Geochemistry of Li in fluids from volcanic areas in Japan as a tool for understanding the connection between conver- The first goal of Li isotopic study in the “modern gent margin fluid flow and seismicity. In a young analytical era” was to use the system to better volcano on the mainland Nishio et al. identified understand marine geochemistry (Chan and an isotopically light fluid component that was Edmond 1988). At that time this goal was present in samples collected from the primary motivated somewhat by improvements in earthquake epicenters. They explained the light understanding of the dynamics of Li the element isotopic composition of these samples (δ7Li as in the oceans (e.g., Edmond et al. 1979; Delaney low as −5.2 ‰)asreflecting Li derived from the and Boyle 1986; Stoffyn-Egli and Mackenzie lower crust. Alternately, applying the reasoning 1984). It is instructive to look back at what Chan of Verney-Carron et al. (2011), these low δ7Li and Edmond called a preliminary report of Li in values might reflect diffusive exchange at mod- the oceans and see, on the strength of a small erate temperature. However, further studies by number of what are, by modern standards, rather Nishio et al. (2015), applied to fluids from off- imprecise data, how lasting their interpretations shore mud volcanoes in the Nankai accretionary were. Although we have advanced far in terms of complex, appear to substantiate the connection application of proxies in the oceans, their 1988 between fluid flow regimes (deep versus shallow) study mapped out in a very compelling fashion and seismic activity. the framework of much that would follow. Laboratory experiments on basalt leaching at near-surface conditions (temperature, pH) from Verney-Carron et al. (2011) may be important in 6.5.1 The Marine Li Cycle future interpretation of hydrothermal fluids in some systems. Unlike experiments in which The Li dissolved in modern seawater is isotopi- secondary minerals crystallize, saturation was not cally heavier than the primary inputs (rivers and allowed to take place in the Verney-Carron et al. hydrothermal systems), which requires that sink experiments, and isotopic fractionation between (s) exist that preferentially remove light Li on starting basalt glass and fluid enriched the latter short time scales. The seawater Li budget is in light and not heavy Li. The results were con- controlled largely by a small number of major sistent with the greater diffusivity of 6Li, and factors. The combination of hydrothermal sys- although isotopically unimportant at low tem- tems and processes that act at lower temperatures peratures, the authors suggested that in natural generally exert major controls on this budget. hydrothermal systems the diffusive fractionation Based on modeling of ocean crust samples needs to be considered. Indeed they from ODP Hole 504B, Chan et al. (2002a) sug- re-interpreted data from Wimpenny et al. (2010a) gested that the upper 1.8 km of the basaltic crust 6.5 Marine Geochemistry of Li 177 is, in bulk, a sink for marine Li, albeit a minor been investigated by Murphy et al. (2014) for one. This is consistent with the experimental Kalix and Råne rivers in northern Sweden. The results of James et al. (2003). In contrast, marine water of Bothnian Bay is isotopically akin to sediments have been shown to be a viable Li seawater but has much lower salinity (*3 vs. source, liberating Li even at sub-hydrothermal *35 psu for seawater) and has been proposed as temperatures (James et al. 2003). isotopically heavy end member to river waters Calculations appear to restrict the role of with δ7Li as low as +21.7 ‰ for low-salinity water–sediment interaction on the sea floor (dif- samples. A non-conservative behavior of Li in fusion and adsorption, collectively) to only a estuarine systems, involving chromatographic minor sink of marine Li (Zhang et al. 1998). It is effects from suspended river particles may have interesting to note the correlation between the large consequences for global Li input to oceans. seawater Li isotopic record of Misra and Froelich Pogge von Strandmann et al. (2008) suggested (2012) with the pore water data of Zhang et al. that such behavior over time would result in a (1998). Despite the presumably significant con- secular effect of the magnitude that has been trols on pore water chemistry by water–sediment measured in Cenozoic foraminifera (e.g., interaction processes, the data from ODP Site Hathorne and James 2006). 918 reproduce the notable c. 15 Ma δ7Li low. The thorough treatment of river sediments by Wimpenny et al. (2010b) placed two major 6.5.2 Secular Variation of Li constraints on marine Li. Although they noted in Seawater that the majority of Li export in western Green- land is in suspended and not dissolved form, the 6.5.2.1 Experiments and Observations extant data indicate that only a fraction of this Li If Li isotopes can be useful proxies of, for will reach the open ocean. Furthermore, their example, continental weathering intensity, it is data indicate that glacial sources are minor critical to find ways of unlocking temporal globally (<6 % of the flux from rivers during the records. Early measurement of Li in three marine last glacial maximum). ferromanganese crusts gave seawater-like values Pogge von Strandmann et al. (2008) verified, (δ7Li * +33.2 ‰; You and Chan 1996) for two in looking at estuaries, that although dissolved Li of the samples, giving some encouragement of Li behaves conservatively in these environments as a paleocean chemistry tool. However, rapid Li (Stoffyn-Egli 1982), suspended particles may be diffusion was subsequently determined to essen- subject to continued chemical weathering in the tially nullify this use (Henderson and Burton water column. They estimated that a significant 1999). This finding was upheld by a later proportion of riverine Li may fail to reach the investigation of globally-dispersed ferroman- oceans because of estuarine processes. This ganese deposits (Chan and Hein 2007). amount (15–25 % of the global flux) would Marriott et al. (2004b) demonstrated that cal- measurably affect the river/hydrothermal budget cite and aragonite fractionate Li differently, both of marine Li. The rapid decrease in Li abundance elementally and isotopically. Inorganically crys- in coastal seawater from riverine input has been tallized aragonite showed a δ7Li offset of −12 to further demonstrated by Brunskill et al. (2003). −11 ‰ from coexisting seawater-like fluid, From sampling in the Gulf of Papua over 6 years, whereas calcite grown the same way was offset they showed that (i) Li abundance increases with only by −3to−2 ‰. The Ca site in aragonite is increasing salinity in dissolved loads but not in ninefold coordinated, as opposed to a sixfold site particulate matter, and (ii) as much as *90 % of in calcite (Reeder 1983), although Okumura and riverine Li input is deposited in shallow shelf Kitano (1986) suggested that the difference region and on shelf slope whereas only *10 % between these sites is unlikely to drive this iso- of Li is transported further to deep sea. Behavior topic difference. They instead suggested that the of Li in low-salinity estuarine environment has interstitial incorporation of Li in calcite yields 178 6 The Surficial Realm: Low Temperature Geochemistry of Lithium less isotopic preference in calcite than in arago- 20 mg of foraminifera for a single sample and nite under equilibrium conditions. Košler et al. (2001) used an average of 6.8 mg In some cases, the applicability of these per sample. The methods of You and Chan experiments to nature seems clear. For example, (1996), Hall et al. (2005), and Hathorne and Sturchio and Chan (2003) analyzed travertine James (2006) all used approximately 2.5 mg of with δ7Li value 11 ‰ lighter than coexisting foraminifera per sample. The most recent bulk spring water and Tomascak et al. (2003) found foraminifera study using bulk methods was that that spring-precipitated “carbonate” (presumably of Misra and Froelich (2012), in which 1–2mg calcite) had δ7Li approximately 2 ‰ lighter than of foraminifera was used per sample. contemporaneous lake water. Foraminifera The data of Košler et al. (2001) showed no reported by Marriott et al. (2004b) carried with dependency on Li isotopes from variable size of them the expected * −3 ‰ offset relative to foraminifera. Hathorne and James (2006) pro- modern seawater; similarly, foraminifera reported vided further support for that relationship, as well by Košler et al. (2001) ranged in offset from −5to as a hint that individual foraminiferal species do −1 ‰. However, not all studies of foraminiferal not differently fractionate Li isotopes. They also δ7Li report an offset from modern seawater in demonstrated that fractionation is not influenced core-top samples. Hathorne and James (2006), by seawater temperature, consistent with labora- Hall et al. (2005) and Misra and Froelich (2012) tory experiments (Marriott et al. 2004b). all report modern foraminifera with broadly sea- Hoefs and Sywall (1997) were the first to water Li. For calcite, the offset is minor relative to present a large marine carbonate data set, the scale of marine Li isotopic variation (see including a total of 75 analyses of foraminifera, below). Although kinetic effects may be respon- bivalves, carbonate oozes, and limestones. The sible for the lack of consistency in the modern data set, however, faces several problems that foraminiferal data set (i.e., crystallization rate: may relate to the analytical technique, the large Marriott et al. 2004a; Gabitov et al. 2011), the sample size, or a combination of these. The effect has yet to be specifically addressed. cardinal trouble is the lack of modern seawater Vigier and Goddéris (2015) suggest that vital values for any of the modern samples that were effects inforaminifera may play arole in generating analyzed: their modern foraminifera, for exam- some of the short time-span isotopic variations in ple, ranged in δ7Li from +7 to +20 ‰, very far the Cenozoic record of Misra and Froelich (2012). below the modern value. The down-core varia- However, this has yet to be rigorously tested. tions they saw in foraminifera and oozes were Broadly speaking, Li isotopes in the carbonate inconsistent and difficult to reconcile with the skeletons of marine organisms do not appear to be known behavior of Li in the oceans. It is possible dominated by the vital effects which are manifest in that diagenesis could have played a role with stable isotopic systems of biologically mediated some samples. Ullmann et al. (2013) provide elements, such as C and O (i.e., Duplessy et al. evidence for effects of diagenesis on fossil 1970; McConnaughey 1989), making Li a poten- belemnites (calcite), and suggest that in the case tially sensitive tracer in foraminifera. That is not to of their sample set diffusion is not an important say that Li plays no role biologically (see Sect. 6.6), cause of isotopic alteration. Instead the primary but, at least in higher animals, this role is complex means of corrupting original isotopic signatures (Schrauzer 2002) and requires specific study. is from fractionation during recrystallization. Until recently, analytical limitations made the Although some of the foraminiferal data of study of small, monospecific foraminiferal sam- Košler et al. (2001) from seafloor (core top) ples impossible. Advances in analytical tech- samples were equivalent to modern seawater, niques have opened up new possibilities in Li about 30 % were significantly isotopically offset isotopic studies of foraminifera. Hoefs and (up to 20 ‰ heavier). None of the four bulk Sywall (1997) used 20–100 mg of carbonate for foraminiferal samples from You and Chan a single analysis. Marriott et al. (2004b) needed (1996), all from the Pleistocene, had modern 6.5 Marine Geochemistry of Li 179 seawater Li isotopes; δ7Li ranged from as low as an increase in marine δ7Li from a minimum +19 ‰ (c. 12 ‰ lighter than modern seawater) around +26 ‰ at 16 Ma to the modern value by to as high as +42 ‰ (c. 11 ‰ heavier than around 6 Ma. Misra and Froelich (2012) expan- modern seawater). As the Li concentration of ded the marine record back beyond the K–Pg the final, cleaned carbonate samples were within boundary (c. 66 Ma; Kuiper et al. 2008) with the ordinary range for marine carbonates foraminifera taken from seven different sites and (*1 ppm), it is clear that the results of these 10 specific locations. This large data set pro- studies cannot be ascribed to silicate or ferro- duced an amazingly concise record which manganese contamination. However, neither of showed both correlations with and distinctions these studies focused on the interpretation of from the marine records of O, Sr, and Os iso- these samples, so the data are difficult to more topes. From their data it appears that during the deeply interpret. bulk of the Cenozoic the marine δ7Li was lower All of the foraminifera analyzed by Hall et al. than today, reaching a nadir in the Paleocene. (2005), ranging from 0.4 to 44 ka, turned up Significantly, there is excellent agreement modern seawater values within uncertainty, between the data sets of Hathorne and James indicating a lack of glacial/interglacial Li iso- (2006) and Misra and Froelich (2012). Aside topic offset. The lack of significant Li isotopic from the upturn in δ7Li in the mid-Miocene, the variation during the Pleistocene is consistent with major change in marine Li captured in the data of the results from foraminifera samples of both Misra and Froelich (2012) is at the K–Pg Hathorne and James (2006) and Misra and boundary, where late Cretaceous Li appears close Froelich (2012). The consensus is thus that, in to modern (δ7Li * +31 ‰), and within *1 Myr the absence of complicating effects unrelated to it drops to * +23 ‰. They also contend that the isotopic composition of the water from which their record documents “pulsed” change in sea- the minerals grew, marine carbonates faithfully water δ7Li through the Cenozoic, although to record seawater δ7Li (Fig. 6.10). unambiguously understand the absolute nature of This conclusion has been used in three studies the secular variation will require a more contin- to date to examine secular variation in marine Li uous data set. isotopes. Hathorne and James (2006) analyzed Instead of foraminifera, Pogge von Strand- foraminifera from sediments in Atlantic and mann et al. (2013) examined bulk shallow marine Pacific cores dating back 18 Ma. The data show carbonate, targeting specific time–stratigraphic

Fig. 6.10 Cenozoic marine δ7Li records from foraminifera. Data from Misra and Froelich (2012) and Hathorne and James (2006) 180 6 The Surficial Realm: Low Temperature Geochemistry of Lithium horizons represented in three locations in Europe. isotopes in sediments from different same-age The c. 1.2 Myr interval examined was around the sites (that have certainly experienced distinct Cenomanian–Turonian boundary (c. 93.5 Ma) post-depositional histories) provides confidence and recorded δ7Li variation from c. +7 to c. that the role of diagenesis is minimized +30 ‰. In all of these sections the onset of (Hathorne and James 2006; Pogge von Strand- oceanic anoxia (“Ocean Anoxic Event 2”) mann et al. 2013). appears to be marked by a rapid and enormous Even in the methods that permit analysis of decrease in marine δ7Li (by 10–20 ‰). Coupled relatively small bulk samples, there are concerns with other data, Pogge von Strandmann et al. about micro-scale complications that may go (2013) interpreted this drop in δ7Li to reflect the unseen by these techniques. As a result, in situ weathering of subaerially exposed basaltic ter- analytical methods present many benefits. ranes that were generated globally just prior to the Although the generally low Li concentration in onset of anoxia. Presuming these sediments marine carbonate presents challenges to high retained their original Li isotopic signatures, the precision measurement of these materials, as we large Li isotopic variations over a geologically have seen, the scale of natural variability may be short period of time requires that the marine res- so great as to tolerate even rather low precision. idence time has changed over time from its cur- The first study to use such techniques on rent value of c. 1 Myr (Stoffyn-Egli and foraminifera was that by Vigier et al. (2007). Mackenzie 1984), owing to reorganization of the Their study showed the importance of the rigor- hydrologic cycle. ous cleaning of foraminiferal tests to eliminate Laboratory studies of Li incorporation in materials included during shell growth with inorganic carbonate precipitates have been con- spurious Li isotopic characteristics. The authors ducted. Some caution was indicated by the suggested that such contamination was not pre- results of Gabitov et al. (2011). Although their sent in the coral samples they evaluated. To that laboratory-grown marine aragonite showed bulk end, Rollion-Bard et al. (2009) measured Li solid–fluid offsets similar to those measured by isotopes in various corals (natural and Marriott et al. (−7.7 to −10.5 ‰), they found laboratory-cultured) to better understand the substantial variation in δ7Li within individual intricacies of Li in the skeletal components of parts of the texturally-heterogeneous aragonite these animals. From their results combined with matrix. However, these all overlapped within 2σ earlier coral analyses (Marriott et al. 2004b)a uncertainty. few conclusions could be made. First, it appears It is unclear why some studies of modern that, unlike B isotopes, there is no significant marine carbonates report values inconsistent with internal difference in Li isotopes relative to the contemporaneous seawater. The effects of dia- morphology of the skeletal element (dense zones genesis and diffusion of Li in a sediment pore versus fibrous). Second, among the small number network are concerns in dealing with sub-sea of species examined, there was an apparent offset floor materials. Misra and Froelich (2012) pro- in δ7Li between those inhabiting deep water and vided a convincing argument that their samples those living in shallow water, in this case were not influenced in such ways. Same-age approximately 3 ‰. Marriott et al. (2004a) noted samples taken from cores from different regions ac.2‰ difference in the two species of shallow of the oceans and those representing a range in water coral analyzed. water depth returned identical data within Vigier et al. (2015) applied the in situ ana- uncertainty, effectively ruling out diagenetic lytical approach to a series of cultured for- control of Li signatures. Also, analysis of coex- aminiferal tests (genus Amphistegina). Their isting pore waters showed no co-variation of δ7Li experiments looked to deconvolve effects of values with the foraminifera, as would be the variable seawater temperature, pH, and dissolved case if Li diffusion through the sediment column inorganic carbon (DIC) on both Li isotopes and was significant. In general, the consistency in Li Li/Ca. The major finding that samples cultured in 6.5 Marine Geochemistry of Li 181 high DIC water carry elevated δ7Li permits rely more on the change in the Li flux from determination of secular variation of DIC. When continental weathering, ultimately dictated by the applied to late Quaternary bulk core-top samples variable rate of soil formation (i.e., their model from Hall et al. (2005) the initial set of experi- matched the Misra and Froelich marine Li curve mental results, conducted over five different DIC with a decrease of river δ7Li through the Ceno- concentrations, when coupled with ice core pCO2 zoic). Their calculations presumed the modern data, produced pH variations that were within dominance of Li in suspended particles over uncertainty of those calculated over the same dissolved load in rivers to have persisted through time period from other geochemical proxies. time. The Vigier and Goddéris (2015) model is 6.5.2.2 Interpretation and Modeling unlike that of Wanner et al. (2014) in that it took Substantial sets of secular marine Li data (e.g., a broader view of weathering as a continent-scale Misra and Froelich 2012; Hathorne and James process. Vigier and Goddéris (2015) substanti- 2006) have provided the basis for a new surge in ated the interpretation that globally thick soil interpretive studies, as the community grapples accumulations around 60 Ma argue against low with the meaning of seawater Li isotopic change seawater δ7Li at that time, as suggested by Misra over time. Lithium isotopes have the major and Froelich (2012). The Vigier and Goddéris advantage over the radiogenic Sr isotope system (2015) model was the first to explicitly link in that silicates dominate the budget for Li, marine Li isotopic variations to carbon cycle whereas both silicates and carbonate contribute change, which appears to afford it an advantage importantly to the Sr that is liberated during in terms of constraining variables in the series of continental weathering. Since continental calculations involved. weathering is tied to atmospheric CO2, under- None of the modeling studies independently standing marine Li should, in principle, shed light on the question of causation: models strengthen the understanding of the carbon cycle. attempt to calculate fluxes that satisfy measured The extent to which this can be applied quanti- isotopic proxy results, but do not specify the tatively has been the subject of significant driver (atmospheric temperature, tectonics, ongoing scrutiny, as models are constructed with hydrology, etc.). Clearly this is an area of rapid varying intrinsic assumptions. evolution which will benefit significantly from The model of Li and West (2014) explained future empirical studies which add to the set of marine Li variation like Misra and Froelich high quality and high resolution Li isotopic data (2012) by balancing changes in the deep sea from the oceans over time. Although the utility of sinks and their inherent fractionation signatures Li isotopic data in examining paleoclimate with changes in river water Li isotopes. They questions appears strong, the exact shape of its suggested the mean global δ7Li of rivers at c. application continues to be sketched out. 60 Ma may have been as low as +13 ‰, con- sistent with the modeling of Wanner et al. (2014). Misra and Froelich (2012) and Froelich 6.6 Biological and Anthropogenic and Misra (2014) suggested that this value could Li Isotopic Fractionation have been as low as +3 ‰ heading into the Paleocene–Eocene thermal maximum, arguing 6.6.1 Lithium in Plants and Animals for a predominance of congruent weathering. Vigier and Goddéris (2015) argued that the Kabata-Pendias (2011) lists the ranges in Li secular variation in marine Li isotopes does not abundances for many plant families worldwide in require the δ7Li of rivers to change. Although order to link them with specific lithological secondary minerals fractionate Li isotopes during substrate and to assess more useful information weathering reactions, Vigier and Goddéris pre- on individual families. These ranges are, in many ferred the Li isotopic composition of seawater to cases, quite large (from tens of ppb to tens ppm 182 6 The Surficial Realm: Low Temperature Geochemistry of Lithium dry weight). It is clear, however, that some plants properties of the wood, source lithology, climate may accumulate Li when grown in arid areas. conditions, or other parameters. Although there is some information available on It has long been known that Li is a Li abundances in specific materials (e.g., leaves, non-essential element for the human body. In needles, etc.; Wu et al. 1997; Sucharová and fact, it may be quite toxic when administered in Suchara 2006), there are few studies of the tox- larger doses (Aral and Vecchio-Sadus 2008). icity of Li to plants when supplied in overabun- Alexander et al. (1982) supplied artificially 6Li- dance (e.g., Naranjo et al. 2003; Kalinowska and 7Li-enriched lithium to mice and noted et al. 2013). substantially deteriorated behavior as well as Lithium isotopic studies of vegetation mate- increased lethality for 6Li-enriched dosing. These rials are, expectedly, even more limited in num- results were also stressed by tests applying acute ber and the extent. It may come to pass that Li overdosing with 6Li- or 7Li-enriched substances isotopes could enable a better understanding of that, again, resulted in dramatically different the transfer of water from the soil environment to biological response for the two respective iso- plant tissues, as Li is mobile in aqueous solutions topes. Conversely, unchanged motor properties and its isotopes may respond differentially to and recovery to full health were superior for changing chemical–physical conditions, but this 7Li-doped mice. These effects may perhaps result has yet to be realized empirically. Lemarchand from different ionic radii and masses of the et al. (2010) found rather limited variations in respective isotopes during transfer through the both Li contents (7.3–154 μmol/kg) and isotopic living body, and perhaps higher efficacy of 6Li compositions (< –4.3 ‰) for spruce needles, uptake by erythrocytes via membrane transport branches, and roots. The Li abundance in plant (Hughes and Birch 1992). Stoll et al. (2001) also tissues was about three orders of magnitude found that for longer treatment periods, more lower than soils the spruce grows on. The indi- serious malfunctions of the kidney system follow vidual δ7Li values did not vary greatly between dosing with 6Li-rich substances, and that 7Li is above ground (needles, branches) and below more benign in this respect. However, the ground (roots) biomass, and were in the range of enhanced toxicity of 6Li compared to 7Li is soil solutions. Thus, biological uptake from soil accompanied by somewhat better therapeutic solutions may have little effect on the isotopic effects. Clearly, carefully controlled medical balance of Li. Nevertheless, Schmitt et al. (2012) experiments could provide important information pointed out the potential utility of combined on the mobility and isotopic fractionation of Li in studies of Li isotopes with other the human body. Such experiments may serve to biologically-sensitive stable isotopic systems improve our knowledge of the effects of dosing (Mg, Ca) in studies of the dynamics of dissolved Li to patients suffering from manic-depressive components in soils. illness (also known as bipolar disorder) as well as The findings of Lemarchand et al. (2010) were the consequences of long-term treatment and its confirmed by the pilot study of Myška et al. withdrawal. (2011), who analyzed Li contents and isotopic The indications that Li is not a biologically compositions in spruce (roots, stem wood, young mediated element do not necessarily spell an end needles, old needles) and beech (roots, stem to the chapter of Li and biology. Vigier et al. wood), as well as in the underlying soils (2015) suggest, based on data from foraminifera, (>5 ppm; δ7Li < −2 ‰). Sub-ppm Li abundances that biological calcification is accompanied by a (0.07–0.44 ppm) were found for all tree com- fractionation distinct from purely inorganic ponents but large δ7Li variations were discovered crystallization. This fractionation, which favors (−5 to +13 ‰), in particular for spruce. How- 7Li incorporation in shell calcite, may relate to ever, further work is required to constrain the key the kinetics of the biochemical reaction. In a variables that produce this range: taxonomy, somewhat less direct approach, Pogge von 6.6 Biological and Anthropogenic Li Isotopic Fractionation 183

Strandmann et al. (2010) made the astute obser- water analyzed by Wang et al. (2015) was not vation that dissolved Li in river water relates (in extraordinarily isotopically heavy (δ7Li = principle, at least) to weathering intensity. Since +40.8 ‰), but permitted the determination that biomass has a potentially significant role in some parts of the Changjiang River system had mineral breakdown in soils, biological activity significant (c. 15 %) anthropogenic contribu- may indeed help establish the link to Li isotopes tions. The results of Kısakűrek et al. (2005) also in studies of weathering. hinted at an anthropogenic component in some Tibetan river waters that bore anomalously, if not definitively, heavy isotope signatures, but this 6.6.2 Tracer Applications in Water hypothesis was not carried further in their study. Treatment Another tactic for the use of Li isotopes in and Agriculture environmental studies draws its basis in the work of Fritz and Whitworth (1994), who suggested The use of Li isotopes as tracers in both surface that membrane filtration would yield significant and groundwater studies has only begun to show isotopic effects. These effects would certainly promise, which seems counterintuitive, given the apply to anthropogenic filtration, but perhaps significant potential for the development of iso- also to fluid movement through impermeable topically distinct, traceable reservoirs in these sediments in nature. Kloppmann et al. (2008a) settings. The study of Qi et al. (1997) showed found that during the desalination process of massive (c. 400 ‰) isotopic shifts present in seawater and brackish water, (i) Li abundance certain anthropogenically-manufactured materi- decreased by up to four orders of magnitude from als. That materials can become highly isotopi- *180 ppb down to 0.01 ppb, and (ii) measurable cally enriched has been demonstrated by the Li isotopic effects were imparted due to lower generation of percent-level fractionation by ion retardation of 6Li in the reverse osmosis mem- exchange chromatography (Taylor and Urey branes, with δ7Li values lower by 12 ‰ in stages 1938; Oi et al. 1991). Indeed, one of the banes of of desalination associated with decreased pH. It Li isotopic analysis is the very fact that *100 % was also stressed that the extent and direction of yield must be assured in the ion exchange pro- Li isotopic fractionation may hinge on membrane cess in order for isotopic measurements to be type and operating conditions (Kloppmann et al. meaningful (see Chap. 2). Tiny fractions of such 2008a). isotopically anomalous Li would be easily The impact on water quality of wastewater traceable signals of specific pollutants surface injection has been evaluated by Kloppmann et al. and ground waters. (2008b), who tested blending of wastewater Some studies have begun to use Li as an recharge into natural aquifers under controlled environmental tracer, though, and to useful conditions. Whilst there was no significant effect effect. The best example to date comes from a on both Li concentration and Li isotopic com- study of waters from a peat bog from southern position in the injection–recharge zone, apart France, where Négrel et al. (2010) found from minor fractionation associated with clay anomalous δ7Li in groundwater samples that mineral precipitation, significant mixing takes were best explained by mixing between normal place with brackish aquifer water loads from a aquifer Li and isotopically heavy-enriched Li wider area of the test site. Further experimental deriving from fertilizer. The analyses of soil work by Kloppmann et al. (2009) revealed a amendments in this study showed substances steady-state decrease in Li contents, likely due to with δ7Li as high as +215 ‰, whereas the pre- sorption paralleled by decrease in δ7Li. Thus, a dicted end member composition generating the sink for heavy Li must exist within the aquifer. isotopically anomalous water had δ7Li of at least Kloppmann et al. (2009) suggest that Li becomes +1250 ‰ (i.e., with absolute 7Li/6Li ratio more a non-conservative monitor in these dynamic than twice that of the L-SVEC standard!). Waste regimes. Water management becomes an 184 6 The Surficial Realm: Low Temperature Geochemistry of Lithium increasingly important topic with time consider- radially away from ground zero (from 48 to ing decreased availability of drinking water in >74 m), whereas distance and 152Eu activities were many populated areas, and hence Li isotopic effectively anti-correlated (Bellucci et al. 2013). It applications in this field may play a more sig- may well be that the time for implementing the Li nificant role in the future. isotopic anomalies was too short, or that neutron consumption by the reactions was too large to induce measurable variations, or that the samples 6.6.3 Nuclear Forensics did not contain a large enough fraction of materials from the detonating device. Further analyses of One goal of nuclear forensics is to identify trinitite glass, considered as remnants of the blast nuclear material through isotopic and elemental tower and the bomb itself, are warranted. properties in parallel with chemical and physical characteristics. In the past, Li was used exten- sively in certain nuclear facilities owing to the fact that 6Li has a very high cross section for References thermal neutrons (*940 barn; the barn is gener- ally considered as the probability of a neutron Alexander GJ, Lieberman KW, Okamoto M, Stokes PE, collision with the target nucleus, and is defined as Triana E (1982) Lithium toxicology: effect of isotopic −28 2 composition on lethality and behavior. Pharmacol 10 m , roughly corresponding to cross-section Biochem Behavior 16:801–804 7 area of a uranium nucleus). Contrary to this, Li Aral H, Vecchio-Sadus A (2008) Toxicity of lithium to has very small thermal neutron cross section humans and the environment–a literature review. – (*0.045 barn) and hence is of little use in the Ecotoxicol Environ Safety 70:349 356 Araoka D, Kawahata H, Takagi T, Watanabe Y, nuclear industry. These bifurcated properties Nishimura K, Nishio Y (2014) Lithium and strontium combined with the predominance of 7Li in nature isotopic systematics in playas in Nevada, USA: resulted in vast amounts of 7Li-enriched residues constraints on the origin of lithium. Miner Deposita – after the isolation of the sought-after 6Li. Such 49:371 379 Bellucci JJ, Simonetti A, Wallace C, Koeman EC, volumes of non-natural Li materials were then Burns PC (2013) Isotopic fingerprinting of the World’s commercially distributed for laboratory and first nuclear device using post-detonation materials. industrial use. No records of which we are aware Anal Chem 85:4195–4198 exist to estimate how much of these materials was Benton LD, Ryan JG, Savov IP (2004) Lithium abun- dance and isotope systematics of forearc serpentinites, distributed or to where these volumes ultimately Conical Seamount, Mariana forearc: insights into the went. That such materials were distributed, mechanics of slab-mantle exchange during subduc- however, is not in doubt (e.g., Qi et al. 1997). tion. Geochem Geophys Geosys 5, paper num- The light isotope 6Li was also used as a fusion ber Q08J12. doi:10.1029/2004GC000708 Berger G, Schott J, Guy C (1988) Behavior of lithium, fuel in the form of lithium deuteride in order to rubidium, and cesium during basalt glass and olivine boost thermonuclear reaction yields because 6Li dissolution and chlorite, smectite and zeolite precipitation can produce tritium. Although 7Li has somewhat from seawater: experimental investigations and modeling ° – different physical properties, it may breed 6Li between 50 and 300 C. Chem Geol 71:297 312 Berlo K, Blundy J, Turner S, Cashman K, through an (n,2n) reaction which is quite effec- Hawkesworth C, Black S (2004) Geochemical pre- tive. Such secondary 6Li then produces further cursors to volcanic activity at Mount St. Helens, USA. tritium and thus, the yields of tritium, fusion and Science 306:1167–1169 fission efficiency increase significantly. Bernal NF, Gleeson SA, Dean AS, Liu XM, Hoskin P (2014) The source of halogens in geothermal fluids To our knowledge, no tests of the effect of from the Taupo Volcanic Zone, North Island, New nuclear blasts on the Li isotopic compositions have Zealand. Geochim Cosmochim Acta 126:265–283 been made so far. Preliminary data from Alam- Boschetti T, Etiope G, Pennisi M, Millot R, Toscani L (2013) Boron, lithium and methane isotope composi- ogordo (Magna and Simonetti, unpublished data 7 6 tion of hyperalkaline waters (Northern Apennines, 2013) show no distinctive Li/ Li ratios between Italy): terrestrial serpentinization or mixing with near-ground zero samples and those collected brine? Appl Geochem 32:17–25 References 185

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Geochim Chem Geol 284:26–34 Cosmochim Acta 62:2437–2450 Summation: What Have We Learned and Where Can We Go? 7

7.1 Retrospective long-lived (recycled components) or may be shallow level in origin (magma chamber con- tamination). There is still a level of uncertainty in In the 2004 review of the state of Li isotope the interpretation of Li isotope variability in research by Tomascak, a series of conclusions MORB samples. However, the global average were put forth. Some of these were actually appears robust and, ironically, recapitulates the suggestions of where Li isotopes had either not estimates from the very first MORB study (Chan been effectively applied or where the existing and Edmond 1988). data failed to completely assess a topic. In light The potential influence of diffusion on the of the studies that have followed, many areas of isotopic compositions of high temperature progress are apparent, and the 2004 list can be materials has become apparent in the intervening re-evaluated. Below we recount the concluding years (see Chap. 4). The results of laboratory points from Tomascak (2004) and indicate where experiments and computational studies serve to progress has been made (i.e., where understand- temper (and in some cases invalidate) many of ing now prevails over uncertainty) as well as the original interpretations from igneous and where new questions have emerged and where metamorphic rocks. Future natural studies need some of the larger gaps remain in the field. to take the constraints imposed by diffusional processes into consideration to avoid erroneous • “There is no consensus on the specific mantle conclusions. On the other hand, substantial Li isotopic composition. Currently reported improvement of in situ analytical capacity has values for normal MORB suggest a range opened the door to use diffusion-induced frac- (δ7Li = +1.5 to +5.6), but no single study has tionation in productive ways, for example in shed light on whether this range is an intrinsic quantifying rates of geological processes. feature of the upper mantle, to what extent Although the specter of diffusion may hang cryptic contamination plays a role, or if bulk heavily over many studies of the mantle and samples can only define an average for a crust, there are still fundamental lessons to be reservoir with broad heterogeneity.” learned in translating laboratory observations into natural settings. Between a richer set of data for MORB and more thorough study of direct mantle samples, • “Given the fractionation of Li isotopes by there is now clear consensus on the Li isotopic processes near the Earth’s surface, fraction- composition of the upper mantle. Sam- ated materials should be returned to the pling MORB at different spatial scales has yiel- mantle throughout geologic time. There is ded suggestions of heterogeneity that may be

P.B. Tomascak et al., Advances in Lithium Isotope Geochemistry, 191 Advances in Isotope Geochemistry, DOI: 10.1007/978-3-319-01430-2_7 © Springer International Publishing Switzerland 2016 192 7 Summation: What Have We Learned and Where Can We Go?

apparent correlation between mantle hetero- radiogenic isotopes. Whether carefully selected geneity in radiogenic isotopes and Li. Thus samples can defeat such problems remains to be far the incorporation of material with demonstrated. δ7Li < MORB, as predicted based on studies of subduction zone lavas and eclogites, is not • “The extent to which isotopic fractionation evident in oceanic lavas. An initial attempt to occurs at sub-1000 °C temperatures in mag- track this recycling process over Earth history matic and regional metamorphic systems is with peridotite xenoliths suggested an unstudied in detail, but limited natural and increase in δ7Li by *4 ‰ since the experimental data in granitic systems suggest mid-Archean.” that measurable fractionation may occur under these conditions. This fractionation As indicated in Chaps. 4 and 5, the impact of might be applied in studies of magmatic and diffusion on understanding the isotopic evolution ore evolution, either for thermometry or of metamorphic rocks has advanced in funda- mineral deposit exploration.” mental ways since the 2004 summary of the field. Earlier inferences of extreme Li isotopic deple- Laboratory experiments in the appropriate tion during prograde metamorphism have been temperature range for the rigorous study of significantly curtailed. Even so, traceable Li crustal metamorphism are challenging, but isotopic signatures are clearly developed in the Chap. 4 maps out the current state of under- crust which should be applicable to the study of standing. Existing studies have not yielded a mantle heterogeneity. clear vision of the utility of Li isotopes in For example, the “HIMU” signature of many studying metamorphic terranes, although some OIB groups appears to carry an imprint of heavy have begun to clarify the conditions in nature in Li (Nishio et al. 2005), indicating the survival of which they can help us understand such pro- crustally-fractionated material in at least one part cesses (e.g., Romer and Meixner 2014). of the upper mantle smorgasbord of recycled Although potentially quite complex in terms of components enumerated primarily through the interplay between diffusion and heteroge- radiogenic isotopes. The capacity to measure neous fluid flow, the capacity to approach Li isotopic variations to higher and higher precision isotope analysis on varying spatial scales may during the last decade has made constraining the yet allow valuable constraints on the histories of nature of less isotopically distinct deeply-recycled rocks in these settings. components feasible. Nevertheless, the isotopi- cally heavy HIMU component is currently the • “Clastic sediments are reservoirs of informa- only convincing Li isotopic fingerprint of tion about weathering processes, but are suf- broad-scale mantle heterogeneity so far revealed. ficiently complex that no study has yet to Formation of crustal silicate reservoirs with realize their potential. Despite a number of predominantly light Li isotopic composition, initial reports of relatively isotopically heavy coupled with quantification of Li in the modern samples, the majority of data for clastic sed- mantle, implies secular Li isotopic fractionation imentary rocks have an average δ7Li * 0 ‰, of the mantle. This was envisaged by Bouman equivalent to the estimated average isotopic et al. (2000) and Elliott et al. (2004), although the composition of the continental crust. results of Halama et al. (2007, 2008) do not indicate a fundamental change in mantle Li since Abundant data now exist for both terrigenous the late Archean. The potential impact of diffu- and marine sediment, as well as for sedimentary sional fractionation on the interpretation of rocks. The processes by which sedimentary rocks mantle rocks as recorders of mantle Li isotopes develop their Li isotopic signatures are now rel- over time present challenges to the realization of atively well documented, as is the average iso- a mantle evolution diagram like those used with topic composition of the crust itself. The capacity 7.1 Retrospective 193 of Li isotopes to constrain the incorporation of hence an evolution of the continents toward subducted sediment in the sources of arc magmas lower δ7Li over Earth history.” (e.g., Tang et al. 2014) is among the most com- pelling recent findings. Some of the greatest advances in the use of Li isotopes to better understand the Earth system • “Although ocean water is well mixed with have been in paleoceanography. Some rather respect to Li, and many laboratories measure thrilling results (e.g., Misra and Froelich 2012) Li isotopes in seawater, the range in values expand on interpretations derived from estab- reported in the literature (δ7Li = +29.3 to lished oceanic geochemical tracers (e.g., +33 ‰) cast some uncertainty on this reser- 87Sr/86Sr) and pave the way for future studies of voir. As precision of analytical methods Li as a proxy for global change. Exactly how Li improves, a check of the viability of the will be best used, though, is controversial. “seawater standard” should be carried out.” Studies have approached the problem from the oceans (in marine sediments or foraminifera) and The isotopic composition of Li in the modern from the continents (dissolved and particulate ocean is now well-characterized, through several constituents of river systems and fluvial excellent studies. The existence of a widely sediments). available certified seawater standard for Li iso- There remains, however, considerable uncer- topes (IRMM BCR-403; Millot et al. 2004)isa tainty on the interpretation of driving forces for reality. these isotopic records. The extent to which Li isotopes can be tied to continental weathering, • “Initial experiments on the viability of marine atmospheric temperature, and/or greenhouse gas (as well as lacustrine) carbonates as geo- fluctuation is hotly debated (no pun intended). chemical proxies have produced encouraging Needless to say there is abundant room for results toward the goal of tracing temporal important contributions in this area of research, changes in the Li isotopic composition of the from natural studies (including studies extending Earth’s surface. Inorganic and biogenic car- records further back in time), laboratory experi- bonate precipitation appears to favorably ments, and computational models. incorporate light Li, but this fractionation is not temperature sensitive. A better understanding • “Consensus should be reached on the inter- of metabolic effects of invertebrate species will pretation of Li isotope data for chondritic be necessary. More laboratory experiments meteorites. Can bulk planetary Li isotopic and controlled natural studies should address compositions be accurately estimated from the prospective of Li isotopes in assessing meteorites, or do these objects preserve only compositional changes in waters over time. parent body near-surface or metamorphic Modern analytical techniques, permitting pre- histories? Are either of these possibilities cise analysis of small (≤1 ng Li) samples, will viable? Detailed, high precision studies permit the examination of detailed populations should permit this assessment.” of these low concentration materials.” “Weathering releases isotopically heavy Li to Despite exciting new data from meteorites, the near-surface environment, regardless of major questions persist in cosmochemical appli- the intensity of chemical weathering. Much of cations of Li isotopes. The study of Li isotopes in this Li goes to stream flow, contributing to the components of individual meteorites appears to riverine average δ7Li * +24 ‰. The detailed be an essential step toward better defining both histories of weathering products on the con- the nature and evolution of these materials. tinents—sediments and soils—can be com- Whether the complex thermal and/or aqueous plex, but studies of simple systems suggest histories of certain meteorites do not allow for Li retention of isotopically light material, and isotopes to fully record the entire lifetime of 194 7 Summation: What Have We Learned and Where Can We Go? these materials remains to be seen, but it appears out. Spatially-resolved analytical techniques that coupling Li with other novel stable isotope finally are producing accurate and in many cases systems (Mg, Si, Fe, Cd) offers potential for very precise Li isotopic data. reaching heretofore unattainable understanding of the evolution of the early Solar System. Although a more comprehensive set of anal- 7.3 Final Words yses now exists for lunar materials, we are only beginning to understand the Li isotopic system- Chapter 4 lays out in great detail the theoretical atics of our nearest celestial neighbor. Like with and laboratory basis for understanding Li ele- asteroidal meteorites, dissected planetary mate- mental and isotopic variations in natural materi- rials need to be examined in order to fully als. Because of the fast diffusion of Li compared appreciate what Li isotopes have to tell about the to other elements, and because of the difference histories of these rocks and their source planets. in rates of diffusion of the Li isotopes, several Speculations into the capacity of Li isotopes to cautions resonate from this chapter. Experiments address questions of water in planetary mantles and computations give very clear boundary require further detailed investigation. conditions on time- and length-scales of diffusion and must be taken into consideration in the interpretation of data from natural materials. 7.2 Analytical Prospects Even when in situ analysis permits the evaluation of zoning within crystals, elemental and isotopic Progress since 2004 on the precise and accurate zonation have been misinterpreted. In light of the analysis of Li isotopes makes the measurement a complex effects of diffusion on Li and its iso- much more routine effort in a much greater topes the admonitions of Sect. 4.3.4 need to be diversity of laboratories around the globe today. kept in mind, especially, but not exclusively, for The development of very high precision analyt- the study of high temperature materials. ical techniques (e.g., external precision better Although the prospects for exciting discoveries fi than ±0.4 ‰) has opened up whole new areas of in the eld of Li isotopes appear certain, future study. Recent advances using less costly instru- work must attentively abide by the constraints mentation have paved the way for measurements imparted by theory and laboratory studies in at intermediate precision (±c. 1 ‰), particularly order for authentic progress to be made. applicable for studies requiring large data throughput, and making a substantial array of References scientifically important work feasible for many more laboratories world-wide. In bulk analysis it Bouman C, Elliott TR, Vroon PZ, Pearson DG (2000) Li is still clear that emphasis must be maintained on isotope evolution of the mantle from analyses of demonstrating the efficacy of chemical separa- mantle xenoliths. J Conf Abstr 5:239 tion, but this stricture appears to be well known Chan L-H, Edmond JM (1988) Variation of lithium now, and experimenters are watchful. isotope composition in the marine environment: a preliminary report. 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Earth Planet Sci Lett 254:77 89 Halama R, McDonough WF, Rudnick RL, Bell K (2008) on high spatial resolution (e.g., to examine evi- Tracking the lithium isotopic evolution of the mantle dence of diffusional processes) can be carried using carbonatites. Earth Planet Sci Lett 265:726–742 References 195

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