Molybdenum Stable Isotope Composition of Meteorites: Constraints on Planetary Core Formation 155 VI–1

Research Collection

Doctoral Thesis

Stable isotope behaviour during planetary differentiation Implications of mass dependent Fe, Si and Mo isotope fractionation between metal and silicate liquids

Author(s): Hin, Remco Christiaan

Publication Date: 2012

Permanent Link: https://doi.org/10.3929/ethz-a-007621154

Rights / License: In Copyright - Non-Commercial Use Permitted

This page was generated automatically upon download from the ETH Zurich Research Collection. For more information please consult the Terms of use.

ETH Library DISS. ETH NO. 20825

STABLE ISOTOPE BEHAVIOUR DURING PLANETARY DIFFERENTIATION Implications of mass dependent Fe, Si and Mo isotope fractionation between metal and silicate liquids

A dissertation submitted to

ETH ZURICH

for the degree of

Doctor of Sciences

presented by

Remco Christiaan Hin

Master of Science in Geosciences of Basins and Lithosphere, Vrije Universiteit Amsterdam

born on December 1, 1983

citizen of The Netherlands

accepted on the recommendation of

Prof. Dr. Max W. Schmidt Prof. Dr. Bernard Bourdon Prof. Dr. Tim Elliott

2012

(Panta rhei, ouden menei)

“And now that I have unfolded my sail at open sea and go with the wind – there is nothing in this whole world that remains.

Everything flows, each thing shapes and passes by.

Time too passes by in a continuous motion like a river that cannot stop its stream just as a running hour cannot stand still; as water pushes water forward while being pushed forward and pushing forward itself, as such time runs forward and chases itself and renews itself; what has been, is now past, and now is, what has not been; every moment changes.”

Pythagoras explains Herakleitos views, as passed on to us through Ovidius’ Metamorphosae. (Personal translation of the Dutch translation by M. D’Hane-Scheltema.) Table of Contents

Abstract/Zusammenfassung 3

Chapter I. Introduction 9 I–1. Planetary accretion and core formation 10 I–1.1 The formation of the Solar System 10 I–1.2. Differentiation of planetary objects 11 I–2. Fractionation of isotopes 15 I–2.1 Mass dependent fractionation of stable isotopes 15 I–2.2 Theoretical background of stable isotope fractionation 16 I–3. Objectives of this thesis 17

Chapter II. Methodology 19 II–1. Experimental 20 II–1.1. Starting mixtures 20 II–1.2. Piston cylinder experiments 20 II–2. From a quenched experiment towards chemical dissolution 32 II–3. Isotopic analyses 34 II–3.1. Ion exchange chemistry 35 II–3.2. MC-ICPMS 37 II–4. Failed Silicon isotope analyses by laser ablation 44

Chapter III. Experimental evidence for the absence of iron isotope fractionation betweenmetal and silicate liquids at 1 GPa and 1250-1300 °C and its cosmochemical consequences 47 III–1. Introduction 48 III–2. Methods 50 III–2.1. Experimental methods 50 III–2.2. Analytical methods 53 III–3. Results 56 III–3.1 Textures and elemental compositions of the experimental run products 56 III–3.2. Isotopic composition of the experimental liquids 60 III–4. Attainment of equilibrium 62 III–5. Discussion 67 III–5.1. Comparison to previous studies 67 III–5.2. Implications for Fe isotope variability in natural systems 69 III–6. Conclusions 74 III–7. Appendix 75

Chapter IV. Experimental determination of the Si isotope fractionation factor between liquid metal and liquid silicate 89 IV–1. Introduction 90 IV–2. Methods 92 IV–2.1. Experimental methods 92 IV–2.2. Analytical methods 94 IV–3. Results 97 IV–3.1 Textures and elemental compositions 97 IV–3.2. Isotopic compositions 100 IV–4. Equilibrium isotope fractionation 102 IV–5. Discussion 108 IV–5.1. Comparison with previous studies 108 IV–5.2. Implications for core formation 110 IV–6. Conclusion 112

Chapter V. Experimental evidence for Mo isotope fractionation between metal and silicate liquids 121 V–1. Introduction 122 V–2. Methods 124 V–2.1. Experimental methods 124 V–2.2. Analytical methods 125 V–3. Results 129 V–3.1 Textures and elemental compositions 130 V–3.2. Isotopic compositions 134 V–4. Discussion 135 V–4.1. Fractionation at 1400 °C 135 V–4.1. Fractionation at 1600 °C 142 V–5. Implications for core formation 142 V–6. Conclusions 145 V–7. Appendix 146 V–7.1. Mo concentration analyses by laser ablation ICPMS 150 V–7.1. Double spike calibration 151

Chapter VI. Molybdenum stable isotope composition of meteorites: Constraints on planetary core formation 155 VI–1. Introduction 156 VI–2. Analytical methods 158 VI–2.1. Sample preparation and Mo separation 158 VI–2.2. Mass spectrometry and data reduction 159 VI–3. Results 160 VI–3.1. Mo concentrations 160 VI–3.2. Mo isotope data and δ98/95Mo values 160 VI–4. Discussion 166 VI–4.1. Stable isotope composition of the inner solar system and bulk planetary bodies 166 VI–4.2. Comparison of experimental and observed Mo stable isotope fractionation during metal segregation 167 VI–5. Conclusions 172

Chapter VII. Conclusions 173 VII–1. Summary of the main conclusions 174 VII–2. General conclusions and implications 175 VII–2.1. Stable isotope fractionation 175 VII–2.2. Conditions of core formation 177

Bibliography 179

Appendix AI. 190 Appendix AII. 194 Appendix AIII. 196

Acknowledgements 204 Abstract/Zusammenfassung — Page 3

ABSTRACT

Stable isotope fractionation may inform on the conditions and chemical consequences of core-mantle differentiation in planetary objects. Equilibrium fractionation of stable isotopes, however, decreases strongly with increasing temperature and at very high temperatures becomes difficult to observe. With the advent of mass spectrometry, though, stable isotope fractionation has in many cases become observable during high temperature processes such as core formation. Nevertheless, the magnitude and direction of equilibrium stable isotope fractionation during such processes still remain largely unknown. In this study, I have experimentally determined the Fe, Si and Mo isotope fractionation factors between liquid metal and liquid silicate. I used these factors to interpret variations in stable isotope compositions in natural rocks and uncover their implications for planetary differentiation processes. Experiments were performed in a centrifuging piston cylinder in talc/pyrex assemblies at a pressure of 1 GPa. Elemental tin was used to lower the melting temperatures of iron-based alloys to below 1500 °C. Silicate compositions were adapted to reach an appropriate liquidus. The centrifugation segregated the liquid metal and the liquid silicate, enabling analyses of bulk pieces of metal and silicate that were free of cross contamination. The bulk pieces of metal and silicate were cleaned and crushed prior to ion exchange chemistry to separate the element of interest from its matrix. Analyses were then performed on a multi-collector inductively coupled plasma mass spectrometer. Experiments for Fe isotope fractionation were run at temperatures of 1250-1300 °C. The analyses demonstrate that 8 of the 10 experiments equilibrated in a closed isotopic system. Statistically significant iron isotope fractionation between the quenched metals and silicates was absent in 9 of the 10 experiments and all 10 experiments yield an average fractionation factor of 0.01 ± 0.04‰. The presence or absence of carbon or sulphur did not affect this result. At low pressures, Fe isotopes thus do not fractionate during metal-silicate segregation under magmatic conditions. This implies that the 0.07 ±�� 0.02‰ heavier composition of bulk magmatic iron meteorites relative to the average of bulk ordinary/ carbonaceous chondrites cannot result from equilibrium Fe isotope fractionation during core segregation. The up to 0.5‰ lighter sulphide than metal fraction in iron meteorites and in one ordinary chondrite can only be explained by fractionation during subsolidus processes. Silicon isotope fractionation between liquid metal and liquid silicate was studied at 1450 °C and 1750 °C. Metal is consistently enriched in light isotopes relative to the silicate, yielding average metal-silicate fractionation factors of -1.48 ± 0.08‰ and -1.11 ± 0.14‰ at 1450 and 1750 °C, respectively. These results are unaffected by the presence or absence of carbon. The temperature Page 4 — Abstract/Zusammenfassung dependence of equilibrium Si isotope fractionation between metal and silicate 30 × 6 2 liquids can therefore be described as Δ SiMetal-Silicate = -4.47(±0.31) 10 /T . Using a bulk silicate Earth δ30Si value of -0.29 ± 0.01‰, this temperature dependence can be used to calculate δ30Si values for Earth’s core. For metal-silicate equilibration temperatures of 2500-3500 K, the Si isotope composition of Earth’ core is -1.01 ± 0.05‰ to -0.66 ± 0.03‰. Mass balance calculations indicate that for those core compositions, the Bulk Earth has a δ30Si of -0.38 ± 0.02‰ or -0.33 ± 0.01‰, respectively, if Earth’s core is assumed to contain 6 wt% Si, or -0.40 ± 0.02‰ or -0.35 ± 0.02‰, respectively, for 8 wt% Si in the Earth’s core. The reported Si isotope compositions of enstatite chondrites are significantly lighter (-0.62 ± 0.05‰) and they are therefore unlikely to represent the Bulk Earth Si isotope composition. A comparison with the reported Si isotope composition of other chondrites is unfortunately hindered by the dispute about these Si isotope compositions. Molybdenum isotope compositions in the silicate were 0.193 ± 0.030‰ and 0.120 ± 0.020‰ heavier than in the metal at 1400 and 1600 °C, respectively. This fractionation is independent of the presence or absence of carbon and tin as well as of oxygen fugacity in the range ΔIW-1.79 to ΔIW+0.47. Equilibrium Mo isotope fractionation between liquid metal and liquid silicate can therefore 98/95 × 5 2 be described as Δ MoMetal-Silicate = -4.80(±0.55) 10 /T . The experiments performed at 1600 °C furthermore demonstrate that non-equilibrium isotope fractionation results from highly reactive capsule material. Rapid dissolution of 98/95 capsule-derived MgO into the silicate melt has led to Δ MoMetal-Silicate of -0.027 ± 0.036‰, clearly resulting from disequilibrium isotope fractionation. Initial determination of Mo isotope compositions of meteorites as well as terrestrial and lunar basalts indicates that equilibrium metal-silicate segregation may have occurred at ~1800 °C in the angrite parent body and at ~2100 °C in the Earth-Moon system. These new results suggest that core formation in the Earth did not occur at the base of a deep magma ocean, but rather took place at a shallower depth, either during descent of metal droplets through the magma ocean or by metal-silicate equilibration in Earth’s precursor bodies. Overall, the findings in this study demonstrate that stable isotope fractionation can occur during metal-silicate differentiation events. It may therefore constitute a tool to further constrain models of core formation. In this respect, Mo isotopes imply that models of high pressure and high temperature equilibrium core formation at the base of a deep magma ocean may be oversimplified. Finally, it appears that a factor two increase in valence state of an element in the silicate liquid may increase stable isotope fractionation between metal and silicate liquids by an order of magnitude, which appears a stronger effect than that caused by a change from VI-fold to IV-fold coordination state. This furthermore emphasises that equilibrium fractionation of stable isotopes is not solely a function of element mass. Abstract/Zusammenfassung — Page 5

ZUSAMMENFASSUNG

Die stabile Isotopenfraktionierung gibt Informationen über die Umstände und chemischen�� Folgen von Kern-Mantel Differen�z��ierungspro�z��essen in pla- netaren Objekten. Gleichgewichtsfraktionierung stabiler Isotope nimmt mit steigender Temperatur stark ab, bis die Fraktionierung schwer messbar wird. Aufgrund Verbesserungen in der Massenspektrometrie kann heute jedoch die Isotopenfraktionierung in Hoch-Temperatur Prozessen - wie zum Beispiel der Kernbildung - in vielen Fällen gemessen werden. Trotzdem ist das Ausmass und die Richtung der Gleichgewichtsfraktionierung stabiler Isotope während solcher Prozesse grösstenteils immer noch unbekannt. In dieser Studie habe ich an Hand von Experimenten die Fe, Si und Mo Isotopenfraktionierungsfaktoren zwischen Metal- und Silikatschmelzen bestimmt. Diese Faktoren habe ich benutzt, um Unterschiede in der stabilen Isotopenzusammensetzung in Gesteinen zu erklären und ihre Auswirkung auf planetarische Differenzierung zu ermitteln. Die Experimente wurden mit Talk/Pyrex Bauteilen in einem zentrifugieren- den Stempel-Zylinder bei einem Druck von etwa 1 GPa ausgeführt. Elementares Zinn wurde benutzt, um den Schmelzpunkt von auf Eisen basierten Legierungen bis unter 1500 °C zu senken. Gleichzeitig wurde die Silikat Zusammensetzung angepasst, um den gewünschten Liquidus zu erreichen. Die Zentrifugierung trennte die Metall- und Silikatschmelzen, wodurch die zwei Phasen ohne Risiko von Kreuzverunreinigungen analysiert werden konnten. Die zwei Phasen wurden gesäubert und pulverisiert bevor das gewünschte Element durch Ionenaustausch- chemie von seiner Matrix getrennt wurde. Die Proben wurden anschliessend auf einem Massenspektrometer gemessen. Die Experimente für Fe Isotopenfraktionierung wurden bei Temperaturen von 1250 bis 1300 °C�� ausgeführt. Die Analysen z�eigen,�� dass in 8 von 10 Experi- menten das Gleichgewicht in einem isotopisch geschlossenen System erreicht wurde. In 9 von 10 Experimenten gab es keine statistisch bedeutende Fe Iso- topenfraktionierung zwischen den abgeschreckten Metallen und Silikaten. Die 10 Experimente ergaben einen durchschnittlichen Fraktionierungsfaktor von 0.01 ± 0.04‰. Die An- oder Abwesenheit von Kohlenstoff oder Schwefel hatte keinen Effekt auf die Fraktionierung. Folglich findet unter magmatischen Bedingungen und bei tiefen Drücken während der Metall-Silikat Differenzie- rung keine Fe Isotopenfraktionierung statt. Das bedeutet, dass die 0.07 ± 0.02‰ schwerere Zusammensetzung magmatischer Eisenmeteorite im Vergleich zu Chondriten nicht durch Gleichgewichtsfraktionierung von Fe Isotopen während der Kernbildung erklärt werden kann. Die bis zu 0.5‰ leichtere sulfidische Zusammensetzung verglichen mit metallischen Teilen in Eisenmeteoriten und Chondriten können nur durch Fraktionierung während subsolidus Prozessen erklärt werden. Page 6 — Abstract/Zusammenfassung

Die Si Isotopenfraktionierung zwischen Metall- und Silikatschmelzen wurde bei Temperaturen von 1450 und 1750 °C untersucht. Metall war im Vergleich zu Silikat immer mit den leichten Isotopen angereichert und ergab durchschnittliche Fraktionierungsfaktoren von -1.48 ± 0.08‰ und -1.11 ± 0.14‰ fü��r 1450 be�z��ie- hungsweise 1750 °C. Die An- oder Abwesenheit von Kohlenstoff hatte keinen Einfluss auf die Ergebnisse. Die Temperaturabhängigkeit der Si Isotopenfraktio- nierung zwischen Metall- und Silikatschmelzen kann demnach folgendermassen 30 × 6 2 ausgedrückt werden: Δ SiMetal-Silicate = -4.47(±0.31) 10 /T . Diese Temperaturab- hängigkeit kann dazu verwendet werden, den d30Si Wert des Erdkernes zu berechnen, wenn man einen Wert von -0.29 ± 0.01‰ für die silikatische Erde annimmt. Für Metall-Silikat-Gleichgewichtstemperaturen von 2500-3500 K beträgt die Si Isotopenzusammensetzung des Erdkernes -1.01 ± 0.05‰ bis -0.66 ± 0.03‰. Massenbilanzberechnungen deuten darauf hin, dass für diese Kernzu- sammensetzung die Gesamterde ein δ30Si von -0.38 ± 0.02‰ oder -0.33 ± 0.01‰ hat, wenn der Erdkern 6 wt% Si enthält, beziehungsweise -0.40 ± 0.02‰ oder -0.35 ± 0.02‰ wenn der Erdkern 8 wt% Si enthält. Die publizierten Si Isotopen- zusammensetzungen von Enstatit-Chondriten sind wesentlich leichter (-0.62 ± 0.05‰) und repräsentieren daher eher nicht die Si Isotopenzusammensetzung der Gesamterde. Ein Vergleich zu die publizierten Si Isotopenzusammensetzung von anderen Chondriten ist wegen der Kontroverse um diese Si Isotopenzusam- mensetzungen nicht möglich. Molybdän Isotopenzusammensetzungen in Silikaten waren 0.193 ± 0.030‰ und 0.120 ± 0.020‰ schwerer bei 1400 beziehungsweise 1600 °C als im Metall. Diese Fraktionierung ist unabhängig von der An- oder Abwesenheit von Kohlenstoff und Zinn, sowie von der Sauerstofffugazität zwischen ΔIW-1.79 und ΔIW+0.47. Die Fraktionierung von Mo Isotopen zwischen Metall- und 98/95 Silikatschmelzen kann wie folgt beschrieben werden: Δ MoMetal-Silicate = -4.80(±0.55)×105/T2. Ferner haben die Experimente bei 1600 °C gezeigt, dass die Ungleichgewichtsfraktionierung in den jeweiligen Experimenten auf das stark reagierende Kapselmaterial zurückzuführen ist. Sich rasch in Silikatschmelze 98/95 auflösendes MgO der Kapsel verursachte Δ MoMetal-Silicate von -0.027 ± 0.036‰, was eindeutig der Ungleichgewichtsfraktionierung zuzuschreiben ist. Erste Mo Isotopenzusammensetzungen von Meteoriten sowie terrestrischen und lunaren Basalten deuten darauf hin, dass die Gleichgewichtstrennung zwischen Metall und Silikat in Angrite Mutterplanetoiden bei ~1800 °C und im Erde-Mond System bei ~2100 °C stattfand. Diese neuen Resultate zeigen, dass die Kernbildung nicht am Boden eines tiefe Magmaozeans stattfand sondern in weniger tiefen Bereichen entweder in der Form von Metalltröpfchen, die durch den Magmaozean wanderten, oder innerhalb von Planetoiden. Insgesamt zeigt diese Studie, dass die Fraktionierung stabiler Isotope während der Metall-Silikat-Differenzierung möglich ist. Stablile Isotopenfraktionierung kön��nte daher bunut�z��t werden um Modelle des Kernbildungsvorgangs �z��u verbes- Abstract/Zusammenfassung — Page 7 seren. Die Mo Isotope deuten darauf hin, dass solche Modelle unter Annahme von Gleichgewichtsbedingungen bei höheren Temperaturen und Drücken am Boden eines tiefe Magmaozeanes bisher zu stark vereinfacht wurden. Zudem haben die Resultate gezeigt, dass eine Verdopplung des Oxidierungsgrad eines Elements in der Silikatschmelze in einer Verzehnfachung der Isotopenfraktio- nierung zwischen Metall- und Silikatschmelze resultiert, was einen grösseren Einfluss als der Koordinierungsgrad scheint zu sein. Dies verdeutlicht letztlich, dass Gleichgewichtsfraktionierung von stabilen Isotopen nicht nur vom Gewicht des Elements abhängt.

CHAPTER I.

INTRODUCTION Page 10 — Chapter I

This thesis concerns the experimental investigation of equilibrium, mass dependent fractionation of stable isotopes between liquid metal and liquid silicate. Liquid metal and liquid silicate are thought to be the dominant phases during core-mantle segregation in rocky (terrestrial) planetary bodies. This process leads to the formation of metallic cores and silicate mantles in terrestrial planets and their smaller precursor bodies (planetesimals and planetary embryos). It is the largest chemical and physical event in the history of terrestrial bodies. On Earth, core formation eventually led to a magnetic field that protects us from harmful cosmic radiation. Core formation is also for a large part responsible for the rareness of precious metals like gold, because most of the mass of such siderophile elements is hidden from us in Earth’s core. Differentiation into a metallic core and silicate mantle occurs in the early history of planetary objects, i.e. in roughly the first 100 My after formation of the Solar System 4.567 Ga. This differentiation is normally studied by comparing undifferentiated meteorites (chondrites) to differentiated meteorites or samples from the terrestrial mantle/crust. Such comparisons have evolved from mineralogical to chemical, the latter focussing on elemental abundances. In the last decade, relative abundances of stable isotopes have emerged as a new tool to investigate core segregation (and many other processes in geosciences). Hitherto, it remains unclear how to link the measured relative abundances of stable isotopes to geological processes, as there is only limited experimental and theoretical understanding of how stable isotopes behave during geological processes. Therefore, my study is a further step to better understand the behaviour of stable isotopes during geological processes. Because it is such an important event in the history of various planetary bodies, I have chosen to apply my study to the process that differentiates planetary bodies into a metallic core and silicate mantle. A more personal motivation for that application is that a large scale process that occurred almost 4.5 billion years ago triggers my imagination: I enjoy trying to understand/visualise how that process may have occurred, and I enjoy to search for facts that can serve as the fundament for that understanding and to force my imagination to stick to those facts. I hope that upon reading my work, you will enjoy it as much as I did while performing it, and that you find it brings us a small, but significant step forward in understanding “stable isotope behaviour during planetary differentiation”.

I–1. PLANETARY ACCRETION AND CORE FORMATION I–1.1. The formation of the Solar System The Solar System formed by collapse of a dense molecular cloud (De Pater and Lissauer, 2001). Such clouds typically have a few thousand molecules per 3 cm , have a temperature of 10-30 K, and consist mainly of H2 and He with small amounts of molecules containing H, C, N and O. Usually, such clouds are stable Chapter I — Page 11 in the sense that their kinetic (thermal) energy balances their gravitational energy. When either their cores reach too high densities or, more commonly, a trigger such as (shock) compression occurs, these clouds start to collapse towards the initially densest centre. This process is self-amplifying due to the continuously increasing density and gravitation in the core of the cloud, nearly all mass being attracted into the collapsing core. The Sun contains over 99.8% of the mass of the Solar System. Mainly due to increased pressure and transformation of gravitational energy into kinetic (thermal) energy, temperatures inside the growing Sun raised sufficiently high (>107 K) for nuclear reactions. Over 98% of the angular momentum in the Solar System is in the remaining ≤0.2% of mass. This rotational energy has prevented that mass from accreting into the Sun and formed a flattened (proto-planetary) disk around the equatorial plane of the growing Sun. The heat produced by this flattening and particularly the radiative heat of the Sun is thought to have eventually created temperatures of a few thousand K close to the Sun and ~100 K at 10 AU. This is the main reason that the inner part of the Solar System consists of rocky (‘terrestrial’) planets, while the outer part consists of gas giants. Upon cooling of the gas in the inner parts of the proto-planetary disk, condensation temperatures are reached and micrometre particles form. As originally detailed in Grossman (1972), Al2O3 and CaTiO3 are the first condensates. So-called Ca- and Al-rich inclusions (CAI’s) are therefore considered to be the oldest particles in the Solar System, dated to be 4.567 Ga (Amelin et al., 2002). The particles that formed by condensation, together with very few interstel- lar particles (‘pre-solar grains’), rapidly accreted to form ~10 kilometre sized bodies. The evolution stage from micrometre particles to 10 km sized bodies is poorly understood, but afterwards growth continues by gravitational attraction of surrounding material towards the ≥10 km sized body. Such bodies are thought to still exist in the asteroid belt between Mars and Jupiter, and samples of them occur on Earth as meteorites called chondrites. The final evolution of the Solar System towards its present state encompasses the accretion of the various bodies into the eight planets with their moons, and various asteroids. The formation of the planets involved a stage of run-away growth into planetary embryos (~Mars size) and a final stage with large-scale impacts, such as the Giant Impact that is thought to have formed Earth’s Moon. I–1.2. Differentiation of planetary objects Once bodies reach sizes over 10 km (‘planetesimals’), they may retain sufficient heat to differentiate into a metallic core and silicate mantle due to melting (Hevey and Sanders, 2006). The heat required for this melting can come from conversion of gravitational energy into thermal energy, from impact heating, and from decay of abundant short-lived nuclides early in the Solar System, most notably the decay of 26Al to 26Mg (De Pater and Lissauer, 2001; Ghosh et al., Page 12 — Chapter I

2006). Samples of differentiated bodies occur on Earth in the form of achondritic meteorites originating from the silicate portions of such bodies, and iron meteorites originating from their metallic cores. Numerical modelling and radiogenic age determinations indicate that the above processes from molecular cloud collapse to the growth of the first planetesimals lasted about 1 My (De Pater and Lissauer, 2001; Kleine et al., 2009; Krot et al., 2009). Age determinations also indicate that the formation of planetesimals was a somewhat complicated process that continued for at least 4 My (Krot et al., 2009). Due to higher material densities and shorter orbital periods, material close to the Sun accreted faster than material at larger heliocentric distances (Ghosh et al., 2006). As a consequence, there were planetesimals that had accreted when abundant heat was still produced by decay of 26Al. These bodies melted (nearly) completely and had therefore differentiated before the precursor bodies of chondrites formed (Kleine et al., 2009). Chondrites, and more specifically the carbonaceous type CI chondrites, are considered to represent the average composition of the Solar System (Zanda, 2004). The CI chondrites have a one-to-one correlation of elemental abundances with the solar photosphere (normalised to 106 Si atoms), except for the volatile noble gases, H, He, C and O. This observation forms the basis of the ‘chondritic model’: the assumption that terrestrial planetary bodies consist of Solar System material with an average bulk composition similar to the Solar System. Excep- tions to this rule are volatile elements that may be lost by volatilisation during planetary accretion and, therefore, do not occur in Solar System abundances. The chondritic model and the cosmochemical classification of elements form the framework of virtually all studies of core segregation in planetary bodies. The cosmochemical classification (Palme and Jones, 2003) groups elements according to their condensation temperatures as well as according to their preference for metal or silicate phases. Based on these properties, we predict which elements should occur in chondritic relative abundances in a body (i.e. the refractory elements) and which elements would be fractionated between metal cores and various silicate reservoirs (e.g. mantle and crust). Compositions of silicate samples can then be compared to CI chondritic compositions to estimate Bulk Silicate compositions of differentiated planetary objects, i.e. homogenising the various silicate reservoirs. Many such attempts have been made for the Earth (Palme and O’Neill, 2003) and they serve as a basis for estimates of the composition of Earth’s core (McDonough, 2003). This led Ringwood (1959) to propose that Si may be present in the core in wt% levels in addition to a Fe90Ni10 alloy. For bodies such as Mars or the various parent bodies of meteorites, core composition estimates are scarcer because of the limited number of samples relative to the Earth. Core composition estimates also require more detailed knowledge than cosmochemical classification. Experiments determining element distribution between Chapter I — Page 13 metal (liquid) and silicate (liquid) have greatly contributed to our knowledge of core formation (e.g. Wood et al., 2006). The distribution coefficients determined in such studies not only constrain which elements may be extracted from the silicate mantle into the metallic core. They also establish pressure, temperature, composition and oxygen fugacity conditions under which these elements are extracted from the mantle as well as the magnitude of extraction. Element distribution coefficients between metal and silicate can therefore be used to constrain hypotheses on the composition of metal cores, but also the conditions of core segregation (Righter and Drake, 1996; Righter, 2003; Wood et al., 2006). Core segregation in Earth is described with equilibrium and disequilibrium models (Figure I–1). The latter assume that the material that accreted to Earth changed composition in the course of accretion and core formation (Wänke, 1981). Equilibrium models are easier to test against data. Moreover, the recent models explain features of the Bulk Silicate Earth (BSE) that formerly required disequilibrium models, most notably the higher than expected abundances of certain siderophile elements. Employing experimentally determined distribution coefficients, equilibrium models explain present day element abundances by high pressure core segregation in a progressively oxidising Earth (Wade and Wood, 2005).

a b

Equilibrated Molten metal droplets Completely mantle molten mantle Metal pond

Metal Unequilibrated Solid diapir metal blobs mantle

Metal Metal core core

Figure I–1. Sketches that represent two models of core formation. Panel (a) suggests that impacting material emulsifies and that small sinking droplets of metal equilibrate with the molten silicate. At the top of the solid mantle, the metal droplets then pond until sufficient mass allows for large metal diapirs to sink to the core without further equilibration with the surrounding silicate (after Wood et al., 2006.) Panel (b) shows an alternative model in which large impactors melt the entire silicate mantle of the impacted body. The molten metal of the impactor has such high velocities that it does not emulsify, but rapidly merges with the core of the impacted body. In this scenario, no re-equilibration occurs between metal and silicate in the impacted body. Metal-silicate equilibrium conditions are thus inherited from the smaller impactor. Page 14 — Chapter I

A recurring issue with equilibrium models, however, is the fact that physical models predict large impactor bodies to merge their core with Earth’s core without re-equilibration. The large impacts in the final stage of accretion of terrestrial planets involves so much energy that the impactor core would not emulsify, a requirement for fast equilibration with the Earth’s molten silicate mantle. In this case, an equilibrium core segregation model would physically not be valid. Rudge et al. (2010), however, showed that elemental abundances of siderophile elements in the BSE can be equally well explained by equilibration of all accreting material as by equilibration of only part (~40%) of it. This implies that conditions of core segregation at lower pressures (and temperatures) in impactors may be inherited in Earth, and, hence, that the high pressures and temperatures (20-60 GPa and ~2500-3000 °C) derived from equilibrium models may not have been a precondition for core segregation on Earth. Stable isotope fractionation, as opposed to elemental distribution, between metal and silicate (liquids) may provide further insights into core segregation in planetary objects. As detailed in the next section, mass dependent stable isotope fractionation between two phases strongly depends on temperature. Further- more, sudden changes in fractionation may appear due to pressure related phase changes. As such, mass dependent stable isotope fractionation may help dis- criminating between equilibrium and disequilibrium models of core segregation. Williams et al. (2006), for instance, hypothesized that Fe isotopes may imply the presence of the high pressure phase perovskite (>20 GPa) during core segregation on Earth. By contrast, Moynier et al. (2011) suggested that Cr isotope compositions imply an inheritance of relatively low temperatures of core segregation in smaller impactors. To fully extract the information stable isotopes may provide, we need to understand their fractionation behaviour. The contrasting hypotheses of Williams et al. (2006) and Moynier et al. (2011) are both based on stable isotope compositions of meteorites and terrestrial samples. It requires theoretical understanding of stable isotope fractionation and experimentally determined magnitudes and directions of fractionation (‘fractionation factors’) to interpret those data. While theoretical predictions of fractionation factors can be made, so far they are fairly restricted in number and usually can only be made for solid material (Polyakov et al., 2007). Furthermore, the accuracy of such predictions has rarely been verified by experimental determinations of fractionation factors. It is therefore the purpose of this thesis to experimentally determine stable isotope fractionation factors between liquid metal and liquid silicate in order to use them for the interpretation of stable isotope compositions of meteorites and terrestrial samples in a framework. By doing so, I aim to shed further light on the process of core segregation. Chapter I — Page 15

I–2. FRACTIONATION OF ISOTOPES I–2.1. Mass dependent fractionation of stable isotopes Isotopes can be fractionated both independent of and dependent on their masses. Mass-independent fractionation is well known from radiogenic ingrowth of a single isotope of an element due to decay of another element. There are also other forms of mass-independent fractionation of isotopes, for instance by prefer- ential production of specific isotopes of an element in supernovae or other stellar environments (‘nucleosynthesis’) or by interaction of isotopes with cosmic rays. These types of isotopic fractionation are not the topic of this thesis, which studies mass dependent fractionation. This type of fractionation is a function only of the masses of the isotopes and involves two sub-types: kinetic and equilibrium fractionation. Kinetic mass dependent fractionation of isotopes can qualitatively be understood to occur because it requires more energy to transport a heavy mass than a light mass. For a given amount of energy, heavy isotopes will thus travel slower than light isotopes, meaning that the residue is relatively enriched in heavy isotopes if a transfer process is interrupted prior to completion. The specific type of fractionation I have studied is equilibrium mass dependent isotope fractionation (in the remainder I will often refer to this with the slightly simpler term ‘stable isotope fractionation’). This occurs when heavy isotopes relative to light isotopes prefer one phase more than a second phase. Rudge et al. (2009) described the process of mass dependent isotope fractionation mathematically as: α ii XXmi = ⋅ jjXX (I.1) O m j where i and j are the isotopes of element X and m are their isotopic masses. By convention, isotope j in these ratios is usually the heavy isotope. The factor α is the mass fractionation factor. The subscript 0 indicates the initial isotopic ratio prior to fractionation. Somewhat unfortunate, the symbol α is also used to describe equilibrium mass dependent isotope fractionation between two phases A and B, mathematically defined as: i X j X A α AB− = (I.2) i X j X B The convention in equation (1.2) is the most commonly used in studies of stable isotope fractionation. This α��-notation,�� though,�� is mainly used by theoreti- cians; geochemists mostly use the symbol Δ for the fractionation factor, which is simply the difference between the measured isotopic compositions of the two Page 16 — Chapter I phases A and B (see more about the expressions of isotopic compositions in Chapter X). It is this Δ��-notation ��that I will use throughout this thesis. For geological purposes, α does generally not deviate from 1.00 by more than 0.01, such that α and Δ can be simply related as: ij/ij // ij 1000ln()αAB−−=D=−XXXAB ddAB (I.3) where δi/jX stands for the measured isotopic composition given as the ratio of isotopes i and j of element X. I–2.2. Theoretical background of stable isotope fractionation In principle, stable isotope fractionation can occur between any phases, and therefore as a consequence of any (geo)chemical process. The preference of heavy isotopes relative to light ones for one phase (A) compared to another (B) arises from differences in the bond stiffness between the two phases A and B. Furthermore, this preference is dependent on the relative masses of isotopes and is proportional to the inverse of the square of temperature, without a first order pressure dependence. Schauble (2007) summarised the general mathematical relationship as: Dm DF ln()α ≈ AB− (I.4) AB− mT22

in which ΔFA-B refers to the difference in force constants (i.e. bond stiffness) between phases A and B. The term Δm refers to the difference in masses of the isotopes. The temperature T is in K. Equation (1.4) results from simplifications of thermodynamic and quantum mechanical relations used to calculate equilibrium constants for isotope exchange reactions (Bigeleisen and Mayer, 1947). For detailed derivations of those equations and their physicochemical background, I recommend the mainly qualitative descriptions of Bigeleisen (1965) as well as the more quantitative works of Schauble (2004), Urey (1947) and Bigeleisen and Mayer (1947). I have made a summary of the key points in Appendix A1. It follows from equation (1.4) that stable isotope fractionation may always occur, except when there is no mass difference in the isotopes or when there is no difference in the bond stiffness between the two phases. Smaller differences in isotopic mass or bond stiffness lead to smaller fractionation. In general, heavy isotopes tend to concentrate in the phase with the stiffest bonds. There are some qualitative rules of thumb that are useful for estimating the magnitude of the difference in bond stiffness between two phases of interest (Schauble, 2004). The most important ones for geological materials state that the stiffest bonds tend to occur for i) elements with a high valence state, ii) low coordination numbers, and iii) for transition elements with low-spin electronic configurations. Furthermore, in a general classification of bond types, covalent bonds are stronger than ionic bonds while metallic bonds are weakest. Increases in the element mass and in Chapter I — Page 17 the temperature of fractionation lead to decreases in the fractionation. This is the reason that until the mid-1990’s stable isotope fractionation was almost only detectable in low temperature environments (i.e. below ~200 °C) and for light elements (e.g. H, C, O), which have relatively large differences between their isotopic masses. With the progress in mass spectrometry, measurement precision has improved to levels at which fractionation at temperatures exceeding 1000°C can be detected for isotope systems of ‘non-traditional’ elements, i.e. elements with masses exceeding those of S. The detectability of stable isotope fractionation enables the use of stable isotope fractionation for (geo)chemical, and thereby geological, processes. This requires full knowledge of the direction and magnitude of stable isotope fractionation that a (geo)chemical process yields.

I–3. OUTLINE OF THIS THESIS

I have chosen to investigate three different isotope systems: Fe, Si and Mo. Isotopic compositions of a variety of meteorites and terrestrial samples were available for Fe and Si, and experimental constraints are necessary to interpret these data (see for instance a discussion about Fe isotopes in EPSL, vol. 256). Molybdenum was chosen as a third element because mass balance calculations, and their inherent uncertainties, are generally not required to interpret Mo isotope compositions of natural silicate samples. In addition, Mo has a high and possibly variable valence state: both Mo4+ and Mo6+ may occur in equilibrium with metal. Finally, these three elements cover a range of atomic masses from 28.0855 atomic mass units (amu) for Si to 95.96 amu for Mo. They also cover a range of valence states in the silicate from 2+ (Fe) to 4+ (Si and Mo) to possibly 6+ (Mo). The remainder of this thesis thus consists of three main chapters describing, respectively, Fe, Si and Mo isotope fractionation between liquid metal and liquid silicate, and their implications for core segregation in planetary objects. An additional collaborative chapter presents Mo isotope compositions of a variety of meteorites and terrestrial samples. Prior to these chapters, however, I will describe all the methodology involved in this thesis.

CHAPTER II.

METHODOLOGY Page 20 — Chapter II

In this work, I have combined methods of experimental petrology and isotope geochemistry. These techniques will be described in the current chapter to provide the details necessary for any future research following up on this work. In addition, this chapter contains information about failed methodologies that may once prove useful. This chapter follows the set-up of my study, starting with the experimental part, followed by elemental analyses and preparation for treatment in a clean chemistry laboratory before I describe the ion exchange chemistry and isotopic analyses in detail. I hope it is educational in the parts where you have no experience.

II–1. EXPERIMENTAL II–1.1. Starting mixtures All experiments are based on synthetic powders as starting mixtures. The silicate components have been prepared from (usually 99.99+%) purified oxide powders, except for the elements Ca, K and Na. These elements are available as carbonates and were first mixed with SiO2 (and Al2O3) powder to create molar ratios of CaO, K2O and Na2O to SiO2 (to Al2O3) equivalent to those of the minerals wollastonite, orthoclase and albite. These powder mixtures were then placed in a furnace at 400 °C and slowly heated to 1500 °C (wollastonite) and ~1100 °C (the two feldspars). The fused material was then crushed and stored for further use in oxide mixtures representing silicate melts, which in my study always (and only) contained additional SiO2, Al2O3, MgO and Fe2O3. Metal components are all available as element powders so that they can simply be weighed in the desired proportions prior to homogenisation. Such homogenisation is generally done with a pestle and mortar. Because metals can be very soft and easily smear out under the pestle, I always first mixed the harder oxides before adding the softer metals. After gentle and patient stirring under acetone or ethanol, a fine and homogeneous powder forms, although the homogeneity decreases by density separation during evaporation of the excess acetone or ethanol. The resulting heterogeneity is, however, easily homogenised again with a short step of dry stirring. Most of the various starting mixtures are presented in the relevant chapters and in Appendix AII. II–1.2. Piston cylinder experiments II–1.2.1. General The technique to generate high pressures and temperatures to simulate conditions below the Earth’s surface is straightforward. Temperatures of up to 2500 °C are generated by running an electric current of up to ~400 A through a graphite cylinder (‘furnace’) inside which the sample powder is positioned inside a container (‘capsule’). The resistance of the graphite furnace then produces the heat, a process no different than the process that produces heat in a standard light Chapter II — Page 21

bulb. Quenching is done by thermocouple wire in cutting power. Pressures ≥ graphite furnace Al2O3 or Mullite sleeve 0.5 GPa are produced by Talc pumping oil into a hydraulic P M M P cylinder reservoir with a yr g g yr O O large surface which in turn e e x x pressurizes a small surface (or volume), the latter containing the sample capsule. Corundum Pressure transfer onto the sample capsule inside the Capsule furnace takes place through 36.0 3.7 Sample material that becomes very ~10.5 soft at the temperatures Metal and pressures of interest, jacket yielding approximately hydrostatic pressures. In my 6 mm Cement study, these materials were always talc and pyrex/silica glass (see Figure II–1 for a sketch of an assembly). A piston cylinder generates pressures of 0.5-4 GPa and allows for relatively 14 mm large capsules. The pressure transfer in this apparatus Figure II–1. Design of assembly. Dimensions are in takes place through a millimetre. conically shaped ‘bridge’ and a tungsten-carbide (WC) piston that is in contact with the furnace and other assembly materials surrounding the capsule. A sketch of a Boyd and England type (Boyd and England, 1960) end-loaded piston cylinder is shown in Figure II–2. The term end-loaded refers to an additional axial pressure that is exerted on the WC core material column, but that is not transferred onto the sample itself. This ‘end-load’ serves to compress the WC also axially, lateral support being provided by the pressure vessel, thus counteracting the pressure exerted onto the sample. Non-end-loaded piston cylinders are consequently used until lower pressures (< 2 GPa). A special type of piston cylinder at ETH Zurich is the centrifuging piston cylinder, details of which can be found in Schmidt et al. (2006). It is a min- iaturised version of a non-end-loaded piston cylinder that is mounted into a rotating table (Figure II–3). Its usage is a little more elaborate than that of a standard version, because oil cannot be pumped in during centrifugation to adjust Page 22 — Chapter II

Figure II–2. Design of a Boyd and England-type piston metal frame cylinder. Inside the metal frame, the parts with solid, dark grey cooling water fill are made of tungsten-car- plate bide. The bomb has a white to bomb with dark grey gradient fill. Parts

water

cooling WC core with a sand pattern fill are non- piston (WC) moveable. Moveable parts are moved up by pumping hydraulic bridge oil in the reservoir underneath

(piston) ram with WC- them, while they are moved

hydraulic oil pushing piece downwards by pumping oil in end-load ram the reservoir along their sides. Pressure is calculated with a conversion factor between the ram and the piston (surface

(end-load) hydraulic oil of ram divided by surface of piston) with a correction for friction determined from pressure calibration. (Courtesy of Peter Ulmer.) 0 10 20 cm pressures. In contrast, it is advisable to heat only during centrifugation, such that high currents are only pulsed through the slip rings when rotating. As the material compacts due to the exerted pressure, particularly after the talc and pyrex/silica glass soften, pressure is lost during heating. A multi-step oil pumping/heating procedure is thus required for experiments in the centrifuging piston cylinder. All the experiments I analysed for their isotopic compositions have been run in this centrifuging piston cylinder, because the analyses require metal and silicate powders that are free of cross contamination. As the experiments were designed to contain liquid metal and liquid silicate, which have densities of ~7500 and ~3000 kg m-3, respectively, Figure II–3. Picture of the centrifuging piston near-perfect separation can be cylinder at ETH Zurich. Highlighted are the min- achieved during the experi- iaturised non-end-loaded piston cylinder (similar ment by centrifugation (Figure to the design in Figure 2, but without the end-load II–4a). Cross contamination, ram) and the counter-mass. Both are mounted on if any, only consists of tiny the rotating table. Chapter II — Page 23

Figure II–4. Back-scatter electron image of a centrifuged experiment (a) compared to a static experiment (b) under identical conditions. White reflections are metal, grey are silicate and black are graphite or laromin epoxy. droplets of the contaminating phase (< 2 μm in diameter) remaining in the main pool of one of the phases, such that the analysed silicate and metal phases are >99.9% pure. This contrasts with a conventional static experiment that can lead to phase separation as poor as in Figure II–4b. Such experiments would require careful hand picking to separate the metal and silicate phases, an approach unlikely to lead to purities as high as by centrifugation. Richter et al. (2008) have found that temperature gradients can lead to similar isotopic fractionation in a single phase as investigated here between two phases. Temperatures in a cylindrical graphite furnace vary between the cold ends in contact with metal from the stack and a hotspot region near the centre. Because temperature gradients are smallest in the hotspot, I have carefully investigated where exactly in the furnace the hotspot was located, so as to place my samples there. In order to determine the location of the hotspot, I have calibrated the temperature along the long axis of the graphite furnace following the technique described in Watson et al. (2002). They found that when MgO and Al2O3 are in direct contact, spinel will grow according to the reaction:

MgO + Al23 O MgAl24 O (II.1) This growth is diffusion limited and therefore strongly dependent on temperature, a relation that Watson et al. (2002) empirically determined to be:

11 48865 D=X 8.58 × 10 ⋅exp  − −2.08 ⋅P ⋅ t (II.2) T in which DX is the thickness of the spinel growth layer in μm, T the temperature in Kelvin, P the pressure in GPa and t the time in seconds. The design of the assembly for this temperature calibration and its results are presented in Figure II–5. From the regression equation, it appears that for the length of capsules used in my study, the maximum temperature difference is 9 °C between the hotpot and the coldest end of the capsule. Page 24 — Chapter II

Figure II–5. Plot of temperature against distance from the top of the capsules during the experi- 1465 ment. Temperatures (°C) were calculated from the thickness 1460 of the spinel growth layer with 1455 equation (II.2). The regression equation was then used to 1450 calculate the 6-9 °C temperature difference between the coldest 1445 and hottest parts of a sample (°C) Temperature 1440 -06 2 -02 assuming a length for the sample T = -2.0 10 d + 1.48 10 d + 1431 R2 = 0.97 3.4 mm during compression 1435 (i.e. approximately 10% shorter 0 1 2 3 4 5 6 7 8 than the original length of 3.7 Distance from top of capsule mm prior to compression as during experiment (mm) presented in Figure 1).

II–1.2.2. Strategy for experiments involving metal and silicate liquids and isotope fractionation A number of specific details concerning the design of the assemblies had to be investigated. These mainly concern sample interactions with the capsule material and compositions of the starting mixtures. With respect to the former, the presence of both liquid metal and liquid silicate limits suitable materials for capsules. Any metal capsule will chemically interact with the liquid metal, while generally oxides or silicates will interact with the silicate liquid. It was furthermore important to avoid contamination of the sample with the element of interest from the capsule; e.g. with an SiO2 capsule for Si isotopes. As presented in Appendix AIII, I have tried a number of different capsule materials. Among these were crushable ZrO2 (two types with porosities of ~30% and ~10%), Al2O3 (porosity <10%), graphite from various sources, SiO2 glass and single crystal MgO. All of these capsules, however, were placed in metal jackets – usually Pt or Au50Pd50 depending on temperatures – because they can be welded. This ‘double capsule’ technique ensures that the samples are chemically isolated from its surroundings, except for hydrogen atoms that are small enough to diffuse through metal.

The experiments with ZrO2 were unsuccessful (RH01 and RH26 in Appendix AIII), because it had too high porosity, resulting finally in partial capsule collapse due to too much permeation of the sample liquids into the capsule pores (Figure

II–6). This was further enhanced by reaction of the silicate liquid with the ZrO2, a process probably partly responsible for continuous pathways even under 1 GPa pressure. Al2O3 capsules had lower porosities and seemed more robust, but they are so hard that they are difficult to produce. Although this material is generally very reactive with silicate liquids and initially led to formation of Al2O3 crystals, Chapter II — Page 25 this can be avoided with shorter run durations. Tests with simple plugs as bottom and top sealings resulted in too much leaking along the contact between plugs and tube, indicating that more complex sealing is necessary (RH29 and RH34 in Appendix AIII). My final choices fell on graphite and, depending on the isotopic system Figure II–6. Back-scatter electron image of a failed under consideration, SiO2 glass and single crystal experiment in a ZrO2 capsule. Due to the high porosity MgO. Although the pores of the capsule, the liquid sample permeated into the capsule that then collapsed towards the central part. of graphite capsules seem All metal in the sample has disappeared because it to close well under 1 GPa alloyed with the metal jacket, which forms an infinite pressure, particularly the reservoir compared to the volume of the sample. The low viscosity metal liquids zone between the sample and the former capsule-sam- were still able to permeate ple contact is a mixture between ZrO2 grains from the into the graphite capsule capsule and silicate glass from the sample. along grain boundaries. This, in fact, initially led to severe problems with capsule collapses (Figure II–7a-c), and I experimented with various sources of graphite (Appendix AIII). Although I have not made it my scientific purpose to identify the exact reason for this problem, it seems related to the permeability of the graphite. As a rule of thumb, the various tests I performed imply that the permeability increases with decreasing grain size, although there was at least one exception to this pattern among my tests. A piece of graphite with measured low gas permeability (kindly donated by Colin Maden) yielded satisfactory results. Professor Bernard J. Wood (University of Oxford), however, brought my attention to the graphite produced by Morgan Industrial Carbon, which provided the best results (Figure II–7d) and has been the material of choice for all later experiments in graphite capsules (Appendix AIII). All three capsule materials mentioned above, however, react with the metal liquid (graphite) or silicate liquid (SiO2 and MgO). Graphite dissolves in the metal, but its solubility is limited to about 5 wt% depending on pressure (P), temperature (T), composition (X), and oxygen fugacity (fO2) conditions (Bouchard and Bale, 1995). Because the presence of carbon in a metal is considered to affect elemental partitioning (Jana and Walker, 1997), I have chosen to also perform experiments in SiO2 glass and single crystal MgO capsules in addition to those Page 26 — Chapter II

Figure II–7. Back-scatter electron images of experiments with various types of graphite. Brightest reflections are metal, darkest reflections are graphite or epoxy. Experiments presented in panels (a) and (b) were performed in the same type of graphite. Its permeability was so high that particularly in panel (b) the original sample area is barely distinguishable from the capsule. The graphite used in the experiment presented in panel (c) is a large improvement, but the best results occurred from runs in the graphite used in the experiment presented in panel (d), in which the black reflections inside the grey silicate glass are epoxy, not graphite. in graphite capsules. The reactivity of SiO2 glass capsules is relatively limited compared to MgO, SiO2 contents increasing from 50 wt% in the starting mixture to 60 wt% in the run product. Reaction progress can furthermore be kept minimal when run durations are kept to a minimum. In order to use relatively low temperatures (~1300-1400 °C) to maximise isotopic fractionation required modifications to the compositions of starting mixtures compared to a binary Fe-Ni alloy and peridotite in metal-silicate equilibration experiments. Lowering the liquidus of the silicate can be achieved by increasing the silica and alkali contents in the starting mixture. For the Fe-metal, there are few elements that lower the melting point, alloy with iron in liquid form and allow for the desired range of oxygen fugacity, i.e. close to the iron- wüstite (IW) buffer. An attempt with Zn (melting temperature of 420 °C at 1 bar) failed, because too much Zn oxidised, which led to the formation of Zn-rich Chapter II — Page 27 spinel crystals. Tin (melting temperature of 232 °C) remained as best choice, and all experiments below the melting temperature of pure Fe (1540 °C) have been performed including Sn. II–1.2.3. Fe isotope specific settings For the study of Fe isotope fractionation, experiments were performed at

1250-1300 °C in graphite and SiO2 glass capsules. The use of MgO capsules was avoided, because these would lead to FeO-MgO exchange rendering an uncontrollable, open system with respect to Fe (isotopes). Starting mixtures with approximately 50 wt% metal were prepared including and excluding S, as well as with and without an isotope tracer. Sulphur-bearing compositions were run in graphite capsules, because of liquid immiscibility in the ternary Fe-C-S system (Corgne et al., 2008). Further details are provided in chapter III. II–1.2.4. Si isotope specific settings For experiments performed to study Si isotope fractionation, the main purpose was to find a composition that would render sufficient metallic Si for analyses. The relative amount of metal was therefore larger with ~75 wt% than in experiments performed for Fe isotope fractionation. Additionally, the experiments had to be more reducing and starting mixtures were therefore based on the starting and equilibrated compositions of mixture/experiment MK15 and MKX presented in Table 1 and 3 in Kilburn and Wood (1997). Two starting mixtures were prepared, because the relatively large isotopic fractionation of Si enabled run temperatures (1750 °C) above the liquidus of pure Fe (see chapter IV for more details). The ability to perform experiments at 1750 °C presented a Figure II–8. Back-scatter electron image of problem, though. Pyrex becomes an experiment that was run at 1750 °C with too weak and too reactive as an a talc/silica glass assembly in the centrifug- assembly part at such high tem- ing piston cylinder following the multi-step peratures, and therefore it is heating and re-pumping procedure. A crack as it occurred in 75% of the 1750 °C experi- recommended to use silica glass ments that followed that procedure is visible instead. This material, however, in the upper left corner of the image. The con- has a much higher softening tem- sequence of the crack is a pathway for liquid perature of ~1500 °C compared sample to reach the Pt jacket. Due to the high to ~800 °C for pyrex. I experi- temperature, this leads to Pt entering the mentally found that softening sample as can be seen from the (Pt-bearing) metal dispersed in the silicate glass. Page 28 — Chapter II starts around 1200 °C. In centrifuge experiments therefore, I initially pumped to about 0.4 GPa, heated to 1200 °C to soften the silicate glass, quenched and then re-pumped to ~1.2 GPa prior to heating to the desired run temperature. Either the higher temperature or the properties of silica glass, however, led to a 75% failure rate due to cracking of the inner capsule perpendicular to its long axis (Figure II–8). The cracking occurred with both graphite and MgO capsules, but was never observed when using pyrex instead of silica glass sleeves. Neither slow cooling instead of quenching after softening, nor pressurising to 0.25 instead of 0.4 GPa led to successful experiments. Only pumping directly to ~1 GPa and omitting the softening step has led to an acceptable failure rate of about 20%, albeit with consequently lower and less predictable final run pressures. Finally, the use of stepped furnaces (necessary for temperatures >1600 °C in the centrifuge) often leads to shearing where the step in wall thickness occurs. This can lead to temperature instabilities, and stepped furnaces are therefore also to be avoided when possible.

II–1.2.5. Mo isotope specific settings The main issue with experiments performed to investigate Mo isotope fractionation was the content of Mo in the silicate phase. Because Mo is siderophile, the vast majority (> 99%) of all bulk Mo resides in the metal phase. Silicate phases only contain traces (ppm level) of Mo, as found by analysis of several experiments with laser ablation ICP-MS following standard protocols. Because at least 60 ng Mo is required for a single isotopic analysis (see section II–3.2. below for more details), 10 mg of silicate powder must be recovered for a silicate with 6 ppm Mo. Recovering more than 10 mg, however, is difficult from a typical experiment. In addition, 60 ng is a minimum, because for statistical reasons > 3 analyses of one sample are preferable. To reach relatively high contents of Mo in the silicate, I have therefore used about 10 wt% Mo in the metal phase.

In addition, I used single crystal MgO capsules instead of SiO2 glass capsules, as the solubility of elements with valence states ≥ 3+, such as Mo, increases with depolymerisation in the silicate liquid (Ellison and Hess, 1986; Walter and Thibault, 1995; Hillgren et al., 1996). Finally, the oxygen fugacity of the experiments was kept higher than 2 log units below the iron-wüstite buffer (ΔIW-2). This was done because at more reducing conditions, Mo becomes too siderophile and even with 10 wt% Mo in the metal, Mo contents of the silicate would not reach 5 ppm. Experiments between ΔIW-1 and ΔIW-2 have been performed in capsules with an inner and outer diameter of 2.8 mm and 4.5 mm, respectively, compared to the standard sizes presented in Figure II–1. The experiments were performed at 1400 °C and 1600 °C and various starting mixtures were prepared to vary oxygen fugacity (by varying Fe2O3 contents) and compositional parameters (Sn). Further details about published experiments are provided in chapter V. Chapter II — Page 29

30 MoS2 Mo (at.%)20 10 Fe

S 1540°C δ-Fe 90 δ-Fe + L 80 1150°C 70 2 60 + L γ-Fe S (at.%) 1 50 L 40 2 1080°C α-Fe + MoS 30 γ-FeS 20 + 2 L L 10 (g) γ-FeS + L + S Fe 2 FeS (l)

+ S S δ-Fe Fe + L γ-FeS+ 2 1540°C 990°C 2 FeS (s) FeS + S S 2 (l) (g) 2

γ (s) -Fe γ-FeS α-FeS+ FeS 2 + S 90 + S β + S -FeS+ FeS 2 2

2 2 β 2

FeS + MoS 3

α-Fe + γ-FeS MoS S 80 2 1750°C

α MoS MoS L Mo L + MoS L (g) 70 1150°C α-Fe + β-FeS 2 (eutectic?) 2

α-Fe + α-FeS 60 3 L + S 3 S MoS 2 S 2 50 Fe Mo L S (at.%) L + Mo L 3 40

2 S

λ + α− + λ

-Fe

γ 30 +

1540°C + Mo -Fe + MoS

Literature data for 1 atm. Fe

µ + α− + µ + L

-Fe 20 δ

α− Mo 1550°C Mo

Experiment at 1600°C, 1 GPa µ

R 665°C Mo R

µ R

1450°C 10

µ µ

Mo R

Μο + µ + Μο

+ R

σ Mo

Μο + σ + Μο

1600°C Mo

2620°C

L 90

Mo + L + Mo

Mo 70

Fe (at.%) Fe

Mo 2620°C

Figure II–9. Phase relations in the ternary system Fe-Mo-S. Only the three binaries and one thermal section through the ternary are known. Data inside the ternary are from an experiment I performed at 1600 °C. Tie lines connect the phases I observed in the experiment, the open star represents the bulk composition. Binaries are from the Landolt- Börnstein database, the ternary section from Villars et al. (1995). Page 30 — Chapter II

In addition to the experiments presented in chapter V, I performed a number of experiments with an S-bearing metal to investigate the influence of S on Mo isotope fractionation. For the metal component, I used a ternary Fe-S-Mo system. Consequently, the experiments were performed in MgO instead of graphite capsules. These experiments are presented here, because none of them were sufficiently successful to be included in chapter V. The reason for the low success rate is the lack of information for the Fe-S-Mo system: only the three binary phase diagrams and one ternary section exist (Figure II–9). I performed several experiments at 1400 °C and 1600 °C with various compositions aimed to be near the binary Fe-S eutectic. However, metal crystals were stable at 1400 °C as well as 1600 °C in experiments with a metal starting composition by weight of ~Fe67S26Mo7 (Figure II–10a). Adjusting the metal starting composition to the S-rich metal liquid in these experiments subsequently resulted in two immiscible metal/sulphide liquids at 1400 °C (Figure II–10b). Another attempt to reach a single metal liquid by adjusting the metal component of the starting mixture to one of the two immiscible liquids again resulted in two immiscible metal liquids, albeit with different compositions (Figure II–10c). It seems therefore that the ternary Fe-S-Mo system is a complex system whose eutectics/peritectics

Figure II–10. Back-scatter electron images of experiments with metal liquids in the Fe-Mo-S compositional system. Panel (a) shows the four phases that occurred in an experiment at 1600 °C with their approximate compositions by mass. Panel (b) shows a centrifuged experiment performed at 1400 °C with a Mo-poorer starting composition. This produced two immiscible liquids. Possibly due to the presence of small amounts of Pt, the experiment shown in panel (c) has also produced two liquids, one of which is richer in Mo and poorer in S than the Mo-rich phase shown in panel (b). Chapter II — Page 31 are deeply entrenched in the liquidus surface. Compositional changes during an experimental run due to redox reactions can shift the bulk metal system into another stability field. I have summarised the results of the 1400 °C experiments in a provisional ternary diagram (Figure II–11).

S 90 80 70 2 60 + L S (at.%) 1 50 L 40 1080°C 30 γ-FeS 20 + 2 L L 10 (g) γ-FeS + L + S Fe 2 FeS (l)

+ S S δ-Fe Fe + L γ-FeS+ 2 1540°C 990°C 2 FeS (s) FeS + S S 2 (l) (g) 2

γ (s) -Fe γ-FeS α-FeS+ FeS 2 + S 90 + S β + S -FeS+ FeS 2 2

2 2 β 2

FeS + MoS 3

α-Fe + γ-FeS MoS S 80 2 1750°C

α MoS MoS L Mo L + MoS L 1150°C (g) 70 2 α-Fe + β-FeS (eutectic?) 2

α-Fe + α-FeS 60 3 L + S 3 S MoS 2 S 2 L 50 Fe Mo L S (at.%) L + Mo L 3 40

2 S

λ + α− + λ

-Fe

γ 30 +

1540°C + Mo -Fe + MoS

µ + α− + µ + L

-Fe 20

Simplified binaries for 1 atm. δ

α− Mo 1550°C Mo

Experiments at 1400°C, 1 GPa µ

R 665°C Mo R

µ R

1450°C 10

µ µ

Mo R

Μο + µ + Μο

+ R

σ Mo

Μο + σ + Μο

1600°C Mo

2620°C

L 90

Mo + L + Mo

Mo 70

Fe (at.%) Fe

Mo 2620°C

Figure II–11. Phase relations in the ternary system Fe-Mo-S with updated information at 1400 °C. Open stars indicate bulk compositions, tie lines connect the stable phases. Apparently there is only a narrow field close to the Fe-S binary that results in a completely molten system at 1400 °C. Page 32 — Chapter II

II–2. FROM A QUENCHED EXPERIMENT TOWARDS CHEMICAL DISSOLUTION

After quenching of an experiment, preparatory procedures are carried out. Most important among these are electron imaging and microprobe analyses that serve both as a quality check of the experiments (e.g. no cracks in the capsule) and as a basis for mass balance calculations after isotopic analyses. Prior to microprobe analyses, however, the capsule is removed from the assembly and embedded in epoxy. This is then ground on SiC grinding paper until the sample material is exposed. Because the silicate glass is brittle and mostly cracked, I always impregnated the sample with the low viscosity epoxy Laromin prior to final grinding to remove approximately a third of the capsule radius. The exposed material is subsequently polished with 3 or 1 μm diamond paste and coated with carbon. Coated samples were analysed with an electron microprobe to examine textures and determine elemental compositions. For details of this technique, see Reed (1993). If experiments were performed as tests, I analysed them semi- quantitatively by EDS with a scanning electron microprobe (SEM). For good signal intensities, the rule of thumb is that the acceleration voltage of the incident electron beam should be twice that of the excitation potential of the electrons that produce the X-rays that are analysed. For my samples, the K-shell of Fe had the highest excitation potential of about 7.1 keV and I therefore always used an acceleration voltage of 15 keV. To avoid diffusion of elements in the glasses due to the intensity of the incident electron beam, silicate glasses were analysed with a beam current of 6-7 nA. Metals, on the other hand, were analysed with a current of 20 nA. Beam diameters were 1-10 μm for silicates and 10 μm for metals and between 7 and 16 spots were analysed on each phase. For each session, pure metal standards were analysed, relative to which the metal samples were analysed. The more common silicate and oxide standards used for silicate analyses were only standardised when necessary, i.e. when their compositions were found to deviate by more than one per cent from the previous standard analysis. After microprobe analyses, samples were prepared for dissolution. First, I removed excessive epoxy with a diamond saw. Subsequently, I used a diamond wire saw with a diameter of 220 or 300 μm to cut the samples at the contact between metal and silicate, and at the sample-capsule contacts at the top and bottom as well as the sides of the capsule (Figure II–12). Capsule material thus remained on one side of the sample, and this I removed by grinding on ~30 μm

SiC (for Fe isotopes) or 30 μm Al2O3 grinding paper (Si and Mo isotopes). The same type of grinding paper was also used to clean each side that was cut with the diamond wire saw to remove any contamination by the sawing procedure. I used Al2O3 instead of SiC grinding paper for Si as well as Mo because Mo readily forms carbides and might therefore be present as contaminant in the SiC paper. Chapter II — Page 33

This method results in one clean piece of metal and one clean piece of silicate, each compris- ing a volume of approximately 1-2 mm3. The final preparatory step I performed consisted of crushing the metal and silicate pieces with a pestle and mortar. For samples in preparation for Fe isotope analyses, I used the agate pestle and mortar for common use in the laboratory. For Si isotopes, Figure II–12. Representation of the removal of the sample from the capsule. Areas where the I purchased two alumina sets sample/capsule is cut with the diamond wire saw to prevent cross contamina- are indicated. The remaining pieces of metal and tion: one for the low Si content silicate are then further cleaned with abrasive (~5 wt%) metallic samples paper and crushed prior to dissolution in a clean and a separate one for the high chemistry laboratory. Si content (~25 wt%) silicate samples. I expected the concentrations of Mo in the silicates for Mo isotope analyses to be below 100 ppm. Therefore, I crushed these samples in the alumina pestle and mortar that had previously only been used for the silicate glasses of which I analysed their Si isotope composition, which contained < 100 ppb Mo. All crushed material was transferred from the mortar into a glass or polypropylene vial for transfer into a clean chemistry laboratory.

Table II–2. Anion exchange protocol for Fe. 1 ml AG1-X4, 200-400 mesh. Table II–1. Material cleaning for Fe chemistry. Beakers were Purpose Acid type Volume (ml)

cleaned by filling them with Pre-clean ~2.8 M HNO3 2x2 acid, closing them and putting Equilibration 6 M HCl 2x0.5 them on hotplates at 130°C. Load 6 M HCl ~0.5 Type Beakers Rinse 6 M HCl 6x0.5 Collect 14 M HNO3 x 1 M HCl 7x0.5 Clean prior H O rinse x ~2.8 M HNO3 2x1.5 2 MQ to storage in 0.05M HCl 0.05 M HCl 0.5 Page 34 — Chapter II

II–3. ISOTOPIC ANALYSES

Isotopic analyses Table II–3. Material cleaning for Si chemistry. consist of two procedures: ion exchange Type Beakers Columns chemistry to separate at ~24°C in 6 M HCl x the element of PP-beakers H2OMQel x interest from all other H O x elements present, and at ~80°C in 2 MQ 7 M HNO x actual analysis in a glass beakers; 3 all chemicals H2OMQ rinse x multi-collector mass p.a. grade spectrometer. The H2OMQ x ion exchange proce- 6 M HCl-1 M HF (48 h) x dures are different for at ~130°C in H2OMQ rinse x every element, but teflon beakers; H O (24 h) x the actual analyses of all chemicals 2 MQel H O rinse x the elements in my p.a. grade 2 MQ study all took place H2OMQel (24 h) x on a multi-collector inductively coupled plasma mass spectrometer (MC-ICPMS). More specifically, Table II–4. Cation exchange protocol for Si. 1 ml AG50W-X8, 200-400 mesh. the MC-ICPMS was a large geometry type Nu1700 at ETH Zurich. Part of the Mo Volume Purpose Acid type isotope analyses, though, were performed (ml) on a Neptune Plus at the University of H O ~1.5 Münster, produced by Thermo Fisher 2 3 M HCl ~1.5 Scientific. Both ion exchange and MC-ICPMS 6 M HCl ~1.5 14 M HNO ~1.5 analyses are performed in liquid state and Pre-clean 3 all samples therefore have to be dissolved. 6 M HCl ~1.5 The crushed samples are weighed in 3 M HCl ~1.5 PFA beakers, which are highly resistant 6 M HCl ~1.5 to acids, and then dissolved with acids. 3 M HCl ~1.5 For metals, usually 12 M HCl with a Equilibration H2O 4x~1.5 bit of 14 M HNO3 are sufficient for dissolution. Silicates, on the other hand, Load+collect1 ~0.01 M HCl ~0.5 2 have such strong bonds that they require Collect H2O 2x1 a more aggressive combination of 24 M HF and 14 M HNO . After dissolution, 1 The exact acid molarity and volume in 3 this step are sample dependent. the samples are dried and treated further 2 The indicated 2x1 ml is used for silicates; with acids depending on the requirement for metals this step consitst of 4x1 ml. of the specific ion exchange procedures. Chapter II — Page 35

However, the dissolution procedure depends on the element of interest: the above details are valid for most elements analysed by MC-ICPMS, but they are different for Si isotopes, because Si becomes volatile below a critical Si/HF ratio. Samples for Si isotopes were therefore treated differently, as described in chapter IV. II–3.1. Ion exchange chemistry The separation of elements occurs through ion exchange, which relies on the variable distribution coefficients (or adsorption coefficients) between an acid and grains of a solid substance. The solid substance is usually an organic polymer (‘resin’) that has either cations (e.g. H+) or anions (e.g. Cl-) on exchangeable sites in its structure. If the distribution coefficient (KD) of an element or molecule - (e.g. FeCl2 ) is larger than 1, it will replace the cation or anion on its site and stick to the resin. The resin is therefore placed in a column, such that the acid can be washed through to separate elements with KD’s >> 1 from those with KD’s << 1. The KD’s are dependent on type of acid, acid molarity and type of resin. The elution of elements as a function of acid volume can further be varied in a limited range by varying the volume of resin, the grain size of the resin (usually between 40 and 150 μm), and Table II–5. Material cleaning for Mo chemistry. the aspect ratio of Pipet Columns for the column. All of Type Beakers these parameters also tips TRU-spec influence the elution 3% RBS 50 x x x flux and thereby H2OMQ rinse x x x ~12 h at the time an elution 6 M HCl x x x procedure takes. ~80°C in glass beakers; H O rinse x x x To prevent any 2 MQ all chemicals 7 M HNO x x x contamination of the p.a. grade 3 element of interest H2OMQ rinse x x x by the environment, H2OMQ x x all chemistry proce- 3:1 12 M HCl- dures are carried out 14 M HNO 3 x in an ultra-clean lab- (p.a. grade, ≥3 oratory. Furthermore, d at ~140°C)

(non-disposable) H2OMQ rinse x Beakers filled equipment is cleaned 5:1 24 M HF- and closed on 14 M HNO prior to use and each hotplate 3 x sample is treated with (1x dist., ≥6 d separate equipment at ~170°C) H O rinse x to prevent cross- 2 MQ H O (≥1 d, contamination of 2 MQel x samples. Details of ~130°C) the ion exchange and H2OMQel rinse x Page 36 — Chapter II cleaning procedures for Fe, Si and Mo are provided in Table II–1 to II–6 and in chapters III to V. Table II–6. Ion exchange protocol for Mo. mM = 0.5 mM HCl-0.5 mM HF. a. Cation exchange; 3 ml AG50W-X8 200- b. Anion exchange; 2 ml AG1-X8 200- 400 mesh 400 mesh Volume Volume Purpose Acid type Purpose Acid type (ml) (ml) 6 M HCl ~8 6 M HCl-1 M HF ~9

1 M HF ~8 H2O ~9 6 M HCl ~8 3 M HNO (2x) ~9 Pre-clean Pre-clean 3 1 M HF ~8 H2O ~9 6 M HCl ~8 1 M HCl ~9 1 M HF ~8 0.1 M HCl ~9 Equilibration 1 M HCl-0.1 M HF 2x3 Equilibration 1 M HCl 2x4 Load+collect 1 M HCl-0.1 M HF 1 Load 1 M HCl 1 Collect 1 M HCl-0.1 M HF 4 Collect 1 M HCl 1 6 M HCl ~8 Collect 1 M HCl 2x6 Clean prior to 1 M HF ~8 3 M HNO (2x) ~9 storage in mM Clean prior to 3 mM ~8 storage in mM mM ~9 c. Anion exchange; 2 ml AG1-X8 200-400 d. Cation exchange; 1 ml TRU- mesh Spec Volume Volume Purpose Acid type Purpose Acid type (ml) (ml)

6 M HCl-1 M HF ~9 H2O ~7 mM 2x2 1 M HNO ~7 Pre-clean 3 6 M HCl-1 M HF ~9 0.1 M HNO ~7 Pre-clean1 3 mM 2x2 H2O ~7

6 M HNO3-0.2 M HF ~9 1 M HCl ~7

3 M HNO3 (2x) ~9 0.1 M HCl ~7 Equilibration 1 M HF 2x3 Equilibration 1 M HCl 2x1.5 Load 1 M HF 6 Load 1 M HCl 1 Rinse 1 M HF 2x6 Rinse 1 M HCl 6 Rinse Ti/Zr/ 6 M HCl-1 M HF 2x8 Collect 0.1 M HCl 6.5 Hf/W 1 Rinse H2O 2 This procedure is repeated twice.

Collect 3 M HNO3 6.5 Clean prior to 6 M HCl-1 M HF ~9 storage in mM mM ~9 Chapter II — Page 37

II–3.2. MC-ICPMS Precise analysis of isotopic ratios relies mainly on two basic principles: ionisation of the element to be analysed and separation of its isotopes by travel through a magnetic field (Figure II–13). Once the samples have been cleared of all elements except the element of interest, the weakly acidic measurement solution is aspirated from a vial. The weak acid is then normally evaporated in a desolvat- ing unit (DSN-100) and separated from the precipitated element aerosols by a membrane. The aerosols are then transferred with a noble gas flow (usually Ar) to a plasma. That plasma is created by collisions of Ar with free electrons in an electromagnetic field generated by a high-frequency alternating electric current. Temperatures in such plasma reach up to 10000 K, under which the aerosols of sample material disintegrate into single atoms that are then ionised, i.e. they lose an electron. Aided by the directed Ar flow and the lower (vacuum) air pressure behind the plasma, the ions then travel through cones with narrow holes to let a narrow beam of ions enter an electric field with various high voltages. The ions convert that potential energy into kinetic energy and are directed into a narrow beam of ions travelling at the same velocities before they Table II–7. Cup configuration Fe isotope analyses. enter a magnetic field (Figure Cup L7 L6 L5 L4 L3 L2 L1 Ax II–13). In the magnetic field, - - 52Cr - - - 54Fe - ion travel paths are deflected as a function of their mass- Cup H1 H2 H3 H4 H5 H6 H7 H8 to-charge ratio, which is - - - - 56Fe 57Fe - - effectively a function of mass because the vast majority Table II–8. Cup configuration Si isotope analyses. of atoms will have lost only one electron. The separated Cup L7 L6 L5 L4 L3 L2 L1 Ax isotope trajectories are finally 28 29 directed into collector cups - - - Si - - Si - where the charge built up by Cup H1 H2 H3 H4 H5 H6 H7 H8 the entered ions is converted - - - - - 30Si - - into an electric current that is measured as output simultaneously on Table II–9. Cup configuration Mo isotope analyses. all cups. Cup L7 L6 L5 L4 L3 L2 L1 Ax Specific settings 90 92 94 95 96 for MC-ICPMS Nu1700 - Zr Mo - Mo Mo - Mo apply to each Neptune Plus - 91Zr 92Mo 94Mo 95Mo 96Mo element of interest, Cup H1 H2 H3 H4 H5 H6 H7 H8 e.g. the strength of Nu1700 - 97Mo - 98Mo 99Ru 100Mo - - the magnetic field Neptune Plus 97Mo 98Mo 99Ru 100Mo - Page 38 — Chapter II

Sample solution Desolvator

102 Pa 10-2 Pa 10-6 Pa 10-7 Pa ElectroStatic Lens Lens Analyser plates stack 1 stack 2 (HV) (HV) Focussing lenses 3

Rotary Turbo pumps pump Monitor plate

Cone Light isotope enriched

Heavy isotope enriched to collectors

Light isotope enriched Magnet

Quadrupole focussing lenses

Collectors

Resistors (typically 1011 Ω) Figure II–13. Schematic representation of isotopic analyses by multi-collector inductively coupled plasma mass spectrometry (MC-IPCMS). The dissolved sample is aspirated from a beaker and sprayed into a desolvator with a nebuliser. The dry aerosols are then transported into the plasma, where they disintegrate and ionise. Out of the cloud of ions, heavy isotopes pass preferentially through the series of cones compared to light isotopes, causing an instrumental mass bias (see inset). After acceleration and focussing of the ion beam, the mass separation of the isotopes occurs in the magnetic field. The various beams are finally focussed into the collector cups, in which the positively charged ions build up an electric potential and generate an electric current that is measured as output value. The ratio of two simultaneously measured electric currents from two different cups then forms the (raw) isotopic ratio. (Figure modified after Nu Instruments Ltd.) Chapter II — Page 39 has to be changed for each element depending on the masses of its isotopes. Cup configurations of the MC-ICPMS are presented for each element in Table II–7, 8 and 9. Further details are provided in chapters III to V. For each element, I used an ASX-100 (Cetac) auto-sampler and a PFA nebuliser (~140 μl min-1 uptake rate) to aspirate the samples and transfer them into a DSN-100 desolvator attached to the Nu1700. In all cases, each analysis consisted of 36 cycles of 5 s integration each. Some of the Mo analyses were performed on a Neptune Plus at the University of Münster. This MC-ICPMS had a sensitivity that was about 3 times higher than that of Nu1700 at ETH Zurich, meaning that 4 times less sample material was consumed for a single analysis (see chapter V for more details). This was required for the small amounts of Mo (<100 ng) recovered from experiments performed at oxygen fugacities

II–3.2.1. Standard-sample bracketing versus double spike Each mass spectrometer internally fractionates isotopes according to their mass. In MC-ICPMS this occurs because heavy isotopes are concentrated in the centre of the plasma relative to light isotopes (Figure II–13). The ions in the central part preferentially pass through all cones and the final measurement output is thus biased to heavier isotope ratios than the true one of a sample. The fractionation is mass dependent and can thus be described identically to equation (I.1) (Albarede et al., 2004). This instrumental fractionation must be corrected for. In studies of mass dependent fractionation, a common correction technique is to bracket a sample analysis with analyses of a standard. The assumption is then made that the standard and sample are equally fractionated by the instrument, which holds true as long as the elemental compositions of the sample solution are identical; i.e. acid type and molarity as well as concentration of the element of interest (Schoenberg and von Blanckenburg, 2005; Hughes et al., 2011). Because instrumental fractionation is fairly constant with time in MC-ICPMS, the sample composition is then simply given as the fractionation relative to the bracketing standard analyses (in per mil deviation): i X j X d ij/ X =sample −⋅1 1000 (3) i X j X standard where i and j refer to the isotopes of element X. For standard-sample-standard bracketing, the ratio of i/jX of the standard is taken as the average of the ratio determined directly before and after a sample analysis. I have followed this standard-sample-standard bracketing procedure for Fe and Si isotope analyses. A second approach to correct for instrumental fractionation is to add a double Page 40 — Chapter II

Figure II–14. Schematic representation of double spike analyses (after Rudge et al., 2009). Squares indicate that compositions are known, stars represent unknowns. Two single spikes of known composition (S and S ) are mixed to form S D.S. S 1 2 1 2 a double spike (D.S.). The D.S. is then mixed in unknown molar proportion p with a sample (Sa) p of unknown composition to form an unknown mixture (M). That mixture (M) is analysed, but the output of the analysis (m) is fractionated M β m relative to the true composition M by an instrumental mass bias that can be described with a 1-p mass dependent fractionation law with instrumental fractionation factor β. Furthermore, Sa α Std the sample composition can be calculated with the assumption that it is fractionated with the same mass dependent fractionation law (now with natural fractionation factor α) relative to a standard (Std) of known composition. spike to the measurement solution. This is possible only when the element of interest has a minimum of four isotopes, because three isotope ratios are required to unravel the instrumental fractionation from the ‘natural’ fractionation, i.e. the sample composition relative to a standard. The procedure relies on the known composition of both a double spike and a standard, the latter having a natural (i.e. unspiked) isotopic composition. With mass dependent fractionation laws that describe natural and instrumental fractionation, and with mass balance equations, it is then possible to solve iteratively for the natural mass dependent fractionation factor (α; Figure II–14). The isotopic composition of the sample in δ units is then related to α and the isotopic masses (m): m dαij/ X =−⋅1000 ln(i ) (4) m j The mathematical background for this double spike deconvolution is given in Rudge et al. (2009). It should be noted however, that this procedure can only be used for analysis of mass dependent fractionation, because of the intrinsic assumption that the natural fractionation is mass dependent. Without this assumption, the double spike deconvolution consists of a task that is not analytically solvable (see Appendix A in Rudge et al., 2009), unless another law can be written to estimate the three sample ratios (see Figure II–15 for a graphical representation). In the study I have performed, I used the double spike procedure for Mo isotope analyses, because it has additional advantages compared to standard- sample bracketing. First, for a trace element like Mo impurities of other elements that occurred in tens of weight per cent in the sample can always remain after ion Chapter II — Page 41

80 D.S.

Mo (‰) 40 m 100/95 δ 20 M

0 4 Std 3 40 δ 98/95 2 Sa 30 Mo (‰) 20 1 10 δ97/95 Mo (‰) 0 0 Figure II–15. Graphic representation of the double spike deconvolution (after Siebert et al., 2001). Three isotopic ratios are required (Mo was chosen as an example). As in Figure 14, squares (cubes) are knowns, while stars are unknowns. The exponential mass dependent fractionation laws form straight lines in this plot for variable values of the instrumental and natural fractionation factors β and α, respectively. Furthermore, mixing of a sample (Sa) and double spike (D.S.) forms a straight line in this plot for various mixing proportions. By initially estimating a natural fractionation factor, using the known standard (Std) and D.S. compositions, a flat surface can then be drawn between the D.S. composition and the line that represents possible sample compositions. The intersection of the line for possible ‘true’ (M) compositions of the analysed (m) mixture with that surface then provides a first estimate of the instrumental fractionation factor and the true composition (M) of the mixture. Extrapolating the line between D.S. and M then provides a first estimate of the natural fractionation factor and the sample composition (Sa). After a few iterations, the values for the instrumental and natural fractionation factors, and thereby M and Sa, do not change anymore and the true composition of the sample as well as the true D.S.-sample mixing proportion (p) has been determined. exchange chemistry. Such impurities, however, would make the measurement solution slightly different from the standard solution, thereby causing different instrumental fractionation for the two solutions; an effect referred to as ‘matrix effect’. Furthermore, by adding the double spike to the sample prior to dissolution, any mass dependent fractionation during ion exchange chemistry as a result of yields that are <100%, is corrected for in the double spike deconvolution. Therefore, the ion exchange procedures could be focused on purification without compromise. Page 42 — Chapter II

Figure II–16. Contoured field of internal analytical errors (1SD) 98/95 for δ Mo as a function of the 1 97 proportion of Mo in the double 0.9 0.13‰ spike (D.S.) and the proportion Mo D.S. 0.8

of D.S. in the D.S. – sample 100 mixture. Errors are modelled 0.7 Mo-

97 0.6 with the error propagation 0.10‰ model in the ‘double spike 0.5

toolbox’ described in Rudge et Mo in

97 0.4 al. (2009). The brightest area 0.3 indicates the smallest errors with the cross indicating the 0.2 optimum D.S. composition and 0.1

molar D.S. proportion in the Proportion of 0 D.S. – sample mixture. Contours 0 0.2 0.4 0.6 0.8 1 are isolines for errors in steps Proportion of D.S. in D.S. − sample mixture of 0.001‰. The smallest error is 0.10‰ 1SD, i.e. 0.032‰ 2SE for 40 scans. The compositions and calibration scheme for the Mo double spike and standard NIST SRM3134 are given in Table II–10. The double spike was mixed after dissolution of a 97Mo and 100Mo metal powder purchased from the Oak Ridge National Laboratory (ONRL). These two isotopes were chosen because (i) they have relatively low isotopic abundances (~9.6% each), (ii) there is only isobaric interference from 100Ru (isotopic abundance of 12.6%), (iii) it results in small intrinsic errors on the quoted δ98/95Mo (Rudge et al., 2009), and (iv) several other studies had already published δ98/95Mo compositions. It is not necessary to calibrate each of the two single spikes relative to a gravimetric standard before mixing them to yield the double spike, but (Christoph Burkhardt and) I have done this (Table II–10). In doing so, it turned out that the 100Mo spike was contaminated with 5.7 at% ‘natural’ Mo compared to the composition provided by the ORNL. The optimum double spike mixture was then calculated with the ‘double spike toolbox’ (Rudge et al., 2009) and the correct abundances of the two single spikes (Figure II–16). The subsequent calibration of the double spike was initially performed against an in-house gravimetric Mo standard (Alfa Aesar), but later adjusted to the NIST SRM3134 standard in an attempt to establish a common international reference standard. The Alfa Aesar Mo standard had a δ98/95Mo composition of -0.17‰ relative to NIST SRM3134. The calibration of the double spike was repeated on the Neptune Plus in Münster against the NIST SRM3134 standard. After calibration on each MC-ICPMS, the accuracy was verified with runs of five different double spike-standard mixtures varying in double spike proportion from 0.18 to 0.80 (Figure II–17). Finally, the actual internal errors were found to compare well to those predicted from intrinsic error Chapter II — Page 43

0.3 0.2 NIST SRM3134 Standard with Neptune 2SD = 0.025‰ 0.1 0 -0.1 Mo (‰)

98/95 -0.2 δ -0.3 -0.4 Alfa Aesar Mo Standard with Nu1700 2SD = 0.119‰ -0.5 0 0.2 0.4 0.6 0.8 1 Proportion of D.S. in D.S. − standard mixture Figure II–17. Mo isotope composition of reference standards in double spike (D.S.) – standard mixtures plotted against proportion of D.S. in the mixtures. Circles are for runs performed on the Neptune Plus in Münster, diamonds for runs performed on the Nu1700 in Zürich. The former runs analysed mixtures with the NIST SRM3134 standard, which is the reference standard relative to which all Mo samples are reported in this thesis. The runs on Nu1700 analysed mixtures with an in-house Alfa Aesar Mo standard, which has a δ98/95Mo composition of -0.17‰ relative to the SRM3134 standard.

0.30 0.27 14 V total beam (Nu1700) 0.24 25 V total beam (Neptune) 0.21 0.18 0.15 Μο (‰, 1SD) 0.12 98/95 0.09 in δ

Internal analytical error 0.06 0.03 0 0 0.2 0.4 0.6 0.8 1 Proportion of D.S in D.S. − sample mixture Figure II–18. Internal analytical error (1SD) for δ98/95Mo as a function of the proportion of double spike (D.S.) in the D.S. – sample mixtures. The curves are produced with the error propagation model of the ‘double spike toolbox’ (Rudge et al., 2009): the broken curve is for a total Mo beam of 14V applicable to the D.S. – standard mixtures run on Nu1700, the solid curve is for a total Mo beam of 25 V applicable to the D.S. – standard mixtures run on the Neptune Plus. Circles represent the internal errors of the runs on the Neptune, diamonds the ones on Nu1700. Page 44 — Chapter II propagation in Rudge et al. (2009) (Figure II–18). Deconvolution of all data was performed with the geometrically motivated iteration described in Siebert et al. (2001). It produced identical results to deconvolution with the ‘double spike toolbox’, but the procedure of Siebert et al. (2001) was more convenient for use with the Nu Plasma software. Table II–10. Details of the Mo double spike calibration. Abundances presented here were determined from repeated runs of pure solutions of the NIST SRM3134 and each of the spike solutions on the Neptune Plus in Münster. The NIST SRM3134 runs were internally normalised to a 97Mo/95Mo ratio of 0.602083 (Lu and Masuda, 1994). The spike runs were standard-sample-standard bracketed, where the standards were also internally normalised to determine the instrumental fractionation factor. The average instrumental fractionation factor of the bracketing standard runs was then applied to a spike run for calculation of mass bias corrected ratios and, finally, abundances. Values in brackets are 2SD uncertainties and refer to the last decimal place.

97Mo-100Mo NIST Abundance 97Mo spike 100Mo spike D.S. SRM3134 92 0.002101(±6) 0.013346(±5) 0.00906(±3) 0.147285(±27) 94 0.001861(±10) 0.006956(±5) 0.00502(±2) 0.092109(±6) 95 0.004787(±1) 0.011929(±3) 0.00921(±2) 0.158920(±3) 96 0.012390(±2) 0.013017(±6) 0.01277(±2) 0.166735(±5) 97 0.942028(±14) 0.008014(±3) 0.36345(±6) 0.095683(±2) 98 0.033763(±3) 0.024577(±4) 0.02807(±2) 0.242292(±17) 100 0.003070(±4) 0.922160(±2) 0.57243(±16) 0.096976(±18) Atomic mass (g mol-1) 96.910904(±3) 99.57391(±32) 98.5606(±7) 95.9417(±2) Concentration (μg g-1)1 147.21(±4) 85.73(±12) 0.6983 14.6 1 Concentrations of the D.S. and NIST SRM3134 determined gravimetrically. Those of the single spikes were determined by calibration, following these steps: 5 solutions with different spike/standard ratios were prepared and analysed. The spike/standard ratios of these solutions were calculted from mass bias corrected (through standard-sample-standard bracketing, as for pure spike runs) ratios, standard abundances, and previously determined spike abundances. Subsequently, spike concetrations were calculated from gravimetrically determined standard concentrations and weighing data for each of the solutions.

II–4. FAILED SILICON ISOTOPE ANALYSES BY LASER ABLATION

In addition to solution MC-ICPMS, I have also performed analyses of Si isotope compositions by laser ablation MC-ICPMS in collaboration with Caroline Fitoussi. These analyses were performed at ETH Zurich in Hönggerberg with both an Excimer (~20 nano-second pulse) and an in-house femto-second (fs) laser system. The Excimer laser was operated at ~17 J cm-2, while the fs laser was Chapter II — Page 45 operated at about 0.15 J cm-2. The fs system was mostly used, because the shorter energy pulse should improve ablation of metal, in which electronic heat dissipa- tion occurs on the scale of about 1 ns. Tests with 800 and 266 nm wavelengths for the fs-laser yielded more satisfactory results with 266 nm wavelength. With that wavelength, Si isotope compositions of silicate standards were closest to the compositions determined by solution MC-ICPMS. However, not all standards produced accurate results with this method. Our hypothesis was that this may be due to the relatively large wavelength: matrix effects in silicates generally decrease with reducing laser wavelength down to at least 193 nm (Guillong et al., 2003; Kroslakova and Gunther, 2007). A comparison was therefore made of particle size distributions produced by laser ablation with a ns-193 nm laser (Excimer) and ablation with a fs-266 nm laser. This comparison seemed to confirm that the fs-laser produced larger aerosols (Figure II–19), although the results are not unequivocal (cf. NIST610). Another method to verify whether particle sizes are a cause of inaccurate analyses consists of varying the temperature of the plasma, which can be done by varying the argon flux towards the plasma (a higher flux leading to lower plasma temperatures). Relatively high plasma temperatures lead to a more efficient ionisation of larger aerosols compared to lower plasma temperatures. If aerosols are too large for full ionisation, the residue of the aerosol will be enriched in heavy isotopes, causing the measurements to be biased towards light isotopic compositions. At low plasma temperatures, this effect should thus be stronger than at

0.20 0.20 a b 0.18 NIST610 glass 0.18 NIST610 glass 0.16 ns-193 nm Opx 0.16 fs-266 nm Opx Cpx Cpx 0.14 San Carlos olivine 0.14 San Carlos olivine 0.12 Metal (from RH59) 0.12 Metal (from RH59) 0.10 0.10 0.08 0.08 0.06 0.06 Volume fraction Volume 0.04 fraction Volume 0.04 0.02 0.02 0 0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1 Particle diameter (µm) Particle diameter (µm) Figure II–19. Size distributions of particles produced by laser ablation. The distributions are based on measurements with an optical particle counter and are given for various phases. As can be seen in panel (a), particles produced by ablation of silicates with an Excimer laser produces tight distributions at particle sizes <200 nm, while a metal produces significant amounts of larger particles up to 600 nm. Panel (b) shows that ablation with 266 nm femto-scond laser produces metal particles that have a size distribution that is more or less identical to those of silicates. Those distributions, however, show that many more larger particles (up to almost 1000 nm) are produced during ablations with this type of laser. Page 46 — Chapter II higher temperatures. If the size distribution of aerosols varies among samples and standards, the isotopic compositions of standard-sample-standard bracketed analyses should then vary as a function of plasma temperature. The results of analyses with the Excimer laser as a function of argon flux are presented in Figure II–20. The tests were performed on an orthopyroxene (opx), a clinopyroxene (cpx) and a San Carlos olivine grain. These three should all have isotopic compositions of about -0.29‰ according to solution MC-ICPMS analyses. Variations of several per mil occurred (Figure II–20) upon varying the argon flux from 0.95 to 0.5 l min-1. Furthermore, the composition of San Carlos olivine did not converge to the ‘true’ value of -0.29‰, although the opx and cpx analyses did seem to converge to accurate values below an argon flux of about 0.7 l min-1. After these tests, I decided to focus on solution MC-ICPMS for Si isotope analyses.

0.5 0 -0.5 -1.0 Composition of all three silicates determined by -1.5 solution analyses

Si (‰) -2.0 30 δ -2.5 -3.0 orthopyroxene clinopyroxene -3.5 San Carlos olivine -4.0 0.4 0.5 0.6 0.7 0.8 0.9 1 Ar flux towards plasma (l min-1)

Figure II–20. Standard-sample-standard bracketed Si isotope compositions of three silicates plotted against Ar flux towards the plasma. The argon flux is a proxy for the temperature of the plasma, lower Ar fluxes indicating higher plasma temperatures. All analyses were performed with an Excimer laser. The negative correlation between Si isotope composition and Ar flux suggests that not all aerosols that enter the plasma are fully evaporated and ionised under all conditions; i.e. at too high Ar fluxes (too low plasma temperatures), particles are partially evaporated leaving behind residues enriched in heavy isotopes that are not analysed. Unfortunately, though, the San Carlos olivine did not appear to stabilise within the analytical uncertainty of the composition as determined by solution analyses. CHAPTER III.

EXPERIMENTAL EVIDENCE FOR THE ABSENCE OF IRON ISOTOPE FRACTIONATION BETWEEN METAL AND SILICATE LIQUIDS AT 1 GPA AND 1250-1300 °C AND ITS COSMO- CHEMICAL CONSEQUENCES1

1 This chapter has been published as: R. C. Hin, M. W. Schmidt, B. Bourdon, 2012. Experimental evidence for the absence of iron isotope fractionation between metal and silicate liquids at 1 GPa and 1250-1300 °C and its cosmochemical consequences. Geochimica et Cosmochimica Acta 93, 164-181. Page 48 — Chapter III

ABSTRACT

Iron isotope fractionation during metal-silicate differentiation has been proposed as a cause for differences in iron isotope compositions of chondrites, iron meteorites and the bulk silicate Earth. Stable isotope fractionation, however, rapidly decreases with increasing temperature. We have thus performed liquid metal – liquid silicate equilibration experiments at 1250-1300 °C and 1 GPa to address whether Fe isotope fractionation is resolvable at the lowest possible temperatures for magmatic metal-silicate differentiation. A centrifuging piston cylinder apparatus enabled quantitative metal-silicate segregation. Elemental tin or sulphur was used in the synthetic metal-oxide mixtures to lower the melting temperature of the metal. The analyses demonstrate that 8 of the 10 experimental systems equilibrated in a closed isotopic system, as was assessed by varying run durations and starting Fe isotope compositions. Statistically significant iron isotope fractionation between quenched metals and silicates was absent in 9 of the 10 experiments and all 10 experiments yield an average metal-silicate fractionation factor of 0.01±0.04‰, independent of whether graphite or silica glass capsules were used. This implies that Fe isotopes do not fractionate during low pressure metal-silicate segregation under magmatic conditions. In large bodies such as the Earth, fractionation between metal and high pressure (>20 GPa) silicate phases may still be a possible process for equilibrium fractionation during metal-silicate differentiation. However, the 0.07±0.02‰ heavier composition of bulk magmatic iron meteorites relative to the average of bulk ordinary/ carbonaceous chondrites cannot result from equilibrium Fe isotope fractionation during core segregation. The up to 0.5‰ lighter sulphide than metal fraction in iron meteorites and in one ordinary chondrite can only be explained by fractionation during subsolidus processes.

III–1. INTRODUCTION

The segregation of a metallic core from a silicate mantle is the major physical and chemical differentiation process within terrestrial planets. Yet there is still debate about the conditions under which core formation took place on Earth (Wade and Wood, 2005; Rudge et al., 2010; Rubie et al., 2011) and what conditions are necessary to allow core segregation within planetary embryos. Because Fe and other siderophile elements have different oxidation states and bonding environments in metal compared to silicate liquids (or minerals), it has been proposed that there is potential for mass dependent stable isotope fractionation between metal and silicate phases (Poitrasson et al., 2005; Weyer et al., 2005; Schoenberg and von Blanckenburg, 2006; Williams et al., 2006). Temperatures well exceeding 1000 °C necessary for core segregation were long considered too high to record any detectable isotope fractionation, but with the progress Chapter III — Page 49 of mass spectrometry, non-traditional stable isotope fractionations (e.g. Fe, Mg, Si, Cr) have been measured in samples that are thought to have undergone only high-temperature processes (Zhu et al., 2002; Williams et al., 2004; Poitrasson and Freydier, 2005; Georg et al., 2007; Weyer and Ionov, 2007; Wiechert and Halliday, 2007; Fitoussi et al., 2009; Ziegler et al., 2010; Moynier et al., 2011). Mass dependent stable isotope fractionation may thus provide a new tool to decipher core formation conditions and its chemical consequences. For this purpose, a full knowledge of isotopic fractionation factors during core segregation processes is required. Although the theory of mass dependent stable isotope fractionation is quite well understood (Bigeleisen and Mayer, 1947; Urey, 1947; Schauble, 2004), the direction and especially the magnitude of such fractionations during geological processes remain poorly constrained. Theoretical predictions are hitherto limited to crystalline materials and can only be performed for elements and systems whose vibrational energies can be determined. Furthermore, the predictions of Fe isotope fractionation factors involving magnetite, for instance, have varied by a factor of two despite that they have been calculated by similar approaches (Polyakov and Mineev, 2000; Polyakov et al., 2007). In addition to theoretical predictions, therefore, experimental determinations of mass dependent stable isotope fractionation factors for minerals and in particular for melts are required to quantify isotope fractionation processes. The iron isotope system has probably been the most studied non-traditional stable isotope system, because of the high abundance and variable oxidation states of Fe in the Solar System. Several studies have found variations in Fe isotope compositions between various meteorite groups and terrestrial rocks, and between various phases in these samples (Poitrasson et al., 2004; Weyer et al., 2005; Schoenberg and von Blanckenburg, 2006; Williams et al., 2006; Craddock and Dauphas, 2011). In particular, differences between chondrites, iron meteorites, and Bulk Silicate Earth (BSE) estimates have led several of these studies to hypothesize that isotopic fractionation between metallic and ferrous Fe during differentiation of planetary objects into a metallic core and silicate mantle may explain the observed differences. Theoretical estimates of Fe isotope fractionation between metallic Fe and olivine, however, suggest that equilibrium fractionation between these phases at temperatures above 1000 °C should not be detectable (Polyakov and Mineev, 2000). As such theoretical estimates are limited to crystalline phases, it remains unclear whether they are directly applicable to liquid-liquid or liquid-mineral interactions that could take place during core formation processes. Poitrasson et al. (2009) experimentally determined the Fe isotope fractionation between liquid metal and liquid silicate and concluded that no detectable fractionation occurred at temperatures of 1750 and 2000 °C, and pressures up to 7.7 GPa. Carbon was present in most of their experiments. A few weight per cent of this light element may affect the elemental partitioning between metal and silicate Page 50 — Chapter III phases (Jana and Walker, 1997), but its effect on isotopic fractionation has not yet been investigated. Furthermore, temperatures of 1750-2000 °C might be relatively high for metal-silicate segregation in planetary embryos or planetesimals. In this context, the observed differences in Fe isotope compositions, particularly between chondrites and iron meteorites, may therefore still be explained by Fe isotope fractionation during metal-silicate segregation. To constrain whether metal-silicate segregation under any magmatic condition can lead to detectable variations in Fe isotope compositions, we have determined the equilibrium Fe isotope fractionation factor between liquid metal and liquid silicate at 1300 °C and 1 GPa, making use of a centrifuging piston cylinder for improved segregation of metal and silicate liquids. The results complement the work of Poitrasson et al. (2009), and are discussed in light of core segregation on the iron meteorite parent bodies and Earth.

III–2. METHODS III–2.1. Experimental methods Optimal conditions to obtain and measure isotope fractionation between a silicate and metal melt are achieved at the lowest possible experimental temperatures and by a perfect experimental segregation of the silicate and metal melts. To prevent mixed signatures of metal and silicate phases during isotopic analyses, we have used the centrifuging piston cylinder at ETH Zurich (Schmidt et al., 2006) allowing more effective segregation of the metal and silicate liquids compared to emulsions generally obtained from static piston cylinder experiments. Furthermore, metal and silicate melt compositions were designed to work at the lowest possible temperature where a depolymerized silicate melt would coexist with a Fe-dominated metallic melt. For the silicate melt we have thus used a non-primitive basaltic composition to be run in graphite capsules and a quartz saturated andesite composition to be run both in graphite and silica capsules. For the metal melt, we have lowered the melting point of pure Fe by admixing ~25 wt% Sn (Okamoto, 1993). Experiments were performed between 0.5 and 1.3 GPa and 1250-1300 °C with three different synthetic starting mixtures (Table III–1). Mix1 is a sulphur- free, Fe-Sn-Mo metal mixture homogenised with an oxide mixture of basaltic composition. Mix2a is an S-free, Fe-Sn-Mo metal mixture homogenised with an 54 andesitic oxide mixture that was enriched with a Fe2O3 isotope tracer (98.37 at% 54Fe, 1.55 at% 56Fe, 0.07 at% 57Fe, 0.008 at% 58Fe, Oak Ridge National Laboratory Batch 154891). Mix2b consists of the same andesitic, 54Fe-enriched oxide mixture as Mix2a homogenised with a 70:30 mixture of elemental Fe and S by weight, which is close to the eutectic composition at 1 GPa. An S-bearing mixture was created because theoretical calculations indicate that the sign of Fe Chapter III — Page 51

isotope fractionation between silicate and Table III–1. Compositions (wt%) troilite should be opposite compared to frac- of start mixtures. The parameters tionation between S-free metal and silicate fM and fSil are the weight fraction of metal and silicate, respectively, in (Polyakov and Mineev, 2000; Polyakov et the bulk starting mix. al., 2007). Isotopic fractionation of relatively heavy Start Start Start elements is thought to be mainly driven by Mix1 Mix2a Mix2b differences in oxidation state and bonding environment (Schauble, 2004). The Metal amplitude of the fractionation factor is then Fe 73.7 73.7 70.1 mainly controlled by temperature. As the Sn 24.6 24.7 expected Fe fractionation factors are small, Mo 1.72 1.60 we optimised our system such that the lowest possible temperatures for magmatic S 29.9 processes could be used without compro- Silicate mising on oxidation state and bonding environment. Hence, the silicate as well as SiO 49.8 59.9 59.9 2 metal compositions are not exactly those 16.7 12.1 12.1 Al2O3 of differentiating planets. However, small MgO 7.63 6.13 6.13 to negligible effects are to be expected by Fe O 9.87 8.95 8.95 ~11 mol% Sn in the metal and by using an 2 3 andesitic instead of an ultramafic silicate CaO 10.6 8.97 8.97 melt. In both cases the oxidation state and 3.59 2.98 2.98 Na2O bonding environment remains identical to 1.83 1.00 1.00 K2O pure Fe-melts and ultrabasic primordial f 0.5155 0.5949 0.5696 melts and any effect of relatively small M compositional variations would be expected f 0.4845 0.4051 0.4304 Sil to be much less than the generally dominant effect of temperature. We have not introduced C or S in the first series of experiments, as these chemical components may lead to structurally different bonding environments for Fe, in particular in the metal phase. As detailed below, we have varied C- and S-contents as independent variables through direct comparison of C-free systems with C-bearing systems, as well as S-free with S-bearing systems. 54 Mixtures 2a and 2b were spiked with a Fe2O3 isotope tracer to induce a detectable difference in the isotopic ratios of the starting Fe metal and oxide. This approach provides additional confidence to the assessment of isotopic equilibration compared to a time series, because the induced difference needs to adjust to equilibrium conditions. Moreover, the use of an isotope tracer leads to a mass independent offset in the isotopic compositions of the Fe metal and oxide, and, as explained by Shahar et al. (2008), this mass independent difference has to be erased through isotopic homogenisation irrespective of whether any mass Page 52 — Chapter III dependent fractionation occurs. The experiments with Mix1 were performed as a time series to determine the run duration required for isotopic equilibrium. These experiments were carried out in graphite capsules with 3.5 mm outer diameter, 2.0 mm inner diameter and 3.7 mm inner length. Carbon saturation, however, has been found to possibly affect elemental distribution between liquid metal and silicate (Jana and Walker, 1997), yet its effect on isotopic fractionation remains unknown. Thus, experiments with Mix2a were run both in graphite and silica glass capsules (of identical dimensions as above) to investigate potential effects of carbon saturation on Fe isotope fractionation. For the experiment with the S-bearing Mix2b, a silica glass capsule was used instead of a graphite capsule. This was done because C-rich and S-rich metallic liquids are immiscible in the Fe-C-S system at pressures below 5 GPa, while the C-free Fe-S system is miscible in the compositional range of our experiment (Corgne et al., 2008). All experiments were performed in talc/pyrex assemblies with inner MgO parts and mullite thermocouple sleeves. Graphite and silica glass capsules were welded into tin can shaped Pt or Au50Pd50 jackets (4.0 mm outer diameter, ~10.5 mm outer length). To prevent oxidation or melting of the metallic Sn powder before the experiments, starting mixtures were not dried directly prior to welding of the metallic jackets. Karl-Fischer Titration analyses on two test experiments suggest that up to ~1 wt% of H2O may be present in the bulk experimental products. Two experiments (RH21, RH44) were run in a static, Boyd and England type (Boyd and England, 1960) end-loaded piston cylinder, all other experiments in the single stage centrifuging piston cylinder. Centrifugation at 600g should lead to a sink velocity of minimum 1 mm h-1 for metal droplets of 5 μm diameter in our compositional systems, meaning such droplets should take about 2.5 h to sink into the main metal pool in an experiment centrifuged at 600g. For an imaginary fractionation factor of 0.10‰ between metal and silicate, and for the chemical compositions of our metal and silicate phases, no less than 10 vol% metal would be required as contamination in the analysed silicate phase to shift the silicate composition to within 0.04‰ (i.e. within analytical uncertainty) of the metal composition. The centrifugation ensures that the silicate glass is sufficiently clean to prevent such contamination. The centrifuging as well as static piston cylinders have a 14 mm diameter bore and are run with identical talc-pyrex assemblies. The friction correction for this assembly was calibrated at the quartz-coesite transition (Bose and Ganguly, 1995) to 10%. In the centrifuging piston cylinder, oil pressure of the hydraulic ram cannot be adjusted during centrifuging and it is advisable to only pass heating current through the slip rings when centrifuging. Through a time-consuming multi-step heating/oil pressure adjustment procedure, final run pressure can be well controlled. However, we have skipped this time-consuming procedure, Chapter III — Page 53 because pressure variations of ±0.4 GPa have been found to be insignificant for Fe isotope fractionation by theoretical calculations (Polyakov, 2009). Minimal temperature gradients are important for experimental investigation of isotopic fractionation, because thermal diffusion can enrich the colder end of a sample in heavy isotopes, fractionating for instance Fe isotopes in basaltic liquids by 1.1‰/100 °C/atomic mass unit (Richter et al., 2009). A spinel crystallization experiment following Watson et al. (2002) was performed inside a Pt capsule of identical dimensions to those used in this study to quantify the temperature gradient inside a sample. The results suggest that by centring the sample in the hotspot the temperature difference between the infinitesimally small hot centre and the cold ends of the sample is 6-9 °C in our experiments (see Appendix A). The difference in average temperatures between metal and silicate, however, are much smaller because the metal and silicate phases do not comprise infinitesimally small volumes at the centre and the cold ends of the capsule. They rather comprise approximately 30 and 70 vol% of the sample area, respectively. A side effect of centring the relatively large capsules in the hotspot is that the thermocouple tip, separated from the sample by the graphite/silica and Pt capsule lids and a corundum disc, is at approximately 50 °C less than the temperature in the centre of the sample (see also Appendix A). We have therefore applied a +50 °C correction to the temperatures recorded by the thermocouples and temperature indications in this manuscript refer to corrected ‘sample temperatures’. Experimental runs were quenched by switching off power, and the recovered capsules were embedded and polished for microprobe analyses. After microprobe analysis, samples were cut at the metal-silicate contact with a 220 μm diameter diamond wire saw. The same saw was then used to remove the sample from the capsule material by cutting each phase on three sides. The capsule material on the remaining side was removed by grinding and polishing with SiC paper. To ensure the purity of each sample, the outer 100-300 μm rims were also removed by this procedure. By doing so, for each experiment remained a clean, single piece of silicate glass and a clean, single piece of quenched metal (each 1-2 mm3). These pieces were gently crushed under ethanol prior to Fe isotope chemistry procedures. III–2.2. Analytical methods Microprobe analyses were carried out with a Jeol JXA 8200 equipped with five spectrometers. Glasses were analysed with a current of 5-6 nA and a beam diameter of 1 or 10 μm, while metals were analysed with a 20 nA current and a 10 μm beam diameter. Pure Fe, Sn and Mo metals were used as standards for the metals, while natural and synthetic minerals and oxides were used as standards for the silicate glasses, with exception of SnO2 that was also standardised to the metallic Sn standard for silicate analyses. Carbon contents in metals were calculated by difference to 100% and amount to 1.1 to 3.1 wt%. Typically 7-10 spots Page 54 — Chapter III were analysed in each phase. Iron isotope chemistry and analyses followed the procedure of Kiczka et al. (2010) and Mikutta et al. (2009), respectively. Samples were weighed in 7 ml

PFA beakers pre-cleaned with pro analysi 14M HNO3 on a hotplate (130 °C) for >48 hours. Reagents were prepared by diluting distilled 14.3M HNO3 and double distilled 12M HCl with ultrapure H2O (≥18.3MΩcm). Silicates were first dissolved in 2:2:1 mixtures of 14M HNO3, 12M HCl, and 23M HF, respectively, while metals were dissolved in 1:1 mixtures of 14M HNO3 and 12M HCl. After this first dissolution step, the solutions were evaporated and the residue was taken up in 14M nitric acid twice prior to re-dissolution in 1 ml 6M HCl. These solutions were again evaporated and the residue re-dissolved in 6M HCl for loading approximately 100 μg of Fe in circa 0.5 ml onto the columns. The AG1-X4 anion exchange resin in the columns was pre-cleaned with 2 steps of 2 ml ~2.8M HNO3 and conditioned with 2 steps of 0.5 ml 6M HCl prior to sample loading. The column was first washed with 6 steps of 0.5 ml 6M HCl, and the Fe was eluted with 7 steps of 0.5 ml 1M HCl. Column calibrations showed that the presence of Sn did not affect this elution scheme as no Sn was observed in the Fe fractions. Approximately 40-60% of Sn was recovered in 3 cleaning steps of 0.5 ml ~2.8M HNO3, thus we additionally cleaned the columns with 2 steps of 1.5 ml ~2.8M HNO3 prior to storage of the columns. No adjustments were made to the analytical protocol for samples containing 2- sulphur. Pribil et al. (2010) found that S-rich samples (SO4 /Cu >50) required an additional clean-up step in their chemistry procedure for Cu, which is very similar to the procedure for Fe. However, S/Fe weight ratios in the samples processed in this study do not exceed 0.5. Furthermore, aliquots of IRMM-014 and our 2- 2- in-house ETH Fe salt solutions have been doped with SO4 at SO4 /Fe weight ratios of ~0.7 and processed through the same chemical protocol. The resulting isotopic compositions of these standards were within uncertainties equal to the undoped solutions, which suggests that the presence of S at such low S/Fe ratios does not require modifications in the chemical procedure. The same conclusion was reached previously by Williams et al. (2006). The Fe was analysed as 100 to 150 μg L-1 solutions in 0.05M HCl on a multi- collector inductively coupled plasma mass spectrometer (MC-ICPMS; Nu1700, Nu Instruments Ltd, Wrexham, UK). Samples were aspirated from 2 ml polypropylene vials with a PFA nebuliser with ~140 μL min-1 aspiration rate in an ASX-100 auto-sampler (Cetac), and transferred to a DSN 100 desolvation system. The large geometry Nu1700 MC-ICPMS can operate at a mass resolution of >2500 (m/Δm), which allows to fully resolve argide (ArN+, ArO+, ArOH+, ArC+) and chloride (37Cl16OH) interferences on Fe peaks. 52Cr was monitored for interference of 54Cr on 54Fe, but intensity ratios of 54Cr/54Fe were always < 5×10-5. Samples were bracketed with the international reference standard IRMM-014 and typically 3 to 6 standard-sample-standard brackets (SSB) were analysed for Chapter III — Page 55 each sample. One analysis consisted of a single block of 36 cycles with an integration time of 5 s per cycle. All beams were collected in Faraday cups with 1011 Ω resistors and the intensity on mass 56 was ~6×10-11 A. Iron isotope compositions are reported as the per mil deviations of 56Fe/54Fe ratios of samples from the 56Fe/54Fe ratio of the international reference standard IRMM-014: 56Fe 54 Fe d 56Fe =sample −⋅1 1000 56 Fe (III.1) 54Fe IRMM −014 Fractionation factors (Δ) are the difference in the δ56Fe composition between two phases A and B such that a negative A - B fractionation factor means that phase B has a heavier Fe isotope composition than phase A: 56 56 56 D=−FeAB− dd FeA FeB (III.2) Repeated analyses of the in-house ETH Fe salt standard over the 22 month course of this study resulted in an average δ56Fe value of -0.71‰ (n=123) with 2SD long-term reproducibility of 0.10‰ calculated from single standard-sample-standard brackets as: 2 ()xx− 22SD = ⋅ ∑ i (III.3) n −1

Where xi stands for a single measurement on a single sample, x for the average of several measurements on a single sample and n for the number of measurements on a single sample. These values are identical to those reported by Kiczka et al. (2010): -0.71±0.10‰ (n=216). Test runs with Ni-doping in single scans and in peak-jump mode did not lead to any improvement of the precision. Experiments with higher signal intensities by switching the collector for mass 56 to a 20 V range improved the reproducibility of the δ57Fe values by approximately a factor of 2, but did not significantly improve the reproducibility of our δ56Fe values. In addition, these changes in beam intensities did not lead to shifts away from the mass dependent fractionation line in δ57Fe- δ56Fe plots. Therefore, no adjustments were made to the original run protocols described above and in Mikutta et al. (2009). In this manuscript, errors are presented as the 95% confidence interval (95% c.i.) of the number (n) of repeated standard-sample-standard brackets: − 2 ∑()xxi (III.4) n −1 95% ci. . = ⋅t0.95 n in which t0.95 refers to a two-tailed 95% confidence test. Its values were obtained from Miller and Miller (2010). Page 56 — Chapter III

III–3. RESULTS

Melt compositions in terms of major elements are presented in Table III–2. Run conditions and isotopic compositions of the silicate and metal melts as well as starting materials are given in Table III–3. III–3.1. Textures and elemental compositions of the experimental products Centrifugation forced the dense metal liquids to the outer bottom side of the capsules (Figure III–1). In most experiments, surface tension prevented small drops (usually less than 4 in a polished section, forming less than 2 vol.% of the liquids) of either metal or silicate liquid to be released from the capsule walls and they therefore did not collect in the metal or silicate liquid pools, respectively (Figure III–1a). These drops, however, were removed upon cutting the sample prior to chemical separation procedures. Only in the S-bearing experiment RH73 centrifugation was not sufficient for an almost full segregation of the two liquids. However, sufficient volume of metal-free silicate glass and silicate-free metal quench was collected for isotopic analysis. In the time series experiments, segregation did not succeed fully (~95 vol.% of the metal liquid collected in the bottom pool) because the temperature of 1250 °C did not exceed the liquidus of the silicate phase. In these experiments, a thin layer of <20 μm sized olivine crystals formed in the silicate melt just above the silicate-metal melt interface. The crystals prevented full segregation of the two liquids (Figure III–1b) and thus retained a few small droplets of metal melt. This interface layer was always <200 μm wide and was thus directly removed when the metal-silicate interface was sawed with the diamond wire saw. In addition, occasional unmixing of liquid Sn-rich and liquid Fe-rich melts resulted in the formation of a rim of pure Sn around the metal phase (Figure III–1c). Although ternary Fe-Sn-C phase diagrams do not exist for appropriate temperature ranges, binary Fe-Sn and Fe-C phase diagrams do not have two-liquid immiscibility gaps at the pressures, temperatures and compositions of our experiments (Okamoto, 1993). Furthermore, centrifugation did not lead to segregation of the two metal alloys, although this would be expected due to the density difference between pure Sn and Fe-Sn alloys. We thus interpret this unmixing to occur upon quench and because of the Fe-free nature of the occasional Sn rims we ignored this feature. In all other experiments, the run temperature exceeded the silicate liquidus and the silicate liquid quenched into a glass free of micrometre sized metal blobs, indicating that centrifuging completely segregated the metal and silicate (Figure III–1d; compare also the silicate glass details in Figure III–1b, c and e). The metal liquids always showed quench textures with usually <10 μm wide dendritic phases (Figure III–1f) with variable Fe/Sn or Fe/S ratios. Such quench Chapter III — Page 57

Figure III–1. Back-scatter electron images of selected experiments. Panel a shows an experiment of the time series performed with Mix1 in a graphite capsule. The temperature of 1250 °C in these experiments did not exceed the liquidus. Therefore a thin layer of small (<20 μm) olivine crystals accumulated during the experiment at the contact between metal and silicate liquids. Panel b shows a detail of the metal-silicate contact area, while panel c highlights the exsolved Sn rim around the metal liquid in the same experiment. Panel d shows an experiment in a silica glass capsule at 1300 °C. This temperature exceeded the silicate liquidus, leading to quantitative separation of the metal and silicate liquids into two pools. A detail of the contact area between the silicate glass capsule wall and silicate melt shows the disintegration of the capsule material (panel e). Panel f shows a detail of typical metal quench textures with relatively high Fe/Sn (dark reflections) and relatively low Fe/Sn (light reflections). Page 58 — Chapter III textures are characteristic in metals and we have not attempted to analyse the quench features but used a 10 μm beam diameter to obtain average compositions (Table III–2). Experiments performed with Mix1 have resulted in FeO contents in the basaltic glass of ~17.0 wt%, up from 8.88 wt% in the starting oxide mixture.

Furthermore, about 1.7 wt% SnO2 is dissolved in the silicate melt. Consequently, all other oxide concentrations in the silicate melt decreased slightly with respect to the starting material. The use of graphite capsules has led to the dissolution of ~3 wt% carbon in the metal liquids, calculated from the difference of the total to 100 wt% (Table III–2). Experiments with Mix2a in graphite capsules led to ~ 14 wt% FeO in the andesite glass (8.05 wt% in the starting material) and oxidation of Sn to form about 1.6 wt% SnO2. In silica glass capsules, these increases were smaller and the andesitic melt contained 10.6 wt% FeO and ~0.8 wt% SnO2. The increase of FeO contents in the andesitic melt to 11.0 wt% in S-bearing Mix2b was similar to that found in experiments with Mix2a in silica glass capsules. An advantage of the use of graphite and silica glass capsules is that diffusive exchange of metallic Fe or Fe2+ with those capsule materials is minimal compared to other materials. Microprobe analyses within 80 μm of the capsule- sample interface show that after 7-8 hours run duration no more than 0.5 wt% FeO diffused into the capsule walls. We furthermore consider this percentage a maximum value, because small scale percolation of silicate or metal melt into the disintegrating capsule wall (see Figure III–1e) can lead to excitation of electrons in the Fe-richer sample during spot analyses near the capsule-sample interface with mixed analysis of capsule wall and sample material as a consequence. Oxygen fugacities relative to the iron-wüstite buffer (ΔIW) were calculated from X ⋅γ D=⋅IW2 log FeO FeO (III.5) X Fe⋅γ Fe

where Xi is the molar fraction and γi the activity coefficient of species i. Following Holzheid et al. (1997), γFeO was taken to be 1.7. Activity coefficients for metallic Fe were calculated with the Wagner epsilon approach as modified by Ma (2001). Interaction parameters were taken from the Steelmaking Data Sour- cebook (1988). Calculated ΔIW varied between -0.47 and -0.90 for experiments in graphite capsules and between -1.44 and -1.96 in silica glass capsules (Table III–2), in accordance with a lesser oxidation of Fe and Sn in the latter. The higher oxygen fugacity conditions in graphite capsules compared to silica glass capsules can be related to the presence of C. The general increase of FeO contents with respect to the starting material in all experiments may be caused by dissociation of H2O, which was present because the starting material could not be dried at high temperature prior to welding of the metallic jackets. The redox dissociation of H2O liberates O2 that oxidises metallic Fe to form FeO. The presence of C can affect the dissociation reaction by the formation of methane, which leads to CH4 Chapter III — Page 59

Table III–2. Compositions (wt%) of the run products of selected experiments. Errors (2SD) refer to the reproducibility of usually 7-10 different analysis spots. Experimental parameters and capsule types are also given: P = pressure, T = temperature, t = run duration, fO2 = oxygen fugacity relative to the iron-wüstite buffer (see text for details); n.d. = not determined.

Start Mix1 Start Mix2a Start Mix2b RH30 2SD RH43 2SD RH48 2SD RH57 2SD P (GPa)1 1.03 0.05 0.79 0.01 1.49 0.06 0.91 0.12 T (°C) 1250 1300 1300 1300 t (h) 3.0 3.9 8.1 6.0 -0.59 -0.86 -1.46 -1.96 fO2 (ΔIW)

Capsule type Graphite Graphite SiO2 glass SiO2 glass

Metal Si 0.04 0.02 0.02 0.02 0.04 0.02 n.d. n.d. Fe 69.9 1.1 74.3 4.2 76.2 1.3 66.3 5.1 Sn 25.7 1.3 21.1 4.0 22.0 1.5 n.d. n.d. Mo 1.62 0.08 1.81 0.31 1.81 0.09 n.d. n.d. S n.d. n.d. n.d. n.d. n.d. n.d. 33.1 6.4 C-calc2 2.7 2.7 Total 97.3 97.3 100.1 99.4

Silicate 44.3 2.1 55.0 1.1 63.0 1.2 60.6 1.9 SiO2 14.2 1.3 10.5 1.4 8.82 0.44 10.3 0.6 Al2O3 MgO 7.17 1.66 5.67 0.92 4.40 0.39 4.59 0.26 FeO 16.8 2.1 13.9 1.5 10.6 1.0 10.9 1.0 CaO 9.39 0.11 8.35 0.41 6.68 0.41 7.15 0.19 3.04 0.47 2.70 0.31 1.83 0.19 1.88 0.12 Na2O 1.44 0.42 0.90 0.17 0.64 0.17 0.68 0.12 K2O 1.80 0.15 1.77 0.21 0.79 0.30 n.d. n.d. SnO2 Total 98.1 0.9 98.8 0.8 96.7 0.6 96.1 1.9

1 Errors refer to the time dependent variation of the pressure in centrifuge experiments. They are calculated as the 2SD of the pressure registrations in 30 s intervals in our log files (see text for more explanation). 2 Carbon contents are calculated as the difference of the totals from 100%. Page 60 — Chapter III solubility in liquid basalt of ~100 ppm at the oxygen fugacities and pressures of our experiments (Ardia et al., 2011). Approximately 0.7 wt% H2O would have to dissociate to explain the observed increases in FeO and SnO2 contents in experiments with Mix2a. Raman spectroscopic analyses of a test experiment with Mix2a revealed the presence of CH4 in the silicate glass. However, H2 was not observed which indicates that it must have been lost from the system, most likely by diffusion through the Pt jacket. III–3.2. Isotopic composition of the experimental liquids Two aliquots of the metallic Fe powder used in all starting mixtures have 56 been determined to have a δ Fe of 0.47±0.04‰ (Table III–3). The Fe2O3 powder used in Mix1 has a δ56Fe of 0.57±0.03‰, yielding a calculated bulk δ56Fe of 0.48±0.04‰ that is identical within error to the measured value of 0.45±0.05‰. The Fe isotope compositions of the corresponding metal and silicate phases of the run product of the time series experiments carried out with Mix1 are within uncertainty identical to each other, except for the 16 hour run of sample RH21 (Figure III–2a). The metal fraction of that sample is 0.07±0.05‰ heavier than its silicate fraction. The δ56Fe values of all metal and silicate phases vary from 0.31±0.01‰ to 0.53±0.03‰. All compositions plot within error of the terrestrial mass dependent fractionation line in a plot of δ57Fe versus δ56Fe (Figure III–2b). 56 57 54 The δ Fe and δ Fe of the Fe2O3 powder spiked with the Fe2O3 tracer are -3.44±0.09‰ and -3.16±0.14‰, respectively. Mix2a thus had a calculated bulk δ56Fe of 0.23±0.04‰ which compares well to the measured value of 0.31±0.09‰. Despite the difference in starting Fe isotope composition between metallic Fe and Fe-oxide, the run products of all experiments performed with Mix2a in both graphite and SiO2 glass capsules are identical within uncertainties (Figure III–2c). The compositions cluster between δ56Fe values of 0.15±0.03‰ and 0.32±0.03‰. These data would not plot on the terrestrial mass dependent fractionation line due to the addition of 54Fe as isotope tracer. A secondary mass dependent fractionation line with a slope identical to the terrestrial one (i.e. 1.475 in Figure III–2d; (Young et al., 2002) has therefore been anchored to the measured bulk δ56Fe and δ57Fe values of the spiked starting powder Mix2a. Except for the natural isotopic composition of the Fe powder used in the starting mixtures, all data plot on this secondary mass dependent fractionation line (Figure III–2d). Sulphur bearing Mix2b was mixed from the metal powder and the spiked Fe-oxide, identical to Mix2a. It had a calculated bulk δ56Fe of 0.20±0.04‰ and resulted in a measured δ56Fe value of 0.15±0.04‰, which is within error of the calculated bulk composition. One experiment with this starting material resulted in an Fe isotope composition of 0.12±0.10‰ and 0.19±0.06‰ for the metal and silicate phase, respectively (Figure III–2c). Both values plot within uncertainty of the secondary mass dependent fractionation line mentioned above (Figure III–2d). Chapter III — Page 61

0.7 a Mix1 1.2 b Mix1 0.6 Fe2O3 1.0

Fe 0.8 0.5 Mix1Calc

Mix1Meas

Fe (‰) 0.6 56 Fe (‰) δ

0.4 57 δ 0.4 1250°C 0.3 1.0 GPa Metal 0.2 Silicate Bulk (calculated) 0.2 0 0 2 4 6 8 10 12 14 16 0.2 0.3 0.4 0.5 0.6 0.7 Time (hours) δ56Fe (‰)

0.6 c 1300°C Mix2a/b 1.0 GPa 0.5 Fe Metal Silicate Bulk (calculated) 1.2 d 0.4 Mix2a/b Mix2a Meas 1.0 0.3

Mix2aCalc 0.8 Unspiked 0.2 Fe Mix2b Fe (‰) Calc

56 0.6 δ 0.1 Mix2bMeas Fe (‰) 57 δ

Graphite Graphite SiO2 SiO2 0.4 capsule capsule capsule capsule

0.2 -3.4 SiO2 capsule Fe2O3 -3.5 0 0 0.1 0.2 0.3 0.4 0.5 0.6 0 2 4 6 8 δ56Fe (‰) Time (hours)

Figure III–2. Iron isotope compositions plotted against run duration of the experiments (a, c), and three-isotope plots (b, d). Panel a shows the Fe isotope compositions of experiments performed with Mix1 (all in graphite capsules). Circles and diamonds represent the compositions of silicate and metal phases, respectively, whereas squares represent the bulk compositions of the experiments, calculated from the silicate and metal phases (see text for calculation details). Note that the positions of the calculated bulk compositions correspond to experimental run durations, while the silicate and metal compositions are shifted for readability. The compositions of the metallic Fe and Fe-oxide used in the starting mix are given as upward and downward pointing triangles, respectively. The solid line and shade indicate the measured bulk composition of Mix1 and its uncertainty, respectively. The calculated bulk composition of Mix1 is shown as a broken line, its propagated uncertainties by the stippled lines. Panel c shows the Fe isotope compositions for Page 62 — Chapter III experiments performed with the spiked starting materials (Mix2a and Mix2b; symbols as in panel a). The experiment performed with S-bearing Mix2b was run for a duration of 6 hours. The measured and calculated compositions of that mixture, as well as their (propagated) uncertainties, are plotted from 5 to 7 hours to distinguish them from Mix2a, with which the four other experiments were performed. With exception of the 16 h run duration experiment in panel a, all metal and silicate pairs presented in panels a and c are within error of each other. The bulk compositions of 8 of the 10 experiments are within error of both measured and calculated starting mixture compositions. All errors are 95% confidence intervals. Panels b and d correspond to the analyses presented in panels a and c, respectively. Their symbols are also identical, with exception of squares: in panels b and d they represent the measured compositions of the starting mixtures. The solid lines indicate the theoretical equilibrium mass dependent fractionation lines (Young et al., 2002). As expected, the natural isotopic composition of the Fe metal powder used in the start mixes plots off the secondary mass dependent fractionation line in panel d. All other analyses are within uncertainty (95% c.i.) of their mass dependent fractionation lines.

III–4. ATTAINMENT OF EQUILIBRIUM

To determine an equilibrium isotope fractionation factor, one should assess whether equilibrium has been reached in the experiments. Poitrasson et al. (2009) found that run durations of ~100 s were sufficient to reach Fe isotope equilibrium in their experiments at 2000 °C. They concluded that this is in agreement with elemental diffusion data for Fe2+ of Guillot and Sator (2007), which result in diffusion distances similar to the dimensions of their capsules within 100 s. The experimental temperatures of 1250-1300 °C in our experiments are much lower and we thus performed a time series to determine the run duration required for isotopic equilibrium. Using diffusion data from Guillot and Sator (2007), Fe diffusion in silicate melts over the ~3.5 mm length of our samples should take about 5 hours, while diffusion in Fe metal should be more than an order of magnitude faster following diffusion coefficients of Dobson (2002). Neverthe- less, the results of nearly all of our time series experiments (0.5, 1.6, 3.0 and 12.0 hours) did not yield any variation in the Fe isotope fractionation factor, suggesting that isotopic equilibrium was reached within 30 minutes instead of 5 hours. It remains unclear why the experiment with longest run duration (16 h) has a metal composition that is slightly heavier (0.07±0.05‰) than the silicate it equilibrated with. These time series experiments were performed with Mix1, in which the difference in the Fe isotope composition between the metal and silicate melts was barely resolvable (i.e. 0.1‰). To verify that run durations of 30 minutes indeed suffice to achieve Fe isotope equilibrium at 1300 °C, run durations of experiments performed with Mix2a were also varied. This starting powder had a difference of 3.91±0.10‰ between the metallic Fe and Fe-oxide. After 25 minutes, the shortest of our run durations, the difference in isotopic composition between the metal and silicate phase was no longer resolvable. Chapter III — Page 63

2 (‰) 95% c.i.

M-Sil Fe (‰) 56 Δ 4 4 8 4 3 4 6 6 n 16 12 123

= fractionation factor between metal = fractionation M-Sil 0.25 0.14 0.15 0.14 0.14 0.10 0.13 0.14 0.11 0.13 0.09 0.13 0.28 0.06 (‰) 95% c.i. Fe 56

Fe 57 0.13 0.09 0.66 0.90 0.65 0.68 0.59 0.43 0.35 0.39 0.62 0.71 -1.06 -3.16 -2.98 (‰) δ

0.10 0.03 0.04 0.04 0.03 0.05 0.04 0.09 0.09 0.04 0.04 0.04 0.03 0.04 (‰) 95% c.i.

Fe 56 0.09 0.05 0.47 0.57 0.45 0.48 0.31 0.23 0.15 0.20 0.53 0.46 -0.71 -3.44 -3.35 (‰) δ

type Capsule

(g) RCF

t (h)

2 O f (ΔIW)

T (°C) 1

P (GPa) = oxygen fugacity to relative the buffer iron-wüstite (see text for details), t = run duration, and RCF = Metal Silicate 2 5 3 4 6 - spiked - calc. - spiked, 3 3 3 O O O 2 2 2 in-house Fe Salt Fe in-house BHVO-2 BCR-2 powders Starting Fe Fe Mix1 Start calc. Mix1 Start Fe Fe Mix2a Start calc. Mix2a Start Mix2b Start calc. Mix2b Start Samples (Mix1) Natural RH21 Sample Experimental conditions and isotopic compositions of starting material and run products. Experimental parameters as well Experimental conditions and isotopic compositions of starting material run products. Experimental Table III–3. Table = for metal and silicate values. P bulk values instead of in rows for calculated given in rows as fractionation factors are metal-silicate fO = temperature, T pressure, Δ (g); constant gravitation terrestrial the of multiple a as given Force, Centrifugal Relative and silicate; n.d. = not determined. Page 64 — Chapter III

2 0.05 0.08 0.02 0.06 0.10 0.11 0.06 (‰) 95% c.i.

M-Sil Fe 0.07 0.00 0.02 0.06 0.08 0.03 -0.01 (‰) 56 Δ 4 4 6 6 6 6 3 3 3 4 8 8 n

0.24 0.08 0.15 0.07 0.19 0.10 0.16 0.16 0.15 0.14 0.09 0.14 0.08 0.12 0.06 0.11 0.14 0.11 0.13 (‰) 95% c.i.

Fe 57 0.63 0.59 0.62 0.59 0.47 0.40 0.46 0.47 0.53 0.48 0.69 0.50 0.66 0.37 0.36 0.37 0.41 0.41 0.41 (‰) δ

0.03 0.05 0.06 0.04 0.02 0.01 0.02 0.03 0.05 0.03 0.02 0.10 0.02 0.11 0.03 0.10 0.05 0.04 0.05 (‰) 95% c.i.

Fe 56 0.52 0.42 0.42 0.42 0.33 0.31 0.33 0.35 0.36 0.35 0.45 0.39 0.44 0.23 0.15 0.22 0.24 0.21 0.24 (‰) δ

C C C C C C C type Capsule

- (g) 606 703 517 579 752 402 RCF

t 0.5 3.0 1.6 0.4 3.9 (h) 12.0 16.0

2 O f n.d. n.d. -0.47 -0.59 -0.90 -0.72 -0.86 (ΔIW)

T 1250 1250 1250 1250 1250 1300 1300 (°C) 1

P 1.00 (GPa) 0.88±0.07 0.93±0.02 0.93±0.04 0.49±0.17 0.89±0.05 0.71±0.01 Bulk calc. Metal Silicate Bulk calc. Metal Silicate Bulk calc. Metal Silicate Bulk calc. Metal Silicate Bulk calc. Metal Silicate Bulk calc. Metal Silicate Bulk calc. RH23 RH28 RH30 RH33 (Mix2a) Spiked RH42 RH43 Sample Table III–3, continued. Table Chapter III — Page 65 , respec - 3 O Errors for 2 2 0.06 0.06 0.12 (‰) 95% c.i.

to tracer Fe M-Sil 3 Fe O -0.05 -0.03 -0.07 (‰) 2 56

Δ 7 8 7 4 4 8 n      Fe 56 ⋅ Sil Fe Sil f 0.06 0.10 0.07 0.07 0.07 0.10 0.18 0.09 0.11 (‰) () 95% c.i.

Fe Fe 56 57 Sil 0.60 0.47 0.47 0.52 0.47 0.27 0.41 0.28 0.46 (‰) δ Fe () 56 Sil ci 95% .95% . 0.03 0.05 0.04 0.05 0.04 0.10 0.06 0.09 0.05 ) values are errors based on the 2SD uncertainties of EPMA ) values are errors based on the 2SD uncertainties of EPMA (‰)     Fe 95% c.i. f

Fe Sil Fe 56 0.32 0.27 0.24 0.27 0.24 0.12 0.19 0.13 0.27 Sil (‰) Fe δ () er f    2 2 2 SiO SiO SiO type 22 Capsule     

Fe 56

- dd (g) 290 ⋅++⋅ 1230 RCF M Fe M f () t 8.1 6.0 4.0 22 (h) Fe 56

M 2 dd Fe ()

f O 56 -1.46 -1.96 -1.44 M

(ΔIW) dd ci Fe -spiked are calculated from the weight mixing ratio 3652:1 of natural Fe 56 3 Sil 95% .95% . T O     2 1300 1300 1300 (°C) 22 ci 1 M Fe M Fe P ff () 1.00 + (GPa) 22 er f 1.34±0.05 0.91±0.12    Fe 56 M      dd ()() Fe values for =+⋅ 57 ci Metal Silicate Bulk calc. Silicate Bulk calc. Silicate Bulk calc. Metal Metal 95% .95% . 6 Fe and δ = 56 ci ci RH44 Sample RH48 S (Mix2b) incl. Spiked, RH73 95% .95% . . 95% . . 95% .

Errors are 2SD and refer to the time dependent variation of the pressure in centrifuge experiments. They are calculated as the 2SD of pressure Errors are 2SD and refer to the time dependent variation of pressure in centrifuge experiments. Propagated according to the following equation: The uncertainties given for this standard are 2SD. Calc. refers to calculation of a bulk isotopic composition from the analysed metal and silicate compositions by mass balance. The calculated δ Experiments RH21 and RH44 were run in a static piston cylinder press, not the centrifuging press. these calculated values are propagated according to the following equation, in which er( tively. Errors were not estimated here. tively. Table III–3, continued. Table 1 registrations in 30 s intervals our log files (see text for more explanation). 2 3 4 5 6 analyses: Page 66 — Chapter III

The fact that equilibrium is reached much faster than expected from calculated diffusion distances could be explained by the grain size and homogenisation of the starting mixtures. Upon reaching the liquidus of both metal and silicate phases, initial droplets of a few tens of μm in size require diffusion timescales two orders of magnitude shorter than is the case for the entire dimensions of the capsules, a feature that furthermore implies that isotopic equilibrium was reached before mechanical segregation of the metal and silicate melts. Neverthe- less, the glasses of experiments with run durations of ≤1.5 h are chemically not as homogeneous as those with longer run durations (e.g. FeO decreases from ~20 wt% close to the metal-silicate contact to ~17.5 wt% farther from the contact in sample RH28; see Table III–A.1 at the end of this chapter). This suggests that Fe isotope equilibrium may be reached faster than elemental equilibrium. Such decoupling of elemental and isotopic diffusion has been experimentally observed for both major and trace elements (Lesher, 1990). Van der Laan et al. (1994) explained the decoupling of isotopic and elemental diffusion by pointing out that concentration gradients may not equal activity gradients and that in systems with a large concentration but small activity gradient, isotopes diffuse faster than elements. Given the FeO concentration gradient near the metal-silicate contact, it may have been that our experiments have represented such systems. Care should be taken to draw such conclusions, however, because the redox contrast at the metal-silicate contact can affect a simple diffusion process. Furthermore, we have performed bulk isotopic analyses, while elemental analyses were performed in-situ. This means that a small isotopic gradient may have been present that did not affect our bulk analyses beyond analytical uncertainties. The constancy of the Fe isotope fractionation factor between liquid metal and silicate suggests that equilibrium has been reached in our experiments. True equilibrium, however, is easiest to assess if the system is closed. To evaluate the latter, the bulk Fe isotope compositions of each starting mixture have been measured (Table III–3) and the bulk compositions of each experimental run have been calculated from the measured metal and silicate compositions by mass balance: 56 M 56 Sil 56 dFeBulk=⋅ f Fe dd FeM +⋅ f Fe FeSil (III.6)

where fFe is the mass fraction of elemental Fe determined from the metal/ silicate mass mixing ratio and the concentration of Fe in the metal and silicate phase. The super- and subscripts M and Sil stand for the metal and silicate phase, respectively. As a result, the measured and the reconstructed bulk compositions of all mixtures are within uncertainties identical in 8 of our 10 experiments

(Figure III–2). Moreover, the spiked Fe2O3 of Mix2a and Mix2b was obtained 54 from an extreme mass mixing ratio of “natural” Fe2O3 to tracer Fe2O3 of 3652:1, yet its measured and calculated Fe isotope compositions are also identical within uncertainties. Only the bulk compositions of our experiments RH28 and RH30 (0.5 and 3.0 Chapter III — Page 67 h run durations, respectively) are not within uncertainty identical to their bulk starting composition. This could possibly be related to a small layer of olivine crystals and metal droplets at the contact between the main metal and silicate pools (Figure III–1a-c). This layer was removed and not analysed. To account for the deviation in bulk Fe composition, the ~200 μm slice with 30% olivine crystals, would have to have an Fe isotope composition of ~4.5‰. Assuming compositions of the metal and silicate melts in this slice identical to those of the analysed pieces, the olivine in the slice would have to be >14‰ heavier than the two other phases, an extreme fractionation that appears highly unlikely. An alternative explanation could lie in diffusional losses of Fe to the capsule wall, but this should cause an effect opposite to the observed one, as such kinetic fractionation should cause the loss of light Fe isotopes by their faster diffusion (Richter et al., 1999). Finally, Fe that migrated into the graphite capsule walls could have resulted in the formation of carbides (Lord et al., 2009), and the loss of heavy isotopes in experiments RH28 and RH30 could potentially be explained by equilibrium fractionation with iron carbide. Nevertheless, in this case all graphite experiments should have shown isotopic open system behaviour, which is not observed. Interestingly, the Fe isotope fractionation between the metal and silicate phases in these two experiments is smaller than uncertainties and excluding their fractionation factors from the average of all 10 experiments leads to 0.01±0.05‰ compared to 0.01±0.04‰ previously. Generally, we could envisage a scenario where the rate of equilibration between the phases of interest (two liquids in this case) exceeds the rate of loss of isotopes to the capsule, enabling equilibrium fractionation despite very minor Fe isotope losses. In summary, although two experiments have shown unexplainable, very minor open system behaviour, 8 of our 10 experiments show no loss of Fe isotopes during the experiments. As these 8 experiments fulfil the criteria of a closed system for Fe isotopes, and as all their run durations were sufficiently long, we conclude that Fe isotope equilibrium has been attained in all those experiments.

III–5. DISCUSSION III–5.1. Comparison to previous studies This study demonstrates that equilibrium fractionation of Fe isotopes between liquid metal and liquid silicate is smaller than present analytical uncertainties. The average of all 10 fractionation factors determined in this study is 0.01±0.04‰ (95% confidence interval). This result is not in conflict with the equilibrium fractionation factor of 0.4±0.3‰ (2σ, recalculated to 56Fe) at 1500°C found in Roskosz et al. (2006). Their study was specifically designed to investigate kinetic fractionation of Fe isotopes by extracting Fe from a silicate liquid by means of a Pt wire and Roskosz et al. (2006) underline that their value for equilib- Page 68 — Chapter III rium fractionation should be considered a first approximation. The experimental work by Poitrasson et al. (2009), designed to study equilibrium fractionation, showed an absence of equilibrium Fe isotope fractionation up to a pressure of 7.7 GPa. The lowest temperature at which Poitrasson et al. (2009) investigated Fe isotope fractionation between liquid metal and liquid silicate was 1750 °C. The data presented here demonstrate that even at temperatures of 1250-1300 °C, the lowest possible temperatures for magmatic processes involving metal and silicate liquids in planetary bodies, such fractionation is absent. Polyakov and Mineev (2000) calculated Fe isotope fractionation factors as a function of temperature on the basis of Mössbauer Doppler shift measurements on 57Fe in various phases. Such spectroscopic measurements are limited to crystalline phases. Their Fe-metal - olivine couple represents the closest analogue to liquid metal - liquid silicate systems. Polyakov and Mineev (2000) predicted α-Fe to be 0.038‰ heavier (re-calculated to 56Fe values) than olivine at 1250 °C. Another potentially relevant modelled fractionation factor can be found in Polyakov et al. (2007), who predicted α-Fe to be heavier by 0.008‰ than wüstite at 1250°C. Polyakov et al. (2007) also calculated fractionation factors for troilite that could be taken as an analogue for the S-bearing Fe-metal melt of our study. The predicted troilite - olivine fractionation factor at 1250 °C is -0.037‰ (i.e. troilite should be 0.037‰ lighter than olivine), which is within the analytical uncertainties in this study. Instead, troilite should be 0.069‰ lighter than wüstite (Polyakov et al., 2007), which is outside our analytical uncertainties. The latter modelled difference is thus the only theoretically predicted fractionation factor (Polyakov and Mineev, 2000; Polyakov et al., 2007) that is not within error of our experimental results. Shahar et al. (2008) experimentally determined a magnetite - fayalite Fe isotope fractionation factor of 0.28±0.04‰ at 800 °C that agrees well with the modelled fractionation factor of 0.23‰ (fayalite and magnetite data respectively from (Polyakov and Mineev, 2000; Polyakov et al., 2007). Schuessler et al. (2007) found that their experimentally determined troilite - silicate melt fractionation factor of -0.35±0.04‰ at 840 to 1000 °C agrees with a theoretically modelled fractionation factor for a solid-solid analogue consisting of a mixture of ferrous and ferric iron bearing minerals yielding a Fe2+/ΣFe ratio of 0.38±0.02. Although some theoretically predicted fractionation factors have varied significantly (c.f. magnetite in Polyakov and Mineev, 2000; Polyakov et al., 2007) the combined results of Schuessler et al. (2007), Poitrasson et al. (2009), and of our study strongly suggest that the theoretical approach of Polyakov et al. (2007) yields realistic fractionation factors for Fe isotopes and that these factors can be reasonably well extended to liquid-liquid systems. Furthermore, differences of a factor of two in theoretically modelled fractionation factors as for magnetite may often not be analytically resolvable for high temperature (>1000 °C) processes because the fractionation factors are so much smaller than current analytical Chapter III — Page 69 uncertainties. III–5.2. Implications for Fe isotope variability in natural systems The absence of equilibrium Fe isotope fractionation between liquid metal and liquid silicate at temperatures of 1250-1300 °C has implications for the interpretation of Fe isotope compositions measured over the last decade. Chon- drites have a significantly lighter Fe isotope composition than iron meteorites and terrestrial ultramafic and mafic rocks (Poitrasson et al., 2005; Weyer et al., 2005; Schoenberg and von Blanckenburg, 2006; Williams et al., 2006). These studies have hypothesized among other possibilities that core segregation was the process fractionating Fe isotopes between the metallic core and the silicate mantle of planetary bodies. III–5.2.1. Meteorites A compilation of all available stable Fe isotope data on meteorites (Figure III–3; Table III–A.2) shows that the bulk iron meteorite data tend to have almost 0.1‰ heavier compositions than bulk chondrite data do. To test whether Fe isotope differences among meteorites are statistically significant, we employed Monte Carlo t-test simulations. Monte Carlo t-test simulations assign random variables to the data of a population, forming a normal distribution defined by the analytical uncertainty, and determine the probability of a statistically significant difference between two population means. Such simulations (10 000 runs with a 1SD analytical uncertainty of 0.05‰) indicate that the probability is 2% that the mean Fe isotope compositions of ordinary and carbonaceous chondrites are statistically distinct at the 95% confidence level, yet there is a 79% probability that enstatite chondrites are distinct from the ordinary and carbonaceous chondrites. Furthermore, Monte Carlo t-tests comparing the population of ordinary/ carbonaceous chondrites to the bulk Fe isotope compositions of magmatic iron meteorites resulted in a 100% probability that these two populations are statistically distinct. The average bulk Fe isotope composition of the magmatic iron meteorites is +0.04±0.01‰, while the average ordinary/carbonaceous chondrite composition is -0.03±0.01‰ (errors are 95% confidence interval). This suggests a difference between bulk magmatic iron meteorites and ordinary/carbonaceous chondrites of 0.07±0.02‰ (propagated 95% confidence intervals), which is larger than the analytical uncertainty of 0.04‰ in this study. Furthermore, despite the fact that metal-silicate differentiation events in planetesimals may have taken place at temperatures as low as ~1300 °C, it is unlikely that the average core segregation temperature was that low (Righter and Drake, 1996; Yoshino et al., 2003; Wood et al., 2006; Ricard et al., 2009). More specifically, bodies that have segregated a metal core within a few million years after Solar System formation, such as the parent bodies of iron meteorites (Markowski et al., 2006), are likely to have Page 70 — Chapter III

EL Bulk Enstatite Chondrites EH

C2 CV Carbonaceous Bulk CO Chondrites CM CB CI Sulphide LL LL Metal H L

LL Ordinary Bulk Chondrites L

IVA Sulphide IIIAB IIAB IVB IVA IIIAB IIIF Metal IIE Magmatic IID Iron Meteorites IIAB IC Unclassified IVB IVA Bulk IIIAB IID IIAB IC Sulphide IIICD IAB Non-Magmatic IIICD Metal IAB Iron Meteorites IIICD Bulk IAB

-0.6 -0.4 -0.2 0 0.2 0.4 δ56Fe (‰)

Figure III–3. Literature compilation of iron isotope compositions of chondrites and iron meteorites. Symbol colours are alternated for different sub-groups of meteorites, e.g. H-, L-, and LL-group ordinary chondrites. The names of these sub-groups are also given (for references and names of individual samples, see Table EA2). Circles refer to bulk compositions, diamonds to metal compositions, and squares to sulphide compositions. Error bars on individual data points are 2SD. The two lines with shades show the average bulk composition including 95% confidence interval of ordinary/carbonaceous chondrites (-0.03±0.01‰), and of magmatic iron meteorites (+0.04±0.01‰). Chapter III — Page 71 reached temperatures of core formation >1300 °C due to radiogenic heating by short-lived 26Al (Ghosh et al., 2006). Such a temperature of 1300 °C should hence be considered as a minimum temperature. The lack of detectable equilibrium Fe isotope fractionation at those temperatures therefore indicates that the observed difference of 0.07±0.02‰ between iron meteorites and ordinary/ carbonaceous chondrites cannot originate from equilibrium fractionation of iron isotopes during metal-silicate segregation in the iron meteorite parent bodies. If the observed Fe fractionation would be derived during magmatic core segregation processes, kinetic fractionation between metal and silicate would thus remain as the only possibility that could yield the observed Fe fractionation between iron meteorites and ordinary/carbonaceous chondrites. In addition to the difference between bulk ordinary/carbonaceous chondrites and bulk magmatic iron meteorites, Figure III–3 shows that the sulphides (troilite) in iron meteorites as well as in chondrites have significantly lighter Fe isotope compositions than the metal fractions. There is considerable spread in the compositions of the sulphides, but they can be up to ~0.5‰ lighter than the metal fractions. Williams et al. (2006) and Williams and Archer (2011) found that this difference was largest in the iron meteorites they analysed and argued that most likely ~0.5‰ is the equilibrium Fe isotope fractionation factor between metal and sulphide minerals at the closure temperature of ~500 °C for Fe diffusion (Dean and Goldstein, 1986), while the observed variability could reflect incomplete isotopic equilibrium during cooling of the iron meteorite (parent) bodies. The work of Polyakov et al. (2007) followed by experimental studies of Poi- trasson et al. (2009) and our study show that equilibrium Fe isotope fractionation between metal and silicate as well as between metal and troilite at temperatures >1300 °C is smaller than analytical uncertainties of 0.04‰. These results agree with the hypothesis of Williams et al. (2006) and Williams and Archer (2011) that differences of up to ~0.5‰ between metal and troilite most likely originate from subsolidus processes. Such a conclusion is further substantiated by data from the Parnallee chondrite (Needham et al., 2009) (see Figure III–3), which is characterised by a striking similarity between the Fe isotope composition of its sulphides with that of iron meteorites. However, the equilibrium fractionation factor between metal and troilite reaches 0.5‰ only at 350 °C (Polyakov et al., 2007), which is 150 °C below the closure temperature of Fe diffusion in metal. This offset cannot be explained by analytical uncertainties of 0.1‰ and therefore argues against equilibrium fractionation of Fe isotopes between metal and troilite in iron meteorites being the single process leading to the observed fractionation. Williams et al. (2006) argued that solid state kinetic fractionation can only occur when coupled diffusion takes place, i.e. when Fe diffusing from metal to sulphide is replaced by another element that is transported in the opposite direction. They then calculated that δ62Ni/58Ni values >40‰ should result from kinetic fractionation of Fe isotopes during diffusion from metal to sulphide. Page 72 — Chapter III

Because such values have not been observed, Williams et al. (2006) ruled out kinetic fractionation as an alternative to (partial) equilibrium fractionation to explain the differences in Fe isotope compositions between troilite and metal. However, it should be considered that such calculations may include too many simplifications, as phase relations also suggest that Ni concentrations in troilite decrease from 0.5 to 0.2 wt% upon cooling from 900 to 450 °C (Buchwald, 1977). In a state of equilibrium, this would lead to Fe-Ni exchange and thereby complicate calculations for kinetic fractionation, especially when considering temperature fluctuations in meteorites (and other rocks) due to thermal events. Needham et al. (2009) found that unequilibrated metamorphic Type 3 ordinary chondrites are significantly heavier than metamorphic Type 4-6 ordinary chondrites, an observation that indicates that thermal/metamorphic events can affect Fe isotope compositions. We therefore conclude that the majority of Fe isotope variations observed in meteorites most likely originates from non-magmatic metamorphic processes, but that it remains at present open whether those processes resulted in kinetic or (partial) equilibrium fractionation of Fe isotopes, or both. III–5.2.2. Earth Estimates of the Fe isotope composition of the Bulk Silicate Earth (BSE) vary from 0.09±0.05‰ (Beard et al., 2003) to 0.04±0.03‰ (Dauphas et al., 2010) yielding a significantly heavier Fe isotope composition than the ordinary/carbonaceous chondrite average of -0.03±0.01‰ (95% confidence level; Table III–A.2, and Figure III–3). Experimental studies indicate that the Fe-Ni metal segregating into the Earth’s core has probably last equilibrated at temperatures of 2500-3000 °C and at pressures of 25-50 GPa (Righter, 2003; Wood et al., 2006; Rubie et al., 2011). Our study, combined with the results of Poitrasson et al. (2009), demonstrates that terrestrial core formation cannot have produced detectable equilibrium Fe isotope fractionation at pressures ≤7.7 GPa. Furthermore, the heavier composition of BSE relative to chondrites contrasts with predictions from theoretical models that metal phases should preferentially incorporate heavy Fe isotopes compared to Fe2+-bearing silicate phases such as olivine (Polyakov and Mineev, 2000). Because of the uptake of Fe3+ in perovskite (McCammon, 1997), equilibration of metal phases with the high pressure silicate phase perovskite has been speculated to change both the magnitude and the direction of Fe isotope fractionation (Williams et al., 2006). Theoretical predictions by Polyakov (2009) suggest 0.062±0.012‰ lighter metal than Fe3+-free post-perovskite at 130 GPa and 4000 K. Polyakov concluded that this fractionation can explain the difference in Fe isotope composition between chondrites and estimates of the BSE. However, this implies that post-perovskite chemically interacted with the rest of the BSE to homogenise its Fe isotope composition. The following mass balance calculations, however, suggest that simple mixing of a post-perovskite bearing layer and the rest of the BSE increases the composition of the BSE by less than 0.005‰ compared to the Bulk Earth composition: Chapter III — Page 73

D56 =Fe D +d56 +−Fe dd56 −56 (III.7) FeBSE−− BE f ppv( ppv M FeBE )()1 fppv FeBE FeBE

where Δppv-M is the equilibrium fractionation factor between post-perovskite Fe and metal (0.062‰), f ppv the mass fraction of Fe accommodated in a layer of 56 post-perovskite at the base of Earth’s mantle, and δ FeBE the Fe isotope composition of the Bulk Earth taken to be identical to the average of carbonaceous/ ordinary chondrites. In this calculation, it is assumed that no fractionation takes place between metal and silicate at the relevant temperatures. Core segregation occurs while the Earth growths by accretion, but the post-perovskite layer only develops at the very end of this stage when pressures at the base of the mantle reach the post-perovskite stability field. It is thus assumed that the post-perovskite layer equilibrates with a core with an Fe isotope composition identical to the Bulk Earth. The thickness of the post-perovskite layer has been assumed to be 250 km based on the low pressure limits of the post-perovskite stability field (Tateno et al., 2009) and the mass of the post-perovskite layer has been estimated from density data derived from the Preliminary Reference Earth Model using a core-mantle boundary depth of 2891 km beneath Earth’s surface. Finally, Fe concentrations of the post-perovskite layer were taken to be identical to the BSE values (6.26 wt%), and all Fe has been assumed to reside in post-perovskite. A recent experimental investigation by Williams et al. (2012) indicates that Fe3+-bearing perovskite could be minimum 0.43‰ (95% confidence, recalculated to 56Fe) heavier than metal at 1850°C, which would resolve the issue of a heavier instead of lighter BSE Fe isotope composition relative to chondrites. These results seem to contrast the theoretical predictions of Polyakov (2009), who predicted Fe2+-bearing perovskite to be about 0.03‰ lighter than metal at lower mantle pressures and at temperatures of ~1850°C. The cause for the discrepancy between the results of Williams et al. (2012) and Polyakov (2009) could arise from the comparison of ferric and ferrous iron bearing perovskites between the two studies. As Earth is the largest body of the terrestrial planetary objects in the Solar System, it is expected to have experienced the highest temperatures of metal-silicate segregation (Righter and Drake, 1996; Righter, 2003), leading to the smallest stable isotope fractionation among planetary objects. It has remained puzzling that the BSE and lunar mantle Fe isotope compositions are the only known silicate portions of a planetary object that seem to deviate from the ordinary/carbonaceous chondrite Fe isotope composition (Poitrasson et al., 2004; Schoenberg and von Blanckenburg, 2006). The data of Williams et al. (2012) indicate that the expected small isotopic fractionation at Earth’s high temperatures of metal-silicate segregation associated with its large mass may be counterbalanced for Fe isotope fractionation by phase changes that accompany the high pressures in Earth. Earth would then remain the only terrestrial body in the Solar System that may have recorded equilibrium fractionation of Fe isotopes Page 74 — Chapter III during core segregation. Alternatively, other processes that are not necessarily high temperature magmatic planetary differentiation processes may be required to explain the relatively heavy BSE Fe isotope composition.

III–6. CONCLUSIONS

At temperatures above 1250 °C and pressures of 1 GPa, equilibrium mass dependent fractionation of Fe isotopes between liquid metal and liquid silicate is not detectable at current analytical precisions. The implication of this result is that any variation in Fe isotope compositions of natural rocks cannot be explained by equilibrium fractionation during magmatic metal-silicate differentiation events. Therefore, either processes unrelated to magmatic metal-silicate differentiation, or kinetic fractionation between metal and silicate liquids may potentially explain the observations that Fe isotope compositions of iron meteorites are 0.07±0.02‰ heavier. Mass balance calculations imply that the relatively heavy composition of the BSE cannot be explained by fractionation between metal and post-perovskite. However, Earth may be the only terrestrial body of the Solar System in which metal-silicate segregation has resulted in equilibrium Fe isotope fractionation, which would be a consequence of the presence of the ferric iron bearing high pressure phase perovskite (Williams et al., 2012). The up to 0.5‰ lighter Fe isotope compositions of sulphides relative to metal in iron meteorites cannot be explained by equilibrium fractionation between these two phases because closing temperatures of Fe diffusion are too high to reach such fractionation of 0.5‰. The differences in metal and sulphide compositions are thus probably due to (kinetic) fractionation at subsolidus temperatures during cooling and/or during thermal events.

Acknowledgements. Felix Oberli is greatly thanked for his education in the mass spectrometry laboratory and for keeping Nu1700 in good shape. Bruno Zürcher was indispensible for maintenance and repairs of the centrifuging piston cylinder. Jan G. Wiederhold has been of great value for the chemical and analytical parts of this work. Discussions with Herbert Palme during his stay at ETH were fruitful, educational and pleasant. We also thank Franck Poitrasson, Mathieu Roskosz and an anonymous reviewer for their constructive comments that improved the original version of this manuscript. Chapter III — Page 75

III–7. APPENDIX

Magnesia (MgO) and alumina (Al2O3) are very common ceramics used in experimental assemblies. Watson et al. (2002) found that when MgO and Al2O3 are in contact, spinel forms according to the reaction

MgO + Al2O3 ↔ MgAl2O4 They empirically determined the temperature and pressure dependence of this reaction to be:

11 48865 D=X 8.58 × 10 ⋅exp  − −2.08 P ⋅ t (III.A1) T in which ΔX is the thickness of the spinel growth layer in μm, T the temperature in Kelvin, P the pressure in GPa and t the time in seconds. With this equation, the temperature distribution in an experimental assembly can be determined from the thickness variation of the thermocouple wire graphite spinel growth layer. To determine talc the temperature distribution inside a Pyrex Pyrex capsule of our experiments, we ran MgO MgO a spinel crystallisation experiment inside a Pt capsule of similar dimensions as used in the experiments in this study. Hence, the thermocouple Pt capsule was also placed in a similar position as in the experiments presented in this manuscript (Figure III–4). The MgO sleeve was manufactured by 10.2 6.4 Norton Ceramics (now Saint-Gobain cement

Advanced Ceramics), the Al2O3 rod corundum and the thermocouple were Degussit AL23 (manufacturer: Frialit-Degus- sit Technical Ceramics). The experiment was run in a static, end-loaded piston cylinder at a temperature (thermocouple reading) of 1400 °C and a pressure of 1.7 GPa. The thickness of the spinel growth layer was determined Degussit AL23 (= ~95% dense Al O ) with a Jeol JSM-6390LA scanning 2 3 electron microscope and converted Figure III–4. Sketch of the assembly used to temperature with equation III.A1. for the spinel growth experiment. The total The temperature calculated in this uncompressed length of the assembly is manner at the thermocouple position 36.06 mm. The length of the capsule and the

Al2O3 bar inside it are given in millimetre. Page 76 — Chapter III was 1407 °C, which falls within the 10 °C uncertainty of the thermocouple reading. The results of the axial temperature distribution inside the capsule are presented in Figure III–5. The regression line (R2 = 0.969) has been used to calculate differences in the temperature in the coldest and hottest ends of the sample area of our capsules (3.7 mm long prior to compression, assumed to be compressed to 3.4 mm during an experiment). Assuming perfect centring of the sample area on the hotspot to a precision of 1 μm, the theoretical temperature difference is 6 °C between a 1 μm thick slice in the hotspot and one at the top or bottom of the capsule. Taking a more realistic uncertainty in the positioning of the capsule of ±0.5 mm, the temperature difference can increase to 9 °C. Fur- thermore, while the temperature at the thermocouple position calculated from the spinel growth layer (1407 °C) agrees well with the temperature indicated by the Eurotherm controller, the temperature calculated in the hotspot inside the capsule is 1458 °C. This relatively large difference arises from the length of the capsules used in this study and the +50 °C correction for “sample temperatures” reported in this study is based on this difference between the modelled temperatures for the thermocouple position and the hotspot.

1465 1460 1455 1450 1445 Temperature (°C) Temperature 1440 T = -2.0 10-06d2 + 1.48 10-02d + 1431 R2 = 0.97 1435 0 1 2 3 4 5 6 7 8 Distance from top of capsule during experiment (mm)

Figure III–5. Axial temperature distribution inside a capsule. Temperatures are determined from the thickness of the spinel growth layer with equation A1. They are plotted as a top-to-bottom profile along the axis of the capsule. The inner wall of the Pt capsule has been chosen as the reference point for the distance d (μm). The Pt capsule wall was protected from the Al2O3 rod by a ~1 mm thick MgO wafer (see Figure III–4). Capsules in experiments presented in the current study were always centred on the hotspot (i.e. top of the parabola). This yields a temperature difference of 6-9 °C between the hottest and coldest 1 μm thick slice of the sample, as can be calculated from the regression equation in the bottom of the figure. Chapter III — Page 77

Table III–A.1 Compositions (wt.%) of all run products. As Table I-2 in manuscript; ce = silicate analysis in centre of glass, co = silicate analysis near contact with metal.

Start Mix1 Start Mix2a RH28 2SD RH28 2SD RH30 2SD RH33 2SD RH42 2SD ce co ce P (GPa) 1.04 0.04 1.04 0.04 1.03 0.05 0.55 0.02 0.99 0.05 T (°C) 1250 1250 1250 1250 1300 t (h) 0.5 0.5 3.0 1.6 0.4 Capsule Graph- Graph- Graph- Graph- Graph- type ite ite ite ite ite

Metal Si 0.05 0.03 0.05 0.03 0.04 0.02 0.03 0.02 0.03 0.03 Fe 70.9 1.8 70.9 1.8 69.9 1.1 73.7 0.9 71.4 1.7 Sn 24.2 1.5 24.2 1.5 25.7 1.3 20.9 0.8 23.9 1.6 Mo 1.67 0.34 1.67 0.34 1.62 0.08 4.24 0.28 1.50 0.24 S n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. C calc.2 3.1 3.1 2.7 1.1 3.1 Total 96.9 0.8 96.9 0.8 97.3 0.5 98.9 0.5 96.9 0.5

Silicate 43.9 41.1 44.3 2.1 45.5 0.5 55.4 SiO2 14.6 14.2 14.2 1.3 14.7 0.2 11.5 Al2O3 MgO 7.18 7.95 7.17 1.66 6.57 0.30 4.83 FeO 17.4 20.0 16.8 2.1 16.3 0.7 13.9 CaO 9.47 9.7 9.39 0.11 5.36 0.51 8.36 2.93 2.63 3.04 0.47 3.17 0.38 2.87 Na2O 1.50 1.12 1.44 0.42 1.57 0.08 0.92 K2O 1.58 1.82 1.80 0.15 1.59 0.17 1.37 SnO2 Total 98.5 98.6 98.1 0.9 94.7 1.2 99.1 Page 78 — Chapter III

Table III–A.1, continued.

Start Mix2a Start Mix2b RH42 2SD RH43 2SD RH441 2SD RH48 2SD RH73 2SD co P (GPa) 0.99 0.05 0.79 0.01 1.00 1.49 0.06 0.91 0.12 T (°C) 1300 1300 1300 1300 1300 t (h) 0.4 3.9 4.0 8.1 6.0

Capsule Graph- Graph- SiO2 SiO2 SiO2 type ite ite glass glass glass

Metal Si 0.03 0.03 0.02 0.02 0.03 0.03 0.04 0.02 0.00 0.00 Fe 71.4 1.7 74.3 4.2 73.3 1.5 76.2 1.3 66.3 5.1 Sn 23.9 1.6 21.1 4.0 25.1 1.4 22.0 1.5 n.d. n.d. Mo 1.50 0.24 1.81 0.31 1.65 0.04 1.81 0.09 n.d. n.d. S n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. 33.1 6.4 C calc.2 3.1 2.7 Total 96.9 0.5 97.3 0.6 100.1 0.5 100.1 0.5 99.4 0.9

Silicate 52.5 55.0 1.1 62.4 1.1 63.0 1.2 60.6 1.9 SiO2 8.56 10.5 1.4 8.92 0.34 8.82 0.44 10.3 0.6 Al2O3 MgO 8.72 5.67 0.92 4.14 0.15 4.40 0.39 4.59 0.26 FeO 16.4 13.9 1.5 10.6 0.7 10.6 1.0 10.9 1.0 CaO 8.73 8.35 0.41 6.54 0.36 6.68 0.41 7.15 0.19 2.09 2.70 0.31 1.60 0.13 1.83 0.19 1.88 0.12 Na2O 0.60 0.90 0.17 0.57 0.09 0.64 0.17 0.68 0.12 K2O 1.71 1.77 0.21 0.79 0.23 0.79 0.30 n.d. n.d. SnO2 Total 99.3 98.8 0.8 95.6 1.5 96.7 0.6 96.1 1.9

1 Experiment RH44 was run in a static piston cilinder press, not in the centrifuging piston cilinder press. 2 Carbon contents are calculated as the difference of the totals from 100%. Chapter III — Page 79 Reference Horn et al. (2006) Horn et al. (2006)* Horn et al. (2006) Horn et al. (2006) Schoenberg and Von Blanckenburg (2006) Blanckenburg Von and Schoenberg et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams (2006) Blanckenburg Von and Schoenberg et al. (2005) Weyer et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams (2006) Blanckenburg Von and Schoenberg et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams (2006) Blanckenburg Von and Schoenberg et al. (2006) Williams et al. (2006) Williams 0.11 0.07 0.10 0.10 0.10 0.10 0.07 0.10 0.10 0.10 0.07 0.10 0.10 0.10 0.10 0.10 0.07 0.10 2SD Fe 57 0.20 0.20 0.24 0.15 0.12 0.05 0.13 0.20 0.18 0.15 0.16 -0.02 -0.56 -0.41 -0.21 -0.25 -0.40 -0.58 δ 0.06 0.06 0.06 0.06 0.05 0.07 0.07 0.07 0.07 0.05 0.04 0.07 0.07 0.07 0.05 0.07 0.07 0.07 0.07 0.07 0.05 0.07 0.07 2SD Fe 56 0.11 0.07 0.14 0.14 0.16 0.10 0.19 0.08 0.06 0.09 0.14 0.12 0.07 -0.06 -0.06 -0.07 -0.38 -0.28 -0.14 -0.17 -0.27 -0.39 δ 0.14-0.34 1 1 1 Specs Bulk Bulk Kamacite Taenite Bulk Metal Metal Sulphide Metal Bulk Bulk Metal Sulphide Sulphide Bulk Metal Sulphide Metal Metal Sulphide Bulk Metal Sulphide Sub-type IA IA IA IA IA IA IA IA IAB IAB IAB IAB IAB IAB IAB IAB IAB IIICD IIICD IIICD IIICD IIICD IIICD Name Toluca Toluca Toluca Toluca Toluca Toluca Toluca Toluca Caddo County Canon Diablo Canon Diablo Canon Diablo Canon Diablo Canon Diablo Odessa Odessa Odessa Mundrabilla Mundrabilla Mundrabilla Nantan Seelägsen Seelägsen Literature compilation Fe isotope data. Literature

Type Iron meteorites Table III–A.2 Table Page 80 — Chapter III Reference et al. (2006) Williams Poitrasson et al. (2005)** Schoenberg and Von Blanckenburg (2006) Blanckenburg Von and Schoenberg et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams (2006) Blanckenburg Von and Schoenberg Horn et al. (2006) Horn et al. (2006) Poitrasson et al. (2005)** (2006) Blanckenburg Von and Schoenberg et al. (2006) Williams et al. (2006) Williams 0.11 0.12 0.04 0.07 0.10 0.10 0.10 0.10 0.10 0.12 0.10 0.10 0.10 0.07 0.10 0.07 0.10 0.10 2SD Fe 57 0.19 0.07 0.02 0.09 0.17 0.18 0.17 0.17 0.22 0.28 0.31 0.01 0.32 0.12 0.18 0.06 0.08 0.13 0.23 -0.55 -0.56 δ 0.08 0.03 0.18 0.05 0.20 0.05 0.07 0.07 0.07 0.07 0.07 0.07 0.08 0.07 0.07 0.07 0.05 0.07 0.05 0.07 0.05 0.07 0.07 2SD Fe 56 0.13 0.06 0.12 0.08 0.12 0.12 0.12 0.15 0.19 0.21 0.01 0.22 0.08 0.12 0.03 0.06 0.03 0.05 0.03 0.16 -0.37 -0.29 -0.38 δ Specs Sulphide Bulk Bulk Metal Metal Metal Metal Metal Metal Sulphide Metal Metal Metal Bulk Bulk1 Bulk Bulk Bulk Metal Sulphide Sub-type IIICD IIICD IC IC IIA IIA IIA IIA IIA IIA IIA IIA IIA IIAB IIAB IIAB IIAB IIAB IIAB IIAB Name Seelägsen z ewell Ta bulk, non-magmatic Average Average metal (solution), non-magmatic Average sulphide (solution), non-magmatic Average Arispe Arispe Bennet County Coahuila Coahuila Gressk Gressk Gressk Guadalupe y Calvo Murphy Negrillos Locust Grove North Chile North Chile North Chile North Chile North Chile North Chile

Type Iron meteorites Table III–A.2, continued. Table Chapter III — Page 81 Reference et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams Poitrasson et al. (2005)** et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams (2006) Blanckenburg Von and Schoenberg (2006) Blanckenburg Von and Schoenberg et al. (2006) Williams 0.11 0.12 0.10 0.13 0.10 0.10 0.10 0.09 0.10 0.10 0.10 0.10 0.10 0.10 0.10 0.10 0.10 0.10 0.10 0.10 0.07 0.07 0.10 2SD Fe 57 0.11 0.22 0.13 0.19 0.19 0.13 0.14 0.21 0.27 0.01 0.14 0.17 0.04 0.01 0.09 0.02 0.15 -0.47 -0.46 -0.06 -0.05 -0.60 -0.59 δ 0.08 0.07 0.09 0.07 0.07 0.07 0.06 0.07 0.07 0.07 0.07 0.07 0.07 0.07 0.07 0.07 0.07 0.07 0.07 0.07 0.05 0.05 0.07 2SD Fe 56 0.15 0.09 0.13 0.13 0.07 0.09 0.09 0.14 0.18 0.01 0.09 0.12 0.03 0.01 0.05 0.02 0.10 -0.32 -0.31 -0.04 -0.03 -0.41 -0.40 δ Specs Sulphide Sulphide Metal Metal Metal Metal Bulk Metal Metal Metal Metal Metal Silicate Silicate Metal Metal Metal Metal Sulphide Sulphide Bulk Bulk Metal Sub-type IIAB IIAB IIB IIB IIB IIB IID IID IID IID IID IIE IIE IIE IIIF IIIF IIIA IIIA IIIA IIIA IIIAB IIIAB IIIB Name North Chile North Chile Mount Joy Navajo Sao Juiao de Moreira Alin Sikhote Carbo Carbo Carbo Carbo Hraschina Watson Watson Watson Clark County Neslon County Merceditas Merceditas Merceditas Merceditas Cape York Mountain Ruffs Bear Creek

Type Iron meteorites Table III–A.2, continued. Table Page 82 — Chapter III Reference et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams Poitrasson et al. (2005)** et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams (2006) Blanckenburg Von and Schoenberg et al. (2006) Williams et al. (2006) Williams (2006) Blanckenburg Von and Schoenberg (2006) Blanckenburg Von and Schoenberg et al. (2006) Williams Poitrasson et al. (2005)** et al. (2006) Williams et al. (2006) Williams et al. (2006) Williams Horn et al. (2006) Horn et al. (2006)* 0.10 0.10 0.10 0.12 0.10 0.10 0.13 0.04 0.10 0.10 0.10 0.07 0.12 0.10 0.07 0.07 0.10 0.13 0.10 0.10 0.10 2SD Fe 57 0.11 0.01 0.14 0.13 0.01 0.02 0.10 0.16 0.01 0.17 0.08 0.09 0.12 0.06 0.04 0.15 0.14 -0.18 -0.05 -0.07 -0.12 δ 0.07 0.07 0.07 0.08 0.07 0.07 0.09 0.03 0.07 0.07 0.07 0.05 0.08 0.07 0.05 0.05 0.07 0.08 0.07 0.07 0.07 0.13 0.05 2SD Fe 56 0.11 0.01 0.09 0.09 0.01 0.01 0.07 0.03 0.12 0.05 0.07 0.07 0.04 0.03 0.10 0.09 0.07 0.03 0.03 -0.12 -0.03 -0.05 -0.08 δ Specs Sulphide Metal Metal Sulphide Metal Metal Metal Bulk Metal Metal Sulphide Bulk Metal Metal Bulk Bulk Metal Bulk Metal Metal Metal Bulk Bulk Sub-type IIIB IIIB IIIB IIIB IIIB IVA IVA IVA IVA IVA IVA IVA IVA IVB IVB IVB IVB IVB IVB IVB IVB ? ? Name Bear Creek Chupaderos Grant Grant Mount Edith Bushman Land Duchesne Duel Hill 1854 Duel Hill 1854 Gibeon Gibeon Hvi z o Pa Yanhuitlan Chinga Hoba Santa Clara Valley Tawallah Tlacotepec Tlacotepec Tlacotepec Range Warburton Glennormiston Glennormiston

Type Iron meteorites Table III–A.2, continued. Table Chapter III — Page 83 Reference (2006) Blanckenburg Von and Schoenberg (2006) Blanckenburg Von and Schoenberg Needham et al. (2009) (2006) Blanckenburg Von and Schoenberg Craddock and Dauphas (2010) Craddock and Dauphas (2010) Poitrasson et al. (2005)** Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Needham et al. (2009) (2006) Blanckenburg Von and Schoenberg Needham et al. (2009) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Needham et al. (2009) Needham et al. (2009) 0.07 0.07 0.07 0.06 0.05 0.06 0.09 0.05 0.07 0.07 0.09 0.05 2SD Fe 57 0.11 0.05 0.01 0.02 0.13 0.02 0.02 0.08 0.02 0.01 0.04 -0.11 -0.03 -0.04 -0.01 -0.02 -0.07 -0.03 δ 0.05 0.05 0.04 0.12 0.35 0.10 0.12 0.29 0.04 0.05 0.04 0.03 0.04 0.06 0.03 0.05 0.04 0.05 0.04 0.06 0.03 0.09 0.04 2SD Fe 56 0.03 0.03 0.04 0.10 0.05 0.10 0.01 0.01 0.01 0.02 0.08 0.01 -0.22 -0.25 -0.03 -0.02 -0.01 -0.07 -0.01 -0.18 -0.07 -0.14 -0.02 δ Specs Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Metal Bulk Sub-type ? ? H3.8 H3-H5 H4 H4 H4 H4 H4 H4 H5 H5-H6 H6 H6 H6 H6 H/L3.9 Name Glennormiston Pinon bulk, magmatic Average metal (solution), magmatic Average sulphide (solution), magmatic Average bulk, all irons Average metal (solution), all irons Average sulphide (solution), all irons Average Dhajala Dar al Gani 300 Bielokrynitschie Bielokrynitschie Forest Vale Ochansk Ochansk Kesen Allegan Djoumine Kernouve Kernouve Kernouve Kernouve Bremervoorde

Type Iron meteorites Ordinary chondrites Table III–A.2, continued. Table Page 84 — Chapter III Reference Needham et al. (2009) Needham et al. (2009) Needham et al. (2009) Poitrasson et al. (2005)** Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Needham et al. (2009) Needham et al. (2009) Craddock and Dauphas (2010) Craddock and Dauphas (2010) (2006) Blanckenburg Von and Schoenberg (2006) Blanckenburg Von and Schoenberg Craddock and Dauphas (2010) Needham et al. (2009) Needham et al. (2009) Needham et al. (2009) Needham et al. (2009) Needham et al. (2009) Needham et al. (2009) Needham et al. (2009) Needham et al. (2009) Needham et al. (2009) 0.11 0.08 0.05 0.04 0.05 0.05 0.07 0.07 0.05 2SD Fe 57 0.00 0.02 0.02 0.06 0.02 -0.00 -0.00 -0.05 -0.02 δ 0.07 0.06 0.07 0.07 0.06 0.03 0.04 0.07 0.06 0.03 0.03 0.05 0.05 0.03 0.10 0.06 0.08 0.09 0.06 0.04 0.04 0.06 0.06 2SD Fe 56 0.15 0.03 0.01 0.04 0.02 0.20 0.00 0.01 0.03 -0.11 -0.01 -0.00 -0.01 -0.02 -0.02 -0.10 -0.04 -0.01 -0.03 -0.04 -0.37 -0.18 -0.14 δ Specs Metal Metal Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Metal Bulk Bulk Bulk Metal Metal Sulphide Sulphide Sulphide Sub-type H/L3.9 L3.2 L3.6 L3.7 L4 L4 L4 L4 L5 L5 L5 L5 L5 L6 L6 L/LL4 LL3.4 LL3.6 LL3.6 LL3.6 LL3.6 LL3.6 LL3.6 Name Bremervoorde 86018 LEW Khohar Me z ö-Madaras Bald Mountain Baratta Dalgety Downs Saratov Elenovka Farmington Farmington Homestead z erait Mt. Ta Harleton Barwell Bjurbole Chainpur Parnallee Parnallee Parnallee Parnallee Parnallee Parnallee

Type Ordinary chondrites Table III–A.2, continued. Table Chapter III — Page 85 Reference Needham et al. (2009) (2006) Blanckenburg Von and Schoenberg Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Needham et al. (2009) Craddock and Dauphas (2010) Needham et al. (2009) Needham et al. (2009) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Needham et al. (2009) Needham et al. (2009) Craddock and Dauphas (2010) Craddock and Dauphas (2010) 0.07 0.04 0.04 0.05 0.09 0.04 0.09 0.05 0.07 0.07 0.05 2SD Fe 57 0.07 0.00 0.10 0.04 0.02 0.02 0.02 0.08 0.02 -0.01 -0.18 -0.15 -0.14 -0.05 δ 0.06 0.05 0.04 0.04 0.03 0.06 0.07 0.04 0.10 0.06 0.06 0.03 0.05 0.08 0.03 0.04 0.03 0.10 0.17 0.13 2SD Fe 56 0.03 0.01 0.04 0.02 0.01 0.07 -0.11 -0.11 -0.11 -0.02 -0.00 -0.13 -0.01 -0.07 -0.04 -0.03 -0.01 -0.02 -0.03 -0.02 δ Specs Sulphide Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Metal Bulk Bulk Bulk Bulk Bulk Bulk Bulk Sub-type LL3.6 LL4 LL4 LL4 LL4 LL4 LL4 LL4 LL5 LL5 LL5 LL5 LL5 LL6 LL6 LL6 LL6 Name Parnallee Dar al Gani 298 Hamlet Kelly Kelly Kelly Soko Banja Soko Banja Oliven z a Oliven z a Paragould Paragould Tuxtuac Mangwendi Dhurmsala Saint-Séverin Saint-Séverin bulk Average metal (solution) Average bulk+metal Average

Type Ordinary chondrites Table III–A.2, continued. Table Page 86 — Chapter III Reference (2006) Blanckenburg Von and Schoenberg Craddock and Dauphas (2010) Poitrasson et al. (2005)** Craddock and Dauphas (2010) Poitrasson et al. (2005)** (2006) Blanckenburg Von and Schoenberg Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Poitrasson et al. (2005)** (2006) Blanckenburg Von and Schoenberg (2006) Blanckenburg Von and Schoenberg Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) et al. (2008) Teng (2006) Blanckenburg Von and Schoenberg 0.11 0.07 0.07 0.07 0.06 0.07 0.04 0.04 0.07 0.07 0.05 0.05 0.07 0.18 0.07 0.04 0.04 0.05 0.06 0.05 0.05 0.07 2SD Fe 57 0.00 0.02 0.02 0.01 0.01 0.02 0.02 -0.11 -0.23 -0.03 -0.04 -0.02 -0.04 -0.08 -0.01 -0.04 -0.04 -0.01 -0.00 -0.01 -0.00 -0.01 δ 0.05 0.05 0.04 0.07 0.07 0.05 0.03 0.03 0.05 0.04 0.03 0.03 0.04 0.09 0.05 0.03 0.03 0.03 0.03 0.03 0.03 0.06 2SD Fe 56 0.01 0.00 0.00 -0.17 -0.03 -0.03 -0.02 -0.03 -0.06 -0.03 -0.01 -0.02 -0.01 -0.07 -0.02 -0.01 -0.00 -0.01 -0.02 -0.01 -0.01 -0.01 δ Specs Bulk or matrix? Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Sub-type CB CI CI CI CM2 CM2 CM2 CM2 CM2 CO3 CO3 CO3.2 CV3 CV3 CV3 CV3 CV3 CV3 CV3 CV3 CV3 CV3 Name Bencubbin-matrix Ivuna Orgueil Orgueil Murchinson Murchinson Murchinson Murchinson Mighei Lancé Lancé Kainsa z Allende Allende Allende Allende Allende Allende Allende Allende Allende Axtell -

Type Carbona ceous chondrites Table III–A.2, continued. Table Chapter III — Page 87 Reference Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) (2006) Blanckenburg Von and Schoenberg Craddock and Dauphas (2010) Craddock and Dauphas (2010) (2006) Blanckenburg Von and Schoenberg Craddock and Dauphas (2010) Craddock and Dauphas (2010) Poitrasson et al. (2005)** (2006) Blanckenburg Von and Schoenberg Craddock and Dauphas (2010) Poitrasson et al. (2005)** Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) (2006) Blanckenburg Von and Schoenberg Craddock and Dauphas (2010) 0.10 0.09 0.08 0.07 0.04 0.06 0.07 0.07 0.07 0.05 0.07 0.07 0.09 0.05 0.05 0.07 0.13 0.07 0.05 0.16 0.05 2SD Fe 57 0.04 0.00 0.02 0.01 0.06 0.02 0.01 0.03 0.03 0.06 0.05 0.01 -0.02 -0.06 -0.09 -0.01 -0.02 -0.08 -0.01 -0.02 δ -0.03 -0.19 0.07 0.06 0.06 0.05 0.04 0.07 0.07 0.05 0.05 0.05 0.03 0.06 0.05 0.06 0.03 0.03 0.05 1.03 0.03 0.03 0.08 0.03 2SD Fe 56 0.03 0.01 0.03 0.02 0.00 0.03 0.01 0.05 0.04 0.00 -0.03 -0.06 -0.05 -0.01 -0.02 -0.01 -0.00 -0.04 -0.01 -0.04 δ -0.02 -0.14 Specs Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Sub-type CV3 CV3 CV3 CV3 C2 C2 EH3 EH3 EH3 EH4 EH4 EH4 EH4 EH4 EH4 EH5 EH5 EH5 EH5 EL6 EL6 Name Grosnaja Viagarano Viagarano SAU (unnamed) Lake Tagish Lake Tagish bulk Average Sahara Sahara 97072 Qing z hen Abee Abee Adhi Kot Indarch Indarch Indarch St. Marks St. Marks Saint-Sauveur Saint-Sauveur Atlanta Blithfield -

Type Carbona ceous chondrites Enstatite chondrites Table III–A.2, continued. Table Page 88 — Chapter III Reference Craddock and Dauphas (2010) (2006) Blanckenburg Von and Schoenberg Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) Craddock and Dauphas (2010) 0.05 0.08 0.04 0.04 0.05 0.10 0.05 0.05 0.04 0.05 0.05 2SD Fe 57 0.18 0.02 0.10 0.04 0.14 0.09 0.04 0.14 0.10 0.02 0.01 0.02 0.02 -0.03 -0.02 δ 0.03 0.06 0.04 0.04 0.03 0.07 0.03 0.03 0.04 0.04 0.04 0.10 0.09 2SD Fe 56 0.12 0.01 0.06 0.03 0.09 0.06 0.01 0.08 0.06 0.02 0.07 0.13 -0.00 -0.01 -0.03 δ Specs Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Bulk Sub-type EL6 EL6 EL6 EL6 EL6 EL6 EL6 EL6 EL6 EL6 EL6 Name Kuil Daniel’s Eagle Eagle Happy Canyon Hvittis Hvittis Ilafegh 009 Jajh deh Kot Lalu Khairpur Pillistfer Yilmia bulk Average bulk ordinary/carbonaceous chondrites Average ΔIron meteorites (bulk)-Chondrites ΔIron meteorites (metal)-Chondrites Type Enstatite chondrites = fs-LA-ICP-MS Table III–A.2, continued. Table 1 * = in Horn et al. (2006) quoted as unplished data from Schoenberg. ** = Errors are 2SE rather than 2SD. give respectively the average, 2SD and 95% confidence intervals. Averages Individual errors are 2SD; He z el et al. (2010) measurements are not included because of poor uncertainties due to low sample weights. CHAPTER IV.

EXPERIMENTAL DETERMINATION OF THE SI ISOTOPE FRACTIONA- TION FACTOR BETWEEN LIQUID METAL AND LIQUID SILICATE1

1 This chapter will be submitted to Earth and Planetary Sciences Letters as: R. C. Hin, C. Fitoussi, M. W. Schmidt, B. Bourdon. Experimental determination of the Si isotope fractionation factor between liquid metal and liquid silicate. Page 90 — Chapter IV

ABSTRACT

The conditions of core formation and the abundance of light elements in Earth’s core remain debated. Silicon isotope fractionation may shed new light on these debates. We present experimentally determined Si isotope fractionation factors between liquid metal and liquid silicate for 1450 °C and 1750 °C. These data have been used to calibrate the temperature dependence of Si isotope fractionation. Experiments were performed in a centrifuging piston cylinder at a pressure of 1 GPa, employing both graphite and MgO capsules. Tin was used to lower the melting temperature of the metal alloys for experiments performed at 1450 °C. An alkaline fusion technique was employed to dissolve silicate as well as metal phases prior to ion exchange chemistry and mass spectrometric analysis. The results show that metal is consistently enriched in light isotopes relative to the silicate, yielding average metal-silicate fractionation factors of -1.48 ± 0.08‰ and -1.11 ± 0.14‰ at 1450 °C and 1750 °C, respectively. The temperature dependence of equilibrium Si isotope fractionation between metal 30 × 6 2 and silicate can therefore be described as Δ SiMetal-Silicate = -4.47(±0.31) 10 /T . Using a bulk silicate Earth δ30Si of -0.29 ± 0.01‰ (95% confidence interval), the determined temperature dependence of Si isotope fractionation can be used to calculate δ30Si values of -1.01 ± 0.05‰ and -0.66 ± 0.03‰ for Earth’s core using metal-silicate equilibration temperatures of 2500 and 3500 K, respectively, during core formation on Earth. Mass balance calculations indicate that the Si isotope composition of the Bulk Earth must be -0.38 ± 0.02‰ or -0.33 ± 0.01‰ for 2500 or 3500 K if Earth’s core is assumed to contain 6 wt% Si, or -0.40 ± 0.02‰ or -0.35 ± 0.02‰, respectively, for 8 wt% Si in the Earth’s core. These values are difficult to match with reported Si isotope compositions of enstatite chondrites, but a comparison with the Si isotope compositions of other chondrites is currently hindered by the dispute about these chondrite compositions.

IV–1. INTRODUCTION

The segregation of the metallic core from the silicate mantle is the largest chemical and physical differentiation process in Earth’s history. In terms of its major components, the chemical composition of the Earth’s upper mantle is quite well known and often assumed to be identical to that of the lower mantle (Palme and O’Neill, 2003). Seismological observations indicate that the core’s density is about 7% lower than an Fe-Ni alloy at the appropriate pressures and temperatures (Anderson and Isaak, 2002; Birch, 1952). Although several elements have been proposed to explain this density deficit (e.g. H, C, O, S, Si), Si has received considerable attention because several per cent of Si in Earth’s core could explain the low Mg/Si ratio of the Bulk Silicate Earth (BSE) relative to carbonaceous chondrites (Ringwood, 1959). However, the contents of light elements in Earth’s Chapter IV — Page 91 core remain debated (McDonough, 2003; Poirier, 1994; Wood et al., 2006). The temperature, pressure and oxygen fugacity conditions under which core segregation took place could help to constrain the nature of the light elements in Earth’s core. These conditions have mainly been constrained to a range of 2500-3500 K, 20-60 GPa and about 1-4 log units below the iron-wüstite buffer, based on elemental distribution coefficients that were determined with (liquid) metal – (liquid) silicate equilibration experiments (Corgne et al., 2008; Mann et al., 2009; Righter, 2003; Siebert et al., 2011; Wade and Wood, 2005). However, those conditions have been subject to a longstanding debate and are, for instance, dependent on whether equilibrium or disequilibrium accretion models are considered (Rubie et al., 2011; Wood et al., 2006). In this respect, Rudge et al. (2010) have shown that elemental partitioning is only mildly sensitive to the degree of final metal-silicate equilibration during core formation. Mass dependent fractionation of stable isotopes, however, is strongly dependent on temperature and may therefore be a new tool to help decipher the conditions under which metal- silicate segregation occurred in planetary bodies. Mass dependent fractionation of stable isotopes occurs when the zero- point vibrational energies of isotopes vary significantly between two phases (Bigeleisen and Mayer, 1947; Schauble, 2004; Urey, 1947). Such variability is mainly dependent on the bond stiffness and isotopic fractionation decreases with the inverse of the square of temperature. The stiffness of bonds is mainly a function of bond type, valence state, coordination state and bonding partner (Schauble, 2004). Because the fractionation is mass dependent, relative mass differences between the isotopes of an element are also an important factor determining the magnitude of fractionation. A relatively light element such as Si is therefore more likely to fractionate to analytically resolvable magnitudes than a heavier element. The only experimental study of Si isotope fractionation between metal and silicate has been performed by Shahar et al. (2011; 2009). Shahar et al (2009) found that liquid metal was at least 2‰ lighter in δ30Si isotope composition than liquid silicate at 1800 °C. In a follow up study, Shahar et al. (2011) determined that similar fractionation occurred between liquid metal and olivine. They concluded that equilibrium Si isotope fractionation between metal and silicate 30 × 6 2 can be described as Δ SiMetal-Silicate = -7.45 ± 0.41 10 /T . Combining their results with the Si isotope compositions of chondrites and the BSE (Armytage et al., 2011; Chakrabarti and Jacobsen, 2010; Fitoussi et al., 2009; Georg et al., 2007; Savage et al., 2010), Shahar et al. (2009) concluded that their work supports the presence of several weight per cent of Si in Earth´s core in an equilibrium model of core formation. The difference in Si isotope compositions between chondrites and the BSE, however, is disputed as particularly the chondrite compositions are not well defined. This is partly related to the fact that Si isotope compositions were found Page 92 — Chapter IV to be difficult to analyse in mafic rocks/minerals, perhaps due to matrix effects arising from the ion exchange chemistry or sample introduction procedures during analyses (Armytage et al., 2011; Chakrabarti and Jacobsen, 2010; Fitoussi et al., 2009; Georg et al., 2007). We have determined the equilibrium Si isotope fractionation between liquid metal and liquid silicate, making use of a centrifuging piston cylinder at ETH Zurich for segregation of metal and silicate phases. Our results will be discussed in the context of core formation models in the early Earth.

IV–2. METHODS IV–2.1. Experimental methods Our approach to determine Si isotope fractionation factors between liquid metal and liquid silicate was to enhance the stable isotope fractionation. We have therefore performed experiments at relatively low temperatures for metal-silicate segregation in a centrifuging piston cylinder that ensures perfect metal-silicate separation and thus allows for bulk sample chemistry prior to mass spectrometric analysis. The metal had a Fe based composition, to which Sn was added to lower the Table IV–1. Compositions (wt%) of liquidus of the metal alloy below 1500 °C starting mixtures. fM and fSil are the (Okamoto, 1993). Basaltic compositions mass fractions of metal and silicate, were chosen for the silicate phase. Two respectively. synthetic starting mixtures were prepared: Start Mix1 Start Mix2 Mix1 consisted of a basaltic oxide mixture homogenised with a Fe-Sn-Si-Mo metal Metal powder; Mix2 had a nearly identical com- Fe 82.4 92.9 position to Mix1, except that the metal Sn 10.4 component was Sn-free (Table IV–1). Both mixtures were prepared with approx- Si 5.68 5.65 imately 75 wt% metal and 25 wt% oxide Mo 1.53 1.48 component. Theoretically, mass dependent stable Silicate SiO 49.6 49.5 isotope fractionation should not be 2 14.6 14.6 sensitive to the exact composition of the Al2O3 phases involved, because the valence and MgO 15.4 15.5 coordination state of atoms as well as type Fe2O3 5.05 5.13 of bonds are the main parameters that CaO 13.1 13.0 control fractionation (Schauble, 2004). In Na O 1.80 1.80 addition, Poitrasson et al. (2009) and Hin 2 K O 0.49 0.49 et al. (2012) have found that neither the 2 presence of carbon nor that of sulphur in fM 0.7586 0.7584 the metal component influenced Fe isotope 0.2414 0.2416 fSil Chapter IV — Page 93 fractionation between liquid metal and liquid silicate at temperatures as low as 1250 °C. We have nevertheless verified the influence of Sn and C by direct comparison of Sn-bearing with Sn-free (Mix1 and Mix2), as well as C-bearing with C-free systems as detailed below. A time series was performed at temperatures of 1450 °C to determine the run duration required for Si isotope equilibrium in our experiments. Shahar et al. (2009) found that 30 minutes run duration did not suffice to fully reach equilibrium at 1800 °C, so we varied run duration in four experiments stepping from 0.5 to 1.5, 4.0 and 22.0 hours. These experiments were performed in graphite capsules with 3.5 mm outer diameter, 2.0 mm inner diameter and 3.7 mm inner length. An MgO capsule was used in one experiment (4 h run duration) at 1450 °C to directly compare C-bearing and C-free systems. In addition to experiments at 1450 °C, we have performed experiments at 1750 °C. This enabled melting a Sn-bearing as well as Sn-free metal alloy. These additional experiments also better constrain the temperature dependence of isotopic fractionation, which is theoretically expected to decrease as a function of the square of temperature (Bigeleisen and Mayer, 1947; Schauble, 2004; Urey, 1947). Again, both graphite and MgO capsules were used for these experiments. All capsules were welded into tin can shaped Pt jackets (4.0 mm outer diameter, ~10.5 mm outer length). To prevent oxidation or melting of the Sn powder before the experiments, filled capsules were only dried for 30 minutes at 110 °C prior to welding. The welded capsules were positioned in the piston cylinder graphite furnace such that temperature gradients in the sample were minimised, following the protocol detailed in Appendix A in Hin et al. (2012). Temperatures given here are ‘sample temperatures’ that include a +50 °C correction of the temperatures recorded by the thermocouple. Because the metal and silicate phases each comprise roughly 50 vol% of the sample, the careful positioning of the samples in the hotspot results in identical average temperatures in the metal and silicate phases. Talc/pyrex assemblies with cylindrical furnaces with MgO spacers and mullite thermocouple sleeves were used for experiments at 1450 °C. Talc/SiO2 glass assemblies were used with stepped furnaces and 99% Al2O3 thermocouple sleeves for experiments at 1750 °C. One experiment (RH27) was run in a static, Boyd and England type (Boyd and England, 1960) end-loaded piston cylinder. All other experiments were run in a single stage centrifuging piston cylinder (Hin et al., 2012; Schmidt et al., 2006). As this apparatus is limited to a heating current of 380 A, we have used the better heating efficiency of stepped furnaces for experiments at 1750 °C. For the compositional systems and temperatures employed here, centrifugation at 400g produces a sinking velocity ≥1 mm h-1 for metal droplets with diameters as small as 2 μm, using silicate melt viscosities for olivine tholeiite from Kushiro et al. (1976). Under these conditions, such droplets would have sunk through the full length of the silicate melt within 1.5 h at 400g. Page 94 — Chapter IV

After quenching by switching off power, all samples were embedded in epoxy and polished for microprobe analyses. The samples were subsequently cut at the metal-silicate contact and taken out of the capsules with a 220 or 300 μm diameter diamond wire saw. This procedure removed any contamination within about ~300 μm from the metal-silicate and sample-capsule contact areas. The bulk pieces of silicate and metal were polished with 30 μm Al2O3 polishing paper to remove any surface contamination such that a single piece of clean, metal- free silicate glass (generally between 1.7 and 4 mg) and a single piece of clean, silicate-free metal (4-15 mg) remained. These pieces were finally crushed under acetone with separate alumina pestles and mortars for metals and silicates to prevent cross contamination. IV–2.2. Analytical methods IV–2.2.1. Elemental compositions Polished and C-coated samples were analysed for their elemental compositions with a Jeol JXA 8200 electron microprobe equipped with five spectrometers. An acceleration voltage of 15 kV was employed to analyse silicates with a 6 nA beam intensity and a 1-10 μm beam diameter. Metals were analysed with a higher intensity of 20 nA and a beam diameter of 10 μm. Analyses were performed relative to pure metal standards for metal compositions, and relative to natural and synthetic silicate and oxide standards for silicate compositions. On average, 10 spots were analysed on silicates and 14 spots on metals.

IV–2.2.2. Ion exchange chemistry We employed an alkaline fusion technique for dissolution of our samples (Georg et al., 2006). Approximately 1 mg of crushed silicate sample was weighed in silver crucibles. A sodium hydroxide pellet was added to the sample. Disso- lution then took place by fusion of the NaOH and sample mixture in a muffle furnace at ~730 °C for 10 minutes. The cooled samples were subsequently transferred with the silver crucible into pre-cleaned Savillex PFA beakers containing a volume of ultra-pure H2O (MQ-water; ≥183 kΩ m). After ultrasonic treatment (~5 min) and equilibration on a hotplate overnight, the silver crucibles were removed from the solution and rinsed with MQ-water. Sample solutions were then acidified to a pH of 2.0-2.3 prior to loading on 2 ml polypropylene columns filled with 1 ml Biorad AG50W-X8 200-400 mesh cation exchange resin in H+ form. This resin was pre-cleaned with 3M and 6M

HCl and one step of 14M HNO3 before it was equilibrated with MQ-water to produce a neutral pH. About 0.5 ml of sample was loaded onto a column and eluted with 2 ml MQ-water. The collected 2.5 ml of Si solution was subsequently diluted with MQ-water to generate solutions with 600-700 ppb Si. Chemical and analytical protocols for metal samples are based on those described above for silicate samples. Some modifications were required because Chapter IV — Page 95 of the lower Si concentrations and higher Fe/Si in the metals relative to silicates. Sample solutions were obtained according to the procedure described above, but the solutions were acidified with HCl to render a pH of ~1.0 instead of 2.0-2.3. This stronger acidification prevents precipitation of iron hydroxides, but it required column elution with 4 ml instead of 2 ml MQ-water, probably due to + formation of H5SiO4 during column elution in the more acidic environment. Fitoussi et al. (2009) observed that aliquots of NBS-28 that were acidified to a pH of ~1.0 yielded δ30Si compositions of -0.3 to -0.4‰ relative to non-acidified NBS-28 solutions (i.e. pH of 2.0-2.3). To analyse the standards under the same conditions as the samples, we therefore acidified aliquots of solutions of the NBS-28 and ‘Diatomite’ standards to a pH of ~1.0 prior to batches of ion exchange chemistry involving metal samples (see explanation below).

IV–2.2.3. Mass spectrometric analyses Analyses were performed on a Nu Plasma 1700 multi-collector inductively coupled plasma mass spectrometer (MC-ICPMS) at ETH Zurich following a protocol described in Georg et al. (2006) and Fitoussi et al. (2009). A Cetac ASX-100 auto-sampler was used to aspirate the samples from 7 ml PFA beakers with a PFA nebuliser (~80 μm min-1 uptake rate) and transfer them into a DSN-100 desolvation system. The MC-ICPMS was operated at a mass resolution of >2000 (m/Δm, 10% valley convention) to fully resolve all isobaric interferences. All three isotopes of Si were collected simultaneously. Sample analyses were bracketed with analyses of the international Si isotope reference standard NBS-28. For each dissolved sample, at least 2 aliquots of the solution were passed through ion exchange chromatography procedures, while commonly 6 to 8 standard-sample-standard brackets were analysed for a single aliquot. A single analysis consisted of 36 cycles with a 5 s integration time, usually with an intensity of ~6×10-11 A on mass 28. Silicon isotope compositions are reported as per mil deviations of the sample 30Si/28Si ratio from the 30Si/28Si ratio of the NBS-28 standard: 30Si 28 Si d 30Si =sample −⋅1 1000 (IV.1) 30Si 28Si NBS −28 Uncertainties for sample isotopic compositions are reported as the 95% confidence interval (95% c.i.) of the number (n) of repeated standard-sample-standard brackets: SD (IV.2) 95% ci. . = ⋅ t0.95 n

where t0.95 refers to a two-tailed 95% confidence test (values from Miller and Page 96 — Chapter IV

Miller, 2010). SD is the standard deviation calculated as: 2 ∑()xxi − SD = (IV.3) n −1

in which xi and x represent a single measurement on a single sample and the average of n measurements on a single sample, respectively. The internation- ally accepted secondary standard ‘Diatomite’ was analysed repeatedly in each measurement session and all data collected in a session were discarded when the Diatomite standard did not have a δ30Si of 1.26 ± 0.02‰ (95% c.i.) (Reynolds et al., 2007). The average δ30Si of the Diatomite standard obtained during our study is 1.24 ± 0.02‰ (95% c.i., n = 76; Table IV–2). Differences in the isotopic compositions of two phases are presented as the fractionation factor (Δ��) between phases A and B, such that a negative fractionation factor indicates a lighter Si isotope composition for phase A: 30 30 30 D=−SiAB− dd Si A Si B (IV.4)

IV–2.2.4. Accuracy of metal analyses The pH of NBS-28 and Diatomite solutions was carefully matched with the pH of metal sample solutions by acidification prior to batches of ion exchange chemistry. This led to Diatomite δ30Si compositions of 1.22 ± 0.03‰ (95% c.i., n = 117; Table IV–2), which is within analytical uncertainty identical to its true composition. This means that the acidification together with the modification of the elution volume does not lead to any deviation in the secondary standard.

Table IV–2. Isotopic compositions of silicon-oxide and metal used in starting mixtures.

δ30Si 95% c.i. δ29Si 95% c.i. n Powder (‰) (‰) (‰) (‰) Diatomite 1.24 0.02 0.64 0.01 76 Diatomite (a)1 1.22 0.03 0.62 0.01 117

SiO2 -2.12 0.04 -1.09 0.02 42

1:1 SiO2 : Fe94Si6 -1.90 0.04 -0.98 0.03 9

1:1 SiO2 - Fe94Si6 (a) -1.91 0.06 -0.99 0.03 16

Fe94Si6 (a) -0.38 0.04 -0.19 0.03 19 2 -0.29 0.64 -0.22 0.32 Fe94Si6 calc.

1 a indicates that the acidified method was used, which was designed for analyses of metals (see text for details). 2 calc. indicates that the value was calculated by mass balance from the

SiO2 and 1:1 SiO2-Fe94Si6 mixture. Chapter IV — Page 97

We have further investigated the validity of the modifications to the ion exchange and analysis protocols for metal samples compared to the original protocols by directly comparing the two methods. For this purpose, we have made a 1:1 mixture of the SiO2 and Fe83Si17 (mixed with Fe to form Fe94Si6) powders used for Mix1 and Mix2 in our experiments. This mixture was analysed with both the conventional and the modified protocols, which resulted in identical compositions of -1.90 ± 0.04‰ and -1.91 ± 0.09‰ (95% c.i.), respectively (Table IV–2).

Additionally, the pure SiO2 was analysed following the conventional method for silicate samples, while the pure metal was analysed with the modified protocol

(Table IV–2). Using the pure SiO2 composition of -2.12 ± 0.04‰ and an average of -1.91 ± 0.06‰ for the 1:1 mixture, the metal composition can be calculated by mass balance. This calculation yielded a metal composition in the 1:1 mixture of -0.29 ± 0.64‰ (propagated 95% c.i.), which is very close to the measured value of -0.38 ± 0.04‰. The above findings imply that the modifications we made to our analytical methods for metal samples provide accurate results.

IV–3. RESULTS

Elemental concentrations are presented in Table IV–3 and Table IV–A.1 (at the end of this chapter), while isotopic compositions as well as details of experimental run conditions are given in Table IV–4. IV–3.1. Textures and elemental compositions All experiments performed at 1450 °C exceeded the liquidus of the silicate and metal phase. The metal liquid was forced to the bottom part of the capsule, leading to an almost complete segregation of the metal and silicate melts (Figure IV–1). Only a small number of metal droplets of few μm diameter remained in the silicate melt at the capsule contact, held by surface forces (Figure IV–1a, b). Following Stokes’ law, metal droplets with diameters ≥2 μm that were not affected by capsule sticky forces should have travelled the full length of the silicate phase in all experiments as an effect of centrifugation and thus should have collected in the main pool of metal melt. Image area analyses indicate that droplets with diameters <2 μm constitute <0.1 vol% of the silicate melt pool. Experiments performed in MgO capsules at 1450 °C show quench needles of olivine as well as small (< 15 μm) euhedral grains of olivine at the contact between the silicate phase and the MgO capsule (Figure IV–1c). These can be attributed to dissolution of MgO from the capsule into the silicate liquid, the basaltic liquid being undersaturated in MgO. The enrichment of MgO in the silicate melt at the contact with the capsule thus stabilises olivine and at the same time depolymerises the basaltic melt, leading to growth of olivine during quenching. As presented in Figure IV–1d and 1e, faster dissolution of MgO from the Page 98 — Chapter IV capsules in experiments at 1750 °C resulted in much larger and more abundant quench needles and euhedral grains of olivine than in experiments at 1450 °C. Furthermore, the segregation of metal in experiments at 1750 °C was found to be more difficult due to the presence of stable olivine grains (Figure IV–1e). The higher temperatures and larger instability of the assembly occasionally led to

Figure IV–1. Back-scattered electron images of selected experimental products. Panel (a) shows the segregated metal and silicate pools in an experiment in graphite at 1450 °C. Tiny bright spots at the silicate- capsule contact are a few tiny metal droplets that have not segregated. As shown in panel (b), sub-micrometre sized metal droplets also occur at larger distance from the capsule wall. In MgO capsules, crystallisation of euhedral olivine as well as formation of olivine needles upon quenching is much less at 1450 °C (panel d) than at 1750 °C (panel e). In the latter experiments, the large euhedral olivine crystals can prevent full segregation of metal and silicate liquids (panel e). Panel (f) shows a crack in the capsule after a run at 1750 °C. As a consequence, Pt-bearing metal occurs in the sample. The metal that is not well segregated from the silicate centrifugation was removed from the silicate by hand picking prior to chemistry procedures. Chapter IV — Page 99

Table IV–3. Elemental compositions (wt%) of selected experimental products. Pressures (P), temperatures (T), run durations (t) and oxygen fugacity (fO2) of the experiments are indicated as well. Uncertainties can be relatively high due to quench heterogeneity in all metals as well in silicates from experiments performed at 1750 °C.

Start Mix1 Start Mix2 RH31 2SD RH51 2SD RH66 2SD RH65 2SD P (GPa)1 0.93 0.01 0.83 0.08 0.63 0.27 0.54 0.12 T (°C) 1450 1450 1750 1750 t (h) 0.52 4.02 1.52 1.52 fO2 (ΔIW) -4.11 -4.53 -4.79 -4.82 Capsule Graph- MgO MgO MgO type ite

Metal Fe 81.7 0.7 85.0 3.8 80.4 17.2 93.5 1.4 Sn 9.12 1.54 8.84 4.82 13.4 20.0 n.d. n.d. Si 5.02 0.27 5.56 0.24 4.74 1.23 4.93 0.14 Mo 1.45 0.35 1.33 0.22 1.65 1.09 1.53 0.48 Pt 0.04 0.05 0.03 0.08 0.25 0.42 0.03 0.05 Total 97.3 0.9 100.7 1.1 100.5 1.7 100.0 1.2 C (calc)2 2.7

Silicate

SiO2 57.65 0.96 48.43 0.54 34.90 1.36 35.40 1.88

Al2O3 13.44 1.76 12.29 0.33 11.15 2.80 9.68 4.08 MgO 15.25 1.66 23.72 0.40 36.47 2.98 39.13 5.76 FeO 0.16 0.19 0.12 0.09 0.08 0.09 0.11 0.06 CaO 12.5 0.3 12.1 0.2 12.3 1.4 10.8 3.1

Na2O 1.62 0.13 1.60 0.12 1.77 0.14 1.54 0.49

K2O 0.43 0.05 0.46 0.05 0.48 0.08 0.42 0.15

SnO2 0.02 0.03 0.01 0.04 0.02 0.02 n.d. n.d. Total 101.1 0.8 98.7 0.8 97.1 0.5 97.1 0.6

where Xi and γi are the molar fractions and activity coefficients of i, respectively, in the metal or silicate melt. Holzheid et al. (1997) observed that the activity coefficient of FeO does not significantly depend on the silicate composition in the temperature range 1300-1600 °C, at least for FeO contents up to ~12 wt% and MgO contents up to ~20 wt%. For these calculations, γFeO was therefore taken to be 1.7 after Holzheid et al. (1997), although our experiments performed at 1750 °C, particularly those performed in MgO capsules, fall outside the calibrated temperature and composition range. Values for γFe were calculated with a modified Wagner epsilon approach (Ma, 2001) using interaction parameters from the Steelmaking Data Sourcebook (1988). Oxygen fugacities were found to vary from ΔIW -4.5 to -3.2.

IV–3.2. Isotopic compositions 30 The δ Si compositions of the SiO2 and Fe83Si17 as sources for the starting mixtures have been measured to be -2.12 ± 0.04 and -0.38 ± 0.04‰ (95% confidence intervals), respectively (Table IV–2 and Table IV–4). Mass balance calculations using these compositions yield a calculated bulk composition for Mix1 and Mix2 of -1.36 ± 0.03‰, which is within analytical uncertainty of the measured bulk composition of -1.32 ± 0.09‰ for Mix1. Mix2, however, has a heavier measured composition of -1.03 ± 0.08‰. Chapter IV — Page 101

Experiments performed at 1450 °C in graphite have resulted in metal melt δ30Si compositions between -1.98 ± 0.08‰ and -2.63 ± 0.05‰ (Figure IV–2a). Silicate melt compositions in these experiments vary from -0.57 ± 0.12‰ to -0.79 ± 0.03‰. The metal and silicate melt compositions of the 1450 °C experiment in an MgO capsule (RH51) fall in the same range with -2.25 ± 0.04‰ and -0.77 ± 0.04‰, respectively. Experiments performed at 1750 °C have metal melt compositions from -2.19 ± 0.05‰ to -1.70 ± 0.08‰ for MgO capsules (Figure IV–2b). The corresponding silicate melts have compositions between -0.93 ± 0.04‰ and -0.61 ± 0.14‰. The experiment in a graphite capsule at 1750 °C (RH58) resulted in similar metal and silicate melt compositions of -1.92 ± 0.08‰ and -0.78 ± 0.04‰, respectively. The metal fractions of experiments at 1750 °C are thereby lighter than the silicate fractions they equilibrated with, which was also observed in the experiments at 1450 °C. Taking into account the 29/28Si ratios determined for all samples, δ30Si compositions of all phases can be plotted against their δ29Si compositions (Figure IV–3). All these data plot within uncertainty of the mass dependent fractionation line.

0 0 a Metal b Bulk Sn-free (MgO capsule) -0.4 Silicate -0.4 -0.8 -0.8

-1.2 -1.2 Si (‰)

30 -1.6 -1.6 δ -2.0 -2.0

-2.4 -2.4 1450 °C 1750 °C -2.8 -2.8 0 1 2 3 4 20 24 0 1 2 Run duration (h) Run duration (h) Figure IV–2. Silicon isotope compositions of experimental run products and bulk materials. Panel (a) shows data for 1450 °C, while data for 1750 °C are presented in panel (b). Measured δ30Si compositions of silicates are shown as circles, metals as diamonds, and the calculated δ30Si compositions of the bulk run products as squares. Open symbols are for experiments performed in graphite capsules, closed symbols for experiments in MgO capsules. Half-filled symbols in panel (b) are for Sn-free experiments in MgO capsules. The line represents the measured bulk δ30Si composition of Mix1, where the shade indicates the uncertainty. All errors are 95% confidence intervals. Page 102 — Chapter IV

-0.5 Si (‰) 29 δ -1.0

-1.5 -3.0 -2.5 -2.0 -1.5 -1.0 -0.5 0 δ30Si (‰) Figure IV–3. Plot of δ29Si against δ30Si compositions. All data plot within uncertainty of the curve that represents the terrestrial, equilibrium mass dependent fractionation line (Young et al., 2002). Symbols for silicates and metals are as in panel (a). The square and up- and downward pointing triangles represent, respectively, the measured bulk composition of Mix1, the measured composition of the metal used for starting mixtures, and the measured composition of the SiO2 used for starting mixtures. Error bars in both panels are 95% confidence intervals.

IV–4. EQUILIBRIUM ISOTOPE FRACTIONATION

The δ30Si isotope composition of the silicate melt has changed by approximately +1.3 to +1.5‰ from -2.12‰ in the starting material to about -0.6 to -0.8‰ in all experiments. The metal melt compositions have changed in the opposite direction: from -0.38‰ to at least -1.7‰. This indicates that the metal and silicate melts have chemically interacted. The resulting lighter metal than silicate composition is expected from theoretical calculations and is overall consistent with previous investigations (Shahar et al., 2011; Shahar et al., 2009; Ziegler et al., 2010). Bulk Si isotope compositions of each experiment were calculated from the measured metal and silicate compositions by mass balance to assess whether the experiments were closed systems with respect to Si isotopes (Table IV–4): 30 M 30 Sil 30 dSiBulk=⋅ f Si dd SiM +⋅ f Si SiSil (IV.6)

where fSi refers to the mass fraction of Si, derived from the bulk metal and silicate fractions of the starting mixtures and the Si contents of the metal and silicate phases determined by electron microprobe analyses. The sub- and superscripts M and Sil indicate metal and silicate, respectively. Bulk mass fractions of metal and silicate, determined by regression analyses with major element cations, indicate that the bulk metal fraction in experiments performed at 1450 °C has changed by <2 wt% compared to the starting mixture. For experiments Chapter IV — Page 103

2 (‰) 0.11 0.11 0.09 95% c.i.

3 (‰) 0.15 BulkStart-Final -0.19 -0.16 = oxygen fugacity 2 Si 30 Δ

2 (‰) 0.14 0.06 0.15 95% c.i.

M-Sil Si (‰) -1.38 -1.84 -1.49 30 Δ n 30 18 16 24 16 31 18 13 10 42 19

(‰) 0.03 0.04 0.05 0.02 0.02 0.05 0.03 0.03 0.04 0.06 0.02 0.02 0.03 0.02 0.05 0.03 95% c.i.

Si 29 (‰) δ -0.58 -1.22 -0.75 -1.00 -0.32 -0.39 -0.29 -0.59 -1.37 -1.04 -0.49 -0.70 -0.65 -0.70 -1.09 -0.19

(‰) 0.11 0.06 0.08 0.03 0.08 0.03 0.12 0.06 0.05 0.09 0.08 0.03 0.03 0.04 0.09 0.04 95% c.i.

Si 30 (‰) δ -0.60 -1.13 -2.34 -1.47 -1.98 -1.16 -2.63 -0.79 -0.57 -2.05 -1.36 -1.03 -1.32 -1.36 -0.38 -2.12

C C C type Capsule

(g) 434 702 RCF

t (h) 1.5 0.5 22.0

2 O f -4.11 -3.66 -4.03 (ΔIW)

T (°C) 1450 1450 1450 1

P 1.00 (GPa) 0.89±0.03 0.93±0.01 Isotopic compositions and experimental run parameters. P = pressure (GPa), T = temperature ( °C), fO = temperature (GPa), T = pressure run parameters. P Isotopic compositions and experimental 5 Bulk calc. Metal Metal Silicate Metal Silicate Bulk calc. Silicate Bulk calc. Metal 4 6 6 Si 2 94 RH37 RH32 RH31 RH27 Samples Mix1, Sn-bearing Start Mix2 calc. Start Mix1 calc. Start Mix2 SiO Start Mix1 Sample Fe Starting powders Table IV–4. Table to relative the buffer iron-wüstite (ΔIW), t RCF = = centrifugal in run relative force multiples duration of in the hour, gravity constant analyses, capsule type C = graphite. (g), n = number of repeated Page 104 — Chapter IV

2 (‰) 0.11 0.11 0.10 0.09 0.10 0.09 95% c.i.

3 (‰) 0.11 0.06 0.05 BulkStart-Final -0.06 -0.14 -0.26 Si 30 Δ

2 (‰) 0.08 0.06 0.09 0.08 0.06 0.16 95% c.i.

M-Sil Si (‰) -1.58 -1.48 -1.15 -0.93 -1.27 -1.09 30 Δ n 32 57 32 52 41 69 33 39 24 18 14

(‰) 0.02 0.03 0.02 0.02 0.02 0.04 0.02 0.05 0.03 0.02 0.04 0.02 0.03 0.03 0.04 0.07 0.03 95% c.i.

Si 29 (‰) δ -0.37 -0.70 -1.16 -0.39 -0.70 -1.02 -0.41 -0.67 -0.89 -0.42 -0.60 -1.13 -0.46 -0.73 -0.86 -0.29 -0.52

(‰) 0.03 0.04 0.04 0.04 0.02 0.08 0.04 0.06 0.06 0.04 0.05 0.05 0.04 0.03 0.08 0.14 0.07 95% c.i.

Si 30 (‰) δ -0.76 -1.38 -2.25 -0.77 -1.37 -1.92 -0.78 -1.26 -1.75 -0.82 -1.18 -2.19 -0.93 -1.43 -1.70 -0.61 -1.06

C C type MgO MgO MgO MgO Capsule

(g) 293 486 245 354 355 697 RCF

t (h) 4.0 4.0 1.5 1.5 1.5 0.5

2 O f -3.23 -4.53 -4.24 -4.79 -4.82 -5.14 (ΔIW)

T (°C) 1450 1450 1750 1750 1750 1750 1

P (GPa) 0.84±0.04 0.83±0.08 0.58±0.21 0.63±0.27 0.54±0.12 0.64±0.07 Silicate Bulk calc. Metal Silicate Bulk calc. Metal Silicate Bulk calc. Metal Silicate Bulk calc. Metal Silicate Bulk calc. Metal Silicate Bulk calc. Errors are 2SD and refer to the time dependent variation of pressure in centrifuge experiments. Sample RH51 RH58 RH66 Mix2, Sn-free RH65 RH67 1 They are calculated as the 2SD of pressure registrations in 30 s intervals our log files (see text for more explanation). Table IV–4, continued. Table Chapter IV — Page 105

Footnotes Table IV–4, continued: 2 Propagated according to the following equation:

30 2230 95% ci. .=  95% ci. .ddSi + 95% ci. . Si ()M ()Sil 3 This refers to the difference in δ30Si of the bulk starting material (measured) and the bulk δ30Si calculated from the metal and silicate analyses after the experiments. The measured starting composition of starting Mix1 is used for all these calculations because of the deviant measurement for starting Mix2 (see text). 4 This metal composition was created by mixing pure Fe metal powder with Fe83Si17 metal powder (purchased from Goodfellow) 5 Calc. refers to calculation of a bulk isotopic composition from the analysed metal and analysed silicate compositions by mass balance. Errors for these calculated values are propagated from errors for metal and silicate compositions as well as weighing uncertainties for mass fractions. 6 Experiment RH27 was run in a static piston cylinder press, not in the centrifuging piston cylinder press. at 1750 °C in MgO capsules, the bulk metal fraction has decreased by about 6 wt% to approximately 69 wt% due to dissolution of capsule-derived MgO. A comparison of the bulk Si isotope compositions of each experiment with that of the starting mixtures indicates that not all experiments are balanced for Si isotopes (Figure IV–2a, b). Six of the nine experiments have bulk δ30Si isotope compositions that differ from the bulk starting compositions by 0.11 ± 0.09‰ to 0.26 ± 0.11‰ (95% c.i.). Both light and heavy isotopes appear to have been lost (Table IV–3). A cause for the differences in bulk δ30Si isotope compositions between starting mixtures and experimental run products seems difficult to find, as previously observed in Hin et al. (2012) for Fe isotopes. The absence of Si isotope mass balance seems rather random: heavy as well as light isotopes are lost, and loss occurs in graphite as well as MgO capsules and at 1450 °C as well as 1750 °C (Table IV–3). There is no apparent system in this behaviour, which implies that it may not be related to formation of up to 7 vol% olivine in MgO capsules or to the involvement of SiC as an additional phase in graphite capsules. Moreover, phase relations in the ternary system Fe-Si-C indicate that SiC is not stable at concentrations below ~17 wt% Si at atmospheric pressure (Lacaze and Sundman, 1991) or below ~12 wt% Si at 8 GPa (Kocherahinski and Kulik, 1996) compared to 4.74-5.83 wt% Si in our experiments at 1 GPa. Oxygen fugacities below ΔIW-7.4 (1450 °C) or ΔIW -5.9 (1750 °C) are required to form SiC by reduction of silicate (Ulmer et al., 1998). The only cause that may explain the observed open system behaviour seems inhomogeneity in the metal and silicate fractions by poor mechanical mixing. The aliquots loaded in the capsules may thus contain different metal fractions. Our previous experience with mechanical mixing during the preparation of starting powders indicates that this usually works well, even for material with density differences as large as for metal and silicate. However, mechanical heterogeneity in the starting mixtures may be confirmed by the dif- Page 106 — Chapter IV ference between the calculated and measured bulk compositions of Mix2. The fractionation factors between metal and silicate melt vary little at each temperature (Table IV–4 and Figure IV–4). The experiments performed at 1450 °C are within analytical uncertainty identical to each other, except for experiment RH31. The run duration of this experiment was shortest with 30 minutes. Silicon diffusivities in our metal and silicate melts at 1450 °C are in the order of 2×10-9 (Saito and Maruya, 1958) and 1×10-10 m2 s-1 (Guillot and Sator, 2007), respectively, leading to diffusion lengths of about 400 μm in the diffusion limiting silicate liquid after 30 minutes. We therefore conclude that this experiment had not yet reached equilibrium. Mass dependent stable isotope fractionation between metal and silicate melt is independent of presence or absence of C, as indicated by a direct comparison of C-bearing and C-free systems (RH37 and RH51, respectively). Similar observations were previously made for Si by Shahar et al. (2011; 2009) and for Fe by Hin et al. (2012). Excluding the experiment that had not yet reached equilibrium, we determine an average equilibrium metal-silicate melt fractionation factor 30 (Δ SiMetal-Silicate) of -1.48 ± 0.08‰ (95% c.i.) at 1450 °C. Experiments performed at 1750 °C scatter more in their fractionation factors

-0.6

MgO capsule -0.8 Graphite capsule

-1.0

-1.2

(‰) Average 1750°C

-1.4 Metal−Silicate Si

30 -1.6 Average 1450°C ∆ -1.8

-2.0

-2.2 0 1 2 3 4 20 24 Run duration (h) Figure IV–4. Silicon isotope fractionation factors between metal and silicate plotted against experimental run duration. Open symbols represent experiments in graphite capsules. Closed symbols are for experiments in MgO capsules, while half-filled symbols represent Sn-free in MgO capsules. Only the experiment run for 0.5 hours at 1450 °C is significantly different from the indicated averages. Chapter IV — Page 107

(Table IV–4 and Figure IV–4). This scatter may be related to greater difficulty of these experiments, as detailed in section 3.1. A direct comparison between Sn-bearing and Sn-free systems does not lead to a clear conclusion. The best comparison should be between experiments RH65 and RH66, both of which were run in MgO capsules for 1.5 h. The Sn-free experiment (RH65) has a 0.33 ± 0.10‰ smaller fractionation factor than the Sn-bearing one (RH66), implying that the presence of Sn possibly reduces the fractionation of Si isotopes between metal and silicate liquids. Unfortunately, however, experiment RH66 had a crack in the capsule (Figure IV–1e), while of all experiments RH65 appears to have suffered the largest loss of Si from the bulk system. A comparison between RH65 and RH58 (Sn-bearing, 1.5 h in graphite) implies that there would not be an effect of Sn on Si isotope fractionation, because these two experiments have within analytical uncertainty identical fractionation factors. Furthermore, the fractionation factor experiment RH67 (Sn-free, 0.5 h in MgO) is smaller than that of RH65, but its analytical uncertainty is so large that it is within uncertainty of all experiments performed at 1750 °C. The comparison between Sn-free and Sn-bearing systems therefore remains inconclusive. An average of all four 30 experiments yields an equilibrium Δ SiMetal-Silicate of -1.11 ± 0.14‰. The average fractionation factors at the two experimental temperatures can be used to anchor the functional form of Si isotope fractionation between metal and silicate melt. Following Bigeleisen and Mayer (1947) and Urey (1947), mass dependent stable isotope fractionation can be described with the simplified equation:

Dm − DF D≈ij AB− ⋅1000 (IV.7) AB− 2 mmij T

in which ΔA-B is the fractionation factor of an isotope ratio (as ‰ deviation from a reference ratio) between phases A and B, m the atomic mass of the isotopes i and j of the element of interest, Δm the difference in masses of the isotopes i and j used for the isotope ratio, ΔFA-B the difference in force constants acting on the isotopes in phases A and B, and T the temperature in K. In this formalism, temperature is thus the only variable that controls fractionation between two phases A and B, meaning that the right-hand side of equation (IV.7) can be simplified to C D= (IV.8) AB− T 2 30 where C is a constant. The Δ SiMetal-Silicate fractionation factors determined above then yield a value of C in equation (IV.8) equal to -4.40(±0.24)×106 (95% c.i.) for 1450 °C and -4.54(±0.57)×106 (95% c.i.) for 1750 °C, or an average of -4.47(±0.31)×106. The values at 1450 and 1750 °C are within uncertainty identical, i.e. the results of our experiments concur with the functional form of equation (IV.7) (Figure IV–5). Page 108 — Chapter IV

IV–5. DISCUSSION

-4.0

-3.5 -4.47(±0.31) 106/T2

-3.0

-2.5 (‰)

-2.0 Metal−Silicate Si

30 -1.5 ∆

-1.0

-0.5

0 800 1000 1200 1400 1600 1800 2000 2200 2400 2600 Temperature (°C) Figure IV–5. Temperature dependence of Si isotope fractionation. The two diamonds indicate the average fractionation factors determined for 1450 °C and 1750 °C. The stippled line is the error envelope for the temperature curve (95% confidence interval). IV–5.1. Comparison with previous work Our study demonstrates that there is measurable Si isotope fractionation between equilibrated metal and silicate liquids. The fractionation factor of -1.11 ± 0.14‰ at 1750 °C as well as the average value of -4.47 ± 0.31×106 for the constant describing the temperature dependence can directly be compared to the works of Shahar et al. (2011; 2009). Those two studies presented data for 1800 °C instead of 1750 °C and should therefore have found fractionation factors at their temperature of 1800 °C that are about 0.05‰ smaller than those found in 30 our study at 1750 °C. However, Shahar et al. (2009) found a Δ SiMetal-Silicate of -1.87 ± 0.14‰ (2SE), which implies a 0.76 ± 0.16‰ larger difference between metal and silicate melt than our result for 1750 °C. In a more comprehensive study, Shahar et al. (2011) reported a fractionation factor of -1.77 ± 0.32‰ (2SE) at 1800 °C and concluded that this confirmed the earlier work of Shahar et al. (2009). Consequently, there is nearly a factor two difference in the constant that Chapter IV — Page 109 describes the temperature dependence: -7.45(±0.41)×106 in Shahar et al. (2011) compared to our value of -4.47(±0.31)×106 (95% c.i.). Despite these differences, our study as well as those of Shahar et al. (2011; 2009) indicate that the presence or absence of C does not affect equilibrium stable isotope fractionation beyond analytical uncertainties. This is in accordance with the independence of Fe isotope fractionation on C or S (Hin et al., 2012; Poitrasson et al., 2009) as well as with theoretical considerations that valence state, coordination state and type of bond determine equilibrium mass dependent fractionation. There are several potential explanations for the deviations between the studies of Shahar et al. (2011; 2009) and our study. In their first study, Shahar et al. (2009) attempted to equilibrate liquid metal with liquid silicate at 1 GPa and 1800 °C in graphite capsules. However, their Si isotope fractionation factor did not stabilise with increasing experimental run duration. In addition, their experiments were not balanced for Si isotopes, because they showed an apparent loss of light Si isotopes that grew to ~0.35‰ on δ30Si with increasing run duration from one to 30 minutes. In their follow-up study, Shahar et al. (2011) attempted to solve these issues with longer run durations up to 60 minutes at 1800 °C and with experiments at temperatures of 2000 and 2200 °C at a pressure of 7 GPa. They employed both graphite and MgO capsules, but they did not report their laser ablation MC-ICPMS analyses on experiments in graphite capsules because according to the authors these analyses proved untrustworthy, possibly due to matrix effects. Their experiments in MgO capsules, however, crystallised large volumes of olivine, so that Si isotope fractionation factors had to be determined between metal liquid and olivine. As indicated in Shahar et al. (2011), Si self-diffusion is slow in olivine and this could have led to 0.3 μm diffusion distances under their experimental run conditions. This raises some doubts about how representative their average Si isotope fractionation factors are for equilibrium fractionation. Nevertheless, a similar temperature dependence of -7.64(±0.47)×106/T2 for Si isotope fractionation was reported by Ziegler et al. (2010). They performed solution and laser ablation MC-ICPMS analyses of Si isotope compositions in metal and silicate fractions from the aubrites (enstatite achondrites) Mount Egerton and Norton County. The final causes for the deviations between the results of Ziegler et al. (2010) and Shahar et al. (2011; 2009) compared to our results remain unclear. Another comparison of our work could be made with the molecular dynamics calculations presented in Georg et al. (2007) and Shahar et al. (2009). These calculations, performed with solid Fe3Si and forsterite or ringwoodite as analogues for metal and silicate melts, originally resulted in a value of about -5.5×106 to -5.8×106 for the constant in equation (IV.8). These values are lower than the value of -7.45(±0.41)×106 determined experimentally in Shahar et al. (2009), but only slightly higher than our experimentally determined value of -4.47(±0.31)×106. However, Shahar et al. (2009) indicated that the uncertainties Page 110 — Chapter IV of molecular dynamics calculations are rather large, and therefore anchored them to their experimentally determined value.

IV–5.2. Implications for core formation Several studies have concluded that the Mg/Si ratio of the bulk silicate Earth (BSE) is 1.045 ± 0.014, which is higher than the value of 0.90 ± 0.03 of the Solar System as derived from carbonaceous chondrites that could represent the Bulk Earth composition (Palme and O’Neill, 2003). Loss of Si relative to Mg by volatilisation during Earth’s accretion (Ringwood, 1979) as well as by incor- poration of Si into the Earth’s core (Ringwood, 1959) have been proposed as potential explanation for this difference. Experimental determination of Si partitioning between metal and silicate liquids has shown that Si could be the most abundant light element in Earth’s core (Corgne et al., 2008; Kilburn and Wood, 1997; Mann et al., 2009), explaining partly its ~7% density deficit (Anderson and Isaak, 2002; Birch, 1952) as well as the relatively high Mg/Si ratio of the BSE. 30 The Si isotope composition of the core (δ SiC) as well as the Bulk Earth 30 (δ SiBE) can be calculated using the Si isotope fractionation factor between metal and silicate liquids determined in our study, if the δ30Si of the BSE is known: 30 CC30 30 ddSiBE= f Si ⋅()() SiC +−1 fSi ⋅ d SiBSE (IV.9) with 30 30 30 ddSiC = D+ SiMetal− Silicate SiBSE (IV.10) in which the sub- and superscripts BE and C refer to the Bulk Earth and core, respectively. The BSE Si isotope composition seems well defined from measurements on basalts, peridotites and pyroxenites, with average δ30Si values of -0.29 ± 0.02‰ (95% c.i.) in Fitoussi et al. (2009), -0.29 ± 0.01‰ in Savage et al. (2010) and -0.32 ± 0.03‰ in Armytage et al. (2011). Notably, Chakrabarti and Jacobsen (2010) present a significantly different BSE Si isotope composition of -0.38 ± 0.02‰, but the composition they report for the BHVO standard is different from all other studies that are in agreement, which raises doubt about the accuracy of this data set. Averaging the data of Fitoussi et al. (2009), Savage et al. (2010) and Armytage et al. (2011) results in a BSE δ30Si isotope composition of -0.29 ± 0.01‰ (95% c.i.). In equation (IV.10), the metal-silicate fractionation factor is temperature dependent and the temperature of Si isotope equilibration during core formation on Earth probably varied and is uncertain. A temperature of 2500 to 3500 K for metal-silicate equilibration during core segregation was determined from models 30 based on elemental partitioning data (Wade and Wood, 2005). Using a δ SiBSE of -0.29 ± 0.01‰, we calculate a δ30Si of -1.01 ± 0.05‰ and -0.66 ± 0.03‰ for Earth’s core at temperatures of 2500 and 3500 K, respectively (Table IV–5). The Si concentration in the core remains the only unknown in calculating the Bulk Earth Si isotope composition with equation (IV.9). Earth’s density deficit, Chapter IV — Page 111 however, would not allow for Table IV–5. Calculated Si isotope composi- 30 more than 12.3 wt% Si (Anderson tions for the Earth’s core (δ SiC) and the Bulk Earth (δ30Si ), given metal-silicate equilibra- and Isaak, 2002). Moreover, con- BE sidering the volatility trend and a tion temperatures of 2500 and 3500 K. Two scenarios are presented for the Bulk Earth, variety of elements in a chondritic one assuming a Si concentration of 6 wt% in reference model, compensation of Earth’s core and one assuming 8 wt%. Com- 30 the core density deficit requires positions are calculated for a δ SiBSE of -0.29 about 6 wt% Si (McDonough, ± 0.01‰. 2003). Furthermore, about 7-8 wt% of Si in the Earth’s core com- 2500 K 3500 K 30 -1.05 ± 0.05 -0.66 ± 0.03 pensates for the high Mg/Si of δ SiC (‰) 30 the BSE relative to carbonaceous δ SiBE 6 wt% (‰) -0.38 ± 0.02 -0.33 ± 0.01 chondrites (Allègre et al., 1995; 30 -0.40 ± 0.02 -0.35 ± 0.02 δ SiBE 8 wt% (‰) Palme and O’Neill, 2003). Con- siderations of densities and compressional sounds velocities of various Fe alloys with e.g. Si indicate that 8 wt% Si in Earth’s core would be an upper limit (Badro et al., 2007; Sherman, 1997). We can therefore assume an Si content of 6-8 wt% for Earth’s core to determine the δ30Si composition for the Bulk Earth. For 6 wt% Si in Earth’s core we thus calculate a Bulk Earth Si isotope composition of -0.38 ± 0.02‰ or -0.33 ± 0.01‰ for equilibration temperatures of 2500 or 3500 K, respectively. If an Si content of 8 wt% is taken for the Earth’s core, the calculated δ30Si of the Bulk Earth becomes -0.40 ± 0.02‰ or -0.35 ± 0.02‰ for 2500 or 3500 K, respectively (Table IV–5). The calculated Bulk Earth δ30Si values can be compared to the Si isotope compositions of chondrites. Unfortunately, the average composition of chondrites is disputed. Fitoussi et al. (2009) reported carbonaceous chondrites to be on average 0.06 ± 0.04‰ heavier than the average δ30Si isotope composition of ordinary chondrites (-0.36 ± 0.02‰ compared to -0.42 ± 0.03‰, respectively; 95% confidence intervals, see Table IV–A.2 at the end of this chapter). One enstatite chondrite was found to have a δ30Si of -0.55 ± 0.04‰. Using a larger set of samples, Armytage et al. (2011) and Fitoussi and Bourdon (2012) confirmed the relatively light Si isotope composition of enstatite chondrites (-0.62 ± 0.05‰, 95% c.i.). However, the averages of Armytage et al. (2011) for ordinary and carbonaceous chondrites of -0.46 ± 0.03‰ and -0.48 ± 0.04‰, respectively, are indistinguishable. Chakrabarti and Jacobsen (2010) made similar observations: a relatively light δ30Si for one enstatite chondrite (-0.56 ± 0.05‰) and indistinguishable ordinary and carbonaceous chondrites with a total average of -0.40 ± 0.02‰ (95% c.i.). Of the various Si isotope compositions reported for chondrites, only the carbonaceous chondrite average of Fitoussi et al. (2009) is indistinguishable from the Bulk Earth compositions calculated for equilibration temperatures of 3500 K. If Si isotopes equilibrated during core formation at 2500 K, the calculated Page 112 — Chapter IV composition of the Bulk Earth would be indistinguishable from the ordinary and carbonaceous chondrite averages of Fitoussi et al. (2009) as well as from the ordinary/carbonaceous chondrite average of Chakrabarti and Jacobsen (2010). If the ordinary/carbonaceous chondrite average of Armytage et al. (2011) is considered representative of the Bulk Earth Si isotope composition, metal-silicate equilibration temperatures of ≤2100 K are required to match the calculated Bulk Earth δ30Si value. Such temperatures are lower than those of 2500-3500 K determined from core formation models for Earth based on elemental partitioning data (Corgne et al., 2008; Mann et al., 2009; Wade and Wood, 2005). If the ordinary/ carbonaceous chondrite average of Armytage et al. (2011) represents the Bulk Earth, this could thus indicate that Si isotope fractionation in Earth records a temperature of metal-silicate equilibration averaged over various temperature conditions. Such an interpretation, however, relies on the use of 95% confidence intervals as calculated by equation (IV.2). Given the dispute about the chondrite compositions, it might be more appropriate to use 2SD uncertainties. In that case, all the above mentioned δ30Si values for chondrites are indistinguishable of the calculated Bulk Earth compositions for 2500 as well as 3500 K, with exception of the enstatite chondrite compositions that would still require temperatures below 2200 K if 2SD uncertainties are used for both the BSE and the enstatite chondrite averages (-0.29 ± 0.08‰ and -0.60 ± 0.09‰, respectively). Although enstatite chondrites are thus unlikely to represent the Bulk Earth Si isotope composition, as previously concluded by Fitoussi and Bourdon (2012), a better determination of ordinary and carbaceous chondrite Si isotope compositions is required to determine whether Si isotopes agree with core formation on Earth in a chondritic reference model at temperatures ≥2500 K.

IV–6. CONCLUSION

Our experimental study, employing a centrifuging piston cylinder, quantifies equilibrium fractionation of Si isotopes between liquid metal and liquid silicate at pressures of 1 GPa. At temperatures of 1450 °C, metal liquid is 1.48 ± 0.08‰ lighter than silicate liquid, while this difference is 1.11 ± 0.14‰ at 1750 °C. 2 30 These observations concur with the theoretical 1/T dependence: Δ SiMetal-Silicate (‰) = -4.47(±0.31)×106/T2. The relation between the Si isotope fractionation factor and temperature can be used to calculate Bulk Earth Si isotope compositions, using the δ30Si of the Bulk Silicate Earth and assuming 6 or 8 wt% Si in Earth’s core. For metal-silicate equilibration temperatures of 2500 and 3500 K, we determine Bulk Earth Si isotope compositions of, respectively, -0.38 ± 0.02‰ and -0.33 ± 0.01‰ (95% c.i.) for 6 wt% Si in Earth’s core, or -0.40 ± 0.02‰ and -0.35 ± 0.02‰ 8 wt% Si. Although enstatite chondrites probably have too light Si isotope compositions to be representative of such Bulk Earth values, a further comparison of the calcu- Chapter IV — Page 113 lated Bulk Earth values with reported chondrite compositions is limited due to the variation in reported chondrite compositions.

Acknowledgements We are greatly indebted to Bruno Zürcher and Felix Oberli for their maintenance works on the centrifuging piston cylinder and Nu1700, respectively. Page 114 — Chapter IV 3.3 1.5 n.d. 0.07 0.37 1.81 0.13 2SD 1 n.d. 0.64 0.51 92.7 5.20 1.81 0.05 99.8 1750 -5.14 MgO RH67 1.7 0.27 17.2 20.0 1.23 1.09 0.42 2SD 1 1.5 0.63 80.4 13.4 4.74 1.65 0.25 1750 -4.79 MgO 100.5 RH66 1.4 1.2 n.d. 0.12 0.14 0.48 0.05 2SD 1 1.5 n.d. 0.54 93.5 4.93 1.53 0.03 1750 -4.82 MgO 100.0 RH65 4.9 2.5 0.21 9.29 1.86 1.89 0.82 2SD - ite 1.5 2.6 0.58 82.8 5.86 5.83 1.46 1.51 97.4 1750 -4.24 RH58 Graph 3.8 1.1 0.08 4.82 0.24 0.22 0.08 2SD 4.0 0.83 85.0 8.84 5.56 1.33 0.03 1450 -4.53 MgO 100.7 RH51 1.4 0.6 0.04 2.68 0.45 0.95 2SD - ite 4.0 0.8 n.d. 0.84 82.3 8.19 5.45 3.25 99.2 1450 -3.23 RH37 Graph 2.0 0.7 0.03 2.42 0.35 0.53 2SD - ite 1.5 1.2 n.d. 0.89 81.7 8.86 5.28 2.96 98.8 1450 -3.66 RH32 Graph 0.7 0.9 0.01 1.54 0.27 0.35 0.05 2SD - ite 2.7 0.93 0.52 81.7 9.12 5.02 1.45 0.04 97.3 1450 -4.11 RH31 Graph 1.7 0.2 1.44 0.35 0.10 0.56 2SD - Elemental compositions (wt%) of all experimental products. Pressures (P), temperatures (T), run durations (t) and oxygen (P), temperatures Pressures products. compositions (wt%) of all experimental Elemental ) of the experiments are indicated as well. ) of the experiments are ite 22 1.0 3.9 2 79.5 6.92 5.31 1.34 3.06 96.1 1450 -4.03 RH27 Graph (ΔIW) 2 O Table IV–A.1 Table fugacity (fO P (GPa) P (°C) T t (h) f Capsule type Metal Fe Sn Si Mo Pt Total C (calc) Chapter IV — Page 115 1.3 0.4 n.d. 0.79 2.86 3.69 0.05 0.23 0.12 2SD 1 n.d. 0.07 10.7 1.45 0.38 97.0 37.74 10.34 36.30 RH67 1.4 0.5 1.36 2.80 2.98 0.09 0.14 0.08 0.02 2SD 1 0.08 12.3 1.77 0.48 0.02 97.1 11.15 34.90 36.47 RH66 3.1 0.6 n.d. 1.88 4.08 5.76 0.06 0.49 0.15 2SD 1 n.d. 0.11 9.68 10.8 1.54 0.42 97.1 35.40 39.13 RH65 0.3 0.6 0.64 0.34 0.35 0.12 0.10 0.03 0.04 2SD 0.14 13.1 1.59 0.43 0.01 99.4 56.15 13.71 14.31 RH58 0.2 0.8 0.54 0.33 0.40 0.09 0.12 0.05 0.04 2SD 0.12 12.1 1.60 0.46 0.01 98.7 48.43 12.29 23.72 RH51 0.8 0.67 0.55 0.16 0.26 0.55 0.08 0.10 0.02 2SD 0.11 7.67 1.58 0.46 0.01 94.8 56.93 13.79 14.20 RH37 0.8 0.46 1.25 0.93 0.30 0.42 0.13 0.06 0.03 2SD 0.09 7.00 1.61 0.47 0.01 95.0 57.77 13.48 14.58 RH32 0.3 0.8 0.96 1.76 1.66 0.19 0.13 0.05 0.03 2SD 0.16 12.5 1.62 0.43 0.02 57.65 13.44 15.25 101.1 RH31 0.7 3.4 2.11 0.58 0.55 0.19 0.23 0.04 0.03 2SD 0.14 12.3 1.55 0.43 0.02 99.2 56.51 13.80 14.41 RH27 3 2 2 O O 2 O 2 2 Based on a weighted average (weight from estimated volume fraction) of quenched olivine and glass Table IV–A.1, continued. Table Silicate SiO Al MgO FeO CaO Na K SnO Total 1 Page 116 — Chapter IV Si) 29 0.04 0.03 0.03 0.04 0.03 0.04 0.04 0.10 2SD (δ Si) 30 0.06 0.07 0.07 0.05 0.05 0.04 0.06 0.17 2SD (δ 2 3 3 1 3 4 1 12 18 19 12 25 15 15 n= n= n= n= n= n= n= n= n= n= n= n= n= n= Comments Reference Fitoussi et al. (2009) Fitoussi et al. (2009) Fitoussi et al. (2009) Fitoussi et al. (2009) Fitoussi et al. (2009) Fitoussi et al. (2009) Fitoussi et al. (2009) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Armytage et al. (2011) Armytage et al. (2011) Chakrabarti and Jacobsen (2010) Chakrabarti and Jacobsen (2010) Chakrabarti and Jacobsen (2010) Chakrabarti and Jacobsen (2010) c.i. 0.01 95% 0.03 0.02 0.02 0.04 0.02 0.02 0.04 0.04 0.06 0.05 0.04 0.03 0.08 0.03 0.03 0.02 0.02 0.03 0.02 0.03 0.03 0.05 2SD/SE Si 29 -0.16 -0.12 -0.16 -0.12 -0.15 -0.15 -0.13 -0.13 -0.16 -0.16 -0.17 -0.17 -0.18 -0.20 -0.21 -0.20 -0.16 -0.18 -0.19 -0.20 -0.19 -0.17 δ c.i. 0.02 95% 0.11 0.04 0.04 0.04 0.05 0.03 0.02 0.06 0.08 0.07 0.14 0.06 0.06 0.16 0.05 0.04 0.04 0.06 0.05 0.04 0.06 0.10 2SD/SE Si 30 -0.30 -0.26 -0.31 -0.26 -0.31 -0.28 -0.24 -0.29 -0.31 -0.29 -0.33 -0.35 -0.34 -0.37 -0.39 -0.40 -0.30 -0.35 -0.37 -0.41 -0.39 -0.33 δ Literature compilation terrestrial silicates and meteorites. compilation terrestrial Literature Name WR18 RH07 RH08 K1 K2 K3 Montana del Fuego Kilbourne Hole San Carlos C235A H93 X1 H93 X8 BD730 BD1542 J30 C235A San Carlos olivine ALM2 SC Ol SC Opx SC Cpx Peridotites Average Table IV–A.2 Table Type Peridotites Chapter IV — Page 117 Si) 29 0.11 0.02 0.06 0.05 0.07 0.07 2SD (δ Si) 30 0.11 0.04 0.06 0.07 0.26 0.19 2SD (δ 4 2 3 14 14 14 n= n= n= n= n= n= Comments Reference Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Fitoussi et al. (2009) Fitoussi et al. (2009) Fitoussi et al. (2009) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) c.i. 0.07 95% 0.04 0.06 0.02 0.06 0.01 0.04 0.03 0.04 0.05 0.03 0.04 0.03 0.02 0.06 0.01 0.02 0.06 0.03 0.04 0.02 0.01 0.06 2SD/SE Si 29 -0.16 -0.16 -0.21 -0.18 -0.15 -0.16 -0.17 -0.15 -0.14 -0.13 -0.16 -0.12 -0.13 -0.13 -0.13 -0.16 -0.12 -0.13 -0.17 -0.15 -0.14 -0.18 δ c.i. 0.14 95% 0.11 0.07 0.06 0.10 0.02 0.04 0.04 0.09 0.10 0.07 0.09 0.01 0.05 0.07 0.03 0.05 0.02 0.02 0.04 0.03 0.02 0.07 2SD/SE Si 30 -0.30 -0.29 -0.39 -0.33 -0.29 -0.32 -0.32 -0.27 -0.27 -0.28 -0.30 -0.24 -0.26 -0.23 -0.27 -0.29 -0.25 -0.25 -0.33 -0.25 -0.28 -0.30 δ Name BD744 J25 J27 Pyroxinites Average A127 D22-5 BHVO-2 Koolau BHVO-1 BHVO-2 BCR-1 D1-5 D5-5 D15-1 D22-1 D26-5 D27-5 R82-1 R94-2 VE23-5 VE32-3 MD57 D13-7 MD34 D4 Table IV–A.2, continued. Table Type Pyroxinites Terrestrial basalts Page 118 — Chapter IV Si) 29 0.11 0.11 0.04 0.04 0.07 0.08 0.07 0.18 0.14 0.16 0.19 0.10 0.08 0.05 0.04 0.10 0.12 2SD (δ Si) 30 0.11 0.08 0.12 0.19 0.15 0.27 0.29 0.40 0.19 0.14 0.13 0.09 0.08 0.26 0.25 0.19 0.15 2SD (δ 15 15 13 15 14 80 51 64 15 23 18 22 4 28 26 15 14 n= n= n= n= n= n= n= n= n= n= n= n= n= n= n= n= n= Comments Reference Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Savage et al. (2010) Armytage et al. (2011) Armytage et al. (2011) Armytage et al. (2011) Armytage et al. (2011) Chakrabarti and Jacobsen (2010) Chakrabarti and Jacobsen (2010) Chakrabarti and Jacobsen (2010) Chakrabarti and Jacobsen (2010) c.i. 0.01 0.01 95% 0.01 0.01 0.02 0.02 0.02 0.02 0.02 0.02 0.05 0.02 0.02 0.01 0.02 0.02 0.02 0.03 0.03 0.04 0.05 2SD/SE Si 29 -0.15 -0.12 -0.14 -0.13 -0.14 -0.13 -0.14 -0.16 -0.15 -0.14 -0.16 -0.16 -0.16 -0.20 -0.19 -0.17 -0.20 -0.15 -0.16 δ c.i. 0.02 0.01 95% 0.02 0.03 0.03 0.05 0.04 0.03 0.04 0.05 0.05 0.03 0.03 0.02 0.04 0.05 0.05 0.05 0.04 0.09 0.10 2SD/SE Si 30 -0.30 -0.24 -0.28 -0.23 -0.28 -0.25 -0.27 -0.32 -0.28 -0.28 -0.30 -0.32 -0.31 -0.40 -0.38 -0.34 -0.42 -0.29 -0.31 δ Name MD57 D10-4 MD34 D6 MD57 D10-1 D04-54 D10-10 URA6 GUG13 ALAM5 SS15.1 A127 D-5 P-13 R937 BHVO-1 and -2 CHEPR25 MAR BCR-1 BHVO-1 basalts terr. Average Table IV–A.2, continued. Table Type Terrestrial basalts all Average Chapter IV — Page 119 Si) 29 2SD (δ Si) 30 2SD (δ Comments Reference c.i. 95% 2SD/SE Si 29 δ c.i. 0.02 0.03 0.02 0.01 95% 0.06 0.08 0.06 0.08 2SD/SE Si 30 -0.29 -0.32 -0.38 -0.29 δ Name Table IV–A.2, continued. Table Type individual: Averages, Fitoussi et al. (2009) Armytage et al. (2011) Chakrabarti and Jaconbsen (2010) Savage et al. (2010)

CHAPTER V.

EXPERIMENTAL EVIDENCE FOR MO ISOTOPE FRACTIONATION BETWEEN METAL AND SILICATE LIQUIDS1

1 This chapter has been submitted to Earth and Planetary Sciences Letters as: R. C. Hin, C. Burkhardt, M. W. Schmidt, B. Bourdon, T. Kleine. Experimental evidence for Mo isotope fractionation between metal and silicate liquids. Page 122 — Chapter V

ABSTRACT

Stable isotope fractionation of siderophile elements may inform on the conditions and chemical consequences of core-mantle differentiation in planetary objects. The extent to which Mo fractionates during such high temperature processes, however, was so far unexplored. We have therefore investigated equilibrium fractionation of Mo isotopes between liquid metal and liquid silicate to evaluate the potential of Molybdenum isotopes as a new tool to study core formation. We have performed experiments at 1400 and 1600 °C in a centrifuging piston cylinder, which perfectly segregated the metal and silicate fractions. Tin was used to lower the melting temperature of the Fe-based metal alloys, while variable Fe-oxide contents were used to vary oxygen fugacity in graphite and MgO capsules. Isotopic analyses were performed using a double spike technique. In experiments performed at 1400 °C, the δ98/95Mo of silicate is 0.193 ± 0.030‰ (95% confidence interval) heavier than that of metal. This fractionation is not significantly affected by the presence or absence of carbon, nor by the presence or absence of Sn. Molybdenum isotope fractionation is furthermore independent of oxygen fugacity in the range IW-1.79 to IW+0.47, which are plausible values for core formation. Experiments at 1600 °C show that, at equilibrium, the δ98/95Mo of silicate is 0.120 ± 0.020‰ heavier than that of metal. Equilibrium Mo isotope fractionation between liquid metal and liquid silicate as a function of 98/95 × 5 temperature can therefore be described as Δ MoMetal-Silicate = -4.80(±0.55) 10 / T2. The experiments performed at 1600 °C furthermore demonstrate that disequilibrium isotope fractionation results from highly reactive capsule material. Rapid 98/95 dissolution of capsule-derived MgO into the silicate melt has led to Δ MoMetal- Silicate of -0.027 ± 0.036‰, clearly due to disequilibrium isotope fractionation. Our equilibrium experiments, however, show that Mo isotope fractionation may be resolvable to temperatures of about 2500 °C, rendering Mo isotopes a powerful new tool to investigate the conditions of core formation.

V–1. INTRODUCTION

The segregation of a metallic core from a silicate mantle is arguably the largest differentiation event in growing planetary objects. Core formation in the Earth (and other terrestrial planets) may have involved merging of metal cores derived from already differentiated objects added during accretion (Karato and Murthy, 1997; Rubie et al., 2011; Taylor et al., 1993). Alternatively, it may have been associated with metal-silicate equilibration in a deep magma ocean (Righter and Drake, 1996; Wood et al., 2006). In the former case, the chemical signatures of core formation would reflect those of the relatively low temperatures and pressures characteristic for the planetesimals and embryos constituting the building blocks of the Earth. In the latter case, they would be the result of metal- Chapter V — Page 123 silicate separation at the high pressures and temperatures prevailing in the deep Earth. Most of the information about the conditions of core formation has come from studies of elemental distributions combined with experimental studies of element partitioning between metal and silicate liquids (Righter and Drake, 1996; Wood et al., 2006). As a result of these studies, equilibrium models of core formation in growing planets have become favoured over disequilibrium models, because the observed abundances of siderophile elements in the Earth’s mantle match those predicted for metal-silicate equilibration at high temperatures and pressures in a deep magma ocean quite well (Corgne et al., 2008; Righter et al., 2010; Wood et al., 2006). However, more recently it was shown that disequilibrium models provide an equally good match to the observed siderophile element abundances, indicating that elemental partitioning studies alone cannot constrain the degree of metal-silicate equilibration during core formation (Kleine and Rudge, 2011; Rudge et al., 2010). Rubie et al. (2011) concluded that a model with incomplete re-equilibration of the accreting material best reproduces the element budget of the silicate Earth. In recent years, stable isotopes have emerged as a tool to constrain the conditions of core formation (Fitoussi et al., 2009; Georg et al., 2007; Moynier et al., 2011; Polyakov, 2009; Williams et al., 2006). Mass dependent equilibrium fractionation of isotopes occurs due to differences in the bond stiffness of isotopes between various phases (Bigeleisen and Mayer, 1947; Urey, 1947). The magnitude of fractionation decreases strongly with increasing temperature and temperatures during metal-silicate segregation have to be high, well exceeding 1000 ��°C.�� Nev- ertheless, resolvable metal-silicate isotopic fractionation may occur due to the relatively large differences in bonding environment and oxidation state between metal and silicate phases (Zhu et al., 2002). In principle, fractionation factors could be calculated from vibrational frequencies of the isotopes in both phases (Bigeleisen and Mayer, 1947; Schauble, 2004; Urey, 1947), but these frequencies remain poorly constrained for most elements and phases. Thus, experimental calibration of the magnitude and direction of mass dependent isotope fractionation between metal and silicate is the best method to develop stable isotopes as tracers of core formation. Molybdenum is an element of interest for core formation. It is refractory under reducing conditions (Palme and O’Neill, 2003) and therefore unlikely to have been affected by isotopic fractionation by condensation/evaporation during heating events. Its siderophile behaviour suppresses the necessity for mass balance calculations, because bodies with core mass fractions as small as 5% will have over 90% of their Mo in the metallic core. As such, the isotopic effect is quantitatively transferred to the silicate portion, making analyses of silicate samples sufficient to study metal-silicate fractionation. We present the first experimental study of Mo isotope fractionation between liquid metal and liquid silicate. To determine whether there is resolvable fraction- Page 124 — Chapter V ation, we have performed experiments at temperatures of 1400 °C and 1600 °C at a pressure of 1 ��GPa in a centrifuging piston cylinder, which ensured perfect separation of metal from silicate prior to chemical processing for isotopic analyses. Although no data are available yet on natural rocks, our results indicate that Mo isotopes constitute a suitable tool to investigate conditions of core formation in planetary bodies.

V–2. METHODS V–2.1. Experimental methods In the silicate Earth, Mo is classified as a moderately siderophile element, but its metal/silicate liquid partitioning may exceed 1000 at oxygen fugacities relevant for core formation (Holzheid et al., 1994; Wade et al., 2012). Such an elemental distribution implies that a silicate melt contaminated by ~0.03 vol% metal consists of an approximately 1:1 mixture by mass of metallic Mo and Mo-oxide. Segregation of the metal and silicate liquids prior to isotopic analysis is therefore essential for a reliable interpretation of the isotopic data. In order to achieve this, we have used a centrifuging piston cylinder at ETH Zurich, whose acceleration results in a perfect density segregation of the synthetic metal liquid from the synthetic silicate liquid during the experiments. Equilibrium isotope fractionation decreases strongly as a function of temperature, favouring low temperatures to analytically resolve fractionation. We have thus performed experiments at 1400 °C and 1600 °C, at pressures of 1 GPa. The melting temperature of Fe-Mo alloys (>1500 °C) was lowered with Sn to <1400 °C (Okamoto, 1993). Both Sn-bearing and Sn-free experiments were performed at 1600 ��°C to ��control for potential effects of Sn on the Mo isotope fractionation factors (Table V–1, cf. mixtures B and F). In addition, graphite as well as MgO capsules were employed to establish whether carbon affects Mo isotope fractionation, because element partitioning may be affected by carbon (Jana and Walker, 1997). Run durations were varied to assess isotopic equilibration, while oxygen fugacity was varied to assess the potential effect of Mo4+/Mo6+ in the silicate melt on isotopic fractionation. We varied oxygen fugacity by varying the concentration of Fe2O3 in the oxide component of the starting material from 6.42 to 44.7 wt% (Table V–1, mixtures B, C, D). These runs were all performed at 1400 °C. A consequence of the siderophile character of Mo is that high Mo concentrations in the metal are required to yield measureable quantities of Mo in the silicate. We have therefore doped the metal component of our starting mixtures with 5-10 wt% Mo and prepared bulk starting mixtures with approximately 65 wt% silicate component. Under natural conditions, however, Mo concentrations in metals are generally below 50 ppm. We have therefore performed one experiment with a starting mixture with 1.25 wt% Mo in the metal component to verify Chapter V — Page 125

Table V–1. Compositions (wt%) of starting mixtures. Weight fractions (f) of metal and silicate (subscripts M and Sil, respectively) are given at the bottom.

A B1 B2 C D E F

Metal Fe 66.9 72.0 72.0 71.3 71.2 77.0 92.3 Sn 27.2 20.1 20.3 19.1 20.7 21.7 Mo 5.83 7.82 7.76 9.59 8.09 1.25 7.65

Silicate

SiO2 47.9 49.0 55.2 52.8 47.9 49.5 54.2

Al2O3 4.51 2.92 2.10 8.61 1.47 1.87 2.26 MgO 4.43 3.51 1.37 8.49 0.99 6.93 1.45

Fe2O3 30.1 32.5 31.3 6.42 44.7 31.9 32.2 CaO 12.0 11.0 9.06 18.5 4.08 8.88 8.90

Na2O 1.02 1.03 0.68 4.16 0.89 0.68 0.67

K2O 0.00 0.00 0.29 1.02 0.00 0.30 0.28

fM 0.3449 0.3556 0.3494 0.3321 0.3475 0.3501 0.3445 0.6551 0.6444 0.6506 0.6679 0.6525 0.6499 0.6555 fSil that, as expected, the high concentrations of Mo in our experiments did not affect isotopic fractionation (Table V–1, mixture E). All starting mixtures were prepared using the same metallic Mo powder and all experiments were performed with talc/pyrex assemblies in a 14 mm bore centrifuging piston cylinder. B-type thermocouples were protected by mullite (1400

°C) or Al2O3 (1600 °C) sleeves, while crushable MgO was used as spacers. The graphite and MgO capsules were welded into Au50Pd50 or Pt outer capsules with an outer diameter of 4.0 mm. Temperature calibrations in the employed assemblies indicate 6-9 °C temperature differences between a micrometre slice in the hottest and coldest region of the sample (Hin et al., 2012). The volume of metal liquid should thus have had an average temperature that was at maximum 3.5 °C lower than the average temperature of the silicate melt. Further technical specifi- cations and details of experimental protocols for the centrifuging piston cylinder at ETH Zurich are provided in Schmidt et al. (2006) and Hin et al. (2012). V–2.2. Analytical methods V–2.2.1. Elemental compositions Elemental concentrations were analysed in polished and carbon-coated samples. These analyses were performed with a Jeol JXA 8200 electron micro- Page 126 — Chapter V probe equipped with 5 spectrometers. An acceleration voltage of 15 kV was employed for all analyses. Silicates were analysed with a 6 nA current and a beam diameter of 1-10 µm, while metals were analysed with a 20 nA current and a beam diameter of 10 µm. Standards consisted of pure metals for metal compositions, and natural and synthetic silicates and oxides for silicate compositions, except for SnO2 that was analysed relative to a metallic Sn standard. Approxi- mately ten spots were analysed on each phase. Molybdenum contents in the silicate melt were always below microprobe detection limits. Molybdenum concentrations were therefore determined by isotope dilution (see below). In addition, selected samples were analysed by laser ablation ICPMS. Details for these procedures are provided in Appendix A.

V–2.2.2. Isotopic compositions After determination of elemental compositions, metal and silicate in each sample were separated and removed from the capsule with a diamond wire saw (300 µm wire diameter). A previously unused diamond wire was employed to prevent contamination. Furthermore, each analysed phase was carefully ground into a hexagonal prism on 30 µm alumina grinding paper to remove any contamination from sawing. This also removed tiny metal droplets that had stuck to the capsule walls during centrifugation. The clean, bulk pieces of silicate or metal were sonicated in acetone for 5 minutes. They were then crushed under acetone with separate Al2O3 pestles and mortars for metal and silicate that had previously not been used to crush Mo-bearing materials. A final inspection with a binocular microscope ensured that no visible metal contamination occurred in the silicate phase. All samples were weighed in pre-cleaned 15 ml Savillex beakers and were digested in 2:1 mixtures of 14 M HNO3 and 24 M HF (silicate samples and starting materials) or in 12 M HCl with a few drops of 14 M HNO3 (metals). A 100Mo-97Mo double spike (see electronic Appendix A for calibration details) was added to silicate samples prior to digestion, while starting materials and metal samples were digested before the double spike was added to an aliquot. The spiked aliquots were equilibrated overnight on a hotplate (130 °C). After digestion and spike-sample equilibration, all samples were dried down and treated with 1:1 mixtures of 14 M HNO3 and 30% H2O2 to remove any remaining organic material. A four-stage ion exchange procedure subsequently separated Mo from its matrix. Three of the four stages are identical to the protocol of Burkhardt et al. (2011), except that resin volumes for the first and second stages were scaled down from 14 and 7 ml to 3 and 2 ml, respectively, as our samples never exceeded masses of 20 mg. The first of these three stages employed a cation resin (Dowex AG50W-X8), onto which the samples were loaded in 1 ml 1 M HCl – 0.1 M HF. Chapter V — Page 127

The loaded solution was collected together with an additional 4 ml 1 M HCl – 0.1 M HF. This stage separates Mo from a large portion of its matrix. The second stage employed anion resin (Biorad AG1-X8). Samples were loaded in 6 ml 1 M HF, followed by rinsing with 12 ml 1 M HF, 16 ml 6 M HCl – 1 M HF and 2 ml ultra-pure H2O to remove remaining matrix (particularly Fe and Zr contami- nants). Molybdenum was then eluted with 6.5 ml 3 M HNO3. A TRU-spec resin was finally used to ensure that any remaining contamination by Zr and Ru were removed. Samples were loaded on the 1 ml TRU Spec resin in 1 ml 1 M HCl and rinsed with additional 6 ml 1 M HCl before Mo was eluted with 6.5 ml 0.1 M HCl. An additional ion exchange step was placed between the original first and second stages to remove Sn. This was done by eluting Mo from an anion resin (AG1-X8) with 1M HCl, making use of the different distribution coefficients of Mo and Sn under these conditions (Mizuike, 1983). The resin was equilibrated twice with 4 ml 1 M HCl prior to loading of the samples in 1 ml 1 M HCl. This 1 ml of sample was not collected after elution, but the following 13 ml 1 M HCl were collected and then dried to prepare for the next stages described above. Molybdenum blanks for the full procedure generally varied between 100 and 250 pg, which is negligible given that about 103 times higher amounts of Mo were analysed. The isotopic analyses were performed on two different multi-collector inductively coupled plasma mass spectrometers (MC-ICPMS). ). At ETH Zurich, a Nu Plasma 1700 equipped with a DSN-100 desolvator and PFA nebuliser was employed. On-peak baselines were collected for 60 s prior to each analysis. All seven isotopes of Mo were collected in 40 scans of 5 s integration together with 90Zr and 99Ru to monitor interferences, which were always insignificant. Double spike – sample mixtures were typically analysed as solutions of 75-100 ng total Mo per millilitre, yielding total Mo intensities of approximately 14×10-11 A at an uptake rate of ~140 µl min-1. This protocol consumed about 45-60 ng total Mo (corresponding to 23-30 ng sample Mo) per analysis. Experiments run at an oxygen fugacity more than one log unit below the iron-wüstite buffer yielded too low amounts of Mo for multiple analyses on the Nu1700 at ETH Zurich. They were therefore analysed on a ThermoScientific Neptune Plus at the Westfälische Wilhelms-Universität Münster. Measurement protocols were identical to those described for the Nu1700, except that 91Zr instead of 90Zr was measured to monitor Zr interferences and that on-peak back- grounds were collected for 30 s prior to each analysis. Furthermore, normal H cones were used and samples were introduced with a Cetac Aridus II desolvator after aspiration with a PFA nebuliser. Solutions of 50 ng total Mo per millilitre were analysed (consuming ~24 ng total Mo per analysis), which yielded a total Mo intensity of about 14×10-11 A at an uptake rate of ~100 µl min-1. Data reduction was performed for all analyses with the geometrically motivated Page 128 — Chapter V double spike deconvolution scheme of Siebert et al. (2001), using the evalua- tion software of Nu Instruments Ltd. The molar spike proportion determined with this scheme was used to calculate Mo concentrations of all silicate samples.

Moreover, the natural fractionation factor of a sample (αSample) obtained from the deconvolution was used to calculate the reported δ98/95Mo. This was done after correction for short-term fluctuations by subtracting the meanα�� in�� a given meas- Mean urement session for the reference standard NIST SRM 3134 (αSRM 3134 ):  98/95 Mean m98 (V.1) dMo =1000( ααSample− SRM 3134 )ln m95 98 95 where m98 and m95 are the atomic masses of Mo and Mo, respectively. Ana- lytical uncertainties on sample compositions are reported as the 95% confidence interval (95% c.i.) of the number (n) of repeated sample analyses: SD (V.2) 95% ci. . = ⋅ t0.95 n

in which t0.95 refers to a two-tailed 95% confidence test (Miller and Miller, 2010). SD is the standard deviation calculated as

2 ∑ ()xx− SD = i (V.3) n −1

where xi and x refer to a single measurement on a sample and the average of n measurements on a sample, respectively. Long-term reproducibility (2SD) over the course of this study was evaluated on the NIST SRM 3134 standard as well as on aliquots of BHVO-2 powder and a solution of molybdenite sample AJ011 (Greber et al., 2011). Spiked runs of the NIST SRM 3134 standard yielded an average δ98/95Mo of -0.002 ± 0.085‰ (2SD) on the Nu1700 (n = 159) and -0.003 ± 0.049‰ on the Neptune Plus (n = 131). Averages for BHVO-2 on the Nu1700 and Neptune Plus are, respectively, -0.068 ± 0.130‰ (n = 16 on 5 aliquots) and -0.060 ± 0.052‰ (n = 15 on 2 aliquots), while those for AJ011 are 0.703 ± 0.121‰ (n = 6 on 2 aliquots) and 0.697 ± 0.057‰ (n = 12 on 2 aliquots). These latter values are within analytical uncertainty identical to those of Greber et al. (2011) when corrected to the same reference standard. In addition, we analysed an aliquot of the JMC Mo standard from the University of Bern, which gave an average of -0.260 ± 0.017‰ (95% c.i.), which compares well to the value of -0.25 ± 0.09‰ (2SD) as determined by Greber et al. (2012). Fractionation factors (��Δ��) between metal and silicate are presented as differences between the bulk isotopic composition of the starting mixture and the analysed silicate composition. This is done because elemental Mo concentrations in the metal exceed those in the silicate by a factor ~1500, i.e. >99.9% of the bulk Mo resides in the metal phase. In this scenario, a fractionation of > 9‰ between metal and silicate would be required to shift the metal by 0.05‰ relative to the Chapter V — Page 129 bulk composition, while a fractionation of 0.2‰ would shift the metal <0.002‰ relative to the bulk. Metal isotopic compositions are thus not analytically resolvable from isotopic compositions of bulk starting mixtures, such that 98/95 98/95 98/95 98/95 DMoMetal−− Silicate =D=−MoBulk Silicate ddMoBulk MoSilicate (V.4)

V–3. RESULTS

Major element compositions of the melts are presented in Table V–2 and in the Appendix to this chapter in Table V–A.1. Elemental Mo contents in silicates are presented in Table V–3 together with isotopic compositions as well as all details of experimental run conditions.

Figure V–1. Back-scatter electron images of selected experiments. Panel (a) shows an experiment in an MgO capsule at 1600 °C with euhedral olivine crystals (a few hundred µm in size). Quench needles (olivine) in the silicate are visible. Panel (b) shows an example of an experiment in an MgO capsule at 1400 °C. At this temperature, these experiments lack significant euhedral olivine crystals and quench needles are barely visible. Panel (c) presents an example of an experiment performed in a graphite capsule, where surface forces often retained few metal droplets attached to the capsule wall. These droplets, however, were always removed with grinding paper. Panel (d) shows crushed material of the remaining silicate, which shows no indication of any metal presence. Page 130 — Chapter V

V–3.1. Textures and elemental compositions In general, experiments were performed at temperatures exceeding the liquidi of metal alloys as well as the silicate mixture in both graphite (Figure V–1a) and MgO (Figure V–1b) capsules. The silicate liquidus was not reached in experiments performed at 1600°C in MgO capsules, where rapid dissolution of MgO from the capsule led to a significant shift in melt composition and to olivine saturation (Figure V–1c). Centrifugation has perfectly separated the metal liquids from the silicate liquids, except for a few small metal droplets that stuck to the capsule walls, in particular in graphite capsules (Figure V–1a). These droplets were removed by grinding the exterior portions of the silicate glass cylinders (see section 2.2.2.). Detailed inspection by SEM indicated that no tiny (0.1-10 µm) metal droplets are present in the quenched silicate melt (Figure V–1d). Molybdenum contents in metal alloys have increased by 5-45%, while Sn

(only added in metal form) oxidised to form 0.21 to 4.36 wt% SnO2 in the silicate melt. In one experiment (RH76) a separate ~Pd40Sn60 alloy formed in addition to a relatively Sn-poor Fe-Mo alloy (Table V–A.1 in the Appendix to this chapter). Mo was below microprobe detection limits (~0.03 wt%) in the Pd-Sn alloy, and hence the formation of the Pd-Sn alloy could not have modified the Mo isotope composition of the Fe-Mo-Sn alloy. The formation of such a Pd-Sn alloy in a single experiment is readily explained by diffusive exchange of the Fe-Sn-Mo alloy with the Au-Pd capsule.

Except for SnO2, most oxides in the silicate melt from experiments in MgO capsules have been diluted by MgO dissolution from the capsule. Only FeO contents in experiments performed in MgO capsules with starting mixture C have increased from 5.78 wt% to about 8.4 wt%, despite the increase in MgO contents from 8.49 wt% to about 15 wt%. As to be expected, MgO dissolution at 1600 °C has been much more extensive than at 1400 °C: after 1.1 hours MgO contents in the silicate melt in experiment RH80 had increased from 1.37 wt% to 24.4 wt%, which compares to an increase to 16.0 wt% after 4.0 hours at 1400 °C in experiment RH77. Oxygen fugacity was calculated relative to the iron-wüstite buffer (ΔIW): X ⋅γ D=⋅IW2 log FeO FeO (V.5) X Fe⋅γ Fe

where XFeO and XFe are the molar fractions of FeO and Fe in the silicate and metal melts, respectively, while γFeO and γFe are their respective activity coefficients. Activity coefficients for metallic Fe were calculated with the modified Sn Wagner epsilon approach of Ma (2001). The interaction parameter ε Mo was determined to be -4.94 at 1873 K based on the Sn-free and Sn-bearing experiments RH102 and RH103, respectively, again following the approach of Ma (2001). All other interaction parameters were taken from the Steelmaking Data Chapter V — Page 131 4.6 0.6 n.d. n.d. 0.02 4.22 0.09 2SD 0.8 n.d. n.d. 0.95 89.2 9.79 0.10 99.3 1600 -0.73 MgO Start Mix F RH86 0.8 1.1 n.d. n.d. 0.09 0.76 0.05 2SD 1.2 5.2 n.d. n.d. 0.10 0.06 85.4 9.21 0.04 94.8 1600 Start Mix F RH102 Graphite 5.1 5.3 0.4 n.d. 0.05 0.13 0.05 2SD 1.5 n.d. 0.89 72.4 26.2 1.66 0.04 1400 -0.52 MgO 100.3 RH99 Start Mix E 0.5 0.5 0.4 n.d. 0.04 0.43 0.03 2SD 1.7 5.1 n.d. 11.8 0.95 0.47 73.8 9.04 0.01 94.9 1400 Start Mix D RH101 Graphite 1.5 0.4 1.5 0.7 n.d. 0.04 0.04 2SD 1.5 n.d. 0.98 68.9 21.1 10.2 0.06 1400 -1.68 MgO 100.5 RH97 Start Mix C 4.8 4.2 1.2 n.d. 0.06 1.83 0.12 2SD 2.2 n.d. 0.91 67.8 22.8 9.81 0.09 1400 -0.64 MgO 100.6 RH82 Start Mix B2 = oxygen fugacity relative to the iron-wüstite buffer (ΔIW), t = run duration in hour. to the iron-wüstite = oxygen fugacity relative 1.7 1.1 0.4 n.d. 2 0.08 1.00 0.13 2SD 1.5 4.3 n.d. 0.91 0.14 73.2 12.5 9.86 0.14 95.7 1400 RH72 Start Mix B1 Graphite Elemental compositions (wt%) of selected samples with their experimental run conditions. P = pressure (GPa), pressure = P run conditions. experimental their samples with selected of (wt%) compositions Elemental 2 1 (ΔIW) 2 O P (GPa) P (°C) T t (h) f Capsule type Metal Fe Sn Mo Pt Pd Total C (calc) Table V–2. Table (°C), fO = temperature T Page 132 — Chapter V 0.9 4.0 5.1 0.7 1.19 0.74 0.24 0.17 0.02 2SD 38.5 2.45 18.3 30.4 9.62 0.86 0.33 0.01 100.4 Start Mix F RH86 0.6 0.6 0.9 0.12 0.09 0.19 0.10 0.03 0.00 2SD 46.9 1.97 1.32 40.4 8.05 0.64 0.23 0.00 99.5 Start Mix F RH102 5.6 4.8 1.7 11.3 0.11 0.94 6.92 0.43 0.77 2SD 42.4 1.65 12.5 32.5 7.91 0.66 0.27 0.94 98.8

RH99 Start Mix E 0.8 1.6 1.3 0.12 0.10 0.42 0.12 0.02 0.39 2SD 39.9 1.22 0.85 49.0 3.39 0.83 0.01 4.06 99.2 Start Mix D RH101 0.9 0.4 0.7 0.8 0.21 0.39 0.26 0.15 0.10 2SD 45.8 7.25 14.4 8.71 16.7 3.66 0.92 0.26 97.8 RH97 Start Mix C 5.7 8.9 5.6 1.6 0.68 4.88 0.34 0.12 0.84 2SD 46.0 1.67 14.3 28.6 7.27 0.56 0.24 1.14 99.7 RH82 Start Mix B2 1.1 0.6 1.1 0.13 0.18 0.26 0.13 0.01 0.95 2SD 42.8 2.47 3.10 37.1 9.55 0.97 0.00 3.03 99.0 RH72 Start Mix B1 3 2 2 O O 2 O 2 2 Errors refer to the time dependent variation of pressure in centrifuge experiments. of the totals from 100%. Carbon contents are calculated as the difference Table V–2, continued.. Table Silicate SiO Al MgO FeO CaO Na K SnO Total 1 They are calculated as the 2SD of pressure registrations in 30 s intervals our log files (see text for more explanation). 2 Chapter V — Page 133

Sourcebook (1988). Holzheid et al. (1997) determined that the activity coefficient of FeO in silicate liquid is constant at 1.7 between 1300 and 1600 °C for broadly basaltic compositions with up to 20 wt% MgO and 12 wt% FeO. They also determined an FeO activity coefficient of 1.7 in an experiment with 26.0 wt% FeO and 4.87 wt% MgO, but some of our experiments have up to 49.0 wt % FeO, falling outside the range of Holzheid et al. (1997). In absence of better approximations, we have kept γFeO constant at 1.7 in equation (V.5). Calcu- lated ΔIW thus varied from -1.05 with starting mixture C to +0.47 with starting mixture D in graphite capsules, and from -1.79 with starting mixture C to -0.62 with starting mixture B2 in MgO capsules (Table V–3 and Table V–A.1). We note, however, that Richardson (1956) found that FeO activities in CaO-FeO-

SiO2 melts approach Raoult’s law behaviour at 50 mol% FeO. Taking an activity coefficient of 1.0 instead of 1.7 for our most oxidised experiments decreases the calculated oxygen fugacity (ΔIW) from +0.46 to 0.00. Molybdenum contents in the silicate melts vary from 6.09 to 13.3 ppm at ΔIW below -1.0, and up to 119 ppm at ΔIW +0.47. Anomalously high and variable Mo contents were observed for experiments in MgO capsules at 1600 °C (ΔIW about -0.7). They varied from 108 to 193 ppm after run durations of ~0.75 hours, and from 246 to 496 ppm after run durations of 1.1 hours (Table V–3).

Figure V–2. Molybdenum isotope compositions for starting mixtures and Starting mixture experimental run products. The star Metal represents the average composition of Silicate the starting mixture relative to which (C capsule) metal-silicate fractionation factors Silicate (MgO capsule) are calculated. Metal run products 1600°C are presented as diamonds adjacent to 1400°C silicates (circles for graphite capsules, squares for MgO capsules) they equilibrated with. All analysed metals are within uncertainty identical to the bulk composition of starting materials. Open symbols of experimental run products refer to experiments performed at 1400 °C, closed symbols to experiments at 1600 °C. Half-filled symbols represent Sn-free experiments at 1600 °C. Averages (with shades as 95% confidence intervals) are given for 1400 °C and 1600 °C, where -0.6 -0.5 -0.4 -0.3 -0.2 -0.1 0 two different averages are given for the δ98/95Mo (‰) latter due to disequilibrium fractionation in MgO capsules (grey symbols; see section 4.2.). Page 134 — Chapter V

V–3.2. Isotopic compositions V–3.2.1. Starting material and metal melts A total of three aliquots from starting mixtures B1 and F were analysed and yielded an average δ98/95Mo starting composition of -0.389 ±�� 0.020‰ (95% confidence interval). Equilibrated metal fractions from experiments RH55, RH76 and RH77 have δ98/95Mo compositions of -0.374 ± 0.042‰, -0.359 ± 0.056‰, and -0.382 ± 0.025‰ (95% c.i.), respectively, indistinguishable from the starting composition (Figure V–2 and Table V–3). V–3.2.2. Silicate melt from experiments at 1400 °C Six experiments were performed with starting mixtures B1 and B2 at variable run durations from 0.5 to 7.0 hours. Experiment RH71 with a run duration of 0.5 hours resulted in a silicate δ98/95Mo of -0.277±0.044‰ (95% c.i.). Silicates from the other five experiments are all within analytical uncertainty of each other, but generally slightly heavier than RH71, with δ98/95Mo between -0.179 ± 0.043‰ and -0.244 ± 0.098‰ (Figure V–3a and Table V–3). The two experiments performed in MgO capsules (RH77 and RH82) have Mo isotope compositions indistinguishable from those performed in graphite capsules (Figure V–3a and Table V–3). 0.2 0.2 a Graphite capsule b Graphite capsule 0.1 MgO capsule 0.1 MgO capsule (‰) 0 0 -0.1 Average 1600°C -0.1 Metal-Silicate

Mo -0.2 -0.2

98/95 Average 1400°C

∆ -0.3 -0.3 [Mo] = 1.66 wt% -0.4 -0.4 Metal 0 1 2 3 4 5 6 7 8 -2.5 -2.0 -1.5 -1.0 -0.5 0 0.5 1.0 ∆ Run duration (h) log fO2 ( IW) Figure V–3. Molybdenum isotope fractionation factors between metal and silicate plotted against experimental run duration (panel a) and oxygen fugacity (panel b). Circles represent experiments in graphite capsules, squares MgO capsules. Errors are 95% confidence intervals. Panel (a) shows that at 1400 °C (open symbols) the shortest run duration of 0.5 h had probably not fully equilibrated. Experiments with longer run duration are within uncertainties identical, independent of whether they were performed in graphite or MgO capsules. Closed symbols in panel (a) represent Sn-bearing experiments at 1600 °C, while half-filled symbols refer to Sn-free experiments at 1600 °C. Experiments at 1600 °C in MgO capsules are in grey, because they represent disequilibrium fractionation (see section 4.2.). Panel (b) shows that Mo isotope fractionation factors are independent of oxygen fugacity in the range from ΔIW -1.79 to +0.47. All experiments presented in panel (b) were performed at 1400 °C. The experiment with 0.5 h run duration was not plotted here, because it probably represents disequilibrium fractionation. Chapter V — Page 135

To assess the effects of oxygen fugacity on δ98/95Mo, experiments with variable oxygen fugacity between ΔIW -1.79 and +0.47 were performed. The δ98/95Mo of the silicate melts from these experiments vary between -0.141 ± 0.032‰ and -0.294 ± 0.022‰ (95% c.i.), but they do not correlate with oxygen fugacity (Figure V–3b). Again, no resolvable differences in δ98/95Mo occurred in experiments performed in graphite or MgO capsules (Figure V–3b and Table V–3). The experiment performed with the lowest Mo content in the metal (1.66 wt%; experiment RH99) resulted in a silicate δ98/95Mo of -0.104 ± 0.017‰. V–3.2.3. Silicate melt from experiments at 1600 °C Sn-bearing as well as Sn-free experiments were performed at 1600 °C in both graphite and MgO�� capsules. Among both types of capsules, the Mo isotope composition of silicate melts from Sn-bearing and Sn-free experiments are within analytical uncertainty identical. However, there is a systematic offset between the δ98/95Mo compositions of silicates derived from graphite capsules and those derived from MgO capsules: the former vary from -0.260 ± 0.012‰ to -0.275 ± 0.013‰, while the latter vary from -0.333 ± 0.097‰ to -0.386 ± 0.069‰ (95% c.i.; Figure V–3a and Table V–3).

V–4. DISCUSSION

Analyses performed on metals from three experiments show that their δ98/95Mo compositions are within uncertainty identical to the average bulk Mo isotope composition of the starting mixtures, as expected from mass balance considerations. This confirms that isotopic differences between the bulk starting mixtures and the silicate melts recovered from the experiments can be equated to differences in isotopic compositions of the equilibrated metal and silicate phases (equation V.4). As explained above, Mo isotope fractionation factors between metal and silicate liquids are thus obtained by subtracting silicate compositions from the composition of the bulk starting mixture. V–4.1. Fractionation at 1400 °C V–4.1.1. Time series The results of the experimental time series imply that a run duration of 0.5 hour is not sufficient to reach isotopic equilibrium at 1400°C . The metal-silicate fractionation factor of RH71 is smaller by 0.071 ± 0.055‰ (95% c.i.) than the average of the other five experiments with run durations of 1.5 to 7.0 hours. Based on diffusion coefficients of approximately 1.5 ´10-11 m2 s-1 for Ti and Zr in haplobasaltic melt at 1400 °C (LaTourrette et al., 1996), we estimate that the diffusion length scale of Mo is <170 µm in the silicate melt in experiment RH71. Given the maximum length of the silicate melt fractionation of ~3 mm, it is conceivable that Mo isotopes had not fully equilibrated in this experiment. Page 136 — Chapter V 2

[Mo] (ppm) 6 6 8 6 6 n 16 15 12 35

(‰) 95% c.i.

M-Sil Mo (‰) 98/95 Δ

(‰) 0.130 0.052 0.121 0.057 0.017 0.020 0.042 0.056 0.025 95% c.i.

Mo (‰) 0.703 0.697 98/95 -0.068 -0.060 -0.260 -0.389 -0.374 -0.359 -0.382 δ

C C type MgO Capsule

(g) 487 393 515 RCF

t 4.0 7.0 4.0 (h)

2 O f 0.10 0.13 -0.70 (ΔIW)

T (°C) 1350 1400 1400 = oxygen fugacity relative to the iron-wüstite buffer (ΔIW), t = run duration in hour, RCF = relative relative = RCF hour, in duration run = t (ΔIW), buffer iron-wüstite the to relative fugacity oxygen = 2 1

P (GPa) A B1 B2 mix Start Isotopic compositions and details of experimental run conditions. Zurich = performed measurements on the Nu1700 at 3

Table V–3. Table pressure = P Münster. in Wilhelms-Universität Westfälische at Plus Neptune the on performed measurements = Münster Zurich; ETH (°C), fO temperature = (GPa), T analyses, capsule type C = graphite. in multiples of the gravity constant (g), n = number repeated centrifugal force Sample Standards BHVO-2 (Zurich) BHVO-2 (Münster) AJ011 (Zurich) AJ011 (Münster) JMC Bern Starting composition Metals RH55 M RH76 M RH77 M Chapter V — Page 137 2

119 11.0 52.5 73.3 74.9 79.1 71.2 9.61 58.2 47.6 6.09 7.61 13.3 [Mo] (ppm) 4 9 8 6 4 5 6 7 7 5 6 6 6 n

(‰) 0.040 0.048 0.028 0.041 0.100 0.037 0.029 0.047 0.049 0.048 0.029 0.036 0.026 95% c.i.

M-Sil Mo (‰) -0.112 -0.181 -0.182 -0.184 -0.145 -0.247 -0.171 -0.209 -0.195 -0.203 -0.094 -0.216 -0.285 98/95 Δ

(‰) 0.035 0.044 0.019 0.036 0.098 0.032 0.021 0.043 0.045 0.044 0.022 0.030 0.017 95% c.i.

Mo (‰) 98/95 -0.207 -0.277 -0.207 -0.204 -0.244 -0.141 -0.218 -0.179 -0.194 -0.185 -0.294 -0.173 -0.104 δ

C C C C C C C type MgO MgO MgO MgO MgO MgO Capsule

(g) 487 514 393 742 748 515 764 764 739 489 739 1130 1140 RCF

t 4.0 0.5 1.5 4.0 7.0 1.7 1.7 4.0 2.2 1.5 1.5 4.0 1.5 (h)

2 O f 0.10 0.16 0.14 0.14 0.13 0.47 -1.05 -0.70 -0.64 -1.68 -1.71 -1.79 -0.52 (ΔIW)

T (°C) 1350 1400 1400 1400 1400 1400 1400 1400 1400 1400 1400 1400 1400 1

P (GPa) 0.79±0.10 1.06±0.10 0.91±0.08 0.73±0.17 0.82±0.07 0.97±0.03 0.95±0.04 0.90±0.05 0.91±0.06 0.98±0.04 1.00±0.07 0.91±0.04 0.89±0.05 E A C C C C D B1 B1 B1 B1 B2 B2 mix Start Table V–3, continued. Table Sample Silicates 1400°C RH55 Sil RH71 Sil RH72 Sil RH74 Sil RH76 Sil RH94 Sil RH101 Sil RH77 Sil RH82 Sil RH97 Sil RH98 Sil RH104 Sil RH99 Sil Page 138 — Chapter V 2

496 193 246 108 52.6 50.3 43.7 [Mo] (ppm) 6 6 6 6 3 6 6 n

(‰) 0.110 0.023 0.024 0.028 0.065 0.099 0.072 95% c.i.

M-Sil Mo (‰) -0.114 -0.116 -0.129 -0.034 -0.055 -0.017 -0.003 98/95 Δ

(‰) 0.012 0.013 0.019 0.061 0.097 0.109 0.069 95% c.i.

Mo (‰) 98/95 -0.260 -0.275 -0.272 -0.355 -0.333 -0.372 -0.386 δ

C C C type MgO MgO MgO MgO Capsule

(g) 849 908 893 860 1070 1863 1871 RCF

t 1.7 1.4 1.2 1.1 0.7 1.1 0.8 (h)

2 O f 0.12 0.06 0.06 -0.70 -0.62 -0.72 -0.73 (ΔIW)

T (°C) 1600 1600 1600 1600 1600 1600 1600 1

P (GPa) 0.98±0.08 0.95±0.17 0.95±0.09 1.00±0.04 1.06±0.02 0.98±0.04 0.95±0.02 F F F B2 B2 B2 B2 mix Start Errors are 2SD of the pressure variation during experimental run, based on automatic log files with recordings every 30s. Concentrations determined from double spiked isotope ratio analyses. Relative uncertainties (2SD) are better than 2% and mainly dependent standard (lot 602332B). JMC Bern = Johnson Matthey ICP Table V–3, continued. Table Sample 1600°C RH92 Sil RH103 Sil RH102 Sil RH80 Sil RH87 Sil RH81 Sil RH86 Sil 1 2 on weighing uncertainties of the small silicates sample masses (between 6 and 18 mg). 3 Chapter V — Page 139

This resulted in kinetic enrichment of the silicate melt in light Mo isotopes by oxidation of metallic Mo and we have consequently excluded experiment RH71 from calculations of averages. Experiments run with a duration ≥1.5 hour, however, do not vary in Mo isotope fractionation factors. We therefore conclude that these experiments display equilibrium fractionation of Mo isotopes, the silicate being on average 0.183 ± 0.026‰ (95% c.i.) lighter than metal. Furthermore, experiments performed in graphite and MgO capsules yielded identical results, implying that the presence of carbon has no resolvable effect on Mo isotope fractionation. A similar observation was previously made by Hin et al. (2012) for Fe isotopes. V–4.1.2. Fractionation as a function of oxygen fugacity In silicate melts, Mo begins to change valence state from 4+ to 6+ at an oxygen fugacity around ΔIW -1 (Farges et al., 2006; Holzheid et al., 1994; O’Neill and Eggins, 2002). Structural analyses by XANES imply that the change in valence state is accompanied by a structural modification from octahedral Mo4+ to tetra- hedral Mo6+ coordination (Farges et al., 2006). As detailed in Schauble (2004), equilibrium mass dependent isotope fractionation is not only a function of mass and temperature, but also of bonding environment, the heavy isotopes concen- trating in phases with stiffer bonds. In general, the stiffest bonds form around atoms with the lowest coordination and the highest valence state. Because of the bonding environment changes Mo may undergo around ΔIW -1, we have varied oxygen fugacity to investigate whether changes in Mo valence and coordination state affect the fractionation of Mo isotopes. The range of oxygen fugacities at reach for our purpose was limited by the amount of Mo in the silicate at low oxygen fugacity and by the stability of Fe-based metal at high oxygen fugacity. As experiments in MgO and graphite capsules yielded identical results (see section 4.1.1.), we used both MgO and graphite capsules to vary oxygen fugacity between ΔIW -1.79 and +0.47. Following O’Neill and Eggins (2002), Mo4+ should decrease from ~80% to ~23% of the total Mo from ΔIW -1.79 to +0.47, respectively, for the silicate melt compositions of our experiments (Figure V–4a). These percentages are, however, sensitive to uncertainties in Gibbs free energies, mainly for the formation of liquid MoO2. A 10% increase in the Gibbs free energy of the reaction, for instance, would predict approximately 10% lower Mo4+ fractions in our samples (68% and 14% for ΔIW -1.79 to +0.47, respectively). Despite the variation in oxygen fugacity and the expected corresponding change in valence state, Mo isotope fractionation factors between metal and silicate in our experiments do not vary as a function of oxygen fugacity (Figure V–3b). This could indicate that i) Mo isotope fractionation between liquid metal and liquid silicate is insensitive to the bonding environment in the silicate liquid, or ii) the bonding environment of Mo has not changed significantly in our exper- Page 140 — Chapter V iments. The latter can be investigated by regressing oxygen fugacity against elemental Mo distribution. Molybdenum distribution (and isotope fractionation) between metal and silicate occurs by the reaction n Mo+ O MoO (V.i) 4 2 n/2 where n indicates the valence state of Mo. The equilibrium constant (K) for this reaction can be written as []MoOn/2 (V.6) K = n 4 []Mo fO2 where f refers to fugacity, while rectangular brackets indicate activities, which are identical to the product of molar fractions (X) and activity coefficients (γ) for each species. The latter can be calculated with solution models, such as those of Ma (2001) for metal activity coefficients. Equation (V.6) can be re-written as []Mo n (V.7) logK=−− log logfO2 []MoOn/2 4 []Mo It follows from equation (V.7) that in log versus log fO2 space, []MoOn/2 straight lines with slopes of -1.0 or -1.5 correspond to Mo valence states of 4+ or 6+, respectively. A gradual change of valence state should, therefore, result in a curve that changes its slope from -1.0 to -1.5 from low to high oxygen fugacity. Such a gradual change in slope was observed by O´Neill and Eggins (2002). For our experiments, the logarithms of the ratios of molar Mo activity and molar MoOn/2 fraction versus log fO2 fall on a straight line with a slope of -1.0 if a constant FeO activity coefficient of 1.7 is used for all experiments. This implies a dominant Mo valence state of 4+ in the silicate melt of all experiments

(Figure V–4b). This conclusion does not change when molar MoOn/2 fractions are converted to molar MoOn/2 activities with activity coefficients calculated with the model of Wade et al. (2012), which accounts for variable SiO2, MgO, Al2O3 and CaO contents in the silicate melt. In fact, the variation in Mg/Si ratios of 0.03-0.45 in experiments at 1400 °C has little effect on the Mo distribution: MgO capsules yielded ~35% lower Mo contents in the silicates compared to graphite capsules (Table V–3), but none of these data deviate from the trend (Figure V–4b). In contrast to previous work (Farges et al., 2006; Holzheid et al., 1994; O’Neill and Eggins, 2002), however, we do not observe a change in valence in the silicate melt from Mo4+ to Mo6+ around ΔIW -1. A possible cause for this discrepancy could be the choice for a constant FeO activity coefficient of 1.7. Based on Richardson (1956), it might be more appropriate to decrease the activity coefficient to ~1.3 and ~1.0 for experiments with ~35 and ~49 wt% FeO. Re-plotting the data in Figure V–4b with oxygen fugacities calculated with these lower FeO activity coefficients would allow for some Mo6+ (Figure V–A.3 in the Appendix to this chapter). Chapter V — Page 141

1 6.5 Theoretical slope Graphite capsule 0.9 a 1400°C b for 4+ valence ) 1600°C 6.0 MgO capsule

6+ 0.8 fO2 experiments n/2 5.5 Theoretical slope 0.7 for 6+ valence

+ Mo 0.6

MoO 5.0 Mg/Si ~ 0.45 4+ 0.5 /X Mg/Si ~ 0.6, 4.5 0.4 Mo 45 min Mg/Si /(Mo ~ 0.09 4+ 0.3 4.0 Mg/Si ~ 0.7,

log a 60 min Mo 0.2 3.5 Depolymerization effect 0.1 of MgO capsules at 1600°C 0 3.0 -6 -5 -4 -3 -2 -1 0 1 2 3 4 5 6 -2.5 -2.0 0.50-0.5-1.0-1.5 1.0 ∆ log fO2 ( IW) log fO2 (∆IW) Figure V–4. Determination of Mo valence state as a function of oxygen fugacity. Panel (a) shows how Mo4+/(Mo4+ + Mo6+) should vary in our experiments with oxygen fugacity according to the work of O’Neill and Eggins (2002).The equation in their Table 8 was used to calculate Mo-oxide activity coefficients for the silicate composition we employed, following which their equation 19 was used to produce the curves for 1400 °C (solid line) and 1600 °C (broken line). The stars on the horizontal axes indicate the oxygen fugacity determined for our experiments, while the arrows indicate the expected Mo4+/(Mo4+ + Mo6+). Panel (b) shows elemental partition coefficients as a ratio of Mo metal activity over molar content of Mo-oxide. Symbols are identical to Figures 3 and 4, except for the additional circles with crosses inside, which represent experiments that were not analysed for isotopic compositions. Their silicate Mo contents were determined by LA-ICPMS (see Appendix). A regression through the data derived from experiments at 1400 °C gives a +0.16 slope of −0.92−0.10, which is within uncertainty of the indicated theoretical slope of -1.0 for Mo4+ (solid line). The theoretical slope for Mo6+ is shown for reference. Note that the data for experiments in MgO capsules at 1600 °C (grey symbols) deviate from the correlation displayed by all other data, apparently as a function of run duration and Mg/Si, while experiments in MgO capsules at 1400 °C do not display this behaviour. Analytical uncertainties are always less than twice the symbol sizes. [An alternative for panel (b) using different FeO activity coefficients is presented as Figure V–A.3 in the Appendix.]

The absence of a correlation of Mo isotope fractionation with oxygen fugacity indicates that all experiments at 1400 °C can be used to calculate an average Mo isotope fractionation factor between metal and silicate. In addition to experiment RH71 with 0.5 hour run duration, two samples are outliers compared to all other experiments, although we do not have an explanation for this (Figure V–2). Experiment RH98, for instance, was performed under conditions identical to experiments RH97 and RH104, yet its fractionation factor is smaller: -0.094 ± 0.029‰ compared to -0.203 ± 0.048‰ and -0.216 ± 0.036‰, respectively (Table V–3). Because the somewhat different results for experiments RH98 and RH99 remain puzzling, we have only excluded experiment RH71 from the average Mo isotope fractionation factor between metal and silicate at 1400 °C: -0.193 ± 0.030‰ (95% c.i.). Page 142 — Chapter V

V–4.2. Fractionation at 1600 °C In contrast to the results at 1400 °C, there is a resolvable difference at 1600 °C in Mo isotope fractionation between experiments in graphite and MgO capsules. This is ultimately the consequence of kinetic fractionation with a silicate melt whose composition evolved with time in the MgO capsules. The experiments in graphite capsules yield an average fractionation factor of -0.120 ± 0.020‰ compared to -0.027 ± 0.036‰ in MgO capsules (95% c.i.). The Mo contents of silicate melts from MgO capsules are 2 to 10 times higher compared to those from graphite capsules. These Mo contents are anomalously high for the oxygen fugacity at which they were performed and they plot off the slope for Mo4+ (Figure V–4b). Moreover, the Mo concentrations in silicate melts of the four experiments in MgO capsules increase with run duration and co-vary with increasing Mg/Si ratios (Figure V–4b). The solubility of high valence elements (such as Mo) in silicate melts increases with melt depolymerisation (Ellison and Hess, 1986; Hillgren et al., 1996). The silicate melts in MgO capsules depolymerise with run duration as a consequence of progressive dissolution of the capsule with time. Consequently, the solubility of Mo also increases progressively during the experiment. This explains the observed dependence of Mo concentrations in silicate melts (run in MgO capsules) on MgO concentration (see Figure V–A.4 in the Appendix at the end of this chapter), which increases with temperature and run duration. The progressively increasing solubility of Mo in the silicate melt leads to progressive oxidation of metallic Mo. If the oxidation rate exceeds the rate of isotopic equilibration, kinetic Mo isotope fractionation is expected to occur. As light isotopes diffuse faster than heavy isotopes (Richter et al., 2009; Schauble, 2004), this scenario fits the observation that the silicates from MgO capsules are on average lighter than those recovered from graphite capsules. We therefore conclude that imperfect equilibration of the silicate melt at 1600 °C in MgO capsules yielded disequilibrium Mo isotope fractionation. These observations indicate that care should be taken when experimentally studying equilibrium isotope fractionation in capsules that are reactive with the studied phases; at least for high valence elements in capsules that are very reactive with a silicate melt. The equilibrium Mo isotope fractionation factor between liquid metal and liquid silicate at 1600 °C is, in our study, represented by the average fractionation factor observed in graphite capsules: -0.120 ± 0.020‰ (95% c.i.). As expected, this is smaller than the fractionation factor of -0.193 ± 0.030‰ observed at 1400 °C.

V–5. IMPLICATIONS FOR CORE FORMATION

The experiments show that the presence or absence of C or Sn in the metal does not affect Mo isotope fractionation between liquid metal and liquid silicate. Chapter V — Page 143

Variations in silicate composition also do not seem to affect Mo isotope fractionation, at least in the range from e.g. about 40-48 wt% SiO2, 7.9-49 wt% FeO and 0.85-16 wt% MgO. Additionally, we did not observe an oxygen fugacity dependence of the Mo isotope fractionation factor to an oxygen fugacity as high as ΔIW +0.47, which may allow for the presence of some Mo6+. This oxygen fugacity is in excess of what is expected for core formation in many planetesimals or planetary embryos (Righter and Drake, 1996; Wade and Wood, 2005). Hence, the Mo isotope fractionation factors determined here are applicable to the entire oxygen fugacity range of core formation. Bigeleisen and Mayer (1947) and Urey (1947) showed that equilibrium, mass dependent fractionation of isotopes decreases as a function of the square of temperature. Based on their work, the Mo isotope fractionation factor between metal and silicate melts can be described as

98/95 DDmF98−− 95 Metal Silicate D≈Mo −  ⋅1000 Metal Silicate 2 (V.8) mm98 95 T

98 95 where m98 and m95 are the atomic masses of Mo and Mo, respectively, Δm the difference between the isotopic masses, ΔFMetal-Silicate the difference in force constants of the bonds of the isotopes in the metal and silicate melts, and T the temperature in K. Implementing our average equilibrium fractionation factors for 1400 °C (-0.193 ± 0.030‰) and 1600 °C (-0.120 ± 0.020‰) in equation (V.8), we determine that the Mo isotope fractionation factor between liquid metal and liquid silicate can be described as −±×4.80( 0.55) 105 D 98/95 = MoMetal− Silicate 2 T (V.9) As can be seen in Figure V–5, equilibrium Mo isotope fractionation between metal and silicate may be analytically resolvable to over 2500 ��°C,�� with fractionation decreasing to ~0.05‰ between 2500 °C and 3000 °C. Molybdenum isotopes may therefore constitute a new tool to study the conditions of metal-silicate equi- libriation and core formation in planetary objects. It may be particularly useful for objects for which only silicate samples are available, because in many cases > 90% of a bulk body’s Mo will reside in the core. As Mo is furthermore a refractory element, this enables the use of chondrites as an analogue for both bulk and core compositions of planetary objects. Temperatures of metal-silicate segregation on the parent bodies of various achondrites could hence be estimated from chondrite and achondrite Mo isotope compositions with the above determined temperature dependence. Preliminary Mo isotope data for chondrites, angrites, and terrestrial and lunar basalts, for instance, suggest that there may indeed be variations in δ98/95Mo that are related to core formation (Burkhardt et al., 2012). Thus, ultimately it may be possible to determine the temperature conditions of core formation through the measurement of Mo isotope compositions. Such con- Page 144 — Chapter V straints would be particularly important for interpreting the significance of Hf-W isotope systematics for inferring the timing of core formation in the Earth. The calculated Hf-W ages are highly dependent on the assumed degree of metal- silicate equilibration (Halliday, 2004; Kleine et al., 2009; Rudge et al., 2010). A large degree of disequilibrium core formation would imply lower metal-silicate equilibration temperatures than a large degree of metal-silicate equilibration in a deep terrestrial magma ocean. Thus, determining the temperature of metal-silicate equilibration between the Earth’s mantle and core can provide the necessary constraint to distinguish between equilibrium and disequilibrium Hf-W models for core formation in the Earth. A large advantage of using Mo isotopes in this context, compared with the approach based on siderophile element contents, is that Mo isotopes are unambiguously dependent on temperature. Whether Mo isotope fractionation is pressure dependent, which is relevant for large bodies, remains to be investigated.

-0.50

-0.45

-0.40 -4.80(±0.55) 105/T2 -0.35

-0.30 (‰) -0.25

Metal-Silicate -0.20 Mo

98/95 -0.15 Graphite capsules ∆ -0.10

-0.05

0.05 800 1000 1200 1400 1600 1800 2000 2200 2400 2600 Temperature (°C) Figure V–5. Temperature dependence of equilibrium Mo isotope fractionation between liquid metal and liquid silicate. The triangles refer to average equilibrium fractionation factors for 1400 °C and 1600 °C with 95% confidence intervals as error bars (see section 4 for determination of averages). The solid line shows the temperature evolution of the fractionation factor; the stippled lines represent the error envelope (95% c.i.), 5 98/95 which reflects the uncertainty (0.55×10 ) of the empirical constant in Δ MoMetal-Silicate = -4.80(±0.55)×105/T2. This uncertainty was calculated by propagating the uncertainties of the average fractionation factors for 1400 and 1600 °C into equation (V.8). Chapter V — Page 145

V–6. CONCLUSIONS

Our study establishes that: 1. There is equilibrium fractionation of Mo isotopes between liquid metal and liquid silicate. At 1400 °C, the δ98/95Mo value of the silicate melt is 0.193 ± 0.030‰ (95% confidence interval) heavier than that of the metal melt, while it is 0.120 ± 0.020‰ heavier at 1600 °C. The temperature dependence of the Mo isotope fractionation factor (Δ98/95Mo) between liquid metal and liquid silicate 98/95 × 5 2 can thus be described as Δ MoMetal-Silicate = -4.80(±0.55) 10 /T (95% c.i.). 2. The addition of C or Sn to the compositional system does not affect Mo isotope fractionation between metal and silicate liquids. Furthermore, Mo isotope fractionation also appears independent of variations in silicate composition for basanitic to basaltic melt compositions. This suggests that mass dependent equilibrium fractionation of Mo isotopes might be largely independent of composition, at least for Fe-based metal and primitive silicate melts. 3. Mo isotope fractionation between metal and silicate liquids is constant over a range of oxygen fugacities from ΔIW -1.79 to +0.47. This is probably because Mo does not substantially change valence state in our experiments and appears to be dominantly present as Mo4+ throughout the entire range of oxygen fugacities, although Mo6+ may be present at the highest oxygen fugacities. 4. Experiments performed at 1600 °C in MgO capsules resulted in silicate melt compositions that were on average 0.027 ± 0.036‰ (95% c.i.) heavier than metal. These experiments suffered from kinetic enrichment of the silicate in light Mo isotopes, which yielded disequilibrium fractionation due to progressively increasing Mo solubility, which in turn was a consequence of depolymerisation of the silicate liquid by capsule-derived MgO. This effect of rapid dissolution of MgO from the capsule underlines that care should be taken when investigating equilibrium isotope fractionation in capsule material that is reactive with the sample. 5. Mass dependent Mo isotope fractionation constitutes a new tool to study the conditions of core formation in planetary bodies. Uncertain mass balance calculations are most likely not required for this due to the siderophile character of Mo, which reduces >90% of all Mo to a metallic form. This enables direct comparison of Mo isotope compositions of silicate samples to the average of chondrites. The temperatures of core formation implied by stable isotope fractionation may further constrain models of core formation, as for instance required for interpreting the Hf-W isotope record of the Earth in terms of accretion and core formation timescales.

Acknowledgements We thank Felix Oberli for the ‘Thermo2Nu’ program and maintenance of the Nu1700. Mario Fischer-Gödde was of great help for the analyses performed on the Neptune Plus in Münster. Page 146 — Chapter V

Ben Reynolds is greatly thanked for his help with the double spike deconvolution. Bruno Zürcher was indispensable for maintenance on the centrifuging piston cylinder. We appreciate the donation of an aliquot of sample AJ011 by Nicolas Greber.

V–7. APPENDIX

Table V–A.1 Elemental compositions (wt%) of selected samples with their experimental run conditions. P = pressure (GPa), T = temperature (°C), fO2 = oxygen fugacity relative to the iron-wüstite buffer (ΔIW), t = run duration in hour. Start Mix A Start Mix B1 RH55 2SD RH71 2SD RH72 2SD RH74 2SD RH76 2SD P (GPa)1 0.79 0.10 1.06 0.10 0.91 0.08 0.73 0.17 0.82 0.07 T (°C) 1400 1400 1400 1400 1400 t (h) 4.0 0.5 1.5 4.0 7.0

fO2 (ΔIW) 0.10 0.16 0.14 0.14 0.13 Capsule Graph- Graph- Graph- Graph- Graph- type ite ite ite ite ite

Metal Fe 71.6 3.0 70.9 3.7 73.2 1.7 75.4 3.7 81.1 0.9 Sn 15.8 4.1 14.5 2.8 12.5 1.1 8.79 5.19 1.46 0.34 Mo 8.38 1.05 10.4 2.0 9.86 1.00 10.7 3.1 11.6 0.7 Pt n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. Pd n.d. n.d. 0.07 0.14 0.14 0.13 0.33 1.05 0.23 0.15 Total 95.8 1.2 95.8 0.9 95.7 0.4 95.2 1.2 94.4 0.5 C (calc)2 4.2 4.2 4.3 4.8 5.6

Silicate

SiO2 42.9 0.7 43.0 1.7 42.8 1.1 43.1 0.9 43.6 0.4

Al2O3 3.92 0.26 2.45 0.24 2.47 0.13 2.43 0.12 2.46 0.10 MgO 3.92 0.20 3.05 0.21 3.10 0.18 3.12 0.13 3.14 0.11 FeO 32.8 0.9 36.7 0.7 37.1 0.6 37.1 1.3 38.7 0.5 CaO 10.6 0.1 9.61 0.34 9.55 0.26 9.56 0.38 9.76 0.20

Na2O 0.96 0.11 0.98 0.18 0.97 0.13 0.98 0.13 0.98 0.07

K2O 0.00 0.01 0.00 0.02 0.00 0.01 0.00 0.01 0.00 0.01

SnO2 3.57 0.23 3.04 1.27 3.03 0.95 2.40 0.44 0.73 0.21 Total 98.6 0.8 98.8 0.8 99.0 1.1 98.7 2.4 99.4 0.8 Chapter V — Page 147

Table V–A.1, continued. Start Mix B2 Start Mix C RH77 2SD RH82 2SD RH94 2SD RH97 2SD RH98 2SD P 0.90 0.05 0.91 0.06 0.97 0.03 0.98 0.04 1.00 0.07 (GPa)1 T (°C) 1400 1400 1400 1400 1400 t (h) 4.0 2.2 1.7 1.5 1.5 fO2 -0.70 -0.64 -1.05 -1.68 -1.71 (ΔIW) Capsule Graph- MgO MgO MgO MgO type ite

Metal Fe 64.0 10.1 67.8 4.8 70.0 1.9 68.9 1.5 68.6 2.0 Sn 23.7 8.8 22.8 4.2 15.5 2.0 21.1 0.4 21.4 1.1 Mo 9.65 1.89 9.81 1.83 10.2 1.1 10.2 1.5 9.90 0.89 Pt n.d. n.d. 0.09 0.12 0.05 0.05 0.06 0.04 0.08 0.06 Pd 2.00 2.19 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. Total 99.4 0.8 100.6 1.2 96.0 0.8 100.5 0.7 100.2 0.9 C 4.0 (calc)2 Silicate

SiO2 45.5 4.8 46.0 5.7 48.6 0.7 45.8 0.9 45.2 0.9

Al2O3 1.78 0.79 1.67 0.68 7.73 0.38 7.25 0.21 7.19 0.23 MgO 16.0 11.1 14.3 8.9 8.01 0.34 14.4 0.4 14.8 0.6 FeO 27.2 2.5 28.6 5.6 9.59 0.32 8.71 0.39 8.38 0.49 CaO 8.22 5.57 7.27 4.88 17.3 0.4 16.7 0.7 16.6 0.9

Na2O 0.63 0.36 0.56 0.34 3.94 0.20 3.66 0.26 3.66 0.21

K2O 0.25 0.08 0.24 0.12 0.98 0.07 0.92 0.15 0.91 0.15

SnO2 0.74 0.51 1.14 0.84 0.74 0.31 0.26 0.10 0.22 0.07 Total 100.4 1.3 99.7 1.6 96.9 1.0 97.8 0.8 97.0 1.1 Page 148 — Chapter V

Table V-A.1 continued Start Mix C Start Mix D Start Mix E Start Mix B2 RH104 2SD RH101 2SD RH99 2SD RH80 2SD RH87 2SD P 0.91 0.04 0.95 0.04 0.89 0.05 1.00 0.04 1.06 0.02 (GPa)1 T (°C) 1400 1400 1400 1600 1600 t (h) 4.0 1.7 1.5 1.1 0.7 fO2 -1.79 0.47 -0.52 -0.70 -0.62 (ΔIW) Capsule Graph- MgO MgO MgO MgO type ite

Metal Fe 68.9 2.1 73.8 0.5 72.4 5.1 67.4 10.1 66.5 9.6 Sn 20.5 1.3 9.04 0.43 26.2 5.3 23.0 11.2 22.6 8.3 Mo 10.3 0.9 11.8 0.5 1.66 0.13 9.63 3.04 10.9 2.3 Pt 0.09 0.09 0.01 0.03 0.04 0.05 0.13 0.24 0.09 0.10 Pd n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. Total 100.0 0.9 94.9 0.4 100.3 0.4 100.2 1.1 100.3 1.1 C 5.1 (calc)2

Silicate

SiO2 46.2 0.5 39.9 0.8 42.4 5.6 37.6 1.5 40.6 2.7

Al2O3 7.14 0.22 1.22 0.12 1.65 0.94 1.79 0.55 1.89 0.80 MgO 16.4 1.0 0.85 0.10 12.5 11.3 24.4 6.2 18.3 5.2 FeO 7.89 0.36 49.0 1.6 32.5 4.8 26.6 6.4 28.7 7.8 CaO 16.8 1.2 3.39 0.42 7.91 6.92 7.75 1.63 8.93 1.17

Na2O 3.64 0.12 0.83 0.12 0.66 0.43 0.65 0.20 0.65 0.10

K2O 0.86 0.12 0.01 0.02 0.27 0.11 0.28 0.08 0.24 0.07

SnO2 0.21 0.09 4.06 0.39 0.94 0.77 1.07 0.46 1.27 0.24 Total 99.1 0.7 99.2 1.3 98.8 1.7 100.1 0.9 100.7 0.9 Chapter V — Page 149

Table V-A.1 continued Start Mix B2 Start Mix F RH92 2SD RH103 2SD RH81 2SD RH86 2SD RH102 2SD P (GPa)1 0.98 0.08 0.95 0.17 0.98 0.04 0.95 0.02 0.10 0.09 T (°C) 1600 1600 1600 1600 1600 t (h) 1.7 1.4 1.1 0.8 1.2 fO2 0.12 0.06 -0.72 -0.73 0.06 (ΔIW) Capsule Graph- Graph- Graph- MgO MgO type ite ite ite

Metal Fe 73.0 1.1 73.2 1.2 92.8 4.0 89.2 4.6 85.4 0.8 Sn 12.3 1.2 12.7 0.6 0.00 0.00 0.00 0.00 0.00 0.00 Mo 9.67 0.88 9.43 0.86 7.50 3.68 9.79 4.22 9.21 0.76 Pt 0.04 0.05 0.03 0.04 0.11 0.13 0.10 0.09 0.04 0.05 Pd n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. Total 95.3 0.9 95.6 0.7 100.4 0.9 99.3 0.6 94.8 1.1 C (calc)2 4.7 4.4 5.2

Silicate

SiO2 47.7 2.2 48.2 2.1 37.7 1.6 38.5 0.9 46.9 0.6

Al2O3 1.83 0.10 1.77 0.15 2.64 0.68 2.45 1.19 1.97 0.12 MgO 1.25 0.13 1.21 0.08 18.5 4.8 18.3 4.0 1.32 0.09 FeO 34.5 1.8 33.8 1.3 30.8 5.0 30.4 5.1 40.4 0.6 CaO 7.91 0.21 8.00 0.19 9.54 1.48 9.62 0.74 8.05 0.19

Na2O 0.62 0.11 0.64 0.10 0.94 0.29 0.86 0.24 0.64 0.10

K2O 0.24 0.07 0.24 0.04 0.33 0.11 0.33 0.17 0.23 0.03

SnO2 4.36 0.64 4.39 0.76 0.00 0.00 0.01 0.02 0.00 0.00 Total 98.4 0.9 98.2 0.7 100.3 0.7 100.4 0.7 99.5 0.9

V–7.1. Mo concentration analyses by laser ablation ICPMS Molybdenum concentrations in silicates of selected samples were analysed by laser ablation ICPMS (LA-ICPMS; Table V–A.2). The purpose of these analyses was to investigate the feasibility of isotopic analyses, but the data obtained here are provided for comparison with concentrations obtained from double spiked isotope ratio analyses (MC-ICPMS). The samples RH75, RH91 and RH93 were analysed by LA-ICPMS, but not by MC-ICPMS. Samples that are presented in Table V–3 of the main manuscript but lacking in Table V–A.2 were only analysed by MC-ICPMS. Analyses by LA-ICPMS were Table V–A.2 LA-ICPMS concentration data performed with a Geolas excimer for Mo. Except RH86 and RH87, all data were laser with a wavelength of 193 internally normalised to CaO concentrations nm, coupled to a quadrupole determined by EPMA. RH86 and RH87 were ICPMS (Perkin Elmer Elan 6100 internally normalised to FeO, because of heterogeneity in other oxide concentrations due DRC). The laser was operated at to large olivine quench needles. n = number of 15 Hz with a spot size of ~40 µm repeated analyses. and an energy fluence of ~12 J cm-2. Ablated material was trans- [Mo] (ppm) [Mo] (ppm) Sample SD n 1 ported from the sample cell to the LA-ICPMS MC-ICPMS RH55 44.3 2.4 4 52.5 ICP torch by an N2 flux. The NIST SRM610 glass standard was RH77 42 20 3 58.2 analysed before and after each RH86 127 48 4 108 sample to correct for mass bias RH87 269 108 4 193 drift. All analyses were internally normalised to CaO concentra- RH92 45.4 5.4 3 52.6 tions determined by electron RH94 6.1 0.8 3 9.61 microprobe analyses, except for RH97 6.3 5.3 3 6.09 samples RH86 and RH87. For RH98 7.3 4.0 3 7.61 their analyses, internal normali- RH99 11.3 7.7 3 11.0 sation to FeO concentrations was RH75 5.1 0.6 3 n.d. found to be most satisfactory due to heterogeneity in other oxide RH93 111.4 0.5 3 n.d. concentrations as a result of large 1 Relative uncertainty (SD) better than 1%. olivine quench needles. As can be seen in Table V–A.2 and Figure V–A.1, average Mo concentrations determined by LA-ICPMS generally match well with bulk Mo concentrations determined by MC-ICPMS. This confirms that the careful metal-silicate separation procedure involving centrifugation, sawing, polishing, crushing and microscopic inspection produced metal-free silicate pieces for analyses by MC-ICPMS. Chapter V — Page 151

103 V–7.2. Double spike Graphite capsule calibration MgO capsule We have added a 100Mo- 97Mo double spike to our 2 samples prior to digestion 10 to correct for instrumental mass bias as well as potential mass dependent isotope fractionation during 101 chemical separation. Both 97 100

Mo-enriched and Mo- [Mo] MC−ICPMS (ppm) enriched metal powders were purchased from the 100 Oak Ridge National Labo- 100 101 102 103 100 ratory (Batch number Mo [Mo] LA−ICPMS (ppm) 159992, Batch number 97 Mo 159791). After dis- Figure V–A.1 Silicate melt Mo concentrations solution, their isotopic determined by isotope ratio analyses plotted against compositions and concen- concentrations determined by LA-ICPMS. The trations were determined methods agree well. Error bars for LA-ICPMS data relative to an in-house are 1SD; those for concentrations determined by gravimetric Mo standard MC-ICPMS are consistently smaller than symbol sizes. (Mo metal powder from Alfa Aesar; LOT number C24P28). The optimum double spike mixture was then determined from the abundances of the single spikes using the ‘double spike toolbox’ of Rudge et al (2009). The gravimetrically prepared double spike was subsequently analysed for its isotopic composition and corrected for instrumental mass bias by bracketing with an Mo standard of natural isotopic composition. The instrumental fractionation factor in the bracketing standards was determined by internal normalisation to a 97Mo/95Mo ratio of 0.602083 (Lu and Masuda, 1994). The isotopic composition of the same Mo standard was also internally normalised to provide the isotopic composition of the reference standard in the double spike deconvolution. This calibration was initially performed on Nu1700 at ETH Zurich against the in-house Mo standard (Alfa Aesar). In an attempt to establish an international reference standard (Greber et al., 2012), however, the calibration was repeated on a Neptune Plus (Münster) against the NIST SRM 3134 Mo standard relative to which all previously performed analyses have been recalculated. The results of these calibration procedures have been verified with double spike – standard mixtures with various proportions of double spike ranging from 0.185 to 0.806, with 0.477 as an optimum spike proportion. No trends of isotopic Page 152 — Chapter V compositions are observed as a function of double spike proportion, confirming the accuracy of the calibration (Figure V–A.2).

0.3

0.2 NIST SRM3134 Standard with Neptune 0.1 2SD = 0.025‰

-0.1 Mo (‰) 98/95

δ -0.2

-0.3

-0.4 Alfa Aesar Mo Standard with Nu1700 2SD = 0.119‰ -0.5 0 0.2 0.4 0.6 0.8 1 Proportion of D.S. in D.S. − standard mixture Figure V–A.2 Plot of δ98/95Mo against spike proportion in various double spike – standard mixtures. The isotopic compositions are shown for mixtures with two different standards: the in-house Alfa Aesar Mo standard on Nu1700 (Zurich; diamonds) and NIST SRM 3134 on Neptune Plus (Münster; circles). Isotopic compositions are independent of spike proportions in either data set, confirming the accuracy of the double spike and standard calibrations. Individual error bars are 2SE internal errors.

6.5 Theoretical slope for 4+ valence Graphite capsule 6.0 MgO capsule

5.5 Theoretical slope for 6+ valence n/2

5.0 Mg/Si ~ 0.45 MoO /X

Mo 4.5 Mg/Si ~ 0.6, 45 min

log a Mg/Si ~ 0.7, 4.0 60 min Mg/Si ~ 0.09 Depolymerization effect 3.5 of MgO capsules at 1600°C

3.0 -2.5 -2.0 -1.5 -1.0 -0.5 0 0.5 1.0 log fO2 (∆IW) Figure V–A.3 Determination of Mo valence state as a function of oxygen fugacity; as Figure V–4b in main manuscript. The data are plotted here with variable activity coefficients after Holzheid et al. (1997) and Richardson (1956): 1.7 for FeO contents <30 wt%, 1.3 for FeO between 30 and 40 wt%, and 1.0 for FeO contents >40 wt%. This approach suggests that Mo6+ is present above ΔIW -0.5. Chapter V — Page 153

600 Figure V–A.4 Molybdenum 500 content against MgO content. The presented data are from experiments 400 performed in MgO capsules at an oxygen fugacity of ΔIW -0.64 to 300 -0.73. Open symbols are for experiments performed at 1400 °C, closed Mo (ppm) 200 symbols for 1600 °C. All errors 100 (2SD) are smaller than twice the symbol sizes. 0 12 14 16 18 20 22 24 26 MgO (wt%)

CHAPTER VI.

MOLYBDENUM STABLE ISOTOPE COMPOSITION OF METEORITES: CONSTRAINTS ON PLANETARY CORE FORMATION1

1 This work was part of the PhD thesis of C. Burkhardt at ETH Zurich (PhD thesis number 20454) and this updated version of his original chapter 4 (part B) is presented here because of its connec- tion to chapter V in the current thesis. An extended version of this chapter will be published as: C. Burkhardt, R. C. Hin, T. Kleine. Molybdenum stable isotope composition of meteorites: Constraints on planetary core formation. Page 156 — Chapter VI

ABSTRACT

Significant Mo isotope fractionation between liquid metal and liquid silicate as determined in an accompanying study reveal the potential of Mo isotopes to constrain the conditions of core formation in planetesimals and terrestrial planets. Here, we present Mo stable isotope data for a variety of meteorites, including chondrites, iron meteorites, eucrites, angrites and a martian meteorite as well as lunar and terrestrial samples. Chondrites and iron meteorites define a common δ98/95Mo of –0.17 ± 0.02 ‰ relative to the NIST SRM 3134 Mo standard. The Mo isotope compositions of the silicate samples are heavier to variable degrees, which at least in part reflects Mo stable isotope fractionation during core formation. However, as is evident from the heavy Mo isotope composition of one angrite (NWA 4801) and one shergottite (DaG 476), other processes such as terrestrial weathering have overprinted the Mo isotope signatures of core formation. Nevertheless, when excluding samples whose Mo isotope signatures were modified during weathering, the temperatures of metal-silicate equilibration during core formation can be estimated. Assuming metal-silicate equilibrium and instantaneous core formation, temperatures between ~2100 °C (Earth and Moon) and ~1800 °C (angrites) are obtained, consistent with metal-silicate equilibration in a magma ocean. These new results suggest that core formation in the Earth did not occur at the base of a deep magma ocean, but rather took place at a shallower depth, either during descent of metal droplets through the magma ocean or by metal-silicate equilibration in Earth’s precursor bodies. The heavier Mo isotope composition of the eucrites reflect either core formation by percolation at lower temperatures (~1000 °C) or additional post-core formation Mo isotope fractionation within the eucrite parent body. This study demonstrates that once the different processes that can fractionate Mo isotopes at high temperatures are fully understood, Mo isotopes provide a powerful tool to investigate the conditions of core formation in planetary bodies.

VI–1. INTRODUCTION

The major chemical differentiation process in terrestrial planets and in many planetesimals is the separation of a metallic core from a silicate mantle. Core formation requires large-scale melting to facilitate efficient metal-silicate separation and metal segregation. The energy required for melting has been provided either by the decay of short-lived, now-extinct 26Al or by the release of gravitational energy during collisional accretion, or both (Rubie et al., 2007). Core formation results in a strong chemical fractionation of siderophile elements, which preferentially partition into the core, from lithophile elements, which remain in the mantle. The magnitude of this chemical fractionation depends, for a given element, on the pressure, temperature and oxygen fugacity under which metal and silicate separate (Wood et al., 2006). Thus, investigating the distri- Chapter VI — Page 157 bution of siderophile elements in the mantle of a differentiated planetary body can help to constrain the conditions under which the metal core segregated. For instance, the observation that the Ni and Co partition coefficients converge at high pressures to values consistent with the observed Ni and Co abundances in the Earth’s mantle has led to the idea that core formation in the Earth occurred under elevated pressures and temperatures in a deep magma ocean (Li and Agee, 1996). Current models of terrestrial planet formation predict that the Earth (and other terrestrial planets) accreted by collision of several Moon- to Mars-sized planetary embryos (Chambers, 2004; Stevenson, 2008). These embryos had already undergone core formation, raising the question as to whether their metal cores first re-equilibrated with the Earth’s mantle before entering Earth’s core or, alternatively, directly merged with Earth’s core without any prior equilibration with Earth’s mantle. Which of these two scenarios is more appropriate for the Earth is currently unclear, but is critical for understanding core formation in the Earth. Furthermore, application of the Hf-W system to constrain the timing of core formation critically depends on the degree of equilibrium during metal-silicate separation (Kleine et al., 2004; Rudge et al., 2010). Independent constraints on the physico-chemical conditions of core formation are therefore required. Such constraints can in principal be obtained from experimental element partitioning studies in conjunction with observed depletion of siderophile elements in the Earth’s mantle (Wood et al., 2006). However, the abundances of siderophile elements in the Earth’s mantle, at least for the Earth, have been shown to be equally consistent with both equilibrium and disequilibrium core formation (Rudge et al., 2010). In recent years, mass dependent isotope fractionation during metal-silicate separation has emerged as a new tool to study the conditions of core formation. Silicon and Cr isotopes appear to fractionate as a result of metal-silicate separation and have been used to argue for partitioning of Si and Cr into the Earth’s core (Armytage et al., 2011; Fitoussi et al., 2009; Georg et al., 2007; Moynier et al., 2011). Since the degree of isotope fractionation in general depends on the temperature of metal-silicate equilibration, such isotope studies can provide important new information regarding the conditions of core formation. We have recently performed liquid-metal liquid-silicate experiments, which demonstrate that Mo isotopes fractionate during metal-silicate equilibration (chapter 5). These experimental results indicate that, with current analytical precision, metal-silicate fractionation may be detectable up to metal-silicate equilibration temperatures of ~2500 °C. Here we present the first stable Mo isotope data for meteorites and terrestrial mantle-derived rocks, obtained by the double spike technique. These data are used to assess the extent of Mo isotope fractionation in meteorites, and to explore the potential of Mo isotopes as a new tracer of core formation in planetesimals and terrestrial planets. Page 158 — Chapter VI

VI–2. ANALYTICAL METHODS VI–2.1. Sample preparation and Mo separation The samples investigated for this study include four chondrites [the carbonaceous chondrites Allende (CV3) and Murchinson (CM2); the ordinary chondrite Kernouvé (H6); and the enstatite chondrite Daniel’s Kuil (EL6)], two iron meteorites [Gibeon (IVA); Cape of Good Hope (IVB)], two angrites (D’Orbigny, NWA 4801), and two eucrites (Dhofar 007; Millbillillie). In addition, one martian meteorite (DaG 476) as well as several terrestrial rock standards (BHVO-2, BIR-1, W-2a, DNC-1) and two lunar samples (low-Ti Mare basalt 12004; KREEP-rich highland breccia 68115) were analyzed. All samples were obtained as small chips and were cleaned carefully with abrasive paper and by ultrasonication in distilled water and ethanol. Three of the samples are meteorites found in the desert (Dhofar 007, DaG476, NWA4801) and show signs of terrestrial alteration, such as oxidation veins along cracks. When possible, obviously altered parts of these samples were avoided. For NWA 4801, however, this was not possible, because all of the available material was weathered. In an attempt to minimize the effects of terrestrial weathering, only grain-sizes larger than 40 μm were used for this sample. All the silicate samples were powdered in an agate mortar or in an agate mill. Except for Murchinson, all the chondrites are aliquots from the powders analyzed previously for mass independent (i.e., nucleosynthetic) Mo isotope anomalies (Burkhardt et al., 2011). The KREEP-rich sample is a left-over from an earlier W isotope study (Touboul et al., 2007), for which some metal was separated from the powder. All silicate samples were weighed into 60 ml Savillex beakers, spiked with 100 97 a Mo- Mo tracer, and digested in HCl (iron meteorites) or HF-HNO3-HClO4 (6:3:1) (silicate samples) at 180°C on a hotplate for several days. This procedure ensured complete spike-sample equilibration. After digestion, samples were dried down and re-dissolved in HNO3-H2O2 to attack any remaining organics. Molyb- denum was separated from the sample matrix using a three-stage ion exchange procedure employing both anion and cation exchange resins (Burkhardt et al., 2011). Molybdenum (together with Ti, W, HFSE and Ru) was first separated from major elements using a cation exchange column (Bio-Rad AG 50W-X8) and 1 M HCl-0.1 M HF (Patchett and Tatsumoto, 1980). Further separation of Mo from the remaining elements was achieved in HCl-HF media using an anion exchange resin (Bio-Rad AG 1-X8). Any remaining Fe, Ni and a large part of the Ru was removed from the resin with 1 M HF, while Ti, Zr, Hf and W were eluted using different HCl–HF mixtures (Kleine et al., 2004). Finally Mo together with some remaining Ru was eluted using 3 M HNO3. The remaining Ru was removed from the Mo fraction using TRU-Spec resin and 7 M HNO3, followed by elution of Mo with 0.1 M HNO3. This last step was repeated using 1 M HCl and 0.1 M HCl instead of 7 M HNO3 and 0.1 M HNO3, resulting in Ru/Mo and Zr/Mo < Chapter VI — Page 159

3×10-5 for most samples. Total procedural blanks ranged from 0.3 to 0.5 ng Mo and were negligible for all samples, for which ~30-100 ng Mo were analyzed. Total chemistry yields were 60–85 %. VI–2.2. Mass spectrometry and data reduction Molybdenum isotope measurements were performed using a Nu Instru- ments Plasma 1700 MC-ICP-MS at ETH Zurich and a Thermo Neptune Plus MC-ICPMS at the University of Münster. On both instruments all seven Mo isotopes were measured simultaneously in static mode, as well as 90Zr (Zurich) or 91Zr (Münster) and 99Ru to monitor potential isobaric interferences. On the Nu 1700, samples were introduced via a DSN-100 desolvator equipped with a PFA nebulizer with an uptake rate of ~140 μl min-1. For a normalized 100 μl min-1 uptake rate this setup resulted in a sensitivity of ~180 V per ppm Mo. Here, measurements were usually performed with 70-100 ppb solutions. For measurements with the Neptune, a Cetac AridusII desolvater equipped with a PFA nebulizer having an uptake rate of ~80 μl min-1 was used. For a normalized 100 μl min-1 uptake rate this setup resulted in a sensitivity of ~430 V per ppm Mo, so that measurements were made at lower concentrations (30–70 ppb Mo) than on the Nu 1700. The Mo isotope measurements consisted of 30 s (Neptune Plus) or 60 s (Nu1700) baseline integrations (made on-peak using the same 0.5 M

HNO3–0.05 M HF solution also used for the samples) and 40 isotope ratio measurements of 5 s each. To monitor any potential drift in the measured Mo isotope ratios, the sample measurements were interspersed with measurements of the SRM 3134 Mo standard. For the measurements on the Nu 1700, data reduction and calculation of δ98/95Mo values was performed on-line following the geometrically motivated iteration procedure for the double spike inversion of Siebert et al. (2001). Data reduction of the data obtained on the Neptune was made off-line using the “double spike toolbox” (Rudge et al., 2009), which employs MATLAB’s non- linear equation solving routine for the iteration. To validate that both routines give consistent results, some of the Nu 1700 data were reduced with the “double spike toolbox” and some of the Neptune data were converted and re-analyzed off-line with the Nu 1700 software. Both methods yield identical results to within 0.005‰, even for samples with relatively large interference corrections. The Mo isotope data are reported as δ98/95Mo values, which have been calculated from αSample as given from the double spike inversion. The mean α of Mean bracketing double-spiked-SRM standard runs (αSRM 3134 ) was subtracted from the value for αSample to correct for short term drift in the Mo isotope measurements:  98/95 Mean m98 (VI.1) dMo =1000( ααSample− SRM 3134 )ln m95 98 95 where m98 and m95 are the atomic weights of Mo and Mo. The external Page 160 — Chapter VI reproducibility of the SRM standard measurements over the course of this study is 0.085‰ (2SD) for the Nu 1700 and 0.049‰ (2SD) for the Neptune.

VI–3. RESULTS VI–3.1. Mo concentrations The Mo concentration and isotopic data are reported in Table VI–1 and the δ98/95Mo values are plotted on Figure VI–1. The Mo concentrations in the meteorites vary from ~5-22 ppm in the iron meteorites, to ~1.1-1.4 ppm in the chondrites, to values of a few tens of ppb in the extraterrestrial mantle-derived samples. Among the latter, angrites tend to have the highest Mo concentrations of ~100-140 ppb, while the eucrite Millbillillie exhibits the lowest concentration of only ~7 ppb Mo. The Mo concentrations obtained in the present study agree to within ~5% with the values reported in Burkhardt et al. (2011) for the same powders/samples. For the angrite NWA4801 a ~30% lower concentration was determined here compared to our previous work (Burkhardt et al., 2011) and this most likely reflects the fact that here the grain size fraction <40 μm was excluded from the analysis. The Mo concentrations determined here for three eucrites show a similar range to that obtained in a previous study on other eucrite samples (Newsom, 1985). The two lunar samples have Mo concentrations at the lower end of values reported previously (Newsom, 1986). However, the Mo concentration of 44 ppb obtained for the KREEP-rich highland breccia 68115 is only a lower limit because from this sample metal was removed in a previous study (Touboul et al., 2007). For low-Ti mare basalt 12004 the Mo concentration is about three times lower than that determined by Neal (2001) for another fragment of this rock, presumably due to sample heterogeneities. We have also analyzed four international rock standards and for three of them (BIR-1, DNC-1, W-2a) report the first high-precision Mo concentration data. The BHVO-2 basalt standard was digested five times and each of these digestions was analyzed multiple times. The measured Mo concentrations range from 3.84 to 5.46 ppm, indicating that the BHVO-2 sample powder is heterogeneous with regard to Mo. The average concentration of 4.55 ppm obtained from the five independent digestions agrees with literature values (~4 ppm) as given in the GeoReM database. VI–3.2. Mo isotope data and δ98/95Mo values The Mo isotope data are given as δ98/95Mo values calculated relative to the 98/95 SRM 3134 Mo standard, where δ Momeasured represents the measured isotope 98/95 fractionation as obtained from the double spike inversion, while δ Mocorrected is the purely mass dependent isotope fractionation obtained after correction for mass independent (i.e., nucleosynthetic) isotope anomalies. The double spike inversion of a sample run iteratively calculates a fractionation factor α relative Chapter VI — Page 161 to the SRM 3134 standard by assuming that the difference in the isotopic composition of the sample and the standard is entirely mass dependent. However, for

Table VI–1. Molybdenum concentration and isotopic composition of meteoritic and terrestrial samples. Moa δ98/95Mo c δ98/95Mo d Sample Type Nb measured corrected [ng/g] [‰] [‰] Chondrites Murchinson CM 1193 5 -0.20 ± 0.03 -0.04 ± 0.05 NWA2090 CO 1257 6 -0.15 ± 0.04 -0.13 ± 0.05 Allende CV 1439 4 -0.26 ± 0.06 -0.18 ± 0.07 Gujba (metal) CB 5273 6 -0.26 ± 0.04 -0.21 ± 0.04 Kernauvé OC 1319 5 -0.17 ± 0.04 -0.16 ± 0.04 Daniel’s Kuil EC 1103 5 -0.19 ± 0.02 -0.18 ± 0.02

Iron meteorites Odessa IAB 6863 6 0.22 ± 0.02 0.22 ± 0.03 Negrillos IIAB 6425 6 -0.17 ± 0.02 -0.14 ± 0.02 Henbury IIIAB 5739 6 -0.15 ± 0.04 -0.13 ± 0.04 Gibeon IVA 5472 5 -0.17 ± 0.04 -0.16 ± 0.04 Cape of Good Hope IVB 22470 5 -0.25 ± 0.06 -0.20 ± 0.06 Chondrites+irons mean -0.20 ± 0.02 -0.17 ± 0.02

Eucrites Millbillillie mmict. eucrite 7.3 6 0.16 ± 0.08 0.16 ± 0.08 Stannern mmict. eucrite 11 4 -0.42 ± 0.28 -0.42 ± 0.28 Dhofar 007 cm. eucrite 83 4 0.26 ± 0.04 0.26 ± 0.04 Eucrites mean 0.20 ± 0.06 0.20 ± 0.06 Angrites NWA4801 Angrite 100 2 0.49 ± 0.01 0.49 ± 0.01 D’Orbigny Angrite 136 6 -0.05 ± 0.04 -0.05 ± 0.04

Mars DAG467 Shergottite 53 2 0.43 ± 0.06 0.43 ± 0.06

Moon 12004.142 Low -Ti Marebasalt 25 2 -0.09 ± 0.03 -0.09 ± 0.03 68115.112 KREEP-rich breccia 44 2 -0.10 ± 0.03 -0.10 ± 0.03 Moon mean -0.09 ± 0.03 -0.09 ± 0.03 Page 162 — Chapter VI

Table VI–1, continued.

a 98/95 c 98/95 d Mo δ Momeasured δ Mocorrected Sample Type Nb [ng/g] [‰] [‰]

Terrestrial basalts BHVO-2 Hawaii basalt 4545 24 -0.06 ± 0.03 -0.06 ± 0.03 BIR-1 Iceland ol-tholeite 32 3 -0.11 ± 0.03 -0.11 ± 0.03 DNC-1 N. Carolina ol-dolerite 121 3 -0.18 ± 0.06 -0.18 ± 0.06 W-2a Virgina diabase 393 2 -0.05 ± 0.06 -0.05 ± 0.06 Terrestrial basalts mean -0.10 ± 0.10 -0.10 ± 0.10 e MoS2 Bern Molybdenite 6 0.70 ± 0.06 a Mo concentration as determined with the double spike. b Number of isotopic analyses. c Measured Mo isotopic mass fractionation relative to the NIST SRM 3134 Mo standard (aliquots are provided upon request). Uncertainties are two-sided 95% Student-t distributions (t0.95,n-1σ/√n) for group means and samples measured more than twice and 2SD for samples measured twice. d Mo isotope 98/95 98/95 mass fractionation corrected for nucleosynthetic effects using δ Mocorrected = δ Momeasured – 96 ε Mo(7/5)nucleosynthetic × 0.066 and nucleosynthetic anomaly data from Burkhardt et al. (2011). Errors are propagated assuming a 20% uncertainty on the correction. e Molybdenite AJ011 from the Bern 98/95 98/95 group (Greber et al., 2011). δ MoJMC Bern = δ MoNIST SRM 3134 + 0.25 ‰ (Greber et al., 2012). samples having mass independent isotopic anomalies relative to the standard, such as iron meteorites and chondrites (Burkhardt et al., 2011; Dauphas et al., 2002) this routine will lead to false α values and thus to incorrect δ98/95Mo. The 98/95 δ Momeasured values, therefore, must be corrected for nucleosynthetic isotope effects to obtain the purely mass dependent Mo isotope fractionation. The effects of nucleosynthetic Mo isotope anomalies on the measured δ98/95Mo values can be corrected using two different approaches, both of which require knowledge of the isotopic composition of the unspiked sample. For most of the samples analyzed in this study this information is available from our previous Mo isotope study (Burkhardt et al., 2011). The effects of nucleosynthetic isotope anomalies can then be corrected by using these Mo isotope compositions of the samples instead of that of the SRM 3134 standard in the double spike inversion. This procedure will provide δ98/95Mo values corrected for nucleosynthetic isotope anomalies. Alternatively, the measured δ98/95Mo values obtained from the double spike inversion using the SRM 3134 standard composition can subsequently be corrected using the independently measured nucleosynthetic isotope anomaly of each sample. In the first approach the Mo isotopic data from Burkhardt et al. (2011) were internally normalized to 97Mo/95Mo = 0.602083 (Lu and Masuda, 1994), the same ratio that was also used to calculate the relative Mo isotope abundances of the SRM 3134 standard. The Mo isotope abundances thus obtained for each Chapter VI — Page 163 sample were then used in the double spike inversion instead of the abundances of the SRM 3134 standard. Figure VI–2 illustrates the results obtained from this inversion procedure for the IVB iron meteorite Cape of Good Hope. The blue line in Figure VI–2a shows the δiMo values relative to the SRM 3134 standard as obtained from using the internally normalized values in the inversion. These δiMo values include both, mass dependent and mass independent Mo isotope variations. The red curve is the pure mass independent part as obtained from the unspiked run for normalization to 97Mo/95Mo (Burkhardt et al., 2011). The true mass dependent fractionation is then obtained as the difference between the blue and the red curve, represented as the black line in Figure VI–2b. Also shown in Figure VI–2b are the δiMo values calculated in the double spike inversion with reference to the composition of the SRM 3134 standard (green line). For Cape of Good Hope, the difference between these two results corresponds to ~0.05‰ δ98/95Mo, indicating that for this sample the measured δ98/95Mo of –0.25‰ needs 98/95 to be corrected upwards to δ Mocorrected = –0.20‰ (see Table VI–1).

Mars DAG 476 shergottite Dhofar 007 cm. Eucrites Stannern Millbillillie mmict. Angrites D’Orbigny NWA 4801 Moon 68115.112 KREEP-rich Breccia 12004.142 Low-Ti Mare Basalt W-2a Terrestrial DNC-1 basalts BIR-1 BHVO-2 Odessa IAB Negrillos IIAB Henbury IIIA Iron Gibeon IVA meteorites Cape of Good Hope IVB + Daniels Kuil EL6 Chondrites Kernouvé H6 Gujba (metal) CB Allende CV3 NWA 2090 CO Murchinson CM2 ( ) -0.8 -0.6 -0.4 -0.2 0 0.2 0.4 0.6 δ98/95Mo (‰)

Figure VI–1. Mo isotope fractionation of planetary materials relative to the SRM 3134 Mo standard after correction for nucleosynthetic anomalies. Uncertainties on single data point and group means as given in Table 4.1, except for the angrites where the shaded area represents total observed variation. Page 164 — Chapter VI

The measured δ98/95Mo values can also be corrected for nucleosynthetic isotope anomalies after the double spike inversion. The nucleosynthetic Mo isotope anomalies in bulk meteorites are caused by a heterogeneous distribution of s-process Mo (Burkhardt et al., 2011; Dauphas et al., 2002). We, therefore, 98/95 modeled the changes in δ Momeasured as a function of the size of the nucleosynthetic anomalies by adding or subtracting s-process Mo (Arlandini et al., 1999) from the Mo isotope abundances of the SRM 3134 standard. From this experiment we obtained the empirical relationship 98/95 98/95 96 δ Mocorrected = δ Momeasured – ε Mo(7/5)nucleosynthetic × 0.066 (VI.2) 96 96 where ε Mo(7/5)nucleosynthetic is the nucleosynthetic anomaly for ε Mo (normalized to 97Mo/95Mo). For Cape of Good Hope, this equation provides exactly the same δ98/95Mo value as obtained above (corresponding to a 0.05 ‰ correction of the measured δ98/95Mo value). Table VI–1 summarizes both δ98/95Mo values obtained by using the SRM 3134 98/95 98/95 composition in the double spike inversion (δ Momeasured), and δ Mo corrected 98/95 for mass independent Mo isotope variations (δ Mocorrected). Among the chondrites and iron meteorites analysed for the present study, the smallest correction of 0.01‰ was necessary for the enstatite chondrite Daniel’s Kuil, while the largest correction of 0.16‰ was applied to the Murchinson data. After correction, all the chondrites (except Murchinson) and magmatic iron meteorites have indistinguishable δ98/95Mo values averaging at –0.17 ± 0.02‰ (95 % confidence interval). Only Murchinson and the non-magmatic iron meteorite Odessa exhibit 98/95 a δ Mocorrected value different from the other chondrites and iron meteorites. The relatively heavy isotopic composition of the IAB iron meteorite Odessa might be related to its non-magmatic origin. It was suggested that non-magmatic iron meteorites may have formed during impacts (Choi et al., 1995). The heavy Mo isotope composition of Odessa could then, for instance, expose its origin as an evaporation residue. More data on non-magmatic iron meteorites are required to substantiate this interpretation, but for present purposes we have excluded Odessa from calculating a mean δ98/95Mo of chondrites and iron meteorites. The difference for Murchinson most likely reflects sample heterogeneity, because the Murchinson powder used in the present study is different from that used to measure nucleosynthetic Mo isotope anomalies (Burkhardt et al., 2011). Note that for the other chondrites the same powders have been used for both measurements. Measured nucleosynthetic isotope anomalies in Murchinson can vary on the sampling scale due to a heterogeneous distribution of isotopically diverse, presolar components, and are also sensitive to the digestion method, some of which may leave certain presolar components incompletely dissolved (Burkhardt 98/95 et al., 2012; Burkhardt et al., 2011). The comparatively high δ Mocorrected value for Murchinson, therefore, most likely is due to an overcorrection for this sample [i.e., the powder used for determining δ98/95Mo has a smaller nucleosynthetic Chapter VI — Page 165 isotope anomaly than that reported in Burkhardt et al. (2011) for another powder of Murchinson]. Alternatively to sample heterogeneity, the composition of Murchinson could be part of a trend in the Mo isotope compositions of the carbonaceous chondrites, progressing from heavy to light in the order CM, CO, CV, CB. This trend resembles the volatility trend in the carbonaceous chondrites, CM being the least volatile depleted of the four groups mentioned above. However, with the current data set, the trend is not resolvable given the uncertainty in the accuracy of the correction for Murchinson. Re-measuring Murchinson for its nucleosynthetic as well as mass dependent Mo isotope composition on a single powder should resolve this issue. Furthermore, we will analyse a CI chondrite, because CI chondrites are classified as least volatile element depleted. Currently, however, we exclude Murchinson from calculating the mean δ98/95Mo of chondrites and iron meteorites. In contrast to the chondrites and iron meteorites, no resolvable nucleosynthetic Mo isotope anomalies have been observed in angrites and martian meteorites (Burkhardt et al., 2011). Moreover, although no eucrites and lunar samples have yet been investigated for nucleosynthetic Mo isotope anomalies, no significant mass independent Mo isotope variations are expected for these samples (Burkhardt et al., 2011). Therefore, no correction had to be applied for these samples. Most of the terrestrial samples (excluding DCN-1), the two lunar 0.6 0.6 a b mass dependent + 0.4 mass independent 0.4 fractionation 0.2 0.2 0 0 true Mo (‰) Mo (‰) i i fractionation δ mass independent δ -0.2 effect only -0.2 measured -0.4 -0.4 fractionation 92 94 95 96 97 98 100 92 94 95 96 97 98 100 iMo iMo Figure VI–2. Effect of nucleosynthetic anomalies on stable Mo isotopic fractionation exemplified for the IVB iron meteorite Cape of Good Hope. The broken line in Figure 4.2a represents the δiMo values relative to the SRM 3134 standard as obtained from using the internally normalized values in the inversion, including mass dependent and mass independent Mo isotope variations. The stippled curve is the pure mass independent part as obtained from the unspiked run for normalization to 97Mo/95Mo (Burkhardt et al., 2011). The true mass dependent fractionation is the difference between the broken line and the stippled curve, represented as the solid line in Figure 4.2b. Also shown in Figure 4.2b are the δiMo values calculated in the double spike inversion referenced to the composition of the SRM 3134 standard (broken line). For Cape of Good Hope the difference between these two results corresponds to ~0.05‰ δ98/95Mo, indicating that for this sample the 98/95 98/95 measured δ Mo of –0.25‰ needs to be corrected upwards to δ Mocorrected = –0.20‰. Page 166 — Chapter VI samples and one angrite exhibit similar δ98/95Mo values ranging between -0.11‰ and -0.05‰. The other angrite investigated for this study as well as the shergottite exihibit more positive δ98/95Mo values between 0.16‰ and 0.49‰. Two of the three eucrites have similarly heavy compositions of about 0.16‰ and 0.26‰, while the third one (Stannern) has a very light δ98/95Mo of -0.42 ± 0.28‰. Among the silicate samples analyzed in the present study, the terrestrial dolerite DCN-1 shows the lowest δ98/95Mo of -0.18‰, indistinguishable from the value of chondrites and iron meteorites. All other silicate samples have δ98/95Mo values more positive than the chondrites and iron meteorites.

VI–4. DISCUSSION

The new Mo isotope data presented here provide clear evidence for mass dependent Mo isotope fractionation in meteorites. Although the data set is still too small, such that a comprehensive assessment of the different processes causing Mo isotope fractionation cannot yet be made, the silicate samples investigated so far exhibit a heavy Mo isotope composition compared to chondrites and iron meteorites. This is consistent with results from liquid-metal liquid-silicate exper- 98/95 iments, which reveal significant Mo isotope fractionation of Δ Mometal-silicate = -0.19 ± 0.03‰ at 1400 °C and 1 GPa (chapter 5). Thus at equilibrium conditions 98/95 the silicates are isotopically heavier than the metal, described as Δ MoMetal-Silicate = -4.80(±0.55) × 105 / T2, indicating that the Mo isotope fractionation observed in the meteorites may at least in part be a consequence of core formation. To assess the significance of the Mo isotope data with regard to core formation requires knowledge of the Mo isotope composition of the bulk, undifferentiated planetary body, and of the Mo isotope composition representative of the entire mantle. In the following sections these two constraints will be derived from the Mo isotope data, and the significance of these data for constraining the conditions of core formation will be discussed by comparing the observed Mo isotope fractionation between metal and silicates to that determined in metal-silicate separation experiments. VI–4.1. Stable Mo isotope composition of the inner solar system and bulk planetary bodies The stable Mo isotope composition of bulk, undifferentiated planetary bodies most likely is identical to that of chondrites. All the chondrites analyzed in the present study (excluding Murchinson) have a common δ98/95Mo, regardless of oxidation state and petrologic type. The δ98/95Mo of the chondrites is also indistinguishable from that of the iron meteorites. This is not surprising, because as a siderophile element most of the Mo will reside in the core after planetary differentiation. For instance, a body with a core of only 5% of its total mass and with ~6 ppm Mo in the metallic core and 50 ppb Mo in the silicate mantle will have Chapter VI — Page 167 over 85% of its bulk Mo in the core, causing the metal to be only 0.03‰ lighter than the bulk. Consequently, the stable Mo isotope composition of the metal core can be considered indistinguishable from that of the bulk, undifferentiated body. The common δ98/95Mo of -0.17 ± 0.02‰ of chondrites and iron meteorites, therefore, provides strong evidence that this value represents the stable Mo isotope composition of bulk, undifferentiated planetary bodies. The homogeneous stable Mo isotope composition of different chondrite groups contrasts with mass dependent isotope variations reported among chondrites for volatile [e.g., Zn (Luck et al., 2005)], moderately volatile [e.g., Cu, Cr (Bishop et al., 2012; Moynier et al., 2011)], and refractory [Ca, (Simon and DePaolo, 2010), Sr (Moynier et al., 2010)] elements. These isotopic variations have been ascribed to fractionation by condensation/evaporation processes in the nebula and subsequent incomplete mixing/homogenization of the fractionated nebular reservoirs and their products, particularly CAI (Luck et al., 2005; Moynier et al., 2011; Simon and DePaolo, 2010). The absence of stable Mo isotopic variations among different chondrite classes thus indicates that either the Mo isotopes have not been fractionated by these processes, or that the Mo concentrations in their products are too small to significantly contribute to the Mo isotope signature of the bulk meteorite. Which of these scenarios is responsible for the homogeneous stable Mo isotope composition among chondrites will be investigated in a future study by analyzing the isotopic signatures of different chondrite components. For the following discussion only the bulk planetary stable Mo isotope composition is relevant, which we assume to be the same as that of chondrites and iron meteorites (δ98/95Mo = -0.17 ± 0.02‰) for all material accreted in the inner solar system, including the precursor material of the Earth, Moon and Mars. VI–4.2. Comparison of experimental and observed Mo stable isotope fractionation during metal segregation All the silicate samples analyzed for the present study (except the terrestrial dolerite DCN-1) exhibit Mo isotope compositions heavier than those of chondrites. This heavier Mo isotope composition may reflect stable isotope fractionation during core formation, but may also be due to other processes such as magmatism or low-temperature alteration during terrestrial weathering. Deter- mining the δ98/95Mo value representative of the bulk silicate part of a planetary body, therefore, requires distinguishing Mo stable isotope fractionation caused by core formation from those related to post-core formation processes. The δ98/95Mo representative for the bulk silicate portion of a differentiated planetary body in conjunction with the experimental determination of the equilibrium fractionation between liquid metal and liquid silicate (see chapter 5) can be used to constrain the temperature of metal-silicate equilibration. This approach assumes complete metal-silicate equilibrium and quantitative metal segregation during core formation. The latter is required because if significant Page 168 — Chapter VI amounts of metal remain in the mantle after core formation, the Mo isotopic composition of the mantle will be dominated by that of the metal, which contains most of the Mo, and hence will be chondritic. At least for the terrestrial planets, efficient metal segregation almost certainly occurred. This is because accretion of the terrestrial planets is thought to have involved several giant impacts, which released sufficient energy to cause global melting, facilitating efficient metal- silicate separation and metal segregation in a magma ocean (Rubie et al., 2007; Stevenson, 2008). The temperature of metal-silicate equilibration can be calculated from the 5 98/95 Mo isotope data using T = √(4.80±0.55×10 / Δ Mobulk-silicate; see chapter V), 98/95 98/95 98/95 where Δ Mobulk-silicate = δ Mochondrites – δ Mosilicate. This approach assumes no pressure dependence, which is expected from theoretical considerations (Bigeleisen and Mayer, 1947; Schauble, 2004). VI–4.2.1. Earth and Moon The terrestrial samples have δ98/95Mo values ranging from the chondritic composition of about –0.18‰ (Dolerite DNC-1) to values that are ~0.1‰ heavier than the composition of chondrites and iron meteorites. The average of the terrestrial samples (excluding DNC-1) is δ98/95Mo = –0.07 ± 0.06‰ (2SD), which is resolvably heavier than chondrites. From the available data it is difficult to assess as to whether this difference is the sole result of core formation or reflects isotope fractionation during later processes. Large mass dependent Mo isotope fractionation in terrestrial samples have been reported primarily in low-temperature environments (Anbar, 2004), but the extent of Mo isotope fractionation during high-temperature magmatic processes is not yet known. The data presented here demonstrate that there are mass dependent Mo isotope variations of at least up to 0.1‰ among terrestrial igneous rocks, but whether this is the result of igneous processes or rather reflects isotope fractionation during low-temperature alteration of the basalts is unclear. Precisely defining δ98/95Mo value characteristic of the bulk silicate Earth will require a better understanding of any Mo stable isotope fractionation during igneous processes as well as a more comprehensive set of data covering the major silicate reservoirs of the Earth. For the purpose of the present study, we will assume that the δ98/95Mo value of the bulk silicate Earth is given by the average of the three basalts BHVO-2, BIR-1 and W-2a. 98/95 98/95 Their δ Mo translates to a Δ Mobulk-silicate of -0.11 ± 0.06‰, which, using the experimentally derived relationship between temperature and Mo stable isotope fractionation, corresponds to a temperature of metal-silicate equilibration of 1987 ± 792 °C (Table VI–2). If, alternatively, the δ98/95Mo value of the bulk silicate Earth is similar to that of the dolerite samples DNC-1, then there is no resolvable difference relative to chondrites, which would correspond to a metal-silicate equilibration temperature of more than ~2500 °C. The two lunar samples analyzed for the present study have indistinguishable Chapter VI — Page 169

Mo isotope compositions averaging at a δ98/95Mo of –0.09 ± 0.03‰ (2SD). Given that these two samples derive from different lunar geochemical reservoirs and were formed by different magmatic processes on the Moon, their indistinguishable δ98/95Mo suggests that Mo isotope fractionation during igneous processes on the Moon were minor to absent. The δ98/95Mo of the two lunar samples can thus be assumed to represent that of the bulk silicate Moon, corresponding to a 98/95 Δ Mobulk-silicate of –0.09 ± 0.03‰ relative to chondrites. This would translate to a metal-silicate equilibration temperature of 2274 ± 603 °C (Table VI–2). However, the interpretation of this temperature estimate with regard to lunar core formation is complicated. It could in principle reflect the temperature of core formation in the Moon, but the Moon did not form out of material with chondritic abundances of siderophile elements. The giant impact hypothesis of lunar origin rather predicts formation of the Moon predominantly out of silicate mantle material either from the impactor or from the proto-Earth. The precursor material of the Moon, therefore, must already have already undergone core formation and, depending on the temperature of the metal-silicate separation, might have had a fractionated Mo isotope composition. Thus, one endmember model to interpret the lunar Mo isotope data is that they reflect core formation in the precursor bodies of the Moon and not in the Moon itself. However, numerical simulations of the Moon-forming impact suggest that some metal of the impactor remixed with the mantle material and ultimately was included into the Moon (Canup, 2004). This material most likely had chondritic δ98/95Mo, was strongly enriched in Mo and thus may have dominated the Mo isotope composition of the proto-Moon. Therefore, the alternative endmember model to interpret the lunar Mo isotope data is that they solely reflect core formation in the Moon. An important observation from the Mo isotope data for lunar and terrestrial samples is that all these samples (except dolerite sample DNC-1) have indistinguishable δ98/95Mo values averaging at –0.08 ± 0.05‰ (2SD). Although this set of data is too limited to firmly establish the Mo isotope composition characteristic for the bulk silicate Table VI–2. Mo isotope fractionation and inferred Earth and Moon, a common core formation temperatures. δ98/95Mo of the lunar and ter- 98/95 a b restrial samples analyzed for Planetary body Δ MoBulk-Silicate Tmetal-silicate eq. the present study suggests a [‰] [°C] 98/95 common Δ Mobulk-silicate of Earth -0.094 ± 0.065 1987 ± 792 –0.09 ± 0.05‰ for the lunar Moon -0.074 ± 0.034 2274 ± 603 and terrestrial mantles. This corresponds to a metal-sili- Angrite PB -0.112 ± 0.048 1797 ± 459 cate equilibration temperature Eucrite PB -0.364 ± 0.061 875 ± 117 of 2131 ± 750 °C for both the a Isotope fractionation between bulk (chondrites and iron Earth and Moon. However, meteorites mean) and planetary silicates. b Liquid metal liquid silicate equilibrium temperature calculated using core formation in the Earth 5 98/95 T = √(4.80[±0.55]×10 / Δ Mobulk-silicate). Page 170 — Chapter VI and Moon is not expected to have involved metal-silicate equilibration at the same temperature, because core formation in the Earth is thought to have taken place at the base of a deep magma ocean, i.e., much deeper and hence at higher temperatures than in the Moon. The Mo isotope data may, therefore, be inter- preted to indicate that metal-silicate equilibration during core formation in the Earth occurred at a shallower depth and not at the base of a deep magma ocean. It is conceivable that small metal droplets descending through the magma ocean equilibrated at some intermediate depths and not at the base of the magma ocean. Alternatively, the metal cores of newly accreted impactors did not fully emulsify in the terrestrial magma ocean but merged with the Earth’s core without re-equilibration in the magma ocean. The Mo isotope signature of the Earth’s mantle would in this case reflect the conditions of core formation in the impactors and not of those in the Earth. VI–4.2.2. Angrites and eucrites The Mo isotope data for angrites and eucrites are more variable than those of the lunar and terrestrial samples, suggesting that in addition to core formation other processes were important in defining the Mo isotope composition of these samples. Among the angrites and eucrites investigated for this study, D’Orbigny has the lowest δ98/95Mo of –0.05 ± 0.04‰, slightly heavier than but not resolvable from the values obtained for lunar and terrestrial samples. Two of the three eucrites analyzed for this study have a heavier Mo and barely overlapping isotope composition with δ98/95Mo values of +0.16 ± 0.08‰ for Millbillillie and +0.26 ± 0.04‰ for the cumulate eucrite Dhofar 007. This may suggest stable Mo isotope fractionation during magmatism or later metamorphism in the eucrite parent body, but may equally well reflect the effects of terrestrial weathering. Given that low-temperature processes such as weathering have been shown to fractionate Mo isotopes (Goldberg et al., 2009), the eucrite with the heavier Mo isotope composition – Dhofar 007, a desert find – may have been affected by weathering in the desert. That weathering in the desert can affect the stable Mo isotope composition of meteorites is most obviously shown by two other desert meteorites analyzed in the present study, which have δ98/95Mo values of about +0.49‰ (angrite NWA 4801) and about +0.43‰ (shergottite DaG 476), the heaviest Mo isotope compositions observed here for meteorites. Such large Mo isotope fractionations have so far only been observed for low-temperature environments and thus at least for the angrite NWA 4801 almost certainly reflect Mo isotope fractionation during terrestrial weathering. It is noteworthy that this sample also shows clearly visible signs of terrestrial alteration that could not be removed during sample preparation. Therefore, until more Mo isotope data for eucrites and angrites will be available, we will restrict the discussion of the Mo isotope fractionation as a result of core formation to the two samples with lowest δ98/95Mo, i.e., D’Orbigny and Millbillillie. Chapter VI — Page 171

98/95 D’Orbigny exhibits a Δ Mobulk-silicate of –0.12 ± 0.04‰, which translates to a metal-silicate equilibration temperature of 1797 ± 459 °C. This temperature is well above the peridotite liquidus for even the largest asteroids (Walter, 2003) and is thus consistent with core formation by efficient metal-silicate separation and metal segregation during large-scale melting. This in turn is consistent with the early timing of core formation as deduced from Hf-W data, indicating that at the time of parent body accretion 26Al was sufficiently abundant to cause global melting (Kleine et al., 2012), and with the presence of a core-dynamo for at least ~10 My (Weiss et al., 2008). The Mo isotopic composition of the eucrite Millbillillie is heavier than that of D’Orbigny and corresponds to a metal-silicate equilibration temperature of 875 ± 117 °C. This temperature estimate overlaps with the Fe-FeS eutectic melting temperature of ~1000 °C, the minimum temperature required for melting of metal and, hence, for core formation. Taken at face value, the inferred temperature would indicate only limited melting on the eucrite parent body during metal segregation. It has been shown that at low pressures (i.e., in asteroids) and fO2 conditions >IW-2 even small amounts of melt are sufficient to form intercon- nected networks and initiate percolative core formation (Terasaki et al., 2008). However, quantitative metal extraction in such a scenario seems unlikely and given the geochemical evidence for efficient metal segregation in the eucrite parent body (Holzheid and Palme, 2007), core formation should have occurred at higher temperatures, at which a significant fraction of the silicates was also molten. The rather low temperature of 875 ± 117 °C derived from the Mo isotope data for Millbillillie, therefore, may not only reflect metal-silicate separation but also other processes than core formation that fractionated the Mo isotopes. Since this meteorite is an observed fall, terrestrial weathering cannot account for its heavy Mo isotope composition. This suggests that processes during magmatism and thermal metamorphism in the eucrite parent body might be responsible for the heavy Mo isotope composition of Millbillillie, possibly related to the formation of native Fe metal during eucrite magmatism and/or later thermal metamorphism. A more complicated origin of the Mo isotope composition of eucrites is also suggested by the composition of the third eucrite (Stannern) of -0.42 ± 0.28‰, which is the lightest meteorite composition observed so far. Clearly, more data are needed to assess the different processes that can have fractionated Mo isotopes in the eucrites. VI–4.2.3. Mars To investigate any potential Mo stable isotope fractionation during core formation in Mars, the shergottite DaG 476 was analyzed. This sample together with the angrite NWA 4801 exhibits the highest δ98/95Mo among the meteorites 98/95 investigated for this study. Its Δ Mobulk-silicate of –0.61 ± 0.06‰ would translate to an apparent metal-silicate equilibration temperature of only 626 ± 73 °C. Such Page 172 — Chapter VI a low temperature cannot be a signature of core formation. The heavy Mo isotope composition of DaG 476 must rather reflect Mo isotope fractionation during other processes, most likely during terrestrial weathering (as has also been argued for the angrite NWA 4801). Once more Mo isotope data for other martian samples are available, it will be possible to constrain the origin of the heavy Mo isotope composition of DaG 476 more tightly.

VI–5. CONCLUSIONS

We have demonstrated that Mo isotopes are fractionated by high-temperature processes and can provide important new constraints on planetary differentiation and core formation. Our new data show that chondrites and iron meteorites have indistinguishable Mo isotope compositions and define the stable Mo isotope composition of bulk, undifferentiated planetary bodies in the inner solar system. By contrast, all the investigated silicate samples (except the terrestrial dolerite DNC-1 and eucrite Stannern) are isotopically heavier, at least in part as a result of Mo isotope fractionation during metal-silicate equilibration. Inferred metal- silicate equilibration temperatures range from ~1800 °C for the angrite parent body to ~2100 °C for the Earth and Moon, consistent with the high temperature required for core formation. The metal-silicate equilibration temperature calculated for the Earth suggests that metal-silicate equilibration did not occur at the base of a deep magma ocean but rather took place at an intermediate depth or, alternatively, reflects the signature of core formation in Earth’s precursor bodies. This implies that core formation occurred at lower pressures than inferred in some studies using the partitioning of siderophile elements under high pressures. The variable Mo isotope compositions of the silicate samples investigated for the present study demonstrate that processes other than core formation might fractionate Mo isotopes in meteorites. This complicates the interpretation of the Mo isotope data in terms of metal-silicate equilibration temperatures, and requires a more detailed understanding of the different processes leading to mass dependent Mo isotope fractionations. Future studies will have to identify these processes and better constrain the Mo isotope composition of the bulk silicate portions of asteroids and terrestrial planets. The present study demonstrates that, once this is established, Mo isotopes can provide important new constraints on planetary core formation. CHAPTER VII.

CONCLUSIONS Page 174 — Chapter VII

In this thesis you may have read that detectable stable isotope fractionation occurs between metal and silicate liquids, but not for all elements. Isotopic fractionations vary strongly, but the variations are not a straightforward function of the element masses. Below, I will briefly summarise the main results my experiments and their implications, before I finish with some more general conclusions.

VII–1. SUMMARY OF THE MAIN CONCLUSIONS

Iron isotopes do not fractionate between liquid metal and liquid silicate at 56 temperatures above 1250 °C. The average fractionation factor Δ FeMetal-Silicate was 0.01 ± 0.04‰ (95% confidence interval), which is independent of the presence or absence of carbon or sulphur. The direct implication of this observation is that any variation in Fe isotope compositions among meteorites cannot have been caused by equilibrium metal-silicate melt fractionation of Fe isotopes during magmatic core-mantle segregation. The observed difference in Fe isotope compositions of 0.07 ± 0.02‰ between ordinary/carbonaceous chondrites and iron meteorites, for instance, can only be explained by kinetic fractionation during metal-silicate segregation or by processes unrelated to core formation. Earth is the only known body in which Fe isotopes might have fractionated significantly between equilibrated metal and silicate during core segregation. This could be due to the presence of Fe3+ in perovskite, a high pressure phase occurring in Earth, but not in Mars or smaller bodies, as these are not large enough to stabilise that phase. Silicon isotopes fractionate between metal and silicate melts by -1.48 ± 0.08‰ (on δ30Si) at 1450 °C and by -1.11 ± 0.14‰ at 1750 °C, the silicate liquid having a heavier composition than the metal liquid. Carbon does not affect Si isotope fractionation. The temperature dependence of Si isotope fractionation 30 × 6 2 can be described as Δ SiMetal-Silicate = -4.47(±0.31) 10 /T . Together with a bulk silicate Earth δ30Si value of -0.29 ± 0.01‰, this temperature dependence can be used to calculate δ30Si values for Earth’s core. For metal-silicate equilibration temperatures of 2500-3500 K, the Si isotope composition of Earth’s core is -1.01 ± 0.05‰ to -0.66 ± 0.03‰. Mass balance calculations indicate that for those core compositions, the Bulk Earth has a δ30Si of -0.38 ± 0.02‰ or -0.33 ± 0.01‰, respectively, if Earth’s core is assumed to contain 6 wt% Si, or -0.40 ± 0.02‰ or -0.35 ± 0.02‰, respectively, for 8 wt% Si in the Earth’s core. These calculated Bulk Earth Si isotope compositions can be compared to the reported Si isotope compositions of chondrites, but the latter remain disputed (see chapter IV, section IV.5.2.). Enstatite chondrites have reported δ30Si values of -0.62 ± 0.05‰ and are therefore unlikely to represent the Bulk Earth Si isotope composition. A conclusion about whether Si isotopes concur with high temperature (≥2500 K) core formation in a chondritic reference model, however, requires better determination of the Si isotope compositions of chondrites. Chapter VII — Page 175

Molybdenum isotope compositions are heavier by 0.193 ± 0.030‰ (on δ98/95Mo) in silicate than in metal melt equilibrated at 1400 °C. Combining this with a difference of 0.120 ±�� 0.020‰�� at 1600 °C�� yields a temperature depend- 98/95 × 5 2 ence of Δ MoMetal-Silicate = -4.80(±0.55) 10 /T . This fractionation is valid for silicates with dominantly Mo4+, which is the case for silicate melts in (partially) molten bodies during core formation. The fractionation is furthermore not affected by the presence or absence of carbon or tin, or by variations in silicate melt compositions from basanitic to basaltic. This establishes Mo isotopes as a tool to investigate temperature conditions during metal-silicate equilibration and segregation in a variety of planetary bodies. Analyses on meteorites indicate that if single stage models of core formation are employed, the Earth/Moon system formed its core at ~2100 °C. Core formation in the angrite parent body seems to have taken place at ~1800 °C. Particularly the temperature for core formation on Earth is at the lower end of temperatures deduced from models based on elemental distribution. The calculated temperature of ~2100 °C rather suggests that Mo isotopes record an average temperature of metal-silicate equilibration during multiple accretion events, or a mixture of conditions of core formation in Earth as well as in planetesimals and other precursor bodies prior to their accretion to Earth. The isotopic analyses of natural rocks, however, also indicate that Mo isotopes may be fractionated by other processes than metal-silicate differentiation. A temperature of ~875 °C for core formation on the eucrite parent body, for instance, is too low for melting of the metal or silicate fractions. The potential of fractionation during subsolidus processes therefore needs to be further investigated.

VII–2. GENERAL CONCLUSIONS AND IMPLICATIONS VII–2.1. Stable isotope fractionation As mentioned above, I have varied the compositions of the metal as well as silicate melts to determine whether they affect stable isotope fractionation between these phases. Compositional variations did not influence the fractionation of any of the three isotopic systems I investigated (Table VII–1). This is in marked contrast to elemental distribution, which can strongly vary as a function of silicate or metal melt composition. In my opinion, this contrast is readily understandable from a theoretical perspective: for isotopic fractionation only the bond stiffness is important, which is similar in any compositional system as long as the bonding partner(s) remain(s) identical and as long as compositional changes do not lead to changes in the structural environment, such as valence and coordination state. Elemental distribution, on the other hand, depends on the available sites in a crystal or melt structure or on the affinity of one element for another. It is thus dependent on composition. Conceptually, the above mentioned difference is probably best explained with Page 176 — Chapter VII

Table VII–1. Summary of compositional effects on Fe, Si and Mo isotope fractionation between metal and silicate liquids. Andesitic basalt Basanitic C S Sn to basalt to basaltic Fe No effect No effect - No effect - Si No effect - No effect - No effect Mo No effect - No effect - No effect the example of Mo. This is an element with high valence (4+), which does not allow for Mo-O-Si bonds and hence needs a relatively large amount of non- bridging oxygen ions in silicate melts. The concentration of Mo in a silicate melt (and therefore its distribution between metal and silicate melts) is critically dependent on the availability of non-bridging oxygen ions and, hence, a strong function of melt polymerisation. Whether a silicate melt is highly polymerised or highly depolymerised, however, does not change the fact that Mo will always be bonded to those oxygen ions. Consequently, the bond stiffness that determines its vibrational energy and thereby its isotopic fractionation remains the same. As such, isotopic fractionation is unaffected by compositional changes that affect elemental distribution. A second general observation can be made from the results of my study: equilibrium stable isotope fractionation is not a simple function of the molar mass of elements. The isotopes of Mo fractionated between metal and silicate liquids, while the isotopes of the lighter element Fe did not fractionate. The parameter ΔF in the following equation (also see Chapter I, equation I.4) is fairly unconstrained: Dm DF ln()α ≈ AB− AB− mT22 It is qualitatively known that ΔF�� is mainly affected by valence states, coordination states and electronic spin configurations. If we isolate ΔF in the three isotopic systems I have studied and combine them with the available information about bonding environment (Table VII–2), we can speculate about the relative influences of valence and coordination state on ΔF. It appears that changes in coordination state affect stable isotope fractionation by less than 30% (cf. Si and Mo in Table VII–2), while a factor two increase in valence state may cause an order of magnitude increase in the isotopic fractionation (compare Fe with Si and Mo in Table VII–2). Nevertheless, this consideration is based on a limited number of data and ignores the changes in bond stiffness in the metal phases. Although the latter may vary less than bond stiffness in silicates, this issue certainly needs a more thorough investigation. If the above observation is accurate, the implications are quite important: at temperatures above 1250 °C isotopes will not fractionate between two phases if the isotopes are about as heavy as Fe (unless the mass differences between the isotopes are much larger than for Fe) and if the Chapter VII — Page 177 element has a difference in valence of two units or less between the two phases. Consequently, elements like Cu, Zn, Ni and probably Cr would not significantly fractionate their isotopes between metal and silicate. Furthermore, significant equilibrium fractionation is unlikely to occur during partial melting of silicate at these temperatures: valence state differences will be even smaller between a silicate crystal and a silicate melt, even if e.g. Fe2+ is strongly fractionated into the solid (olivine, orthopyroxene, etc) with respect to Fe3+. As mentioned, though, this topic certainly needs to be investigated in more detail.

Table VII–2. Comparison of values for the parameter ΔF between three isotopic systems with variable bonding environments. Valence state Coordination state References for structural ΔF Metal-Silicate (in silicate) (in silicate) environment Fe <1.40×105 2+ dominantly V-fold Jackson et al. (2005) Si -1.88(±0.13)×106 4+ IV-fold Warren and Biscoe (1938) This study and Farges et al. Mo -1.49(±0.17)×106 4+ VI-fold (?) (2006)

VII–2.2. Conditions of core formation The results of my work indicate that stable isotope fractionation between metal and silicate liquids can occur at detectable levels during core formation. Interestingly, our initial data set of Mo isotopes seems to indicate a discrepancy between the temperatures of core formation in Earth as determined by isotopic versus elemental distributions. While most studies of elemental distributions indicate that core formation in Earth occurred at temperatures above 2500 °C, Mo isotope analyses rather imply temperatures below 2500 ��°C.�� This discrepancy could be related to the recent conclusion that core formation models based on elemental distributions are not fully sensitive to the degree of equilibration (Rudge et al., 2010). The Mo isotope data may thus record conditions of core formation integrated over events with variable temperatures (and depth) of metal-silicate equilibration. They may even record conditions of core formation in precursor bodies that did not re-equilibrate their metallic core with silicate in Earth upon accretion. Such a scenario would be a step forward in resolving the dispute between end-members of equilibrium and disequilibrium models of core formation in the sense that it would show that a combined scenario occurred. Fe and Mo isotope analyses on meteorites also indicate that other (probably subsolidus) processes fractionate stable isotopes, thereby complicating the interpretation of these data. These complications will have to be deciphered in future work to establish stable isotope fractionation as a powerful tool to investigate conditions of core formation and to further constrain core formation models.

Bibliography — Page 179

BIBLIOGPRAHY

Albarede F., Telouk P., Blichert-Toft J., Boyet M., Agranier A., and Nelson B., 2004. Precise and accurate isotopic measurements using multiple-collector ICPMS. Geochim. Cosmochim. Acta 68, 2725-2744. Allègre C. J., Poirier J. P., Humler E., and Hofmann A. W., 1995. The Chemical Composition of the Earth. Earth Planet. Sci. Lett. 134, 515-526. Amelin Y., Krot A. N., Hutcheon I. D., and Ulyanov A. A., 2002. Lead isotopic ages of chondrules and calcium-aluminum-rich inclusions. Science 297, 1678-1683. Anbar A. D., 2004. Molybdenum stable isotopes: Observations, interpretations and directions. Rev Mineral Geochem 55, 429-454. Anderson O. L. and Isaak D. G., 2002. Another look at the core density deficit of Earth’s outer core. Phys. Earth Planet. In. 131, 19-27. Ardia P., Withers A. C., Hirschmann M. M., and Hervig R. L., 2011. Methane solubility under reduced conditions in a haplobasaltic liquid, applicable to degassing of magma oceanGoldschmidt. Mineralogical Magazine. Arlandini C., Kappeler F., Wisshak K., Gallino R., Lugaro M., Busso M., and Straniero O., 1999. Neutron capture in low-mass asymptotic giant branch stars: Cross sections and abundance signatures. Astrophys. J. 525, 886-900. Armytage R. M. G., Georg R. B., Savage P. S., Williams H. M., and Halliday A. N., 2011. Silicon isotopes in meteorites and planetary core formation. Geochim. Cosmochim. Acta 75, 3662-3676. Badro J., Fiquet G., Guyot F., Gregoryanz E., Occelli F., Antonangeli D., and d’Astuto M., 2007. Effect of light elements on the sound velocities in solid iron: Implications for the composition of Earth’s core. Earth Planet. Sci. Lett. 254, 233-238. Beard B. L., Johnson C. M., Skulan J. L., Nealson K. H., Cox L., and Sun H., 2003. Application of Fe isotopes to tracing the geochemical and biological cycling of Fe. Chem. Geol. 195, 87-117. Bigeleisen J. and Mayer M. G., 1947. Calculation of equilibrium constants for isotopic exchange reactions. J. Chemical Physics 15, 261-267. Bigeleisen J., 1965. Chemistry of Isotopes. Science 147, 463-471. Birch F., 1952. Elasticity and constitution of the Earth’s interior. J. Geophys. Res. 57, 227-286. Bishop M. C., Moynier F., Weinstein C., Fraboulet J. G., Wang K., and Foriel J., 2012. The Cu isotopic composition of iron meteorites. Meteorit. Planet. Sci 47, 268-276. Bose K. and Ganguly J., 1995. Quartz-Coesite transition revisited: Reversed experimental determination at 500-1200 °C and retrieved thermochemical properties. Am. Miner. 80, 231-238. Page 180 — Bibliography

Boyd F. R. and England J. L., 1960. Apparatus for phase-equilibrium measurements at pressures up to 50 kilobars and temperatures up to 1750 °C. J. Geophys. Res. 65, 741-748. Buchwald V. F., 1977. The mineralogy of iron meteorites. Philosophical Trans- actions of the Royal Society of London Series A - Mathematical and Physical Sciences 286, 453-491. Burkhardt C., Kleine T., Oberli F., Pack A., Bourdon B., and Wieler R., 2011. Molybdenum isotope anomalies in meteorites: Constraints on solar nebula evolution and origin of the Earth. Earth Planet. Sci. Lett. 312, 390-400. Burkhardt C., Hin R. C., and Kleine T., 2012a. Molybdenum stable isotope constraints on planetary core formationEuropean Mineralogical Conference, Frankfurt. Burkhardt C., Kleine T., Dauphas N., and Wieler R., 2012b. Origin of isotopic heterogeneity in the solar nebula by thermal processing and mixing of nebular dust. Earth Planet. Sci. Lett. Canup R. M., 2004. Simulations of a late lunar-forming impact. Icarus 168, 433-456. Chakrabarti R. and Jacobsen S. B., 2010. Silicon isotopes in the inner Solar System: Implications for core formation, solar nebular processes and partial melting. Geochim. Cosmochim. Acta 74, 6921-6933. Chambers J. E., 2004. Planetary accretion in the inner Solar System. Earth Planet. Sci. Lett. 223, 241-252. Choi B. G., Ouyang X. W., and Wasson J. T., 1995. Classification and Origin of Iab and IIIcd Iron-Meteorites. Geochim. Cosmochim. Acta 59, 593-612. Corgne A., Keshav S., Wood B. J., McDonough W. F., and Fei Y. W., 2008. Metal-silicate partitioning and constraints on core composition and oxygen fugacity during Earth accretion. Geochim. Cosmochim. Acta 72, 574-589. Craddock P. R. and Dauphas N., 2011. Iron isotopic compositions of geological reference materials and chondrites. Geostand Geoanal Res 35, 101-123. Dauphas N., Marty B., and Reisberg L., 2002. Molybdenum evidence for inherited planetary scale isotope heterogeneity of the protosolar nebula. Astrophys. J. 565, 640-644. Dauphas N., Teng F. Z., and Arndt N. T., 2010. Magnesium and iron isotopes in 2.7 Ga Alexo komatiites: Mantle signatures, no evidence for Soret diffusion, and identification of diffusive transport in zoned olivine. Geochim. Cosmochim. Acta 74, 3274-3291. De Pater I. and Lissauer J., 2001. Planetary sciences. Cambridge University Press, Cambridge. Dean D. C. and Goldstein J. I., 1986. Determination of the interdiffusion coefficients in the Fe-Ni and Fe-Ni-P systems below 900 °C.Metall. Trans. A 17, 1131-1138. Dobson D. P., 2002. Self-diffusion in liquid Fe at high pressure. Phys. Earth Bibliography — Page 181

Planet. In. 130, 271-284. Ellison A. J. and Hess P. C., 1986. Solution behavior of +4 cations in high silica melts - Petrologic and geochemical implications. Contrib. Mineral. Petrol. 94, 343-351. Farges F., Siewert R., Brown G. E., Guesdon A., and Morin G., 2006. Structural environments around molybdenum in silicate glasses and melts. I. Influence of composition and oxygen fugacity on the local structure of molybdenum. Can. Min. 44, 731-753. Fitoussi C., Bourdon B., Kleine T., Oberli F., and Reynolds B. C., 2009. Si isotope systematics of meteorites and terrestrial peridotites: implications for Mg/Si fractionation in the solar nebula and for Si in the Earth’s core. Earth Planet. Sci. Lett. 287, 77-85. Fitoussi C. and Bourdon B., 2012. Silicon Isotope Evidence Against an Enstatite Chondrite Earth. Science 335, 1477-1480. Georg R. B., Reynolds B. C., Frank M., and Halliday A. N., 2006. New sample preparation techniques for the determination of Si isotopic compositions using MC-ICPMS. Chem. Geol. 235, 95-104. Georg R. B., Halliday A. N., Schauble E. A., and Reynolds B. C., 2007. Silicon in the Earth’s core. Nature 447, 1102-1106. Ghosh A., Weidenschilling S. J., McSween Jr. H. Y., and Rubin A., 2006. Aster- oidal heating and thermal stratification of the asteroid belt. In: Lauretta D. S. and McSween H. Y. J. Eds.), Meteorites and the Early Solar System II. University of Arizona Press, Tuscon. Goldberg T., Archer C., Vance D., and Poulton S. W., 2009. Mo isotope fractionation during adsorption to Fe (oxyhydr)oxides. Geochim. Cosmochim. Acta 73, 6502-6516. Greber N. D., Hofmann B. A., Voegelin A. R., Villa I. M., and Nagler T. F., 2011. Mo isotope composition in Mo-rich high- and low-T hydrothermal systems from the Swiss Alps. Geochim. Cosmochim. Acta 75, 6600-6609. Greber N. D., Siebert C., Nägler T. F., and Pettke T., 2012. d98/95Mo values and molybdenum concentration data for NIST SRM 610, 612 and 3134: Towards a common protocol for reporting Mo data. Geostand Geoanal Res. Grossman L., 1972. Condensation in Primitive Solar Nebula. Geochim. Cosmochim. Acta 36, 597-&. Guillong M., Horn I., and Gunther D., 2003. A comparison of 266 nm, 213 nm and 193 nm produced from a single solid state Nd : YAG laser for laser ablation ICP-MS. J. Anal. At. Spectrom. 18, 1224-1230. Guillot B. and Sator N., 2007. A computer simulation study of natural silicate melts. Part I: Low pressure properties. Geochim. Cosmochim. Acta 71, 1249-1265. Halliday A. N., 2004. Mixing, volatile loss and compositional change during impact-driven accretion of the Earth. Nature 427, 505-509. Page 182 — Bibliography

Hevey P. J. and Sanders I. S., 2006. A model for planetesimal meltdown by 26Al and its implications for meteorite parent bodies. Meteorit. Planet. Sci 41, 95-106. Hillgren V. J., Drake M. J., and Rubie D. C., 1996. High pressure and high temperature metal-silicate partitioning of siderophile elements: The importance of silicate liquid composition. Geochim. Cosmochim. Acta 60, 2257-2263. Hin R. C., Schmidt M. W., and Bourdon B., 2012. Experimental evidence for the absence of iron isotope fractionation between metal and silicate liquids at 1 GPa and 1250–1300 °C and its cosmochemical consequences. Geochim. Cosmochim. Acta 93, 164-181. Holzheid A., Borisov A., and Palme H., 1994. The Effect of Oxygen Fugacity and Temperature on Solubilities of Nickel, Cobalt, and Molybdenum in Silicate Melts. Geochim. Cosmochim. Acta 58, 1975-1981. Holzheid A., Palme H., and Chakraborty S., 1997. The activities of NiO, CoO and FeO in silicate melts. Chem. Geol. 139, 21-38. Holzheid A. and Palme H., 2007. The formation of eucrites: Constraints from metal-silicate partition coefficients.Meteorit. Planet. Sci 42, 1817-1829. Hughes H. J., Delvigne C., Korntheuer M., de Jong J., Andre L., and Cardinal D., 2011. Controlling the mass bias introduced by anionic and organic matrices in silicon isotopic measurements by MC-ICP-MS. J. Anal. At. Spectrom. 26, 1892-1896. Jana D. and Walker D., 1997. The impact of carbon on element distribution during core formation. Geochim. Cosmochim. Acta 61, 2759-2763. Karato S. and Murthy V. R., 1997. Core formation and chemical equilibrium in the Earth - I. Physical considerations. Phys. Earth Planet. In. 100, 61-79. Kiczka M., Wiederhold J. G., Frommer J., Kraemer S. M., Bourdon B., and Kretzschmar R., 2010. Iron isotope fractionation during proton- and ligand- promoted dissolution of primary phyllosilicates. Geochim. Cosmochim. Acta 74, 3112-3128. Kilburn M. R. and Wood B. J., 1997. Metal-silicate partitioning and the incom- patibility of S and Si during core formation. Earth Planet. Sci. Lett. 152, 139-148. Kleine T., Mezger K., Munker C., Palme H., and Bischoff A., 2004. 182Hf-182W isotope systematics of chondrites, eucrites, and martian meteorites: Chronol- ogy of core formation and early mantle differentiation in Vesta and Mars. Geochim. Cosmochim. Acta 68, 2935-2946. Kleine T., Touboul M., Bourdon B., Nimmo F., Mezger K., Palme H., Jacobsen S. B., Yin Q. Z., and Halliday A. N., 2009. Hf-W chronology of the accretion and early evolution of asteroids and terrestrial planets. Geochim. Cosmochim. Acta 73, 5150-5188. Kleine T. and Rudge J. F., 2011. Chronometry of meteorites and the formation of the Earth and Moon. Elements 7, 41-46. Bibliography — Page 183

Kleine T., Hans U., Irving A. J., and Bourdon B., 2012. Chronology of the angrite parent body and implications for core formation in protoplanets. Geochim. Cosmochim. Acta 84, 186-203. Kocherahinski Y. A. and Kulik O. G., 1996. Equilibrium phase diagrams and manufacture of synthetic diamonds. Powder Metallurgy and Metal Ceramics 35, 470-483. Kroslakova I. and Gunther D., 2007. Elemental fractionation in laser ablation- inductively coupled plasma-mass spectrometry: evidence for mass load induced matrix effects in the ICP during ablation of a silicate glass. J. Anal. At. Spectrom. 22, 51-62. Krot A. N., Amelin Y., Bland P., Ciesla F. J., Connelly J., Davis A. M., Huss G. R., Hutcheon I. D., Makide K., Nagashima K., Nyquist L. E., Russell S. S., Scott E. R. D., Thrane K., Yurimoto H., and Yin Q. Z., 2009. Origin and chronology of chondritic components: A review. Geochim. Cosmochim. Acta 73, 4963-4997. Kushiro I., Yoder H. S., and Mysen B. O., 1976. Viscosities of Basalt and Andesite Melts at High-Pressures. J. Geophys. Res. 81, 6351-6356. Lacaze J. and Sundman B., 1991. An assessment of the Fe-C-Si system. Metall. Trans. A 22, 2211-2223. LaTourrette T., Wasserburg G. J., and Fahey A. J., 1996. Self diffusion of Mg, Ca, Ba, Nd, Yb, Ti, Zr, and U in haplobasaltic melt. Geochim. Cosmochim. Acta 60, 1329-1340. Lesher C. E., 1990. Decoupling of chemical and isotopic exchange during magma mixing. Nature 344, 235-237. Li J. and Agee C. B., 1996. Geochemistry of mantle-core differentiation at high pressure. Nature 381, 686-689. Lord O. T., Walter M. J., Dasgupta R., Walker D., and Clark S. M., 2009. Melting in the Fe-C system to 70 GPa. Earth Planet. Sci. Lett. 284, 157-167. Lu Q. and Masuda A., 1994. The isotopic composition and atomic weight of Molybdenum. Int J Mass Spectrom 130, 65-72. Luck J. M., Ben Othman D., and Albarede F., 2005. Zn and Cu isotopic variations in chondrites and iron meteorites: Early solar nebula reservoirs and parent-body processes. Geochim. Cosmochim. Acta 69, 5351-5363. Ma Z. T., 2001. Thermodynamic description for concentrated metallic solutions using interaction parameters. Metall. Mater. Trans. B 32, 87-103. Mann U., Frost D. J., and Rubie D. C., 2009. Evidence for high-pressure core- mantle differentiation from the metal-silicate partitioning of lithophile and weakly-siderophile elements. Geochim. Cosmochim. Acta 73, 7360-7386. Markowski A., Leya I., Quitte G., Ammon K., Halliday A. N., and Wieler R., 2006. Correlated helium-3 and tungsten isotopes in iron meteorites: Quan- titative cosmogenic corrections and planetesimal formation times. Earth Planet. Sci. Lett. 250, 104-115. Page 184 — Bibliography

McCammon C., 1997. Perovskite as a possible sink for ferric iron in the lower mantle. Nature 387, 694-696. McDonough W. F., 2003. Compositional model for the Earth’s core. In: Carlson R. W. (Ed.), The Mantle and Core. Elsevier-Pergamon, Oxford. Mikutta C., Wiederhold J. G., Cirpka O. A., Hofstetter T. B., Bourdon B., and Von Gunten U., 2009. Iron isotope fractionation and atom exchange during sorption of ferrous iron to mineral surfaces. Geochim. Cosmochim. Acta 73, 1795-1812. Miller J. N. and Miller J. C., 2010. Statistics and chemometrics for analytical chemistry. Prentice Hall, Harlow. Mizuike A., 1983. Enrichment techniques for inorganic trace analysis Springer, Berlin. Moynier F., Agranier A., Hezel D. C., and Bouvier A., 2010. Sr stable isotope composition of Earth, the Moon, Mars, Vesta and meteorites. Earth Planet. Sci. Lett. 300, 359-366. Moynier F., Yin Q. Z., and Schauble E., 2011. Isotopic evidence of Cr partitioning into Earth’s core. Science 331, 1417-1420. Neal C. R., 2001. Interior of the Moon: The presence of garnet in the primitive deep lunar mantle. J. Geophys Res.-Planets 106, 27865-27885. Needham A. W., Porcelli D., and Russell S. S., 2009. An Fe isotope study of ordinary chondrites. Geochim. Cosmochim. Acta 73, 7399-7413. Newsom H. E., 1985. Molybdenum in Eucrites - Evidence for a Metal Core in the Eucrite Parent Body. J. Geophys. Res. 90, C613-C617. Newsom H. E., 1986. Constraints on the origin of the Moon from the abundance of Molybdenum and other siderophile elements. Proceedings of the Confer- ence “Origin of the Moon”, 206-229. O’Neill H. S. C. and Eggins S. M., 2002. The effect of melt composition on trace element partitioning: an experimental investigation of the activity coefficients of FeO, NiO, CoO, MoO2 and MoO3 in silicate melts. Chem. Geol. 186, 151-181. Okamoto H., 1993. Phase diagrams of binary iron alloys. ASM International, Materials Park, Ohio. Palme H. and Jones A., 2003. Solar System abundances of the elements. In: Davis A. M. (Ed.), Meteorites, Comets, and Planets. Elsevier-Pergamon, Oxford. Palme H. and O’Neill H. S. C., 2003. Cosmochemical estimates of mantle composition. In: Carlson R. W. (Ed.), The Mantle and Core. Elsevier-Pergamon, Oxford. Patchett P. J. and Tatsumoto M., 1980. A Routine High-Precision Method for Lu-Hf Isotope Geochemistry and Chronology. Contrib. Mineral. Petrol. 75, 263-267. Poirier J. P., 1994. Light elements in the Earths outer core - a critical review. Phys. Earth Planet. In. 85, 319-337. Bibliography — Page 185

Poitrasson F., Halliday A. N., Lee D. C., Levasseur S., and Teutsch N., 2004. Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms. Earth Planet. Sci. Lett. 223, 253-266. Poitrasson F. and Freydier R., 2005. Heavy iron isotope composition of granites determined by high resolution MC-ICP-MS. Chem. Geol. 222, 132-147. Poitrasson F., Levasseur S., and Teutsch N., 2005. Significance of iron isotope mineral fractionation in pallasites and iron meteorites for the core-mantle differentiation of terrestrial planets. Earth Planet. Sci. Lett. 234, 151-164. Poitrasson F., Roskosz M., and Corgne A., 2009. No iron isotope fractionation between molten alloys and silicate melt to 2000 °C and 7.7 GPa: Experi- mental evidence and implications for planetary differentiation and accretion. Earth Planet. Sci. Lett. 278, 376-385. Polyakov V. B. and Mineev S. D., 2000. The use of Mössbauer spectroscopy in stable isotope geochemistry. Geochim. Cosmochim. Acta 64, 849-865. Polyakov V. B., Clayton R. N., Horita J., and Mineev S. D., 2007. Equilibrium iron isotope fractionation factors of minerals: Reevaluation from the data of nuclear inelastic resonant X-ray scattering and Mössbauer spectroscopy. Geochim. Cosmochim. Acta 71, 3833-3846. Polyakov V. B., 2009. Equilibrium iron isotope fractionation at core-mantle boundary conditions. Science 323, 912-914. Pribil M. J., Wanty R. B., Ridley W. I., and Borrok D. M., 2010. Influence of sulfur-bearing polyatomic species on high precision measurements of Cu isotopic composition. Chem. Geol. 272, 49-54. Reed S. J. B., 1993. Electron microprobe analysis. Cambridge University Press, Cambridge. Reynolds B. C., Aggarwal J., Andre L., Baxter D., Beucher C., Brzezinski M. A., Engstrom E., Georg R. B., Land M., Leng M. J., Opfergelt S., Rodushkin I., Sloane H. J., van den Boorn S. H. J. M., Vroon P. Z., and Cardinal D., 2007. An inter-laboratory comparison of Si isotope reference materials. J. Anal. At. Spectrom. 22, 561-568. Ricard Y., Sramek O., and Dubuffet F., 2009. A multi-phase model of runaway core-mantle segregation in planetary embryos. Earth Planet. Sci. Lett. 284, 144-150. Richardson F. D., 1956. Activities in ternary silicate melts. T Faraday Soc 52, 1312-1324. Richter F. M., Liang Y., and Davis A. M., 1999. Isotope fractionation by diffusion in molten oxides. Geochim. Cosmochim. Acta 63, 2853-2861. Richter F. M., Watson E. B., Mendybaev R. A., Teng F. Z., and Janney P. E., 2008. Magnesium isotope fractionation in silicate melts by chemical and thermal diffusion. Geochim. Cosmochim. Acta 72, 206-220. Richter F. M., Dauphas N., and Teng F. Z., 2009a. Non-traditional fractiona- Page 186 — Bibliography

tion of non-traditional isotopes: Evaporation, chemical diffusion and Soret diffusion. Chem. Geol. 258, 92-103. Richter F. M., Watson E. B., Mendybaev R., Dauphas N., Georg B., Watkins J., and Valley J., 2009b. Isotopic fractionation of the major elements of molten basalt by chemical and thermal diffusion. Geochim. Cosmochim. Acta 73, 4250-4263. Righter K. and Drake M. J., 1996. Core formation in Earth’s Moon, Mars, and Vesta. Icarus 124, 513-529. Righter K., 2003. Metal-silicate partitioning of siderophile elements and core formation in the early Earth. Ann. Rev. Earth Planet Sci. 31, 135-174. Righter K., Pando K. M., Danielson L., and Lee C. T., 2010. Partitioning of Mo, P and other siderophile elements (Cu, Ga, Sn, Ni, Co, Cr, Mn, V, and W) between metal and silicate melt as a function of temperature and silicate melt composition. Earth Planet. Sci. Lett. 291, 1-9. Ringwood A. E., 1959. On the Chemical Evolution and Densities of the Planets. Geochim. Cosmochim. Acta 15, 257-283. Roskosz M., Luais B., Watson H. C., Toplis M. J., Alexander C. M. O., and Mysen B. O., 2006. Experimental quantification of the fractionation of Fe isotopes during metal segregation from a silicate melt. Earth Planet. Sci. Lett. 248, 851-867. Rubie D. C., Nimmo F., and Melosh H. J., 2007. Formation of Earth’s core. In: Stevenson D. J. (Ed.), Evolution of the Earth. Elsevier-Pergamon, Oxford. Rubie D. C., Frost D. J., Mann U., Asahara Y., Nimmo F., Tsuno K., Kegler P., Holzheid A., and Palme H., 2011. Heterogeneous accretion, composition and core-mantle differentiation of the Earth. Earth Planet. Sci. Lett. 301, 31-42. Rudge J. F., Reynolds B. C., and Bourdon B., 2009. The double spike toolbox. Chem. Geol. 265, 420-431. Rudge J. F., Kleine T., and Bourdon B., 2010. Broad bounds on Earth’s accretion and core formation constrained by geochemical models. Nat. Geosci. 3, 439-443. Saito T. and Maruya K., 1958. Diffusion of silicon and manganese in liquid iron: I Diffusion in liquid iron saturated with carbon. The 70th report of the Research Institute of Mineral Dressing and Mettalurgy, 259-268. Savage P. S., Georg R. B., Armytage R. M. G., Williams H. M., and Halliday A. N., 2010. Silicon isotope homogeneity in the mantle. Earth Planet. Sci. Lett. 295, 139-146. Schauble E. A., 2004. Applying stable isotope fractionation theory to new systems. Rev Mineral Geochem 55, 65-111. Schauble E. A., 2007. Role of nuclear volume in driving equilibrium stable isotope fractionation of mercury, thallium, and other very heavy elements. Geochim. Cosmochim. Acta 71, 2170-2189. Schmidt M. W., Connolly J. A. D., Gunther D., and Bogaerts M., 2006. Element Bibliography — Page 187

partitioning: The role of melt structure and composition. Science 312, 1646-1650. Schoenberg R. and von Blanckenburg F., 2005. An assessment of the accuracy of stable Fe isotope ratio measurements on samples with organic and inorganic matrices by high-resolution multicollector ICP-MS. Int. J. Mass Spectrom. 242, 257-272. Schoenberg R. and von Blanckenburg F., 2006. Modes of planetary-scale Fe isotope fractionation. Earth Planet. Sci. Lett. 252, 342-359. Schuessler J. A., Schoenberg R., Behrens H., and von Blanckenburg F., 2007. The experimental calibration of the iron isotope fractionation factor between pyrrhotite and peralkaline rhyolitic melt. Geochim. Cosmochim. Acta 71, 417-433. Shahar A., Young E. D., and Manning C. E., 2008. Equilibrium high-temperature Fe isotope fractionation between fayalite and magnetite: An experimental calibration. Earth Planet. Sci. Lett. 268, 330-338. Shahar A., Ziegler K., Young E. D., Ricolleau A., Schauble E. A., and Fei Y. W., 2009. Experimentally determined Si isotope fractionation between silicate and Fe metal and implications for Earth’s core formation. Earth Planet. Sci. Lett. 288, 228-234. Shahar A., Hillgren V. J., Young E. D., Fei Y. W., Macris C. A., and Deng L. W., 2011. High-temperature Si isotope fractionation between iron metal and silicate. Geochim. Cosmochim. Acta 75, 7688-7697. Sherman D. M., 1997. The composition of the Earth’s core: constraints on S and Si vs. temperature. Earth Planet. Sci. Lett. 153, 149-155. Siebert C., Nagler T. F., and Kramers J. D., 2001. Determination of molybdenum isotope fractionation by double-spike multicollector inductively coupled plasma mass spectrometry. Geochem. Geophy. Geosy. 2, art. no.-2000GC000124. Siebert J., Corgne A., and Ryerson F. J., 2011. Systematics of metal-silicate partitioning for many siderophile elements applied to Earth’s core formation. Geochim. Cosmochim. Acta 75, 1451-1489. Simon J. I. and DePaolo D. J., 2010. Stable calcium isotopic composition of meteorites and rocky planets. Earth Planet. Sci. Lett. 289, 457-466. Steelmaking T. J. S. f. t. P. o. S. a. t. N. C. o., 1988. Part 2: Recommended values of activity and activity coefficients, and interaction parameters of elements in iron alloys. Gordon and Breach Science Publishers, Montreux. Stevenson D. J., 2008. A planetary perspective on the deep Earth. Nature 451, 261-265. Tateno S., Hirose K., Sata N., and Ohishi Y., 2009. Determination of post-perovskite phase transition boundary up to 4400 K and implications for thermal structure in D ‘’ layer. Earth Planet. Sci. Lett. 277, 130-136. Taylor G. J., Keil K., Mccoy T., Haack H., and Scott E. R. D., 1993. Asteroid dif- Page 188 — Bibliography

ferentiation - Pyroclastic volcanism to magma oceans. Meteoritics 28, 34-52. Terasaki H., Frost D. J., Rubie D. C., and Langenhorst F., 2008. Percolative core formation in planetesimals. Earth Planet. Sci. Lett. 273, 132-137. Touboul M., Kleine T., Bourdon B., Palme H., and Wieler R., 2007. Late formation and prolonged differentiation of the Moon inferred from W isotopes in lunar metals. Nature 450, 1206-1209. Ulmer G. C., Grandstaff D. E., Woermann E., Gobbels M., Schonitz M., and Woodland A. B., 1998. The redox stability of moissanite (SiC) compared with metal-metal oxide buffers at 1773 K and at pressures up to 90 kbar. Neues Jb Miner Abh 172, 279-307. Urey H. C., 1947. The thermodynamic properties of isotopic substances. J. Chem. Soc., 562-581. Van der Laan S., Zhang Y. X., Kennedy A. K., and Wyllie P. J., 1994. Compari- son of element and isotope diffusion of K and Ca in multicomponent silicate melts. Earth Planet. Sci. Lett. 123, 155-166. Villars P., Prince A., and Okamoto H., 1995. Handbook of ternary alloy phase diagrams. ASM International. Wade J. and Wood B. J., 2005. Core formation and the oxidation state of the Earth. Earth Planet. Sci. Lett. 236, 78-95. Wade J., Wood B. J., and Tuff J., 2012. Metal-silicate partitioning of Mo and W at high pressures and temperatures: Evidence for late accretion of sulphur to the Earth. Geochim. Cosmochim. Acta 85, 58-74. Walter M., 2003. Melt extraction and compositional variability in mantle lithosphere. In: Holland D. and Turekian K. K. Eds.), The mantle and core. Elsevier-Pergamon, Oxford. Walter M. J. and Thibault Y., 1995. Partitioning of Tungsten and Molybdenum between Metallic Liquid and Silicate Melt. Science 270, 1186-1189. Wänke H., 1981. Constitution of terrestrial planets. Philos T Roy Soc A 303, 287-302. Watson E. B., Wark D. A., Price J. D., and Van Orman J. A., 2002. Mapping the thermal structure of solid-media pressure assemblies. Contrib. Mineral. Petrol. 142, 640-652. Weiss B. P., Berdahl J. S., Elkins-Tanton L., Stanley S., Lima E. A., and Carporzen L., 2008. Magnetism on the Angrite Parent Body and the Early Differentia- tion of Planetesimals. Science 322, 713-716. Weyer S., Anbar A. D., Brey G. P., Munker C., Mezger K., and Woodland A. B., 2005. Iron isotope fractionation during planetary differentiation. Earth Planet. Sci. Lett. 240, 251-264. Weyer S. and Ionov D. A., 2007. Partial melting and melt percolation in the mantle: The message from Fe isotopes. Earth Planet. Sci. Lett. 259, 119-133. Wiechert U. and Halliday A. N., 2007. Non-chondritic magnesium and the origins of the inner terrestrial planets. Earth Planet. Sci. Lett. 256, 360-371. Bibliography — Page 189

Williams H. M., McCammon C. A., Peslier A. H., Halliday A. N., Teutsch N., Levasseur S., and Burg J. P., 2004. Iron isotope fractionation and the oxygen fugacity of the mantle. Science 304, 1656-1659. Williams H. M., Markowski A., Quitte G., Halliday A. N., Teutsch N., and Levasseur S., 2006. Fe isotope fractionation in iron meteorites: New insights into metal-sulphide segregation and planetary accretion. Earth Planet. Sci. Lett. 250, 486-500. Williams H. M. and Archer C., 2011. Copper stable isotopes as tracers of metal-sulphide segregation and fractional crystallisation processes on iron meteorite parent bodies. Geochim. Cosmochim. Acta 75, 3166-3178. Williams H. M., Wood B. J., Wade J., Frost D. J., and Tuff J., 2012. Isotopic evidence for internal oxidation of the Earth’s mantle during accretion. Earth Planet. Sci. Lett. 231-322, 54-63. Wood B. J., Walter M. J., and Wade J., 2006. Accretion of the Earth and segregation of its core. Nature 441, 825-833. Yoshino T., Walter M. J., and Katsura T., 2003. Core formation in planetesimals triggered by permeable flow.Nature 422, 154-157. Young E. D., Galy A., and Nagahara H., 2002. Kinetic and equilibrium mass- dependent isotope fractionation laws in nature and their geochemical and cosmochemical significance.Geochim. Cosmochim. Acta 66, 1095-1104. Zanda B., 2004. Chondrules. Earth Planet. Sci. Lett. 224, 1-17. Zhu X. K., Guo Y., Williams R. J. P., O’Nions R. K., Matthews A., Belshaw N. S., Canters G. W., de Waal E. C., Weser U., Burgess B. K., and Salvato B., 2002. Mass fractionation processes of transition metal isotopes. Earth Planet. Sci. Lett. 200, 47-62. Ziegler K., Young E. D., Schauble E. A., and Wasson J. T., 2010. Metal-silicate silicon isotope fractionation in enstatite meteorites and constraints on Earth’s core formation. Earth Planet. Sci. Lett. 295, 487-496. Page 190 — Appendix AI.

Appendix AI.

In general, stable isotope fractionation arises from the desire of a system to minimise its energy. On the atomic scale, this involves minimising the (potential) energy of an atom on its site when it is bound to other atoms; the term zero- point energy refers to the minimum (potential) energy. Figure I.1a schematically shows the zero-point energy as a function of interatomic distance in a diatomic molecule. At large distance, the atoms move independently of one another, but when they come closer they attract each other, losing potential energy, until they get too close and repel one another again. This potential energy, however, is not a continuum, but is quantised; i.e. only discrete levels of energy can be occupied. The lowest level (zero-point energy) is equal to ½hν, where h is Planck’s constant and ν the vibrational frequency: 1 A ν = (AI.1) 2π m in which A is a factor that is only dependent on force constants, and m is the mass of the atom. The zero-point energy of an atom thus decreases when its mass increases, or, in other words, heavy isotopes have a lower zero-point energy than light isotopes. Figure I.1. Schematic represen- a Dissociated atoms tation of the energy of atoms in diatomic molecules. Panel (a) shows the potential energy as a function of interatomic EL E H distance. The energy of disso- Touching 0 ciated atoms decreases once atoms Quantised E levels they are close enough to be attracted to each other, until the ½hν L energy rapidly increases due to ν ½h H Potential energy repellent forces. The energy the atoms can obtain is not con- Interatomic distance (~10-10 m) tinuous but quantised, meaning that only discrete energy levels can be obtained. The minimum 0 b energy level (E , ‘zero-point energy’) is given by ½hν, Rotational meaning that heavy atoms have lower zero-point energy than Vibrational light atoms. Panel (b) gives the three types of motions that Vibrational are used to describe the total Rotational energy of an atom: vibrational, rotational and translational Translational energy. Appendix AI. — Page 191

Urey (1947) and Bigeleisen and Mayer (1947) showed that for isotopic exchange reactions, the equilibrium constant (K) can be determined by calculating the energy changes associated with the reactions. The energy of a molecule can be calculated from so-called partition functions (Q), which essentially sum the probabilities (over all energy levels) that a molecule occupies a certain energy level. Hence, we can write an isotope exchange reaction between two molecules (or phases) A and B with the light and heavy isotopes 1 and 2, respectively:

aA12+=+ bB aA 21 bB Because isotope exchange reactions don’t involve changes in molar volumes, the Helmholtz free energy and the Gibbs free energy are identical for these reactions. It follows therefore, that: G= RTln() Q (AI.2) And because of the familiar relation: D=G RTln() K (AI.3) the equilibrium constant of the isotope exchange reaction can be expressed as: a Q A2 Q A1 (AI.4) K = b Q B2 Q B1 where the powers a and b can be set to 1 assuming that only one atom is exchanged in its molecule/phase. The partition function (Q) is actually the summation of the partition functions for the vibrational, rotational and translational energies of the molecule (Fig. 1b). As indicated in Urey (1947) and Bigeleisen and Mayer (1947), however, with exception of H atoms, the rotational and translational energies cancel out for ratios of partition functions of a single Q molecule (e.g. A2 ) when temperatures exceed ~100 K. The partition function Q A1 is therefore effectively the partition function for vibrational energies, and it is sometimes referred to as the reduced partition function, which is often indicated with the symbol β instead of Q2 /Q1. Both classical papers then derive an equation for the reduced partition function of polyatomic molecules: u 2i −u 2 1i Q σ u2 ee1− 21= i (AI.5) ∏ −u u 22ii Quσ i 1− e 12 1i e 2 where σ stands for the symmetry number of the molecule, while: hν u = i (AI.6) i kT Page 192 — Appendix AI.

where h is again Planck’s constant, k Boltzmann’s constant and T the temperature in K. Bigeleisen and Mayer (1947) further simplified this equation by defining uu= +D u and by realising that Du is small for every scenario 12ii i i except for exchange of H and D, and that ui itself is small in case T > ~100 K: σ Q Du 2 22=1 + ∑ i (AI.7) σ11Q i 24 It follows from equation (AI.1) that: 1 Dm D=ν 2 (AI.8) ∑ i 2 A i 4π mmij Substitution of equation (AI.7) into (AI.8) yields: σ Q DmA 22=+⋅ (AI.9) 1 C 2 σ11Q mmij T with: 2 h 1 C =  (AI.10) k 96π 2 Symmetry numbers are often ignored for two reasons: i) the ratios of symmetry numbers are approximately 1 because the symmetry of a molecule is not much affected by isotopic substitution, and ii) their exact values are very difficult to determine accurately. Therefore, the definition: σ Q 22 σ 11Q A (AI.11) α AB− = σ Q 22 σ 11Q B simplifies to: Q 2 Q1 A (AI.12) α AB− ≈≈K Q 2 Q 1 B It thus follows that: DmA DDmA mDA (AI.13) ln()α ≈ ln11 +⋅ C −ln +⋅ C ≈⋅C AB− AB− mm T2mm T 22 mm T ij ABij ij This equation is the source for equation (I.4) in chapter I of this thesis (Intro- duction), in which the symbol A in equation (AI.16) is simply re-named as F, 2 and where mimj is simplified as m because for non-traditional isotopes mass 2 differences are small enough that mimj ≈ m . Furthermore, for qualitative or semi- quantitative purposes, the constant C can be ignored. Appendix AI. — Page 193

Equations (AI.5) and (AI.6) form the basis for theoretical calculations of stable isotope fractionation factors, because they only require two input parameters: vibrational frequencies (ν) and symmetry numbers. As mentioned above, symmetry numbers are often ignored, which is a major source of uncertainty in theoretical calculations of stable isotope fractionation factors. Vibrational frequencies can either be directly calculated from, for instance, Raman, Mössbauer, or Inelastic Nuclear Resonant X-ray spectroscopy, where sources of uncertainty mainly arise from difficulties to measure certain frequencies. Page 194 — Appendix AII. Appendix AII. 15 6.13 8.95 8.97 2.98 1.00 70.14 29.86 59.91 12.06 0.5696 0.4304 0.1931 14 5.83 4.51 4.43 1.02 30.11 66.93 27.24 47.90 12.03 0.3449 0.6551 0.2902 13 4.92 6.92 8.79 7.64 8.91 3.43 1.04 69.09 19.06 56.47 13.72 0.3579 0.6421 0.2726 12 8.51 9.69 4.02 11.82 49.42 16.54 0.1371 11 1.40 6.95 81.61 10.05 0.0754 10 1.60 6.13 8.95 8.97 2.98 1.00 73.73 24.67 59.91 12.06 0.5949 0.4051 0.8409 9 1.97 51.83 15.30 16.71 14.18 0.1169 1.62 5.67 5.22 1.79 0.45 8(II) 82.24 10.47 49.55 14.55 15.24 13.20 0.7473 0.2527 0.2643 8 1.53 5.68 5.05 1.80 0.49 82.39 10.40 49.59 14.58 15.38 13.10 0.7499 0.2501 0.5268 7 5.10 7.33 0.94 45.57 25.19 15.88 0.2546 6 1.43 5.22 7.19 5.64 1.10 72.72 25.85 70.32 10.53 0.4270 0.5730 0.7167 5 1.72 7.63 9.87 3.59 1.83 73.69 24.59 49.79 16.73 10.55 0.5155 0.4845 0.9686 4 1.61 6.88 9.63 3.67 1.38 57.68 40.71 46.79 15.14 16.51 0.3272 0.6728 0.8166 3 1.68 4.56 6.58 5.10 2.02 0.70 57.68 40.65 70.70 10.33 0.3418 0.6582 0.7569 2 1.61 7.60 9.86 3.59 1.83 57.68 40.71 49.99 16.62 10.52 0.3647 0.6353 0.7326 1 1.31 8.97 9.05 9.96 3.98 2.04 11.74 50.98 35.97 48.04 17.97 0.3672 0.6328 0.7998 1 3 3 3 2 O O O 2 O 2 2 O 2 2 The working names of these starting mixtures were either ‘PwdM-Sil-N’ or ‘Pwd-N’ (where N in both cases stands for the number or ‘Pwd-N’ The working names of these starting mixtures were either ‘PwdM-Sil-N’ Appendix AII. Compositions (wt%) of starting mixtures. Appendix Name 1 given above), depending on whether it was a metal-silicate mixture or not. Fe Metal Sn Mo Si Zn S Ni SiO Silicate Al MgO Fe FeO CaO Na K B Metal fraction Silicate fraction Total weight (g) Total Appendix AII. — Page 195 30a 1.07 7.59 9.46 1.81 1.05 11.32 73.25 25.68 55.19 13.57 0.3485 0.6515 0.1908 29 1.97 2.99 4.88 2.86 11.12 83.00 17.00 61.24 14.94 28 8.09 1.47 0.99 4.08 0.89 71.21 20.70 47.88 44.68 0.3475 0.6525 0.1269 27 7.82 2.92 3.51 1.03 11.02 72.05 20.14 49.02 32.50 0.6444 0.4587 0.3556 26a 1.25 1.87 6.93 8.88 0.68 0.30 77.01 21.74 49.46 31.89 0.6499 0.2176 0.3501 25a 2.24 7.46 9.21 1.82 1.07 11.42 84.56 13.20 55.35 13.66 0.3471 0.6529 0.1082 24a 7.59 6.68 6.49 1.66 0.95 66.71 25.70 55.14 15.67 13.41 0.1196 0.6529 0.3471 23a 7.63 2.26 1.30 9.05 0.68 0.29 68.30 24.08 55.16 31.25 0.1171 0.3495 0.6505 22 7.65 2.26 1.45 8.90 0.67 0.28 92.35 54.23 32.20 0.6555 0.2364 0.3445 21 9.59 8.61 8.49 6.42 4.16 1.02 71.27 19.14 52.80 18.48 0.6679 0.3701 0.3321 20 7.89 2.69 9.15 9.66 5.91 5.03 75.52 13.90 51.61 18.63 0.3383 0.6617 0.2825 19a 7.76 2.10 1.37 9.06 0.68 0.29 71.99 20.25 55.24 31.26 0.6506 0.4669 0.3494 19 7.82 2.92 3.51 1.03 11.02 72.05 20.14 49.02 32.50 0.3556 0.6444 0.4587 18 7.83 2.90 3.56 1.02 11.03 71.89 20.28 48.99 32.50 0.3274 0.6726 0.3054 17 6.04 93.96 0.0572 16 1.48 5.65 5.13 1.80 0.49 92.87 49.53 14.55 15.47 13.04 0.7584 0.2416 0.4409 1 3 3 3 2 O O O 2 O 2 2 O 2 2 The working names of these starting mixtures were either ‘PwdM-Sil-N’ or ‘Pwd-N’ (where N in both cases stands for the number or ‘Pwd-N’ The working names of these starting mixtures were either ‘PwdM-Sil-N’ Appendix AII continued. Compositions (wt%) of starting mixtures. Appendix 1 given above), depending on whether it was a metal-silicate mixture or not. Name Metal Fe Sn Mo Si Zn S Ni Silicate SiO Al MgO Fe FeO CaO Na K B Silicate fraction weight (g) Total Metal fraction Page 196 — Appendix AIII.

Appendix AIII.

2 The test than in RH01, but still too much infiltration. 2 glass capsule (donated by Bernie Wood). Too much leaking Too Wood). glass capsule (donated by Bernie 2

Comments with crushable Zirconia as inner capsule: too much infiltration of melt, Test all Fe in Pt jacket + not reducing enough for Zn metal, so wt% levels of ZnO Many tests with graphite capsules: various sources of used, but all resulted in too much infiltration of melt with capsule collaps as a consequence. with a quartz had too much leaking at the flat disk lid. See more comments in text, section II.2.2. Industrial Carbon, test pieces donated First success. Graphite source: Morgan Later ordered from company as product code M/09FG07 Wood. by Bernie (November 2009, contact person Edwin Haas). www.morgancarbon.com.au with SiO Test by opx. (See Figure a in this at the flat disk lid; olivine crystels overgrown Appendix.) III of thesis. Published in chapter Raman analyses and later Succes, but not published. Used e.g. for III of thesis. mentioned in chapter of which are capsules analyses, results Analyses not satisfactory for publication in basalt liquid. Soret diffusion (matrix issue). with much denser ZrO Test of thesis. IV Published in chapter III of thesis. Published in chapter with dense alumina capsule. Crystallised corundum and leaking along Test plugs. (Fig. b) that stabilised spinel.

8 24 16 20 22 ~6 0.5 (h) ~24 ~20 ~48 8+4 Duration T 1150 (°C) 1250 1200 1200 1200 1200 1250 1375 1200 1400 1200

) 2 3

2 2 -C; O 50 2 glass 2 Pd Pt-C Pt-C Pt-C Pt-C Pt-C Pt-C Pt-C; 50 Pt-ZrO Pt-ZrO Pt-Al Capsule Au (1 x Pt-SiO Pt-SiO

S S S S S S S S S C S+C Static (S) Centrifuge (C)/

-5 -5 -8 -5 -1 -5 -6 -5 -1-4 M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil Pwd-7 Pwd-7 mixture Starting Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Si Fe Fe Fe Fe Fe Fe Fe Fe Fe Fe system Isotopic Appendix AIII. List of experiments. Note that temperatures given here are thermocouple temperatures; they do not include thermocouple temperatures; are given here AIII. List of experiments. Note that temperatures Appendix images of failed experiments that do not appear in this Back-scatter electron for sample temperatures. a correction Appendix, can be found in chapter II. Name RH01 RH02- 20 RH21 RH22 RH23 RH24 RH25 RH26 RH27 RH28 RH29 Appendix AIII. — Page 197

(Fig. d)

-doped. Published in chapter III of thesis. -doped. Published in chapter 3 O 2 Fe Comments III of thesis. Published in chapter of thesis. IV Published in chapter T that 20°C because I realised increased T of thesis; IV Published in chapter gradient as I thought initially; T as much with increased doesn’t increase probably in RH36. tested this later III of thesis. Published in chapter This time corundum saturated basalt as start with dense alumina capsule. Test composition. Still leaking along plugs; probably works well with proper capsules, but these will be difficult to machine (dense alumina is very hard), so I recommend them as last alternative. (Fig. c) Prepared for analyses, but in the end only analysed by fs-LA-MC-ICPMS calibration (spinel crystallisation method) with capsule inside Temperature III of thesis. Results published in chapter assembly. of thesis. Published in chapter IV what was in capsule. Never ran this experiment. Forgot with quite coarse grained, very low gas permeability graphite (donated by Test Colin Maden). Quite succesfull, but not quite as good graphite from Morgan Industrial Carbon. Successful test with single crystel MgO capsule. with ultrapure (free of ‘dust filling’) graphite, donated by Colin Maden. Test Unsuccesful: same infiltration problem as before. Reaction with ‘dust filling’ apparently not the issue. (Fig. e) 54

3 4 4 ~4 0.5 1.5 1.5 3.5 (h) ~62 ~24 ~0.5 ~4.5 Duration T (°C) 1200 1375 1395 1200 1350 1395 1400 1395 1250 1395 1250 1250 3 S.C. O 2 Pt Pt Pt-C Pt-C Pt-C Pt-C Pt-C Pt-C Pt-C Pt-C Pt-Al Capsule Pt-MgO

S S S S S S C C C C C C Static (S) Centrifuge (C)/

-5 -8 -8 -5 -8 -5 -8 -5 -10 - M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil Pwd-7 Pwd-9 mixture Starting Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd - Si Si Si Si Si Fe Fe Fe Fe Fe Fe system Isotopic Appendix AIII continued. List of experiments. Appendix Name RH30 RH31 RH32 RH33 RH34 RH35 RH36 RH37 RH38 RH39 RH40 RH41 RH42 Page 198 — Appendix AIII. content. Only a bit too much Sn in metal. Analysed by LA- content. Only a bit too much Sn in metal. 3 glass softening (has higher temperature than 2 O 2 , but lowered it. Should have added it as NiO, would stabilise 2 O -doped. Published in chapter III of thesis. -doped. Published in chapter III of thesis. -doped. Published in chapter III of thesis. -doped. Published in chapter f 3 3 3 O O O 2 2 2 with higher Fe 2 Fe Fe Fe O Comments 54 54 metal liquids separated by wafer: for laser ablation tests with different Two matrices. well Worked much plagioclase crystallisation. Too in basalt liquid. Soret diffusion in RH25, so less water this time? (Fig. f) Analyses not satisfactory for publication (matrix in basalt liquid. Soret diffusion issue). 54 Failed because capsule cracked along cleavage of MgO, probably during quenching after SiO pyrex). (Fig. g) high current during experiment. Gave up on Too Again attempt for Soret diffusion. it is apparently written in the stars that I shall not get publishable Soret diffusion: results for this. of thesis. IV Published in chapter Added Ni to reach to find good chemical system for Mo isotope project. Test higher olivine. Fairly succesful test for chemical system Mo isotope project. Reached higher f ICPMS for Mo content in silicate. at ~1680°C to see whether issue with RH49 was due only the high Test temperature, or rather due to the quenching as initially hypothesi z ed.

4 4 2 8 4 ~5 ~5 ~6 ~4 0.3 1.5 (h) ~3.5 Duration T 1150 (°C) 1250 1250 1395 1250 1280 1250 1680 1395 1290 1300 1680

- - S.C. -C S.C. S.C. 50 50 50 Pd Pd glass glass Pd 2 2 50 50 Pt-C Pt-C Pt-C Pt-C Pt-C Pt-C 50 Capsule Au Au SiO SiO Pt-MgO Pt-MgO Au Pt-MgO

S S S S S S S S C C C C Static (S) Centrifuge (C)/

-8 + -8 -8 -10 -10 -10 -13 -14 3 17 O 2 M-Sil Si M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil 83 Fe Pwd-7 Pwd-11 Pwd-12 Pwd-9+ mixture Fe Starting Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Si Si Si Si Fe Fe Fe Fe Fe Fe Mo Mo system Isotopic Name RH43 RH44 RH45 RH46 RH47 RH48 RH49 RH50 RH51 RH52 RH53 RH54 Appendix AIII continued. List of experiments. Appendix Appendix AIII. — Page 199 sleeve to accomodate 3 O 2 glass softening. 2 glass softening (has higher temperature than 2 sleeve around 5 mm Pt capsule always shows faulting. Probably requires larger volume for analysis. Graphite capsule always collapses requires larger 3 2

O O 2 f Comments small amount of metal to see whether Mo isotope project Succesful test for centrifugation. Separated metal and silicate separation after is sufficient for of thesis. V analysed and published in chapter Succesful test with S-bearing composition. Analyses not satisfactory for publication (matrix issue). of thesis. IV Published in chapter metal liquids separated by wafer: for laser ablation tests with different Two matrices. Failed because capsule cracked in right angle to its long axis, probably during quenching after SiO pyrex). Repeat of RH60. Failed again for same reason, despite that I cooled slowly (instead of quenching) after SiO Unsuccesfull test with 4.5 mm instead of 3.5 (O.D.) graphite capsule, because lower and Al Al volume of graphite is too much compacted for large compaction. Succesful test with chemical composition for Mo isotope project in an MgO capsule. Analysed by LA-ICPMS for Mo content in silicate. Repeat of RH62 with same result. (Fig. h)

4 4 ~6 ~7 ~4 ~4 1.5 1.5 1.5 (h) 0.25 Duration T (°C) 1300 1050 1200 1700 1525 1700 1700 1300 1300 1300

- - - S.C. 50 50 50 S.C. Pd Pd Pd glass glass 2 2 50 50 50 Pt-C Pt-C Pt-C Pt-C Pt-C Pt-C MgO Capsule Au Au Au SiO SiO Pt-MgO

S S S S C C C C C C Static (S) Centrifuge (C)/

-8 -14 -15 -15 -16 -16 -14 -18 -18 M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil Pwd-17 mixture Starting Pwd-11+ Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Si Si Si Si Fe Fe Mo Mo Mo Mo system Isotopic Name RH55 RH56 RH57 RH58 RH59 RH60 RH61 RH62 RH63 RH64 Appendix AIII continued. List of experiments. Appendix Page 200 — Appendix AIII.

2 experiments. Too low Mo contents in silicate glass. Too experiments. low Mo contents in silicate glass, but would Too experiments. 2 2 O O f f -doped. Published in chapter III of thesis. -doped. Published in chapter 3

O 2

Fe Comments of thesis. Successful this time because I skipped SiO IV Published in chapter glass than usual). Simply pumped pressure softening step (hence somewhat lower a bit Stepped furnaces do have a tendency to shear pressure. cold to desired at the step. of thesis. IV Published in chapter of thesis. IV Published in chapter Homogenisation of potential metal standard for Si isotope work. Eventually not used because LA-MC-ICPMS failed. remember why it failed. Failed repeat of RH57. I don’t for lower Test of thesis. V Published in chapter of thesis. V Published in chapter 54 of thesis. V Published in chapter for lower Test be sufficient with 4.5 instead of 3.5 mm (O.D.) inner capsule and slightly more sensitive machine. of thesis. V Published in chapter of thesis. V Published in chapter

4 6 4 4 8 4 ~6 1.5 1.5 0.5 0.5 1.5 (h) ~0.1 Duration T (°C) 1700 1700 1700 1500 1200 1300 1350 1350 1250 1350 1350 1350 1350

- - - - -C -C -C -C S.C. S.C. S.C. -C 50 50 50 50 50 50 50 50 50 S.C. Pd Pd Pd Pd glass glass glass C Pd Pd Pd Pd Pd 2 2 2 50 50 50 50 50 50 50 50 50 MgO Capsule Au Au Au Au SiO SiO SiO Au Pt-MgO Pt-MgO Pt-MgO Au Au Au Au

S C C C C C C C C C C C C Static (S) Centrifuge (C)/

-16 -16 -19 -19 -15 -19 -19 -15 -20 -21 -19a -8(II) M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil mixture SRM891 Starting Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Si Si Si Si Fe Fe Mo Mo Mo Mo Mo Mo Mo system Isotopic Appendix AIII continued. List of experiments. Appendix Name RH65 RH66 RH67 RH68 RH69 RH70 RH71 RH72 RH73 RH74 RH75 RH76 RH77 Appendix AIII. — Page 201 / 4+ to (hopefully) vary Mo 2 O f content for higher 3 O 2

o6+. Comments to see Westrenen) van Wim (and Tronche Unsuccesful experiment for Elodie whether when sample Lost ocean. magma lunar from by density separates really plagiclase polishing. AuPd from the Unsuccesful repeat of RH77, because it had leaked a bit Too capsule. low centrifugation. of thesis. V Published in chapter of thesis. V Published in chapter of thesis. V Published in chapter to see Westrenen) van Wim (and Tronche Unsuccesful experiment for Elodie whether Above liquidus. plagiclase really separates by density from lunar magma ocean. Failed test with S-bearing system for Mo isotope project. Metal not above liquidus. Unsuccesful because metal and silicate not well separated. Stepped furnace sheared again.Stopped using stepped furnaces. of thesis. V Published in chapter of thesis. V Published in chapter Failed test with S-bearing system for Mo isotope project. Metal still not above liquidus. Unsuccesful because silicate not chemically homogeneous. Succesful test with higher Fe M Lost by electrostatics in chemistry.

6 2 1 1 2 6 1 1 ~4 ~2 (h) 1.25 ~0.75 ~0.75 ~0.75 Duration T (°C) 1240 1350 1550 1550 1350 1240 1350 1550 1550 1550 1550 1550 1350 1550 S.C. S.C. S.C. S.C. S.C. S.C. S.C. S.C. S.C. Pt-C Pt-C Pt-C Pt-C Pt-C Capsule Pt-MgO Pt-MgO Pt-MgO Pt-MgO Pt-MgO Pt-MgO Pt-MgO Pt-MgO Pt-MgO

S S C C C C C C C C C C C C Static (S) Centrifuge (C)/

-27 -19a -22a -19a -22a -19a -19a -23a -19a -24a -19a -22a M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil Elodie Elodie Tronche Tronche mixture Starting Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd - - Mo Mo Mo Mo Mo Mo Mo Mo Mo Mo Mo Mo system Isotopic Name RH78 RH79 RH80 RH81 RH82 RH83 RH84 RH85 RH86 RH87 RH88 RH89 RH90 RH91 Appendix AIII continued. List of experiments. Appendix Page 202 — Appendix AIII.

Comments of thesis. V Published in chapter Lost by electrostatics in chemistry. of thesis. V Published in chapter Failed test with S-bearing composition for Mo isotope project. Flat disk used as lid and leaked again along lid. with S-bearing composition for Mo isotope project. Fairly succesful: metal Test liquid, but with two-liquid immiscibility. of thesis. V Published in chapter of thesis. V Published in chapter of thesis. V Published in chapter Analysed, but finally not published. of thesis. V Published in chapter of thesis. V Published in chapter of thesis. V Published in chapter of thesis. V Published in chapter Failed attempt to adjust chemical composition based on RH25a get a single metal liquid. experiments. with B2O3-bearing for Si diffusion Test

1 4 52 1.5 1.5 1.5 1.5 1.5 3.5 (h) 1.75 1.75 1.25 ~1.5 ~1.5 ~1.5 Duration T (°C) 1550 1350 1350 1350 1350 1350 1350 1350 1350 1350 1550 1550 1350 1350 1060 S.C. S.C. S.C. S.C. -C S.C. S.C. S.C. S.C. 50 Pd Pt-C Pt-C Pt-C Pt-C Pt-C Pt-C 50 Capsule Pt-MgO Pt-MgO Pt-MgO Pt-MgO Au Pt-MgO Pt-MgO Pt-MgO Pt-MgO

S S S C C C C C C C C C C C C Static (S) Centrifuge (C)/

-21 -21 -21 -28 -21 -27 -29 -19a -26a -22a -19a -25a -25a -25a -30a M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil M-Sil mixture Starting Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Pwd Si Mo Mo Mo Mo Mo Mo Mo Mo Mo Mo Mo Mo Mo Mo system Isotopic Appendix AIII continued. List of experiments. Appendix Name RH92 RH93 RH94 RH95 RH96 RH97 RH98 RH99 RH100 RH101 RH102 RH103 RH104 RH105 RH106 Appendix AIII. — Page 203

Appendix AIII continued. Back-scatter electron images. Page 204 — Acknowledgements

ACKNOWLEDGEMENTS

I suppose the end of a PhD thesis, generally culminating in the writer’s first book(let), is a customary moment to become sentimental and think about the important things in life. I cannot, however, in all honesty say that writing this thesis has felt like a four year job of blood, sweat and tears that justifies sentimental behaviour. On the other hand, as some people who know me very well can confirm, I tend to be a bit of philosopher and thus think about the important things in life even when it is not customary to do so. I find that the latter justifies acknowledging my parents and my sister above anyone else who appears in this text. They have been a great support throughout my life, in good and bad times, and they have always set good examples. It is up to others to judge, but I like to think they could have done a much worse job raising me. And Saskia, thanks for being the older one, the pioneer who walked all the paths before me so I knew what to expect.

Specifically for this PhD thesis, I have to express my gratitude to my supervisor, Max W. Schmidt, for giving me the space to do what I wanted or what I thought was right to do. Having been given so much freedom felt good. The best education is furthermore to find things out for oneself, and I feel I have been able to do that. When I had my interview here, I had no idea what Max talked about when he showed me the experimental equipment during my lab tour. About two or three weeks after I started, in her PhD defence talk someone presented a data table to show that she had calibrated acoustic waves through WC for her experiments; I wondered for about 10 seconds why for goodness sake she needed to calibrate a toilet, until I suddenly understood that WC stood for tungsten carbide, one of the materials the experimental equipment is made of. Now, four years later, I feel piston cylinders have no secrets for me anymore. I owe this for a large part to the freedom Max gave me. I also want to thank my co-supervisor, Bernard Bourdon. Although we did not discuss very often, the times that we did were truly great for me: I always left his office, or wherever we discussed, with the feeling I had made a step forward and, even more importantly, with more motivation and more ideas than I had entered our discussion with. I much appreciated Bernard’s ability to actually listen to what I wanted to say, think about it and then reply to it, never afraid to reply with a question or a comment that was clearly open to further discussion. In addition, I suppose this is also the place to apologise for any inappropriately long e-mails I wrote to him or for subjects I wanted to discuss that had less to do with science than we both wished. There were moments where I used him as a mentor, despite the fact that he did not sign up for that job. Yet he never complained. Many thanks go to Tim Elliott for his positive response to the request to read Acknowledgements — Page 205 my entire thesis and come to Zurich to comment it. I also much appreciate it that my sister (specialised in ancient history rather than geosciences) spent some of her holiday time to pay back a four year old debt by carefully reading my thesis in order to pick out spelling mistakes and to improve its legibility. Although they were not directly involved in my PhD, I would also like to express my gratitude to Gareth Davies and Wim van Westrenen at the Vrije Uni- versiteit Amsterdam. Their great education and their enthusiasm have had a large influence on my decision to continue to work in science and I am still grateful for all guidance they offered me during my studies in Amsterdam. I have enjoyed good collaborations with Caroline Fitoussi and particularly Christoph Burkhardt. Christoph, thanks for sharing your knowledge and the good and open collaboration on our molybdenum isotope work. During my stay here, I feel I have made many friends, but one of them stands out: Mattia Pistone. Despite that we are in many respects opposite characters, we have been great friends from the start until the end. Thanks, Mattia, for the chats at the end of our writing work, thanks for the times we spoke about life, thanks for regularly going somewhere together to meet others without either of us having the slightest idea where that somewhere actually was, and, above all, thanks for all the fun moments we shared anywhere on this globe. Although she arrived a year later than Mattia and I, Rita Seifert has become an equally good friend: thanks for simply being such a good friend, and for the skiing, the hiking, the dinners, and for sharing some beautiful moments with Mattia and me on Hawaii and in Italy. I am afraid there are too many people to name here for contributing to the good time I had in Zurich. Of all the people I have learned to know in my life, Paola Ardia probably takes best care of her friends; I am glad to be one them. Together with ‘the French girls’ (Jessica Langlade, Marion Louvel, Marion Campani; and Pinar Ozfirat as semi-French) she took me up in the ETH community from the first day I arrived. I am grateful for how open the five of you were to my presence, even when I just sat with you and had nothing to say. Thanks also to Esther Schwarzenbach with whom I did not just share office, but also a telephone for most of my stay here. Although in the first few weeks we barely spoke and I only learned that there is another language besides Dutch that cannot be any good for the health of one’s throat, we have become close friends. While I write these words, I am about to visit her in Blacksburg, Virginia, for a relaxing holiday! Many thanks also to Ute Mann, for being able to speak about anything and for all the education in the piston lab, sometimes by frustrating and time consuming trial-and-error (sorry for the too optimistic back-packing trip in Wallis, by the way!); and to Tamara Baumberger, for the fun during multiple occasions, not the least the tennis evenings with Esther, Mark, Arno, Greg and Florian. I cannot in good conscience finish my PhD thesis without acknowledging all Page 206 — Acknowledgements the girls I shared office with over the past four years: thanks for the good working atmosphere! I very much appreciated the presence of Claudia Büchel, who organises a lot in our institute including putting me in that office with all those girls, and Lydia Zehnder, who has been a great and very friendly help for plenty of things related to the labs! In those laboratories, I could not have done my work properly without Bruno Zürcher, Urs Graber, Thomas Good and Andreas Jallas for the experimental labs, and particularly Felix Oberli in the ICP lab. They always repaired or produced what I needed, when I needed it and they always delivered the best- of-quality, even though I requested some of them to repeatedly do some very frustrating jobs. Carmen Sanchez-Valle was a great skiing partner and somebody I could rely on for some good advice when I needed it. Having Peter Ulmer next door is rather amazing: he seems to know everything and seems able to explain it to everyone. I also want to thank Jamie Connolly and Christian Liebske for the cycling that Jamie and I should have started six months after I arrived instead of six months before I left. Thanks also, Jamie, for letting me know that my sometimes lazy attitude is simply a consequence of the second law of thermody- namics (minimisation of energy) and for telling me that the best way to finish a PhD is to cycle more. Last-but-not-least I have a desire to show my appreciation to two things, instead of persons. First and foremost, this is my bicycle, because it has always been and continues to be the one thing I can always rely on to clear my mind. In addition, when I did not need it to clear my mind, it has given me plenty of thoughts, many of which I think were truly brilliant; even it is just because the vast majority of them were gone with the wind before I could realise they were probably not so brilliant after all. My last gratitude then goes to the Swiss Alps. Although this thesis did not result from four years of blood, sweat and tears, there were certainly stressful and difficult moments. After five or six days of hard work, though, one day in the sun in the mountains completely recharged my batteries. Finally, Dutch people go to the mountains when they have holiday; simply because we do not have any mountains of our own. So many thanks to the Swiss Alps, for despite the difficult times they never stopped giving me the feeling this PhD was after all just one big holiday. If only I had spent more time enjoying their beauty.