# The variability of ruthenium in chromite from chassignite and olivine-phyric shergottite meteorites: New insights into the behavior of PGE and sulfur in Martian magmatic systems

Raphael J. Baumgartner, Marco L. Fiorentini, David Baratoux, Ludovic

Ferrière, Marek Locmelis, Andrew Tomkins, Kerim A. Sener

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## To cite this version:

Raphael J. Baumgartner, Marco L. Fiorentini, David Baratoux, Ludovic Ferrière, Marek Locmelis, et al.. The variability of ruthenium in chromite from chassignite and olivine-phyric shergottite meteorites: New insights into the behavior of PGE and sulfur in Martian magmatic systems. Meteoritics and Planetary Science, 2017, 52, pp.333-350. 10.1111/maps.12795. insu-03661380

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# The variability of ruthenium in chromite from chassignite and olivine-phyric shergottite meteorites: New insights into the behavior of PGE and sulfur in Martian magmatic systems

Raphael J. BAUMGARTNER<fnref n="1" />*, Marco L. FIORENTINI<fnref n="1" />, David BARATOUX<fnref n="2" />,<fnref n="3" />, Ludovic FERRIÈRE<fnref n="4" />, Marek LOCMELIS<fnref n="5" />, Andrew TOMKINS<fnref n="6" />, and Kerim A. SENER<fnref n="7" />

**Abstract**–The Martian meteorites comprise mantle-derived mafic to ultramafic rocks that formed in shallow intrusions and/or lava flows. This study reports the first in situ platinum-group element data on chromite and ulvöspinel from a series of dunitic chassignites and olivine-phyric shergottites, determined using laser-ablation ICP-MS. As recent studies have shown that Ru has strongly contrasting affinities for coexisting sulfide and spinel phases, the precise in situ analysis of this element in spinel can provide important insights into the sulfide saturation history of Martian mantle-derived melts. The new data reveal distinctive differences between the two meteorite groups. Chromite from the chassignites Northwest Africa 2737 (NWA 2737) and Chassigny contained detectable concentrations of Ru (up to ~160 ppb Ru) in solid solution, whereas chromite and ulvöspinel from the olivine-phyric shergottites Yamato-980459 (Y-980459), Tissint, and Dhofar 019 displayed Ru concentrations consistently below detection limit (<42 ppb). The relatively elevated Ru signatures of chromite from the chassignites suggest a Ru-rich (~1–4 ppb) parental melt for this meteorite group, which presumably did not experience segregation of immiscible sulfide liquids over the interval of mantle melting, melt ascent, and chromite crystallization. The relatively Ru-depleted signature of chromite and ulvöspinel from the olivine-phyric shergottites may be the consequence of relatively lower Ru contents (<1 ppb) in the parental melts, and/or the presence of sulfides during the crystallization of the spinel phases. The results of this study illustrate the significance of platinum-group element in situ analysis on spinel phases to decipher the sulfide saturation history of magmatic systems.

## INTRODUCTION

The Martian surface is dominated by mantle plume-related, mafic to ultramafic rocks (e.g., McSween et al. 2008; Baratoux et al. 2011; Grott et al. 2013). Detailed information on their chemical and petrographical

characteristics is derived from the analysis of Martian meteorites, which are categorized into three major groups, i.e., the basaltic, olivine-phyric, and lherzolitic shergottites; the clinopyroxenitic nakhlites; and the dunitic chassignites (commonly referred to as SNC meteorites). In addition, there are several ungrouped

<footnote n="1">¹School of Earth and Environment, Centre for Exploration Targeting, ARC Centre of Excellence for Core to Crust Fluid Systems, The University of Western Australia, 6009 Crawley, Australia</footnote>

<footnote n="2">²Géosciences Environnement Toulouse, CNRS, IRD and University of Toulouse, 31400 Toulouse, France</footnote>

<footnote n="3">³Institut Fondamental d'Afrique Noire Cheikh Anta Diop, 5005 Dakar, Senegal</footnote>

<footnote n="4">⁴Natural History Museum Vienna, 1010 Vienna, Austria</footnote>

<footnote n="5">⁵Department of Geosciences and Geological and Petroleum Engineering, Missouri University of Science & Technology, Rolla, Missouri 65409, USA</footnote>

<footnote n="6">⁶School of Geosciences, Monash University, 3800 Melbourne, Australia</footnote>

<footnote n="7">⁷Matrix Exploration Pty Ltd, 6112 Armadale, Australia</footnote>

<footnote>*Corresponding author. E-mail: raphael.baumgartner@research.uwa.edu.au</footnote>

<footnote>(Received 07 March 2016; revision accepted 08 October 2016)</footnote>
specimens, such as the cumulitic orthopyroxenite Allan Hills 84001 (Mittlefehldt 1994), and the polymict regolith breccias Northwest Africa 7533 and Northwest Africa 7034 (Agee et al. 2013; Humayun et al. 2013). While the shergottites presumably emplaced as (shallow) intrusions or surficial lava flows (e.g., Greshake et al. 2004; Bridges and Warren 2006; Papike et al. 2009; Balta et al. 2015), chassignites and nakhlites are likely to represent various parts of a layered intrusion or a stack of thick lava flows (Floran et al. 1978; Friedman-Lentz et al. 1999; Treiman 2005; McCubbin et al. 2013). Existing whole-rock platinum-group element (PGE: Os, Ir, Ru, Rh, Pt, and Pd) data from Martian meteorites provide unique insights into the early differentiation history of Mars. Owing to their highly siderophile and chalcophile nature (with metal-silicate and sulfide-silicate distribution coefficients in excess of $10^4$; e.g., Mann et al. 2012; Patten et al. 2013; Mungall and Brenan 2014), the overall elevated PGE concentrations in the Martian meteorites are thought to be the result of equilibrium partitioning during the early differentiation of Mars into a metallic core and a silicate mantle, and/or the potential addition of PGE via the accretion of a late chondritic veneer following the cessation of core formation (e.g., Kong et al. 1999; Jones et al. 2003; Brandon et al. 2012; Righter et al. 2015).

The study of PGE in mantle rocks and derived partial melts not only gives insights into the early accretion and differentiation history of a planet, but can provide additional information on the evolution of silicate mantle reservoirs and derived partial melts, as demonstrated for the Earth (e.g., Oshin and Crocket 1982; Rehkämper et al. 1999; Lorand et al. 1999; Burton et al. 2002; Barnes and Fiorentini et al. 2008; Baumgartner et al. 2013). In the absence of metallic iron, the PGE are essentially controlled by the stability of sulfides and platinum-group minerals, PGM (Lorand and Luguet [2016] and references therein). Furthermore, the PGE have strongly contrasting affinities for silicate and oxide phases, in particular with respect to chromite. For instance, the Pd-subgroup PGE (PPGE: Rh, Pt, and Pd) have an overall low preference to associate with chromite, whereas the Ir-subgroup PGE (IPGE: Os, Ir, and Ru) generally form discrete PGM inclusions (e.g., Augé 1985; Lee and Tredoux 1986; von Gruenewaldt and Merkle 1995; Melcher et al. 1997; Fiorentini et al. 2004; Finnigan et al. 2008; Kapsiotis et al. 2011) and/or partition as solid solution in chromite (e.g., Righter et al. 2004; Locmelis et al. 2011, 2013; Brenan et al. 2012; Pagé et al. 2012; Pagé and Barnes 2016). The study by Brandon et al. (2012) reports subtle positive correlations between MgO, $Cr_2O_3$, and respective IPGE whole-rock concentrations among a limited number of Martian meteorites. This trend suggests a broad

chromite-silicate control on PGE during melt genesis and/or differentiation. The potential influence of sulfide fractionation or accumulation may be visible in distinctive deviations in whole-rock PGE abundances. However, such interpretations may not be valid given that variations in PGE are masked/mimicked by analytical barriers and experimental issues (e.g., nugget effect; see Meisel and Horan 2016) in the whole-rock analysis of the small sample masses used in the study of Martian meteorites (Jones et al. 2003; Brandon et al. 2012).

The detailed characterization of the mineralogical hosts of PGE could provide important information on many unknown characteristics regarding the PGE abundances and sulfide genesis in the Martian meteorites. This study thus represents the first step toward quantifying the IPGE and Rh concentrations in chromite and ulvöspinel from olivine-phyric shergottites and chassignites, using laser-ablation ICP-MS (LA-ICP-MS) analysis. The set of investigated meteorite samples includes the dunitic chassignites NWA 2737 and Chassigny, as well as the incompatible element depleted olivine-phyric shergottites Y-980459, Tissint, and Dhofar 019. We show for the first time that chromite in the chassignites contains detectable Ru concentrations in solid solution, whereas PGE contents in chromite and ulvöspinel in the olivine-phyric shergottites are generally below detection limit. To evaluate the potential causes of these differences, we review the known magmatic factors that control the partitioning of PGE in chromite from magmatic systems on Earth, and test whether similar mechanisms may be responsible for the varied Ru signature in the analyzed Martian spinels. In particular, by combining our results with the known contrasting affinity of Ru for sulfide and chromite (Locmelis et al. 2013), our data provide new constraints on the sulfide saturation history and the PGE inventory of the melts parental to the investigated Martian meteorites, in particular with respect to chassignites.

## METHODOLOGY

### Sample Preparation and Basic Textural Investigations

Two polished thin sections of Tissint (NHM-Vienna N-9404) and Chassigny (NHM-Vienna N-3255) were provided by the Natural History Museum Vienna (Austria). Polished mounts prepared with fragments of the Tissint meteorite were provided by Prof. H. Chennaoui Aoudjehane from the Hassan II University in Casablanca (Morocco). The set of olivine-phyric shergottites was complemented by chips of Y-980459 (provided by Mihiko Hirai from the National Institute of Polar Research in Tokyo, Japan) and Dhofar 019
(purchased from the Hupé Collection), which were prepared as polished mounts. A polished slice of NWA 2737 purchased from the Lucy Cullen Collection was prepared as a polished thick section. The chemical, mineralogical, and petrographic properties of chromite and ulvöspinel were examined using reflected light and secondary electron microscopy. Backscattered electron (BSE) imagery was carried out using a JEOL JSM 6610-LV at the Natural History Museum in Vienna (Austria), and with a TESCAN VEGA 3 electron microscope installed at the Centre for Microscopy, Characterisation and Analysis (CMCA), at The University of Western Australia in Perth (Australia).

## Electron Probe Microanalysis

The spinel major and minor element chemistry was determined using a JEOL 8530F field-emission gun electron probe microanalyzer installed at the CMCA in Perth. The determined elements include Ti, Al, Cr, V, Mg, Fe, Ni, and Mn. The following natural and synthetic standard reference materials were used: olivine (for Mg), pyrope (for Al), rutile (for Ti), synthetic chromium oxide (for Cr), spessartine (for Mn), magnetite (for Fe and V), and metallic nickel (for Ni). The microprobe was operated at 15 kV acceleration voltage, 15 nA beam current, and peak counting times ranging between 15 and 20 s for each element. Probe for electron probe microanalysis (EPMA) from the Probe Software Inc. was used to reduce the data. The detection limits for the minor elements V and Ni were ~200 ppm and ~170 ppm, respectively. The Fe²⁺ and Fe³⁺ proportions were determined by stoichiometric calculations using the equation of Droop (1987).

## Laser Ablation ICP-MS

Time-resolved LA-ICP-MS spot and traverse analyses were performed on 52 chromite grains from the chassignites (46 in NWA 2737 and 6 in Chassigny), as well as on 40 chromite grains (8 in Y-980459, 14 in Tissint, and 18 in Dhofar 019) and 22 ulvöspinel grains (all in Dhofar 019) from the selected olivine-phyric shergottites. These analyses were carried out at the University of Québec at Chicoutimi (Canada) using a New Wave Research 193 nm YAG UV laser coupled to a Thermo X7 ICP-MS. A helium gas mixed with argon was used as the carrier medium. Chromite and ulvöspinel were analyzed for the following masses: ²⁵Mg, ³⁴S, ⁴⁴Ca, ⁴⁷Ti, ⁵¹V, ⁵³Cr, ⁵⁵Mn, ⁵⁷Fe, ⁵⁹Co, ⁶⁰Ni, ⁶³Cu, ⁶⁶Zn, ⁶⁹Ga, ⁷¹Ga, ⁷⁴Ge, ⁸⁸Sr, ⁹⁵Mo, ¹⁰¹Ru, ¹⁰²Ru, ¹⁰³Rh, ¹⁰⁴Ru, ¹⁰⁵Pd, ¹¹⁵In, ¹¹⁸Sn, ¹²⁵Te, ¹²⁸Te, ¹⁸⁵Re, ¹⁸⁹Os, ¹⁹⁰Os, ¹⁹¹Ir, ¹⁹²Os, and ¹⁹³Ir. The analyses were conducted using 5 mJ pulse⁻¹ laser power, 15 Hz

repetition rate, and 5 µm s⁻¹ scan speed at traverse analyses. A beam diameter of 73 µm was chosen. The gas blank was collected for ~10–30 s prior and after laser ablation. The data reduction was done using the Iolite software (Paton et al. 2011).

⁵⁷Fe was chosen as the internal reference element. In the absence of any matrix-matched reference material for spinel, the certified reference material Laflamme Po-727 (synthetic FeS doped with ~40 ppm of each PGE; Wilson et al. 2002) was used for quantification of selected PGE (i.e., ¹⁰¹Ru, ¹⁰³Rh, ¹⁸⁹Os, and ¹⁹³Ir). The GSE-1G standard (synthetic basaltic glass standard with ~400 ppm Cr and ~450 ppm Ti; Guillong et al. 2005) was used for the quantification of ⁵³Cr and ⁴⁷Ti. The analysis of PGE was monitored by repeatedly analyzing the inhouse reference material JB-MSS-5 (synthetic FeS doped with ~20–65 ppm of each PGE; e.g., Dare et al. 2010; Patten et al. 2013) during each analytical run. The detection limits for the respective PGE in each individual quantification were calculated from the gas blank and the apparent sensitivity of the reference material following the equation given in Longerich et al. (1996), based on which we determined average detection limits for the entire analytical session (i.e., 40 ppb for Os, 15 ppb for Ir, 42 ppb for Ru, and 7 ppb for Rh). Ruthenium was the only PGE above detection limit in spinel from the analyzed meteorites, as discussed below. Potential interferences between ⁶¹Ni⁴⁰Ar species and ¹⁰¹Ru were monitored through the analysis of a NiFeS₂ free from Ru. However, corrections for ¹⁰¹Ru were insignificant due to the low Ni contents detected in chromite and ulvöspinel (usually below the detection limit of ~170 ppm in microprobe analysis), as well as the low natural abundance of ⁶¹Ni of which only low fractions will form ⁶¹Ni⁴⁰Ar species. Determinations on the effect of ⁶³Cu⁴⁰Ar on ¹⁰³Rh were waived given the overall lack of detectable Rh in the analyzed spinel crystals. Following interference corrections, the repeated analyses for Ru in the JB-MSS-5 reference material reproduce the certified Ru concentrations with a relative difference of better than ~5%.

The quantification of a suitable internal standard element for LA-ICP-MS analysis is impeded by the chemical zonation of chromite from the chassignites, as well as the intimate chromite-ulvöspinel associations in the shergottites. While spinel in the latter specimens generally lack detectable PGE concentrations (see below), iron is chosen for chromite in the chassignites by taking advantage of a distinctive negative correlation between Cr and Fe contents in EPMA (coefficients of determination, R², >0.9). Data processing in Iolite provides elemental ratios between unknown analyte (e.g., ⁵³Cr) and the selected internal standard element (⁵⁷Fe), which in the case of the examined spinels can be
used to constrain the bulk Fe content at laser spots. This approach was tested by comparing calculated and measured (by EPMA) Fe contents of several monomineralic and rather unzoned chromite and ulvöspinel grains. The internally derived Fe contents are uniformly offset, which likely resulted from the uncertain and/or heterogeneously distributed Cr concentrations in the GSE-1G standard (Guillong et al. 2005), or a matrix effect in the ablation process. However, after correcting the ⁵³Cr/⁵⁷Fe ratios in Iolite by a uniform iteratively constrained correction factor, the calculated Fe contents at respective laser spots reproduce the EPMA data within ~5% difference.

Through time-resolved LA-ICP-MS analysis it is possible to distinguish whether a certain element occurs in a mineral inclusion or as solid solution. Flat spectra are commonly attributed to PGE occurring in solid solution in the host crystal lattice, whereas the presence of distinctive peaks implies the presence of discrete PGE-rich inclusions in the form of PGM and/or sulfides. Understanding whether PGE occur as solid solution or as inclusions is pivotal in the interpretation of the LA-ICP-MS data and to model the PGE history of a magmatic system (Locmelis et al. 2011, 2013; Pagé et al. 2012; Park et al. 2012; Pagé and Barnes 2016). The identification of PGM inclusions is commonly limited by the spatial resolution of the analysis. The background signal and the presence of spurious particles, the latter of which can come from the transport tubing or the ablation cell, may aggravate the recognition of PGM. In traverse analysis, we determined their presence and size on the basis of the defined scan speed and ablation repetition rate (i.e., 5 µm s⁻¹ and 15 Hz, respectively), which resulted in laser pulse counts every 0.33 µm along the ablation lines. In order to avoid false identification, we require PGM to show three consecutive laser pulse counts two or more orders of magnitude higher than the count rates in the gas blank, which allows us to identify inclusions larger than 1 µm. The mineralogical association of PGE-rich inclusions was qualitatively assessed by monitoring elements that are incompatible in spinel, such as ³⁴S and ²⁹Si as proxies for sulfides and silicates.

## PETROGRAPHY AND PETROGENESIS OF CHROMITE AND ULVÖSPINEL

The petrography of spinel in the analyzed olivine-phyric shergottites and chassignites was documented and interpreted in previous studies, such as Beck et al. (2006) and Treiman et al. (2007) for NWA 2737, Floran et al. (1978) for Chassigny, Greshake et al. (2004) for Y-980459, Gattacceca et al. (2013) and Balta et al. (2015)

for Tissint, as well as Taylor et al. (2002) and Goodrich et al. (2003) for Dhofar 019. We complement these studies by documenting the mineralogy of chromite and ulvöspinel further analyzed with high EPMA and LA-ICP-MS analysis. In the chassignites, chromite occurs within olivine (both as individual and inclusion-hosted crystals), but most commonly occurs as euhedral to subhedral crystals along olivine crystal boundaries (Fig. 1a) in conjunction with variable amounts of mesostasis silicates, oxides, and sulfides (pyrrhotite ± pentlandite). Chromite is more abundant in NWA 2737 (~2 vol%) than in Chassigny (~1 vol%). The chromite grains interstitial to olivine are up to ~300 µm in size in NWA 2737, and ~150 µm in Chassigny. In both samples, chromite included within olivine was too small (~50 µm in diameter) for LA-ICP-MS analysis.

In the olivine-phyric shergottite Y-980459, spinel occurs as euhedral to anhedral chromite associated with pyroxene and olivine (mega-) crystals in an aphyric groundmass (Fig. 1b). The groundmass contains dendritic to chain-like olivine and pyroxene, skeletal iron oxides, and dispersed bleby sulfides composed of pyrrhotite and minor pentlandite exsolutions. Tissint and Dhofar 019 lack the glassy interstitial material and spinel in the form of chromite and ulvöspinel occurs in a coarse matrix composed of olivine, phyric pyroxene, maskelynite, and minor amounts of sulfides (pyrrhotite ± pentlandite and Cu-sulfides), as shown in Fig. 1c. In both samples, chromite can be present as monomineralic grains, but mostly occurs in association with ulvöspinel. In Tissint, ulvöspinel solely occurs as rims around chromite cores (Fig. 1c), whereas in Dhofar 019, ulvöspinel is much more abundant and either forms rims around chromite, or exsolutions in chromite (Fig. 1d). The grain sizes vary between the analyzed specimens. Y-980459 contains chromite usually smaller than 50–60 µm, with rare grains up to ~100 µm in diameter. Consistently, only a limited number of chromite grains could have been analyzed with LA-ICP-MS at ~73 µm laser diameter. Tissint and Dhofar 019 frequently contain larger spinel grains of up to 150 µm in diameter. In all samples, pyroxene and olivine can be host to euhedral to subhedral chromite, but the small grain sizes (<50 µm) did not allow for LA-ICP-MS analysis.

Chromite and ulvöspinel in the chassignites and olivine-phyric shergottites is interpreted to be of magmatic origin. In the chassignites, chromite and olivine presumably accumulated from a variably evolved, incompatible element (e.g., light rare earth elements, LREE) enriched basaltic precursor melt (e.g., Filiberto 2008; He et al. 2013). Magma emplacement may have happened as a shallow slow-cooling magma body or as a stack of thick lava flows (Monkawa et al. 2004; Mikouchi et al. 2005; Beck et al. 2006; McCubbin
<figure><img src="image_3.png" /><figcaption>Fig. 1. Backscattered electron (BSE) images showing chromite and ulvöspinel in the examined chassignites and olivine-phyric shergottites. a) Cluster of large chromites in the chassignite NWA 2737, which occur interstitial to olivine and associated with various amounts of mesostasis minerals, such as clinopyroxenes and sulfides. b) Euhedral chromites in the olivine-phyric shergottite Y-980459 associated with the aphyric groundmass. Note the small chromite inclusion in pyroxene at the bottom of the picture. c, d) Spinel in the olivine-phyric shergottites Tissint and Dhofar 019. Chromite and ulvöspinel can occur as single crystals, but mostly form composite grains with sharp contacts (a) or gradual intergrowths (d) between chromite-rich cores and ulvöspinel rims. Chr = chromite; Cpx = clinopyroxene; Ol = olivine; Pl = plagioclase; Po = pyrrhotite; Px = unclassified pyroxene; Spl = spinel; Usp = ulvöspinel.</figcaption></figure>

et al. 2013). Previous studies reconciled the varied spinel mineralogy in the investigated olivine-phyric shergottites with a magmatic origin, in which chromite represents an early population of spinel that crystallized at high temperature from an incompatible element depleted melt. In Y-980459, the association of large chromite and olivine crystals (Fig. 1b) embedded in the aphyric groundmass suggests two-stage magmatic cooling. Some chromites may have crystallized with olivine in deep reservoirs, migrated with the magma to the surface, and may have been set in the quenched groundmass during

rapid cooling in a (thin) lava flow (e.g., Greshake et al. 2004). In Tissint and Dhofar 019, ulvöspinel is thought to have been formed relatively late during slow cooling, presumably in a shallow magmatic intrusion (e.g., Goodrich et al. 2003; Balta et al. 2015).

## MAJOR AND MINOR ELEMENT CHEMISTRY OF SPINEL

The major and minor element concentrations of chromite and ulvöspinel analyzed in this study are
<table><caption>Table 1. Composition of spinel in the chassignites NWA 2737 and Chassigny, and the olivine-phyric shergottites Y-980459, Tissint, and Dhofar 019.</caption><thead><tr><th rowspan="3"></th><th colspan="4">Chr—NWA 2737</th><th colspan="4">Chr—Chassigny</th><th colspan="4">Chr—Y-980459</th></tr><tr><th colspan="4">n = 39</th><th colspan="4">n = 41</th><th colspan="4">n = 27</th></tr><tr><th>Min.</th><th>Max.</th><th>Avg.</th><th>SD</th><th>Min.</th><th>Max.</th><th>Avg.</th><th>SD</th><th>Min.</th><th>Max.</th><th>Avg.</th><th>SD</th></tr></thead><tbody><tr><td>TiO<sub>2</sub><br>(wt%)</td><td>1.09</td><td>2.00</td><td>1.49</td><td>0.22</td><td>1.64</td><td>3.99</td><td>2.40</td><td>0.52</td><td>0.42</td><td>1.10</td><td>0.57</td><td>0.20</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>6.19</td><td>10.93</td><td>8.07</td><td>1.20</td><td>6.65</td><td>12.30</td><td>9.86</td><td>1.30</td><td>5.60</td><td>8.52</td><td>6.35</td><td>0.82</td></tr><tr><td>Cr<sub>2</sub>O<sub>3</sub></td><td>49.33</td><td>57.18</td><td>53.64</td><td>2.23</td><td>43.01</td><td>54.77</td><td>48.43</td><td>2.93</td><td>54.59</td><td>61.78</td><td>60.11</td><td>1.82</td></tr><tr><td>V<sub>2</sub>O<sub>3</sub></td><td>0.16</td><td>0.25</td><td>0.20</td><td>0.02</td><td>0.28</td><td>0.40</td><td>0.35</td><td>0.03</td><td>0.38</td><td>0.71</td><td>0.49</td><td>0.09</td></tr><tr><td>MgO</td><td>4.47</td><td>6.52</td><td>5.48</td><td>0.45</td><td>2.99</td><td>4.11</td><td>3.44</td><td>0.21</td><td>6.30</td><td>7.46</td><td>6.80</td><td>0.26</td></tr><tr><td>FeO</td><td>27.49</td><td>31.71</td><td>29.27</td><td>1.15</td><td>31.37</td><td>36.52</td><td>33.73</td><td>1.23</td><td>22.26</td><td>24.50</td><td>23.09</td><td>0.53</td></tr><tr><td>MnO</td><td>0.39</td><td>0.53</td><td>0.46</td><td>0.03</td><td>0.41</td><td>0.58</td><td>0.48</td><td>0.04</td><td>0.31</td><td>0.49</td><td>0.38</td><td>0.04</td></tr><tr><td>NiO</td><td>&lt;0.02</td><td>0.05</td><td>0.01</td><td>0.02</td><td>&lt;0.02</td><td>0.09</td><td>0.01</td><td>0.02</td><td>&lt;0.02</td><td>0.06</td><td>0.02</td><td>0.02</td></tr><tr><td>Fe#</td><td>0.68</td><td>0.77</td><td>0.72</td><td>0.02</td><td>0.80</td><td>0.85</td><td>0.83</td><td>0.01</td><td>0.62</td><td>0.66</td><td>0.65</td><td>0.01</td></tr><tr><td>Cr#</td><td>0.75</td><td>0.86</td><td>0.82</td><td>0.03</td><td>0.71</td><td>0.85</td><td>0.77</td><td>0.03</td><td>0.82</td><td>0.88</td><td>0.86</td><td>0.02</td></tr><tr><td>100<br>× Fe<sup>3</sup>/R<sup>3+</sup></td><td>5.14</td><td>9.47</td><td>6.54</td><td>0.88</td><td>5.39</td><td>10.19</td><td>7.42</td><td>1.20</td><td>0.24</td><td>5.60</td><td>1.45</td><td>1.07</td></tr><tr><td>100<br>× V/(Cr+Al)</td><td>0.26</td><td>0.43</td><td>0.33</td><td>0.04</td><td>0.47</td><td>0.77</td><td>0.62</td><td>0.08</td><td>0.57</td><td>1.10</td><td>0.76</td><td>0.15</td></tr><tr><td>fO<sub>2</sub> (ΔQFM)</td><td>0.7</td><td>0.0</td><td></td><td></td><td>-0.1</td><td>-1.3</td><td></td><td></td><td>-0.4</td><td>-3.1</td><td></td><td></td></tr></tbody></table>

Oxygen fugacity ($fO_2$) is calculated using the qualitative V valence-state oxybarometer by Papike et al. (2004). The parenthesized 100 × V/(Cr+Al) values and related $fO_2$ estimates for chromite in Dhofar 019 exclude the analyses that were potentially affected by ulvöspinel co-analysis (see text). avg. = average; Chr = chromite; Usp = ulvöspinel; min. = minimum; max. = maximum; n = number of analyses; SD = standard deviation; Cr# = Cr/(Cr+Al); Fe# = $Fe^{2+}$/(Fe$^{2+}$+Mg); 100 × $Fe^{3+}$/R$^{3+}$ = 100 × $Fe^{3+}$/(Fe$^{3+}$+Cr$^{3+}$+Al$^{3+}$); ΔQFM = $fO_2$ given as differences in log units relative to the quartz-fayalite-magnetite buffer (QFM).

summarized in Table 1, and illustrated in Fig. 2. The new data fall well into the range of those reported in the literature; see Treiman et al. (2007) for NWA 2737, Floran et al. (1978) for chassigny, Greshake et al. (2004) for Y-980459, Gattacceca et al. (2013) for Tissint, and Goodrich et al. (2003) for Dhofar 019. The chromites from NWA 2737 have Cr# values (with Cr# = Cr/[Cr+Al]) between 0.75 and 0.86, Fe# values (with Fe# = $Fe^{2+}$/[Fe$^{2+}$+Mg]) between 0.68 and 0.77, and 100 × $Fe^{3+}$/R$^{3+}$ values (with R$^{3+}$ = $Fe^{3+}$ + $Cr^{3+}$+Al$^{3+}$) that fall in the range of 5.14 and 9.47. Chromite in Chassigny shows broadly similar Cr# values (0.71–0.85; Fig. 2a), distinctly higher Fe# values (0.80–0.85; Fig. 2a), and slightly higher 100 × $Fe^{3+}$/R$^{3+}$ values (5.39–10.19; Fig. 2c). In both samples, the TiO$_2$ contents positively correlate with the Fe# values (Fig. 2b), and are relatively lower in NWA 2737 (1.09–2.00 wt%) than in Chassigny (1.64–3.99 wt%).

The chromites from Y-980459 have Cr# values of 0.82–0.88, Fe# values of 0.62–0.66, 100 × $Fe^{3+}$/R$^{3+}$ values of 0.24–5.60, and 0.42–1.10 wt% TiO$_2$. The chromites from Tissint have broadly similar Cr# values (0.81–0.85), but distinctly higher Fe# values (0.64–0.88) than Y-980459. The 100 × $Fe^{3+}$/R$^{3+}$ values (1.22–3.37) and TiO$_2$ contents (0.62–1.45 wt%) are elevated. Chromite contained in Dhofar 019 shows the lowest

Cr# values (0.65–0.85; Fig. 2a) and highest Fe# values (0.73–0.91; Fig. 2a). Similarly, the measured TiO$_2$ contents (0.84–9.08 wt%) and calculated 100 × $Fe^{3+}$/R$^{3+}$ values (4.63–17.68) are the highest among the analyzed shergottite specimens (Figs. 2b and 2c, respectively). However, it is noted that some of the analyses with extremely high 100 × $Fe^{3+}$/R$^{3+}$ values and TiO$_2$ concentrations in Dhofar 019 (Fig. 2b) may correspond to (sub-) micron scale chromite and ulvöspinel intergrowths not resolvable with BSE imagery, rather than extreme $Fe^{3+}$ and Ti enrichments along the Fe-Ti trend previously reported in Goodrich et al. (2003).

In Tissint and Dhofar 019, ulvöspinel crystals display a broad continuum of chemical compositions, which are reflected by decreasing Cr# values along with increasing Fe# values (Fig. 2a), TiO$_2$ contents (Fig. 2b), and 100 × $Fe^{3+}$/R$^{3+}$ values (Fig. 2c). In Tissint, ulvöspinel shows 48.92–56.09 wt% FeO and 18.17–30.45 wt% TiO$_2$. In Dhofar 019, ulvöspinel shows similar FeO (45.69–54.97 wt%), but lower TiO$_2$ contents (11.73–17.93 wt%). In ulvöspinel from all samples, the 100 × $Fe^{3+}$/R$^{3+}$ values fall between 8.32 and 38.04.

The redox conditions during crystallization of chassignites and olivine-phyric shergottites were determined and intensively discussed in previous studies
<table><thead><tr><th colspan="4">Chr—Tissint<br>n = 28</th><th colspan="4">Usp—Tissint<br>n = 25</th><th colspan="4">Chr—Dhofar 019<br>n = 19</th><th colspan="4">Usp—Dhofar 019<br>n = 35</th></tr><tr><th>Min.</th><th>Max.</th><th>Avg.</th><th>SD</th><th>Min.</th><th>Max.</th><th>Avg.</th><th>SD</th><th>Min.</th><th>Max.</th><th>Avg.</th><th>SD</th><th>Min.</th><th>Max.</th><th>Avg.</th><th>SD</th></tr></thead><tbody><tr><td>0.62</td><td>1.45</td><td>0.80</td><td>0.21</td><td>18.17</td><td>30.45</td><td>22.33</td><td>2.43</td><td>0.84</td><td>9.08</td><td>3.32</td><td>2.56</td><td>11.73</td><td>17.93</td><td>15.18</td><td>1.60</td></tr><tr><td>6.54</td><td>8.42</td><td>7.28</td><td>0.38</td><td>2.49</td><td>3.86</td><td>3.30</td><td>0.37</td><td>5.65</td><td>13.19</td><td>7.32</td><td>1.92</td><td>3.07</td><td>13.54</td><td>4.65</td><td>2.26</td></tr><tr><td>53.83</td><td>59.56</td><td>57.16</td><td>1.49</td><td>6.57</td><td>22.54</td><td>15.98</td><td>3.77</td><td>35.69</td><td>56.23</td><td>47.85</td><td>7.29</td><td>18.23</td><td>29.32</td><td>23.92</td><td>3.33</td></tr><tr><td>0.31</td><td>0.79</td><td>0.48</td><td>0.10</td><td>0.36</td><td>0.62</td><td>0.46</td><td>0.07</td><td>0.43</td><td>0.83</td><td>0.66</td><td>0.11</td><td>0.46</td><td>0.95</td><td>0.63</td><td>0.12</td></tr><tr><td>2.24</td><td>7.17</td><td>4.86</td><td>1.46</td><td>1.14</td><td>2.29</td><td>1.78</td><td>0.27</td><td>2.47</td><td>5.32</td><td>3.80</td><td>0.93</td><td>2.20</td><td>4.48</td><td>2.74</td><td>0.50</td></tr><tr><td>23.87</td><td>31.79</td><td>27.36</td><td>2.40</td><td>48.92</td><td>56.09</td><td>52.04</td><td>2.02</td><td>28.44</td><td>42.95</td><td>34.67</td><td>4.50</td><td>45.69</td><td>54.97</td><td>50.33</td><td>2.53</td></tr><tr><td>0.25</td><td>0.50</td><td>0.38</td><td>0.06</td><td>0.40</td><td>0.68</td><td>0.59</td><td>0.60</td><td>0.33</td><td>0.53</td><td>0.44</td><td>0.07</td><td>0.49</td><td>0.70</td><td>0.60</td><td>0.05</td></tr><tr><td>&lt;0.02</td><td>0.05</td><td>0.02</td><td>0.02</td><td>&lt;0.02</td><td>0.12</td><td>0.03</td><td>0.03</td><td>&lt;0.02</td><td>0.08</td><td>0.02</td><td>0.02</td><td>&lt;0.02</td><td>0.13</td><td>0.05</td><td>0.03</td></tr><tr><td>0.64</td><td>0.88</td><td>0.75</td><td>0.07</td><td>0.92</td><td>0.96</td><td>0.94</td><td>0.01</td><td>0.73</td><td>0.91</td><td>0.82</td><td>0.08</td><td>0.73</td><td>0.92</td><td>0.87</td><td>0.05</td></tr><tr><td>0.81</td><td>0.85</td><td>0.84</td><td>0.01</td><td>0.64</td><td>0.82</td><td>0.76</td><td>0.04</td><td>0.65</td><td>0.85</td><td>0.81</td><td>0.05</td><td>0.65</td><td>0.86</td><td>0.80</td><td>0.04</td></tr><tr><td>1.22</td><td>3.37</td><td>2.65</td><td>0.62</td><td>8.32</td><td>33.09</td><td>18.91</td><td>6.67</td><td>4.63</td><td>17.68</td><td>9.98</td><td>4.57</td><td>20.29</td><td>38.04</td><td>26.90</td><td>5.19</td></tr><tr><td>0.62</td><td>1.31</td><td>0.76</td><td>0.17</td><td>1.55</td><td>3.88</td><td>2.47</td><td>0.60</td><td>0.78</td><td>1.75<br>(1.16)<br>-1.4</td><td>1.11</td><td>0.27</td><td>1.53</td><td>3.93</td><td>2.30</td><td>0.63</td></tr><tr><td>-0.7</td><td>-4.2</td><td></td><td></td><td></td><td></td><td></td><td></td><td>-4.7<br>(-3.4)</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr></tbody></table>

<figure><img src="image_4.png" /><figcaption>Fig. 2. Chemical composition of chromite and ulvöspinel in the chassignites (NWA 2737 and Chassigny) and olivine-phyric shergottites (Y-980459, Tissint and Dhofar 019) in (a) the plot Fe# (Fe²⁺/[Fe²⁺+Mg]) versus Cr# (Cr/[Cr+Al]), (b) the plot Cr# versus TiO₂, and (c) the ternary diagram Cr-Al-Fe³⁺. (Color figure can be viewed at wileyonlinelibrary.com.)</figcaption></figure>

(e.g., Herd et al. 2002; Goodrich et al. 2003; McKay et al. 2004; Herd 2006; Beck et al. 2006; Treiman et al. 2007; Balta et al. 2015; among others). Qualitative estimates are derived using the V valence-state oxybarometer proposed by Canil (2002), and revisited by Papike et al. (2004) and Herd (2006), in which the linear relationship between fO₂ and chromite 100 × V/(Cr+Al) atomic values determined by EPMA readily allows us to approximate oxygen fugacity during crystallization (hereafter given by differences in log

units of oxygen fugacity, fO₂, relative to the quartz-fayalite-magnetite buffer, QFM). Our chromite data (Table 1) suggest more oxidizing conditions in the crystallization of NWA 2737 (fO₂ ~QFM to ~QFM+0.7) relative to Chassigny (fO₂ ~QFM-1.3 to ~QFM-0.1). By way of comparison, Beck et al. (2006) calculated with the same method comparable values (fO₂ = QFM+0.6 and QFM-1.3 for NWA 2737 and Chassigny, respectively). Treiman et al. (2007) utilized the olivine-pyroxene-spinel oxygen barometer by Wood
(1991) and found similar $fO_2$ in Chassigny ($fO_2$ ~QFM~1.3), but slightly lower $fO_2$ in NWA 2737 ($fO_2$ ~QFM~0.5). The mineral data for chromite in the three analyzed olivine-phyric shergottites give a broad continuum of 100 × V/(Cr+Al) values (Table 1). Our qualitative results ($fO_2$ ~QFM~4.0 to ~QFM~0.5; Table 1) broadly align with the values found in previous studies ($fO_2$ ~QFM~4 to ~QFM~2; see McKay et al. 2004; for Y-980459, Balta et al. 2015; for Tissint and Goodrich et al. 2003, for Dhofar 019), with redox conditions that are relatively lower than in the chassignites.

## DISTRIBUTION AND CONCENTRATION OF OS, IR, RU, AND RH IN SPINEL

Representative spectra of time-resolved LA-ICP-MS traverse and spot analyses on chromite in the chassignites, as well as on chromite and ulvöspinel in the olivine-phyric shergottites, are illustrated in Fig. 3. The PGE-bearing inclusions usually consist of randomly distributed PGM composed of one or more IPGE ± Rh. Signals of Pt and Pd were usually absent. While no preferred occurrence and elemental association of PGM was observed, their overall composition could suggest that Os-Ir-Ru alloys of variable composition are the main PGM species in the analyzed meteorites. Rare polymineralic assemblages of silicates and PGE-bearing sulfides enriched in Ru, Rh ± Pt, and Pd were observed in ulvöspinel from the olivine-phyric shergottites only. Aside from the PGE-bearing inclusions, chromite and ulvöspinel from the chassignites and olivine-phyric shergottites showed Os, Ir, and Rh signals usually not distinguishable from the background (Figs. 3c and 3d). Notable exceptions were the slightly increased count rates of Os and Ir in some chromites from the chassignite NWA 2737 (Fig. 3a), but their origin remains uncertain given the insufficiently low signal to background ratios. The majority of chromites from the chassignites, as opposed to chromite and ulvöspinel in the olivine-phyric shergottites, showed Ru counts distinctly above the background levels (Figs. 3a and 3b). These signals are interpreted to reflect the presence of Ru in solid solution.

Quantitative analysis for Ru, Rh, Os, Ir, Ti, and Cr (Table 2) was restricted to intervals free of apparent silicate and PGE-bearing inclusions (Figs. 3a-c), and/or randomly distributed PGE spikes (Fig. 3d) that may come from PGM microcrysts or contamination during analysis. The quantifications for Rh, Os, and Ir are generally below the limit of detection (Rh = 7 ppb, Os = 40 ppb, Ir = 10 ppb; average values) for all spinel

analyses in the chassignites and olivine-phyric shergottites. The same is true for Ru analysis (average detection limit of 42 ppb) on chromite and ulvöspinel from the olivine-phyric shergottites. In contrast, chromite from the chassignites (42 out of 46 in NWA 2737, and 4 out of 6 in Chassigny) yielded concentrations ranging between the respective limits of detection and up to ~160 ppb (Table 2). The majority of analyzed chromites contained ~60–120 ppb Ru (Fig. 4), with apparent correlations with Cr and Ti (Figs. 5a and 5b, respectively).

## DISCUSSION AND IMPLICATIONS

### Constraints on the Partitioning of Ru in Chromite

The partitioning of the IPGE Os, Ir, and Ru into magmatic chromite (expressed by the distribution coefficient $D_{IPGE}^{chromite-melt}$) is controlled by the magma type and mode of emplacement (Pagé et al. 2012; Pagé and Barnes 2016), the presence of immiscible sulfides (Fiorentini et al. 2008; Locmelis et al. 2011, 2013), and the local redox conditions during crystallization (Capobianco and Drake 1990; Capobianco et al. 1994; Righter et al. 2004; Brenan et al. 2012; Pagé et al. 2012; Park et al. 2012). Under the redox conditions typical for terrestrial magmatic systems ($fO_2$ ~QFM~2 to ~QFM+2), the partitioning of Ru in chromite is generally found to positively correlate with oxygen fugacity. This is explained by a combined effect that arises due to the stability of different ionic species with increasing oxygen fugacity (i.e., $Ru^{2+} \rightarrow Ru^{3+}$), and their contrasting compatibility in octahedral sites in the spinel structure (Brenan et al. 2012). Oxidizing conditions also result in increased proportions of $Fe^{3+}$ in crystallizing chromite; exerting a positive effect on the partitioning of Ru that preferentially substitutes for $Fe^{3+}$ (Brenan et al. 2012).

The experimental study by Brenan et al. (2012) calculated $D_{Ru}^{chromite-melt}$ <100 for $Fe^{3+}$-poor spinel crystallizing at $fO_2$ <QFM+2. Following the unnaturally high-oxygen fugacity experimental studies by Righter et al. (2004), the partitioning of Ru at $fO_2$ >QFM+2 could be one or more orders of magnitude higher. The distribution coefficients calculated based on whole-rock analysis and in situ LA-ICP-MS analysis on chromite from natural samples further underpins this view. Chromite from relatively reduced ($fO_2$ <QFM+1) magmatic rocks on Earth, such as komatiites, komatiitic basalts, ferropicrites, picrites, boninites, and tholeiites, can contain >100 ppb Ru (Locmelis et al. 2011; Brenan et al. 2012; Pagé et al. 2012; Pagé and Barnes 2016), with the maximum of 600 ppb Ru found in komatiite-hosted chromites from the Central Lapland greenstone
<figure><img src="image_5.png" /><figcaption>Fig. 3. Time-resolved laser-ablation ICP-MS analyses on selected spinels in the Martian meteorites (see the Methodology section for analytical parameter). a, b) Traverse analyses along chromite crystals in the chassignites NWA 2737 and Chassigny. c, d) Spot analyses on chromite and ulvöspinel in the olivine-phyric shergottites Tissint and Dhofar 019. Domains free of conspicuous silicate and discrete PGE-rich phases (gray fields) were integrated for PGE quantification. The Os peak in (b) can be a submicron scale PGM or spurious particles released from the transport tubing or the ablation cell during analysis. Signal intensities (y-axis) below 10² are zero. (Color figure can be viewed at wileyonlinelibrary.com.)</figcaption></figure>

belt (Locmelis et al. 2013). The calculated distribution coefficients in this case fall in the ~50–150 range. Lastly, Fe³⁺-rich chromite from the oxidized (fO₂ ~QFM+2.5) basalts at the Ambae volcano, Vanuatu, was found to contain up to ~100 ppb Ru, with a calculated Dₘₚᵢₘ-ešt of ~2450 (Park et al. 2012). Because our study is the first that systematically measured Ru

concentrations in magmatic ulvöspinel, and given the overall absence of detectable Ru in this spinel species, an estimation of the partitioning of Ru is not straightforward. However, the overall Fe³⁺-rich nature of this spinel species in the shergottites (Table. 1), as well as the preferred substitution of Ru³⁺ for Fe³⁺, suggests that the partitioning of Ru is likely to be
<table><caption>Table 2. Results of LA-ICP-MS analysis on chromite from the chassignites NWA 2737 and Chassigny.</caption><thead><tr><th>Analysis</th><th>Fe (wt%)</th><th>Cr (wt%)</th><th>1 SE</th><th>Ti (wt%)</th><th>1 SE</th><th>Ru (ppb)</th><th>1 SE</th><th>d.l. (ppb)</th></tr></thead><tbody><tr><td colspan="9">NWA 2737</td></tr><tr><td>N1</td><td>21.5</td><td>35.6</td><td>0.1</td><td>0.65</td><td>0.04</td><td>103</td><td>12</td><td>22</td></tr><tr><td>N2</td><td>21.5</td><td>35.8</td><td>0.1</td><td>0.69</td><td>0.04</td><td>105</td><td>11</td><td>45</td></tr><tr><td>N3</td><td>21.4</td><td>35.9</td><td>0.1</td><td>0.68</td><td>0.04</td><td>105</td><td>11</td><td>30</td></tr><tr><td>N4</td><td>21.9</td><td>35.1</td><td>0.1</td><td>0.83</td><td>0.05</td><td>&lt;d.l.</td><td>—</td><td>93</td></tr><tr><td>N5</td><td>21.3</td><td>36.1</td><td>0.1</td><td>0.66</td><td>0.03</td><td>94</td><td>12</td><td>45</td></tr><tr><td>N6</td><td>21.7</td><td>35.4</td><td>0.1</td><td>0.76</td><td>0.05</td><td>61</td><td>12</td><td>53</td></tr><tr><td>N7</td><td>21.5</td><td>35.8</td><td>0.1</td><td>0.71</td><td>0.07</td><td>120</td><td>19</td><td>79</td></tr><tr><td>N8</td><td>21.2</td><td>36.2</td><td>0.1</td><td>0.73</td><td>0.05</td><td>115</td><td>14</td><td>79</td></tr><tr><td>N9</td><td>21.4</td><td>35.8</td><td>0.1</td><td>0.63</td><td>0.03</td><td>88</td><td>9</td><td>12</td></tr><tr><td>N10</td><td>21.9</td><td>35.1</td><td>0.2</td><td>0.77</td><td>0.08</td><td>85</td><td>15</td><td>14</td></tr><tr><td>N11</td><td>21.7</td><td>35.4</td><td>0.2</td><td>0.72</td><td>0.06</td><td>98</td><td>15</td><td>36</td></tr><tr><td>N12</td><td>21.3</td><td>36.1</td><td>0.1</td><td>0.69</td><td>0.04</td><td>89</td><td>9</td><td>42</td></tr><tr><td>N13</td><td>21.2</td><td>36.3</td><td>0.1</td><td>0.61</td><td>0.03</td><td>102</td><td>12</td><td>66</td></tr><tr><td>N14</td><td>21.3</td><td>36.0</td><td>0.1</td><td>0.61</td><td>0.03</td><td>97</td><td>9</td><td>12</td></tr><tr><td>N15</td><td>22.2</td><td>34.6</td><td>0.2</td><td>0.98</td><td>0.05</td><td>&lt;d.l.</td><td>—</td><td>97</td></tr><tr><td>N16</td><td>21.0</td><td>36.6</td><td>0.1</td><td>0.63</td><td>0.05</td><td>92</td><td>9</td><td>13</td></tr><tr><td>N17</td><td>20.9</td><td>36.9</td><td>0.2</td><td>0.76</td><td>0.10</td><td>77</td><td>15</td><td>10</td></tr><tr><td>N18</td><td>20.7</td><td>37.6</td><td>0.1</td><td>0.69</td><td>0.07</td><td>128</td><td>18</td><td>9</td></tr><tr><td>N19</td><td>21.2</td><td>36.2</td><td>0.1</td><td>0.66</td><td>0.06</td><td>104</td><td>13</td><td>15</td></tr><tr><td>N20</td><td>21.4</td><td>35.9</td><td>0.2</td><td>0.69</td><td>0.07</td><td>96</td><td>24</td><td>18</td></tr><tr><td>N21</td><td>21.2</td><td>36.2</td><td>0.1</td><td>0.68</td><td>0.05</td><td>110</td><td>13</td><td>16</td></tr><tr><td>N22</td><td>21.1</td><td>36.4</td><td>0.1</td><td>0.64</td><td>0.04</td><td>159</td><td>18</td><td>5</td></tr><tr><td>N23</td><td>21.7</td><td>35.3</td><td>0.1</td><td>0.67</td><td>0.04</td><td>72</td><td>12</td><td>12</td></tr><tr><td>N24</td><td>21.4</td><td>35.9</td><td>0.1</td><td>0.70</td><td>0.04</td><td>97</td><td>10</td><td>11</td></tr><tr><td>N25</td><td>21.1</td><td>36.5</td><td>0.1</td><td>0.65</td><td>0.03</td><td>114</td><td>14</td><td>28</td></tr><tr><td>N26</td><td>21.2</td><td>36.2</td><td>0.1</td><td>0.67</td><td>0.04</td><td>102</td><td>9</td><td>16</td></tr><tr><td>N27</td><td>21.2</td><td>36.2</td><td>0.1</td><td>0.67</td><td>0.04</td><td>100</td><td>13</td><td>22</td></tr><tr><td>N28</td><td>22.2</td><td>34.6</td><td>0.1</td><td>0.95</td><td>0.05</td><td>&lt;d.l.</td><td>—</td><td>38</td></tr><tr><td>N29</td><td>22.1</td><td>34.8</td><td>0.1</td><td>0.77</td><td>0.06</td><td>71</td><td>14</td><td>23</td></tr><tr><td>N30</td><td>21.8</td><td>35.2</td><td>0.1</td><td>0.70</td><td>0.05</td><td>65</td><td>13</td><td>42</td></tr><tr><td>N31</td><td>21.6</td><td>35.6</td><td>0.2</td><td>0.68</td><td>0.04</td><td>76</td><td>17</td><td>39</td></tr><tr><td>N32</td><td>22.6</td><td>34.0</td><td>0.2</td><td>1.15</td><td>0.07</td><td>43</td><td>13</td><td>31</td></tr><tr><td>N33</td><td>21.7</td><td>35.5</td><td>0.1</td><td>0.80</td><td>0.05</td><td>54</td><td>12</td><td>29</td></tr><tr><td>N34</td><td>22.1</td><td>34.8</td><td>0.2</td><td>0.95</td><td>0.07</td><td>23</td><td>11</td><td>15</td></tr><tr><td>N35</td><td>21.1</td><td>36.5</td><td>0.1</td><td>0.64</td><td>0.03</td><td>74</td><td>18</td><td>21</td></tr><tr><td>N36</td><td>21.5</td><td>35.7</td><td>0.1</td><td>0.71</td><td>0.04</td><td>108</td><td>22</td><td>67</td></tr><tr><td>N37</td><td>21.4</td><td>35.9</td><td>0.1</td><td>0.65</td><td>0.04</td><td>68</td><td>17</td><td>23</td></tr><tr><td>N38</td><td>21.1</td><td>36.5</td><td>0.1</td><td>0.55</td><td>0.03</td><td>137</td><td>22</td><td>35</td></tr><tr><td>N39</td><td>21.0</td><td>36.8</td><td>0.2</td><td>0.56</td><td>0.03</td><td>59</td><td>17</td><td>30</td></tr><tr><td>N40</td><td>21.4</td><td>36.0</td><td>0.1</td><td>0.54</td><td>0.03</td><td>81</td><td>18</td><td>29</td></tr><tr><td>N41</td><td>21.1</td><td>36.5</td><td>0.2</td><td>0.58</td><td>0.04</td><td>70</td><td>12</td><td>15</td></tr><tr><td>N42</td><td>20.9</td><td>36.9</td><td>0.1</td><td>0.57</td><td>0.03</td><td>107</td><td>17</td><td>13</td></tr><tr><td>N43</td><td>21.0</td><td>36.8</td><td>0.1</td><td>0.60</td><td>0.02</td><td>103</td><td>11</td><td>28</td></tr><tr><td>N44</td><td>20.8</td><td>37.1</td><td>0.1</td><td>0.64</td><td>0.04</td><td>97</td><td>10</td><td>14</td></tr><tr><td>N45</td><td>21.4</td><td>35.9</td><td>0.1</td><td>0.64</td><td>0.04</td><td>&lt;d.l.</td><td>—</td><td>72</td></tr><tr><td>N46</td><td>21.4</td><td>36.0</td><td>0.1</td><td>0.70</td><td>0.03</td><td>92</td><td>15</td><td>27</td></tr><tr><td colspan="9">Chassigny</td></tr><tr><td>C1</td><td>24.8</td><td>36.7</td><td>0.1</td><td>1.02</td><td>0.05</td><td>67</td><td>9</td><td>12</td></tr><tr><td>C2</td><td>24.9</td><td>36.2</td><td>0.1</td><td>0.95</td><td>0.04</td><td>92</td><td>12</td><td>71</td></tr><tr><td>C3</td><td>25.4</td><td>35.0</td><td>0.1</td><td>1.14</td><td>0.09</td><td>&lt;d.l.</td><td>—</td><td>61</td></tr><tr><td>C4</td><td>25.8</td><td>33.9</td><td>0.1</td><td>1.20</td><td>0.06</td><td>&lt;d.l.</td><td>—</td><td>98</td></tr><tr><td>C5</td><td>26.9</td><td>31.6</td><td>0.1</td><td>1.74</td><td>0.08</td><td>15</td><td>5</td><td>15</td></tr><tr><td>C6</td><td>25.0</td><td>36.1</td><td>0.1</td><td>1.00</td><td>0.05</td><td>57</td><td>8</td><td>23</td></tr></tbody></table>

d.l. = detection limit; 1 SE = one standard error.
<figure><img src="image_6.png" /><figcaption>Fig. 4. Frequency diagram of Ru in chromite from the chassignites NWA 2737 and Chassigny. The vertical dashed line represents the average detection limit (i.e., 42 ppb) of all chromite analyses. *n* = number of analyses.</figcaption></figure>

higher ($D_{Ru}^{\text{ulvospinel-melt}} >150$) than for Fe$^{3+}$-poor chromite.

Sulfide immiscibility requires particular attention when interpreting Ru hosted in solid solution or as PGM inclusions in chromite. The extremely chalcophile behavior of Ru ($D_{Ru}^{\text{sulfide-melt}} \sim 10^4$ to $10^5$; Patten et al. 2013; Mungall and Brenan 2014) suggests that sulfide liquids presumably have an overriding control and concentrate this element approximately a hundred to thousand times more efficiently than co-crystallizing Fe$^{3+}$-poor chromite ($D_{Ru}^{\text{chromite-melt}} \sim 50-150$ at $fO_2 < \text{QFM}+2$). Consistently, the studies by Fiorentini et al. (2008) and Locmelis et al. (2011, 2013), which investigated the relationship between Ru and chromite in various terrestrial komatiites, komatiitic basalts, and ferropicrites, both in the whole-rock and at the mineral scale, showed that the solid solution signature of Ru in chromite may be utilized as a petrogenetic indicator for sulfide saturation in silicate melts.

## The Partitioning of Ru in Martian Chromites: Inferences on the Role of Sulfides, Magma Fractionation, and Ru in the Parental Melt

### Chassignites

The stability of sulfides in the formation of chassignites remains poorly understood, but can be addressed when evaluating the concentrations of Ru in

solid solution in the contained chromites. Sulfide blebs are minor constituents and occur in conjunction with mesostasis silicates interstitial to cumulate olivines and chromites (Fig. 1a) (cf. Lorand et al. 2012), or are rarely found in minor proportions in olivine-hosted silicate melt inclusions (Lorand et al. 2012). Both features can be explained by olivine and chromite crystallization in sulfide undersaturated systems. However, Ding et al. (2015) numerically modeled the fate of sulfur during crystallization of Martian magmas and found that the bulk sulfur concentrations in the chassignites (i.e., <300 ppm S in Chassigny; Lodders 1998; Franz et al. 2014) could reflect contribution from cumulate sulfides. Given the contrasting affinity of Ru for chromite and sulfide, it is possible to test whether the Ru concentrations detected in the analyzed chromites indicate their crystallization in the presence or absence of immiscible sulfides. Righter et al. (2015) modeled the potential PGE inventory of Martian mantle melts based on results from melting experiments and published partition coefficients. Their results suggest that sulfide undersaturated silicate melts could contain ~1–10 ppb Ru, whereas those saturated in sulfide may contain one order of magnitude lower concentrations. Considering chromite-melt partitioning of Ru appropriate for the crystallization conditions in the chassignites ($D_{Ru}^{\text{chromite-melt}} \sim 50-150$ at $fO_2 < \text{FMQ}+2$; Locmelis et al. 2011; Brenan et al. 2012; Pagé et al. 2012; Pagé and Barnes 2016), and the strong affinity of Ru for sulfide liquid, the Ru content in the parental melt must have been implausibly high (i.e., >100 ppb Ru) to account for the Ru concentrations observed in chromite from NWA 2737 (i.e., up to ~160 ppb Ru; Table 2). Similar calculations for chromite crystallization in the absence of sulfide require only ~1–4 ppb Ru to be present in the parental melt, which is in agreement with the theoretical model on the Ru inventory of Martian mantle-derived melts that are undersaturated in sulfide (Righter et al. 2015).

Among the analyzed chromites, we note a trend of progressively decreasing Ru with decreasing Cr (Fig. 5a; linear correlation coefficient, *R*, ~0.65), which is accompanied by an increase in Ti concentrations (Fig. 5b) and $100 \times \text{Fe}^{3+}/\text{R}^{3+}$ values (Fig. 2c). Such a trend was never observed before in the in situ analysis of Ru in chromite from terrestrial magmatic chromites (see Locmelis et al. 2011, 2013; Pagé et al. 2012; Park et al. 2012; Pagé and Barnes 2016; for comparison). The attainment of sulfide saturation during crystallization of chromite is unlikely to account for the observed variations as it would account for a sharp rather than a gradual reduction in Ru concentrations. On the other hand, the partitioning of Ru may attenuate gradually toward more reducing conditions. However, the study
<figure><img src="image_7.png" /><figcaption>Fig. 5. Concentrations of Ru, Ti, and Cr in chromite from the chassignites NWA 2737 and Chassigny. Data points lying on the x-axis represent analyses with Ru below the respective detection limits (Table 2). The horizontal dashed line represents the average detection limit (42 ppb Ru) for all chromite analyses in this study. It is noted that Ru positively correlates with Cr (a) and inversely correlates with Ti (b). The error bars indicate ±1 standard error. (Color figure can be viewed at wileyonlinelibrary.com.)</figcaption></figure>

by Brenan et al. (2012) predicts the partitioning of Ru to remain relatively constant over the interval of calculated oxygen fugacities in the chassignites (i.e., fO₂ ~QFM-1.3 to ~QFM+0.6; this study). We argue that the strong affinity of Ru for chromite itself could have been the main factor responsible for the observed Ru variations. For instance, by assuming ~1.5 ppb Ru in the parental melt, and that chromite is the sole phase that scavenges Ru at solid solution levels (Dₘ Chromite-melt ~100; Brenan et al. 2012; Pagé et al. 2012; Pagé and Barnes 2016), the concentrations in the evolving melt component would drop to ~0.5 ppb Ru within 1 wt% of chromite fractionation. Over the same interval, the Ru concentrations in chromite would decrease from initially ~150 ppb to ~50 ppb, the latter of which is close to the average detection limit (i.e., 42 ppb) of LA-ICP-MS analysis in this study. These calculations suggest that sulfide saturation was not achieved during olivine-chromite accumulation, and that the mesostasis sulfides, although occasionally in intimate contact with chromite (Fig. 1a) (cf. Lorand et al. 2012), could come from another source, such as trapped intercumulus melt during crystal settling, and/or late-stage migration/percolation of externally derived evolved melt components.

The calculations presented above provide crucial information on the potential Ru inventory and sulfide saturation history of the chassignite precursor melt during melt genesis in the mantle, ascent, and

emplacement. Whole-rock measurements for PGE are restricted to Chassigny (Jones et al. 2003). However, this set of data does not allow estimation of the Ru content of the melt parental to chassignites, given that potential sample heterogeneities and the small sample masses used in the analysis of PGE could have resulted in measured concentrations that do not accurately represent the true whole-rock abundances (usually referred to as nugget effect; see Meisel and Horan 2016). In addition, the fractionation and accumulation of chromite and olivine, both of which can take up Ru, likely have led to bulk concentrations in the chassignites that strongly differ from their parental melt. Our result of ~1–4 ppb thus represents a unique estimate for the Ru inventory of the primary mantle melt from which the chassignites formed, although we note that the fractionation of chromite, sulfide, or PGM prior to melt emplacement could have reduced the Ru tenor inherited from mantle melting. However, silicate and oxide fractionation may have been very limited given that the melt parental to the chassignites was shown to be very similar to the near-primary mantle melt parental to the Adirondack-class basalts at Gusev Crater (Filiberto 2008; Filiberto et al. 2008; He et al. 2013). Moreover, ~1–4 ppb Ru broadly overlaps with the calculated Ru inventory (i.e., ~1–10 ppb) of Martian mantle melts undersaturated in sulfide (Righter et al. 2015). These considerations support the inference that the parental melts did not
experience considerable segregation of sulfides prior to emplacement (i.e., during melt genesis and/or ascent), and that the calculated ~1–4 ppb Ru could represent concentrations inherited from mantle melting.

## Olivine-Phyric Shergottites

The analyzed olivine-phyric shergottites (Y-980459, Tissint, and Dhofar 019) display strong similarities, the most prominent of which are (1) broadly comparable mineralogical and petrographic features (e.g., Greshake et al. 2004; Papike et al. 2009; Balta et al. 2015), and (2) relative depletions in incompatible elements (Taylor et al. 2002; Shirai and Ebihara 2004; Balta et al. 2015), which were potentially inherited from the same or a very similar mantle source (e.g., Nyquist et al. 2001; Symes et al. 2008). Nevertheless, specific features could suggest that they are unrelated to each other, but derived from a common mantle source, and/or mimic petrographic commonalities by similar modes of magma emplacement and crystallization. These distinctive characteristics include variations in crystallization ages between 470 and 570 Ma based on the Sm-Nd and Rb-Sr isotopic system (e.g., Brennecka et al. [2014] and references therein), and subtle variations in chemical, mineralogical, and petrographic characteristics (see discussions in Papike et al. [2009] and Balta et al. [2015]). Despite the above uncertainties on the genetic relationships, it is reasonable to question why chromite and ulvöspinel from all analyzed shergottite specimens completely lack detectable concentrations of Ru in solid solution. Chromite in the studied olivine-phyric shergottites crystallized at lower oxygen fugacities (fO₂ between ~QFM-4 and ~QFM-2; e.g., Herd et al. 2002; Goodrich et al. 2003; Balta et al. 2015; this study) than in the chassignites (fO₂ between ~QFM-1.3 and ~QFM+0.7; Beck et al. 2006; Treiman et al. 2007; this study). These differences are reflected in relatively lower 100 × Fe³⁺/R³⁺ values among chromite from the olivine-phyric shergottites (Table 1), and could have influenced the uptake of Ru in chromite and ulvöspinel. However, over the interval of fO₂ and 100 × Fe³⁺/R³⁺ values calculated for chromite in both meteorite groups, Dₘₚᵢₘᵢ-eₘ₁t is predicted to remain constant within an approximated interval of 50–150 (Locmelis et al. 2011; Brenan et al. 2012; Pagé et al. 2012; Pagé and Barnes 2016). Accordingly, it appears unlikely that differences in the local redox conditions can solely account for the absence of detectable Ru concentrations in chromite from the olivine-phyric shergottites.

The mantle melts parental to Tissint and Dhofar 019 presumably experienced mineral fractionation prior to emplacement and crystallization (e.g., Taylor et al. 2002; Balta et al. 2015). We have shown that the continuous sequestration of Ru by crystallizing

chromite could have gradually decreased the Ru concentrations in the melt parental to the chassignites. Similarly, chromite fractionation (or alternatively PGM) prior to emplacement of Tissint and Dhofar 019 could have been a process that scavenged Ru from the melt. Alternatively, Ru may have been deficient due to melting of a mantle relatively depleted in PGE, or the sequestration of Ru by sulfides during melt genesis, ascent, and/or emplacement. The Dhofar 019 specimen requires particular attention in this respect as it shows anomalously low total PGE concentrations (Brandon et al. 2012) and a pronounced depletion in Ni (Taylor et al. 2002). These characteristics were explained by the early segregation of sulfides during melt genesis or ascent (Baumgartner et al. 2015; Righter et al. 2015), and such a process could indeed account for the absence of detectable Ru in chromite from Dhofar 019. The same arguments hold true for the absence of Ru in ulvöspinel, which formed late in the crystallization sequences from evolved melt components possibly saturated in sulfide.

The proposed mechanism of silicate-oxide fractionation is not adequate to explain the overall absence of Ru in chromite from the Y-980459 specimen, from which laboratory melting experiments suggest that it represents a primary mantle melt composition (Musselwhite et al. 2006). The PGE content and sulfide saturation history in the parental melt must be addressed. Brandon et al. (2012) observed total Ru concentrations in Y-980459 (~21–23 ppb Ru) one order of magnitude higher than in comparable olivine-phyric shergottites (usually ~1–5 ppb Ru). Although elevated whole-rock PGE concentrations were found in three sample aliquots, their results could be erroneous or nonrepresentative given the nugget effect in the analysis of small sample masses. Measured bulk sulfur concentrations of Y-980459 (<1500 ppm; Shirai and Ebihara 2004; Franz et al. 2014) are much lower than the capacity of the parental melt to dissolve sulfur (~5000 to ~3500 ppm for anhydrous melting of Y-980459 at 1 to 2.5 GPa, and ~4000 ppm S at crustal pressure regimes; Righter et al. 2009; Ding et al. 2014). This difference possibly reflects melt emplacement at sulfide undersaturated conditions, but its cause can be reconciled with various factors, such as deficient sulfur in the mantle, the release of sulfur-rich volatile species during melt ascent and/or emplacement, or sulfide saturation (Righter et al. 2009; Gaillard et al. 2013; Ding et al. 2015), the latter of which would have scavenged PGE from the melt.

Given the large uncertainties on sulfide genesis and PGE in Y-980459, an explanation for the absence of detectable Ru in chromite from this meteorite type
remains speculative. Yet, in the case that the elevated total PGE concentrations measured in Y-980459 are realistic, and that the parental melt emplaced in a sulfide undersaturated state, it is enigmatic why chromite, which is among the early phases to have crystallized and scavenged IPGE from the melt, is not the host of Ru in solid solution. The key to such a conundrum could lie in the apparent multistage magmatic history of Y-980459 (Greshake et al. 2004; Usui et al. 2008), at which chromite may have started to crystallize from a sulfide-bearing melt at depth, migrated upwards with the ascending magma, and settled in the rapidly cooled aphyric groundmass (Fig. 1b). Similarly, PGM, or alternatively PGE-rich sulfide-vapur phase assemblages (see Mungall et al. 2015) not at equilibrium with the sulfide undersaturated carrier melt (i.e., due to sulfide saturation at depth and the increasing sulfur capacity of silicate melts with decompression) could have been transported and deposited with the Y-980459 specimen. Such processes could account for (1) the elevated whole-rock PGE contents of Y-980459, (2) the apparent sulfide undersaturated signature of the precursor melt, and (3) the absence of detectable Ru in solid solution in chromite found in this study. Further detailed studies are required to decipher the mineralogical hosts of PGE in this meteorite specimen.

## SYNTHESIS AND CONCLUSION

In situ LA-ICP-MS analysis for the platinum-group elements Os, Ir, Ru, and Rh was carried out for the first time on magmatic spinel grains from a set of chassignite (NWA 2737 and Chassigny) and olivine-phyric shergottite meteorites (Y-980459, Tissint, and Dhofar 019). These analyses reveal the frequent occurrence of PGE-bearing inclusions, which comprise discrete Os, Ir, and Ru-rich PGM, and minor PGE-rich sulfides. Aside from PGE-rich inclusions, chromite and ulvöspinel from the olivine-phyric shergottites yielded PGE concentrations consistently below the detection limit, whereas chromite contained in the chassignite generally contained detectable Ru contents in solid solution (up to ~160 ppb Ru). Our data record the first evidence of the partitioning of Ru into the chromite lattice in a rock from a planetary body other than Earth.

The partitioning of Ru in chromite relies critically on the redox conditions and the stability of sulfides during crystallization, with sulfides predicted to have an overriding control on the partitioning of Ru in chromite over the range of oxygen fugacities recorded from the chassignites and olivine-phyric shergottites. These

characteristics allowed us to utilize concentrations of Ru in chromite as a proxy for the stability of sulfides and the Ru inventory in their parental melts. It is argued that the elevated Ru concentrations detected in chromite from the chassignites require their fractionation from a primitive and overall Ru-rich (~1–4 ppb) mantle melt that did not experience sulfide saturation over the interval of mantle melting, melt ascent, and chromite crystallization.

The absence of detectable Ru concentrations in chromite and ulvöspinel from the analyzed olivine-phyric shergottites can be explained by a variety of potential processes. The silicate melts parental to Tissint and Dhofar 019 presumably experienced silicate-oxide fractionation prior to emplacement, which could have efficiently scavenged and sequestered Ru from the melt. Moreover, Dhofar 019 shows chemical evidence for sulfide saturation and segregation prior to emplacement. Y-980459 is viewed to reflect a primary mantle melt with high bulk PGE, but low bulk sulfur abundances. The lack of detectable Ru concentrations in chromite from this specimen is difficult to reconcile with (1) the compatibility of Ru in Fe³⁺-rich chromite, and (2) the contention that Y-980459 emplaced at sulfide undersaturated conditions.

This work not only represents the first step toward characterization of the mineralogical hosts of platinum-group elements in Martian meteorites, but demonstrates that LA-ICP-MS analysis of Ru in magmatic chromite yields important quantitative information about the Ru inventory of the silicate melt parental to the Martian meteorites, and can shed light on the stability of sulfides in their genesis. This approach is particularly important in the study of planetary material such as Martian meteorites, where the suite of available sample material is small, and accurate extrapolation of precursor melt characteristics based on whole-rock characterization of sulfur and Ru is limited.

*Acknowledgments*—R. Baumgartner acknowledges the receipt of an UWA International Postgraduate Research Scholarship (IPRS). The research was financially supported by the ARC Centre of Excellence for Core to Crust Fluid Systems (CCFS), Matrix Exploration Pty Ltd and the French “Programme National de Planétologie” (INSU). M. Fiorentini acknowledges support from the Australian Research Council through Linkage Project LP120100668 and the Future Fellowship Scheme (FT110100241). D. Baratoux acknowledges support from the Programme National de Planétologie, France, and the Institut de Recherche pour le Développement for his current 4 years visiting appointment at the Institut Fondamental
d'Afrique Noire Ch. Anta Diop (Senegal). We thank H. Chennaoui Aoudjehane (Hassan II University in Casablanca, Morocco) and M. Hirai (National Institute of Polar Research in Tokyo, Japan) for providing samples for this study. F. Nemeth (University of Western Australia in Perth, Australia) is acknowledged for his help with sample preparation. M. Roberts and S. Caruso (University of Western Australia in Perth, Australia) are thanked for assistance with EPMA analysis. S. Mehdi, M. Prašek, and D. Savard (Université du Québec à Chicoutimi in Chicoutimi, Canada) are acknowledged for technical support with LA-ICP-MS analysis. We thank J. Brenan and C. Herd for helpful reviews, and A. Ruzicka for his insightful comments and editorial handling. This is contribution 873 from the ARC Centre of Excellence for Core to Crust Fluid Systems (http://www.ccfs.mq.edu.au).

*Editorial Handling*—Dr. Alexander Ruzicka

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