Non-chondritic iron isotope ratios in planetary mantles as a result of core formation

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1 In the format provided by the authors and unedited. S. M. Elardo and A. Shahar SUPPLEMENTARY INFORMATION DOI: /NGEO2896 Non-chondritic iron isotope ratios in planetary mantles as a result of core formation Calculation of the δ 57 Fe of Bulk Planetary Mantles The δ 57 Fe of the bulk silicate portions of Earth, the Moon, Mars and Vesta were calculated using Equation (2) in the main text. The experientially determined Δ 57 Fe Core-Mantle at 1850 C is calculated from a linear regression of the data in Figure 2 and includes our experimental data from Ni-bearing experiments the S-bearing experiments of Shahar et al. (2015). Both sets of experiments were included in the regression because the data sets overlap within analytical error. The regression did not include Ni- and S-free experiments in Figure 2. The regression resulted in the following relationship: Δ 57 Fe Core-Mantle at 1850 C = at. % Ni+S Fe Alloy , R 2 = 0.92 The full range of compositional parameters used to model the δ 57 Fe of planetary mantles is listed in Extended Table 3. The minimum and maximum for each planetary body represent the parameters that result in the minimum and maximum Δ 57 Fe Core-Mantle at 1850 C for that body and thus the least negative and most negative δ 57 Fe. Those minimum and maximum estimates are the upper and lower boundary curves in each plot in Fig. 3 in the main text. Our estimates of the core-mantle distribution of Fe, core Ni content, and mantle Fe content for Earth were taken from McDonough (2003). For the Moon, we use the core mass fraction of 2.4% from Steenstra et al. (2016b) and the Fe content of the bulk lunar mantle from the Lunar Primitive Upper Mantle composition of Longhi (2006) to calculate the core-mantle Fe distribution for each estimate of the composition of the lunar core. Estimates of the composition of the lunar core are from Righter and Drake (1996), Newsom (1984), Righter (2002), Ringwood and Seifert (1986), O Neill (1991) and the S content from Weber et al. (2011). For Mars, we use the core mass fraction of 20.6 % from Dreibus and Wänke (1985) and the Fe content of the bulk martian mantle from Lodders and Fegley (1997) to calculate the core-mantle Fe distribution for each estimate of the composition of the martian core. Estimates of the composition of the martian core are from Dreibus and Wänke (1985), Lodders and Fegley (1997), Morgan and Anders (1979), Bertka and Fei (1997), Sanloup et al. (1999), Johnston and Toksöz (1977), and Anderson (1972). For Mars, we also include the average composition of magmatic Fe meteorites from Mittlefehldt et al. (1998), which is the martian minimum estimate. This average would be within the range already established for the other planets, so including it for those would be redundant. For Vesta, we use the core mass fraction of 18% from Russell et al. (2012) and the Fe content of the bulk vestian mantle from Ruzicka et al. (1997) to calculate the core-mantle Fe distribution for each estimate of the composition of the vestian core. Estimates of the composition of the vestian NATURE GEOSCIENCE Macmillan Publishers Limited, part of Springer Nature. All rights reserved.

2 core are from Dreibus and Wänke (1980), Morgan et al. (1978), Hertogen et al. (1977) and Righter and Drake (1996). The extrapolation of the δ 57 Fe at 1850 C for each planetary body to higher and lower temperatures was done by scaling the Δ 57 Fe Core-Mantle according the 1/T 2 law of equilibrium mass dependent isotope fractionation, wherein fractionation goes to 0 at infinite T. The core formation temperature ranges for each planetary body were taken from literature estimates. The range from 2700 to 3600 C for Earth come from the estimates of Corgne et al. (2009), Siebert et al. (2011; 2013). The range from 1777 to 3077 C for the Moon comes from the estimates of Righter and Drake (1996) and Steenstra et al. (2016b). The range 1627 to 2027 C for Mars comes from the estimates of Righter and Drake (1996) and Righter and Chabot (2011). The range from 1452 to 1727 C for Vesta comes from the estimates of Righter and Drake (1996), Pringle et al. (2013), and Steenstra et al. (2016a). All of these calculations assume that the core formation process was homogeneous. This simplification is necessary in order to make a calculation of the bulk δ 57 Fe of planetary mantles, but we recognize that the actual process of core formation was likely heterogeneous (Wood et al., 2006; Rubie and Jacobson, 2015). Does magma ocean crystallization and mantle partial melting fractionate Fe isotopes in the absence of Fe 3+? We propose a model wherein the δ 57 Fe of bulk planetary mantles are light due to core formation, but are made heterogeneous with regions becoming progressively heavier during silicate differentiation due to mineral melt fractionation resulting in an enrichment of heavy Fe in the melt phase. This would also act to create basaltic partial melts that are heavier than their mantle source regions. There is convincing evidence that this process occurs on Earth where magmatic systems are relative oxidized and the presence of Fe 3+ creates differences in average Fe bonding environment between minerals and basaltic melt, resulting in fractionation. However, there is much debate about whether this mineral melt fractionation occurs in reduced magmatic systems such as those found on Mars, the Moon and Vesta. Below we discuss the lines of reasoning that suggest it does occur. 1) Mineral melt fractionation in terrestrial magmatic systems The clearest sample evidence of mineral melt fractionation during mantle partial melting on Earth comes from samples of MORBs and abyssal peridotites. In this tectonic setting, we can directly measure the δ 57 Fe of mantle partial melts and xenoliths of the mantle from which those melts were extracted. MORBs have an average δ 57 Fe of ± (Teng et al., 2013), which is distinctly heavier than their source region. Abyssal peridotites have an average measured δ 57 Fe of 0.01 ± 0.02 ; however Craddock et al. (2013) argued that after the effects of seafloor alteration are accounted for, the average of these samples is 0.04 ± They also calculated that the effects of ~10 15% partial melting on the residual peridotite mantle would be too small to resolve within the scatter of the abyssal peridotite dataset, so the measured composition of these samples is a reasonable approximation of the oceanic mantle. There are some studies that have found a heavier δ 57 Fe for some peridotites and this has been used to argue that Earth s mantle may be isotopically heavy (e.g., Poitrasson et al., 2013). There is almost certainly some

3 heterogeneity in Earth s upper mantle (Williams and Bizimis, 2014); however most studies infer a bulk mantle value close to the chondritic value (e.g., Weyer and Ionov, 2007; Craddock et al., 2013), and the global average δ 56 Fe is ± (Dauphas et al., 2017). A possible complication in interpreting these data is that the Mg# s of most MORBs indicate they are not primary magmas and have undergone some degree of fractional crystallization since extraction from their source regions. In a study of the Kilauea Iki lava lake samples, Teng et al. (2008) showed that fractional crystallization of olivine resulted in heavy isotope enrichment in the residual melt, a result supported by measurements of samples from the Red Hill intrusion by Sossi et al. (2012). This suggests that the δ 57 Fe of MORBs may be heavier than their primary parental magmas, thus inflating estimates of the partial melting fractionation. However, Teng et al. (2013) showed that the δ 57 Fe of MORBs does not vary in a statistically significant way over the relatively narrow range in MgO content of the basalts, indicating that the amount of olivine fractionation that occurred since source region extraction was minor enough to not measurably affect their δ 57 Fe. Therefore, the average MORB δ 57 Fe of ± is a valid representation of the composition of melts extracted from the oceanic mantle. Accordingly, we can calculate a terrestrial mantle partial melting fractionation factor for Fe isotopes. Using the seafloor alteration corrected value of 0.04 ± 0.04 for abyssal peridotites from Craddock et al. (2013) yields a partial melting fractionation of 0.11 ± Although it is not yet know what the relative contributions are from various Fehosting phases (i.e., olivine, pyroxenes, sulfides, oxides, garnet) in the mantle to the overall fractionation factor, we can establish that mantle melts become measurably enriched in heavy Fe isotopes through the melting process under terrestrial conditions. 2) Nuclear Resonant Inelastic X-ray Scattering (NRIXS) measurements by Dauphas et al. (2014) In a study that aimed to determine the effects of melt structure and redox state on Fe isotope fractionation, Dauphas et al. (2014) used the NRIXS technique to determine force constants of Fe bonds in quenched basaltic, dacitic, and rhyolitic glasses and Fo 82 olivine. They applied these data to address the enrichment of heavy Fe isotopes in MORBs relative to their source materials, abyssal peridotites. This δ 57 Fe offset is roughly 0.11 ± 0.04 (Fig. 1 in main text; Weyer and Ionov, 2007; Craddock et al., 2013; Teng et al., 2013). Their results indicate redox state can only account for ~1/3 of the observed fractionation between MORBs and their source peridotites, leaving a roughly 0.07 ± 0.03 offset that is presumably attributed to differences in the bonding environment of Fe 2+ between basaltic melts and the minerals. This is the first direct evidence that mineral melt equilibrium in reduced planetary magmatic systems that contain no Fe 3+ will result in fractionation of Fe isotopes. 3) Large Fe isotope variations in lunar basalts

4 The measurements of Dauphas et al. (2014) indicate that only 1/3 of the terrestrial mantle partial melting fractionation is due to the effects of Fe 3+ is supported by δ 57 Fe data for lunar basalts, which due to the reducing conditions of lunar magmatic systems, do not contain Fe 3+. Lunar mare basalts and volcanic glasses, which represent partial melts of the compositionally heterogeneous lunar mantle, range in δ 57 Fe from to 0.35 (Poitrasson et al., 2004; Weyer et al., 2005). These basalts show a trend in δ 57 Fe that correlates with bulk rock TiO 2 content and to a lesser extent MgO (Extended Figures 3 and 4). TiO 2 is incompatible during LMO crystallization up to the point of ilmenite saturation at ~95% solid (Snyder et al., 1992), so TiO 2 in lunar magmas is often used as a general indicator of when during LMO crystallization a basalt s source region formed. The clear trend of increasing δ 57 Fe with TiO 2 content strongly argues for mineral melt fractionation during crystallization the lunar magma ocean. Sample data from martian basaltic shergottites also supports mineral melt fractionation in martian systems. The redox state of martian magmas are significantly more reduced than terrestrial magmas and though some martian magmas contain Fe 3+, many do not and Fe 2+ dominates in all (Herd et al., 2001; 2002; Herd, 2003). However, shergotties show variations in δ 57 Fe that correlate with MgO content (Sossi et al., 2016), indicating a mineral melt fractionation that measurably affects the δ 57 Fe of martian magmas. Mineral melt fractionation factors for most silicate minerals and Fe-bearing oxides are not known; however it is likely that orthopyroxene, pigeonite, clinopyroxene, chromite, ilmenite, and possibly troilite are the minerals hosting most Fe in the lunar mantle, in addition to olivine (Snyder et al., 1992; Elardo et al., 2011). The NRIXS measurements of Dauphas et al. (2014) suggest that the Fe 2+ partitioning between Fo 82 olivine and basaltic melt does not fractionate Fe isotopes, but a wider compositional range in olivine needs investigation before conclusions can be drawn. Also, it is likely that pyroxenes are the dominant host of Fe in the source regions of lunar mare basalts and volcanic glasses (Hess, 2000; Shearer et al., 2006) and the Fe-hosting M-sites in pyroxenes become increasingly distorted with increasing Ca content (e.g., Cameron and Papike, 1981). This distortion directly affects Fe bonding environment and bond lengths, and therefore pyroxenes are a plausible candidate for a phase that fractionates Fe isotopes in Fe 3+ -free systems, though they need not be the only one. Estimating the δ 57 Fe of Planetary Mantles from the Compositions of Planetary Samples: A Comparison to the Predictions from Experiments In order to compare the prediction of negative δ 57 Fe values in planetary mantles from our experiments and mass-balance calculations, we can apply the estimated reduced mantle melting partial melting isotopic difference, or Δ 57 Fe Melt-Mantle, of 0.07 ± 0.03 that we discuss above to the compositions of basalts from the Moon, Mars and Vesta. Choosing samples to make this estimate is non-trivial. Lunar basalts span a wide range in compositions from to 0.35, so we discuss below our rationale for choosing a sample that will give a best estimate for a primitive source region in the Moon. The range in δ 57 Fe for Vesta and Mars is more restricted. This could be the result of sampling bias, relatively less mantle heterogeneity than in the Moon, or both. Mantle convection and mixing will act to erase the original heterogeneity created by magma ocean crystallization. This has almost certainly happened to some degree in Mars and the

5 Moon, although the very wide compositional range in major and trace elements in lunar basalts shows this mixing was inefficient. In the main text we use the range in δ 57 Fe of shergottites and non-cumulate eucrites to estimate their source region compositions. The basaltic shergottite range in δ 57 Fe is to which yields estimated mantle δ 57 Fe of to assuming Δ 57 Fe Melt-Mantle of 0.07 (see above). This estimate is necessarily first order because of uncertainties in the style, degree, and temperature of partial melting and the effects of varying degrees by fractional crystallization and crystal accumulation. Overall, the shergottite data are consistent with derivation from isotopically light source regions that are broadly consistent in composition with the predictions of our experimental results. The range in δ 57 Fe for non-cumulate eucrites is to This range yields an estimated mantle δ 57 Fe range from to using the Δ 57 Fe Melt-Mantle of The range in the eucrite dataset is subject to the same caveats of the martian dataset, but is broadly consistent with the predictions of our experimental results. The δ 57 Fe of the Lunar Mantle: Which Samples Provide the Best Estimate? Evidence from lunar samples and compositional models indicate the lunar mantle is extremely heterogeneous with respect to composition and mineralogy (e.g., Snyder et al., 1992; Shearer et al., 2006; Elardo et al., 2011; 2015) and this extends to Fe isotopes. Additionally, there are no lunar mantle xenoliths yet identified in the Apollo and meteorite sample collections (Shearer et al., 2015a). Therefore, it becomes problematic to select a basaltic sample which can be used to calculate a best estimate of the δ 57 Fe of the bulk lunar mantle. Such a sample may not exist because due to the heterogeneity created by lunar magma ocean (LMO) crystallization there may not be a region of the lunar mantle that is representative of the bulk mantle. However, within the currently available lunar sample collection, we argue the Apollo 15 green glasses are the most appropriate samples for this estimate for a few reasons. The Apollo 15 green volcanic glasses, and group C glasses especially, have the highest Mg# of any magmatic liquid composition, indicating they are derived from melting of one of the most lunar mantle primitive source regions yet sampled. They also have very low abundances of incompatible trace elements compared to other lunar basalts and have relative small negative Eu anomalies (Shearer and Papike, 1993). These geochemical characteristics indicate the source regions of the Apollo 15 green glasses crystallized from the LMO earlier than the source regions of other lunar magmas. The low-ti and high-ti mare basalt groups both have higher abundances of incompatible trace elements and deep, negative Eu anomalies, with the latter indicating that a least a component of source regions formed after plagioclase saturation in the LMO. Crystallization models of the LMO indicate that plagioclase does not begin to crystallize until roughly 80% solid (Snyder et al., 1992). Therefore, the source regions (or at least a component of the source regions) for the low-ti and high-ti mare basalt groups formed after roughly 80% of the magma ocean underwent fractional crystallization. This makes these groups poor choices for estimating the bulk mantle δ 57 Fe as the potential effects of Fe isotope fractionation during LMO crystallization will be more enhanced in these samples. Wang et al. (2015) argued that the very light δ 57 Fe (-0.58 ) of lunar dunite is a direct reflection of the composition of the deep lunar mantle; however it is

6 unlikely that this is the case. There is no evidence that dunite is a lunar mantle xenolith as argued by Wang et al. (2015). Rather, isotopic, mineralogical, geochemical and trace element evidence shows that is a magmatic cumulate rock and a member of the Mg-suite (Dymek et al., 1975; Ryder, 1992; Shearer and Papike, 2005; Elardo et al., 2011; Shearer et al., 2015b). The magma that crystallized was very likely derived from a source region that contained a component of the earliest dunitic cumulates from the LMO, but also components of the lateformed plagioclase-rich lunar crust (72415 contains plagioclase, also ruling out a high-pressure deep mantle origin) and the highly fractionated KREEP reservoir. Furthermore, textural and mineral compositional analysis of 7215 shows that the rock underwent very slow cooling, reequilibration, and possibly metasomatic alteration by sulfur- and/or phosphorus-rich fluids (Shearer et al., 2015b). Therefore, its δ 57 Fe is not suitable for extrapolation to the lunar mantle. Teng et al. (2008) showed that, at least under terrestrial redox conditions, the δ 57 Fe of a magma increases with increasing degrees of fractional crystallization. Furthermore, the fact that the source of the Apollo 15 green glasses appears to be the most primitive yet sampled also means that the temperature at which that source crystallized from the LMO and the temperature reached during melting of the parental magmas to the glasses will be higher than for other basalts and their source regions. Higher temperatures favor smaller isotopic fractionations, so this indicates the source of the green glasses is more likely to record a δ 57 Fe similar to the bulk lunar mantle than other mare basalt source regions. A potential complication in using the δ 57 Fe of the green glasses to estimate the δ 57 Fe of their mantle source region is that many of the lunar volcanic glasses have thin coatings that are thought to have formed from condensing vapor during the eruption that produced them. These coatings are rich in volatile elements such as S, Zn, Cl, and K (Butler and Meyer, 1976). Ding et al. (1983) attributed light S isotopic compositions in the orange glasses to the preferential degassing of light sulfur and its recondensation onto the surfaces of the glasses during the eruption process. Moynier et al. (2006) attributed the isotopically light δ 57 Fe of in the Apollo 17 orange glasses to light Fe isotopes this vapor condensate. They also found light isotopic compositions in for Zn and Cu as well. If true, this processes could artificially lower the isotopic composition of the Apollo 15 green glasses and shift our mantle estimate to lower values. However, vapor condensation is unlikely to have altered the Fe isotopic composition of either glass suite for a number of reasons. Butler and Meyer (1976) studied the compositions of coatings on various types of lunar glass beads, including the Apollo 15 green and 17 orange, and found detectable Fe in the coating of only a single bead. This shows there is not enough Fe present in these coatings to affect the isotopic compositions measured for the whole bead. The Apollo 17 orange volcanic glasses that Moynier et al. (2006) measured contain 23 wt. % FeO (Shearer and Papike, 1993), whereas they have only ~141 ppm Zn (Morgan et al., 1974). Therefore, it is highly unlikely that the thin coating of condensate on the glass beads surfaces would affect the measured composition of the orange glass for a major element like Fe. Zn isotopic compositions, however, with 3 orders of magnitude less Zn in the glasses, would certainly be affected by such a process. Furthermore, Poitrasson et al. (2004) measured the Apollo 17 orange glass as well and obtained a δ 57 Fe of 0.013, in stark contrast to Moynier et al. (2006) value of Therefore, we believe the measured δ 57 Fe of the Apollo 15 green glasses is likely to robust and reflect the Fe isotopic composition of a primitive, relatively unfractionated mantle melt.

7 The measured composition of the Apollo 15 green glasses is (Poitrasson et al., 2004). This value yields estimated mantle δ 57 Fe of to using the Δ 57 Fe Melt-Mantle of 0.07 (see above). This range is consistent with the δ 57 Fe of the bulk lunar mantle predicted by our experimental results. We can also apply the mantle melt fractionation factor to the low-ti and high-ti mare basalts to obtain estimates of the δ 57 Fe of their source regions, which formed after extensive lunar magma ocean crystallization and were likely fractionated to values heavier than the bulk silicate Moon. The range in δ 57 Fe for low-ti mare basalts is to This range yields estimated mantle δ 57 Fe of to using the Δ 57 Fe Melt-Mantle of The range in δ 57 Fe for high-ti mare basalts is to This range yields a range in source region δ 57 Fe, using the Δ 57 Fe Melt-Mantle of 0.07, of to These ranges for both mare basalt groups are subject to the same caveats discussed above, and also likely represent the inefficient mixing of the lunar mantle during mantle overturn and any subsequent convection that occurred after lunar magma ocean crystallization. Extended Data Figure 1: The Δ 57 Fe Core-Mantle as a function of experimental duration in Nifree experiments. This time series shows that the Δ 57 Fe Core-Mantle is relatively constant in experiments ranging from 0.5 to 3 hours, suggesting that isotopic equilibrium has been closely approached in these experiments. Extended Data Figure 2: The Δ 56 Fe of metal and silicate fractions of all experiments. The Δ 56 Fe value represents deviation from mass-dependent behavior. For a metal and silicate fraction to be in isotopic equilibrium, they must possess the same Δ 56 Fe, indicating they lie on the same mass-dependent Fe isotope fractionation line. All experiments presented here are within isotopic equilibrium within error. Deviations in the Δ 56 Fe of the bulk experiment were achieved by spiking the peridotite starting materials in 54 Fe such that the starting peridotite and metal were out of isotopic equilibrium (three isotope exchange method, see text). Extended Figure 3: The iron isotope compositions of lunar mare basalts and volcanic glasses vs. bulk rock TiO 2 content. Solid line is a linear regression of the dataset and the dashed lines denote the 95% confidence interval of the regression. Extended Figure 4: The iron isotope compositions of lunar mare basalts and volcanic glasses vs. bulk rock MgO content. Solid line is a linear regression of the dataset and the dashed lines denote the 95% confidence interval of the regression. References Cited Anderson, D. L., (1972) Internal Constitution of Mars. Journal of Geophysical Research 77, Bertka, C. M. and Fei, Y., (1997) Mineralogy of the Martian interior up to core-mantle boundary pressures. Journal of Geophysical Research 102,

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12 Supplementary Table 1: Chemical Compositions of Experimental Run Products Expt. # Phase T. ( C) P. (GPa) Duration (h) Starting Metal SiO 2 Al 2 O 3 FeO MgO CaO NiO Fe Ni C corr Total PC 1350 Silicate None S. D PC 1375 Silicate S. D Metal Fe S. D PC1373 Silicate S. D Metal Fe S. D PC1364 Silicate S. D Metal Fe S. D PC1376 Silicate S. D Metal Fe S. D PC 1361 Silicate bdl S. D Metal Fe 0.94 Ni S. D PC 1448 Silicate bdl S. D Metal Fe 0.94 Ni S. D PC 1487 Silicate S. D Metal Fe 0.75 Ni S. D PC 1488 Silicate S. D Metal Fe 0.85 Ni S. D

13 Supplementary Table 2: Iron Isotopic Compositions of Geostandards, Starting Materials, and Experimental Run Products Sample Phase n δ 56 Fe ( ) 2*S.E. δ 57 Fe ( ) 2*S.E. Δ 56 Fe ( ) 2*S.E. Δ 57 Fe Core-Mantle ( ) 2*S.E. Geostandards AGV Accepted Values* BIR Accepted Values* BHVO Accepted Values* Starting Materials Alfa Aesar Fe Metal Peridotite Mix Experiments PC 1375 Silicate Metal PC1373 Silicate Metal PC1364 Silicate Metal PC1376 Silicate Metal PC 1361 Silicate Metal PC 1448 Silicate Metal PC 1487 Silicate Metal PC 1488 Silicate Metal *Values from Craddock and Dauphas (2010)

14 Supplementary Table 3: Values used to model the δ 57 Fe of planetary mantles Core Mass Fe wt. % Ni wt. % S wt. % Δ 57 Fe Core-Mantle Fraction f Fe Mantle Mantle Core Core at. % Ni Core at. % S Core at 1850 C Ref. Comments Earth Used for all Earth estimates Earth Minimum Earth Maximum The Moon , 3 Used for all Moon estimates Moon Minimum , 9 Moon Maximum Mars , 11 Used for all Mars estimates Mars Minimum a Mars Maximum Vesta , 19 Used for all Vesta estimates Vesta Minimum Vesta Maximum a = The "Mars Minimum" values are calculated assuming 8.4 wt. Ni in the core, the average values of magmatic Fe meteorites References [1] McDonough, 2003 [2] Steenstra et al., 2016, [3] Longhi, 2006 [4] Weber et al. 2011, [5] Righter and Drake, 1996 [6] Newsom, 1984 [7] Righter, 2002 [8] Ringwood and Seifert, 1986 [9] O'Neill, 1991 [10] Dreibus and Wänke, 1985 [11] Lodders and Fegley, 1997 [12] Mittlefehldt et al., 1998, [13] Morgan and Anders, 1979 [14] Bertka and Fei, 1997 [15] Sanloup et al., 1999 [16] Johnston and Toksöz, 1977 [17] Anderson, 1972 [18] Russell et al., 2012 [19] Ruzicka et al., 1997 [20] Dreibus and Wänke, 1980 [21] Morgan et al., 1978 [22] Hertogen et al., 1977

15 0.2 Ni-free series at 1850 C Δ 57 Fe Core-Mantle ( ) Time (hours)

16 PC1364 Graphite Capsule With Varying Ni No S PC1375 Δ 56 Fe ( ) PC1373 PC1376 PC1488 PC PC1487 PC Graphite Capsule No Ni or S Circles - Silicate Fractions Squares - Metal Fractions

17 0.4 Lunar Mare Basalts and Volcanic Glasses 0.3 δ 57 Fe ( ) Low-Ti Mare Basalts High-Ti Mare Basalts A15 Green Glass A17 Orange Glass Bulk Rock TiO 2 (wt. %)

18 0.3 Lunar Mare Basalts and Volcanic Glasses 0.4 Low-Ti Mare Basalts High-Ti Mare Basalts A15 Green Glass A17 Orange Glass δ 57 Fe ( ) Bulk Rock MgO (wt. %)