<figure><img src="image_1.png" /></figure>

HAL
open science

# Oxidized conditions in iron meteorite parent bodies

Pierre Bonnand, A. Halliday

► To cite this version:

Pierre Bonnand, A. Halliday. Oxidized conditions in iron meteorite parent bodies. Nature Geoscience, 2018, 11 (6), pp.401 - 404. 10.1038/s41561-018-0128-2. hal-01833374

HAL Id: hal-01833374

https://uca.hal.science/hal-01833374v1

Submitted on 2 Oct 2025

HAL is a multi-disciplinary open access
archive for the deposit and dissemination of sci-
entific research documents, whether they are pub-
lished or not. The documents may come from
teaching and research institutions in France or
abroad, or from public or private research centers.

L'archive ouverte pluridisciplinaire HAL, est
destinée au dépôt et à la diffusion de documents
scientifiques de niveau recherche, publiés ou non,
émanant des établissements d'enseignement et de
recherche français ou étrangers, des laboratoires
publics ou privés.

<figure><img src="image_2.png" /></figure>

Distributed under a Creative Commons Attribution 4.0 International License
# Oxidized conditions in iron meteorite parent bodies

P. Bonnand<fnref n="1" />,²* and A. N. Halliday<fnref n="1" />,³

Magmatic iron meteorites are thought to originate from the cores of differentiated planetary embryos, but it is unclear where they formed in the circumstellar disk. Oxygen fugacity ($fO_2$) is thought to have varied radially across the disk, affecting planetary accretion and core formation. Metal-silicate equilibration experiments show that the chromium concentration in metallic melt increases with decreasing $fO_2$. Chromium can therefore be used to reconstruct $fO_2$ during metal segregation. Here, we report mass-dependent fractionation of stable chromium isotopes measured for several groups of magmatic iron meteorites. We find that chromium concentrations are variable and most samples are isotopically heavy compared with the bulk silicate Earth and chondrites. Although it is evident that the chromium isotopes were systematically affected by fractional crystallization, we use the mass-dependent fractionation behaviour to derive initial chromium concentrations for the planetary embryo cores. We find that the cores were diverse in $fO_2$ and demonstrate that core formation in some planetary embryos occurred under relatively oxidizing conditions ($fO_2$ relative to iron-wüstite (IW) buffer ($\Delta$IW) = -0.95 ± 0.15). It is difficult to reconcile this with the view that iron meteorite parent bodies formed closer to Earth and subsequently migrated outwards. Instead, we suggest that these parent bodies formed at greater distances, beyond Mars.

Many elements (referred to as 'siderophile') are partially or strongly partitioned into core-forming metals and metallic liquids, and are therefore depleted in the silicate mantle of differentiated planetary bodies relative to chondrites. The degree of partitioning of moderately siderophile elements between metallic iron and iron alloy liquids on the one hand and silicates and silicate liquids on the other is dependent on physical and chemical parameters such as temperature, pressure and oxygen fugacity ($fO_2$). Constraining these parameters is crucial for understanding the physical and chemical gradients during the early stages of planetary differentiation. Such reactions have been extensively studied by (1) examining the composition of natural samples, such as the Earth's mantle and iron meteorites, and (2) conducting experiments (for example, refs. <fnref n="1" />,²). However, the conditions and mechanisms that govern these reactions in planets and planetary embryos are debated. The $fO_2$ in the solar nebula is highly heterogeneous with chondrites believed to have formed over a wide range of $fO_2$ (relative to iron-wüstite ($\Delta$IW)), from -4 for enstatite chondrites to 0 for carbonaceous chondrites³,⁴. It has been proposed that the solar nebula was zoned in terms of $fO_2$, with regions closer to the Sun depleted in oxygen and thus relatively reduced. It is still unclear where the iron meteorite parent bodies formed⁵,⁶ and whether they formed under reduced or oxidized conditions.

## Chromium concentrations as a redox proxy

Chromium is a moderately siderophile element. The integrated model distribution coefficient for chromium between metal and silicate melt (DCR$_{metal-silicate}$) consistent with its overall levels of depletion in the bulk silicate Earth is 2.5–3.5 (refs ²,⁷,⁸). Chromium partitioning is strongly affected by $fO_2$, but unaffected by the iron-to-nickel ratio in the iron melts, major element concentration in the silicate melt and pressure⁹⁻¹⁴. This makes studying the chromium concentration in the metal fraction a good proxy for $fO_2$. As illustrated in Fig. 1, the chromium concentration in the metal fraction during metal-silicate

equilibrium experiments increases with decreasing $fO_2$. To calculate the chromium concentrations shown in Fig. 1, several assumptions were made. First, only low-pressure experiments ($\le 3$ GPa)¹¹ were used because high-pressure experiments are not applicable to iron meteorite parent bodies. Second, two chromium concentrations are presented in Fig. 1: 2,500 µg g⁻<fnref n="1" /> (similar to Earth's mantle) and 3,200 µg g⁻<fnref n="1" /> (similar to ordinary chondrites). It is important to note that varying the chromium concentration in the silicate fraction does not dramatically change the relationship between the chromium concentration in the metal and $fO_2$. The correlation between the chromium concentration in the metal and $fO_2$ suggests that the chromium concentration in iron metal could be an effective proxy for determining $fO_2$ during metal segregation. Determining the initial chromium concentration in the iron liquid formed during the differentiation of planetary embryos would allow us to assess one of the key parameters ($fO_2$) controlling chemical reactions during core formation.

## Chromium concentrations in magmatic iron meteorites

Magmatic iron meteorites provide unique insights into core formation processes both on Earth and smaller planetary embryos (for example, ref. ¹⁵). Based on chemical heterogeneities, iron meteorites have been divided into a number of groups, each of which reflects a unique parent body¹⁵. Chemical trends (usually with nickel or gold concentrations) have been explained by fractional crystallization of solid iron during cooling of the liquid. Several models have been proposed to explain the chemical variations¹⁵, with certain elements such as iridium viewed as highly compatible and others such as gold being highly incompatible. Chromium is poorly understood because experiments provide evidence that it should behave incompatibly during iron liquid crystallization with a distribution coefficient between iron solid and iron liquid (DCR$_{solid-liquid}$) of ~0.62 (ref. ¹⁶). However, chromium concentrations in the three main groups (IIAB, IIIAB and IVA) decrease with increasing gold concentration, which

<footnote n="1">¹Department of Earth Sciences, University of Oxford, Oxford, UK. ²Laboratoire Magmas et Volcans, Université Clermont Auvergne, CNRS, Clermont-Ferrand, France. ³The Earth Institute, Columbia University, New York, NY, USA. *e-mail: pierre.bonnand@uca.fr</footnote>
<figure><img src="image_3.png" /><figcaption>**Fig. 1 |** Chromium concentrations in the metal fraction during metal-silicate equilibrium experiments versus $fO_2$ (ΔIW). The experiments were performed in graphite capsules at temperatures ranging from 1,850 to 2,600 K and pressures from 1.5 to 3 GPa. The chromium concentrations in the metal fractions at equilibrium were calculated assuming an Earth's mantle chromium concentration of 2,500 µg g⁻¹ (circles) and an LL chondrite chromium concentration of 3,200 µg g⁻¹ in the silicate fraction. The solid and dashed lines are linear regressions for the respective chromium concentrations in the silicate fraction. Partition coefficients from ref. ¹¹. R² is the coefficient of determination obtained in the experimental study¹¹.</figcaption></figure>

means that chromium behaves compatibly ($DCr_{solid-liquid} > 1$) during fractional crystallization of iron liquid (Fig. 2)¹⁷⁻¹⁹. In group IVA, for example, the chromium concentration decreases from 600 µg g⁻¹ to <100 µg g⁻¹ while the gold concentration increases from 0.6 µg g⁻¹ to 1.5 µg g⁻¹ (Fig. 2)¹⁸. One of the proposed explanations is the early crystallization of chromite in the crystallization sequence¹⁶. Given the high variability of chromium content in iron meteorites, it is difficult to assess the chromium concentration in the parent iron liquid before fractional crystallization. Mass-dependent stable isotope variations in iron meteorites have the potential to provide new constraints and have previously been reported for iron²⁰, zinc²¹, copper²² and nickel²³,²⁴. Here, we report a detailed study of mass-dependent chromium isotope variations aimed at constraining the behaviour of chromium in planetary cores. The variations are reported in per mil (‰) deviations expressed as $\delta^{53}\text{Cr} = ((^{53}\text{Cr}/^{52}\text{Cr})_{\text{sample}}/(^{53}\text{Cr}/^{52}\text{Cr})_{\text{NIST979}} - 1) \times 1,000$, where $(^{53}\text{Cr}/^{52}\text{Cr})_{\text{sample}}$ and $(^{53}\text{Cr}/^{52}\text{Cr})_{\text{NIST979}}$ are the chromium isotopic compositions of the sample and the isotopically certified chromium reference material NIST SRM979, respectively.

## Chromium isotopes

Mass-dependent chromium isotope variations have been studied recently in a range of terrestrial and extraterrestrial samples. Much of the variability is associated with oxidation-reduction at low temperatures (for example, ref. ²⁵) with relatively little associated with igneous processes except during metasomatism and/or re-fertilization²⁶,²⁷. However, it has been shown that chromium isotopes are fractionated during fractional crystallization on the Moon²⁸. It has also been reported that the silicate Earth is isotopically fractionated by terrestrial core formation relative to chondritic values²⁹,³⁰, but more recent work has shown that this is not the case³¹,³². Metal-silicate piston cylinder experiments performed at 1.5 GPa and ~1,925 K³¹ show that there is no equilibrium isotopic fractionation between metal and silicate, consistent with the more recent measurements of chondrites. On this basis, it seems most likely that metallic cores have chondritic chromium isotopic compositions.

<figure><img src="image_4.png" /><figcaption>**Fig. 2 |** Concentrations of gold and chromium in iron meteorites. Chromium concentrations versus gold concentrations in iron meteorites from the three main groups (IIAB (diamonds), IIIBAB (circles) and IVA (squares)). Data from refs ¹⁷⁻¹⁹.</figcaption></figure>

We analysed 18 iron meteorites from five groups (IC, IIAB, IIIBAB, IIIF and IVA). We determined trace element concentrations, together with chromium isotopic compositions. Each sample was measured twice—once for radiogenic ($\epsilon^{53}\text{Cr}$) and nucleosynthetic ($\epsilon^{54}\text{Cr}$) anomalies and cosmic ray spallation processes (both $\epsilon^{53}\text{Cr}$ and $\epsilon^{54}\text{Cr}$), and once for their mass-dependent isotopic variations (Supplementary Table 1). As previously reported, there is a strong correlation between $\epsilon^{53}\text{Cr}$ and $\epsilon^{54}\text{Cr}$, and these variations are attributed to spallation³³,³⁴,³⁵ (Supplementary Fig. 1). Mass-dependent variations, investigated using a double-spike technique, were corrected for mass-independent variations (see Methods). Except for one sample, all the iron meteorites are characterized by heavy $\delta^{53}\text{Cr}$ values compared with the bulk silicate Earth and chondrites, ranging from +0.22 ± 0.10 ‰ to +2.85 ± 0.10 ‰ (Fig. 3). They are, to date, the heaviest isotopic compositions measured in high-temperature geological systems. The five studied groups present large intragroup chromium isotope variations. For example, $\delta^{53}\text{Cr}$ in group IIIBAB samples ranges from +0.89 ± 0.10 ‰ to +2.78 ± 0.10 ‰ (Fig. 3). There is a strong correlation between the chromium isotopic composition and chromium concentration in group IVA. This correlation is not present for the other groups, in which the variation in the chromium concentration is more limited (10 to 100 µg g⁻¹). Chromium isotopes are also correlated with nickel and gold concentrations (Supplementary Fig. 3) for the IVA group.

The chromium concentrations in iron meteorites are highly variable (10 to 600 µg g⁻¹; Fig. 2) and although there is a covariation with gold concentrations, the processes responsible for the chromium concentration variations are debated¹⁶. It is difficult to determine an initial chromium concentration in the iron liquid at equilibrium with silicate during metal segregation based on iron meteorite samples strongly affected by fractional crystallization. A direct estimate of the $fO_2$ during core formation is therefore intractable. Mass-dependent chromium isotope fractionations can help resolve this problem. As chromium isotopes do not fractionate during metal-silicate equilibrium, hence presumably during core formation³¹, the chromium isotopic composition of the initial iron liquid at equilibrium with a silicate mantle should be chondritic. As shown in Fig. 3, chromium isotopes in iron meteorites are highly variable and this indicates that they have been affected by a second process after core formation. The compatible behaviour of chromium cannot be explained by iron solid crystallizing out of iron liquid; therefore, the negative correlation between chromium isotopic composition and concentration in the IVA group (Fig. 3) suggests the role of another phase that is isotopically light. The lack of correlation in other groups could be explained by different crystallization histories of each parent body and the absence of the specific phase responsible for chromium removal in group IVA.
<figure><img src="image_5.png" /><figcaption>**Fig. 3** | Chromium isotopic compositions versus chromium and nickel concentrations in iron meteorites. **a**, Chromium isotopic composition versus chromium concentration. The grey horizontal bar is the bulk silicate Earth (BSE) and chondritic value<sup>26,31,32</sup>. Two models are presented. The equilibrium model was calculated with an isotopic fractionation ε of −2.25‰, and the kinetic fractional model was calculated with an ε of −0.8‰. For the kinetic model, 'solid inst.' is the instantaneous chromium isotopic composition of the crystallizing solid, while 'solid acc.' is the chromium isotopic composition of the accumulated solid. **b**, Chromium isotopic composition versus nickel concentration. The solid line is the linear regression through the IVA dataset. Error bars represent 2 s.d.</figcaption></figure>

Using both chromium concentrations and chromium isotopes in iron meteorites, we can determine the chromium concentration in the original iron liquid in group IVA formed at equilibrium with a silicate phase. A Rayleigh fractionation model (kinetic reactions) and an equilibrium fractionation model, fitted to the data to calculate the initial chromium concentration in the metal (see Supplementary Information) are presented in Fig. 3a. Our models assume that the initial δ<sup>53</sup>Cr value of the iron liquid before fractional crystallization starts is similar to the chondrite composition of δ<sup>53</sup>Cr = −0.12 ± 0.10 ‰ (refs <sup>24,29,30</sup>). It is interesting to note that the IIIF iron meteorite lies close to the chromium isotopic composition calculated for the crystallizing phase. The equilibrium and kinetic fractionation models presented in Fig. 3 give an initial chromium concentration in the iron liquid in the IVA group of 550 ± 50 µg g<sup>−1</sup> and 850 ± 150 µg g<sup>−1</sup>, respectively. This initial chromium concentration is small compared with the assumed chromium concentration in the Earth's core of about 7,500 µg g<sup>−1</sup> (ref. 7). If we assume the aforementioned chromium concentrations in the initial IVA iron liquid (between 500 and 1,000 µg g<sup>−1</sup>), an fO<sub>2</sub> value (ΔIW) of −0.95 ± 0.15 can be calculated using the correlations presented in Fig. 1. This provides evidence that iron meteorites formed from a reservoir that was relatively oxidized compared with Earth, but similar to the fO<sub>2</sub> associated with chondrite formation (for example, ref. 3). The relatively high fO<sub>2</sub> obtained for iron meteorite parent bodies is consistent with accretion near the asteroid belt. It is difficult to reconcile this with models proposing that iron meteorite parent bodies formed in the vicinity of the terrestrial planets and migrated outwards<sup>5</sup>. The similarity in fO<sub>2</sub> between IVA iron meteorites' parent body and ordinary chondrites

is also supported by similar oxygen and nucleosynthetic ε<sup>54</sup>Cr isotopic compositions<sup>36,37</sup>.

Several crystallizing phases could be responsible for the chromium isotope variations observed in iron meteorites. Two main chromium-bearing minerals can explain the compatible behaviour of chromium; namely, the chromium oxide chromite (FeCr<sub>2</sub>O<sub>4</sub>) and the chromium sulphide daubréelite (FeCr<sub>3</sub>S<sub>4</sub>). Both have been described in iron meteorites<sup>18,38,39</sup>. Arguments can be made in favour of each of these phases being responsible for the removal of chromium. Experimental petrology provides evidence that in the presence of sulphur, chromium becomes more incompatible during fractional crystallization<sup>16</sup>. It is therefore difficult to explain the decrease in chromium early in the crystallization sequence with sulphides crystallizing. However, chromite is believed to be a very early phase in the crystallization history of iron liquids represented by some iron meteorites. Therefore, it has been proposed in the literature that chromite may be the best explanation for the compatible behaviour of chromium during fractional crystallization. The isotopic variation could be due to equilibrium fractionation between oxides and metallic liquids. However, the chromium isotopic compositions of terrestrial chromites are similar to bulk mafic/ultramafic rocks and not very light, as would be required to generate the trends in Fig. 3 (ref. 40). Furthermore, ab initio calculations show that chromite and daubréelite would be isotopically heavy compared with iron liquids<sup>29</sup>, whereas CrS would be isotopically light. Thus, if the isotopic variations observed are due to equilibrium reactions, CrS is a better candidate than chromites. However, fractional crystallization is rarely an equilibrium reaction and the observed variations could be due to the necessary oxidation reactions required to form oxides (or sulphides) in the iron liquid environment. Indeed, during core formation processes, chromium exists in three oxidation states: Cr<sup>0</sup> in the metal fraction, Cr<sup>2+</sup> in CrS, and Cr<sup>3+</sup> in chromite and daubréelite. In this case, the fractionation would follow a kinetic reaction and could be easily explained by a Rayleigh fractionation model (Fig. 3). Further work is needed to evaluate the degree of isotopic fractionation between metallic liquids, sulphides and oxides and to investigate the small number of negative δ<sup>53</sup>Cr values in our dataset, hence the apparent absence of cumulate-rich compositions.

## Implications for parent body formation

Large mass-dependent chromium isotopic variations have been discovered in a group of 18 iron meteorite samples. Nearly all are characterized by heavy δ<sup>53</sup>Cr relative to chondrites and this indicates that chromium isotopes are fractionated during fractional crystallization of the iron liquid. Given the chromium behaviour during fractional crystallization, the observed fractionation is probably due to the crystallization of chromium-bearing phases such as chromites or sulphides. The mechanism responsible for chromium isotopic fractionation could be equilibrium fractionation between iron liquid and sulphides or kinetic fractionation during oxidation of chromium from Cr<sup>0</sup> to Cr<sup>2+</sup> or Cr<sup>3+</sup>.

The correlation between chromium concentrations and chromium isotopes in group IVA facilitates determination of the initial chromium concentration of the iron liquid during metal segregation. The relationship between the chromium concentration and fO<sub>2</sub> then defines the initial fO<sub>2</sub> of the IVA group. This fO<sub>2</sub> is relatively oxidized and has a value (ΔIW) of −0.95 ± 0.15. The lack of correlation between chromium isotopes and chromium concentration in other groups (IIAB and IIIBAB) does not permit determination of the initial fO<sub>2</sub> for these bodies. However, assuming similar reactions and therefore fractionation pathways, it is possible to say that these bodies also formed under relatively oxidized conditions.

It has been proposed in the literature that the solar nebula was zoned in terms of fO<sub>2</sub> with reduced conditions near the Sun. With an fO<sub>2</sub> just below the IW buffer, this suggests that the planetesimals
sampled by the iron meteorites formed between Mars and the asteroid belt. It is also consistent with the proposed chemical link between ordinary chondrites and IVA iron meteorites' parent body based on similar oxygen and nucleosynthetic $e^{54}Cr$ and molybdenum isotopic compositions.

## Methods

Methods, including statements of data availability and any associated accession codes and references, are available at https://doi.org/10.1038/s41561-018-0128-2.

Received: 14 June 2017; Accepted: 13 April 2018;
Published online: 7 May 2018

## References

1. Newsom, H. E. in *Origin of the Earth* (eds Newsom, H. E. & Jones, J. H.) 273–288 (Oxford Univ. Press, New York, 1990).

2. Wood, B. J., Wade, J. & Kilburn, M. R. Core formation and the oxidation state of the Earth: additional constraints from Nb, V and Cr partitioning. *Geochim. Cosmochim. Acta* **72**, 1415–1426 (2008).

3. Brett, R. & Sato, M. Intrinsic oxygen fugacity measurements on 7 chondrites, a pallasite and a tekite and the redox state of the meteorite parent bodies. *Geochim. Cosmochim. Acta* **48**, 111–120 (1984).

4. Cartier, C. et al. Experimental study of trace element partitioning between enstatite and melt in enstatite chondrites at low oxygen fugacities and 5 GPa. *Geochim. Cosmochim. Acta* **130**, 167–187 (2014).

5. Bottke, W., Nesvorny, D., Grimm, R., Morbidelli, A. & O'Brien, D. Iron meteorites as remnants of planetesimals formed in the terrestrial planet region. *Nature* **439**, 821–824 (2006).

6. Kruijer, T. S., Burkhardt, C., Budde, G. & Kleine, T. Age of Jupiter inferred from the distinct genetics and formation times of meteorites. *Proc. Natl Acad. Sci. USA* **114**, 6712–6716 (2017).

7. Allègre, C., Poirier, J., Humler, E. & Hofmann, A. The chemical composition of the Earth. *Earth Planet. Sci. Lett.* **134**, 515–526 (1995).

8. McDonough, W. & Sun, S. The composition of the Earth. *Chem. Geol.* **120**, 223–253 (1995).

9. Gessmann, C. & Rubie, D. The origin of the depletions of V, Cr and Mn in the mantles of the Earth and Moon. *Earth Planet. Sci. Lett.* **184**, 95–107 (2000).

10. Righter, K. Prediction of metal-silicate partitioning coefficients for siderophile elements: an update and assessment of PT conditions for metal-silicate equilibrium during accretion of the Earth. *Earth Planet. Sci. Lett.* **304**, 158–167 (2011).

11. Siebert, J., Corgne, A. & Ryerson, F. J. Systematics of metal-silicate partitioning for many siderophile elements applied to Earth's core formation. *Geochim. Cosmochim. Acta* **75**, 1451–1489 (2011).

12. Tuff, J., Wood, B. J. & Wade, J. The effect of Si on metal-silicate partitioning of siderophile elements and implications for the conditions of core formation. *Geochim. Cosmochim. Acta* **75**, 673–690 (2011).

13. Fischer, R. A. et al. High pressure metal-silicate partitioning of Ni, Co, V, Cr, Si, and O. *Geochim. Cosmochim. Acta* **167**, 177–194 (2015).

14. Clesi, V. et al. Effect of H₂O on metal-silicate partitioning of Ni, Co, V, Cr, Mn and Fe: implications for the oxidation state of the Earth. *Geochim. Cosmochim. Acta* **192**, 97–121 (2016).

15. Goldstein, J. I., Scott, E. R. D. & Chabot, N. L. Iron meteorites: crystallization, thermal history, parent bodies, and origin. *Chem. Erde Geochem.* **69**, 293–325 (2009).

16. Chabot, N. L., Saslow, S. A., McDonough, W. F. & Jones, J. H. An investigation of the behavior of Cu and Cr during iron meteorite crystallization. *Meteorit. Planet. Sci.* **44**, 505–519 (2009).

17. Wasson, J. Trapped melt in IIAB irons: solid/liquid elemental partitioning during the fractionation of the IIAB magma. *Geochim. Cosmochim. Acta* **63**, 2875–2889 (1999).

18. Wasson, J. & Richardson, J. Fractionation trends among IVA iron meteorites: contrasts with IIAB trends. *Geochim. Cosmochim. Acta* **65**, 951–970 (2001).

19. Wasson, J. T., Huber, H. & Malvin, D. J. Formation of IIAB iron meteorites. *Geochim. Cosmochim. Acta* **71**, 760–781 (2007).

20. Williams, H. M. et al. Fe isotope fractionation in iron meteorites: new insights into metal-sulphide segregation and planetary accretion. *Earth Planet. Sci. Lett.* **250**, 486–500 (2006).

21. Bridgestock, L. J. et al. Unlocking the zinc isotope systematics of iron meteorites. *Earth Planet. Sci. Lett.* **400**, 153–164 (2014).

22. Bishop, M. C. et al. The Cu isotopic composition of iron meteorites. *Meteorit. Planet. Sci.* **47**, 268–276 (2012).

23. Cook, D. L. et al. Mass-dependent fractionation of nickel isotopes in meteoritic metal. *Meteorit. Planet. Sci.* **42**, 2067–2077 (2007).

24. Gall, L., Williams, H. M., Halliday, A. N. & Kerr, A. C. Nickel isotopic composition of the mantle. *Geochim. Cosmochim. Acta* **199**, 196–209 (2017).

25. Ellis, A. S., Johnson, T. M. & Bullen, T. D. Chromium isotopes and the fate of hexavalent chromium in the environment. *Science* **295**, 2060–2062 (2002).

26. Schoenberg, R., Zink, S., Staubwasser, M. & von Blanckenburg, F. The stable Cr isotope inventory of solid Earth reservoirs determined by double spike MC-ICP-MS. *Chem. Geol.* **249**, 294–306 (2008).

27. Xia, J. et al. Chromium isotope heterogeneity in the mantle. *Earth Planet. Sci. Lett.* **464**, 103–115 (2017).

28. Bonnand, P., Parkinson, I. J. & Anand, M. Mass dependent fractionation of stable chromium isotopes in mare basalts: implications for the formation and the differentiation of the Moon. *Geochim. Cosmochim. Acta* **175**, 208–221 (2016).

29. Moynier, F., Yin, Q.-Z. & Schauble, E. Isotopic evidence of Cr partitioning into Earth's core. *Science* **331**, 1417–1420 (2011).

30. Schiller, M., Van Kooten, E., Holst, J. C., Olsen, M. B. & Bizzarro, M. Precise measurement of chromium isotopes by MC-ICP-MS. *J. Anal. At. Spectrom.* **29**, 1406–1416 (2014).

31. Bonnand, P., Williams, H. M., Parkinson, I. J., Wood, B. J. & Halliday, A. N. Stable chromium isotopic composition of meteorites and metal-silicate experiments: implications for fractionation during core formation. *Earth Planet. Sci. Lett.* **435**, 14–21 (2016).

32. Schoenberg, R. et al. The stable Cr isotopic compositions of chondrites and silicate planetary reservoirs. *Geochim. Cosmochim. Acta* **183**, 14–30 (2016).

33. Birck, J.-L. & Allègre, C. J. in *Isotopic Ratios in the Solar System* 21–25 (Cepadues-Editions, Toulouse,1985).

34. Shima, M. & Honda, M. Distribution of spallation produced chromium between alloys in iron meteorites. *Earth Planet. Sci. Lett.* **1**, 65–74 (1966).

35. Qin, L., Alexander, C. M. O., Carlson, R. W., Horan, M. F. & Yokoyama, T. Contributors to chromium isotope variation of meteorites. *Geochim. Cosmochim. Acta* **74**, 1122–1145 (2010).

36. Clayton, R. & Mayeda, T. Oxygen isotope studies of achondrites. *Geochim. Cosmochim. Acta* **60**, 1999–2017 (1996).

37. Trinquier, A., Birck, J.-L. & Allegre, C. J. Widespread Cr-54 heterogeneity in the inner solar system. *Astrophys. J.* **655**, 1179–1185 (2007).

38. Wasson, J., Lange, D., Francis, C. & Ulff-Moller, F. Massive chromite in the Brenham pallasite and the fractionation of Cr during the crystallization of asteroidal cores. *Geochim. Cosmochim. Acta* **63**, 1219–1232 (1999).

39. Isa, J., Ma, C. & Rubin, A. E. Joegoldsteinite: a new sulfide mineral (MnCr₂S₄) from the Social Circle IVA iron meteorite. *Am. Mineral.* **101**, 1217–1221 (2016).

40. Farkas, J. et al. Chromium isotope variations (δ⁵³/⁵²Cr) in mantle-derived sources and their weathering products: implications for environmental studies and the evolution of δ⁵³/⁵²Cr in the Earth's mantle over geologic time. *Geochim. Cosmochim. Acta* **123**, 74–92 (2013).

## Acknowledgements

This research was supported by STFC grant ST/M001318/1. We thank M. Boyet and A. Bouhifd for commenting on an earlier version of the manuscript. We also thank the Natural History Museum (London) and Smithsonian for access to samples.

## Author contributions

P.B. performed all the measurements. P.B. and A.N.H. wrote the manuscript.

## Competing interests

The authors declare no competing interests.

## Additional information

*Supplementary information* is available for this paper at https://doi.org/10.1038/s41561-018-0128-2.

Reprints and permissions information is available at www.nature.com/reprints.

Correspondence and requests for materials should be addressed to P.B.

Publisher's note: Springer Nature remains neutral with regard to jurisdictional claims in published maps and institutional affiliations.
## Methods

**Sample preparation.** Up to 3 g of iron meteorite sample was dissolved in 6 M HCl. The fully dissolved sample was evaporated to dryness and then re-dissolved in 6 M HCl. Two aliquots of this solution, each containing ~2 µg of chromium, were dried and dissolved in 7 M HCl. To obtain accurate mass-dependent chromium isotopic compositions, one aliquot was analysed for its radiogenic and nucleosynthetic chromium isotopic composition. The other aliquot was spiked with the requisite amount of double spike ($^{50}$Cr and $^{54}$Cr) and analysed for its mass-dependent chromium isotopic composition. Spiking occurred before any separation chemistry and the unspiked and spiked samples were processed in the same way. The chromium separation protocol involves three columns and has already been described in detail$^{28,31,41}$. A summary of the method used in this study is given below.

The first column allowed the complete removal of iron from the samples. A disposable Bio-Rad Poly-Prep column was filled with up to 8 ml of AG1 X8 and the sample was loaded in 7 M HCl. The chromium was eluted, leaving iron on the column. After this first column, the samples were dried down and dissolved in 6 M HCl. The second column separated chromium from the main cations. Bio-Rad Poly-Prep 10 ml columns were filled with 1 ml of AG50 X8 (200–400 mesh chloride-form resin) and the samples were loaded and eluted in 0.5 M HCl. The third column was a clean-up column to ensure that the chromium fraction was free of isobaric interferences (titanium, vanadium and iron). The chromium fraction was dried down and dissolved in 0.5 M HNO$_3$. A Bio-Rad Poly-Prep column was filled with 1 ml of AG50 X8 200–400 mesh resin. The sample was loaded in 4 ml 0.5 M HNO$_3$. The column was then washed with 7 ml 0.5 M HCl and chromium was eluted with 6 ml 3 M HCl. The total procedural yield was about 80%. The procedural blank was 0.25 ng, which is negligible compared with the amount (2 µg) of chromium processed through chemistry.

**Mass spectrometry.** Chromium isotopic compositions were determined using a Thermo Fisher Scientific Triton thermal ionisation mass spectrometer in static multicollection mode at the University of Oxford (United Kingdom). Chromium samples were loaded onto outgassed, zone-refined Re filaments in 1 µl 6 M HCl, along with 3 µl of a mixture of silicic and boric acids. Amplifier gains were measured daily and the baseline was measured after each block of cycles, with amplifier rotation after each block. Samples and standards were run with a typical beam current of 6 × 10$^{-11}$ A of $^{52}$Cr for 54 blocks of 10 cycles (integration time: 8.4 s). Interferences were monitored on $^{56}$Fe, $^{49}$Ti and $^{51}$V. Corrections for isobaric interferences were always negligible. For unspiked runs, chromium isotope ratios were corrected for mass fractionation using an exponential law and $^{50}$Cr/$^{52}$Cr = 0.051859 (ref. $^{42}$). Radiogenic and nucleosynthetic variations are reported using the equation: $\epsilon^{52}Cr = (^{52}Cr/^{52}Cr_{sample}/^{52}Cr/^{52}Cr_{NIST979} - 1) \times 10,000$ where $^{52}Cr/^{52}Cr_{sample}$ is the mass fractionation corrected isotope ratios ($^{52}Cr$ is either $^{53}$Cr or $^{54}$Cr).

The accuracies for $\epsilon^{53}$Cr and $\epsilon^{54}$Cr were assessed as 2 standard deviations (s.d.) of NIST 979 processed through chemical purification, and corresponded to 0.00 ± 0.10 and 0.00 ± 0.19, respectively (2 s.d., $n = 15$)$^{31}$. The external reproducibility for $\epsilon^{53}$Cr and $\epsilon^{54}$Cr was assessed by 2 s.d. of the measurements of the Arispe iron meteorite sample, the values being 17.71 ± 0.15 and 69.98 ± 0.69, respectively (2 s.d., $n = 10$; Supplementary Table 1). The accuracy of the mass-dependent variation ($\delta^{53}$Cr) was assessed by repeat measurements of a NIST SRM979 chromium standard and was $-0.005 \pm 0.009 \, \text{‰}$ (2 s.d., $n = 15$)$^{31}$. The external reproducibility was assessed with repeat measurements of the Arispe sample to be 2.85 ± 0.10 ‰ (2 s.d., $n = 10$; Supplementary Table 1). The uncertainties associated with the mass-dependent $\delta^{53}$Cr reported in Supplementary Table 1 incorporated errors propagated from the unspiked and spiked analyses, as well as the uncertainties of the spallation correction described below.

The mass-dependent chromium isotope deconvolution procedure used in this study allows the correction for mass fractionation that occurs during chemical separation and mass spectrometry. As described in refs $^{31,32}$, extraterrestrial samples do not lie on a single mass fractionation trend in four-isotope space such that we need to have two isotope measurements (unspiked and spiked) to derive the stable isotope composition of chromium isotopes. For the iron meteorite samples, the unspiked runs were internally normalized using an exponential law to the canonical $^{50}$Cr/$^{52}$Cr ratio of 0.051859 (ref. $^{42}$) to reveal any mass-independent $\epsilon^{53}$Cr and $\epsilon^{54}$Cr variations (Supplementary Table 1). The unspiked run was then combined with the spiked run for each sample in the double-spike calculation and this allowed determination of the mass-dependent chromium stable isotope composition. However, it has been demonstrated that in iron-rich samples chromium isotopes can be strongly affected by spallation, with production rates for $^{50}$Cr, $^{52}$Cr, $^{53}$Cr and $^{54}$Cr in the proportions 0.2:1:1:1, respectively$^{33,35}$. This indicates

that during internal normalization of the unspiked runs, an artificial bias is created. To correct for this bias, the mass-dependent values obtained after double-spike deconvolution were corrected assuming a production rate for the four chromium isotopes of 0.2:1:1:1. We used the following equation.

$$ \delta^{53}Cr_{corr} = \delta^{53}Cr - (\epsilon^{53}Cr \times -0.00152 + 0.0004) $$

where $\delta^{53}$Cr is the value obtained after deconvolution of the spiked run associated with the mass-bias-corrected unspiked run ($\epsilon^{53}$Cr). The equation above was determined analytically using the International Union of Pure and Applied Chemistry natural abundancies for chromium isotopes ($^{50}$Cr: 0.043452; $^{52}$Cr: 0.837895; $^{53}$Cr: 0.095006; and $^{54}$Cr: 0.023647). The spallation effect was then modelled by adding the required amount for each isotope using the production rate proportions of 0.2:1:1:1. Each run was then fractionated to simulate an unspiked measurement and deconvolved by internal normalization using an exponential law and the $^{50}$Cr/$^{52}$Cr ratio of 0.051859 to obtain an $\epsilon^{53}$Cr value. Using a Monte Carlo simulation, a double-spike run was then simulated and deconvolved using the previously determined $\epsilon^{53}$Cr value. Three different calculations were performed with the double-spike runs (Supplementary Table 2); one without accounting for the spallation effects (column a), one using the unfractionated chromium isotopic composition after spallation effects were added (column b) and one using the mass-bias-corrected $\epsilon^{53}$Cr value (column c). It is clear that there is an artificial bias during the double-spike deconvolution procedure (Supplementary Fig. 1). The $\delta^{53}$Cr values obtained after correction for spallation effects are strongly dependent on the production rates for each isotope used in the calculations. Using a production rate of 0.3:1:1:1 instead of 0.2:1:1:1 affects the samples with large spallation effects. For all IVA samples except Altonah (0.12 ‰), the error introduced by a such an incorrect production rate changes the values by less than 0.05 ‰, as illustrated in Supplementary Fig. 1. This is small compared with the large variations observed in the samples.

The uncertainties associated with the measurements also account for the fact that an exponential fractionation law is assumed for our samples. Exponential laws are used to describe kinetic fractionation, whereas equilibrium fractionation follows a power law. This produces an inaccuracy of $-0.009 \, \text{‰}$ per 1 ‰ on $\delta^{53}$Cr (refs $^{26,41}$). Propagation of all these uncertainties, together with the analytical uncertainty, indicates that the external reproducibility of the iron meteorite measurements for $\delta^{53}$Cr is 0.1 ‰ (2 s.d.), which is still small compared with the observed variations.

**Major and trace elements.** Major and trace element concentrations were determined by inductively coupled plasma mass spectrometry (ELEMENT; Thermo Fisher Scientific) at the University of Oxford. The instrument was fitted with an Aridus 1 and a 100 µl min$^{-1}$ nebulizer. Major elements analyses were externally calibrated against mixed synthetic solutions. The concentrations ranged from 100 ng g$^{-1}$ to 5 µg g$^{-1}$ for iron, 1 ng g$^{-1}$ to 100 ng g$^{-1}$ for cobalt, and 10 ng g$^{-1}$ to 1 µg g$^{-1}$ for nickel. Trace element analyses were standardized using the standard addition method at the beginning of each analytical run. To this end, synthetic solutions were prepared in a clean laboratory using ultrapure multi elemental standard solutions. The standard solution was then added to one sample in five concentrations (1 pg g$^{-1}$, 10 pg g$^{-1}$, 100 pg g$^{-1}$, 1 ng g$^{-1}$ and 10 ng g$^{-1}$). The final solutions were diluted to 10 µg g$^{-1}$ iron in solution. The level of iron was kept as constant as possible to avoid matrix effects. Limited loss of signal due to high total dissolved solids was observed during the run. The external reproducibility was assessed with repeat measurements of Arispe and was about 5%.

**Data availability.** The data supporting the findings of this study are available in the Supplementary Information.

## References

41. Bonnand, P., James, R. H., Parkinson, I. J., Connelly, D. P. & Fairchild, I. J. The chromium isotopic composition of seawater and marine carbonates. *Earth Planet. Sci. Lett.* **382**, 10–20 (2013).

42. Shields, W., Murphy, T., Catanzaro, E. & Garner, E. Absolute isotopic abundance ratios and atomic weight of a reference sample of chromium. *J. Res. NBS A Phys. Chem.* **70**, 193–197 (1966).

43. Bonnand, P., Parkinson, I. J., James, R. H., Karjalainen, A.-M. & Fehr, M. A. Accurate and precise determination of stable Cr isotope compositions in carbonates by double spike MC-ICP-MS. *J. Anal. At. Spectrom.* **26**, 528–535 (2011).