# Silicon in iron meteorite metal

Andreas Pack, Ingo Vogel, Claire Rollion-Bard, Béatrice Luais, Herbert Palme

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Andreas Pack, Ingo Vogel, Claire Rollion-Bard, Béatrice Luais, Herbert Palme. Silicon in iron meteorite metal. Meteoritics and Planetary Science, 2011, 46 (10), pp.1470-1483. 10.1111/j.1945-5100.2011.01239.x. insu-03619289

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# Silicon in iron meteorite metal

Andreas PACK<fnref n="1" />,<fnref n="2" />,<fnref n="3" />*, Ingo VOGEL<fnref n="4" />, Claire ROLLION-BARD<fnref n="1" />, Béatrice LUAIS<fnref n="1" />, and Herbert PALME<fnref n="4" />,<fnref n="5" />

*(Received 08 November 2010; revision accepted 28 June 2011)*

**Abstract**–We report Si concentrations in the metal phases of iron meteorites. Analyses were performed by secondary ion mass spectrometry using a CAMECA 1270 ion probe. The Si concentrations are low (0.09–0.46 µg g⁻<fnref n="1" />), with no apparent difference in concentration between magmatic and nonmagmatic iron meteorites. Coexisting kamacite and Ni-rich metal phases have similar Si contents. Thermodynamic calculations show that Fe,Ni-metal in equilibrium with silicate melts at temperatures where metal crystallizes should contain approximately 100 times more Si than found in iron meteorites in this work. The missing Si may either occur as tiny silicate inclusions in metal or it may have diffused as Si-metal into surrounding silicates at low temperatures. In both cases, extensive low-temperature diffusion of Si in metal is required. It is therefore concluded that low Si in iron meteorites is a result of subsolidus reactions during slow cooling.

## INTRODUCTION

### Iron Meteorites

Iron meteorites are predominantly composed of Fe,Ni alloys with minor inclusions of troilite (FeS), schreibersite (Fe,Ni)₃P, graphite, cohenite (FeNi)₃C and/or silicates. Based on trace element systematics, iron meteorites are classified as magmatic or nonmagmatic (Wasson et al. 1980; Wasson and Wang 1986; Goldstein et al. 2009). Magmatic iron meteorites (Wasson 1985; Goldstein et al. 2009 suggested the term “fractionally crystallized”) have complex trace element patterns, which can best be explained by fractional crystallization of slowly cooling liquid metal masses including variable amounts of S in cores of approximately 5–1000 km sized planetesimals (Goldstein et al. 2009). Nonmagmatic iron meteorites (groups IAB [with subgroups, see Wasson and Kallmeyn 2002] and IIE) have largely unfractionated trace element patterns and have not experienced

fractional crystallization (Wasson et al. 1980; Wasson and Wang 1986; Choi et al. 1995; Goldstein et al. 2009). Goldstein et al. (2009) suggested the term “silicate-bearing” for these groups. In this contribution, we use the traditional terms “magmatic” and “nonmagmatic.”

The thermal history of iron meteorites can be derived from the metal structure. Metallographic cooling rates (i.e., subsolidus cooling rates) between <10 and 17000 K Ma⁻<fnref n="1" /> have been proposed (Goldstein et al. 2009; therein table 2). Advances in the models have led to higher cooling rates. The magmatic IIIAB iron meteorites are suggested to have cooled at 60–340 K Ma⁻<fnref n="1" /> (Yang and Goldstein 2006). The high cooling rate is suggested to be the result of a partial loss of the insulating silicate mantle of the IIIAB parent body. Older data for IIAB suggested 0.8–10 K Ma⁻<fnref n="1" /> (Randich and Goldstein 1978). Application of more advanced models for the metal structure, however, may lead to 1–2 orders of magnitude higher cooling rates (Goldstein et al. 2009). The highest cooling rates were

<footnote n="1">¹Centre de Recherches Pétrographiques et Géochimiques (CRPG-CNRS), 15 rue Notre Dame des Pauvres, F-54501 Vandœuvre-lès-Nancy, France</footnote>

<footnote n="2">²Institut für Mineralogie, Universität Hannover, Callinstraße 3, D-30167 Hannover, Germany</footnote>

<footnote n="3">³Geowissenschaftliches Zentrum, Abteilung Isotopengeologie, Georg-August-Universität, Goldschmidtstraße 1, D-37077 Göttingen, Germany</footnote>

<footnote n="4">⁴Insitut für Geologie und Mineralogie, Universität Köln, Zülpicher Straße 49b, D-50674 Köln, Germany</footnote>

<footnote n="5">⁵Senckenberg Forschungsinstitut und Naturmuseum, Senckenberganlage 25, D-60325 Frankfurt am Main, Germany</footnote>

<footnote>*Corresponding author. E-mail: apack@uni-goettingen.de</footnote>
observed in IVA and IVB iron meteorites. Their parent asteroids may have retained only a very thin insulating silicate mantle during solidification (Asphaug et al. 2006; Yang et al. 2007).

## Si in Iron Meteorites and Its Relationship with Redox State

The first report of Si concentrations in iron meteorites is from Goldschmidt (1937) on three iron meteorites with values between 20 and 50 µg g⁻¹ (see Wai and Wasson 1969). Later and using the electron microprobe, Wai and Wasson (1969) reported Si measurements on 19 iron meteorites with 17 meteorites having <30 µg g⁻¹.

Two iron meteorites have 8000 µg g⁻¹ Si (Tucson, ungrouped, 9.45 wt% Ni) and 25,000 µg g⁻¹ Si (Horse Creek, ungrouped, 3.8 wt% Ni). In extension to their work, Wai and Wasson (1970) reported 1400 µg g⁻¹ Si in the Nedagolla ungrouped Ni-rich iron meteorite. Nedagolla was suggested to have been remelted after its formation (Miyake and Goldstein 1974). No clear relationship between Cr and Si concentration in iron meteorite metal was observed. Such a correlation is expected as the concentrations of both elements would increase in parallel with decreasing fO₂. A relationship between fO₂ and Si concentration is obvious in the Tucson iron meteorite in which Si-rich metal (0.8 wt%) occurs along with highly reduced silicates (olivine with fO₀.₀₈–0.₂₂, and orthopyroxene with fS₀.₄, [Wai and Wasson 1969; Nehru et al. 1982]). Highly reduced silicates and Si-rich metal are also known from enstatite chondrites. A relationship of the Horse Creek meteorite with enstatite chondrites was therefore proposed (Wai and Wasson 1969; Nehru et al. 1982).

## Aim of This Study

In conventional models, it is assumed that iron meteorites formed by gravitational separation of immiscible liquid metal or metal-sulfide in partially or fully molten planetesimals at temperatures exceeding approximately 1570 K. If thermodynamic equilibrium between metal and silicates is assumed during separation, some Si must have entered the metal phase. The amount of Si partitioning into metal under equilibrium conditions can be calculated, if the relevant thermodynamic data on the stability of SiO₂ in silicates and the activity coefficients of Si in Fe,Ni alloys are known. Data on the activities of Si in Fe,Ni alloys as function of temperature, oxygen fugacity and composition have been determined experimentally by Vogel et al. (2006).

By using these data, Si concentrations expected from high-temperature metal-silicate equilibration can be

calculated and compared with the Si contents measured in the metal phases of iron meteorites as reported in this article. However, as iron meteorites have undergone prolonged annealing at low temperatures developing coarse-grained textures and subsolidus exsolution of kamacite from taenite, subsolidus redistribution of Si between coexisting metal phases and/or silicates may have also affected the Si distribution in planetary cores.

## ANALYTICAL PROCEDURES

### Sample Preparation, Backscattered Electron, and X-Ray Imaging

All iron meteorite samples were ground and polished with Si-free materials. Alumina sand paper was used for grinding. The final polish was conducted with 3 and 1 µm diamond suspensions. Samples were carefully cleaned with ethanol before they were coated with C for X-ray mapping and electron probe microanalyses. Backscattered electron (BSE) imaging was conducted to search for silicate inclusions. Nickel-Kα X-ray element distribution maps were obtained using a JEOL 8900RL electron microprobe (University of Cologne). These maps were used for selecting areas for analysis and for phase identification (low-Ni kamacite, Ni-rich regions [former taenite], Ni-free troilite and Ni-rich schreibersite).

### Electron Microprobe Analyses

The chemical composition of metal, sulfide, and phosphide was analyzed using a JEOL 8900RL electron microprobe (University of Göttingen). The probe was operated with a current of 20 nA and an acceleration voltage of 15 kV. We used pure metal standards for Fe, Co, and Ni; apatite for P; and sphalerite for S. The peaks (Kα) were analyzed for 20 s and the backgrounds for 10 s on each side of the peaks. The Fe concentrations were used as internal standard for the ion probe measurements. Sulfides and phosphides were analyzed for phase identification only.

### Secondary Ion Mass Spectrometry

After electron microprobe work, we have removed the C coating by polishing with 1 µm diamond suspension. Samples were cleaned and coated with Au for Secondary ion mass spectrometry (SIMS) analyses. Concentrations of Si were determined by means of SIMS. Analyses were performed using the CAMECA 1270 ion microprobe (CRPG, Nancy). We have analyzed Si and Fe on ²⁸Si⁺ and ⁵⁴Fe⁺ peaks at a mass resolution of *m*/Δ*m* ≈ 3500. Separation of ⁵⁶Fe⁺⁺ from ²⁸Si⁺
requires a mass resolution of approximately 2950. Intensities of $^{28}$Si$^{+}$ were normalized to the intensity of $^{54}$Fe$^{+}$. Count rates on $^{28}$Si$^{+}$ in iron meteorite metal were typically approximately 10 cps µg$^{-1}$ Si (cps: counts per second). The SIMS background was negligible. The spot sizes were 15–20 µm. The relative sensitivity factor for $^{28}$Si/$^{54}$Fe was determined on standard metals (Fig. 1). We used three steel standards of the Bundesanstalt für Materialforschung (BAM, Berlin, Germany): 184-1 (0.218 wt% Si, 3.32 wt% Ni), 096-1 (0.263 wt% Si, 0.02 wt% Ni), and 182-1 (0.368 wt% Si, 0.152 wt% Ni). These standards were complemented by two Fe,Ni alloys from Vogel (2005): 030203B (0.082 wt% Si, 37.3 wt% Ni) and 120303B (0.055 wt% Si, 10 wt% Ni). These samples were prepared by melting appropriate metal mixtures at reducing conditions. We have chosen alloys with low Si concentrations (<0.5 wt%), but variable Fe/Ni-ratios, to study the influence of Ni content (matrix) on the sensitivity factor for Si.

The relative sensitivity factor for Si is within error independent of the Fe/Ni-ratio, at least within the range of Ni contents of iron meteorites, from 0 to 37 wt% (Fig. 1). The intercept of the regression line is within error zero (−0.001 ± 0.004). The errors were calculated using Gaussian error propagation including counting errors and the uncertainty intrinsic to the calibration line. We report 2σ errors throughout the manuscript. Each analysis was carefully checked for surface contamination. We identified contamination by a systematic decrease in $^{28}$Si/$^{54}$Fe-ratios with time and a linear correlation between the $^{27}$Al/$^{54}$Fe- and the $^{28}$Si/$^{54}$Fe-ratios during sputtering. Contaminated samples were excluded from the data set as they would give erroneously high Si concentrations. The source of the surface contamination remains unknown. Although we assume that the data reported here are accurate, small amounts of surface contamination (e.g., from the rims of the sputtering pits) may have remained unidentified on basis of the stability of the $^{28}$Si/$^{54}$Fe- and $^{27}$Al/$^{54}$Fe-ratios. Certainly, however, concentrations were not higher than reported in this study.

## Thermodynamic Calculation of Silicate Partial Melts

The compositions of partial mantle melts were calculated using MELTS (Ghiorso and Sack 1995; Asimov and Ghiorso 1998) release 5.0 (MacBook Pro, OS X 10.6). For sake of simplification, it is referred to MELTS only without reference to the original literature.

## SAMPLES

We have studied the Si concentrations on polished sections of 15 iron meteorites. Magmatic iron meteorites

<figure><img src="image_2.png" /><figcaption>Fig. 1. Plot of the measured $^{28}$Si/$^{54}$Fe intensity ratios versus Si/Fe concentration ratios in Fe-based alloys. Ni concentrations are indicated for each sample (circles: BAM standards 184-1, 096-1, 182-1; squares: secondary reference samples 120303B, 030203B). The slope is a measure for the relative sensitivity factor with an uncertainty of about 15%. The error envelope has been calculated using ISOPLOT (Ludwig 2003).</figcaption></figure>

include Guadalupe y Calvo, Sao Juliao de Moreira, Walker County, San Martin specimen of North Chile (IIABs), Perryville, Ballinoo (IICs), and Cape York (Agpalilik specimen; IIIBA). The nonmagmatic iron meteorites include Landes, Toluca, Canyon Diablo, Seeläsgen (IABs), Watson, Mont Dieu, Miles (IIEs). Wiley is now classified as an ungrouped iron meteorite (Wasson, personal communication). Except for San Martin, we studied the same samples as Luais (2007).

## RESULTS

The results are displayed in Table 1. Except for Landes, no silicate inclusions have been observed by BSE imaging and reflected light microscopy in the studied sections. The size of the ion microprobe spots did not allow obtaining unequivocal taenite analyses. The Ni-rich regions, however, were likely submicron intergrowths of metal phases (Goldstein et al. 2009). The Si concentrations in magmatic and nonmagmatic iron meteorites vary within a narrow range, between 0.09 and 0.31 µg g$^{-1}$ (Fig. 2). The lowest Si concentration is observed in Seeläsgen (IAB). The highest concentration (0.46 µg g$^{-1}$) is measured in Wiley (ungrouped). No systematic difference in Si concentration is found between kamacite and Ni-rich metal regions (see data for Cape York and Toluca in Table 1). All Si data obtained for a single magmatic iron meteorite are within the error intervals of the mean value. A slightly larger scatter is observed for nonmagmatic iron meteorites and the ungrouped iron meteorite Wiley.
<table><caption>Table 1. Results of EPMA (Fe, Co, Ni) and SIMS (Si) analyses on 15 magmatic and nonmagmatic iron meteorites. Major elements are used for phase identification. For schreibersite, we have used the metal calibration line. Therefore, the Si number given has to be regarded as semiquantitative.</caption><thead><tr><th>Meteorite</th><th>Type</th><th>Text.</th><th>Phase</th><th>Analysis</th><th>Fe<br>wt%</th><th>Co<br>wt%</th><th>Ni<br>wt%</th><th>Si<br>µg g<sup>-1</sup> 2σ</th></tr></thead><tbody><tr><td rowspan="2">Magmatic iron meteorites<br>Guadalupe y Calvo</td><td rowspan="2">IIAB</td><td rowspan="2">H</td><td rowspan="2">kam</td><td>@1</td><td>94.1</td><td>0.45</td><td>5.2</td><td>0.23 ± 0.05</td></tr><tr><td>@2</td><td></td><td></td><td></td><td>0.16 ± 0.02</td></tr><tr><td rowspan="2">Sao Juliao de Moreira</td><td rowspan="2">IIAB</td><td rowspan="2">Ogg</td><td rowspan="2">kam</td><td>@1</td><td>94.2</td><td>0.51</td><td>5.5</td><td>0.19 ± 0.02</td></tr><tr><td>@2</td><td></td><td></td><td></td><td>0.16 ± 0.02</td></tr><tr><td rowspan="2">North Chile (San Martin)</td><td rowspan="2">IIAB</td><td rowspan="2">H</td><td rowspan="2">kam</td><td>@2</td><td>93.9</td><td>0.45</td><td>5.6</td><td>0.24 ± 0.04</td></tr><tr><td>@4</td><td></td><td></td><td></td><td>0.24 ± 0.02</td></tr><tr><td rowspan="2">Walker County</td><td rowspan="2">IIAB</td><td rowspan="2">H</td><td rowspan="2">kam</td><td>@1</td><td>93.5</td><td>0.46</td><td>5.3</td><td>0.16 ± 0.05</td></tr><tr><td>@2</td><td></td><td></td><td></td><td>0.12 ± 0.01</td></tr><tr><td rowspan="4">Perryville</td><td rowspan="4">IIC</td><td rowspan="4">Opl</td><td>kam</td><td>@1</td><td>91.6</td><td>0.59</td><td>7.1</td><td>0.29 ± 0.04</td></tr><tr><td>kam/Ni-rich</td><td>@2</td><td></td><td></td><td>0.31 ± 0.04</td></tr><tr><td>kam/Ni-rich</td><td>@3</td><td></td><td></td><td>0.31 ± 0.03</td></tr><tr><td>kam</td><td>@4</td><td></td><td></td><td>0.22 ± 0.02</td></tr><tr><td rowspan="3">Ballinoo</td><td rowspan="3">IIC</td><td rowspan="3">Opl</td><td>kam</td><td>@1</td><td>91.8</td><td>0.61</td><td>7.0</td><td>0.22 ± 0.03</td></tr><tr><td>kam</td><td>@2</td><td></td><td></td><td>0.17 ± 0.02</td></tr><tr><td>kam/Ni-rich</td><td>@3</td><td></td><td></td><td>0.17 ± 0.02</td></tr><tr><td rowspan="2">Wiley</td><td rowspan="2">ungr.</td><td rowspan="2">Opl</td><td rowspan="2">kam</td><td>@1</td><td>91.2</td><td>0.87</td><td>7.2</td><td>0.31 ± 0.03</td></tr><tr><td>@3</td><td></td><td></td><td></td><td>0.46 ± 0.04</td></tr><tr><td rowspan="4">Cape York (Agpalilik specimen)</td><td rowspan="4">IIIAB</td><td rowspan="4">Om</td><td>Ni-rich</td><td>@1</td><td>82.1</td><td>0.35</td><td>17.0</td><td>0.11 ± 0.03</td></tr><tr><td>kam</td><td>@2</td><td>92.0</td><td>0.54</td><td>7.1</td><td>0.15 ± 0.01</td></tr><tr><td>kam</td><td>@4</td><td>92.0</td><td>0.54</td><td>7.1</td><td>0.14 ± 0.02</td></tr><tr><td>schreib</td><td>@3</td><td>51.4</td><td>0.11</td><td>32.9</td><td>1.78 ± 0.08</td></tr><tr><td rowspan="4">Nonmagmatic iron meteorites<br>Landes</td><td rowspan="4">IAB</td><td rowspan="4">O</td><td rowspan="4">kam</td><td>@2</td><td>92.3</td><td>0.46</td><td>6.3</td><td>0.27 ± 0.04</td></tr><tr><td>@3</td><td></td><td></td><td></td><td>0.20 ± 0.02</td></tr><tr><td>@4</td><td></td><td></td><td></td><td>0.10 ± 0.01</td></tr><tr><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td rowspan="6">Toluca</td><td rowspan="6">IAB</td><td rowspan="6">Og</td><td>Ni-rich</td><td>@1</td><td>84.1</td><td>0.28</td><td>14.6</td><td>0.19 ± 0.02</td></tr><tr><td>Ni-rich</td><td>@2</td><td>80.2</td><td>0.24</td><td>18.2</td><td>0.16 ± 0.02</td></tr><tr><td>kam</td><td>@3</td><td>91.7</td><td>0.51</td><td>6.9</td><td>0.21 ± 0.02</td></tr><tr><td>kam</td><td>@4</td><td>91.7</td><td>0.51</td><td>6.9</td><td>0.16 ± 0.02</td></tr><tr><td>Ni-rich</td><td>@5</td><td>83.5</td><td>0.26</td><td>15.0</td><td>0.20 ± 0.03</td></tr><tr><td>Ni-rich</td><td>@6</td><td>84.1</td><td>0.28</td><td>14.6</td><td>0.17 ± 0.03</td></tr><tr><td rowspan="4">Canyon Diabolo</td><td rowspan="4">IAB</td><td rowspan="4">Og</td><td rowspan="4">kam</td><td>@7</td><td>91.7</td><td>0.51</td><td>6.9</td><td>0.11 ± 0.01</td></tr><tr><td>@2</td><td>92.5</td><td>0.51</td><td>6.5</td><td>0.18 ± 0.03</td></tr><tr><td>@3</td><td></td><td></td><td></td><td>0.19 ± 0.03</td></tr><tr><td>@4</td><td></td><td></td><td></td><td>0.29 ± 0.03</td></tr><tr><td></td><td></td><td></td><td></td><td>@7</td><td></td><td></td><td></td><td>0.10 ± 0.02</td></tr><tr><td rowspan="2">Seeläsgen<br>Miles</td><td rowspan="2">IAB</td><td rowspan="2">Og</td><td rowspan="2">kam</td><td>@1</td><td>92.8</td><td>0.47</td><td>6.4</td><td>0.09 ± 0.01</td></tr><tr><td>@1</td><td>92.5</td><td>0.46</td><td>6.8</td><td>0.10 ± 0.01</td></tr><tr><td rowspan="2">Watson</td><td rowspan="2">IIE</td><td rowspan="2">–</td><td rowspan="2">kam</td><td>@1</td><td>92.3</td><td>0.47</td><td>7.2</td><td>0.11 ± 0.01</td></tr><tr><td>@2</td><td></td><td></td><td></td><td>0.12 ± 0.02</td></tr><tr><td rowspan="2">Mont Dieu</td><td rowspan="2">IIE</td><td rowspan="2">–</td><td rowspan="2">kam</td><td>@1</td><td>91.6</td><td>0.46</td><td>7.3</td><td>0.10 ± 0.01</td></tr><tr><td>@2</td><td></td><td></td><td></td><td>0.12 ± 0.01</td></tr></tbody></table>

Co and Ni concentrations for the North Chile (San Martin) meteorite are from Wasson et al. (1989), Fe was calculated by difference.

kam = kamacite, Ni-rich = Ni-rich metal region, former taenite; schreib = schreibersite; H = hexahedrite; O = octahedrite (pl = plessitic; m = medium; g = coarse; gg = very coarse).

## DISCUSSION

### Comparison with Literature Data

Secondary ion microprobe is appropriate for the determination of Si concentrations in Fe,Ni-metal in the

lower µg g⁻¹ range. Our data are in agreement with earlier findings of <20 µg g⁻¹ Si in iron meteorite metal (Wai and Wasson 1969, 1970).

The Si concentrations are not correlated with the presence of silicate inclusions. Nonmagmatic iron meteorites contain abundant silicate inclusions, but show
<figure><img src="image_3.png" /><figcaption>Fig. 2. Variation diagrams showing the Si concentrations in metal of 15 different iron meteorites. (A) Magmatic iron meteorites and Wiley (ungrouped), (B) nonmagmatic iron meteorites. The errors are 2σ.</figcaption></figure>

the same Si concentration as magmatic iron meteorites
with and without silicate inclusions.

The data on kamacite and Ni-rich regions (former
taenite) from Cape York and Toluca (Table 1) suggest
that the partition coefficient of Si between kamacite and
taenite is about one. This conclusion is in agreement with
Raghavan (2003), who reported a Si α-Fe,Ni/γ-Fe,Ni
partitioning coefficient of 1.3 at 1273 K in Fe,Ni alloys
with 5–13 wt% Ni.

Thermodynamic Modeling

Temperature Range

To assess the significance of our data, we have
calculated the Si concentration in solid metal cores of
small planetesimals. We assume that the core metal
formed by crystallization from metal that segregated at
temperatures between 1773 and 1260 K. The upper limit
is the liquidus temperature of Fe,Ni alloys (Predel
2010a). The lower limit is the Fe,Ni-(Fe,Ni)S eutectic
temperature (Raghavan 2004; Predel 2010b).

Core Composition and Mass Balance

The composition of the bulk core can be expressed as
function of FeO in the silicate by mass balance, assuming
that the iron meteorite parent body had a composition of
anhydrous CI1 meteorites (Lodders 2003). We use a CI1
composition to start with solar relative proportions of the

major elements. The calculations would give a very
similar result with the composition of other chemically
primitive meteorites. For simplification, we consider only
(O), Na, Mg, Al, Si, (S), K, Ca, Ti, Cr, Mn. Iron is
distributed between core (Fe⁰) and mantle (FeO). The
amount of O varies as we vary the Fe/FeO-ratio of the
bulk model asteroid. Variable bulk FeO contents of the
mantle between 5 and 25 wt% are considered. Nickel is
assumed to occur only in the core, considering its very
high metal/silicate partition coefficient at low pressures
(Kegler et al. 2008). Sulfur is varied between 0 and
17.5 wt% in the bulk core. Sulfur is assumed to be
present in the liquid portion of the core. By mass balance,
we calculate the composition of the core as a function of
FeO in the mantle.

Structure of the Parent Body

In calculating the Si content in metal phase, we
consider two endmember models: (A) the equilibrium
model and (B) the internal oxidation model.

Case A: *Equilibrium Model*: In this model (Fig. 3A), solid metal forms by fractional crystallization from liquid metal. The instantaneously formed solid metal is in equilibrium with the liquid metal. The liquid metal, in turn, is in equilibrium with the surrounding silicate mantle. In this scenario, newly formed solid metal and silicate mantle are in thermodynamic equilibrium via the liquid metal. We will show below that the solid metal core cannot equilibrate with surrounding silicates on a scale larger than a few meters after crystallization is completed. This leads to a series of chemically zoned metals.

Case B: *Internal Oxidation Model*: This model (Fig. 3B) suggests that the first metal crystallizes at the core-mantle boundary as planetesimals cool from the outside. The solid metal does not sink toward the center of the planetesimal (as it does in case A). Instead, the solid metal forms a shell that separates the liquid metal from the silicate. After formation of a shell of solid metal, no further equilibration with silicates is possible. The time for completion of the fractional crystallization of the core is much shorter than the time for diffusion of Si and O through a solid metal layer. Thus, the evolving core behaves as a closed system. As in the equilibrium model, metal fractionally crystallizes from the liquid. The total amount of Si in the metal is determined by the initial liquid metal–liquid silicate equilibrium. The Si concentration of the solid metal at the time of crystallization is controlled by the solubility of Si in the liquid metal and the solid/liquid metal Si partitioning coefficient.

Thermodynamic Modeling: Case A (Equilibrium Model)

We first model the Si concentration in solid metal
that formed at the center of the metallic core (Fig. 3A).
<figure><img src="image_4.png" /><figcaption>Fig. 3. Sketch illustrating the parent body models (A: Case A and B: Case B).</figcaption></figure>

We calculated the oxygen fugacity ($f_{O_2}$) prevailing during metal-silicate equilibrium for two different cases. In the first case, we assumed that the mantle was entirely molten. In this case, the FeO content of the bulk mantle and the Fe content of the metal phase determine $f_{O_2}$. In the second case, we assume that the silicate mantle is only partially molten and that the $f_{O_2}$ is determined by the FeO content of the melt and the Fe concentration in the core metal. The composition of the partial mantle melt is calculated using the MELTS program (Fig. 4). With increasing degree of partial mantle melting the FeO content of the partial melt approaches the total silicate melt FeO-content. At least about 20% melting is required for effective core formation (Taylor 1992). At lower melt fractions, metal cannot effectively separate. In our system, 20% melting is reached at $T \ge 1430$ K (Fig. 4A). At 1790 K, between 55 and 75 wt% of the mantle is molten. The mantle reaches the solidus at approximately 1270 K.

The oxygen fugacity in the interior of the planetesimal is buffered by the Fe-FeO equilibrium according to the net reaction:

$$Fe_s^{metal} + \frac{1}{2}O_2 \rightleftharpoons FeO_1^{silicate} \quad (1)$$

with

$$K_1 = \frac{a_{FeO}^{silicate}}{a_{Fe}^{metal} \sqrt{f_{O_2}}} \quad (2)$$

with $a_i^j$ being the activity of component $i$ in phase $j$ and $a_i^j = X_i^j \gamma_i^j$. The reaction constant $K_1$ is identical to $K_f$ for formation of liquid FeO from the elements in the reference state (Chase 1998) with:

$$\log K_1 = \log K_f^{FeO_1} = 13.074 \times \frac{10^3}{T} - 2.6807 \quad (3)$$

The Fe activity coefficient $\gamma_{Fe}^{metal}$ for solid Fe,Ni-metal with > 90 wt% Fe is approximately one (Swartzendruber

<figure><img src="image_5.png" /></figure>

<figure><img src="image_6.png" /><figcaption>Fig. 4. (A) Plot of the melt fraction versus $T$ for bulk FeO in the mantle between 5 and 25 wt%. The dashed line indicates the minimum degree of melting (20%) that is required for metal segregation (Taylor 1992). (B) Plot of the FeO concentration in the partial mantle melt as function of temperature and for starting compositions with bulk silicate FeO contents from 5 to 25 wt% as indicated in the figure. At high melt fractions, partial melt FeO approaches bulk FeO. Melt compositions were calculated using MELTS ($p = 0.1$ kbar). The FeO contents (5–25 wt%) of the bulk mantle are indicated.</figcaption></figure>

et al. 1991). The activity coefficient $\gamma_{FeO}^{silicate}$ for silicate melts with a variety of compositions has been determined experimentally by Holzheid et al. (1997) to be $1.7 \pm 0.22$ (1σ). It is independent of the composition of the silicate liquid. Combining Equation 2 with Equation 3 gives the expression for the oxygen buffer:

$$\log f_{O_2} = -2 \underbrace{\log K_1}_{\text{Chase (1998)}} + 2\log \underbrace{\frac{\overbrace{X_{FeO}^{silicate}}^{\text{bulk mantle, MELTS}}}{\overbrace{X_{Fe}^{metal}}^{\text{mass balance}}} + \underbrace{0.46}_{\text{Holzheid et al. (1997)}} \quad (4)$$

At high temperatures (> 1600 K), both approaches (bulk FeO versus FeO in partial mantle melt) give a similar $f_{O_2}$ (Fig. 5). At $T = 1300$ K, equilibrium $f_{O_2}$ becomes 1.5–2 log-units lower for the case, in which we use the FeO concentration of the partial mantle melt to
calculate $f_{O_2}$ (Fig. 5). This is related to the decreasing FeO concentration in the mantle melt at $T < 1600$ K (Fig. 4). It has to be noted that the FeO concentrations in the residual melt at low degree of melting as given by MELTS are lower than experimental data (10 kbar) for peridotite melting (Baker et al. 1995; Hirschmann et al. 1998). The difference may arise from the difference in $p$ and uncertainties in the thermodynamics of MELTS. The resulting differences in the concentrations of Si in metals are small and have no effect on the general conclusions of the study.

The partitioning of Si between silicate and metal is expressed in the form of reaction:

$$Si_s^{metal} + O_2 \rightleftharpoons SiO_2^{silicate} \quad (5)$$

with

$$K_2 = \frac{a_{SiO_2}^{silicate}}{X_{Si}^{metal} \gamma_{Si}^{metal} f_{O_2}} \quad (6)$$

The value of $K_2$ is identical to $K_f$, which is the tabulated equilibrium constant for the formation reaction of silica from the elements in their reference states. We adopt the values for $K_f$ for liquid silica from Chase (1998) with:

$$\log K_2 = \log K_f^{SiO_2} = 47.713 \times \frac{10^3}{T} - 9.4696 \quad (7)$$

Equations 5 and 6 show that the solubility of Si in the metal is a function of the Si activity coefficient $\gamma_{Si}^{metal}$, oxygen fugacity $f_{O_2}$, and silica activity $\gamma_{SiO_2}^{silicate} \times X_{SiO_2}^{silicate}$ in the mantle.

We assume that the silica activity in the silicate melt is constant at 0.5. MELTS gives silica activities in the range of $0.3 \le a_{SiO_2}^{silicate} \le 0.6$ with most data lying between 0.4 and 0.5. Combining Equations 6 and 4 and inserting the Gibbs equation for $K_2$ in Equation 6 gives

$$\log X_{Si}^{metal} = 2\log K_1 - \log K_2 + 2\log \frac{X_{Fe}^{metal}}{X_{FeO}^{silicate}} - \underbrace{2\log\gamma_{FeO}^{silicate}}_{0.46} - \log\gamma_{Si}^{metal} + \log a_{SiO_2}^{silicate} \quad (8)$$

The last unknown (apart from $X_{Si}^{metal}$) is the activity coefficient $\gamma_{Si}^{metal}$ of Si in Fe,Ni-metal. It has been experimentally determined by Vogel et al. (2005, 2006) that $\gamma_{Si}^{metal}$ is a function of temperature and Ni content of the metal. Experiments were conducted for metal compositions between 0 and 100 wt% Ni and a temperature range of 1523–1673 K. Regression analysis of all experiments gives

$$\log\gamma_{Si}^{metal} = -0.17 - \frac{5193K}{T} - 1.22X_{Si}^{metal} \quad (9)$$

<figure><img src="image_7.png" /></figure>

<figure><img src="image_8.png" /><figcaption>Fig. 5. Plot of the calculated $f_{O_2}$ (on log-scale relative to the iron/wüstite [IW] buffer) at the metal to silicate interface versus $T$, with silicates either completely (A) or partially (B) molten. As in Fig. 4, in both cases, the FeO percentages indicate the FeO content of the bulk mantle.</figcaption></figure>

The uncertainty in $\gamma_{Si}^{metal}$ is within the range of ±15%. As an example, $\gamma_{Si}^{metal}$ is 0.00045 ($T = 1673$ K, $X_{Ni} = 0.06$). The combination of Equation 8 with Equation 9 allows calculating the concentration of Si in solid Fe,Ni-metal that coexists with an entirely (Fig. 6A) and a partially molten silicate mantle (Fig. 6B).

At a crystallization temperature of 1773 K, solid Fe,Ni-metal should contain between approximately 6 and approximately 300 µg g⁻¹ Si (Fig. 6). Low Si is expected for oxidized parent bodies (i.e., high FeO in the mantle) and high Si for reduced parent bodies (i.e., low FeO in the mantle). The equilibrium Si concentration decreases with decreasing $T$. The decrease in Si with decreasing $T$ is more pronounced for models with FeO of the bulk mantle composition (Fig. 6A). As asteroidal crust (e.g., eucrites) and mantle (ureilites) rocks contain considerable amounts of FeO, iron meteorites probably contained much less than 300 µg g⁻¹ Si when they crystallized. Decrease in crystallization temperature due to S, P, and C in the host metal has, in addition, lowered
<figure><img src="image_9.png" /><figcaption>Fig. 6. Plot of the Si concentration in solid Fe,Ni-metal in equilibrium with silicates for a CI1 chondrite parent body with variable FeO in the bulk silicate fraction (FeO concentrations are indicated). The gray shaded area outlines the observed Si concentrations in iron meteorite metal. (A) The FeO concentration is fixed at 5, 10, 15, 20, and 25 wt%. (B) The FeO concentration is given by the composition of the partial mantle melt.</figcaption></figure>

the initial Si content. Thus, an initial Si content of < 10 µg g⁻¹ is plausible.

Several uncertainties are inherent to the prediction of Si concentrations. The activity coefficient of Si in Fe,Ni-metal has experimentally been determined for a T range of 1523–1673 K (Vogel et al. 2005, 2006). The mean deviation of measured γ<sub>Si</sub><sup>metal</sup> data from the fit (Equation 9) is ±15% relative. Including the uncertainty of the experiments and the uncertainty intrinsic to the low-T extrapolation, we estimate the total uncertainty of γ<sub>Si</sub><sup>metal</sup> of ±30% relative. For the silicate phase, we assume an activity of a<sub>SiO<sub>2</sub></sub><sup>silicate</sup> = 0.5. We assume that this value is within ±50% correct. We have taken an FeO activity coefficient of 1.7 (Holzheid et al. 1997). O'Neill and Eggins (2002) reported γ<sub>FeO</sub><sup>silicate melt</sup> of 1–2. This transforms to an uncertainty in f<sub>O₂</sub> of 0.3 log-units and to an uncertainty in Si in metal of a factor of 2. This is, however, much lower than the uncertainty that is due to the choice of the f<sub>O₂</sub> buffer. Using the composition (FeO)

of a partial mantle melt gives by up to two orders of magnitude higher Si concentrations when compared with the model that assumes that the mantle is totally molten (Fig. 6). Even the large uncertainty intrinsic to the f<sub>O₂</sub> buffer model, however, is not a reasonable explanation for the observed low Si concentrations in iron meteorite metal.

### Thermodynamic Modeling: Case B (Internal Oxidation Model)

In this model, solid Fe,Ni-metal forms as a thin layer enclosing the liquid metal core (Fig. 3B). The metal prevents Si exchange between the liquid metal and the outer liquid (and solid) silicates. As heat can be effectively transported through the solid metal layer, the core will cool by fractional crystallization of the enclosed metal liquid. The total amount of Si in the metal-sulfide core was established by metal-silicate equilibrium at high temperatures before the system was isolated and removed from contact with liquid silicates.

The initial O concentration is controlled by the silicate/metal equilibrium. For simplification, we assume logf<sub>O₂</sub> = IW – 2. The solubility of O in liquid Fe that is saturated in FeO (logf<sub>O₂</sub> = IW) is a function of T with

$$ \log_{10}[O] = -\frac{6380}{T} + 1.765 \quad (10) $$

(Turkdogan 1980; [O] gives the concentration in wt%). We use Equation 10 to calculate the bulk O concentration in the liquid metal before formation of the isolating solid metal shell. At T = 1790 K and logf<sub>O₂</sub> = IW – 2 160 µg g⁻¹ O are dissolved in the metal. The effect of S on the activity coefficient of O in liquid Fe is neglected. The solubility of O and Si are coupled with:

$$ \log_{10}[\text{Si}] + 2\log_{10}[\text{O}] = -\frac{30410}{T} + 11.59 \quad (11) $$

The values in brackets are concentrations in wt% (Turkdogan 1980; for liquid Fe with trace concentrations of Si and O). The solubility product (Equation 11) decreases with decreasing T. At 1800 K, no S and logf<sub>O₂</sub> = IW – 2, a Si concentration of 193 µg g⁻¹ is calculated for the liquid metal. Using a Si solid/liquid partition coefficient of 0.5 (Ghosh 2001), a concentration of 80 µg g⁻¹ is calculated for the first solid metal. This is in excellent agreement with the Si concentration calculated for solid metal (T = 1790 K) in the equilibrium model (Fig. 6; logf<sub>O₂</sub> = IW – 2 corresponds to approximately 10 wt% FeO in mantle, Fig. 5).

During fractional crystallization O, Si, and S will partition between solid and liquid metal. For the
modeling, we assume a solid/liquid partition coefficient for O and S of zero. For Si, a solid/liquid partition coefficient of 0.5 was determined by Ghosh (2001) for Fe, Si-systems with low Si concentrations. The influence of alloying metals (Ni) and nonmetals (O, C, S, P) has not yet been experimentally investigated. For the influence of S on the Si partitioning, we use the data for Ge as proxy. Chabot (2004) gives a value of 0.66 for the Ge solid/liquid partition coefficient for S free systems, which is similar to 0.5 reported by Ghosh (2001) for Si. Chabot (2004) demonstrated that $D_{Ge}^{solid/liquid}$ increases with increasing S content. For reasons of the lack of experimental data, we assume that Si behaves similarly. The S concentration in the residual melt increases with increasing degree of crystallization *f*. Hence, $D_{Ge}^{solid/liquid}$ can be expressed as a function of *f* (Fig. 7).

The Si concentration in the fractionally crystallizing solid metal is controlled by the *T*-dependent Si solubility in the liquid (Equation 11) and the S-dependent Si solid/liquid partition coefficient (Fig. 7; the S concentration in the liquid is function of *f*). The crystallization *T* is function of S content (C and P are neglected here) and is derived from the Fe,Ni,S-phase diagram (we use a Ni concentration of approximately 8 wt%; Raghavan 2004). The fractional crystallization model hence allows expressing the Si concentration as a function of *f* and initial S concentration in the liquid (Fig. 8).

The calculations show that Si concentrations in the instantaneously forming metal are > 10 µg g⁻¹ for S contents between 0 and 17.5 wt% (Fig. 8). The initial Si concentrations are generally lower for S-rich systems because of the lower crystallization temperatures. The model shows that the Si concentrations are largely independent of the degree of crystallization, except for cases with > 10 wt% S. The increase in Si with increasing *f* is related to the strong increase in $D_{Ge}^{solid/liquid}$ in S-rich systems (Fig. 7).

The modeling, however, must be substantiated by experimental data on the effect of S on the solid/liquid partitioning of trace amounts of Si in Fe,Ni(S) alloys. Also, the effect of other nonmetals (C, P, O) should be investigated in experiments.

## Why Is the Si So Low in Iron Meteorite Metal?

Both models (Case A and B) do not explain the observed low and uniform Si concentrations in iron meteorite metal.

For the equilibrium model (Case A), our calculations show that Si concentrations should range from approximately 1 to 10 µg g⁻¹ in iron meteorite metal (Fig. 6). The bulk Ni concentrations of the analyzed iron meteorites range from 5.4 to 11.5 wt% Ni (Wasson 1969; Luais 2007; Wasson et al. 2007). We assume that the

<figure><img src="image_10.png" /></figure>

Fig. 7. Plot of the Si partition coefficient (using Ge as proxy, Chabot 2004) as a function of degree of crystallization *f*. The numbers indicate bulk S contents in wt%. The S solid/liquid partition coefficient is assumed to be zero.

<figure><img src="image_11.png" /></figure>

Fig. 8. Plot of the Si concentration in the instantaneously forming solid metal as a function of the degree of crystallization *f* and the initial S content (0–17.5 wt%) of the bulk metal core. The numbers indicate the bulk S concentrations. The shaded area outlines the observed minimum and maximum Si concentrations in the iron meteorite metal.

bulk compositions of the iron meteorites are roughly representative of the bulk composition (Fe,Ni) of the cores of their parent asteroids. This assumption is reasonable for bulk core compositions with < 20 wt% S with $D_{Ni}^{solid/liquid} \approx 1$ (Jones and Drake 1983). Higher Ni contents in the core imply more oxidized iron, assuming a chondritic Fe/Ni-ratio for the bulk planet. Thus, the spread in Ni contents of 5.4–11.5 wt% leads to a corresponding range in FeO contents of < 5 to approximately 23 wt% FeO in the mantles.

If Si in iron meteorite metal resulted from metal-silicate equilibration during core crystallization, Ni should be inversely correlated with Si (Fig. 6), as high Ni contents in the cores indicate oxidizing conditions during metal-silicate equilibration. A decrease in Si with increasing Ni is, however, not observed (Fig. 9). In fact, the highest Si concentrations were analyzed in Wiley,
<figure><img src="image_12.png" /><figcaption>Fig. 9. Plot of Si concentrations (means) versus bulk Ni concentration of magmatic (solid circles) and nonmagmatic (solid squares) iron meteorites. The bulk Ni data are taken from Luais (2007) and Wasson (1969).</figcaption></figure>

which is an ungrouped iron meteorite with a bulk Ni concentration of 11.5 wt% (Luais 2007).

The absence of a negative correlation between Si and Ni could suggest that iron meteorites crystallized at different temperatures or that Si was redistributed after crystallization of the metal. Different crystallization temperatures for each iron meteorite are difficult to envisage as Ni does not greatly affect the solidus temperature in the compositional range of interest (Swartzendruber et al. 1991).

A decrease in crystallization temperature, however, could be due to S or other nonmetals in the metal melt (Predel 2010b). For type IIAB iron meteorites, Chabot (2004) estimated 17 wt% S in the parent melt. This decreases the solidus temperature from approximately 1790 K to approximately 1600 K (Predel 2010b). The decrease in *T*, however, is not sufficient to explain the observed Si concentrations and would lead to 0.5–90 µg g⁻¹ Si in the metal (Fig. 6). The number given by Chabot (2004) was criticized by Wasson et al. (2007), who suggest a lower content of only 7.5 wt% S in the parent metal melt of the IIAB irons. This would lead to a decrease in the solidus temperature to approximately 1670 K and would not explain the low Si concentrations. Because of the low Ni concentrations in the type IIAB iron meteorites (5.2–6.5 wt%; Wasson 1969; Wasson et al. 2007), the FeO content of the corresponding mantle was likely <5 wt%, fairly reducing conditions for metal separation. In a parent body with 5 wt% FeO in the silicate mantle, 30 and 110 µg g⁻¹ Si are expected for the two models (Fig. 6). This is much higher than observed in the group IIAB iron meteorite metal (Fig. 2). The group IIIAB iron meteorites contain 7.4–10.7 wt% Ni in the metal (Wasson and Kimberlin 1967), corresponding to 18.6 wt% FeO in the mantle. For the

group of IIIAB iron meteorites, a S content of 12 wt% was suggested for the parent melt (Chabot 2004). This leads to a decrease in solidus temperature to approximately 1620 K (Predel 2010b). The corresponding Si concentrations would be 1–2 µg g⁻¹ (Fig. 6), which is by one order of magnitude higher than observed in iron meteorite metal (Fig. 2).

Enrichment of S in the residual metal melt leads to a further decrease in crystallization temperature. If such a process is responsible for the observed low Si concentrations, it can be tested with the set of IIAB meteorites analyzed in this study. Four different IIAB iron meteorites span nearly the entire range in Ir, from 0.01 to 44 µg g⁻¹ (Wasson et al. 2007). The crystallization sequence based on Ir among the analyzed IIAB iron meteorites is: Guadalupe y Calvo (42.6 µg g⁻¹ Ir) followed by North Chile (3.43 µg g⁻¹ Ir), Walker County (3.27 µg g⁻¹ Ir), and Sao Juliao (0.012 µg g⁻¹ Ir). The corresponding degrees of crystallization were estimated to be 2–48% (Wasson et al. 2007). Although these meteorites cover the whole range of Ir concentrations, no systematic variation of Si with Ir is observed.

In model Case A, we have considered equilibration between solidifying metal and liquid silicates where the communication between solid metal and liquid silicates is achieved through the liquid portion of the core. The activity coefficient of Si in liquid Fe,Ni,C,S,P alloys has not yet been studied in detail. As mentioned above, for cast iron, Ghosh (2001) reported a Si γ-Fe/liquid-Fe partition coefficient of 0.5. This implies that liquid Fe in equilibrium with silicate has twice the Si content of the solid metal. This number will decrease with increasing S content at increasing degree of crystallization, as discussed before.

The observed Si concentrations, however, are also lower than predicted by the Case B model (Fig. 8).

## Is the Subsolidus Equilibration Iron Meteorite Metal with Silicates Responsible for the Observed Low Si?

The slow cooling of magmatic iron meteorites within the range of <1 to >1000 K Ma⁻¹ allowed extensive subsolidus reactions that are, for example, visible in the form of Widmannstätten patterns (Hopfe and Goldstein 2001; Haack and McCoy 2007; Goldstein et al. 2009). The high Si concentrations in two iron meteorites reported by Wai and Wasson (1969, 1970) and discussed above are found in two meteorites with very high cooling rates. Apparently, slow cooling allows removal of Si from the metal, while high cooling rates lead to Si retention during cooling.

Subsolidus equilibration of metal with silicate (i.e., loss of Si from metal through diffusion) may have caused the observed low Si concentrations in iron
meteorite metal. We modeled the subsolidus diffusional exchange between metal and silicate using diffusion coefficients for α- and γ-Fe from LeClaire and Neumann (1990). The diffusion calculations were conducted numerically for spheroidal silicate inclusion surrounded by metal. The starting temperature was 1773 K with exponential cooling at rates of 10 and 100 K $Ma^{-1}$ at $T = 700$ K and a final temperature of 200 K. The T-dependency of the diffusion coefficients were taken from LeClaire and Neumann (1990). The γ- to α-transition was assumed to occur instantaneously at $T = 1273$ K (LeClaire and Neumann (1990). The computations were run for $t_{max} = 10^9$ yr. The length scales for diffusional exchange are in the range up to a few meters at maximum. Hence, low Si concentrations in iron meteorite metal can be explained by subsolidus exchange with silicate on the meter scale. Re-equilibration of km-sized metal cores would require billions of years and could not have occurred.

More importantly, large-scale diffusion of Si in a planetary core would contradict the observed chemical trends (e.g., Ir versus Au) within individual magmatic iron meteorite groups. These trends have been established by fractional crystallization with only little disturbance after solidification. If Si had been redistributed over large distances (e.g., through enhanced diffusion along grain boundaries), simultaneous diffusion of Ir and other trace elements would have erased variations in concentrations of trace elements in different members of a group of iron meteorites. Thus, the constancy of Si in iron meteorites, in particular, in different meteorites of a single group contrasts with the large variations of Ni, Ir, Au etc. in individual members of a single group of iron meteorites. This cannot be the result of differences in diffusion coefficients, e.g., high for Ir and low for Si. Although Righter et al. (2005) reported a much lower diffusion coefficient for Ir compared with Si in Fe metal, other elements with variable concentrations in iron meteorites have diffusion coefficients that are comparable (e.g., Ga) to or higher (e.g., Au) than Si.

One feasible model to explain subsolidus equilibration between Fe,Ni-metal and silicate is that metal occurred in the form of metal pockets that were surrounded by silicates in the parent body (raisin-bread model, e.g., Goldstein and Axon 1973). Such a scenario has also been proposed by Roskosz et al. (2006) to explain the re-equilibration of Fe isotopes between metal and silicate. The size of the pockets is limited by Si diffusion to the meter scale. In such a scenario, silicates have acted as barriers for simultaneous redistribution of siderophile elements between individual pockets.

A further possibility is that Si exsolved from the solid metal as silicate or $SiO_2$, which now occur as

inclusions in iron meteorites. Subsolidus exsolution of silica by oxidation of formerly dissolved Si during metamorphism has been reported from chondrite metal (Zanda et al. 1994). A similar process may have operated in iron meteorites. Buchwald (1975, 1977) reported the occurrence of silicate inclusions in a number of iron meteorites (see Iron Meteorites section). Many silicate inclusions, however, may simply have been overlooked in polished sections of magmatic iron meteorites due to their rarity (Wasson, 2006, personal communication).

Except for the silicate-rich Landes meteorite, none of the meteorite sections that we studied contain silicate inclusions. The probability of finding a silicate inclusion is estimated for an iron meteorite with initially $10 \mu g g^{-1}$ Si in the Fe,Ni-alloy. Oxidation of the entire Si results in the formation of $0.002 \text{ wt}\% \text{ }SiO_2$, which corresponds to $0.006 \text{ vol}\% \text{ }SiO_2$ ($\rho_{metal} = 7.8 \text{ g cm}^{-3}$, $\rho_{SiO_2} = 2.65 \text{ g cm}^{-3}$). This transfers to a probability of only one $0.25 \text{ cm}^2$ inclusion in an area of $0.4 \text{ m}^2$. The apparent absence of inclusions in the studied sections and the absence of reports of silica inclusions in magmatic iron meteorites thus could indicate that inclusions are in the cm-range, i.e., their detection in a section is highly unlikely. Some inclusions may also have been lost during cutting and polishing. Alternatively, initial Si concentrations have been < $10 \mu g g^{-1}$, i.e., the concentration of inclusions is even lower. At $1 \mu g g^{-1}$ Si, a $4 \text{ m}^2$ section would have to be investigated to find a single $0.25 \text{ cm}^2$ inclusion.

Wasson et al. (2006) discussed the formation of large silica inclusions in the Gibeon IVA iron meteorite by precipitation from SiO gas. This process could potentially have operated on scales larger than 20 m. As we have not analyzed IVA iron meteorites for Si, we cannot conclude that the proposed process is the cause for the observed low Si concentrations.

Precipitation of Si in C-rich inclusions has not been investigated. The partition coefficient of Si between cementite ($Fe_3C$) and α-Fe is < 1 (Al-Salman et al. 1979). We therefore conclude that cohenite ($Fe,Ni)_3C$ is not a likely host of Si in iron meteorites. Although schreibersite contains approximately 10 times more Si than coexisting kamacite (Table 1, note that the schreibersite analysis is semi-quantitative only), there is not enough Si in schreibersite to account for the missing Si in iron meteorites.

## S in Iron Meteorite Metal—Analogous to Si?

We have shown here that Si concentrations in iron meteorite metal are much lower than expected from metal/silicate equilibration at the time of their crystallization. We suggested that subsolidus equilibration with silicate
(surrounding mantle or inclusions) is the cause for the low Si contents. A similar process may have established the S concentration in iron meteorite metal.

Two S concentration measurements of iron meteorite metal were published in the literature. In an article reporting S isotope analyses, Rumble et al. (1993) found 1.02 µg g⁻¹ S in Gibeon (IVA) and 2.6 µg g⁻¹ in Canyon Diablo (IAB) metal. Most iron meteorites contain variable amounts of troilite inclusions that represent trapped melts incorporated during solidification of the metal. Fe,Ni-metal in equilibrium with FeS (T = 1270 K), however, should contain approximately 280 µg g⁻¹ S (Hansen and Anderko 1958). As in the case of Si, the discrepancy between expected and measured S concentration can be explained by low-temperature solid-state exsolution of sulfide or sulfur from metal.

However, as for the case of sulfide inclusions, a majority of silicate inclusions in iron meteorites are unlikely to have formed by exsolution of dissolved Si. This is especially true for the abundant silicate inclusions in nonmagmatic iron meteorites. Subsolidus precipitation of sulfide and silica, however, may have occurred at the interfaces between pre-existing, trapped inclusions and metal.

## SUMMARY AND CONCLUSIONS

In this work, we demonstrate that SIMS is an appropriate tool to determine Si in metal down to the ng g⁻¹ concentration range. We found that iron meteorite metal contains between 0.09 and 0.46 µg g⁻¹ Si. No differences between magmatic and nonmagmatic groups and between coexisting kamacite and Ni-rich fine-grained assemblages were observed.

Thermodynamic calculations based on metal-silicate equilibration and mass balance at temperatures where liquid Fe,Ni,S-metal separates from silicates predict higher and variable Si concentrations in iron meteorite metal. The rather uniform Si concentrations in iron meteorite metal are in stark contrast to the large variations in Ni, Ir and other trace elements in individual members of a single group of iron meteorites. For example, correlations of Si with Ni expected from our calculations were not observed.

We suggest that the low Si in metal of iron meteorites has been established by subsolidus equilibration between metal and silicate at temperatures below 1270 K.

We discuss two different scenarios that may explain the low Si concentrations:

1. Silicon was lost from metal by diffusion into silicates. Diffusion calculations show that this model can only be applied to metal bodies less than a few meters in size. Equilibration of larger, kilometer-

sized asteroidal cores is impossible as diffusion rates of Si are too low. Also, diffusional exchange of Si over kilometer distances would contradict the siderophile element (e.g., Ir) heterogeneity that is observed in iron meteorites of a single group. This scenario implies a parent asteroid with small metal pockets surrounded by silicate.

2. Alternatively, Si may have been removed from metal by formation of scattered silicate inclusions. Silicate inclusions have been reported from a variety of iron meteorites, including magmatic iron meteorites. The absence of silicate inclusions in our samples and the small number of inclusions in magmatic iron meteorites may be explained by the presence of large (cm-sized) inclusions. The probability of cross-cutting such a large inclusions in a section is very low and may explain the few reports of inclusions.

Possibly, however, a combination of these possibilities is required to explain the observed low Si in iron meteorite metal.

*Acknowledgments*—We are indebted to the staff of the CRPG during the 2-yr stay of A. P. in Nancy. M. Chaussidon, E. Deloule, N. Vigier, G. Libourel, and B. Marty are thanked for discussions and their help with the ion microprobe work. This manuscript benefited from reviews by J. T. Wasson, J. I. Goldstein, and N. Chabot (AE). The project was made possible through financial funding by the Deutsche Forschungsgemeinschaft (DFG; PA 346/34-1 [H. P.] and PA-909/1-1 and PA-909/2-1, Emmy-Noether Program [A. P.]).

*Editorial Handling*—Dr. Nancy Chabot

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