# Silicon isotope variations in the inner solar system: Implications for planetary formation, differentiation and composition

Thomas Zambardi, Franck Poitrasson, Alexandre Corgne, Merlin Méheut, Ghylaine Quitté, Mahesh Anand

► To cite this version:

Thomas Zambardi, Franck Poitrasson, Alexandre Corgne, Merlin Méheut, Ghylaine Quitté, et al.. Sil- icon isotope variations in the inner solar system: Implications for planetary formation, differentiation and composition. Geochimica et Cosmochimica Acta, 2013, 121, pp.67-83. 10.1016/j.gca.2013.06.040 . hal-04749541

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1 Silicon isotope variations in the inner solar system:

2 Implications for planetary formation, differentiation and composition

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4 Thomas Zambardi<sup>a,b*</sup>, Franck Poitrasson<sup>a</sup>, Alexandre Corgne<sup>c,d</sup>, Merlin Méheut<sup>a</sup>, Ghylaine

5 Quitté<sup>c</sup> and Mahesh Anand<sup>f,g</sup>

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7 * Corresponding author:

8 Tel: 00337.86.27.29.84; e-mail: zambardi@illinois.edu

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10 <sup>a</sup> Géosciences Environnement Toulouse, CNRS – Université de Toulouse – IRD, 14 avenue
Edouard Belin, 31400 Toulouse, France.

12 <sup>b</sup> Department of Geology – Natural History Building, University of Illinois at Urbana-
Champaign, 1301 W. Green Street, 61801 Urbana, IL, USA.

14 <sup>c</sup> Institut de Recherche en Astrophysique et Planétologie, CNRS – Université de Toulouse, 14
avenue Edouard Belin, 31400 Toulouse, France.

16 <sup>d</sup> Instituto de Geociencias, Universidad Austral de Chile, Casilla 567, Valdivia, Chile

17 <sup>e</sup> Laboratoire de Géologie de Lyon : Terre, Planètes, Environnement, CNRS, ENS de Lyon,
Université Lyon 1, 46 allée d'Italie, 69364 Lyon, France.

19 <sup>f</sup> Department of Physical Sciences, The Open University, Milton Keynes, MK7 6AA, UK.

20 <sup>g</sup> Department of Mineralogy, The Natural History Museum, London SW7 5BD, UK.

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**Abstract**

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Accurate and precise Si isotope measurements were obtained using magnesium doping and high-resolution plasma source mass spectrometry for samples representative of the Earth, as well as lunar samples, meteorites from Mars (SNC), eucrites, a howardite, carbonaceous chondrites (CC), ordinary chondrites (OC) and enstatite chondrites (EC). Our data confirm that significant Si isotope fractionations exist among the inner solar system planetary bodies.

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They show that the Earth and the Moon share the same Si isotopic composition, which is heavier than all other measured bodies, in agreement with most of previous studies. At the other end of the spectrum, enstatite chondrites have the lightest Si isotope compositions. In order to precisely estimate the amount of Si that may have entered the Earth's core, we developed a refined model of Si partitioning based on continuous planetary accretion that takes into account the likely variations in *T*, *P* and *fO*₂ during the Earth's accretion, as well as isotopic constraints involving metal-silicate partitioning derived from both experimental and natural sample data sets.

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Assuming that the difference between the isotopic signature of the bulk silicate Earth (BSE) and chondrites solely results from Si isotope fractionation during core formation, our model implies that at least ~12wt% Si has entered the Earth's core, which is greater than most of the estimates based on physical constraints on core density or geochemical mass balance calculations.

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This result leads us to propose two hypotheses to explain this apparent contradiction:
1) At least part of the Earth's building blocks had a Si isotope composition heavier than that observed in chondrites (i.e., $\delta^{30}\text{Si} > -0.39\‰$).
2) If on the contrary the Earth accreted only from material having chondritic $\delta^{30}\text{Si}$, then an additional process besides mantle-core differentiation is required to generate a stronger isotope fractionation and lead to the observed heavy isotope composition of the bulk silicate Earth. It may be the loss of light Si isotopes

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47 during partial planetary vaporization in the aftermath of the Moon-forming giant impact. This
48 process, which may have affected metallic cores, required a thorough isotopic re-equilibration
49 between core and silicate to explain the similar heavy isotope composition of the silicate
50 portions of the Earth and the Moon.

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52 Keywords: silicon isotopes, terrestrial planets, core formation, inner solar system
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# 1. INTRODUCTION

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The core of the Earth has been shown to have a 5-10% density deficit relative to a pure Fe-Ni alloy (Birch, 1968; Boehler, 2000; Anderson and Isaak, 2002), suggesting significant incorporation of light elements such as C, O, S and/or Si. Given the superchondritic Mg/Si ratio of the silicate Earth and assuming that terrestrial planets have a bulk composition similar to those of chondrites (e.g., Jagoutz et al., 1979; Allègre et al., 1995; McDonough and Sun, 1995), Si has been considered among the most likely candidate light elements to occur in the core (e.g., MacDonald and Knopoff, 1958; Allègre et al., 1995). However, the Si contribution to the light element budget of the Earth's core remains poorly constrained, reflecting the conflicting evidence from cosmochemical, experimental and theoretical studies concerning which light element(s) is/are present (e.g., Poirier, 1994; Allègre et al., 1995; Sherman, 1997; Kilburn and Wood 1997; Gessmann et al., 2001; Li et al., 2001; Alfè et al., 2002; Rubie et al., 2004; Asahara et al., 2007; Badro et al., 2007; Côté et al., 2008; Corgne et al., 2009; Antonangeli et al., 2010). The determination of the Si isotope signatures (from here on expressed as permil $\delta^{30}$Si) of chondritic and Earth materials has been proposed to provide additional constraints on the amount of Si present in the Earth's core (e.g., Georg et al., 2007): for a given Si isotopic fractionation between the Earth's core and the Earth's mantle, the difference in $\delta^{30}$Si between the silicate portion of the Earth and the bulk Earth is indeed proportional to the fraction of total Si going into the core. Assuming that the mean Earth signature can be estimated by measuring chondritic meteorites, the difference in isotopic composition between rock samples representative of the bulk silicate Earth and chondritic materials ($\Delta^{30}$Si$_{BSE-Bulk Earth} = \Delta^{30}$Si$_{BSE-Chond.}$) can thus be used to estimate the Si content of the core.

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With the exception of Chakrabarti and Jacobsen (2010a), all recent studies indicate that chondritic Si isotope compositions are clearly distinct from the Earth's mantle signature
(Georg et al., 2007; Fitoussi et al., 2009; Zambardi et al., 2009; Ziegler et al., 2010; Armytage et al., 2011; Fitoussi and Bourdon, 2012). All these studies consider that the $\Delta^{30}\text{Si}_{\text{BSE-Chond.}}$ values directly result from the Si incorporation into the Earth's core during accretion. However two issues remain: (1) the measured $\Delta^{30}\text{Si}_{\text{BSE-Chond.}}$ values vary from one study to another, which leads to significant variations in estimated core contents of Si and (2) the Si fractionation during terrestrial accretion must be modeled more realistically in order to better understand the relationship linking $\Delta^{30}\text{Si}_{\text{BSE-Chond.}}$ to the Si content of the Earth's core. To further explore the usefulness of Si isotopes in this context, and in particular to be able to compare the Si-core content inferred from isotope data with other independent constraints, it is crucial to address these two points.

In addition to potential analytical bias during isotopic measurements that may be adequately addressed through interlaboratory comparison of geostandard analyses (Reynolds et al., 2007; Zambardi and Poitrasson, 2011), part of the uncertainty in estimating the Si content of the core resides in the choice of which chondritic reference to use as a proxy for the bulk Earth. This is because different chondrites types do not necessarily share the same Si isotope signature, with carbonaceous and ordinary chondrites showing $\delta^{30}\text{Si}$ values notably higher than enstatite chondrites (Georg et al., 2007; Fitoussi et al., 2009; Armytage et al., 2011; Savage and Moynier, 2013). Furthermore, the nature of the terrestrial building blocks remain hypothetical, and different Earth forming models have been proposed, requiring the use of variable chondritic sources (e.g., Allègre et al., 1995, 2001; Palme and O'Neill, 2003; Trinquier et al., 2007; Carlson et al., 2007; Leya et al., 2008; Simon and DePaolo, 2010; Javoy et al., 2010; Schönbächler et al., 2010).

Last but not least, another potential source of uncertainty in estimating the Si content of the Earth's core comes from the use of single-stage core formation models (Georg et al., 2007; Shahar et al., 2009; Ziegler et al., 2010; Armytage et al., 2011) that involve fixed
conditions of temperature (T), pressure (P) and oxygen fugacity ($fO_2$) throughout this protracted event. Such models neglect the continuous variations of these three parameters during core-mantle differentiation, while they are known to significantly affect Si metal-silicate partitioning (e.g., Gessmann et al., 2001; Corgne et al., 2008). Furthermore, T variations are also known to affect isotope fractionation during metal-silicate melt partitioning (e.g., Shahar et al., 2009, 2011; Ziegler et al., 2010). For these reasons, it is more appropriate to consider continuous core formation models (Wade and Wood, 2005; Corgne et al., 2008; Chakrabarti and Jacobsen, 2010a).

To address these issues, we have produced new Si isotope measurements for a number of terrestrial and lunar rocks, as well as Martian meteorites, eucrites, a howardite and different types of chondrites. Our analytical approach combines Mg-doping and high-resolution mass spectrometry, which improves both data accuracy and reproducibility (Zambardi and Poitrasson, 2011). A $\Delta^{30}\text{Si}_{\text{BSE-Chond.}}$ value is inferred from our data set, and incorporated into a continuous accretion elemental partitioning model (Wade and Wood, 2005; Corgne et al. 2008). Our model also considers the constraints on Si isotope metal-silicate fractionation (Ziegler et al., 2010; Shahar et al., 2009, 2011). This approach allows us to refine the estimates of Si incorporated into the Earth's core as a function of the selected chondrite type(s) that contributed matter to the growing Earth. We eventually compare the expected Si-core concentrations inferred from our model to those derived from purely physical constraints and/or geochemical mass balance calculations, and then discuss the range of possible processes and materials involved in terrestrial planet formation during the early stages of the solar system.

## 2. SAMPLES AND METHODS

This study focuses on the Si isotope compositions of silicate samples spanning a large
range of igneous rock compositions from the inner solar system. They include samples from the Earth (11 peridotites, 4 basalts, 1 gabbro, 1 diorite, 1 dolerite and 1 komatiite), Moon (4 basalts, and 1 norite), Mars (5 shergottites, 4 nakhlites and 1 chassignite), as well as HEDs (6 eucrites and 1 howardite), 6 ordinary chondrites, 3 carbonaceous chondrites and 2 enstatite chondrites. Each sample group consists of a variety of rocks that reflect as much as possible the geochemical diversity of their respective parent bodies. The full sample set is listed in Table 1.

For meteorites and lunar samples, the aliquots used for the analyses were taken from at least 1g of well-homogenized bulk-rock powders (or at least 0.8g for Martian meteorites) to ensure that the sample is representative of the bulk rock as far as possible even for these rare samples. For the more abundant terrestrial rock samples, powders were prepared by crushing more than 100g of starting material. Since a complete description of the analytical procedure is already given in Zambardi and Poitrasson (2011), only the main steps are summarized hereafter.

The sample solutions were prepared using an alkali fusion method adapted from Georg et al. (2006). Silicon purification was achieved by column chromatography using Biorad AG50W-X12 cationic resins. Typical Si recovery was between 95 and 100% after the decomposition and purification steps. Procedural blanks for Si were ~30ng, corresponding to less than 0.2% of the total amount of Si analyzed per sample and were thus negligible. All measurements were performed at the Observatoire Midi-Pyrénées (Toulouse, France) and at the Open University (Milton Keynes, United Kingdom) using a Neptune (Thermo) MC-ICP-MS in medium or high-resolution mode, with wet plasma conditions and slightly acidic HCl solutions (0.05M for both samples and standard). The instrumental mass bias drift was corrected simultaneously using the sample standard bracketing (SSB) technique and by adding magnesium (Mg) as an internal standard in both standard and samples. Measurements
of $^{25}$Mg/$^{24}$Mg ratios were performed in the dynamic mode. The mass bias drift correction was then applied using the Russel's exponential law (Russel et al., 1978). All the data are reported in Table 1 as $\delta^{30}$Si and $\delta^{29}$Si (‰) relative to the NIST international standard reference material SRM 8546 (NBS-28) following the relation:

$$ \delta^{xx} \text{Si} = \left( \frac{\left( \frac{^{xx}\text{Si}}{^{28}\text{Si}} \right)_{\text{Sample}} - 1}{\left( \frac{^{xx}\text{Si}}{^{28}\text{Si}} \right)_{\text{NBS-28}}} \right) \times 1000 \quad (1) $$

where the superscript xx denotes the Si isotope with either mass number 30 or 29.

The long-term external reproducibility for $\delta^{30}$Si and $\delta^{29}$Si was given by the standard deviation of 54 measurements of BHVO-2 over three years. It yielded 0.07‰ and 0.04‰ (2SD) for $\delta^{30}$Si and $\delta^{29}$Si, respectively. In the following, uncertainties of planetary and meteorite parent body means are computed as two standard error of the mean (2SE) rather than two standard deviation (2SD) since the purpose is to indicate how well we know the mean rather than to give a measure of the spread of a set of individual measurements around the mean (see e.g., Miller and Miller, 1993). Noteworthy, these two standard error uncertainties represent a direct proxy of t-tests (see e.g., Poitrasson et al., 2004, 2005). All literature data have been recomputed this way to be comparable. However, whether computed as 2SE or 2SD, it is important to keep in mind that these uncertainty estimates only include random errors but not systematic ones potentially associated to analytical or sampling biases.

## 3. RESULTS

All data are given in Table 1. Bulk rock $\delta^{30}$Si values for inner solar system bodies range from -0.19‰ to -0.64‰. Comparison between $\delta^{30}$Si and $\delta^{29}$Si in a three-isotope plot shows that these Si isotope fractionations are mass dependent (Fig. 1), in agreement with previous data (e.g., Douthitt, 1982; Molini-Velsko, 1986; Ding 1996; Georg et al., 2007; Fitoussi et al., 2009; Savage et al., 2010; Ziegler et al., 2010; Armytage et al., 2011).
Terrestrial mafic rocks yield δ³⁰Si ranging from -0.19‰ to -0.33‰ (Table 1; Fig 2), which includes the smaller variation range displayed by lunar mafic samples (δ³⁰Si ranging from -0.24‰ to -0.29‰) and therefore suggests the two bodies share similar isotopic features.

Meteoritic samples including chondrites, rocks thought to originate from Mars and 4-Vesta all yield lighter δ³⁰Si values, ranging from -0.33‰ to -0.64‰ (Table 1; Fig. 2). Among chondrites, enstatite chondrites clearly display the lightest isotope compositions (δ³⁰SiEC = -0.59‰ and -0.64‰) relative to both ordinary chondrites (δ³⁰SiOC = -0.47 ± 0.01‰ (2SE, n=6)) themselves not statistically different from carbonaceous chondrites (δ³⁰SiCC = -0.43 ± 0.05‰ (2SE, n=3)). As the signatures of CC and OC overlap within uncertainties, their averaged signature is reported in the following as δ³⁰SiO/CC = -0.46 ± 0.02‰ (2SE, n=9). We note that Savage and Moynier (2013) found significant differences between EL and EH chondrites, with EH being enriched in light isotopes compared to EL. Although the present study does not reproduce such an extensive E-chondrite set, it should be noticed that our two EH-chondrites do not yield comparable negative values. In fact, the two EH we measured rather plot within the range of the EL chondrites analyzed by Savage and Moynier (2013).

This issue thus requires additional measurements of E-chondrites to be solved.

A review of previously published data reveals significant differences between laboratories, even for the same extra-terrestrial samples and terrestrial geostandards (Georg et al., 2007; Fitoussi et al., 2009; Zambardi et al., 2009; Ziegler et al., 2010; Chakrabarti and Jacobsen, 2010a; Armytage et al., 2011, 2012; Zambardi and Poitrasson, 2011; Fitoussi and Bourdon 2012; Savage et al., 2012, Savage and Moynier, 2013, see Fig. 3). These variations may have multiple origins that remain complex to identify with certainty. Among them, sample heterogeneity is a plausible explanation, especially for rare samples like meteorites and lunar rocks given the limited amount of sample available for analyses. Armytage et al. (2011) previously noticed the effect of sample heterogeneity for the Martian meteorite
202 Governador Valadares, with two different preparations producing a 0.1‰ difference in δ³⁰Si.
203 In the present study, sample heterogeneity was avoided as much as possible by powdering at
204 least 0.8g of bulk SNC or 1g of bulk chondrites, achondrites and lunar samples, prior to
205 sample preparation. Furthermore, several methodological aspects that may influence the final
206 results are still a subject of debate in the literature (e.g., Zambardi and Poitrasson, 2011).
207 These include for example the use of HCl or HNO₃ sample matrices (e.g. Cardinal et al.,
208 2003; van den Boorn et al., 2009), removal of S from samples (van den Boorn et al., 2009),
209 the use of a collision cell (Chakrabarti and Jacobsen, 2010a), operating a MC-ICP-MS under
210 dry or wet plasma conditions and more broadly, the contrasted architectures of the plasma
211 source mass spectrometer used. Altogether, the present data are well consistent with those of
212 Armytage et al., (2011), especially for extra-terrestrial Si isotopic compositions (Fig. 4).

213 Given that the models estimating the planetary Si-core contents are highly sensitive to
214 the Δ³⁰SiBSE-Chond. value (see details in section 4.3), it is extremely important to estimate each
215 planetary δ³⁰Si average as accurately and precisely as possible. In this study, the analytical
216 protocol includes the use of Mg as an internal standard. This technique has been demonstrated
217 to correct for isotope fractionation caused by rapid mass bias variations and matrix effects
218 (Cardinal et al., 2003; Engström et al., 2006; Zambardi and Poitrasson, 2011). Combined with
219 high-resolution mass spectrometry, it helps to significantly improve the reproducibility and
220 accuracy of Si isotope measurements.
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## 4. DISCUSSION

### 4.1. δ³⁰Si variability among planetary bodies

222 Planetary δ³⁰Si averages calculated from our new results as well as from recent
223 literature data are presented in Figure 4. For terrestrial samples, the effect of magmatic
224 differentiation has been shown to increase the amount of heavy Si isotopes in the most
differentiated materials (Douthitt 1982; Ding et al., 1996; Savage et al., 2011). These previous
results suggest that, assuming a homogeneous mantle source (see Savage et al., 2010), only
mafic and ultramafic materials should be considered as representative of the bulk silicate parts
of the solar system planetary bodies. We thus computed the isotopic composition of the bulk
silicate Earth (BSE) following this rule. This yields for our measurements δ³⁰SiBSE = -0.27 ±
0.01‰ (2SE, n=19; see Table 1, Fig.4), in excellent agreement with values recently proposed
by Savage et al. (2010, 2011), and Fitoussi and Bourdon (2012), but heavier by ~0.1‰ than
that of Georg et al. (2007) and Chakrabarti and Jacobsen (2010a).

Four basalts and a norite sample allowed us to estimate the isotopic composition of the
lunar mantle (BSMo). We found δ³⁰SiBSMo = -0.27 ± 0.02‰ (2SE, n=5), which is identical to
the Si isotope signature of the BSE (Fig. 4), and in agreement with the average computed by
Georg et al. (2007), Armytage et al. (2012) and Fitoussi and Bourdon (2012). However this
signature is significantly heavier than the signature of a lunar breccia (-0.45 ± 0.05‰, 2SE,
n=1) measured by Chakrabarti and Jacobsen (2010a).

Among the Martian meteorites studied here, the Los Angeles shergottite represents the
most differentiated and felspar-rich basaltic rock (Xirouchakis et al., 2002; Papike et al.,
2009). Like differentiated terrestrial samples, this more felsic sample was found to have a
heavier Si isotopic signature (Table 1). Los Angeles was thus omitted from the bulk silicate
Mars (BSMa) calculation with only less evolved rocks (meaning all samples but Los Angeles)
used to estimate the Martian mantle δ³⁰Si. We thus derive a bulk silicate Mars signature
(δ³⁰SiBSMa) of -0.49 ± 0.03‰ (2SE, n=9), significantly lighter than both the δ³⁰SiBSE and
δ³⁰SiBSMo (Table 1, Fig. 4). Since Nakhlites may contain up to 1wt% of secondary minerals
(e.g., smectite) (Bunch and Reid, 1975), their original isotopic signature could be shifted
toward lower δ³⁰Si values, given the generally accepted effects of low-temperature alteration
on Si isotopes (e.g., Ziegler et al., 2005a). However, incorporation of 1% of smectite with a
252 light δ³⁰Si signature of -2‰ (e.g., Ziegler et al., 2005b) would induce a maximum shift of
253 0.02‰ on bulk Nakhlites samples, which does not lead to a significantly different δ³⁰SiBSMa
254 value. Our δ³⁰SiBSMa value is in excellent agreement with the -0.48‰ ± 0.06 (2SE, n=5)
255 estimate proposed by Armytage et al. (2011), is identical within uncertainties with the result
256 of Georg et al. (2007) and again significantly differs from the value reported by Chakrabarti
257 and Jacobsen, 2010a).

258 The study of HED meteorites helps to constrain the bulk isotope composition of their
259 parent body, most probably asteroid 4-Vesta (Consolmagno and Drake, 1977). Eucrites are
260 thought to be fragments of the asteroid's upper crust, whereas diogenites probably originated
261 from deeper levels. Howardites may be mixed breccias of eucrites and diogenites (e.g.,
262 Mittlefehldt, 1994). The lack of a resolvable difference in δ³⁰Si between the howardite and
263 eucrites measured in the present study suggests a homogeneous Si isotope composition for the
264 silicate portion of the HED parent body, in contrast with Li isotope compositions (Magna et
265 al., 2006). The HED signature of -0.41 ± 0.03‰ (2SE, n=7) that we obtained is in agreement
266 with the data from Fitoussi et al. (2009), Armytage et al. (2011) and Chakrabarti and Jacobsen
267 (2010a), but differs from data by Georg et al. (2007).

268 Lastly, the chondrite averages often vary among authors (Fig. 4), notably for the O/C
269 groups. Given that the ordinary chondrites do not seem to show much isotopic differences
270 among their types (e.g., H, L or LL; Table 1; Armytage et al., 2011), we tentatively attribute
271 these discrepancies to inter-laboratory biases.

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273 **4.2. Planetary bulk Si isotopic compositions and chemical heterogeneity of the inner**
274 **solar system**

275 As previously shown in most of the recent studies, the lunar mantle and BSE have
276 similar δ³⁰Si, both have significantly heavier isotope compositions than Mars, 4-Vesta and
277 chondrite parent bodies (e.g., Georg et al., 2007; Armytage et al., 2011, 2012, Fitoussi and Bourdon 2012, see Fig. 4). Similarities between the Earth and the Moon were also observed for other isotope systems such as O (e.g., Wiechert et al., 2003; Spicuzza et al., 2007; Hallis et al., 2010), Cr (e.g., Lugmair and Shukolyukov, 1998), Ti (Zhang et al., 2012), Li (e.g., Magna et al., 2006), W (e.g., Touboul et al., 2007) and Sr (Moynier et al., 2010), thereby supporting a genetic link between the Moon and the Earth's mantle. Assuming that chondrites are the building blocks from which the terrestrial planets formed (e.g., Jagoutz et al., 1979; Allègre et al., 1995; McDonough and Sun, 1995), the heavy Si isotope composition of the Earth's mantle relative to chondrites is therefore not representative of the bulk Earth. As previously discussed in the literature, incorporation of isotopically light Si into the Earth's core is a possible explanation for this "heavy mantle" (Georg et al., 2007; Shahar et al., 2009; Fitoussi et al., 2009; Ziegler et al., 2010; Chakrabarti and Jacobsen, 2010a; Armytage et al., 2011; Fitoussi and Bourdon, 2012).

290 The δ³⁰Si signatures of Mars and Vesta do not show such a shift with respect to chondrites. A possible explanation (further discussed in section 4.4) is that these bodies have not segregated a significant amount of Si into their core reflecting their relatively small size and different redox conditions of formation, in agreement with models of planetary differentiation (e.g., Wade and Wood, 2005; Corgne et al., 2008).

295 To identify the process that led to the heavy Si isotope composition of the Earth's mantle, and eventually better constrain the amount of Si segregated into the Earth's core, an important objective is to obtain a precise value of the bulk-Earth signature. However current geochemical models indicate that the Earth may have accreted from C-chondrites (e.g., Allègre et al., 1995, 2001; Palme and O'Neill, 2003), E-chondrites (e.g., Javoy, 1995; Trinquier et al., 2007; Javoy et al., 2010), O-chondrites (e.g., Carlson et al., 2007; Leya et al., 2008; Simon and de Paolo, 2010) or mixed types of chondrites (e.g., Schönbächler et al., 2010).
302 2010; Fitoussi and Bourdon 2012). Given that chondrites types do not all share the same Si
303 isotope signature (Fig. 4), the bulk-Earth signature may then vary from -0.62‰ assuming a
304 pure E-chondrites end-member, to -0.46‰ considering a pure O/C-chondrites end-member.

305 Given that the HED parent body and Mars did not experience significant core
306 partitioning of Si during metal segregation (e.g., Wade and Wood, 2005; Corgne et al., 2008),
307 their mantle signature should then correspond to that of the bulk planetary body. Since the Si
308 isotopic signature for HED samples is indistinguishable within uncertainties from that of O/C-
309 chondrites, we infer the bulk composition of their parent body is consistent with a starting
310 material dominated by O/C-chondrite-like materials. In contrast, our data show that the
311 Martian isotopic composition appears slightly but significantly lighter than 4-Vesta. While
312 this result is compatible with an O/C-chondrite-dominated Mars composition, it may betray
313 the contribution of a significant amount of E-chondrites as well. A simple mass balance
314 calculation can be used to give a rough estimate of the proportions of different chondrite types
315 in the starting composition of Mars. Doing so leads to a range of 0 to 40wt% (with an average
316 at 20wt%) E-chondrites for Mars' bulk composition. Interestingly, Sanloup et al. (1999) and
317 Mohapatra and Murty (2003) proposed even larger contributions of E-chondrites for Mars
318 building blocks (~70wt% E-chondrites) based on O isotope data. Despite the uncertainties of
319 such models, these data suggest that E-chondrites could have significantly contributed to
320 building Mars but only played a limited role in contributing to the HED parent body.
321 Consequently, the possible distinct nature of the building blocks of 4-Vesta and Mars would
322 argue in favor of the existence of chemical heterogeneities in the early solar system. While
323 still a matter of debate (e.g., Lewis, 1974; Morgan and Anders, 1980; Palme 2000; Taylor and
324 Scott, 2003; Poitrasson et al., 2005), the existence of a temperature gradient in the solar
325 nebula may have induced a chemical gradient following early condensation, with more
326 oxidized and volatile-rich materials (C-chondrites-like) occurring preferably farther from the
Sun and more reduced materials (E-chondrites-like) occurring in the inner parts of the protoplanetary accretion disk.

329

### 4.3. Constraints on the Earth's core composition

Using the signature of the BSE estimated above ($\delta^{30}\text{Si} = -0.27\text{‰}$) and assuming the Earth accreted from chondritic materials, the difference in isotopic composition between the bulk silicate Earth and its chondritic building blocks can be calculated using $\Delta^{30}\text{Si}_{\text{BSE-Chond.}} = \delta^{30}\text{Si}_{\text{BSE}} - \delta^{30}\text{Si}_{\text{Chond.}}$. Assuming that this difference is due to metal-silicate partitioning during mantle-core differentiation, $\Delta^{30}\text{Si}_{\text{BSE-Chond.}}$ can be used to constrain both the Si concentration and isotopic composition of the core. However, the conditions under which mantle-core differentiation occurred must be known. Temperature ($T$) is a critical parameter for isotope fractionation and typical equilibrium fractionations scale with $1/T^2$ for high temperatures (Bigeleisen and Mayer, 1947). It is therefore generally accepted that the temperature dependence of the fractionation law can be written as:

$$ \epsilon_{\text{Si}}(T) = \delta^{30}\text{Si}_{\text{Silicate}(T)} - \delta^{30}\text{Si}_{\text{Metal}(T)} = \frac{B \cdot 10^6}{T^2} \qquad (2) $$

where $\epsilon_{\text{Si}}$ denotes the fractionation factor of Si given in permil and defined as $\epsilon = 1000 \times \ln(\alpha)$.

The value of the constant $B$ has been estimated using either natural samples (7.64 ± 0.47, Ziegler et al., 2010) or from metal-silicate fractionation experiments (7.45 ± 0.41, Shahar et al., 2011). As mantle-core differentiation occurred under high-pressure conditions, pressure ($P$) may also have influenced Si fractionation, in particular through changes in the silicate melt structure. However, theoretical calculations performed by Shahar et al. (2009) suggest that the effect of $P$ is only secondary. In the following, we will therefore consider that the isotope fractionation during accretion and core formation is only temperature dependent and
follows Equation (2).

Most recent work (Georg et al., 2007; Fitoussi et al., 2009; Shahar et al., 2009; Armytage et al., 2011; Fitoussi and Bourdon, 2012) has modeled Si isotope fractionation during mantle-core differentiation as a steady-state process at constant T, which gives the following relation:

$$ \varepsilon_{\text{Si}}(T) = \frac{\Delta^{30}\text{Si}_{\text{BSE-Chond.}}}{F_{\text{Core}}} \qquad (3) $$

where $F_{\text{Core}}$ represents the Si mass fraction of the Earth present in the core. The concentration of Si in the core ($C_{\text{Si}}^{\text{Core}}$) is then derived from:

$$ C_{\text{Si}}^{\text{Core}} = \left( \frac{m_{\text{BSE}}}{m_{\text{Core}}} \right) \cdot \left( \frac{F_{\text{Core}}}{1 - F_{\text{Core}}} \right) \cdot C_{\text{Si}}^{\text{BSE}} \qquad (4) $$

knowing that $m_{\text{BSE}}/m_{\text{Core}}$, the BSE/core mass ratio, is 2.1 (see e.g., Yoder, 1995) and that the concentration of Si in BSE ($C_{\text{Si}}^{\text{BSE}}$) is 21wt% (see e.g., McDonough and Sun, 1995; O'Neill and Palme, 1998; Palme and O'Neill, 2003).

It should be noted that, in an effort to account for varying conditions of core-mantle differentiation, some authors have also considered two-stages models (Fitoussi et al. 2009), each stage corresponding to a steady-state process. The major problem with the single- or two-stage(s) approach is that geochemical and dynamical constraints indicate that core formation was a continuous process rather than a discrete one (e.g., Wetherill, 1986; Canup and Asphaug, 2001; Halliday, 2008; Wood and Halliday, 2010). It is therefore more appropriate to use continuous core formation models such as the ones proposed by Wade and Wood (2005) and Corgne et al. (2008). In these models based on metal-silicate partitioning of siderophile elements, T, P and oxygen fugacity ($fO_2$) are allowed to continuously evolve as core formation proceeds. As these parameters have a major influence on elemental Si metal-silicate partitioning (e.g., Corgne et al., 2008), it is crucial to consider their effects in details to obtain the most precise estimate of the Si content in the Earth's core. We thus use an approach
similar to that of Chakrabarti and Jacobsen (2010a), but in the present study all parameters (*T*, *P* and *fO*₂) are constrained by siderophile element abundances in the BSE and their metal-silicate partitioning behavior.

### 4.4. Model of continuous core formation

Continuous core formation is modeled as follows:

a. Core formation takes place during Earth's accretion, with final metal-silicate equilibration at the base of a homogeneous magma ocean.

b. Growth of the Earth is linearly modeled in 100 steps. At each step, added metal represents 32.3% of the added mass and is equilibrated with the total mass of the mantle. The metal then migrates to the core without further re-equilibration.

c. The pressure of final equilibration at the base of the magma ocean is assumed to increase proportionally to the pressure at the core-mantle boundary (CMB) of the growing Earth. The pressure at the base of the magma ocean is fixed at 40% of the pressure at the CMB for nickel (Ni) partitioning data (essentially pressure-dependent) to satisfy the well-constrained Ni content of the BSE (Corgne et al., 2008). This means that, at the end of accretion, the pressure at the CMB is ~135 GPa while the pressure at the base of the magma ocean is ~54 GPa.

d. The peridotite solidus as defined by Fiquet et al. (2010) provides the variation of *T* at the base of the magma ocean during accretion (hence during core formation (Fig. 5a)). Average *T* and *P* conditions of core formation are about 2900 K and 33 GPa, respectively.

Siderophile element abundances in the BSE have been shown to be consistent with a continuous accretion model (Wade and Wood, 2005; Corgne et al., 2008; Wood et al., 2008).

In particular, nickel and vanadium provide tight constraints on the required conditions of *T*, *P*
and fO₂ during accretion and core formation, as shown in Fig. 5b. Variations of the core/BSE Si elemental ratio during accretion can be calculated with the T-P-fO₂ conditions and the parameterization of the Si metal-silicate partition coefficient (D<sub>Si</sub>) given by Corgne et al. (2008):

$$ \log D_{\text{Si}} = 2.32(\pm 0.3) - \frac{21800}{T} - \frac{11(\pm 33) \cdot P}{T} - 2\log\left(\frac{x_{\text{FeO}}^{\text{Silicate}}}{x_{\text{Fe}}^{\text{Metal}}}\right) - \log \gamma_{\text{Si}}^{\text{Metal}} \quad (5) $$

where T is the temperature (K), P is the pressure (GPa), x<sub>FeO</sub><sup>Silicate</sup> is the molar fraction of FeO in the silicate mantle, x<sub>Fe</sub><sup>Metal</sup> is the molar fraction of Fe in the metallic droplet, and γ<sub>Si</sub><sup>Metal</sup> is the activity coefficient in the metallic phase calculated from Dresler (1989). As shown in Figure 5b for a fO₂ path satisfying siderophile element constraints (from Wood et al., 2008), the core/BSE concentration ratio of Si is poorly modeled due to large uncertainties associated with parameterizing the partitioning (Eq. 5). It is thus impossible to derive an accurate Si content of the core using this approach (Corgne et al., 2008). Values inferred this way span a huge and partly unrealistic range from 1 to 50wt%, considering 1SD uncertainty on the regression coefficients. As shown below, isotopic constraints on BSE and chondrites derived in section 3 can greatly reduce the range of estimates for the Si content of the core.

In the following, we present the procedure used to derive the isotopic signatures of core and BSE. At step *i*, the isotope signature of the added metal (AM) subsequently going to the core (δ³⁰Si<sub>AM(i)</sub>) can be calculated knowing the BSE isotope signature and the metal-silicate fractionation law at the base of the magma ocean:

$$ \delta^{30}\text{Si}_{\text{AM(i)}} = \delta^{30}\text{Si}_{\text{BSE(i)}} - \varepsilon_{\text{Si}}(T_{(i)}) \quad (6) $$

From mass-balance considerations, the following relationship can be derived between the isotope signatures at step *i* and step *i+1*:

$$ \delta^{30}\text{Si}_{\text{BSE(i)}} \cdot m\text{Si}_{\text{BSE(i)}} + \delta^{30}\text{Si}_{\text{Chond.}} \cdot m\text{Si}_{\text{Chond.}} = \delta^{30}\text{Si}_{\text{BSE(i+1)}} \cdot m\text{Si}_{\text{BSE(i+1)}} + \delta^{30}\text{Si}_{\text{AM(i+1)}} \cdot m\text{Si}_{\text{AM(i+1)}} \quad (7) $$

where mSi<sub>BSE(i)</sub>, mSi<sub>AM(i)</sub> and mSi<sub>Chond.</sub> are the Si mass in BSE and added metal at step *i*, and
the added Si mass from chondrites at each step, respectively. Combining (6) and (7) yields:

$$425 \qquad \delta^{30}\text{Si}_{\text{BSE}(i+1)} = \frac{\delta^{30}\text{Si}_{\text{BSE}(i)} \cdot m\text{Si}_{\text{BSE}(i)} + \varepsilon_{\text{Si}}(T_{(i+1)}) \cdot m\text{Si}_{\text{AM}(i+1)} + \delta^{30}\text{Si}_{\text{Chond.}} \cdot m\text{Si}_{\text{Chond.}}}{m\text{Si}_{\text{BSE}(i+1)} + m\text{Si}_{\text{AM}(i+1)}} \qquad (8)$$

Since we are interested in isotope signatures with respect to the initial chondritic material, $\delta^{30}\text{Si}_X$ can be replaced by $\Delta^{30}\text{Si}_{X-\text{Chond.}}$ in equation (8), with $\Delta^{30}\text{Si}_{X-\text{Chond.}} = \delta^{30}\text{Si}_X - \delta^{30}\text{Si}_{\text{Chond.}}$, giving:

$$429 \qquad \Delta^{30}\text{Si}_{\text{BSE-Chond.}(i+1)} = \frac{\Delta^{30}\text{Si}_{\text{BSE-Chond.}(i)} \cdot m\text{Si}_{\text{BSE}(i)} + \varepsilon_{\text{Si}}(T_{(i+1)}) \cdot m\text{Si}_{\text{AM}(i+1)}}{m\text{Si}_{\text{BSE}(i+1)} + m\text{Si}_{\text{AM}(i+1)}} \qquad (9)$$

The value of $\Delta^{30}\text{Si}_{\text{BSE-Chond.}}$ at the end of core formation is calculated using Eq. (9) with the initial conditions $\Delta^{30}\text{Si}_{\text{BSE-Chond.}(0)} = 0$ and $m\text{Si}_{\text{BSE}(0)} = 0$ and the final condition $m\text{Si}_{\text{BSE}(100)}$ given by current BSE concentration models (21wt% Si, e.g., McDonough and Sun, 1995; O'Neill and Palme, 1998), and using Eq. (5) to derive Si mass partitioning between the added metal ($m\text{Si}_{\text{AM}}$) and silicate ($m\text{Si}_{\text{BSE}}$). The isotopic signature of the core at step $i$ is then given by the equation:

$$436 \qquad \Delta^{30}\text{Si}_{\text{Core-Chond.}(i)} = \frac{\sum_{j=1}^{i} \Delta^{30}\text{Si}_{\text{AM-Chond.}(j)} \cdot m\text{Si}_{\text{AM}(j)}}{\sum_{j=1}^{i} m\text{Si}_{\text{AM}(j)}} \qquad (10)$$

Figure 5c presents the theoretical variations of $\Delta^{30}\text{Si}_{\text{BSE-Chond.}}$ during accretion for a typical oxygen fugacity path satisfying siderophile element constraints, and with the fractionation law given by Ziegler et al. (2010) (Eq. (2) with B=7.64). Given the large uncertainties on elemental partitioning which are incorporated into the model (see for instance Fig. 5b), the predicted difference in isotopic signature between BSE and chondrites is also highly uncertain, ranging between 0.03‰ and 0.5‰ in the example considered here (Fig. 5c). To circumvent this difficulty, we calculated the variation during accretion of the $\Delta^{30}\text{Si}_{\text{BSE-Chond.}(i)/F_{\text{Core}(i)}}$ ratio, where $F_{\text{Core}(i)}$ is the Si fraction of the Earth contained in the core at step $i$. The following relationship links $F_{\text{Core}(i)}$ to the partition coefficient $D_{\text{Si}}$:
$$ \frac{1}{F_{\text{Core}(i)}} = \frac{1}{D_{\text{Si}(i)}} \frac{m_{\text{BSE}}}{m_{\text{Core}}} + 1 \qquad (11) $$

where $m_{\text{BSE}}/m_{\text{core}}$, the BSE/core mass ratio, equals 2.1 (e.g. Yoder, 1995). For a single step steady-state process (a single-stage of core formation at a fixed $T$), $\Delta^{30}\text{Si}_{\text{BSE-Chond.}}/F_{\text{Core}}$ is simply equal to $\varepsilon_{\text{Si}}$ (Eq. 2). Considering a continuous accretion scenario, the $\Delta^{30}\text{Si}_{\text{BSE-Chond.}(i)}/F_{\text{Core}(i)}$ ratio (Fig. 6) can be seen as the integrated fractionation factor during accretion, up to step $i$ ($\Delta^{30}\text{Si}_{\text{BSE-Chond.}(100)}/F_{\text{Core}(100)} = \varepsilon_{\text{Si}}$) and taking into account all relevant effects (effect of $T$ on isotope fractionation, $T-P-fO_2$ effects on elemental Si partitioning, effect of isotope dilution in the mass-increasing BSE). An important outcome of the model is that the value of $\Delta^{30}\text{Si}_{\text{BSE-Chond.}(100)}/F_{\text{Core}(100)}$ ($= \varepsilon_{\text{Si}}$) at the end of accretion varies little ($0.88 \pm 0.06$), regardless of the $fO_2$ path considered (all paths span a realistic range (see insert of Fig. 6), while remaining consistent with siderophile element constraints). To evaluate the total uncertainty on $\varepsilon_{\text{Si}}$ inferred from the model, it is important to also take into account the uncertainty on $T$ through Earth's accretion. Given that $T$ is constrained by the peridotite solidus (Fiquet et al., 2010), we can evaluate the impact of the $T$ uncertainty on the final $\varepsilon_{\text{Si}}$ (or $\Delta^{30}\text{Si}_{\text{BSE-Chond.}(100)}/F_{\text{Core}(100)}$) using the uncertainties from their experiments ($\pm 150$K, see Fiquet et al., 2010). This leads to an additional $-0.053/+0.047$ uncertainty (approximated to $\pm 0.05$), resulting in a final estimate of $\Delta^{30}\text{Si}_{\text{BSE-Chond.}(100)}/F_{\text{Core}(100)} = 0.88 \pm 0.11$. In a steady state model at constant $T$ (Eq. (3)), this $\Delta^{30}\text{Si}_{\text{BSE-Chond.}(100)}/F_{\text{Core}(100)}$ value would correspond to core formation at a fixed $T$ of 2940K, close to the 3000K value proposed by Ziegler et al. (2010).

In the approach detailed above, we have used the fractionation law of Ziegler et al. (2010). More generally, it can be shown that $\Delta^{30}\text{Si}_{\text{BSE-Chond.}}/F_{\text{Core final}}$ is proportional to $B$ (as in Equation (3)), yielding:

$$ \Delta^{30}\text{Si}_{\text{BSE-Chond.}}/F_{\text{Core final}} = (0.115 \pm 0.008) * B \qquad (12) $$
The integrated fractionation factor can then be used together with the isotope signatures of BSE and chondrites to estimate $F_{\text{Core}}$, and hence the amount of Si present in the core, using Eq. (4). Thus we derive a minimum estimated Si core content of $12_{-2}^{+4}$ wt% using O/C-chondrites as building material for the Earth ($\Delta^{30}\text{Si}_{\text{BSE-O/CC}} = 0.19$), and a maximal $29_{-9}^{+21}$ wt% Si using E-chondrites ($\Delta^{30}\text{Si}_{\text{BSE-EC}} = 0.35$). Any mix of chondrite types as starting materials will thus result in a $0.19 < \Delta^{30}\text{Si}_{\text{BSE-Chond}} < 0.35$, and therefore produce an estimated Si-core concentration in-between 12 and 29wt% (Fig. 7). The corresponding $\delta^{30}\text{Si}$ isotope signature of the core is then approximately $-1.1 \pm 0.2\%$ according to Eq. (10).

## 4.5 Application of the model to literature data

Applying our model to previously published data sets from various laboratories (Georg et al., 2007; Fitoussi et al., 2009; Chakrabarti and Jacobsen 2010a; Armytage et al., 2011, 2012; Fitoussi and Bourdon 2012, Savage and Moynier 2013) yields estimated Si-core concentrations varying between about 1 and 13wt% using O/C-chondrites, and between 11wt% and 35wt% using E-chondrites (see Table 2 for details). Since the model is very sensitive to the $\Delta^{30}\text{Si}_{\text{BSE-Chond}}$ value, it is extremely important to find the most accurate estimates of both $\delta^{30}\text{Si}_{\text{BSE}}$ and $\delta^{30}\text{Si}_{\text{Chond}}$. In order to improve the statistical robustness of our model, below we computed a $\Delta^{30}\text{Si}_{\text{BSE-Chond}}$ that incorporates the previously published data. Given the significant offsets in both the Georg et al. (2007) and Chakrabarti and Jacobsen (2010a) results from all other recently published data including ours, we omit the former from the calculation. Thus using data from Fitoussi et al., (2009), Savage et al., (2010), Armytage et al., (2011), Fitoussi et Bourdon (2012), Savage and Moynier (2013) and this study, we estimate an average $\delta^{30}\text{Si}_{\text{BSE}} = -0.28 \pm 0.01\%$, (2SE, n=62), $\delta^{30}\text{Si}_{\text{O/CC}} = -0.46 \pm 0.01\%$, (2SE, n=37), and $\delta^{30}\text{Si}_{\text{EC}} = -0.64 \pm 0.03\%$, (2SE, n=26). This result in a $\Delta^{30}\text{Si}_{\text{BSE-O/CC}} = 0.18$, which leads to minimum estimate of $12_{-3}^{+3}$ wt% Si expected in the core, and a $\Delta^{30}\text{Si}_{\text{BSE-EC}} =$
495 0.36, leading to a maximum estimate of $31_{-10}^{+19}$ wt% Si expected in the core (see Table 2, Fig. 7). It is worth noting that, although absolute values are apparently shifted, data from Georg et al. (2007) lead to similar results ($\Delta^{30}\text{Si}_{\text{BSE-O/CC}} = 0.20$ and Si-core = 13wt%, and/or $\Delta^{30}\text{Si}_{\text{BSE-EC}} = 0.31$ and Si-core = 24wt%, see Table 2) on average.

499

500 **4.6. Did the Earth accrete from chondrites only?**

501 Assuming that mantle-core differentiation is the sole process responsible for Si isotope
502 fractionation during Earth's formation, our model suggests that the minimum expected Si core
503 concentration is ~12wt%, using most of the available isotopic data (Table 2, Fig. 7). Such a
504 value exceeds most of the estimates based on the expected density of Fe-Si alloys at core
505 conditions (e.g., Sherman, 1997; Alfè et al., 2002; Dobson et al., 2003; Badro et al. 2007;
506 Brosh et al., 2009; Antonangeli et al., 2010; Ricolleau et al., 2011; Bouhifd et al., 2012) or
507 geochemical mass balance calculations (Allègre et al., 1995; 2001; Javoy et al., 2010). The
508 latter usually limit the Si core content at between 1 and 7wt% (in most cases between 2 and
509 5wt%), with the exception of a few models (Ricolleau et al., 2011) that predict higher Si core
510 contents, up to 11wt%.

511 Consequently, if the actual Si concentration of the Earth's core does not exceed 5 to
512 7wt%, and if mantle-core differentiation is effectively the sole process responsible for the
513 heavy isotopic signature of the BSE, then our model does not argue in favor of an Earth
514 purely made of chondrites, but rather implies that the Earth has at least partly accreted from
515 non-chondritic materials. Such a scenario is worth being considered further since it has
516 already been suggested by several other studies on the basis of, for example, non-chondritic
517 $^{142}$Nd/$^{144}$Nd ratio (Caro et al., 2008, 2010). Hence, assuming the Earth has formed from non-
518 chondritic materials, these latter should have had a heavier Si isotope composition than that of
519 chondrites (i.e., $\delta^{30}\text{Si} > -0.39\text{‰}$).
520

## 4.7. Silicon isotope fractionation by vaporization during a giant interplanetary impact?

Preferential loss of Si to outer space during the Earth's accretion may provide an alternative explanation for both the superchondritic terrestrial Mg/Si ratio and the Si isotope fractionation observed between the BSE and chondrites. A study by Knight et al. (2009) shows that Si vaporization from calcium-aluminum-rich inclusions (CAIs) in meteorites results in very strong Si isotope fractionation. These authors found a constant fractionation factor $\alpha^{29}\text{Si} = 0.98985 \pm 0.00044$ for the 1600-1900°C temperature range, with the gas phase being enriched in light isotopes. Considering a ~2000K Si degassing source (e.g., Poitrasson et al., 2004; Pahlevan and Stevenson, 2007), data from Knight et al. (2009) suggest that even small amounts of vaporized Si can significantly fractionate Si isotopes. Following a simple Rayleigh distillation model, we can estimate that loss of between 0.9 and 1.6wt% of the bulk Earth's Si loss through volatilization can explain the isotope fractionation we measured between the BSE and chondrites, without the need of Si isotope fractionation during mantle-core differentiation (Fig. 8A). If one now considers a chondritic Earth with 7wt% Si core content (which can be translated into a 0.12‰ decrease of the starting $\Delta^{30}\text{Si}_{\text{BSE-Chond}}$ via mantle-core differentiation; see Fig. 7), a loss of 0.4 to 1.1wt% bulk Earth's Si is sufficient to reconcile our partitioning model with isotopic observations (Fig. 8B). Interestingly, this result is almost identical to the Fe loss required through partial vaporization (between 0.5 and 1wt%) to explain the observed Fe isotope differences between the Earth-Moon couple and the other solar system planetary bodies (Poitrasson et al., 2004).

The most widely accepted model for lunar formation involves a violent impact between the proto-Earth and a Mars-sized body (e.g., Canup and Asphaug 2001; Poitrasson et al., 2004; Touboul et al., 2007). In this scenario, the ejected materials should have suffered significant vaporization during and after the impact, accompanied by isotope fractionation.
Elliptical trajectories of the ejecta (Cameron, 2000; Canup and Asphaug, 2001) would have facilitated the loss of light isotopes escaping from the Earth-Moon system (Poitrasson et al., 2004). This event would have been followed by a thorough homogenization process between the material that subsequently re-accreted to form the Earth and the Moon, yielding undistinguishable Si isotope compositions for the two bodies (Nimmo and Agnor, 2006; Pahlevan and Stevenson, 2007, 2011; Touboul et al., 2007). Noteworthy, the gas pressure during CAI formation was most likely different – and lower – than in the aftermath of the impact, so that the fractionation factor provided by the CAI data is possibly too high for the conditions of Earth accretion (e.g., Young et al., 2002). Unfortunately, the “true” Si isotope fractionation factor during large impacts is unknown, hence this model should be considered as an end-member scenario.

While Fe and Si show significant mass-dependent isotopic differences between the Earth-Moon system, Mars and 4-Vesta (e.g., Poitrasson et al., 2004; Georg et al., 2007), this is not the case for elements such as O, K and Mg (e.g., Humayun and Clayton, 1995; Wiechert et al., 2001; Bourdon et al., 2010; Teng et al., 2010). The lack of observable mass-dependent interplanetary isotope fractionation may result from the large fractionations observed within planets for O, or because of analytical issues for K (see Poitrasson et al., 2004). A possible explanation for the difference between Mg on one hand and Fe and Si on the other hand could be the fact that Si and Fe occur in significant quantities in the Earth's core (e.g., Shahar et al., 2009, 2011; Corgne et al., 2009; Ricolleau et al., 2011). One could thus hypothesize that the elemental losses and isotopic fractionation in Si and Fe preferentially result from the partial vaporization of metallic core components in the aftermath of the giant impact (see Poitrasson et al., 2004). Such a scenario suggests that the ejected Si-rich core particles have lost light Si isotopes to space before re-accretion. These core particles should then have acquired a heavier isotopic signature and subsequently re-equilibrated with
570 mantle components within the turbulent accretion disk of the new Earth-Moon system (e.g.,
571 Pahlevan and Stevenson, 2007). This process may therefore have ultimately led to an
572 isotopically heavier mantle signature. This hypothesis is in agreement with numerical
573 simulations of interplanetary impacts that show significant ejection of metallic core particles
574 (e.g., Cameron, 2000). Additionally, experiments have demonstrated that the vaporization
575 flux of elements is much more effective if they are initially present in metal rather than in
576 oxide form (Wang et al., 1994).

577 The feasibility of this process can be evaluated by mass balance calculations.
578 Assuming the Earth's core contains 7wt% Si prior to the impact, and considering the need to
579 vaporize 1.6% (worst possible case, i.e., using only E-chondrites, see Fig. 8) of the bulk
580 Earth's Si preferentially from the core, this means that ~10% of the Si present in the core has
581 to be lost in the aftermath of the impact, which will eventually drive the Si-core concentration
582 to 6.2wt% after re-accretion. This concentration remains very close to the initially assumed
583 7wt% Si-core content. If correct, this evaluation thus suggests that only marginal amounts of
584 Si are lost from the core after the impact, given the current uncertainties on the Si-core
585 content. This result also suggests that the minimum amount of bulk core that needs to be
586 processed in such a case is 10%. Although no upper limit can be given, this still implies that a
587 relatively limited core proportion may have been processed through this mechanism. In
588 summary, the combination of a volatilization-driven isotopic fractionation process and
589 mantle-core differentiation can produce the observed heavier Si isotope signature of the
590 silicate Earth and Moon relative to chondrites.

591

592 **5. CONCLUSIONS**

593 Significant differences in Si isotope composition were found among planetary bodies
594 of the inner solar system, in agreement with most previous studies. The bulk silicate Earth and
595

the bulk silicate Moon have the same Si isotope composition, both are isotopically heavier
than O/C-chondrites, E-chondrites, achondrites from 4-Vesta and Martian meteorites. This
argues in favor of a common origin for the Earth and the Moon.

596

597

598

Coupling Si isotope constraints with a realistic continuous accretion core formation
model allows us to estimate the amount of Si incorporated into the Earth's core. Assuming
that mantle-core differentiation is the sole process that fractionates Si isotopes, our model
suggests that the Earth's core has incorporated at least ~12wt% Si. While these results are
consistent with Si being a major light element in the Earth's core, this refined model requires
higher Si core contents than permitted by geophysical and experimental constraints.

599

600

601

602

603

604

We propose two explanations to reconcile the geochemical, dynamical and
geophysical data:

605

606

1) The Earth has – at least partly – accreted from non-chondritic materials, the latter
having on average heavier Si isotopic signature than chondrites (i.e., δ³⁰Si > -0.39‰). Such
materials would be required to reduce the isotopic difference Δ³⁰SiBSE-Bulk Earth to 0.12‰ or
less, hence limiting the Si core concentration to a reasonable value (i.e., < 7wt%).

607

608

609

610

2) If instead the Earth accreted only from chondritic materials and the Si core content
is <12wt% as dictated by geophysical constraints, then mantle-core differentiation cannot be
the sole mechanism from which planetary Si isotopes variations originate. An additional
process is then required, particularly if the Earth accreted significant amount of E-chondrite
material during its formation. We propose that volatilization-driven isotopic fractionation in
the aftermath of the Moon's forming giant impact could be this process. This mechanism
could have preferentially affected the metallic core, which explains the lack of isotopic
variations observed for lithophile elements, such as Mg, among planetary bodies. This
scenario requires a final isotopic re-equilibration on a planetary scale, leading to the heavier
isotopic compositions of the lunar and terrestrial mantles.

611

612

613

614

615

616

617

618

619
Finally, it is clear from our modeling that the mantle-core differentiation process alone cannot explain the observed terrestrial planet Si isotope compositions. Our current level of knowledge however, does not allow yet favoring one of the two hypotheses (i.e., non-chondritic Earth and/or impact generated Si vaporization) proposed in this study.

624

625 Acknowledgements

This study would not have been possible without the allocation by the NASA's CAPTEM committee of lunar rocks samples to FP. The Natural History Museum of London and Meteorite Working Group is gratefully acknowledged for the provision of Martian meteorites to MA, as well as the Museum National d'Histoire Naturelle de Paris for providing eucrites samples to GQ. We also thank the Museum d'Histoire Naturelle de Toulouse for providing aliquots of Orgueil and St Sauveur meteorites and the Field Museum of Chicago for a piece of meteorite Abee. Michel Gregoire is thanked for providing part of the terrestrial peridotite samples. Carole Boucayrand, Manuel Henry, Sébastien Gardoll and Jonathan Prunier are gratefully acknowledged for assistance during clean-lab work, as well as Jérôme Chmeleff, Rémi Freydier and Manuela Fehr for analytical assistance on Neptune spectrometers. Craig Lundstrom is warmly thanked for fruitful discussions as well as for English language corrections. MA thanks RCUK for a research fellowship during which this collaborative project was carried out. This work has been funded by a grant from the Programme National de Planétologie (PNP) of CNRS-INSU to FP. Anonymous reviewers are thanked for in depth reviews that helped improving both the clarity and the quality of this manuscript. Associate Editor Mark Rehkämper is also acknowledged for his work on this manuscript and helpful comments.

643
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903
904 **Figure Captions:**

905 Fig. 1: Three-isotope plot showing δ<sup>29</sup>Si vs. δ<sup>30</sup>Si for individual analyses of all the samples
906 from this study. The dashed line represents the theoretical mass dependent fractionation slope
907 defined as: δ<sup>29</sup>Si = δ<sup>30</sup>Si × 0.5092. The long-term reproducibility (2SD) is shown in the
908 bottom left quadrant of the figure.

909
910 Fig. 2: δ<sup>30</sup>Si values relative to NBS-28 for terrestrial and extraterrestrial samples.
911 Uncertainties for individual samples are given as 2SE, computed as 2SD/√(n-1) × t, where t is
912 the Student's correcting factor (Platzner et al., 1997). BSMa stands for bulk silicate Mars, and
913 BSMo means bulk silicate Moon. All data are from Table 1.

914
915 Fig. 3: Comparison of published δ<sup>30</sup>Si data relative to NBS-28 for samples from planetary
916 bodies in the inner solar system. Uncertainties are given as 2SE, computed as 2SD/√(n-1) × t,
917 where t is the Student's correcting factor (Platzner et al., 1997). AGV-2 data is from
918 Zambardi and Poitrasson, (2011). C.C. = carbonaceous chondrites, E.C. = enstatites
919 chondrites, R.M. = reference materials.

920
921 Fig. 4: Comparison of δ<sup>30</sup>Si averages relative to NBS-28 for planetary bodies in the inner
922 solar system from this and previous work. All uncertainties are given as ± 2SE.

923
924 Fig. 5: Model of continuous core formation during Earth's accretion (Wade and Wood, 2005;
925 Corgne et al., 2008; Wood et al., 2008). (a) Variations of *T* of last equilibration (at the base of
926 the magma ocean) and FeO content of the BSE during accretion. The *T* profile is defined by
927 the peridotite solidus (Fiquet et al., 2010). FeO variations correspond to the three-step profile
928 proposed by Wood et al. (2008). Such temperatures and FeO variations lead to modeled
929 depletions of moderately siderophile elements in the BSE consistent with observations (as
930 shown for Ni and V in (b)). (b) Modeled core/BSE variations during accretion for Ni, V and
931 Si using P-T-FeO conditions defined in (a) and parameterizations for metal-silicate
932 partitioning from Corgne et al. (2008) for Ni and Si and Wood et al. (2008) for V. Note the
933 large range of final Si core/BSE ratios (yellow path). (c) Expected variations of Δ<sup>30</sup>Si<sub>BSE-Chond</sub>.
934 showing an important level of uncertainty at the end of accretion, due to large uncertainties in
935 elemental metal-silicate partitioning. Note that the prediction is independent from the type of
936 chondrite chosen as reference.
937

938

939

940

941

942

943

944

945

946

947

948

949

950

951

952

953

954

955

956

957

958

959

960

961

962

963

964

965

966

967

968

969

Fig. 6: Variations of the $\Delta^{30}\text{Si}_{\text{BSE-Chond}}/F_{\text{Core}}$ ratio during continuous core formation for several FeO$_{\text{BSE}}$ paths (insert, same abscissa) satisfying constraints from moderately siderophile elements. The blue dashed curve is the three-step model proposed by Wood et al. (2008) and considered in Fig. 5. Taking into account the uncertainty on elemental metal-silicate Si partitioning, $\Delta^{30}\text{Si}_{\text{BSE-Chond}}/F_{\text{Core}}$ lies along a restricted range at the end of accretion (0.88 ± 0.11), regardless of the chosen FeO path.

Fig. 7: Predicted Si core (wt.%) concentrations as a function of measured $\Delta^{30}\text{Si}_{\text{BSE-Chond}}$. The model (thick curve) is computed using $\Delta^{30}\text{Si}_{\text{BSE-Chond}}/F_{\text{Core}} = 0.88$ (see text and Fig. 6). Thin curves show the ± 0.11 variability of $\Delta^{30}\text{Si}_{\text{BSE-Chond}}/F_{\text{Core}}$, which includes the uncertainties on both fO$_2$ and T variations during the Earth accretion. Experimental constraints (Shahar et al., 2009) and metal-silicate Si isotope partitioning in natural samples (Ziegler et al., 2010) were also taken into account. Any mix of chondrite (whatever their type) shall lead to a $\Delta^{30}\text{Si}_{\text{BSE-Chond}}$ that falls on the red section of the curves. The model indicates that the initial $\Delta^{30}\text{Si}_{\text{BSE-Chond}}$ should have been 0.12 if one assumes a maximum possible Si core concentration of 7wt% after mantle-core differentiation, which is below the minimum value given by O/C-chondrites (0.18). Hence, mantle-core differentiation alone cannot explain the observed different Si isotope composition between chondrites and the bulk silicate Earth if the accreted materials had the composition of chondrites.

Fig. 8: A) Rayleigh models predicting the variations of $\delta^{30}\text{Si}_{\text{Chond}}$ during a vaporization process that may have occurred along with giant interplanetary impact. The equation is defined as $\delta^{30}\text{Si}_{\text{Solid residue}} = \delta^{30}\text{Si}_{\text{Chond}} \times (\alpha^{30}\text{Si} - 1) \times \ln(f)$, where $\alpha$ is the fractionation factor of Si and $f$ is the fraction of Si remaining in the residual solid. The computation involves a Si isotope fractionation factor during vaporization $\alpha^{29}\text{Si} = 0.98985 \pm 0.00044$ (Knight et al., 2009), which has been converted into $\alpha^{30}\text{Si} = 0.97893 \pm 0.00088$, assuming the fractionation process is mass dependent. Starting $\delta^{30}\text{Si}_{\text{Chond}}$ values were calculated from sample measurements and considered as $\delta^{30}\text{Si}_{\text{Bulk Earth}}$. $\delta^{30}\text{Si}_{\text{Chond}}$ values were calculated using O-chondrites (blue), or O/C-chondrites (red). Solid lines represent the corresponding Rayleigh models. The paths delimited by dashed lines takes into considerations both the uncertainties on $\delta^{30}\text{Si}_{\text{Chond}}$ averages and $\alpha^{30}\text{Si}$. The intersection between the model and the black solid line ($\delta^{30}\text{Si}_{\text{BSE}}$ value) gives the fraction Si remaining in the Earth residue at the end of the
970 vaporization process (i.e., when $\Delta^{30}\text{Si}_{\text{BSE-Chond.}} = 0$). These 2 models do not consider Si isotope
971 fractionation from mantle-core differentiation, and therefore are considered as an extreme
972 case. B) The same Rayleigh models assuming that the mantle-core differentiation fractionated
973 Si isotopes prior to vaporization, and the Earth's core incorporated 7wt% Si (thus leading to a
974 0.12‰ decrease of the $\Delta^{30}\text{Si}_{\text{BSE-Chond.}}$ value (Fig. 7) translated on the figure into a 0.12‰ shift
975 on the starting materials). According to these latter models, between 0.4 and 1.1wt% on
976 average of Si loss by vaporization is enough to reconcile a chondritic Earth with isotopic
977 observations while keeping a reasonable amount of Si in the core.
978

978

979 Table Captions:

980 Table 1: Silicon concentration and isotope data for the Earth, Moon and meteorite samples.
981 Uncertainties are given as ± 2SE, computed as $2\text{SD}/\sqrt{(n-1) \times t}$, where t is the Student's
982 correcting factor (Platzner et al., 1997) and where n is the number of replicate analyses of a
983 sample. With the exception of Enstatite chondrites, uncertainties on planetary average values
984 are also given as 2SE, computed using the same relation, where n means the number of
985 individual samples (see text for justification). **Given the E-chondrite dataset only gathers 2
986 samples, a propagated uncertainty calculation was adopted for these data instead of the 2SE
987 calculation. °Data from Zambardi and Poitrasson (2011). *Sample not included in the bulk-
988 silicate-body average calculation. For averages, N means the number of samples.
989

989

990 Table 2: Si-core concentrations inferred from our model on all available $\Delta^{30}\text{Si}_{\text{BSE-Chond.}}$ data
991 from recent literature and this study. All $\Delta^{30}\text{Si}_{\text{BSE-Chond.}}$ uncertainties were propagated from the
992 2SE of both $\delta^{30}\text{Si}_{\text{BSE}}$ and $\delta^{30}\text{Si}_{\text{Chond.}}$. All extreme Si-core concentrations include uncertainties
993 from both the model and the $\Delta^{30}\text{Si}_{\text{BSE-Chond.}}$ estimates. *Data computed using the $\delta^{30}\text{Si}_{\text{BSE (Mean)}}$
994 from Savage et al. (2010), see Table 2 in Armytage et al. (2011). **Data computed using an
995 averaged E-chondrites (EH and EL) $\delta^{30}\text{Si}_{\text{EC}} = -0.67 \pm 0.04$ (2SE, n=13) from Savage and
996 Moynier (2013). ***Data computed from the average of all available analysis of terrestrial
997 and chondrites samples in recent literature, with the exception of Georg et al. (2007) and
998 Chakrabarti and Jacobsen (2010a) data. The $\delta^{30}\text{Si}_{\text{BSE}}$ also includes all data from Savage et al.
999 (2010).
$\delta^{30}\text{Si}$

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

$\delta^{29}\text{Si}$
Chondrites

<figure><img src="image_3.png" /></figure>

<figure><img src="image_4.png" /></figure>

Ordinary chondrites

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

Carbonaceous chondrites

Howardite

<figure><img src="image_6.png" /></figure>

Eucrites

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

HED

<figure><img src="image_8.png" /></figure>

SNC

<figure><img src="image_9.png" /></figure>

Moon

Mare basalts

Norite (Nr)

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

Earth (BSE)

Dunite (Dn)

Komatiite (Km)

Peridotites

Basalt trachyandesite (Bt)

Diorite (Dr)

Dolerite (Dl)

Gabbro (Gb)

Basalts

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

-0.7

-0.6

-0.5

-0.4

-0.3

-0.2

-0.1

$\delta^{30}\text{Si}$ (‰)
<figure><img src="image_12.png" /></figure>

C.C.

E.C.

Eucrites

SNC

Moon

Terrestrial RM
<figure><img src="image_13.png" /></figure>

$\delta^{30}\text{Si} \text{‰}$
<figure><img src="image_14.png" /></figure>

<figure><img src="image_15.png" /></figure>

<figure><img src="image_16.png" /></figure>
<figure><img src="image_17.png" /></figure>
<figure><img src="image_18.png" /></figure>

Expected Si (wt%) in the Earth's core
A

Mass fraction Si remaining in solid Earth

<figure><img src="image_19.png" /></figure>

<figure><img src="image_20.png" /></figure>

B

Mass fraction Si remaining in solid Earth

<figure><img src="image_21.png" /></figure>

<figure><img src="image_22.png" /></figure>
<table><thead><tr><th>Samples</th><th>Nature</th><th>Location</th><th>SiO<sub>2</sub> (wt%)</th><th>δ<sup>30</sup>Si (‰)</th><th>2SE</th><th>δ<sup>29</sup>Si (‰)</th><th>2SE</th><th>n# analyses</th></tr></thead><tbody><tr><td colspan="9"><strong>Terrestrial mafic rocks (BSE)</strong></td></tr><tr><td>°BCR-2</td><td>Basalt</td><td>Columbia river, USA</td><td>54.0</td><td><strong>-0.19</strong></td><td>0.03</td><td><strong>-0.11</strong></td><td>0.03</td><td>6</td></tr><tr><td>°BHVO-2</td><td>Basalt</td><td>Hawaï, USA</td><td>49.9</td><td><strong>-0.27</strong></td><td>0.01</td><td><strong>-0.14</strong></td><td>0.01</td><td>54</td></tr><tr><td>°BIR-1</td><td>Basalt</td><td>Reikjavik, Iceland</td><td>48.0</td><td><strong>-0.27</strong></td><td>0.04</td><td><strong>-0.12</strong></td><td>0.02</td><td>6</td></tr><tr><td>°PMS</td><td>Microgabbro</td><td>Pitscurrie, UK</td><td>46.9</td><td><strong>-0.29</strong></td><td>0.03</td><td><strong>-0.14</strong></td><td>0.02</td><td>6</td></tr><tr><td>°OU-2</td><td>Dolerite</td><td>Belford, UK</td><td>51.0</td><td><strong>-0.25</strong></td><td>0.02</td><td><strong>-0.11</strong></td><td>0.02</td><td>6</td></tr><tr><td>°DR-N</td><td>Diorite</td><td>Neuntelstein, France</td><td>52.9</td><td><strong>-0.26</strong></td><td>0.02</td><td><strong>-0.12</strong></td><td>0.03</td><td>6</td></tr><tr><td>°ES-9104</td><td>Basalt Trachyandesite</td><td>Massif d'Estrel, France</td><td>52.3</td><td><strong>-0.25</strong></td><td>0.02</td><td><strong>-0.11</strong></td><td>0.02</td><td>12</td></tr><tr><td>GM92-234</td><td>Peridotite</td><td>Kerguelen Island</td><td>42.2</td><td><strong>-0.33</strong></td><td>0.05</td><td><strong>-0.16</strong></td><td>0.02</td><td>3</td></tr><tr><td>GR97-38</td><td>Peridotite</td><td>Kerguelen Island</td><td>44.6</td><td><strong>-0.27</strong></td><td>0.04</td><td><strong>-0.12</strong></td><td>0.02</td><td>3</td></tr><tr><td>GR97-52</td><td>Peridotite</td><td>Kerguelen Island</td><td>44.6</td><td><strong>-0.31</strong></td><td>0.07</td><td><strong>-0.14</strong></td><td>0.04</td><td>3</td></tr><tr><td>JGM-92-1c</td><td>Peridotite</td><td>Kerguelen Island</td><td>42.4</td><td><strong>-0.29</strong></td><td>0.01</td><td><strong>-0.14</strong></td><td>0.02</td><td>6</td></tr><tr><td>°JP-1</td><td>Peridotite</td><td>Horoman, Japan</td><td>42.4</td><td><strong>-0.29</strong></td><td>0.06</td><td><strong>-0.13</strong></td><td>0.03</td><td>6</td></tr><tr><td>°KNO3</td><td>Peridotite</td><td>Cameroon</td><td>45.0</td><td><strong>-0.23</strong></td><td>0.06</td><td><strong>-0.11</strong></td><td>0.04</td><td>6</td></tr><tr><td>MG91-3</td><td>Peridotite</td><td>Kerguelen Island</td><td>39.6</td><td><strong>-0.30</strong></td><td>0.06</td><td><strong>-0.16</strong></td><td>0.04</td><td>3</td></tr><tr><td>NK-01</td><td>Peridotite</td><td>Cameroon</td><td>43.5</td><td><strong>-0.29</strong></td><td>0.02</td><td><strong>-0.14</strong></td><td>0.01</td><td>3</td></tr><tr><td>PREM89-1</td><td>Peridotite</td><td>South Africa</td><td>40.9</td><td><strong>-0.28</strong></td><td>0.04</td><td><strong>-0.14</strong></td><td>0.04</td><td>3</td></tr><tr><td>PREM90-24</td><td>Peridotite</td><td>South Africa</td><td>40.9</td><td><strong>-0.30</strong></td><td>0.05</td><td><strong>-0.15</strong></td><td>0.02</td><td>3</td></tr><tr><td>°DTS-2b</td><td>Dunite</td><td>Bellingham, USA</td><td>39.4</td><td><strong>-0.23</strong></td><td>0.05</td><td><strong>-0.10</strong></td><td>0.04</td><td>6</td></tr><tr><td>°WITS-1</td><td>Komatiite</td><td>Barberton, South Africa</td><td>43.3</td><td><strong>-0.26</strong></td><td>0.04</td><td><strong>-0.11</strong></td><td>0.02</td><td>9</td></tr><tr><td><strong>Average BSE</strong></td><td></td><td></td><td></td><td><strong>-0.27</strong></td><td><strong>0.01</strong></td><td><strong>-0.13</strong></td><td><strong>0.01</strong></td><td>N=19</td></tr><tr><td colspan="9"><strong>Lunar samples</strong></td></tr><tr><td>10049</td><td>High Ti mare basalt</td><td>1969, Mare Tranquillitatis</td><td>42.0</td><td><strong>-0.28</strong></td><td>0.03</td><td><strong>-0.13</strong></td><td>0.02</td><td>7</td></tr><tr><td>12054</td><td>Ilmenite basalt</td><td>1969, Oceanus Procellarum</td><td>42.2</td><td><strong>-0.29</strong></td><td>0.05</td><td><strong>-0.13</strong></td><td>0.01</td><td>9</td></tr><tr><td>14053</td><td>Low Ti mare basalt</td><td>1971, Fra Mauro</td><td>46.3</td><td><strong>-0.27</strong></td><td>0.07</td><td><strong>-0.10</strong></td><td>0.02</td><td>6</td></tr><tr><td>15555</td><td>Olv. normative mare basalt</td><td>1971, Hadley Apennines</td><td>44.6</td><td><strong>-0.28</strong></td><td>0.04</td><td><strong>-0.13</strong></td><td>0.02</td><td>8</td></tr><tr><td>77215</td><td>Cataclastic norite</td><td>1972, Taurus Littrow</td><td>51.2</td><td><strong>-0.24</strong></td><td>0.04</td><td><strong>-0.10</strong></td><td>0.03</td><td>6</td></tr><tr><td><strong>Average BSMo</strong></td><td></td><td></td><td></td><td><strong>-0.27</strong></td><td><strong>0.02</strong></td><td><strong>-0.12</strong></td><td><strong>0.02</strong></td><td>N=5</td></tr><tr><td colspan="9"><strong>Carbonaceous chondrites</strong></td></tr><tr><td>NWA-1757</td><td>CO3</td><td>Find 2001, Morocco</td><td>34.1</td><td><strong>-0.45</strong></td><td>0.04</td><td><strong>-0.21</strong></td><td>0.03</td><td>6</td></tr><tr><td>NWA-4425</td><td>CK3</td><td>Find 2006, Algeria</td><td>33.4</td><td><strong>-0.41</strong></td><td>0.01</td><td><strong>-0.20</strong></td><td>0.04</td><td>3</td></tr><tr><td>Orgueil</td><td>CII</td><td>Fall 1864, France</td><td>20.1</td><td><strong>-0.44</strong></td><td>0.03</td><td><strong>-0.23</strong></td><td>0.02</td><td>12</td></tr><tr><td colspan="9"><strong>Ordinary chondrites</strong></td></tr><tr><td>Acfer-380</td><td>H5</td><td>Find 2002, Algeria</td><td>36.2</td><td><strong>-0.49</strong></td><td>0.01</td><td><strong>-0.25</strong></td><td>0.02</td><td>3</td></tr><tr><td>NWA-2258</td><td>L5</td><td>Find 2001, Algeria</td><td>35.6</td><td><strong>-0.47</strong></td><td>0.06</td><td><strong>-0.23</strong></td><td>0.03</td><td>6</td></tr><tr><td>NWA-4278</td><td>L6</td><td>Find 2005, Morocco</td><td>40.3</td><td><strong>-0.46</strong></td><td>0.02</td><td><strong>-0.21</strong></td><td>0.02</td><td>6</td></tr><tr><td>NWA-4260</td><td>L6</td><td>Find 2003, Algeria</td><td>34.9</td><td><strong>-0.47</strong></td><td>0.04</td><td><strong>-0.20</strong></td><td>0.02</td><td>12</td></tr><tr><td>Mocs</td><td>L6</td><td>Fall 1882, Romania</td><td>38.4</td><td><strong>-0.46</strong></td><td>0.06</td><td><strong>-0.21</strong></td><td>0.03</td><td>6</td></tr><tr><td>Tanezrouft-065</td><td>L4</td><td>Find 2002, Algeria</td><td>39.6</td><td><strong>-0.48</strong></td><td>0.01</td><td><strong>-0.24</strong></td><td>0.05</td><td>3</td></tr><tr><td><strong>Average O/C-Chond.</strong></td><td></td><td></td><td></td><td><strong>-0.46</strong></td><td><strong>0.02</strong></td><td><strong>-0.22</strong></td><td><strong>0.02</strong></td><td>N=9</td></tr><tr><td colspan="9"><strong>Enstatite chondrites</strong></td></tr><tr><td>Abee</td><td>EH</td><td>Fall 1952, Canada</td><td>41.1</td><td><strong>-0.64</strong></td><td>0.04</td><td><strong>-0.33</strong></td><td>0.07</td><td>3</td></tr><tr><td>St sauveur</td><td>EH</td><td>Fall 1914, France</td><td>35.4</td><td><strong>-0.59</strong></td><td>0.05</td><td><strong>-0.30</strong></td><td>0.01</td><td>3</td></tr><tr><td><strong>Average E-Chond.</strong></td><td></td><td></td><td></td><td><strong>-0.62</strong></td><td><strong>0.04**</strong></td><td><strong>-0.32</strong></td><td><strong>0.03**</strong></td><td>N=2</td></tr><tr><td colspan="9"><strong>Mars (SNC)</strong></td></tr><tr><td>Chassigny</td><td>Chassignite, dunite</td><td>Fall 1815, France</td><td>36.0</td><td><strong>-0.47</strong></td><td>0.06</td><td><strong>-0.21</strong></td><td>0.02</td><td>7</td></tr><tr><td>DaG-476</td><td>Shergottite, Ol-phyric basalt</td><td>Find 1999, Lybia</td><td>45.4</td><td><strong>-0.43</strong></td><td>0.03</td><td><strong>-0.20</strong></td><td>0.03</td><td>5</td></tr><tr><td>Los Angeles*</td><td>Shergottite, Pxi-phyric basalt</td><td>Find 1980, CA USA</td><td>48.6</td><td><strong>-0.33</strong></td><td>0.06</td><td><strong>-0.16</strong></td><td>0.02</td><td>5</td></tr><tr><td>SaU-005</td><td>Shergottite, Ol-phyric basalt</td><td>Find, Oman</td><td>46.3</td><td><strong>-0.45</strong></td><td>0.02</td><td><strong>-0.21</strong></td><td>0.02</td><td>5</td></tr></tbody></table>
<table><tbody><tr><td>Shergotty</td><td>Shergottite, Px-phyric basalt</td><td>Fall 1865, India</td><td>49.5</td><td><strong>-0.45</strong></td><td>0.06</td><td><strong>-0.21</strong></td><td>0.05</td><td>5</td></tr><tr><td>Zagami</td><td>Basalt Shergottite</td><td>Fall 1962, Nigeria</td><td>49.7</td><td><strong>-0.49</strong></td><td>0.05</td><td><strong>-0.21</strong></td><td>0.06</td><td>7</td></tr><tr><td>Governador</td><td>Nakhlite, cpxnite</td><td>Find 1958, Brasil</td><td>48.0</td><td><strong>-0.51</strong></td><td>0.07</td><td><strong>-0.23</strong></td><td>0.04</td><td>5</td></tr><tr><td>Lafayette</td><td>Nakhlite, cpxnite</td><td>Find 1931, USA</td><td>46.5</td><td><strong>-0.53</strong></td><td>0.06</td><td><strong>-0.25</strong></td><td>0.03</td><td>8</td></tr><tr><td>MIL-03 346</td><td>Nakhlite, cpxnite</td><td>Find 2003, Antarctica</td><td>49.1</td><td><strong>-0.49</strong></td><td>0.06</td><td><strong>-0.21</strong></td><td>0.02</td><td>5</td></tr><tr><td>Nakhla</td><td>Nakhlite, cpxnite</td><td>Fall 1911, Egypt</td><td>48.0</td><td><strong>-0.56</strong></td><td>0.07</td><td><strong>-0.25</strong></td><td>0.02</td><td>5</td></tr><tr><td><strong>Average BSMa</strong></td><td></td><td></td><td></td><td><strong>-0.49</strong></td><td><strong>0.03</strong></td><td><strong>-0.22</strong></td><td><strong>0.01</strong></td><td>N=9</td></tr><tr><td><strong>4-Vesta (HED)</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>ALHA-78132</td><td>Eucrite</td><td>Find 1978, Antarctica</td><td>49.9</td><td><strong>-0.42</strong></td><td>0.03</td><td><strong>-0.20</strong></td><td>0.04</td><td>6</td></tr><tr><td>Bereba</td><td>Eucrite</td><td>Fall 1924, Burkina Faso</td><td>52.7</td><td><strong>-0.38</strong></td><td>0.04</td><td><strong>-0.15</strong></td><td>0.03</td><td>6</td></tr><tr><td>Bouvante</td><td>Eucrite</td><td>Find 1978, France</td><td>50.1</td><td><strong>-0.40</strong></td><td>0.02</td><td><strong>-0.19</strong></td><td>0.03</td><td>6</td></tr><tr><td>Juvinas</td><td>Eucrite</td><td>Fall 1821, France</td><td>50.6</td><td><strong>-0.41</strong></td><td>0.03</td><td><strong>-0.18</strong></td><td>0.02</td><td>7</td></tr><tr><td>NWA-1777</td><td>Eucrite</td><td>Find 2002, Algeria</td><td>49.9</td><td><strong>-0.46</strong></td><td>0.02</td><td><strong>-0.21</strong></td><td>0.02</td><td>6</td></tr><tr><td>Pasamonte</td><td>Eucrite</td><td>Fall 1933, NM USA</td><td>49.3</td><td><strong>-0.38</strong></td><td>0.04</td><td><strong>-0.17</strong></td><td>0.03</td><td>6</td></tr><tr><td>Le Teilleul</td><td>Howardite</td><td>Fall 1845, France</td><td>49.9</td><td><strong>-0.45</strong></td><td>0.05</td><td><strong>-0.20</strong></td><td>0.02</td><td>6</td></tr><tr><td><strong>Average 4-Vesta</strong></td><td></td><td></td><td></td><td><strong>-0.41</strong></td><td><strong>0.03</strong></td><td><strong>-0.19</strong></td><td><strong>0.01</strong></td><td>N=7</td></tr></tbody></table>
<table><caption>Table 2</caption><thead><tr><th></th><th>Δ<sup>30</sup>Si<sub>BSE-O/CC</sub></th><th>Si-core conc. (wt.%)</th><th>Δ<sup>30</sup>Si<sub>BSE-EC</sub></th><th>Si-core conc. (wt.%)</th></tr></thead><tbody><tr><td>Georg et al. (2007)</td><td>0.20 ± 0.05</td><td>13<sup>+9</sup><sub>-5</sub></td><td>0.31 ± 0.04</td><td>24<sup>+14</sup><sub>-7</sub></td></tr><tr><td>Fitoussi et al. (2009)</td><td>0.11 ± 0.04</td><td>6<sup>+5</sup><sub>-3</sub></td><td>0.27 ± 0.05</td><td>20<sup>+13</sup><sub>-7</sub></td></tr><tr><td>Chakrabarti and Jacobsen (2010a)</td><td>0.02 ± 0.04</td><td>1<sup>+3</sup><sub>-1</sub></td><td>0.18 ± 0.05</td><td>12<sup>+8</sup><sub>-5</sub></td></tr><tr><td>Armytage et al. (2011)</td><td>0.15 ± 0.05<br>0.19 ± 0.01*</td><td>9<sup>+7</sup><sub>-5</sub><br>12<sup>+4</sup><sub>-2</sub></td><td>0.31 ± 0.09<br>0.35 ± 0.08*</td><td>24<sup>+26</sup><sub>-11</sub><br>29<sup>+29</sup><sub>-12</sub></td></tr><tr><td>Fitoussi and Bourdon (2012)</td><td>-</td><td>-</td><td>0.34 ± 0.05</td><td>28<sup>+17</sup><sub>-9</sub></td></tr><tr><td>Savage and Moynier (2013)</td><td>-</td><td>-</td><td>0.39 ± 0.04**</td><td>35<sup>+24</sup><sub>-11</sub></td></tr><tr><td>This study</td><td>0.19 ± 0.01</td><td>12<sup>+4</sup><sub>-2</sub></td><td>0.35 ± 0.05</td><td>29<sup>+21</sup><sub>-9</sub></td></tr><tr><td><strong>All data***</strong></td><td>0.18 ± 0.01</td><td>12<sup>+3</sup><sub>-3</sub></td><td>0.36 ± 0.04</td><td>31<sup>+19</sup><sub>-10</sub></td></tr></tbody></table>