# Calcium isotopic fractionation in mantle peridotites by melting and metasomatism and Ca isotope composition of the Bulk Silicate Earth

Jin-Ting Kang, Dmitri Ionov, Fang Liu, Chen-Lei Zhang, Alexander V.

Golovin, Li-Ping Qin, Zhao-Feng Zhang, Fang Huang

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## To cite this version:

Jin-Ting Kang, Dmitri Ionov, Fang Liu, Chen-Lei Zhang, Alexander V. Golovin, et al.. Calcium isotopic fractionation in mantle peridotites by melting and metasomatism and Ca isotope composition of the Bulk Silicate Earth. Earth and Planetary Science Letters, 2017, 474, pp.128-137. 10.1016/j.epsl.2017.05.035. hal-01686768

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Distributed under a Creative Commons Attribution 4.0 International License
Calcium isotopic fractionation in mantle peridotites by
melting and metasomatism and Ca isotope composition of
the Bulk Silicate Earth

Jin-Ting Kang<sup>a,b</sup>, Dmitri A. Ionov<sup>c</sup>, Fang Liu<sup>b,d</sup>, Chen-Lei Zhang<sup>b,d</sup>,

Alexander V. Golovin<sup>e,f,g</sup>, Li-Ping Qin<sup>a</sup>, Zhao-Feng Zhang<sup>b,*</sup>, and Fang Huang<sup>a,*</sup>

a. CAS Key Laboratory of Crust-Mantle Materials and Environments, School of Earth and Space Sciences, University of Science and Technology of China, Hefei 230026, China.

b. State Key Laboratory of Isotope Geochemistry, Guangzhou Institute of Geochemistry, the Chinese Academy of Sciences, Guangzhou 510640, China.

c. Géosciences Montpellier, Université de Montpellier, Montpellier 34095, France

d. University of Chinese Academy of Sciences, Beijing 100049, China

e. Sobolev Institute of Geology and Mineralogy, Siberian Branch Russian Academy of Sciences, Koptyuga 3, Novosibirsk 630090, Russian Federation

f. Novosibirsk State University, Pirogova 2, Novosibirsk 630090, Russian Federation

g. Diamond and Precious Metal Geology Institute, Siberian Branch Russian Academy of Sciences, Yakutsk 677007, Russian Federation

18

19

20

21 5494 words, 7 figures, 2 Tables and 1 Supplementary material

22

23 *Correspondence and requests for materials should be addressed to Fang Huang

24 (fhuang@ustc.edu.cn) or Zhao-Feng Zhang (zfzhang@gig.ac.cn).
25 **ABSTRACT**

26 To better constrain the Ca isotopic composition of the Bulk Silicate Earth (BSE)
27 and explore the Ca isotope fractionation in the mantle, we determined the Ca isotopic
28 composition of 28 peridotite xenoliths from Mongolia, southern Siberia and the
29 Siberian craton. The samples are divided in three chemical groups: (1) fertile,
30 unmetasomatized lherzolites (3.7–4.7 wt.% Al₂O₃); (2) moderately melt-depleted
31 peridotites (1.3–3.0 wt.% Al₂O₃) with no or very limited metasomatism (LREE-
32 depleted cpx); (3) strongly metasomatized peridotites (LREE-enriched cpx and bulk
33 rock) further divided in subgroups 3a (harzburgites, 0.1–1.0% Al₂O₃) and 3b (fertile
34 lherzolites, 3.9–4.3% Al₂O₃). In Group 1, δ⁴⁴/⁴⁰Ca of fertile spinel and garnet peridotites,
35 which experienced little or no melting and metasomatism, show a limited variation from
36 0.90 to 0.99‰ (relative to SRM 915a) and an average of 0.94±0.05‰ (2SD, n=14),
37 which defines the Ca isotopic composition of the BSE. In Group 2, the δ⁴⁴/⁴⁰Ca is the
38 highest for three rocks with the lowest Al₂O₃, i.e. the greatest melt extraction degrees
39 (average 1.06±0.04‰, i.e. ~0.1‰ heavier than the BSE estimate). Simple modeling of
40 modal melting shows that partial melting of the BSE with 10³lnαperidotite-melt ranging
41 from 0.10 to 0.25 can explain the Group 2 data. By contrast, δ⁴⁴/⁴⁰Ca in eight out of nine
42 metasomatized Group 3 peridotites are lower than the BSE estimate. The Group 3a
43 harzburgites show the greatest δ⁴⁴/⁴⁰Ca variation range (0.25–0.96‰), with δ⁴⁴/⁴⁰Ca
44 positively correlated with CaO and negatively correlated with Ce/Eu. Chemical
45 evidence suggests that the residual, melt-depleted, low-Ca protoliths of the Group 3a
harzburgites were metasomatized, likely by carbonate-rich melts/fluids. We argue that such fluids may have low (≤0.25‰) δ<sup>44/40</sup>Ca either because they contain recycled crustal components or because Ca isotopes, similar to trace elements and their ratios, may be fractionated by kinetic and/or chromatographic effects of melt percolation in the mantle. The δ<sup>44/40</sup>Ca in Group 3b lherzolites (0.83–0.89‰) are lower than in the BSE as well, but the effects of metasomatism on δ<sup>44/40</sup>Ca are smaller, possibly because of the high Ca contents in their protoliths and/or smaller δ<sup>44/40</sup>Ca differences between the protoliths and metasomatic agents. The BSE estimates based on fertile peridotites in this study fall in the δ<sup>44/40</sup>Ca ranges for oceanic and continental basalts, various meteorites (achondrites; carbonaceous, ordinary and enstatite chondrites), Mars, and the Moon. These results provide benchmarks for the application of Ca isotopes to planet formation, mantle evolution, and crustal recycling.

58

59 **Keywords:** Ca isotopes; bulk silicate Earth; partial melting; metasomatism; lithospheric mantle; peridotite xenolith

60

61
62

## 1. Introduction

64

Calcium is an alkaline earth element with six stable isotopes: <sup>40</sup>Ca (96.941%), <sup>42</sup>Ca

(0.657%), <sup>43</sup>Ca (0.135%), <sup>44</sup>Ca (2.086%), <sup>46</sup>Ca (0.004%) and <sup>48</sup>Ca (0.187%). Due to its

high abundance in the Earth (e.g. McDonough and Sun, 1995; Rudnick and Gao, 2003)

and large relative mass-differences (~20% between <sup>48</sup>Ca and <sup>40</sup>Ca), Ca isotopes can be

applied to studies of a number of fundamental geochemical processes. Because Ca is

highly refractory with 50% condensation temperature (T<sub>c</sub>) of 1517 K (Lodders, 2003),

Ca isotope data on meteorites can be used to gain insights into planetary building blocks

and the heterogeneity of the proto-planetary disk (e.g. Huang et al., 2012; Schiller et

al., 2015; Simon and DePaolo, 2010; Valdes et al., 2014). Furthermore, because Ca

isotopes can be strongly fractionated during low-temperature biological and

surficial/supergene geological processes (e.g. Fantle and Tipper, 2014), they have the

potential to trace recycled crustal materials in the deep Earth (e.g. DePaolo, 2004;

Huang et al., 2011).

The mantle is the Earth's largest reservoir for Ca. Its Ca isotopic composition can

be roughly assessed using oceanic basalts as 'probes' assuming that they represent the

average isotopic signature of their mantle source regions, as was previously done for

other metal stable isotope systems (e.g. Zn in Chen et al., 2013; Mg in Teng et al., 2007).

The δ<sup>44/40</sup>Ca values [δ<sup>44/40</sup>Ca (‰) = (<sup>44</sup>Ca/<sup>40</sup>Ca)<sub>sample</sub>/(<sup>44</sup>Ca/<sup>40</sup>Ca)<sub>NIST SRM 915a-1</sub>] reported
for oceanic basalts range from 0.62 to 1.54‰ with an average of 0.90±0.28‰ (2SD, n=61) (Fig. 1; Amini et al., 2009; DePaolo, 2004; Holmden et al., 2010; Hindshaw et al., 2013; Huang et al., 2010; Huang et al., 2011; Jacobsen et al., 2015; Skulan et al., 1997; Simon and DePaolo, 2010; Valdes et al., 2014). The large δ<sup>44/40</sup>Ca range for the basalts may indicate either broad Ca isotope variations in the mantle or, alternatively, that the basalts do not mirror the δ<sup>44/40</sup>Ca of their mantle sources due to isotope fractionation during partial melting and magma differentiation and/or addition of external components during magma emplacement. The average δ<sup>44/40</sup>Ca for the basalts is about 0.1‰ lower than the “upper mantle” value estimated by Huang et al. (2010) based on analyses of clinopyroxene (cpx) from two mantle xenoliths. However, due to the scarcity of data on δ<sup>44/40</sup>Ca in peridotites and the broad δ<sup>44/40</sup>Ca variations in the oceanic basalts, it is not clear yet if there is an isotopic offset between the basalts and the mantle. Even less certain may be attempts to infer the composition of the Bulk Silicate Earth (BSE) from data on oceanic basalts because their parental magmas are generated not from pristine but from depleted mantle possibly containing enriched components (e.g. Hofmann, 2014).

Recent studies of pyroxenes in mantle xenoliths from the US and China found considerable inter-mineral Ca isotopic fractionation, e.g. a broad range (0 to 1.1‰) of δ<sup>44/40</sup>Ca between cpx and orthopyroxene (opx) (Huang et al., 2010; Kang et al., 2016). Whole-rock δ<sup>44/40</sup>Ca estimated for the peridotites from the pyroxene data range from 0.7 to 1.1‰, but no whole-rock Ca isotope analyses of mantle xenoliths have been
104 published as yet. $\delta^{44/40}$Ca for three peridotite reference materials (PCC-1, BM90/21-G
105 and DTS-1) were reported by several studies and show ~0.5‰ variation (1.01–1.49‰)
106 (e.g. Amini et al., 2009; Feng et al., 2016; He et al., 2016;), yet it is not clear to what
107 extent Ca isotopes in these materials were affected by late-stage processes like crustal
108 contamination or low-temperature alteration. In addition, no Ca isotope data have been
109 reported so far for garnet-facies peridotites, which are common at depths >60 km in the
110 lithospheric mantle. Therefore, a sufficient number of appropriate peridotites need to
111 be analyzed to precisely characterize the Ca isotopic composition of the upper mantle
112 and the BSE.

113 Here, we report Ca isotope data on 28 well-characterized mantle peridotites from
114 spinel and garnet depth facies and from cratonic and off-craton mantle. These samples
115 represent major mantle rock types including (1) fertile spinel and garnet lherzolites, (2)
116 melt-depleted, unmetasomatized lherzolites and harzburgites, and (3) strongly
117 metasomatized refractory and fertile peridotites. We seek to constrain: (a) the Ca
118 isotope composition of the BSE; (b) the range of Ca isotope variation in the continental
119 lithospheric mantle; (c) the mechanisms responsible for the Ca isotope variations.

120 Our results show that fertile peridotites have a consistent $\delta^{44/40}$Ca of 0.94±0.05‰
121 (2SD, n=14), which we propose as the BSE value. We further find that melt extraction
122 at high melting degrees (>25%) may slightly elevate the $\delta^{44/40}$Ca values of peridotites
123 while metasomatism may decrease their $\delta^{44/40}$Ca values, particularly in low-Ca rocks.

124
125

## 2. Geological background and sample description

127

128 The samples were selected from well-studied collections of mantle xenoliths from central and northeast Asia (see KML map) to represent the range of modal and chemical compositions of peridotite in the continental lithospheric mantle (CLM). The xenoliths are large, modally homogeneous, and show no or negligible alteration and/or contamination by host magma based on petrographic and chemical data. Locality, rock type, bulk-rock contents of Al₂O₃ and CaO, and data sources for each sample are given in Table 2; their CaO, Al₂O₃ and FeO contents are plotted in Fig. 2.

135 The majority of the xenoliths (23 out of 28) are from alkali basaltic pyroclastic rocks in three volcanic fields in central Asia: (1) Vitim east of Lake Baikal in Russia, (2) Tariat in central Mongolia and (3) Dariganga in southeast Mongolia. Three out of four Vitim xenoliths are fertile (3.7–4.6 wt.% Al₂O₃; 3.5–3.6 wt.% CaO) coarse garnet lherzolites hosted by 16 Ma old tuffs; sample 621-16 is a cpx-bearing garnet-spinel harzburgite (1.3 wt.% Al₂O₃, 1.07 wt.% CaO; ~3% cpx) from the 0.8 Ma old Yaksha volcano (Ionov, 2004; Ionov et al., 1993, 2005a). Seventeen out of 18 Tariat xenoliths are from the 0.5 Ma old Shavaryn-Tsaram eruption center (labeled S, ST and STZ), one (H-25) is from the Holocene Haer volcano (Ionov, 2007). The Shavaryn-Tsaram site is famous for the abundance of large, fertile lherzolites brought up from a broad depth range (45–75 km) between the crust-mantle boundary and the garnet-spinel peridotite
146 facies (Ionov et al., 1998; Press et al., 1986). Out of the 18 Tariat samples in this study, 147 ten are fertile, cpx-rich spinel lherzolites, one (ST53389) is fertile garnet-spinel 148 lherzolite, five range from low-cpx spinel lherzolites to harzburgites, and two more 149 (4230-16 and 4230-19) are modally metasomatized fertile spinel lherzolites that contain 150 ~2% of disseminated, texturally equilibrated phlogopite (phl) (Ionov et al., 1997). The 151 single Dariganga xenolith (BN-8) in this study is a spinel harzburgite from the Barun- 152 Nart eruption center (Ionov et al., 1992; Wiechert et al., 1997).

153 The other five samples are from the Udachnaya-East kimberlite in the central 154 Siberian craton, Russia including a coarse garnet harzburgite (419/09), a cpx-bearing 155 (107/03) and two apparently cpx-free (105/03, 137/08) spinel harzburgites, and a 156 sheared garnet lherzolite (U85) (Doucet et al., 2012, 2013; Ionov et al., 2010).

157 In addition to the usual petrographic classification based on mineral assemblages 158 and modal proportions, we identify 3 chemical groups among our samples based on 159 their major oxide contents as well as rare earth element (REE) patterns (Table 2). REE 160 patterns of peridotites are sensitive indicators of either melt extraction (depletion in 161 light REE, LREE) because mineral/melt partition coefficients (K<sub>d</sub>) for the highly 162 incompatible LREE are an order of magnitude lower than for the moderately 163 incompatible heavy REE (HREE) (see Ionov et al., 2002 for references), or of 164 metasomatism (LREE-enrichment) because metasomatic media are LREE-enriched. 165 This is also the reason why some LREE may reside not in cpx, the main REE mineral 166 host in spinel peridotites, but also in interstitial micro-phases (e.g. Hiraga et al., 2004),
such that a bulk rock may have higher LREE/HREE than its cpx. Importantly, the behavior of Ca in the mantle is similar to that for the HREE, but distinct from the behavior of the LREE. In particular, if a rock is LREE-enriched, it does not necessarily imply that the metasomatism also affected its Ca budget. The share of Ca hosted by each mineral in our samples calculated from CaO measured in WR and in the minerals, and modal abundances are given in S-Table 1 of the Supplementary Material. These estimates show in particular that metasomatism may affect more the Ca budget of initially strongly melt-depleted, low-Ca harzburgites (e.g. by depositing interstitial materials) than that of high-Ca, cpx-rich lherzolites.

Group 1 rocks are fertile (3.7–4.7 wt.% Al₂O₃, 3.2–4.0 wt.% CaO), cpx-rich (14–20%) unmetasomatized lherzolites from Vitim and Tariat, including four garnet-bearing (9-13%), which have nearly flat whole-rock (WR) REE patterns and cpx depleted LREE (Fig. 3a). These 14 samples are viewed as pristine (unmelted) mantle rocks or residues of low-degree partial melting. Over 90% of Ca in these rocks are hosted by the cpx (and garnet if present), 3-8% by opx and <2% by olivine.

Group 2 consists of five moderately melt-depleted (1.3–3.0 wt.% Al₂O₃, 0.75–2.3 wt.% CaO) lherzolites and harzburgites from Tariat, Vitim and Dariganga with slightly to moderately LREE-enriched WR patterns (but very low LREE well below those in primitive mantle) and LREE-depleted cpx patterns (Fig. 3b). The LREE depletion in the cpx indicates that this main carrier phase of Ca in the peridotites has not been significantly affected by metasomatism while the bulk-rocks may have been affected
188 by ingress of LREE-bearing fluids, shortly before or during transport in host magma,
189 along grain boundaries (e.g. Wiechert et al., 1997).

190 Group 3 includes two sub-groups of strongly metasomatized peridotites. Group 3a
191 has six ‘anhydrous’ harzburgites from Tariat and Udachnaya, typically very low in Ca,
192 with LREE-enriched WR as well as LREE-enriched and strongly fractionated cpx REE
193 patterns (Fig. 3c). No cpx has been identified in thin sections in two of these samples,
194 but it is possible that fine-grained cpx (or other Ca-rich phases) may be present in
195 interstitial materials; only 14-24% of WR Ca is hosted by opx and olivine.

196 Group 3b combines two phl-spinel lherzolites from Tariat and a sheared garnet
197 lherzolite from Udachnaya; all of them are ‘fertile’ in terms of major oxide contents
198 (3.9–4.3 wt.% Al₂O₃; 3.1–3.9 wt.% CaO), but unlike Group 1 xenoliths, they have
199 LREE-enriched WR and cpx patterns (Fig. 3d) with typical (La/Nd)N≥1 or (La/Ce)N ≥1
200 in the cpx [the subscript N indicates normalization to primitive mantle (McDonough
201 and Sun, 1995)] due to equilibration with LREE-rich silicate melts (Ionov and Hofmann,
202 1995; Ionov et al., 2002). Phl is very low in Ca, with cpx (plus garnet in U85) hosting
203 ≥95% of Ca.

204

205 **3. Analytical methods**

206

207 Chemical purification and mass spectrometric analyses of Ca were done at the State

208 Key Laboratory of Isotope Geochemistry, Guangzhou Institute of Geochemistry (GIG),
209 Chinese Academy of Sciences (CAS). Ca isotope analyses were done on a thermal
210 ionization mass-spectrometer (Thermo Triton). Instrumental mass dependent
211 fractionation was corrected by the double-spike method as described in Heuser et al.
212 (2002). Detailed information about the chemical procedures and mass spectrometry can
213 be found in the Supplementary Material.

214 The internal precision (reproducibility) of δ<sup>44/40</sup>Ca values in this study is ±0.08‰
215 (2SE) based on 3-4 repeated measurements of each sample solution. Two standard
216 deviations of average δ<sup>44/40</sup>Ca in SRM915a measured in the course of this study were
217 ±0.13‰ (n=48), which represents long-term external precision. The values for
218 SRM915a, IAPSO Seawater, and three USGS basaltic reference samples (BCR-2,
219 BHVO-2, and BIR-1) analyzed in the same session, as well as literature data for the
220 same materials are reported in Table 1. The mean values of δ<sup>44/40</sup>Ca obtained for
221 SRM915a and IAPSO Seawater are 0.02±0.13‰ (2SD, n=48) and 1.83±0.13‰ (2SD,
222 n=47), respectively; those for BCR-2, BHVO-2 and BIR-1 are 0.82±0.09‰ (2SD,
223 n=12), 0.80±0.10‰ (2SD, n=16) and 0.84±0.09‰ (2SD, n=9), respectively. All these
224 results are in agreement with data in the literature (e.g. Amini et al., 2009; Feng et al.,
225 2016; He et al., 2016; Magna et al., 2015; Huang et al., 2010; Valdes et al., 2014), which
226 demonstrates the robustness of our analytical methods. In addition, full duplicates of 4
227 samples (S-14, S-62, U85 and UV-105/03) obtained by digestion of duplicate batches
228 of rock powder show good reproducibility within the analytical errors (Table 2).

229
230 **4. Results**

231

232 Ca isotopic compositions are reported in Table 2 and shown in Fig. 4. The 14 Group

233 1 rocks (fertile, unmetasomatized spinel and garnet lherzolites) have a narrow range of

234 δ<sup>44/40</sup>Ca from 0.90±0.07‰ (2SE, n=3) to 0.99±0.08‰ (2SE, n=3), with an average of

235 0.94±0.01‰ (n=14, 2SE). Two moderately melt-depleted (2.2–3.0 wt.% Al<sub>2</sub>O<sub>3</sub>, 1.8–2.3

236 wt.% CaO) Group 2 lherzolites (S-16 and S-22) have δ<sup>44/40</sup>Ca of 0.97±0.09‰ and

237 0.98±0.07‰ (n≥3, 2SE), close to the upper limit of the range for Group 1. By contrast,

238 the three strongly melt-depleted (1.3–1.7 wt.% Al<sub>2</sub>O<sub>3</sub>, 0.8–1.3 wt.% CaO) Group 2

239 harzburgites from 3 localities (BN-8, 621-16 and S-62) show distinctively heavier

240 δ<sup>44/40</sup>Ca values (1.03–1.09‰) with an average of 1.06±0.04‰ (n=3, 2SE).

241 The Group 3 (metasomatized) rocks have a broad δ<sup>44/40</sup>Ca ranging from 0.96±0.08‰

242 (n=3, 2SE) to 0.25±0.01‰ (n=3, 2SE), which extends to much lighter values than for

243 samples in the other two groups (with no or only minor metasomatism). The fertile

244 Group 3b rocks (two phl-spl lherzolites from Tariat and a sheared garnet lherzolite from

245 Udachnaya) have δ<sup>44/40</sup>Ca from 0.83±0.04‰ to 0.89±0.07‰ (n=3, 2SE), defining a

246 tight cluster just below the range of unmetasomatized rocks (Fig. 4). By contrast, the

247 δ<sup>44/40</sup>Ca range in the six ‘anhydrous’ Group 3a harzburgites from Udachnaya and Tariat

248 (0.25–0.96‰) is much greater and includes four samples with significant low δ<sup>44/40</sup>Ca

249 (0.25–0.72‰).

250
251 **5. Discussion**

252

253 **5.1. Assessing mantle Ca isotope compositions via analyses of bulk rocks vs. minerals**

254 The pioneering study of Huang et al. (2010) reported Ca isotope analyses of cpx
255 and opx from two peridotite xenoliths from southwestern USA. They found that acid-
256 leached and unleached cpx in one xenolith had similar δ<sup>44/40</sup>Ca, but unleached opx
257 yielded a lower average δ<sup>44/40</sup>Ca than its leached fractions (1.21‰ vs. 1.40‰). Huang
258 et al. (2010) attributed the lower δ<sup>44/40</sup>Ca in the unleached opx to Ca contribution from
259 hypothetical, presumably post-eruption, low-δ<sup>44/40</sup>Ca carbonates precipitated on the opx
260 surface and used only data on leached pyroxenes to constrain Ca isotope composition
261 of “upper mantle”. Yet, because no mineralogical or trace element data on the xenolith
262 were reported by Huang et al. (2010), it is difficult to ascertain the origin of the Ca-
263 bearing materials in the mineral separates.

264 It is well known, however, that mantle peridotites may contain cryptocrystalline
265 inter-granular materials of mantle origin rich in incompatible elements (e.g. Hiraga et
266 al., 2004). Furthermore, metasomatized mantle rocks, including Group 3 samples in
267 this study, may contain various interstitial micro-phases, from the common amphibole
268 and phl (e.g. Ionov et al., 1997) to more rare phosphates, carbonates, alkali feldspar, Ti-
269 rich oxides precipitated from percolating melts and fluids (e.g. Ionov et al., 1999, 2006).
270 For this reason, we chose to analyze bulk rock xenoliths, with the assumption that their
271 interstitial components are mainly of mantle origin. To minimize contributions from
272 non-mantle materials we used cores of well-studied large and fresh xenoliths with no
273 evidence for intrusion of host basalt and no or minimal post-eruption alteration.

274 The first advantage of this approach is that all Ca-bearing components of bulk rocks
275 are taken into account including those from the low-Ca olivine and potentially Ca-rich,
276 but low-volume interstitial materials. Another one is that our Ca isotope data are direct
277 bulk-rock measurements whereas values based on pyroxene analyses are estimates
278 calculated using modal abundances of Ca-rich minerals, which is an indirect method to
279 infer bulk-rock values. The latter is the case for the “upper mantle” estimates in Huang
280 et al. (2010), which are based on approximate models for modal opx and cpx in the
281 mantle. Finally, estimates based on cpx analyses are impossible for rocks, like
282 harzburgite, that contain no optically recognizable cpx (e.g. two samples from this
283 study), in cases if Ca-bearing minerals are too small for mineral picking, heterogeneous,
284 or when their mineral proportions cannot be established reliably.

285

286 5.2 Ca isotopic composition of the BSE

287 The stable isotope composition of the BSE serves as a benchmark to evaluate the
288 directions and degrees of isotope fractionation in terrestrial rocks and geochemical
289 reservoirs. The BSE, which can be viewed as equivalent to the primitive mantle, is the
290 silicate fraction of the Earth that remained after the separation of metallic core and loss
291 of volatiles early in its history. It later fractionated to form the crust and the
292 complementary melt-depleted asthenosphere and lithospheric mantle. Pioneering stable
isotope studies often use data on basalts to estimate BSE values, but it is not known to what extent basaltic magmas represent the primitive mantle because: 1) partial melting may cause isotope fractionation; 2) sources of oceanic basalts vary widely; 3) magmas may fractionate and/or be contaminated by other materials upon transport and emplacement (e.g. Weyer and Ionov, 2007; Williams and Bizimis, 2014).

A more direct approach in constraining the Ca isotope composition of the BSE is to look at pristine mantle peridotites that suffered no, or only little, fractionation by melting or contamination. This approach was previously used to estimate the abundances of refractory lithophile elements including Ca in the primitive mantle, i.e. the equivalent of the BSE (McDonough and Sun, 1995). In this study, Group 1 peridotites meet these criteria, with fertile major element compositions (Fig. 2) and nearly flat or slightly LREE-depleted patterns (Fig. 3a; Ionov 2004; Ionov and Hofmann, 2007; Press et al., 1986), and thus are appropriate to estimate Ca isotope composition of the BSE. Notably, our Vitim samples were previously used to constrain Li, Mg, V, and Zn isotopic compositions of the BSE and yielded uniform isotope values within a given isotope system (Doucet et al., 2016; Pogge von Strandmann et al., 2011; Prytulak et al., 2013). Here, we estimate the δ<sup>44/40</sup>Ca of BSE to be 0.94±0.05‰ (2SD) as a mean of average (3–4 duplicates) analyses of the 14 Group 1 lherzolites. Importantly, δ<sup>44/40</sup>Ca in the four garnet and garnet-spinel lherzolites match those of spinel lherzolites, show no differences between the spinel and garnet facies mantle (Fig. 4).

Magna et al. (2015) suggested that because the Ca-O bonds in garnet are shorter
314 than in cpx, and because heavy isotopes prefer stronger bonding environments, garnets
315 may have higher $\delta^{44/40}$Ca than the coexisting cpx. This argument, however, if valid,
316 only concerns inter-mineral isotope fractionation, but does not imply that garnet
317 peridotites in general have heavier Ca isotopes than garnet-free peridotites because
318 there is no net mass transfer during the transformation of spinel to garnet peridotites
319 thus precluding Ca isotope fractionation.

320

321 **5.3 Ca isotope fractionation during mantle melting**

322 Calcium is preferentially partitioned into melt and is depleted in residues during
323 melting of peridotite. Cpx, the most important silicate host of Ca in the fertile upper
324 mantle, with ~20% CaO, is exhausted at ≥25% melting at 0–2 GPa (e.g., Pickering-
325 Witter and Johnston, 2000). The effects of melt extraction on the Ca isotopic
326 composition of melting residues are examined here by comparing the Group 1 fertile
327 rocks with the Group 2 melt-depleted, unmetasomatized peridotites.

328 While the two moderately melt-depleted (2.2–3.0 wt.% Al₂O₃, 1.8–2.3 wt.% CaO)
329 Group 2 lherzolites show $\delta^{44/40}$Ca values (0.97±0.09‰ to 0.98±0.07‰) near the upper
330 limit of the Group 1 range, the three strongly melt-depleted (0.75–1.3 wt.% Al₂O₃,
331 0.8–1.7 wt.% CaO) Group 2 harzburgites have considerably higher $\delta^{44/40}$Ca, 1.03±0.06
332 to 1.09±0.06‰. These harzburgites are residues of much greater melt-extraction
333 degrees than the lherzolites, i.e. ~25–30% vs. <20% of isobaric batch melting at 4 GPa
334 (Fig. 2). Furthermore, because the LREE-depleted cpx patterns in the Group 2
335 harzburgites (Fig. 3b) indicate little (if any) metasomatic influence (Ionov et al., 2005, 336 Press et al., 1986), the higher δ<sup>44/40</sup>Ca most likely reflect the effects of mantle partial 337 melting. Given that the fractionation scale of Ca isotopes during partial melting is 338 relatively small, partial melting-induced Ca isotopic shift can only be distinguished in 339 highly melt-depleted peridotites (25–30% melting, Fig. 2b) with the analytical 340 uncertainty in this study (~0.13‰).

341 Earlier studies have suggested that cpx (CaO ~20.0 wt.%) may have lighter Ca 342 isotopic composition than coexisting olivine (CaO <0.1 wt.%), opx (CaO <1.5 wt.%), 343 and garnet (CaO ~3-5 wt.%) (Feng et al., 2014; Huang et al., 2010; Kang et al., 2016; 344 Magna et al., 2015). Because modal abundances of cpx decrease whereas those of 345 olivine and opx increase in residues at higher melting degrees, the residues may be 346 progressively enriched in heavy Ca isotopes with the exhaustion of cpx, which is 347 generally consistent with our data for the Group 2 and Group 1 rocks.

348 To provide quantitative constraints on the fractionation factor between peridotite 349 and melt (αperidotite-melt), modal batch melting modeling similar to that in Weyer and 350 Ionov (2007) and Doucet et al. (2016) was applied. In this model, CaO variations in 351 residues and partial melts were estimated with the method of Niu (1997) initially 352 applied to abyssal peridotites. Details of the modeling are given in Supplementary 353 Material. As shown in Fig. 4 and 5, the modeling suggests that the variations of 354 δ<sup>44/40</sup>Ca in the melt-depleted peridotites can be explained by melting of fertile mantle 355 at 3 GPa with αperidotite-melt ranging from 0.99975 to 0.99990 (10<sup>3</sup>lnαperidotite-melt ≈
0.10–0.25). However, this estimate is approximate because the fractionation factor may vary with cpx consumption and also depend on P-T conditions, like for Zn isotopes (Doucet et al., 2016).

## 5.4 Ca isotopic composition of metasomatized peridotites

The notion of mantle metasomatism includes a broad range of phenomena taking place in melt-depleted residues due to reaction with migrating melts or fluids. Initially, such residues are depleted in incompatible elements and may contain no or only small amounts of Ca-rich minerals (cpx, garnet), but interaction with metasomatic media may cause important changes in their chemical (cryptic), and possibly modal compositions (patent or modal) (e.g. Bodinier et al., 1990). Cryptic metasomatism produces enrichments in incompatible elements, notably in LREE over MREE-HREE (Fig. 3c-d) without apparent changes in phase assemblage. Modal metasomatism, in addition, precipitates new phases, like water-bearing amphibole and phl (e.g. Ionov et al., 1997) as well as Ca-rich cpx, carbonates, phosphates, and alkali feldspar (e.g. Ionov et al., 1999, 2006).

In general, the peridotites affected by mantle metasomatism in this study tend to have lower $\delta^{44/40}$Ca values than fertile (0.90–0.97‰) or melt-depleted (0.97–1.09‰) rocks, which we attribute to reaction of the latter with lighter, Ca-rich fluids/melts. In detail, out of nine metasomatized samples in this study, the three Group 3b fertile lherzolites show a tight $\delta^{44/40}$Ca range from 0.83‰ to 0.89‰, which is only slightly
377 lighter than the BSE. By contrast, the six Group 3a harzburgites display a much broader
378 $\delta^{44/40}$Ca range from 0.25‰ to 0.96‰. The $\delta^{44/40}$Ca in the metasomatized rocks may be
379 controlled by (a) the nature of metasomatic media, (b) degree and type of metasomatism,
380 (c) the composition of the rocks that undergo metasomatism, and (d) diffusion-related
381 isotope fractionation during melt percolation. These factors are examined below for
382 subgroups 3a and 3b.

383

384 **5.4.1. Very light Ca isotope compositions in Group 3a**

385 Highly refractory, olivine-rich rocks, like the Group 3a harzburgites in this study,
386 are most likely to be intruded and metasomatized by migrating melts (e.g. Frey and
387 Green, 1974) because they are most permeable for porous melt flow (e.g. Toramaru and
388 Fujii, 1986), in particular for carbonate-rich melts (Hunter and McKenzie, 1989).
389 Furthermore, because such refractory residues are low in Ca, metasomatic additions of
390 Ca or equilibration with percolating melts may modify their Ca isotopic compositions
391 much more than those of Ca-rich fertile peridotites. As a result, the Ca isotope
392 compositions of such strongly metasomatized, refractory peridotites may mainly reflect
393 the nature of the metasomatic media.

394 The six Group 3a harzburgites (four from Udachnaya and two from Tariat) show a
395 positive correlation of $\delta^{44/40}$Ca with CaO (Fig. 4) and negative correlations with
396 LREE/MREE ratios (e.g. Ce/Eu, Fig. 6), which we view as evidence for links of their
397 Ca isotopic composition with metasomatism. Notably, the most refractory Udachnaya
peridotite in this study, cpx-free KC-137/08, has the lightest $\delta^{44/40}$Ca of 0.25‰ and a very high Ce/Eu of 81, the latter identical to Ce/Eu for an average of fresh Udachnaya kimberlites (Kamenetsky et al., 2012).

Primary kimberlite magmas are CO₂-rich and appear to be generated from rocks enriched in carbonates and incompatible elements by sub-lithospheric fluids (e.g. le Roex et al., 2003). Kamenetsky et al. (2012) argued that the parental melt of the Udachnaya kimberlite (and possibly of Type I kimberlites worldwide) was a halogen-rich Na-Ca carbonatite, which progressively assimilated silicate components to form kimberlite magmas. All these melts intruded host mantle and metasomatized it. The Udachnaya samples in this study contain no kimberlite veins and related phl-cpx pockets, hence their budgets of incompatible components are related to metasomatism in the mantle and not to ingress of kimberlite to xenoliths during their transport (Agashev et al., 2013; Doucet et al., 2013). Carbonate-rich media were also inferred as likely metasomatic agents for Tariat harzburgites, which often contain microcrystalline apatite (Ionov, 2007; Ionov et al., 2006).

It was recently suggested, based on studies of mantle-derived rocks from eastern China affected by Meso-Cenozoic subduction, that subduction-related melts or fluids could impart low-$\delta^{44/40}$Ca signatures to peridotites because some recycled marine carbonates have lower $\delta^{44/40}$Ca (-1.09 to 1.81‰) with an average of 0.61‰, Fantle and Tipper, 2014) than the upper mantle (Kang et al., 2016). There is no evidence however that the mantle beneath Udachnaya (≥1000 km from the margins of the Siberian craton),
419 or beneath central Mongolia, could be directly affected by subduction-derived fluids
420 either at the time of the eruption or earlier in its history. Alternatively, recycled
421 carbonates with light Ca-isotope signatures could be first subducted to the deep mantle
422 and to be later brought up by mantle upwelling to source regions of intra-plate magmas
423 (Ionov et al., 2006; Moyen et al., 2016).

424 As shown in Fig. 6, the correlation between δ<sup>44/40</sup>Ca and Ce/Eu in the Group 3a
425 samples cannot be explained by binary mixing of refractory peridotites and a single
426 carbonate melt end-member (find the modeling parameters in S-Table 2 of
427 supplementary material). This feature, however, is consistent with the complex
428 evolution of Udachnaya kimberlite magmas, which is related to reactions of the parental
429 alkali-CO<sub>2</sub>-halogen-rich fluids with the host mantle to gradually increase its silicate
430 contents. Hence, both carbonatic melts and related metasomatic media could have a
431 large range of δ<sup>44/40</sup>Ca and Ce/Eu. Likewise, Ionov et al. (2010) and Agashev et al.
432 (2013) argued that Udachnaya peridotites experienced multiple stages of enrichments,
433 first by carbonatite then by OIB-like silicate liquids. It is thus possible that the
434 relationship between Ce/Eu and δ<sup>44/40</sup>Ca in the Group 3a rocks reflects a net result of
435 multiple metasomatic events by evolving CO<sub>2</sub>-rich media, shown as the three-end-
436 member mixing model in Fig. 6.

437 An additional mechanism to yield light Ca isotopic values in metasomatized rocks
438 could be kinetic isotope fractionation previously proposed for stable isotopes of other
439 elements in mantle peridotites, e.g. Fe (Weyer and Ionov, 2007) and Li (Rudnick and
Ionov, 2007). Because lighter isotopes diffuse faster they may be locally enriched at a
metasomatic front or in melt-rock reactions. Such ‘chromatographic’ effects were
earlier invoked to explain the decoupling of trace elements (e.g. Bodinier et al., 1990,
Ionov et al., 2002a) and Sr and Nd isotopes (Ionov et al., 2002b) in enriched peridotites.

444

445 5.4.2. Moderately-light Ca isotopic composition of Group 3b samples

The metasomatic effects on Ca isotopes in the three patently metasomatized Group
3b lherzolites (two phl-bearing spinel lherzolites from Tariat and a sheared garnet
lherzolite from Udachnaya) are uniformly minor, with only slightly lighter δ<sup>44/40</sup>Ca
(0.83–0.89‰) than the BSE value, in spite of evidence for equilibration with
metasomatic media like LREE-enriched, concave down REE patterns (Fig. 3d) and
texturally equilibrated phl in the Tariat xenoliths (Ionov et al., 2002a; Ionov and
Hofmann, 1995). This difference may be due either to low mass ratios of Ca in the
metasomatic media to Ca in the host peridotites with high CaO (3.1–3.9 wt.% in Group
3b rocks) or alternatively, much smaller differences between Ca isotopic composition
of the pre-metasomatic rocks and the metasomatic media.

456

5.5 Relationships between δ<sup>44/40</sup>Ca of the mantle and oceanic basalts

A compilation of published Ca isotope data on basalts (Fig. 1) shows a large range
from 0.62 to 1.54‰ with an average of 0.90±0.28‰ (n=61, 2SD) (data from Amini et
al., 2009; Huang et al., 2010, 2011; Holmden et al., 2010; Hindshaw et al., 2013;
461 Jacobsen et al., 2015; Skulan et al., 1997; Simon and DePaolo, 2010; Valdes et al., 2014).
462 A recent study by Zhu et al. (2016) shows that δ<sup>44/40</sup>Ca in normal and depleted MORB
463 from the Pacific Ocean range from 0.75 to 0.88‰ with an average of 0.83 ± 0.07‰
464 (2SE, n=13). Such variations can be explained by equilibrium Ca-isotope fractionation
465 during mantle melting (Fig. 5). Much more fractionated δ<sup>44/40</sup>Ca in oceanic basalts
466 (0.62–1.54‰, 60 samples, Figs. 1 and 5) can be due to several factors. Interaction with
467 seawater, which has a high δ<sup>44/40</sup>Ca value of ~1.82‰ (e.g. Fantle and Tipper, 2014 and
468 references therein), may elevate their δ<sup>44/40</sup>Ca values (e.g., John et al., 2012). Further,
469 crustal materials like sediments and carbonates may have much lighter Ca isotopic
470 composition than the BSE (Fantle and Tipper, 2014 and references therein).
471 Incorporation of such isotopically lighter materials in the source regions of basaltic
472 magmas may lower their δ<sup>44/40</sup>Ca values, in which case Ca isotopes in basalts could
473 serve as recycling tracers (e.g. Huang et al., 2011). On the other hand, the effects of
474 partial melting should also be taken into account. Our modeling suggests that melts
475 derived from pristine mantle may have δ<sup>44/40</sup>Ca from 0.79 to 0.94‰ depending on
476 melting degrees and fractionation factors (Fig. 5). Thus, we suggest that δ<sup>44/40</sup>Ca<0.79‰
477 in basalts may be seen as indication of the contribution of recycled isotopically lighter
478 materials in their mantle sources.
479

480 5.6 Interplanetary comparisons

481 Our new estimate of the δ<sup>44/40</sup>Ca in the BSE may help evaluate likely building
blocks of the Earth. As shown in Fig. 7, achondrites and ordinary chondrites have $\delta^{44/40}$Ca ranging from 0.88 to 1.13‰ (e.g., Huang et al., 2016; Simon and DePaolo, 2010; Valdes et al., 2014), i.e. close to the BSE estimate (0.94±0.05‰) obtained in this study. The Ca isotopic composition of the majority of Martian meteorites is similar to our BSE estimate as well (Magna et al., 2015; Simon and DePaolo, 2010). By comparison, carbonaceous chondrites display a broad $\delta^{44/40}$Ca range of 0.27–1.19‰ (Huang et al., 2016; Simon and DePaolo, 2010; Valdes et al., 2014; Amsellem et al. 2017), with CM, CV3, CR2 and CI isotopically lighter than the BSE and most CO similar to the BSE. Such variations may be caused by disequilibrium condensation of the solar nebula or be related to the Calcium-Aluminium-rich Inclusions, CAI's (Simon and DePaolo, 2010). In contrast, some enstatite chondrites have heavier Ca isotopic compositions (0.90–1.51‰) than the BSE, casting doubts on their role as possible Earth's building blocks. Valdes et al. (2014) attributed this signature to variations in modal proportions of oldhamite, which is enriched in heavy Ca isotopes relative to silicate phases of enstatite chondrites. More work is needed to understand the implications of Ca isotope compositions of the rocky planets and meteorites.

The Earth-Moon system has been proposed to form by collision of the proto-Earth with a Mars-sized impactor named Theia (e.g. Cameron, 2001). The Ca isotopic compositions of several lunar samples reported by Simon and DePaolo (2010) and Valdes et al. (2014) are consistent with the BSE estimate in this study (Fig. 6). Such a consistent Ca isotopic signature is reminiscent of the identical O and Ti isotopic
503 compositions of terrestrial and lunar samples (e.g., Wiechert et al., 2001), which
504 supports the hypothesis that the majority of the Moon-forming materials came from the
505 proto-Earth rather than Theia (e.g. Zhang et al., 2012).

506

507 **6. Summary and Conclusions**
508

508

509 We report whole-rock Ca isotope data for three groups of well-studied peridotite
510 xenoliths from central Asia and the Siberian craton to constrain Ca isotopic composition
511 of the BSE and explore Ca isotope fractionation in the upper mantle.

512 (1) Fertile spinel and garnet lherzolites (Group 1) show identical δ<sup>44/40</sup>Ca, their
513 average defines a δ<sup>44/40</sup>Ca for BSE of 0.94±0.05‰ (2SD, n=14).

514 (2) The average δ<sup>44/40</sup>Ca (1.07±0.04‰, n=3) of harzburgites without significant
515 metasomatism (Group 2) is slightly higher than the BSE value, suggesting that partial
516 melting may induce observable Ca isotopic fractionation in refractory peridotites and
517 in complementary melts at melting degrees ≥20%. A modal batch-melting model with
518 α<sub>peridotite-melt</sub> ranging from 0.99975 to 0.99990 can explain our data.

519 (3) Ca isotope variations in metasomatized peridotites are largely controlled by the
520 nature of metasomatic media, degree and type of metasomatism, and the Ca contents of
521 the initial rocks. Metasomatized, low-Ca harzburgites (Group 3a) have a broad range
522 of δ<sup>44/40</sup>Ca from 0.25±0.01‰ to 0.96‰±0.08‰ that are negatively correlated with
523 LREE/MREE ratios. Their light Ca isotope compositions may be imparted by a range
524 of metasomatic media from carbonatites to proto-kimberlitic liquids, possibly
525 containing recycled crustal components. δ<sup>44/40</sup>Ca in metasomatized fertile peridotites
526 (Group 3b) are only slightly lighter (0.83–0.89‰) than in the BSE, possibly due to low
527 mass ratios of Ca in the metasomatic media to Ca in the host peridotites or smaller Ca
528 isotope differences between these two components.

529

530 Acknowledgements

531 This work was supported by the Strategic Priority Research Program (B) of the
532 Chinese Academy of Sciences (XDB18000000), the National Science Foundation of
533 China (41325011, 41573017) and the 111 project. DAI acknowledges funding from the
534 French CNRS (INSU-PNP annual grants in 2010-17 for mantle studies), AVG was
535 supported by Russian State Assignment Project ± 0330-2016-0006. We thank Hongli
536 Zhu with discussion and the help of analysis, Maria C. Valdes for data on some basalts,
537 Gui-Qin Wang for help with isotope analyses and Chen Zhu and Anne Hereford for help
538 with polishing English.

539

540 Appendix A. Supplementary material

541 Supplementary material related to this article can be found on-line at
542 http://dx.doi.org/10.1016/j.epsl.2017.xx.xxx.

543

544
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720 **Figure Captions**

721 **Fig. 1.** A compilation of available Ca isotopic data for terrestrial basalts and peridotites.

722 The red dashed line is the average of oceanic basalt of 0.90‰. The blue vertical

723 band represents the δ<sup>44/40</sup>Ca of BSE estimated in this study. Data for basalts are

724 from Skulan et al. (1997), Amini et al. (2009), Huang et al. (2010, 2011),

725 Holmden et al. (2010), Simon and DePaolo (2010), Hindshaw et al. (2013),

726 Valdes et al. (2014) and Jacobsen et al. (2015). Data for peridotites are from

727 Amini et al. (2009) and Kang et al. (2016). The error bars are 2SE. The grey (left)

728 and yellow (right) lines represent the average δ<sup>44/40</sup>Ca of marine carbonates and

729 seawater. Note that marine carbonates have a broad range of δ<sup>44/40</sup>Ca from -1.09

730 to 1.81‰ (Fantle and Tipper, 2014 and references therein). For color codes refer

731 to the online version of the paper.

732 **Fig. 2.** Major element relations of whole-rock peridotites. Black dashed lines: isobaric

733 batch melting residues at 2, 4 and 6 GPa (Herzberg, 2004); continuous red lines:

734 residues of polybaric fractional melting at 2–0, 3–0, 5–1, and 7–2 GPa

735 (Herzberg, 2004); red dotted lines: 20%, 30% and 40% melting degrees. See text

736 for references for the BSE composition. The major element data of peridotites

737 are from the literature (listed in Table 2) and unpublished data of DA Ionov and

738 RW Carlson (for Tariat xenoliths). For color codes refer to the online version of

739 the paper.

740 **Fig. 3.** Primitive mantle-normalized (McDonough and Sun, 1995) REE patterns of
741 typical samples in Groups 1, 2, 3a and 3b: open symbols, clinopyroxene (cpx),
742 filled symbols, whole-rocks.

743 **Fig. 4.** Plots of δ<sup>44/40</sup>Ca vs. CaO. The grey field represents the BSE value estimated in
744 this study (0.94±0.05‰, 2SD). Ca isotopic variation in melting residues is
745 calculated for pure batch melting assuming fractionation factor α<sub>peridotite-melt</sub>
746 ranging from 0.99975 to 0.99990. For color codes refer to the online version of
747 the paper.

748 **Fig. 5.** Comparison of mantle peridotites with oceanic basalts (OIB and MORB) on a
749 plot of δ<sup>44/40</sup>Ca vs. CaO. Grey field represents the BSE value from this study.
750 Modeling results for residues and partial melts of pure batch melting of the BSE
751 are shown assuming α<sub>peridotite-melt</sub> ranging from 0.99975 to 0.99990. Literature
752 data of peridotites are from Amini et al (2009) and Kang et al (2016). OIB data
753 are from Amini et al (2009), Huang et al (2011), Valdes et al (2014) and Jacobsen
754 et al (2015). The range of δ<sup>44/40</sup>Ca for the Pacific MORB is from Zhu et al (2016).
755 Some oceanic basalt data plotted in Fig. 4 are not included in this figure because
756 their CaO contents are not published.

757 **Fig. 6.** A plot of δ<sup>44/40</sup>Ca vs. Ce/Eu for peridotites in this study. Simple binary mixing
758 between carbonate melts and refractory mantle (blue curve) cannot explain the
759 variations of Ca isotopes and Ce/Eu in Udachnaya harzburgites. A mixing model
760 for three end-members, carbonate melt, OIB-like silicate melt and refractory
761 mantle, is applied to examine the δ<sup>44/40</sup>Ca and Ce/Eu variations. The percentages
besides the curve denote the contribution of metasomatism agents to the REE

The parameters and references for the model are given in supplementary

material. For color codes refer to the online version of the paper.

**Fig. 7.** A compilation of Ca isotopic data for primitive meteorites, lunar samples, and Martian meteorites. The gray vertical band represents the $\delta^{44/40}\text{Ca}$ of the BSE, i.e. $0.94\pm0.05\‰$ (2SD, n=14) estimated in this study. Data sources: primitive meteorite data, Simon and DePaolo (2010) and Valdes et al. (2014), lunar samples, Valdes et al. (2014), and Martian meteorites, Simon and DePaolo (2010) and Magna et al. (2016). Achondrite data include those for aubrite, angrite, diogenite, and eucrite. Carbonaceous chondrite data include those for the CM, CV, CO, CR and CI types. Ordinary chondrite data include those for LL and L. Enstatite chondrite data include those for the EH and EL types.
<figure><img src="image_3.png" /></figure>
<figure><img src="image_4.png" /></figure>

<figure><img src="image_5.png" /></figure>
<figure><img src="image_6.png" /><figcaption>779 Fig. 3</figcaption></figure>

780

781
<figure><img src="image_7.png" /><figcaption>782 Fig. 4</figcaption></figure>

783
<figure><img src="image_8.png" /></figure>
786 Fig. 6

787

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

788

789
BSE

Achondrite

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

Carbonaceous Chondrite

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

Ordinary Chondrite

<figure><img src="image_12.png" /></figure>

Enstatite Chondrite

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

<figure><img src="image_14.png" /></figure>

CR

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

CV

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

CI

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

CO

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

CM

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

EL

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

EH

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

<figure><img src="image_22.png" /></figure>

Moon

<figure><img src="image_23.png" /></figure>

Mars

$\delta^{44/40}$Ca(‰)
<table><caption>792 Table 1. Ca isotopic compositions of SRM915a, seawater, and USGS standards.</caption><thead><tr><th>Samples</th><th>References</th><th>δ<sup>44/40</sup>Ca</th><th>2SD</th><th>n</th></tr></thead><tbody><tr><td rowspan="3">SRM 915a</td><td><strong>This study</strong></td><td><strong>0.02</strong></td><td><strong>0.13</strong></td><td><strong>48</strong></td></tr><tr><td>Valdes et al. (2014)</td><td>0.00</td><td>0.17</td><td>74</td></tr><tr><td>Huang et al. (2010)</td><td>0.04</td><td>0.13</td><td>45</td></tr><tr><td rowspan="2">Seawater</td><td><strong>This study</strong></td><td><strong>1.83</strong></td><td><strong>0.12</strong></td><td><strong>47</strong></td></tr><tr><td>Huang et al. (2010)</td><td>1.90</td><td>0.12</td><td>16</td></tr><tr><td rowspan="3">BCR-2</td><td><strong>This study</strong></td><td><strong>0.82</strong></td><td><strong>0.09</strong></td><td><strong>12</strong></td></tr><tr><td>Amini et al. (2009)</td><td>0.81</td><td>0.17</td><td>6</td></tr><tr><td>Valdes et al. (2014)</td><td>0.88</td><td>0.32</td><td>3</td></tr><tr><td rowspan="3">BHVO-2</td><td><strong>This study</strong></td><td><strong>0.80</strong></td><td><strong>0.10</strong></td><td><strong>16</strong></td></tr><tr><td>Magna et al. (2015)</td><td>0.90</td><td>0.11</td><td>5</td></tr><tr><td>Valdes et al. (2014)</td><td>0.88</td><td>0.04</td><td>13</td></tr><tr><td rowspan="3">BIR-1</td><td><strong>This study</strong></td><td><strong>0.84</strong></td><td><strong>0.09</strong></td><td><strong>9</strong></td></tr><tr><td>Amini et al. (2009)</td><td>0.77</td><td>0.34</td><td>14</td></tr><tr><td>Valdes et al. (2014)</td><td>0.90</td><td>0.08</td><td>4</td></tr></tbody></table>

793
<table><caption>794 Table 2. Ca isotopic compositions of peridotites.</caption><thead><tr><th rowspan="2">Sample</th><th rowspan="2">Locality</th><th rowspan="2">Rock</th><th rowspan="2">CaO<br>wt.%</th><th rowspan="2">Al<sub>2</sub>O<sub>3</sub><br>wt.%</th><th colspan="5">δ<sup>44/40</sup>Ca</th><th rowspan="2">Mean</th><th rowspan="2">2SD</th><th rowspan="2">2SE</th><th rowspan="2">n</th><th rowspan="2">Data<br>source<sup>c</sup></th></tr><tr><th>1<sup>a</sup></th><th>2</th><th>3</th><th>4</th><th>R<sup>b</sup></th></tr></thead><tbody><tr><td colspan="16" style="text-align: center; font-weight: bold;">Group 1 fertile, unmetasomatized peridotites</td></tr><tr><td>Ts-53389</td><td>Tar</td><td>Gar-Sp Lh</td><td>3.82</td><td>4.70</td><td>0.99</td><td>0.89</td><td>0.95</td><td></td><td></td><td></td><td>0.94</td><td>0.10</td><td>0.06</td><td>3</td><td>(3)</td></tr><tr><td>S-17</td><td>Tar</td><td>Sp Lh</td><td>3.89</td><td>4.65</td><td>0.94</td><td>0.96</td><td>0.92</td><td></td><td></td><td></td><td>0.94</td><td>0.04</td><td>0.02</td><td>3</td><td>(1)</td></tr><tr><td>313-8</td><td>Vit</td><td>Gar Lh</td><td>3.53</td><td>4.58</td><td>1.05</td><td>0.92</td><td>1.00</td><td></td><td></td><td></td><td>0.99</td><td>0.13</td><td>0.08</td><td>3</td><td>(4)</td></tr><tr><td>S-4</td><td>Tar</td><td>Sp Lh</td><td>3.92</td><td>4.48</td><td>0.93</td><td>0.98</td><td>0.93</td><td></td><td></td><td></td><td>0.95</td><td>0.06</td><td>0.03</td><td>3</td><td>(1)</td></tr><tr><td>313-102</td><td>Vit</td><td>Gar Lh</td><td>3.56</td><td>4.48</td><td>0.92</td><td>0.85</td><td>1.02</td><td>0.90</td><td></td><td></td><td>0.92</td><td>0.14</td><td>0.08</td><td>4</td><td>(4)</td></tr><tr><td>stz-1</td><td>Tar</td><td>Sp Lh</td><td>4.01</td><td>4.40</td><td>0.95</td><td>0.93</td><td>0.92</td><td></td><td></td><td></td><td>0.93</td><td>0.03</td><td>0.02</td><td>3</td><td>(2)</td></tr><tr><td>S-37</td><td>Tar</td><td>Sp Lh</td><td>3.25</td><td>4.39</td><td>0.91</td><td>0.94</td><td>0.94</td><td></td><td></td><td></td><td>0.93</td><td>0.03</td><td>0.02</td><td>3</td><td>(1)</td></tr><tr><td>S-1</td><td>Tar</td><td>Sp Lh</td><td>3.65</td><td>4.27</td><td>1.00</td><td>0.87</td><td>0.89</td><td></td><td></td><td></td><td>0.92</td><td>0.14</td><td>0.08</td><td>3</td><td>(1)</td></tr><tr><td>S-14</td><td>Tar</td><td>Sp Lh</td><td>3.40</td><td>4.20</td><td>0.99</td><td>0.98</td><td></td><td>0.90</td><td></td><td></td><td>0.96</td><td>0.10</td><td>0.06</td><td>3</td><td>(1)</td></tr><tr><td>S-21</td><td>Tar</td><td>Sp Lh</td><td>3.24</td><td>4.09</td><td>0.93</td><td>0.87</td><td>0.97</td><td></td><td></td><td></td><td>0.92</td><td>0.10</td><td>0.06</td><td>3</td><td>(1)</td></tr><tr><td>stz-2</td><td>Tar</td><td>Sp Lh</td><td>3.32</td><td>4.07</td><td>0.96</td><td>0.90</td><td>0.84</td><td></td><td></td><td></td><td>0.90</td><td>0.12</td><td>0.07</td><td>3</td><td>(1)</td></tr><tr><td>S-15</td><td>Tar</td><td>Sp Lh</td><td>3.40</td><td>4.00</td><td>0.88</td><td>0.98</td><td>0.83</td><td></td><td></td><td></td><td>0.90</td><td>0.15</td><td>0.09</td><td>3</td><td>(1)</td></tr><tr><td>S-2</td><td>Tar</td><td>Sp Lh</td><td>3.26</td><td>3.86</td><td>0.91</td><td>0.99</td><td>0.94</td><td></td><td></td><td></td><td>0.95</td><td>0.08</td><td>0.05</td><td>3</td><td>(1)</td></tr><tr><td>313-104</td><td>Vit</td><td>Gar Lh</td><td>3.49</td><td>3.66</td><td>0.98</td><td>0.90</td><td>1.03</td><td></td><td></td><td></td><td>0.97</td><td>0.13</td><td>0.08</td><td>3</td><td>(4)</td></tr><tr><td colspan="16" style="text-align: center; font-weight: bold;">Group 2 moderately refractory peridotites with LREE-depleted cpx</td></tr><tr><td>S-16</td><td>Tar</td><td>Spl Lh</td><td>2.34</td><td>3.04</td><td>0.89</td><td>0.94</td><td>1.07</td><td>0.99</td><td></td><td></td><td>0.97</td><td>0.15</td><td>0.09</td><td>4</td><td>(1)</td></tr><tr><td>S-22</td><td>Tar</td><td>Spl Lh</td><td>1.77</td><td>2.24</td><td>0.93</td><td>0.95</td><td>1.05</td><td></td><td></td><td></td><td>0.98</td><td>0.13</td><td>0.07</td><td>3</td><td>(1)</td></tr><tr><td>S-62</td><td>Tar</td><td>Sp Hz</td><td>1.32</td><td>1.74</td><td>1.05</td><td>1.07</td><td></td><td>1.15</td><td></td><td></td><td>1.09</td><td>0.11</td><td>0.06</td><td>3</td><td>(1)</td></tr><tr><td>621-16</td><td>Vit</td><td>Gar Hz</td><td>1.07</td><td>1.31</td><td>1.18</td><td>1.02</td><td>1.05</td><td>1.03</td><td></td><td></td><td>1.07</td><td>0.15</td><td>0.09</td><td>4</td><td>(1)</td></tr><tr><td>BN-8</td><td>Dar</td><td>Spl Hz</td><td>0.75</td><td>1.26</td><td>1.08</td><td>1.03</td><td>0.98</td><td></td><td></td><td></td><td>1.03</td><td>0.10</td><td>0.06</td><td>3</td><td>(1)</td></tr><tr><td colspan="16" style="text-align: center; font-weight: bold;">Group 3a refractory peridotites with carbonatite-type metasomatism</td></tr><tr><td>419/09</td><td>Ud</td><td>Gar Hz</td><td>1.37</td><td>1.00</td><td>0.77</td><td>0.77</td><td>0.82</td><td></td><td></td><td></td><td>0.79</td><td>0.06</td><td>0.03</td><td>3</td><td>(6)</td></tr><tr><td>S-29</td><td>Tar</td><td>Spl Hz</td><td>0.44</td><td>0.90</td><td>0.73</td><td>0.69</td><td>0.75</td><td></td><td></td><td></td><td>0.72</td><td>0.06</td><td>0.04</td><td>3</td><td>(1)</td></tr><tr><td>105/03</td><td>Ud</td><td>Spl Hz</td><td>0.33</td><td>0.67</td><td>0.54</td><td>0.70</td><td>0.53</td><td>0.61</td><td>0.66</td><td></td><td>0.61</td><td>0.15</td><td>0.09</td><td>5</td><td>(5)</td></tr><tr><td>H-25</td><td>Tar</td><td>Spl Hz</td><td>0.41</td><td>0.62</td><td>0.54</td><td>0.56</td><td>0.56</td><td></td><td></td><td></td><td>0.55</td><td>0.02</td><td>0.01</td><td>3</td><td>(1)</td></tr><tr><td>107/03</td><td>Ud</td><td>Spl Hz</td><td>0.54</td><td>0.55</td><td>0.88</td><td>1.01</td><td>0.99</td><td></td><td></td><td></td><td>0.96</td><td>0.14</td><td>0.08</td><td>3</td><td>(5)</td></tr><tr><td>KC-137/08</td><td>Ud</td><td>Spl Hz</td><td>0.26</td><td>0.09</td><td>0.25</td><td>0.25</td><td>0.26</td><td></td><td></td><td></td><td>0.25</td><td>0.01</td><td>0.01</td><td>3</td><td>(5)</td></tr><tr><td colspan="16" style="text-align: center; font-weight: bold;">Group 3b fertile, silicate-melt metasomatized peridotites</td></tr><tr><td>4230-16</td><td>Tar</td><td>Phl lh</td><td>3.85</td><td>4.32</td><td>0.94</td><td>0.91</td><td>0.83</td><td></td><td></td><td></td><td>0.89</td><td>0.11</td><td>0.07</td><td>3</td><td>(8)</td></tr><tr><td>U85</td><td>Ud</td><td>Gar Lh</td><td>3.35</td><td>3.96</td><td>0.81</td><td>0.82</td><td></td><td>0.87</td><td></td><td></td><td>0.83</td><td>0.06</td><td>0.04</td><td>3</td><td>(7)</td></tr><tr><td>4230-19</td><td>Tar</td><td>Phl lh</td><td>3.11</td><td>3.92</td><td>0.92</td><td>0.90</td><td>0.82</td><td></td><td></td><td></td><td>0.88</td><td>0.11</td><td>0.06</td><td>3</td><td>(8)</td></tr></tbody></table>

795

796 a. The numbers (1 to 4) represent repeated analyses of the same solution.

797 b. R, full-procedure duplicate (repeated dissolution and analysis)
798 c. References to data sources: (1) Unpublished data of R. Carlson and D. Ionov; (2)

799 Press et al. (1986); (3) Ionov et al. (1998); (4) Ionov et al. (2004); (5) Doucet et al.

800 (2012); (6) Doucet et al. (2013); (7) Ionov et al. (2010); (8) Ionov and Hofmann

801 (2007).

802 The order of samples in each group is that of decreasing Al₂O₃ contents (roughly

803 corresponds to increasing degrees of melt extraction by partial melting).

804 Location abbreviations: Tar, Tariat; Vi, Vitim; Dar, Dariganga; Ud, Udachnaya.

805 Rock type abbreviations: Lh, lherzolite; Hz, harzburgite; Gar, garnet; Sp, spinel; Phl,

806 phlogopite.

807