The calcium isotope systematics of Mars

Earth and Planetary Science Letters 430 (2015) 86–94

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Earth and Planetary Science Letters www.elsevier.com/locate/epsl

The calcium isotope systematics of Mars Tomáš Magna a,b,∗ , Nikolaus Gussone a , Klaus Mezger a,c a b c

Institut für Mineralogie, Universität Münster, Corrensstr. 24, D-48149 Münster, Germany Czech Geological Survey, Klárov 3, CZ-118 21 Prague, Czech Republic Institut für Geologie, Universität Bern, Baltzerstr. 1+3, CH-3012 Bern, Switzerland

a r t i c l e

i n f o

Article history: Received 15 April 2015 Received in revised form 13 August 2015 Accepted 14 August 2015 Available online xxxx Editor: B. Marty Keywords: calcium isotopes Mars mantle crust carbonate low-temperature fractionation

a b s t r a c t New Ca isotope data from a suite of Martian meteorites provide constraints on the Ca isotope composition of the Martian mantle and possible recycling of surface materials back into the mantle. A mean δ 44/40 Ca of 1.04 ± 0.09h (2SD) is estimated for the Martian mantle which can also be taken as an approximation for Bulk Silicate Mars. This value is identical with the estimates for Bulk Silicate Earth, and the inner Solar System planets can therefore be considered homogeneous with respect to Ca isotopes. The Ca isotope composition of two Martian dunites varies by ∼0.3h despite strong chemical and mineralogical similarities and this difference can be caused by the presence of carbonate, probably of pre-terrestrial origin, implying a protracted period of the existence of CaCO3 -rich fluids and sufficient amounts of CO2 on the surface of Mars. The variability of δ 44/40 Ca within the groups of shergottites and nakhlites (clinopyroxene cumulates) cannot be related to partial melting and fractional crystallization in any simple way. However, there is no necessity of incorporating surface lithologies with isotopically light Ca into the mantle sources of Martian meteorites. These inferences are consistent with the absence of large scale crust–mantle recycling and thus of plate tectonics on Mars. © 2015 Elsevier B.V. All rights reserved.

1. Introduction Mars and Earth are rock-dominated planets, having broadly similar bulk compositions, with Mars being enriched in volatile elements (e.g., S, Se, K, Pb) compared to Earth (e.g., Dreibus and Wänke, 1987; McLennan, 2003). As a consequence of its smaller size (∼11% of the terrestrial mass), Mars can lose most of its internally generated heat by conduction and appears not to show evidence for plate tectonic activity. Its major geologic activity ceased early in its history (Grott et al., 2013; Mezger et al., 2013) while Earth’s has been continuing vigorously. The study of magmatic rocks from Mars and their chemical and isotope signatures can therefore provide important insights into early mantle differentiation processes of planetary bodies (e.g., Kleine et al., 2009). Stable isotope variations of major elements (e.g., Mg, Si, Fe, Ca) in magmatic systems have long been below analytical resolution of mass spectrometry techniques to become useful geochemical tools and only more recently have high-precision studies found that the isotopes of major elements such as Mg, Si, Fe and Ca fractionate during high-temperature magmatic processes

*

Corresponding author at: Czech Geological Survey, Klárov 3, CZ-118 21 Prague, Czech Republic. Tel.: +420 2 5108 5331; fax: +420 2 5181 8748. E-mail address: [email protected] (T. Magna). http://dx.doi.org/10.1016/j.epsl.2015.08.016 0012-821X/© 2015 Elsevier B.V. All rights reserved.

(e.g., Amini et al., 2009; Georg et al., 2007; Teng et al., 2010; Williams et al., 2004). These elements are among the most abundant in the mantles and/or cores of rocky planets, planetesimals and moons, and therefore play key roles in deciphering planetary rock cycles, core–mantle segregation, phase transitions and melting processes. The early studies of mass-dependent Ca isotope fractionation focused on biological cycling (Skulan and DePaolo, 1999; Zhu and Macdougall, 1998) as well as high temperature condensation– evaporation processes in the early Solar System (e.g., Niederer and Papanastassiou, 1984; Russell et al., 1978), while more recent Ca isotope research focused on low-temperature geochemical and biogeochemical cycling, such as the oceanic Ca cycle (Farkaš et al., 2007), bio-mineralization (Böhm et al., 2006; Gussone et al., 2007), carbonate diagenesis (Fantle and DePaolo, 2007; Teichert et al., 2009) or forest ecosystems (Holmden and Bélanger, 2010). These observations demonstrate the potential utility of Ca isotopes as tracer for geological processes. Calcium has a high abundance, lithophile character and plays an essential role in a variety of geological and biological processes, including phenomena such as crystallization of major rock-forming minerals, metamorphism, hydrothermal processes, weathering, diagenesis, bio-mineralization, plant growth and the global carbon cycle. Calcium isotope fractionation during inorganic mineral precipitation is expected to be small (<2h), although systematic δ 44/40 Ca variations in different

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igneous rocks were reported (Amini et al., 2009) and are further supported by the observations of systematic Ca isotope fractionation between orthopyroxene and clinopyroxene in mantle xenoliths, the extent of which is suggested to be compositionally dependent (Feng et al., 2014). These observations indicate that resolvable Ca isotope fractionation may indeed occur during hightemperature igneous processes such as partial melting and fractional crystallization (e.g., Huang et al., 2010b, 2011). Recent data of Valdes et al. (2014) do not imply systematic differences in δ 44/40 Ca among melts derived from different chemical reservoirs in the Earth’s mantle (i.e., HIMU, DMM, EM). Notably, most basalts analyzed by Valdes et al. (2014) and Jacobson et al. (2015) have δ 44/40 Ca values below that inferred for the Earth’s upper mantle which is estimated to be 1.05 ± 0.04h (2σ ; Amini et al., 2009; Huang et al., 2010b). Other processes that may cause Ca isotope variability in igneous lithologies include fluid–rock interactions (John et al., 2012) as well as reprocessing of low-temperature mineralization products such as Ca-carbonates and Ca-sulfates in near-surface environments (e.g., Gussone et al., 2003; Huang et al., 2011; Marriott et al., 2004). The Ca isotope variability in Martian meteorites is still poorly understood, but scant available data imply similar Ca isotope compositions to those of the Earth’s mafic lithologies (Simon and DePaolo, 2010; Valdes et al., 2014). Here, we aim to contribute further to the understanding of geochemical history of Mars through the Ca isotope analyses of SNC (Shergotty–Nakhla–Chassigny) meteorites and Allan Hills (ALH) 84001 orthopyroxenite, and to constrain the δ 44/40 Ca of the Bulk Silicate Mars (BSM), compare it with that of the Earth, and evaluate the processes leading to Ca isotope variations among Martian samples. 2. Samples and methodology Calcium isotope compositions were obtained for 23 Martian meteorites. Following petrologic classification for shergottites, four basaltic shergottites (Shergotty, Zagami, Los Angeles, Northwest Africa [NWA] 856), six olivine-phyric shergottites (Elephant Moraine [EETA] 79001 lithology A, Larkman Nunatak [LAR] 06319, Yamato [Y] 980459, NWA 1068, NWA 6162, Sayh al Uhaymir [SaU] 005), one olivine–orthopyroxene-phyric shergottite (Roberts Massif [RBT] 04262), two lherzolitic shergottites (ALH 77005, Y-000097), and one diabasic shergottite (NWA 5990) were analyzed. Six clinopyroxene-bearing nakhlites (Nakhla, Lafayette, Miller Range [MIL] 03346, Y-000593, NWA 817, NWA 5790), two dunitic chassignites (Chassigny, NWA 2737), and ALH 84001 were also measured for Ca isotope compositions. Fusion crusts were removed and only interior portions of the samples were processed. Powdered whole-rock samples were dissolved in a mixture of ultrapure concentrated HNO3 –HF (1/6 v/v) in Teflon® screwtop vials; dried residues were refluxed repeatedly with concentrated HNO3 and equilibrated in 6M HCl. An aliquot corresponding to ∼1.8 μg Ca was taken from the stock solution, which was checked for the absence of precipitates in order to exclude laboratory-induced Ca isotope fractionation, and spiked with a mixed 42 Ca–43 Ca tracer prior to ion-exchange separation. Calcium was purified using 100-μl columns packed with MCI-gel (Mitsubishi Chemicals) and 1.8N HCl following the method described in Teichert et al. (2009). Calcium isotope ratios were measured using a Triton T1 thermal ionization mass spectrometer, housed at the Universität Münster, following the protocols of Gussone et al. (2011). The Ca isotope data in this study is reported in per mil deviation from the NIST SRM 915a reference material (Eisenhauer et al., 2004) and calculated as δ 44/40 Ca (h) = [(44 Ca/40 Ca)sample /(44 Ca/40 Ca)SRM915a − 1] × 1000. Three to four analyses of SRM 915a material per sample wheel were performed with unknown samples. The δ 44/40 Ca of unknown samples were

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normalized to the average SRM 915a value of the session. The reliability of the analytical technique was assessed through the analysis of reference rocks and replicate analyses of unknown samples (see Table 1). Replicate analyses of unknown samples reproduced generally within better than ±0.10h (2SD). 3. Results The stable Ca isotope compositions of Martian meteorites are listed in Table 1 along with CaO contents and plotted in Fig. 1. Age-corrected Ca isotope compositions (from 40 K decay to 40 Ca) differ less than 0.01h from the uncorrected values and make these corrections negligible (see Kreissig and Elliott, 2005, and Caro et al., 2010, for the effects of radiogenic 40 Ca ingrowth). This minor correction is in large part due to young magmatic ages of most meteorites from this study combined with low K/Ca < 0.03, with the exception of chassignites having K/Ca of 0.06–0.07. The overall δ 44/40 Ca range is ca. 0.4h (from 0.71 to 1.14h) and does not correlate with CaO contents or petrography in a simple way, perhaps with the exception of slightly larger δ 44/40 Ca variation with increasing modal pyroxene (mostly clinopyroxene; orthopyroxene is dominant in ALH 8400 and is a minor phase in EETA 79001A and LAR 06319; Fig. 2a). Basaltic shergottites span the range in δ 44/40 Ca with Shergotty showing the heaviest and Zagami the lightest Ca isotope composition of the whole Martian suite. The results for Zagami differ from those given by Farkaš et al. (2009) and Simon and DePaolo (2010) outside the analytical uncertainty (Fig. 1) but this may reflect lithological heterogeneity of Zagami (McCoy et al., 1999), also noted for recent Li data obtained from the same aliquot (Magna et al., 2015). The rather narrow range of δ 44/40 Ca of 0.91 ± 0.03h (2SD) for the enriched (based on chondrite-normalized LaN /SmN ; Anders and Grevesse, 1989) olivine(–orthopyroxene)-phyric shergottites (LAR 06319, NWA 1068, RBT 04262) closely resembles that of basaltic shergottites (0.94 ± 0.35h, 2SD) indicating that incompatible element enrichment/depletion (Fig. 2b) is not paralleled by differences in Ca isotope compositions. These findings are underscored by intermediate and depleted shergottites having a mean δ 44/40 Ca of 1.03 ± 0.09h (2SD) and 0.93 ± 0.24h, respectively. Nakhlites show a mean δ 44/40 Ca of 0.88 ± 0.19h (2SD), mimicking the δ 44/40 Ca range found for shergottites, but at significantly higher CaO contents due to large amounts of modal clinopyroxene. Two dunites (Chassigny, NWA 2737) reveal a distinct δ 44/40 Ca that differs by ca. 0.3h despite broadly similar petrography and chemistry, with the exception of the presence of carbonate (∼1%) in NWA 2737 compared with only traces of carbonate in Chassigny. The orthopyroxenite ALH 84001 has a δ 44/40 Ca of 1.06h, indistinguishable from that reported by Simon and DePaolo (2010) within the uncertainty. 4. Discussion 4.1. Calcium isotope composition of Mars and significance for the inner Solar System planets In order to derive an estimate for the δ 44/40 Ca of BSM, several factors must be considered: (i) The ∼2.4 wt.% CaO in BSM (McSween, 2003) is significantly less than ∼3.6 wt.% CaO in Bulk Silicate Earth (McDonough and Sun, 1995). From the average mineral assemblage of the Earth’s upper mantle (Salters and Stracke, 2004) and considering high-temperature Ca isotope fractionation for the orthopyroxene– clinopyroxene pair of 44/40 Caopx–cpx = ∼0.4–0.8h in mantle peridotites (e.g., Amini et al., 2009; Feng et al., 2014), Huang et al. (2010b) calculated δ 44/40 CaBSE = 1.05 ± 0.04h (2SD). Due to its low CaO content, olivine can be considered a subordinate Caphase (De Hoog et al., 2010) despite high modal abundance in

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Table 1 Calcium contents and isotope compositions in Martian meteorites. Petrography

CaO (wt.%)

δ 44/40 CaSRM915a (h)

2SE

2SD

n

olivine-phyric basalt (NASA subsample 47) basalt (USNM 7058) basalt olivine-phyric basalt olivine–orthopyroxene-phyric basalt (NASA subsample 45) basalt (USNM 6676) basalt

6 .5 10.0 10.2 7. 9 5 .7 9 .6 10.5

0.93 0.99 0.91 0.90 0.90 1.14 0.71 0.97 0.87

0.14 0.02 0.06 0.02 0.07 0.10 0.01

0.20 0.03 0.08 0.02 0.10 0.14 0.01 0.05 0.14

2 2 2 3 2 2 2

Intermediate shergottites ALHA 77005 lherzolite (NASA subsample 223) EETA 79001A olivine-phyric basalt (NASA subsample 644) Y-000097 lherzolite (NIPR subsample 71)

3 .2 7.3 4 .2

1.05 1.05 0.98

0.01 0.03 0.07

0.02 0.05 0.10

2 2 2

Depleted shergottites NWA 5990 NWA 6162 SaU 005 Y-980459

diabasic basalt olivine-phyric basalt olivine-phyric basalt (MPI 1522/2) olivine-phyric basalt (NIPR subsample 53)

6 .5 5 .1 5 .7 6. 8

0.95 0.90 0.80 1.08

0.04 0.01 0.06 0.05

0.06 0.02 0.09 0.08

2 2 2 3

Nakhlites Lafayette MIL 03346 Nakhla (F)

clinopyroxenite (USNM 1505) clinopyroxenite (NASA subsample 193) clinopyroxenite

13.4 15.0 14.7

0.07 0.04 0.01

clinopyroxenite clinopyroxenite clinopyroxenite (NIPR subsample 121)

0.07 0.04 0.05

0.09 0.06 0.02 0.14 0.10 0.06 0.07

3 2 2

NWA 817 NWA 5790 Y-000593

0.94 0.89 0.81 0.89 0.79 0.82 1.04

2 2 2

Chassignites Chassigny (F) NWA 2737

dunite (NHMV subsample H 3398) dunite

0.66 0.84

1.06 0.74

0.10 0.07

0.15 0.10

2 2

Orthopyroxenites ALH 84001

orthopyroxenite (NASA subsample 413)

1 .8

1.06 1.01

0.07

0.10 0.05

2

11.3

0.90 0.83 0.87 1.15

0.05 0.07 0.03 0.03

0.11

5

0.07

4

Enriched shergottites LAR 06319 Los Angeles 001 NWA 856 NWA 1068 RBT 04262 Shergotty (F) Zagami (F)

[1] [2]

[2] 13.1 14.0 14.9

[1] Reference rocks BHVO-2

Hawaiian basalt [3] [4]

JP-1

peridotite

0.56

“F” next to meteorite name denotes an observed fall. The degree of enrichment/depletion is based on chondrite-normalized LaN /SmN values (>0.7 for enriched, 0.3–0.7 for intermediate and <0.3 for depleted shergottites). n – number of individual analyses. [1] Simon and DePaolo (2010); [2] Farkaš et al. (2009); [3] Amini et al. (2009); [4] Valdes et al. (2014). CaO contents are taken from literature (see Supplementary information).

the mantle. However, it may still represent a non-negligible Ca carrier, particularly in deep mantle regions dominated by a lowCa orthopyroxene–olivine assemblage where the solubility of Ca in olivine increases (Köhler and Brey, 1990). Garnet may be present as a minor phase in the Martian mantle (Draper et al., 2003) but its effect on the δ 44/40 CaBSM (and δ 44/40 CaBSE , by inference) remains unconstrained due to the lack of experimental data. Similarly, the effect of perovskite, which possibly formed above the Martian core–mantle boundary (Bertka and Fei, 1997) is yet unconstrained. (ii) The extent to which clinopyroxene, orthopyroxene, olivine, and garnet (and possibly perovskite in the deep mantle) fractionate Ca isotopes critically depends on the Ca–O bond length in the respective phase (see Huang et al., 2010b). Moreover, lower Ca/Mg in orthopyroxene leads to heavier Ca isotope compositions compared with orthopyroxene having high Ca/Mg, probably related to Ca–Mg substitution (Feng et al., 2014). By inference, very low Ca/Mg in olivine could generate similar effects. Indeed, Ca-rich olivine (monticellite), Ca-doped enstatite and Ca-garnet (grossular) all show shorter weighted mean Ca–O distances (<2.40 Å; Beran et al., 1996; Nestola and Tribaudino, 2003; Sharp et al., 1987) than clinopyroxene (>2.43 Å; Cameron et al., 1973). These crystal-

lographic data combined with the existing Ca isotope compositions of orthopyroxene and clinopyroxene suggest that the Ca isotope composition of olivine, and possibly also garnet, may be generally heavier than that of clinopyroxene. These theoretical considerations are in line with the observation that δ 44/40 Ca of Chassigny (olivine cumulate) is identical within analytical uncertainty to that of ALH 84001 (orthopyroxene cumulate), both specimens being isotopically heavier than most Martian basalts, rich in clinopyroxene. (iii) Smith et al. (1983) inferred from CaO contents in olivine grains of Chassigny that this meteorite may represent a product of relatively shallow-level crystallization from a parent magma, but its δ 44/40 Ca of 1.06h is similar to that inferred for pristine terrestrial ultramafic rocks (Amini et al., 2009). The assumed shallow origin thus imposes little, if any, change to the primordial δ 44/40 Ca of deep mantle olivine-dominated lithologies, and probably the Martian mantle itself. From these above inferences and combined with the finding that lherzolitic shergottites (here represented by ALH 77005 and Y-000097) and olivine-phyric shergottite Y-980459 (considered to reflect the mantle source of depleted shergottites; Usui et al., 2008) most closely reflect the Martian mantle composi-

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Fig. 1. The δ 44/40 Ca versus CaO contents in Martian meteorites from this study. Nakhlites have very high CaO concentrations due to high modal augite whereas olivine-dominated chassignites and ALH 84001 are low in CaO contents due to low CaO in olivine and orthopyroxene. Incompatible trace element-enriched shergottites have in general slightly higher CaO abundances than intermediate and depleted shergottites; in particular, basaltic shergottites tend to higher CaO, which is also reflected in their rather high modal plagioclase. Different aliquots of Zagami reveal varying δ 44/40 Ca: S – Simon and DePaolo (2010); F – Farkaš et al. (2009); Z – this study. This difference implies that the measured values may not be representative of the bulk sample/melt and, by inference, may not reflect the average value of the source. Data for chondrites are taken from elsewhere (see Supplementary information). CaO contents in bulk Earth’s continental crust and upper mantle are from Rudnick and Gao (2014) and McDonough and Sun (1995), respectively; CaO content of the bulk Martian crust and bulk silicate Mars is from Taylor and McLennan (2009) and McSween (2003), respectively. The δ 44/40 Ca range of terrestrial basalts and Ca isotope composition of terrestrial upper mantle is from Huang et al. (2010b).

tion (Nyquist et al., 2009), an initial estimate for the δ 44/40 CaBSM of 1.04 ± 0.11h (2SD) is derived. The similarity in incompatible trace elements and Sm–Nd systematics between lherzolitic shergottites and olivine-phyric EETA 79001A (Nyquist et al., 2009) allows integrating its Ca isotope composition into the calculated mean δ 44/40 Ca of 1.04 ± 0.09h (2SD), which is adopted here for the BSM. Interestingly, this value is indistinguishable from that found for the olivine-dominated meteorite Chassigny as well as orthopyroxenedominated ALH 84001. Thus, the similarity in δ 44/40 Ca between Chassigny–ALH 84001 and lherzolitic shergottites–Y-980459–EETA 79001A versus resolvedly lower δ 44/40 Ca found for most other shergottites (enriched and depleted) and nakhlites, combined with the essentially identical estimated mean δ 44/40 CaBSM (this study) and δ 44/40 CaBSE (Huang et al., 2010b), implies a broadly similar sense of Ca isotope fractionation between major mineral phases in the mantles of Mars and Earth. The similar stable isotope compositions of mantles of the Earth and Mars were also noted for Mg (Teng et al., 2010), whilst Si (Zambardi et al., 2013), Fe (Poitrasson et al., 2004) and O (Clayton and Mayeda, 1996) isotope compositions of Martian meteorites are different. For extra-terrestrial materials other than those derived from Mars, Ca isotope data are scarce and only a more detailed assessment of Ca isotope signatures in chondritic and achondritic meteorites may provide further constraints on Ca isotope systematics of the inner Solar System planetary bodies. Limited Ca isotope data for Moon and Vesta (Bermingham et al., 2011; Simon and DePaolo, 2010; Valdes et al., 2014) appear to be similar to the δ 44/40 Ca range observed for Earth and Mars. On the contrary, significant δ 44/40 Ca heterogeneity exists among chondrites (Fig. 1) but the origin of these Ca isotope variations remains elusive (Huang et al., 2012; Simon and DePaolo, 2010; Valdes et al., 2014) and a large uncertainty in the proportion of

Fig. 2. The δ 44/40 Ca versus modal pyroxene (a) and chondrite-normalized (Anders and Grevesse, 1989) LaN /SmN values (b) in Martian meteorites from this study. In (a), mostly pigeonite and augite are considered; orthopyroxene is dominant only in ALH 84001 and subordinate orthopyroxene is reported for a few Martian meteorites. In (b), no particular difference is found for δ 44/40 Ca despite the large range in LaN /SmN values. Modal mineralogy and trace element abundances are taken from the literature (see Supplementary information).

chondritic/non-chondritic materials contributing to the mass of the Earth and Mars must be considered (Warren, 2011). The heterogeneous Ca isotope compositions between different chondrite classes are also evident from the large range in δ 44/40 Ca of calcium– aluminum-rich inclusions (Farkaš et al., 2010; Huang et al., 2010a; Niederer and Papanastassiou, 1984; Simon et al., 2009). The contrasting homogeneity in δ 44/40 Ca of inner Solar System planets could imply efficient attenuation of such nucleosynthetic anomalies prior to the onset of accretion of terrestrial planets as suggested by Simon et al. (2009) although 48 Ca anomalies were recently reported for various chondrites and achondrites (Dauphas et al., 2014). 4.2. Calcium isotope perspective on pre-terrestrial low-temperature alteration The δ 44/40 Ca values of Chassigny and NWA 2737 differ by ∼0.3h despite sharing strong chemical and mineralogical similarities (Beck et al., 2006). Whilst there is no evidence for distinc-

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tive mineralogy of the two chassignites, petrographic investigations revealed that alteration products (i.e., carbonates) formed in NWA 2737 (Beck et al., 2006; Treiman et al., 2007) compared with only trace amounts of Ca–Mg-rich carbonates in Chassigny (Wentworth and Gooding, 2004). Considering the low CaO content of chassignites (<0.84 wt.%), even a minute amount of secondary Ca-rich phase may thus contribute significantly to the Ca isotope composition of the bulk rock, irrespective whether pre-terrestrial or terrestrial in origin. Notably, some carbonates in NWA 2737 are clearly pre-terrestrial, as evidenced by shock-induced dislocations (Beck et al., 2006), low terrestrial exposure of NWA 2737 to hot desert weathering processes (<10 kyr; Meyer, 2012), and high levels of Xe from the Martian atmosphere (Marty et al., 2006). These latter authors suggested that NWA 2737 is among the least terrestrially altered SNC meteorites recovered to date. The Ca isotope signature of inorganic carbonate is generally lighter compared to the fluid from which it precipitated, with 44/40 Ca values (defined as δ 44/40 Cacarb − δ 44/40 Cafluid ) ranging from ∼0 to about −2h (cf. Gussone et al., 2003; Lemarchand et al., 2004; Marriott et al., 2004). Parameters that influence the extent of Ca isotope fractionation include temperature (Gussone et al., 2003; Marriott et al., 2004), growth rate (Lemarchand et al., 2004), and crystal structure of the carbonate minerals (Gussone et al., 2005, 2011). The δ 44/40 Ca of carbonate in NWA 2737 can be approximated by simple mass balance calculation, as exemplified by John et al. (2012), using Eq. (1) and Eq. (2), respectively. A conceptual mixing model was thus constructed in order to test whether the ∼0.9 modal% of CaCO3 phase can quantitatively explain the δ 44/40 Ca difference between NWA 2737 and Chassigny (Fig. 3):

δ 44/40 CaWR = (1 − a) × δ 44/40 Casilicate + a × δ 44/40 Cacarb

(1)

or,

 δ 44/40 Cacarb = 1/a × δ 44/40 CaWR

 + (a − 1) × δ 44/40 Casilicate ,

Fig. 3. A schematic model describing the Ca isotope composition in chassignites. The > 0.3h difference in δ 44/40 Ca is most likely related to the presence of a CaCO3 phase in NWA 2737; on Earth, Ca-carbonate usually carries light Ca isotope values relative to magmatic rocks (see main text). If mixing between a carbonate-rich reservoir and a ‘near-parental’ Shergotty-like end member is considered in order to account for light Ca isotope signature in other basaltic shergottites, then such mixing may not easily explain the dataset. Instead, an isotopically light silicate reservoir with some resemblance to Zagami (and perhaps isotopically even lighter and chemically more evolved composition) is required to reproduce low δ 44/40 Ca without significant increase in CaO contents that would parallel the addition of carbonate. This component could be of crustal origin or could have formed from an evolved mantle magma.

(2)

where a is the proportion of carbonate-bound Ca. Taking into account the similar mineralogy of chassignites, it can be assumed that δ 44/40 Ca values of the silicate fraction are identical for both meteorites. Two scenarios are then calculated for a: (1) ∼0.9 modal% CaCO3 contributes 0.56 wt.% CaO to the bulk 0.84 wt.% CaO and the remaining 0.28 wt.% CaO is bound in the silicate fraction. Solving Eq. (1) yields δ 44/40 Cacarb = 0.59h, i.e., 0.47h lighter than the silicate fraction. This is similar to the extent of carbonate–silicate fractionation (44/40 Cacarb–sil ) observed for terrestrial settings (John et al., 2012). However, this assumption may overestimate the amount of Ca from the carbonate phase, since the calculated CaO content of 0.28 wt.% contributed by the silicate fraction is lower than the ‘magmatic CaO’ content of Chassigny (0.66 wt.%). (2) The silicate fraction of NWA 2737 resembles Chassigny both in Ca isotope composition (δ 44/40 Ca = 1.06h) as well as CaO content (0.66 wt.%). Then, 0.28 wt.% CaO (or, 20%) of the whole-rock Ca in NWA 2737 is contributed by CaCO3 . Solving Eq. (1) yields δ 44/40 Cacarb of −0.55h, i.e., 44/40 Cacarb–sil = −1.61h. This approximation may underestimate the Ca contribution from the carbonate if pure calcite is considered, but it appears permissible that these above approximations delimit the range of possible Ca isotope compositions of Martian carbonates; therefore, the effect of such mixing between silicate Ca and inorganic carbonate Ca may be discernible (Fig. 3). The calculated 44/40 Cacarb-sil values between −0.5 and −1.6h are in the range of experimental and field observations of Ca isotope fractionation during CaCO3 precipitation in terrestrial hydrous

systems (Brown et al., 2013; Gussone et al., 2005, 2003, 2011; John et al., 2012; Lemarchand et al., 2004; Marriott et al., 2004; Reynard et al., 2011; Tang et al., 2008). Alternatively, the light δ 44/40 Ca in NWA 2737 might originate from yet unspecified alteration on Mars by fluids that carried a relatively light Ca isotope signature, possibly from the dissolution of Ca-rich carbonates and/or sulphates (Holmden, 2009). 4.2.1. Calcium isotope evidence for the absence of plate tectonics on Mars? Although the extent of Ca isotope fractionation between silicate rocks and secondary carbonates cannot be constrained more precisely for Mars from the current dataset, the low δ 44/40 Ca of NWA 2737 proves that whatever the origin of carbonate on Mars, it likely reflects precipitation from solute sources that must have been saturated with respect to CaCO3 . Therefore, enough CO2 must have been present either in the atmosphere or dissolved in percolating fluids (Grott et al., 2011; Stanley et al., 2011) in order to form carbonates. Abundant surface exposures of carbonates were identified indicating that, during its history, Mars experienced mild climatic conditions at near-neutral pH and had enough atmospheric CO2 to precipitate carbonates (Morris et al., 2010). This is consistent with a low formation temperature of ∼18 ◦ C determined for carbonate in ALH 84001 (Halevy et al., 2011). Carbonate precipitation in ALH 84001 post-dates crystallization of the whole rock (Borg et al., 1999; Lapen et al., 2010) but bulk δ 44/40 Ca of ALH 84001 does not differ from that estimated for the Martian mantle (Table 1). The post-magmatic formation of carbonates could provide further constraints on the timing of degassing events. Sufficient CO2 levels were present in the Martian atmosphere during the mid/late Noachian (∼3.9 to 3.7 Ga) after the initial loss of primordial CO2 through degassing, escape and atmosphere erosion (e.g., Kass and Lung, 1995). Because the magmatic age of Chassigny and NWA 2737 is significantly younger (∼1.38 Gyr; Misawa et al., 2005), it allows to be suggested that

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explain the Ca isotope variations for Hawaiian basalts most likely does not apply to Mars. The lack of Martian meteorites with progressively low δ 44/40 Ca values thus hints at a negligible input of surface lithologies back into the mantle and argues against major recycling of crust into the mantle, and large-scale plate tectonics on Mars in general. 4.3. Further ramifications of Ca isotope variability in SNC meteorites

Fig. 4. Calcium isotope compositions versus Sr/Nb in the Martian suite. No systematic variations are observed for different classes of Martian meteorites (cf. Hawaii; Huang et al., 2011). Lafayette is the most equilibrated of all nakhlites in this study, with the highest Sr/Nb ratio. The other samples show lower δ 44/40 Ca and Sr/Nb, and may represent more evolved (and shallower seated) parts of the nakhlite pile (e.g., Day et al., 2006). Trace element abundances are taken from the literature (see Supplementary information).

conditions aiding in carbonate formation were maintained during a substantial part of Martian history. This, in turn, could provide indirect evidence for continuous magmatic degassing and/or percolation of CaCO3 -saturated fluids, at least in localized areas. Importantly, Ca-sulfates (i.e., anhydrite, gypsum and bassanite) were detected in various places on Mars and in Martian meteorites (e.g., Clayton and Mayeda, 1988; Squyres et al., 2012). Their role for the Ca isotope inventory of Mars remains unclear although terrestrial analogs carry isotopically light Ca (Holmden, 2009), consistent with experimental results revealing Ca isotope fractionation during precipitation of gypsum and anhydrite from aqueous fluid (44/40 Camin–fluid from −0.8 to −2.2h; Harouaka et al., 2014; Hensley, 2006). Moreover, accessory phosphates (apatite, merillite, whitlockite), carrying predominantly light Ca in terrestrial settings (Farkaš et al., 2010), are rather uncommon in Martian meteorites and their presence does not correlate with bulk δ 44/40 Ca. For example, these accessory phases are absent in Zagami, which has the lowest δ 44/40 Ca of the whole suite. Huang et al. (2011) inferred from light Ca isotope compositions in Hawaiian lavas that mixing of small (∼4%) amounts of recycled ancient marine carbonates can account for correlated behavior of Ca and Sr isotopes, consistent with other indicators of carbonate presence such as increased Sr/Nb and Ba/Y, and with progressively lighter Ca isotope compositions in those lavas that were likely contributed by larger volumes of carbonates. The current data set for Martian meteorites shows systematics between Sr/Nb and δ 44/40 Ca that are dissimilar to the Hawaii suite (Fig. 4) but we note that Huang et al. (2011) investigated samples from one volcanic center while Martian meteorites represent a range of possible mantle sources spanning a range in magmatic ages more than 1 Gyr. Although our results are in favor of isotopically light Ca in non-silicate surface materials and perhaps support Ca isotope dichotomy similar to that observed for Earth, there are likely no suitable physical conditions (Halevy and Schrag, 2009) for the presence and/or recycling of large amounts carbonate at any stage during the petrogenesis of Martian rocks. This, in turn, can be explained by the absence of plate tectonic processes on Mars (Breuer and Spohn, 2003; Grott et al., 2013) that would ultimately feed the source regions of the deep mantle with ample quantities of isotopically light Ca. Therefore, the model of Huang et al. (2011) to

Several further implications are apparent from Ca isotope systematics in Martian meteorites (Fig. 1). First, the mean δ 44/40 Ca of three olivine(–orthopyroxene)-phyric enriched shergottites is not different from that of the enriched basaltic shergottites implying that modal olivine imposes a negligible effect on the bulk δ 44/40 Ca due to its low CaO content. This is consistent with the findings for olivine-rich cumulates, such as lherzolitic shergottites and Chassigny, which show limited Ca isotope variability despite different levels of olivine accumulation, implying very limited isotope fractionation by incorporation of Ca into olivine. Second, no relationship is observed for δ 44/40 Ca versus Mg# (molar Mg/(molar Mg + molar Fe)) for the Martian sample suite (not shown), suggesting that extreme degrees of magmatic differentiation such as experienced by the Los Angeles shergottite (Rubin et al., 2000) do not impart measurable variations to δ 44/40 Ca. These observations are further tied to the lack of systematic variations between δ 44/40 Ca and incompatible element indices of magmatic differentiation (Fig. 2b). And third, the generally light Ca isotope signature in shergottites, with the exception of the most primitive shergottites (Y-980459, lherzolitic shergottites) is similar to the values observed for terrestrial basalts from various geotectonic settings (Amini et al., 2009; DePaolo, 2004; Jacobson et al., 2015; Skulan et al., 1997; Valdes et al., 2014). Valdes et al. (2014) noted that lower δ 44/40 Ca in terrestrial basalts relative to the estimated BSE may result from igneous processes but recycling of weathered residues and carbonates should be equally considered (Fantle and Tipper, 2014) because Earth’s continental crust carries isotopically light Ca relative to mafic silicates (e.g., Amini et al., 2009; De la Rocha and DePaolo, 2000; John et al., 2012; Marshall and DePaolo, 1989; Skulan et al., 1997; Valdes et al., 2014). Mantle plumes on Earth produce high-temperature, low-degree partial melts (e.g., Klein and Langmuir, 1987; Prytulak and Elliott, 2007) that commonly contain recycled oceanic lithosphere (Sobolev et al., 2007). They have probably been active on Earth at least since the Archean (e.g., Tomlinson and Condie, 2001) and are still an important source of basalts today. Their incompatible element characteristics are different from those of large-scale mantle-derived melts at mid-ocean ridges (e.g., Hofmann, 2014). Mantle plumes are considered to have been the major source of volcanism on Mars throughout its history (Grott and Breuer, 2010; Hynek et al., 2011) although it is still unclear how the assumed Martian plumes could survive for a long time on a largely inactive planet (see Reese et al., 2004 for impact-triggered melting scenarios). It is notable that there is no systematic difference in Ca isotope compositions between shergottites and nakhlites despite largely different La/Sm ratios (Fig. 2b) and petrogenetic histories (e.g., Blinova and Herd, 2009; Debaille et al., 2009; Papike et al., 2009). Moreover, Zr/Hf ratios are broadly similar between shergottites and nakhlites without resolved variations in δ 44/40 Ca (Fig. 5) while they differ in major mantle-derived reservoirs on Earth (e.g., David et al., 2000; Münker et al., 2003). These coupled variations may reflect a variable extent of partial melting in the stability field of majorite garnet (Debaille et al., 2009; Draper et al., 2003; Weyer et al., 2003) but the lack of Ca in majorite requires at least a two-stage evolution with additional fractional crystallization of clinopyroxene (David et al., 2000; Hart and Dunn, 1993)

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Juraj Farkaš for comments on an earlier draft of the manuscript. Detailed and constructive reviews by three anonymous referees and editorial handling by Bernard Marty are greatly acknowledged. This work was funded by the Helmholtz Association through the research alliance HA 203 “Planetary Evolution and Life”. T.M. is grateful for a partial support through the Czech Science Foundation grant 13-22351S. Appendix A. Supplementary material Supplementary material related to this article can be found online at http://dx.doi.org/10.1016/j.epsl.2015.08.016. References

Fig. 5. The δ 44/40 Ca versus Zr/Hf in Martian meteorites. Higher Zr/Hf ratios are in general associated with higher modal proportion of plagioclase. Chondritic Zr/Hf of 34.3 is from Münker et al. (2003); the mean δ 44/40 Ca of the Martian mantle is derived in Section 4.1. Trace element abundances are taken from the literature (see Supplementary information).

with a predominantly light Ca isotope signature (Feng et al., 2014; Huang et al., 2010b). 5. Conclusions The range in δ 44/40 Ca (∼0.4h) found for the compositionally variable suite of Martian meteorites (shergottites, nakhlites, ALH 84001) could reflect processes of magmatic differentiation and/or percolation by low-temperature Ca-rich solute sources although the former process cannot be accounted for easily at present. While mixing with chemically more evolved surface lithologies cannot be excluded, the current Ca isotope data set, coupled with incompatible element systematics, does not require the incorporation of surface carbonate into the source regions of Martian meteorites, thereby precluding recycling of surface lithologies back into the Martian mantle and thus large-scale plate tectonics. The Ca isotope compositions of SNC meteorites are consistent with experimental predictions of Ca isotope fractionation between major mineral phases; e.g., δ 44/40 Ca of olivine is predicted to be similar to that of orthopyroxene although this relationship has yet to be investigated in a greater detail for a larger suite of Martian meteorites and their major mineral phases. The δ 44/40 Ca difference of ∼0.3h between the two chassignites is most likely related to a presence of small amounts of late-formed carbonate, likely preterrestrial in origin, suggesting a prolonged existence of Ca-rich fluids and a CO2 -rich atmosphere on Mars. The role of sulfates cannot be constrained with the available suite. A mean δ 44/40 Ca of 1.04 ± 0.09h (2SD) is estimated for the Martian mantle which can also be considered as an approximation for the Bulk Silicate Mars. This estimate is identical to that derived for the Bulk Silicate Earth at ∼1.05h (Huang et al., 2010b) and is consistent with the assertion of Simon and DePaolo (2010) who proposed similar bulk Ca isotope compositions of the inner Solar System planets. Acknowledgements We are grateful to NASA, Tim McCoy (Smithsonian Institution), Franz Brandstätter (NHM Vienna), Hideyasu Kojima (NIPR) and Jutta Zipfel (Senckenberg Forschungsinstitut und Naturmuseum Frankfurt) for sample allocations. We thank Jan Cempírek, Albert Jambon and Stephan Klemme for additional information, and

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