High temperature strontium stable isotope behaviour in the early solar system and planetary bodies

Earth and Planetary Science Letters 329–330 (2012) 31–40

Contents lists available at SciVerse ScienceDirect

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

High temperature strontium stable isotope behaviour in the early solar system and planetary bodies B.L.A. Charlier a,⁎, G.M. Nowell b, I.J. Parkinson a, S.P. Kelley a, D.G. Pearson b, c, K.W. Burton a, b a b c

CEPSAR, The Open University, Walton Hall, Milton Keynes MK7 6AA, United Kingdom Department of Earth Sciences, University of Durham, South Road, Durham DH1 3LE, United Kingdom Department of Earth and Atmospheric Sciences, University of Alberta, Edmonton, Canada T6G 2E3

a r t i c l e

i n f o

Article history: Received 11 July 2011 Received in revised form 13 February 2012 Accepted 14 February 2012 Available online 22 March 2012 Editor: R.W. Carlson Keywords: strontium isotopes 88 Sr/86Sr Zr mass-dependent fractionation MC–ICP-MS

a b s t r a c t This study presents comprehensive strontium stable isotope (88Sr/ 86Sr) data, measured by multiple-collector inductively-coupled plasma mass spectrometry (MC–ICP-MS), for a suite of carbonaceous chondrites, differentiated meteorites, lunar, martian and terrestrial samples. Carbonaceous chondrites comprise a mixture of refractory inclusions and chondrules that have light δ88Sr values between − 0.35 to + 0.05‰ and matrix material that possesses a heavy δ88Sr composition of 0.65‰, confirming the results of earlier studies. Consequently, bulk carbonaceous chondrites are relatively heterogeneous in composition ranging from + 0.12 to + 0.35‰, most likely reflecting clast–matrix variability. The light δ88Sr compositions of the refractory inclusions are consistent with mass dependent fractionation of other refractory elements (Ca and Eu) and are most likely produced by non-equilibrium fractionation (undercooling in the nebula gas) during condensation of hibonite from the solar nebula (Simon and DePaolo, 2010. Stable calcium isotopic composition of meteorites and rocky planets. Earth Planet. Sci. Lett. 289, 457–466). Carbonaceous chondrites, angrites and martian meteorites have indistinguishable compositions at the level of analytical uncertainty of this study. However, statistical analysis indicates that melts derived from the Earth's mantle have heavier δ88Sr values than bulk carbonaceous chondrites and martian meteorites, but compositions indistinguishable from eucrites (δ88Sr = + 0.26 ± 0.12‰). Moreover, terrestrial basalts and andesites have restricted δ88Sr values (+ 0.30 ± 0.07‰), suggesting that mantle melting delivers rather homogenous melts to the Earth's surface with respect to δ88Sr. In contrast, glasses from evolved terrestrial rocks and lunar basalts extend to very light δ88Sr values ~−0.20‰. The Sr stable isotope composition covaries with europium anomaly (Eu/Eu*), as an index of plagioclase fractionation, and δ88Sr can be successfully modelled by the heavy isotopes of Sr being preferentially partitioned into plagioclase with a fractionation factor of ~1.0007 for 88Sr/86Sr. Our results demonstrate that Sr stable isotopes may be significantly fractionated at high temperatures and their measurement can provide insights into planetary evolution and magmatic processes. © 2012 Elsevier B.V. All rights reserved.

1. Introduction Strontium is a refractory lithophile element and has four stable isotopes, 84Sr (0.56%), 86Sr (9.86%), 87Sr (7.00%) and 88Sr (82.58%). However, 87Sr is also produced by the long-lived radioactive decay of 87 Rb, and this parent–daughter isotope system has been widely utilised over the past 40 years for geochronology and as a geochemical tracer. Conventional radiogenic Sr isotope measurements made using thermal ionisation mass spectrometry (TIMS) are corrected for instrumental mass fractionation assuming a value for the 88Sr/86Sr ratio of 8.375209, and this internal normalisation permits the measurement

⁎ Corresponding author. Tel.: + 44 1908 652558; fax: + 44 1908 655151. E-mail address: [email protected] (B.L.A. Charlier). 0012-821X/$ – see front matter © 2012 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2012.02.008

of 87Sr/ 86Sr ratios to a very high precision. To some degree, normalisation to a fixed 88Sr/ 86Sr ratio assumes that this ratio is constant between natural samples (although the choice of the actual figure used is to some extent arbitrary), and by definition normalisation using a mass dependent isotopic fractionation law removes the signature of any mass-dependent variation in the 88Sr/ 86Sr ratio present in natural samples. Early work on refractory calcium–aluminium rich inclusions (CAIs) in carbonaceous chondrites, showed that they preserve anomalous Sr isotope compositions (Papanastassiou and Wasserburg, 1978). However, this study only reported Sr isotope variations attributable to non-mass dependent processes which are most-likely due to a deficit in the least abundant isotope 84Sr, a p-process nucleus, resulting from production at distinct nucleosynthetic sites (Birck, 2004; Papanastassiou and Wasserburg, 1978). In principle, stable isotope fractionation of a heavy element such as Sr is expected to be relatively small, particularly

32

B.L.A. Charlier et al. / Earth and Planetary Science Letters 329–330 (2012) 31–40

at high temperatures. However, it has long been known that for some meteorites, there exist significant and measurable variations in the 88 Sr/ 86Sr isotope composition. The pioneering work of Patchett (1980a, b) using double spike techniques and TIMS, demonstrated that some chondrules and refractory inclusions in the carbonaceous chondrite “Allende” preserve covariations between 88Sr/ 86Sr and 84Sr/ 86 Sr consistent with the expected effects of mass-dependent fractionation, with CAIs and chondrules being depleted in the lighter isotopes of Sr by up to ~ 3‰ compared to the bulk meteorites and terrestrial samples. The variations observed by Patchett (1980a, b) for CAIs and chondrules were dismissed by some workers because over-spiking of the samples may have propagated significant uncertainties into the data (Niederer and Papanastassiou, 1984). The enrichment in light isotopes is indeed difficult to explain since refractory inclusions preserve evidence for evaporative loss and recondensation (Richter et al., 2002; Shahar and Young, 2007), and mass dependent theory predicts that the condensed phase should be enriched in heavy isotopes. Such enrichment in the refractory phase has been observed for a number of elements during experimental evaporation of material with a solar composition (e.g. Davis et al., 1995; Esat et al., 1986; Wang et al., 2001). Nevertheless, Ca (Niederer and Papanastassiou, 1984; Simon and DePaolo, 2010), Eu (Moynier et al., 2006), Zn (Luck et al., 2005) and Cd (Wombacher et al., 2008) all show enrichments in light isotopes in refractory inclusions, attributed to kinetic effects (Richter, 2004) or electromagnetic separation (Moynier et al., 2006). Over recent years, with the advent of multi-collector inductivelycoupled plasma mass-spectrometry (MC–ICP-MS), Sr stable isotope variations have been observed in a range of terrestrial materials (Aggarwal, 2006; De Souza et al., 2007; Fietzke and Eisenhauer, 2006; Halicz et al., 2008; Krabbenhöft et al., 2010; Ohno and Hirata, 2006), with differences that are easily resolved at current analytical precision. In particular, the fractionation of Sr stable isotopes accompanying the precipitation of calcium carbonate (Fietzke and Eisenhauer, 2006; Rüggeberg et al., 2008) highlights the potential of Sr stable isotopes as a tracer of weathering processes or changes in carbonate sedimentation (Krabbenhöft et al., 2010). Until recently there had been no attempt to duplicate the results of Patchett (1980a, b), or indeed to investigate the possibility of mass-dependent fractionation in differentiated meteorites or other terrestrial planets. However, the study of Moynier et al. (2010) confirmed that chondrules and refractory inclusions in Allende do indeed have light 88Sr/86Sr isotope compositions. Moreover, they also reported data for terrestrial igneous rocks (see also Ohno and Hirata, 2006), martian and lunar meteorites, and differentiated basaltic meteorites that indicated little 88Sr/ 86Sr isotope variation at the level of precision of that study (±50 ppm). This study presents comprehensive 88Sr/ 86Sr data for carbonaceous chondrites (and chondrules and refractory inclusions), eucrites (thought to be derived from the asteroid 4-Vesta), angrites, martian and lunar rocks, and terrestrial igneous rocks (that range from primitive mid-ocean ridge basalts to highly evolved continental rhyolites). These data, like those of Moynier et al. (2010), confirm the variations in Sr stable isotope composition for CAIs and chondrules reported previously (Patchett, 1980a, b), albeit at a reduced magnitude. These results also show that Mars, along with 4-Vesta and the angrite parent body (as sampled here), possess 88 Sr/ 86Sr ratios that are indistinguishable from those of bulk carbonaceous chondrites at the resolution of our measurements. However, mantle-derived terrestrial igneous rocks have 88Sr/ 86Sr ratios that are subtly, although statistically significantly, heavier than bulk carbonaceous chondrites, while more evolved rhyolite glass samples and lunar basalts possess significantly lighter stable isotope compositions, demonstrating that high-temperature magmatic processes are capable of generating measurable variations in Sr stable isotopes.

2. Analytical techniques 2.1. Sample preparation For bulk chondrites, eucrites, angrites, martian rocks and MORB samples, small fragments ranging from ~100 mg to >2 g were cleaned ultrasonically several times alternately in MilliQ water (resistivity > 18 MΩ) and ultra-pure methanol. After drying thoroughly in an oven at 100 °C, chips were crushed in a scrupulously cleaned agate mortar in a filtered air environment. Lunar rocks and terrestrial rock standards were supplied as powders and underwent no further treatment prior to dissolution. The rhyolite glass sample P1209 (Oruanui eruption, Taupo volcano) was obtained using standard heavy liquid separation techniques. For the component parts of Allende and Vigarano, as well as for rhyolite glasses YTT-33 (Youngest Toba tuff) and BFC-138 (Fish Canyon tuff), samples were milled from cleaned, polished thick sections using a Merchantek Micromill following the protocols detailed by Charlier et al. (2006). 2.2. Sample digestion and chemical separation Sample digestion was achieved with the addition of ~ 1 ml 29 M HF + trace 10 M HNO3 to the sample in a 3.5 ml Savillex beaker, which was then sealed and placed on a hotplate (ca. 120 °C) overnight. Excess HF and volatile SiF4 were removed by subsequent evaporation of the sample solution. Following evaporation, the sample was sequentially dried and taken up again in 6 M HCl, then 10 M HNO3 and ultimately equilibrated with 3 M HNO3 in preparation for column separation. Carbonaceous chondrites generally required several sustained periods at the 10 M HNO3 stage to ensure full dissolution of the carbonaceous component. Separation of Sr was carried out using cleaned Sr Spec extraction chromatographic resin (Horwitz et al., 1992) (particle size 50–100 μm) supplied by Eichrom Technologies Inc (Darien, Il.) using the separation scheme detailed by Charlier et al. (2006). Total procedural blanks were routinely b20 pg Sr, which can be considered negligible in comparison to the ≫100 ng of Sr used for sample analysis. Dissolved sample amounts ranged from 0.86 mg (Vigarano CAI), to 37 mg (rock standards). 2.3. Mass spectrometry Samples and standards were analysed as ~200 ppb Sr solutions to which a 600 ppb Zr single element standard was added just prior to analysis. Isotope ratios were determined using Thermo-Fisher Neptune MC–ICP-MS instruments at the University of Durham and later at the Open University. Amplifier gains were determined at the start of each day, and 120 s long baselines were measured prior to each analysis. Each analysis was carried out in static multi-collection mode and comprised 1 block of 50 cycles each with a 4-second integration. All data were processed offline for mass bias and interfering element corrections. The analytical method used was identical to that used by Charlier et al. (2006) for the analysis of Rb isotopes, as follows. The measured 88Sr/ 86Sr (after correction for 86Kr interference on mass 86Sr by monitoring mass 83Kr) was used to calculate the mass bias corrected 90Zr/ 91Zr in a Zr-doped NBS-987 (strontium) solution. 88 Sr/ 86Sr is the most appropriate ratio to measure because 84Sr has a low abundance and is difficult to measure by MC–ICP-MS due to a significant 84Kr interference. Based on the 90Zr/ 91Zr value determined in this fashion, the mass bias corrected 88Sr/ 86Sr was then determined in Zr-doped unknown Sr samples. We found that by plotting ln[ 90Zr/ 91 Zr] versus ln[ 88Sr/ 86Sr], the standard data clustered into distinct groups for each analytical session. From this we infer that the relative difference in mass bias coefficients for Sr and Zr is constant within each analytical session, but varies from session to session. For this reason, we chose to determine the 90Zr/ 91Zr for each day of analysis,

B.L.A. Charlier et al. / Earth and Planetary Science Letters 329–330 (2012) 31–40

rather than use a fixed value for this ratio. Despite a limited variation in mass bias within individual analytical sessions, regression of the ln [ 90Zr/ 91Zr] and ln[ 88Sr/ 86Sr] data yield slopes consistent with both exponential fractionation and 90Zr/ 91Zr values from the literature (Sahoo and Masuda, 1997). The 90Zr/ 91Zr determined over the five analytical sessions used here (4.5869 ± 7 (2 s.d.)), is identical to that previously determined by Charlier et al. (2006) (4.5870 ± 12). 87Sr/ 86 Sr ratios in NBS 987 determined by the Zr normalisation technique versus conventional within-run internal normalisation (e.g., with 86Sr/ 88 Sr= 0.1194) yield results that are within uncertainty, indicating the Zr doping technique is robust. 2.4. Standard data To express variations in the isotopic composition of Sr relative to NBS 987, the standard delta per mil notation (δ 88Sr ‰) is used: 2
88

 3 86 Sr= Sr samp −15  1000 δ Sr ð%oÞ ¼ 488 86  Sr= Sr NBS− 987 88

ð1Þ

where 88Sr/ 86SrNBS − 987 = 8.375209 ( 86Sr/ 88SrNBS − 987 ≡ 0.1194). To assess the long-term external precision of the technique, NBS 987 standards that were not included in calculating the daily 90Zr/ 91 Zr ratio, were also run in each of the analytical sessions. Over the five analytical sessions this yielded an 88Sr/ 86Sr value of 8.37517 ± 0.00053 (2 s.d., n = 66) equivalent to a δ 88Sr of −0.005 ±0.063‰. We also duplicated measurements of the rock standards BHVO-1, BIR-1 and AGV-1 and these all reproduced within 2 s.d. analytical uncertainty. The possibility of column fractionation influencing isotopic fractionation was carefully considered, and a series of experiments was run to address this issue (see Supplementary material). 3. Results 3.1. Extraterrestrial samples 3.1.1. Carbonaceous chondrites To assess variability in the stable Sr isotope composition in the early solar system we have analysed a suite of primitive carbonaceous chondrites, and the data are reported in Table 1 and presented in Fig. 1b. These include bulk-rock analyses of representative meteorites from four (CI, CM, CO and CV) of the seven carbonaceous chondrite classes. We have also analysed CAIs, chondrules, a dark inclusion and matrix from Allende and a CAI from Vigarano. These data allow us to study isotopic fractionation between the different components of Allende and assess whether there are differences between oxidised (Allende) and reduced (Vigarano) CV3s. Moreover, our Allende data can be compared with those of Moynier et al. (2010) and Patchett (1980a, b) (Fig. 1a). The δ 88Sr values for bulk carbonaceous chondrites range from + 0.13 ± 0.06 for Alais to +0.35 ± 0.10‰ for Vigarano, which are distinctly heavier than the other samples. There is no obvious systematic variation between the different classes of carbonaceous chondrites and the population has a mean of +0.20 ± 0.14‰ (compared to + 0.19 ± 0.20‰ for the samples studied by Moynier et al. (2010)). The CAIs, chondrules and dark inclusions in Allende all have distinctly light δ 88Sr values, compared to the bulk data. Of these components, the CAIs have the lightest values with δ 88Sr from −0.36 ± 0.14 to −0.07 ± 0.08‰, but with no obvious difference between coarse and fine grained CAIs, and a dark inclusion sampled from within a CAI has a value of 0.05 ± 0.14‰. The two chondrules measured have slightly heavier values of −0.10 ± 0.08 and +0.14 ± 0.08‰. Matrix material was carefully hand-picked to avoid contamination, although it is possible that small fragments of CAI and chondrule were analysed. In contrast to all of the other component parts of Allende the

33

Table 1 Stable Sr isotope composition of carbonaceous chondrites (and component parts), achondrites, martian, lunar and terrestrial samples. MORB samples MD-37, MD-57 and EW-93 were those used by Gannoun et al. (2007); BFC138, Fish Canyon Tuff (Bachmann et al., 2002); P1209, Taupo rhyolite glass (Wilson et al., 2006); YTT30, Youngest Toba Tuff (Jones, 1993). All other terrestrial samples are commercially available bulk rock standards (USGS and GSJ). Uncertainties on the isotope ratio represent 2 s.e. on individual measurements. Dupl. and trip. represent repeated measurements of independent dissolutions. Rock type

Sample

88

Sr/86Sr

CI1 CM2 CM2 CM2 CO3 CO3 CV3 CV3 CV3 CV3 CV3 CV3 CV3 CV3 CV3 CV3 CV3 CV3 CV3 CV3 CV3 CV3 CV3 CV3 Eucrite Eucrite Eucrite Eucrite Eucrite Eucrite Eucrite Eucrite Eucrite Eucrite Eucrite Eucrite Eucrite Angrite Angrite Angrite Angrite Angrite Martian Martian Martian Martian Lunar Lunar Lunar Lunar Lunar Lunar Lunar Lunar MORB MORB MORB Basalt Basalt Basalt Basalt Basalt Basalt Basalt Andesite Andesite

Alais Murchison Nogoya Cold Bokkeveld Kainsaz Ornans Bali Kaba Mokoia Allende Allende (dupl.) Vigarano Allende—Fine-grained CAI Allende—Fine-grained CAI Allende—Coarse-grained CAI Allende—Coarse-grained CAI Allende—Coarse-grained CAI Allende—Coarse-grained CAI Allende—Coarse-grained CAI Allende—Dark inclusion in CAI Allende—Chondrule Allende—Chondrule Allende—Matrix Vigarano—CAI Camel Donga Dhofar 007 Ibitera Jonzac Jonzac (dupl.) Juvinas Serra de Magé Millbillillie Pasamonte Sioux County Stannern Talampaya Moore County Sahara 99555 Angra Dos Reis D'Orbigny D'Orbigny (dupl.) D'Orbigny (trip.) Nahkla Zagami Shergotty Los Angeles 15555 15555 (dupl.) 67955 71501 NWA032 NWA3160 68501 12032 MD-37 MD-57 EW-93 DNC-1 BIR-1 BIR-1 (dupl.) BHVO-1 BHVO-1 (dupl.) JB-2 W2 AGV-1 AGV-1 (dupl.)

8.3763 8.3766 8.3774 8.3768 8.3776 8.3768 8.3766 8.3763 8.3761 8.3774 8.3771 8.3782 8.3731 8.3746 8.3726 8.3731 8.3722 8.3731 8.3725 8.3756 8.3764 8.3744 8.3806 8.3757 8.3774 8.3768 8.3772 8.3783 8.3775 8.3767 8.3777 8.3773 8.3769 8.3772 8.3780 8.3781 8.3772 8.3776 8.3775 8.3766 8.3768 8.3770 8.3767 8.3767 8.3766 8.3763 8.3775 8.3764 8.3767 8.3765 8.3764 8.3752 8.3765 8.3769 8.3776 8.3777 8.3778 8.3774 8.3776 8.3777 8.3777 8.3779 8.3784 8.3773 8.3779 8.3778

2 s.e.

δ88Sr (‰)

2 s.e.

0.0005 0.0005 0.0003 0.0004 0.0002 0.0002 0.0003 0.0003 0.0003 0.0007 0.0007 0.0008 0.0008 0.0007 0.0008 0.0006 0.0011 0.0007 0.0005 0.0012 0.0007 0.0007 0.0006 0.0007 0.0005 0.0004 0.0003 0.0003 0.0008 0.0009 0.0004 0.0005 0.0006 0.0006 0.0005 0.0004 0.0003 0.0004 0.0005 0.0007 0.0009 0.0005 0.0010 0.0013 0.0008 0.0011 0.0011 0.0009 0.0009 0.0004 0.0008 0.0014 0.0013 0.0009 0.0002 0.0002 0.0004 0.0005 0.0009 0.0012 0.0007 0.0009 0.0009 0.0010 0.0007 0.0009

0.13 0.17 0.26 0.20 0.28 0.19 0.17 0.14 0.10 0.26 0.23 0.35 − 0.25 − 0.07 − 0.31 − 0.25 − 0.36 − 0.26 − 0.32 0.05 0.14 − 0.10 0.65 0.05 0.26 0.19 0.23 0.37 0.27 0.17 0.30 0.25 0.20 0.24 0.33 0.35 0.23 0.29 0.27 0.16 0.19 0.22 0.17 0.17 0.16 0.13 0.27 0.14 0.18 0.15 0.14 0.00 0.15 0.21 0.29 0.29 0.32 0.27 0.28 0.29 0.29 0.32 0.39 0.25 0.32 0.32

0.06 0.07 0.03 0.05 0.03 0.02 0.04 0.03 0.04 0.09 0.09 0.10 0.10 0.08 0.09 0.07 0.14 0.08 0.06 0.14 0.08 0.08 0.08 0.09 0.06 0.04 0.04 0.04 0.10 0.11 0.04 0.06 0.07 0.07 0.06 0.05 0.04 0.05 0.06 0.08 0.11 0.06 0.11 0.15 0.10 0.13 0.13 0.11 0.11 0.05 0.09 0.17 0.15 0.10 0.03 0.03 0.04 0.06 0.11 0.15 0.08 0.11 0.10 0.12 0.08 0.11

(continued on next page)

34

B.L.A. Charlier et al. / Earth and Planetary Science Letters 329–330 (2012) 31–40

Table 1 (continued) Rock type

Sample

88

Sr/86Sr

Quartz-latite Granodiorite Granodiorite Granite Granite Rhyolite glass Rhyolite Rhyolite glass Rhyolite glass Rhyolite glass K-feldspar

QLO-1 GSP-1 GSP-1 (dupl.) G2 G2 (dupl.) RGM-1 JR-2 BFC138 P1209 YTT30 SRM607

8.3769 8.3777 8.3776 8.3779 8.3783 8.3750 8.3767 8.3766 8.3757 8.3736 8.3746

2 s.e.

δ88Sr (‰)

2 s.e.

0.0005 0.0007 0.0006 0.0004 0.0009 0.0013 0.0010 0.0004 0.0006 0.0011 0.0005

0.20 0.29 0.29 0.32 0.37 − 0.03 0.18 0.16 0.05 − 0.19 − 0.07

0.06 0.08 0.07 0.05 0.11 0.15 0.12 0.05 0.07 0.13 0.06

matrix is isotopically heavy with a δ 88Sr of + 0.65 ± 0.08‰; this is in fact the heaviest extraterrestrial or terrestrial silicate material that we have measured. Our data concur with those of Moynier et al. (2010), who also found that the matrix was the heaviest material (+0.66 ± 0.04‰). We have only analysed one CAI from Vigarano, but in accordance with the Allende data, it is significantly lighter than the bulk composition with a value of + 0.05 ± 0.09‰.

88

8.350

8.360

a

8.380

8.390

Matrix Coarse CAIs

3.2. Terrestrial samples

Fine CAIs

In order to assess the stable Sr isotope composition of melts generated from the Earth's mantle and the effects of fractional crystallisation, we have analysed a suite of igneous rocks that range in composition

Chondrules -3

-2

3.1.3. Lunar samples We have analysed a range of lunar samples that include three basalts, three regolith samples and an anorthosite (Table 1 and Fig. 2). The regolith and anorthosite samples have a very restricted compositional range between +0.15± 0.05‰ to +0.21 ± 0.10‰. In contrast, the lunar basalts have a wider range of compositions, extending from + 0.27 ± 0.13‰ for the primitive basalt sample 15555 to 0.00 ± 0.17‰ for the most evolved lunar basalt NWA3160. 3.1.4. Martian meteorites In this study we have determined the composition of four martian meteorites, three basaltic shergottites and one nahklite. The four samples encompass a limited composition range of +0.13 ±0.13‰ to +0.17 ± 0.15‰ and are identical at the level of analytical uncertainty (Table 1 and Fig. 2). By comparison Moynier et al. (2010) analysed three shergottites that gave a slightly higher average value of +0.25±0.07.

Sr/86Sr 8.370

3.1.2. Achondrites We have analysed a suite of achondrites (Table 1, Fig. 2). Within this suite we have analysed eleven eucrites (thought to have originated from 4-Vesta), and three angrites: Sahara 99555, Angra Dos Reis and D'Orbigny. The eucrites define a normally distributed population that has an average value of + 0.26 ± 0.12‰. By comparison Moynier et al. (2010) analysed one eucrite and one diogenite that gave an average value of +0.33 ± 0.04‰. The first two angrites have indistinguishable values of +0.29 ± 0.05 and +0.27 ± 0.06‰ respectively, whereas D'Orbigny has a weighted mean value of +0.20 ± 0.04‰ (2 s.d., MSWD = 0.73, n = 3 on independent dissolutions).

-1

0

1 88Sr/86Sr

8.370

8.372

8.374

8.376

8.378

8.380

8.370

8.382

8.372

8.374

8.376

8.378

b

Murchison Nogoya

CM2

Cold Bokkeveld Kainsaz - CO3.2 Ornans - CO3.4

Eucrites

Bali - CV Kaba Mokoia Allende bulk

8.380 Camel Donga Dhofar 007 Ibitera Jonzac Jonzac Juvinas Serra de Magé Milibillillie Pasamonte Sioux County Stannern Talampaya Moore County

Alais - CI

CV3

Vigarano bulk Allende fine CAI

Sahara 9555 Angra Dos Reis Allende coarse CAI

Angrites D'Orbigny

Allende dark inclusion

Nahkla Zagami Shergotty Los Angeles

Allende single chondrule Allende matrix Vigarano CAI

-0.5

-0.3

-0.1

0 0.1

0.3

0.5

Mars

15555

0.7

67955 71501 NWA 032 NWA 3160 68501 12032

δ 88Sr (‰) Moon Fig. 1. a) Compilation of δ88Sr in carbonaceous chondrites and their component parts from the study of Patchett (1980a, b) analysed by double-spike TIMS and Moynier et al. (2010) analysed by MC–ICP-MS. To allow comparison with our data, Patchett's NBS 987 data have been re-normalised to 88Sr/86SrNBS 987 = 8.375209, which produce a shift to 0.4‰ lighter values for all of his data. Fields represent data for CAIs, chondrules and matrix. b) Plot of δ88Sr for bulk carbonaceous chondrites and their component parts analysed by MC–ICP-MS (this study). Error bars represent 2 s.e. uncertainties on individual measurements. The thick line is the best estimate for the terrestrial mantle based on an average of our new data and those of Moynier et al. (2010).

-0.5

-0.3

-0.1 0

0.1

0.3

0.5

δ 88Sr (‰) Fig. 2. Plot of δ88Sr in achondrites, martian and lunar samples analysed by MC–ICP-MS in this study. Error bars represent 2 s.e. uncertainties on individual measurements and the thick line is the best estimate for the terrestrial mantle.

B.L.A. Charlier et al. / Earth and Planetary Science Letters 329–330 (2012) 31–40

35

ratio for a given 88Sr/86Sr ratio (known as 87Sr/86Sr*) for marine Sr data. While this is useful for modelling Sr isotope data in 86Sr–87Sr–88Sr space, it is potentially confusing if workers re-normalise their TIMS data 88 to a different Sr/86Sr ratio. Therefore, we prefer to indicate a mean mantle value for δ88Sr of +0.29 (based on this work and the data of Moynier et al., 2010), in Figs. 1–3 so that it is clear what the relation of the various planetary bodies are to the composition of the Earth's mantle.

from primitive basalts to glasses from highly evolved rhyolites (Table 1 and Fig. 3). The majority of the samples are international rock standards but are supplemented by three MORB samples and rhyolitic glasses from the Toba Tuff, the Fish Canyon Tuff and the Taupo Volcanic Zone. Our data can be simply divided into two groups. The majority of samples, which include all the basalts, one each of the andesite, granodiorite and granite samples are all enriched in 88Sr relative to NBS 987 and have a restricted compositional range, with δ88Sr of +0.30± 0.07‰ (2 s.d.). In contrast, the more evolved samples that include all of the rhyolite glasses and a quartz latite have distinctly lighter δ 88Sr values spanning from + 0.20 ± 0.06 to − 0.19 ± 0.13‰. We also analysed the K-feldspar standard NBS 607, which has a light δ88Sr value of −0.07± 0.06‰.

4.2. Strontium stable isotope variations in planetary materials A large body of work now exists on the composition of stable isotopes (e.g., Li, O, Ca, Mg and Fe) in chondrites, the Earth, Moon, Mars and 4-Vesta. Our new data adds further information to the debate on the origin of compositional differences and similarities between these planetary materials. The variations in δ 88Sr between the bulk compositions of each of these groups is limited, however there is enough variation relative to the analytical uncertainty to make some first order observations. We use the student t-test to make a statistical analysis of the difference between the various groups of data. In some cases 2 s.e. uncertainties on individual analyses are larger than the stated external precision for the NBS 987 standard, and in order to derive useful statistical information these uncertainties must be incorporated into the t-test. This can be accomplished by using a Monte Carlo simulation and we use a methodology similar to that used by De Souza et al. (2007). We here compare the following data sets: basaltic terrestrial rocks, carbonaceous chondrites, Martian meteorites, eucrites and the angrites. We illustrate below in Section 4.3 that the evolved terrestrial samples have been affected by high-level fractionation processes. We have not included our lunar data in this analysis, because of the inference that these too are affected by fractionation and mixing processes. For each data set we take the uncertainty on each individual sample and generate 5000 data points for each analysis that have a Gaussian distribution around its mean value. Thus, for example, the population for the carbonaceous chondrites is composed of 5000 data sets each containing 12 data points, with each point part of a Gaussian distribution around its mean value. The student t-test is then undertaken on each of the 5000 data sets that are generated and the mean probability calculated between the various populations.

4. Discussion 4.1. Note on the NBS 987 standard and the normalisation of stable Sr data It is interesting to note that the majority of the terrestrial silicate samples analysed in this study and those reported by Moynier et al. (2010) have values that are heavier than the NBS 987 standard. This standard is supplied as strontium carbonate, but it is not clear whether the original sample was a natural carbonate or if the Sr was derived from a silicate source and the Sr carbonate synthesised in the laboratory. NBS 987 has an 87Sr/ 86Sr ratio of ~0.71025, which is consistent with the Sr being derived from a crustal source. Furthermore, it is now well known that low temperature precipitation of carbonate can significantly fractionate the Sr isotopes, with the carbonate having lighter δ 88Sr than the fluid from which it was precipitated (e.g., Fietzke and Eisenhauer, 2006), and so the light δ 88Sr value may well have been generated during preparation of the standard. Several other non-traditional stable isotope systems (Ca, Fe) are referenced to standards that have compositions comparable to mantle or terrestrial composition, (e.g. δ 44Ca of the silicate Earth ~0‰; Simon and DePaolo, 2010). While it might be tempting to re-normalise our data to a notional terrestrial composition, there is a large body of radiogenic 87Sr/86Sr data that has been corrected to 88Sr/86Sr ratio of 8.375209. In principle re-normalising the data would require recalculating the 87Sr/ 86Sr ratio to be self-consistent. This approach has been utilised by Krabbenhöft et al. (2010), who compute the 87Sr/86Sr

88Sr/86Sr

8.370

8.372

8.374

8.376

8.378

8.380 MD-37 MD-57 EW-93 DNC-1

Less evolved

MORB glasses

BIR-1 BHVO-1 JB-2 W-2 AGV-1

Earth

QLO-1 GSP-1 G2

RGM-1 JR-2 Fish Canyon Tuff bulk Taupo rhyolite glass Youngest Toba Tuff bulk SRM 607 K-fsp Rb-Sr std

More evolved -0.5

-0.3

-0.1

0 0.1

0.3

0.5

δ 88Sr (‰) Fig. 3. Plot of δ88Sr in terrestrial igneous rocks with samples becoming more evolved down the figure. Error bars represent 2 s.e. uncertainties on individual measurements and the thick line is the best estimate for the terrestrial mantle.

36

B.L.A. Charlier et al. / Earth and Planetary Science Letters 329–330 (2012) 31–40

An obvious limitation is that the results from the student t-test are only meaningful if the populations being tested are representative (i.e. that we capture the true variance for each of the populations). While our stable Sr data set is not exhaustive, we have incorporated the data of Moynier et al. (2010) into our analyses and this currently provides our best means to assessing the variations in stable Sr isotopes between different planetary bodies. Results from the t-test on the non-evolved terrestrial samples are presented in Table 2. Several conclusions can be reached from these data. 1). Melts from the Earth's mantle are isotopically heavier than carbonaceous chondrites and Mars at 99% and 90% confidence levels respectively, indicating that the Earth has the most distinct bulk δ 88Sr composition of the solar system material we have measured. 2). Eucrites are isotopically heavier than Mars at ~90% confidence levels, but are indistinguishable from the Earth. 3). Carbonaceous chondrites, angrites and martian samples have compositions that are the same at the level of analytical uncertainty. Lunar data are hard to assess, but the lunar basalt that is least affected by plagioclase fractionation (15555) has a δ 88Sr value of +0.26 which overlaps with the values for terrestrial basalts. Our statistical analyses suggests that mantle-derived melts from the Earth have the heaviest δ 88Sr values of all the material we have analysed and are consistently heavier than carbonaceous chondrites (see also Moynier et al., 2010). However, the combined data from this study and that of Moynier et al. (2010) would indicate that carbonaceous chondrites are heterogeneous and it is possible that some of the lighter δ 88Sr values are because we have preferentially sampled CAIs. Although there is a correlation with CAI abundance and δ 88Sr in the data of Moynier et al. (2010), it would require ~10% of a CAI with a δ 88Sr of − 1‰ to shift the mean value of the carbonaceous chondrites to lighter values by 0.1‰. Therefore we suggest that there is a small but significant difference between carbonaceous chondrites and the terrestrial mantle, but this needs to be tested in more detail by analysing bulk carbonaceous chondrite powders that are more homogeneous. The combined data from this study and those of Moynier et al. (2010) raise several key questions that require more detailed studies of stable strontium isotopes at higher precision. The origin of the Earth having a slightly heavier δ 88Sr signature than carbonaceous chondrites could be due to several processes. It is possible that there was a small amount of volatile loss of light Sr during planetary formation, and this is consistent with our data for other differentiated planetary bodies such as Mars and 4-Vesta. Alternatively, heavy Sr may have been recycled into the mantle at subduction zones since modern seawater has a δ 88Sr +0.38‰ and hydrothermal fluids at ridge crests have values that lie between 0.26 and 0.38‰ (Fietzke and Eisenhauer, 2006; Krabbenhöft et al., 2010; Parkinson et al., 2008). Strontium may be lost from the subducted slab via dehydration reactions, thus making the mantle wedge heavier in a similar manner to that envisaged for Li isotopes (Elliott et al., 2004). In contrast, the Earth is statistically indistinguishable from both the enstatite or ordinary chondrite data of Moynier et al. (2010), which may be coincidence or provide further evidence for an enstatite–chondrite origin for the Earth (e.g., Javoy et al., 2010). Higher-precision determinations of Table 2 p-Values calculated from a student t-test for the planetary and meteoritic groups analysed in this study. Terrestrial samples only include mantle derived melts. The student t-test takes into account the uncertainties on individual measurements within each group by using the Monte-Carlo simulations described in the text.

Earth Eucrite Carb. chond. Mars Angrites

Earth

Eucrite

Carb. chond.

Mars

Angrites

1 0.497 0.005 0.052 0.132

0.497 1 0.023 0.107 0.249

0.005 0.023 1 0.631 0.489

0.052 0.107 0.631 1 0.542

0.132 0.249 0.489 0.542 1

δ 88Sr would help to further resolve and explain some of the differences between planetary bodies that we have identified in this study. 4.3. Evolved magmas and the role of feldspar The majority of terrestrial rocks we have measured have a fairly uniform δ 88Sr composition of + 0.30 ± 0.07‰ (2 s.d.). Most of these are basaltic to andesitic in composition suggesting that mantle melting delivers melts to the Earth's surface that have restricted stable Sr isotope compositions, and that fractionation of olivine and clinopyroxene have little effect on δ 88Sr. By contrast, highly evolved melts that range in composition from andesites to high-silica rhyolites have δ 88Sr values that extend from +0.19 to − 0.19‰, and represent the only terrestrial silicates to have light δ 88Sr values. These melts are characterised by decreasing Sr and increasing Rb with increasing SiO2, consistent with extensive plagioclase and potassium feldspar fractionation, with δ 88Sr correlating negatively with Rb/Sr ratios. Therefore, there are two possible explanations for the decrease in δ 88Sr. First, fractionation is coupled with crustal assimilation, (e.g., assimilation fractional crystallisation (AFC)), such that the melts assimilate progressively more crust with light δ 88Sr. Second, that feldspar preferentially partitions heavy Sr isotopes, producing melts with increasing light δ 88Sr with increasing fractionation. Both of these models are relatively simple to test and can be modelled quantitatively. 4.3.1. Assimilation fractional crystallisation As part of this study we routinely analysed all the samples for the radiogenic 87Sr/ 86Sr ratio, which can be used as a monitor of crustal assimilation. For the limited sample suite we have analysed, there is no systematic relationship between the initial 87Sr/ 86Sr ratio and δ 88Sr, with an r 2 of 0.1, with only the Young Toba Tuff sample having a relatively elevated 87Sr/ 86Sr ratio of 0.719 (Jones, 1993) and a light δ 88Sr. Analysis of the evolved terrestrial lavas in this study is not the ideal way to assess the effects of crustal contamination on stable Sr isotopes, and this would be better accomplished by a systematic study of a single magmatic system that has been demonstrably contaminated. However, our first order observation is that crustal contamination is not the dominant driver for δ88Sr becoming progressively lighter in the evolved samples. 4.3.2. Fractional crystallisation modelling High-temperature fractionation of Sr isotopes during fractional crystallisation could occur due to equilibrium or kinetic fractionation, with the heavier Sr isotopes preferentially partitioned into plagioclase or potassium feldspar relative to the melt. There are no published equilibrium or kinetic fractionation factors for Sr, but it is possible to estimate their approximate values. For equilibrium fractionation, a useful starting point is to use published values for elements of similar mass to Sr. At ~730 °C (an approximate temperature for granite petrogenesis), Ge, Br and Ru compounds yield fractionation factors (f) of 1.0002–1.0010 per atomic mass unit (amu) (Schauble, 2004), equivalent to or 1.0004–1.0020 for 88Sr/ 86Sr ratios. For kinetic fractionation factors we can use the simple empirical relationship derived for diffusion-related fractionation by Richter et al. (2006) to estimate the possible fractionation factor, f, f ¼

  m88 β m86

ð2Þ

where m88 and m86 are the masses of the Sr isotopes and β is an experimentally derived exponent. Other group 2 metals, such as Mg and Ca, have β values of 0.05 and 0.075 respectively, whereas Ge (with a mass similar to Sr) has a β values b0.025. Using Eq. (2) and β values of 0.025–0.075 yield kinetic fractionation factors of 1.0006–1.0017, which are likely to be extreme estimates for the kinetic fractionation factor.

B.L.A. Charlier et al. / Earth and Planetary Science Letters 329–330 (2012) 31–40

We have analysed a selection of evolved rocks from a variety of localities rather than a suite of co-magmatic rocks from a single locality. This makes modelling slightly more problematic because each of the melts may have had different initial Sr and Rb contents. A simple way to deal with this problem is to model europium anomalies (Table 3) as a proxy for plagioclase fractionation, because it is reasonable to assume that primary melts have Eu/Eu* values of 1 ± 0.05 and the modelling is not sensitive to knowing absolute REE concentrations in the primary melts. This approach has the additional advantage of allowing an assessment of differences between terrestrial and lunar samples, because the Eu plagioclase/melt partition coefficients are likely to be higher in the more reduced Moon. However, modelling the Sr isotope ratios does require some estimate of the initial Sr concentration in the melt. We use values of 400–600 ppm based on published data for Taupo (Wilson et al., 2006), the Fish Canyon Tuff (Bachmann et al., 2002) and Toba (Jones, 1993), and 250 ppm for the lunar basalts (Schnetzler et al., 1972). The Sr partition coefficients are derived from the temperature and compositional dependent expression for plagioclase/melt partitioning of Blundy and Wood (1991). For the terrestrial samples, we assume a mean temperature of 775 °C and a plagioclase composition of An35 based on estimates from geothermometry and mineral chemistry for evolved rhyolitic systems (e.g., Bachmann et al., 2002; Sutton, 1995), which yields a partition coefficient of ~7.4 for plagioclase. For lunar samples we use 1000 °C and a plagioclase composition of An85, the average anorthite content of the three lunar basalt modelled here (Crawford, 1973; Fagan et al., 2001; Zeigler et al., 2006), which yields a partition coefficient of ~ 1.5 for plagioclase. There are likely to be differences in the REE partitioning between the lunar basaltic melts and evolved terrestrial melts, due to the more reduced nature of the Moon and the large temperature differences. For the lunar data we utilise measured REE plagioclase/glass partitioning data from 15555 (Schnetzler et al., 1972) and for the rhyolites, plagioclase/glass and sanidine/glass partition data from the Fish Canyon Tuff (Bachmann et al., 2002; Charlier, unpublished data). We have considered sanidine because it is a common phase in many of the rhyolites, although its REE partitioning data are not significantly different to plagioclase for the Fish Canyon Tuff. Concentrations of trace elements in the melt (CL) are calculated using the fractional crystallisation equation

CL ¼ C0 F

ðD−1Þ

ð3Þ

where C0 is the initial melt composition, F is the melt fraction and D is the partition coefficient. The δ 88Sr of the liquid (δ 88SrL) is then calculated using 88

δ SrL ¼ ½δ88 Sr0 þ 1000F L

15555 NWA032 NWA3160 MD-37 MD-57 EW-93 DNC-1 BIR-1 BHVO-1 JB-2 W2 AGV-1 QLO-1 GSP-1 G2 RGM-1 JR-2 BCF138 P1209 YTT30

ð4Þ

4.4.1. Previous studies Previous studies of non-radiogenic Sr isotopic variations in primitive meteorites are limited, but can be separated into those that resolve non0.6 0.5 0.4

Eu/Eu* 0.905 0.442 0.163 1.004 0.974 0.934 1.083 1.149 0.997 0.965 0.980 0.920 0.910 0.398 0.767 0.504 0.074 0.726 0.593 0.342

−1000

4.4. Carbonaceous chondrites

0.2 0.1 0 -0.1

Basalts and andesites

‘Anomalous’ terrestrial samples GSP-1

G-2

JR-2

0.3

δ 88Sr (‰)

Lunar basalt Lunar basalt Lunar basalt MORB MORB MORB Basalt Basalt Basalt Basalt Basalt Andesite Quartz-latite Granodiorite Granite Rhyolite Rhyolite Rhyolite Rhyolite Rhyolite

Sample

ðf −1Þ

where δ 88Sr0 is the initial stable Sr isotope composition and FL is the fraction of Sr remaining in the melt. Fig. 4 is a plot of δ 88Sr versus Eu/Eu* and illustrates that melts that have experienced plagioclase ± potassium feldspar fractionation lie on three distinct trends that evolve from basaltic and andesitic terrestrial rocks which have δ 88Sr of 0.30 ± 0.07‰ and Eu/Eu* of 0.92–1.13. Evolved terrestrial rhyolite glass samples define a curvi-linear trend that extends to a δ 88Sr of − 0.19‰ and a Eu/Eu* of 0.34. In contrast, lunar basalts range from + 0.26 to 0‰, but extend to a much lower Eu/Eu* value of 0.16, consistent with the increased compatibility of Eu. Three anomalous terrestrial samples, (granite G-2, granodiorite GSP-1 and a very Sr-depleted rhyolite, JR-2) lie on a much flatter trend that reaches a δ88Sr value of only +0.19‰ but with an extremely low Eu/Eu* of 0.07, indicating extensive plagioclase±potassium feldspar fractionation with only minor stable Sr isotope fractionation. Given that these are bulk samples (as opposed to glass separates), we cannot rule out that they may contain significant amounts of xenocrystic feldspar. The terrestrial rhyolite glass samples and lunar basalts have been modelled using the partitioning data described above with the only variable being the fractionation factor for Sr isotopes. Best fits of the two data sets yields a f value for 88Sr/86Sr of 1.000598 (β = 0.026) for the terrestrial rhyolites and 1.000823 (β = 0.036) for the lunar samples. This is on the lower end of the range estimated from theoretical studies, but the difference is hardly surprising, as there is no experimental data for Sr fractionation. Critically, the modelling indicates that extensive fractionation of a phase in which Sr is compatible, in this case plagioclase (±sanidine in the terrestrial samples), can significantly change the stable isotope composition of Sr in the evolving melt fraction. This is therefore a viable mechanism for generating significant δ88Sr variations in highly evolved magmatic systems on the Earth and may provide insights into the evolution of the lunar crust.

Table 3 0.5 0.5 Eu/Eu* for lunar and terrestrial rocks. Eu/Eu* is calculated as EuN / (SmN × GdN ) or 0.666 0.333 EuN / (SmN × TbN ) depending on whether Gd data were available. Rock type

37

x

x

x

x

x

x x x x

x

x

x

x

x

x

x

x

BFC-138

x

x

x x

P1209

x

x

Lunar x 0.2 samples

0.4 x

0.6 RGM-1

0.8

1

1.2

Eu/Eu*

x

-0.2 x

-0.3 -0.4

YTT-30

x

Rhyolite glasses

Terrestrial Lunar

Fig. 4. δ88Sr versus Eu/Eu* in terrestrial magmas and lunar basalts. Three model curves have been generated using Eqs. (3) and (4). The rhyolites and lunar basalts have been modelled using f values of 1.000598 and 1.000823 respectively for the Sr isotopic fractionation, but different Sr, Eu, Sm, and Gd partition coefficients (see text for details). Eu/Eu* data from Borg et al. (2009), Charlier et al. (2007), Gannoun et al. (2007), Jones (1993), Potts et al. (1992), Schnetzler et al. (1972), Sutton (1995) and Zeigler et al. (2006).

38

B.L.A. Charlier et al. / Earth and Planetary Science Letters 329–330 (2012) 31–40

mass dependent nucleosynthetic variations, (Andreasen and Sharma, 2007; Papanastassiou and Wasserburg, 1978), and those that resolve mass-dependent stable isotope variations, (Moynier et al., 2010; Patchett, 1980a, b). The double-spike Sr isotope data of Patchett (1980a, b) measures the true composition of all four Sr isotopes without any internal normalisation. Therefore, in principle, both mass dependent variations in 84 Sr/ 86Sr and 88Sr/ 86Sr as well as nucleosynthetic 84Sr/ 86Sr can be determined. However, the data of Patchett (1980a, b) are not of particularly high precision (2σ uncertainties of 0.15–0.4‰ on 88Sr/ 86Sr ratios and 0.4–0.9‰ on 84Sr/ 86Sr ratios), not particularly accurate (NBS 987 run as an unknown gives a δ 88Sr of 0.4‰ ± 0.28) and so only large depletions 84Sr/ 86Sr would be resolved. The 84Sr/ 86Sr and 88 Sr/ 86Sr data define a broad negative trend consistent with mass dependent fractionation, with no evidence of 84Sr depletions and with much of the CAI and chondrule data having negative δ 88Sr values. Niederer and Papanastassiou (1984) have suggested that much of the data of Patchett (1980a, b) can be dismissed because the samples were chronically over-spiked. As part of this study we have assessed the effects of the over-spiking using a similar methodology to that described in Galer (1999). Our calculations indicate that although all of the samples were over-spiked, the primary effect that this causes is a significant error magnification during double-spike deconvolution, making the data less precise. However, we would suggest that the lightest δ88Sr values (−3.2 to −2‰) could well be inaccurate. These data are the most over-spiked, are poorly reproduced and lie off the mass fractionation trend. The most likely reason for this is a differential blank problem, rather than over-spiking per se, since the spiked-runs have compositions far displaced from the natural blank composition. We conclude, that although Patchett's data are of low precision, the broad correlation between 84Sr/ 86Sr and 88Sr/86Sr (with slope ~−1) is consistent with considerable mass dependent fractionation and that his data suggest that CAIs and chondrules have δ 88Sr values that are light, extending to values as low as ~−1‰. This is further supported by the data of Moynier et al. (2010), who find δ 88Sr values as low −1.73‰ in some chondrules. 4.4.2. New data and the origins of δ 88Sr fractionation It is useful to compare our new bulk carbonaceous chondrite data with those of Moynier et al. (2010) and Patchett (1980a, b), although there are large analytical uncertainties associated with the latter. However, our new higher-precision data do confirm that CAIs and chondrules in Allende have light δ 88Sr values, albeit with a more restricted range than that reported by either Patchett (1980a, b) (see Fig. 1a) or Moynier et al. (2010). The critical question is why do CAIs and chondrules have light δ88Sr? Moynier et al. (2010) have suggested that redistribution of Sr between matrix and chondrules during aqueous alteration produces light chondrules and a complementary matrix with heavy δ88Sr values. Low Sr concentrations (~10 ppm) in the matrix and chondrules make them susceptible to alteration, although it would be useful to test the redistribution model in a carbonaceous chondrite other than Allende. CAIs have much higher Sr concentrations and here we focus on the processes that could produce mass-dependent fractionation in the CAIs, such as solid-state diffusional exchange, or possible fractionation processes in the solar nebula such as evaporation/condensation, non-equilibrium fractionation (e.g., Simon and DePaolo, 2010) or electromagnetic effects (e.g., Moynier et al., 2006). 4.4.3. Solid-state diffusional exchange It is now well established that the lighter isotopes of an element diffuse faster than heavier isotopes, which can generate significant kinetic mass dependent fractionation. Recent work has demonstrated that the amount of isotopic fractionation is a function of the fractionation factor and the concentration gradient (Parkinson et al., 2007; Richter et al., 2003). To generate negative δ88Sr values requires enrichment of Sr at the rim of the CAI or chondrule, so that 86Sr preferentially

diffuses into the grain. Patchett (1980a) reports a rim Sr concentration that is twice that of the core for a CAI from Allende, and we have unpublished laser ablation ICP-MS data which supports these observations. We have modelled both the Sr isotopic profiles and bulk mineral compositions for a variety of concentration gradients and f values (see Parkinson et al., 2007 for methodology). Using f values calculated from Eq. (2), which are likely to be too high (see Section 4.3.2.), would require concentration gradients greater than 10 to generate the light δ 88Sr values in the CAIs that we report. We conclude that diffusional exchange is unlikely to generate the observed δ 88Sr in the CAIs, unless f values are markedly higher than we have estimated, or that concentration gradients were appreciably higher (>10) at some point in the history of the CAIs. 4.4.4. Fractionation processes in the solar nebula Evaporation/condensation processes were suggested by Patchett (1980a, b) as a viable mechanism to explain the light δ 88Sr values in the CAIs. Essentially, the enrichment of the light isotopes of Sr is by condensation from a gas kinetically enriched in the lighter isotopes. First order modelling of this process, using the Rayleigh fractionation equation (e.g., Moynier et al., 2010), indicates that for elements that have vapour/solid fractionation factors close to unity, which is likely to be the case for Sr; ~63% condensation is required before any of the solids are enriched in the lighter isotope. However, this would produce large amounts of solid that have δ 88Sr of +2 to +6‰ and no material with such heavy Sr isotope compositions has been measured so far. Additionally, Mg and Si, two elements that are slightly less refractory than Sr, generally have isotopically heavy compositions in CAIs consistent with an origin by condensation from light isotope depleted gas, and are in accordance with experimental condensation studies (Richter, 2004). Therefore, simple evaporation/condensation processes are difficult to reconcile with our stable Sr isotope data and the isotopes of other moderately refractory elements such as Eu and Ca (Moynier et al., 2006; Simon and DePaolo, 2010). In order to explain the light δ 44Ca values observed in CAIs, Simon and DePaolo (2010) developed a non-equilibrium condensation model that produces kinetic mass fractionation during condensation. The model utilises recent advances in our understanding of isotopic fractionation during transport of vapour phase components to a growing condensate grain (Richter et al., 2002, 2006). Condensation of CAIs in the solar nebula is likely to take place at low pressures (b10 − 3 bar) such that the transport of chemical species in the gas phase is dependent on the mean velocity. Thus, the relative fluxes of isotopes to the surface of a condensate grain is given by αν ¼

rffiffiffiffiffiffiffi m1 m2

ð5Þ

where m1 and m2 are the true masses of an isotope of an element, where in the case of Sr, m1 = 85.9092672 and m2 = 87.9056188. At equilibrium the flux of isotopes lost from the condensate surface by evaporation is balanced by condensation, so that the ratio is determined by the equilibrium fractionation factor (αeq); at solar nebula condensation temperatures αeq will be equal to one and so no fractionation will occur. By definition, equilibrium conditions require that the partial pressure of the gas phase condensing (Pi) is equal to the equilibrium vapour saturation pressure (Pi,eq-T) for a given species. However, isotopic fractionation will occur if a non-equilibrium process occurs, such as undercooling. In this case, Pi > Pi,eq-T, and the amount of fractionation (αc) can be determined using the following equation
 α eq P i−T 2 =P i;eq−T 1 
 αc ¼
α eq =α kin P i−T 2 =P i;eq−T 1 −1 þ 1

ð6Þ

B.L.A. Charlier et al. / Earth and Planetary Science Letters 329–330 (2012) 31–40

where at solar nebula gas pressures (b10 − 3 bar) αkin = αν and T2 is the temperature that condensation occurs below the equilibrium condensation temperature T1 (Simon and DePaolo, 2010). The model that Simon and DePaolo (2010) developed utilises the fact that Ca condensation in CAIs is controlled by the formation of hibonite, and thermodynamic calculations allow the ratio Pi − T2/Pi, eq − T1 to be computed for different degrees of undercooling. Strontium is geochemically similar to Ca, forming divalent ions that are partitioned in Ca-bearing phases, and has a similar condensation temperature (Tc), with 50% Tc of 1464 versus 1517 °C for Ca (Lodders, 2003). Critically, Sr condensation is controlled by hibonite and titanite (Lodders, 2003), and Sr concentrations in CAIs are consistent with this. For example, Patchett (1980a, b) reports strontium concentrations in CAIs that are commonly over 100 ppm. In addition, ion probe determinations of Sr in hibbonite also indicate supra-chondritic concentrations (Ireland et al., 1988, 1992), and we have unpublished LA–ICP-MS data for hibonite in an Allende CAI with up to 600 ppm Sr. It is reasonable to assume that condensation of hibonite under non-equilibrium conditions such as undercooling, could produce isotopic fractionation of any Sr incorporated into the phase. Calculations using Eqs. (5) and (6) and the equations for condensing hibonite during undercooling from a nebula gas derived by Simon and DePaolo (2010) indicate that light 88Sr/86Sr can be incorporated into hibonite. Taking the lightest δ44Ca value in a CAI from Niederer and Papanastassiou (1984) of −4.2‰, we predict a corresponding δ88Sr value of~−1‰. In fact, the model calculations produce curves that are related to the δ44Ca curves of Simon and DePaolo (2010) by multiplying the δ 44Ca by a factor approximated by (1 − αν, Ca)/(1 − αν, Sr). Published data indicate that δ 88Sr values of ~ −1‰ and possibly lighter (see Section 4.4.1.) are found in CAIs, consistent with a non-equilibrium origin for the light isotope enrichment of refractory elements. It is interesting to note that modelling of Eu, another element controlled by hibonite condensation, would produce a δ153Eu value of ~−0.6‰, consistent with the data of Moynier et al. (2006). Other models have been suggested to explain the enrichment of light isotopes of refractory elements such as Sr and Eu in CAIs. Moynier et al. (2006), suggest that one way to produce a region of the solar nebula that is isotopically light in these elements is that Sr and Eu are magnetically separated within the ionised fraction of the nebula gas. This may occur because Sr and Eu have similar ionisation potentials (5.69 and 5.67 eV respectively), which are significantly lower than other refractory elements such as Mg, Fe and Si. By contrast, Ca has a slightly higher ionisation potential of 6.11 eV, while rare earth elements (REE) span values from 5.43 to 6.25. Therefore, a critical test of both non-equilibrium condensation and electromagnetic separation models would be to analyse Ca, Sr and Eu stable isotopes as well as REE patterns in the same CAI material. This would allow an assessment of chemical and physical conditions during formation of the host meteorites, and the relationship between these meteorites and the refractory element composition of the Earth. 5. Conclusions 1. We have demonstrated that 88Sr/ 86Sr ratios can be measured by MC–ICP-MS with Zr doping to a precision of 0.06‰, which is ten times smaller than the total range in 88Sr/ 86Sr discovered in extraterrestrial and terrestrial silicate samples. This indicates that stable Sr isotopes are significantly fractionated at high temperatures and their measurement by MC–ICP-MS can provide insights into planetary evolution and magmatic processes. 2. Carbonaceous chondrites are heterogeneous in composition and are composed of CAIs and chondrules (−0.36 to +0.1‰), and matrix material with heavy δ88Sr of +0.6‰, with the bulk compositions (~+0.2‰) between these components. The matrix may represent the record of a complementary reservoir to the chondrules due to redistribution of Sr isotopes during alteration. However, the

39

light δ 88Sr values of the CAIs are a primary feature which is most plausibly produced by non-equilibrium fractionation of Sr isotopes during condensation of hibonite from the solar nebula. 3. Rocks derived from the Earth's mantle are the heaviest and most homogeneous bulk-silicates we have measured with a δ 88Sr of + 0.29 ± 0.07‰ and suggest that mantle melting generates melts that are homogenous in composition. The Earth is statistically distinct from carbonaceous chondrites, but indistinguishable from Eucrites and enstatite/ordinary chondrites indicating that the origins of δ88Sr values in the Earth needs further investigation. 4. Evolved terrestrial, and basaltic lunar rocks have δ 88Sr that range to light (~−0.20‰) values. δ 88Sr co-varies with Eu/Eu*, an index of plagioclase (±sanidine) fractionation, and can be successfully modelled by the heavy isotopes of Sr being preferentially partitioned into plagioclase with a fractionation factor of ~ 1.0007 for 88 Sr/ 86Sr. 5. Higher-precision determinations of δ 88Sr by MC–ICP-MS would certainly help to further resolve and explain some of the differences between planetary bodies that we have identified in this paper. However, measurement of all three Sr isotope ratios (i.e., 88Sr/86Sr, 87 Sr/86Sr and 84Sr/ 86Sr) by high-precision double-spike TIMS methods is necessary to resolve the timing and contribution of nucleosynthetic and mass dependent inputs into the evolution of Sr in the early solar system. Acknowledgements BLAC was supported by a NERC fellowship and a Leverhulme Grant to SPK. Isotope facilities at the Open University are supported by NERC and the Open University. We thank Manuela Fehr for assistance using the Open University MC-ICP-MS instrument, and Colin Wilson for his comments on an earlier version of this manuscript. Reviews by Justin Simon and Frédéric Moynier and editorial comments by Rick Carlson considerably improved the manuscript. Appendix A. Supplementary data Supplementary data to this article can be found online at doi:10. 1016/j.epsl.2012.02.008. References Aggarwal, J.K., 2006. Mass-dependent strontium isotopic fractionation. Geochim. Cosmochim. Acta 70, A3. Andreasen, R., Sharma, M., 2007. Mixing and homogenization in the early solar system: clues from Sr, Ba, Nd, and Sm isotopes in meteorites. Astrophys. J. 665, 874–883. Bachmann, O., Dungan, M.A., Lipman, P.W., 2002. The Fish Canyon magma body, San Juan Volcanic Field, Colorado; rejuvenation and eruption of an upper crustal batholith. J. Petrol. 43, 1469–1503. Birck, J.-L., 2004. An overview of isotopic anomalies in extraterrestrial materials and their nucleosynthetic heritage. Rev. Mineral. Geochem. 55, 25–64. Blundy, J.D., Wood, B.J., 1991. Crystal–chemical controls on the partitioning of Sr and Ba between plagioclase feldspar, silicate melts, and hydrothermal solutions. Geochim. Cosmochim. Acta 55, 193–209. Borg, L.E., Gaffney, A.M., Shearer, C.K., DePaolo, D.J., Hutcheon, I.D., Owens, T.L., Ramon, E., Brenneck, G., 2009. Mechanisms for incompatible-element enrichment on the Moon deduced from the lunar basaltic meteorite Northwest Africa 032. Geochim. Cosmochim. Acta 73, 3963–3980. Charlier, B.L.A., Ginibre, C., Morgan, D., Nowell, G.M., Pearson, D.G., Davidson, J.P., Ottley, C.J., 2006. Methods for the microsampling and high-precision analysis of strontium and rubidium isotopes at single crystal scale for petrological and geochronological applications. Chem. Geol. 232, 114–133. Charlier, B.L.A., Bachmann, O., Davidson, J.P., Dungan, M.A., Morgan, D.J., 2007. The upper crustal evolution of a large silicic magma body: evidence from crystal-scale Rb–Sr isotopic heterogeneities in the Fish Canyon magmatic system, Colorado. J. Petrol. 48, 1875–1894. Crawford, M.L., 1973. Crystallization of plagioclase in mare basalts. Proc 4th Lunar Sci. Conf. Geochim. Cosmochim. Acta (suppl. 4), 705–717. Davis, A.M., Hashimoto, A., Clayton, R.N., Mayeda, T.K., 1995. Isotopic and chemical fractionation during evaporation of CaTiO3. XXVI Lunar and Planetary Science Conference, LPI, Houston, TX, p. 317. De Souza, G.F., Reynolds, B.C., Bourdon, B., 2007. Evidence for stable strontium isotope fractionation during chemical weathering. Geochim. Cosmochim. Acta 71, A220.

40

B.L.A. Charlier et al. / Earth and Planetary Science Letters 329–330 (2012) 31–40

Elliott, T., Jeffcoate, A., Bouman, C., 2004. The terrestrial Li isotope cycle: light-weight constraints on mantle convection. Earth Planet. Sci. Lett. 220, 231–245. Esat, T.M., Spear, R.H., Taylor, S.R., 1986. Isotope anomalies induced in laboratory distillation. Nature 319, 576–578. Fagan, T.J., Taylor, G.J., Keil, K., Bunch, T.E., Wittke, J.H., Korotev, R.L., Jolliff, B.L., Gillis, J.J., Haskin, L.A., Jarosewich, E., Clayton, R.N., Mayeda, T., Fernandes, V.A., Burgess, R., Turner, G., Eugster, O., Lorenzetti, S., 2001. Northwest Africa 032: product of lunar volcanism. Meteor. Planet. Sci. 37, 371–394. Fietzke, J., Eisenhauer, A., 2006. Determination of temperature-dependent stable strontium isotope (88Sr/86Sr) fractionation via bracketing standard MC–ICP-MS. Geochem. Geophys. GeosysT. 7. doi:10.1029/2006GC001243. Galer, S.J.G., 1999. Optimal double and triple spiking for high precision lead isotopic measurement. Chem. Geol. 157, 255–274. Gannoun, A., Burton, K.W., Parkinson, I.J., Alard, O., Schiano, P., Thomas, L.E., 2007. The scale and origin of the osmium isotope variations in mid-ocean ridge basalts. Earth Planet. Sci. Lett. 259, 541–556. Halicz, L., Segal, I., Fruchter, N., Stein, M., Lazar, B., 2008. Strontium stable isotopes fractionate in the soil environments? Earth Planet. Sci. Lett. 272, 406–411. Horwitz, E.P., Chiarizia, R., Deitz, M.L., 1992. A novel strontium-selective extraction chromatographic resin. Solvent Extr. Ion Exch. 10, 313–336. Ireland, T.R., Fahey, A.J., Zinner, E.K., 1988. Trace-element abundances in hibonites from the Murchison carbonaceous chondrite: constraints on high-temperature processes in the solar nebula. Geochim. Cosmochim. Acta 52, 2841–2854. Ireland, T.R., Zinner, E.K., Fahey, A.J., Esat, T.M., 1992. Evidence for distillation in the form of HAL and related hibonite inclusions. Geochim. Cosmochim. Acta 56, 2503–2520. Javoy, M., Kaminski, E., Guyot, F., Andrault, D., Sanloup, C., Moreira, M., Labrosse, S., Jambon, A., Agrinier, P., Davaille, A., Jaupart, C., 2010. The chemical composition of the Earth: enstatite chondrite models. Earth Planet. Sci. Lett. 293, 259–268. Jones, S, 1993. Mechanisms of large silicic magma chamber zonation: the Youngest Toba Tuff, Sumatra. Unpublished Ph.D. Thesis, The Open University. Krabbenhöft, A., Eisenhauer, A., Böhm, F., Vollstaedt, H., Fietzke, J., Liebetrau, V., Augustin, N., Peucker-Ehrenbrink, B., Müller, M.N., Horn, C., Hansen, B.T., Nolte, N., Wallmann, K., 2010. Constraining the marine strontium budget with natural strontium isotope fractionations (87Sr/86Sr*, δ88/86Sr) of carbonates, hydrothermal solutions and river waters. Geochim. Cosmochim. Acta 74, 4097–4109. Lodders, K., 2003. Solar system abundances and condensation temperatures of the elements. Astrophys. J. 591, 1220–1247. Luck, J.M., Othman, D.B., Albarede, F., 2005. Zn and Cu isotopic variations in chondrites and iron meteorites: early solar nebula reservoirs and parent–body processes. Geochim. Cosmochim. Acta 69, 5351–5363. Moynier, F., Bouvier, A., Blichert-Toft, J., Telouk, P., Gasperini, D., Albarède, F., 2006. Europium isotopic variations in Allende CAIs and the nature of mass-dependent fractionation in the solar nebula. Geochim. Cosmochim. Acta 70, 4287–4294. Moynier, F., Agrainer, A., Hezel, D.C., Bouvier, A., 2010. Sr stable isotope composition of Earth, the Moon, Mars, Vesta and meteorites. Earth Planet. Sci. Lett. 300, 359–366. Niederer, F.R., Papanastassiou, D.A., 1984. Ca isotopes in refractory inclusions. Geochim. Cosmochim. Acta 48, 1279–1293. Ohno, T., Hirata, T., 2006. Stable isotope geochemistry of strontium using MC–ICP-MS. Geochim. Cosmochim. Acta 70, A453. Papanastassiou, D.A., Wasserburg, G.J., 1978. Strontium isotopic anomalies in the Allende meteorite. Geophys. Res. Lett. 5, 361–376. Parkinson, I.J., Hammond, S.J., James, R.H., Rogers, N.W., 2007. High-temperature lithium isotope fractionation: insights from lithium isotope diffusion in magmatic systems. Earth Planet. Sci. Lett. 257, 609–621.

Parkinson, I.J., Burton, K.W., Mokadem, F., 2008. High-precision measurement of radiogenic and stable Sr isotopes: applications to the marine Sr record. EOS, Trans. AGU 89 (53) Fall Meet. Suppl. Abstract B14B-08. Patchett, P.J., 1980a. Sr isotopic fractionation in Ca–Al inclusions from the Allende meteorite. Nature 283, 438–441. Patchett, P.J., 1980b. Sr isotopic fractionation in Allende chondrules: a reflection of solar nebular processes. Earth Planet. Sci. Lett. 50, 181–188. Potts, P.J., Tindle, A.G., Webb, P.C., 1992. Geochemical Reference Materials. CRC Press, p. 313. Richter, F.M., 2004. Timescales determining the degree of kinetic isotope fractionation by evaporation and condensation. Geochim. Cosmochim. Acta 68, 4971–4992. Richter, F.M., Davis, A.M., Ebel, D.S., Hashimoto, A., 2002. Elemental and isotopic fractionation of type B calcium-, aluminium-rich inclusions: experiments, theoretical considerations, and constraints on their thermal evolution. Geochim. Cosmochim. Acta 66, 521–540. Richter, F.M., Davis, A.M., DePaolo, D.J., Watson, E.B., 2003. Isotope fractionation by chemical diffusion between molten basalt and rhyolite. Geochim. Cosmochim. Acta 67, 3905–3923. Richter, F.M., Mendybaev, R.A., Christensen, J.N., Hutcheon, I.D., Williams, R.W., Sturchio, N.C., Beloso Jr., A.D., 2006. Kinetic isotopic fractionation during diffusion of ionic species in water. Geochim. Cosmochim. Acta 70, 277–289. Rüggeberg, A., Fietzke, J., Liebetrau, V., Eisenhauer, A., Dullo, W.-C., Freiwald, A., 2008. Stable strontium isotopes (δ88/86Sr) in cold-water corals—a new proxy for reconstruction of intermediate ocean water temperatures. Earth Planet. Sci. Lett. 269, 570–575. Sahoo, S.K., Masuda, A., 1997. Precise measurement of zirconium isotopes by thermal ionisation mass spectrometry. Chem. Geol. 141, 117–126. Schauble, E.A., 2004. Applying stable isotope fractionation theory to new systems. Rev. Mineral. Geochem. 55, 65–111. Schnetzler, C.C., Philpotts, J.A., Nava, D.F., Schuhmann, S., Thomas, H.H., 1972. Geochemistry of Apollo 15 Basalt 15555 and Soil 15531. Science 175, 426–428. Shahar, A., Young, E.D., 2007. Astrophysics of CAI formation as revealed by silicon isotope LA–MC–ICP-MS of an igneous CAI. Earth Planet. Sci. Lett. 257, 497–510. Simon, J.I., DePaolo, D.J., 2010. Stable calcium isotopic composition of meteorites and rocky planets. Earth Planet. Sci. Lett. 289, 457–466. Sutton, A.N., 1995. Evolution of a large silicic magmatic system: Taupo volcanic centre, New Zealand. PhD thesis, The Open University, Milton Keynes, UK, 416pp. Wang, J., Davis, A.M., Clayton, R.N., Mayeda, T.K., Hashimoto, A., 2001. Chemical and isotopic fractionation during the evaporation of the FeO–MgO–SiO2–CaO–Al2O3– TiO2 rare earth elements melt system. Geochim. Cosmochim. Acta 65, 479–494. Wilson, C.J.N., Blake, S., Charlier, B.L.A., Sutton, A.N., 2006. The 26.5 ka Oruanui eruption, Taupo volcano, New Zealand: development, characteristics and evacuation of a large rhyolitic magma body. J. Petrol. 47, 35–69. Wombacher, F., Rehkamper, M., Mezger, K., Bischoff, A., Munker, C., 2008. Cadmium stable isotope cosmochemistry. Geochim. Cosmochim. Acta 72, 646–667. Zeigler, R.A., Korotev, R.L., Irving, A.J., Jolliff, B.L., Kuehner, S.M., Hupé, A.C., 2006. Petrography and composition of lunar basaltic meteorite NWA 3160. Lunar Planet. Sci. Conf. XXXVII, Abstract No. 1804. Lunar Planetary Institute, Houston (CD-ROM).