high temperature silicon isotope geochemistry

Lithos 190–191 (2014) 500–519

Contents lists available at ScienceDirect

Lithos

j ourna l homepage: www.e lsev ie r .com/ locate / l i thos

Invited review article

High temperature silicon isotope geochemistry

Paul S. Savage a,b,⁎, Rosalind M.G. Armytage c, R. Bastian Georg d, Alex N. Halliday e

a Department of Earth and Planetary Sciences, Washington University in St. Louis, One Brookings Drive, St. Louis, MO 63130, USAb Department of Earth Sciences, Durham University, Science Labs, Durham DH1 3LE, UKc Department of Earth and Atmospheric Sciences, University of Houston, 312 Science and Research 1, Houston, TX 77204, USAd Water Quality Centre, Trent University, 1600 West Bank Drive, Peterborough, Ontario K9J 7B8, Canadae Department of Earth Sciences, University of Oxford, South Parks Road, Oxford OX1 3AN, UK

⁎ Corresponding author. Tel.: +1 314 935 5619; fax: +E-mail address: [email protected] (P.S. Savage).

0024-4937/$ – see front matter © 2014 Elsevier B.V. All rihttp://dx.doi.org/10.1016/j.lithos.2014.01.003

a b s t r a c t
a r t i c l e i n f o
Article history:Received 17 October 2013Accepted 15 January 2014Available online 23 January 2014

Keywords:Silicon isotopesBulk silicate EarthIgneous processesSi in Earth's coreMC-ICP-MS

Silicon (Si) is the defining element of silicate reservoirs yet, despite its dominance inmajor Earth processes, there isstill no clear understanding of howmuch is hosted in Earth's core, how the enriched continental crust forms or evenif the crust is isotopically different from themantle because of a long history ofweathering, erosion and subduction.With the advent of multiple collector inductively coupled plasma mass spectrometry it has become relativelystraightforward to explore small (100 ppm level) mass dependent variations in Si isotopic composition resultingfrom high temperature fractionation and to develop new isotopic fingerprints of magmatic processes and sourceregions. This paper reviews the technique developments, the new data and the veracity of current interpretations.Only a small Si isotopic effect is associated with basalt formation via mantle melting. However, there now iscompelling evidence, based on a considerable number of samples (N100), that the silicate Earth is isotopicallyfractionated by 100–200 ppm per amu to a heavy composition relative to that of chondrites and also all differen-tiated stony meteorites. This could reflect variability in the circumstellar disc, but this is not well supportedby data for enstatite chondrites which are isotopically light. The most plausible current explanation is that Si isa light element in Earth's core and that differences in the bond stiffness between silicate- and metal-hosted Siresulted in substantial fractionation. Interestingly, the Moon has the same Si isotope composition as Earth'smantle, which is hard to explain unless the Moon's atoms were mainly derived from Earth.Differentiated magmatic sequences such as those of Hekla, Iceland display a systematic relationship betweenisotopic composition and Si content. More complex magmatic suites, such as I- and S-type granites, reveal arange of isotopic compositions not well correlated with chemical composition. Similar effects are found inlower crustal granulite facies xenoliths. Nonetheless the overall composition of the continental crust is onlyslightly heavy relative to the mantle; in fact the effect is barely resolvable. Therefore, the amounts of surfacedissolved heavy Si removed to the mantle via weathering over time have been small or these losses have beenbalanced by subduction of isotopically light clay.

© 2014 Elsevier B.V. All rights reserved.

Contents

1. Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5011.1. Geochemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 501

2. Nomenclature and analytical techniques . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5012.1. Nomenclature and theory . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5012.2. Analytical techniques . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 502

2.2.1. Background . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5022.2.2. Alkali fusion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5022.2.3. High resolution MC-ICP-MS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5022.2.4. In-situ Si isotope analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 503

2.3. External standards . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5033. The silicon isotope composition of the bulk silicate earth . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 504

3.1. Background . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 504

1 314 935 7361.

ghts reserved.

http://crossmark.crossref.org/dialog/?doi=10.1016/j.lithos.2014.01.003&domain=pdf

http://dx.doi.org/10.1016/j.lithos.2014.01.003

mailto:[email protected]

http://dx.doi.org/10.1016/j.lithos.2014.01.003

http://www.sciencedirect.com/science/journal/00244937

501P.S. Savage et al. / Lithos 190–191 (2014) 500–519

3.2. The silicon isotope composition of BSE as estimated using MC-ICP-MS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5043.3. A robust estimate for the Si isotope composition of BSE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 505

4. Silicon isotopes and Earth's accretion and core formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5054.1. Background . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5054.2. The silicon isotopic composition of the BSE is heavy compared to that of meteorites . . . . . . . . . . . . . . . . . . . . . . . . . . 5064.3. The silicon isotopic composition of the BSE was fractionated by core formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5074.4. Estimating δ30Sichondrite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5074.5. Silicon in the Earth's core . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5084.6. Silicon isotopes and the formation of the Moon . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 510

5. The behaviour of Si isotopes during magmatic differentiation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5115.1. Background . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5115.2. Silicon isotope fractionation in rocks from Hekla volcano, Iceland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5115.3. The “igneous array” for silicon isotopes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5125.4. Intermineral Si isotope fractionation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5125.5. The possibility of high temperature kinetic silicon isotope fractionation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 513

6. Silicon isotopes and the formation and composition of continental crust . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5146.1. Background . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5146.2. The “igneous” composition of the continental crust . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5146.3. The silicon isotopic composition of granites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5146.4. The silicon isotope composition of the upper continental crust . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 516

7. Summary and future work . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 516Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 517References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 517

1. Introduction

The field of silicon (Si) isotope geochemistry has grown rapidly overthe past decade following the development of multiple collector induc-tively coupled plasma mass spectrometry (MC-ICP-MS; Walder andFreedman, 1992). This facilitated high precision isotopic measurementsof many elements in natural samples (e.g. Johnson et al., 2004). Withthe associated improvements in precision it also became possible toexplore minor isotopic fractionations in nature such as those typicallyproduced at small but measurable levels in high temperature processes,like crystal-silicatemelt differentiation. In this respect it is hard to imag-ine any element more fundamental to understanding Earth processesand therefore more suitable for stable isotope studies than Si.

1.1. Geochemistry

Silicon (atomic weight = 28.08553) is the third most abundantelement on Earth (~16.1 wt.% Si; McDonough, 2003). It is lithophile,and behaves moderately incompatibly during mantle melting, suchthat both the oceanic and continental crust are more enriched in Sithan the mantle. Silicon has one valence state (Si4+) and does notform volatile compounds readily in terrestrial systems. In the silicateEarth, native Si is uncommon; instead Si is almost ubiquitously bondedto O in the form of the SiO44− tetrahedron. As well as being a major cat-ion in the silicate Earth, Si is important in the biosphere, diatomsrepresenting almost half of the ocean's primary productivity (Nelsonet al., 1995). It is also a major product of erosion of the continents,which feeds the oceanic biosphere. Therefore, it is not surprising thatthe earliest applications of Si isotopes, which were within reach of lowprecisionmethods, were focused on the Earth's supergene environment(c.f. De la Rocha, 2006; Ziegler et al., 2005a, 2005b).

Silicon has three stable isotopes with the following abundances:28Si: 92.23%; 29Si: 4.68%; and 30Si: 3.09%. Relatively large mass differ-ences of 3.5% and 7% between the isotopes imply the possibility of mea-surable stable isotope fractionation resulting from low temperaturechemical and biological processes. Per mil (1‰ = 0.1%) level isotopicvariations have therefore long been known to occur as a result of lowtemperature processes and have been reported in:

• chemical weathering and formation of terrigenous sediment(e.g. Georg et al., 2006a; Georg et al., 2009; Pogge von Strandmannet al., 2012; Ziegler et al., 2005a, 2005b)

• precipitation of secondary silicate minerals such as clays and lowtemperature silica (Basile-Doelsch et al., 2005; Opfergelt et al., 2011,2012)

• biological uptake of Si from seawater (Cardinal et al., 2007; De LaRocha, 2003, 2006; De La Rocha et al., 1998; Egan et al., 2012;Hendry et al., 2010; Reynolds et al., 2006).

However, MC-ICP-MS renders it possible to also explore smallerhigh temperature isotopic fractionations of Si, which is a subject thathas really only burgeoned in the past 10 years.

Silicon isotope fractionation, like that of any other element,is relatively limited at high temperatures (see O'Neil, 1986 andreferences therein). The potential for isotopic fractionation is furtherlimited by silicon's low volatility, single valence state and invariantbonding environment. Early Si isotope studies, utilising gas sourcemass spectrometry (GS-MS), struggled to measure any high-temperature Si isotope fractionation outside of analytical error (approx-imately ±0.4‰; Douthitt, 1982; Reynolds and Verhoogen, 1953). Theadvent and development of MC-ICP-MS (Walder and Freedman,1992) and in particular high resolution MC-ICP-MS (Halliday et al.,2011) facilitated high precision determination of Si isotopic composi-tions. However, this was not enough; novel sample preparation andintroduction techniques (Georg et al., 2006b; see Section 2.2) werealso necessary. The result has been improved analytical precisionfor natural samples to levels at which high-temperature isotopicvariations of ≤0.1‰ in Si extracted from silicate matrices can bemeasured, opening up, for the first time, a new area for geochemicalinvestigation.

The results so far have already afforded important new insightsinto the origin of the Earth, its core and Moon, as well as identifyinga potentially powerful new tracer of magmatic fractionation andmetasedimentary contributions to magma sources. The last review ofSi isotopes was made over 15 years ago, by Ding et al. (1996). As such,it is timely that we here summarise these latest breakthroughs inhigh-temperature Si isotope variations in particular.

2. Nomenclature and analytical techniques

2.1. Nomenclature and theory

As with the vast majority of stable isotope systems, mass dependantSi isotope variations are reported using delta notation, defined as the

502 P.S. Savage et al. / Lithos 190–191 (2014) 500–519

deviation of a sample's isotopic ratio from that of a standard (in per mil,‰), as such:

δ30Si ¼30Si=28Sisample30Si=28Sistandard

−1

!� 1000; δ29Si ¼

29Si=28Sisample29Si=28Sistandard

−1

!� 1000

ð2:1Þ

The standard reference used for reporting Si isotopic data is the silicasand reference material RM 8546, distributed by the National Institute ofStandards and Technology (NIST). For historical reasons, the equivalentname used bymost of the literature for this standard is NBS28; this prac-tice is continued throughout this work. Before NBS28 was established,the Caltech Rose Quartz Standard (RQS) was used (e.g. Douthitt, 1982).Unfortunately, the isotopic composition of RQS is poorly constrainedagainst NBS28 and so quantitative comparison of older data is not ideal,although the most recent estimate indicates that both standards haveapproximately equal isotopic compositions (Georg et al., 2007).

In terrestrial samples it is valid to assume that isotopic fractionation ofsilicon's three stable isotopes will be a mass-dependent process, hencethe extent of fractionation of the 30Si/28Si ratio will be greater than thatof the 29Si/28Si purely as a function of the mass difference of the ratios.As such, the δ30Si value will be ~2 times that of the δ29Si value, andleads to Si isotopic data being commonly represented using only δ30Si.Mass dependence can be described more accurately in terms of the frac-tionation factors:

α29=28 ¼ α30=28

� �β ð2:2Þ

whereαx/28 is defined as the ratio of the Si isotopic ratios of two phases Aand B (where x = 29 or 30), as such:

αx=28 ¼xSi=28SiAxSi=28SiB

¼ 1000þ δxSiA1000þ δxSiB

: ð2:3Þ

Here, the exponent β is defined in terms of the atomic masses ofeach isotope and with differences for kinetic and equilibrium fraction-ation (Young et al., 2002). On a three isotope plot (a plot of δ30Si vs.δ29Si), mass dependent isotopic data will fall on a straight line with aslope of βeq or βkin, depending on whether fractionation is a result ofequilibrium or kinetic processes. For Si, equilibrium isotope fraction-ation, βeq = 0.5178 and for kinetic fractionation, βkin = 0.5092. Theseslopes are almost identical for the range of Si isotopic fractionationassociated with high temperature processes, such that current analyticalprecision is insufficient to distinguish between the two.

2.2. Analytical techniques

2.2.1. BackgroundThe isotopic analysis of Si is not a novel area of research, and a few

pioneering studies were made as early as the 1950s (e.g., Reynoldsand Verhoogen, 1953). These studies utilised gas source mass spec-trometry (GS-MS),whereby the Si is analysed by themass spectrometerin the form of the volatile SiF3+ species. Up until the end of the 20th cen-tury, analysis by GS-MS techniques had generatedmost of the Si isotopedata, from the lunar studies in the 1970s (e.g. Epstein and Taylor, 1970)to the biogenic silica studies of De La Rocha (2003), including three ofthe most cited works in the field (Ding et al., 1996; Douthitt, 1982;Molini-Velsko et al., 1986). The precision in these studies was on theorder of ±0.15 to ±0.50‰, which is insufficient to resolve muchhigh temperature Si isotopic variation. Recent method improvements(e.g. Brzezinski et al., 2006) have improved precision to better than±0.15‰ using GS-MS, however, these still involve the use of extremelyhazardous, fluorine-based, chemicals. The advent of MC-ICP-MS and itsapplication to the Si isotope systemmeans that, nowadays, the analysis

of Si isotopes is semi-routine and can be made at much improved de-grees of precision.

The theory of sample preparation and general application of MC-ICP-MS to stable isotope analysis is eloquently described elsewhere (Albarèdeand Beard, 2004; Halliday et al., 2011; Weyer and Schwieters, 2003) andis beyond the scope of this review. Instead we will briefly describe thespecific methods typically involved in solution MC-ICP-MS Si isotopeanalysis and concentrate on two important aspects of the analytical tech-niques that havemade high precision Si isotopemeasurement of igneousmaterial possible: the sample digestion procedure and high-resolutionMC-ICP-MS.

2.2.2. Alkali fusionTypically, digestion of silicate material utilises hydrofluoric acid

(HF), which readily attacks silicates and produces the volatile com-pound SiF4. This means that Si can easily be lost from the sample,which can induce artificial isotopic fractionation (Georg et al., 2006b).The other disadvantage of including HF during sample processing isthat it can suppress the ionisation efficiency of Si in the inductivelycoupled plasma source. However, HF can be avoided by dissolvingsilicates in a strong alkaline solution, for example using an alkali fusionmethod. This solutionwasfirst proposed for Si isotope analysis byGeorget al. (2006b), though alkali fusion is not a novel technique for process-ing silicate samples (Potts, 1987). Briefly, the sample and alkali flux areplaced in a crucible at a high heat (N700 °C). At these temperatures theflux melts to produce a hyper-alkaline liquid that completely dissolvessilicate material. When this cools, it forms a water-soluble meta-stablesilicate, which can be dissolved in a weakly acidic solution.

Most Si isotope studies since have utilised such a procedure to avoidthe use of HF; for instance, Armytage et al. (2011, 2012), Fitoussi et al.(2009, 2012), Pringle et al. (2013a), Savage et al. (2010, 2011, 2012),Zambardi and Poitrasson (2011a, 2011b), and many others closelyfollow the methods described by Georg et al. (2006b). Abraham et al.(2008) use a similarmethod but with a different flux and Si purificationtechniques. Finally, Chakrabarti and Jacobsen (2010), Kempl et al.(2013), and van Den Boorn et al. (2006, 2007, 2009) all use a methodthat involves lower temperature alkali digestion in Parr bombs,but the theory is the same. Some workers do still use HF dissolution(e.g. Ziegler et al., 2010) but are careful to limit the amount of acid used.

2.2.3. High resolution MC-ICP-MSDue to the high ionisation efficiency of the Ar-based plasma on an

MC-ICP-MS instrument, a small fraction (usually b1%) ofmost elementsin a sample solution will get transferred into themass analyser throughan inefficient cone interface. There is an associated very strong (percentlevel) isotopic fractionation ormass bias generated by space charge dur-ing extraction of the ions from the plasma through the cone interface.This needs to be carefully corrected in order to get accurate data. Eachelement behaves differently in a plasma and it has been known sincebefore the advent of MC-ICP-MS that admixing other elements canaffect the size of signal (increased or decreased) from the element ofinterest (or the “analyte”). Furthermore, the presence of a matrix canalter the measured isotopic ratio, presumably by changing the spacecharge. Therefore, purification of the element of interest before analysisis necessary in order to eliminate matrix effects and so that the massbias from the standards and the sample is the same. For Si isotopes, thisis typically achieved using ion exchange chromatography, utilising arobust one-step column procedure with cation exchange resin whichresults in quantitative Si yields with complete removal of all cationicmatrix elements (Georg et al., 2006b). Subsequent to purification, Sistable isotope measurement is either performed using a simplestandard-sample bracketing protocol (e.g. Georg et al., 2007; Savageet al., 2010, 2011, 2012 etc.) or standard-sample in tandem with Mgdoping (e.g. Zambardi and Poitrasson, 2011a; Zambardi et al., 2013) tocorrect for instrumental mass bias, using NBS28 (see above) as thebracketing standard.

Table 1BHVO estimates.

δ30Si 2 s.d.a δ29Si 2 s.d.a n

BHVO-1Abraham et al. (2008) −0.33 0.17 −0.17 0.10 11Georg et al. (2009) −0.29 0.16 −0.17 0.11 87Savage et al. (2010) −0.27 0.09 −0.15 0.04 192Armytage et al. (2011) −0.30 0.15 −0.16 0.08 117Hughes et al. (2011) −0.32 0.10 −0.16 0.06 8Mean −0.30 0.05 −0.16 0.02 5

BHVO-2Abraham et al. (2008) −0.29 0.27 −0.17 0.13 8Fitoussi et al. (2009) −0.32 0.15 −0.16 0.15 14Chakrabarti and Jacobsen (2010) −0.42 0.15 −0.20 0.11 14Savage et al. (2010) −0.27 0.10 −0.14 0.05 59Zambardi and Poitrasson (2011a) −0.27 0.08 −0.14 0.05 42Armytage et al. (2011) −0.28 0.14 −0.15 0.08 223Savage et al. (2011) −0.29 0.09 −0.14 0.08 188Armytage et al. (2012) −0.28 0.14 −0.15 0.08 26Pringle et al. (2013a) −0.28 0.12 −0.15 0.09 26Savage et al. (2013a) −0.29 0.10 −0.14 0.09 44Savage and Moynier (2013) −0.28 0.09 −0.15 0.06 44Mean −0.30 0.09 −0.15 0.04 11Mean (without C&J, 2010)b −0.28 0.03 −0.15 0.02 10

n = number of sample repeats. The values in bold are the best estimates (mean andassociated precision) of the BHVO standards, calculated using available literature data.

a In the published work the uncertainties have been quoted in different ways from±1s.d. to ±2 s.e. Here ±2 s.d. has been applied uniformly for consistency.

b Here, the italics denote the best estimate for BHVO-2 calculated without using thedata from Chakrabarti and Jacobsen (2010). The significant discrepancy between theBHVO-2 value of Chakrabarti and Jacobsen (2010) and all the other studies listed aboveis discussed in the text. Note the significant increase in precision when this value is re-moved from the average.


A major problem associated with the application of MC-ICP-MS tomany isotope systems, not least Si, is that of polyatomic interferences.Even though samples are chemically purified before analysis, elementspresent in the ambient atmosphere, the sample solution and the plasmagas supply (e.g. C, O, N, Ar) become ionised and can combine to formmolecules that are, nominally, of the same mass as the isotope(s) of in-terest. Thesemolecules aremeasured as extra ion beam signal, resultingin an overestimation of certain isotopes. This can have a significantanalytical effect, especially for the minor isotope beams. Moleculessuch as 28SiH+, 12C16O+, 14N2

+ and 14N16O+ are all such examples inthe case of Si. In particular, the large interference of 14N16O+ on 30Si+

(the least abundant isotope) meant that early MC-ICP-MS studiescould only accurately measure 29Si/28Si (e.g. Cardinal et al., 2003).This problem was solved by the introduction of “high-” and “medium-”resolution instruments.

Despite being notionally similar in mass, polyatomic interferencesare almost always slightly heavier than the isotope of interest. Suchinterferences can therefore be resolved by reducing the mass/chargewindow of the ion beam for admittance to the detector at one time,thereby improving the “mass resolution” of the instrument (c.f. Weyerand Schwieters, 2003). This can be achieved, firstly, by physicallymasking the ion beams with slits that define the egress of a smallerspread in mass/charge ratio to the detector. This means that, at thefocal plane of a mass spectrometer, the interfering beams will be sepa-rated from the Si beams. Narrowing the beam reduces the sensitivityof the instrument so a compromise must be reached between massresolution and sensitivity. The Nu Instruments NuPlasma HR (and NuPlasma II) and Thermo-Finnigan Neptune (and Neptune Plus) are allsuch instruments that employ this technique. A second way of increas-ing the mass resolution further without additional loss of sensitivity isto increase the geometry of the instrument. By doubling the lengththe ion beamhas to travel, the separation of two beams can be essential-ly doubled. An example of this sort of instrument is the Nu InstrumentsNu1700. Amore comprehensive description of high-resolutionMC-ICP-MS is provided in Halliday et al. (2011), needless to say that withoutthe above-mentioned improvements in instrument mass resolution,accurate MC-ICP-MS analysis of all three Si isotopes would not befeasible.

2.2.4. In-situ Si isotope analysisSo far we have concentrated on solution MC-ICP-MS, but analysis of

Si isotopes is also achievable with in-situ techniques. Initial studiesutilised Secondary Ion Mass Spectrometry (SIMS; Heck et al., 2011;Robert and Chaussidon, 2006), but more recently laser ablation–MC-ICP-MS has been employed (e.g. Shahar and Young, 2007; Steinhoefelet al., 2011; Ziegler et al., 2010). The advantage of such techniques isthat they allow investigation of Si isotope variations at high spatialresolution, for instance in finely laminated siliceous sediments such ascherts (Robert and Chaussidon, 2006) or banded iron formations(Heck et al., 2011). They also provide a means of analysing smalldisseminated phases such as metal and silicate phases in enstatitemeteorites (Ziegler et al., 2010). The drawback with both SIMS andLA–MC-ICP-MS methods is that their associated precisions (±0.2to 0.3‰; 2 s.d.) have often not been as good as that achievable withsolution based methods. Recent improvements in laser technologyhave increased precision on Si isotopic measurements, which are nowapproaching those attainable by solution MC-ICP-MS (e.g. Steinhoefelet al., 2011), although little data pertinent to the scope of this reviewis so far available.

2.3. External standards

The analysis of well-characterised standards as a way of assessingmethod accuracy and precision is commonplace in both elemental andisotopic studies. For Si isotopes, where double-spike correction isimpossible, this technique is the usual method by which data are

demonstrated to be robust and comparable to results generated byother laboratories. This is particularly important for studies concerninghigh-temperature Si isotope variations, where the amount of fraction-ation can be close to analytical precision; a discrepancy of ±0.1‰ canhave significant scientific ramifications.

Reynolds et al. (2007) were the first to respond to the growingneed for established Si isotopic standards. Those characterised wereIRMM-018 (δ30Si =−1.65± 0.22‰; 2 s.d.; a SiO2 standard distributedby Institute for Reference Materials and Measurements, Belgium),Big Batch (δ30Si=−10.48± 0.54‰; 2 s.d.; a highly fractionated chem-ically prepared silica powder prepared at the University of California,Santa Barbara) and diatomite (δ30Si = +1.26 ± 0.20‰; 2 s.d.; chemi-cally purified natural diatomite, again prepared at UCSB). Further tothis, Abraham et al. (2008) analysed the USGS rock powders BHVO-1and BHVO-2 for Si isotopes (respectively δ30Si = −0.33 ± 0.13‰and −0.29 ± 0.27‰; 2 s.d.). In practice, the most commonly utilisedstandards from this suite are diatomite and BHVO-1/-2, and the latteris most useful for assessing total method accuracy for analysis of silicaterocks. This is because BHVO-1 and -2 are natural basaltic rock samplesand, as such, they have chemical matrices similar to the samples ofinterest.

Since the publication of Abraham et al. (2008), many Si isotope stud-ies of silicate rocks have utilised the BHVO standards; Table 1 provides alist of these measurements. It is clear from these data that the methodprecision provided by Abraham et al. (2008) of δ30Si = ±0.13 and±0.27‰ (2 s.d.) is not an accurate reflection of how well-characterisedthese standards are now. For BHVO-1, taking ameanof the 5 values avail-able in the literature gives a value of δ30Si = −0.30 ± 0.05‰ (2 s.d.).For BHVO-2, the same exercise gives δ30Si = −0.30 ± 0.09‰ (2 s.d.);this value is identical to that for BHVO-1 (this is expected, as the sam-ples are taken from the same lava flow) but the error is almost twiceas large. Inspecting the data in Table 1, there is one outlier (δ30Si =−0.42 ± 0.15‰; 2 s.d.) given by the study of Chakrabarti andJacobsen (2010). There are possible analytical issues for the offset inthe data of Chakrabarti and Jacobsen (2010), as discussed in Armytage

Table 2BSE estimates.

δ30Si 2 s.d.a δ29Si 2 s.d.a n

StudyGeorg et al. (2007)b −0.38 0.13 −0.20 0.09 12Fitoussi et al. (2009) −0.29 0.06 −0.15 0.04 10Chakrabarti and Jacobsen (2010)b −0.38 0.06 −0.19 0.02 8Savage et al. (2010) −0.29 0.08 −0.15 0.05 35Armytage et al. (2011) −0.32 0.08 −0.16 0.04 6Zambardi et al. (2013) −0.27 0.07 −0.13 0.04 19

Averagesc

Mantle average −0.30 0.09 −0.15 0.05 32−0.31 0.09 −0.15 0.05 33

MORB average −0.27 0.06 −0.14 0.03 20−0.30 0.13 −0.16 0.07 25

OIB average −0.31 0.04 −0.16 0.05 13−0.32 0.10 −0.17 0.06 18

Other basalt sensu lato −0.28 0.04 −0.13 0.04 13−0.28 0.04 −0.13 0.03 13

BSE average −0.29 0.07 −0.15 0.05 78−0.30 0.10 −0.16 0.06 89

n = number of samples analysed. The values in bold are the best estimates (mean andassociated precision) for the Si isotope composition of bulk silicate Earth (BSE).

a In the published work the uncertainties have been quoted in different ways from±1 s.d. to ±2 s.e. Here ±2 s.d. has been applied uniformly for consistency.

b The studies of Georg et al. (2007) and Chakrabarti and Jacobsen (2010)find somewhat lighter BSE estimates compared to the other studies listed in this table —

possible reasons for this discrepancy are discussed in the text.c The data from the studies marked ‘b’ are not included in the new BSE estimate provided

by this paper; in the interests of transparency, we have recalculated the estimates withoutexcluding these studies, these are the values given in italics. See text for discussion.


et al. (2011). Removing their data from the BHVO-2mean gives a valueof δ30Si=−0.28± 0.03‰ (2 s.d.); the error is now three times smallerand the value is still statistically identical to that of BHVO-1. Given howwell-characterised these samples are now, we would recommend theuse of the means provided in Table 1 as new literature values for bothstandards.

3. The silicon isotope composition of the bulk silicate earth

3.1. Background

Over 90% of Earth's Si is in the mantle. All of the Si present in newlyformed oceanic and continental crust is derived from this reservoir;thus, the Si that ultimately enters the hydrological and biological cyclesas a result of weathering is alsomantle-derived. Therefore, providing anaverage Si isotope composition for the Bulk Silicate Earth (BSE) isimportant for establishing a global terrestrial baseline, to which allother Si reservoirs, from meteorites to terrigenous sediment, can becompared. This section details the previous attempts to estimate thevalue for the BSE and then recent data are combined to provide a robustestimate.

The first estimate for BSE was by Douthitt (1982) whose averageδ30Si value of−0.4± 0.3‰ (2 s.d.) was calculated using predominantlyisland arc basalt (IAB) data from the Marianas. This was reappraised byDing et al. (1996) who gave a lighter, but within uncertainty identical,estimate of−0.61± 0.55‰. Another dataset that hinted at the compo-sition of BSE was that of Ziegler et al. (2005a), which provided a set ofanalyses from an un-weathered basaltic flow from Hawaii (δ30Si =−0.5 ± 0.3‰; 2 s.d.). Thus, until 2007, all that could confidently besaid about the Si isotopic composition of BSE was that it probably laysomewhere between δ30Si = −1.0 and −0.1‰. This is a wide rangeof uncertainty given the likely limits of high temperature Si isotopicfractionation and it was unclear whether this reflected the lower preci-sion of these data or whether widespreadmantle isotope heterogeneityexisted.

3.2. The silicon isotope composition of BSE as estimated using MC-ICP-MS

The first “modern” estimate for BSE, usingMC-ICP-MS instrumenta-tion, was that of Georg et al. (2007). This paper noted that primitivemeteorites were, on average, resolvably lighter than terrestrial rockswith respect to Si isotopes. The significance of this offset is discussedin Section 4; what is important here is to note that their BSE estimate,of δ30SiBSE = −0.38 ± 0.13‰ (2 s.d.; n = 12; Table 2) is more precisebut almost identical to Douthitt's (1982) value. A notable aspect ofthis dataset is that there still appeared to be resolvable Si isotope varia-tions between individual ultramafic andmafic rocks. Given their limitedterrestrial sample set, however, it was unclear whether the 0.20‰ scat-ter in the Georg et al. (2007) data was analytical or represented mantleheterogeneity.

A subsequent paper by Fitoussi et al. (2009) provided a heavier andmore precise δ30SiBSE value of δ30Si = −0.29 ± 0.06‰ (2 s.d.; n = 10;Table 2). This work saw no evidence for resolvable isotopic differencesbetween mantle derived samples. The possible difference betweenthese two more “modern” BSE estimates (~0.1‰) appears small, butsuch an offset can have implications in the field of high temperatureSi isotope geochemistry, as shall be discussed later. For a brief time,the exact Si isotopic composition of BSE became a hotly debated issue.

Savage et al. (2010) went someway to resolving this with a detailedstudy of many more (35) samples from varied tectonic and geologicalsettings. These included various mantle rocks (spinel and garnetlherzolite, harzburgite, dunite, websterite), sampled as both xenolithsand obducted terranes. Also, a number of mid-ocean ridge basalts(MORB) were analysed. These samples came from all major ocean ba-sins, thereby providing an opportunity to test for possible geographicalbias, and alsowhether mantle melting fractionates Si isotopes. Finally, a

set of island arc basalts (IAB) and basaltic andesites was analysed, fromthe Marianas and South Sandwich arcs, where variable sedimentaryinput could potentially alter the Si isotopic composition of the subarcmantle.

The group data averages from Savage et al. (2010) were as follows:δ30Siultramafic = −0.33 ± 0.08‰, δ30SiMORB = −0.27 ± 0.06‰ andδ30SiIAB=−0.28± 0.06‰ (2 s.d.). All these averages are similar, imply-ing that there currently is no resolvable Si isotopic difference betweenultramafic rocks and basaltic lavas; as such, no significant fractionationoccurs during mantle melting (Fig. 1). There is also no evidence forsignificant isotopic variations related to geographical location, ageor chemistry. From their analyses, Savage et al. (2010) estimated theSi isotopic composition of the BSE to be δ30Si = −0.29 ± 0.08‰ (2s.d.; Table 2), very similar to the previous estimate of Fitoussi et al.(2009).

A number of subsequent Si isotope studies have provided their ownestimates for the BSE, and these are listed in Table 2. Although theseoverlap within uncertainty, they fall into two δ30SiBSE groups, oneclose to −0.4‰ (Chakrabarti and Jacobsen, 2010; Georg et al., 2007),the other nearer −0.3‰ (Armytage et al., 2011; Fitoussi et al., 2009;Savage et al., 2010; Zambardi et al., 2013). One reason for the disparitycould be the choice of samples, in particular the use of mineral separateanalyses by Chakrabarti and Jacobsen (2010) and Georg et al. (2007), incalculating the BSE value. However, for both studies, removing themineral separate analyses from their calculations still results in relative-ly light estimates. Therefore, analytical issues, rather than sample selec-tion, are more likely to be biassing the data of these two studies. It hasalready been noted that the work of Chakrabarti and Jacobsen (2010)showed a −0.1‰ offset between their BHVO-2 value and other litera-ture values (Table 1). Recalculating their data using this offset gives aBSE estimate in line with the rest of the studies in Table 2. For thestudy of Georg et al. (2007), much of the data (both terrestrial andextra-terrestrial) appear to be offset to lighter isotopic values (by~0.1‰). Reanalysis of the same samples (Savage et al., 2010) and eventhe same sample aliquots (Armytage et al., 2011) gave systematicallyheavier values than found by Georg et al. (2007). Therefore, it appearsthat there may have been some bracketing standard issue or systematic

Fig. 1. Histograms (top panel) and probability density functions (bottom panel) of the var-ious ultramafic andmafic datasets used to calculate the new estimate for Bulk Silicate Earth(BSE) in this study. New BSE estimate plotted as dashed line, with grey box representing the2 s.d. uncertainty on the estimate (s.l. = sensu lato). Data are from Abraham et al. (2008),Armytage et al. (2011), Fitoussi et al. (2009), Savage et al. (2010, 2011, 2013a), Savage andMoynier (2013) and Zambardi et al. (2013).


error affecting the data of this first MC-ICP-MS study. Nevertheless, allthe estimates listed in Table 2 are within error; that is, at 2 standarddeviations, all BSE estimates are equivalent.

3.3. A robust estimate for the Si isotope composition of BSE

Reliable data from the literature forwhole rock ultramafic or basaltic(sensu lato) samples from the studies of Abraham et al. (2008),Armytage et al. (2011, 2012), Fitoussi et al. (2009), Savage et al.(2010, 2011, 2013a), Savage and Moynier (2013) and Zambardi et al.(2013) have here been combined to derive a new best estimate forδ30SiBSE. Repeat analyses of the same sample (i.e. BHVO-2, BCR-1,San Carlos) from separate studies were averaged before the data wereused, so as not to bias the overall mean. In addition to ultramafic andMORB groups, there is now a collection of ocean island basalt (OIB)data, taken from Iceland (Savage et al., 2011; Zambardi et al., 2013),the Cameroon line (Armytage et al., 2011) and Hawaii (Fitoussi et al.,2009; BHVO-1 and -2). As ocean island volcanoes potentially samplethe lower mantle, further constraints on Si isotope mantle heterogene-ity can thus be made. A final difference from the study of Savage et al.(2010) is that island arc samples have been subsumed into an “otherbasalt” group, which includes data for continental flood basalts andrift zone basalts, along with various hypabyssal and intrusive maficlithologies (Savage et al., 2010, 2011, 2013a; Zambardi et al., 2013).

The data for each group are plotted in Fig. 1, as well as thenormal distributions of each data set. The averages of each group areas follows: δ30Siultramafic = −0.30 ± 0.09‰, δ30SiMORB = −0.27 ±0.06‰, δ30SiOIB = −0.31 ± 0.04‰, δ30Siother basalt = −0.28 ± 0.04‰(2 s.d.; Table 2). As is shown in Fig. 1 and Table 2, there is no resolvableSi isotopic difference between any of the groups, despite the samplesbeing taken from many various and varied tectonic settings. Thesedata are thus used to calculate a new Si isotopic composition for BSE,of δ30SiBSE =−0.29± 0.07‰ (2 s.d.; n= 78). This value is representedin Fig. 1 as the dashed vertical line, with the precision shown by the greyband; except for some ultramafic samples, all data plot within the band;the outliers are typically melt-depleted (dunite, websterite) and maynot be representative of mantle compositions. The new δ30SiBSE is iden-tical to that of Armytage et al. (2011), Fitoussi et al. (2009), Savage et al.(2010), and Zambardi et al. (2013), but with slightly improved preci-sion. Data from Chakrabarti and Jacobsen (2010) and Georg et al.(2007) were not used in this first estimate, given their possible analyt-ical discrepancies, as detailed above. A separate calculation using thesedata has beenmade; these are the values in italics in Table 2. Using suit-able data from these two studies does not significantly alter the aver-ages of each group, due in part to the fact that a large proportion oftheir isotope data was from mineral separates and so was not utilised.As such, the BSE average value is also not significantly different(δ30Si = −0.30 ± 0.10‰, 2 s.d.), but the precision is slightly poorer.Given this loss in precision, we recommend the use of δ30SiBSE =−0.29 ±0.07‰ (2 s.d.).

4. Silicon isotopes and Earth's accretion and core formation

4.1. Background

Improving our understanding of the origin and earliest history of theEarth has long been a major goal of isotope geochemistry. Of particularinterest are the nature of terrestrial core formation and the origin of theMoon. Silicon isotope measurements have proved to be immenselyvaluable in both these areas.

Earth's accretion and core formation are believed to have beenprotracted, with the bulk of the planet formed during energetic colli-sions with other proto-planetary objects (Raymond et al., 2009). Thisaddition of mass and energy provides heat to generate metallic ironand facilitate further core growth such that a fraction of the core mayhave formed by core–coremerging rather than equilibrium segregation(Halliday, 2004). It has been known for more than 50 years, since thework of Francis Birch (1964), that the density of the outer core is toolow for it to be made only of an iron–nickel alloy; some other lighterelement(s) must also be present. Since then a great deal of effort, notonly in the area of experimental petrology (e.g. Corgne et al., 2008;Gessmann et al., 2001; Wade and Wood, 2005), but also in bulk Earthelemental composition estimation (e.g. Allègre et al., 1995; Drake andRighter, 2002; McDonough, 2003), has been expended on investigatingwhat those light elements could be andwhat their concentrationsmighttell us about the conditions under which Earth first differentiated. Asexplained below, silicon isotopes have been insightful in this regard.

The formation of the Moon is thought to be the result of the finalmajor stage in Earth's accretion following a collision with another plan-et (Canup, 2008) sometimes called “Theia” (Halliday, 2000). Evidence tosupport the Giant Impact origin of the Moon (and indeed the overallaccretion of Earth via such oligarchic growth processes) would be isoto-pic fractionation in semi refractory elements that would only boil atextremely high temperatures. Poitrasson et al. (2004) first argued forsuch an effect in Fe isotopes. Silicon isotopes have again been decisivein this area. Even more significantly, high precision Si isotopes forlunar samples have been able to shed new light on the nature of theGiant Impact itself and provide powerful constraints that are hard toreconcile with some current dynamic simulations.

-0.80 -0.70 -0.60 -0.50 -0.40 -0.30 -0.20

Ziegler et al. (2010)HF dissolutionNeptunem/Δm =12000

Georg et al. (2007)Ag crucible NaOH fusionNu1700Mass resolution 2000

Armytage et al. (2011)Ag crucible NaOH fusionNuPlasmam/Δm =3500

Fitoussi et al. (2009)Ag crucible NaOH fusionNu1700Mass resolution 2000

Chakrabarti and Jacobsen (2010)Teflon bomb NaOH fusionGV Isoprobem/Δm =5000

Fig. 2. The range of published Si isotope compositions of CM2 carbonaceous chondrite“Murchison”, illustrating the possible isotope heterogeneities within undifferentiatedmeteorite groups and also the possible analytical discrepancies related to various tech-niques and instrumentation (see text for discussion, s.l. = sensu lato). Error barsare 2 s.e. (standard error of themean), the vertical line and associated grey box representsour best estimate and 2 s.d. uncertainty for the average composition of bulk Earth, definedby ordinary and carbonaceous chondrites (Section 4.4).


4.2. The silicon isotopic composition of the BSE is heavy compared to thatof meteorites

Georg et al. (2007) measured an offset between the silicon isotopiccomposition of BSE and allmeteorites, including carbonaceous, ordinaryand enstatite chondrites, martian meteorites, howardites, eucrites anddiogenites (HED, thought to come from asteroid 4-Vesta) and ureilites.There were several Si isotope studies published on meteorites beforethis important finding (e.g. Molini-Velsko et al., 1986) but none weresufficiently precise to resolve a BSE-meteorite difference. Since thework of Georg et al. (2007), subsequent studies have not all agreedabout the exact composition of meteorites and the magnitude of thedifference relative to the BSE. A striking example is that of the carbona-ceous chondrite Murchison, which has been analysed by more groupsthan any other meteorite, but yields very different results betweengroups (Fig. 2); such variability could be a result of different analyticalsetups. It could, also, be explained by Si isotopic heterogeneity withindifferent components of carbonaceous chondrites (see Hezel et al.,2010, and references therein).

Table 3Comparison of recent δ30Si studies.

Meteorite average(‰),±2 s.d.a

n

Georg et al. (2007) −0.58 ± 0.11 22Fitoussi et al. (2009) −0.41 ± 0.12 9Fitoussi and Bourdon (2012)Chakrabarti and Jacobsen (2010) −0.42 ± 0.13 10Armytage et al. (2011) −0.48 ± 0.13 36Armytage et al. (2012)Savage and Moynier (2013) −0.49 ± 0.08 5

Zambardi et al. (2013) −0.46 ± 0.05 9

n = number of samples analysed.a In the published work the uncertainties have been quoted in different ways from ±1 s.d.b The studies of Georg et al., 2007 and Fitoussi et al. (2009) only contain lunar basalts, Chakr

highland samples (ferroan anorthosites), picritic glasses and lunar basalts and Zambardi et al.c As only one sample was measured, the error represents ±2 s.e. for the sample, rather thand Weighted averages for the two different "types" of enstatite chondrites, EH and EL.

All groups agree, however, that enstatite chondrites have the lightestδ30Si of all the chondrites (Fitoussi and Bourdon, 2012; Savage andMoynier, 2013; see Table 3, Fig. 3) and all workers (Armytage et al.,2011; Fitoussi et al., 2009; Georg et al., 2007; Zambardi et al., 2013)also find some difference between the BSE and meteorites, with theexception of Chakrabarti and Jacobsen (2010). It is still unclear as towhy Chakrabarti and Jacobsen (2010) did not measure an offset but,as has already mentioned, their BSE estimate is lighter and may be bi-assed by analytical issues. Fitoussi et al. (2009)was the only study to re-port differences between ordinary and carbonaceous chondrites butsubsequent studies with many more samples (Armytage et al., 2011)have been unable to reproduce this (Fig. 3).

The solar system is isotopically heterogeneous with respect to anumber of elements, such as O, Cr, Ti and Ni (Clayton, 1993; Regelouset al., 2008; Trinquier et al., 2007, 2009). Therefore, the lack of composi-tional parity between the Si isotopic compositions of chondrites andBSE might relate to such effects. However, this is unlikely for a numberof reasons. Firstly, Earth's composition can be explained as a resultof mixing of different proportions of chondritic components with di-verse O, Cr and Ti compositions (e.g., Lodders, 2000). In the case of Sihowever, there are no plausible solar system reservoirs with such aheavy composition as measured for BSE (some calcium aluminiumrefractory inclusions are extremely heavy in Si; Davis et al., 1990;however, chemically, it makes little sense that bulk Earth is madefrom such refractory material; Wänke and Dreibus, 1988). Theclosest chondritic isotopic matches to Earth's O, Cr and Ti are enstatiteand ordinary chondrites (Clayton, 1993; Trinquier et al., 2007, 2009;Zhang et al., 2012). Ordinary chondrites are like other meteorites interms of Si isotopes and enstatite chondrites are isotopically verylight, not heavy (Fitoussi and Bourdon, 2012; Savage and Moynier,2013; Ziegler et al., 2010).

Furthermore, solar system O, Cr, Ti and Ni isotope heterogeneities areoften related to mass-independent processes, and are measurable on abulk meteorite scale. However, Pringle et al. (2013b) show that thereare no measureable mass-independent Si isotope anomalies (withincurrent method precision) in a wide variety of extra-terrestrial bulk me-teorite samples, when analysed relative to a terrestrial standard. This im-plies that the solar system was well-mixed with respect to Si isotopesbefore planet formation began.

Therefore, unless the inner solar system material from which Earthaccreted bears no resemblance to the material currently falling toEarth as meteorites, which cannot be reconciled with dynamic simula-tions, the Si isotopic composition of the BSE has been altered subse-quent to formation.

Enstatite chondrite(‰)

n Lunar averageb

(‰)n

−0.69 ± 0.03c 1 −0.31 ± 0.07 4−0.55 ± 0.04c 1−0.59 ± 0.10 3 −0.29 ± 0.05 6−0.56 ± 0.05c 1 −0.45 ± 0.05c 1−0.63 ± 0.07 3

−0.29 ± 0.08 24−0.67 ± 0.21 13EH:−0.77 ± 0.08d 6EL:−0.59 ± 0.09d 7−0.61 ± 0.07 2 −0.27 ± 0.04 5

to ±2 s.e. Here ±2 s.d. has been applied uniformly for consistency.abarti and Jacobsen's (2010) sample is lunar breccia, Armytage et al. (2012) include lunar(2013) included main mare basalt, with one cataclastic norite.2 s.d. on the population.

Fig. 3.Histograms and probability density functions summarising the variousmeteorite Siisotope data available in the literature, used to calculate the new estimate for bulk Earth,referred to as δ30Simeteorite* in this work (Section 4.4). Our new estimate of δ30Simeteorite*

is plotted as the broad dashed line, with light grey box representing the 2 s.d. uncertaintyon the estimate (thedotted line and dark grey box are the BSE estimate, see Fig. 1). The toppanel describes undifferentiatedmeteorite data, themiddle panel the differentiated stonymeteorites and the bottom panel shows lunar data. Data are from Armytage et al. (2011,2012), Fitoussi and Bourdon (2012), Fitoussi et al. (2009), Pringle et al. (2013a), Savageand Moynier (2013) and Zambardi et al. (2013).


4.3. The silicon isotopic composition of the BSE was fractionated by coreformation

Given the above arguments, the BSE can only be isotopically heavyrelative to chondrites if it is residual to another reservoir of light Si.Either, this reservoir is material that has been lost to space and othersolar system bodies, or, it is present in Earth's core. Poitrasson et al.(2004) found a small (0.03‰ per amu) effect in Fe isotopes interrestrial (and lunar) basalts relative to those from Mars and Vesta,and raised the possibility that this might relate to fractionation by lossto space during accretion. Iron is not very refractory (TC = 1337 K)and so might lose light isotopes during vaporisation during a GiantImpact or during earlier stages of accretion when Earth was smallerand material was more readily lost. O'Neill and Palme (2008) presentevidence for such losses.

However, such isotopic effects should be more pronounced in other(lighter) elements of similar volatility but greater relative mass

difference between isotopes. This is seen in small solar system objects,such as calcium aluminium-rich inclusions (CAIs), whereby both Mgand Si compositions are significantly enriched in the heavier isotopesrelative to chondritic. To give further support to this being as a resultof evaporative processes,where there isMg and Si isotope data for a sin-gle (zoned) CAI, these data are positively correlated (Shahar and Young,2007). However, for larger bodies where escape velocities are muchhigher, studies have failed to find a difference between BSE andMartianmeteorites or eucrites in, for instance, Li and Mg isotopes (Bourdonet al., 2010; Magna et al., 2006; Schiller et al., 2010; Seitz et al., 2007;Teng et al., 2010; Wiechert and Halliday, 2007). The Li and Mg isotopiccompositions of chondrites and the BSE have been the matter ofsome debate but the most recent studies of Pogge von Strandmannet al. (2011) and Lai et al. (2013) appear to show, atmost, a small differ-ence. Finally, there is some evidence that the heavier Fe isotopic compo-sition of basalts (with respect to chondrites) is due to partial meltingeffects, and that the composition of BSE is actually identical to chon-drites (e.g. Craddock et al., 2013). Therefore, a more likely explanationof the heavy Si isotope composition of BSE is that isotopically light Sipartitioned into the core, as proposed by Georg et al. (2007).

The possibility that Si is one of the “light elements” present inthe Earth's core has been argued by a number of workers on the basisof cosmochemical models (Allègre et al., 1995; McDonough, 2003).This is corroborated by experimental studies that show that, withincreasing pressure and temperature, and decreasing oxygen fugacity,Si exhibits increasingly siderophile behaviour (Gessmann et al., 2001;Wade andWood, 2005). Hence, independent evidence for the presenceof Si in the core would indicate that the early Earth's mantle must havebeen reducing at some time, and that core segregation occurred at thebase of a deep magma ocean (Wade and Wood, 2005).

One of the key arguments for a non-zero ΔBSE-chondrite being theresult of high pressure terrestrial core formation is that achondrites,differentiated meteorites from low pressure planetary bodies such asMars and Vesta, have approximately the same Si isotopic compositionas chondrites with only the Moon sharing the same δ30Si as BSE(Fig. 3). An exception to this are the aubrites, a group of meteoritesthought to have differentiated from enstatite chondrite-precursor ma-terial; given the enrichment of light Si isotopes in enstatite chondrites,it is perhaps not surprising that the aubrites have comparable Si isotopecompositions (Fig. 3). The conditions under which Si will partition intothe metal phase during metal–silicate differentiation are not expectedto have occurred on a planetary body the size of Mars and smaller(e.g. Gessmann et al., 2001; Kilburn and Wood, 1997; Wade andWood, 2005), except in the case of extremely low oxygen fugacities(Ziegler et al., 2010). Intriguingly, a recent paper by Pringle et al.(2013a) showed a small but resolvable enrichment in the heavier Siisotopes in meteorites from Vesta (HED), relative to chondrites, whichthey interpreted as evidence that Vesta's core contains ~1 wt.% Si;if this is correct, then the asteroid must have accreted at extremelylow oxygen fugacities (ΔIW ∼−4, where ΔIW is the oxygen fugacityexpressed as log units from the iron-wüstite buffer).

Fig. 3 shows the distribution of Si data from chondrites, achondritesand the Moon, with repeat analyses of the same samples from differentstudies being averaged out. As the studies of Chakrabarti and Jacobsen(2010) and Georg et al. (2007) were excluded from the earlier estimateof δ30SiBSE they have not been included in these compilations either. Thesimilarity of the various differentiated bodies and chondrites suggests,with the exception of enstatite meteorites, that the region of the innersolar system where all this material was accreted was relatively wellmixed with respect to Si isotopes.

4.4. Estimating δ30Sichondrite

In order to constrain ΔBSE-chondrite and hence the amount of Si inthe core, it is necessary to determine δ30SiBulk Earth. This might be as-sumed to be the same as δ30Sichondrite. However, with subtle variations


between different Si isotopic datasets, the composition of the bulk Earthreservoir is less than straightforward. For some mass independent iso-tope anomalies, enstatite chondrites appear the best match for proto-Earth material (see above), although one of the few common findingsbetween the recent Si isotopic studies points to enstatite chondriteslying at the extreme light end of the different meteorite groups. Indeed,this observation and its attendant high percentages of Si in the core ledFitoussi and Bourdon (2012) to argue on the basis of Si isotopic data thatenstatite chondrites could not represent the material that accreted toform Earth. Following this, Savage and Moynier (2013) performed thefirst extensive study on the enstatite chondrites (and achondrites).They found a significant difference between EH and EL chondrites,with the EH chondrites enriched in the lighter Si isotopes relative toEL (Table 3). These workers attributed the overall light Si enrichmentto the presence of isotopically light Si in the metal phase of enstatitechondrites (which formed under extremely reducing conditions). Theoffset between the two groups (EH and EL) was explained by the factthat EH chondrites have both more modally abundant metal and, also,higher Si concentrations in the metal phase. This was confirmed byanalysing metal-free enstatite chondrite powders, which had Si isotopecompositions indistinguishable from those of ordinary and carbona-ceous chondrite. As to why previous workers had not identified suchan offset before can be explained by sample heterogeneity; small differ-ences in metal abundance in a sample chip will have a large effect on itsSi isotope composition, due to the extremely light (~−5‰; Ziegler et al.,2010) composition of such metal — the same can be intimated regard-ing an aliquot of powder. Nevertheless, the argument against enstatitechondrites being included in an estimate of δ30SiBulk Earth is strong.

Fitoussi et al. (2009) used only carbonaceous chondrites in estimatingδ30Sichondrite, as the refractory lithophile element ratios in the Earth arebest explained by carbonaceous chondrites being Earth's building blocks(Palme and O'Neill, 2003). However, theirs was the only study to reportsmall differences between carbonaceous chondrites and other meteoritegroups, allowing for such selection. Georg et al. (2007) used an average ofall the chondrites, including enstatite chondrites as the average for bulkEarth. Armytage et al. (2011) calculated an estimate that they denote“δ30Simeteorite*” instead of δ30Sichondrite, which includes achondrite databut excludes enstatite meteorites based on the arguments presentedabove. Using these three approaches, the Si isotopic data from Fig. 3,would give δ30Sicarb. chondrite = −0.46 ± 0.10‰, δ30Sichondrite =−0.52 ± 0.22‰, and δ30Simeteorite* = −0.46 ± 0.09‰ (all ±2s.d.) forthe best estimates of the silicon isotopic composition of bulk Earth.The difference between these estimates is not significant andwell with-in the internal precision on individual samples (although the widerange of enstatite chondrite compositions explains the poorer precisionon δ30Sichondrite). Our preferencewould be for δ30Simeteorite*, as it is basedon more samples than δ30Sicarb. chondrite (64 rather than 14), and doesnot include the enstatite chondrites with their significantly lighterδ30Si compositions. Note that this average uses literature HED dataalthough Pringle et al. (2013a; Section 4.3) demonstrate that they aresomewhat enriched in the heavier Si isotopes — the offsets are smallhowever, and removing this group from the average does not affectthe final value, nor the precision.

4.5. Silicon in the Earth's core

Although Armytage et al. (2011), Fitoussi et al. (2009) andGeorg et al.(2007) all argue that the most likely cause of a non-zero Δ30SiBSE-meteorite

is Si entering the metal phase during core formation, translating this intoa well-defined compositional model for the core is harder.

The percentage of Si in the core can be calculated simply by massbalance and regardless of the mechanism or degree of equilibriumachieved, as follows:

δ30Simeteorite ¼ fδ30SiBSE þ 1− fð Þδ30Sicore ð4:1Þ

where f is the fraction of Si in the silicate Earth. If the fractionationfactor for Si isotopes between metal and silicate (ε — Eq. (2.2)) andΔ30SiBSE-meteorite = δ30SiBSE − δ30Simeteorite are substituted into Eq. (4.1):

Δ30SiBSE‐meteorite ¼ ε 1− fð Þ: ð4:2Þ

The percentage of Si in the core (X) is then given by calculating the Si“deficit” in BSE:

X ¼ 100 MBSE=MCoreð Þ c= fð Þ–c½ � ð4:3Þ

whereMBSE andMcore are themass fractions of BSE and the core, respec-tively, and c is themeasured Si fraction in BSE (0.212; Palme andO'Neill,2003). Combining Eqs. (4.2) and (4.3):

X ¼ 100 MBSE=MCoreð Þ c=1−Δ30SiBSE‐meteorite=ε� �

−ch i

: ð4:4Þ

The three variables in Eq. (4.4) (X, ε andΔSiBSE-meteorite) can all be in-dependently constrained. The range of X can be found either throughpartitioning experiments, or constraints from chondrite models ofbulk Earth chemistry; ε can be arrived at either through first principleslattice dynamical models or can be experimentally determined, whileΔSiBSE-meteorite is found by measuring the δ30Si of appropriate samples.

It has been known for some time that the fractionation factor for iso-tope exchange reactions is related to temperature by A / T2 (Bigeleisenand Mayer, 1947). Shahar et al. (2011) experimentally found themetal–silicate 30Si/28Si fractionation factor temperature dependence(in Kelvin) to be

ε ¼ 7:45� 106 � 0:41=T2 ð4:5Þ

which agrees well with previous experimental, empirical and theoreti-cal calculations (Georg et al., 2007; Shahar et al., 2009; Ziegler et al.,2010). Taking Δ30SiBSE-meteorite = 0.17 and 10 wt.% Si in the core as aworking maximum for X, Eq. (4.4) gives a minimum value for ε of0.936. From Eq. (4.5), the maximum temperature of core formationis therefore 2822+77

–79 K, which is in the range of the expected temper-ature (e.g. Gessmann and Rubie, 2000; Wade and Wood, 2005) formetal–silicate equilibration in the terrestrial magma ocean.

Errors onΔ30SiBSE-meteoritemean that the isotope data cannot be usedto put strong constraints on the temperature of core formationalthough, from partitioning studies (e.g. Wade and Wood, 2005) fixingthe maximum temperature of core formation to the peridotite liquidusis a good assumption. At core formation pressures of 40 GPa this gives3000 K which equates to ε = 0.828. This can be used along with themeasured Δ30SiBSE-meteorite to constrain the amount of Si in the core.There is still, however, considerable uncertainty for the amount of Siin the core resulting from errors in Δ30SiBSE-meteorite. Based on therange of published values the core could contain anywhere from 0 to44 wt.% Si, though the higher values are implausible based on chemicaland physical constraints. What is becoming more apparent, with in-creasing data, is that the simple, “single-stage”, core formation model,which is described above, appears to give estimates of X that all tendto be on the high side. For instance, utilising our new δ30Simeteorite* andδ30SiBSE values from this study gives X = 11.9+6.9

–5.0 wt.% Si. Morelikely, core formation occurred over a range of ambient pressure (P),temperature (T) and oxygen fugacity (fO2) conditions, somore compre-hensive modelling of Δ30SiBSE-meteorite can be used to place bounds onthe critical parameters.

Assuming, however, thatΔ30SiBSE-meteorite is solely the result of plan-etary differentiation, after identifying the crucial parameter(s) affectingΔ30SiBSE-meteorite, we can use inverse modelling techniques to placebounds around those very parameter(s). Detailed descriptions ofthe following model are beyond the scope of this review, however,examples for models of continuous core formation related to Si isotopefractionation can be found in Ricolleau et al. (2011), Zambardi et al.(2013), and Ziegler et al. (2010). The partitioning of Si into metal can

Fig. 4. Predicted relationship between the Si deficiency in Earth's mantle (i.e. the amountof Si that partitioned into Earth's core) and the degree to which Earth's mantle was morereducing relative to current day mantle. This last parameter is defined as gradient fO2 interms of log units (ΔIW) below modern day mantle oxidation state. Note that predictedSi deficiencies are invariant (~2.5 wt.%) if Earth's mantle started at ≤2 log units belowcurrent fO2 conditions.

Fig. 5. Modelled relationship between Si isotope fractionation caused by Si partitioninginto the core (Δ30SiBSE) and degree to which Earth's mantle was more reducing relativeto current day mantle (as in Fig. 4). Note that, in contrast to the amount of Si partitioninginto the core which is predicted to be invariant below gradient fO2 ~2 (Fig. 4), Si isotopefractionation continues to vary below this point, as a function of the T dependence onthe isotope fractionation factor. The letters associated with the various Δ30SiBSE valuesare in reference to various literature estimates of this value, as follows: a) Ziegler et al.(2010); b) Georg et al. (2007); c) Armytage et al. (2011); d) Fitoussi et al. (2009);e) Chakrabarti and Jacobsen (2010). Our model struggles to predict the small Δ30SiBSEvalues measured by refs d and e.


be described as a function of the P–T–fO2 in themantle during accretion(Corgne et al., 2008; Wade and Wood, 2005). High-P/T petrologyprovides the experimental framework for the mathematical expressionof that complex process, which enables us to formulate a mathematicalmodel to simulate Si partitioning during core formation in an accretingplanet. Although an abstract representation of reality, such models canprovide important insights into processes related to planetary accretion.Some model parameters may be fixed to comply with experimentalconstraints and model outputs can then be compared to actual dataobservation in an effort to understand how a systemmust have changedin order to create the observed data. The partitioning of Si is described bythe partition coefficient,which is the ratio of Si concentration in themetalto the Si concentration in the silicate:

DSi ¼ CSimetal=C

Sisilicate: ð4:6Þ

However, the partitioning of Si between metal and silicates is notas precisely constrained as is for other elements, such as nickel (Ni).Therefore, it is helpful to fix the Si partitioning to comply with thepartitioning of other siderophile elements, such as Ni and Cr. This reducesthe numbers of variables for Si partitioning by fixing the evolution ofP and T and the modern day oxygen fugacity in the mantle to complywith siderophile element partitioning. Expressing elemental partitioningas a function of P, T and fO2, we obtain:

logxmetalm

xsilicatem

!¼ aþ b

Tþ cP

Tþ d

nbot

− v4ΔIWþ 2log

γsilicateFeO

γmetalFe

!−log

γmetalm

γmetalFe

� �v=2 !

ð4:7Þ

And, to a good approximation:

Dm ¼ 10log

xmetalm

xsilicatem

� �ð4:8Þ

In Eqs. (4.7) and (4.8), m represents either Si or Ni, ΔIW is the oxygenfugacity (fO2) expressed as log units from the iron-wüstite buffer, xi isthe mole fraction of species i in the specified phase and γi is the activitycoefficient of species i in the specified phase. The regression parametersfor Si and Ni (a, b, c, d and v in Eq. (4.7)) are given elsewhere (Corgneet al., 2008; Rudge et al., 2010).

The model continuously recalculates Si partitioning and isotopefractionation as a function of evolving P, T and fO2 in an accretingEarth. With the evolution of P, T and modern day fO2 being fixed tocomply with siderophile element partitioning, the only variable forthe Si partitioning is the oxygen fugacity at the beginning of accretion,and thus the resulting gradient in fO2 to reach amodern daymantle ox-idation state. Therefore, the Si deficiency (i.e. the fraction of Si removedfrom BSE into the core) must also be a function of fO2 during accretion.Fig. 4 shows the Si deficiency as a function of fO2 gradient during accre-tion. A gradient of 3 in thisfigure indicates that accretion startedwith anoxygen fugacity 3 log units (ΔIW) belowmodern daymantle. The largerthe gradient, the larger the fraction of Si removed into the core andhence the Si deficiency in BSE (Fig. 4). It is interesting to note that theSi deficiency becomes independent of fO2 when the gradient is lessthan 2, i.e. fO2 began at ΔIW = −2 below modern day mantle. Thishas direct implications for the amount of Si in the core and BSE, asboth remain approximately the same for a wide range of fO2 gradients.However, the Si isotope composition provides a higher level of informa-tion and allows us to further restrict the range of possible oxygenfugacity gradients. The isotope fractionation during metal–silicatepartitioning is sensitive to temperature (Shahar et al., 2011) so thatfinal Δ30SiBSE-meteorite and Si deficiency are decoupled (Fig. 5). Here wecan see that, compared to Si deficiency (Fig. 4), the finalΔ30SiBSE-meteorite

is also sensitive to fO2gradients below 2, owing to the additional

temperature dependency on isotope fractionation. Here it also becomesapparent that Δ30SiBSE-meteorite is not necessarily a measure for the

amount of Si in the core and why single stage fractionation models(e.g. Armytage et al., 2011; Fitoussi et al., 2009; Georg et al., 2007) arefalling short of simulating Si isotope fractionation during core forma-tion. It should also be noted that the model can produce different Sicontents in the core for exact the same value of Δ30SiBSE-meteorite.While the Si isotope composition alone cannot represent the amountof Si in the core, it provides another level of information and allowsfor a stricter assessment of partitioning parameter.

Assuming the composition of bulk Earth to be close to that ofcarbonaceous chondrites (CV-type) we can run the model and solvefor the amount of Si in the core and the resulting Δ30SiBSE-meteorite, as a

Fig. 6. Predicted relationship between Si content in the core versus Si isotope fractionationcausedby Si partitioning into the core (Δ30SiBSE). References as in Fig. 5, themodel predictsbetween ~11 and 5wt.% Si in the core can explain the isotopic offset between BSE and bulkEarth (chondrite). Our new estimate of δ30Simeteorite* gives a Si core content of 6.2 wt.%(red dashed line).

Fig. 7. Range of literature Si isotope data available for various lunarmaterials. Colours referto various studies, sample names are given on the right adjacent to the data points. Averageδ30Silunar is shown as black line and grey box; the average composition including the studiesof Chakrabarti and Jacobsen (2010) and Georg et al. (2007) is shown by the blue dotted(average) and dashed (2 s.d. uncertainty) lines. Both averages are identical to BSE.


function of the starting fO2 (Fig. 6). Considering published estimates forΔ30SiBSE-meteorite, we can estimate that the core contains between ~10and 5.4 wt.% Si, with our new estimate for Δ30SiBSE-meteorite* (0.17‰)corresponding to a Si concentration in the core of 6.2 wt.%, which is inaccord to geophysical estimates (e.g. Anderson and Isaak, 2002) and ex-perimental estimates (Corgne et al., 2008; Wade and Wood, 2005).

4.6. Silicon isotopes and the formation of the Moon

The Moon is thought to have formed as a result of collision betweenthe proto-Earth and another planet “Theia”, and this has been successfullymodelled with smooth particle hydrodynamic (SPH) models (e.g. Canup,2008). One of the aspects of most of these simulations is that it is neces-sary to source the majority (70–80%) of the Moon-forming materialfrom the impactor. However, another constraint on the Moon's origin isthat lunar rocks have Δ17O compositions that are indistinguishable fromthe terrestrial fractionation line (Spicuzza et al., 2007; Wiechert et al.,2001). Variations in Δ17O are thought to be linked to accretion sourceregions in the early solar nebula (e.g. Clayton, 1993), and it is highlyunlikely that such a large, late impactor as Theia would have an identicalΔ17O composition to the proto-Earth.

In order to reconcile the SPH models and the oxygen isotopeevidence, Pahlevan and Stevenson (2007) proposed homogenisationacross a vapour cloud immediately following the giant impact. Howev-er, the processes of liquid–vapour exchange and reservoir separationthat would be involved in such a homogenisation could potentiallylead to mass dependent isotopic fractionations between the Earth andMoon. By modelling rainout from this vapour cloud, Pahlevan et al.(2011) predicted an offset of Δ30Si ~0.14‰ between the Moon andBSE. With the current high precision MC-ICP-MS techniques discussedearlier, such an offset should be resolvable.

Unlike the BSE and meteorite data, the variation between differenthigh precision Si isotope studies of lunar materials is more limited(Fig. 7, Table 3). The sample population includes ferroan anorthosites(FANs), picritic glasses, and high and low Ti-basalts. The FANs representsamples from the lunar highland crust, the high and low Ti basalts arethought to result from distinct source regions in the mantle based ontheir εNd and εHf trends (e.g. Unruh et al. 1984), while the picriticglasses are products of highly effusive eruptions from relatively deepsources in the lunar mantle (Head and Wilson, 1992). Despite theirvaried petrology, no isotopic differences can be resolved betweenthese different lithological types in δ30Si (Fig. 7). This limited distribu-tion in δ30Si rules out any fractionation related to lunar magma ocean(LMO) formation, mantle melting and differentiation, or source regionswith varying δ30Si at the current levels of precision. It is likely that thehomogeneity of the Si isotope composition of lunar mantle sampled sofar is the result of the initial genesis of the Moon. The lack of variationrelated to LMO formation, and the flotation of a plagioclase crust are

somewhat surprising, giving the significant inter-mineral Si isotopefractionation observed in terrestrial settings (e.g. Savage et al., 2011,Section 6.4). Unfortunately the dataset for relevant mineral separatesis still somewhat small, thus limiting our ability to fully addressthis issue.

Fig. 3 clearly shows that the Moon is the only body so far measuredin the inner solar system that has a BSE-like Si isotopic composition. Theaverage Si isotopic composition based on the data from Armytage et al.(2012), Chakrabarti and Jacobsen (2010), Fitoussi and Bourdon (2012),Georg et al. (2007), and Zambardi et al. (2013) is δ30Si = −0.30 ±0.09‰ (2 s.d.). As with earlier estimates, when excluding the data ofChakrabarti and Jacobsen (2010) and Georg et al. (2007) the averagebecomes δ30Si =− 0.29 ± 0.07‰ (2 s.d.). Both compositions are iden-tical to that defined as the “best” BSE composition earlier in Section 3.3,making the Earth–Moon system unique for Si isotopes in the innersolar system.

This compositional similarity has several implications for the earlyhistory of the Earth–Moon system. An impactor the size of Mars ismost likely to share the δ30Si composition of δ30Simeteorite* rather thanδ30SiBSE, as the conditions underwhich Si exhibits siderophile behaviour


(Gessmann et al., 2001) are thought not to have existed on smallerplanetary bodies. There are currently no convincing arguments thatthe Earth and Theia had similar accretionary histories such that thebulk silicate Earth and Theia should end upwith the same isotopic com-position. The simplest explanation of the Earth–Moon δ30Si similaritywould therefore be that the Moon predominantly formed from BSEmaterial, after Si segregated into the core. This would mean that the Siisotopic data, like oxygen, are in conflict with the SPH models thatconstrain Theia to be the primary source of theMoon-formingmaterial.The most recent models by Reufer et al. (2012) source at least 45% ofthe Moon from the impactor. Based on mass balance calculations,with these new constraints, the Moon should have a composition ofδ30Si = −0.37‰. This composition lies at the margins of the 2 s.d.error envelope for our lunar averages (Fig. 7). Therefore, to be consistentwith any of the giant impact models it is necessary for Si isotopes to behomogenised in the aftermath of the lunar forming collision. However,the Si isotopic data are also in opposition to the isotopic fractionationpredicted by Pahlevan et al. (2011) based on the oxygen isotopehomogenisation processes in the proto-lunar disc (Pahlevan andStevenson, 2007). As the Si isotopic composition of the Moon does notshow evidence of either a major contribution from Theia, and is notcompatible with the homogenisation processes proposed to explainthe identical Δ17O of the Earth–Moon system, the current Si dataset isnot consistent with the current Moon-forming giant impact models.These conclusions from Si are also echoed by data for O (Wiechertet al., 2001), Ti (Zhang et al., 2012) and W (Touboul et al., 2007),which are much heavier than Si yet also have identical BSE and Moonisotope compositions. An answer to this dilemmamay lie inmore recentsimulations with a different solution to the angular momentum andwhich facilitate either a small amount of material being derived fromTheia (Ćuk and Stewart, 2012) or an identical mix of material from theEarth and Theia (Canup, 2012).

5. The behaviour of Si isotopes during magmatic differentiation

5.1. Background

Since advances in instrumentation and analytical techniquesallowed for the investigation of small, high temperature variationsin Si isotopes, much research interest has centred on the Si isotopicvariations between BSE and bulk Earth/chondrites, as detailed above.Savage et al. (2010) indicated that there is limited Si isotope fraction-ation during mantle melting to form basalt; however, at that timethere was no specific study of how Si isotopes behaved during furthermagmatic differentiation toward rhyolitic compositions, that is, towardthe composition of the continental crust. Placing constraints on thisprocess was important: firstly, the continental crust is themajor sourceof riverine and oceanic silica (Derry et al., 2005), hence providing theisotopic composition of this reservoir is important for Si isotope studiesconcerning these reservoirs; secondly, and more generally, if Si isotopefractionation does take place as a result of magmatic differentiation,identifying the cause(s) was fundamental to the developing under-standing of the Si isotope system.

The effect of magmatic differentiation on various other “non-traditional” stable isotope systems has been investigated, with mixedresults. There is good evidence that Fe isotopes are affected bymagmat-ic differentiation, whereby more evolved (Si-rich) lithologies are moreenriched in the heavier Fe isotopes (e.g. Poitrasson and Freydier,2005; Schuessler et al., 2009; Teng et al., 2008). However, this fraction-ation is possibly linked to the redox sensitivity of Fe. For monovalentelements, such as Mg and Li, magmatic differentiation appears to havelittle effect on isotopic composition (e.g. Schuessler et al., 2009; Tenget al., 2007; Tomascak et al., 1999). This might suggest that Si isotopeswould be equally unaffected by igneous processes.

However, some evidence from early Si isotope data suggests thatfractionation can occur as a result of magmatic differentiation. In the

widely-cited Si isotope compilations of Ding et al. (1996) and Douthitt(1982), both showed that silica-rich (rhyolites, granites, dacites) lithol-ogies have, on average, heavier Si isotope compositions than primitive(ultramafic and basaltic) material. Unfortunately, by utilising graniticsamples, where petrogenesis may involve melting non-igneous sourcerocks, the effect of magmatic differentiation was difficult to deconvolve.Also, the large spread of data and relatively large error bars meant thatfurther constraints could not be made.

There is also theoretical evidence to suggest that resolvable equilibri-um Si isotope fractionation occurs between co-existing mineral phases,suggesting that e.g. fractional crystallisation will affect bulk rock Si iso-tope compositions. Following the techniques established by Bigeleisenand Mayer (1947), Grant (1954) calculated, from first principles, the Siisotope fractionation factors for a number of co-existing mineral pairs.He showed that fractionation was dependent on the vibrationalfrequencies of the Si-O bonds in a silicate; in particular, increasingthe degree of polymerisation results in stiffer Si\O bonds, so a morepolymerised (and hence, more Si-rich) phase is preferentially enrichedin the heavier Si isotopes. Although the sense of fractionation calculatedby Grant (1954) agrees with subsequent empirical data, the predictedmagnitudes of fractionation are much larger (up to Δ30Si ~1–5‰between mineral phases) than that measured in even the earliest Siisotope measurements (cf. Reynolds and Verhoogen, 1953).

Over fifty years later, this topic would be reappraised with a moresophisticated theoretical approach by Méheut et al. (2009), whoused density functional theory (DFT; see Capelle, 2006, for an eloquentintroduction) to approximate fractionation factors. As well as calculat-ing much smaller (and more likely) degrees of fractionation betweencrystal phases (Δ30Si b 0.5‰), these workers also demonstrated thatthe relationship between degree of polymerisation and Si isotopecomposition is not as simple as proposed by Grant (1954). Specifically,they showed that other network-forming cations in a silicate structure,such as Al, should alter the vibrational frequency of crystallographicsites, ultimately affecting degree of fractionation — meaning that themost polymerised silicate phase will not necessarily have the heaviestisotopic composition. Given the theoretical and possible empiricalevidence, it was apparent that Si isotopes may be fractionated bymagmatic differentiation. However, the predicted degrees of fraction-ation were relatively small, so it was only with the application of highresolution MC-ICP-MS that such phenomena could be adequatelyinvestigated.

5.2. Silicon isotope fractionation in rocks from Hekla volcano, Iceland

Savage et al. (2011) were the first to implicitly investigate the effectof magmatic differentiation on Si isotopes, principally using a set ofsamples taken from the volcano of Hekla. Hekla is an active fissurevolcano in the SW of Iceland (see Höskuldsson et al., 2007 and refer-ences therein) which is, in many ways, ideal for the study of magmaticdifferentiation. Firstly, it produces eruptive material with a wide rangeof silica contents (46 to 72 wt.% Si; basalt, basaltic andesite, andesite,dacite and rhyolite) from a cogenetic source. Second, the lavas areerupted through relatively young and unaltered mafic crust. Third,Hekla had been the focus of a number of previous geochemical studies,including Sr, Th, O, Fe and Li isotope analysis (Schuessler et al., 2009;Sigmarsson et al., 1992), which already placed a number of constraintson petrogenesis. In particular, these studies showed that magmaticevolution from basalt to rhyolite is not simple; fractional crystallisation,partial melting of crust and magma mixing all dominate at differentpoints in the suite, meaning that all such processes could be investigat-ed in terms of their effects on Si isotopes.

The Si isotope compositions of the Hekla samples vary from δ30Si =−0.34 to−0.14‰ (Fig. 8a), and define a wider range than that seen inmafic and ultramafic rocks analysed by MC-ICP-MS (Section 3, Fig. 1).Crucially, the isotope data exhibit strong positive correlations withSiO2 contents (R2 = 0.84; Fig. 8a) as well as other proxies for

a) Hekla

b) Various localities

Fig. 8. Silicon isotopic fractionation due tomagmatic differentiation— top panel shows thesample data from Hekla volcano, Iceland, as a function of SiO2 content and the apparentlylinear increase in δ30Si over the range of differentiation, over both fractional crystallisationand magma mixing regimes. Bottom panel shows other various “fresh” igneous samples,taken from various localities, which plot on the same array as the Hekla data. This seem-ingly consistent behaviour was termed the “igneous array” by Savage et al. (2011).Data taken from Savage et al. (2011, 2013a).


differentiation (i.e. Ba content, Mg#). This was the first conclusive em-pirical evidence that equilibrium Si isotope fractionation occurred as aresult of magmatic differentiation.

Two things are notable about this behaviour. Firstly, the sense of frac-tionation, whereby the more evolved lithologies are more enriched inthe heavier Si isotopes (relative to more primitive lithologies), agreeswith earlier theoretical studies and less precise empirical evidence(Section 5.1). Secondly, the Si isotope compositions become progressive-ly heavier over the whole range of increasing SiO2 contents (Fig. 8a).Because the Hekla lavas do not represent a simple differentiation suite,the Si isotope compositions evolve when fractional crystallisation andpartialmelting/magmamixing processes are the dominant differentiationmechanisms (Fig. 8a). These data appear then to indicate that afundamental property of the system, rather than one specific process,is controlling the Si isotope composition. Given the theoretical evidencepresented above (Grant, 1954;Méheut et al., 2009), Savage et al. (2011)suggested that themost likely cause of fractionation is the differences ininternal vibrational frequencies of the Si\O bonds in coexisting phases.During fractional crystallisation and/or partialmelting therewill always

be variations of this property between solid(s) and melt, as a result ofdifferences in polymerisation degree and presence of other network-forming cations in coexisting and newly forming mineral/melt phases.This discriminating factor therefore appears to be central to the behav-iour of the Si isotopic system.

5.3. The “igneous array” for silicon isotopes

For the relationship between SiO2 content and δ30Si to be trulyfundamental to the Si isotope system, it should be observable else-where. As well as the Hekla samples, Savage et al. (2011) analysed a se-lection of igneous samples from elsewhere on Iceland, the Afar Rift zonein Ethiopia and the USGS standard collection. When plotted againstSiO2, all of these samples show an increase in the heavier Si isotopesas a function of increasing silica content. This is also true for a selectionof granulite facies xenoliths from the lower continental crust, whoseprotoliths were equilibrium melt assemblages (Savage et al., 2013a).More significantly, these data are all collinear with the Hekla samples(Fig. 8b) — i.e., they fall on the same array. Part of this phenomenoncan be explained by the fact that the starting composition for all ofthese materials is that of BSE, which is homogeneous with respect toSi isotopes (Section 3). However, for all the samples to exhibit essential-ly the same degree of Si isotope fractionation away from BSE, regardlessof tectonic setting, is good evidence that such fractionation is a funda-mental property of the Si isotope system.

It appears then, that Si isotopes behave predictably duringmagmaticdifferentiation. As such, given a sample's silica content, its Si isotopecomposition can be estimated to a fairly accurate degree. This empiricalformula, determined by Savage et al. (2011), was termed the “igneousarray” for Si isotopes (Fig. 8b), and is defined as such:

δ30Si ‰ð Þ ¼ 0:0056� SiO2 wt:%ð Þ–0:567 �0:05;2� s:e:of theregressionð Þ:ð5:1Þ

For such a relationship to hold, a sample must be mantle derived,and represent an equilibrium melt assemblage, i.e. the formula willnot work for cumulate material (see Section 5.4).

5.4. Intermineral Si isotope fractionation

Even though the Si\O bond frequency is most likely to be a majorcontrol over the Si isotope composition of various silicate phases,there must be some mass flux to generate isotopic fractionation. Theseprocesses, in igneous systems, are most likely to be either fractionalcrystallisation during cooling, or partialmelting (and restite formation).

During crystallisation, to first order, more mafic (less polymerised)phases will form first, and during melting the more felsic (morepolymerised) phases will melt first. Table 4 provides a selection ofmineral-melt Si isotope fractionation factors for both mafic (theSkaergaard layered intrusion, Greenland) and felsic (granites from SWEngland and Australia) lithologies. Both olivine and clinopyroxene arepreferentially enriched in the lighter isotopes— fractional crystallisationof one or both of these phases will drive the melt toward heavierisotopic compositions (following the trend seen on Hekla and else-where, see Fig. 8), leaving isotopically light cumulate material.Conversely, partial melting of the same material will preferentiallycause melting of the plagioclase and clinopyroxene, leaving an olivine-rich restite (see Fig. 9a) — again, creating a melt that is preferentiallyenriched in the heavier isotopes relative to the solid. This simple de-scription effectively describes the phenomena of Si isotope fractionationduring magmatic differentiation.

The fractionation factors in Table 4 also illustrate that intermineral Siisotope fractionation is not completely controlled by relative polymeri-sation degree; although quartz seems always to be enriched in theheavier isotopes, and olivine in lighter isotopes, clinopyroxene is

Table 4Mineral-melt fractionation factors.

Δ30Si(mineral-melt) (‰)

Mafic (Skaergaard layered intrusion)Savage et al. (2011)

Olivine −0.15Clinopyroxene −0.21Plagioclase 0.11

Felsic (Granites — SW England, Australia)Georg (2006) Savage et al. (2012)

Quartz 0.04 0.05Plagioclase −0.05Muscovite 0.01 −0.02K-feldspar 0.02 0.06Tourmaline 0.17Biotite −0.28

a)

b)

Fig. 9. Silicon isotope variation in equilibrium melt assemblages compared to cumulateassemblages — samples are all granulite facies xenoliths, analysed by Savage et al.(2013a). In the top panel, samples that are representative of melts plot on the igneousarray (as in Fig. 8) whereas cumulate samples do not. In the bottom panel, the samedata are plotted against normative diopside content, and the cumulate samples showa good negative trend, indicating that the Si isotope composition of such samples iscontrolled by mineralogical abundances.


almost always lighter than coexisting olivine (see also Chakrabarti andJacobsen, 2010; Georg et al., 2007; Savage et al., 2011). As noted byMéheut et al. (2009), polymerisation is not the only fundamental con-trol over fractionation factor— both the v1 vibrational frequency (pulsa-tion along the Si\O bond) and presence of other network-formingcations (such as Al) exert strong controls. As a result, their calculationspredict that clinopyroxene should be isotopically lighter than olivine,as many researchers have measured. Such effects may also explainwhy, in Table 4, biotite is much lighter than muscovite, and tourmalineis much heavier than quartz.

As far as the “igneous array” is concerned, this means that thepredictable relationship between SiO2 and δ30Si will only hold whereequilibrium melt assemblages are concerned. Cumulates, made upof mineral phases that are not in equilibrium, will not exhibit such a re-lationship. This fact was demonstrated by Savage et al. (2013a) whoanalysed the Si isotope compositions of two suites of granulite facies(lower crustal) xenoliths from Queensland, Australia. The first suite,from McBride (c.f. Rudnick and Taylor, 1987) contains predominantlymetaigneous rocks whose protoliths represent equilibrium melt-derived material — these samples plot on the “igneous array” (Figs. 8b,9a), as expected. Samples from the second suite, from Chudleigh (c.f.Rudnick et al., 1986) are all metaigneous rocks with cumulate protoliths.These samples have a similar range of Si isotope composition to theMcBride suite, but do not exhibit a relationship with SiO2 contents, orany other proxy for magmatic differentiation (Fig. 9a). Instead, theyshow good relationships with mineralogical indices, for instant Mg#,Eu anomaly, normative diopside contents (Fig. 9b), implying that themodal abundance of various mineral phases is the controlling factor onthe Si isotope composition of the Chudleigh xenoliths. The negative rela-tionship in Fig. 9b also confirms the sense of isotopic fractionationbetween differing phases as described above: cumulates with moremafic phases have lighter Si isotope compositions that those with morefelsic phases.

A final point to make is that fractionation factors depend on the rela-tive bond strengths of the various crystal phases as well as the coexistingmelt. This means that specific phases will not have fixed mineral-meltfractionation factors. An example is plagioclase feldspar (Table 4). Inmafic systems it tends to have a positive Δ30Simineral-melt value, whereasin felsic systems it is enriched in the lighter isotopes.

5.5. The possibility of high temperature kinetic silicon isotope fractionation

Stable isotope fractionation as a result of kinetic, diffusion-related,processes in high temperature systems has been reported by a numberof experimental and theoretical studies (see Richter et al., 2009a, for areview). As the degree of fractionation, in most cases, can be muchlarger than that generated by equilibrium processes at equivalent

temperatures, here we review the current knowledge of kinetic Si iso-tope fractionation during chemical and thermal diffusion.

Chemical diffusion (chemical flux over a compositional gradient)has been shown to generate fractionation in a number of isotope sys-tems, such as Li, Mg and Ca, in both experimental and natural studies(Dauphas et al., 2010; Richter et al., 2003;Watkins et al., 2011). Howev-er, it does not appear that Si isotopes are affected in the sameway. UsingGeO2 as a proxy for silica, Richter et al. (1999) show that there is noresolvable isotope fractionation during diffusion experiments. Even ifGe is not an adequate proxy, Richter et al. (1999) note that silica gradientsnecessary to induce the degree of chemical diffusionnecessary for isotopefractionation rarely exist in nature.

In contrast to chemical diffusion, thermal (Soret) diffusion has beenshown, experimentally, to result in kinetic Si isotope fractionation


(Richter et al., 2009b). In all cases, the cooler region of the thermalgradient is preferentially enriched in the heavier Si isotopes,with some-what larger magnitudes of fractionation than those generated by equi-librium processes (Fig. 10). It is, however, unclear whether suchsignatures should be detectable in natural settings. For instance, the ef-fect Soret diffusion has on Si isotopes is shown to be much smaller thanthe Mg and Ca isotope systems (Richter et al., 2009b, Fig. 10) — both ofwhich show only veryminor isotopic variations in natural high temper-ature systems (Marshall and DePaolo, 1989; Teng et al., 2007), andthese have been attributed to equilibrium, not kinetic, processes. It isalso unclear how the experimental settings, made on chemically homo-geneous samples, can be applied to magma chambers where processessuch as fractional crystallisation, assimilation and even chemical diffu-sion will complicate the system. Finally, the timescales necessary togenerate significant Si isotope fractionation in natural settings, as a re-sult of thermal diffusion, are probably on a scale of millions of years(Lundstrom, 2009) — which would limit these effects to slow-coolingbodies, such as granite plutons. Nevertheless, there is also some limitedevidence that the Si isotope composition of granite plutons may evolveas a result of thermal gradients (Lundstrom, 2009), but the effects arenot predicted to be widespread in the Si isotope system, nor easilydeconvolved from other causes of fractionation.

Regarding extraterrestrial samples, there is good evidence for Siisotope fractionation by kinetic processes, specifically in calcium–

aluminium rich inclusions (CAIs), as noted in Section 4.3. Calcium–

aluminium rich inclusions are sub-millimetre to centimetre scale ob-jects formed in the early solar systemwhich show significant Si isotopicfractionations (Clayton et al., 1988) relative to other high temperaturematerials such as bulk meteorites or terrestrial igneous rocks. IgneousCAIs (i.e. those whose textures imply that they cooled from melts)tend to be enriched in the heavier isotopes of Si with δ30Si compositionsup to ~14‰, though the average is closer to ~3‰ (Clayton et al., 1988).These large variations in Si isotope composition are consistent withexperimental and theoretical work (Davis et al., 1990; Grossman et al.,2000; Richter et al., 2002; Wang et al., 1999, 2001) on kinetic isotopeeffects during evaporation in a primordial solar nebula. The directionand magnitude of fractionation indicate that igneous CAIs are evap-orative residues of a melt, with the isotopic fractionation followingRayleigh-like distillation (Davis et al., 1990). Recent in-situ work on Siisotopes in igneous CAIs (Bullock et al., 2013; Shahar and Young,2007) have found uniform (heavy) isotopic compositions, with the

Fig. 10.Kinetic isotope fractionation (via Soret diffusion) of various stable isotope systems,based on experimental data and figures in Richter et al. (2009b). In all systems, thecold end of the experiment becomes enriched in the heavier isotopes. The curves haveno scientific basis, they are fitted second (Si) and third (Mg, Ca) order polynomials. TheΩ values are defined as the fractionation (‰) per 100 °C per atomic mass unit differencebetween the isotopes (Richter et al., 2009b). Silicon isotopes show a much more limitedisotope fractionation than Mg and Ca (also Fe and O, not shown), but the magnitudesare still large compared to equilibrium isotope fractionation.

exception of the outer b200 μm, pointing to multiple melting eventsin these objects' histories. Some of the finer grained CAIs in Claytonet al. (1988) have light δ30Si compositions down to ~−3‰, which hasbeen suggested to be the result of condensation from a depleted vapour(Clayton et al., 1988; Davis et al., 1990).

Chondrules, which are the dominant constituent (60–80%) of themost common class of chondrite, show limited range (δ30Si −1 to1‰) in Si isotopic composition relative to CAIs (Clayton et al., 1991;Georg et al., 2007; Hezel et al., 2010). This restricted range points to sup-pression of kinetic isotope fractionation during the formation of chon-drules, as similarly limited isotopic variations have been observedin chondrules for other elements such as K (Alexander et al., 2000),Fe (Alexander and Wang, 2001; Zhu et al., 2001) and Mg (Galy et al.,2000). In parallel with the current lack of consensus on the locationand mechanism(s) of chondrule formation, there is no clear paradigmfor the origin of stable isotope signatures in chondrules.

6. Silicon isotopes and the formation and composition ofcontinental crust

6.1. Background

The continental crust is the ultimate source of silica to thesupracrustal cycle and to the riverine and oceanic Si budget (Tréguerand De La Rocha, 2013). The application of Si isotopes to such systemscan provide insights into continental weathering (Georg et al., 2006a,2009; Pogge von Strandmann et al., 2012; Ziegler et al., 2005a, 2005b),ocean palaeotemperature (Robert and Chaussidon, 2006) and oceanicprimary productivity (De la Rocha 2003; Hendry et al., 2010). For allof these systems, a global baseline is required. Given that magmatic dif-ferentiation affects the Si isotope composition of igneous rocks(Section 6), there is no a priori reason to assume that the continentalcrust has an identical Si isotope composition to that of BSE. Also,magmatic differentiation is not the only process that occurs during con-tinental crust petrogenesis. The following section describes how pro-cesses such as sedimentary anatexis and partial melting of pre-existingmetabasaltic material affect Si isotopes during the petrogenesisof continental crust, as well as the efforts to calculate a robust isotopecomposition for this important silicate reservoir.

6.2. The “igneous” composition of the continental crust

Using Eq. (5.1), aswell as the average SiO2 content of the continentalcrust (60.6 wt.%; Rudnick and Gao, 2003) an “igneous” composition forcontinental crust can be deduced. This value, δ30Si = −0.23 ± 0.05‰(2 s.e.), is enriched in the heavier isotopes compared to BSE and reflectsthe andesitic composition of this reservoir. However, this value repre-sents the composition of crust generatedwholly by equilibrium igneousprocesses. This is a very simplistic average: first, it does not take intoaccount the presence of sedimentary material in the crust, which definesa much wider range of isotopic compositions (e.g. Basile-Doelsch, 2006).Also, even though Si-rich igneous material comprises a significantamount of the continental crust, it is principally granitoid rocks, notrhyolites and andesites, which dominate the budget.

6.3. The silicon isotopic composition of granites

Granites (sensu lato) make up over half of themass of the continen-tal crust (Wedepohl, 1995), but are too Si-rich to be in equilibriumwiththemantle (Rudnick, 1995). Formation of much of the continental crustcannot, therefore, be a result of simple mantle melting. A number ofother processes, including fractional crystallisation, partial melting ofpre-existing mafic crust and sediment anatexis can all affect granitepetrogenesis.

In particular, the anatexis of sedimentary lithologies should generatedata that donot lie on the Si isotope “igneous array” (Section 5.3). This is


becauseweathering tends to enrich secondary minerals in the lighter Siisotopes relative to igneousmaterial (Opfergelt et al., 2012; Ziegler et al.,2005a, 2005b). Hence, partialmelting of pelitic sediment, or contamina-tion of mantle-derived melt by such material, should result in theformation of a crust which has a lighter Si isotope composition thanpurely “igneous” (e.g. I-type granites and rhyolites) material. In thisway, Si isotopes have the potential to differentiate between granitesthat have sedimentary or igneous affinities, thus allowing us to identifymantle-derived and hence “new” volumes of continental crust.

Douthitt (1982) tested this hypothesis by analysing a selection ofI- (Igneous, Infracrustal) and S-type (Sedimentary, Supracrustal) gran-ites from the Lachlan Fold Belt, Australia (see Chappell andWhite, 2001,for definitions). Their data span a wide range of isotopic compositions(δ30Si = −0.8 to +0.3‰) but no systematic variations between thetwo granite types were noted. However, it was always possible thatthe large errors (up to ±0.6‰) associated with these data could beobscuring significant variations. Subsequent work using MC-ICP-MSfurther investigated this field, but these studies were either limited inscope (Georg, 2006, presented the analysis of one S- and one I-typegranite) and/or the data remain unpublished (in conference abstracts,e.g. Zambardi and Poitrasson, 2011b). In both of these cases, however,the S-type samples analysed displayed a resolvably lighter isotopiccomposition than the I-type, consistent with the hypothesis that S-type granites form from sedimentary material preferentially enrichedin the lighter Si isotopes.

Fig. 11. Granite Si isotope data plotted versus SiO2 (top panels), age corrected 87Sr/86Sr and ASIplot on the “igneous array” indicating an igneous protolith for these samples (they have similimplying these samples derived from more isotopically variable protoliths. The good negatsedimentary protoliths, indicate that the S-type granites petrogenesis involved anatexis of isot

The study of Savage et al. (2012) analysed a broad selection of I-, S-and A-type (anorogenic granite; Loiselle and Wones, 1979) samples, aswell granitic hypabyssal lithologies (such as aplites and pegmatites;London, 2005) from Australia and England. Their granite data rangefrom δ30Si = −0.40 to −0.11‰, much more limited than recorded byDouthitt (1982); the large range displayed by the latter study is mostlikely a result of their larger analytical uncertainties. However, thedata from Savage et al. (2012) define a broader range compared toother high-Si igneous rocks, such as rhyolites and dacites (Section 5;Fig. 8) In particular, the S-type granite samples define the broadestrange of data and have, on average, the lightest Si isotope composition.This deviation away from the predicted “igneous” value can be seen on aplot of δ30Si vs. SiO2 against the “igneous array” (Fig. 11a & b). Whereasboth I- andA-type granites plot on or around the array, all but one of theS-type granites plot below the array, many to much more negativevalues thanwould be predicted in the absence of a sedimentary compo-nent. Further evidence that source variation, in particular sediment con-tamination or anatexis, is causing the Si isotope variability in granitescomes from the good negative correlations between initial 87Sr/86Srand ASI (alumina saturation index) values in the same samples(Fig. 11c & d). Both are good indicators of evolved crustal componentsin granites, hence, the very negative Si isotope compositions (whichare in most cases S-type samples) can be linked to the involvement ofpelitic material in themagma source of granites. The linear relationshipalso indicates that more weathered/evolved sediment will impart a

(alumina saturation index)— data taken from Savage et al. (2012). I- and A-type granitesar Si isotope compositions to rhyolites), whereas the S-type samples are more scattered,ive trends versus Sr isotopes and ASI, which are both proxies for “enriched” crustal oropically light (clay-rich) material.


more negative Si isotope signature to its partialmelt, and eventually to agranite formed from such material.

It is important to note, however, that the Si isotope system does notdifferentiate between I- and S-type granites as well as Sr isotopes or ASIindex. This is particularly evident for samples from the Lachlan Fold Belt(LFB), Australia, which is the type locality for these granite affinities(Chappell and White, 2001). The question is whether this is a generalissue with the Si isotope system, or whether it is a property of the LFBsamples.

A caveat with the hypothesis that Si isotopes will differentiatebetween granite affinities is that the sedimentary material from whichthe granites are formed is isotopically distinct from that of igneous ma-terial. For the LFB granites, thismay not be the case. Savage et al. (2012)analysed a set of Ordovician turbidite samples, cited bymanyworkers tobe the sedimentary source for LFB granites (e.g. Keay et al., 1997) andshowed them to be unfractionated relative to igneous rocks. Hence,any amount of these sediments in the magmatic source should notresult in a Si isotope signature different to that of I-type granites. How-ever, the fact that some LFB S-type granites are resolvably lighter couldindicate the presence of an as-yet unanalysed, isotopically light, sedi-mentary source. Furthermore, Zambardi and Poitrasson (2011b)analysed a selection of S-type granites sourced from France, the UKand Iran, and showed that they are, on the whole, resolvably differentfrom A- and I-type granites. The overlap between I- and S-type granitesmay be simply evidence that granites in the LFB are representative of acompositional continuum between igneous and sedimentary sources,rather than of distinct sources (e.g. Collins, 1998).

6.4. The silicon isotope composition of the upper continental crust

Granite petrogenesis includes both sedimentary and igneoussources, so they are arguably most representative of the lithologiespresent in the upper continental crust. Savage et al. (2012) thereforecalculated the Si isotope composition of the upper continental crustby taking an average of their granite data, giving a value of δ30Si =−0.23 ± 0.15‰ (2 s.d.). This estimate is identical to the value calculat-ed using the igneous array (Section 6.1) although the larger errors areconsistent with the wider spread of data when granites are concerned.Also, the “igneous” estimate uses the average SiO2 content of the conti-nental crust (60.6 wt.%), rather than the upper continental crust, whichis more Si-rich (~66 wt.%; Rudnick and Gao, 2003) and would give acomposition of δ30Si ~−0.20‰. The slightly lighter granite averagereflects the incorporation of isotopically light sedimentary materialinto the granite source region, and is arguably a more robust value.In a subsequent paper, Savage et al. (2013b) analysed a large rangeof upper crustal sedimentary lithologies, including shales and loess.Although these samples span amuch larger spread of Si isotope compo-sitions than igneous material, a weighted lithological upper crustalaverage gives a composition of δ30Si =−0.25 ± 0.16‰ (2 s.d.), identi-cal to that of the granitic average and also very close to that of BSE. Thereasons for this are three-fold: first, any weighted lithological estimateof the upper crust will be dominated by granites (sensu lato) whichcomprise ~50% of this reservoir (Wedepohl, 1995); second, the Si isoto-pic compositions of many upper crustal lithologies, such as loess, aredominated by the high modal abundances of detrital quartz, which isigneous in origin and should therefore be similar to granites; and finally,the light Si isotope enrichment that is associated with neoformation ofclay minerals during weathering is effectively balanced out by a) theslight enrichment to heavier Si isotope compositions as a result ofmagmatic differentiation and b) the incorporation of isotopicallyheavy biogenic Si (such as opals and chert) into the crustal stratigraphy(see Savage et al., 2013b for further information). The striking similaritybetween BSE and the crust implies that the amount of surface dissolved,isotopically heavy, Si removed to the oceans via weathering and subse-quently to the mantle via subduction over time is small, or such losseshave been balanced by subduction of isotopically light clay.

7. Summary and future work

In the past few years our understanding of the Si isotopic composi-tion of the Earth has changed dramatically, providing us with importantnew insights into major Earth reservoirs and some powerful new linesof evidence constraining Earth processes.

The first, without question, is the realisation that the mantle isessentially homogeneous with respect to Si isotopes (Savage et al.,2010). This was ambiguous until precise methods were developed andeven early attempts left the issue unresolved (Georg et al., 2007). Nowthere is no clear evidence for discrete isotopic domains, reservoirs oreven enriched streaks that reflect recycled heavy Si from subductedArchaean cherts, for example. The mantle is well mixed with respectto Si isotopes.

The second important discovery is that Si isotopes are not greatlyfractionated by partial melting of the Earth's mantle. There is some ef-fect but it is small. More precisemethods are needed to fully and reliablyresolve these effects; however, it is clear that comparisons betweenthe Si isotopic compositions of basalts from different planets probablyreflect the parent bodies themselves. This is not the case for all stableisotope systems, Fe isotopes for example (e.g. Halliday, 2013 and refer-ences therein).

Perhaps the biggest discovery has been that meteorites, includingbasalts from other planets, are not as isotopically heavy as terrestrialbasalts and mantle peridotites (e.g. Armytage et al., 2011; Fitoussiet al., 2009; Georg et al., 2007; Zambardi et al., 2013). With the excep-tion of one study (Chakrabarti and Jacobsen, 2010), a significant andsystematic difference has been found between the silicate Earth andmeteorites. This could reflect isotopic diversity in the circumstellardisc from which the planets formed; however, all meteorites appear tohave the same Si isotope composition, with the exception of enstatitechondrites which are, isotopically, even lighter, not heavy. Amore likelyexplanation, consistent with experiments, theory, outer core densityand a non chondritic Si/Mg in the primitivemantle is that the Si isotopiccomposition of the silicate Earth has been fractionated bymetal–silicateequilibria. As such, Si isotopes provide powerful new evidence that Si isone of the light elements in the core. This process can be modelledthrough accretion but in future requires more sophistication basedon further experiments, because Earth clearly grew and evolved as itscore formed resulting in changing pressure, temperature and oxygenfugacity.

This distinctive isotopic composition of the silicate Earth has onlybeen replicated in samples from the Moon (e.g. Armytage et al., 2012).This is an important new constraint on the origin of the Moon. Nearlyall Giant Impact models using smooth particle hydrodynamic codegenerate the majority of the material in the Moon from the impactor.Previously it has been shown that O (Wiechert et al., 2001) and Tiisotopes (Zhang et al., 2012) are the same in the Earth and Moon. Thiscould reflect the possibility that the proto Earth and impacting planetTheia were formed at the same heliocentric distance. However, the Siand W (Touboul et al., 2007) isotope similarity can only plausibly beexplained if the atoms in the Moon were derived from, or equilibratedwith, those in the silicate Earth (Pahlevan and Stevenson, 2007).

The continental crust is an enriched reservoir, the andesitic compo-sition ofwhichhas been amatter of interest for decades given thatman-tle melting today largely generates basalt, even in arcs. One model thathas been proposed is that weathering followed by recycling of materialinto the mantle by subduction has gradually changed the compositionof the residual crust (Albarède and Michard, 1986). Silicon isotopescould record this effect because the composition is known to be frac-tionated by clay formation duringweathering. The study of differentiat-ed magmatic sequences formed away from mature continental crustdemonstrates that Si isotopes are slightly and systematically fractionat-ed to heavier compositions in rhyolites by removal of isotopically lightolivine and pyroxene (Savage et al., 2011; Zambardi and Poitrasson,2011b). This simple fractionation trend can be used to calculate an


average isotopic composition of continental crust based purely on itschemistry. Sedimentary rocks are highly variable in Si isotopic composi-tion because of fractionation during low temperature chemicalweathering and/or biological utilisation. It is no surprise therefore thatgranites, many of which have melted or assimilated sediments, havemore variable Si isotopes for a given silica content. What is surprisingis that an estimate for the average isotopic composition of the continen-tal crust based on lithologies looks so similar to that based on Si contentand closed system fractional crystallisation.Weathering appears to havehad no currently resolvable effect on the bulk isotopic composition ofthe continental crust and it is only slightly heavy relative to the mantle.Therefore, there is no vestige of a signature of continental compositionthat results from long term weathering and mantle recycling. Sucheffects must be at best small.

Future research will need to explore the isotopic fractionation thattakes place duringmelting, core formation and differentiation in furtherdetail. Clearly, calibration of these effects with rigorous experimentsthrough a range of parameter space is key to further utilisation of thetechnique and its application to other solar system samples. Regardingtechnique development, more work is needed to control mass bias toachieve further improved sample–standard bracketing that is lessmatrix sensitive. The utilisation of the method to solve new scientificproblemswill require techniques that arewell understood, shared openlyand not plagued with the same inter-laboratory differences that havebeen seen in the first few years of using MC-ICPMS for this element.

Acknowledgements

The authors are very grateful to Tim Horscroft for inviting the reviewand for being so patient during its protracted writing. The authors wouldalso like to extend their thanks to the two anonymous reviewers whosecomments greatly improved this manuscript and, also, to Andrew Kerrfor astute editorial assistance. PS and RA both extend their gratitude toHelenWilliams andKevin Burton for sound academic advice and support,as well as all colleagues at the Department of Earth Sciences at Oxford,where their initiation into the field of Si isotopes occurred, and wheremany of the data quoted in this review were generated. Support forsilicon isotope research at Oxford University has been providedby grants to ANH from STFC, NERC and ERC. PS would also like toacknowledge the financial support of the McDonnell Center for SpaceSciences and the Marie Curie IOF fellowship programme.

References

Abraham, K., Opfergelt, S., Fripiat, F., Cavagna, A.-J., de Jong, J.T.M., Foley, S.F., André, L.,Cardinal, D., 2008. δ30Si and δ29Si determinations on USGS BHVO-1 and BHVO-2 ref-erence materials with a new configuration on a Nu Plasma Multi-Collector ICP-MS.Geostandards and Geoanalytical Research 32, 193–202.

Albarède, F., Beard, B., 2004. Analytical methods for non-traditional isotopes. In: Johnson,C.M., Beard, B.L., Albarède, F. (Eds.), Reviews in Mineralogy and Geochemistry, 55.Mineralogical Society of America, pp. 113–152.

Albarède, F., Michard, A., 1986. Transfer of continental Mg, S, O and U to the mantlethrough hydrothermal alteration of the oceanic crust. Chemical Geology 57, 1–15.

Alexander, C.M.O., Wang, J., 2001. Iron isotopes in chondrules: implications for the role ofevaporation during chondrule formation. Meteoritics and Planetary Science 36,419–428.

Alexander, C.M.O., Grossman, J.N., Wang, J., Zanda, B., Bourot-Denise, M., Hewins, R.H.,2000. The lack of potassium-isotopic fractionation in Bishunpur chondrules. Meteoriticsand Planetary Science 35, 859–868.

Allègre, C.J., Poirier, J.-P., Humler, E., Hofmann, A.W., 1995. The chemical composition ofthe Earth. Earth & Planetary Science Letters 134, 515–526.

Anderson, O.L., Isaak, D.G., 2002. Another look at the core density deficit of Earth's outercore. Physics of the Earth and Planetary Interiors 131, 19–27.

Armytage, R.M.G., Georg, R.B., Savage, P.S., Williams, H.M., Halliday, A.N., 2011. Siliconisotopes in meteorites and planetary core formation. Geochimica et CosmochimicaActa 75, 3662–3676.

Armytage, R.M.G., Georg, R.B., Williams, H.M., Halliday, A.N., 2012. Silicon isotopesin lunar rocks: implications for the Moon's formation and the early history of theEarth. Geochimica et Cosmochimica Acta 77, 504–514.

Basile-Doelsch, I., 2006. Si stable isotopes in the Earth's surface: a review. Journal ofGeochemical Exploration 88, 252–256.

Basile-Doelsch, I., Meunier, J.D., Parron, C., 2005. Another continental pool in the terrestrialsilicon cycle. Nature 433, 399–402.

Bigeleisen, J., Mayer, M.G., 1947. Calculation of equilibrium constants for isotopicexchange reactions. Journal of Chemical Physics 15, 261–267.

Birch, F., 1964. Density and composition of mantle and core. Journal of GeophysicalResearch 69, 4377–4388.

Bourdon, B., Tipper, E.T., Fitoussi, C., Stracke, A., 2010. Chondritic Mg isotope compositionof the Earth. Geochimica et Cosmochimica Acta 74, 5069–5083.

Brzezinski, M.A., Jones, J.L., Beucher, C.P., Demarest, M.S., Berg, H.L., 2006. Automateddetermination of silicon isotope natural abundance by the acid decomposition ofcesium hexafluosilicate. Analytical Chemistry 78, 6109–6114.

Bullock, E.S., Knight, K.B., Richter, F.M., Kita, N.T., Ushikubo, T., MacPherson, G.J., Davis,A.M., Mendybaev, R.A., 2013. Mg and Si isotopic fractionation patterns in types B1and B2 CAIs: implications for formation under different nebular conditions. Meteoriticsand Planetary Science 48, 1440–1458.

Canup, R.M., 2008. Lunar-forming collisions with pre-impact rotation. Icarus 196,518–538.

Canup, R.M., 2012. Forming a moon with an Earth-like composition via a giant impact.Science 338, 1052–1055.

Capelle, K., 2006. A bird's-eye view of density-functional theory. Brazilian Journal of Physics36, 1318–1341.

Cardinal, D., Alleman, L.Y., De Jong, J., Ziegler, K., André, L., 2003. Isotopic composition ofsilicon measured by multicollector plasma source mass spectrometry in dry plasmamode. Journal of Analytical Atomic Spectrometry 18, 213–218.

Cardinal, D., Savoye, N., Trull, T.W., Dehairs, F., Kopczynska, E.E., Fripiat, F., Tison, J.-L.,André, L., 2007. Silicon isotopes in spring Southern Oceandiatoms: large zonal changesdespite homogeneity among size fractions. Marine Chemistry 106, 46–62.

Chakrabarti, R., Jacobsen, S.B., 2010. Silicon isotopes in the inner Solar System: implica-tions for core formation, solar nebular processes and partial melting. Geochimica etCosmochimica Acta 74, 6921–6933.

Chappell, B.W., White, A.J.R., 2001. Two contrasting granite types: 25 years later.Australian Journal of Earth Sciences 48, 489–500.

Clayton, R.N., 1993. Oxygen isotopes in meteorites. Annual Review of Earth and PlanetarySciences 21, 115–149.

Clayton, R., Hinton, R., Davis, A., 1988. Isotopic variations in the rock-forming elements inmeteorites. Philosophical Transactions of the Royal Society of London Series a-Mathematical Physical and Engineering Sciences 325, 483–501.

Clayton, R.N., Mayeda, T.K., Goswami, J.N., Olsen, E.J., 1991. Oxygen isotope studies ofordinary chondrites. Geochimica et Cosmochimica Acta 55, 2317–2337.

Collins, W.J., 1998. Evaluation of petrogenetic models for Lachlan Fold Belt granitoids:implications for crustal architecture and tectonic models. Australian Journal ofEarth Sciences 45, 483–500.

Corgne, A., Keshav, S., Wood, B.J., McDonough, W.F., Fei, Y., 2008. Metal–silicatepartitioning and constraints on core composition and oxygen fugacity during Earthaccretion. Geochimica et Cosmochimica Acta 72, 574–589.

Craddock, P.R., Warren, J.M., Dauphas, N., 2013. Abyssal peridotites reveal the near-chondritic Fe isotopic composition of the Earth. Earth and Planetary Science Letters365, 63–76.

Ćuk, M., Stewart, S.T., 2012. Making the moon from a fast-spinning earth: a giant impactfollowed by resonant despinning. Science 338, 1047–1052.

Dauphas, N., Teng, F.-Z., Arndt, N.T., 2010. Magnesium and iron isotopes in 2.7 Ga Alexokomatiites: mantle signatures, no evidence for Soret diffusion, and identificationof diffusive transport in zoned olivine. Geochimica et Cosmochimica Acta 74,3274–3291.

Davis, A.M., Hashimoto, A., Clayton, R.N., Mayeda, T.K., 1990. Isotope mass fractionationduring evaporation of Mg2SiO4. Nature 347, 655–658.

De La Rocha, C.L., 2003. Silicon isotope fractionation by marine sponges and thereconstruction of the silicon isotope composition of ancient deep water. Geology31, 423–426.

De La Rocha, C.L., 2006. Opal-based isotopic proxies of paleoenvironmental conditions. Glob-al Biogeochemical Cycles 20 (4), GB4S09. http://dx.doi.org/10.1029/2005GB002664.

De La Rocha, C.L., Brzezinski, M.A., DeNiro, M.J., Shemesh, A., 1998. Silicon-isotope compo-sition of diatoms as an indicator of past oceanic change. Nature 395, 680–683.

Derry, L.A., Kurtz, A.C., Ziegler, K., Chadwick, O.A., 2005. Biological control of terrestrialsilica cycling and export fluxes to watersheds. Nature 433, 728–731.

Ding, T., Jiang, S., Wan, D., Li, Y., Li, J., Song, H., Liu, Z., Yao, X., 1996. Silicon IsotopeGeochemistry. Geological Publishing House, Beijing.

Douthitt, C.B., 1982. The geochemistry of the stable isotopes of silicon. Geochimica etCosmochimica Acta 46, 1449–1458.

Drake, M.J., Righter, K., 2002. Determining the composition of the Earth. Nature 416,39–44.

Egan, K.E., Rickaby, R.E.M., Leng, M.J., Hendry, K.R., Hermoso, M., Sloane, H.J., Bostock, H.,Halliday, A.N., 2012. Diatom silicon isotopes as a proxy for silicic acid utilisation:a Southern Ocean core top calibration. Geochimica et Cosmochimica Acta 96, 174–192.

Epstein, S., Taylor Jr., H.P., 1970. 18O/16O, 30Si/28Si, D/H, and 13C/12C studies of lunar rocksand minerals. Science 167, 533–535.

Fitoussi, C., Bourdon, B., 2012. Silicon isotope evidence against an enstatite chondriteearth. Science 335, 1477–1480.

Fitoussi, C., Bourdon, B., Kleine, T., Oberli, F., Reynolds, B.C., 2009. Si isotope systematics ofmeteorites and terrestrial peridotites: implications for Mg/Si fractionation in the solarnebula and for Si in the Earth's core. Earth and Planetary Science Letters 287, 77–85.

Galy, A., Young, E.D., Ash, R., O'Nions, R.K., 2000. The formation of chondrules at high gaspressures in the solar nebula. Science 290, 1751–1753.

Georg, R.B., 2006. Geochemistry of Stable Silicon Isotopes measured by High-ResolutionMulti-Collector Inductively-Coupled Plasma Mass Spectrometry (HR-MC-ICP-MS).(PhD thesis) ETH, Zurich.

http://refhub.elsevier.com/S0024-4937(14)00007-3/rf0005






















































































http://dx.doi.org/10.1029/2005GB002664
































Georg, R.B., Reynolds, B.C., Frank, M., Halliday, A.N., 2006a. Mechanisms controlling thesilicon isotopic compositions of river waters. Earth and Planetary Science Letters249, 290–306.

Georg, R.B., Reynolds, B.C., Frank, M., Halliday, A.N., 2006b. New sample preparationtechniques for the determination of Si isotopic compositions usingMC-ICPMS. ChemicalGeology 235, 95–104.

Georg, R.B., Halliday, A.N., Schauble, E.A., Reynolds, B.C., 2007. Silicon in the Earth's core.Nature 447, 1102–1106.

Georg, R.B., Zhu, C., Reynolds, B.C., Halliday, A.N., 2009. Stable silicon isotopes of ground-water, feldspars, and clay coatings in the Navajo Sandstone aquifer, Black Mesa,Arizona, USA. Geochimica et Cosmochimia Acta 73, 2229–2241.

Gessmann, C.K., Rubie, D.C., 2000. The origin of the depletions of V, Cr and Mn in themantles of the Earth and Moon. Earth and Planetary Science Letters 184, 95–107.

Gessmann, C.K., Wood, B.J., Rubie, D.C., Kilburn, M.R., 2001. Solubility of silicon in liquidmetal at high pressure: implications for the composition of the Earth's core. Earthand Planetary Science Letters 184, 367–376.

Grant, F.S., 1954. The geological significance of variations in the abundances of theisotopes of silicon in rocks. Geochimica et Cosmochimica Acta 5, 225–242.

Grossman, L., Ebel, D.S., Simon, S.B., Davis, A.M., Richter, F.M., Parsad, N.M., 2000. Majorelement chemical and isotopic compositions of refractory inclusions in C3 chondrites:the separate roles of condensation and evaporation. Geochimica et CosmochimicaActa 64, 2879–2894.

Halliday, A.N., 2000. Terrestrial accretion rates and the origin of the Moon. Earth andPlanetary Science Letters 176, 17–30.

Halliday, A.N., 2004. Mixing, volatile loss and compositional change during impact-drivenaccretion of the Earth. Nature 427, 505–509.

Halliday, A.N., 2013. Earth science: small differences in sameness. Nature 497, 43–45.Halliday, A.N., Stirling, C.H., Freedman, P.A., Oberli, F., Reynolds, B., Georg, R.B., 2011. High

precision isotope ratio measurements using multiple collector inductively coupledplasma mass spectrometry. In: Houk, R., Matthews, D. (Eds.), Elemental, Isotopicand Inorganic Analysis by Mass Spectrometry. Encyclopaedia of Mass Spectrometry,vol. 5. Elsevier.

Head, J.W., Wilson, L., 1992. Lunar mare volcanism: stratigraphy, eruption conditions, andthe evolution of secondary crusts. Geochimica et Cosmochimica Acta 56, 2155–2175.

Heck, P.R., Huberty, J.M., Kita, N.T., Ushikubo, T., Kozdon, R., Valley, J.W., 2011. SIMSanalyses of silicon and oxygen isotope ratios for quartz from Archean andPaleoproterozoic banded iron formations. Geochimica et Cosmochimica Acta 75,5879–5891.

Hendry, K.R., Georg, R.B., Rickaby, R.E.M., Robinson, L.F., Halliday, A.N., 2010. Deep oceannutrients during the Last Glacial Maximum deduced from sponge silicon isotopiccompositions. Earth and Planetary Science Letters 292, 290–300.

Hezel, D.C., Needham, A.W., Armytage, R., Georg, B., Abel, R.L., Kurahashi, E., Coles, B.J.,Rehkämper, M., Russell, S.S., 2010. A nebula setting as the origin for bulk chondruleFe isotope variations in CV chondrites. Earth and Planetary Science Letters 296,423–433.

Höskuldsson, Á., Óskarsson, N., Pedersen, R., Grönvold, K., Vogfjöro, K., Ólafsdóttir, R.,2007. The millennium eruption of Hekla in February 2000. Bulletin of Volcanology70, 169–182.

Hughes, H.J., Delvigne, C., Korntheuer, M., de Jong, J., André, L., Cardinal, D., 2011. Control-ling themass bias introduced by anionic and organic matrices in silicon isotopic mea-surements by MC-ICP-MS. Journal of analytical atomic spectrometry 26, 1892–1896.

Johnson, C.M., Beard, B.L., Albarède, F., 2004. Geochemistry of non-traditional stable isotopes.Reviews in Mineralogy and Geochemistry, 55. Mineralogical Society of America.

Keay, S., Collins, W.J., McCulloch, M.T., 1997. A three-component Sr–Nd isotopic mixingmodel for granitoid genesis, Lachlan fold belt, eastern Australia. Geology 25, 307–310.

Kempl, J., Vroon, P.Z., Zinngrebe, E., van Westrenen, W., 2013. Si isotope fractionationbetween Si-poor metal and silicate melt at pressure–temperature conditions relevantto metal segregation in small planetary bodies. Earth and Planetary Science Letters368, 61–68.

Kilburn, M.R., Wood, B.J., 1997. Metal–silicate partitioning and the incompatibility of Sand Si during core formation. Earth and Planetary Science Letters 152, 139–148.

Lai, Y.-J., Pogge von Strandmann, P.A.E., Elliott, T., Russell, S., Brooker, R., 2013. MineralogicalMagazine 77 (5), 1539.

Lodders, K., 2000. An oxygen isotope mixing model for the accretion and composition ofrocky planets. Space Science Reviews 92, 341–354.

Loiselle, M.C., Wones, D.R., 1979. Characteristics and origin of anorogenic granites.Geological Society of America Abstracts with Programs 11, 468.

London, D., 2005. Granitic pegmatites: an assessment of current concepts and directionsfor the future. Lithos 80, 281–303.

Lundstrom, C., 2009. Hypothesis for the origin of convergent margin granitoidsand Earth's continental crust by thermal migration zone refining. Geochimica etCosmochimica Acta 73, 5709–5729.

Magna, T., Wiechert, U., Halliday, A.N., 2006. New constraints on the lithium isotopecompositions of the Moon and terrestrial planets. Earth and Planetary Science Letters243, 336–353.

Marshall, B.D., DePaolo, D.J., 1989. Calcium isotopes in igneous rocks and the origin ofgranite. Geochimica et Cosmochimica Acta 53, 917–922.

McDonough, W.F., 2003. Compositional model for the Earth's core. In: Holland, H.D.,Turekian, K.K. (Eds.), Treatise on Geochemistry. Pergamon, Oxford, pp. 547–568.

Méheut, M., Lazzeri, M., Balan, E., Mauri, F., 2009. Structural control over equilibriumsilicon and oxygen isotopic fractionation: a first-principles density-functional theorystudy. Chemical Geology 258, 28–37.

Molini-Velsko, C., Mayeda, T.K., Clayton, R.N., 1986. Isotopic composition of silicon inmeteorites. Geochimica et Cosmochimica Acta 50, 2719–2726.

Nelson, D.M., Tréguer, P., Brzezinski, M.A., Leynaert, A., Queguiner, B., 1995. Productionand dissolution of biogenic silica in the ocean: revised global estimates, comparison

with regional data and relationship to biogenic sedimentation. Global BiogeochemicalCycles 9, 359–372.

O'Neil, J.R., 1986. Theoretical and experimental aspects of isotopic fractionation. In: Valley,J.W., Taylor Jr., H.P., O'Neil, J.R. (Eds.), Stable Isotopes in High Temperature GeologicalProcesses. Reviews in Mineralogy, 16. Mineralogical Society of America, pp. 1–40.

O'Neill, H.S.C., Palme, H., 2008. Collisional erosion and the non-chondritic composition ofthe terrestrial planets. Philosophical Transactions of the Royal Society A: Mathematical,Physical and Engineering Sciences 366, 4205–4238.

Opfergelt, S., Georg, R.B., Burton, K.W., Guicharnaud, R., Siebert, C., Gislason, S.R., Halliday,A.N., 2011. Silicon isotopes in allophane as a proxy for mineral formation in volcanicsoils. Applied Geochemistry 26, S115–S118.

Opfergelt, S., Georg, R.B., Delvaux, B., Cabidoche, Y.-M., Burton, K.W., Halliday, A.N., 2012.Silicon isotopes and the tracing of desilication in volcanic soil weathering sequences,Guadeloupe. Chemical Geology 326, 113–122.

Pahlevan, K., Stevenson, D.J., 2007. Equilibration in the aftermath of the lunar-forminggiant impact. Earth and Planetary Science Letters 262, 438–449.

Pahlevan, K., Stevenson, D.J., Eiler, J.M., 2011. Chemical fractionation in the silicate vaporatmosphere of the Earth. Earth and Planetary Science Letters 301, 433–443.

Palme, H., O'Neill, H.S.C., 2003. Solar system abundances of the elements. In: Holland, H.D.,Turekian, K.K. (Eds.), Treatise on Geochemistry, Vol. 2. Pergamon, Oxford, pp. 1–38.

Pogge von Strandmann, P.A.E., Elliott, T., Marschall, H.R., Coath, C., Lai, Y.-J., Jeffcoate, A.B.,Ionov, D.A., 2011. Variations of Li andMg isotope ratios in bulk chondrites andmantlexenoliths. Geochimica et Cosmochimica Acta 75, 5247–5268.

Pogge von Strandmann, P.A.E., Opfergelt, S., Lai, Y.-J., Sigfússon, B., Gislason, S.R., Burton,K.W., 2012. Lithium, magnesium and silicon isotope behaviour accompanyingweathering in a basaltic soil and pore water profile in Iceland. Earth and PlanetaryScience Letters 339–340, 11–23.

Poitrasson, F., Freydier, R., 2005. Heavy iron isotope composition of granites determinedby high resolution MC-ICP-MS. Chemical Geology 222, 123–147.

Poitrasson, F., Halliday, A.N., Lee, D.-C., Levasseur, S., Teutsch, N., 2004. Iron isotopedifferences between Earth, Moon, Mars and Vesta as possible records of contrastedaccretion mechanisms. Earth and Planetary Science Letters 223, 253–266.

Potts, P.J., 1987. A Handbook of Silicate Rock Analysis. Blackie, Glasgow.Pringle, E.A., Savage, P.S., Badro, J., Barrat, J.-A., Moynier, F., 2013a. Redox state during core

formation on asteroid 4-Vesta. Earth and Planetary Science Letters 373, 75–82.Pringle, E.A., Savage, P.S., Jackson, M.G., Barrat, J.-A., Moynier, F., 2013b. Si isotope

homogeneity of the solar nebula. The Astrophysical Journal 779, 123.Raymond, S.N., O'Brien, D.P., Morbidelli, A., Kaib, N.A., 2009. Building the terrestrial

planets: constrained accretion in the inner Solar System. Icarus 203, 644–662.Regelous, M., Elliott, T., Coath, C.D., 2008. Nickel isotope heterogeneity in the early Solar

System. Earth and Planetary Science Letters 272, 330–338.Reufer, A., Meier, M.M.M., Benz, W., Wieler, R., 2012. A hit-and-run giant impact scenario.

Icarus 221, 296–299.Reynolds, J.H., Verhoogen, J., 1953. Natural variations in the isotopic constitution of

silicon. Geochimica et Cosmochimica Acta 3, 224–234.Reynolds, B.C., Frank, M., Halliday, A.N., 2006. Silicon isotope fractionation during nutrient

utilization in the North Pacific. Earth and Planetary Science Letters 244, 431–443.Reynolds, B.C., Aggarwal, J., André, L., Baxter, D., Beucher, C., Brzezinski, M.A., Engström, E.,

Georg, R.B., Land, M., Leng, M.J., Opfergelt, S., Rodushkin, I., Sloane, H.J., Van DenBoorn, S.H.J.M., Vroon, P.Z., Cardinal, D., 2007. An inter-laboratory comparison of Siisotope reference materials. Journal of Analytical Atomic Spectrometry 22, 561–568.

Richter, F.M., Liang, Y., Davis, A.M., 1999. Isotope fractionation by diffusion in moltenoxides. Geochimica et Cosmochimica Acta 63, 2853–2861.

Richter, F.M., Davis, A.M., Ebel, D.S., Hashimoto, A., 2002. Elemental and isotopic fraction-ation of type B calcium-, aluminum-rich inclusions: experiments, theoretical consid-erations, and constraints on their thermal evolution. Geochimica et CosmochimicaActa 66, 521–540.

Richter, F.M., Davis, A.M., DePaolo, D.J., Watson, E.B., 2003. Isotope fractionation bychemical diffusion between molten basalt and rhyolite. Geochimica et CosmochimicaActa 67, 3905–3923.

Richter, F.M., Dauphas, N., Teng, F.-Z., 2009a. Non-traditional fractionation of non-traditional isotopes: evaporation, chemical diffusion and Soret diffusion. ChemicalGeology 258, 92–103.

Richter, F.M., Watson, E.B., Mendybaev, R., Dauphas, N., Georg, B., Watkins, J., Valley, J.,2009b. Isotopic fractionation of the major elements of molten basalt by chemicaland thermal diffusion. Geochimica et Cosmochimica Acta 73, 4250–4263.

Ricolleau, A., Fei, Y., Corgne, A., Siebert, J., Badro, J., 2011. Oxygen and silicon contents ofEarth's core from high pressure metal–silicate partitioning experiments. Earth andPlanetary Science Letters 310, 409–421.

Robert, F., Chaussidon, M., 2006. A palaeotemperature curve for the Precambrian oceansbased on silicon isotopes in cherts. Nature 443, 969–972.

Rudge, J.F., Kleine, T., Bourdon, B., 2010. Broad bounds on Earth's accretion and coreformation constrained by geochemical models. Nature Geosciences 3, 439–443.

Rudnick, R.L., 1995. Making continental crust. Nature 378, 571–578.Rudnick, R.L., Gao, S., 2003. Composition of the continental crust. In: Holland, H.D.,

Turekian, K.K. (Eds.), Treatise on Geochemistry. Pergamon, Oxford, pp. 1–64.Rudnick, R.L., Taylor, S.R., 1987. The composition and petrogenesis of the lower crust:

a xenolith study. Journal of Geophysical Research 92, 13981–14005.Rudnick, R.L., McDonough, W.F., McCulloch, M.T., Taylor, S.R., 1986. Lower crustal xeno-

liths from Queensland, Australia: evidence for deep crustal assimilation and fraction-ation of continental basalts. Geochimica et Cosmochimica Acta 50, 1099–1115.

Savage, P.S.,Moynier, F., 2013. Silicon isotopic variation in enstatitemeteorites: clues to theirorigin and Earth-forming material. Earth and Planetary Science Letters 361, 487–496.

Savage, P.S., Georg, R.B., Armytage, R.M.G., Williams, H.M., Halliday, A.N., 2010. Siliconisotope homogeneity in the mantle. Earth and Planetary Science Letters 295,139–146.





































































































































































Savage, P.S., Georg, R.B., Williams, H.M., Burton, K.W., Halliday, A.N., 2011. Silicon isotopefractionation during magmatic differentiation. Geochimica et Cosmochimica Acta 75,6124–6139.

Savage, P.S., Georg, R.B.,Williams, H.M., Turner, S., Halliday, A.N., Chappell, B.W., 2012. Thesilicon isotope composition of granites. Geochimica et Cosmochimica Acta 92,184–202.

Savage, P.S., Georg, R.B., Williams, H.M., Halliday, A.N., 2013a. Silicon isotopes in granulitexenoliths: insights into isotopic fractionation during igneous processes and thecomposition of the deep continental crust. Earth and Planetary Science Letters 365,221–231.

Savage, P.S., Georg, R.B., Williams, H.M., Halliday, A.N., 2013b. The silicon isotopecomposition of the upper continental crust. Geochimica et Cosmochimica Acta 109,384–399.

Schiller, M., Handler, M.R., Baker, J.A., 2010. High-precision Mg isotopic systematics ofbulk chondrites. Earth and Planetary Science Letters 297, 165–173.

Schuessler, J.A., Schoenberg, R., Sigmarsson, O., 2009. Iron and lithium isotope systematicsof the Hekla volcano, Iceland — evidence for Fe isotope fractionation during magmadifferentiation. Chemical Geology 258, 78–91.

Seitz, H.-M., Brey, G.P., Zipfel, J., Ott, U., Weyer, S., Durali, S., Weinbruch, S., 2007. Lithiumisotope composition of ordinary and carbonaceous chondrites, and differentiatedplanetary bodies: bulk solar system and solar reservoirs. Earth and Planetary ScienceLetters 260, 582–596.

Shahar, A., Young, E.D., 2007. Astrophysics of CAI formation as revealed by silicon isotopeLA–MC-ICPMS of an igneous CAI. Earth and Planetary Science Letters 257, 497–510.

Shahar, A., Ziegler, K., Young, E.D., Ricolleau, A., Schauble, E.A., Fei, Y., 2009. Experimentallydetermined Si isotope fractionation between silicate and Fe metal and implicationsfor Earth's core formation. Earth and Planetary Science Letters 288, 228–234.

Shahar, A., Hillgren, V.J., Young, E.D., Fei, Y., Macris, C.A., Deng, L., 2011. High-temperatureSi isotope fractionation between ironmetal and silicate. Geochimica et CosmochimicaActa 75, 7688–7697.

Sigmarsson, O., Condomines, M., Fourcade, S., 1992. A detailed Th, Sr and O isotope studyof Hekla: differentiation processes in an Icelandic Volcano. Contributions to Mineralogyand Petrology 112, 20–34.

Spicuzza, M.J., Day, J.M.D., Taylor, L.A., Valley, J.W., 2007. Oxygen isotope constraints onthe origin and differentiation of the Moon. Earth and Planetary Science Letters 253,254–265.

Steinhoefel, G., Breuer, J., von Blanckenburg, F., Horn, I., Kaczorek, D., Sommer, M., 2011.Micrometer silicon isotope diagnostics of soils by UV femtosecond laser ablation.Chemical Geology 286, 280–289.

Teng, F.-Z., Wadhwa, M., Helz, R.T., 2007. Investigation of magnesium isotope fraction-ation during basalt differentiation: implications for a chondritic composition of theterrestrial mantle. Earth and Planetary Science Letters 261, 84–92.

Teng, F.-Z., Dauphas, N., Helz, R.T., 2008. Iron isotope fractionation during magmaticdifferentiation in Kilauea Iki lava lake. Science 320, 1620–1622.

Teng, F.-Z., Li, W.-Y., Ke, S., Marty, B., Dauphas, N., Huang, S., Wu, F.-Y., Pourmand, A., 2010.Magnesium isotopic composition of the Earth and chondrites. Geochimica etCosmochimica Acta 74, 4150–4166.

Tomascak, P.B., Tera, F., Helz, R.T., Walker, R.J., 1999. The absence of lithium isotopefractionation during basalt differentiation: newmeasurements by multicollector sectorICP-MS. Geochimica et Cosmochimica Acta 63, 907–910.

Touboul, M., Kleine, T., Bourdon, B., Palme, H., Wieler, R., 2007. Late formationand prolonged differentiation of the Moon inferred from W isotopes in lunar metals.Nature 450, 1206–1209.

Tréguer, P.J., De La Rocha, C.L., 2013. The world ocean silica cycle. Annual Review ofMarine Science 5, 477–501.

Trinquier, A., Birck, J.-L., Allègre, C.J., 2007. Widespread 54Cr heterogeneity in the innersolar system. Astrophysical Journal Letters 655, 1179–1185.

Trinquier, A., Elliott, T., Ulfbeck, D., Coath, C., Krot, A.N., Bizzarro, M., 2009. Origin of nucle-osynthetic isotope heterogeneity in the solar protoplanetary disk. Science 324,374–376.

Unruh, D.M., Stille, P., Patchett, P.J., Tatsumoto, M., 1984. Lu-Hf and Sm-Nd evolution inlunar mare basalts. Journal of Geophysical Research: Solid Earth 89 (2), 459–477.

van Den Boorn, S.H.J.M., Vroon, P.Z., Van Belle, C.C., Van Der Wagt, B., Schwieters, J.,Van Bergen, M.J., 2006. Determination of silicon isotope ratios in silicate materialsby high-resolution MC-ICP-MS using a sodium hydroxide sample digestion method.Journal of Analytical Atomic Spectrometry 21, 734–742.

van den Boorn, S.H.J.M., van Bergen, M.J., Nijman, W., Vroon, P.Z., 2007. Dual role ofseawater and hydrothermal fluids in Early Archean chert formation: evidence fromsilicon isotopes. Geology 35, 939–942.

van Den Boorn, S.H.J.M., Vroon, P.Z., Van Bergen, M.J., 2009. Sulfur-induced offsets in MC-ICP-MS silicon-isotope measurements. Journal of Analytical Atomic Spectrometry 24,1111–1114.

Wade, J., Wood, B.J., 2005. Core formation and the oxidation state of the Earth. Earth andPlanetary Science Letters 236, 78–95.

Walder, A.J., Freedman, P.A., 1992. Isotopic ratio measurement using a double focusingmagnetic sector mass analyser with an inductively coupled plasma as an ion source.Journal of Analytical Atomic Spectrometry 7, 571–575.

Wang, J., Davis, A., Clayton, R., Hashimoto, A., 1999. Evaporation of single crystalforsterite: evaporation kinetics, magnesium isotope fractionation, and implicationsof mass-dependent isotopic fractionation of a diffusion-controlled reservoir.Geochimica et Cosmochimica Acta 63, 953–966.

Wang, J., Davis, A.M., Clayton, R.N., Mayeda, T.K., Hashimoto, A., 2001. Chemical and iso-topic fractionation during the evaporation of the FeO–MgO–SiO2–CaO–Al2O3–TiO2

rare earth element melt system. Geochimica et Cosmochimica Acta 65, 479–494.Wänke, H., Dreibus, G., 1988. Chemical composition and accretion history of terrestrial

planets. Philosophical Transactions of the Royal Society of London. Series A, Mathe-matical and Physical Sciences 325, 545–557.

Watkins, J.M., DePaolo, D.J., Ryerson, F.J., Peterson, B.T., 2011. Influence of liquid structureon diffusive isotope separation in molten silicates and aqueous solutions. Geochimicaet Cosmochimica Acta 75, 3103–3118.

Wedepohl, K.H., 1995. The composition of the continental crust. Geochimica etCosmochimica Acta 59, 1217–1232.

Weyer, S., Schwieters, J.B., 2003. High precision Fe isotope measurements with high massresolution MC-ICPMS. International Journal of Mass Spectrometry 226, 355–368.

Wiechert, U., Halliday, A.N., 2007. Non-chondritic magnesium and the origins of the innerterrestrial planets. Earth and Planetary Science Letters 256, 360–371.

Wiechert, U., Halliday, A.N., Lee, D.-C., Snyder, G.A., Taylor, L.A., Rumble, D., 2001. Oxygenisotopes and the Moon-forming giant impact. Science 294, 345–348.

Young, E.D., Galy, A., Nagahara, H., 2002. Kinetic and equilibriummass-dependent isotopefractionation laws in nature and their geochemical and cosmochemical significance.Geochimica et Cosmochimica Acta 66, 1095–1104.

Zambardi, T., Poitrasson, F., 2011a. Precise determination of silicon isotopes in silicate rockreferencematerials byMC-ICP-MS. Geostandards andGeoanalytical Research 35, 89–99.

Zambardi, T., Poitrasson, F., 2011b. A combined Earth–Moon Si isotopes model to trackrocks petrogenesis. Mineralogical Magazine 75, 2243.

Zambardi, T., Poitrasson, F., Corgne, A., Méheut, M., Quitté, G., Anand, M., 2013. Siliconisotope variations in the inner solar system: implications for planetary formation,differentiation and composition. Geochimica et Cosmochimica Acta 121, 67–83.

Zhang, J., Dauphas, N., Davis, A.M., Leya, I., Fedkin, A., 2012. The proto-Earth as a signifi-cant source of lunar material. Nature Geoscience 5, 251–255.

Zhu, X., Guo, Y., O'Nions, R.K., Young, E., Ash, R., 2001. Isotopic homogeneity of iron in theearly solar nebula. Nature 412, 311–313.

Ziegler, K., Chadwick, O.A., Brzezinski, M.A., Kelly, E.F., 2005a. Natural variations of δ30Siratios during progressive basalt weathering, Hawaiian Islands. Geochimica etCosmochimica Acta 69, 4597–4610.

Ziegler, K., Chadwick, O.A., White, A.F., Brzezinski, M.A., 2005b. δ30Si systematics in a gra-nitic saprolite, Puerto Rico. Geology 33, 817–820.

Ziegler, K., Young, E.D., Schauble, E.A., Wasson, J.T., 2010. Metal–silicate silicon isotopefractionation in enstatite meteorites and constraints on Earth's core formation.Earth and Planetary Science Letters 295, 487–496.




























































































































high temperature silicon isotope geochemistry

Documents