Iron, zinc, magnesium and uranium isotopic fractionation ... - CiteSeerX

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Geochimica et Cosmochimica Acta 97 (2012) 247–265 www.elsevier.com/locate/gca

Iron, zinc, magnesium and uranium isotopic fractionation during continental crust differentiation: The tale from migmatites, granitoids, and pegmatites Myriam Telus a,⇑, Nicolas Dauphas a, Fre´de´ric Moynier b, Francßois L.H. Tissot a, Fang-Zhen Teng c, Peter I. Nabelek d, Paul R. Craddock a, Lee A. Groat e a b

Origins Laboratory, Department of the Geophysical Sciences and Enrico Fermi Institute, The University of Chicago, Chicago, IL 60637, USA Department of Earth and Planetary Sciences and McDonnell Center for the Space Sciences, Washington University in St. Louis, St. Louis, MO 63130, USA c Department of Geosciences & Arkansas Center for Space and Planetary Sciences, University of Arkansas, Fayetteville, AR 72701, USA d Department of Geological Sciences, University of Missouri-Columbia, Columbia, MO 65211, USA e Department of Earth and Ocean Sciences, University of British Columbia, Vancouver, BC, Canada V6T 1Z4 Received 27 September 2011; accepted in revised form 14 August 2012; Available online 25 August 2012

Abstract The causes of some stable isotopic variations in felsic rocks are not well understood. In particular, the origin of the heavy Fe isotopic compositions (i.e., high d56Fe values, deviation in & of the 56Fe/54Fe ratio relative to IRMM-014) of granites with SiO2 > 70 wt.% compared with less silicic rocks is still debated. It has been interpreted to reflect isotopic fractionation during late stage aqueous fluid exsolution, magma differentiation, partial melting, or Soret (thermal) diffusion. The present study addresses this issue by comparing the Fe isotopic compositions of a large range of differentiated crustal rocks (whole rocks of migmatites, granitoids, and pegmatites; mineral separates) with the isotopic compositions of Zn, Mg and U. The samples include granites, migmatites and pegmatites from the Black Hills, South Dakota (USA), as well as I-, S-, and A-type granitoids from Lachlan Fold Belt (Australia). The nature of the protolith (i.e., I- or S-type) does not influence the Fe isotopic composition of granitoids. Leucosomes (partial melts in migmatites) tend to have higher d56Fe values than melanosomes (melt residues) indicating that partial melting of continental crust material can possibly fractionate Fe isotopes. No clear positive correlation is found between the isotopic compositions of Mg, U and Fe, which rules out the process of Soret diffusion in the systems studied here. Zinc isotopes were measured to trace fluid exsolution because Zn can easily be mobilized by aqueous fluids as chloride complexes. Pegmatites and some granitic rocks with high d56Fe values also have high d66Zn values. In addition, high-SiO2 granites show a large dispersion in the Zn/Fe ratio that cannot easily be explained by magma differentiation alone. These results suggest that fluid exsolution is responsible for some of the Fe isotopic fractionation documented in felsic rocks and in particular in pegmatites. However, some granites with high d56Fe values have unfractionated d66Zn values and were presumably poor in fluids (e.g., A-type). For these samples, iron isotopic fractionation during magma differentiation is a viable interpretation. Equilibrium Fe isotopic fractionation factors between silicic melts and minerals remain to be characterized to quantitatively assess the role of fractional crystallization on iron isotopes in granitoids. Ó 2012 Elsevier Ltd. All rights reserved.

⇑ Corresponding author. Present Address: Hawai‘i Institute of

Geophysics and Planetology, School of Ocean and Earth Science and Technology, University of Hawai‘i at Ma¯noa, Honolulu, HI 96822, USA. Tel.: +1 239 247 0122; fax: +1 808 956 6322. E-mail address: [email protected] (M. Telus). 0016-7037/$ - see front matter Ó 2012 Elsevier Ltd. All rights reserved. http://dx.doi.org/10.1016/j.gca.2012.08.024

1. INTRODUCTION With its complex redox chemistry, iron is a key element in biological and geochemical processes and its isotopic variations can be used to probe these interactions. Initial

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studies found all terrestrial igneous rocks to have homogeneous Fe isotopic composition (Beard and Johnson, 1999, 2004; Beard et al., 2003), suggesting that Fe isotopes could be discriminate indicators of biological activity and low temperature aqueous processes, which can significantly fractionate Fe isotopes (Johnson et al., 2004; Dauphas and Rouxel, 2006; Anbar and Rouxel, 2007). However, growing evidence suggests that igneous processes such as partial melting, magmatic differentiation, and low temperature aqueous fluid exsolution can also impart measurable iron isotopic variations among bulk igneous rocks (Williams et al., 2004, 2005, 2009; Poitrasson and Freydier, 2005; Weyer et al., 2005; Dauphas and Rouxel, 2006; Poitrasson, 2006; Schoenberg and von Blanckenburg, 2006; Schuessler et al., 2007, 2009; Weyer and Ionov, 2007; Heimann et al., 2008; Teng et al., 2008; Dauphas et al., 2009, 2010). This opens the prospect of using iron isotopes as petrogenetic tracers of mantle and crustal differentiation. Poitrasson and Freydier (2005) observed that several granites were isotopically heavier than the average d56Fe composition of igneous rocks and identified a positive correlation between d56Fe and SiO2 wt.%. They suggested that fluids exsolved from the magma were responsible for removing isotopically light Fe from the crystallizing melt, resulting in a heavy granitic magma body. Heimann et al. (2008) found that magnetite from evolved granitic rocks had high d56Fe values while Fe-silicates had d56Fe values close to the average d56Fe composition of igneous rocks, which they attributed to isotopic exchange between magnetite and fluid rich in ferrous chloride. Thus, these authors also concluded that fluid exsolution was responsible for high d56Fe values in evolved granites and rhyolites. Schuessler et al. (2009) found similar iron isotopic trends for Fe isotopes, but did not detect measurable variation in the isotopic composition of Li, which is easily mobilized by fluids. Based on the absence of a correlation between Li and Fe isotopic compositions, they ascribed these trends to fractional crystallization of a Fe-bearing phase that preferentially incorporated light Fe isotopes; thus, continually enriching the residual melt in the heavier isotopes of iron. Similar results were found for a basaltic system at Kilauea Iki lava lake (Teng et al., 2008). In the latter study, the light Fe isotopic composition reflected diffusion-driven kinetic isotope fractionation during Fe–Mg exchange in olivine (Teng et al., 2011; Sio et al., 2011). The present study addresses the range of hypotheses put forth to explain these iron isotope variations by combining Fe, Zn, Mg and U isotopic data of migmatites, granitoids, and pegmatites to develop a more complete view of Fe isotopic fractionation during the formation and differentiation of felsic rocks. Particular attention is directed towards the following questions: Does partial melting fractionate iron isotopes? This has been documented in mantle peridotites and arc lavas through correlations between d56Fe values and proxies of degree of partial melting (Weyer and Ionov, 2007; Dauphas et al., 2009). It is not known however whether iron isotopes are fractionated in the process of generating felsic melts. Migmatites may provide the opportunity to directly answer this question, as they represent highly metamorphosed

rocks arrested in the process of generating granitic melts (Brown, 1994). Does the nature of the protolith influence the iron isotopic composition of felsic rocks? It was suggested that during partial melting of the mantle, the iron isotopic composition of magmas was related to the degree of partial melting, the Fe3+/Fe2+ ratio, and oxygen buffering capacity (Dauphas et al., 2009). Is that true for felsic rocks? Do granites that were derived from sources with elevated Fe3+/Fe2+ ratios have higher d56Fe? To address this question, a comparison is made between S- and I-type granitic rocks (Chappell and White, 2001) that are derived from sedimentary and igneous sources, which differ in their mineral assemblages and Fe3+/Fe2+ ratios (Flood and Shaw, 1975; Whalen and Chappell, 1988). Is magmatic differentiation (fractional crystallization and restite unmixing), fluid exsolution or Soret (thermal) diffusion responsible for the correlation between iron isotopic composition and silica content? Magmatic differentiation should be supported by fractional crystallization models (Schoenberg and von Blanckenburg, 2006). The fluid exsolution hypothesis (Poitrasson and Freydier, 2005; Heimann et al., 2008) requires that significant amounts of isotopically light iron be removed from the magma body. Pegmatites, which represent the latest stages of granitic magma evolution, should be particularly sensitive to such processes (Walker et al., 1989). Another means of testing the fluid exsolution hypothesis is to study the isotopic compositions of elements, such as Zn, that are mobile in such fluids (Nakano and Urabe, 1989; Shinohara, 1994), and to search for correlations with iron isotopic variations. This is tested by comparing the Zn and Fe isotopic compositions of felsic rocks. Finally, an alternative model to the standard view of granite formation was presented by Lundstrom (2009), who suggested that some granites form by a top-down thermal migration process. The model predicts that non-traditional stable isotope systems such as Fe, Mg and U should show signatures of thermal diffusion. To address this hypothesis, the isotopic compositions of Mg, U and Fe were measured in the same granitic rocks. Magnesium, U and Fe isotopes show strong and correlated mass dependent fractionation during experiments of thermal diffusion in silicate melts (Richter et al., 2008, 2009; Huang et al., 2010; Lacks et al., 2012), thus providing a straightforward means of identifying Soret effect in geological processes (Dauphas et al., 2010). 2. SAMPLE SELECTION 2.1. I-, S- & A- type granites: bulk rocks and mineral separates To determine whether the Fe isotopic compositions of granites depend on their source material, granitic rocks from the Silurian-Devonian Lachlan Fold Belt (LFB, Southeastern Australia) and Proterozoic Harney Peak Granite (Black Hills, South Dakota) were analyzed for their Fe isotopic compositions. These samples have been well characterized (Hine et al., 1978; Chappell, 1984; Nabelek

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et al., 1992; Chappell and White, 1993; Sha and Chappell, 1999) and are divided into three groups according to their source rock (Chappell and White, 1974; Loiselle and Wones, 1979). I-type granites (nine samples; I1, I2, I3, I4, I5, I6, I7, I8, and I9) are derived from partial melting of infracrustal metaigneous rocks. They are hornblende-bearing and have relatively high Ca and Na contents. S-type granites (11 samples; S1, S2, S3, S4, S5, S6, S7, S8, S9, HP49-A, and HP21-C) are derived from partial melting of supracrustal sedimentary source material that has been exposed to chemical weathering. They are generally more felsic than I-types and have low Ca and Na content relative to I-type granites due to aqueous removal of these elements during weathering of feldspars into clay minerals (Chappell and White, 2001). Another important difference between Iand S- type granites is that they tend to differ in their Fe3+/ Fe2+ ratios. S-type granites generally have lower Fe3+/ (Fe3++Fe2+) ratios (0.16) than I-types (0.30) (Whalen and Chappell, 1988). The metasedimentary sources of Stype granites are generally rich in graphite (Flood and Shaw, 1975), which explains the more reducing conditions of their formation (French and Eugster, 1965). Classification of A-type granite (two samples; A1 and A2) is more complex as it also involves considerations of the tectonic setting (Loiselle and Wones, 1979). Whalen et al. (1987), Eby (1990) and Bonin (2007) provide detailed descriptions of the chemical and petrographic characteristics of A-type granites. They form in anorogenic environments, often in tensional regimes, are mildly alkaline, and their parental magmas are inferred to have been relatively anhydrous. There is considerable uncertainty regarding their origin and a unique petrogenetic model may not be appropriate. Some of the major models that have been proposed include melting of relatively anhydrous granulitic meta-igneous restite left from a prior episode of melt extraction (Collins et al., 1982; Clemens et al., 1986; Whalen et al., 1987), melting of granulitic crust metasomatized by mantle-derived alkaline fluids (Martin, 2006), derivation from basaltic magma by magmatic differentiation (Loiselle and Wones, 1979; Turner et al., 1992), and partial melting of crustal tonalitic-granodioritic rocks (Anderson, 1983; Creaser et al., 1991). The chemical compositions of most samples studied here are available in the literature (Hine et al., 1978; Chappell, 1984; Chappell and White, 1992, 1993; Nabelek et al., 1992; King et al., 1997; Sha and Chappell, 1999). However, when not available (four out of 22 samples studied; I4, I6, I7 and S7), the chemical compositions were measured at Service d’Analyse des Roches et des Mine´raux, Nancy, France, using established protocols (Carignan et al., 2001). The SiO2 concentrations of bulk samples range from 56 to 77 wt.%. In order to better understand the processes that control Fe isotopic fractionation in felsic rocks, mineral separates (magnetite, biotite and hornblende) from LFB granitoids (samples I1, I2, I3, I5 and A2) were also measured. 2.2. Migmatites and pegmatites Migmatites are generally found in high-grade metamorphic terrains and are made up of leucosomes (L) that are

249

dominated by large quartz crystals and are thought to represent partial melts, and melanosomes (M) that are dominated by biotite and sillimanite and are thought to represent the residues of partial melting. Although not sampled for this study, migmatites also consist of non-migmatized (non-separated) portions, called mesosomes, which consist of varying portions of quartz, biotite, sillimanite and plagioclase. Schists, which are compositionally very similar to mesosomes, were analyzed for this study. These schists are associated with both migmatites and granites of the Black Hills, South Dakota (Nabelek, 1999). Migmatites are important for studying melt segregation and processes governing granite genesis and differentiation. They represent rocks in which melt has segregated on a local scale but has not escaped from the system. Melting in migmatites begins at grain boundaries and as melting progresses, melt pockets combine to form a network of melt channels, which subsequently can migrate to produce granitic bodies (Sawyer, 1991, 1994; Brown, 1994, 2002). A total of 11 migmatite components (five leucosome– melanosome pairs and one unpaired leucosome) and two schists from the Black Hills region of South Dakota were studied. Melanosomes are very similar in chemical composition, while leucosomes show more variations (e.g., 115-1L and 131-1-L have very high concentrations of K2O and Na2O and 118-1-L has high concentrations of CaO and P2O5 compared to other leucosomes). The bulk SiO2 content ranges from 72 to 79 wt.% for leucosomes, from 47 to 62 wt.% for melanosomes and from 63 to 69 wt.% for schists. The FeO concentrations range from 0.3 to 0.6 wt.% for leucosomes, and from 8 to 17 wt.% for melanosomes. Schists fall in between with 5.5 wt.% FeO. Pegmatites are coarse-grained rocks, typically associated spatially and compositionally with granites. Large quartz, feldspar and muscovite crystals make up the bulk of most pegmatites. They have been interpreted as the products of extreme crystal fractionation of granitic magmas (London, 1996, 2005). They provide insight to the final stages of granite differentiation, which is heavily influenced by fluid exsolution processes (Norton and Redden, 1990). Such processes have been suggested as the major mechanism for producing variations in the Fe isotopic composition of evolved felsic rocks (Poitrasson and Freydier, 2005; Heimann et al., 2008). To address the influence of fluid exsolution on the Fe isotopic composition, we studied four samples from simple pegmatites (81BH 5-1, 81BH 6-3, 81BH 6-4, WC-9) and three wall zone samples from the more complex, zoned, Tin Mountain pegmatite (81BH9-2, 81BH10-3, 82BH431), which are closely associated (12 km southwest) with the Proterozoic Harney Peak Granite and migmatites from the Black Hills (Walker, 1984; Teng et al., 2006). The Tin Mountain pegmatite consists of five zones (wall zone, first intermediate zone, second intermediate zone, third intermediate zone, and core), which are characterized by their distinct mineral assemblages, with the wall zone crystallizing first. Oxygen isotopic thermometry yields crystallization temperatures ranging from >600 °C in the outer regions to 500 °C in the core (Walker et al., 1986), while some temperatures as low as 350 °C have been reported (Sirbescu

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and Nabelek, 2003). Trace-element data (Rb, Li, Sr, and Ba) suggest two possible origins for the Tin Mountain Pegmatite: low degree partial melting of metasedimentary rocks that experienced moderate fractional or equilibrium crystallization, or derivation from the Harney Peak Granite through multiple stages of crystal-liquid fractional crystallization (Walker et al., 1989). Pegmatites from Little Nahanni, Logan Mountains, Northwest Territories, Canada (ca. 82 Ma) were also studied (six samples; P389704, P389910, P389913, P389726, P389739, and P389734). They have been characterized in detail by Groat et al. (2003). These samples are rare-element pegmatites, which generally have lower crystallization temperatures compared to simple and zoned pegmatites. The nearest granitic intrusion is 4.7 km2, 84–87 Myr and 15 km away. Rare-element granitic pegmatites typically contain high concentrations of H2O, Li, B, F, and P. They solidify under chemical and physical conditions that differ from predictable paths of crystallization for normal felsic melts. The mechanism for the formation of these pegmatites is not well understood, but extreme fractional crystallizaˇ erny´, 1985, 1991). tion is the preferred interpretation (C Chemical evolution of these pegmatites shows systematic trends of Mg and Fe decreasing, while Al, Mn, Rb and Cs increase from wall to core (Groat et al., 2003). Little Nahanni pegmatites are divided into spodumene-bearing and spodumene-free, but apart from this distinction, they are similar in their major element composition to each other and to pegmatites from the Black Hills. The SiO2 concentrations of all pegmatites measured for this study range from 68 to 74 wt.% and FeO concentrations range from 0.3 to 2.2 wt.%. The major element compositions for all samples analyzed in this study are compiled in Table 1. A total alkali vs. SiO2 diagram (Fig. 1) provides an illustration of the compositional diversity of the samples studied. 3. ANALYTICAL METHODS Samples were crushed to a fine powder in an agate mortar. Mafic minerals from samples I1, I2, I3, I5 and A2 were separated using a combination of magnetic separation, manual extraction and density separation for Fe isotopic analysis. Mineral mode compositions were estimated by analyzing thin sections of these rocks with an optical microscope. A chart was used as a visual aid to help determine the percentages of the Fe-bearing minerals in each thin section. The modal abundances range from 1–3& for magnetite, 6–20% for biotite and 10–18% for hornblende. These numbers are rough estimates. The samples were crushed, sieved under water, and fractions between 125 and 250 lm were collected. Magnetite was separated by covering a hand magnet with weighing paper and rastering it across the grains. A Frantz magnetic separator was used on each sample to separate mafic from non-mafic mineral fractions. Biotite was then separated from the mafic minerals by simply pouring out the sample on a sheet of paper and continuously tapping and rotating the paper at various angles until biotite was all that remained since its sheet silicate crystal structure allows it to stick to the sheet of paper.

Pure hornblende was collected through heavy liquid separation and was then handpicked from the high-density fraction under an optical microscope. 3.1. Sample preparation and instrument analysis The samples were prepared for iron isotope analyses according to the procedure described in Dauphas et al. (2004, 2009). Sizes of the sample powder aliquots ranged from 10 to 20 mg for geostandards, from 20 to 30 mg for plutonic rocks, from 21 to 35 mg for migmatites, from 18 to 70 mg for pegmatites, and from 2 to 9 mg for the mineral separates. Blanks and sets of geostandards were measured together with samples. The blank contribution (10 lg). Isotope ratios are expressed as relative deviations, in d-notation, from the Zn JMC-Lyon standard as: h i di Zn ¼ ði Zn=64 ZnÞsample =ði Zn=64 ZnÞJMCLyon  1  103 ;

Table 1 Iron isotope and major element compositions of geostandards, granitoids, pegmatites and migmatites. Sample type

Name

d56Fe(&)

d57Fe(&)

Averages of replicate analyzes*

Major element compositions (wt.%)

d56Fe(&)

SiO2

Al2O3

Fe2O3

FeO

MnO

MgO

CaO

Na2O

d57Fe(&)

P2O5

H2O+

H2O

CO2

References

1.41

0.99

0.28

1.03

0.18

0.28

2.74

3.43

0.58

0.19

1.26

0.09

0.20

4.32

2.63

2.99

0.53

0.11

0.95

0.13

0.08

0.08 2.85

0.38 6.05

3.63 2.60

4.71 2.01

0.06 0.54

0.00 0.11

1.20

0.20

0.10

0.15 0.06 0.07

8.20 0.59 2.40

7.18 2.56 3.74

2.04 4.81 2.65

2.09 2.57 4.08

0.84 0.42 0.88

0.24 0.10 0.24

1.02

0.20

0.32

3.40

0.09

2.19

4.72

2.46

2.74

0.59

0.11

0.68

0.16

0.31

0.59 0.81

3.48 4.02

0.06 0.11

1.76 2.13

0.95 2.37

1.49 1.80

3.73 3.60

0.54 0.68

0.13 0.15

1.28 1.98

0.14 0.26

0.13 0.13

14.70 14.36 14.25 13.41 14.00 13.08

0.68 0.81 0.93 0.44 1.60 0.16

4.03 3.36 3.10 1.47 0.98 2.27

0.07 0.06 0.07 0.04 0.03 0.05

2.22 1.95 1.68 0.71 0.32 0.61

2.26 1.97 3.11 1.70 0.80 1.20

1.92 1.99 2.23 2.63 3.17 2.58

3.60 3.94 3.77 4.63 4.36 4.60

0.64 0.57 0.53 0.22 0.17 0.36

0.15 0.16 0.13 0.08 0.27 0.17

1.51 1.65 1.28 0.74

0.22 0.23 0.12 0.14

0.12 0.12 0.29 0.11

0.90

0.16

0.09

68.05

14.41

0.86

3.57

0.06

2.05

2.68

2.03

3.58

0.61

0.15

1.21

0.15

0.24

0.227 ± 0.034 0.467 ± 0.032 0.270 ± 0.055

73.60 77.00 72.40

12.72 11.83 16.10

1.99 0.40 –

0.82 1.05 1.31

0.04 0.04 0.13

0.20 0.04 0.15

0.83 0.61 0.57

3.62 3.06 7.31

4.11 4.98 0.11

0.38 0.15 0.02

0.08 0.02 0.22

0.73 0.47

0.29 0.18

0.29 0.07

Sha and Chappell (1999) Sha and Chappell (1999) Sha and Chappell (1999) This study Sha and Chappell (1999) This study This study Chappell and White, 1992 Chappell and White (1992) Chappell (1984) Chappell and White (1993) Chappell (1984) Hine et al. (1978) Chappell (1984) Chappell (1984) This study Chappell and White (1992) Sha and Chappell (1999) King et al. (1997) King et al. (1997) Nabelek et al. (1992)

0.240 ± 0.041

0.368 ± 0.047

75.10

13.50

–

1.52

0.02

0.41

0.76

2.81

3.96

0.16

0.18

Nabelek et al. (1992)

81BH 5-1 81BH 6-3 81BH 6-4 WC-9

0.270 ± 0.054 0.151 ± 0.054 0.290 ± 0.054 0.209 ± 0.054

0.330 ± 0.105 0.262 ± 0.105 0.377 ± 0.105 0.305 ± 0.105

74.45 72.21 71.99 73.34

15.85 9.40 15.10 12.60

1.52 1.03 0.12 0.40

– – – 0.01

0.04 0.13 0.00 0.00

0.45 0.00 0.38 0.36

0.56 0.40 0.09 2.83

5.75 – 2.42 6.90

1.39 4.08 10.10 0.06

0.06 0.01 – 0.50

0.09 – –

Walker Walker Walker Walker

81BH 9-2 81BH 10-3 82BH 43-1

0.197 ± 0.054 0.221 ± 0.054 0.385 ± 0.054

0.268 ± 0.105 0.375 ± 0.105 0.562 ± 0.105

74.40 68.32 70.78

13.54 17.54 15.67

– – –

2.20 1.54 0.51

0.05 0.26 0.05

0.02 0.01 0.01

0.64 0.93 0.61

2.68 9.13 8.01

2.22 0.39 2.01

0.02 0.02 0.08

0.56 1.12 0.50

Walker et al. (1989) Walker et al. (1989) Walker et al. (1989)

P389734 P389910

0.189 ± 0.039 0.061 ± 0.039

0.285 ± 0.060 0.101 ± 0.059

72.31 70.76

16.73 17.56

0.27 0.28

0.26 0.26

0.23 0.15

0 0.01

0.08 0.61

3.24 4.77

3.92 3.58

0.45 0.64

1.03 0.78

I1

0.123 ± 0.020

0.184 ± 0.026

60.13

16.42

2.04

3.73

0.11

3.27

5.91

3.94

Granodiorite, Yurammie

I2

0.081 ± 0.021

0.120 ± 0.038

65.27

15.22

1.76

2.83

0.09

2.01

4.40

Granodiorite, Glenbog

I3

0.112 ± 0.021

0.167 ± 0.028

67.39

14.48

1.58

3.02

0.09

1.72

Adamellite, Wallagaraugh Tonalite, Jindabyne

I4 I5

0.193 ± 0.021 0.067 ± 0.020

0.287 ± 0.030 0.129 ± 0.026

74.87 62.29

12.97 16.46

1.04 1.85

0.60 3.41

0.05 0.10

Qtz. Diorite, Clear Hills Adamellite, Eugowra ___, Lysterfield

I6 I7 I8

0.064 ± 0.021 0.121 ± 0.022 0.125 ± 0.020

0.109 ± 0.030 0.194 ± 0.041 0.193 ± 0.025

56.20 69.51 63.73

13.13 15.34 15.14

8.18 2.53 0.87

6.23 1.22 3.77

Dacite, Kadoona

I9

0.112 ± 0.020

0.147 ± 0.031

66.17

14.52

1.80

Granodiorite, Cooma Granodiorite, Cowra

S1 S2

0.099 ± 0.022 0.097 ± 0.020

0.133 ± 0.034 0.114 ± 0.029

72.00 67.87

13.72 14.59

Granodiorite, Jillamatong Adamellite, Minnegans Granodiorite, Dalgety Adamellite, Numbla ___, Koetong Granite, Strathbogie

S3 S4 S5 S6 S7 S8

0.114 ± 0.020 0.155 ± 0.033 0.113 ± 0.020 0.139 ± 0.021 0.112 ± 0.022 0.152 ± 0.020

0.150 ± 0.029 0.215 ± 0.059 0.182 ± 0.026 0.185 ± 0.038 0.170 ± 0.032 0.219 ± 0.025

67.68 68.72 68.21 73.48 73.65 73.65

Dacite, Hawkins

S9

0.072 ± 0.021

0.122 ± 0.030

Granite, Watergums Granite, Mumbella Granite, Black Hills, S. Dakota, USA Granite, Black Hills, S. Dakota, USA

A1 A2 HP49-A

0.148 ± 0.022 0.305 ± 0.022 0.161 ± 0.052

HP21-C

K2O

Pegmatites Simple Pegmatites Black Hills, S. Dakota, USA Black Hills, S. Dakota, USA Black Hills, S. Dakota USA Black Hills, S. Dakota, USA Tin Mountain Pegmatites Black Hills, S. Dakota, USA Black Hills, S. Dakota USA Black Hills, S. Dakota, USA Rare-element Pegmatites Little Nahanni, Canada Little Nahanni, Canada

0.053 ± 0.027

0.096 ± 0.042

(1984) (1984) (1984) (1984)

M. Telus et al. / Geochimica et Cosmochimica Acta 97 (2012) 247–265

TiO2

Granitoids Tonalite, Tuross Head

Groat et al. (2003) Groat et al. (2003) (continued on next page)

251

252

Table 1 (Continued) Sample type

Name

d56Fe(&)

d57Fe(&)

Averages of replicate analyzes*

Major element compositions (wt.%)

d56Fe(&)

d57Fe(&)

SiO2

Al2O3

Fe2O3

FeO

MnO

MgO

CaO

Na2O

K2O

TiO2

P2O5

H2O+

H2O

CO2

References

Little Nahanni, Canada

P389913

0.046 ± 0.039 0.160 ± 0.039

0.091 ± 0.060 0.225 ± 0.059

0.137 ± 0.027

0.206 ± 0.042

67.16

19.42

0.46

0.32

0.10

0.10

1.07

5.41

3.20

0.74

1.30

Little Nahanni, Canada

P389704

0.115 ± 0.039 0.202 ± 0.041

0.185 ± 0.060 0.317 ± 0.047

0.155 ± 0.027

0.261 ± 0.036

74.32

16.71

0.15

0.13

0.14

0.01

0.42

3.66

2.32

0.00

0.26

(Spodumene-bearing) Little Nahanni, Canada

P389726

0.118 ± 0.036 0.083 ± 0.039

0.184 ± 0.055 0.113 ± 0.059

0.067 ± 0.027

0.105 ± 0.040

70.67

17.24

0.27

0.26

0.21

0.04

0.63

4.66

3.22

0.91

0.91

Groat et al. (2003)

(Spodumene-bearing) Little Nahanni, Canada

P389739

0.052 ± 0.036 0.215 ± 0.052

0.097 ± 0.055 0.307 ± 0.055

71.77

17.36

0.28

0.26

0.16

0.00

0.65

4.17

2.26

0.89

0.92

Groat et al. (2003)

115-1L

0.246 ± 0.031

0.372 ± 0.045

72.46

14.85

0.68

0.01

0.22

0.40

1.62

8.43

0.06

0.41

0.59

115-1M

0.138 ± 0.031

0.205 ± 0.045

61.86

20.67

7.65

0.11

2.55

0.18

0.23

4.10

0.69

0.14

1.25

118-1L

0.272 ± 0.029

0.416 ± 0.050

74.91

14.55

0.39

0.10

0.07

4.85

0.06

0.23

0.01

3.84

0.33

118-1M

0.077 ± 0.029

0.153 ± 0.050

59.24

13.92

12.53

0.15

4.23

0.36

0.12

5.67

1.35

0.27

0.88

131-1B-L

0.273 ± 0.031

0.395 ± 0.045

76.40

13.82

0.55

0.01

0.15

1.00

3.64

3.24

0.05

0.09

0.55

Nabelek (1999) Nabelek (1999) Nabelek (1999) Nabelek (1999) Nabelek (1999)

131-1B-M

0.279 ± 0.041 0.237 ± 0.031

0.403 ± 0.047 0.329 ± 0.045

61.75

20.21

8.30

0.08

2.90

0.11

0.32

4.00

0.93

0.06

0.12

127-1L

0.355 ± 0.028

0.517 ± 0.052

79.03

18.59

0.39

0.01

0.04

0.26

0.42

0.67

0.02

0.16

0.54

127-1M

0.302 ± 0.028

0.456 ± 0.052

46.86

16.83

16.82

0.17

5.39

0.06

0.16

7.64

2.20

0.09

1.17

129-1B-L

0.283 ± 0.028

0.389 ± 0.052

79.45

16.52

0.54

0.01

0.16

0.17

0.67

2.32

0.05

0.04

0.40

129-1B-M

0.186 ± 0.028

0.270 ± 0.052

55.56

20.65

9.62

0.11

3.59

0.33

0.87

5.19

1.09

0.10

1.52

118-2

0.480 ± 0.029

0.738 ± 0.050

78.74

18.35

0.80

0.02

0.22

0.42

0.00

0.37

0.06

0.37

0.65

84-1

0.146 ± 0.031

0.253 ± 0.060

69.34

14.67

5.14

0.06

1.75

0.65

1.73

3.57

0.7

0.16

1.24

157-1

0.158 ± 0.028

0.213 ± 0.052

63.89

17.39

6.14

0.08

2.15

0.76

1.90

4.31

0.8

0.17

1.44

Groat et al. (2003) 0.71

Groat et al. (2003)

Melanosomes, Black Hills, S. Dakota, USA Leucosomes, Black Hills, S. Dakota, USA Melanosomes, Black Hills, S. Dakota, USA Leucosomes, Black Hills, S. Dakota, USA Melanosomes, Black Hills, S. Dakota, USA Leucosomes, Black Hills, S. Dakota, USA No melanosome pair Schists, Black Hills, S. Dakota, USA Schists, Black Hills, S. Dakota, USA

0.275 ± 0.025

Notes: Error bars on d56Fe and d57Fe are 95% confidence intervals. * Replicates are for newly digested samples.

0.399 ± 0.032

Nabelek (1999) Nabelek (1999) Nabelek (1999) Nabelek (1999) Nabelek (1999) Nabelek (1999) Nabelek (1999) Nabelek (1999)

M. Telus et al. / Geochimica et Cosmochimica Acta 97 (2012) 247–265

(Spodumene-bearing) Migmatites Leucosomes, Black Hills, S. Dakota, USA Melanosomes, Black Hills, S. Dakota, USA Leucosomes, Black Hills, S. Dakota, USA Melanosomes, Black Hills, S. Dakota, USA Leucosomes, Black Hills, S. Dakota, USA

M. Telus et al. / Geochimica et Cosmochimica Acta 97 (2012) 247–265 12 Na 2 O + K 2 O (wt.%)

Foidite

Trachyte

Granitoids I-type S-type A-type

Trachyandesite

8

Basaltic trachyandesite

Pegmatites Rhyolite

4

Basalt

Basaltic andesite

Andesite

Migmatites Leucosomes Melanosomes Schists

Dacite

0 45

55

65

75

85

SiO2 (wt.%) Fig. 1. Total alkali vs. SiO2 wt.% illustrates the compositional range of samples analyzed for this study. Dashed lines connect leucosomes to their associated melanosomes.

where i refers to mass 66 or 68. External reproducibilities (2SD) of ±0.09& for d66Zn and ±0.3& for d68Zn are estimated from 10 replicate measurements of one sample that had been split into 10 separate aliquots and measured independently by MC-ICP-MS (Herzog et al., 2009). Magnesium isotopic analyses were performed on a Nu Plasma MC-ICP-MC at the Isotope Laboratory of the University of Arkansas, Fayetteville, following established procedures (Yang et al., 2009; Li et al., 2010; Teng et al., 2010b). The digested samples were taken up in 1 M HNO3 for column purification. Separation of Mg was achieved by cation exchange chromatography with BioRad 200–400 mesh AG50 W-X8 resin in 1 M HNO3. This step was repeated twice for each sample. The procedural blank was less than 10 ng (i.e., less than 0.1% of Mg loaded on the column). Geostandards were processed with samples to verify that the column purification does not introduce spurious effects (Teng et al., 2010a). Magnesium isotopic compositions were analyzed by the standard bracketing method and data are reported in standard d-notation in permil relative to DSM3, a synthetic standard: h i di Mg ¼ ði Mg=24 MgÞsample =ði Mg=24 MgÞDSM3  1  1000; where i refers to mass 25 or 26. The long-term precision, based on replicate analyses of synthetic solutions, minerals and rock standards, is ±0.07& (2SD) (Yang et al., 2009; Teng et al., 2010a). Samples were prepared for uranium isotope analyses according to the procedure outlined by Tissot and Dauphas (2011, 2012) and Weyer et al. (2008). The samples were double-spiked prior to digestion using commercially available IRMM-3636 (49.51% of 236U and 50.46% of 233 U). The amount of spike added is such that the ratio Uspike/Usample is 3% for each sample. In granitic rocks, U can be hosted in accessory minerals that are notoriously difficult to digest, such as zircons. To ensure that the U isotope measurements are representative of the bulk rocks, the samples were digested in several steps: (1) two 24 h digestions in HF/HNO3 2:1 on hot plates, (2) a 5 day digestion in Parr Bombs in HF/HNO3 2:1 at 180 °C, and (3) two

253

24 h digestions in HCl/HNO3 2:1 on hot plates. After drying, the samples were dissolved in HNO3 3 M to be loaded on U/Teva resin (Eichrom). Most matrix elements were eluted in 40 ml HNO3 3 M. Neptunium and Th were subsequently removed with 20 ml HCl 5 M + 1 M oxalic acid (H2C2O4) followed by 10 ml HCl 5 M to rinse off the oxalic acid. Uranium was finally eluted in 25 ml HCl 0.05 M and the procedure was repeated twice to ensure complete matrix removal. All U isotope measurements were performed on the Neptune MC-ICPMS at the Origins Laboratory of the University of Chicago, using a desolvating nebulizer Aridus II. Uranium isotope analyses were done after the instrument was upgraded with an OnTool Booster 150 Jet pump (Pfeiffer). Enhanced signal stability was achieved by placing a spray chamber between the Aridus II and the MC-ICPMS. Correction of isotopic mass fractionation introduced during chemical separation and mass spectrometry was done using a 233U/236U double spike (IRMM3636). An additional correction was applied by bracketing samples with standard measurements also spiked with IRMM-3636 at the same level as the samples. The procedural blank was estimated to be 0.02–0.05 ng U (0.01% of sample uranium). Contributions of the 238U tail onto 236 U, 235U and 234U were estimated to be, respectively, 0.6  106, 0.25  106 and 0.1  106 of the 238U intensity and were corrected for. The 238U tail corrections were significant only for 234U and 236U. Ratios are reported in d238U notation relative to the U standard CRM-112a (also named SRM960, NBL112-a or CRM-145). h i d238 U ¼ ð238 U=235 UÞsample =ð238 U=235 UÞCRM112a  1  103 : The typical precision on d238U is ±0.05& (2SD). Two granite geostandards were analyzed (G-2 and JG-1; Table 2) and show reasonable agreement with the data reported by Weyer et al. (2008). Comparison of a large record of geostandards analyzed at the Origins Laboratory shows no systematic bias when compared to literature data. 3.2. Replicates and geostandards Samples that were especially difficult to dissolve were replicated to ensure reproducibility of d56Fe measurements. Replicates were done mostly for pegmatite samples. Out of five replicate analyses, only one (P389704) has a replicate value that does not agree within error (i.e., d56Fe = +0.202 ± 0.041& vs. +0.118 ± 0.036&). The reason for this is unclear but could be due to incomplete digestion. The averages of the replicate analyses are used in the figures (Table 1). Accuracy and precision of sample preparation and isotopic measurements were also evaluated by analyzing several international standards, such as AC-E, BE-N, BHVO-1, BHVO-2, DNC-1, DR-N, GA, and GSN. BHVO-1 was measured during each run and the replicate compositions agree within analytical uncertainty. The measured d56Fe values of the standards (Table 2) agree within error with previously published values (Craddock and Dauphas, 2010 and references therein).

254

Table 2 Iron , Zn, Mg and U isotopic compositions of geostandards. Averages of replicate analyses* Name

d56Fe (&)a

d57Fe (&)

Granite, Ailsa Craig island, Scotland Basalt, Essey-la-Cote, France Basalt, Hawaii, USA

AC-E

BHVO-2 DNC-1

0.311 ± 0.027 0.355 ± 0.065 0.162 ± 0.039 0.135 ± 0.031 0.091 ± 0.028 0.104 ± 0.029 0.133 ± 0.052 0.133 ± 0.031 0.139 ± 0.036 0.137 ± 0.033 0.120 ± 0.052 0.065 ± 0.039

0.471 ± 0.037 0.318 ± 0.025 0.479 ± 0.035 0.556 ± 0.115 0.242 ± 0.060 0.146 ± 0.024 0.210 ± 0.036 0.466 ± 0.09 0.849 ± 0.30 0.192 ± 0.045 0.150 ± 0.040 0.118 ± 0.013 0.170 ± 0.019 0.141 ± 0.044 0.162 ± 0.055 0.194 ± 0.045 0.192 ± 0.055 0.194 ± 0.046 0.173 ± 0.055 0.294 ± 0.09 0.601 ± 0.30 0.106 ± 0.059

DR-N GA

0.126 ± 0.039 0.202 ± 0.059 0.137 ± 0.031 0.195 ± 0.045

GSN MDO-G

0.152 ± 0.041 0.239 ± 0.047

Basalt, Hawaii, USA Dolerite, North Carolina, USA Diorite, Neuntelsein, France Calc-alkaline granite, Vosges Mountains,France Granite, France Trachyte, Mont Dore, France Serpentinite, ____ Pitscurrie Microgabrro, Scotland Sill Dolerite, England synthetic standard Mg:Fe:Al:Ca:Na:K:Ti = 1:1:1:1:1:1:0.1 Olivine, Kilbourne Hole, New Mexico, USA Granite, Rhode Island, USA Sori Granodiorite, Japan a

BE-N BHVO-1

d56Fe (&)

d57Fe (&)

d66Zn (&)b

d68Zn (&)

d26Mg (&)

U (ppm)

d238U (&)d

0.288 ± 0.09 0.502 ± 0.30 0.267 ± 0.09 0.585 ± 0.30

0.344 ± 0.09 0.773 ± 0.30

UB-N PM-S

0.380 ± 0.09 0.726 ± 0.30 0.444 ± 0.09 0.936 ± 0.30

WS-E DSM-3

0.173 ± 0.09 0.372 ± 0.30

KH setp22 G2 JG-1

d25Mg (&)c

0.01 ± 0.07 0.01 ± 0.09

0.18 ± 0.07 0.33 ± 0.09 1.86 ± 0.01 0.11 ± 0.10 4.00 ± 0.02 0.30 ± 0.04

Error bars on d56Fe and d57Fe are 95% confidence intervals. External reproducibilities (2SD) of ± 0.09& for d66Zn and ±0.3& for d68Zn are estimated from 10 replicate measurements of one sample that had been split into 10 separate aliquots and measured independently by Herzog et al. (2009). c External reproducibilities (2SD) of 0.06& for d25Mg and 0.07& for d26Mg is based on replicate analyses of synthetic solutions, mineral and rock standards reported in Yang et al. (2009) and Teng et al. (2010a). 2SD = 2 times the standard deviation of the population of 4 repeat measurements of a sample solution. d External reproducibilities (2SD) of 0.05& for d238U is based on numerous (>250) replicate measurements of the bracketing standard spiked at the same level as the samples. * Replicates are for newly digested sample. b

M. Telus et al. / Geochimica et Cosmochimica Acta 97 (2012) 247–265

Sample type

M. Telus et al. / Geochimica et Cosmochimica Acta 97 (2012) 247–265

56

δ Febulk (‰)

0.5

Granitoids I-type S-type A-type Pegmatites Black Hills Little Nahanni Migmatites Leucosomes Melanosomes Schists

0.3

0.1

-0.1 45

MORBs

Literature data Granitoids Volcanic Rocks

55

65 SiO2 (wt.%)

75

85

Fig. 2. d56Fe vs. SiO2 wt.% of migmatites, granitoids and pegmatites measured in this study compared to results of previous studies for silicic plutonic and volcanic rocks (Poitrasson and Freydier, 2005; Heimann et al., 2008; Sossi et al., 2012). The shaded area represents the d56Fe value of terrestrial basalts (MORBs) of +0.110 ± 0.018& (Dauphas et al., 2009). Granitic rocks become enriched in 56Fe with increasing SiO2 wt.%, which is in agreement with previous work. The d56Fe values of leucosomes are consistently heavier than average granites and MORBs, while the d56Fe values of melanosomes are scattered. Dashed lines connect leucosomes to their associated melanosomes. Schists are slightly heavier than MORBs, but similar in composition to average granitic rocks.

4. RESULTS 4.1. Iron isotopic compositions of felsic rocks The Fe isotopic compositions of 22 granitic rocks (I-, S-, and A-type), 11 migmatite components (six leucosomes and five melanosomes), and 13 samples of pegmatites are reported in Table 1 and plotted in Fig. 2. The results for geostandards are reported in Table 2 and 14 mineral separates (five biotite, four amphibole, and five magnetite) are reported in Table 3. All samples follow mass-dependent fractionation, i.e., d57Fe = (+1.49 ± 0.09)  d56Fe. 4.1.1. Granitoids The d56Fe values of granitic rocks measured in this study range from +0.064 ± 0.021& to +0.305 ± 0.022& (95% confidence intervals). The range of d56Fe values for I-type granitoids is from +0.064 ± 0.021& to +0.193 ± 0.021&. S-type granitoids have a similar range of d56Fe values (from +0.072 ± 0.021& to +0.240 ± 0.041&). Poitrasson and Freydier (2005) and Heimann et al. (2008) found that plutonic and volcanic rocks with >70 wt.% SiO2 had high d56Fe values. This study obtained similar results (Fig. 2) although some granitoids between 70 and 75 wt.% SiO2 did not show the expected enrichment in heavy Fe isotopes. The d56Fe values of most of the granitic rocks measured in this study lie within the range of MORBs (+0.110 ± 0.018&; Dauphas et al., 2009) (Fig. 2). 4.1.2. Pegmatites The iron isotopic compositions of pegmatites from the Black Hills are heavier than the average igneous rock composition with d56Fe values ranging from +0.151 ± 0.054&

255

to +0.385 ± 0.054& (Fig. 2). The “rare-element” pegmatites from Little Nahanni have a range of d56Fe values between 0.067 ± 0.027& and +0.215 ± 0.052&. It comprises the only pegmatite sample with a negative d56Fe value. They have been classified as spodumene-bearing and non-spodumene-bearing (Groat et al., 2003; see Table 1), but this classification does not seem to influence the d56Fe composition. Rare-element granitic pegmatites, like those from Little Nahanni, contain high concentrations of volatile elements, such as H2O, Li, B, F, and P (Groat et al., 2003), which may explain the scatter in their d56Fe compositions (Fig. 2). Pegmatites from both the Black Hills and Little Nahanni show no obvious correlation between d56Fe and SiO2 wt.%, or any other major oxides or trace elements. 4.1.3. Migmatites The Fe isotopic compositions of thirteen migmatite components from the Black Hills, South Dakota were measured (Fig. 2): six leucosomes, five melanosomes and two schists. The d56Fe values of the leucosomes range from +0.246 ± 0.031& to +0.480 ± 0.029&. Melanosomes have d56Fe compositions ranging from +0.077 ± 0.029& to +0.302 ± 0.028&. The two schists have an intermediate d56Fe composition of +0.153 ± 0.021&. Ten of the 13 migmatite components are leucosomes–melanosomes pairs (i.e., the leucosomes and melanosomes are spatially related to each other). The differences between d56Fe values for associated melanosomes and leucosomes range from +0.038 ± 0.056& to +0.196 ± 0.058&. The Fe isotopic compositions of leucosomes are systematically heavier than melanosomes compositions. Three of the five pairs do not overlap in composition. Leucosome d56Fe values overlap with the heaviest granite and pegmatites. Unlike granitoids and pegmatites, none of the leucosome Fe isotopic compositions overlap with MORBs. Melanosome Fe isotopic compositions vary widely, some are indistinguishable from MORBs and others are comparable to the heaviest pegmatites and granitoids. Schists are slightly heavier than MORBs (Fig 2). 4.1.4. Mineral separates Magnetite was found to have the highest d56Fe composition with values that range from +0.214 ± 0.026& to +0.645 ± 0.026& (Fig. 3). Without considering A2, which has an anomalously high d56Fe value for biotite (+0.249 ± 0.029&), the range for biotite is from 0.022 ± 0.023& to +0.056 ± 0.029&. The fractionation between magnetite and biotite (D56Femgt-bt) ranges between +0.234 ± 0.032& and +0.396 ± 0.037&. Hornblende d56Fe values range from 0.001 ± 0.023& to +0.076 ± 0.029& and are very similar in compositions to biotite. Iron isotopic fractionation between biotite and amphibole (D56Febt-amph) ranges between 0.048 ± 0.041& and +0.030 ± 0.041&. The iron isotope measurements of mineral separates agree with previous results of Heimann et al. (2008). To assess whether all of the major Fe-bearing minerals in each sample were accounted for, the bulk d56Fe value calculated from mineral mode analysis and the Fe isotopic

256

Table 3 Fe isotopic compositions of mineral separates from granitoids. Name

d56Fe(&) (mineral)

d57Fe(&) (mineral)

d56Fe(&) bulk (measured)

Density* (gm/ cm3)

Fe Concentration (g/g)

Mineral mode (%)**

Fraction of Fe (mineral)

d56Fe(&) bulk (calculated)

Amphibole, Tonalite Biotite, Tonalite Magnetite, Tonalite Amphibole, Granodiorite Biotite, Granodiorite Magnetite, Granodiorite Amphibole, Granite Magnetite, Granite Amphibole, Tonalite Biotite, Tonalite Magnetite, Tonalite Amphibole, Tonalite Biotite, Tonalite Magnetite, Tonalite

I1-Amph

0.053 ± 0.023

0.086 ± 0.036

0.123 ± 0.02

3.3

0.107

18

0.272

0.147 ± 0.040

I1-Bt I1-Mgt

0.045 ± 0.023 0.278 ± 0.023

0.073 ± 0.036 0.423 ± 0.036

3.0 5.2

0.129 0.636

18 3

0.301 0.427

I2-Amph

0.001 ± 0.023

0.001 ± 0.036

3.3

0.113

10

0.215

I2-Bt

0.022 ± 0.023

0.038 ± 0.036

3.0

0.121

15

0.316

I2-Mgt

0.214 ± 0.026

0.324 ± 0.041

5.2

0.520

3

0.469

A2-Bt

0.249 ± 0.026

0.380 ± 0.041

0.305 ± 0.022

3.0

0.107

6

0.455

0.465 ± 0.037

A2-Mgt I3-Amph

0.645 ± 0.026 0.026 ± 0.029

0.945 ± 0.041 0.051 ± 0.036

0.112 ± 0.021

5.2 3.3

0.444 0.124

1 16

0.545 0.338

0.163 ± 0.050

I3-Bt I3-Mgt

0.056 ± 0.029 0.433 ± 0.029

0.071 ± 0.036 0.642 ± 0.036

3.0 5.2

0.141 0.573

16 2

0.352 0.310

I5-Amph

0.076 ± 0.029

0.116 ± 0.036

3.3

0.123

16

0.440

I5-Bt I5-Mgt

0.028 ± 0.029 0.283 ± 0.029

0.016 ± 0.036 0.409 ± 0.041

3.0 5.2

0.108 0.331

20 1

0.443 0.118

0.081 ± 0.021

0.067 ± 0.02

Notes: Bulk Fe isotopic data calculated using the fraction of Fe and the d56Fe for each mineral. Error bars on d56Fe and d57Fe are 95% confidence intervals. * Averages from Nesse (2000). ** Mineral modes were determined by examining thin sections under an optical microscope and estimating what percentages of the minerals occupy the thin sections.

0.093 ± 0.042

0.079 ± 0.050

M. Telus et al. / Geochimica et Cosmochimica Acta 97 (2012) 247–265

Sample type

M. Telus et al. / Geochimica et Cosmochimica Acta 97 (2012) 247–265 1.0 I-type This Study Amphibole Biotite Magnetite

0.4

Heimann et al. (2008) Amphibole Biotite Magnetite

0.2

0.8

66

0.6

δ Zn (‰)

δ56Fe of mineral separates (‰)

0.8

257

S-type A-type PegmatitesBlack Hills Leucosomes Melanosome

0.6 0.4

MORBs

0.0 -0.2 55

0.2

60 SiO2

65

70

75

0.0 55

80

Fig. 3. d56Fe of mineral separates vs. SiO2 wt.% of bulk rocks. Fesilicates (biotite and amphibole) generally have lighter or indistinguishable iron isotopic compositions compared to MORBs, while magnetite compositions are significantly heavier.

Table 4 Zinc isotopic compositions of granitoids, pegmatites, and migmatites (relative to JMC-Lyon standard). Sample type

60

of bulk rocks (wt.%)

Name

d66Zn (&)

d68Zn (&)

I1 I2 I4 I6 I7 I8 S1 S2 S3 S4 S5 S6 S7 S8 S9 HP49-A HP21-C A1 A2

0.247 ± 0.09 0.192 ± 0.09 0.266 ± 0.09 0.171 ± 0.09 0.119 ± 0.09 0.179 ± 0.09 0.130 ± 0.09 0.122 ± 0.09 0.330 ± 0.09 0.224 ± 0.09 0.325 ± 0.09 0.209 ± 0.09 0.357 ± 0.09 0.355 ± 0.09 0.303 ± 0.09 0.486 ± 0.09 0.342 ± 0.09 0.269 ± 0.09 0.233 ± 0.09

0.473 ± 0.30 0.506 ± 0.30 0.682 ± 0.30 0.317 ± 0.30 0.426 ± 0.30 0.331 ± 0.30 0.357 ± 0.30 0.226 ± 0.30 0.664 ± 0.30 0.506 ± 0.30 0.677 ± 0.30 0.464 ± 0.30 0.755 ± 0.30 0.722 ± 0.30 0.665 ± 0.30 0.918 ± 0.30 0.675 ± 0.30 0.482 ± 0.30 0.462 ± 0.30

81BH 5-1 81BH 9-2 81BH 10-3 82BH 43-1 WC-9

0.673 ± 0.09 0.875 ± 0.09 0.752 ± 0.09 0.631 ± 0.09 0.534 ± 0.09

1.318 ± 0.30 1.655 ± 0.30 1.422 ± 0.30 1.178 ± 0.30 1.026 ± 0.30

115-L 131-1B-L 131-1B-M

0.258 ± 0.09 0.238 ± 0.09 0.179 ± 0.09

0.481 ± 0.30 0.432 ± 0.30 0.357 ± 0.30

Pegmatites

Migmatites

Notes: External reproducibilities (2SD) of ±0.09& for d66Zn and ±0.3& for d68Zn are estimated from 10 replicate measurements of one sample that had been split into 10 separate aliquots and measured independently by Herzog et al. (2009).

compositions of mineral separates were compared with the measured bulk rock compositions. The mineral mode, the

75

80

Fig. 4. d66Zn vs. SiO2 wt.%. There is no clear correlation, but there seems to be more dispersion in the Zn isotopic composition for more silicic magmas. The average pegmatite Zn isotopic composition is heavier than the average value for granitic rocks and migmatites by 0.4&.

Table 5 Magnesium isotope compositions of granitoids (relative to DSM-3 standard). Sample type

Granitoids

65 70 SiO2 (wt.%)

Name

d25Mg (&)

d26Mg(&)

I4 I8 S4 S6 S8 HP21C HP49A A1 A2

0.216 ± 0.070 0.129 ± 0.070 0.123 ± 0.070 0.107 ± 0.070 0.104 ± 0.070 0.121 ± 0.070 0.388 ± 0.070 0.104 ± 0.070 0.014 ± 0.070

0.446 ± 0.092 0.216 ± 0.092 0.251 ± 0.092 0.201 ± 0.092 0.209 ± 0.092 0.218 ± 0.092 0.720 ± 0.092 0.214 ± 0.092 0.026 ± 0.092

Granitoids

Notes: External reproducibilities (2SD) of 0.06& for d25Mg and 0.07& for d26Mg and are based on replicate analyses of synthetic solutions, minerals and rock standards reported in Yang et al. (2009) and Teng et al. (2010a). 2SD = 2 times the standard deviation of the population of four repeat measurements of a sample solution.

density of each mineral and the Fe concentration of each mineral give the fraction of Fe locked in each mineral x, (f Fe_x). A simple mass balance equation, 56Febulk = d56Femgtf Fe_mgt + d56Feamphf Fe_amph + d56Febtf Fe_bt, was used to calculate the d56Fe bulk value from the compositions of the minerals separates (mgt refers to magnetite, amph refers to amphibole, and bt refers to biotite). Most of the calculated bulk d56Fe measurements are within the range of the measured d56Fe values (Table 3). Predicted values that are slightly inconsistent with measured values probably reflect inaccuracy in mineral mode estimation using visual charts (see “Methods” for the technique use to estimate the mineral modes). 4.2. Zinc isotopic compositions of felsic rocks Zinc isotopic compositions were determined for 19 granitoids, three migmatite components, five pegmatites

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M. Telus et al. / Geochimica et Cosmochimica Acta 97 (2012) 247–265 1.0 I-type S-type A-type Pegmatites Black Hills

66

δ Zn (‰)

0.8 0.6

Leucosome Melanosome

Table 6 Uranium concentrations and isotope compositions of granitoids (relative to CRM-112a standard). Sample type

0.2

0.1

0.2

0.3

0.4

0.5

δ56Fe (‰)

Fig. 5. d66Zn vs. d56Fe. Pegmatites and some granitic rocks have high d66Zn and d56Fe values. Since Zn is easily mobilized by aqueous fluids, this correlation suggests that fluid exsolution may be responsible for the iron isotope variations measured in some felsic rocks, as suggested by Poitrasson and Freydier (2005) and Heimann et al. (2008). However, some samples (e.g., A-type granites and migmatites) have high d56Fe values, but unfractionated d66Zn values. For these samples, fractional crystallization is the most likely cause of iron isotopic fractionation.

0.8

26

δ Mg (‰)

0.6

This Study I-type S-type A-type

0.4

Literature I-type S-type A-type

0.2 0.0 -0.2

U(ppm)

d238U(&)

I1 I2 I3 I4 I5 I6 I7 I8 I9

1.518 ± 0.010 2.065 ± 0.015 2.527 ± 0.017 7.109 ± 0.064 1.530 ± 0.007 2.076 ± 0.014 3.731 ± 0.026 7.595 ± 0.063 3.944 ± 0.027

0.363 ± 0.033 0.272 ± 0.033 0.228 ± 0.033 0.250 ± 0.049 0.231 ± 0.030 0.270 ± 0.037 0.510 ± 0.039 0.385 ± 0.039 0.320 ± 0.039

S1 S1* S2 S3 S4 S5 S6 S7 S8 S9

3.455 ± 0.025 3.497 ± 0.025 2.911 ± 0.021 3.433 ± 0.025 3.343 ± 0.023 3.008 ± 0.021 4.272 ± 0.034 7.314 ± 0.070 7.533 ± 0.062 3.881 ± 0.028

0.234 ± 0.049 0.255 ± 0.033 0.306 ± 0.033 0.272 ± 0.049 0.300 ± 0.037 0.237 ± 0.037 0.401 ± 0.039 0.335 ± 0.039 0.285 ± 0.039 0.282 ± 0.039

A1 A2

5.913 ± 0.047 7.208 ± 0.055

0.328 ± 0.049 0.340 ± 0.049

Granitoids

0.4

0.0 0.0

Name

Notes: External reproducibilities (2SD) of 0.05& for d238U is based on numerous (>250) replicate measurements of the bracketing standard spiked at the same level as the samples. * Denotes replicates on a newly digested sample.

MORBs

2.5

Pyroxene

-0.4 65

70 SiO 2 (wt.%)

75

Fig. 6. d26Mg vs. SiO2 wt.% for granitic rocks compared to previous studies by Shen et al. (2009), Li et al. (2010) and Liu et al. (2010). The d26Mg values for the I- and S-type granitoids from this study are indistinguishable from MORBs (0.25 ± 0.07&; Teng et al., 2010a,b). Sample A2 and many other A-type granites (Li et al., 2010) have significantly higher d26Mg values than MORBS. 68

66

Biotite

80

(Table 4, d Zn = (+1.80 ± 0.32)  d Zn), and eight geostandards (Table 2). The range in d66Zn for granitic rocks is between +0.12& and +0.49& (2SD) and the range for pegmatites is from +0.53& to +0.88&. The Zn isotopic compositions of migmatites (average +0.23 ± 0.05&) overlap with those of granitic rocks. Pegmatites are heavier than granitoids and migmatites by an average of +0.4&. There is no clear correlation between d66Zn and SiO2 wt.% as observed for d56Fe (Fig. 4). There is however a relationship between d66Zn and d56Fe for pegmatites (Fig. 5) in so much that samples with high d56Fe values also have high d66Zn values. For granites and migmatites, the systematics is less clear with some samples showing slightly heavy d66Zn values, but for the most part, high d56Fe values are not associated with high d66Zn values.

Olivine

DZn /DFe

-0.6 60

Amphibole

2

Magnetite

1.5

Ilmenite Plagioclase

1 0.5 0 0.01

0.1

1 10 DFe (solid/liquid)

100

1000

Fig. 7. Solid/melt distribution coefficients of Fe and Zn in pyroxene, amphibole, biotite, olivine, magnetite, ilmenite, and plagioclase (Ewart and Griffin, 1994). Zinc is similar to Fe in radius and valence. In addition, fractionation of Fe and Zn in minerals common in granitic rocks such as biotite, amphibole, and oxides is small, so limited fractionation of the Zn/Fe ratio is expected during magmatic differentiation.

4.3. Magnesium isotopic data The magnesium isotopic compositions of eight granitic rocks (two I-types, five S-types and two A-types) were analyzed and are reported in Table 5. Results for the standards

M. Telus et al. / Geochimica et Cosmochimica Acta 97 (2012) 247–265

140

Fractional crystallization of a hydrous dioritic melt

120

Fractional crystallization of an anhydrous dioritic melt

by +0.2& (Fig. 6). These results agree with those reported by Shen et al. (2009), Liu et al. (2010) and Li et al. (2010).

Fluid exsolution

Zn/Fe ( 104, atomic ratio)

160

100

4.4. Uranium isotopic data

80 60 Magma differentiation

40 20 0

55

60

65

70

75

80

SiO2 (wt.%) Fig. 8. Zn/Fe atomic ratio (104) vs. SiO2 wt.% for granitic rocks (data is from GEOROC database, http://georoc.mpch-mainz.gwdg.de/georoc/) and magma differentiation model results (green and purple circles). The dashed line represents the average value for peridotites (Le Roux et al., 2010). Little fractionation is expected between Zn and Fe during magmatic differentiation. However, strong fractionation can occur in exsolved aqueous fluids (Holland, 1972; Khitarov et al., 1982; Urabe, 1987; Simon et al., 2004; Zajacz et al., 2008). The large dispersion in Zn/Fe of highly silicic magmas (i.e., SiO2 > 70 wt.%) is interpreted to reflect Fe–Zn mobilization as chloride complexes during fluid exsolution. Trajectories of Zn/Fe vs. SiO2 wt.% during magma differentiation were calculated using Rhyolite Melts (1300–950 °C; 5 kbars) for fractional crystallization of a hydrous (1.36 wt.% H2O; green circles) and anhydrous dioritic melt (purple circles) with an initial Zn concentration of 50 ppm. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

are reported in Table 2. Expectedly, the data follow massdependent fractionation, d25Mg = (+0.530 ± 0.146)  d26Mg. The average d26Mg compositions for A-, S- and Itypes are 0.094 ± 0.065&, 0.320 ± 0.041&, and 0.331 ± 0.065&, respectively (2SD). They are indistinguishable from MORBs (0.25 ± 0.07&; Teng et al., 2010a), except sample A2 (+0.026 ± 0.092), which is enriched in the heavier isotopes of Mg relative to MORBs

A

259

The uranium isotopic compositions of 22 granitoids (I-, S-, and A-type) are reported in Table 6 and results for the standards are reported in Table 2. The typical uncertainty for measurements of granitic rocks is 0.04& (95% confidence interval). For the most part granitoids show little isotopic variations with most of the samples between 0.4& and 0.2& in d238U. I-, S- and A-type granites have different protoliths and different modes of formation; however there is no significant difference in d238U between these three granite types. Individual granitic rocks show variations in d238U. In particular, one I-type rock (I7) has a very low isotopic composition of 0.510 ± 0.039&. The average U isotopic composition of all the granitoids is 0.305 ± 0.143& (2SD). These isotopic compositions agree well with previously documented U isotopic composition of igneous rocks, including basalts (Weyer et al., 2008; Tissot and Dauphas, 2012). There is no correlation between U isotope variations and proxies of felsic magma evolution, such as SiO2 content. 5. DISCUSSION Previous work has suggested that Fe isotopes fractionate during partial melting (e.g., Weyer and Ionov, 2007; Dauphas et al., 2009), fractional crystallization (e.g., Schoenberg and von Blanckenburg, 2006; Teng et al., 2008; Schoenberg et al., 2009), fluid exsolution (Poitrasson and Freydier, 2005; Heimann et al., 2008), and also as a result of Soret (thermal) diffusion in plutons (Lundstrom, 2009). Here we discuss the possible origins of Fe isotopic variations documented in felsic rocks and how our results can shed new light on this issue. The fractional crystallization model has difficulties explaining why felsic rocks become heavy when a significant Fe-bearing phase crystallizing is an oxide (magnetite or

B

Fig. 9. Zn vs. SiO2 wt.% for granites from the GEOROC database (A) and Zn/Fe ratio vs. Zn (B). The GEOROC Zn data is considered reliable because the dispersion in the data does not depend on the Zn concentration.

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M. Telus et al. / Geochimica et Cosmochimica Acta 97 (2012) 247–265 0.4

Lacks et al. (2012)

0.3

δ 56

0.2

0.1

0.1

56

0.2

A 0.0 -1

-0.5

B

0.5

0

-0.6

26

-0.5

-0.4 δ

δ Mg (‰) 56

26

Fe (‰)

δ Fe (‰)

0.3

0.4

Richter et al. (2009) and Lacks et al. (2012) I-type S-type A-type

56

238

-0.3

-0.2

-0.1

0.0

U (‰)

238

Fig. 10. Expected isotopic fractionation of d Fe–d Mg (A), and d Fe–d U (B) produced by thermal (Soret) diffusion in a silicate melt (dashed lines; Richter et al., 2009 and Lacks et al., 2012). Soret diffusion has been proposed as a possible mechanism for fractionating iron isotopes in felsic rocks (Lundstrom, 2009). The data do not follow the correlations expected for Soret diffusion.

ilmenite). Indeed, magnetite and ilmenite are expected to have heavy rather than light Fe isotopic compositions relative to silicate phases (Polyakov et al., 2007; Shahar et al., 2008; Craddock et al., 2010), which should drive the system towards lower d56Fe values, opposite to what is observed. However, in order to definitely rule out this model, the reduced partition function ratios for iron isotopes in silicate magmas remain to be determined. The Fe isotopic composition of most of the granitic rocks measured in this study did not show significant deviation from the igneous average (+0.09&; Beard et al., 2003) (relative to IRMM-014), except for the most felsic granites (>70 wt.% SiO2; Fig. 2). The d56Fe values of I- and S- type granites overlap, so the redox state of the source rock cannot be the main control on iron isotopic variations in these samples.

melanosome pairs (Fig. 2) suggests that some of the melanosomes have equilibrated with isotopically evolved leucosomes. Re-equilibration has been shown to also affect the major and trace element compositions of leucosomes (Fourcade et al., 1992; Johannes et al., 1995). Despite this possible overprint, the Fe isotopic data preserve evidence that partial melting of crustal lithologies yields melts with fractionated Fe isotopic compositions, consistent with the interpretations of Weyer and Ionov (2007) and Dauphas et al. (2009). A remaining issue is the extent to which fractional crystallization and reequilibration of the leucosomes during cooling affected iron isotopes.

5.1. Iron isotopic fractionation during partial melting

Pegmatites are evolved melts that may be particularly sensitive to fluid exsolution. For example, Shearer et al. (1986) documented extensive interaction between pegmatite-derived fluids and wallrock in the Black Hills, South Dakota. The d56Fe values of pegmatites are similar or heavier than the igneous average (except for sample P389726), indicating that pegmatites are not the predominant reservoir for isotopically light Fe. Examining the fractionation between zinc and iron provides a method of further evaluating the effect of fluid exsolution on Fe isotopes. Zinc is a divalent element that is similar in ionic radius to Fe2+ ˚ ) and hence also similar in behavior. We measured (0.74 A the zinc isotopic compositions of many of our samples and found in pegmatites and some granites that high d56Fe were accompanied by high d66Zn values (Fig. 5), implying that fluid exsolution is indeed an important source of Fe isotopic variations in felsic rocks as suggested by Poitrasson and Freydier (2005) and Heimann et al. (2008). However, the d56Fe values for most granitoids and in particular those of A-type, which formed under relatively dry conditions (Collins et al., 1982), do not correlate with their Zn isotopic compositions. For these samples that do not show fractionated d66Zn values, the most likely cause of iron isotopic fractionation is fractional crystallization. A similar conclusion was reached by Sossi et al. (2012) based

Migmatites are important samples for addressing the question of Fe isotopic fractionation during partial melting of granitic source material because they represent melts arrested in the process of migrating from their source. Although leucosomes form by partial melting, their Fe isotopic compositions can also be affected by fractional crystallization of the partial melts. Leucosomes measured in this study have d56Fe values heavier by +0.2& relative to schists, which are most representative of migmatite protoliths, and also relative to the average composition of bulk granitic rocks. In terms of mass-balance, most of the iron is expected to stay in melanosomes, which are the residues of partial melting and are rich in mafic minerals. Melanosomes are, therefore, expected to have an isotopic composition that is similar to that of the bulk protolith. All melanosomes have lighter Fe isotopic compositions than their leucosome complements (Fig. 2). Of the five leucosome–melanosome pairs that were analyzed, at least two melanosomes have Fe isotopic compositions identical to that of schists, whereas two others have heavy d56Fe values by up to +0.2&. This enrichment is difficult to explain by closed system behavior. Instead, the broadly positive correlation between the Fe isotopic compositions of leucosome–

5.2. Iron isotopic fractionation during fluid exsolution and fractional crystallization

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on the argument that A-type granites were unlikely to have experienced significant iron mobilization by fluid exsolution, yet these granites often have high d56Fe values. Iron isotopic fractionation during magma differentiation is a process that could be tested by modeling. Unfortunately, most equilibrium iron isotopic fractionation factors between crystallizing minerals and melts are unknown, which should be the focus of future studies. Mafic magmas show limited fractionation of Zn from Fe (Le Roux et al., 2010; Lee et al., 2010). Felsic magma differentiation is also not expected to induce large fractionation of Zn/Fe as the minerals involved have similar mineral/melt distribution coefficients (Fig. 7). The average DZn/DFe ratios (where D, the distribution coefficient between the solid and the liquid, is calculated as the ratio of the concentration of the element in the solid and in the liquid, Csolid/Cliquid) for the minerals involved are 0.5 for pyroxene, 0.3 for amphibole, 0.6 for biotite, 0.4 for olivine, 0.5 for magnetite, 0.2 for ilmenite, and 1.0 for feldspar (Ewart and Griffin, 1994). All DZn/DFe ratios (except for feldspar) are lower than one, meaning that the Zn/Fe ratio of the melt should increase during felsic magma differentiation. On the other hand, Zn and Fe can be significantly fractionated by complexation in chlorine-bearing magmatic fluids (Holland, 1972; Khitarov et al., 1982; Urabe, 1987; Simon et al., 2004; Zajacz et al., 2008). Fig. 8 shows a compilation of Zn/ Fe ratios in granitoids from the GEOROC database (http:// georoc.mpch-mainz.gwdg.de/georoc/). As discussed, little fractionation of Zn/Fe ratios is expected during granitic magma evolution and only a small increase in Zn/Fe vs. SiO2 is observed in the data. However, a notable feature is the large dispersion in Zn/Fe ratio of high-SiO2 granites, which spans two orders of magnitude for the most silicic samples (>70 wt.% SiO2). We modeled fractional crystallization of hydrous (1.36 wt.% H2O) and anhydrous felsic melts to examine whether magmatic differentiation could explain the large dispersion in Zn/Fe ratios of granites. We used Rhyolite Melts, a program that enables calculations of fractional crystallization of hydrous silicic melts under a range of pressures and temperatures (Gualda et al., 2012; http:// melts.ofm-research.org/). The initial composition of the melt is assumed to be dioritic (66 wt.% SiO2) with initial Zn and Fe contents of 50 ppm and 2.84 wt.%, respectively. The values for DFe for orthopyroxene, clinopyroxene, magnetite and other oxides were determined by calculating the ratio of the concentration of Fe in the minerals to that in the melt and values for DZn for each mineral were estimated by multiplying the DFe values by published DZn/DFe ratios (Ewart and Griffin, 1994). The concentration of Fe for the melt during fractional crystallization was computed by the Melts program, whereas changes in the concentration of Zn for the melt were determined by numerically modeling the incremental removal of Zn incorporated in the minerals that crystalized from the melt. Model results are plotted in Fig. 8 (green and purple circles) and indicate that fractional crystallization can explain the correlation between Zn/Fe and SiO2 wt.%, but it cannot readily explain the large dispersion of Zn/Fe ratios observed for granites. This dispersion could potentially result from differences in the

261

quality of Zn and Fe analyses reported in the GEOROC database. However, plots of Zn vs. SiO2 content (Fig. 9A) and Zn/Fe vs. Zn (Fig. 9B) indicate that the Zn concentration data appear to be reliable. Thus, the dispersion in Zn/ Fe ratios is likely produced by mobilization of Zn and possibly Fe in exsolved fluids. During late stage crystallization of a granitic melt, Fechlorides can form and subsequently be lost in exsolved fluids (Simon et al., 2004; Zajacz et al., 2008), which could possibly leave the remaining silicate melt enriched in heavy iron isotopes. This mechanism has been proposed as a mechanism for enrichment of heavy Fe isotopes in evolved granitic rocks (Poitrasson and Freydier, 2005; Heimann et al., 2008). Heimann et al. (2008) presented a quantitative assessment of this scenario based on solubility data and predicted equilibrium fractionation factors for iron in minerals and fluids. A significant result of the present study is that in pegmatites and some granitic rocks, Zn isotopes show a moderate correlation with Fe isotopes (Fig. 5). This observation, together with the observed scatter in Zn/Fe ratio for granitic rocks (Fig. 8), suggests that fluid exsolution is responsible in part for the isotopic fractionation of iron in felsic rocks. With the fluid exsolution model, a question is posed regarding the fate of the isotopically light exsolved reservoir. Significant amounts of iron must be removed to explain the heavy iron isotopic composition of pegmatites and some granitic rocks. Such a reservoir has not been conclusively identified but it is worth noting that the Little Nahanni pegmatites, which are from volatile-rich sources, have d56Fe compositions that are highly variable, including light Fe isotopic compositions (0.07&; Fig. 2). Further work on fluid inclusions is needed to ascertain the nature of the putative complementary reservoir to isotopically heavy iron in pegmatites and some granites. 5.3. Demonstration that Soret effect is absent in the studied granitoids from Mg, U and Fe isotope systematics Magnesium isotopic compositions of granitic rocks are important to compare with Fe isotopic data because Mg isotopes do not fractionate during partial melting or magmatic differentiation, but are fractionated significantly during diffusion and weathering (Teng et al., 2007, 2010b; Richter et al., 2008; Shen et al., 2009; Lundstrom, 2009; Liu et al., 2010). The Lachlan Fold Belt granitoids measured in this study are fresh samples and their Mg isotopic compositions were not affected by weathering. We measured Mg isotopic data for granitic rocks with the highest d56Fe values to determine whether Soret diffusion and/or derivation from sedimentary protoliths are influencing the d56Fe composition of these felsic rocks. Little variation in d26Mg among the different types of granitic rocks was found. The one exception is sample A2, which has a d26Mg composition of +0.026 ± 0.092&, significantly heavier than MORBs. Sample A2 also has the highest d56Fe of +0.305 ± 0.022&. Heterogeneous d26Mg values for A-type granites have also been reported by Li et al. (2010). The Mg isotopic variations of A-type granites possibly reflect chemical heterogeneities inherited from the source (Shen et al., 2009; Li et al., 2010; Liu et al., 2010). The reason for hea-

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ment sensitive to fluid exsolution, Zn, in the same samples that were analyzed for iron. Pegmatites and some granitoids show high d56Fe values that are accompanied by high d66Zn values (Fig. 5), so fluid exsolution is most likely responsible for the iron isotope variations measured in these samples. This is supported by the observation that the Zn/Fe ratio shows increasing scatter in more silicic granites, which cannot easily be explained by magma differentiation and could be due to fluid exsolution instead (Fig. 8). 4. Fluid exsolution alone cannot explain the high d56Fe values measured in all granitoids, as many of these samples show no variations in d66Zn. Fractional crystallization is the most likely cause of iron isotope variations in these samples. This is particularly relevant to A-type granites, which presumably formed in relatively dry conditions and are not expected to have experienced significant fluid exsolution. 5. Finally, d26Mg and d238U results for granitic rocks suggest that thermal migration (Soret diffusion), which has been proposed as a mechanism of granitic magma differentiation (Lundstrom, 2009), can be ruled out in our sample set because the d26Mg, d238U and d56Fe do not follow the correlations expected for Soret diffusion (Richter et al., 2009; Lacks et al., 2012) (Fig. 10).

vier Mg and Fe isotopic compositions in this sample is unclear. Soret (thermal) diffusion has been proposed as a mechanism for producing heavy d26Mg and d56Fe compositions in plutons (Lundstrom, 2009). Based on the experimental results from Richter et al. (2009) and Lacks et al. (2012), Fe and Mg isotopes from a source that has experienced Soret diffusion in a thermal gradient should follow the trend illustrated in Fig. 10A; however, the granitic rocks that were measured for both d56Fe and d26Mg do not clearly follow the expected correlation. Further evidence that Soret diffusion played no role in the isotopic fractionation documented here comes from U and Fe isotope measurements. One of our motivations to investigate the U isotopic composition of granitic rocks was to assess whether the U isotopic compositions of zircons could be related to the nature of the protolith (i.e., sedimentary vs. igneous) with potential application to Hadean zircons, which provide the only record of Earth’s evolution during that eon. We did not find any evidence for that; U isotope variations in granites are not related to the nature of their protoliths. Another reason for measuring U isotopes is that Soret diffusion could fractionate U isotopes in silicate melts (Lacks et al., 2012). Such fractionation would produce correlated Fe–U isotopic variations (the slope between d56Fe and d238U in laboratory experiments is 3.18). As shown in Fig. 10B, we do not see any correlation in the sample set studied that would suggest that Soret diffusion was present. If anything, the variations would define a trend that is orthogonal to the trend predicted for Soret diffusion. We conclude, based on the absence of correlation between Mg, U and Fe isotopic compositions, that the granitoids investigated here (I-, S-, and A-type granitoids from Lachlan Fold Belt) were not influenced by Soret effect.

Our results on a limited, albeit representative, sample set support fluid exsolution and fractional crystallization as the primary mechanisms for the d56Fe variations of felsic rocks. Equilibrium fractionation factors between silicate melts and minerals have to be characterized to quantitatively assess the role of fractional crystallization on iron isotopes in granitoids.

6. CONCLUSIONS

ACKNOWLEDGEMENTS

The d56Fe values of granitoids are positively correlated with SiO2 wt.%, consistent with previous studies (Poitrasson and Freydier, 2005; Heimann et al., 2008). Several mechanisms for enriching the melt in heavy Fe isotopes have been proposed including: partial melting, fractional crystallization, fluid exsolution and thermal migration (Soret diffusion). This work sheds new light on the mechanisms governing the isotopic fractionation of iron and other nontraditional stable isotope systems in the crust:

We thank Alfred T. Anderson for valuable assistance with mineral separation; Mark S. Ghiorso and Julia E. Hammer for assisting with the Melts calculations; Richard J. Walker for providing the Black Hills pegmatite samples. We also thank Francis Albarede and Philippe Telouk for the generous access to the Nu Plasma mass-spectrometer in Lyon for the Zn isotopic measurements. Drafts of this manuscript have benefitted from thoughtful comments by Alfred T. Anderson and Cin-Ty A. Lee. Comprehensive reviews by three anonymous referees helped improve the manuscript. This work was supported by NASA (NNX09AG59G, NNX12AH60G), NSF (EAR-0820807) and a Packard Fellowship to N.D., by NSF (EAR-0838227 and EAR1056713) and Arkansas Space Grant Consortium (SW19002) to F.Z.T, and by NASA Grants (NNX12AD88G and NNX12AH70G) to F.M.

1. The enriched Fe isotopic compositions of leucosomes provide evidence that partial melting of the crust can fractionate the isotopes of Fe by at least 0.2& in d56Fe (Fig. 2). However, the extent to which the measured leucosomes were affected by fractional crystallization and reequilibration is unknown. 2. The nature and redox state of the protolith do not appear to influence the d56Fe and d238U compositions of granitic rocks. 3. Poitrasson and Freydier (2005) and Heimann et al. (2008) suggested that exsolved fluids could have removed isotopically light iron and hence explain the heavy iron isotopic composition of some felsic rocks. This idea was tested by measuring the isotopic variations of an ele-

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