Protocrustal evolution of the Nuvvuagittuq Supracrustal Belt as determined by high precision zircon Lu–Hf and U–Pb isotope data

Earth and Planetary Science Letters 428 (2015) 162–171

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

Protocrustal evolution of the Nuvvuagittuq Supracrustal Belt as determined by high precision zircon Lu–Hf and U–Pb isotope data Lars Eivind Augland a,c,∗ , Jean David a,b a

GEOTOP, Université du Québec à Montréal, CP 8888, Succ. Centre-Ville, Montréal, Québec H3C 3P8, Canada Bureau d’Exploration Géologique du Québec, Ministère de l’Énergie et des Ressources Naturelles, Université du Québec à Montréal, Succ. Centre-Ville, Montréal, Québec H3C 3P8, Canada c Department of Geosciences, University of Oslo, P.O. Box 1047 Blindern, N-0316 Oslo, Norway b

a r t i c l e

i n f o

Article history: Received 21 January 2015 Received in revised form 14 July 2015 Accepted 16 July 2015 Available online xxxx Editor: A. Yin Keywords: Nuvvuagittuq Supracrustal Belt Eoarchean Lu–Hf isotopes U–Pb geochronology Hadean

a b s t r a c t The Nuvvuagittuq Supracrustal Belt (NSB) in northern Quebec, Canada, represents one of the oldest known crustal fragments preserved in the Earth’s crust. Its age has, however, been disputed and different authors present crustal formation ages varying from ca. 3.8 Ga to 4.4 Ga (e.g. O’Neil et al., 2012; Guitreau et al., 2013). Here we report new high precision U–Pb geochronological and coupled Lu–Hf isotope data from zircons that reveal the age of the NSB and provide new constraints on the source rocks to this piece of early crust. Two rocks have been analysed, a felsic schist, interpreted to represent a volcanite, and a mylonitic tonalite that is intrusive into the NSB. The felsic schist was emplaced at 3771+5/−3 Ma, dating the formation of the NSB. The Lu–Hf model age indicates that the parental melt to the felsic schist was extracted from the mantle at this age or within a few tens of million years, but no more than 300 m.y., before its extrusion. This mantle extraction age provides an absolute maximum age for the NSB, but the most probable age of its protocrust is ca. 3.8 Ga. The mylonitic tonalite was emplaced at 3667+3/−1 Ma. Its Lu–Hf isotopic composition reveals the presence of an older crustal component than that of the NSB, requiring that its precursor melt formed from a Hadean source that was older than the NSB. Based on recent models for Archean TTG formation, we propose that this source represents re-melted Hadean hydrated crust that must have been translated below the NSB at ca. 3667 Ma. © 2015 Elsevier B.V. All rights reserved.

1. Introduction The Nuvvuagittuq Supracrustal Belt (NSB) is one of a few known Eoarchean or older crustal fragments preserved in the Earth’s crust (Cates and Mojzsis, 2007; O’Neil et al., 2008; David et al., 2009). These crustal fragments provide unique possibilities to study the evolution of the early Earth and to gain insights into processes that were active during their formation, as for example whether subduction processes were involved in the formation of felsic rocks already in the Eoarchean (e.g. Moyen and Martin, 2012; Turner et al., 2014). Crucial to the study of these rock complexes is knowing their timing of formation, specifically their crystallisation and crustal evolution, and the timing and nature of mantle extraction, differentiation and evolution. As the NSB is highly deformed (as is generally the case for Eoarchean rock complexes) and has

*

Corresponding author at: GEOTOP, Université du Québec à Montréal, CP 8888, Succ. Centre-Ville, Montréal, Québec H3C 3P8, Canada. E-mail address: [email protected] (L.E. Augland). http://dx.doi.org/10.1016/j.epsl.2015.07.039 0012-821X/© 2015 Elsevier B.V. All rights reserved.

gone through several cycles of high grade metamorphism (Cates and Mojzsis, 2007; David et al., 2009), obtaining robust and unambiguous geochronological and isotopic information unravelling the different aspects of its evolution have presented a challenge. In recent years there has been a debate on whether (some of) the gneisses of the NSB were actually crystallised in and incorporated into the crust in the Early Hadean (O’Neil et al., 2008, 2012, 2013) or whether these rocks were formed in the Eoarchean (Cates and Mojzsis, 2007; Cates et al., 2013; Guitreau et al., 2013) and rather sampled an enriched mantle source that had differentiated in the early Hadean as has been suggested for example for the Isua Supracrustal Belt (e.g. Rizo et al., 2011). To shed light on this discussion and test models for the presence or absence of Hadean crust in the NSB or in the source to the NSB felsic rocks, we here present new high precision U–Pb isotope dilution thermal ionisation mass spectrometry (ID-TIMS) and coupled Lu–Hf solution inductively coupled plasma mass spectrometry (S-MC-ICPMS) data from two felsic gneisses originally dated by David et al. (2009). The new data from a felsic schist (1091F) interpreted to be an extrusive volcanic rock (David et al., 2009), provide precise age

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Fig. 1. Simplified geological map of the Nuvvuagittuq Supracrustal Belt, modified from David et al. (2009). Sample locations are indicated with arrows.

constraints on the formation of the NSB. Due to the high resolution and precision of the data obtained with our methods of choice, we can also give much more robust estimates than previous studies on when precursor melts to the NSB crust separated from the mantle, and thus contribute to settle the debate for the protolith age of the NSB. The re-analysis of a tonalite intruding the NSB (POR23; David et al., 2009), further allows the investigation of which source(s) were involved in tonalite–trondhjemite– granodiorite (TTG) melt formation and what processes were responsible for crust-generation in the Eoarchean. 2. Geological setting 2.1. Regional geology The NSB (Fig. 1) occupies a ca. 20 km2 area in the ca. 2.8–2.7 Ga Inukjuak Domain in the Superior Province in northwestern Québec, Canada (David et al., 2009). It comprises a tightly to isoclinally folded, highly sheared and metamorphosed volcano-sedimentary unit, including a banded iron formation, that has been refolded in an open, shallowly plunging large scale fold (David et al., 2009). The NSB is surrounded by a ca. 2750 Ma tonalitic gneiss (Simard et al., 2003) and is in its centre intruded by a metagabbro that was metamorphosed at ca. 2693 Ma (David et al., 2009). The volcanosedimentary sequence itself consists of amphibolites, banded iron formation and mafic and ultramafic sills in addition to minor felsic volcanic and sedimentary units as well as tonalitic intrusive sheets (Fig. 1; Cates and Mojzsis, 2007; David et al., 2009). One of these intrusive sheets, a tonalitic mylonite, was dated by David et al. (2009) to 3659 ± 3 Ma. The NSB rocks have been metamorphosed at upper amphibolite facies at ca. 2.70 Ga and were also metamorphosed under unknown conditions at ca. 3.66 Ga and 3.36 Ga as documented by metamorphic zircons (Cates and Mojzsis, 2009; David et al., 2009).

Other 147 Sm–143 Nd-ages from NSB rocks analysed by David et al. (2009) and Roth et al. (2013) indicate ages of 3924 ± 290 Ma and 3869 ± 130 Ma, respectively, and a Lu–Hf regression of Guitreau et al. (2013) gives an age of 3864 ± 70 Ma. Notably, none of these ages corresponds to the 4.4 Ga age proposed by O’Neil et al. (2008, 2011, 2012, 2013). O’Neil et al. (2012) argued that the discrepancies between the ages calculated from the two Sm–Nd isotope systems was a result of partial resetting of the 147 Sm–143 Nd system while the 146 Sm–142 Nd system remained undisturbed. Roth et al. (2013) showed through quantitative isotope modelling that the observed isotopic patterns (i.e. “scatterchrons”) could not be explained by undisturbed 146 Sm–142 Nd systematics. Guitreau et al. (2013) interpreted the positive correlation between 142 Nd/144 Nd and 147 Sm/144 Nd ratios, that was used to calculate the Hadean “scatterchron” (O’Neil et al., 2011), to represent an inherited isotopic signature from a mantle reservoir that had experienced fractionation and subsequent re-fertilisation in the early Hadean. U–Pb geochronology from different TTG rocks interpreted to be intrusive and/or volcanic have indicated formation of the NSB at or before 3750–3880 Ma (Cates and Mojzsis, 2007; David et al., 2009; Darling et al., 2013; O’Neil et al., 2013) and detrital zircons from what has been interpreted as metasedimentary units of 3780 ± 22 Ma, show that the oldest zircon-bearing units of the NSB are most likely not older than 3800 Ma (Cates and Mojzsis, 2007; David et al., 2009; Cates et al., 2013). Later, laser ablation multicollector inductively coupled plasma mass spectrometry (LA-MCICP-MS) U–Pb and Lu–Hf isotope data from TTGs of the NSB reported by O’Neil et al. (2013) were interpreted to indicate the presence of a Hadean crustal source in the host magma to the zircons, again taken as evidence for the presence of Hadean crust in the NSB. 3. Geochronology and isotope geochemistry 3.1. Choice of analytical method in the context of previous work

2.2. Age of the Nuvvuagittuq Supracrustal Belt O’Neil et al. (2008, 2012) observed a correlation between 142 Nd/144 Nd and 147 Sm/144 Nd, which they interpreted as an isochron. The latter gives an age of 4.3–4.4 Ga, which is interpreted by the same authors as the age of some of the NSB mafic and ultramafic rocks. On the other hand the 147 Sm–143 Ndage for the same rocks is 3819 ± 270 Ma (O’Neil et al., 2012).

Archean TTGs are generally accepted to form by melting of hydrated metabasaltic rocks in the garnet stability field (e.g. Moyen and Martin, 2012). Therefore, pre-existing mafic crust must have been involved in the formation of TTGs and thus one can expect to find traces of this crust in the isotopic record of TTGs, or felsic rocks in general, in ancient gneiss terranes such as the NSB. Several whole rock studies have been conducted in the NSB, but

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as the NSB was metamorphosed at ca. 3.66, 3.36 and 2.70 Ga (Cates and Mojzsis, 2009; David et al., 2009; O’Neil et al., 2012, 2013; Darling et al., 2013), the isotope systems have been disturbed, as was recently shown by Roth et al. (2013). Analysis of Lu–Hf in zircon provides a way around such problems because unrecrystallised zircon preserves the initial Hf-isotopic composition of its igneous host rock even through high temperature metamorphism that may lead to significant Pb-loss (Kinny et al., 1991; Kinny and Maas, 2003; Gerdes and Zeh, 2009; Lenting et al., 2010; Guitreau and Blichert-Toft, 2014). Furthermore, Hf is a compatible element in zircon and generally comprises ca. 0.5 to 2% and initial Lu/Hf ratios in zircon are usually very low, thus making it ideal for the determination of the Hf-isotopic composition in its host rock and the source from which it formed (Patchett, 1983; Kinny and Maas, 2003). The fact that zircons can be precisely dated by U–Pb means that the initial Hf-isotopic composition at the right age can also be determined with great precision and certainty. The combined U–Pb and Lu–Hf isotope study of zircons from the NSB by O’Neil et al. (2013) dated (U–Pb) and measured the Hfisotopic compositions of zircons during two different laser-ablation mass spectrometry sessions, respectively. Essential to this analytical approach is that the same zircon domains are analysed for both isotopic systems, i.e. that only the domain giving the specific U–Pb age is analysed for Lu–Hf. It is also of major importance that the data are sufficiently accurate and precise in both isotopic systems to rule out mixed and discordant ages (U–Pb) and false mixing lines due to spread in data (Lu–Hf). Both of these aspects often hamper the validity of the interpretation of data obtained by LA-ICPMS analyses of zircon (e.g. Gerdes and Zeh, 2009; Lenting et al., 2010; Fisher et al., 2014; Guitreau and Blichert-Toft, 2014) and they result mainly from one or a combination of three methodological limitations. The first such limitation is that different volumes of zircon are consumed during the analyses of U–Pb and Lu–Hf thus necessitating assumptions regarding the homogeneity of the two analysed volumes, i.e. inferring that only one and the same zircon domain is measured for both isotopic systems (something that may be difficult to evaluate as the texture study of zircons before analysis is usually based on two-dimensional CLimages). The second limitation is that the analytical precision and uncertainties stemming from correction procedures can make it difficult to evaluate concordance of the obtained U–Pb data. Slight degrees of discordance in zircons that have experienced more than one potential Pb-loss event leads to notorious difficulties in determining the real age of the zircon as the timing of Pb-loss is required for calculating the age. For that type of LA-MC-ICP-MS data, the best approach is to include the range of ages obtained by calculation of upper intercept ages of the earliest and latest potential Pb-loss events (unless the Pb-loss event is well constrained) in the uncertainty of the data, or simply report minimum ages (Halpin et al., 2012). The third limitation is that the precision and accuracy of LA-MC-ICP-MS Lu–Hf isotope data depend on making the appropriate corrections of large isobaric interferences of 176 Yb and 176 Lu on the low-abundance radiogenic 176 Hf isotope (see e.g. discussions in Griffin et al. (2006), Corfu (2007), Blichert-Toft (2008), Guitreau et al. (2012) and Fisher et al. (2014)). For unknown samples where heterogeneities and differing REE concentrations may be present, corrections for isobaric interferences (that in particular are highly sensitive to mass bias corrections, see discussion in Blichert-Toft (2008)) may lead to a spread in data over several epsilon units that can lead to serious misinterpretation of data. To avoid all these analytical and interpretational pitfalls, we have analysed zircons from two samples reported by David et al. (2009) for U–Pb by isotope dilution thermal ionisation mass spectrometry (ID-TIMS) and for Lu–Hf isotope systematics by solution multi-collector inductively coupled plasma mass spectrometry

(S-MC-ICPMS), respectively, obtaining high precision crystallisation ages and directly coupled Hf-isotopic data. 3.2. U–Pb geochronology The U–Pb analyses were conducted by ID-TIMS at GEOTOP, Université du Québec à Montréal. The samples were crushed and zircons were separated for the study by David et al. (2009) and reanalysed here. Zircon grains were selected based on optical characteristics and comparison with CL-imaged zircons. The selected zircons were annealed for ca. 60 h at 900 ◦ C and subsequently chemically abraded (Mattinson, 2005). The chemical abrasion consisted of a one-step dissolution procedure in 48% HF for 14 to 16 h at 195 ◦ C. The abraded grains best suited for analysis were selected based on their morphology, colour and internal textures and washed in dilute HNO3 , de-ionised water and acetone, using an ultrasonic bath, to remove any contamination. Each sample was subsequently weighed on a microbalance and spiked with a 202 Pb–205 Pb–233 U–235 U tracer. The samples were then dissolved in 48% HF and a drop of 8 N HNO3 in Teflon bombs at 195 ◦ C for 5 days. The U and Pb of the dissolved zircons were chemically separated using standard column chemistry (Krogh, 1973). Isotope measurements followed the standard U–Pb ID-TIMS analytical procedures at GEOTOP outlined in detail in Augland et al. (in press). Analytical errors and corrections were incorporated and propagated using an Excel sheet based on the algorithms of Schmitz and Schoene (2007) and ages were calculated using ISOPLOT (Ludwig, 2003). Errors are reported at a 2σ level, not considering decay constant errors (Table 1) and the decay constants of Jaffey et al. (1971) were used in the age calculations. 3.3. Lu–Hf S-MC-ICP-MS The Lu–Hf analyses were conducted by S-MC-ICP-MS at the Department of Geosciences at the University of Oslo following the routines of Patchett and Tatsumoto (1980) and Augland et al. (2012). The washes from the chemical separation of U and Pb of the dissolved zircons were made up to a 1 N HCl–0.1 N HF solution. A small (ca. 10 μg) aliquot for the determination of Lu/Hf ratios was pipetted out and converted to a 2% HNO3 solution (we follow this procedure as it has been shown that there is no Lu–Hf fractionation during Lu and Hf elution in the U–Pb column chemistry [U. Söderlund pers. comm. to F. Corfu, 2004]). The rest of the solutions were run through micro-columns with cation exchange resin AG50W-X8 (repeatedly washed with 48% HF, 6 N HCl and H2 O) to extract Hf and exclude any ions that isobarically interfere with the Hf-isotopes (Yb and Lu). After the ion exchange chromatography the samples were dried down and treated with aqua regia to remove any potential organic material stemming from the resin. The samples were converted to a 2% HNO3 solution and diluted to give a total Hf-concentration of ca. 50 ppb. The chemically separated sample solutions were measured on an Nu Plasma HR MC-ICP-MS fitted with Nu Instruments’ DSN-100 desolvating nebulizer. The chemically treated Hf separates were measured in 2 blocks of 30 cycles with 5 s integration time. Raw data were corrected for mass discrimination using an exponential law, and the mass discrimination factor for Hf was determined assuming 179 Hf/177 Hf = 0.7325 (Patchett and Tatsumoto, 1980). Lu and Yb were monitored through the measurement of 173 Yb and 175 Lu. The measured 173 Yb and 175 Lu were used to generate interferencecorrected 176 Hf/177 Hf ratios reported in Table 2. Note that these corrections are negligible due to the very small Yb and Lu concentrations (176 Yb/177 Hf < 5 × 10−5 , with one exception being as high as 1 × 10−4 [analysis No. 2 of 1091F], and 176 Lu/177 Hf < 4 × 10−6 ). The 176 Lu/177 Hf ratios were measured on the untreated aliquots with a Bruker Aurora M90 ICP-MS, doing 10 scans per replicate

Th/238 U using Th/U [magma] = 3.

Pb/204 Pb = 15.55 ± 0.13%; f

g

Errors are 2-sigma, propagated using the algorithms of Schmitz and Schoene (2007). Calculations are based on the decay constants of Jaffey et al. (1971). 206 Pb/238 U and 207 Pb/206 Pb ages corrected for initial disequilibrium in

230

207

Abbreviations: a.r.: aspect ratio; zr.: zircon; euh.: euhedral; lrg.: large; metam.: metamict; corr.: correlation; coef.: coefficient; disc.: discordance. a Uncertainty of microbalance is 0.001 mg. b Nominal U and total Pb concentrations subject to uncertainty in weight. c Model Th/U ratio calculated from radiogenic 208 Pb/206 Pb ratio and 207 Pb/235 U age. d Pb∗ and Pbc represent radiogenic and common Pb, respectively. e Corrected for fractionation, spike, and common Pb; up to 2 pg of common Pb was assumed to be procedural blank: 206 Pb/204 Pb = 18.60 ± 0.14%; 1-sigma). Excess over blank was assigned to initial common Pb.

0.87 0.88 0.69 0.91 0.92 0.80 0.93 0.76 0.76 0.85 3548.25 3455.59 3378.50 3467.68 3667.10 3660.98 3661.84 3665.78 3652.20 3665.12 0.944 0.933 0.969 0.950 0.956 0.960 0.929 0.968 0.951 0.939 0.11 0.10 0.08 0.12 0.19 0.16 0.13 0.17 0.12 0.13 0.6887 0.6459 0.6260 0.6591 0.7661 0.7574 0.7570 0.7621 0.7542 0.7599 0.15 0.14 0.12 0.16 0.20 0.18 0.16 0.19 0.15 0.16 0.057 0.057 0.044 0.059 0.060 0.052 0.061 0.050 0.050 0.056 03-POR-23 1 1 euh. metam. h. a.r. zr. 2 1 euh. clear h. a.r. zr. 3 1 stubby subrounded zr. 4 1 subhedral clear zr. fragm. 5 1 euh. metam. m. a.r. zr. 6 1 stubby, euh. metam. zr. 7 1 euh. h. a.r. zr. fragm. 8 1 euh. clear stubby zr. 9 1 euh. metam. stubby zr. 10 1 euh. metam. stubby zr.

0.003 0.002 0.005 0.004 0.001 0.002 0.004 0.005 0.006 0.003

209 175 128 104 109 66 51 42 68 63

0.82 0.79 0.48 0.61 0.89 0.89 1.16 0.75 0.86 0.91

192 152 99 88 119 69 55 43 69 66

4.1 9.5 7.2 5.9 5.9 3.2 3.6 2.4 2.6 2.0

140 31 68 59 19 42 60 90 161 95

0.3155 0.2972 0.2828 0.2995 0.3410 0.3396 0.3398 0.3407 0.3377 0.3405

29.96 26.46 24.41 27.22 36.02 35.46 35.47 35.80 35.11 35.68

0.88 1.2 4.1 1.6 1.9 3731.90 3745.2 3771.5 3763.1 3770.6 0.924 0.990 0.955 0.968 0.972 0.11 0.57 0.89 0.36 0.49 0.7382 0.7673 0.7852 0.7847 0.7924 0.14 0.58 0.92 0.40 0.52 0.058 0.082 0.27 0.10 0.12 1091F 1 2 3 4 5

1 1 1 1 1

euh. metam. high a.r. zr. metam. zr. fragm. euh. stubby metam. zr. stubby subrounded metam. zr. stubby euh. metam. zr. prism.

0.003 0.001 0.001 0.001 0.001

173 24 21 117 52

0.49 0.35 0.70 0.62 0.84

166 25 24 123 61

5.2 1.4 1.9 2.0 4.9

95.6 17.3 12 59 11

0.3558 0.3589 0.3652 0.3631 0.3649

36.21 37.97 39.53 39.29 39.87

±f Pb g

206 Pb

207

% errf Pb e 206 Pb

207

Pb∗ d Pbc Th c U

Ub (ppm) Wt.a (mg)

Zr. characteristics An. no.

Table 1 Zircon U–Pb geochronological data.

Compositional parameters

Pbb (ppm)

Pbc d (pg)

Radiogenic isotope ratios

207

Pb e 235 U

% errf

206

Pb e 238 U

% errf

corr. coef.

Isotopic ages

208P

b/204 Pb = 37.8 ± 0.13% (all uncertainties

4.8 7.0 7.3 5.9 0.0 0.7 0.8 0.4 0.8 0.6 3.0 2.6 2.1 3.2 5.2 4.5 3.6 4.8 3.3 3.7 3377.6 3212.0 3133.6 3263.6 3666.7 3634.7 3633.2 3651.9 3622.8 3643.8 1.5 1.4 1.2 1.6 2.0 1.8 1.6 1.9 1.4 1.5 3485.6 3363.9 3284.8 3391.3 3666.9 3651.7 3651.7 3660.9 3641.8 3657.6

4.5 2.0 0.9 0.8 0.2 3.0 16 25 10 14 3564.1 3671 3736 3734 3762 1.4 5.7 9.1 4.0 5.2 3672.3 3719.1 3759.1 3752.9 3767.5

Pb g 203

Pb g 207

235 U

±f

238 U

±f

Disc. %

L.E. Augland, J. David / Earth and Planetary Science Letters 428 (2015) 162–171

165

and 5 replicates per sample. A dwell time of 10 ms was applied between the individual analyses. Four solutions with different Lu/Hf and Yb/Hf ratios gravimetrically prepared from certified mono-element standards with a known concentration were used as bracketing standards. Measured Lu/Hf ratios were corrected for isobaric interferences of Lu and Hf based on measured 172 Yb and 175 Lu and the measured and corrected 175 Lu/177 Hf was used to calculate the 176 Lu/177 Hf ratios reported in Table 2 based on the terrestrial average 176 Lu/175 Lu of 0.02656 (Blichert-Toft et al., 1997). A value of 1.867 × 10−11 a−1 for the decay constant of 176 Lu was used in all calculations (Scherer et al., 2001; Söderlund et al., 2004) and for the calculations of ε Hf we used present-day chondritic 176 Hf/177 Hf = 0.282785 and 176 Lu/177 Hf = 0.0336 (Bouvier et al., 2008). The depleted mantle model of Griffin et al. (2000), modified to the λ176 Lu and the chondritic compositions given in Bouvier et al. (2008) is adopted. This model produces a present day depleted mantle 176 Hf/177 Hf-value of 0.28325 similar to that of average mid-ocean ridge basalt over 4.56 Ga, from chondritic initial hafnium at 176 Lu/177 HfDM = 0.0388 (Andersen et al., 2009; see Section 4.2.1 for discussion of alternative models). The JMC-475 standard of reference (courtesy of Jonathan Patchett, University of Arizona) was run at frequent intervals obtaining 176 Hf/177 Hf = 0.282150 ± 0.000013 (2σ ; 8 analyses), overlapping the certified value of 0.282163 ± 0.000009 (Blichert-Toft et al., 1997). The external 2σ precision is thus considered to be 0.005%, and when internal precision of the average values is lower than this, a 2σ of 0.005% was used in calculation of initial 176 Hf/177 Hf (176 Hf/177 Hf(t) ) and initial εHf (εHf(t) ). The reported 176 Hf/177 Hf(t) and εHf(t) values are calculated at U–Pb emplacement ages of the rocks and are reported in Table 2. The data presented in Table 2 has been normalised to the certified 176 Hf/177 Hf-value of the JMC-475 and errors have been propagated using the algorithms of Ickert (2013). For the calculation of εHf(t) values his model 1 error propagation scheme was used and for the calculation of 176 Hf/177 Hf(t) we used his model 2 error propagation scheme (Ickert, 2013). 3.4. Sample and zircon characteristics and analytical results 3.4.1. Felsic schist – 1091F This sample was collected from a plagioclase–quartz–biotite schist on the southwestern limb of the major synform (Fig. 1; see David et al. (2009) for sample locality GPS co-ordinates). The felsic schist is concordant, up to 0.5 m thick and interlayered and isoclinally folded with amphibolitic schist and a quartz biotite schist interpreted to be a conglomerate (David et al., 2009). No cross-cutting relationships with the surrounding inferred metasupracrustal rocks were observed and the rock has been interpreted to be a metamorphosed felsic volcanic rock. It was originally dated by David et al. (2009) at 3817 ± 16 Ma, but due to large degrees of discordance and scatter this age is not considered very robust. The zircons in this sample form a relatively homogeneous population of prismatic medium to high aspect ratio euhedral zircons with pyramidal terminations. Some stubby prisms are also present. The zircons range from clear reddish to more metamict and slightly altered brownish appearance. Textures revealed by CLimages show well pronounced oscillatory zonation interpreted to be of primary magmatic origin (David et al., 2009). Five single zircons were analysed in this study (Table 1, Fig. 2). U concentrations of the different zircons range from 21 to 173 ppm and Th/U ratios vary between 0.35 and 0.84. The analyses range from concordant to 4.5% discordant, but three of the zircons are less than 1% discordant. The latter three analyses have the highest Th/U ratios. Analyses Nos. 3 and 5, anchored at 0 ± 100 Ma, give an upper intercept age of 3771 ± 2 Ma (2σ , MSWD = 0.16). Using

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Table 2 Lu–Hf data. The data have been normalised to the certified 176 Hf/177 Hf-value of JMC-475: 0.282163 ± 0.000014 (Blichert-Toft et al., 1997). For POR23 the data above the horizontal line are from zircons that are more than 4% discordant and the data below the horizontal lines are from zircons that are less than 0.8% discordant. The average for POR23 is calculated from the data below the horizontal line (analyses Nos. 5–10). Abbreviations: Min.: minimum; Max.: maximum; No.: number; TDM : depleted mantle model age. Age (Ma)

207

Pb/206 Pb-age (Ma)

176

1091F 1 3 4 5 Average

3772 3772 3772 3772

3731.9 3763.1 3771.5 3763.1

POR23 1 2 3 4 5 6 7 8 9 10 Average

3667 3667 3667 3667 3667 3667 3667 3667 3667 3667

3548.2 3455.6 3378.5 3467.7 3667.1 3661.0 3661.8 3665.8 3652.2 3665.1

No.

Hf/177 Hf

2σ

176

0.280378 0.280380 0.280371 0.280404

0.000014 0.000038 0.000056 0.000016

0.280406 0.280410 0.280439 0.280404 0.280388 0.280366 0.280374 0.280377 0.280368 0.280361

0.000014 0.000014 0.000014 0.000014 0.000021 0.000024 0.000014 0.000014 0.000014 0.000014

Lu/177 Hf

2σ

176

0.00050 0.00048 0.00034 0.00055

0.000028 0.000039 0.000022 0.000019

0.280341 0.280345 0.280347 0.280364 0.280349

0.000015 0.000039 0.000056 0.000017 0.000020

0.000409 0.000769 0.000505 0.000529 0.00067 0.000649 0.000553 0.00050 0.000577 0.000697

0.000058 0.000079 0.000026 0.000049 0.00012 0.000039 0.000019 0.00010 0.000023 0.000085

0.280377 0.280355 0.280403 0.280366 0.280340 0.280320 0.280335 0.280341 0.280327 0.280312 0.280329

0.000015 0.000015 0.000015 0.000015 0.000022 0.000025 0.000015 0.000015 0.000015 0.000015 0.000023

the three collinear data points, 2, 4 and 5 an upper intercept age of 3774 ± 5 Ma (2σ , MSWD = 0.0022 lower intercept 1874 ± 290 Ma) is obtained. The slight discordance of analysis No. 3 is considered a result of modern Pb-loss, while analysis No. 1 must have experienced mixed Pb-loss at a time of metamorphism and modern day. There are no known metamorphic events at ca. 1.87 Ga, but Lu–Hf whole rock-garnet isochrones obtained by O’Neil et al. (2012) give overlapping ages of 1655 Ma and younger, probably related to Proterozoic metamorphic events that could have represented events leading to the observed Pb-loss in the zircons dated here. To be sure to include the range of possible ages for the emplacement of 1091F, we use a conservative error estimate and an average of the two upper intercept ages of 3772+5/−3 Ma. This age is also identical within error to the 207 Pb/206 Pb-age of analysis No. 5 of 3771 ± 2 Ma. The Lu–Hf isotope data from the four analysed single zircons all overlap and 176 Hf/177 Hf(t) vary from 0.280341 to 0.280364 with an average of 0.280349 ± 0.000020 (2σ ) considered to represent the magmatic 176 Hf/177 Hf(t) -value, corresponding to an ε Hf(t) -value of 0.53 ± 0.71 (Table 2, Figs. 2 and 3). 3.4.2. Mylonitic tonalite – POR23 This sample was taken from a mylonitic intrusive tonalite sheet from the western flank of the NSB (Fig. 1; see David et al. (2009) for sample locality GPS co-ordinates) originally dated at 3659 ± 3 Ma (David et al., 2009). Zircons from the mylonitic tonalite are generally low to medium aspect ratio, euhedral, light brown prisms. Some are slightly metamict and altered. A few high aspect ratio zircons also occur. CL-images reveal that these grains have oscillatory zonation and in some cases resorption textures typical of magmatic zircon. Ten single zircons were analysed for the present study (Table 1, Fig. 2). The U concentrations vary from 42 to 209 ppm and Th/U ratios range from 0.48 to 1.16. The analyses are up to 7.3% discordant, but 6 of the analyses are concordant or less than 1% discordant. The latter six analyses have somewhat lower U concentrations than the four remaining analyses. One concordant analysis (No. 5) gives a concordia age of 3667.1 ± 0.9 Ma (2σ , MSWD = 0.028). This age is overlapping an upper intercept of collinear data points Nos. 5, 7, 8, 10 of 3669 ± 1 Ma that have a lower intercept of 1310 ± 220 Ma. O’Neil et al. (2012) obtained ages of ca. 1.2 Ga on Lu–Hf garnet – whole rock isochrones from garnet amphibolites

Hf/177 Hf(t)

2σ

ε Hf(t)

2σ

Min. TDM (Ga)

Max. TDM (Ga)

0.25 0.37 0.45 1.05 0.53

0.66 1.4 2.0 0.72 0.71

3.89 3.89 3.88 3.86

4.08 4.07 4.06 3.99

−0.97 −1.75 −0.02 −1.35 −2.29 −3.01 −2.45 −2.25 −2.75 −3.28 −2.67

0.67 0.68 0.66 0.66 0.90 0.96 0.66 0.70 0.66 0.69 0.83

3.86 3.89 3.82 3.87 3.91 3.94 3.92 3.91 3.93 3.96

4.15 4.24 4.05 4.19 4.30 4.37 4.31 4.29 4.35 4.40

from the NSB possibly indicating disturbance of the isotope systems around this time, consistent with Pb-loss at the lower intercept recorded in these zircons. The collinearity could also be accidental (as they have little spread) and instead represent mixed Pbloss (nevertheless, as these data points are so close to concordance, their upper intercept age will give an age very close to the crystallisation age). The best estimate for the magmatic emplacement age of the tonalite is the concordia age of analysis No. 5 with an enlarged uncertainty of the upper age incorporated from the upper intercept age of analyses Nos. 5, 7, 8 and 10: 3667+3/−1 Ma. The discordance in the data set can be explained by a mixture of modern day Pb-loss and late Archean recrystallisation/Pb-loss as all analyses fall within the triangle defined by a discordia line through analysis No. 5, anchored at modern day, and a discordia line anchored at the documented (Cates and Mojzsis, 2009) metamorphic age of 2.7 Ga (Fig. 2). Because the present data (except from analyses Nos. 1–4) are concordant to only slightly discordant, with less spread than the original data of David et al. (2009), the age obtained here is more precise and slightly older than the previously reported age that required a longer extrapolation to yield an upper intercept age. The degree of discordance in the present data set is most probably linked to Pb-loss for the zircons that are less than 1% discordant and metamorphic recrystallisation for the zircons that are more discordant. This interpretation is based on the Hf isotopic data showing that the concordant to less than 1% discordant zircons (Nos. 5–10) cluster at a lower 176 Hf/177 Hf(t) than the zircons that are more discordant (Nos. 1–4; Fig. 2; Table 2; Section 4.2). The Lu–Hf isotope data cluster in two groups (Table 2, Figs. 2 and 3), one consisting of analyses that are more than 4% discordant (Nos. 1–4) and the other of analyses that are less than 1% discordant (Nos. 5–10). The 176 Hf/177 Hf(t) for the zircons that are less than 1% discordant vary between 0.280312 and 0.280341, corresponding to variations in the εHf(t) -value from −3.28 to −2.25. 4. Discussion 4.1. Age constraints from U–Pb data The age of the felsic schist (1091F) refines and confirms the result from David et al. (2009). Due to the careful selection of zircons and the lesser degree of discordance resulting from the chemical

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Fig. 3. ε Hf(t) against age plot. The data points are plotted at their respective crystallisation ages determined by U–Pb geochronology. The vertical yellow dotted lines mark the minimum and maximum calculated depleted mantle model ages for 1091F and the vertical grey dotted lines mark the minimum and maximum calculated depleted mantle model ages for POR23. The light blue bar represents the range of ε Hf(t) values from amphibolites analysed by Guitreau et al. (2013), the yellow star with error bars represents the average and range of ε Hf(t) values of garnet-free amphibolites analysed by O’Neil et al. (2013), and grey bars represent zircon ε Hf(t) values of TTG zircons analysed by O’Neil et al. (2013).

felsic band similar to that from which 1091F was sampled, had intrusive relationships to mafic rocks of the NSB – that also the felsic schist dated here and by David et al. (2009) is of intrusive origin despite the observations reported in David et al. (2009) of this felsic schist being concordant with surrounding mafic and metasedimentary rocks. Due to the high degrees of deformation and the transposed nature of different units in the NSB, it is not possible to determine unequivocally whether the protolith of this felsic schist was a volcanic or an intrusive rock. However, based on the fact that it formed part of a concordantly layered sequence of amphibolites, banded iron formation and a conglomerate of clearly supracrustal origin (Kitayama et al., 2012), we favour the former interpretation. The new protolith age for the mylonitic tonalite (POR23) of 3667+3/−1 Ma largely confirms the age of David et al. (2009) but puts tighter and more robust age constraints on the intrusive event it represents. 4.2. Constraints from Lu–Hf isotope data For the felsic schist (1091F), there is no correlation between Pb/206 Pb ages and the 176 Hf/177 Hf(t) for the individual analyses (Fig. 2; Table 2) and we thus attribute the slight discordance observed in the U–Pb data to have been the result of Pbloss and not metamorphic recrystallisation. This interpretation is also in agreement with the textural information obtained from CL-imaging (David et al., 2009). As there is no spread in the 176 Hf/177 Hf(t) beyond the uncertainties of the individual analyses and no correlation of 176 Hf/177 Hf(t) with 207 Pb/206 Pb ages for the felsic schist (1091F), we interpret the average 176 Hf/177 Hf(t) ratio of 0.280349 ± 0.000020 (2σ ), corresponding to an εHf(t) -value of 0.53 ± 0.71 (2σ ), to represent the magmatic ratio of the volcanic protolith to the felsic schist. Furthermore, the uniformity of the 176 Hf/177 Hf(t) ratios indicates that the parent melt to the volcanic protolith formed from a homogeneous source (Fig. 2; Table 2). For the mylonitic tonalite (POR23), the consistently higher 176 Hf/177 Hf(t) of the >4% discordant zircons (Nos. 1–4) compared to those of the zircons that were <1 % discordant (Nos. 5–10; Fig. 2; Table 2) leads to the conclusion that these zircons had their Lu–Hf isotopic composition altered by post-magmatic metamorphic recrystallisation or zircon growth. The group of concordant to subconcordant zircons have 176 Hf/177 Hf(t) where the lowest value of 0.280312 ± 0.000015 (2σ ), corresponding to an 207

Fig. 2. U–Pb concordia and 207 Pb/206 Pb age vs. 176 Hf/177 Hf(t) diagrams. All analyses are plotted with 2σ absolute errors. Reported ages are calculated excluding decay constant errors. Analysis numbers correspond to those in Tables 1 and 2 and are written next to the individual error ellipses in the U–Pb concordia diagram and below the error bars in the 207 Pb/206 Pb age vs. 176 Hf/177 Hf(t) diagram. Error bars of 207 Pb/206 Pb ages in the 207 Pb/206 Pb age vs. 176 Hf/177 Hf(t) diagram of analyses Nos. 1–4 from POR23 are located within the individual analysis points. The upper error bar of analysis No. 3 is equal to the lower error bar. The 176 Hf/177 Hf(t) values are calculated at the crystallisation ages determined by U–Pb geochronology.

abrasion procedure, much of the scatter observed in the data set of David et al. (2009) is eliminated and a robust and precise age is obtained. Emplacement at 3772+5/−3 Ma of this felsic, presumably volcanic unit provides firm time constraints on the formation of (at least parts of) the NSB that was formed synchronously. This age is in disagreement with the minimum age proposed by Darling et al. (2013) for the NSB of 3800 Ma. The latter authors claim – based on the observation of Cates and Mojzsis (2007) that one

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εHf(t) -value of −3.28 (No. 10), is just overlapping the analysis with the highest value of 0.280341 ± 0.000015 (2σ ), corresponding to an εHf(t) -value of −2.25 (No. 8). There is, however, no correlation between 176 Hf/177 Hf(t) and 207 Pb/206 Pb ages for these zircons (Fig. 2), so the 176 Hf/177 Hf(t) of this group of zircons are considered to represent magmatic values. The slight spread in magmatic 176 Hf/177 Hf(t) observed for the tonalite (POR23) is best interpreted as the result of the parent melt formed from a source with heterogeneous isotopic signatures or by a mixed source with two different isotope signatures. For reasons of clarity and simplicity we will refer to εHf(t) values in the remainder of the text. 4.3. Source age constraints from Lu–Hf isotope data In order to understand the formation history of the NSB and to gain insight into the petrogenesis of the rocks studied here, we will, in this section, discuss and use model ages calculated from the Lu–Hf and U–Pb data (Fig. 3; Table 2). The model ages presented also reveal information on the pre-formation history of the NSB and thus enable the evaluation of geodynamic models for Eoarchean crust-formation. 4.3.1. Choice of source evolution parameters In calculating the Lu–Hf mantle separation model ages for the NSB samples analysed here there are two main considerations to make. The first is which Lu/Hf ratios are to be considered representative for the source of the rocks analysed and the second is which initial Hf-isotopic composition to use for the chondritic uniform reservoir (CHUR; or bulk silicate Earth: BSE; Bouvier et al., 2008; Bizzaro et al., 2012; Iizuka et al., 2015). Amelin et al. (1999) and subsequently Kemp et al. (2010) suggested a 176 Lu/177 Hf of ca. 0.02 as an average value for mafic crust sourcing ca. 4.3 to 3.9 Ga zircons from the Jack Hills conglomerate, whereas Blichert-Toft and Albarède (2008) suggested a TTG source with 176 Lu/177 Hf of less than 0.01 for ca. 4.1 Ga Jack Hills zircons. The estimate of the 176 Lu/177 Hf of the source to the NSB felsic schist and mylonitic tonalite analysed here thus clearly depends on what the source is considered to be (e.g. mafic or tonalitic crust). To be sure to include the possible range of model ages for the NSB zircons analysed here, we have calculated model ages using a maximum 176 Lu/177 Hf of 0.026, corresponding to the average of the enriched low-Ti metamafic rocks of the NSB (Adam et al., 2012; O’Neil et al., 2013; Section 4.3.2), and a minimum equal to that of the lowest recorded 176 Lu/177 Hf for a low-Ti enriched metamafic rock from the NSB of 0.007 (Fig. 3; Guitreau et al., 2013). The enriched low-Ti series of O’Neil et al. (2011) was shown to be a possible source of Archean TTGs by Adam et al. (2012); it should be noted that garnet-free low-Ti amphibolites have 176 Lu/177 Hf of up to 0.029; however, using this value yields model ages older than the Earth for POR23, so we do not consider such a high 176 Lu/177 Hf in our calculation of potential model ages. Present day MORB results from melting of a depleted mantle that has experienced continuous depletion of incompatible elements throughout the Earth’s history, and can thus not be used as a measure of the mantle melting to create Hadean and Eoarchean crust. Furthermore, as noted by Dhuime et al. (2011) and subsequently by Guitreau et al. (2012), it seems that mantle involved in major crustal events, especially manifested by Archean TTG rocks, were of a more enriched source than the depleted mantle generating MORB, questioning the use of the depleted mantle as a general reference when calculating mantle extraction model ages. A further necessary consideration to make is the possibility that the source rocks from which 1091F formed had a near-CHUR Lu/Hf. If that was the case, it would not be possible to calculate a meaningful mantle separation model age as the source rocks would just have evolved along the CHUR evolution line (Fig. 3). However, as seen

from the whole-rock Lu–Hf isotope data of Guitreau et al. (2013) and O’Neil et al. (2013; garnet-free amphibolites), most amphibolites analysed have positive εHf(t) values relative to CHUR (Fig. 3), indicating that they formed from a mantle source that was depleted relative to CHUR. Also the zircons analysed from 1091F have positive εHf(t) values (Fig. 3) so it is sound to assume that the precursor source rocks to 1091F (and POR23) were formed from a mantle source that was depleted relative to CHUR. Taking into account that the mantle generating the precursor melts to the NSB was most likely enriched relative to the depleted mantle, but depleted relative to CHUR, the depleted mantle model ages reported here are considered maximum ages. The choice of CHUR Lu–Hf parameters also changes the calculated model ages and even though Bizzaro et al. (2012) recently proposed a new, lower CHUR initial Hf-isotopic composition, most workers still use the parameters of Bouvier et al. (2008) which has recently been corroborated by Iizuka et al. (2015) which obtained a CHUR initial of 176 Hf/177 Hf of 0.282793 (±0.000018), overlapping that of Bouvier et al. (2008) of 0.282785 (±0.000011) and that in the work of Blichert-Toft and Albarède (1997; 0.282772 ± 0.000029). The choice of these parameters affects the 176 Hf/177 Hf(t) , εHf(t) , and model ages, and is thus a highly important consideration to make when comparing Lu–Hf model ages with model ages from other isotope systems (e.g. Sm–Nd), but for the comparison with other Lu–Hf isotope data the relative differences, similarities and patterns remain the same using either value (as long as the same parameters are used consistently in comparison of data). We thus plot the data using the parameters of Bouvier et al. (2008) to make it easier for the reader to compare with other published data. We also note that the analyses from the felsic schist would plot above the depleted mantle curve using the parameters of Bizzaro et al. (2012), requiring an extremely depleted source which does not seem to be compatible with the compositions of other NSB rocks (e.g. Guitreau et al., 2013). 4.3.2. Age constraints on the NSB from Lu–Hf isotope data The felsic schist (1091F) has minimum and maximum depleted mantle model ages of 3.86 and 4.08 Ga, respectively, assuming a minimum 176 Lu/177 Hf of 0.007 and a maximum 176 Lu/177 Hf of 0.026 of the crust from which it was sourced (Fig. 3). Using the calculated maximum model age of 4.08 Ga leaves ca. 300 m.y., or less assuming a lower 176 Lu/177 Hf of the source crust and/or a less depleted mantle source, between mantle extraction and generation of the melt from which the rock crystallised. Interestingly, the 176 Hf/177 Hf(t) for the felsic schist is 0.280349 ± 0.000020 (2σ ) overlapping with the initial 176 Hf/177 Hf calculated from a global scatterchron of a range of rocks from the NSB (0.280326 ± 0.000026) and interpreted to represent the 176 Hf/177 Hf(t) of the NSB at 3.8 Ga (Guitreau et al., 2013). The overlap in age and 176 Hf/177 Hf(t) for the NSB determined by Guitreau et al. (2013) with that of the felsic schist dated here, indicates that this rock formed part of the supracrustal sequence of the NSB and that it actually dates formation of (at least part of) the NSB. If, on the other hand the felsic schist crystallised from a heterogeneous melt with a mixed source of juvenile melt extracted from the mantle input and NSB crust, the NSB crustal source may have been older. The latter possibility is, however, unlikely as there is no spread in the data from this sample, something that is expected if there was a mixing of two clearly different sources. The 176 Hf/177 Hf(t) , when time corrected, also overlap the range of 176 Hf/177 Hf(t) of NSB amphibolites reported by Guitreau et al. (2013) and O’Neil et al. (2013; garnet free amphibolites), showing that an older source than 3.8 Ga is unlikely (Fig. 3). If one assumes that the NSB and thus the felsic schist formed by melting of an enriched mantle with near-chondritic Lu/Hf as suggested by Guitreau et al. (2012; see also discussion in Dhuime

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et al., 2011), the U–Pb age of 1091F dates, or is only a few tens of millions of years younger than, juvenile crust formation in the NSB and most likely approximates closely the age of the NSB. It should also be mentioned that the average ε Hf(t) -value of the felsic schist overlaps the average ε Hf(t) -value determined by LA-MC-ICP-MS of the oldest zircons analysed by O’Neil et al. (2013; Fig. 3). However, their data show a very large spread, probably due to the problems described in Section 3.1, and any comparison beyond the fact that the ε Hf(t) values overlap is not informative (the same observation is valid for the younger mylonitic tonalite (POR23), so the zircon Lu–Hf isotope data of O’Neil et al. (2013) will not be discussed further). Although, in our view, the most likely interpretation of the field relationships is that the felsic schist (1091F) is a supracrustal rock in the NSB, the alternative interpretation that this rock represents an intrusive sheet into the NSB also needs to be explored. If the felsic schist was an intrusive sheet into the NSB, the fact that it has identical isotopic compositions to the NSB (Fig. 3; Guitreau et al., 2013), indicates that its melt was formed by re-melting of the NSB without any traceable juvenile (mantle) addition. This again most likely precludes any older crustal component in the NSB than the maximum provided by the mantle separation model age of 4.08 Ga assuming 176 Lu/177 Hf of 0.026, and is most consistent with juvenile crust formation at ca. 3.8 Ga considering that the mantle the NSB rocks were sourced from was enriched relative to the depleted mantle (e.g. Guitreau et al., 2012). The mylonitic tonalite (POR23) has minimum and maximum depleted mantle model ages of 3.96 and 4.40 Ga, respectively, assuming a minimum 176 Lu/177 Hf of 0.007 and a maximum 176 Lu/177 Hf of 0.026 (Fig. 3). Interestingly, the younger tonalite (POR23) aged 3667+3/−1 Ma thus has an older model age than the 3772+5/−3 Ma felsic schist interpreted to be of volcanic origin. However, these model ages may not be geologically meaningful considering the possibility that the tonalite might have formed by re-melting of the low-Ti enriched amphibolites with an arc-like geochemistry present in the NSB (Adam et al., 2012). The average ε Hf(t) -value at 3800 Ma of garnet-free low-Ti enriched amphibolites analysed by O’Neil et al. (2013) is 1.06 with a range of values from −2 to 6 (Fig. 3). As the Lu–Hf system is sensitive to garnet growth and the garnet amphibolites of the NSB define an isochron of ca. 2.5 Ga (and in individual cases younger) interpreted to reflect a metamorphic age (O’Neil et al., 2012, 2013), we conclude that the garnet amphibolites have not retained their magmatic Lu– Hf isotope signatures. They are thus not used for comparison with our data. Using either the average 176 Lu/177 Hf-ratio of the low-Ti enriched amphibolites of 0.026 or the minimum value of 0.007, the evolution curve going from the average ε Hf(t) -value of the lowTi enriched amphibolites does not overlap with the ε Hf(t) values of the 3667 Ma mylonitic tonalite (Fig. 3). The same is the case for the amphibolite Lu–Hf isotope data reported by Guitreau et al. (2013; Fig. 3). However, there is an overlap of our data from the mylonitic tonalite with the two lowest ε Hf(t) -values (−1.86 and −2.06) obtained by O’Neil et al. (2013), but as those samples are clearly different from the rest of the low-Ti enriched amphibolites he analysed (which have ε Hf(t) -values of −0.56 to 6.06, comparable to the values obtained by Guitreau et al. (2013)), it does not appear probable that they are representative for the NSB low-Ti amphibolites that could have potentially re-melted to form the tonalite precursor melt. Furthermore, they have 176 Lu/177 Hf of 0.023 and 0.026, respectively, and using these values for the Lu/Hf evolution curve, there is no overlap at a 2σ level with the lowest ε Hf(t) -value of POR23 (an. No. 10). It thus seems clear that the mylonitic tonalite could not have formed by re-melting of the enriched lowTi metamafic rocks of the NSB as indicated by Adam et al. (2012). The mylonitic tonalite could neither have formed by re-melting of older felsic rocks from the NSB, even considering the lowest ex-

169

treme 176 Lu/177 Hf of 0.002 measured in a felsic rock from the NSB, as an evolution curve with this trend going through the data for the felsic schist (1091F) does not overlap with the data from the mylonitic tonalite (POR23; an. No. 8). It should also be noted that the melt experiments of Adam et al. (2012) indicate that metamorphic conditions required to form a melt with the appropriate composition would have been around 2 GPa and 1050 ◦ C, conditions there is no existing geological evidence the NSB experienced. It is thus clear that an older and different crustal source than that present in the NSB must have been involved in the formation of the precursor melt to the mylonitic tonalite (POR23). This precursor melt was then probably contaminated by the isotopically more juvenile NSB crust during emplacement into the NSB crust, explaining the observed spread in the magmatic ε Hf(t) -values for the mylonitic tonalite. The age of the source that was partially melted to form the tonalite precursor melt is dependent on the 176 Lu/177 Hf ratio of the source rocks. If one assumes that the 176 Lu/177 Hf ratio of the source rocks lies between that suggested for a basaltic source to Hadean crust equivalent to that sourcing the Jack Hills Hadean zircons (0.020; Kemp et al., 2010) and that of the average of the enriched low-Ti amphibolites from the NSB (0.026), maximum depleted mantle model ages in the range 4.17–4.40 Ga can be calculated for the source to the mylonitic tonalite. This corresponds to CHUR model ages of 4.0 to 4.25 Ga. It is thus highly likely that Hadean crust older than 4.0 and possibly up to 4.40 Ga was partially molten to eventually crystallise the tonalite (POR23). 4.4. Tectonic setting during formation of the NSB and intrusion of the 3667 Ma mylonitic tonalite Recently, Touboul et al. (2014), based on W isotopes and highly siderophile element (HSE) concentrations in mafic and ultramafic rocks of the NSB, suggested that the precursor melts to the NSB rocks had formed by melting of a fluid-enriched (metasomatised) mantle similar to a typical modern-day supra-subduction mantle. To explain the measured 182 W excess of +6 to +17 ppm and modern-day HSE concentrations in the NSB crust, Touboul et al. (2014) suggested that the re-melted mantle generating the NSB had been infiltrated by fluids derived from crust that had previously been recycled into the mantle by subduction or delamination. Furthermore, Turner et al. (2014) suggested that the NSB stratigraphy of mafic rocks and their geochemistry have great similarities with the stratigraphy and geochemistry of rocks formed during inception of the Izu–Bonin–Mariana arc. Based on this similarity they interpreted the NSB to have formed in a fore-arc setting. The Lu–Hf isotope data from the felsic schist (1091F) do not put any direct constraints on the tectonic environment during formation of the NSB, but they are consistent with its formation from an enriched mantle source. The Lu–Hf isotope data from the ca. 3667 Ma tonalite (POR23) intruding the NSB do, on the other hand, provide important constraints on the tectonic setting during their formation. The negative ε Hf(t) values from this rock requires the presence of an older source involved in the NSB melt-generation and cannot be explained solely by melting of a metasomatised mantle as in the case for the NSB mafic rocks in the model of Touboul et al. (2014). Martin et al. (2014) recently used geochemical and isotope characteristics to show that the most probable source for Archean TTG rocks was melts generated from hydrated oceanic plateau basalt in the garnet stability field (at pressures ≥1 GPa) that interacted with peridotite. The most likely environment where hydrated oceanic plateau basalt can be brought to sufficient depth, partially melt and interact with peridotite is a subduction environment. Whether or not it was oceanic plateau basalt that was the crust being melted to produce the source to the ca. 3667 Ma mylonitic tonalite (POR23) cannot be constrained

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Fig. 4. Schematic model for the geodynamic setting during formation of the ca. 3667 Ma tonalites in the Nuvvuagittuq Supracrustal Belt. The stratigraphy and legend is schematic and the whole stratigraphy of the NSB is not portrayed.

by the Lu–Hf isotope data reported here. The geodynamic setting proposed by Martin et al. (2014) for Archean TTG formation would, however, provide an explanation for the isotopic compositions recorded by the mylonitic tonalite. It also fulfils the requirement of addition of melt generated from older crust (>4.0 Ga) than the NSB itself that must have been introduced by horizontal transport to a position below the NSB in order to explain its intrusive nature into the NSB. We therefore consider the most likely setting during intrusion of the ca. 3667 Ma mylonitic tonalite (POR23) and related rocks to have been a suprasubduction-zone setting (Fig. 4), supporting other observations indicating that subduction was already active in the Eoarchean (e.g. Martin et al., 2014; Turner et al., 2014). 5. Conclusion This contribution reports new high-precision U–Pb geochronology and coupled Lu–Hf isotope data from a felsic schist, interpreted to represent a volcanite, and a mylonitic tonalite from the Nuvvuagittuq Supracrustal Belt (NSB). The new data show that the NSB most likely formed at 3772+5/−3 Ma and that the parental melt to the felsic volcanite was probably extracted from the mantle within a few tens of m.y. within its intrusion considering that its precursor source rocks formed from an enriched mantle source. This indicates that the NSB is not a Hadean sequence as previously suggested. The Lu–Hf isotope data from the 3667+3/−1 Ma mylonitic tonalite intrusive into the NSB show the presence of an older crustal component not present in the NSB and we interpret this source to represent re-melted subducted Hadean (>4.0 Ga) crust, thus supporting the idea that subduction was active in the Eoarchean. Acknowledgements Thanks to T. Andersen and F. Corfu for providing laboratory access at Oslo University and to S.L. Simonsen for technical assistance during the Lu–Hf analytical sessions. The constructive comments provided by two anonymous reviewers greatly improved the paper. The first author was funded by a grant from the Ministère de l’Énergie et des Ressources naturelles, Quebéc. References Adam, J., Rushmer, T., O’Neil, J., Francis, D., 2012. Hadean greenstones from the Nuvvuagittuq fold belt and the origin of the Earth’s early continental crust. Geology 40 (4), 363–366. Amelin, Y., Lee, D.C., Halliday, A.N., Pidgeon, R.T., 1999. Nature of the Earth’s earliest crust from hafnium isotopes in single detrital zircons. Nature 399 (6733), 252–255.

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