U–Pb geochronology of 3810–3630 Ma granitoid rocks south of the Isua greenstone belt, southern West Greenland

Precambrian Research 126 (2003) 235–257

U–Pb geochronology of 3810–3630 Ma granitoid rocks south of the Isua greenstone belt, southern West Greenland J.L. Crowley∗ Department of Earth Sciences, Memorial University of Newfoundland, St. John’s, Canada NF A1B 3X5 Accepted 5 February 2003

Abstract The oldest well preserved granitoid rocks on Earth envelope the Isua greenstone belt, southern West Greenland. The timing and nature of geological events that affected granitoid rocks south of the Isua belt (“southern gneisses”) are refined with new mapping and U–Pb geochronology. These results and previous work allow for a detailed comparison between the southern gneisses and similar granitoid rocks north of the arcuate Isua belt (“northern gneisses”). Zircons from tonalitic gneiss and meta-tonalite that yielded ages of 3810–3795 Ma are interpreted as having grown during igneous crystallisation of their protoliths based on the consistency of ages within and between samples, and the oscillatory internal zoning that is typical of magmatic zircons. Two dated tonalitic rocks are particularly important because they occur as concordant layers within supracrustal rocks: amphibolite with deformed pillow lava structures in the southwestern part of the Isua belt and quartz-rich gneiss (likely derived from a chemical sediment) that forms an enclave in the southern gneisses. Although these layers lack cross-cutting intrusive relationships, they are interpreted as having intruded as sheets into >3800 Ma supracrustal rocks. Most tonalite was converted into gneiss before intrusion of diorite dated at 3658.3±1.2 Ma and deformed again during or shortly after intrusion of 3650–3630 Ma granitic rocks, whereas a few small lenticular zones of meta-tonalite escaped nearly all strain. The timing of amphibolite facies metamorphism may be recorded by zircon and titanite growth in amphibolitic gneiss at 3632–3620 Ma. Other thermal events were recorded by zircon growth at ∼3550 and 2660 Ma, and titanite growth or Pb loss at ∼2650 and 2600 Ma. These new data and interpretations suggest that the southern gneisses underwent many geological events along with the northern gneisses, including >3658 Ma deformation, 3658 Ma diorite intrusion and 3655–3630 Ma crystallisation of granitic rocks during significant metamorphism and deformation. The major differences are that most tonalite in the southern gneisses is 100 my older than the northern tonalite, and the intensity of the Neoarchaean tectonothermal activity increased southward. Any large-scale juxtaposition of the southern gneisses against the northern gneisses, across the Isua belt, likely occurred before intrusion of the 3658 Ma diorite. © 2003 Elsevier Science B.V. All rights reserved. Keywords: Greenland; Archaean; U–Pb geochronology; Zircon; Gneiss; Isua greenstone belt

1. Introduction

∗ Present address: Department of Earth, Atmospheric and Planetary Sciences, Massachusetts Institute of Technology, 77 Massachusetts Avenue, Cambridge, MA 02139, USA. Tel.: +1-617-452-2784; fax: +1-617-253-6735. E-mail address: [email protected] (J.L. Crowley).

Paleoarchaean tonalitic and granitic rocks that lie immediately north and south of the Isua greenstone belt, southern West Greenland, were little affected by deformation and metamorphism recorded by similar rocks elsewhere in the Godthåbsfjord region (Bridgwater and McGregor, 1974; Moorbath et al.,

0301-9268/$ – see front matter © 2003 Elsevier Science B.V. All rights reserved. doi:10.1016/S0301-9268(03)00097-4

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The aim of this study is to better define the sequence of tectonic events in rocks that border the acruate Isua greenstone belt to the southwest, around lake 682 (Fig. 2), by building upon the previous work and recently gained experience with rocks north of the Isua belt (Crowley et al., 2002). Dating of zircon from tonalitic and granitic rocks by the U–Pb isotope dilution method was undertaken with the goal of obtaining precise magmatic ages while avoiding the problems of inheritance and metamictisation. Metamorphic zircon and titanite were dated to determine the timing of thermal events and the degree to which the rocks were affected by Neoarchaean metamorphism. This work has led to a fuller understanding of the geological history and allows for a detailed comparison between rocks south of the Isua greenstone belt, supracrustal rocks within the belt, and granitoid rocks that lie north of the Isua belt.

2. Regional geology Fig. 1. Geological map of the Godthåbsfjord region. Locations of the Paleoarchaean rocks and terrane boundaries are after Allaart (1982) and McGregor et al. (1991), respectively. The inset map of Greenland locates Nuuk and the main regions of Archaean gneiss that escaped pervasive ductile Proterozoic deformation.

1975; Bridgwater et al., 1976; Nutman et al., 1996, 1999, 2002; Friend et al., 2002; Crowley et al., 2002; Figs. 1 and 2). Ancient rocks with preserved primary features are obviously important because they provide some of the few insights we have into the early stages of Earth evolution. For example, recent studies in the vicinity of the Isua greenstone belt have concluded that plate tectonic processes were active in the Paleoarchaean (Hanmer et al., 2002; Nutman et al., 2002). Locally weakly deformed tonalitic rocks south of the Isua greenstone belt have been investigated in considerable detail; protolith ages of the dominant meta-tonalites are ∼3800 and 3700 Ma and a substantial tectonothermal event occurred at 3650–3620 Ma (Nutman et al., 1999, 2002; Friend et al., 2002). However, many aspects of the geological history of these rocks remain poorly known, including the ages of the supracrustal enclaves and minor intrusive phases, the kinematics of deformation, and the timing and nature of tectonothermal events.

The Archaean gneiss complex in southern West Greenland, in the vicinity of Nuuk (Fig. 1), was mostly derived from Meso- and Neoarchaean tonalite and granodiorite, with minor granite and diorite, that were intruded into basaltic volcanic and sedimentary rocks (Bridgwater et al., 1976). In a 25–75 km wide belt that extends for 200 km through Godthåbsfjord (Fig. 1), Neoarchaean tonalitic and granitic rocks were tectonically interleaved with and intruded into Paleoarchaean tonalitic gneisses, and Paleo- and Mesoarchaean supracrustal rocks. This belt contains the Earth’s most extensive, best exposed, and most intensely studied Paleoarchaean rocks (Nutman et al., 1996, and references therein). McGregor (1973) showed that two groups of tonalitic gneisses could be distinguished by the presence or absence of mafic dykes, termed Ameralik dykes, that crystallised between magmatic and deformation episodes. A protolith age of 3900–3600 Ma for gneisses that predated the Ameralik dykes was first established by Rb–Sr and Pb–Pb whole rock dating and U–Pb zircon dating (see Kamber et al., 2001 for a review of the early dating). Supracrustal enclaves within the gneisses were shown to be of similar age or older. Protoliths of the gneisses that postdated the Ameralik dykes were intruded nearly a billion years

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Fig. 2. Geological map showing the northern and southern gneisses (after Nutman, 1986) and location of U–Pb geochronology samples 1–10. A and B are localities of Nutman et al. (1999) referred to in the text.

later. Further mapping and geochronology led to the recognition of three tectonostratigraphic terranes that were assembled by ∼2.72 Ga (Friend et al., 1996, and references therein). These terranes were interpreted as originally independent rafts of continental crust, distinguished by rocks with different protolith ages and metamorphic histories. The Paleoarchaean gneisses and Ameralik dykes are confined to the Akulleq ter-

rane (Fig. 1). Nutman et al. (1996) introduced the term Itsaq Gneiss Complex to include all of the diverse Paleoarchean rocks in the Akulleq terrane. Protoliths of tonalitic rocks in the complex were interpreted, based on SHRIMP U–Pb zircon dating, as having crystallised during numerous events between 3.87 and 3.63 Ga (Nutman et al., 1996, 1999, 2000, 2002). This interpretation for some rocks in outer Godthåbsfjord

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was challenged by Kamber and Moorbath (1998) and Whitehouse et al. (1999), who proposed that zircons yielding dates older than 3.65 Ga are xenocrystic. The largest package of Paleoarchaean supracrustal rocks lies near the ice cap at Isukasia (Figs. 1 and 2). This package, termed the Isua supracrustal belt by some workers (e.g. Bridgwater and McGregor, 1974) and the Isua greenstone belt by others (e.g. Appel et al., 1998; for brevity referred to as the Isua belt), contains Earth’s best preserved >3.7 Ga supracrustal rocks. The main protoliths are basaltic and high-Mg basaltic pillow lava and pillow lava breccia, chert, banded iron formation, and a minor component of clastic sedimentary rocks derived from chert and basaltic volcanic rocks (Appel et al., 1998; Komiya et al., 1999; Myers, 2001, and references therein). Much of the belt was severely metasomatised (Rosing et al., 1996). U–Pb dating suggests that there are at least two unrelated supracrustal packages that formed at >3.79 and ∼3.71 Ga (Nutman et al., 1997, and references therein). The Isua belt forms a semicircle, 25 km in diameter, of steeply dipping schists around a core of mostly tonalitic to granitic rocks termed the “northern gneisses” (Fig. 2). Similar granitoid rocks south of the belt are the “southern gneisses.” Contacts between the Isua belt and the gneisses are generally strongly deformed, locally mylonitic, especially along the inner contact with the northern gneisses. Based on the presence of layers of gneiss within the belt, Bridgwater and McGregor (1974) interpreted the protoliths of the gneisses as having intruded into the belt, but they allowed that some gneisses could be older. Many subsequent workers have also interpreted the protoliths of the gneisses as originally intrusive into the belt; thus, dating of layers within the belt was used to define a minimum age on the Isua belt (Nutman et al., 1996, 1997, 1999, 2002). The northern gneisses have been mapped (Nutman, 1986; Fig. 2), described (Nutman and Bridgwater, 1986), geochemically analysed (Nutman and Bridgwater, 1986) and dated (Moorbath et al., 1972, 1975; Baadsgaard, 1983; Nutman et al., 1996, 2000; Crowley et al., 2002). This work showed that tonalitic gneiss with a protolith age of ∼3700 Ma dominates over ∼3660 Ma diorite and 3655–3640 Ma granite and pegmatite. Although some deformation occurred before 3660 Ma, the main regional event coincided

with intrusion of the granite and pegmatite. A large area of weak regional strain exists in the middle of the northern gneisses. The southern gneisses have also been mapped (Nutman, 1986) and nearly a dozen tonalitic and quartz dioritic rocks collected around lake 682 (Fig. 2) have yielded ages of ∼3820–3795 Ma (Nutman et al., 1993, 1996, 1999, 2002; Friend et al., 2002) and two others were dated at ∼3700–3690 Ma (Nutman et al., 1993, 2002). Mafic and ultramafic rocks similar to those in the Isua belt typically form large enclaves, while rocks similar to the metachert and banded iron formation in the belt form smaller enclaves. One ultramafic enclave was shown to be >3800 Ma based on dating of a cross-cutting tonalitic rock (Friend et al., 2002). The southern gneisses were affected by multiple deformation events, including one that juxtaposed mafic and ultramafic rocks before 3800 Ma (Friend et al., 2002) and a major period of intercalation along mylonite zones at 3650–3600 Ma (Nutman et al., 2002). Finite strain around lake 682 was highly heterogeneous, with strongly deformed gneisses lying adjacent to small lenses of meta-tonalite that preserve igneous textures. Four localities of little strained rocks were studied in detail by Nutman et al. (1999), and two of them were examined during this study. Metamorphic zircon grains and overgrowths were dated at 3650–3620 Ma (Friend et al., 2002). The Itsaq Gneiss Complex was affected by several, strong thermotectonic and magmatic events in the Paleo- and Neoarchaean (e.g. McGregor, 1973; Bridgwater et al., 1976; Friend et al., 1996; Nutman et al., 1996). Near the Isua belt, however, it was recognised that some rocks were only weakly tectonised in the Neoarchaean based on the state of strain in dolerite dykes that were correlated with the Mesoarchaean Ameralik dykes at the coast (Bridgwater and McGregor, 1974; Bridgwater et al., 1976; Nutman and Bridgwater, 1986). Ameralik dykes north of the Isua belt are a few centimetres to ∼100 m wide and up to several kilometres long (Fig. 2). Within the northern gneisses, the dominant east–west and north–south swarms have ophitic and sub-ophitic textures and primary igneous mineralogy. The dykes were intruded after movement between the northern gneisses and the Isua belt, and nearly all deformation in the gneisses (Bridgwater and McGregor, 1974). Ameralik dykes in the southern gneisses are much narrower, shorter and

J.L. Crowley / Precambrian Research 126 (2003) 235–257

less abundant. Although not everywhere deformed, they were generally recrystallised into amphibolites. U–Pb dating of baddeleyite and zircon from the older, east–west swarm in the northern gneisses indicates crystallisation at ∼3465 Ma (White et al., 2000). Although it cannot be concluded that all Ameralik dykes near the Isua belt crystallised ∼3465 Ma, it is known that all dykes are Mesoarchaean because they are cut by Neoarchaean pegmatite. All rocks and events that are postdated by the dykes are assumed to be Paleoarchaean because there is no evidence for deformation, metamorphism, or granitoid magmatism having occurred during intervals between emplacement of different generations of Ameralik dykes.

3. U–Pb geochronology 3.1. Methods Ten rock samples were collected for U–Pb dating during mapping around lake 682 (Fig. 2). Samples were chosen on the basis that they have important and clear structural relationships, and likely to have zircons with relatively simple U–Pb systematics. For example, migmatitic heterogeneous gneisses that probably contain grains with igneous and metamorphic components were avoided in preference to homogeneous gneisses. Grains were separated into 6–12 groups according to morphology, colour and size. Representative grains from each group were mounted in epoxy, polished and imaged with a Cameca SX-50 electron microprobe equipped with a backscattered electron (BSE) detector. Internal zoning patterns were revealed by the various degrees of brightness in the BSE images, which correspond to differences in the average atomic number (these differences in zircon are mainly controlled by Th and U contents). BSE imaging allowed for the identification of grains or parts of grains that are likely to have grown during igneous crystallisation. Such grains from tonalitic rocks have pronounced fine oscillatory concentric zoning, low U concentrations and lack overgrowths formed during a younger period of igneous crystallisation. Some of the grains do have thin overgrowths, but these are interpreted as having formed during metamorphism based on high U concentrations, weak to nonexistent zoning, and the U–Pb dates. In the granitic rocks, zircons that are interpreted

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as having grown during igneous crystallisation of the granite have weak fine oscillatory concentric zoning and high U concentrations. Xenocrystic cores in the granitic rocks are evidenced by low U concentrations. Metamorphic zircons were identified in two samples based on zoning and grain morphology. Grains selected for dating were similar to those identified by the BSE imaging as most likely to have grown during igneous crystallisation of the samples or during metamorphism. Six of the nine zircon-bearing samples yielded concordant or nearly concordant dates. Highly discordant dates were produced from U-rich grains and grains with metamorphic rims. Careful grain selection and air abrasion of the grains (Krogh, 1982) was performed to reduce the amount of discordance that was likely due to Pb loss from metamict zones or cracks and the presence of overgrowths. Crystallisation ages and errors for samples with concordant analyses were assigned based on the weighted average of the 207 Pb/206 Pb dates. Crystallisation ages of samples with discordant analyses were interpreted from the upper intercept of collinear analyses. The discordia line, including the age errors on the intercepts, was calculated with the use of a modified York (1969) regression (Parrish et al., 1987). Age errors are given at the 2σ level. The clearest titanite grains were abraded, and then analysed in fractions composed of 1–7 grains. Routine sample preparation and analytical procedures were similar to those described by Dubé et al. (1996). The only difference was that most zircon grains were dissolved using microcapsules described by Parrish (1987). Pb blanks were 3–20 pg for zircon analyses and 10–20 pg for titanite. U blanks were <1 pg. U–Pb data are given in Table 1 and plotted on concordia diagrams in Fig. 3. The locations of the samples are shown on maps in Figs. 2 and 4, and field relationships are shown in Fig. 5. 3.2. Results 3.2.1. Sample 1—tonalitic gneiss layer within Isua belt Layers of tonalitic rocks are abundant in the southwestern part of the Isua belt (Fig. 5a). A layer that lies near the contact with the southern gneisses was dated at 3791 ± 4 Ma (Nutman et al., 1996). Tonalitic rocks within the belt are useful because their age and structural setting can place age constraints on the generally

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Table 1 U–Pb analytical data Concentrationsb

Analysisa Wt. (␮g)e

U (ppm)

Pb∗ (ppm)

Atomic ratiosc Pbc (pg)

206 Pb/204 Pb

Age (Ma)d 208 Pb/206 Pb

206 Pb/238 U

207 Pb/235 U

Foliated meta-tonalite—sample 2 (JC-331-00; Z1 7 137 134.9 Z2 5 136 134.8 Z3 5 128 126.0 Z4 4 146 145.2 Z5 5 160 157.0 Z6 4 266 242.7 Z7 3 149 146.6

Fig. 3b) at lat. 65◦ 00 38.2 N, long. 50◦ 14 53.2 W 18 2858 0.070 0.7938 ± 0.10 5 6883 0.078 0.7968 ± 0.15 3 10314 0.086 0.7882 ± 0.12 11 2458 0.077 0.7991 ± 0.11 11 3934 0.092 0.7813 ± 0.10 10 4768 0.045 0.7639 ± 0.10 5 4348 0.084 0.7883 ± 0.13

207 Pb/206 Pb

Disc. (%)f

0.24 0.11 0.16 0.16 0.18 0.26

0.37216 0.37340 0.36831 0.37328 0.37223 0.37309

± ± ± ± ± ±

0.03 0.03 0.06 0.03 0.06 0.05

3800.2 3805.3 3784.5 3804.8 3800.5 3804.0

± ± ± ± ± ±

1.1 0.8 1.8 1.1 1.7 1.4

−0.1 0.0 0.9 0.8 0.0 −0.1

40.143 40.801 39.711 40.962 39.176 36.421 39.907

± ± ± ± ± ± ±

0.11 0.15 0.13 0.11 0.11 0.11 0.14

0.36676 0.37140 0.36541 0.37179 0.36365 0.34579 0.36715

± ± ± ± ± ± ±

0.03 0.05 0.03 0.04 0.03 0.03 0.03

3778.1 3797.1 3772.5 3798.7 3765.2 3688.5 3779.7

± ± ± ± ± ± ±

0.8 1.4 0.9 1.3 0.8 0.9 1.0

0.3 0.5 0.7 0.3 1.2 0.8 0.9

Foliated meta-tonalite—sample 3 (JC-330-00; Fig. 3c and d) at lat. 65◦ 00 41.5 N, long. 50◦ 14 49.3 W Z1 7 112 108.7 28 1341 0.064 0.7874 ± 0.10 Z2 8 163 155.2 3 18817 0.051 0.7813 ± 0.12 Z3 4 144 138.5 10 3269 0.079 0.7753 ± 0.15 Z4 4 110 102.5 6 3517 0.059 0.7669 ± 0.13 Z5 4 153 149.6 10 2971 0.086 0.7839 ± 0.10 Z6 3 291 277.8 8 5394 0.047 0.7866 ± 0.11 Z7 3 115 111.6 14 1193 0.080 0.7798 ± 0.18 T1 42 23 12.0 59 520 0.049 0.4976 ± 0.19 T2 (3) 78 23 11.8 61 935 0.025 0.4982 ± 0.12 T3 (5) 97 33 16.9 124 825 0.015 0.5001 ± 0.12

39.534 38.979 38.297 37.198 39.483 39.479 38.827 11.982 12.024 12.126

± ± ± ± ± ± ± ± ± ±

0.10 0.13 0.15 0.14 0.11 0.12 0.18 0.19 0.13 0.13

0.36417 0.36186 0.35827 0.35176 0.36529 0.36399 0.36113 0.17463 0.17504 0.17585

± ± ± ± ± ± ± ± ± ±

0.03 0.03 0.04 0.05 0.03 0.03 0.06 0.07 0.04 0.04

3767.3 3757.7 3742.5 3714.7 3772.0 3766.6 3754.6 2602.6 2606.4 2614.1

± ± ± ± ± ± ± ± ± ±

1.0 0.9 1.1 1.4 1.0 0.9 1.8 2.3 1.4 1.3

0.6 1.0 1.1 1.2 1.1 0.7 1.0 0.0 0.0 0.0

Foliated meta-tonalite—sample 4 (JC-399-00; Fig. 3d and e) at lat. 65◦ 0 49.6 N, long. 50◦ 12 43.0 W Z1 8 124 129.1 11 4407 0.172 0.7851 ± 0.09 Z2 18 93 96.0 5 18069 0.160 0.7846 ± 0.11 Z3 4 118 122.1 14 1494 0.170 0.7822 ± 0.13 Z4 3 104 103.9 7 1984 0.173 0.7631 ± 0.11 T1 22 65 36.0 64 732 0.079 0.5106 ± 0.13 T2 (2) 22 79 43.9 78 752 0.072 0.5140 ± 0.11 T3 (3) 23 118 72.7 124 769 0.094 0.5499 ± 0.16

39.326 39.606 39.148 37.107 12.755 12.943 16.075

± ± ± ± ± ± ±

0.10 0.11 0.13 0.12 0.15 0.13 0.17

0.36328 0.36609 0.36299 0.35266 0.18116 0.18262 0.21201

± ± ± ± ± ± ±

0.03 0.03 0.05 0.04 0.07 0.04 0.04

3763.6 3775.3 3762.4 3718.5 2663.5 2676.8 2921.0

± ± ± ± ± ± ±

0.9 0.9 1.4 1.1 2.2 1.3 1.4

0.7 1.1 1.0 1.7 0.2 0.1 3.3

J.L. Crowley / Precambrian Research 126 (2003) 235–257

Tonalitic gneiss layer within Isua belt—sample 1 (JC-262-00; Fig. 3a) at lat. 65◦ 05 21.3 N, long. 50◦ 10 34.5 W Z1 3 164 182.1 50 572 0.231 0.8046 ± 0.23 41.287 ± Z2 4 336 382.2 53 1271 0.267 0.8046 ± 0.10 41.423 ± Z3 4 132 134.8 3 8780 0.138 0.7891 ± 0.16 40.070 ± Z4 3 166 185.2 4 5606 0.256 0.7957 ± 0.15 40.953 ± Z5 3 76 78.3 10 1374 0.127 0.8033 ± 0.18 41.229 ± Z6 1 81 83.4 8 654 0.118 0.8048 ± 0.26 41.401 ±

207 Pb/206 Pb

Diorite—sample 5 (JC-200-00; Fig. 3d and f) at lat. 65◦ 04 06.7 N, long. 50◦ 11 08.2 W Z1 8 66 65.5 11 2505 0.160 0.7633 Z2 6 90 86.6 19 1447 0.131 0.7637 Z3 6 98 97.3 16 1701 0.168 0.7651 T1 (3) 89 12 12.3 205 228 0.506 0.6792 T2 (5) 129 11 12.3 302 208 0.633 0.6990

35.689 35.682 35.770 27.815 29.524

± ± ± ± ±

0.12 0.15 0.14 0.17 0.14

0.33913 0.33886 0.33906 0.29703 0.30633

± ± ± ± ±

0.03 0.03 0.03 0.05 0.04

3658.8 3657.7 3658.5 3454.8 3502.6

± ± ± ± ±

0.9 0.9 1.1 1.6 1.3

0.1 0.0 −0.1 3.3 2.4

Foliated pegmatite—sample 6 (JC-274-00; Fig. 3g) at lat. 65◦ 03 42.4 N, long. 50◦ 13 44.9 W Z1 117 770 689.4 381 11155 0.047 0.7530 ± 0.09 Z2 32 1059 930.0 231 6718 0.048 0.7392 ± 0.09 Z3 61 718 638.4 93 22318 0.042 0.7503 ± 0.11 Z4 14 784 686.8 46 10636 0.045 0.7397 ± 0.08

34.818 33.960 34.647 33.741

± ± ± ±

0.10 0.10 0.11 0.09

0.33536 0.33320 0.33490 0.33084

± ± ± ±

0.03 0.03 0.03 0.03

3641.8 3631.9 3639.7 3621.0

± ± ± ±

0.8 0.8 0.8 0.7

0.6 1.8 0.9 1.4

Foliated leucogranite—sample 7 (JC-351-00; Fig. 3h) at lat. 65◦ 00 38.3 N, long. 50◦ 15 02.5 W Z1 4 3655 2969.0 220 2623 0.040 0.6961 ± 0.09 Z2 14 1569 1307.4 21 44272 0.035 0.7140 ± 0.10 Z3 5 1581 1324.7 35 10165 0.047 0.7119 ± 0.09 Z4 11 1358 1124.7 12 54850 0.037 0.7092 ± 0.09 Z5 4 1803 1491.1 15 23523 0.039 0.7075 ± 0.10 Z6 4 1374 1082.4 12 19942 0.038 0.6792 ± 0.10

30.340 31.634 31.490 31.348 31.209 29.097

± ± ± ± ± ±

0.10 0.11 0.10 0.10 0.11 0.11

0.31611 0.32135 0.32083 0.32059 0.31995 0.31073

± ± ± ± ± ±

0.03 0.03 0.03 0.03 0.03 0.03

3551.0 3576.3 3573.8 3572.7 3569.6 3524.5

± ± ± ± ± ±

0.9 0.8 0.8 0.8 0.9 0.9

4.1 2.9 3.0 3.3 3.4 5.2

Amphibolitic gneiss—sample 8 (JC-392-00; Fig. 3i) at lat. 65◦ 01 32.2 N, long. 50◦ 13 20.5 W Z1 23 249 226.8 51 5180 0.069 0.7564 ± 0.09 Z2 6 423 382.8 4 31685 0.064 0.7531 ± 0.11 Z3 4 265 241.0 6 8757 0.068 0.7541 ± 0.10

34.736 ± 0.10 34.575 ± 0.12 34.587 ± 0.10

3631.3 ± 0.8 3630.9 ± 0.9 3629.4 ± 0.9

0.0 0.3 0.2

Metasedimentary gneiss—sample 9 (JC-385-00; Fig. 3j) at lat. 65◦ 00 49.6 N, long. 50◦ 12 43.1 W Z1 11 757 640.5 551 724 0.028 0.7324 ± 0.10 Z2 8 1374 693.6 48 7420 0.014 0.4914 ± 0.09 Z3 4 844 689.6 4 38393 0.018 0.7121 ± 0.10 Z4 3 1693 865.1 13 13363 0.015 0.4973 ± 0.10

31.866 12.240 31.006 12.360

± ± ± ±

0.11 0.10 0.11 0.11

long. 50◦ 13 44.6 W 0.5128 ± 0.11 12.949 ± 0.12 0.5048 ± 0.14 12.475 ± 0.14 0.5104 ± 0.10 12.777 ± 0.11

0.33308 ± 0.03 0.33299 ± 0.03 0.33266 ± 0.03 0.31555 0.18064 0.31582 0.18026

± ± ± ±

0.03 0.03 0.03 0.03

0.18314 ± 0.5 0.17925 ± 0.6 0.18156 ± 0.4

3548.3 2658.7 3549.6 2655.3

± ± ± ±

1.0 0.9 0.8 0.9

0.2 3.1 2.3 2.0

2681.5 ± 1.6 2646.0 ± 1.8 2667.1 ± 1.1

0.5 0.4 0.3

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0.11 0.14 0.13 0.16 0.12

Foliated meta-diorite layer—sample 10 (JC-356-00; Fig. 3d) at lat. 65◦ 00 50.1 N, T1 (2) 17 107 62.4 53 1113 0.139 T2 (3) 11 178 99.3 42 1486 0.102 T3 (7) 18 193 107.8 73 1530 0.090

± ± ± ± ±

Map datum for sample locations is WGS 84. a Mineral analysed (T, titanite; Z, zircon), followed by number of grains in each analysis in parentheses (where >1). b Concentration uncertainty varies with sample weight: estimated at >30% for sample weights <5 ␮g, >10% for sample weights <10 ␮g, <10% for sample weights >10 ␮g. Pb∗ , radiogenic Pb; Pbc , total common Pb in analysis corrected for spike and fractionation. c Ratios corrected for spike, fractionation, blank and initial common Pb, except 206 Pb/204 Pb ratio corrected for spike and fractionation only. Errors are 1σ in percent. d Errors are 2σ in Ma. e Weight of zircon grains estimated visually using a microscope. f Disc. = discordance = 100 − (100 × (206 Pb/238 U age)/(207 Pb/206 Pb age)). 241

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Fig. 3. U–Pb concordia plots. Interpreted ages are weighted averages of 207 Pb/206 Pb dates or upper intercepts from modified York (1969) regressions (see text for details). Age errors and ellipses for the analyses are at the 2σ level. Sample numbers are in parentheses. Analytical data are in Table 1.

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Fig. 3. (Continued ).

mafic lithologies that are difficult to date directly. Sample 1 is from a layer of fine-grained, strongly foliated biotite tonalitic gneiss in the southwestern part of the belt (Figs. 2 and 4). It was collected east of a low hill composed of layers of tonalitic gneiss and amphibolite. One of the amphibolite layers contains oblate structures that are interpreted as flattened pillow lava structures (Fig. 4 in Myers, 2001 and locality 6 in Nutman et al., 2002). Contacts between the sampled layer and adjacent amphibolites are covered, but they are considered to be similar to nearby exposed contacts that are concordant with the dominant foliation. Most tonalitic layers range from a few centimetres to a few metres wide, with the largest being ∼100 m. The small grain size and strong foliation suggest that the protolith of sample 1 was strongly deformed.

Tan to colourless zircon grains are 200–350 ␮m long and elongate (aspect ratios of 4:1 to 6:1). BSE imaging shows fine oscillatory concentric zoning (Fig. 6a) throughout all the grains. These grains are interpreted as having grown during igneous crystallisation of the sample. Four grains with low to moderate U concentrations (76–336 ppm) are concordant with 207 Pb/206 Pb dates of 3805–3800 Ma (Table 1; Fig. 3a), and another grain is 0.8% discordant with a 207 Pb/206 Pb date of 3805 Ma. The weighted average of 3803.3 ± 3.1 Ma (mean square of the weighted deviates (MSWD) = 3.5) is the interpreted igneous crystallisation age. The spread in dates is probably due to small amounts of Pb loss that occurred during metamorphic events at ∼3.6 and (or) 2.7 Ga.

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Fig. 4. Map of the northwestern part of the southern gneisses (mostly after mapping during this study and partly after Nutman, 1986), showing the location of the U–Pb samples (sample numbers in parentheses). Mineral abbreviations after Kretz (1983).

3.2.2. Sample 2—foliated meta-tonalite Sample 2 is from weakly foliated biotite meta-tonalite at the base of a ∼2 m high outcrop south of the Isua belt, at locality A of Nutman et al. (1999; Fig. 2). It was collected ∼1 m east of a thin, subvertical, north-striking, meta-dolerite dyke that is thought to be part of the Mesoarchaean Ameralik swarm (Fig. 5b). Nutman et al. (1999) dated meta-tonalite collected a few metres from sample 2 at 3808 ± 4 Ma, and two meta-tonalite samples from frost-heaved boulders that lie ∼20 m to the north were dated at 3818 ± 8 Ma and 3811 ± 6 Ma. Meta-tonalite in the southern gneisses was described by Nutman et al. (1999) as having a relict igneous texture and a weak biotite foliation (Fig. 5c) or, more rarely, no discernible fabric. Primary igneous contacts exist at locality A between weakly foliated and unfoliated leucocratic phases, in-

cluding a ∼5-cm wide trondhjemitic pegmatite vein dated at 3661 ± 8 Ma (Nutman et al., 1999). There is a dramatic increase in strain away from the sample locality; meta-tonalite located ∼5 m to the west is well foliated and contains thin concordant leucocratic layers, and tonalite ∼20 m away has a tightly folded gneissosity. Leucogranite layers located 120 m to the west (sample 7) are strongly foliated and concordant with tectonic layering in the host amphibolitic gneiss, and pegmatite-banded tonalitic gneiss was deformed by west-verging Neoarchaean folds (Fig. 5d). These relationships indicate that the volume of weakly deformed rocks at locality A is small, perhaps a few hundred cubic metres. Colourless zircon grains are 200–350 ␮m long and elongate (3:1 to 4:1). Grains with fine oscillatory concentric zoning are partly surrounded by narrow (up

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Fig. 5. Field relationships. (a) Contact between amphibolite of southwestern part of the Isua belt and homogeneous tonalitic gneiss of the southern gneisses. Scale is an 18-cm long notebook in centre. View to southeast at lat. 65◦ 04 59 N, long. 50◦ 10 31 W (map datum is WGS 84). (b) Sample 2 meta-tonalite was collected at locality A of Nutman et al. (1999). Person for scale on the right. View to south. (c) Weakly foliated meta-tonalite in a low strain zone, near locality B of Nutman et al. (1999). Scale card is 8 cm long. View to north at lat. 65◦ 00 50 N, long. 50◦ 13 46 W. (d) Pegmatite-banded gneissosity in gneiss cut by Mesoarchaean Ameralik dyke that underwent Neoarchaean folding. Contact is marked by dashed line. This strong deformation contrasts with the weak strain at locality A (120 m to the east). Scale is an 18-cm long notebook in centre. View to south at lat. 65◦ 00 37 N, long. 50◦ 15 04 W. (e) Layers of meta-tonalite (sample 4), metasedimentary gneiss (sample 9), pegmatite and amphibolite in an enclave of supracrustal rocks south of the Isua belt. View to west. (f) Sketch of (e). (g) Diorite a few tens of metres north of sample 5 collection site cuts gneissosity in the host tonalitic gneiss. Scale is 5-cm wide lens cap. View to north. (h) The southern end of the diorite layer (sample 5) is a stockwork of dykes that join at high angles. Scale is 70-cm long backpack in centre. View to west at lat. 65◦ 03 54 N, long. 50◦ 10 54 W. (i) Foliated pegmatite a few metres west of sample 6 collection site. Dark layer to left is an Ameralik dyke. View to north. (j) Foliated, concordant leucogranite layers (sample 7) within amphibolitic gneiss. Scale is 40-cm long hammer. View to south.

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Fig. 5. (Continued ).

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Fig. 6. Backscattered electron (BSE) images of representative polished zircon grains show domains of various brightness that correspond to subtle differences in the average atomic number (mainly controlled by U contents). Sample numbers are in parentheses. (a) Sample 1, concentric oscillatory zoning. (b and c) Sample 2, concentric oscillatory zoning with thin, bright (U-rich) rims. (d and e) Sample 3, zircons are similar to those from sample 2. (f and g) Sample 4, zircons are similar to those from sample 2. (h and i) Sample 5, fine to broad zoning parallel to the length of grains. (j) Sample 6, fine oscillatory concentric zoning. (k) Sample 7, weak fine concentric zoning. (l and m) Sample 8, weak patchy concentric zoning. (n) Sample 9, weak fine concentric zoning.

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to 25 ␮m) U-rich rims that were presumably removed during abrasion (Fig. 6b and c). Four grains with moderate U concentrations (136–160 ppm) define a discordia with upper and lower intercepts of 3814.5 + 10.4/−7.5 and 2774 ± 173 Ma, respectively (MSWD = 0.5; Table 1; Fig. 3b). The two least discordant analyses along the discordia have 207 Pb/206 Pb dates of ∼3798 Ma. The upper intercept is interpreted as the igneous crystallisation age, which agrees with SHRIMP ages obtained from three nearby rocks by Nutman et al. (1999). The large uncertainty on the upper intercept age is due to the low angle of intersection between concordia and the discordia line. The lower intercept age could have been the time of metamorphic growth of the U-rich rims. However, because these thin rims were presumably removed during grain abrasion and SHRIMP dating of U-rich rims in nearby tonalitic rocks (Friend et al., 2002) yielded ages of 3640–3620 Ma, it is more likely that the U-rich rims in sample 2 grew in the Paleoarchaean. In this interpretation, the Neoarchaean lower intercept age is the time of Pb loss from metamict zones and cracks. 3.2.3. Sample 3—foliated meta-tonalite Sample 3 is from foliated hornblende-biotite meta-tonalite located ∼120 m north-northeast of sample 2 (Fig. 2). It differs from sample 2 by containing centimetres-wide leucocratic layers and a moderate foliation, yet the zircon grains are similar (Fig. 6d and e). Five grains with moderate U concentrations (112–291 ppm) define a discordia with upper and lower intercepts of 3795.0 + 20.3/−11.2 and 2772 ± 283 Ma, respectively (MSWD = 0.5; Table 1; Fig. 3c). The two least discordant analyses along the discordia have 207 Pb/206 Pb dates of 3767 Ma. The upper intercept is interpreted as the igneous crystallisation age, in agreement with results from sample 2 and the previously dated nearby meta-tonalites. Three titanite fractions of 1–5 tan grains have moderate U concentrations (23–33 ppm) and concordant dates of 2614–2602 Ma (Table 1; Fig. 3d). 3.2.4. Sample 4—meta-tonalite layer within quartz-rich gneiss Sample 4 is from a 5–15-cm wide layer of foliated biotite meta-tonalite that lies between thin (10–20 cm) layers of clinopyroxene-amphibole-quartz gneiss (sample 9; Fig. 2). The quartz-rich gneiss is part of an

∼20 m × 20 m enclave that also includes metre-thick layers of amphibolite and leucocratic amphibolite and thin layers of amphibolite, rusty quartz-rich gneiss, meta-tonalite and pegmatite (Fig. 5e and f). The quartz-rich gneiss is interpreted as being derived from a chemical sediment for reasons given with description of sample 9. Zircon grains are similar to those from samples 1 and 2 (Fig. 6f and g), except there are more cracks and inclusions, and U-rich rims are thicker (up to 50 ␮m). Three grains with moderate U concentrations (104–124 ppm) define a discordia with upper and lower intercepts of 3800.9 + 27.3/−16.7 and 2820 ± 240 Ma, respectively (MSWD = 2.9; Table 1; Fig. 3e). The two least discordant analyses along the discordia have 207 Pb/206 Pb dates of 3764–3763 Ma. The upper intercept is interpreted as the igneous crystallisation age, in agreement with samples 1–3. Two titanite fractions of one and two tan grains have high U concentrations (65 and 79 ppm) and slightly discordant dates with 207 Pb/206 Pb dates of 2677 and 2664 Ma (Table 1; Fig. 3d). Another fraction (T3) composed of three grains is more discordant and has an older 207 Pb/206 Pb date. The significance of this older fraction is given with discussion of titanite from sample 5. 3.2.5. Sample 5—diorite Sample 5 is from a large (25–80 m× ∼ 650 m) layer of hornblende-biotite diorite (Figs. 2 and 4) that is concordant with the regional gneissosity and nearby enclaves of mafic and ultramafic rocks. Despite being broadly concordant, the margins of the layer and narrow dykes that emanate from it cut across the gneissosity in the host tonalitic gneiss at a high angle (Fig. 5g). Diorite dykes are younger than some pegmatite layers in the host gneiss, yet older than others (Fig. 5g and h). The layer terminates southward in a stockwork of dykes that join at high angles (Fig. 5h). This medium-grained rock contains widely scattered feldspar megacrysts, most being 1–2 cm in size (Fig. 5g) with some up to 30 cm. Typically, the megacrysts are equant and the diorite lacks a foliation. However, there is evidence that the diorite was locally deformed, including flattened megacrysts that lie along a moderately well developed foliation defined by aligned matrix plagioclase and hornblende. The foliation is strongest in diorite that lies adjacent

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to sheared Ameralik dykes that are concordant with the foliation, recrystallised into foliated amphibolites, and have pinch and swell structures. Together these observations require that considerable deformation, metamorphism and pegmatite intrusion in the host gneiss predated diorite intrusion, but a minor amount also occurred afterward, probably after emplacement of the Ameralik dykes; there is no evidence for tectonism in the interval between intrusion of the diorite and Ameralik dykes. Colourless zircon grains are 250–400 ␮m long and elongate (4:1 to 6:1). Broad zoning parallels the length of the grains (Fig. 6h and i). Three grains with low U concentrations (66–98 ppm) are concordant with 207 Pb/206 Pb dates of 3659–3658 Ma (Table 1; Fig. 3f). The weighted average of 3658.3 ± 1.2 Ma (MSWD = 0.3) is interpreted igneous crystallisation age. Two titanite fractions of three and five tan grains have low U concentrations (11 and 12 ppm) and dates that are slightly discordant with 207 Pb/206 Pb dates of 3503 and 3455 Ma (Table 1; Fig. 3d). These two fractions and two each from samples 4 and 5 lie in a co-linear array, defining a discordia with upper and lower intercepts of 3623.7 ± 8.5 Ma and 2653.1 ± 5.5 Ma, respectively (MSWD = 0.6). The fact that these define a discordia is interpreted as indicating that all titanite from sample 5 first cooled below the Pb closure temperature or grew at the upper intercept age, and then underwent partial Pb loss and (or) new titanite growth at the lower intercept age, when most grains from samples 4 and 5 underwent complete Pb loss and (or) new titanite growth. Because analyses from sample 5 plot close to the upper intercept, the ∼2653 Ma thermal overprint in its vicinity is thought to have had minor effects. At least one of the three grains in fraction T3 from sample 4 was also not completely overprinted at ∼2653 Ma. 3.2.6. Sample 6—foliated pegmatite Sample 6 is from a layer of strongly foliated pegmatite that lies in a 1 km × 0.1 km zone that is dominated by vertical, metre-scale layers of foliated pegmatite within tonalitic gneiss (Figs. 2 and 4). Structural relationships in this pegmatite-rich zone are used to interpret the timing and intensity of Paleo- and Neoarchaean deformation. The presence of a weak foliation in the eastern margin of the pegmatite zone compared to the strong gneissosity in the host tonalitic gneiss indicates that substantial deformation predated

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pegmatite intrusion. The minor deformation that affected the eastern margin is known to have occurred in the Paleoarchaean because it was postdated by an undeformed pegmatite dyke that was cut by an Ameralik dyke. Pegmatite in the middle of the zone underwent stronger deformation that resulted in a vertical foliation defined by flattened quartz and feldspar and aligned biotite (Fig. 5i). Thin (0.1–0.5 m) pegmatite layers typically lie along the axial surfaces of upright, shallowly north plunging folds. The gneissosity in the host gneiss that is folded certainly formed before pegmatite intrusion, but it is unclear whether the upright folding occurred during pegmatite intrusion, shortly after pegmatite intrusion in the Paleoarchaean, or much later in the Neoarchaean. Evidence for some of the folding being Neoarchaean comes from Ameralik dykes that lie along the axial surfaces of the upright folds and contain an axial planar foliation, and the presence of Ameralik dykes that are locally folded. In summary, sample 6 was collected from pegmatite that crystallised after considerable deformation and metamorphism affected the host gneisses, yet, it was affected by upright folding that increases in intensity westward across the pegmatite zone, at least some of which is Neoarchaean. Reddish-brown zircon grains are 300–650 ␮m long and elongate (3:1) to stubby (2:1) to subequant. Weak fine oscillatory concentric zoning exists throughout the grains (Fig. 6j) and xenocrystic cores were not seen. Three grains with high U concentrations (718–1059 ppm) define a discordia with upper and lower intercepts of 3647.0 ± 2.4 Ma and 1383 ± 170 Ma, respectively (MSWD = 0.1; Table 1; Fig. 3g). The least discordant analysis forming the discordia has a 207 Pb/206 Pb date of 3642 Ma, which is a firm minimum age of igneous crystallisation. The upper intercept is the interpreted igneous crystallisation, but it cannot be ruled out that crystallisation occurred a few million years before that age until additional grains are analysed. Another analysis that lies above the discordia has a different U–Pb history, possibly having lost Pb in the Neoarchaean. 3.2.7. Sample 7—foliated leucogranite Sample 7 is from a strongly foliated, 15–25-cm wide garnet leucogranite layer that lies within amphibolitic gneiss (Fig. 2). It was collected 120 m west of the weakly foliated meta-tonalites at locality A of

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Nutman et al. (1999; see description of sample 2). The strong foliation is defined by highly flattened quartz and feldspar. This layer is parallel to other thin leucogranite layers in the outcrop, all of which are concordant with the gneissosity in the host amphibolitic gneiss (Fig. 5j). The concordant and foliated nature of the leucogranite layers requires that they were intensely deformed, and a cross-cutting nearby Ameralik dyke indicates that this deformation is Paleoarchaean. These relationships require that a strong strain gradient exists in the 120 m interval between this outcrop and the weakly foliated meta-tonalite at locality A. The youngest structures that affected the layer are open to tight, west-verging folds that plunge shallowly to moderately to the south. The involvement of Ameralik dykes in this folding (Fig. 5d) indicates a Neoarchaean age. Reddish-brown zircon grains are 300–500 ␮m long and elongate (5:1) to stubby (2.5:1). Weak fine concentric zoning exists throughout the grains (Fig. 6k) and xenocrystic cores were not seen. Six grains with high U concentrations (1358–3655 ppm) define a discordia with upper and lower intercepts of 3635.6 + 7.7/−6.9 and 2043 ± 78 Ma, respectively (MSWD = 2.0; Table 1; Fig. 3h). The least discordant analysis along the discordia has a 207 Pb/206 Pb date of 3574 Ma. The upper intercept is the interpreted igneous crystallisation age. 3.2.8. Sample 8—amphibolitic gneiss Sample 8 is from garnet-bearing amphibolitic gneiss on a peninsula in lake 682 composed largely of mafic and ultramafic rocks (Fig. 2). Pink zircon grains are 80–300 mm wide, subequant, anhedral with smooth indentations, and have weak patchy concentric zoning (Fig. 6l and m). These characteristics suggest that the grains grew during metamorphism. Three grains with moderate U concentrations (249–423 ppm) have dates that are concordant to 0.3% discordant with 207 Pb/206 Pb dates of 3631–3629 Ma (Table 1; Fig. 3i). The weighted average of 3630.5 ± 1.3 Ma (MSWD = 0.7) is the interpreted metamorphic growth age. 3.2.9. Sample 9—metasedimentary gneiss Sample 9 is from a 10–20-cm wide layer of clinopyroxene-amphibole-quartz gneiss in an enclave within tonalitic gneiss (Fig. 2). It is concordant with other layers in the ∼20 m × 20 m enclave (Fig. 5e

and f), including metre-thick layers of amphibolite and leucocratic amphibolite and thin layers of amphibolite, a rusty quartz-rich gneiss, pegmatite and meta-tonalite dated at 3801 + 27/−17 Ma (sample 4). Sample 9 is composed of 0.2–2-cm wide alternating bands of quartz (∼50% of the rock), amphibole and amphibole + clinopyroxene. There are trace amounts of magnetite. Although there are no preserved primary structures indicative of the protolith, the quartz-rich gneiss is interpreted as having a sedimentary origin, probably a chemical sediment such as chert, because it is mineralogically distinct from the metaplutonic and metavolcanic rocks in the region and it is associated with amphibolites that may have a supracrustal origin. Based on these characteristics, however, it cannot be ruled out that the gneiss was derived from a quartz vein or silicified and metasomatised amphibolite or ultramafic rock. In fact, a vein does emanate from the a gneiss layer into the adjacent leucocratic amphibolite (Fig. 5e and f). The vein is 2 cm wide, up to 2 m long, and is similar to the gneiss. This relationship perhaps indicates that the gneiss was derived from a quartz vein that infiltrated into tonalite and amphibolite. However, this possibility can be ruled out because the thin vein cuts strongly across the foliation in the amphibolite, and thus formed late in the history of these rocks, in contrast to the gneiss from which it emanates that was clearly affected by all of the deformation. The vein is thus interpreted as having formed along a crack that developed during the waning stages of metamorphism, during conditions of high fluid pressure in which silica mobility was increased. Brown zircon grains are 150–300 ␮m wide and stubby (2:1) to subequant. Some have smooth indentations and protrusions, and all have weak fine concentric zoning (Fig. 6n). These characteristics suggest that the grains grew during metamorphism. Some grains contain strong oscillatory zoned cores that are probably detrital grains. Two grains with high U concentrations (757 and 844 ppm) are 0.3 and 2.3% discordant with 207 Pb/206 Pb dates of 3548 and 3550 Ma, respectively (Table 1; Fig. 3j). The weighted average is 3548.9 ± 1.5 Ma (MSWD = 0.8). Two other grains with high U concentrations (1374 and 1693 ppm) are 2.0 and 3.1% discordant with 207 Pb/206 Pb dates of 2655 and 2659 Ma, respectively. These analyses indicate that there were at least two

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periods of metamorphic zircon growth in this gneiss, at ∼3550 and 2660 Ma. 3.2.10. Sample 10—foliated diorite Sample 10 is from a 10–15-cm wide layer of foliated biotite-hornblende diorite that was intruded into weakly foliated meta-tonalite near locality B of Nutman et al. (1999; Fig. 2). At locality B, a diorite dyke cuts across the foliation and thin leucocratic banding in the host meta-tonalite that was dated at 3809±8 Ma and 3806±8 Ma. The diorite is important because it was intruded after much of the deformation that affected the host meta-tonalites. Nutman et al. (1999) tentatively interpreted diorite crystallisation at ∼3760 Ma based on two discordant SHRIMP dates. Because it proved impossible to collect a sample of the diorite at locality B, sample 10 was taken from a similar diorite layer in a hill located 160 m to the north-northwest. The layer is foliated and concordant with the foliation in the host meta-tonalite at the collection site, but locally the layer cuts across the foliation in the host meta-tonalite. Zircon was not dated because BSE imaging showed that the igneous grains were largely replaced by U-rich zircon. Instead, three titanite fractions of 2–7 tan grains with high U concentrations (107–193 ppm) were dated. Two analyses lie near the lower intercept of the discordia between 3624 ± 9 Ma and 2653 ± 6 Ma defined by titanite from samples 4 and 5 (Table 1; Fig. 3d). Another analysis that lies slightly below the discordia is 0.4% discordant with a 207 Pb/206 Pb date of 2646 Ma.

4. Geological history The following sequence of events are proposed based on previous work and the field relationships and data presented above. This sequence is summarised on the timeline in Fig. 7. 4.1. Magmatism The debate over the protolith ages of some tonalitic rocks in the Itsaq Gneiss Complex is centred on whether the oldest parts of typically heterogeneous zircons grew in the protolith magma or are xenocrystic (Nutman et al., 1996, 2000; Kamber and Moorbath, 1998; Whitehouse et al., 1999; Kamber et al., 2001).

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This controversy does not involve tonalitic rocks in the southern gneisses because there is little doubt that the ∼3820–3795 Ma zircon ages obtained by Nutman et al. (1996, 1999, 2002) and Friend et al. (2002) from nearly a dozen samples collected around lake 682 indeed represent the timing of crystallisation of the protolith magmas. Zircons in these well preserved rocks are homogeneous, prismatic grains with fine, concentric, oscillatory zoning that is typical of zircons grown in tonalite magmas. The uniformity of the ages and zoning patterns within and between samples suggests that xenocrysts, either as cores or whole grains, do not exist. Thin U-rich rims are found on many grains, some of which have 3640–3620 Ma ages (Friend et al., 2002) that overlap with the timing of regional metamorphism. Dating of tonalitic rocks by the IDTIMS method during this study (samples 1–4) yielded ∼3810– 3795 Ma ages, in agreement with the SHRIMP ages. Younger tonalitic rocks (∼3700–3690 Ma) are present in the southern gneisses, where they are thought to be volumetrically minor because they are only known to exist in two localities (Nutman et al., 1993, 2002). Diorite is sparse in the southern gneisses, the only known thick layer of which was dated at 3658.3 ± 1.2 Ma (sample 5). Ubiquitous pegmatite and leucogranite dykes and layers mostly crystallised at 3650–3630 Ma (samples 6 and 7; Friend et al., 2002), and one younger pegmatite yielded an age of 3607 ± 5 Ma (Nutman et al., 2002). Older pegmatite has not been dated, but it is known to exist where cut by 3658 Ma diorite. The quiescent Mesoarchaean began with intrusion of the Ameralik dolerite dykes. Ameralik dykes were only cut by Neoarchaean pegmatite dykes and Proterozoic dolerite dykes. 4.2. Metamorphism The southern gneisses were metamorphosed several times, in events that are thought to have occurred at amphibolite facies conditions based on the lack of granulite facies assemblages and the occurrence of hornblende in the mafic enclaves (resulting from Paleoarchaean heating) and in recrystallised Ameralik dykes (resulting from Neoarchaean heating). Evidence for an early high-grade metamorphism comes from the presence of a pegmatite-banded gneissosity that was intruded by 3658 Ma diorite (sample 5; Fig. 5g).

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Fig. 7. Timeline of geological events in the southern gneisses, northern gneisses and Isua belt. Ages are U–Pb SHRIMP and IDTIMS dates, except for the four labelled Pb–Pb dates. Note that the Mesoarchaean era is mostly omitted, and the protolith ages of supracrustal units in the Isua belt are not given.

Minerals that grew during the >3658 Ma event have not been dated. High-grade metamorphism also likely accompanied deformation that developed a foliation in 3650–3630 Ma granitic rocks. Timing of this event is probably recorded by the metamorphic growth of (i) zircon at 3650–3620 Ma meta-tonalites (Friend et al., 2002), (ii) zircon in amphibolitic gneiss at 3630.5 ± 1.3 Ma (sample 8) and (iii) titanite (or complete Pb loss in titanite) at 3623.7 ± 8.5 Ma in diorite (sample 5). Evidence for another period of Paleoarchaean heating is metamorphic zircon growth

in a metasedimentary gneiss at 3548.9 ± 1.5 Ma (sample 9). Neoarchaean metamorphism is recorded by titanite growth and (or) Pb loss at 2653.1 ± 5.5 Ma and zircon growth in a metasedimentary gneiss at ∼2660 Ma (sample 9). Metamorphic zircons from an altered peridotite were dated at ∼2650 Ma (Friend et al., 2002) and zircons from a schist located ∼5 km south of lake 682 grew at 2685 ± 5 Ma (Nutman et al., 2002). The position of titanite analyses along the 3624–2653 Ma discordia (Fig. 3d) is related to

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the amount of ∼2653 Ma titanite growth and (or) Pb loss, which is presumably related to the intensity of the thermal event. The ∼3624 Ma titanite from diorite (sample 5) was only moderately disturbed by Neoarchaean metamorphism, whereas titanite from samples located farther from the Isua belt grew and (or) suffered complete Pb loss at ∼2653 Ma. This southward increase in Neoarchaean heating is also reflected by an increase in the degree of recrystallisation of Ameralik dykes; dykes near the Isua belt locally contain sub-ophitic primary textures, but southward they are hornblende-plagioclase ± chlorite ± epidote amphibolites with strong linear and planar deformation fabrics. Other titanites from the samples that define the 3624–2653 Ma discordia yielded dates that suggest some grains grew and (or) lost Pb slightly before or after 2653 Ma. In addition, younger titanite from a meta-tonalite yielded ∼2600 Ma dates (sample 3). 4.3. Deformation Ductile strain was heterogeneous across the southern gneisses. Parts of the ∼3800 Ma tonalite and its enclaves were sheared and converted into gneiss relatively early and then deformed a few more times, whereas small, scattered lenticular zones of meta-tonalite escaped essentially all deformation. These low-strain zones commonly lie only a few tens of metres from strongly deformed gneisses. Dated samples with clear structural relationships are used to place age constraints on the deformation, and the Ameralik dykes are used to distinguish Paleoarchaean strain from Neoarchaean strain (based on the assumption that all Ameralik dykes are Mesoarchaean for reasons given above). Early deformation is indicated by the intrusion of by ∼3800 Ma tonalite and ∼3645 Ma pegmatite across tectonic contacts between ultramafic and mafic rocks (Friend et al., 2002). Tonalite between lake 682 and the Isua belt was converted into pegmatite-banded gneiss before intrusion of a 3658 Ma diorite (sample 5; Fig. 5g), with the exception of weakly foliated tonalite lying within a few hundred metres of the belt that were never strongly deformed. In a few localities along the northeastern shore of lake 682, the >3658 Ma gneissosity lies in the axial surface of steeply plunging folds that affected a pegmatite-banded fabric. This relationship suggests that the >3658 Ma gneissosity re-

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sulted from transposition of an even older fabric, and thus, there were multiple early deformation events. Deformation after 3658 Ma in the vicinity of the diorite is thought to have been mild because it is generally weakly foliated (Fig. 5g), plagioclase megacrysts are equant, and dykes join at high angles (Fig. 5h). It is likely that any deformation occurred after emplacement of the Ameralik dykes. Other evidence for early deformation is found where the margin of a large 3647.0 ± 2.4 Ma pegmatite body (sample 6) is weakly foliated relative to it host gneiss. The intensity of the foliation in the pegmatite increases westward (Fig. 5i), as does the number of upright, shallowly north plunging folds. These upright folds do not exist between lake 682 and the Isua belt, but are common to the west and south. The age of this folding is difficult to determine. Some of it may have occurred in the Paleoarchaean, but much is thought to be Neoarchaean because foliated amphibolite layers (derived from Ameralik dykes) lie along the axial surfaces of the folds. Clear evidence for Paleoarchaean deformation to the west of lake 682 is found where a strongly foliated, concordant leucogranite layer dated at 3636 + 8/−7 Ma (sample 7; Fig. 5j) lies near cross-cutting Ameralik dykes that were not affected by this severe deformation (Fig. 5d). Spatial variations in Neoarchaean strain are recorded by the Ameralik dykes. Between lake 682 and the Isua belt, there are only a few folded and sheared dykes. Although many dykes in this region are concordant with the dominant gneissosity (Fig. 4), it is likely that they were not transposed into that orientation because they do not contain deformation fabrics and some dykes are not strained where they cut across the gneissosity at a high angle. Neoarchaean strain increases dramatically to the west and south of lake 682. The line drawn by Nutman (1986) and Nutman et al. (1999) showing the northern limit of ubiquitous Neoarchaean high strain lies at lat. 65◦ N to the south of lake 682 and sweeps northward along its west side. North of this line, the gneissosity formed in the Paleoarchaean and was only locally slightly modified by Neoarchaean strain. To the south in a transition zone a few hundred metres wide, the Paleoarchaean gneissosity was largely transposed into a Neoarchaean fabric. Small zones of weakly deformed meta-tonalite are widely scattered between the western arm of lake 682 and lat. 65◦ N. A low strain zone that is a few hundred cubic metres in size exists at locality A of Nutman

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et al. (1999; Fig. 5b), and a larger zone lies near locality B (Fig. 5c). Meta-tonalite in the middle of the low-strain zones is weakly foliated to unfoliated, and is cut by unstrained Paleoarchaean leucocratic veins and Ameralik dykes. The low-strain zones are surrounded by strongly deformed rocks; a foliation was developed in tonalite within a few metres of the middle of the low-strain lens at locality A, and a strong gneissosity was formed in tonalite a few tens of metres away. Leucogranite layers 120 m away (Fig. 5j) are strongly foliated and concordant with tectonic layering in the host amphibolitic gneiss, and nearby pegmatite-banded tonalitic gneiss was folded by Neoarchaean folds (Fig. 5d). 4.4. Age of supracrustal rocks The protolith age of zircon-poor supracrustal rocks in the Isua belt and enclaves in the southern gneisses is best constrained by carefully examining relationships with adjacent dated tonalitic rocks. Amphibolites with pillow lava structures near the southwestern margin of the Isua belt are intercalated with layers of tonalitic rocks, some of which were dated at 3803.3 ± 3.1 Ma (sample 1) and 3791 ± 4 Ma (Nutman et al., 1996). Nutman et al. (1996, 1997, 1999, 2002) reported that 3800 Ma tonalitic sheets cut across layered amphibolite in the belt where they emanate from homogeneous meta-tonalite that borders the belt to the south. This observation led to the interpretation that the tonalitic rocks were intruded into the amphibolite, requiring that this part of the belt is >3800 Ma. Because no map or photographic evidence of this cross-cutting relationship has been provided, other possible interpretations of intercalated layers of tonalite and amphibolite must be considered, including (i) the volcanic protolith of the amphibolite was extruded upon a tonalite basement, and the contact was subsequently isoclinally folded, and (ii) the units were juxtaposed along faults. The former relationship is considered unlikely because hinges of isoclinal folds of tonalite-amphibolite contacts were not found during this study of a well exposed section at the southwestern margin of the belt (Fig. 5a; locality 5 and Fig. 3a in Nutman et al., 2002). The faulting relationship is also considered unlikely because it requires dozens to perhaps hundreds of faults to have juxtaposed the numerous layers, which range in thickness from a few centimetres to a few

metres. This juxtaposition could not have occurred along brittle faults because such structures do not exist at the margins of the layers, and juxtaposition during ductile shearing is ruled out because many of the tonalitic layers are weakly deformed and pillow lava structures are locally preserved in adjacent amphibolite layers. Therefore, the interpretation of Nutman et al. (1996, 1997, 1999, 2002) that the tonalite of the southern gneisses was intruded as sheets into amphibolite of the Isua belt is considered most likely, requiring that volcanic rocks on the southwestern side of the belt were extruded before 3800 Ma. Quartz-rich gneiss (sample 9) in an enclave within the southern gneisses is interpreted as having a chemical sediment protolith based on mineralogy and association with amphibolites that may have a volcanic origin. A concordant, thin layer of foliated metatonalite that lies within the enclave (Fig. 5e and f) was dated at 3801 + 27/−17 Ma (sample 4). Similar to the example given above, there are multiple possible primary relationships between the meta-tonalite and quartz-rich gneiss given the lack of cross-cutting intrusive relationships. However, an intrusive relationship is considered most likely because the meta-tonalite does not appear to have been strongly sheared (required if this was a faulted contact) and fold hinges are not seen in the outcrop (required if this was a folded basement-cover contact). It is thus likely that at least some of the supracrustal rocks in enclaves within the southern gneisses are >3800 Ma. Nutman et al. (2002) showed that part of the largest horizon of Paleoarchaean supracrustal rocks south of the Isua belt (informally termed the Tussaap belt) is also >3800 Ma. Detrital zircons from a quartzite in the Isua belt are mainly ∼3850 and 3810 Ma, with oldest zircon possibly being ∼3900 Ma (Nutman et al., 1997). These results show that at least one quartzite in the Isua belt is no older than ∼3810 Ma and the source rocks were <3900 Ma.

5. Comparison of geological histories The considerable amount of recently obtained data allows for a detailed comparison of the geological histories of the southern gneisses, northern gneisses, and supracrustal rocks in the Isua belt. This comparison, which is summarised in Fig. 7, is required before

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a tectonic setting for the tectonothermal events can be proposed. Three rock types exist in both the northern and southern gneisses: 3700 Ma tonalite, 3658 Ma diorite and 3655–3630 Ma granite. The 3658 Ma diorite in the northern gneisses (Crowley et al., 2002) comprises the Inaluk dyke swarm (Nutman and Bridgwater, 1986), which differs from 3658 Ma diorite in the southern gneisses by lacking feldspar megacrysts and forming narrow isolated dykes rather than a single thick layer. However, mineralogical and age similarities suggest that diorites in the northern and southern gneisses crystallised from related magmas. The only major granitoid rock that has not been found in both the northern and southern gneisses is 3800 Ma tonalite that is exclusive to the southern gneisses and adjacent margin of the Isua belt. A strong link between the northern and southern gneisses is that they were first tectonised before 3658 Ma and substantially metamorphosed and deformed at 3650–3630 Ma. Minerals that grew during the >3658 Ma event in the gneisses have not been dated, but Pb–Pb dating of garnet in a vein from the southwestern part of the Isua belt yielded an age of 3739±21 Ma, interpreted as the time of garnet growth during a hydrothermal–metasomatic event (Frei and Rosing, 2001). Sm–Nd data from garnet in the southwestern part of the belt (Blichert-Toft and Frei, 2001) and Pb–Pb data from magnetite in the northeast part of the belt (Frei et al., 1999a) indicate similar events at 3714 ± 24 Ma and 3691 ± 22 Ma, respectively. Significant tectonothermal activity occurred throughout the region at 3650–3630 Ma, the interpreted time of the main deformation and metamorphism in the northern gneisses based on structural relationships of granite and pegmatite layers (Crowley et al., 2002). Although some parts of the southern gneisses were not deformed at this time, there was metamorphic zircon and titanite growth (Friend et al., 2002, this study).A narrow (80 m) mylonite zone between the northern gneisses and the Isua belt formed in the Paleoarchaean, probably during 3650–3630 Ma regional deformation (Crowley et al., 2002), coincident with the formation of metre-wide mylonite zones in the northern and southern gneisses and in the Isua belt that juxtaposed supracrustal sequences of different ages (Nutman et al., 2002). Evidence for another thermal event before the quiescent Mesoarchaean is provided by metamorphic zircon growth at 3580–3550 Ma in the southern

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gneisses and Isua belt (sample 9; Nutman et al., 2000, 2002). A large-scale fold that is cored by the southern gneisses, with the Isua belt and Tussaap belt on opposing limbs, is thought to have formed at this time (Nutman et al., 2002). The intensity of Neoarchaean tectonism generally increases southward. Titanite data indicate that rocks in the middle of northern gneisses were little affected by Neoarchaean thermal events, whereas titanite near the Isua belt lost considerable Pb and (or) underwent new growth at 2915 ± 32 Ma (Crowley et al., 2002). This age overlaps with two ages from the western side of the belt, a 2948 ± 8 Ma pegmatite dyke (Hanmer et al., 2002) and a Pb–Pb date of 2840 ± 49 Ma from magnetite that is interpreted as the age of metamorphism (Frei et al., 1999a). Neoarchaean titanite from the southern gneisses (∼2650 and 2600 Ma) is considerably younger than titanite from the northern gneisses and slightly younger than the main Neoarchaean metamorphism along the coast south of Nuuk (∼2720–2700 Ma; Friend et al., 1996; Crowley, 2002), although titanite to the south is similar in age (Crowley, 2002). In the Isua belt, Pb–Pb dating of garnet yielded an age of 2661 ± 19 Ma (Frei et al., 1999b). The intensity of Neoarchaean strain increases southward through the southern gneisses, and the Isua belt appears to have been a high strain zone in the Neoarchaean (Nutman, 1986). The steeply southeast-plunging fold that is cored by banded iron formation and metachert near the ice cap is clearly a Neoarchaean structure, and Nutman (1986) suggested that the curvature of the entire belt resulted from Neoarchaean folding.

6. Tectonic setting Crowley et al. (2002) could not propose a tectonic setting for tectonothermal events that affected the northern gneisses because there are no characteristics diagnostic of a particular setting. However, it was noted that the sequence of overprinting geological events is similar to that in younger orogenic belts, especially the coincidence of peak metamorphism, deformation, and anatectic production of granite. The same conclusion could be made for the southern gneisses. Because it is likely that any large-scale juxtaposition across the Isua belt occurred before intrusion of 3700, 3658 and 3655–3630 Ma rocks that are

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common to the northern and southern gneisses, the tectonic setting for <3700 Ma events in the gneisses can be inferred from conclusions made about the Isua belt. Although the timing of tectonothermal events in the Isua belt is poorly known compared with the gneisses, tectonic settings for the belt have recently been proposed. Nutman et al. (2002) considered the large horizontal motions and interleaving that presumably occurred along metre-wide mylonite zones within the Isua belt, within the gneisses, and along some of the contacts between the belt and the gneisses as likely having occurred in the vicinity of plate margins. It was thus concluded that some form of plate tectonics contributed to Paleoarchaean crustal development. Hanmer and Greene (2002) showed that a thrust-nappe stack on the western side of the Isua belt formed during intrusion of 3640 ± 3 Ma tonalite (Hanmer et al., 2002), and concluded that there was a modern structural regime in the Paleoarchaean. A thrust-napped stack has not been identified in the northern or southern gneisses. However, the presence of such a structure could be masked by the monotonous nature of the tonalitic rocks that compose the gneisses; Hanmer and Greene (2002) identified the thrust-nappe stack in the Isua belt amongst heterogeneous lithologies. Future research of the timing and nature of metamorphic and deformational events in the Isua belt will test the hypothesis that they occurred in a plate tectonic setting, and will show whether events in the belt are related to those that affected the surrounding gneisses. Rocks that best preserve the effects of the early (>3658 Ma) tectonothermal events should also be studied futher.

Acknowledgements Field work was supported by the Isua Multidisciplinary Research Project, which was led by P. Appel and funded by the Danish National Science Research Council, the Commission for Scientific Research in Greenland, the Greenland Bureau of Minerals and Petroleum, and the Geological Survey of Denmark and Greenland. Financial support was also provided by the Voisey’s Bay Nickel–Paterson research chair held by J. Myers at Memorial University of Newfoundland and a Natural Sciences and Engineering

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