Journal of Petrology Advance Access published online on June 7, 2007
Journal of Petrology, doi:10.1093/petrology/egm024
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Petrology and Geochronology of Granulites from the McKaskle Hills, Eastern Amery Ice Shelf, Antarctica, and Implications for the Evolution of the Prydz Belt
1Institute of Geomechanics, Chinese Academy of Geological Sciences, Beijing 100081, China2Department of Earth Sciences, The University of Hong Kong, Pokfulam Road, Hong Kong, China3Shrimp Laboratory, Institute of Geology, Chinese Academy of Geological Sciences, Beijing 100026, China
Received February 10, 2006; Revised typescript accepted April 4, 2007
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONREGIONAL FRAMEWORK AND FIELD...SAMPLES AND ANALYTICAL...PETROLOGY AND THERMOBAROMETRYGEOCHRONOLOGYDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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A combined petrological and geochronological study was carried out on mafic granulites and associated felsic gneisses from the McKaskle Hills, eastern Amery Ice Shelf, East Antarctica. Garnet-bearing mafic granulites exhibit reaction textures and exsolution textures that indicate two-stage metamorphic evolution. Thermobarometric estimates from matrix and symplectite assemblages yield peak and retrograde P–T conditions of 9·0–9·5 kbar and 880–950°C and 6·6–7·2 kbar and 700–750°C, respectively. Similar but slightly scattered peak P–T estimates of 7·9–10·1 kbar and 820–980°C are obtained from the core compositions of minerals from felsic para- and orthogneisses. Evidence for the prograde history is provided by muscovite inclusions in garnet from a paragneiss. Sensitive high-resolution ion microprobe U–Pb zircon dating reveals an evolutionary history for the granulites, including a mafic and felsic igneous intrusion at 1174–1019 Ma, sedimentation after 932–916 Ma, and a high-grade metamorphism at 533–529 Ma. In contrast, Sm–Nd mineral–whole-rock dating mainly yields a single age population at
500 Ma. This suggests that the McKaskle Hills form part of the Prydz Belt, and that the relatively high peak P–T conditions and a decompression-dominated P–T path for the rocks resulted from a single Cambrian metamorphic cycle, rather than two distinct metamorphic events as formerly inferred for the granulites from Prydz Bay. The age data also indicate that the Precambrian history of the McKaskle Hills is not only distinct from that of the early Neoproterozoic terrane in the northern Prince Charles Mountains, but also different from that of other parts of the Prydz Belt. The existence of multiple basement terranes, together with considerable crustal thickening followed by tectonic uplift and unroofing indicated by the clockwise P–T–t evolution, suggests that the Prydz Belt may represent a collisional orogen that resulted in the assembly of Gondwana during the Cambrian period. KEY WORDS: Mesoproterozoic basement; Cambrian metamorphism; P–T path; Prydz Belt; East Antarctica
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONREGIONAL FRAMEWORK AND FIELD...SAMPLES AND ANALYTICAL...PETROLOGY AND THERMOBAROMETRYGEOCHRONOLOGYDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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One of the most important developments made in the field of Antarctic geology over the last decade is the recognition and confirmation of late Neoproterozoic–Cambrian (also termed Pan-African) high-grade tectonothermal events in East Antarctica. This has led to the breakdown of the concept of a single continuous late Mesoproterozoic–early Neoproterozoic (i.e. Grenvillian) Circum-East Antarctic Mobile Belt and, therefore, has challenged traditional Rodinia reconstructions (Dalziel, 1991
; Hoffman, 1991; Moores, 1991). Between two major early Neoproterozoic–Cambrian orogenic belt systems in East Antarctica, most geologists have considered the Lützow–Holm Belt and central Dronning Maud Land as a southern continuation of the East African Orogen, which could have resulted from the final collision between East and West Gondwana (e.g. Shiraishi et al., 1994; Shackleton, 1996; Jacobs et al., 1998, 2003; Fitzsimons, 2000). However, the Prydz Belt and Denman Glacier are located in the interior of what was previously considered to be a unified East Gondwana block and preserve polyphase tectonothermal histories of Archaean to early Palaeozoic age. Hence the nature and role of Cambrian tectonism in these areas in the assembly of Gondwana remain a subject for debate (Fitzsimons, 2003; Harley, 2003; Yoshida et al., 2003; Zhao et al., 2003, and references therein).
The eastern Amery Ice Shelf (EAIS) lies near the boundary between the Cambrian high-grade complex of Prydz Bay and the early Neoproterozoic high-grade complex of the northern Prince Charles Mountains (nPCM) (Fig. 1), and could be a key area for understanding the tectonic evolution of the Prydz Belt. Earlier studies suggested that rocks exposed in the Reinbolt Hills were metamorphosed under low- to medium-pressure granulite-facies conditions and were assumed to be related to an early Neoproterozoic tectonothermal event based on the occurrence of a synmetamorphic pegmatite with a U–Pb zircon age of 896 Ma (Grew & Manton, 1981; Nichols & Berry, 1991). However, a Th–U–total Pb monazite age of 536 Ma recently obtained from a sillimanite-bearing pegmatite has led to a suggestion that the area may have been affected by a Cambrian metamorphic event (Ziemann et al., 2005). As no geochronological constraints are available for the metamorphic rocks, it is difficult to determine the timing of regional high-grade metamorphism in the EAIS and to determine whether or not the area has experienced poly-phase metamorphism as inferred from Prydz Bay (Henson & Zhou, 1995; Hensen et al., 1995).
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To investigate the above issues, we undertook a petrological and geochronological study of garnet-bearing mafic granulites and associated felsic gneisses exposed in the McKaskle Hills, northern EAIS. We deduce a clockwise P–T path for these rocks based on mineral reaction textures and thermobarometric calculations and demonstrate that this P–T path is related to a single Cambrian metamorphic event on the basis of sensitive high-resolution ion microprobe (SHRIMP) U–Pb zircon dating and Sm–Nd mineral-whole-rock dating. These new data are then used to discuss the tectonic setting of the EAIS and the evolution of the Prydz Belt in the context of Gondwana assembly.
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REGIONAL FRAMEWORK AND FIELD RELATIONS |
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TOPABSTRACTINTRODUCTIONREGIONAL FRAMEWORK AND FIELD...SAMPLES AND ANALYTICAL...PETROLOGY AND THERMOBAROMETRYGEOCHRONOLOGYDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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The Prince Charles Mountains–Prydz Bay area comprises low- to high-grade metamorphic rocks only within distinct crustal blocks that range in age from Archaean to late Mesoproterozoic–early Neoproterozoic and Cambrian (Fig. 1). Three Archaean blocks, exposed in the Vestfold Hills, Rauer Group and southern Prince Charles Mountains (sPCM), preserve distinct crustal histories and therefore do not form remnants of a single unified craton (Harley, 2003
). The Vestfold Hills are dominated by granulite-facies orthogneiss interlayered with metapelitic supracrustal rocks. Metamorphism, deformation and subsequent mafic–felsic intrusion were dated between 2520 Ma and 2480 Ma (Black et al., 1991; Snape et al., 1997). The Rauer Group is a composite terrane containing Archaean (>3300 Ma and 2840–2800 Ma) tonalitic orthogneisses and late Mesoproterozoic (1030–1000 Ma) mafic to felsic intrusive rocks (Kinny et al., 1993; Harley et al., 1998). Both rock types have been recrystallized during Cambrian high-grade metamorphism (Harley et al., 1998; Kelsey et al., 2003a). The sPCM comprises Archaean (3200–3000 Ma) granitic gneiss basement interleaved with sedimentary rocks of different age groups (Tingey, 1991; Mikhalsky et al., 2001; Boger et al., 2006; Phillips et al., 2006). This basement and at least some of the sedimentary rocks underwent greenschist- to amphibolite-facies metamorphism and deformation prior to 2650 Ma (Boger et al., 2001, 2006).
Late Mesoproterozoic rocks are exposed in the southern sector of the nPCM (i.e. Fisher terrane), where mafic–felsic volcanism and intrusion occurred at 1300–1020 Ma, followed by amphibolite-facies metamorphism at 1020–940 Ma (Beliatsky et al., 1994; Kinny et al., 1997; Mikhalsky et al., 1999). Nd model ages (TDM) for these igneous rocks range from 1780 to 1360 Ma (Mikhalsky et al., 2006). Early Neoproterozoic rocks in the nPCM and adjacent Mawson Coast (i.e. Rayner Complex) are characterized by granulite-facies metamorphism accompanied by widespread charnockitic and granitic magmatism dated at 990–900 Ma (Manton et al., 1992; Kinny et al., 1997; Boger et al., 2000; Carson et al., 2000). The TDM ages of the granitoids range from 2160 to 1600 Ma (Young et al., 1997; Zhao et al., 1997a). Metamorphism is of high-temperature, low- to medium-pressure type (6–8 kbar and 800–880°C), characterized by an isobaric cooling P–T path (Fitzsimons & Harley, 1992, 1994a; Hand et al., 1994; Stephenson & Cook, 1997; Boger & White, 2003). In addition, intraplate deformation and emplacement of minor pegmatite sheets at 550–500 Ma have been observed in some places (Carson et al., 2000; Boger et al., 2002). This may be a response to the intense Cambrian tectonism that occurred in the Prydz Belt.
The Prydz Belt mainly crops out along the Prydz Bay coast (Fig. 1). Equivalently or similarly aged rocks are also exposed in the Grove Mountains (Zhao et al., 2000; Liu et al., 2002) and in the northern part of the sPCM (Boger et al., 2001; Corvino et al., 2005). The high-grade metamorphic rocks in Prydz Bay comprise two lithological associations: mafic–felsic composite orthogneisses and migmatitic paragneisses, which are attributed to the basement and cover sequences, respectively (Fitzsimons & Harley, 1991; Dirks & Hand, 1995; Dirks & Wilson, 1995). Geochronological studies suggest that the basement orthogneisses were formed at 1100 Ma with TDM ages of 2300–1800 Ma (Hensen & Zhou, 1995; Zhao et al., 1995, 2003; Wang et al., 2003), whereas the cover sequence is assumed to have been deposited during the late Mesoproterozoic (Dirks & Wilson, 1995; Carson et al., 1996) or Neoproterozoic (Hensen & Zhou, 1995; Zhao et al., 1995). The peak metamorphic P–T conditions in different localities of the area have been estimated at 6–7 kbar and 760–860°C, followed by a P–T path involving near-isothermal decompression (Motoyoshi et al., 1991; Fitzsimons, 1996; Carson et al., 1997). The major tectonothermal events, including deformation, high-grade metamorphism and emplacement of syn- to post-orogenic granites, took place between 550 Ma and 490 Ma (Zhao et al., 1992, 1997b, 2000; Hensen & Zhou, 1995; Carson et al., 1996; Zhang et al., 1996; Fitzsimons et al., 1997). In addition, an earlier, relatively high-pressure event (9–10 kbar and 850–980°C) is preserved in some areas (Thost et al., 1991; Ren et al., 1992; Tong & Liu, 1997). In rocks from Søstrene Island, this event has been dated at 990 Ma using the Sm–Nd isochron method (Hensen et al., 1995). The rocks from the Grove Mountains record peak P–T conditions (6·1–6·7 kbar and 850°C) similar to those from Prydz Bay (Liu et al., 2003). However, in the rocks of the sPCM, Cambrian peak metamorphic temperatures reach only 650–700°C at 6–7 kbar (Boger & Wilson, 2005).
The principal outcrops in the EAIS area (Manning Nunataks, Reinbolt Hills, Jennings Promontory, McKaskle Hills and Landing Bluff) are dominated by orthopyroxene-bearing orthogneisses interleaved with minor paragneisses, calc-silicate rocks and mafic granulites (Tingey, 1981; Sheraton et al., 1996). Although a few Neoproterozoic–Cambrian ages have been reported for some pegmatites in the Reinbolt Hills (Grew & Manton, 1981; Ziemann et al., 2005) and from granites in Landing Bluff (Tingey, 1991), the timing of regional high-grade metamorphism remains poorly constrained. The McKaskle Hills in the northern EAIS mainly consist of brown, layered or massive orthopyroxene-bearing orthogneisses. These rocks contain variable amounts of biotite, plagioclase, K-feldspar and Fe–Ti oxides (Fig. 2). Garnet, clinopyroxene and amphibole may additionally be present but all three of these minerals never coexist in one sample. Pale garnet-bearing but orthopyroxene-absent orthogneiss and garnet–orthopyroxene-bearing paragneiss are also present in the northern McKaskle Hills. Thin calc-silicate bands comprising garnet, clinopyroxene, scapolite and plagioclase are interleaved with the paragneiss. The original relationships between these metasedimentary rocks and the orthogneisses have been obliterated by deformation and metamorphism. Two-pyroxene mafic granulites are widespread. These rocks mostly occur as irregular blocks, lenses or layers (1–15 m wide) in various orthogneisses. Many layers of mafic granulite define open–tight folds and preserve axial-planar foliation that is parallel to the gneissosity of the host gneisses. Garnet-bearing mafic granulite occurs in garnet-bearing orthogneiss at the northeastern end of the outcrops (70°00'24"S, 72°59'29"E). This unit is relatively small, forming a 1·5 m wide layer that extends more than 10 m along the strike (Fig. 3). In addition, a number of shallowly dipping (10–15°) granitic dykes with widths of several metres occur in the southern McKaskle Hills. These dykes obliquely cut the foliation of country gneisses and are assumed to be Cambrian intrusions.
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SAMPLES AND ANALYTICAL TECHNIQUES |
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TOPABSTRACTINTRODUCTIONREGIONAL FRAMEWORK AND FIELD...SAMPLES AND ANALYTICAL...PETROLOGY AND THERMOBAROMETRYGEOCHRONOLOGYDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Eleven garnet- and orthopyroxene-bearing samples were selected for detailed petrological study. These samples include five mafic granulites (70-11, 80-11, 80-1, 80-2 and 80-6), three felsic paragneisses (70-6, 70-22 and 71-7), and three felsic orthogneisses (72-7, 74-5 and 78-5, Fig. 2). A garnet-bearing mafic granulite (sample 80-1), a garnet-free mafic granulite (sample 72-1), a paragneiss (sample 71-7) and an orthogneiss (sample 72-7) were chosen for SHRIMP U–Pb zircon dating. Two garnet-bearing mafic granulites (samples 70-11 and 80-1) and two paragneisses (samples 70-6 and 71-7) were chosen for Sm–Nd mineral–whole-rock dating. These samples are all fresh and most appear to have avoided partial melting, except for some paragneisses.
The chemical compositions of the constituent minerals were analyzed using a JEOL JXA-8100 wavelength-dispersive electron microprobe at the MOE Key Laboratory of Orogenic Belts and Crustal Evolution of Peking University. The operating conditions were 15 kV accelerating voltage, a 10 nA beam current and a counting time of 10 s for each peak. The beam diameter was set to 2 µm for all minerals except for orthopyroxene lamellae (1 µm) and for micas (5 µm). Natural minerals were used as standards. Ferric iron in garnet, clinopyroxene and orthopyroxene was calculated based on the scheme of Droop (1987). Molecular formulae for amphibole were calculated based on Holland & Blundy (1994), with modifications following Dale et al. (2000). Ferric iron in biotite was estimated based on the assumption of Fe3+/(Fe2+ + Fe3+) = 0.116 (Holdaway et al., 1997). Total iron is reported as ferrous iron in muscovite. Representative mineral analyses are presented in Table 1. Mineral abbreviations are after Kretz (1983), except for amphibole (Amp).
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In an attempt to determine reintegrated pyroxene compositions (i.e. lamellae + host), we employed point reintegrations via spot analyses using a defocused beam (see Liu et al., 2003
). Specifically, the microprobe beam was defocused to 25 µm width, and 20 spots were collected equidistantly across grains containing regularly spaced lamellae. The spot analyses were then averaged to provide a reintegrated composition for each pyroxene.
Zircon analyses were performed using the SHRIMP II ion microprobe at the Beijing SHRIMP Centre, Chinese Academy of Geological Sciences. Zircons were extracted from the samples using conventional techniques, including crushing, sieving, heavy liquid separation and hand-picking. Zircon grains were mounted in an epoxy disc along with the TEMORA zircon standard and polished down to expose their centres. The internal structures of the zircons were characterized by cathodoluminescence (CL) imaging. Instrumental conditions and data acquisition procedures are the same as described by Williams (1998). A primary ion beam of 4·5 nA, 10 kV 
and 20 µm spot were used. The data were collected in sets of five scans per dataset with analyses of the unknown zircons interspaced with analyses of the reference zircon every three analyses. The measured 206Pb/238U ratios were corrected using reference zircon TEMORA (416·75 ± 0·24 Ma; Black et al., 2003). Correction for common Pb was made using the measured 204Pb and assuming a model common Pb composition (Stacey & Kramers, 1975) appropriate to the age of the mineral. The age uncertainties for individual analyses represent 1
, but the weighted mean 206Pb/238U or 207Pb/206Pb ages calculated using the computer program Isoplot (version 2.06; Ludwig, 1999) are quoted at the 95% confidence level.
Mineral–whole-rock Sm–Nd isotopic analyses were conducted at the Institute of Geology and Geophysics, Chinese Academy of Sciences. Garnet, clinopyroxene, orthopyroxene, biotite and plagioclase were separated using a magnetic separator followed by hand-picking under a binocular microscope. The analytical procedures are the same as those reported by Yang et al. (2004). Powder samples were weighed and spiked with 149Sm and 150Nd and then dissolved in acid (HF + HNO3). Dissolution was carried out in Teflon vials at about 100°C for 12 days. Sm and Nd were separated using conventional ion exchange procedures. Mass analysis was performed using a VG 354 multicollector mass spectrometer. Nd isotopic fractionation was corrected assuming 146Nd/144Nd = 0·7219. During the period of data acquisition, the measured isotope ratios for the La Jolla Nd standard was 143Nd/144Nd = 0·511862 ± 0·000007 (2, n = 6). Total procedural blanks were <100 pg. Sm–Nd isochron ages were also calculated using the Isoplot program of Ludwig (1999). Input errors (2) used in age calculations are 147Sm/144Nd = 1% and 143Nd/144Nd = 0·005%. The quoted errors in isochron ages represent two standard deviations (2). Sm–Nd model ages (TDM) were calculated in two ways. The one-stage model age (TDM1) is calculated assuming a linear isotopic evolution of the depleted mantle reservoir from 
Nd(T) = 0 at 4·56 Ga to + 10 at the present. The two-stage model age (TDM2) is obtained assuming that the protolith of the granitic magmas has the Sm/Nd ratio of the average continental crust (Keto & Jacobsen, 1987). Nd isotopic normalization parameters used for the calculation of model ages are the same as those of Liu et al. (2006).
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PETROLOGY AND THERMOBAROMETRY |
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TOPABSTRACTINTRODUCTIONREGIONAL FRAMEWORK AND FIELD...SAMPLES AND ANALYTICAL...PETROLOGY AND THERMOBAROMETRYGEOCHRONOLOGYDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Mineral assemblages and textures
Mafic granulites
Mafic granulites are composed of garnet, clinopyroxene, orthopyroxene, amphibole, intergrown magnetite–ilmenite, apatite and zircon, with or without plagioclase. These minerals (except the accessory phases) form coarse-grained (1–2 mm), near-equigranular textures (Fig. 4a), and are inferred to have equilibrated at peak metamorphic conditions. The mafic granulites can be divided into two subtypes. Type I granulites (samples 70-11 and 80-11) have higher modal abundances of garnet (25–40%), greater modal clinopyroxene than orthopyroxene, and no preserved plagioclase. Type II granulites (samples 80-1, 80-2 and 80-6) contain plagioclase, less garnet (5–10%) and less modal clinopyroxene relative to orthopyroxene. Early amphibole in both subtypes is minor (<5%) and appears in textural equilibrium with the other phases. Garnet contains inclusions of clinopyroxene, orthopyroxene, amphibole, plagioclase (with the exception of sample 70-11), quartz, intergrown magnetite–ilmenite, apatite and less commonly biotite (Fig. 4a). The suite of minerals occurring as inclusions in garnet is very similar to that in the matrix. The exception is quartz, which, although present in garnet, is absent from the matrix. Zircon is observed either as intergranular grains or as inclusions in other minerals.
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In all mafic granulites, garnet porphyroblasts, especially when in contact with clinopyroxene, are partially replaced by fine-grained, radial and vermicular symplectites. The symplectites are typically composed of orthopyroxene and plagioclase, with or without trace amounts of clinopyroxene, amphibole and intergrown magnetite–ilmenite (Fig. 4b and c). These textures are interpreted to have resulted from the following two reactions (Harley, 1988
; Thost et al., 1991):
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Coarse matrix clinopyroxene contains two sets of orthopyroxene lamellae exsolved parallel to the (100) and (001) planes of the host crystals (Fig. 4g and h). Lamellae parallel to (100) planes are long, relatively coarse (maximum up to 2 µm wide) and closely spaced, whereas lamellae parallel to the (001) planes are short (<20 µm in length), narrow (mostly <0·5 µm wide), and very rare. The (100) lamellae are commonly contiguous to symplectic orthopyroxene mantling garnet (see Fig. 4c). Rare Fe–Ti oxide rods usually occur accompanying the orthopyroxene exsolution (Fig. 4h). Coarse matrix orthopyroxene contains very narrow exsolution lamellae (c. 0·2 µm wide) of clinopyroxene as well as Fe–Ti oxides (Fig. 4i and j). Both lamellae minerals can occur on the same (100) plane of the host orthopyroxene. Lamellae of Fe–Ti oxide are discontinuous and occur as parallelogram-like plates 20–80 µm in length and 0·5–1 µm in width.
Felsic paragneisses and orthogneisses
Felsic para- and orthogneisses both contain garnet, orthopyroxene, biotite, plagioclase, K-feldspar, quartz, intergrown magnetite–ilmenite, apatite and zircon, with or without spinel and monazite (Fig. 4k). Paragneisses are distinguished from orthogneisses by having higher modes of garnet (8–20%), orthopyroxene (10–25%) and biotite (mostly >5%) and lower modes of K-feldspar (<2%). Garnet is hypidiomorphic or, in some cases, xenomorphic, with grain sizes ranging from 0·5 to 5 mm. Most garnets contain numerous inclusions of biotite, plagioclase, quartz and intergrowths of magnetite, ilmenite and/or spinel. In sample 70-22, muscovite, which appears to be in textural equilibrium with biotite, plagioclase and quartz, is also observed as inclusions in garnet (Fig. 4l). Orthopyroxene is hypidiomorphic or xenomorphic and may contain inclusions of biotite, plagioclase and quartz. As in the mafic granulites, rare Fe–Ti oxide lamellae are exsolved from orthopyroxene. Spinel occurs in some samples (70-22, 71-7 and 74-5), occurring as blebs or blades within magnetite–ilmenite intergrowths. K-feldspar commonly contains numerous exsolved blebs of albite. In paragneiss samples, late K-feldspar is locally developed along fractures in plagioclase or in intergranular fractures. Vermicular symplectites of quartz and plagioclase also occur around relict biotite or K-feldspar. These might reflect the products of partial melting at or near peak metamorphism.
Mineral compositions
The end-members of garnet were calculated after the scheme of Deer et al. (1992). Garnet in the mafic granulites consists predominantly of almandine (XAlm 0·56–0·58), pyrope (XPrp 0·20–0·23) and grossular (XGrs 0·14–0·20), with minor spessartine (XSps 0·02–0·03) and andradite (XAdr 0·00–0·05) (Fig. 5). Garnet is not significantly zoned in XAlm, XPrp, XGrs and XSps, being characterized by flat profiles across much of a grain, decreasing in XPrp (0·13–0·15) and increasing in XAlm (0·61–0·63) and XSps (0·04–0·05) within 25–50 µm of the rim when in contact with clinopyroxene or symplectite (Fig. 6). Garnet from felsic para- and orthogneisses is dominated by XAlm (0·57–0·65) and XPrp (0·21–0·31), with minor XGrs (0·07–0·10), XSps (0·03–0·05) and XAdr (0·00–0·02). Generally, garnet from paragneisses is more magnesian than that from orthogneisses. Garnet from both of these rock types shows compositional zoning with decreasing XPrp and XGrs and increasing XAlm and XSps from core to rim (Fig. 5). This zoning is more pronounced when garnet is in contact with another ferromagnesian mineral, indicating significant Fe–Mg resetting during cooling.
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The end-members of clinopyroxene were normalized based on wollastonite (Wo) = Ca/(Ca + Mg + Fe2+), enstatite (En) = Ca/(Ca + Mg + Fe2+), and ferrosilite (Fs) = Ca/(Ca + Mg + Fe2+). Clinopyroxene from mafic granulites has compositions in the range of 43–49 mol % Wo, 32–39 mol % En and 15–23 mol % Fs. It contains 1·7–3·9 wt % Al2O3 and XMg [=Mg/(Mg + Fe2+)] ranges from 0·59 to 0·73. Clinopyroxene compositions vary with textural position. Within individual samples, the reintegrated matrix clinopyroxene has a lower concentration of Ca, a higher concentration of Al and a lower XMg when compared with clinopyroxene inclusions in garnet and the symplectic clinopyroxene or clinopyroxene rim adjacent to symplectic orthopyroxene and plagioclase.
Orthopyroxene in the mafic granulites has 1·0–2·5 wt % Al2O3 and XMg ranges from 0·46 to 0·54. Similar to clinopyroxene, orthopyroxene compositions vary with textural position. Generally, the symplectic orthopyroxene has lower Al and Ca contents, whereas orthopyroxene inclusions in garnet have lower Ca contents and XMg relative to the reintegrated matrix orthopyroxene. Orthopyroxene from felsic para- and orthogneisses has 1·3–4·7 wt % Al2O3 and XMg ranges from 0·46 to 0·65. Orthopyroxene in the paragneisses commonly has a higher XMg and Al content than that from the orthogneisses. Individual orthopyroxene grains typically show a decrease of Al and Ca and an increase of XMg from core to rim.
Amphibole in mafic granulites is pargasitic (CaB > 1· 8; NaA + KA > 0·7; Ti < 0·4; Si = 6·07–6·28 p.f.u.) with XMg ranging from 0·53 to 0·64. Amphibole inclusions in garnet have relatively high XMg, whereas the large retrograde amphibole grains have a lower XMg. The XMg of amphibole in the matrix and symplectites is variable. Nevertheless, within a single sample, amphibole in symplectites exhibits both higher Si content and XMg when compared with amphibole in the matrix.
Biotite inclusions in garnet in mafic granulites have TiO2 contents of 1·9–5·0 wt % and XMg of 0·73–0·81. Retrograde biotite has XMg of 0·66–0·69 and a nearly constant TiO2 content of 4·2–4·4 wt %. For felsic para- and orthogneisses, biotite in different samples has variable TiO2 contents (2·5–7·8 wt %) and XMg (0·56–0·78). However, within a given sample, biotite inclusions in garnet have higher TiO2 contents and XMg compared with matrix biotite. No significant core-to-rim variation in composition was observed for matrix biotite grains. Inclusions of muscovite in garnet from paragneiss sample 70-22 show little compositional variation. The Si values range from 3·93 to 3·00 p.f.u., and XMg values from 0·86 to 0·88.
The end-members of plagioclase were normalized based on anorthite (An) = Ca/(Ca + Na) and albite (Ab) = Na/(Ca + Na). Plagioclase from mafic granulites has dominantly labradorite and anorthite compositions depending on textural position. Matrix plagioclase exhibits pronounced compositional zoning, with a homogeneous core (XAn 0·55–0·62) and a distinctly anorthitic rim (XAn 0·90–0·95) (Fig. 7). Plagioclase inclusions in garnet have XAn of 0·60–0·62, similar to the core compositions observed for matrix plagioclase. However, within individual samples, inclusion plagioclases may have slightly higher XAn than matrix plagioclase cores. Symplectic plagioclase has the highest XAn (0·95–0·97). Plagioclase from felsic para- and orthogneisses is andesine to labradorite in composition (XAn 0·36–0·57). Most grains display slight compositional zoning with a rimward enrichment in Ca of less than XAn = 0·04.
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Thermobarometry and P–T path
To evaluate P–T conditions for the different stages of granulite-facies metamorphism in the McKaskle Hills, we employ average P–T calculations using THERMOCALC and thermobarometry based on the retrieval technique. Average P–T calculations followed the method of Powell & Holland (1994
), using an updated and expanded version of the internally consistent thermodynamic dataset (Holland & Powell, 1998) and the software THERMOCALC 3.1. The activities for end-members in garnet, clinopyroxene, orthopyroxene, amphibole, biotite, plagioclase and K-feldspar were calculated using the software AX (Holland, available at http://www.esc.cam.ac.uk/astaff/holland/index.html). The fluid phase was assumed to be an H2O-bearing mixture. The convergence method for the recovery of peak metamorphic P–T conditions (Fitzsimons & Harley, 1994b; Pattison & Bégin, 1994) adjusts Fe–Mg ratios in both garnet and orthopyroxene until the temperature estimates using a Fe–Mg exchange thermometer (e.g. Harley, 1984) and an Al-solubility thermometer (e.g. Harley & Green, 1982) converge. This technique has been successfully used on other granulite terranes (e.g. Bégin & Pattison, 1994; Chacko et al., 1996; Jones & Escher, 2002). Here we employ the RCLC program (Pattison et al., 2003) which uses the TWQ 2.02b thermodynamic dataset (Berman & Aranovich, 1996) with modifications to incorporate the experiments of Aranovich & Berman (1997) on Al solubility of orthopyroxene.
The results of average P–T calculations using THERMOCALC for mafic granulites and felsic gneisses are given in Table 2, and a comparison of uncorrected Fe–Mg, uncorrected Fe–Al and corrected Fe–Mg–Al P–T estimates using RCLC given in Table 3. The core compositions of garnet, amphibole and plagioclase and the reintegrated compositions of clino- and orthopyroxenes for mafic granulites yield average P–T estimates of 883–948°C and 9·0–9·5 kbar for the peak stage, whereas the rim compositions of garnet together with symplectic clinopyroxene (or clinopyroxene rim) + orthopyroxene + amphibole + plagioclase yield average P–T estimates of 710–736°C and 6·4–7·3 kbar for the retrograde stage. In contrast, the corrected Fe–Mg–Al P–T estimates from garnet–orthopyroxene pairs are 763–902°C and 8·3–10·6 kbar for the peak stage, and 698–748°C and 6·6–7·2 kbar for the retrograde stage. The corrected Fe–Mg–Al P–T estimates for the retrograde stage are equivalent to those determined using THERMOCALC, but both retrieval temperatures and pressures for the peak stage (except for sample 80-6) are significantly lower than the THERMOCALC average P–T estimates. As usual, the uncorrected Fe–Mg temperatures are successively lower than the uncorrected Fe–Al and corrected Fe–Mg–Al temperatures, consistent with more rapid diffusion of Fe–Mg vs Al (Fitzsimons & Harley, 1994b; Pattison & Bégin, 1994). However, as many mafic granulites from other areas have shown (Pattison et al., 2003), a reverse order for Fe–Mg, Fe–Al and Fe–Mg–Al temperature estimates is given by the peak mineral assemblages of samples 80-1 and 80-2. The reason for this unreasonable pattern is unclear, but thermodynamic inaccuracy is the most likely single explanation (Pattison et al., 2003).
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Because no retrograde reaction textures occur in the felsic gneisses we can only estimate the peak P–T conditions for the rocks. The average P–T estimates using the core compositions of garnet + orthopyroxene + biotite + plagioclase + K-feldspar + quartz are 698–880°C and 5·9–7·9 kbar, whereas the corrected Fe–Mg–Al P–T estimates are 816–979°C and 7·9–10·1 kbar. In contrast to the peak P–T conditions obtained for mafic granulites, the temperature and pressure estimates for the felsic gneisses determined using THERMOCALC are significantly lower. This suggests that the Fe–Mg compositions of the ferromagnesian minerals used in the THERMOCALC calculations might have been reset during cooling. Furthermore, two samples (70-6 and 72-7) also yield relatively low temperatures and pressures using retrieval thermobarometry.
Considering that the P–T results obtained by RCLC are unreasonable for the matrix assemblage of the mafic granulites and slightly scattered for the felsic gneisses, we prefer the average P–T conditions of 880–950°C and 9·0–9·5 kbar obtained by THERMOCALC from the matrix assemblage of the mafic granulites as the best estimation of the peak metamorphic conditions in the McKaskle Hills. The pre-peak history is indicated by the preservation of quartz inclusions in garnet in the mafic granulites and muscovite inclusions in garnet in the paragneiss. Although P–T conditions for this earlier stage cannot be quantitatively determined as a result of the compositional homogenization of garnet during peak metamorphism, the experimental results suggested that the garnet-in reaction for quartz tholeiites occurred above 7·5–9·0 kbar at 700–800°C (Green & Ringwood, 1967; Rushmer, 1991). The upper stability limit of muscovite is constrained by the muscovite-out reaction, which lies at temperatures not exceeding 650–750°C at 4–9 kbar (Spear et al., 1999; White et al., 2001). The nearly identical P–T estimates of about 700–750°C and 6·6–7·2 kbar obtained by both RCLC and THERMOCALC for symplectites place a good constraint on the retrograde conditions for the mafic granulites. Therefore, the post-peak P–T trajectory of the granulites from the McKaskle Hills involves a decompression of 2–3 kbar concurrent with cooling of 200°C from the metamorphic peak (Fig. 8).
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TOPABSTRACTINTRODUCTIONREGIONAL FRAMEWORK AND FIELD...SAMPLES AND ANALYTICAL...PETROLOGY AND THERMOBAROMETRYGEOCHRONOLOGYDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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SHRIMP U–Pb zircon dating
Sample 80-1: garnet-bearing mafic granulite
Zircons from this sample are colorless or pale yellow and rounded to prismatic in shape, with grain lengths of 70–200 µm. Based on CL images, four distinct domains can be distinguished in these zircon grains. These are banded cores, patchy cores, strongly luminescent rims, and weakly luminescent rims or individual grains. The growth zones in banded cores are broad and straight, and are commonly surrounded by strongly luminescent rims (Fig. 9a). The patchy cores have an extremely low CL response, and in many cases, contain moderately luminescent patches (Fig. 9b and c). Such cores are all embayed by thin or thick, strongly luminescent rims. Amphibole, clinopyroxene, plagioclase and apatite inclusions were observed in both banded and patchy cores. The strongly luminescent rims are commonly homogeneous, but may preserve banded zoning. A more strongly luminescent band commonly occurs on the outer borders of banded or patchy core (Fig. 9a and c). The weakly luminescent rims are very rare, and occur only as overgrowths on banded cores (Fig. 9d). Individual grains, which are nearly isometric and may contain a banded core, are also weakly luminescent, but fir-tree sector zoning is commonly visible in the CL images (Fig. 9e and f).
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Forty-seven analyses of 43 zircon grains were obtained on this sample (Table 4 and Fig. 10a). Apart from two analyses (21.1 and 29.1) that are significantly discordant, banded cores yield a very large age range from 1073 ± 45 Ma (207Pb/206Pb age) to 545 ± 15 Ma (206Pb/238U age), with Th/U ratios ranging from 0·36 to 0·73. A group of nine grains form a concordant population with a weighted mean 207Pb/206Pb age of 1019 ± 33 Ma (MSWD = 0·45). Other banded cores show younger ages that probably reflect Pb loss related to a subsequent thermal event. The data from strongly luminescent rims are mostly discordant and show large errors on 207Pb/206Pb ratios and ages as a result of the small analytical spot size used and the low abundances of U (60–173 ppm) and Th (9–52 ppm). Nearly concordant data from such rims and their enclosed patchy cores give a highly scattered age distribution with a maximum 207Pb/206Pb age of 1012 ± 91 Ma and a minimum 206Pb/238U age of 503 ± 17 Ma. The data from weakly luminescent rims or individual grains are nearly concordant and grouped, yielding 206Pb/238U ages of 561 ± 14 to 500 ± 14 Ma with a mean of 533 ± 9 Ma (n = 15, MSWD = 1·5). Their Th/U ratios are relatively high, ranging from 0·57 to 0·92.
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The broad and straight bands preserved in the zircon cores seem not to be a typical feature of magmatic zircons, although similar zoned patterns have been observed in zircon domains from some gabbros (Rubatto & Gebauer, 2000
). However, the absence of garnet and orthopyroxene inclusions in the banded domains also precludes their formation by metamorphic overgrowth. Thus, the banded cores are likely to be magmatic in origin, but could have been affected by solid-state recrystallization (Hoskin & Black, 2000). In such a case, the concordia age of 1019 ± 33 Ma obtained for these zircon domains is interpreted to be the minimum emplacement age of the mafic granulite protolith. The fir-tree sector zoning preserved in weakly luminescent zircon rims or individual grains is similar to the inner structure of some granulite-facies zircons (Vavra et al., 1996) and interpreted as the result of metamorphic overgrowth. Therefore, the concordia age of 533 ± 9 Ma obtained for such zircon domains is considered to be the age of the Cambrian metamorphic event. The relatively high Th/U ratios for such zircon domains may have resulted from open-system behaviour, breakdown of high Th/U minerals or competition with high-U minerals during metamorphism, and cannot be simply regarded as an indicator of a non-metamorphic origin (e.g. Möller et al., 2003).
The preservation of prismatic zircons suggests that the strongly luminescent rims were formed through recrystallization of previously existing zircons rather than as new overgrowths. The more strongly luminescent bands developed near banded or patchy cores could represent recrystallization fronts, as described by Hoskin & Black (2000), whereas the patchy cores could be the modified and disrupted domains of former igneous zircons. Hence, the age data scattered from late Mesoproterozoic to Cambrian obtained for these zircon domains probably reflect partial or entire resetting of the U–Pb isotopic system during Cambrian recrystallization of pre-existing zircons. They are consequently inferred to provide no meaningful geological data. Alternatively, these ages could represent a late Mesoproterozoic–early Neoproterozoic thermal event at 1000 Ma that has been strongly reworked by a later Cambrian metamorphic event at 530 Ma. However, inferring a metamorphic event without new zircon growth and/or age clustering seems unlikely. Importantly, weakly luminescent rims have never been observed to grow on strongly luminescent domains, suggesting that these two zircon domains were not formed by two separate metamorphic events. Moreover, the presence of Cambrian metamorphic zircons growing directly on the Mesoproterozoic magmatic zircon domains in orthogneiss sample 72-7 (see below) does not support the occurrence of a 1000 Ma thermal event in the area.
Sample 72-1: garnet-free mafic granulite
Garnet-free mafic granulite sample 72-1 is representative of the regional two-pyroxene mafic granulites in the McKaskle Hills. Zircons from this rock are also colorless or pale yellow and rounded to prismatic in shape. They have grain lengths of 50–180 µm. All zircon grains have moderate to weak CL response and preserve banded zoning (Fig. 9g and h). Similar to sample 80-1, this zoning is interpreted to reflect a magmatic origin that may have been affected by solid-state recrystallization. Amphibole, clinopyroxene, plagioclase, quartz and apatite were found as inclusions in these zircons. The bands in the weakly luminescent grains are commonly thickened. Some such grains are partially embayed by moderately luminescent rims (Fig. 9h and i), which are interpreted to have formed during recrystallization of igneous zircon. Local, weakly luminescent overgrowths also occur on some zircon grains (Fig. 9h). These are assumed to be the expression of Cambrian metamorphic growth, but the domains are too thin to be analyzed.
Twenty-five analyses on 21 zircon grains were performed on this sample (Table 4 and Fig. 10b). The banded zircons show a scattered age distribution ranging from 1216 ± 38 Ma (207Pb/206Pb age) to 586 ± 15 Ma (206Pb/238U age), with Th/U ratios ranging from 0·69 to 1·74. However, 10 analyses (207Pb/206Pb ages ranging from 1216 ± 38 Ma to 1077 ± 50 Ma) form a nearly concordant cluster giving a weighted mean 207Pb/206Pb age of 1122 ± 30 Ma (MSWD = 2·3). The remaining 13 analyses show different degrees of discordance produced by radiogenic Pb loss. Two ages from moderately luminescent rims are also discordant and show appreciable Pb loss. When combined, all 25 datasets yield a discordia with an upper intercept age of 1137 ± 46 Ma and a lower intercept age of 560 ± 58 Ma (MSWD = 1·7). The upper intercept age is equivalent within error to the mean age obtained from the banded zircon domains and is interpreted as the minimum emplacement age of the mafic granulite protolith. The inaccurate lower intercept age could reflect the effect of a metamorphic event at 530 Ma as revealed by sample 80-1.
Sample 71-7: paragneiss
Zircons from paragneiss sample 71-7 are pale yellow and rounded in shape and have grain sizes ranging from 30 to 100 µm. Both optical and CL images reveal a typical core–rim structure (Fig. 9j–n) for almost all the zircon grains. All the cores are small and well-rounded. Therefore, they are interpreted to be of detrital origin. These cores have extremely low CL response and appear unstructured or sector-zoned. In addition, an oscillatory-zoned domain is preserved in the inner part of a such core (Fig. 9j), which may represent a relic of an earlier igneous zircon. All the rims grew around a detrital core and show moderate CL response and planar growth banding or fir-tree sector zoning. Some grains have a weakly luminescent mantle and a moderately luminescent rim (Fig. 9k). Radial fractures developed around a detrital core are commonly visible in transmitted light (Fig. 9n). These rims are interpreted to be of metamorphic origin.
Twenty-one U–Pb analyses on 18 zircon grains reveal three age populations (Table 4 and Fig. 10c). The oldest 207Pb/206Pb age of 1096 ± 11 Ma (spot 2.1) was obtained for an oscillatory-zoned domain within a weakly luminescent core, with a Th/U ratio of 0·28. Two other weakly luminescent cores yield 207Pb/206Pb ages of 932 ± 15 Ma (spot 8.1) and 916 ± 19 Ma (spot 5.1). Their Th/U ratios are very low (< 0·10), typical of metamorphic zircon (e.g. Williams & Claesson, 1987; Rubatto & Gebauer, 2000). It should be noted that these two analyses are probably mixed with some rim composition because of the small size of the cores. However, because of the low U concentrations of the rim, the effect of this mixture will be minimal. The remaining six weakly luminescent cores show younger 207Pb/206Pb ages between 536 ± 17 Ma and 445 ± 33 Ma. These grains have Th/U ratios of 0·02–0·07 and show reverse discordance, which is interpreted to be due to high uranium contents (Williams, 1998). The 206Pb/238U ages obtained from the moderately luminescent rims give ages between 556 ± 13 Ma and 508 ± 13 Ma, with a mean of 533 ± 10 Ma (n = 12, MSWD = 1·5). Th/U ratios for these domains range from 0·16 to 1·33. The ages of zircon rims are interpreted to represent the timing of metamorphism of the paragneiss. In contrast, two concordant data for detrital zircon cores are interpreted as the minimum age of early Neoproterozoic metamorphism, whereas other detrital zircon cores may have experienced recrystallization during the Cambrian metamorphism. The age of the oscillatory-zoned zircon domains is in agreement with the ages of igneous events obtained from other samples.
Sample 72-7: orthogneiss
Zircons from orthogneiss sample 72-7 are pale brown and prismatic, and 200–380 µm in length. They all show simple zoning patterns characterized by oscillatory-zoned cores and moderately luminescent overgrowths (Fig. 9o and p). Biotite, K-feldspar, plagioclase, quartz, magnetite and apatite were observed as inclusions in the oscillatory-zoned cores. The overgrowths show planar growth banding and sector zoning and are interpreted to have formed during metamorphism.
The results of 23 analyses on 17 zircon grains form two populations (Table 4 and Fig. 10d). Oscillatory-zoned cores give 207Pb/206Pb ages ranging from 1268 ± 28 Ma to 1005 ± 44 Ma, with Th/U ratios ranging from 0·21 to 0·97. Of these, 12 concordant analyses yield a weighted mean age of 1174 ± 26 Ma (MSWD = 2·7). This age is interpreted as the igneous emplacement age of the orthogneiss protolith. The other five analyses showing younger ages probably reflect Pb loss caused by the subsequent metamorphism. The 206Pb/238U ages obtained for six moderately luminescent overgrowths form a single population (536 ± 14 Ma to 521 ± 14 Ma) with a mean of 529 ± 11 Ma (MSWD = 0·20). These overgrowths have low U but high Th concentrations with high Th/U ratios of 1·58–3·80. The age is interpreted to represent the timing of metamorphic recrystallization of the orthogneiss.
Sm–Nd mineral-whole-rock dating
The coarse-grained garnet, clinopyroxene and orthopyroxene in the mafic granulites and garnet, orthopyroxene, biotite and plagioclase in the paragneisses were separated for Sm–Nd isotopic dating. In the mafic granulites, symplectic clinopyroxene and orthopyroxene are easily separated because of their fine-grained size and intergrowth with plagioclase, whereas amphibole and plagioclase were not selected for isotopic analyses because of their multiple growths. However, garnet crystals commonly contain small amounts of mineral inclusions, of which REE-rich minerals, such as apatite and zircon, could be responsible for the increase of Sm and Nd concentrations and possible modification of Sm/Nd ratios in garnet. However, because the peak metamorphic temperatures are higher than the Nd isotope closure temperatures of the inclusions, the Nd isotopic compositions of the inclusions would be expected to attain isotopic equilibrium with the host garnet. Consequently, an isochron constructed with such garnet and cogenetic phases can still provide a correct age of metamorphism (Jahn et al., 2005).
Sm–Nd isotope data for four samples are presented in Table 5 and further illustrated in Fig. 11. Minerals and whole-rock isochron ages from two mafic granulites and one paragneiss are nearly identical, 500 ± 9 Ma (MSWD = 0·04) for sample 70-11, 499 ± 10 Ma (MSWD = 1·06) for sample 80-1, and 495 ± 6 Ma (MSWD = 0·97) for sample 71-7. The initial 143Nd/144Nd ratios (INd) are 0·511763 ± 0·000022, 0·511481 ± 0·000023 and 0·511491 ± 0·000015, respectively. The data for clinopyroxene from sample 80-1 deviated significantly from the garnet–orthopyroxene–whole-rock trend line and therefore were not involved in the age calculations. The Sm–Nd system for this clinopyroxene seems to have been disturbed during retrogression. The age results of 500–495 Ma are similar to those (515–490 Ma) obtained from granulites in Prydz Bay (Hensen & Zhou, 1995). The whole-rock Sm–Nd results for paragneiss sample 70-6 show a significant deviation from the garnet–orthopyroxene–plagioclase mineral isochron. In addition, the mineral isochron gives a slightly younger age of 457 ± 12 Ma with MSWD = 0·66 and INd = 0·511586 ± 0·000025. This suggests that the minerals may have re-equilibrated at a later time, or, more likely, isotopic disequilibrium between minerals yields an erroneous isochron. Similarly young ages were also reported from Bolingen Islands, Prydz Bay (Hensen & Zhou, 1995).
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The whole-rock analyses for two mafic granulites yield a two-point age of 1027 ± 50 Ma (INd = 0·511113 ± 0·000052). This age is consistent within error with the U–Pb age obtained for the banded zircon domains from sample 80-1, and therefore supports the interpretation that the protoliths of the mafic granulite were emplaced during late Mesoproterozoic time. Two mafic granulites have
Nd(T) values of –3·9 and –4·0, with Nd model ages (TDM1) of 4510 and 1890 Ma. This could suggest that the primary mafic magma was derived from the partial melting of an enriched subcontinental lithospheric mantle. The old (TDM1) age obtained for sample 70-11 may be caused by the unreasonably high 147Sm/144Nd ratio of the rock. Because two paragneisses have 147Sm/144Nd ratios (>0·14) higher than the average crustal ratio of 0·12, it is considered that the two-stage Nd model ages are more realistic for inferring the age of the source. The obtained (TDM2) ages for two samples are 2030 and 1780 Ma, respectively. |
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Peak P–T conditions and P–T–t path of Cambrian metamorphism
Based on geochronological data as well as microstructural and thermobarometric analyses from the Søstrene Island of Prydz Bay, most geologists have accepted that the Prydz Belt is a polymetamorphic belt that was affected by at least two episodes of high-grade metamorphism, in early Neoproterozoic and Cambrian times. The Cambrian metamorphic peak in this belt only reached low- to medium-pressure granulite-facies conditions (6–7 kbar and 760–860°C) (Motoyoshi et al., 1991
; Nichols & Berry, 1991; Thost et al., 1994; Fitzsimons, 1996; Carson et al., 1997), whereas the early Neoproterozoic metamorphism took place at relatively high-pressure and high-temperature conditions (9–10 kbar and 850–980°C) (Thost et al., 1991; Hensen et al., 1995; Tong & Liu, 1997). Moreover, in the Rauer Group, two distinct granulite-facies events have been recognized in the metapelites (Harley, 1987; Sims et al., 1994; Dirks & Wilson, 1995; Harley et al., 1998). It seems that Prydz Bay provides an excellent example showing that mineral reaction textures in a granulite sample may not necessarily reflect a single metamorphic cycle with only one P–T path. However, based on the result of in situ (Th + U)–Pb monazite dating, Kelsey et al. (2003a) have argued that the ultrahigh-temperature metamorphism at 950–1050°C and 9·5–12 kbar and the subsequent overprint at >800–850°C and >7 kbar as recorded in the metapelites of the Rauer Group (e.g. Harley & Fitzsimons, 1991; Harley, 1998; Kelsey et al., 2003b) could have occurred during the same Cambrian event.
Granulites from the McKaskle Hills also preserve two-stage mineral assemblages that recrystallized at 9·0–9·5 kbar and 880–950°C and 6·6–7·2 kbar and 700–750°C, respectively. These P–T estimates seem to be comparable with the P–T conditions of the two metamorphic events in Prydz Bay. However, U–Pb zircon and Sm–Nd mineral–whole-rock data, particularly the growth of new zircon rims on the early Neoproterozoic detrital zircon grains in the paragneiss, clearly suggest that the peak mineral assemblages of granulites from the McKaskle Hills were formed during the Cambrian metamorphism. This implies that the peak P–T conditions of the Cambrian high-grade event in the area are higher than those recorded in much of the Prydz Belt, and the inferred clockwise P–T path for the rocks is related to a single Cambrian metamorphic cycle, rather than two separate metamorphic events. A recent geochronological study of high-grade rocks from the Grove Mountains has revealed that zircon rims with ages of 550–535 Ma grew on magmatic zircon cores with ages of 920–910 Ma (X. Liu et al., 2007), also negating the existence of an early Neoproterozoic metamorphic event in that area. In fact, the recognition of a distinct two-stage decompression in the mafic granulites from Søstrene Island (Thost et al., 1991; Hensen et al., 1995) suggests a different metamorphic evolution compared with other granulites from the Prydz Belt. If so, it is likely that Søstrene Island represents a distinct terrane in the Prydz Belt and that this terrane might have no universal significance. As for the Rauer Group, there is still some possible disagreement in the timing of the ultrahigh-temperature event (Harley, 2003; Hokada et al., 2003; Tong & Wilson, 2006).
There is a difference of 30 Myr between the ages given by the U–Pb zircon and Sm–Nd mineral–whole-rock dating methods for the granulites from the McKaskle Hills. The closure temperature for the zircon U–Pb system is >900°C (Lee et al., 1997; Cherniak & Watson, 2001) although zircon can grow at very low temperatures (
550°C) in the presence of hydrous fluids (Rubatto et al., 1999; Liermann et al., 2002). In contrast, the closure temperature for the garnet Sm–Nd system remains controversial, ranging from 850°C for rapidly cooled terranes (Cohen et al., 1988; Jagoutz, 1988) to 600–650°C for slowly cooled ones (Humphries & Cliff, 1982; Mezger et al., 1992; Burton et al., 1995). This suggests that the age difference of 30 Myr in the present case is likely to result from the difference in closure temperature between the two isotopic systems. Accordingly, the U–Pb zircon age of 530 Ma is interpreted as the approximate age of the peak metamorphism, whereas the Sm–Nd isochron age of 500 Ma is interpreted as a decompressional cooling age. The time interval between the peak metamorphism and the blocking of the garnet Sm–Nd isotopic system defines a cooling rate of 7°C/Ma for granulites from the McKaskle Hills.
Precambrian history and regional correlations
SHRIMP U–Pb zircon dating of mafic granulite and orthogneiss from the McKaskle Hills reveals a mafic–felsic igneous intrusion at 1170–1020 Ma. These late Mesoproterozoic igneous rocks have been attributed to form part of the basement rocks in Prydz Bay (Fitzsimons & Harley, 1991; Dirks & Hand, 1995; Dirks & Wilson, 1995). The interpretation of the deposition ages of the cover sequence, characterized by migmatitic paragneisses in Prydz Bay, is arguable, with one school of thought considering it to be Mesoproterozoic (Dirks & Wilson, 1995; Carson et al., 1996), whereas others have interpreted it to be Neoproterozoic (Hensen & Zhou, 1995; Zhao et al., 1995). A few zircon domains from paragneiss sample 71-7 collected from the McKaskle Hills record an igneous event at 1096 Ma and a metamorphic event at 932–916 Ma. Such zircon domains are clearly of detrital origin as revealed by optical and CL images. This suggests that at least some paragneiss protoliths in the Prydz Belt were deposited during the Neoproterozoic. As the age of one detrital zircon population from the paragneiss is in good agreement with the ages of granulite facies metamorphism in the Rayner Complex (Young & Black, 1991; Boger et al., 2000; Kelly et al., 2002), it is believed that the materials of the paragneiss sequence came from a similar early Neoproterozoic orogen. The comparable TDM ages between the rocks from two areas lends support to this assumption.
The ages of 1170–1020 Ma obtained for the mafic granulites and orthogneisses demonstrate that the basement of the McKaskle Hills differs from the Rayner Complex of the nPCM, where all the igneous and metamorphic rocks were dated at <990 Ma except for a leucosome that yielded an inherited age of 1017 Ma (Boger et al., 2000). This suggests that the Rayner Complex does not extend to the east of the Lambert graben and form part of the Prydz Belt. Within East Antarctica, 1170–1040 Ma volcanic and intrusive events have been reported for the rocks in Dronning Maud Land (Harris et al., 1995; Jacobs et al., 1998, 2003; Paulsson & Austrheim, 2003; Board et al., 2005), the Fisher terrane of the nPCM (Beliatsky et al., 1994; Kinny et al., 1997; Mikhalsky et al., 1999) and the Bunger Hills–Windmill Islands area (Sheraton et al., 1990, 1992; Post et al., 1997). Additionally, 1300 Ma felsic volcanism and intrusion have also been reported for the rocks in the Fisher terrane. These Mesoproterozoic igneous events occurred in different tectonic settings, including a volcanic arc in Dronning Maud Land, an active continental margin in the Fisher terrene, and a post-orogenic environment in the Bunger Hills–Windmill Islands area. Rocks from the McKaskle Hills might also have formed at an active continental margin as inferred from the orthogneisses of the Jennings Promontory, 25 km SW of the McKaskle Hills (Sheraton et al., 1996); however, their low Nd(T) values are distinct from those of juvenile rocks from the Fisher terrane. Furthermore, the above three areas have all experienced late Mesoproterozoic medium- to high-grade metamorphism, although the metamorphism in the different areas was not contemporaneous. Therefore, the McKaskle Hills (and possibly the entire EAIS) might record a Precambrian history different from that of the Mesoproterozoic–Neoproterozoic terranes in East Antarctica and should represent a distinct basement terrane.
Implications for the evolution of the Prydz Belt
U–Pb zircon ages of 533–529 Ma for regional metamorphism in the McKaskle Hills, together with a U–Th–total Pb monazite age of 536 Ma obtained for a pegmatite in the Reinbolt Hills (Ziemann et al., 2005) and a U–Pb zircon age of 500 Ma obtained for a granite at Landing Bluff (Tingey, 1991), suggest that the EAIS is a part of the Prydz Belt. However, the different units in the Prydz Belt seem to record a different earlier tectonic evolution, although all of them have experienced the same intense Cambrian high-grade metamorphic episode. The Larsemann Hills resemble the McKaskle Hills by preserving an igneous event at 1100 Ma, but subsequent metamorphism and granite emplacement at 990–940 Ma have also occurred within the basement in this area (Hensen et al., 1995; Zhao et al., 1995; Wang et al., 2003). In the Grove Mountains, a mafic–felsic intrusive event has been dated at 920–910 Ma, and the detrital zircons from a paragneiss preserve an early Paleoproterozoic age of 2050 Ma (X. Liu et al., 2007). In the Rauer Group and in the northern part of the sPCM, Cambrian high-grade metamorphism was overprinted on the Archaean basement rocks (Kinny et al., 1993; Harley et al., 1998; Boger et al., 2001). The existence of multiple basement terranes in an orogen and the lack of similar terranes outside the orogen is difficult to interpret by an intracontinental reworking model as favored by some workers (e.g. Yoshida, 1995; Wilson et al., 1997; Yoshida et al., 2003).
The shape of the P–T path provides an indication of the tectonic processes involved in the formation and evolution of the granulites (e.g. Harley, 1989). The relatively high peak P–T conditions followed by decompression of 2–3 kbar inferred from the granulites in the McKaskle Hills indicate a considerable crustal thickening subsequently followed by tectonic uplift and unroofing during the Cambrian orogeny. This scenario is consistent with the widespread development of contemporaneous compressional and subsequent extensional deformation in the Prydz Belt (Carson et al., 1995; Dirks & Wilson, 1995; Liu et al., 2006). The intrusion of voluminous syn- to post-orogenic granitoids in the belt, particularly the widespread distribution of intermediate charnockites and sheeted granites in the Grove Mountains, suggests that lithospheric delamination and asthenospheric upwelling could have occurred during post-collisional extension (Liu et al., 2006). Coupled with the juxtaposition of multiple basement terranes, we favor the suggestion that the Prydz Belt represents a collisional orogen that resulted from the assembly of Gondwana during Cambrian times (e.g. Fitzsimons, 2000, 2003; Boger et al., 2001; Zhao et al., 2003).
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TOPABSTRACTINTRODUCTIONREGIONAL FRAMEWORK AND FIELD...SAMPLES AND ANALYTICAL...PETROLOGY AND THERMOBAROMETRYGEOCHRONOLOGYDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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U–Pb zircon and Sm–Nd mineral–whole-rock dating of mafic granulites and felsic gneisses from the McKaskle Hills reveals a peak metamorphic age of 533–529 Ma and a cooling age of
500 Ma, respectively. Hence the area is attributed to a part of the Prydz Belt. Thermobarometric estimates from mafic granulites suggest that the Cambrian metamorphism reached peak pressures of 9·0–9·5 kbar at 880–950°C during crustal thickening, followed by decompression to 6·6–7·2 kbar at 700–750°C as a consequence of orogenic exhumation and unroofing. U–Pb zircon dating also reveals that the basement of mafic granulites and orthogneisses in the McKaskle Hills was formed during a time interval of 1170–1020 Ma, whereas the cover sequence (now paragneisses) was deposited on the basement during the Neoproterozoic. This basement is not only distinct from the nPCM by its old protolith crystallization ages, but also slightly different from that of Prydz Bay by the absence of an early Neoproterozoic metamorphic event. Taking into account multiple basement juxtaposition and a clockwise P–T–t path, we prefer the suggestion that the Prydz Belt represents a collisional orogen that resulted from the assembly of Gondwana during the Cambrian period. |
SUPPLEMENTARY DATA |
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TOPABSTRACTINTRODUCTIONREGIONAL FRAMEWORK AND FIELD...SAMPLES AND ANALYTICAL...PETROLOGY AND THERMOBAROMETRYGEOCHRONOLOGYDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Supplementary data for total microprobe analyses of minerals in this paper are available at Journal of Petrology online.
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ACKNOWLEDGEMENTS |
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We would like to thank Liudong Ren for assistance during field work and for fruitful discussions and valuable comments on the manuscript. Zhuyin Chu and Ji Qiu ably assisted in the Sm–Nd isotopic analyses, and Guiming Shu in the electron microprobe analyses. N. M. Kelly, S. D. Boger, D. E. Kelsey, M. J. Flowerdew and R. W. White are thanked for their detailed and thorough critical reviews, and G. L. Clarke for his editorial comments, which not only greatly improved the manuscript, but also improved the English expression. The field work was carried out during the 2004–2005 Chinese National Antarctic Research Expedition. Logistic support by the Antarctic Administration of China and financial support by the Geological Investigation Project of the Chinese Geological Survey (1212010511505), National Natural Science Foundation of China (No. 40372046) and Programme of Young Scientists of the Ministry of Land and Resources are gratefully acknowledged.
*Corresponding author. Telphone: +86-10-68486756. Fax: +86-10-68422326. E-mail: liuxchqw{at}yahoo.com.cn
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