Journal of Petrology Advance Access published online on November 21, 2006
Journal of Petrology, doi:10.1093/petrology/egl064
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Resetting of the U–Pb Zircon System in Cambro-Ordovician Intrusives of the Deep Freeze Range, Northern Victoria Land, Antarctica
1Dipartimento Di Scienze Della Terra, Università Di Siena, 53100 siena, Italy
2Gemoc Arc National Key Centre, Department of Earth and Planetary Sciences, Macquarie University, Sydney, Nsw 2109, Australia
Received February 25, 2005; Revised typescript accepted September 29, 2006
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ABSTRACT |
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TOP
ABSTRACT
INTRODUCTIONA detailed in situ isotopic (U–Pb, Lu–Hf) and geochemical study of zircon populations in a composite sequence of foliated to massive Cambro-Ordovician intrusions in the Deep Freeze Range (North Victoria Land, Antarctica), has highlighted great complexity in zircon systematics. Zircons in deformed granitoids and tonalites display complex internal textures, a wide spread of concordant U–Pb ages (between 522 and 435 Ma) and unusual trace-element compositions (anomalous enrichment of light rare earth elements, U, Th and Y) within single zircon grains. In contrast, zircons from undeformed samples display a limited range of U–Pb ages and trace-element compositions. Zircons from all age and textural populations in most of the deformed and undeformed samples show a relatively narrow range of
Hf values, suggesting that the Lu–Hf system remained undisturbed. Inferred emplacement ages cover a time interval of about 30 Myr: from 508 to 493 Ma for the oldest strongly foliated synkinematic Howard Peaks megacrystic monzogranites and high-K calc-alkaline mafic to intermediate rocks of the ‘Corner Tonalite’ unit; from about 489 to 481 Ma for the younger massive shoshonitic mafic dyke suite and the high-K calc-alkaline Keinath granite. The observed isotopic and chemical variations in zircon are attributed to a sub-solidus recrystallization under hydrous conditions and varying temperature, in a setting characterized by a transpressional to extensional stress regime. KEY WORDS: Antarctica; Cambro-Ordovician intrusives; Ross Orogen; zircon U–Pb geochronology
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INTRODUCTION |
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Zircon is chemically and physically resistant to modification under most geological conditions; this characteristic makes it one of the most important tools to determine the age, origin and thermal history of crustal igneous and metamorphic rocks. However, many recent studies (e.g. Black et al., 1986
; Pidgeon, 1992; Wayne et al., 1992; Vavra et al., 1996, 1999; Pidgeon et al., 1998; Ashwal et al., 1999; Rubatto et al., 1999, 2001; Schaltegger et al., 1999; Hartmann et al., 2000; Hoskin & Black, 2000; Liermann et al., 2002; Rubatto, 2002; Rubatto & Hermann, 2002; Zeck & Whitehouse, 2002; Tomacheck et al., 2003; Whitehouse & Kamber, 2002; Whitehouse & Platt, 2003; Booth et al., 2005; Giacomini et al., 2005) have shown that a high degree of complexity in the internal textures of zircon grains may be associated with disturbance of the U–Pb system in metamorphic and igneous rocks, and often with changes to the chemical composition of the zircons. These studies clearly demonstrate that zircon does not always behave as a closed system. The interpretation of zircon U–Pb ages may therefore contain ambiguities. A number of processes such as igneous and metamorphic overgrowth, inheritance, annealing after radiation damage, hydrothermal alteration, precipitation and dissolution can severely affect the chemical and isotopic systems, leading to mixtures of age components within single zircon grains. The application of trace-element, U–Pb and Lu–Hf systematics using in situ analytical methods [ion-microprobe, laser ablation microprobe–inductively coupled plasma mass spectometry (LA-ICPMS) and multi-collector-LA-ICPMS)] on discrete domains within the same zircon grains has great potential for distinguishing between magmatic, metamorphic and altered zircons. Indeed, the integrated application of these techniques allows correlations between zircon ages, textures, and elemental and isotopic compositions, which are indispensable in dating specific geological events identified in the zircons’ growth history.
Many researchers have attributed ambiguity in the interpretation of U–Pb zircon dates principally to partial Pb-loss effects, which result in discordant arrays of radiometric ages. Examples of near-complete Pb loss have been ascribed to extreme processes of mylonitization (Wayne & Sinha, 1988) or leaching of metamict grains by fluids under low-temperature conditions (Gebauer & Grunenfelder, 1976). Recently, several cases of completely reset U–Pb isotopic systems, often accompanied by a redistribution of trace elements, have been described (e.g. Ashwal et al., 1999; Hoskin & Black, 2000; Giacomini et al., 2005).
In this paper we present a detailed in situ study of the isotopic (U–Pb and Lu–Hf) and trace-element composition in zircon from nine selected samples of a composite sequence of foliated and massive granitoids and mafic rocks cropping out in the Deep Freeze Range (Northern Victoria Land, Antarctica). The zircons have been analysed by scanning electron microprobe (EMP), LA-ICPMS and MC-LA-ICPMS.
Previous studies of the Wilson Terrane in the Northern Victoria Land constrain the age of intrusive magmatism to between 530 and 480 Ma (Armienti et al., 1990; Tonarini & Rocchi, 1994). However, the intrusive sequence in the Deep Freeze Range is still poorly constrained. Published Rb/Sr whole-rock–biotite pairs and K/Ar data on biotite from both metamorphic and magmatic rocks cropping out in the study area give younger ages between 475 and 449 Ma, which are interpreted as cooling ages (Borsi et al., 1987; Kreuzer et al., 1987; Vita et al., 1991). In addition, the Deep Freeze Range area was affected, during the Ross Orogeny, by ductile deformational events triggered by dextral strike-slip and/or reverse shear zones and several minor local shear bands (Skinner, 1991; Musumeci & Pertusati, 2000). In most of the studied rocks, shearing and deformation produced a pervasive solid-state ductile high-temperature deformation and recrystallization of the magmatic minerals and probably favoured a strong interaction with fluids, as suggested by a common pervasive presence of fluid inclusions. The wide spectrum of U–Pb ages and trace-element compositions in the zircons of some samples indicates that in such a setting, the isotopic and compositional systematics of zircon underwent a strong modification.
This study of zircon populations in selected plutonic rocks focuses on the relationships between syn- and post-magmatic events and the partial resetting of the U–Pb and trace-element chemistry of the zircons. Hf isotopes provide an invaluable tracer in this type of study; they make it possible to unravel the complex history of the zircon populations and help to identify multiple age components.
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GEOLOGICAL BACKGROUND |
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During Late Proterozoic–Early Palaeozoic time, active subduction was occurring along the palaeo-Pacific margin of the East Antarctic Craton. This subduction generated the Late Cambrian–Early Ordovician Ross–Delamerian Orogen, a 4000 km long mobile belt, now exposed on the border of the East Antarctic Craton and in South Australia (Stump, 1995
; Goodge, 2002; Goodge et al., 2004a, 2004b; Wysoczanski & Allibone, 2004, and references therein). The Transantarctic Mountains, generated by uplift of the craton-ward shoulder of the Cretaceous–Cenozoic rift, broadly parallel the older Ross Orogen and expose the crystalline basement.
In Northern Victoria Land (NVL) cratonward-directed subduction resulted in the accretion of the exotic Bowers and Robertson Bay Terranes to the inboard Wilson Terrane, which then represented the early Palaeozoic active margin of the East Antarctic Craton (Bradshaw & Laird, 1983; Kleinschmidt & Tessensohn, 1987; Kleinschmidt et al., 1992; Matzer, 1995; for an updated discussion of the tectonic architecture of the NVL, see Roland et al., 2004; Tessensohn & Henjes-Kunst, 2005). In the Wilson Terrane these accretionary processes generated a wide fold–thrust metamorphic belt and widespread syn-, late- and post-tectonic plutonism between about 530 and 480 Ma (Armienti et al., 1990; Tonarini & Rocchi, 1994, and references therein) represented by the Granite Harbour Intrusives (Gunn & Warren, 1962; Borg et al., 1987; Vetter & Tessensohn, 1987; Armienti et al., 1990; Fenn, 1993; Schüssler et al., 1993; Olesch et al., 1996). The intrusive rocks include both I-type metaluminous and S-type peraluminous calc-alkaline magmas, with variable K enrichment and magmatic arc affinity (Borg et al., 1987; Ghezzo et al., 1987; Vetter & Tessensohn, 1987; Armienti et al., 1990; Biagini et al., 1991a, 1991b; Di Vincenzo & Rocchi, 1999; Dallai et al., 2003; Rocchi et al., 2004). The metamorphic rocks intruded by the GHI can be subdivided into an eastern medium- to high-pressure belt including ultrahigh-pressure eclogitic lenses, and a western low- to medium-pressure–high-temperature belt that is locally in granulite facies (Grew et al., 1984; Talarico & Castelli, 1995; Di Vincenzo et al., 1997; Schüssler et al., 2004; Talarico et al., 2004, and references therein).
In the Deep Freeze Range (Fig. 1), an intermediate belt of the Wilson Terrane, a complex sequence of metaluminous and peraluminous granitoids and mafic intrusive rocks is well exposed and here defined as the Deep Freeze Range Intrusive Complex (DFRIC). Magmatic rocks include plutons and satellite plugs of variable size and degree of foliation, and dyke swarms. Mapping in the Mt. Levick–Browning Pass region has identified four distinct intrusive units: (1) foliated coarse-grained inequigranular granitoids (the Howard Peaks Pluton); (2) a foliated medium-grained gabbro–tonalite–granodiorite sequence (the ‘Corner Tonalite’, Adamson, 1971); (3) unfoliated biotite monzogranites (the Keinath Pluton, Skinner, 1983) and granodiorites; (4) a widespread complex of fine-grained felsic and mafic dykes and of aplite–pegmatites that crosscut the whole basement. The metamorphic country rocks have a complex history characterized by a low- to high-grade metamorphism (with remnants of a granulite-facies complex, Talarico & Castelli, 1995) of low-pressure type with contrasting P–T–t paths (Palmeri, 1997; Talarico et al., 2004, and references therein).
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For this study nine samples from the Howard Peaks Pluton, the ‘Corner Tonalite’ unit, the Keinath Pluton and the unfoliated mafic dyke swarm have been selected.
The Howard Peaks Pluton (HPP; Fig. 1) is the largest intrusion in the region and is a composite pluton. It occurs in scattered outcrops over an area of about 200 km2 between the eastern side of the Priestley Glacier and the Campbell Glacier, in and south of the Tourmaline Plateau. The HPP is composed of dominant coarse-grained porphyritic biotite monzogranites (Fig. 2a) with 5–8 cm K-feldspar megacrysts and rare mafic microgranular enclaves, and minor granodiorites. The HPP corresponds to the ‘Dickason Granite’ of Adamson (1971), but Ghezzo et al. (1987) referred to these rocks as the HPP because Mt. Dickason comprises unrelated migmatitic rocks.
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Intrusive contacts with the medium- to high-grade metamorphic country rocks are sharp and often subvertical. A marked magmatic foliation that dips 60°–80° toward the SW, parallel to the NW–SE-striking regional S2 schistosity, is developed throughout the HPP. This foliation is defined by the preferred orientation of magmatic minerals. A high- to moderate-temperature solid-state ductile deformation is common and often produces an augen-gneiss fabric superimposed on the magmatic fabric; this fabric is more pronounced along discrete shear bands that strike NW–SE. All structural criteria such as deformation and rotation of the feldspar megacrysts (Passchier & Trouw, 2005
) are consistent with a dextral sense of shear.
Preliminary chemical and isotopic analyses (unpublished data) indicate that HPP monzogranites are metaluminous, with SiO2 and K2O contents in the range 68–74% and 5–6%, respectively, high values of large ion lithophile element to high field strength element ratio (LILE/HFSE), a negative Nb–Ta anomaly and Rb/Sr ratios around two. High 87Sr/86Sr ratios at 500 Ma, in the range 0·710–0·716, suggest a crustal origin. The HPP monzogranites share many characteristics with the Abbott Granite in the Northern Foothills (Skinner, 1983; Armienti et al., 1990; Di Vincenzo & Rocchi, 1999), an intrusive unit with a high-K shoshonitic affinity, and with the DV2 monzonitic suite from the Dry Valleys region in southern Victoria Land (Allibone et al., 1993a, 1993b; Cox et al., 2000).
The ‘Corner Tonalite’ (CT; Fig. 1) is a mafic complex mainly cropping out along Black Ridge and the eastern side of the Priestley Glacier. Dominant tonalites [‘Corner Tonalite’ of Adamson (1971)] and quartz diorites are accompanied by minor gabbro and ultramafic rocks. The relatively mafic rocks are locally interlayered with sheets of granodiorite and monzogranite near the margins of the HPP. The rocks are subalkaline with a dominant high-K calc-alkaline affinity. All these intrusives commonly display a steeply dipping NW–SE-striking foliation parallel to the metamorphic foliation in the adjacent country rocks. The foliation is generally magmatic in origin, but along discrete shear bands is commonly overprinted by a solid-state deformation fabric. Where in contact with the Howard Peaks Pluton, the CT shows field evidence of mixing processes that gave rise to a heterogeneous granodioritic facies with large rounded K-feldspar megacrysts.
The Keinath Pluton (KP) is a large (about 50 km2) relatively homogeneous pluton composed of massive equigranular medium-grained biotite monzogranites with abundant mafic microgranular enclaves. Other smaller bodies of similar granitoids crop out in the Cape Sastrugi–Boomerang Glacier area. Petrographic and chemical data (Armienti et al., 1990, and unpublished data) indicate a metaluminous high-K calc-alkaline character similar to the dominant I-type suite in the Wilson Terrane (Borg et al., 1987; Armienti et al., 1990) and to the DV1a suite from the Dry Valleys region (Allibone et al., 1993b; Cox et al., 2000).
The dyke complex is a composite and not cogenetic complex.
Felsic dykes are fairly common, particularly in the medium- to high-grade and granulite-facies metamorphic rocks. Both foliated and unfoliated dykes are present (Musumeci & Pertusati, 2000).
Foliated felsic dykes are mainly represented by peraluminous two-mica ± garnet ± andalusite ± cordierite leucogranites and monzogranites showing concordant (sheet intrusions) or discordant contacts with the metamorphic rocks, sometimes folded, and commonly emplaced along the S2 foliation (‘Synkinematic Peraluminous Leucogranites’, Biagini et al., 1991a). Foliation is related to moderate-temperature solid-state deformation (Musumeci & Pertusati, 2000).
Unfoliated felsic dykes and stocks are common in the metamorphic rocks, typically as satellite plugs of the major plutons. They are equigranular or sometimes porphyritic, fine- to medium-grained metaluminous biotite monzogranites and granodiorites. Undeformed garnet–tourmaline-bearing aplite–pegmatitic veins and dykes crosscut all the basement rocks; they are particularly common in the granulite-facies rocks.
Unfoliated mafic dykes appear as phenocryst-poor, fine-grained massive dark rocks, lacking evidence of deformation. They are mainly microdiorites, quartz microdiorites and microtonalites. The geochemistry of these mafic rocks is characterized by a dominant high-K shoshonitic affinity and trace-element patterns similar to those of the mafic facies of the Vegetation Unit in the Terra Nova Intrusive Complex (Rocchi et al., 2004). All these fine-grained unfoliated rocks show textures that suggest a shallow tensional emplacement regime.
The emplacement of the intrusive sequence in the Deep Freeze Range has been described by Musumeci & Pertusati (2000): a mainly synkinematic emplacement is proposed for most of the intrusive bodies. They were intruded into low- to high-grade metamorphic rocks exploiting orogen-parallel thrust faults and transpressional structures controlled by two major ductile shear zones that run along the Priestley–Boomerang Glaciers and the Campbell Glacier [see also Läufer & Rossetti (2003) and references therein]. These shear zones have been reactivated in Cenozoic time (the ‘Priestley Fault’ and ‘Campbell Fault’) by a regional strike-slip fault system (Salvini et al., 1997; Storti et al., 2006). The Howard Peaks Pluton was emplaced in a push-up zone, as indicated by the sigmoidal pattern of magmatic foliation and the asymmetric shape with tails parallel to the shear zones (Musumeci & Pertusati, 2000). Some bodies of the ‘Corner Tonalite’ unit and many dykes, particularly some peraluminous felsic leucogranites, were also involved in the shear deformation. The emplacement depth of the intrusive sequence is poorly defined: preliminary data from the Al-in-hornblende geobarometer (Anderson & Smith, 1995) suggest an emplacement pressure of about 4 GPa for the CT. Furthermore, the timing of the magmatic and structural events detailed above is poorly constrained.
Previous studies indicated an age of 481 ± 10 Ma (Rb/Sr whole-rock isochron) for peraluminous synkinematic dykes in the Deep Freeze Range (Tonarini & Rocchi, 1994). Rb/Sr whole-rock–biotite ages, and K/Ar ages on biotite (Borsi et al., 1987; Kreuzer et al., 1987; Vita et al., 1991) on samples from the Howard Peak Pluton, the ‘Corner Tonalite’ and some metamorphic rocks from Mt. Emison in the Deep Freeze Range (including a granulite), indicate cooling ages in the range 475–449 Ma.
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ANALYTICAL METHODS |
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To collect a representative zircon population, about 1–2 kg of each sample was processed using a disc mill, shaking table, Frantz magnetic separator and heavy liquids. Representative crystals (about 100 grains) of different morphology and colour were then hand-picked from the final mineral separate, mounted in 25 mm epoxy discs and polished. A CAMEBAX SX50 electron microprobe at the ARC National Key Centre for the Geochemical Evolution and Metallogeny of Continents (GEMOC) was used to analyse the major constituents of zircons (Zr and Si) and selected minor elements (Hf, Y, U and Th), as well to image the internal zoning pattern of the crystal sections using a backscattered electron (BSE) detector. The BSE detector is a light-sensitive diode; when operated at appropriate conditions, it produces images of zircon texture that reflect both variations in the mean atomic number and the cathodoluminescence (BSE/CL images). Operating conditions for the EMP were 15 kV accelerating voltage and a beam current of 20 nA. The sample mounts were cleaned, together with zircon standards, in 2N nitric acid for c. 1 h to remove surface Pb contamination. Zircons were analysed for U, Th and Pb isotopic compositions using two LAMs (266 and 213 nm) attached to an Agilent 4500 series 300 ICPMS system at GEMOC. Samples and standards were ablated in a custom-built sample chamber using He as carrier gas. The ablation in He, rather than Ar, minimizes deposition of ablation products around ablation sites, improves sample transport efficiency, and gives more stable signals and more reproducible U–Pb fractionation (Jackson et al., 2004
). All analyses were performed with the laser focused above the sample (typically 150–250 µm) to minimize dynamic U–Pb fractionation as the laser beam penetrates into the sample (Hirata & Nesbitt, 1995). Constant LAM operating conditions were maintained throughout each run of 20 analyses to ensure constant U–Pb fractionation. Samples and standards were analysed in runs of 20 analyses, including 12 unknowns and eight analyses of the GEMOC/GJ-1 zircon standard (609 Ma). The unknowns included two near-concordant standard zircons, 91500 and Mud Tank, which were analysed in each run as an independent control on reproducibility and instrument stability. Each analysis lasted about 180 s, with gas background measurements being taken over the first 60 s, prior to initiation of ablation. The signals of masses 206Pb, 207Pb, 208Pb, 232Th and 238U were acquired using the instrument's time-resolved analysis data acquisition software. 204Pb was not measured, because of the low count rates and the isobaric interference from Hg that is a contaminant of the Ar supply gas. 235U was calculated from the ratio 235U = 238U/137·88. Raw data were processed using GLITTER, an in-house data reduction program (van Achterbergh et al., 2001). Identical integrated ablation time segments for each sample and standard zircon were selected, thus correcting for the effects of laser-induced fractionation. The contents of common lead were estimated from the three-dimensional discordance pattern in 206Pb/238U–207Pb/235U–208Pb/232Th space and corrected accordingly (Andersen, 2002). Concordia ages were determined using Isoplot 2.32 (Ludwig, 1999). Further details of the analytical technique, instrumental conditions and measurement procedures have been given by Belousova et al. (2001) and Jackson et al. (2004). The precision (1
) of the method (using a spot size of 40 µm and a pulse energy of 0·25 mJ) evaluated on the GJ-1 zircon is close to 1% for the 206Pb/238U and 207Pb/235U ratios. Some age measurements were also performed using a 213 nm laser ablation microprobe, coupled to a magnetic sector ICP-MS at CNR–Istituto di Geoscienze e Georisorse of Pavia, Italy. Ablation of samples and standards was performed in an in-house built ablation cell, using He as the carrier gas and 91500 zircon (1065 Ma) as external standard. The signals of masses 206Pb, 207Pb, 235U and 238U were acquired in magnetic scan mode (B-scan) at a spatial resolution of 40 µm. Accuracy ranges between 1% and 5% and is strictly related to Pb content in zircon. Details of the analytical technique, experimental conditions, data accuracy and precision have been given by Tiepolo (2003).
Trace-element concentrations were obtained on zircons from seven of the selected samples using the LA-ICPMS system at CNR–Pavia. NIST-610 was used as an external standard to correct for mass bias and laser-induced elemental fractionation. The data were normalized for each analysis to the electron microprobe Zr concentration as an internal standard. Precision and accuracy are generally better than 10%, and limits of detection are down to the ppb level for the heaviest elements. A detailed description of the method has been given by Tiepolo et al. (2002).
Hf isotope analyses for the same set of zircons were carried out in situ using two New Wave/Merchantek laser ablation microprobes (193 nm, 213 nm), attached to a Nu Plasma MC-ICPMS system at Macquarie University, Sydney. Operating conditions were 10 Hz repetition rate and energies of 0·6–1·3 mJ/pulse for the 213 nm laser, and 2 Hz repetition rate and energies of 0·03–0·05 mJ/pulse for the Excimer (193 nm) laser. Typical ablation times were 80–120 s, resulting in pits of 40–60 µm and 30–40 µm deep with the 213 and 193 nm lasers, respectively. The methods and analyses of standard solutions and standard zircons have been described by Griffin et al. (2000). Helium was used as carrier gas to transport the ablated sample from the laser-ablation cell to the ICPMS torch, via a mixing chamber. The masses analysed for this work are 172, 175, 176, 177, 178, 179 and 180. Analyses were carried out in static-collection mode and all masses acquired simultaneously in Faraday cups. Data were normalized to 179Hf/177Hf = 0·7325, using an exponential correction for mass bias. Interference of 176Lu on 176Hf was corrected by measuring the intensity of the interference-free 175Lu isotope and using the recommended 176Lu/175Lu = 0·02669 to calculate 176Lu/177Hf. Interference of 176Yb on 176Hf was corrected by measuring the interference-free 172Yb isotope and using 176Yb/172Yb to calculate 176Yb/177Hf. The appropriate value of 176Yb/172Yb = 0·5865 was determined by successively spiking the JMC475 Hf standard (1 ppm solution) with Yb, and iteratively finding the value of 176Yb/172Yb required to yield the value of 176Hf/177Hf obtained on the pure Hf solution. The measured 176Lu/177Hf ratios were then used to calculate initial 176Hf/177Hf ratios. The typical uncertainty on a single analysis of 176Hf/177Hf is ±0·000020, equivalent to an uncertainty of less than one epsilon unit. Mud Tank and 91500 standards were used to check the accuracy and precision of the laser-ablation results within each run.
All analyses were performed directing the laser to specific textural domains and avoiding mixed textural domains and mineral inclusions in the zircon grains.
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SAMPLE CHARACTERISTICS |
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Zircon crystals were analysed from four samples (M14, C1, C10, DF6) selected to be representative of the major foliated porphyritic biotite monzogranitic facies, one sample (C4) of the foliated porphyritic biotite granodiorite from Mills Peak, all from the Howard Peaks Pluton; two samples are a foliated quartz diorite and a tonalite (B2 and MS1) from the ‘Corner Tonalite’ unit; one sample (L1) is a massive biotite monzogranite from the Keinath Pluton; and one is a pyroxene micro-monzodiorite (DR12) from the unfoliated mafic dyke swarm. Locations are given in Fig. 1.
The nine studied samples show different fabrics: M14, C1, C10, DF6, C4 and MS1 show both magmatic foliation and solid-state ductile deformation. Deformation is more pronounced in DF6 and C4, and moderate in MS1. Sample B2 is characterized only by a magmatic foliation. Samples L1 and DR12 are massive. Sample codes, localities and latitude and longitude are given in the Appendix.
Howard Peaks Pluton
Samples M14, C1 and C10 are typical coarse-grained, foliated, megacrystic, biotite monzogranites characterized by the following main mineral phases: variable amounts of perthitic microcline (mainly as 2–8 cm megacrysts), zoned plagioclase (28–42% An), quartz, biotite (XMg = 0·35–0·45) (Fig. 2b); apatite, allanite, ilmenite, monazite, rutile and zircon are common accessory phases. Sample DF6 is a foliated porphyritic monzogranite from the eastern border of the Howard Peaks Pluton. It is petrographically similar to the main monzogranites, but it is a relatively low-silica facies (around 70% SiO2) and contains small amounts of orthopyroxene and slightly higher amounts of biotite. In these rocks deformation and partial recrystallization of the magmatic minerals has produced elongate fine-grained stringers of quartz, plagioclase and biotite that wrap around K-feldspar and plagioclase grains, myrmekite along the margins of igneous K-feldspar, ductile deformation of quartz to form ribbons or aggregates of new grains, and bending of plagioclase crystals, all features consistent with a high- to moderate-temperature solid-state deformation (Paterson et al., 1989; Tribe & D’Lemos, 1996). The common occurrence of ‘chessboard’ subgrain patterns in quartz confirms the high P–T conditions of deformation (Kruhl, 1996). Trails of fluid inclusions, with a variable composition (often CO2-rich), are common in quartz.
Sample C4 is a foliated porphyritic biotite granodiorite with plagioclase and K-feldspar megacrysts, from a stock intruded in the granulite-facies metamorphic rocks at Mills Peaks. In this sample a postkinematic annealing event is overprinted on the gneissic fabric and is characterized by a fine-grained polycrystalline aggregate, lacking evidence of strain and interposed between the feldspar megacrysts. The sample from this outcrop has been included in the HPP, because of its striking fabric similarity to the typical megacrystic monzogranites of the HPP. However, preliminary chemical and isotopic data suggest that it probably belongs to a local, not cogenetic intrusion, coeval with the main HPP.
‘Corner Tonalite’
Samples MS1 and B2 are two foliated hornblende–biotite-bearing plutonic rocks (a tonalite and a quartz diorite) outcropping at the western side of the Howard Peaks intrusion and at Black Ridge, respectively. MS1 is characterized by a pervasive overprinted solid-state deformation; B2 shows no evidence of post-crystallization deformation. They are medium-grained rocks with a colour index of about 30, with abundant green hornblende and minor brown biotite. Plagioclase is strongly zoned and ranges in composition from andesine to labradorite; quartz is interstitial; accessory minerals are apatite, sphene, magnetite and zircon. A few trails of small fluid inclusions occur mainly in sample MS1.
Keinath Pluton
Sample L1 is a massive equigranular medium-grained biotite monzogranite. Plagioclase is a strongly zoned andesine, K-feldspar is a perthitic microcline; accessory minerals are apatite, sphene, magnetite and zircon. The presence of minor epidote and chlorite suggests a moderate subsolidus alteration; a few fluid inclusion trails occur locally in quartz.
Mafic dyke swarm
Sample DR12 is a fine-grained undeformed massive phenocryst-poor (plagioclase) micro-monzodiorite dyke from the western side of the Boomerang Glacier, intruded within strongly foliated and deformed granitoids. It is characterized by dominant euhedral zoned plagioclase (33–48% An), quartz and rare interstitial orthoclase; the mafic phases are biotite, orthopyroxene and clinopyroxene. Apatite is abundant as needles; sphene, zircon and ilmenite are common accessory phases.
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ZIRCON TEXTURES AND U–Pb RESULTS |
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Howard Peaks Pluton
Porphyritic biotite monzogranites (samples M14, C1, C10)
Abundant zircon grains are distributed throughout these samples and commonly are included in the main mineral phases (quartz, feldspar and biotite); they also occur in recrystallized areas. They range in length between about 400 and 50 µm and show variable amounts of apatite and fluid inclusions.
BSE/CL images (Fig. 3a–j) reveal a wide spectrum of internal textures in zircon, not always readily interpretable as primary. Most of the investigated crystals show oscillatory zoning defined by variable degrees of thickness and brightness. Three main groups of textures have been distinguished, which characterize individual domains within zircon grains or the whole grains: fine-scale oscillatory zoning, thick zoning and complex zoning. Fine-scale oscillatory zoning is characterized by thin zones showing low contrast in brightness; thick zoning is defined by thick bright bands alternating with darker zones; complex zoning includes a variety of textures and covers the majority of the grains. Complex textures include alternating convoluted bands distinguished by different BSE/CL intensity, which sometimes preserve ghosts of oscillatory zoning, and unzoned patches with or without small bright patches in BSE/CL images.
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U–Pb results are shown in Fig. 4a–f. Selected U–Pb data are listed in Table 1. The complete dataset is provided as Electronic Appendix 1, available for downloading from http://www.petrology.oupjournals.org. The concordia and weighted average ages are shown at the 95% confidence level. To construct the weighted average age and the probability density plots concordant 207Pb/206Pb ages as calculated by Isoplot 2.32 (probability of concordance >0·001) were used (Ludwig, 1999
).
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In all the analysed samples there is a relationship between texture and the degree of concordance of the U–Pb ages: the domains characterized by thick zoning often show discordant ages.
Inherited components are also often present. They occur mainly as relict cores with different textures, size and morphology; these are not shown in the U–Pb diagrams nor discussed in this paper.
Sample M14 is characterized by a wide scatter of concordant U–Pb ages, between 522 and 435 Ma (Fig. 4a and Table 1). The probability density plot (Fig. 4b) of the concordant ages shows many peaks without clear gaps between them. The spread in concordant ages may be related to a relatively short time interval between the magmatic intrusion and a discrete resetting event. Therefore, the whole U–Pb dataset cannot be here used to infer the emplacement age of the granite; most of the young ages do not reflect the primary magmatic age but represent varying degrees of incomplete or complete resetting of the primary undisturbed igneous zircons. The oldest peak in the probability density plot (Fig. 4b) therefore is used to infer the magmatic age: nine analyses with a mean age of 506 ± 7 Ma (MSWD = 2); the 20 younger ages ranging between 488 and 435 Ma represent reset ages.
In sample C1, U–Pb ages range from 508 to 472 Ma (Table 1, Fig. 4c). In the probability density plot (Fig. 4d), concordant ages define two main peaks. The oldest peak of 18 points with weighted average ages of 499 ± 2·4 Ma (MSWD = 1·3) represents the emplacement age of the sample. A younger peak of 12 points around 484 Ma (Fig. 4d) and a single spot at 472 Ma reflect Pb loss. A relationship between age populations and textural domains is not obvious; the older ages were measured on differently zoned domains (oscillatory and complex zoning).
In sample C10 (Fig. 4e and Table 1) many discordant data (11 analyses) are present, making it difficult to recognize a primary age or a subdivision into different age populations. The relatively large scatter in the Tera–Wasserburg diagram (Fig. 4e) suggests that the data are probably affected by both an uncorrected common lead component and Pb loss. The few concordant data (22) scatter between 505 and 455 Ma (Fig. 4e); the oldest 12 analyses define a peak around 493 ± 3·8 Ma at 95% confidence (MSWD = 1· 3) (Fig. 4f). The older ages are found in zircons showing different textural features: zoned and unzoned domains, and resorbed cores showing convoluted zones. Younger concordant ages between 481 and 455 Ma have been measured on unzoned or weakly zoned domains.
Porphyritic orthopyroxene-bearing monzogranite (sample DF6)
The large number (42 grains) of zircons selected from sample DF6 show grain sizes and morphological features similar to those described for the zircons extracted from the other granitic samples. Inclusions of apatite and fluids are common. As shown in the concordia plot (Fig. 5a), most data (43 of 53 analyses) plot along concordia, defining a wide range of ages between 501 and 448 Ma (Table 1). The probability density plot (Fig. 5b) for the concordant data alone shows three main peaks. The oldest peak defines a plateau around 498 ± 6·6 Ma (four analyses) (MSWD = 1· 3; Fig. 5b), but most of the analyses (39) show relatively ‘young’ ages. No systematic variations can be observed between zoned and unzoned or complexly zoned crystals, but most of the domains characterized by thick zoning and bright-BSE/CL patches show slightly discordant ages.
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Porphyritic biotite granodiorite (sample C4)
Zircons recovered from sample C4 show the same range of textural features and size displayed by the Howard Peaks samples, but with a predominance of domains characterized by complex zoning (Fig. 3j). Fluid inclusions are abundant and apatite inclusions are also present. Crystal shapes vary from elongated and prismatic to anhedral. Oscillatory zoning is common and can either include the entire zircon texture or occur in rims around unzoned or more complexly zoned cores. Bright domains are often present as cores or patches distributed throughout the crystal. Zircons showing oscillatory zoning often display bright thick zones. Finally, more complex features such as internal convoluted surfaces are widespread. Inherited cores, highly variable in size and shape, are also present. They are typically surrounded by weakly zoned or unzoned rims, separated by planar or rounded surfaces with bright bands sometimes present at the core–rim boundary. Textureless inherited grains are also present.
U–Pb results (Table 1, Fig. 5c and d) show a broad spectrum of zircon ages for this sample. No relationship exists between zircon age and morphology, and a subdivision into different age groups based on BSE/CL images is difficult. Concordant ages (Fig. 5c) range from 506 to 450 Ma, but a large number of slightly discordant ages around 470–460 Ma and three younger discordant analyses suggest that a strong Pb-loss event affected this sample. In a probability density plot and a weighted average plot (Fig. 5d) the concordant ages of this group appear to fall into two distinct populations: a minor group of three analyses that may reflect the magmatic age at 500 ± 5·4 Ma (MSWD = 0·87) and a wide cluster of younger reset ages (12 analyses) ranging from 482 to 450 Ma.
Keinath Pluton
Biotite monzogranite (sample L1)
Zircons in the biotite monzogranite L1 concentrate within the main mineral phases: K-feldspar, plagioclase and mainly biotite grains. They are generally euhedral short-prismatic to strongly elongated in shape, have sharp edges and vary in size between a few tens of microns and more than 300 µm. BSE/CL images of most grains show well-developed fine-scale oscillatory growth zoning that covers the whole grain or forms outer rims around homogeneous unzoned cores (Fig. 6a and b). In some zircons the BSE/CL images show bright thick zones, usually located in the inner rim (Fig. 6c). These bright rims often exhibit cracks. Inherited components are rare and usually occur as short-prismatic fractured grains containing bright metamict cores with typically spongy textures. More rarely they form metamict cores inside long-prismatic grains characterized by oscillatory zoning.
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The U–Pb data (Table 1) for the analysed zircons are given in Table 1 and Fig. 5e. They show a range of concordant ages between 495 and 464 Ma. The weighted average age (Fig. 5f) for the main group of U–Pb data (23 analyses) is 481 ± 2·3 Ma (MSWD = 1· 6), interpreted as the emplacement age of the Keinath Pluton. Thick bright zones show a range of ages similar to those obtained for the fine-scale oscillatory zones, except for a single slightly discordant age around 340 Ma. Six younger ages ranging from 470 to 340 Ma reflect lead loss and are not included in the calculation of the magmatic age.
‘Corner Tonalite’
Foliated hornblende–biotite quartz diorite and tonalite (samples B2 and MS1)
Zircons from the B2 quartz diorite are all textureless (Fig. 7e and f). In tonalite sample MS1 zircons are usually unzoned or faintly zoned (Fig. 7a–d). Rounded bright patches occur within some crystals and bright irregular bands are present in a few elongated zoned crystals in sample MS1. Grain shapes vary from euhedral to subhedral. Textures resembling relict cores are absent, and apatite inclusions are common in both samples. In general, textures in the diorite–tonalite zircons are much less complex than those in the granodioritic and granitic zircons (Fig. 3a–j).
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As in the granitic samples, the probability density diagram for sample MS1 (Fig. 8a) shows two populations: an older group of 17 analyses with a weighted average age of 508 ± 3·1 Ma (MSWD = 0·92), and a group of younger ages (11 analyses) that ranges from 489 to 457 Ma (Fig. 8b). The differently zoned domains show the same range in U–Pb ages (Table 1). Sample B2 shows a more restricted range of ages (Fig. 8c) than MS1. The probability density plot (Fig. 8d) exhibits a single population of 25 analyses with a weighted mean age of 497 ± 2·6 Ma (MSWD = 1· 18) if a younger spot related to ancient Pb loss is rejected.
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Dyke swarm
Micro-monzodiorite (sample DR 12)
Zircons from sample DR12 are typically subhedral to euhedral, homogeneous to rarely zoned (oscillatory zoning), with lengths in the range 200–50 µm (Fig. 7g and h). Apatite inclusions are common. Fifteen analyses (Table 1) yield a concordant age (Fig. 8e and f) of 489 ± 3·8 Ma (MSWD = 1· 7).
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ZIRCON TRACE-ELEMENT CHEMISTRY |
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A wide range of trace-element contents, including Y, Nb, Ba, Yb, Hf, Ta, Th, U and rare earth elements (REE), has been measured on zircons that were analysed for both U–Pb and Lu–Hf in seven of the analysed samples (C10, C1, DF6, L1, MS1, B2 and DR12) to relate the trace-element concentrations to the different textural and age domains in zircons. A detailed analysis of the zircon chemistry is required to differentiate magmatic from secondary domains in zircon and to constrain the processes involved in the zircon resetting. Ablation spots, about 20 µm in diameter, were sited near the spots used for U–Pb analyses and on similar morphological domains (Figs 3, 6 and 7).
Representative trace-element compositions are shown in Table 2. (For the complete dataset refer to Electronic Appendix 2, available for downloading from http://www.petrology.oupjournals.org.) A large scatter in trace-element concentrations is found in the zircons from HPP granitoid samples C1, C10 and DF6, with variations over several orders of magnitude in light REE (LREE). There is only limited variability, however, in heavy REE (HREE) concentrations. The elemental abundances (Table 2) range over almost the same order of magnitude in the three samples. A wide variation in Y, U, Th and Hf contents is observed, with ranges between a few tens and thousands of ppm. Sample C1 shows the highest mean values of Y, Hf, 
REE and LREE, and also the highest variability in LREE enrichment. A correlation can be recognized between trace-element variations and textural domains in zircons, with some differences between the three samples. In samples C1 and DF6 the domains characterized by thick zoning show the highest mean values (Table 3) for many of the measured elements. Sample C10 shows a different behaviour; the domains covered by fine-scale zoning have higher mean values for many elements, but again higher mean values of LREE occur in the domains with thick and complex zoning (Table 3).
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The Th/U ratios range from 0·05 to 3 with a strong predominance of values higher than 0·1, typical of igneous zircons; the few low ratios (< 0· 1) were measured in domains with thick zoning in zircons from samples C10 and DF6 (Table 3). However, no clear differences can be outlined for the distinct age populations.
The REE patterns are strongly different in the three textural groups (Fig. 9a–c): the domains showing fine-scale oscillatory zoning typically have low abundances of the LREE and strongly fractionated patterns with well-defined positive Ce and negative Eu anomalies; domains characterized by thick zoning typically show enriched or flat LREE patterns with negative Eu anomalies but variable Ce anomalies; the REE abundances in the complex domains cover a wide range, overlapping with both the fine-scale and thick zoning.
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The increase in the overall REE contents is related to a strong increase in the LREE contents. The LREE enrichment is always positively correlated to most of the trace elements (Y, Th, U, Yb, Nb and Ta in samples C1 and DF6). Th/U ratios are not correlated to
LREE in the three samples.
The zircons of the monzogranite L1 show concentrations of most of the analysed elements (U, Th, Y, Yb, Nb, Ta, Hf, REE; Tables 2 and 3) in the same range as the HPP granitoids. However, the LREE in sample L1 show lower mean values and a distinctly lower variability. According to the BSE/CL images of the zircon textures (Fig. 6a–d), two of the three textural populations observed in the HPP granitoid have been recognized also in sample L1: fine-scale zoning, which constitutes the majority of zircon grains, and thick zoning. Complex domains have not been observed. Fine-scale oscillatory zoning is characterized by LREE-depleted patterns with typical positive Ce and negative Eu anomalies (Fig. 10). The two single analyses of domains with thick zoning show LREE enrichment and higher mean contents of most of the elements (U, Th, Y, Yb, Nb, Ta, Hf, LREE). One small grain, characterized by faint zoning and by a distinctly younger age, is characterized by lower contents of many elements (Y, Nb, Ba, Yb, REE and LREE). The Th/U ratio is in all textural groups always higher than 0·1 and between 0·3 and 0·8.
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The zircons of the amphibole-bearing quartz diorite–tonalite samples (B2 and MS1) show a more restricted variability for most of the analysed elements (U, Th, Y, Nb, Ta; Tables 2 and 3). Mean values of U, Th, Hf, Y, Yb,
REE and LREE in sample MS1 are lower than in sample B2, and Th, U and LREE show a larger variability (Table 3). The REE patterns show LREE depletion, with typical positive Ce and negative Eu anomalies (Fig. 11a and b) but are slightly less steep and more scattered in sample MS1, with larger variations in the LREE. The Th/U ratio is between 0·21 and 0·82.
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The lack of visible textural domains in zircons from sample B2, and the low BSE/CL response of the MS1 zircons, hinders any attempt to relate morphological domains with microchemistry. With regard to the relationship between ‘age populations’ and trace-element variations, different trends are observed in the two samples: a limited LREE variability in sample B2 zircons is related to a restricted range of U–Pb ages, whereas a larger variability in
LREE and other measured elements and lower mean abundances in sample MS1 are connected to a larger spread in the concordant U–Pb ages.
Zircons in the micro-monzodiorite sample (DR12) have trace-element contents and REE patterns similar to those for zircons from sample B2, with slightly higher contents of U, Th, Y, Ta, REE, LREE and Th/U, and the same narrow range of both LREE and HREE as in sample B2 (Tables 2 and 3; Fig. 11c). In sample DR12, as observed for the quartz diorite B2 and partially for the Keinath biotite monzogranite L1, the high proportion of simple zircon textures and the limited variations of the REE are reflected in a tight clustering of the U–Pb data.
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Lu–Hf ISOTOPE DATA |
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Zircons analysed for U–Pb were also analysed for their Hf isotopic composition. The Hf isotopic composition of zircons provides valuable information on the source and age of the protolith of the magma, and is very useful in discriminating between magmatic and inherited components. Ablation spots for the Hf analyses were sited near the spots used for U–Pb isotope analyses (Figs 3, 6 and 7). Table 4 lists the analytical data, as well as initial
Hf values for each grain. Lu concentrations in zircons are very low, making corrections for in situ decay of 176Lu very small. Therefore initial 176Hf/177Hf values are essentially the same as measured 176Hf/177Hf ratios. We report the measured values in Table 4 and Fig. 12. A large number of analyses have been obtained for eight of the nine analysed samples (C10, M14, DF6, C4, MS1, B2, L1 and DR12), but here only values obtained on zircon domains with concordant ages are shown, to clearly relate zircon age and textures to Hf-isotope signature.
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The spread of U–Pb ages seen in most of the analysed rocks is not reflected in the Hf isotope data. The different ‘zircon age populations’ observed in the granitoid and dioritic–tonalitic samples yield Hf-isotope signatures that are indistinguishable from each other, whereas significant differences are found for inherited cores (unpublished data). No correlation is evident between 176Hf/177Hf ratios and textural domains in zircons.
The Howard Peaks granitoid and the quartz diorite and tonalite from the ‘Corner Tonalite’ show the same range of initial 176Hf/177Hf and Hf (Fig. 12 and Table 4). The average 176Hf/177Hf is about 0·2823, and Hf values are in the range between –2·1 and –10·3 in the granitoids. The TDM model ages range between 1·1 and 1·5 Ga, with a mean value of 1·3 Ga. The TDMcrustal model age calculated for the crustal protolith (assuming 176Lu/177Hf = 0·015, equivalent to the average continental crust; Griffin et al., 2000) ranges from 1· 5 to 2·1 Ga, with a mean value of 1· 8 Ga. The Corner quartz diorite and tonalite (B2 and MS1) show average 176Hf/177Hf ratios similar to the granitoids (0·2823) and the same range of TDM and TDMcrustal model ages.
Zircons from the micro-monzodiorite sample DR12 have higher 176Hf/177Hf ratios than zircons from HPP granitoids and CT quartz diorite–tonalite samples, with values ranging from 0·2825 to 0·2824 and Hf between –0·3 and 2·1.
The Keinath monzogranite is characterized by distinctly higher average initial 176Hf/177Hf relative to the HPP granitoids, around 0·2825. Such values correspond to Hf values between –2·69 and 5·96 and a mean Hf of 0·07. The young zircon population has 176Hf/177Hf ratios in the same interval. The TDM model ages range between 0·87 and 1·17 Ga, with a mean value of 1· 08 Ga.
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DISCUSSION |
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The data presented here clearly indicate that in the studied granitic intrusives there is not a simple correlation between isotope data and emplacement age: the wide scatter of U–Pb data along concordia within each sample complicates the inference of the emplacement age and requires a careful evaluation of the data to discriminate possible magmatic ages from overprinted subsolidus resetting. Such discrimination can be made only if the discussion of the U–Pb data is supported by a detailed investigation and comparison of the zircon textures, as previously outlined for the analysed samples, and their trace-element and isotopic composition.
Zircon textures
Zircons from the micro-monzodiorite dyke and the CT quartz diorite–tonalite samples display a simple pattern of textures when compared with zircons from HPP granitoid samples, ranging from unzoned and homogeneous in sample B2 to unzoned to weakly zoned in samples DR12 and MS1. The following discussion will be, therefore, mainly focused on the description of the complex textures in zircons recovered from the HPP granitoid samples.
Zircons from the Keinath L1 biotite monzogranite show a predominance of fine-scale oscillatory zoning and only minor thick bright zones, similar to the thick zoning of the HPP zircons.
LA-ICPMS isotopic ages and BSE/CL imaging allow us to evaluate the roles played by different processes in each textural population of the HPP zircons, separately.
Euhedral oscillatory zoning (fine-scale and thick zoning)
The euhedrally zoned zircons show both thick and fine-scale zoning (Fig. 3a–d). Grains showing thick zoning (Fig. 3d) are characterized by the occurrence of bright zones, related to different U and Th contents in sample DF6 and U, Y, Ta and LREE in sample C10; grains with fine-scale zoning lack very bright zones but can show different degrees of oscillatory zoning. Within grains characterized by thick zoning, radial micro-fractures running perpendicular to the bright zones or fractures concentrically developed along the boundary between the bright and dark domains can occur. Such features are typical of variously metamictized zircons (Corfu et al., 2003, and references therein; Geisler et al., 2003). Domains covered by thick zoning often exhibit discordant ages; domains with fine-scale oscillatory zoning commonly show more concordant ages over a wide interval.
Prismatic forms and euhedral concentric zonation commonly are interpreted as indicators of primary magmatic textures reflecting crystallization from melts (Vavra, 1990; Paterson et al., 1992; Hanchar & Miller, 1993). However, in our opinion this simple interpretation of internal texture is ambiguous. Recent papers have demonstrated that apparently undamaged and euhedral crystals may represent recrystallized zircon. Different types of euhedrally zoned crystals have been recognized by Black et al. (1986) in zircons from metamorphic rocks, where the strength of zoning reflects differently disturbed systems, characterized by different U and/or Th contents. Strongly zoned grains are usually characterized by more discordant data and more pronounced Pb-loss events. Pidgeon et al. (1998) explained the thickening of the primary zones and the formation of trace-element-rich bands in zircons from Archaean granites as due to the progressive migration and reconcentration of trace elements within the zircon crystals. The redistribution of trace elements within zircon grains can provide, according to Pidgeon et al., a mechanism for resetting zircon U–Pb ages independent of radiation damage of the zircon lattice, giving discordant ages. The suggested mechanism for producing such textures could be the instability of zoned zircon caused by the high trace-element contents, which produces a distortion of the lattice (Köppel & Sommerauer, 1974; Sommerauer, 1976). Similar conclusions were reached by Vavra et al. (1999) for the strongly luminescent domains related to primary structures in zircons but, in addition to thermal events, those workers attributed an important role in producing zoning-controlled alteration also to decompressional events.
A different explanation (Hartmann et al., 2000) attributes paired dark–bright bands to diffusion promoted by the interaction of zircons with fluids released during metamorphic events. Such fluid-aided recrystallization could not only produce enhancement of magmatic zones by infiltration and diffusion of elements along euhedral zones, but also produce euhedral zones in originally unzoned or weakly zoned crystals.
It is clear, from this discussion, that whatever the processes that produced zircon recrystallization (solid-state or fluid-aided), euhedral zoning in crystals cannot be always interpreted as a primary texture. Dating of these crystals can sometimes yield a recrystallization age. However, most of the above-mentioned works dealt with zircons from metamorphic rocks, in which the modification of the texture was related to recrystallization under wet or dry conditions during the metamorphism. Still poorly documented is euhedral oscillatory zoning of secondary origin in magmatic rocks (unaffected by metamorphic processes). In the light of this evidence, we interpret the strongly zoned crystals in the studied granitoids as open systems in which significant Pb loss occurred, probably in a multi-stage disturbance event: an initial high-temperature solid-state recrystallization during the slow cooling of the pluton and a progressive low-temperature (probably fluid-assisted) trace-element redistribution. Large BSE/CL bright young cores could represent strongly metamict zones surrounded by outer, unaltered, euhedral rims. Bright thick zones can also, sometimes, be partially metamict, as evidenced by the presence of cracks in BSE/CL images. Zircons with fine-scale oscillatory zoning could be interpreted, on the other hand, as being as old as the strongly zoned ones, but less affected by Pb loss: the oldest ages could represent the real magmatic event, whereas the younger ones, although concordant, could be due to variable Pb loss during a single long-lasting event, which started soon after the magma emplacement.
In conclusion, constraining the magmatic age from zircons with euhedral oscillatory zoning is not always easily done, even in non-metamorphic rocks. Only crystals lacking possible evidence of secondary trace-element redistribution, such as highly luminescent bands, should be considered for discussions of the magmatic age. However, resetting phenomena also may have affected grains covered by fine-scale zoning, making a clear distinction of magmatic from reset zircons difficult on the basis of BSE/CL imaging alone.
Complex zoning
The vast majority of the studied zircons are characterized by very complex textures that show little correlation with the age results. Contorted zones (Fig. 3e–g) or (Fig. e, f and g), characterized by different BSE/CL intensity, are widespread in all the HPP granitoid samples. The curved zones usually are concentrated in the centre of the crystals (Fig. 3f, g and i), but a detailed study of the images shows that they often form protuberances starting at embayments in the crystal surfaces and invading towards the centre, leaving euhedral oscillatory zones undisturbed on the remainder of the crystal edges. Most of the investigated crystals also show BSE/CL bright domains (Fig. 3i), some of which are related to primary features in the crystal (inclusions or almost completely obliterated growth zoning), and sometimes constitute isolated patches not related to crystal terminations or inclusions in the observed two-dimensional section. A few grains also show spongy textures. Furthermore, ghosts of finely banded oscillatory zones (Fig. 3f, g, i and j) can sometimes survive between convoluted bands and spongy domains. Trace-element compositions and REE patterns in the complex domains are widely variable (Tables 2 and 3; Fig. 9a–c), with contents often intermediate or sometimes overlapping both the fine-scale and thick oscillatory zoning.
Textures similar to those described in this study have often been reported, but interpreted in different ways. The curvilinear shape of growth banding has been compared to flow structures by Vavra et al. (1996). In addition, the advance of contorted zones from the surface of the crystals towards the centre has been considered as evidence of resorption, and of the infiltration of hydrous fluids during metamorphism or metasomatism. Similar conclusions were reached by Pidgeon et al. (1998) regarding curved and contorted zoning in zircons of Archaean granites from the Darling Range Batholith. On the other hand, Hoskin & Black (2000), on the basis of the occurrence of ghost zones of magmatic oscillatory zoning within convoluted domains, argued against fluid-assisted recrystallization and favoured solid-state recrystallization. They proposed that the role of inter-granular fluids is critical only in a kinetic sense, providing higher thermal conductivity than that of the surrounding minerals. However, all of these workers agreed on the secondary origin of this kind of texture, and the discussion concerns only the mechanism responsible for the isotopic disturbance.
The wide dispersion of the U–Pb data in the samples from this study can be explained in terms of variable degrees of resetting as seen in crystals characterized by fine-scale oscillatory zoning. It is more difficult to establish the role of solid-state or fluid-assisted recrystallization. The occurrence of ghost zones between convoluted or unzoned patches and of bright domains clearly related to primary features in the crystal seem to point to a solid-state recrystallization. On the other hand, the presence of curved zones, commonly starting at the crystal edge and advancing toward the core, can be explained more easily by interaction of the crystal surface with fluids.
LREE in zircon
A number of studies (Hinton & Upton, 1991; Maas et al., 1992; Barbey et al., 1995; Sutherland et al., 1998; Hoskin & Ireland, 2000; Hoskin et al., 2000; Belousova et al., 2002; Whitehouse & Kamber, 2002; Hoskin & Schaltegger, 2003) of the REE chemistry of zircon have shown that igneous zircons are characterized by steep chondrite-normalized REE patterns with a steep slope from La to Lu, and by prominent positive Ce and negative Eu anomalies. This pattern is observed for a large number of the zircons analysed in this study. However, for many of them, a scatter over three orders of magnitude in the LREE contents also has been observed, coupled with a limited variability in the HREE abundance. A xenotime-type substitution (REE + P = Zr + Si) has been invoked (Speer, 1982; Hinton & Upton, 1991; Hoskin et al., 2000; Finch et al., 2001) as the main mechanism for substitution of quadrivalent Zr by trivalent REE, the charge balance being achieved with the incorporation of P in the Si site. However, as outlined by many researchers (Speer, 1982; Hinton & Upton, 1991; Maas et al., 1992; Hoskin et al., 2000; Finch et al., 2001; Hanchar et al., 2001; Hoskin & Schaltegger, 2003), the atomic ratios of REE:P often deviate from unity, suggesting a more complex charge-balance mechanism. An excess of HREE in HREE-doped synthetic zircons has been explained by the presence of charge compensating species (Mo6+ and Li+) in octahedral and four-coordinated interstitial sites and substitution of Zr in the Si site (Finch et al., 2001; Hanchar et al., 2001). However, the mechanism involved in the incorporation of large amounts of LREE is still poorly understood. The larger ionic radii of LREE cause them to be incompatible in the zircon structure and incorporated only in small amounts; the incorporation of high contents into the triangular dodecahedral site introduces a strong distortion of the zircon lattice (Whitehouse & Kamber, 2002, and references therein). Hoskin & Black (2000) proposed a mechanism of cation expulsion and recrystallization fronts progressing within the zircon lattice during sub-solidus recrystallization under high-temperature conditions to explain the concentration of LREE behind the recrystallization front. In crystals subjected to similar processes, however, REE-enriched bands are usually associated with REE-depleted areas. In the analysed zircons this is true for some grains, but in many others the REE enrichment affects the whole grain. Simple volume-diffusion models can explain the mobility of HREE, but fail to explain LREE enrichments (Cherniak et al., 1997), because they produce higher HREE variations with respect to LREE.
A number of workers (Bea & Montero, 1999; Schaltegger et al., 1999; Rubatto, 2002; Hoskin & Schaltegger, 2003; Whitehouse & Platt, 2003) have recently related the REE patterns of zircons to the crystallization or breakdown of phases coexisting with zircon during metamorphic processes. However, although the behaviour of the HREE partitioned into zircon in equilibrium with garnet is relatively well understood, the behaviour of LREE and their relationship with LREE-rich phases such as monazite and allanite is still unknown. Fluids are important carriers of trace elements in metamorphic processes, but the mobility and immobility of REE during hydrothermal fluid–rock interaction is still a matter of debate. Rolland et al. (2003) described the behaviour of REE in the pre-Alpine granitic basement of Mont Blanc Massif during Alpine fluid-driven shear zone development. They suggested that the changes in REE concentrations can be mainly ascribed to the alteration of pre-existing magmatic REE-bearing minerals during deformation-related fluid–rock interaction and to the precipitation of new REE-bearing minerals. The stability of HREE- or LREE-bearing minerals is dependent on the composition of the fluids. In particular, ligands such as 
, 
and 
and pH are important in controlling the fractionation of LREE from HREE; high sulphur fugacity values at low pH and high contents produce LREE enrichments (Rolland et al., 2003). Geisler et al. (2003) described strong enrichment of LREE in metamict zircons from post-tectonic granites affected by strong hydrothermal alteration related to an intensive reaction with aqueous fluids under low-temperature conditions (120–200°C).
Alternatively, as pointed out by Whitehouse & Kamber (2002), LREE enrichment of zircon can also be an analytical artefact related to the sampling of LREE-enriched phosphate inclusions not visible on the zircon surface. However, the use of time-resolved analytical techniques, as in this study, makes the recognition of such inclusions a trivial matter, if they are larger than a few microns in diameter, and none were observed in the grains described here. However, even submicroscopic inclusions of phosphate phases such as monazite or apatite should be easily detected as they are always associated with high contents of Th and Ca + Sr, respectively (Whitehouse & Kamber, 2002). A positive correlation is present in all the samples between the LREE enrichment and Th. A positive correlation is observed between LREE and Ca for sample C1, but no correlation is present in samples C10 and DF6 (Table 2). This pattern could reflect, according to Whitehouse & Kamber (2002), a contamination by monazite microinclusions in samples C10 and DF6 and by monazite and/or apatite microinclusions in sample C1. However, LREE enrichment is not positively correlated with P in sample C1, which shows the best defined positive trends for both Th and Ca, except for two points (C1-13.1 and C1-20), suggesting that the above-mentioned positive correlations are, in most cases, not related to contamination by monazite and apatite microinclusions.
Lu–Hf isotope data
The relatively narrow interval of 176Hf/177Hf observed in zircons from the HPP granitoids and CT dioritic–tonalitic samples and the lack of correlation between 176Hf/177Hf ratios and textural domains in zircons exclude a restitic origin for the zircon domains characterized by relatively ‘older’ ages and suggest that isotopic resetting affected only the U–Pb system, leaving 176Hf/177Hf ratios undisturbed. Thus the Hf-isotope data strongly support the interpretation of the oldest concordant age populations in each sample as reflecting the intrusion age.
The negative Hf values indicate the involvement of old recycled continental crust, with an age of about 1·8 Ga, in the partial melting events that generated the HPP intrusives. Low 176Hf/177Hf ratios and negative Hf values (Fig. 12 and Table 4) also in the quartz diorite and tonalite samples from the ‘Corner Tonalite’ unit could be related to interaction processes between subduction-related mafic melts and crustal-derived components.
Within single rocks (HPP granitoids) and different age populations the Hf isotope data show small variations (Fig. 12). These ranges in the isotopic composition, which are greater than the analytical uncertainty, can be explained by two mechanisms: (1) mixing between a crustal melt and a more primitive component; (2) a heterogeneous source involved in the melting events. The large range and high values of 87Sr/86Sr (0·710–0·716) and low Nd (–8·9 to –7·5) of the analysed granitoids (unpublished data) are consistent with a composite crustal protolith, probably characterized by large amounts of a sedimentary component. The occurrence of a significant number of old inherited zircons in the granitoids is consistent with this model. On the other hand, the coeval mafic magmatism, which often shows magma-mingling relationships with the adjacent granitoids, could provide a more radiogenic Hf component to the crustal melts through mixing processes.
The isotopic composition of zircons from the micro-monzodiorite sample DR12, higher than that measured in zircons from HPP granitoids and CT quartz diorite and tonalite (Fig. 12), with slightly negative to positive Hf values, is lower than that expected for zircons crystallized from a depleted mantle source, suggesting a derivation of the mafic shoshonitic rocks from an enriched mantle source, or contamination of mantle-derived melts by crustal material.
The Keinath monzogranite has slightly negative to strongly positive Hf values. 176Hf/177Hf ratios extend between the CHUR and the Depleted Mantle evolution lines (Fig. 12). Such values suggest that the Keinath Pluton was derived from a source different from the Howard Peaks granitoids: probably young continental crust with an age of about 1·08 Ga which interacted with a more juvenile mantle component as suggested by the abundant mafic microgranular enclaves.
Geological implications of the isotopic data: the intrusive sequence and the tectono-magmatic evolution
The B2 sample of the ‘Corner Tonalite’ unit, the DR12 micro-monzodiorite dyke and the L1 monzogranite from the Keinath Pluton are all characterized by an undeformed fabric, and show narrow clusters of concordant U–Pb ages giving emplacement ages of 497 ± 2·6, 489 ± 3·8 and 481 ± 2·3 Ma, respectively. Only few analytical points (mainly in the monzogranitic sample L1) give younger ages related to Pb loss. REE patterns for the majority of analysed grains are typical of igneous zircons and the variability of all the trace elements is limited.
The geochronological data obtained on the strongly deformed Howard Peaks granitoids C1, C10, M14, DF6 and C4, and on the deformed MS1 sample from the ‘Corner Tonalite’ unit, show complex age spectra with a large range in the concordant data. Such complexity hinders a ready inference of the true intrusion ages from these samples. Moreover, the frequent overabundance of LREE with respect to HREE and the wide range of U, Th and Y contents exhibited by most of the studied zircons suggest that considerable mobility of these elements occurred after the crystallization of the magmatic zircons. This mobility acted in a complex way, sometimes giving rise to overabundances of LREE and large variations in Y, Th and U contents, more rarely causing the trace elements to be expelled from the zircon lattice in domains belonging to the same age population. Therefore, the best estimate of the age of intrusion cannot always be assessed from the concordant ages measured on zircons characterized by oscillatory zoning and typically magmatic trace-element patterns (steep REE patterns with negative Eu and slightly positive Ce anomalies, Th/U ratios >0·1). Indeed, the resetting of the U–Pb system is not always accompanied by the complete erasing of the magmatic textures, or by a strong modification of the typical compositional signature of the igneous zircons.
We propose that the time span over which crystallization took place can be assessed from the weighted average of the oldest (least affected by resetting of the U–Pb isotopic system) ages: the weighted average ages point to an emplacement age in the range 506–493 Ma for the Howard Peaks Pluton granitoids and of 508–497 Ma for the MS1 and B2 tonalite–quartz diorite samples. Therefore, the emplacement of the ‘Corner Tonalite’ intrusives appears to be coeval with the main HPP monzogranitic rocks. The intrusion of the undeformed felsic and mafic dyke swarm in the Deep Freeze Range postdates the main intrusive sequence and is documented only by the isotopic data for the single DR12 micro-monzodiorite dyke at 489 ± 3·8 Ma. Obviously more data are needed (and are in progress) to better define the emplacement age of the dyke system. The late intrusion of the undeformed Keinath Pluton is well constrained at 481 ± 2·3 Ma.
These results indicate that the emplacement sequence in the Deep Freeze Range took place over a time span of at least 30 Myr, from 508 to 481 Ma, in a stress regime that varied from transpressional to extensional. The coeval intrusion of the dominant high-K megacrystic monzogranitic bodies and the high-K calc-alkaline gabbro–tonalitic ones was followed by the emplacement of the shoshonitic mafic dykes, the felsic undeformed dyke swarm and few late calc-alkaline granitic bodies. Figure 13 illustrates schematically the emplacement sequence compared with the relative volume of the intrusives inferred from field relationships and isotope data. This sequence is consistent with that obtained for the Terra Nova Intrusive Complex, Northern Victoria Land (Armienti et al., 1990; Rocchi et al., 2004) and for the main Cordilleran-style magmatism in the Dry Valleys region in Southern Victoria Land (Encarnacíon & Grunow, 1996; Cox et al., 2000; Allibone & Wysoczanski, 2002). Distinctly younger ages ranging to about 435 Ma, with a large number of ages around 470 and 460 Ma, mainly in the deformed granitoids of the Deep Freeze Range, are thought to reflect minor to complete resetting of the primary zircon ages.
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Isotopic data on zircons obtained in this study indicate that mainly the deformed facies underwent a strong resetting of the U–Pb isotopic system beginning soon after their emplacement, thus producing a continuous array of younger concordant ages (from emplacement ages to 435 Ma). The youngest reset U–Pb ages overlap the K–Ar ages obtained on biotite from different lithotypes (micaschists, gneisses and retrogressed granulites) sampled along a transect between the Priestley–Campbell Glaciers zone in the Deep Freeze Range (Vita et al., 1991
, 1994). The K–Ar data show a weighted average of 455 ± 5·8 Ma, close to the Rb/Sr ages of 453 ± 4 Ma on biotites from the same area (Armienti et al., 1990, and references therein). U–Pb data on monazites, zircons and titanites from migmatites and granulitic gneisses of the same area (Klee et al., 1992; Klee, 1995) give ages between 490 and 464 Ma, similar to the U–Pb ages of this study and suggesting a complex resetting of the U–Pb system also in the host metamorphic rocks, coeval with and subsequent to the emplacement of the granitoids.
Recent petrological and geochronological investigations on the Oates Coast metamorphic rocks (Schüssler et al., 1999, 2004; Henjes-Kunst et al., 2004), forming part of the western low-pressure high-temperature belt in the Wilson Terrane, as well as in the Deep Freeze Range (Grew et al., 1984; Palmeri, 1991, 1997), have provided evidence for a common tectono-metamorphic evolution: a single metamorphic P–T–t loop starting at 500–490 Ma for the granulite-facies metamorphic peak down to widespread retrograde metamorphism under wet high-T conditions between 470 and 465 Ma. The post-granulite-facies metamorphic evolution as proposed by Schüssler et al. (1999, 2004) and Henjes-Kunst et al. (2004), is consistent with the data obtained in this study. However, the age of the granulite-facies metamorphic peak is still uncertain: the Howard Peaks granitoids clearly cross-cut the high-grade granulitic basement along the eastern Mt. Emison–Mt. Queensland border. We think that the 500–490 Ma age in the granulites, assumed as dating the metamorphic peak, could instead represent the age of a thermal overprint during the retrograde metamorphic evolution. The common occurrence of a sequence of intrusive mafic rocks with calc-alkaline to shoshonitic affinity in the Deep Freeze Range and in the Northern Foothills, coeval with or slightly younger (Rocchi et al., 2004) than the emplacement of many granitoid bodies, is clear evidence of an important long-lasting mafic subduction-related magmatic event, responsible for the crustal melting that produced the various granitoid suites and also for the thermal input that allowed the slow cooling of the basement. The same situation is outlined in the Oates Land area, where similar mafic bodies of the Granite Harbour Intrusives are exposed (Schüssler et al., 1993).
Given the above scenario, there still remains the problem of what mechanisms triggered the peculiar complex situation seen in this area: the deformation events, fluid circulation in subsolidus conditions, or a thermal input?
The abundance in the deformed intrusives of secondary trails of fluid inclusions in quartz and of fractures in the main mineral phases filled with secondary minerals (muscovite, calcite, quartz, etc.) confirms the major role played by fluids. The occurrence of an important shear system in this area and the wide distribution of minor local shear bands are probably the driving force behind the deformation and the high degree of fluid–rock interaction in the Howard Peaks intrusives. In such a setting, an extensive, fracture-assisted fluid circulation, under variable T conditions (probably sustained by the late mafic magmatism), was an important agent for the transport of Pb and trace elements from and towards the reset zircon grains. Independent fluid inclusion data for neighbouring migmatitic and granulitic rocks (Frezzotti et al., 1994a, 1994b) indicate that a pervasive circulation of complex H2O–CO2 (±CH4 and N2) fluids occurred during the retrograde amphibolite-facies evolution at temperatures above 400°C, gradually evolving to a H2O (±CO2) dominated low-T system (Frezzotti et al., 1990, and unpublished data; Giorgetti et al., 1996). Data on fluid inclusions in rocks from some foliated peraluminous felsic dykes, and preliminary data from the foliated granitoids, indicate that the same types of fluid flooded these rocks. We suggest that the whole basement in the area, involved in the mega-shear zone, was affected by the same type of fluid–rock interaction process. The origin of the fluids is difficult to assess, as they could be related mainly to the mafic intrusions, as well as to a shallower source.
All the above-described basement features, documented along a linear belt extending from the Northern Foothills and Deep Freeze Range to the Matusevich Glacier area in the Oates Coast Land, are too striking to be unrelated and instead point to a common thermotectonic history of this belt. Indeed, the widespread mafic calc-alkaline to shoshonitic magmatism along this belt, clearly providing a heat source for crustal melting and production of the various granitoid suites, and the common low-P high-T evolution of the metamorphic basement in a tectonic regime evolving from transpressional to extensional, requires peculiar thermotectonic conditions. We argue that such features are consistent with models involving crustal and/or lithospheric extension in a back-arc or intra-arc environment, as proposed by several workers for the contiguous Palaeozoic Delamerian and Lachlan Fold Belts in Australia (Sandiford & Powell, 1986; Turner, 1996; Collins, 2002a, 2002b; Foden et al., 2002).
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CONCLUSIONS |
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In the Deep Freeze Range region the Cambro-Ordovician intrusive complex is constituted by a bimodal magmatic association of granitoids and mafic rocks characterized by different fabrics, that suggest a close correlation between the tectono-metamorphic events and plutonism. A ductile solid-state deformation fabric is common and overprinted on magmatic foliation in many of the intrusive bodies, suggesting a synkinematic emplacement with regional transpressional events. Some intrusions are undeformed: field relationships indicate a younger emplacement age in an extensional regime.
The U–Pb isotopic and trace-element chemistry of zircons from the major intrusive pulses of the Deep Freeze Range have been affected by a strong resetting. The probable mechanism invoked here to explain isotopic and elemental resetting is solid-state recrystallization promoted by continuous intergranular fluid circulation in an open system. Zircon characteristics and rock fabrics are interrelated: resetting is more pronounced in foliated granitoids, which were affected by late- to post-magmatic ductile deformation under subsolidus and wet conditions. The resetting process produced complex internal textures, a wide spread of concordant U–Pb ages and a strong variability of the trace-element compositions, with anomalous enrichments of LREE, U, Th and Y. Zircons in samples with undeformed fabrics exhibit, in contrast, simple internal textures, a restricted range of concordant U–Pb ages and limited variability in trace-element compositions, with REE patterns typical of magmatic undisturbed zircons. The distinct zircon age populations and different textural domains within each grain have similar Hf-isotope compositions, suggesting that the Lu–Hf system remained undisturbed.
A careful examination of the textural, isotopic and geochemical data for the zircon grains allows us to extract the emplacement age of the major intrusions in the Deep Freeze Range that covers a time span of about 30 Myr between 508 and 481 Ma. Zircon data from the undeformed intrusives indicate that the main pervasive shear deformation in this region ended at about 490 Ma.
There appear to be three main processes responsible for the observed isotopic and chemical modification of zircon: (1) extensive fluid circulation; (2) pervasive deformation; (3) protracted cooling of the whole basement. Deformation related to the occurrence of a dextral strike-slip shear zone in this area, coeval with the emplacement of the main synkinematic granitoids, provided variable degrees of fluid access. Fluid circulation probably occurred initially at high-temperature conditions, but proceeded also at low-temperature conditions. The long-lasting intrusive mafic and felsic activity provided a heat source to the whole basement, but this external heat supply alone was unable to modify zircon isotopic compositions, as clearly proven by the youngest undeformed intrusives.
The resetting process strongly affected other isotopic systems, producing a large range of K–Ar and Ar–Ar in biotites from magmatic and metamorphic rocks in the whole Deep Freeze Range (Vita et al., 1991, 1994), and was probably a regional event, as suggested by a similar spread of U–Pb data in zircons from gneissic granitoid plutons in South Victoria Land (Allibone & Wysoczanski, 2002).
Our results suggest that zircon, a key mineral in modern geochronology, is not so stable as commonly assumed, particularly when water–rock interaction is involved, and the modified isotopic compositions may lead to erroneous interpretations of radiometric dating in igneous and metamorphic rocks. However, understanding the mechanisms responsible of the resetting processes is a powerful tool in unravelling metamorphic events and subsolidus phenomena in intrusive rocks.
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SUPPLEMENTARY DATA |
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Supplementary data for this paper are available at Journal of Petrology online.
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APPENDIX: SAMPLE CODES AND LOCALITIES OF SELECTED SAMPLES |
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Selected samples are held at the Earth Science Section of the Museo Nazionale dell’Antartide in Siena (MNA-SI).
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ACKNOWLEDGEMENTS |
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This work was carried out with the financial support of the Italian Programma Nazionale di Ricerche in Antartide. We are grateful to Suzy Elhlou, Norm Pearson and Ashwini Sharma for their constant assistance with electron microprobe and LA-ICPMS analyses at GEMOC. We also thank Massimo Tiepolo for laboratory assistance at LA-ICPMS of CNR–Pavia. Andrew H. Allibone, Paul W. O. Hoskin, two anonymous reviewers and the editor Marjorie Wilson are gratefully acknowledged for their constructive reviews. This is Publication 445 from the ARC National Key Centre for Geochemical Evolution and Metallogeny of Continents (www.es.mq.edu.au/GEMOC/).
*Corresponding author. +39-0577233970. +39-0577233938. bomparola{at}unisi.it
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