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Journal of Petrology Volume 42 Number 8 Pages 1429-1448 2001
© Oxford University Press 2001
U–Pb Zircon Studies of Mid-crustal Metasedimentary Enclaves from the S-type Deddick Granodiorite, Lachlan Fold Belt, SE Australia
1DEPARTMENT OF EARTH SCIENCES, LA TROBE UNIVERSITY, BUNDOORA, VIC. 3086, AUSTRALIA
2DEPARTMENT OF EARTH SCIENCES, MONASH UNIVERSITY, CLAYTON, VIC. 3168, AUSTRALIA
3SCHOOL OF EARTH SCIENCES, CURTIN UNIVERSITY, PERTH, W.A. 6001, AUSTRALIA
Received October 1, 1998; Revised typescript accepted January 22, 2001
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONGEOLOGICAL SETTING AND SAMPLESANALYTICAL PROCEDURESRESULTSDISCUSSIONREFERENCES |
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High-grade metasedimentary enclaves in granites can be used to directly characterize the lower and mid crust. Enclaves in the Silurian S-type Deddick Granodiorite, SE Australia, have chemical and Nd–Sr isotopic compositions broadly similar to Ordovician–Silurian clastic sedimentary rocks throughout the fold belt. SHRIMP U–Pb ages of detrital zircons in the enclaves indicate Early Ordovician maximum depositional ages for their psammitic–pelitic precursors. Patterns of zircon inheritance show the same
500 and 1100–1200 Ma peaks commonly observed in Palaeozoic Gondwana margin sedimentary sequences in eastern Australia, New Zealand and East Antarctica. U–Pb ages for metamorphic zircon rims and Sm–Nd dating of garnet indicate that metamorphism in enclave precursors was coeval with granitic magmatism at 430 Ma. The enclaves represent metamorphic equivalents of the voluminous Lachlan Fold Belt Palaeozoic turbidites. After tectonic transport to the mid crust and attendant amphibolite-facies metamorphism during the Benambran Orogeny (440–430 Ma), these Palaeozoic metasediments became the sources for extensive Silurian magmatism. S-type granite magmas in the eastern Lachlan Fold Belt are therefore derived from Ordovician turbidites with possible contributions from Cambrian volcano-sedimentary (‘greenstones’) sequences and from contemporaneous, mantle-derived magmas. KEY WORDS: enclaves; granites; magma source; zircon U–Pb ages
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTING AND SAMPLESANALYTICAL PROCEDURESRESULTSDISCUSSIONREFERENCES |
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High-grade metasedimentary enclaves are common in S-type granitoids world-wide. Unlike coexisting hornfels xenoliths they are generally more deformed, have high-grade mineral assemblages and lack an obvious source in the immediate wallrocks (Montel et al., 1991
). Enclaves of this type are rare or absent in I-type intrusions, implying a close connection between the S-type host granodiorites and their high-grade enclaves. Most workers therefore agree that such enclaves are derived from crustal levels either within or slightly above the source region of the host S-type magmas (Chen et al., 1989; Barbey, 1991; Fleming, 1996; Maas et al., 1997; Anderson et al.,1998). Available pressure–temperature estimates are in general agreement with this conclusion. Metasedimentary enclaves may therefore be used as probes of the mid crust, in much the same way as high-pressure xenoliths in many continental basalts are used as probes of the deep crust and shallow mantle. As long as the effects of metamorphism and residence in the host magma on enclave compositions are understood, the enclaves can yield uniquely direct information on the age, composition and structural history of the mid crust (e.g. Fleming, 1996; Anderson et al., 1998). This information complements inferences made from surface observations, geophysical data and granite compositions, and provides important constraints for tectonic models.
High-grade metasedimentary enclaves can also be discussed in terms of their relationship to their host granitic plutons. They have been interpreted as restite, as refractory parts of the granite source, or as deep-seated xenoliths (Price, 1983; Chen et al., 1989; Barbey, 1991; Montel et al., 1991). The role of their parent lithologies, if any, in magma production is not clear. Isotopically, metasedimentary enclaves often differ from their hosts although compositions may overlap. Where Nd–Sr isotopic contrasts exist, the enclaves typically have higher 87Sr/86Sr and/or lower 
Nd than the host granites (Clarke et al., 1988; Anderson, 1997; Maas et al., 1997; Waight et al., 2000, 2001). This suggests contributions to the granitic magma from isotopically less evolved source rocks not represented by the metasedimentary enclaves, or from mantle-derived mafic magmas (e.g. Clarke et al., 1993; Moreno-Ventas et al., 1995; Patiño-Douce, 1995; Collins, 1996; Castro et al., 1999). If there are additional components in the granite source, detrital zircon age spectra in the host granite and in the enclaves may differ, and the differences may place age and/or provenance constraints on the source components not represented by the enclaves.
Here we present U–Pb zircon ages for high-grade metasedimentary enclaves from the Deddick Granodiorite, a typical example of the mafic, enclave-rich S-type granites of the Palaeozoic Lachlan Fold Belt in southeastern Australia (e.g. Chappell & White, 1992). Mineral geobarometry for the high-grade enclaves indicates pressures of 0·4–0·5 GPa (15–20 km depth), similar to pressures estimated for early crystallization in the host magma (Maas et al., 1997). The enclaves are therefore considered samples of the mid crust as it existed in this part of the fold belt during Silurian time. The principal tool used here—SHRIMP U–Pb dating of detrital zircons—provides robust estimates of deposition ages for the sedimentary precursors of the enclaves and fingerprints the material through its zircon age pattern (e.g. Williams, 1992; Keay et al., 2000). This evidence is used to examine the nature of the basement underlying the voluminous Palaeozoic turbidites of the Lachlan Fold Belt, a topic that has been controversial for many years. Some workers have proposed a concealed, possibly fragmented, continental basement of Proterozoic (White et al., 1976; Chappell et al., 1988) or Cambrian (Ireland et al., 1998) age, whereas others favor turbidite deposition on a Cambrian oceanic crust (Crook, 1980; Crawford et al., 1984). Resolution of this issue would influence models of granite petrogenesis and palaeogeography, and may help explain why the fold belt lacks metamorphic and structural evidence for major thickening despite massive shortening of the Palaeozoic sedimentary sequences (Gray, 1997).
At present, there is no evidence for a continental basement of any age within the fold belt. For example, the Wagga-Omeo Metamorphic Belt in NE Victoria (Fig. 1) is one of the areas suspected of having an old continental basement and one of the few in which high-grade metamorphic rocks are exposed, yet the gneisses in that area have sedimentary protoliths deposited in the Early Ordovician or later (Keay et al., 2000). Similar results were reported for high-grade metasedimentary enclaves in S-type granites in the same belt (Anderson et al., 1998). Another region possibly underlain by Proterozoic crust is the Kosciusko Batholith (Chappell et al., 1988). The Deddick Granodiorite, which hosts the enclaves studied here, forms part of this batholith. Deep-seated enclaves in this pluton may therefore carry the zircon age signature of Proterozoic crust, if it exists. In fact, zircon age distribution patterns obtained from high-grade enclaves may provide less ambiguous signatures than those from granites, because inherited zircons in granites can originate both from the granite source and from Palaeozoic sediments assimilated by the rising felsic magma. Dating zircons in high-grade enclaves circumvents this problem because the detrital zircons cannot be due to contamination but must be from sedimentary precursors.
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The aims of this study were therefore to (1) use detrital zircon ages to extract information on the tectonic history of the mid crust, (2) examine the nature of the Lachlan Fold Belt ‘basement’, and (3) further explore the relationship between high-grade enclaves and their host granites.
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GEOLOGICAL SETTING AND SAMPLES |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTING AND SAMPLESANALYTICAL PROCEDURESRESULTSDISCUSSIONREFERENCES |
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Host granodiorite
The >100 km2 Deddick Granodiorite (Fig. 1) forms a high-level pluton emplaced into low-grade Late Ordovician and Early Silurian pelitic–psammitic sediments in the southern portion of the Silurian Kosciusko Batholith (Vandenberg, 1999
). Deformation and low-grade metamorphism in the sediments occurred during the Silurian, just before widespread granitic magmatism (Williams, 1992; Maas et al., 1997; Vandenberg, 1999). The pluton is relatively mafic (MgO + FeOt 
6 wt %), with primary magmatic cordierite and garnet and evolved isotopic compositions. It is a typical example of the classic mafic S-type granites of the Lachlan Fold Belt (Chappell & White, 1992) and similar plutons form the major component of the Kosciusko Batholith (Hine et al., 1978). Bulk rock sample DR2/10 is typical of the granodiorite found in extensive exposures along the Snowy River near McKillop’s Bridge (see Maas et al., 1997).
High-grade enclaves
Metasedimentary enclaves from <1 to >20 cm in size make up 5–7% of outcrop and are conspicuous on fresh surfaces. Microgranitoid enclaves of igneous origin are subordinate. The metasedimentary enclaves chosen for study cover the range of lithological character within the enclave suite, from psammo-pelitic gneiss to migmatitic and melt-depleted melanosome rocks (Maas et al., 1997).
Enclave 8249 is an unusually large (50 cm x 10 cm) tabular fragment of quartzo-feldspathic gneiss with prominent alternating planar biotite-rich and alkali feldspar–plagioclase–quartz bands of 0·5–1·0 cm scale. Biotite-rich bands are composed mainly of an aggregate of strongly aligned tabular orange–brown biotite crystals to 1 mm, which ‘pinch and swell’ around lenticular cores of cordierite containing characteristic inclusions consisting of felted fibrous sillimanite and strings of brown–green spinel. These bands contain minor granular alkali feldspar and plagioclase. Zircon is abundant in biotite crystals, which have numerous radiation haloes. Cordierite is often extensively altered to ‘pinite’; biotite shows minor chloritization. Quartzo-feldspathic bands comprise mainly alkali feldspar, plagioclase (An25–40) and quartz (approximately 50:35:15 by volume), showing well-developed sub-equigranular 0·5–2·0 mm granoblastic texture. Feldspars show minor sericitic alteration.
Enclave DR2/3 is a large migmatitic enclave with folded banding, prominent narrow, lenticular leucosome pods up to a few centimetres in length, and quartzo-feldspathic veins of 1–3 mm width. Melanosome material contains 10–15% of equant almandine-rich garnet to 5 mm, usually developing 0·5–1·0 mm cordierite rims, with inclusions of biotite, apatite, zircon and ilmenite. The garnet is set in a matrix of coarse granular to subhedral prismatic cordierite (typically 1–3 mm, occasionally to 5 mm long), biotite plates to 4 mm, and minor alkali feldspar, plagioclase, quartz and ilmenite. Cordierite crystals are charged with dense sheafs of fibrolitic sillimanite and strings of deep green Zn-rich hercynites. Cordierite is partially pinitized, biotite shows minor chloritization, and the feldspars are lightly sericitized. There are rare patches of coarse radiating muscovite and quartz. Zircons again appear to form mainly inclusions in biotite. Leucosome pods consist of irregular to subequant, often strained, quartz crystals to 3 mm (>50%), with less abundant slightly perthitic alkali feldspar of similar grain size, and 10–15% red–brown biotite plates. Veins have similar assemblages, richer in quartz.
Enclave DR4/1 is an example of the dominant ‘cordierite–garnet’ enclaves, which consist mainly of cordierite ± garnet and alkali feldspar, with minor plagioclase (An25–40) and very little quartz. Their mineral assemblages are often very similar to those of melanosomes in banded migmatite enclaves such as DR2/3, but they may be even more aluminous, containing very abundant fibrolitic sillimanite sheafs, Zn-rich hercynite strings and, in the case of DR4/1, corundum crystals up to 5 mm long. These assemblages are residues from significant melt extraction. DR4/1 is patchily banded, with the dominant material consisting of almandine-rich garnets to >5 mm and rare corundum crystals set in a matrix of coarse cordierite with dense sweeping zones of sillimanite and hercynite inclusions, and interstitial granular alkali feldspar. Less voluminous feldspar- and biotite-rich patches or discontinuous bands consist mainly of granoblastic plagioclase and less abundant alkali feldspar, with 10–20% biotite and ilmenite.
Chemical and isotopic compositions
The host granodiorite sample and the enclaves studied here have compositions that are comparable with those reported by Maas et al. (1997). Enclave 8249 (67% SiO2, 13·1% Al2O3, 6·4% FeOt and 1·0% CaO) has major element concentrations similar to those of previously studied psammitic–pelitic enclaves in this and other S-type granites in the area, and plots within or near the field for Ordovician–Silurian clastic sedimentary rocks of the Kosciusko Batholith region (Wyborn & Chappell, 1983). Initial Sr–Nd isotope ratios (87Sr/86Sr = 0·71353, Nd = -10·6) are similar to those of the host granodiorite and are again within the range for Lachlan Fold Belt Ordovician–Silurian sedimentary rocks (see Maas et al., 1997). By contrast, enclaves DR2/3 (50·1% SiO2, 24·5% Al2O3, 8·7% FeOt, 0·8% CaO, 87Sr/86Sr = 0·71768, Nd = -10·9) and DR4/1 (50·2% SiO2, 23·7% Al2O3, 9·8% FeOt, 1·0% CaO, 87Sr/86Sr = 0·71882, Nd = -12·2) have substantially lower SiO2 and higher Al2O3, MgO and FeOt than both enclave 8249 and the host granodiorite, reflecting their more cordierite–garnet-rich mineralogy. Their chemical and isotopic characteristics follow the trends of the gneissic and cordierite–garnet (melanosome) enclaves described by Maas et al. (1997).
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ANALYTICAL PROCEDURES |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTING AND SAMPLESANALYTICAL PROCEDURESRESULTSDISCUSSIONREFERENCES |
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Zircons were separated by conventional heavy liquid and magnetic techniques, mounted in epoxy and polished to expose their mid-sections. U–Th–Pb isotopic analyses were carried out using the SHRIMP II ion-microprobe at Curtin University. 206Pb/238U ages were used for zircons younger than
700 Ma, and 207Pb/206Pb radiogenic Pb model ages were used for all older analytical sites. Common Pb corrections were made by the 208Pb method, or, where Th–Pb systems were disturbed or suspected of being disturbed, the 204Pb method was used (Chen & Williams, 1991). The common Pb composition used was either Broken Hill Pb (where measured 204Pb counts were <6 times the 204Pb counts measured on the zircon standard) or the appropriate common Pb from the Cumming & Richards (1975) growth curve (where measured 204Pb counts were greater). Pb/U ages obtained using the 208Pb or 204Pb common Pb corrections were always within 10 my of the respective ages obtained after correction by the 207Pb method (e.g. Muir et al., 1996). The 204Pb method was used for all analyses older than 700 Ma. Errors for individual analyses are presented as 1
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RESULTS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTING AND SAMPLESANALYTICAL PROCEDURESRESULTSDISCUSSIONREFERENCES |
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Sm–Nd garnet chronometry
Enclave DR4/1 is rich in almandine garnet formed during amphibolite-facies metamorphism (Maas et al., 1997
). Sm–Nd dating was used in an attempt to constrain the age of garnet formation and/or closure. The two-point model age for the whole rock and the garnet (523 ± 110 Ma) is imprecise because of the low Sm/Nd of the garnet separate (Table 1). Two further aliquots of the same garnet separate were therefore leached with strong HCl and HCl–HNO3, respectively, to remove Nd-rich impurities (see Zhou & Hensen, 1995). The residues produced in this way were indeed lower in Nd than the bulk garnet, and Sm/Nd ratios were increased by factors of two and 13 (Table 1). The greatly improved dispersion (Fig. 2) results in a precise isochron age of 434 ± 4 Ma [six data points, mean square weighted deviation (MSWD) 0·54, initial Nd = -12·1 ± 0·2, all errors 95% confidence limits, ISOPLOT 2.96, Ludwig, 1991]. This isochron is strongly controlled by the heavily leached residue of aliquot 3, which has the rare earth element (REE) characteristics of pure garnet (Hickmott et al., 1987). Omitting this data point leaves the age more or less unchanged but results in a larger error (431 ± 19 Ma, MSWD 0·93, Nd = -12·1). This suggests that the source(s) of REE in the leachates, presumably light REE (LREE)-rich inclusions in the garnet, were in isotopic equilibrium with the pure garnet. The 434 ± 4 Ma age is indistinguishable from the U–Pb zircon age for the host granodiorite (see below) and suggests that metamorphic garnet in the enclave precursors formed during the same thermal event. Alternatively, an older garnet age may have been reset during Silurian metamorphism and/or residence in the host magma.
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U–Th–Pb isotope results
Zircon morphology
Zircons separated from the three enclaves and the host granodiorite are generally colourless, vary in size from 50 to 300 µm, and typically show moderate elongation (length to width ratio, l/w, ranging from one to three). A small subset of zircon grains (10–15%) in the host granodiorite shows greater l/w ratios (>4); these are associated with the youngest zircon generation formed from the host melt, whereas inherited grains tend to have smaller l/w ratios. Grain shapes are euhedral or nearly so for 30–50% of grains in the enclaves but slight face angle rounding is observed even on those grains described as euhedral here. Severely rounded shapes are observed for 15–30% of grains, the remainder having transitional properties. This contrasts with the perfectly euhedral outlines of most of the zircon grains from the host granodiorite. Distinct cores are visible optically or in back-scattered electron (BSE) and cathodoluminescence (CL) images in most of the grains. Core shapes in the enclave zircons vary from near euhedral to severely rounded whereas cores in the host are nearly always rounded to subhedral. Where present, inclusions are almost always in the core. Cores are overgrown by one or more later growth stages separated from the core and each other by surfaces often showing evidence of resorption; up to four growth stages may be discerned. CL intensity in cores varies from broad faint striping to oscillatory zoning and highly convoluted, irregular patterns. Zircon rims in the host granodiorite and in the leucosome of migmatite enclave DR2/3 are compositionally zoned, whereas the thin, often discontinuous rims developed on many of the enclave zircons largely lack discernible BSE or CL structure. U–Th–Pb isotope data are reported in Table 2.
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Host granodiorite (DR2/10)
Fifty-four sites on 46 grains were analysed (Table 2a), yielding the typical pattern of mostly concordant to near-concordant U–Pb compositions commonly observed in SHRIMP studies of Phanerozoic rocks (Fig. 3a). Ages cluster at 400–600 Ma and 900–1000 Ma, with a scattering of older ages up to 2300 Ma. The 400–600 Ma cluster can be resolved into several discrete subgroups. Twenty analyses on 18 grains with U–Pb ages between 415 and 442 Ma form the youngest subgroup. These ages were obtained from thick, intensely zoned mantles surrounding distinct cores, and, in eight cases, also from central parts of grains (e.g. g16 and g20). Two grains (g17 and g25) contain coeval, strongly zoned growth stages separated by a discontinuity. U contents are mostly around 300 ppm (107–823 ppm) and Th/U is typically 0·3–0·5 (0·05–0·6). As outlined below, these sites probably represent melt-precipitated zircon formed in the granodiorite. Pooling and exclusion of one outlier (415 Ma) yields a mean of 430 ± 3 Ma (n = 19, 2m, 
2 = 0·73).
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Inherited zircon ages can be grouped into a number of age intervals. Eleven sites on 10 grains yield U–Pb ages in the interval 463–512 Ma. Most of these sites (Th/U mostly 0·1–1, one at 0·004) are on zoned zircon mantles surrounding older cores. Another cluster with ages in the interval 551–704 Ma from 11 sites on 11 grains (Th/U 0·2–2·2, some <0·1) was mostly obtained from distinct cores, but some of the ages again represent zoned overgrowths on older cores. The remainder of sites yields diverse ages from 890 to 2300 Ma on a variety of cores and rounded grains.
Quartzo-feldspathic (psammitic) paragneiss enclave (8249)
Zircon ages in this sample cluster at 460–510 Ma, 800 Ma and 1100–1200 Ma, with a few analyses near concordia from 2000–2800 Ma (Table 2b, Fig. 3b). U–Pb ages for the youngest cluster range from 458 to 507 Ma (22 analyses on 21 grains). These analyses were sited on cores (or grain centres where structure was unclear) or on zoned inner mantles surrounding the cores, within euhedral to subhedral grains. There is no correlation of age with either site type. U and Th/U range from 100 to 691 ppm, and from 0·3 to 1, respectively. Grain b13 was analysed on its core (b13-1) and on a volumetrically minor inner mantle stage (b13-2) with similar ages of 465 and 482 Ma, respectively. This suggests that age differences between inner growth stages are small in at least some cases, consistent with the similarity in ages between cores and mantles in different grains in this age cluster. The zoned inner mantle of grain b35 was analysed at two sites, with identical ages (491 and 494 Ma, the latter not listed in Table 2) and Th/U ratios. Two further analyses made in the central parts of grains b28 (b28-1, 389 Ma) and b27 (b27-1, 408 Ma) yielded ages that are too young to be geologically plausible. Both plot near the concordia and probably belong to the 500 Ma cluster; their young ages are probably due to radiogenic Pb loss.
A histogram of the 450–510 Ma cluster (Fig. 3b) suggests more than one discrete age, consistent with a high 2 value of 3·6 for the entire cluster (average 485 ± 6, 2m). Two peaks (at 480 and 495 Ma) can be resolved visually from Fig. 3b and this is substantiated by mixture modelling (Sambridge & Compston, 1995). A very similar distribution of discrete Cambro-Ordovician and Early Ordovician zircon ages has been observed in enclaves from the Lachlan Fold Belt in central and northeastern Victoria (Anderson et al., 1996). The remaining analyses indicate ages between 548 and 2705 Ma, with most of them between 830 and 1200 Ma. With the exception of two analyses made on mantles, all analytical sites were located on cores or on the central parts of grains without recognizable cores. The older ages in particular were obtained from rounded grains.
Migmatitic enclave (DR2/3)
The unusually large size (30 cm diameter) of this enclave allowed separation of leucosome, melanosome and ptygmatic quartz–feldspar veins by means of a small diamond saw. Despite some mixing of components, zircons recovered from the leucosomes clearly have a larger size range, a higher (40% vs 20–30%) proportion of euhedral grains, and thicker, more strongly zoned rims than the zircons from the other migmatite subcomponents.
Leucosome. Zircon ages for the leucosome are mostly in the range 362–560 Ma, with a scattering of ages from 626 to 2580 Ma (Table 2c, Fig. 4a). U–Pb ages for the outer growth stages of seven euhedral grains range from 362 to 434 Ma. One of these sites is too young (<400 Ma, d7-1) to be geologically plausible and has probably lost Pb. The sites in this group have rather uniform U contents (462–686 ppm) and low Th/U (0·02–0·14, mostly <0·1). Excluding d7-1, an average of 425 ± 6 Ma (2m) is obtained for the remaining six sites. This age is the same within errors as the magmatic age of the host granodiorite. Interestingly, ages in this group were obtained both from the thin outermost growth stages and from structurally older stages. This suggests that age and compositional differences between the outer growth stages are subtle or that the analytical results are dominated by one 425 Ma, U-rich layer.
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Analyses of cores (17 sites on 17 grains) yielded U–Pb ages between 473 and 561 Ma (Fig. 4a). These sites are chemically diverse with 58–1035 ppm U and Th/U between 0·06 and 1·4 (mostly 0·3–0·8). One of these cores (e2-2, 512 Ma) has a near-coeval, faintly zoned rim (e2-1, 473 Ma) with a Th/U (0·26) that is significantly higher than the ratios for other rims on leucosome zircons. No other example of a rim that falls in this age group has been found. Grain e2 is an unusually elongated (l/w >5) grain which would be expected to break during sedimentary cycling unless it was armoured by another mineral. Armouring could explain the absence of
425 Ma low-Th/U rims found on many other grains. Another nine cores, several of them strongly rounded, have ages from 626 to 2580 Ma. U (151–782 ppm) and Th/U (0·1–1) are less variable than in the younger cores. Grain e21 has at least three growth stages and was analysed in two locations. The core (e21-1) yields 2122 Ma whereas an inner mantle (e21-2) yields 657 Ma. Melanosome. Only three U–Pb analyses are available for zircons from the melanosome. These analyses (a14, 15 and 16) were carried out on intensely zoned zircon overgrowing discrete cores and yield U–Pb ages of 483–495 Ma (Table 2c, Fig. 4a). Th/U ratios are between 0·24 and 0·46.
Quartz–feldspar veins. Analyses were made on the cores or centres (where grain structure was not optically discernible) of 18 grains but several analyses are suspect because of poor run quality or inner mantle–rim overlaps. No rims were analysed as these were expected to yield the same age as those in the leucosome. Four cores and the inner mantle of grain f12 yield U–Pb ages between 476 and 519 Ma (Table 2c, Fig. 4a). Another core (f5-1) yields an implausible age of 410 Ma, and probably suffered some Pb loss. The core of grain f12 yields an age of 2010 Ma whereas the ages for 10 other cores range from 1046 to 3120 Ma.
In summary, distinctions between the zircons separated from the migmatite subcomponents are minor and relate mainly to crystal size and shape, i.e. reflecting the effect on zircon growth of more abundant felsic melt in leucosomes. U–Pb age spectra of inherited components do not show obvious differences and it is assumed that the age of 430 Ma measured for zircon rims in the leucosome applies to all three subcomponents.
Cordierite–garnet gneiss (DR4/1)
Zircon ages are predominantly in the range 400–500 Ma with scattered ages out to 2776 Ma (Table 2d, Fig. 4b). U–Pb ages for the rims of six complex grains range between 400 and 452 Ma. These rims have 479–1345 ppm U and low Th/U (<0·1), a signature also found in most of the young zircon rims in the leucosome of enclave DR2/3. Some of the observed age range is almost certainly due to minor Pb loss in some cases, and to overlap of the primary ion beam onto older cores in others. This goes some way towards explaining the large error of the mean for five sites of 425 ± 15 Ma (2m). This age is identical within uncertainties to the more precise ages for the host granodiorite and for the low-Th/U zircon rims in the migmatite leucosome.
Analyses of distinct cores (or central parts of structureless grains) return a variety of ages. A small group of five sites on four grains has U–Pb ages between 481 and 521 Ma, whereas a further 12 analyses on 12 grains have Proterozoic ages ranging from 670 to 2776 Ma, with no particular clustering except for a group of four ages near 1000 Ma. Almost all of these sites have Th/U ratios between 0·3 and 1.
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DISCUSSION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTING AND SAMPLESANALYTICAL PROCEDURESRESULTSDISCUSSIONREFERENCES |
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Crystallization age of the host Deddick Granodiorite
The crystallization age of the host granodiorite is estimated at 430 ± 3 Ma from analyses of euhedral, zoned rims found on almost all zircons from this rock. Dating of such rims in S-type granitoids provides consistent and geologically plausible crystallization ages even where inherited zircon is abundant (e.g. Williams, 1992
; Keay et al., 2000). It is noteworthy that several 430 Ma sites have been found in the centres of grains lacking distinct cores, and in two cases indistinguishable ages were measured on zoned growth stages separated by a discontinuity within the same grain. Apparently, new zircon growth from the granodiorite melt was not restricted to seeds provided by inherited zircon grains, and zircon growth appears to have been punctuated by periods of resorption. The 430 ± 3 Ma zircon age is within error of the 425 ± 2 Ma biotite–whole rock Rb–Sr age for this sample (Maas et al., 1997). It is also consistent with conventional 430 Ma U–Pb ages for monazite and with ionprobe U–Pb zircon ages for other S-type granitoids of the Kosciusko Batholith and the nearby Berridale Batholith (Williams et al., 1983; Chappell et al., 1991; Williams, 1992).
The 430 ± 3 Ma age is early Silurian [Llandovery, 434–425 Ma, time scale for Australia of Young & Laurie (1996)] and can therefore be only marginally younger than the stratigraphic age of the local host rocks, the early Silurian turbidites of the Yalmy Group (Llandovery, Vandenberg, 1999). During the ‘Benambran’ orogenic phase, the Yalmy Group was tectonically interleaved with Ordovician turbidites and black shales, metamorphosed at lower greenschist facies and now forms part of the Yalmy fold and thrust belt. The Deddick Granodiorite and nearby coeval S-type plutons postdate this structure (e.g. Glen & Vandenberg, 1987) and therefore constrain the ‘Benambran’ in this area to a period of a few million years in the Llandovery at 430 Ma (see also Vandenberg, 1999).
Age of high-grade metamorphism and partial melting in the enclaves
Metamorphic mineral assemblages in the Deddick enclaves indicate temperature–pressure conditions during metamorphism of 750–850°C at 0·5 ± 0·1 GPa, conditions similar to those estimated for the granite source (Maas et al., 1997). This suggests that metamorphic recrystallization and partial melting in the enclaves, and the generation of the host granodiorite magma, were part of the same thermal event. Independent support for this correlation is provided by U–Pb ages for thin rims present on many zircons extracted from the enclaves. With the exception of the rims on leucosome zircons, these overgrowths lack the zoning typical of the grain interiors or of igneous rims on zircon grains in the host granodiorite. They have moderate to high U contents and Th/U ratios are low (<0·1), regardless of trace element levels in the grain interior. Although low-Th/U sites have occasionally been observed within zoned zircon, this feature is most pronounced and consistent in zircon rims. Structural uniformity and low Th/U are thought to be typical of zircon formed during metamorphism (Williams & Claesson, 1987; Page & Laing, 1992), and we therefore infer that the low-Th/U rims in the enclaves grew during metamorphism and migmatization in the precursor metasediment before enclave formation. Zircon growth during enclave residence in the host granite magma is a less likely alternative because low-Th/U rims did not form on detrital zircons in similar enclaves collected from the S-type Koetong Adamellite (Anderson et al., 1996).
Low-Th/U rims on zircons from garnet–cordierite gneiss enclave (DR4/1) yield an average 206Pb/238U age of 425 ± 15 Ma. An identical age (425 ± 6 Ma) was observed for prominent low-Th/U rims on leucosome zircons in migmatitic enclave DR2/3. Both ages are similar to the 430 ± 3 Ma age for the host granodiorite, supporting the correlation between metamorphism in the enclaves and host magma generation. The ages also confirm that high-grade metamorphism at depth (15–20 km) and low-grade metamorphism at higher crustal levels within the Yalmy fold and thrust belt were part of the same thermal event (Collins, 1998; Vandenberg, 1999).
Further support for this link comes from the 434 ± 4 Ma Sm–Nd age for metamorphic garnet in enclave DR4/1. Although it is possible that this age simply reflects resetting of a potentially much older garnet Sm–Nd system during residence of the enclave in the host magma (at >750°C), this is considered unlikely. The garnet isochron (excluding the host enclave) has the same initial Nd as the host enclave and the other garnet-bearing enclaves (-12·1 vs -10 to -12) and shows no sign of the raised initial 143Nd/144Nd expected in a reset older garnet (high-) Sm/Nd system.
Depositional age of enclave sedimentary precursors
U–Pb ages for the youngest detrital zircons in any clastic sedimentary rock provide a maximum age for deposition. This concept is applied here to infer upper limits to depositional ages for the enclaves’ clastic precursors. Before this can be done, however, we must first examine if U–Pb systems in the zircons have remained undisturbed. Assessing the extent of Pb loss in a detrital zircon suite is problematic because Pb loss is clearly identifiable only where the analysed sites are strongly discordant or where concordant U–Pb ages are anomalously young. On this basis, we infer that only a few analyses have been affected by Pb loss and may give spuriously young ages. None of the sites on the youngest detrital zircons were strongly discordant and the vast majority belong to sharply peaked distributions, which have similar ages in the four rocks studied here. We will therefore assume that the youngest clearly discernible peaks (excluding rims) in the age distributions represent zircon that was ultimately derived from igneous sources older than, or contemporanous with, sediment deposition.
The age distributions for the youngest detrital zircon ages in all three enclaves peak at around 470–500 Ma, pointing to an Early Palaeozoic age of deposition of their sedimentary precursors. Psammitic gneiss enclave 8249 yields an age of 475 Ma for the youngest cluster of detrital zircon ages (Fig. 5) whereas the youngest major age peaks in the other two enclaves are at 490–500 Ma. This indicates maximum stratigraphic ages in the Early Ordovician and near the Cambrian–Ordovician boundary, respectively. Very similar depositional ages were inferred from zircon studies of enclaves in other S-type granites in northeast and central Victoria (Anderson et al., 1996; Anderson, 1997) and from high-grade metasedimentary rocks in northeast Victoria (Keay et al., 2000). The youngest detrital zircon ages in the host granodiorite also peak at 485 Ma. The presence of these zircons in the granodiorite must at least in part reflect abundant disaggregated enclave material.
Three sites on zircons from the ptygmatic veins in enclave DR2/3 and one site in the cordierite–garnet gneiss have ages around 450 Ma, forming the small hump on the low-age side of the composite peak for these populations (Fig. 5). Taken at face value, these ages would suggest an even lower (Late Ordovician) maximum age for the enclave precursor sediments. Alternatively, the ages may reflect partial resetting of zircon U–Pb systems during residence in the host magma. A similar situation has been reported by Keay et al. (2000) for metasediments in the Wagga-Omeo Metamorphic Belt of northeast Victoria. The youngest major peak of inherited zircon ages in these sediments provides an 465 Ma maximum age for deposition. Rare 430 Ma ages in these rocks were explained as resulting from local resetting of older detrital zircons in the contact aureole of a nearby 430 Ma granite intrusion (Keay et al., 2000). However, the rare 450 Ma ages in the two Deddick enclaves are more readily explained as analytical artefacts: the analysed sites either show minor rim–inner mantle overlaps and/or have high common Pb contents.
Age spectra of inherited zircons
The age spectra in the enclaves show peaks at 470–550 and 900–1150 Ma, and scattered ages to 3100 Ma. Ages between 550 and 900 Ma are rare except for a group of three ages at 800–900 Ma in enclave DR4/1. The host granodiorite shows a similar pattern but contains an additional group of nine ages at 580 Ma.
Peaks in zircon age distributions around 470–600 Ma and 1100 Ma have also been observed in metasedimentary enclaves from other S-type granites in Victoria (Anderson et al., 1996; Anderson, 1997), and they are typical of Palaeozoic sedimentary, metasedimentary, and granitic rocks of the Lachlan Fold Belt (Williams et al., 1990, 1992, 1994; Keay et al., 2000) and other parts of eastern Gondwana preserved in New Zealand and Antarctica (Ireland, 1992; Gibson & Ireland, 1996; Ireland et al., 1998; Veevers, 2000). The sedimentary lithologies represented by the enclaves therefore appear to have the same provenance and may have formed within the same depositional system as the broadly coeval turbiditic sediments exposed throughout southeast Australia, New Zealand’s Western Province and East Antarctica (e.g. Veevers, 2000). The broad chemical and Nd–Sr isotopic similarities between the enclaves and the Ordovician–Silurian Lachlan Fold Belt sediments are consistent with this interpretation (Price, 1983; Chen et al., 1989; Anderson, 1997; Maas et al., 1997). This implies tectonic transport of lateral equivalents of the Lachlan Fold Belt Palaeozoic sedimentary sequences to depths of 15 km before early Silurian granitic magmatism, consistent with evidence for major telescoping and thrust stacking in that period (e.g. Glen & Vandenberg, 1987; Gray, 1997).
The Lachlan Fold Belt basement
Although the Ordovician turbidites and underlying Cambrian volcano-sedimentary sequences have long been related to an oceanic setting for the Lachlan Fold Belt (Crook, 1980), chemical-isotopic data for the voluminous Silurian–Devonian S-type granites appeared to require an additional feldspathic sedimentary source, which was thought to lie below the Ordovician turbidites (e.g. Wyborn & Chappell, 1983). The suggestion that this source could be a Precambrian continental basement (White et al., 1976) appeared to be supported by zircon inheritance and Sr isotopic data for the granites available at the time (Compston & Chappell, 1979; Williams et al., 1983). As pointed out by Collins (1999; see references therein), this ‘required’ continental basement has influenced tectonic models for the fold belt over the last 20 years.
Zircon U–Pb ages for high-grade metasedimentary enclaves in S-type granites are now available from three localities (Cobaw, Koetong and Deddick) in the western, central and eastern Lachlan Fold Belt, respectively (Anderson et al., 1996; Anderson, 1997; this study). All the results imply that Ordovician metasedimentary rocks were (and presumably still are) present within the crustal depth range at which S-type granite magmas originated. There is no evidence from these studies for a Precambrian basement. The mid-crustal Ordovician metasediments represented by the enclaves contain detrital zircons derived from sources ranging in age from Cambro-Ordovician to Archaean, and have evolved Sr–Nd isotopic compositions, similar to what is observed in most S-type granites in southeastern Australia. An older basement may exist below the level sampled by the enclaves, but it is no longer needed to explain the granite data. For example, the granite zircon data have been explained in terms of a largely Ordovician sedimentary source (Williams, 1995), whereas Rossiter & Gray (1996) favoured a Cambrian volcano-sedimentary source similar to the Cambrian ‘greenstone’ sequences of the southern Lachlan Fold Belt in Victoria.
Furthermore, a Proterozoic continental basement would be in conflict with Nd–Sr isotopic data presented here and by Maas et al. (1997). These data show the host Deddick Granodiorite to be isotopically less evolved than the high-grade enclaves. By contrast, Proterozoic basement inliers throughout the Australian continent contain highly radiogenic Sr (calculated for a Silurian age; Gray & Webb, 1995) and any contributions from such a source would cause a much higher 87Sr/86Sr in the granites than is observed.
The Ordovician zircon ages in the enclaves and granites also weaken the case for a 500 Ma old ‘Delamerian-type’ basement (Gibson & Ireland, 1996). Crustal rocks of the Adelaide–Kanmantoo fold belt and the Ross orogen (Fig. 1) were consolidated during the Delamerian–Ross Orogeny at 500 Ma and formed a landmass from which Lachlan Fold Belt turbidites were derived (e.g. Veevers, 2000). Gibson & Ireland (1996) reported that 500 Ma igneous rocks underlie metamorphosed equivalents of the Lachlan Fold Belt turbidites in parts of Fjordland (southwestern New Zealand). The Fjordland igneous rocks are associated with lithologies similar to those in the Kanmantoo fold belt. Gibson & Ireland (1996) therefore proposed that Kanmantoo-type rocks may also form a basement under southeastern Australia. They further suggested that Kanmantoo-type material could be a more suitable source for S-type magmas than the quartz-rich Lachlan Fold Belt turbidites. Support for this argument comes from the similar zircon age spectra found in the Fjordland rocks on the one hand, and in Lachlan Fold Belt sediments and S-type granites on the other (Ireland et al., 1998). Sr–Nd isotopic data also permit such a relationship. Initial 87Sr/86Sr ratios of the Kanmantoo Group and Lachlan Fold Belt turbidites, S-type granites and high-grade enclaves are all similar (e.g. Gray & Webb, 1995) and substantial overlap exists for Nd isotope ratios (Turner et al., 1993). Furthermore, the Kanmantoo (and other rocks of the Adelaide Fold Belt) is chemically diverse (e.g. Turner et al., 1993), with Ca- and/or Na-rich rocks that are rare in Lachlan Fold Belt turbidites.
If this model is correct, the Deddick enclaves may represent such Kanmantoo-type crustal material. However, sediment deposition in the Kanmantoo and underlying Normanville Groups of the Adelaide–Kanmantoo Fold Belt is entirely Cambrian and was terminated by 514 ± 5 Ma, the U–Pb zircon age for syn-orogenic plutonism associated with the Delamerian Orogeny (Foden et al., 1997). Post-orogenic magmatism at 485 Ma marks the end of the Delamerian Orogeny (Turner et al., 1996). The youngest detrital zircons in Kanmantoo-like crust should therefore be no younger than 520 Ma. By contrast, the youngest well-defined detrital age peaks in the three Deddick enclaves occur at 480–490 Ma. Resetting of Cambrian grains by 485 Ma Delamerian post-orogenic intrusive rocks is possible but no such intrusive rocks are known from pre-485 Ma outcrop within the Lachlan Fold Belt, for example in the Cambrian sedimentary sequences of western Victoria. Examination of the zircon age spectra for the Kanmantoo and Fjordland rocks (Gibson & Ireland, 1996; Ireland et al., 1998) is complicated by Pb loss and the age of the youngest detrital grains is not known exactly. On present evidence therefore, it appears unlikely that the enclaves in the Deddick Granodiorite and in other S-type granites are correlative with the Kanmantoo-type rocks. This does, of course, not preclude Kanmantoo-type basement below the level sampled by enclaves and therefore possibly within granite source regions. It also does not preclude derivation of some Lachlan Fold Belt turbidites by erosion of the Delamerian orogen (Gray & Webb, 1995; Ireland et al., 1998); however, this is not the issue here.
The enclave–host granite connection
Most high-grade metasedimentary enclaves in the Deddick Granodiorite show evidence of partial melting and/or melt depletion, and based on the age dating it is inferred that equivalents of these enclaves formed part of the granite magma source. However, as noted by Maas et al. (1997) and in other studies, Ca–Na contents are too low and Nd–Sr isotope data too evolved for the enclaves to represent the only granite source component. Production of the typical Ca-rich S-type magmas appears to require complementary components richer in Ca and Na (Wyborn & Chappell, 1983) and with more primitive Nd–Sr isotopic compositions (McCulloch & Chappell, 1982). Ca–Na-rich Ordovician–Silurian metasediments do occur in the fold belt, e.g. the paragneisses of the Gundowring Terrane in northeastern Victoria (Keay et al., 2000), but they are uncommon elsewhere (Wyborn & Chappell, 1983). Ca–Na-rich, isotopically unevolved contributions to the felsic magmas could also be provided by the Cambrian mafic–intermediate volcanic rocks and associated sediments in the magma source (Collins, 1996; Rossiter & Gray, 1996) and/or by a juvenile, mantle-derived component (Gray, 1984, 1990; Collins, 1996; Keay et al., 1997).
If the enclaves are not representative of the magma source, the host granite may contain detrital zircon components not present in the enclaves. This should show as a gap in the enclave record. In the Deddick Granodiorite, the zircon age spectra in enclaves and host granodiorite are similar, with the exception of a sizeable component (nine grains of 54 analysed) at 580 Ma in the granite; this component is rare in the three enclaves. The geological significance of this observation is uncertain at this stage. It could be a sampling artefact related to the small number of detrital grains (<50) analysed for each enclave. However, the probability of missing a component comprising a fourth of all detrital grains in the host granodiorite sample—if it is indeed present in the enclaves—is very small. The results of Anderson (1997) indicate that 600 Ma detrital grains are present in enclaves from other Victorian S-type granites (Koetong, Cobaw) but they are absent from four samples of metasediment from the Koetong area (Keay et al., 2000). Ordovician turbidites studied by Williams et al. (1994) contain a consistent 500–600 Ma age component, as do (meta-) sedimentary rocks from the Gondwana margin in Antarctica and from New Zealand (Ireland, 1992; Gibson & Ireland, 1996). Clearly, the zircon age patterns in clastic (meta-) sediments show considerable variation in detail (Keay et al., 2000). The 580 Ma gap in the enclave zircon record may therefore be a feature of only local significance.
Implications for S-type magma genesis
On present evidence, S-type granite magmas of the southeastern Lachlan Fold Belt are largely sourced in the Ordovician–Silurian turbidite and, possibly, in Cambrian volcano-sedimentary sequences. A Cambrian volcano-sedimentary ‘greenstone’ source is not needed to explain the detrital zircon ages but it remains attractive as a source of Ca, Na and relatively unevolved Nd and Sr isotopic compositions. In this regard, it is curious that no enclaves have been found that can be linked to the Cambrian greenstones. A Precambrian continental basement source can be ruled out based on the enclave studies, isotopic data for the granites, and geodynamic modelling (e.g. O’Halloran & Rey, 1999).
Mid-Palaeozoic S-type granites and associated volcanic rocks make up >30% of the surface geology in the eastern Lachlan Fold Belt, an unusually high fraction by world standards (e.g. Coney et al., 1990). Crustal melting on this scale would require a combination of hot mid crust and mantle-derived mafic magma input to produce extensive crustal melting zones. Mid-crustal heating may have been a result of underthrusting, crustal thickening and thermal recovery over the 10–20 my before late- to post-Benambran Early Silurian S-type magmatism (e.g. Collins, 1998). Pre-430 Ma, deep underthrusting of Ordovician–Silurian sediments is consistent with seismic data (Gray, 1997), with the chemical, Nd–Sr isotopic and maximum depositional ages for the enclaves at all three sites studied to date, with the Sm–Nd garnet age for Deddick enclave DR4/1, and with the U–Pb ages for metamorphic zircon rims in the Deddick enclaves. The garnet and zircon rim ages indicate only one episode of high-grade metamorphic zircon growth, near 430 Ma, coeval with the time of granite emplacement. This suggests that mid-crustal metamorphic conditions peaked only once following burial and underthrusting, presumably after the time required for thermal recovery and coeval with mafic magma input into the mid crust.
The tectonic setting of Silurian crustal shortening is not clear but a number of workers have suggested a subduction setting (Fergusson & Vandenberg, 1990; Collins & Vernon, 1992; Foster et al., 1999). Mafic magma input has been related to arc or back-arc magmatism (Collins & Vernon, 1992; Soesoo et al., 1997; Collins, 1998; Soesoo and Nicholls, 1999; Scheibner & Veevers, 2000). For example, Collins & Vernon (1992) considered major batholiths such as the Kosciusko Batholith to represent the internal zones of magmatic arcs. Collins (1994) suggested lithospheric delamination as the mechanism driving both deformation and magmatism. Any of these settings would imply mafic magma invasion of the crust, extensive crustal melting in the zones above mafic magma chambers, and almost certainly the formation of hybrid magmas and related cumulates (e.g. Castro et al., 1999). Juvenile contributions to the granite batholiths are present in the form of gabbros with low-K affinities (Cousins et al., 1998) and hybrid mafic enclaves. Their relatively minor volumes in present outcrop may be related to efficient trapping of basaltic magmas in a hot mid crust undergoing massive partial melting. Syn-magmatic juvenile contributions to the S-type granites themselves, however, are not easily gauged and have been controversial, with estimates ranging from nil (Chappell & White, 1992) to 10–40% (based on data of Gray, 1984, 1990; Collins, 1996; Keay et al., 1997). The latter compare well with estimates for the Hercynian S-type granites of Europe (e.g. Barbarin, 2000). Tantalizing hints of an early high-temperature stage in S-type magma production are provided by experimental work (Clemens & Wall, 1981). Melting and crystallization of natural S-type compositions require initial crystallization temperatures >900°C, well in excess of the temperatures estimated from mineral thermometry in slowly cooled granites and high-grade metamorphic rocks in the eastern Lachlan Fold Belt. This suggests that S-types magmas could indeed have evolved from high-temperature, relatively mafic magmas via massive assimilation in the mid crust.
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ACKNOWLEDGEMENTS |
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Tod Waight, Sue Keay and Lian-Kun Sha read earlier versions of this paper. Journal reviewers Richard Price and Richard Arculus provided useful comments and suggestions. A big thanks goes to journal editor R. Arculus for his patience. The Perth SHRIMP consortium kindly provided access to the SHRIMP II at Curtin University. Funding for this research was provided by the Australian Research Council.
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FOOTNOTES |
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*Corresponding author. Telephone: 61-3-9479-1274. Fax: 61-3-9479-1272. E-mail: r.maas{at}latrobe.edu.au
Present address: Department of Earth Sciences, University of Queensland, St. Lucia, Qld. 4072, Australia.
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REFERENCES |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTING AND SAMPLESANALYTICAL PROCEDURESRESULTSDISCUSSIONREFERENCES |
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