SHRIMP and electron microprobe chronology of UHT metamorphism in the Napier Complex, East Antarctica: implications for zircon growth at >1,000 C

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1 Contrib Mineral Petrol (2004) 147: 1 20 DOI /s ORIGINAL PAPER Tomokazu Hokada Æ Keiji Misawa Æ Kazumi Yokoyama Kazuyuki Shiraishi Æ Akira Yamaguchi SHRIMP and electron microprobe chronology of UHT metamorphism in the Napier Complex, East Antarctica: implications for zircon growth at >1,000 C Received: 13 August 2002 / Accepted: 26 November 2003 / Published online: 6 February 2004 Ó Springer-Verlag 2004 Abstract Zircons in ultra-high-temperature (UHT) metamorphosed paragneisses from Mt. Riiser-Larsen in the Napier Complex, East Antarctica, were dated by using ion microprobe (SHRIMP) and electron microprobe (EMP). Both SHRIMP and EMP analyses yield consistent Ma age populations for garnet orthopyroxene-bearing paragneiss and leucosomes enclosed within. The peak UHT event was dated at 2480 Ma by SHRIMP analyses on metamorphic zircons from the garnet orthopyroxene paragneiss and those on magmatic zircons from the leucosomes which are interpreted to be formed at syn-uht. As obtained by SHRIMP, the UHT metamorphic event was terminated no later than 2460 Ma. Minor 2520-Ma SHRIMP age suggests either the onset of prograde metamorphism or another high-grade metamorphic event unrelated to the UHT. EMP analyses on metamorphic zircons from sapphirine quartz and osumilitebearing magnesian paragneisses give c Ma ages. Inherited igneous zircon cores of the magnesian paragneisses yield relatively scattered EMP ages ranging over c Ma, suggesting that igneous materials of these ages sourced the protoliths of the paragneisses and that they were deposited during the interval c Ma. Editorial responsibility: I. Parsons T. Hokada (&) Æ K. Misawa Æ K. Shiraishi Æ A. Yamaguchi National Institute of Polar Research, Kaga, Itabashi, Tokyo, Japan hokada@nipr.ac.jp Tel.: Fax: T. Hokada Æ K. Misawa Æ K. Shiraishi Æ A. Yamaguchi Department of Polar Science, the Graduate University for Advanced Studies, Kaga, Itabashi, Tokyo, Japan K. Yokoyama National Science Museum, Hyakunin-cho, Shinjuku, Tokyo, Japan Introduction Ultrahigh-temperature metamorphism (UHT), with temperatures >900 C, is now recognized as a distinct category of higher-temperature granulite-facies conditions (Spear 1993; Harley 1998), and a greater number of UHT terranes and relatively isolated localities have been recently reported (Harley 1998; and references therein). The Archaean Napier Complex in East Antarctica is known as one of the UHT metamorphic terranes, which is characterized by the occurrence of sapphirine + quartz and osumilite (e.g., Dallwitz 1968; Ellis et al. 1980; Grew 1980; 1982; Motoyoshi and Hensen 1989; Harley and Motoyoshi 2000), and inverted metamorphic pigeonite (e.g., Sandiford and Powell 1986; 1988; Harley 1987). For these characteristic features, the Napier Complex has now become known as the best-preserved and the largest recognized UHT terrane; about a km 2 UHT area has been estimated from the occurrence of diagnostic UHT mineral parageneses (Fig. 1A). The highest-grade metamorphic conditions of the Napier Complex are estimated to be >1,000 C for rocks with sapphirine + quartz, based on experimental results (e.g., Hensen and Green 1973; Bertrand et al. 1991). The stability field of osumilite (Motoyoshi et al. 1993; Audibert et al. 1995; Carrington and Harley 1995a; 1995b) and calculations using geobarometers (Harley 1983; 1985; Ellis and Green 1985; Sandiford 1985) constrain the pressure conditions to be less than 1.1 GPa. Recent estimations of the metamorphic conditions using high-al 2 O 3 orthopyroxene coexisting with sapphirine + quartz (Harley and Motoyoshi 2000) and high-ca ternary mesoperthitic feldspar (Hokada 2001), imply that the peak metamorphic temperature exceeded 1100 C (at GPa). A variety of reaction textures involved in sapphirine and osumilite-bearing rocks and geothermobarometry for garnet orthopyroxene equilibria suggest near isobaric cooling from the peak UHT to amphibolite-facies conditions (e.g., Harley 1983; 1985; Harley 1998; Harley and Hensen 1990).

2 2 Fig. 1 Simplified geological map of the Mt. Riiser-Larsen area of the Napier Complex. Star and circles indicate the localities of the samples used in this study. Inset: Geological outline of the Napier Complex and surrounding area in East Antarctica after Sheraton et al. (1987) and Harley and Hensen (1990). The sapphirine + quartz in isograd defining the >1000 C high-grade region is taken from Harley and Hensen (1990). A Amundsen bay, C Casey bay, N Napier Mountains, T Tula Mountains, S Scott Mountains The earliest recorded history of the Napier Complex started with 3800-Ma tonalitic magmatism at Fyfe Hills, Mt. Sones and Gage Ridge as estimated by SHRIMP zircon analyses (e.g., Black et al. 1986a; Harley and Black 1997). Many other rocks were formed or metamorphosed during the period from mid Archaean to earliest Proterozoic (e.g., Sheraton et al. 1987; Harley and Black 1997). In spite of the large number of geochronological data that have been reported during the last two decades, the age of the UHT event is still in argument; Grew and Manton (1979) suggested that the major high-grade (UHT) event occurred at c Ma based on conventional U Pb geochronology, and this interpretation has been supported by several other geochronological studies (e.g., De Paolo et al. 1982; Grew 1998; Asami et al. 1998; 2002; Carson et al. 2002). On the other hand, older (3070) 2900-Ma UHT metamorphism, deformation and magmatism, followed by slow isobaric cooling to upper-amphibolite conditions at Ma were proposed (Black and James 1983; Black et al. 1983a; 1986b; Sheraton et al. 1987). Based on a SHRIMP study, Harley and Black (1997) concluded that the UHT event occurred no earlier than 2840 Ma and possibly at Ma (Harley et al. 2001) in the Tula and Scott Mountains, and that the pre Ma age referred to another high-grade event that happened only in the Napier Mountains in the north. Zircon and monazite have now been recognized as reliable minerals in dating the age of igneous rocks and high-grade metamorphic rocks. Ion microprobe analyses give ages on the <30-lm crystal domains, which can be a powerful tool for distinguishing multiple thermal events recorded in heterogeneous zircon grains (e.g., Williams 1998). The application of the electron microprobe to Th U Pb mineral dating, especially of monazite (e.g., Suzuki et al. 1991; Montel et al. 1996) and zircon, (Suzuki and Adachi 1991) has been explored successfully in several recent papers and appears to be a very promising technique. This method is less precise than that using an ion microprobe, but has proven to be a valuable tool for geochronologists because of its ability to determine U, Th, and Pb concentrations in domains <10 lm in size, which is considerably smaller than the smallest possible spot size of an ion microprobe. We have carried out ion microprobe (SHRIMP) and electron microprobe (EMP) dating of zircon from the UHT metamorphosed paragneisses from Mt. Riiser- Larsen in the Napier Complex. The UHT samples

3 3 analyzed here preserve anhydrous high-grade mineral assemblages including some of garnet, orthopyroxene, sapphirine, quartz, osumilite, and high-ca ternary feldspar. Leucocratic patches or veins are locally developed in the garnet orthopyroxene-bearing paragneiss, which are considered as an anatectic melt formed syn-uht. Oscillatory-zoned zircons in such a domain are interpreted to document the igneous zircon crystallization during the UHT metamorphism. Based on these geochronological data, we present evidence for zircon growth during the UHT metamorphism, and so constrain the age of the UHT metamorphism in the Napier Complex. Fig. 2 Modes of occurrence and mineral textures of sapphirine quartz osumilite-bearing magnesian paragneisses. Mineral abbreviations are the same as in Table 1. A Field photo of the sapphirine quartz-bearing magnesian paragneiss. Layered structure composed of sapphirine orthopyroxene-rich (blue) and quartzo-feldspathic (white) layers is developed. B Field photo of the osumilite-bearing magnesian paragneiss. Osumilite commonly coexists with garnet and orthopyroxene. C Photomicrograph of sapphirine coexisting with quartz in sapphirine orthopyroxene quartz gneiss (sp ). Plane polarized light. D Photomicrograph of osumilite in garnet orthopyroxene-osumilite gneiss (sp. R2302C). Note that a fine-grained (submicron scale) symplectite (labeled Symp ) of cordierite quartz K-feldspar orthopyroxene replaces osumilite. Setting for optical microscope is plane polarized light Geological outline and sample description Sample localities are shown in Fig. 1. Felsic orthogneiss, mafic granulite and a variety of paragneisses of anhydrous high-grade (UHT) mineral assemblages constitute the area. A massive orthogneiss unit constitutes the eastern part, whereas a layered gneiss unit comprises the central-western part. Apparently the orthogneiss unit overlies the layered gneiss. The layered gneisses are composed of orthopyroxene-bearing felsic orthogneiss, garnet-bearing felsic gneiss, two pyroxene-bearing mafic granulite and a variety of paragneisses. These include garnet gneiss, garnet sillimanite gneiss, garnet orthopyroxene gneiss, and other aluminous, siliceous, ferruginous and magnesian gneisses (e.g., Ishizuka et al. 1998). Magnesian paragneisses occasionally include diagnostic UHT mineral assemblages such as sapphirine + quartz and osumilite. We have analyzed zircons in sapphirine quartz and osumilite-bearing magnesian paragneisses (Fig. 2), garnet orthopyroxene-bearing gneiss and associated patch-vein leucosomes (Fig. 3). Constituent minerals of the samples are listed in Table 1.

4 4 Fig. 3 Modes of occurrence and mineral textures of garnet orthopyroxene-bearing paragneiss and leucocratic patches (leucosomes) enclosed within. Mineral abbreviations are the same as in Table 1. A Quartzo-feldspathic layer (sp. R2301A) composed of garnet, orthopyroxene, quartz (bluish color crystals) and antiperthitic plagioclase (whitish color crystal). B Garnet and quartz-rich siliceous layer including leucocratic feldspathic (perthitic alkali feldspar) domains (sp. R2301B). C Photomicrograph of quartzofeldspathic layer of garnet orthopyroxene gneiss (sp. R2301A). Garnet, orthopyroxene, antiperthitic plagioclase, quartz and ilmenite are the main constituents. Biotite occurs around ilmenite. Setting for optical microscope is plane polarized light. D Photomicrograph of siliceous layer including leucocratic feldspathic domains (sp. R2301B). Zircons in such domains are unusually coarse-grained, up to 500 lm. Optical microscope setting is plane polarized light Sapphirine quartz osumilite-bearing magnesian paragneisses Magnesian paragneisses include anhydrous granulite- UHT mineral assemblages. Sapphirine in direct contact with quartz is observed from three samples (TH , TH , TH : hereafter referred to 10813, and 21326, respectively). Two samples (21326, and R C: hereafter as R2302C) contain osumilite as a major constituent. Sapphirine + quartz is stable at temperatures higher than 1,000 C and osumilite at >900 C with pressures less than 1.0 GPa (Fig. 4). Zircon grains in these magnesian paragneisses were analyzed by EMP. Sapphirine orthopyroxene quartz-plagioclase gneiss (10813) This gneiss consists of a sapphirine orthopyroxenerich layer (1 5 cm in thickness) and a quartzo-feldspathic layer (several centimeters to a few meters) (Fig. 2A). The sapphirine-rich layer is composed mainly of sapphirine, orthopyroxene and quartz, whereas the quartzo-feldspathic layer consists mainly of plagioclase, quartz and minor orthopyroxene. Osumilite, sillimanite, biotite and zircon are usually minor. This sample was collected from the same locality as that studied by Harley and Motoyoshi (2000). They estimated metamorphic temperature as >1,120 C from the Al 2 O 3 content of orthopyroxene (12.8 wt%) coexisting with sapphirine and quartz. Suzuki et al. (2001) reported a 2204± 19-Ma Sm Nd internal isochron age for the same locality.

5 5 Table 1 Constituent minerals of sapphirine-quartz/osumilite-bearing magnesian paragneisses and garnet-orthopyroxene-bearing quartzo-feldspathic paragneisses including leucocratic patches. Mineral abbreviations essentially follow Kretz (1983) except sapphirine and osumilite as follows. Qtz: quartz, Pl: plagioclase, Kfs: alkali feldspar, Grt: garnet, Opx: orthopyroxene, Spr: sapphirine, Os: osumilite, Sil: sillimanite, Crd: cordierite, Bt: biotite, Rt: rutile, Ilm: ilmenite, Zrn: zircon, Mnz: monazite, EMP: zircon chemical dating using electron microprobe, SHRIMP: zircon dating using SHRIMP, +: present, ) minor or local Sample Analysis Qtz Pl Kfs Grt Opx Spr Os Sil Crd Bt Rt Ilm Zrn Mnz Sapphirine quartz/osumilite-bearing magnesian paragneiss EMP ) ) ) ) ) EMP ) ) EMP ) ) R2302C EMP ) + ) ) ) Grt Opx-bearing paragneiss with leucosomes R2301A 1 SHRIMP/EMP ) ) ) ) R2301B 2 SHRIMP/EMP + ) + ) ) ) ) ) ) 1 Quartzo feldspathic layer 2 Siliceous layer including patch-vein leucosomes Fig. 4 Phase relations for UHT magnesian metapelites involving sapphirine and osumilite, compiled and modified after experimental results of Hensen and Green (1973), Bertrand et al. (1991), Motoyoshi et al. (1993), Audibert et al. (1995) and Carrington and Harley (1995a, 1995b). Stability fields of sapphirine + quartz (Spr+Qtz) and osumilite (Os) are distinguished by shading Sapphirine orthopyroxene quartz plagioclase gneiss (20713) The sapphirine orthopyroxene-bearing gneiss occurs as a 2 3-m-thick layer in the layered paragneiss unit which locally includes sapphirine, orthopyroxene, garnet and osumilite. The sample consists of quartz, plagioclase, orthopyroxene and sapphirine. Sillimanite, cordierite, K-feldspar and zircon are accessory minerals. Sapphirine is in contact with quartz and orthopyroxene (Fig. 2C). Sapphirine orthopyroxene osumilite quartz gneiss (21326) This gneiss occurs in layered sequences composed of orthogneisses and paragneisses. Siliceous layers (10 20 m thick) include local sapphirine orthopyroxene osumilite-bearing pods, and are composed of quartz, sapphirine, orthopyroxene and osumilite with minor amounts of rutile and zircon. Sapphirine is in contact with quartz. Osumilite is commonly replaced by a symplectite consisting of very fine-grained cordierite, K-feldspar, quartz and orthopyroxene. Garnet orthopyroxene osumilite quartz gneiss (R2302C) This gneiss occurs intercalated with aluminous paragneisses (Fig. 2B). It consists of garnet, orthopyroxene, osumilite and quartz with minor amounts of rutile, zircon and monazite. Osumilite is commonly replaced by a symplectite consisting of very fine-grained cordierite, K-feldspar, quartz and orthopyroxene (Fig. 2D). Garnet orthopyroxene-bearing paragneiss Garnet orthopyroxene-bearing gneiss (20 30 m thick) is typical of layered paragneisses. It is composed of quartzo-feldspathic layers, siliceous layers, and aluminous

6 6 Fig. 5 Re-integrated one-phase feldspar compositions of antiperthitic plagioclase in a quartzo-feldspathic layer (sp. R2301A) and perthitic alkali feldspar in leucosomes (sp. R2301B) of the garnet orthopyroxene gneiss, modified after Hokada (2001). Ternary feldspar solvus isotherms for 0.8 GPa were calculated using Fuhrman and Lindsley (1988). It should be noted that the estimated temperatures and the position of solvus curves were changed from those in Hokada (2001), because of the correction of interchanged Margules parameters presented in the original paper of Fuhrman and Lindsley (1988) (see Kroll et al. 1993) layers that include sapphirine and osumilite. Feldspathic patch-vein leucosomes occasionally are developed in the siliceous layers. Zircons in both the quartzo-feldspathic layers (sp. R A, hereafter as R2301A, Fig. 3A, C) and the siliceous layers including feldspathic leucosomes (sp. R B, hereafter as R2301B, Fig. 3B, D) were analyzed by SHRIMP and EMP. Quartzo-feldspathic layer in garnet orthopyroxene gneiss (R2301A) The major constituents of this layer are quartz, plagioclase, garnet and orthopyroxene (Fig. 3A, C). Ilmenite, biotite and zircon are usually accessory minerals. Garnet forms porphyroblasts up to 1 cm that occasionally include ilmenite and quartz. Orthopyroxene is also subidioblastic up to 5 mm in diameter. Quartz (2 5 mm in diameter) and plagioclase (1 2 mm in diameter) are mainly xenoblastic. Quartz grains are rarely elongated up to 1 cm in length. Antiperthitic exsolution lamellae (10 20 lm in width) are conspicuous in plagioclase grains. Equilibrium temperature estimated from the integrated composition of the antiperthite is >1,060 C (Fig. 5). Biotite occurs locally at the rim of ilmenite or orthopyroxene grains as a secondary phase. Siliceous layer including leucosomes (R2301B) This siliceous layer is composed mostly of quartz and garnet with a subordinate amount of orthopyroxene (Fig. 3B). Quartz is xenoblastic to rounded (<5 mm in diameter). Garnet forms commonly rounded porphyroblasts up to 1 cm in diameter, occasionally as xenomorphic crystals. Opaque minerals (e.g., ilmenite in the host gneiss) are absent in the siliceous layer. Rutile is an accessory. Patch-vein leucosomes composed mainly of perthitic alkali feldspar and quartz occur rarely in the siliceous layer. The alkali feldspar occurs interstitially among rounded quartz grains and has concave or embayed grain boundaries (Fig. 3B, D). Temperature estimated by integrating perthite composition was >950 C (Fig. 5). Garnet near the leucosomes is also sometimes interstitial to quartz grains. Idioblastic, coarse-grained zircon (up to 400 lm) occurs in or near the leucosomes. They occasionally show oscillatory-zoned structure under microscopy. Coarse-grained and irregular-shaped monazite (up to 500 lm) is commonly associated with zircon. Ion microprobe (SHRIMP) chronology Analytical methods U, Th and Pb isotopes in zircons were analyzed using the sensitive high-resolution ion microprobe (SHRIMP- II) at the National Institute of Polar Research, Tokyo, Japan. Analytical technique essentially follows Williams (1998). A 30-lm-diameter analytical spot was used. Standard zircon SL13 (U=238 ppm) provided by the Australian National University was used for the reference value of U concentration in zircon. Pb/U ratios were corrected for instrumental interelement fractionation using the ratios measured on the standard zircon FC1 (1099 Ma; Paces and Miller 1993). Common Pb corrections were based on the measured 204 Pb. Data reduction and processing were conducted using the computer programs SQUID version1 and ISOPLOT version 2 provided by K.R. Ludwig at Berkeley Geochemistry Center of University of California (Ludwig 2001a, 2001b). Analytical errors shown are at 68% confidence levels (1-sigma uncertainties) for each analysis, whereas those calculated for concordia intercept ages and weighted average ages are at 95% confidence levels (2-sigma uncertainties), including the decay-constant error of concordia curve. SHRIMP ages were obtained from zircon grains mounted on epoxy, which were separated from the quartzo-feldspathic layer (R2301A) and the siliceous layer including leucosomes (R2301B) of the garnet orthopyroxene-bearing paragneiss. Because the leucosomes occur locally in the gneiss, zircon grains in the leucosomes (R2301B) were also analyzed on the normal polished thin section in addition to the separated grains from grounded rock specimens. After polishing, analytical positions of samples and standard zircons (SL13 and FC1) were selected under cathodoluminescence imaging (CLI) to assess their internal structure. Figure 6 represents selected CLI for zircons from garnet orthopyroxene quartzo-feldspathic paragneiss (R2301A) and from siliceous layer enclosing leucosomes (R2301B). Zircons

7 7 Fig. 6 Cathodoluminescence images of selected zircon grains analyzed by SHRIMP. Errors on 207 Pb/ 206 Pb ages are 1-sigma confidence levels from the quartzo-feldspathic paragneiss (R2301A) commonly have a dark CLI-luminescent core with scarce structure and a blight CLI-luminescent structureless rim. The bright-cli rim is too narrow to be analyzed by SHRIMP so we have analyzed dark-cli cores. Zircons from the siliceous layer with leucosomes (R2301B) show either oscillatory zoning or homogeneous domain, both domains were analyzed by SHRIMP. SHRIMP ages for quartzo-feldspathic layer in the garnet orthopyroxene gneiss (R2301A) Zircon grains from the quartzo-feldspathic layer of the gneiss commonly have a relatively structureless core with low luminescence and a thin (<20 lm) brightly-luminescent rim (Fig. 6A). These textural features, lacking euhedral or oscillatory-zoned structure, suggest sub-solidus metamorphic crystal growth or recrystallization. 207 Pb/ 206 Pb ages of the zircons range from 2531 to 2450 Ma, and an error weighted regression (n=40, MSWD=1.9) gave an upper concordia intercept at 2486±11 Ma with an imprecise lower intercept age of 773±410 Ma (Fig. 7A). Probability distribution of 207 Pb/ 206 Pb ages indicates three age populations; weighted averages of 207 Pb/ 206 Pb ages gave 2517±5 Ma (n=8), 2481±3 Ma (n=24) and 2464±3 Ma (n=7; Fig. 7B), but no systematic correlation between zircon age, morphology and chemistry was found. The zircon grains have variable U and Th contents (U: 153 4,054 ppm; Th: ppm; see Table 2), and Th/U ratios range from 0.1 to 1.6.

8 8 Fig. 7 Conventional Wetherill concordia plots (A B) and relative probability 207 Pb/ 206 Pb age distribution curves (C D) for SHRIMP zircon analyses of garnet orthopyroxene gneiss and leucosomes enclosed within. ZS analyses on separated zircon grains. TS analyses on zircons in situ of polished thin-section. Error ellipse of each analysis is 1-sigma confidence level, whereas intercept ages and weighted average ages are of 2-sigma uncertainties SHRIMP ages for siliceous layer and leucosome in the garnet orthopyroxene gneiss (R2301B) Most of zircon grains show oscillatory-zoning and occasionally sector-zoning (Figs. 6B and 8) consistent with the magmatic crystallization in the leucosomes. Only one grain (spot 6.1) and a recrystallized domain (spot 2.1) in the grain mount had no distinct internal visible structure under CLI (Fig. 6B). Eleven spots on 6 zircon grains, all having oscillatory-zoned magmatic features associated with the leucosome domains, were analyzed in situ on the polished thin section. Both grain mount and thin section analyses gave 207 Pb/ 206 Pb ages ranging from 2502 to 2449 Ma (zircon analyses on thin section giving Ma age range), and an errorweighted regression (n=38, MSWD=1.3) gave an upper concordia intercept at 2479±9 Ma, with an imprecise lower intercept age of 196±450 Ma (Fig. 7C). Probability distribution of 207 Pb/ 206 Pb ages indicated a single age peak with younger age tail; the weighted average of 207 Pb/ 206 Pb ages was 2482±2 Ma (n=33, including both grain mount and thin section analyses) for the main age group and 2457±10 Ma (n=4) for the younger tail (Fig. 7D). Analyses of zircons occurring in the leucosomes on thin sections ( Ma) yielded a weighted average 207 Pb/ 206 Pb age of 2481±6 Ma (n=11), which is in good agreement with the 2482±2-Ma age obtained above. Two structureless zircon domains (spots 2.1 and 6.1) were also in the same age bracket within the 2482±2-Ma main age population. The younger c Ma age was obtained from the low- CL core (hence higher-u) (Fig. 6; spot 4.2) of the zoned zircons, and its 2461±10-Ma age was younger than the surrounding zoned domain (2482±8 Ma), suggesting that the c Ma core reflects intracrystalline isotopic disturbance. Most of the zircons were characterized by lower U ( ppm) and higher Th (406 1,693 ppm) in comparison with those in the host quartzo-feldspathic layer (R2301A), representing unusually high Th/U ratios of Two homogeneous domains (2.1 and 6.1) had 1, and ppm Th, respectively; their Th/U ratios of 0.2 to 0.4, were markedly lower than those of zoned domains.

9 9 Table 2 SHRIMP U-Pb zircon analyses of garnet-orthopyroxene gneiss and leucocratic patches from the Mt. Riiser-Larsen U Th Th/U Common 206 Pb/ 238 U 207 Pb/ 206 Pb 207 Pb/ 206 Pb 206/238 Spot (ppm) (ppm) ratio 206 Pb (%) ratio ratio age (Ma) R2301A (Grt-Opx gneiss - zircon separates from quartzo-feldspathic layer) ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± R2301B (Grt Opx gneiss zircon separates from siliceous layer including leucosomes) ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

10 10 Table 2 (Contd.) U Th Th/U Common 206 Pb/ 238 U 207 Pb/ 206 Pb 207 Pb/ 206 Pb 206/238 Spot (ppm) (ppm) ratio 206 Pb (%) ratio ratio age (Ma) ± ± ± ± ± ± R2301B(zircon grains associated with lucosomes on thin section) ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± Errors are 1-sigma uncertainties Electron microprobe (EMP) chemical zircon chronology Analytical method Chemical analyses were made on zircon mineral separates and those in situ in normal polished thin sections using an electron microprobe with a wavelength-dispersive X-ray analytical system (JEOL JXA-8800 M) at the National Science Museum, Tokyo, Japan. Monazite grains from two samples (R2302C and R2301B) were also analyzed for comparison with zircon ages. The analytical conditions were 15kV accelerating voltage, 0.5 la probe current (0.2 la for monazite), 2 lm probe diameter (the estimated area of analyses was <6 lm in diameter) and s counting time for U, Th and Pb, respectively. PRZ corrections (modified ZAF) were applied to the analyses. Standard materials for U, Th and Pb were synthetic c UO 3, ThO 2 and natural crocoite (PbCrO 4 ), respectively. Natural and synthesized minerals were used as standards for other elements. Seven elements (Si, Zr, Y, Hf, U, Th, Pb) were analyzed for zircon and 14 elements (P, Si, La, Ce, Pr, Nd, Sm, Gd, Dy, Y, U, Th, Pb, Ca) for monazite. UMa, ThMa, PbMa lines were used in the U, Th and Pb analyses, respectively, and the spectral interferences of the Th, Y and Zr with the PbMa line, and the Th with the UMa line, were corrected. The theoretical basis of electron microprobe dating follows that of the chemical Th U-total Pb isochron method (CHIME) described by Suzuki et al. (1991) and Suzuki and Adachi (1991). Suzuki et al. (1991) obtained both initial PbO content and age from the regression line on the PbO ThO 2 * diagram (ThO 2 * refers to as the sum of ThO 2 and the ThO 2 equivalent of UO 2 ), assuming that initial PbO is homogenously present in the mineral. Recent studies have shown that the initial Pb is negligible in comparison with radiogenic Pb. Hence in the present study, the ages have been calculated with the Fig. 8 A Photomicrograph of the SHRIMP sample prepared for in situ zircon analyses of leucocratic domains (sp. R2301B). Coarsegrained zircon grains occasionally occur associated with feldspathic patches. B C) Cathodoluminescence images of selected zircon grains analyzed in situ by SHRIMP. Oscillatory-zoned zircons with sector zoning give c Ma with Th/U ratios from 1.0 to 3.0. Errors on 207 Pb/ 206 Pb ages are 1-sigma confidence levels assumption that initial Pb is negligible. Ages of zircons and monazites obtained by EMP were compared with the data by SHRIMP to confirm the applicability of this

11 11 Fig. 9 Backscattered electron image of representative zircon grain (z-11 grain from sp ) along with chemical ages analyzed by electron microprobe (EMP). Each fine line or fine dashed line represents the probability age distribution of each analysis. The total probability age distribution curve is integrated from age and 1-sigma error of each analysis. Best-fit ages for zoned core (heavy dashed line) and overgrowth domain (heavy line) are calculated for the probability curve. Scale for total probability and best-fit curves are reduced for visualization EMP method. Although a few percent shifts are inevitable because of machine drift and standard conditions, ages were consistent with each other. We have checked the accuracy of the chemical ages using monazite and zircon grains from Archaean to Quaternary samples. A 994±5-Ma zircon from Antarctica (K. Shiraishi, unpubl. data) and 3019±4-Ma monazite from Australia (Kiyokawa et al. 2002; the age was obtained from the coexisting zircon), both analyzed by SHRIMP, were used as internal age working standards. These internal age standards were analyzed before and after the analyses of unknown samples to assess the daily drift of the analytical conditions, the daily drift being less than 1% fluctuation. Some other problems of chemical U Th Pb dating using the electron microprobe are summarized in Montel et al. (1996). As the PbO content in zircon is too low to obtain U 7Th Pb ages within reasonable error level from each analysis using the electron microprobe, we have calculated the best-fit age for the sum of probability age distribution curve obtained from the age and 1-sigma error of each analysis (Fig. 9) as an alternative to the calculation of isochron ages as used in Suzuki et al. (1991). The error (1-sigma confidence level) given for each analysis includes instrumental counting statistics only, and was approximately 3% age error at a PbO=0.1 wt% level to 5% age error at a PbO=0.05 wt% level. Least-squares modeling was applied to calculate the best-fit age for the sum of probability distribution curves with an assumption that it gives a single age and that it follows a normal distribution. As can be seen in an example of the analyses on zircon grain z-11 from sapphirine orthopyroxene quartz plagioclase gneiss (sp ) in Fig. 9, a zoned core represents an age range of Ma (errors ranging from ±130 to ±330 Ma). We assumed the probability distribution as one age peak with normal distribution, and obtained the best-fit age of 2771 Ma (1-sigma range of the distribution is ±227 Ma). An overgrowth domain showing an age range of Ma (errors ranging from

12 12

13 13 b Fig. 10 Histograms and relative probability age distribution curves for zircon and monazite chemical ages by electron microprobe analyses. Best-fit ages are calculated for total probability curves Table 3 Selected electron microprobe analyses of zircon (z-11 grain) from sapphirine-orthopyroxene-quartz-plagioclase gneiss (sp ) Sample z-11 wt.% zoned core rim SiO ZrO Y2O HfO UO ThO PbO Total cations (O=4) Si Zr Y Hf U Th Pb Total Age (Ma) Error (Ma) Th/U ±80 to ±160 Ma) gave a best-fit age of 2484 Ma (1- sigma range of the distribution is ±128 Ma). In this way, we calculated the best-fit ages on the accumulated relative probability age distribution curves for total analyses of each sample (Fig. 10). Representative zircon chemical analyses of zircon are shown in Table 3. EMP Chemical ages of Spr-Qtz/Os-bearing magnesian paragneisses Four magnesian paragneiss samples (sp , 21326, 20713, R2302C) yielded similar age populations, the best-fit ages giving c Ma, with relatively minor abundances of older ages up to c Ma (Fig. 10). Fig. 11A J show the backscattered electron images of selected zircon grains from the magnesian paragneiss samples along with the analyzed spots and the calculated ages for each domain. Euhedrally-zoned zircon cores yielded relatively scattered c Ma ages (Fig. 11D, G, H, I). Some zircon grains include relatively heterogeneous patchy domains (Fig. 11C, E, F, H) that yielded c Ma ages. Some of the zircon grains were analyzed in situ in polished thin sections to assess their textural relationships with the other metamorphic minerals: for example zircon Z-201 grain from sapphirine orthopyroxene quartz plagioclase gneiss (sp : Fig. 11A, C) is associated with sapphirine and quartz. Zircon Z-275 from garnet orthopyroxene-osumilite-quartz gneiss (sp. R2302C: Fig. 11B, F) is enclosed within osumilite. A few monazite grains were included in the garnet orthopyroxene osumilite quartz gneiss (sp. R2302C), and they gave 2488 Ma (1-sigma range of the distribution is ±39 Ma; Fig. 10G). EMP Chemical ages of garnet orthopyroxene-bearing paragneiss and leucosomes We analyzed chemical U Th Pb zircon ages of the same samples as those analyzed by SHRIMP. All EMP analyses were carried out on thin sections. Zircons from the quartzo-feldspathic layer (R2301A) yielded a best-fit age of 2470 Ma (1-sigma range of the distribution is ±128 Ma; Fig. 10E), and the three age populations (2520, 2480 and 2460 Ma) obtained by SHRIMP cannot be distinguished in EMP analyses because of their relatively large errors compared with SHRIMP analyses. Zircons from the leucosomes in the siliceous layer (R2301B) had euhedral oscillatory-zoned internal structure rarely accompanied by a thin outer rim (Fig. 11K) which was too narrow (<20 lm) to be analyzed by SHRIMP. Zoned core and outer rim gave bestfit ages of 2461 Ma (1-sigma range of the distribution is ±256 Ma) and 2460 Ma (1-sigma range of the distribution is ±134 Ma), respectively (Fig. 10F). Monazites commonly occur in the leucosomes. They show relatively heterogeneous internal structure, and the radiation damage (metamict state) is conspicuously observed in backscattered electron images. We carefully chose the analytical spots in relatively clear domains to avoid the effect of Pb loss, and obtained an age range of Ma. However, the data do not display a symmetric normal age distribution but suggest an asymmetric probability distribution curve (Fig. 10H), so therefore the best-fit age could not be calculated. Discussion Pre-2600-Ma ages The c Ma ages obtained by EMP analyses on euhedrally-zoned zircon cores of the magnesian paragneiss samples are interpreted as igneous crystallization on the basis of zircon morphology. They suggest that magmatic zircons of these ages were in the source of the protolith of the paragneisses. As described above, each individual zircon grain suggests somewhat different igneous core ages, ranging from c Ma (e.g., grain z-47 in sp. R2302C: Fig. 11G) to c Ma (e.g., grain z-22 in sp : Fig. 11H), although these ages are identical within the 1-sigma range of probability distribution. The ages suggest that multiple igneous sources at c Ma may have supplied the sediments to the precursors of the paragneisses. The sedimentary precursors of the magnesian paragneisses are considered to

14 14 have been deposited after the c Ma magmatism and prior to the c Ma metamorphic event, as will be discussed below. Origin of the patch-vein leucosomes in the garnet orthopyroxene gneiss We suggest that the patch-vein leucosomes (R2301B) in the garnet orthopyroxene gneiss are formed by a magmatic process on the basis of their cross-cutting modes of occurrence in the host gneiss, and the euhedral and oscillatory-zoned (and occasionally sector-zoned) morphology of the zircons (Figs. 6B and 8). These zircon grains commonly have an equidimensional shape, and are not elongated like a typical magmatic zircon. The constituent feldspar in the leucosomes is different from that in the host gneiss; perthitic alkali feldspar constitutes the leucosomes, whereas antiperthitic plagioclase occurs in the host quartzo-feldspathic gneiss (Fig. 5). As was mentioned above, hydrous minerals are typically lacking in the leucosomes except one cordierite grain, which could be hydrous, found at the rim of garnet. Estimated from feldspar thermometry, equilibrium temperatures were >950 C (Fig. 5), suggesting that the leucosomes were recrystallized from melt at syn-uht conditions. The occurrence of these leucosomes is restricted to within the garnet orthopyroxene-bearing gneisses; they are not observed in the neighboring felsic gneiss and mafic granulite layers. We deduce that the leucosomes were formed through partial melting of the garnet orthopyroxene-bearing paragneiss hosting them, and that such magmas were crystallized almost in situ to form patch-vein leucosomes. Nevertheless, the evidence of melting is not obvious in the host garnet orthopyroxene gneiss. Experimental investigations using a mineral mixture of quartz feldspar orthopyroxene (excluding garnet and other minor phases) from the same sample (sp. R2301A) imply that the dry solidus of this quartzo-feldspathic gneiss lies between 1,100 and 1,150 C (Hokada and Arima 2001), which is comparable with the peak metamorphic temperatures attained in the study area. Thus, it is possible for the garnet orthopyroxene gneiss to have been partially molten during the UHT conditions, although the detailed processes of partial melting in the gneiss are too complicated to be examined in detail here Ma ages c Fig. 11 Backscattered-electron images of mineral textures involving zircon grains (A B), and those of selected zircon grains along with intra-grain age distributions (C K). A Zircon grain (z201) associated with sapphirine-quartz assemblage in sapphirine-orthopyroxene-quartz gneiss (sp ). B Zircon enclosed within osumilite in garnet orthopyroxene osumilite gneiss (sp. R2302C). Note that a later fine-grained (submicron scale) symplectite of cordierite quartz K-feldspar orthopyroxene replaces the rim of the osumilite grain. c Z201 from sp (as shown in Fig. 8A) with patchy core, bright (higher-u) mantle and dark (lower-u) rim. The heterogeneous patchy core gave an older age (2588 Ma) than the bright mantle (2459 Ma). The outer dark rim contains too low PbO abundance to obtain reasonable ages. D Z83 from sp with 2723-Ma zoned core, 2530-Ma mantle and 2524-Ma overgrowth domain. E Z86 (sp ) with heterogeneous patchy core, which gave a similar age (2467 Ma) to the overgrowth domain (2464 Ma). F Z275 from sp. R2302C (as shown in Fig. 8b) with euhedral patchy core giving 2509 Ma, bright mantle giving 2469 Ma, and outer dark domain was too low regarding PbO to give a meaningful age. G Z247 from sp. R2302C with euhedrallyzoned core (2975 Ma) replaced by overgrowth domain (2443 Ma). H Z22 from sapphirine orthopyroxene osumiliteapphirine orthopyroxene osumilite quartz gneiss (sp ) with 2650 Ma euhedrally-zoned core, surrounded by 2442-Ma heterogeneous patchy domain. Outer rim gives 2508 Ma. I Z50 from sapphirine orthopyroxene quartz gneiss (sp ) with euhedrally-zoned core of 2892 Ma and an outer rim of 2508 Ma. J Relatively homogeneous grain z25 from sp giving 2485 Ma with local 2288-Ma domain. K Coarse-grained, euhedrally-zoned zircon (z1) from siliceous layer of the garnet orthopyroxene gneiss including leucosomes (sp. R2301B) giving 2482 Ma, with thin outer rim giving 2491 Ma A consistent Ma age cluster was obtained from both SHRIMP and EMP analyses. We found the daily drift of EMP analytical results to be less than 1%, and the best-fit U Th Pb EMP ages ranging Ma are in almost the same age bracket with the SHRIMP concordia intercept ages of 2486±11 Ma (R2301A) and 2479±9 Ma (R2301B). Subhedral or rounded zircons (Fig. 6A) from the garnet orthopyroxene-bearing quartzo-feldspathic paragneiss (R2301A) are interpreted as being formed through metamorphic crystallization on the basis of their morphology. Three age peaks are suggested from the probability distribution of 207 Pb/ 206 Pb ages ranging from 2520 to 2460 Ma, and the most intense 207 Pb/ 206 Pb age population is calculated from the weighted average 207 Pb/ 206 Pb age to be 2480±3 Ma. Magmatic zircons from the syn-uht leucosomes (R2301B) give an weighted average 207 Pb/ 206 Pb age of 2482±2 Ma, which is coincident with that obtained from the host gneiss (R2301A). We have suggested that both the host quartzo-feldspathic gneiss and the leucosomes enclosed within have been metamorphosed or formed at >1,000 C UHT conditions and that the euhedrally-zoned zircon grains associated with the leucosomes were probably crystallized from the UHT melt. The 2482±2 Ma age estimated from igneous zircons from the UHT leucosomes, along with the dominant zircon age population (2480±3 Ma) obtained from the host quartzo-feldspathic gneiss, presumably correspond with the age of the UHT metamorphic event. The weighted average 207 Pb/ 206 Pb SHRIMP age of 2517±5 Ma, which predating the 2480-Ma UHT event, was obtained only from the host gneiss (R2301A). This 2520-Ma age is also interpreted as a metamorphic event on the basis of zircon morphology. Although no inherited metamorphic minerals that might give information of the pre-uht events have been identified in the gneiss, the 2520-Ma zircon age can be interpreted as either the

15 15 onset of prograde zircon recrystallization or another high-grade metamorphic event unrelated to the UHT event. Two episodes of intense ductile deformation (D1 and D2) have been identified as the main fabric-forming events in the Napier Complex (e.g., Black and James 1983; Sheraton et al. 1987). The gneisses studied are heterogeneously deformed and it is not possible to identify deformation features corresponding to D1 or D2. However, there is a possibility that the 2520 and 2480-Ma ages might correspond to the high-grade D1 and D2 events, respectively. The younger age population gave a weighted average 207 Pb/ 206 Pb SHRIMP age of 2464±3 Ma for the garnet orthopyroxene paragneiss (sp. R2301A), which represents the youngest metamorphic age population recorded in the gneiss. The garnet orthopyroxene gneiss occasionally includes biotite, which is understood to be a retrograde phase on the basis of its replacive relationship to orthopyroxene and ilmenite. The UHT gneisses in the Napier Complex have commonly suffered gentle upright deformation corresponding to D3 (Black and James 1983) which, seen on a regional scale (wave length is several hundred meters to kilometer scale) but not recognized on a local outcrop or hand-specimen scale. Sheraton et al. (1987) reported that the D3 deformation is associated with the alignment of biotite. These facts suggest that the zircon recrystallization at 2460 Ma possibly corresponds to the D3 retrograde event associated with external fluid influx. Zircons separated from the sample with leucosomes (R2301B) also suggest a minor group of 2457±10 Ma weighted average 207 Pb/ 206 Pb SHRIMP ages; such younger ages were not obtained from zircons analyzed in situ within the leucosomes using the thin section. Although no hydrous phase has been identified under microscope except one cordierite grain, the rock may have been affected by the retrograde event. Magnesian paragneiss samples also give c Ma U Th-total Pb age populations. Some of the zircon grains occur associated with or enclosed within sapphirine-quartz and osumilite (Fig. 11A, B). These