Two-stage partial melting and contrasting cooling history within the Higher Himalayan Crystalline Sequence in the far-eastern Nepal Himalaya

Lithos 134-135 (2012) 1–22

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Two-stage partial melting and contrasting cooling history within the Higher Himalayan Crystalline Sequence in the far-eastern Nepal Himalaya Takeshi Imayama a,⁎, Toru Takeshita b, Keewook Yi c, Deung-Lyong Cho d, Kouki Kitajima e, Yukiyasu Tsutsumi f, Masahiro Kayama g, Hirotsugu Nishido g, Tasuku Okumura g, Koshi Yagi h, Tetsumaru Itaya g, Yuji Sano i a

Center for Chronological Research, Nagoya University, Nagoya 464-8602, Japan Department of Natural History Sciences, Hokkaido University, Sapporo 060-0810, Japan c Geochronology Team, Korea Basic Science Institute, Chungbuk 363-883, Republic of Korea d Korea Institute of Geoscience and Mineral Resources, Taejon 305-350, Republic of Korea e Department of Geoscience, University of Wisconsin, WI 53706, USA f Department of Geology and Paleontology, National Museum of Nature and Science, Tokyo 169-0073, Japan g Research Institute of Natural Sciences, Okayama University of Science, Okayama 700-0005, Japan h Hiruzen Institute for Geology & Chronology, Okayama 703-8252, Japan i Center for advanced Marine Research, Ocean Research Institute, The University of Tokyo, Tokyo 164-8639, Japan b

a r t i c l e

i n f o

Article history: Received 8 March 2011 Accepted 9 December 2011 Available online 19 December 2011 Keywords: Partial melting Zircon geochronology Nepal Himalaya P–T–t path Large-scale thrust within the HHCS

a b s t r a c t The timing of partial melting and the pressure–temperature (P–T) paths in the High Himalayan Crystalline Sequence (HHCS) in far-eastern Nepal has been investigated using zircon chronology, rare earth element (REE) compositions, and P–T pseudosection analysis. Zircon from migmatites formed during Himalayan thermal events displays inherited magmatic core overgrown by two generations of metamorphic rims. The new rims are distinguished on the basis of their Tertiary ages, low MREE contents, and low Th/U ratios. The inner zircon rims from Sil + Grt + Bt + Kfs + Pl + Qtz and Ky + Sil + Grt + Bt + Ms + Pl + Qtz migmatites at different structural level of the HHCS display ages of c. 33–28 Ma (Early Oligocene) and c. 21–18 Ma (Early Miocene): these rims are characterized by flat MREE to HREE patterns and were overgrown by partial melt through muscovite dehydration melting under the stability of garnet, which occurred at P = c. 7–10 kbar and T = c. 730–780 °C, and at P = c. 8–14 kbar and T = c. 720–770 °C, respectively. The outer zircon rims are relatively enriched in HREE with respect to the inner rims and were overgrown at c. 27–23 Ma (Late Oligocene) and at c. 18–16 Ma (Early Miocene) during melt crystallization accompanying breakdown of garnet at P = c. 4–7 kbar and T = c. 650–725 °C. Early Miocene Ms–Bt leucogranites with two successively overgrown zircon rims at c. 18.3 ± 0.3 Ma and c. 16.3 ± 0.2 Ma were intruded into Early Oligocene migmatite hosts. Microstructural observations and the corresponding P–T conditions associated with the two generations of zircon rims indicate that the Early Oligocene and Early Miocene migmatites show relatively isobaric and nearly isothermal P–T paths during exhumation, respectively. The inferences are consistent with higher average cooling rates for the Early Miocene (c. 30–40 °C/My) than the Early Oligocene (c. 15–25 °C/My) migmatites, inferred from peak-T conditions and FT (c. 6 Ma for both migmatites) and U–Pb zircon ages. The P–T–t paths of the two migmatites indicate that burial of the Early Miocene migmatites has been coeval with exhumation of the Early Oligocene migmatites, implying the formation of large-scale thrust within the HHCS. © 2011 Elsevier B.V. All rights reserved.

1. Introduction The Himalaya orogen is regarded as one of the best examples of an active continental collision belt where Indian and Asian plates have collided since c. 50 Ma (Searle et al., 1997; Zhu et al., 2005). The

⁎ Corresponding author. Tel.: + 81 52 789 2579; fax + 81 52 789 3092. E-mail address: [email protected] (T. Imayama). 0024-4937/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.lithos.2011.12.004

Higher Himalayan Crystalline Sequence (HHCS) is mainly composed of amphibolite to lower granulite facies gneisses, migmatites, and leucogranites. The HHCS is thrust southwards over the low-grade Lesser Himalayan Sequence (LHS) along the Main Central Thrust (MCT) at the base (Fig. 1, Hodges, 2000). The extensional South Tibetan Detachment (STD) separates the HHCS from the overlying Tethys Himalayan Sequence of Early Paleozoic to Early Tertiary age at the top (e.g. Burchfiel et al., 1992). The opposing-sense movements along the MCT and the STD simultaneously operated at least during the Early Miocene, and are believed to have caused the southward extrusion of the entire

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T. Imayama et al. / Lithos 134-135 (2012) 1–22

80oE

70oE

90oE

Nanga Parbat

35oN

Zanskar

TIBET ST

Tethys Himalayan Sequence Higher Himalayan Crystalline Sequence Lesser Himalayan Sequence Sub-Himalayan Sequence Ophiolites STD: South Tibetan Detachment MCT: Main Central Thrust MBT: Main Boundary Thrust MFT: Main Frontal Thrust

D

MC T

30oN

Sutlej Valley Garhwal Kumaun

INDIA 25oN

70oE

30oN

Kailas Indus-Yarlung Suture

Everest Fig. 2

MBT MFT

ca.500km

Kathmandu 80oE

35oN

STD MCT Bhutan

Arun

Sikkim 90oE

25oN

Fig. 1. Simplified geological division of the Himalaya. Box in the map shows location of Fig. 2. Modified after Imayama et al. (2010) and references therein.

HHCS in a framework of midcrustal channel flow (Beaumont et al., 2001, 2004; Grujic et al., 2002; Jamieson et al., 2004; Searle et al., 2003). Within the HHCS in the Himalaya, sillimanite-grade metamorphism overprints on kyanite-bearing rocks: these are generally interpreted as two discrete tectono-thermal events that operated in a specific volume of crust at different times (Hodges, 2000; Godin et al., 2006, and references therein). According to these authors, the radiometric ages in the HHCS indicate that Late Eocene to Oligocene kyanite-grade metamorphism related to crustal thickening (Eohimalayan, M1 stage) was extensively overprinted by an Early to Middle Miocene high-temperature thermal event (Neohimalayan, M2 stage). The later Neohimalayan event was mostly associated with the leucogranite generation (e.g. Searle et al., 2003). Alternatively, Jamieson et al. (2004) regarded these metamorphic episodes as being a single diachronous event within the HHCS. They argued that Eohimalayan and Neohimalayan metamorphisms were recorded by crustal rock volumes that reached peak-P in the early stages of collision and peak-T during exhumation, respectively: these rocks were tectonically assembled by the activities of main structures, such as the MCT and the STD. However, the spatial relationship among the P–T paths of rocks and the relation to the shear zone within the HHCS are still uncertain to constrain these models, and it is necessary to link reliable P–T information to the radiometric ages in the HHCS. Zircon readily (re-)crystallizes during melting and metamorphism, and the U–Pb system that has high-closure temperature is one of the most important geochronometries to constrain the timing of multiple melting and high-temperature metamorphisms in the crust (e.g. Harley et al., 2007). Here we present the U–Pb ages for the polyphase growth in zircons preserved in leucogranites and metamorphic rocks from the lower portion, close to the MCT, to the top of the HHCS in the far-eastern Nepal Himalaya, where the spatial distributions of P–T conditions (Imayama et al., 2010) and the highstrain shear zone (Goscombe et al., 2006) in the HHCS have been well documented. These age data have been integrated with REE pattern of zircon, microstructural observations, and P–T analyses using pseudosection to correlate the ages with specific reactions during melting or high-grade metamorphism. This study presents new findings of the multiple zircon growth in Himalayan rocks that has been remained unexplained. Furthermore, in order to infer the exhumation processes of metamorphic rocks, K-Ar biotite and fission track (FT) zircon ages from these migmatites were also investigated. Finally,

we discuss the tectonic models for the Himalaya orogens, suggesting the multiple melting of the high-grade gneisses and the role of the shear zone played in the formation of the HHCS. 2. Geological setting and previous age constraints Far-eastern Nepal (study area) is located approximately 80 km east of the Everest area and 50 km west of the Sikkim area (Fig. 1). The Tamor–Ghunsa section in far-eastern Nepal is a NE–SW transect, which crosses the northern part of the antiformal Tamor Khola tectonic window of the LHS at the bottom. The lower part of the section corresponds to a brittle–ductile shear zone known as the MCT zone that is bounded by the MCT at the top; the middle to upper part of the section corresponds to the main part of the HHCS (Fig. 2). From the MCT zone to the middle parts of HHCS, the section is characterized by an inverted metamorphic zonation with a sequence of garnet-in, staurolite-in, kyanite-in, sillimanite-in, muscovite-out, and cordierite-in isograds (Fig. 2: Imayama et al., 2010). The LHS is predominantly composed of low-grade metasediments (Taplejung Group: Schelling, 1992), consisting of biotite–muscovite–chlorite phyllites, quartz-rich metasandstones, and minor amphibolites (Fig. 2). The HHCS is mainly composed of kyanite–sillimanite grade paragneisses, migmatitic gneisses, orthogneisses, and subordinate calcareous gneisses and metabasites (Fig. 2). Kyanite-bearing rocks, compared to sillimanite-bearing rocks, are located at the relatively lower structural level of the HHCS. In migmatites that include both kyanite and sillimanite, (fibrolitic) sillimanite directly replaces rim of garnet porphyroblasts including kyanite, indicating that kyanitegrade metamorphism was partly overprinted by sillimanite-grade metamorphism in the HHCS (Imayama et al., 2010). Throughout the Tamor–Ghunsa section, the foliation roughly shows homoclinal structure in both HHCS and LHS, which mostly strikes NW–SE and dips NE at moderate angles (Imayama et al., 2010). The MCT zone has variable thickness up to one to several kilometers between the MCT and lower MCT (Fig. 2; Imayama and Arita, 2008; Schelling, 1992), which consists of garnet ± staurolite schist, mylonitic augen gneisses, and minor amphibolites. The lower MCT (LMCT) is the single brittle thrust fault placed at the base of the Sisne Khola Augen Gneisses (Fig. 2, e.g. Paudel and Arita, 2002; Searle et al., 2008). The MCT zone is characterized by phyllites containing garnet with S-shaped inclusion trails and blastomylonitic

T. Imayama et al. / Lithos 134-135 (2012) 1–22

3

(a)

(b)

Fig. 2. (a) Geological map and (b) cross section along the Tamor–Ghunsa transect of far-eastern Nepal, showing mineral isograds and location of samples used for geochronology. Modified after Schelling (1992), Goscombe et al. (2006), and Imayama et al. (2010). Bt: biotite; Ms: muscovite; Tur: tourmaline; Grt: garnet; Sil: sillimanite; Ky: kyanite; Crd: cordierite; Cpx: clinopyroxene; Chl: chlorite; St: staurolite; Qtz: quartz; Kfs: K-feldspar. Mineral abbreviations after Kretz (1983).

augen gneiss with euhedral to subhedral K-feldspar and albite porphyroclasts (Imayama and Arita, 2008). Asymmetric deformational structures such as a mica fish in phyllites and ribbon quartz in quartz-rich rocks around the MCT show a dominantly top-to-thesouthwest sense of shear. Goscombe et al. (2006) described the brittle–ductile shear zone, well above the MCT, in the middle parts of the HHCS in far-eastern Nepal, which has been called the High Himal Thrust (HHT) (Fig. 2). The HHT consists mainly of mylonitic (garnet)–sillimanite–biotite quartzfelspathic and pelitic gneisses of c. 400 m thick. The HHT essentially worked as a thrust fault, and then it was reactivated as an extensional normal fault, based on the occurrences of normal-sense shear bands consisting of fine-grained sillimanite + biotite ± gedrite overprinting the main foliation formed by the reverse-sense of shearing (Goscombe et al., 2006). The spectrum of the detrital zircon ages for the HHCS has the broad peaks at c. 1.3–0.6 and c. 2.7–2.3 Ga in the metasediments,

and c. 0.5 Ga for inherited zircons in the orthogneisses. In contrast, the LHS mainly yields detrital zircons ages ranging between c. 2.6 and 1.6 Ga with a prominent peak at c. 1.8 Ga (e.g. DeCelles et al., 2000; Martin et al., 2005; McQuarrie et al., 2008). The protoliths of the HHCS are also characterized by the higher εNd (0) values of − 19 to −2, than those (−27 to −16) of the LHS (e.g. Imayama and Arita, 2008; Parrish and Hodges, 1996; Robinson et al., 2001). Therefore, the MCT mapped in eastern Nepal by Schelling (1992) actually corresponds to a large-scale protolith boundary that was preferentially affected by deformation leading to development of the MCT zone. The kyanite-grade metamorphism has been dated by U–Pb monazite ages at c. 32 Ma as a minimum age for the M1 Eohimalayan event, which affected the HHCS in the Everest area (Simpson et al., 2000). Cottle et al. (2009a) have recently reported that early sillimanitegrade metamorphism occurred at least as early as c. 39 Ma in the Everest area, based on U–Pb monazite ages. The U–Th–Pb ages of

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xenotime, monazite, and zircon in the high-grade gneisses in the Everest area constrain the sillimanite-grade metamorphism, referred to as the M2 Neohimalayan event, to have mainly occurred at c. 26–16 Ma under P–T conditions of c. 6–8 kbar and c. 600–750 °C (Catlos et al., 2002; Cottle et al., 2009a; Searle et al., 2003; Simpson et al., 2000; Viskupic et al., 2005). In Sikkim, Sm–Nd ages of garnet cores and rims in kyanite–sillimanite migmatites indicate garnet growth during peak metamorphism at c. 23 Ma and subsequent melting at c. 16 Ma during isothermal decompression at near peak temperatures of c. 750 °C (Harris et al., 2004). The radiometric ages from the LHS and the base of the HHCS along the MCT have been interpreted as metamorphic ages related to thrusting. In the Everest area, initial movement on the MCT is inferred to have occurred at c. 24–17 Ma based on 40Ar/ 39Ar hornblende ages that indicate cooling ages through c. 500 °C (Hubbard and Harrison, 1989) and 208Pb/ 232Th ages of monazite inclusions in syn-kinematic garnet that grew at peak metamorphic temperature (Catlos et al., 2002). The monazite ages in rocks near the MCT become progressively younger southward up to c. 15–10 Ma, suggesting that the movement on the MCT possibly continued until Late Miocene (Catlos et al., 2002). On the other hand, the timing of the extensional movement on the STD in the Everest area is constrained by 40Ar/ 39Ar hornblende ages of c. 20–19 Ma and U–Pb titanite ages of c. 20 Ma from amphibolites below the detachment (Hodges et al., 1992). Based on the U– Th–Pb ages of xenotime, monazite, and zircon from deformed leucogranite sill near the STD, ductile shearing to brittle faulting related to the movement on the STD was active from c. 20 Ma to c. 13 Ma in the Everest area (Cottle et al., 2007, 2009a; Hodges et al., 1998; Leloup et al., 2010; Murphy and Harrison, 1999; Searle et al., 2003). The leucogranites (Jannu Granites, Schelling, 1992) in fareastern Nepal contain the following mineral assemblage: biotite + muscovite + quartz + K-feldspar + plagioclase ± tourmaline ± garnet ± sillimanite (Fig. 2). The U–Th–Pb ages of xenotime, monazite, and zircon from leucogranite plutons mainly range between c. 24–17 Ma in eastern Nepal (Catlos et al., 2002; Harrison et al., 1999; Hodges et al., 1998; Jessup et al., 2008; Murphy and Harrison, 1999; Schärer, 1984; Simpson et al., 2000). Multiple generations of leucogranite sill and dykes are evidenced by the U–Pb age distributions of zircon and monazite at c. 16–12, c. 26–24, and c. 32–30 Ma in the Everest area (Cottle et al., 2007, 2009a; Viskupic et al., 2005).

3. Metamorphic history and sample description along the Tamor– Ghunsa section The P–T conditions associated with the inverted sequence in the Tamor–Ghunsa section have been constrained by average P–T method and geothermobarometry (Imayama et al., 2010). Metamorphic peak-T conditions increase upwards from c. 570 °C to 670 °C from the LMCT to the base of the HHCS, through a significant temperature increase to c. 740 °C near the muscovite-out isograd, and reach roughly isothermal conditions (c. 710–810 °C) in the middle HHCS (Imayama et al., 2010). Metamorphic pressure conditions at peak-T show maximum pressure of c. 11 kbar at the base of the HHCS, which decreases southward to c. 6 kbar in the garnet zone and northward to c. 7 kbar in the Ky ± St zone (Imayama et al., 2010). The high field pressure gradient of c. 1.2–1.6 kbar/km across the MCT could be caused by ductile extrusion along the MCT (Imayama et al., 2010). Subsequently, pressure increases up to secondary peak of c. 10 kbar near the muscovite-out isograd, and then apparently decreases upwards from c. 8–10 to c. 5 kbar across the HHT (Imayama et al., 2010). In the Everest area, the right-way-up isograds, changing from the Sil–Ms/Kfs to the St zones towards the top, have been mapped near the STD, and staurolite schists have been observed as float at the top of the HHCS (Jessup et al., 2008).

Seven samples for zircon U–Pb dating in far-eastern Nepal were collected along the Tamor–Ghunsa section (Fig. 2). Hereafter they will be described from the bottom to the top of the section. The highly sheared mylonitic augen gneiss (M1806) in the MCT zone (Fig. 2) shows the assemblages of biotite + muscovite + augen K-feldspar + plagioclase + quartz with accessory zircon and monazite. A few leucosomes that contain biotite+ muscovite+ quartz + plagioclase ± tourmaline assemblages are found near the MCT, but the evidence for melting is not found in the low-grade LHS rocks. Coarse-grained garnet–mica gneiss (H1205) in the HHCS directly above the MCT (Fig. 2) is collected to investigate the timing of slip along the MCT. Sample H1205 contains garnet, biotite, muscovite, plagioclase, and quartz with minor amounts of tourmaline, ilmenite, rutile, and zircon. According to Imayama et al. (2010), the P–T paths inferred from garnet zoning and pseudosection method indicate that sample H1205 constituting part of hanging wall of the MCT shows a clockwise path, with increasing temperature path (c. 590 °C to 640 °C) during both loading and decompression with the peak pressure of c. 11–12 kbar. The metamorphic evolution is quite different from that recorded by the staurolite-bearing schist in footwall of the MCT, which experienced a nearly isothermal loading path, where the pressure increases from c. 7 to 9–10 kbar was accompanied by a small temperature increase from c. 550 to 620 °C (Imayama et al., 2010). In order to understand the spatial relationship among the P–T–t paths of rocks within the HHCS, three migmatites (samples H2101, H2908, and H2710) at different structural level of the HHCS are collected along this section (Fig. 2). Kyanite–sillimanite migmatitic gneiss (H2101) is a stromatic migmatite that occurs at the middle HHCS near the muscovite-out isograd (Fig. 2). Sample H2101 contains kyanite, (fibrolitic) sillimanite, garnet, biotite, muscovite, plagioclase, and quartz with minor monazite, zircon, rutile, and ilmenite. Kfeldspar is restricted only in leucosome. Kyanite in the migmatites occasionally occurs within a cm-sized leucosome, implying that the melting started within the kyanite stability field. Garnet porphyroblasts include kyanite, biotite, muscovite, quartz, plagioclase, and rutile. Plagioclase locally occurs as antiperthite (Fig. 3a). Fine-grained fibrolite + biotite ± quartz aggregates, which replaced rim of garnet (Fig. 3b) and cut foliation-forming coarse-grained biotite and muscovite (Fig. 3c), suggest that the growth of the fibrolite took place at retrograde stage. Sillimanite migmatitic gneiss (H2908) above the muscovite-out isograd (Fig. 2) in the HHCS contains (fibrolitic) sillimanite, garnet, biotite, plagioclase, K-feldspar, and quartz with minor monazite, zircon, rutile, and ilmenite. Plagioclase locally occurs as antiperthite. Garnet porphyroblasts are euhedral, and include biotite, plagioclase, K-feldspar, and quartz. Coarse-grained sillimanite occurs in the matrix, and is locally replaced by fine-grained fibrolite. Both foliated and massive leucogranites with thickness less than a meter are commonly intercalated with the migmatite at the middle HHCS. Several tens of cm thick leucogranite dyke (H2708) occurs within the sillimanite migmatitic gneiss (H2710) directly below the HHT (Figs. 2 and 3d), and dated to constrain the timing of melting. The leucogranite (H2708) contains biotite, muscovite, K-feldspar, quartz, and plagioclase with minor xenotime, monazite, and zircon. Sillimanite migmatitic gneiss (H2710) contains sillimanite, garnet, biotite, K-feldspar, plagioclase, and quartz with minor monazite, zircon, rutile, and ilmenite. A lamellar intergrowth of quartz and garnet is developed at the outer parts of large garnet (Fig. 3e). Garnet was partly replaced by green biotite and quartz along the fracture in the crystal (Fig. 3e, f), and also fine-grained fibrolite + biotite + quartz aggregates surrounded by recrystallized quartz replaced rim of garnet (Fig. 3f). Further, the intergrowth of plagioclase and biotite is developed at the rim of garnet (Fig. 3g), and zircon often occurs in the replacement texture at garnet rims (Fig. 3g). Retrograde fine-grained muscovite + quartz ± fibrolite aggregates occasionally occur in leucosome, and locally replace K-

T. Imayama et al. / Lithos 134-135 (2012) 1–22

(a) H2101

(b) H2101 Kfs

Pl Ky

Ms

Bt

5

(c) H2101 Fib

Bt Pl Qtz Qtz

Kfs

Bt Fib

Bt

Grt

Bt

Fib

Fib Qtz

Qtz Bt

1mm

0.5 mm

(d)

0.5 mm

(e) H2710

Grt-Cpx-Hbl metabasite

Qtz Bt Pl

Sil-Grt-Bt migmatitic gneiss (H2710)

Qtz Grt

Sil-Bt gneiss

Bt

Bt

Ms-Bt leucogranite (H2708) Bt

Qtz Grt

1mm

50 cm

(f) H2710

(g) H2710

(h) H2710 Pl

Qtz

Bt

Bt

Qtz

Qtz

Bt

Bt

Fib

Kfs

Grt

Bt

Bt Qtz

Bt Pl

Qtz Fib Bt

Ms

Grt

Ilm

Zrn

Grt

Zrn 0.5 mm

1mm

0.5 mm

Fig. 3. Photomicrographs of metapelites and outcrop photograph from the Tamor–Ghunsa section, far-eastern Nepal Himalaya. (a) Antiperthite showing the exsolution of Kfeldspar in sample H2101. Coarse-grained biotite, muscovite, and kyanite also occur. (b) Garnet and (c) coarse-grained biotite in sample H2101 replaced by the fine-grained fibrolite + biotite ± quartz aggregate. (d) Outcrop photograph showing sillimanite migmatitic gneiss H2710 intruded by muscovite–biotite leucogranite dyke H2708. Note that the leucogranite dyke cuts the foliation in the migmatitic gneiss. (e) A lamellar intergrowth of quartz and garnet at the outer part of large garnet (left side), which is partly replaced by green biotite and quartz along the fracture in the crystal (right side) in sample H2710. (f) Fine-grained fibrolite + biotite + quartz aggregates surrounded by recrystallized quartz replacing rim of garnet in sample H2710. (g) Intergrowth of plagioclase and biotite at the reaction rim of garnet in sample H2710. (h) Fine-grained muscovite replacing K-feldspar in matrix in sample H2710. Pl: plagioclase; Fib: Fibrolite; Zrn: zircon. See captions of Fig. 2 for abbreviations of other minerals.

feldspar in the matrix (Fig. 3h). Myrmekite is developed along Kfeldspar grain boundary. Granitic orthogneiss (H90-30) which occurs at the higher structural level of the HHCS (Fig. 2) is intruded into biotite–sillimanite migmatites and quartzo-feldspathic gneisses with minor calcsilicate. Considerable debate has existed regarding whether the origins of the augen gneiss in the MCT zone (H1806) and the granitic gneiss (H90-30) at the upper HHCS are similar or different (e.g. Cottle et al., 2009b), and thus the zircon from these gneisses has been dated to infer their protolith ages. Sample H90-30 includes the assemblages of biotite, muscovite, augen K-feldspar, plagioclase, and quartz with accessory zircon, apatite, and monazite. 4. P–T pseudosection In order to constrain quantitative P–T paths of anatectic rocks in far-eastern Nepal, the P–T conditions at partial melting and formation of retrograde microstructures are inferred from the pseudosection.

The forward modeling using the pseudosection predicts the phase relations depending on a bulk-rock composition, which can constrain the reactions that operated during metamorphism and melting, and permit the comparisons with the observed phase relation (e.g. Powell et al., 1998). Recent improvements in thermodynamic database, in particular solid solution models including silicate-melt, allow the modeling of pseudosection for anatectic rocks (White et al., 2007). The pseudosections for kyanite–sillimanite migmatitic gneiss (H2101) and sillimanite migmatitic gneiss (H2710) from the Tamor– Ghunsa section were calculated in the chemical system NCKFMASTH using the Perplex_X program (Connolly, 2005) with the internally consistent thermodynamic data set (Holland and Powell, 1998; updated in 2002). Bulk rock compositions used in the pseudosection calculations (Fig. 4) were analyzed by X-ray fluorescence analysis (XRF) at Nagoya University. The modeled phase relationship for aluminous metapelites (sample H2101) near muscovite-out isograd is shown in Fig. 4a. The solidus of the system is located at T = c. 650 °C, and the muscovite-out

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(a)

NCKFMASHT (+Pl +Qtz)

14

Na2O CaO K2O FeO MgO Al2O3 SiO2 H2O TiO2 wt% 1.34 0.94 3.61 9.40 2.63 17.8 62.0 1.60 0.87

Thermocalc average

kfs ms grt ru melt

kfs grt ky ru melt (-pl)

kfs ms grt ky ru melt

ms pa bt grt ru bt kfs ms grt ru melt bt ms grt ru melt

Pressure (kbar)

MeltMelt+

KfsKfs+

bt ms grt ru

bt ms grt ky ru

8

bt ms grt bt ms grt ky ru H2O ky ru ilm bt ms g ky ru ilm H2O bt ms grt ky ilm bt ms grt ky st ilm H2O bt ms g st ilm H2O bt ms grt bt ms ky ilm H2O grt st ilm bt ms grt t ky st Ilm s gr O bt milm H2 sil bt ms ky m O st ilm st Il H2 bt ms ky st ilm H2O sil Ilm ms t bt sil s s bt ms sil m bt ilm H2O melt

bt ms grt sil ru ilm melt

bt ms grt ky ru melt

bt kfs ms grt sil ru melt

bt ms grt ky ru ilm melt

kfs grt sil ru melt

bt kfs grt sil ru melt

bt ms grt ky ilm H2O melt grt bt ms elt m sil ilm bt ms grt sil ilm H2O melt bt ms sil ilm melt

bt sil ilm melt

bt kfs grt sil ru ilm melt

bt kfs ms grt sil ilm melt

650

kfs grt sil ru ilm melt

bt kfs grt sil ilm melt

d- crd kfs grt sil ru melt Cr rd+ crd kfs grt sil ilm melt C

bt grt sill ilm melt

bt crd kfs grt sil ilm melt crd kfs grt ilm melt

OpxOpx+

bt crd kfs grt ilm melt

bt crd grt sil ilm melt bt sil bt crd sil ilm melt ilm H2O melt bt crd kfs sil ilm melt bt crd sil ilm H2O melt bt crd kfs ilm melt

H2101

Pentavariant

bt ms grt ky ilm melt

bt ms sil ilm H2O

4 600

Quadrivariant

bt kfs grt ky ru melt

bt ms grt ky ru H2O melt

Grt+ Grt-

6

bt ms grt ru H2O

Trivariant

bt kfs ms grt ky ru melt

bt ms grt ru H2O melt

10

kfs grt ky ru melt

Ms+ MsBt+ Bt-

12

Divariant

700

crd kfs grt opx ilm melt bt crd pl kfs grt opx ilm melt

750

800

850

Temperature (oC)

Thermocalc average

12

lt+ Me ltMe

bt ms grt ru H2O

bt kfs ms grt ru melt

bt ms grt ru H2O melt

bt ms grt ru ilm H2O bt ms grt ru ilm melt bt ms grt ilm H2O melt

Quadrivariant bt kfs grt ru melt (-pl)

bt ms grt ilm melt

bt kfs ms grt ky ru melt

bt kfs ms grt ru ilm melt

bt kfs ms grt ilm melt lt grt me bt ms melt il ilm m rt s ky il g bt ms grt ky s bt ms grt bt m st ilm H2O 2O y st ilm H2O ilm H grt k elt ms rt ky g t m b s bt m s ky H2O ilm bt m O bt ms bt ms ilm H2bt ms ky st ilm ky st ilm H2O bt ms sil ilm H2O H2O melt ilm melt bt kfs sil ilm melt bt ms sil st ilm H2O bt ms fs sil ilm bt kfs ms d k elt bt ms sil H2O melt t cr sil ilm melt b il ilm m ilm H2O s bt ms grt ilm H2O

8

bt kfs g ky ru (-pl)

Pentavariant

bt kfs grt ky ru melt

bt ms grt ru ilm H2O melt

10

Trivariant

bt kfs ms grt ky ru melt (-pl)

rt sg fs m bt k u melt r sil

bt kfs grt sil ru melt

bt kfs grt ru melt

Grt+ Grt-

kfs grt bt kfs grt ru melt (-pl) opx ru melt bt kfs grt opx (-pl) ru melt (-pl) bt kfs grt bt kfs grt opx ru melt bt kfs ms grt kfs grt opx sil ru ilm melt bt kfs grt opx sil ilm melt ru melt ru ilm melt bt kfs grt bt crd kfs grt kfs grt opx ru ilm melt bt kfs grt sil ilm melt ru ilm melt sil ilm melt bt kfs grt bt crd kfs kfs grt opx ilm melt grt ru melt bt kfs grt ilm melt bt crd kfs opx ilm melt grt sil ilm melt crd kfs grt opx ilm melt bt crd kfs bt crd kfs grt grt ilm melt opx ilm melt

6

crd kfs opx ilm melt bt crd kfs opx ilm melt

bt crd kfs ilm melt

H2710 4 600

dCr rd+ C

O Op pxx+

Pressure (kbar)

bt ms grt ru melt

ms pa bt grt ru H2O

Ms + Ms -

ms pa bt grt ru H2O melt

Divariant

kfs grt ru melt (-pl)

bt kfs ms grt ru melt (-pl)

Bt+ Bt-

ms pa bt grt ru melt

Na2O CaO K2O FeO MgO Al2O3 SiO2 H2O TiO2 wt% 0.87 0.77 4.43 8.68 3.29 13.8 65.5 2.02 0.98

KfsKfs+

14

+Qtz)

Als+ Als-

(b) NCKFMASHT (+Pl

650

700

750 o

Temperature ( C)

800

850

T. Imayama et al. / Lithos 134-135 (2012) 1–22

reaction including kyanite occurs at c. 730–780 °C above c. 8 kbar. The peak assemblage garnet + biotite + muscovite + plagioclase + kyanite + quartz + melt + K-feldspar + rutile for H2101 is predicted in a narrow P–T field with steep P/T gradient along muscovite-out line at c. 720–770 °C at c. 8–14 kbar (Fig. 4a), consistent with the average P–T condition at c. 740 ± 50 °C and c. 10.5 ± 2 kbar, as reported by Imayama et al. (2010). These facts suggest that the melt generation for sample H2101 occurred via muscovite dehydration melting (Ms + Pl + Qtz = Ky + Kfs + melt). In contrast, retrograde microstructures in sample H2101 are produced by the back reaction with melt (Grt + Kfs + melt = Bt + Pl + Sil + Qtz), based on the productions of bands of fine-grained fibrolitic sillimanite + biotite ± quartz surrounding garnet porphyroblasts and of antiperthite (Fig. 3a, b, c; e.g. Vernon, 2004; Kriegsman and Álvarez-Valero, 2010). In fact, pseudosection modeling indicates that the disappearances of garnet and K-feldspar occur at T = c. 700 °C below P = c. 6 kbar in the presence of melt under the stability field of sillimanite (Fig. 4a). Coupled with the constraints on minimum pressure by cordierite-in line, retrograde microstructures observed in sample H2101 are interpreted to have been formed at T = c. 650–700 °C at P = c. 4–6 kbar. These facts indicate that the P–T path of sample H2101 experienced a decompression path with fairly large dP/dT slope at the high temperature, right after the P–T conditions at peak-T. Sample H2710 directly below the HHT is a subaluminus metapelite, which bulk composition has relatively lower Al2O3 content, higher SiO2 and K2O contents than those of sample H2101. As a result, calculated P–T pseudosection for sample H2710 shows that aluminosilicate disappears at high temperature above c. 775 °C. Similarly, the biotite-out line moves towards higher temperature side (c. 825–850 °C) with the appearance of orthopyroxene (Fig. 4b), indicating that the biotite dehydration melting for sample H2710 requires higher temperatures than those of sample H2101. The peak assemblage garnet + biotite+ plagioclase + sillimanite + quartz + melt + Kfeldspar + rutile for H2710 is constrained by the P–T field of c. 730–780 °C at c. 7–10 kbar (Fig. 4b), which is in accord with the estimated average P–T conditions of c. 760 ± 60 °C at c. 8 ± 2 kbar within the error (Imayama et al., 2010). Retrograde biotite also occurs as bands of fine-grained fibrolitic sillimanite+ biotite ± quartz and intergrowth of biotite + plagioclase surrounding garnet porphyroblasts in sample H2710. Hence, the back reaction of biotite melting (Grt + Kfs + melt = Bt + Pl + Sil + Qtz) also occurred in sample H2710. Pseudosection modeling indicates that the disappearance of garnet occurs at T = c. 725 °C below P = c. 7 kbar in the presence of melt under the stability field of sillimanite, whereas that of K-feldspar occurs with the reappearance of muscovite (Fig. 4b). The presence of late muscovite such as fine-grained muscovite + quartz ± fibrolite aggregate in leucosomes suggests the P–T path of sample H2710 to have passed through the P–T field of c. 650–700 °C at 4–7 kbar between wet solidus and muscovite dehydration reactions (Fig. 4b), on cooling (Spear et al., 1999). These facts indicate that the P–T path of sample H2710 have experienced a relatively isobaric path, compared to those of sample H2101 during exhumation. 5. Zircon U–Pb ages 5.1. Analytical procedure Zircons from seven samples were separated using standard heavy liquid technique, and then handpicked under a binocular microscope.

7

The zircons from samples were mounted in epoxy resin, polished to expose the grain centers. CL and BSE images were obtained using the scanning electron microscopes JEOL 6610LV at Korea Basic Science Institute (KBSI), SEM-CL JEOL 5410LV at the Okayama University of Science, and JXA-8800 at the Natural Museum of Nature and Science. Mineral inclusions were identified by Oxford energy dispersive X-ray spectrometry (EDX) installed at JEOL 6610LV at KBSI. Zircon U–Pb dating for five samples (H1205, H2101, H2908, H2710, and H2708) was carried out using the sensitive high resolution ion microprobe (SHRIMP) at KBSI. The analytical procedures for SHRIMP dating were based on Williams (1998) and Ireland and Williams (2003). A 15–20 μm spot size was used for all analyses using a 1.5–2 nA negative ion oxygen beam (O2−). The measured 206Pb/238U ratio was calibrated using the standards FC1 zircon (c. 1099 Ma, Paces and Miller, 1993). The 206Pb*/ 238U ages of standard FC1 zircons were 1099 ± 25 Ma (2σ, n = 30) during the course of this work. The data were collected in sets of five scans throughout the mass measurements for each isotope. For the age calculations of cores and rims of zircon yielding the pre-Tertiary and Tertiary, the measured 204Pb/206Pb ratio and 207Pb/206Pb ratio were used for common Pb correction, respectively (Williams, 1998). Data reduction, age calculations, and common Pb corrections were carried out using the Isoplot/Ex 3.41 program (Ludwig, 2003). Both individual and mean 206Pb*/238U ages are quoted at 2σ confidence level. However, populations with mean square weighted deviation (MSWD) much above 2–3 are not statistically homogeneous for sure, and thus the ages for such groups are reported showing the ranges in age in this study. U–Th–Pb analyses of the zircon measured by SHRIMP are listed in Table 1. REE analyses of zircons from samples H2101 and H2710 were also carried out by the SHRIMP at KBSI. The analytical conditions and a spot size of 15–20 μm are the same as those of the U–Pb analyses. The NIST 611 glass standard was used to calibrate relative element sensitivities for the REE analyses of zircon. Detection limits for REE elements are c. 1 ppm. Results of REE contents of zircons are listed in Table 2. Two additional samples (M1806 and H90-30) and one sample (H2710) used for SHRIMP dating at KBSI were analyzed using a Cameca Nano-SIMS NS50 ion microprobe at the Ocean Research Institute of the University of Tokyo. The analytical procedure was the same as those described by Takahata et al. (2008). About 8 nA mass filtered O − primary beam was used to sputter a spot with a diameter of c. 15 μm. Isotopic data for multiple reference zircons indicate that the Nano-SIMS analysis using a 15-μm probe diameter is suitable for ion microprobe U–Pb zircon dating, and the Nano-SIMS 207Pb*/ 206 Pb* ages were comparable to SHRIMP ages (Takahata et al., 2008). The measured 206Pb/ 238U ratio was calibrated using the standards AS3 zircon (c. 1099 Ma, Paces and Miller, 1993). The 206Pb*/ 238 U ages of standard AC3 zircons were 1099 ± 10 Ma (2σ, n = 23) during the course of this work. Common Pb corrections for zircon data were calibrated using the measured 207Pb/ 206Pb ratio. U–Th–Pb analyses of the zircon measured by Nano-SIMS are listed in Table 3. 5.2. Results 5.2.1. Mylonitic augen orthogneiss (M1806) Zircon grains from sample M1806 above the LMCT are euhedral and pinkish white to bronze grains up to 300 μm. Most crystals show a well developed oscillatory zoning in CL image (Fig. 5a). Most grains are fractured and altered, and the primary oscillatory domain was partly modified by the dark-CL domains along the fractures within the crystal (Fig. 5a, b). The presence of faint oscillatory zoning

Fig. 4. P–T pseudosections for (a) aluminous metapelites (sample H2101) and (b) subaluminous metapelites (sample H2710) calculated in NCKFMASH system. The peak and retrograde assemblages are shown by bold, italic and italic letters, respectively. Dashed ellipses indicate the average peak P–T conditions estimated by Imayama et al. (2010). Black and thick arrows indicate P–T paths inferred from this work. Bulk compositions of the samples are shown on the top of each figure. The following solid solution models are used to calculate the pseudosection; muscovite (Chatterjee and Froese, 1975), cordierite (Holland and Powell, 1998), garnet and biotite (White et al., 2000), plagioclase (Newton et al., 1980), K-feldspar (Waldbaum and Thompson, 1968), orthopyroxene (Holland and Powell, 1996) and silicate-melt (Holland and Powell, 2001; White et al., 2001). Quartz is assumed to be in excess. The fluid is considered as pure H2O (aH2O = 1). Ilm: ilmenite; Opx: orthopyroxene; Pa: paragonite; Ru: rutile. See captions of Figs. 2, 3 for abbreviations of other minerals.

8

T. Imayama et al. / Lithos 134-135 (2012) 1–22

Table 1 SHRIMP U–Pb–Th analytical data of zircon. Grain. spota

U (ppm)b

238

U/ Pb

2σ (%)

206 Pb*/238U age (Ma, 2σ)c

207 Pb*/206Pb* age (Ma, 2σ)d

Domain

0.0725 0.0722 0.1637 0.1629 0.2364 0.0719 0.0725 0.0689 0.1053 0.0690 0.0717 0.0956 0.2618 0.0776 0.0996

7.6 2.8 1.2 1.4 1.0 1.0 1.4 4.0 1.4 2.0 1.6 0.6 0.8 11.6 25.0

868 ± 36 949 ± 24 2365 ± 50 2394 ± 52 2706 ± 56 1054 ± 22 999 ± 20 865 ± 26 1411 ± 34 861 ± 22 925 ± 22 1081 ± 22 2451 ± 4 25.2 ± 1.2 19.9 ± 1.2

947 ± 196 896 ± 94 2486 ± 22 2481 ± 22 3093 ± 18 959 ± 52 992 ± 28 811 ± 116 1675 ± 36 911 ± 48 963 ± 38 1482 ± 20 3252 ± 14

Core Core Core Core Core Core Core Core Core Core Core Core Core Rim Rim

1.8 2.0 2.2 3.0 2.4 2.2 2.4 2.2 2.2 2.4 2.8 2.4 3.4 3.2 3.2 3.2 3.6 3.4 3.0 3.2 3.8 3.4 3.4 3.8

0.0614 0.0674 0.0709 0.0720 0.0486 0.0480 0.0475 0.0486 0.0525 0.0476 0.0511 0.0519 0.0471 0.0463 0.0503 0.0478 0.0528 0.0620 0.0980 0.0504 0.0472 0.0545 0.0504 0.0588

1.2 3.8 2.0 0.8 6.2 5.4 7.6 5.0 6.6 6.6 6.4 9.6 6.0 5.2 5.4 6.4 11.8 9.6 6.8 10.2 13.0 10.0 10.8 16.0

565 ± 10 87 ± 1.8 918 ± 20 831 ± 24 20.3 ± 0.4 21.2 ± 0.4 19.0 ± 0.4 20.7 ± 0.4 20.2 ± 0.4 20.5 ± 0.6 19.8 ± 0.6 19.4 ± 0.4 21.2 ± 0.8 19.8 ± 0.6 20.8 ± 0.6 18.9 ± 0.6 16.6 ± 0.6 17.5 ± 0.6 16.8 ± 0.6 18.1 ± 0.6 16.1 ± 0.6 16.4 ± 0.6 16.5 ± 0.6 17.1 ± 0.6

618 ± 44 723 ± 152 1007 ± 64 977 ± 20

Core Core Core Core Inner rim Inner rim Inner rim Inner rim Inner rim Inner rim Inner rim Inner rim Inner rim Inner rim Inner rim Inner rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim

2.8 1.6 3.5 2.4 3.6 3.2 1.6 7.0 3.6 1.8 1.2 4.0 8.4 5.8 1.2 2.6 5.0 3.6 3.8 1.6 1.2 2.4 4.0 3.8 1.2 2.4 3.8 5.8 1.2

0.0686 0.0707 0.1047 0.1030 0.0689 0.0983 0.1001 0.0739 0.0713 0.0752 0.0676 0.0712 0.1293 0.0689 0.0589 0.0696 0.0683 0.0667 0.0694 0.0673 0.0725 0.0987 0.0785 0.0645 0.0565 0.0606 0.0736 0.0601 0.0578

1.4 1.4 1.8 1.4 1.2 1.0 1.0 3.0 1.2 1.2 0.4 1.2 3.0 0.6 5.8 3.2 1.0 3.6 2.0 1.8 1.6 1.0 3.8 4.4 3.6 3.4 7.0 4.6 7.6

876 ± 24 806 ± 12 1639 ± 50 1514 ± 32 866 ± 30 1425 ± 42 1485 ± 22 966 ± 66 904 ± 30 985 ± 16 803 ± 8 795 ± 30 1665 ± 124 696 ± 38 389 ± 4 915 ± 22 581 ± 28 706 ± 24 867 ± 32 616 ± 10 870 ± 10 1462 ± 34 20.1 ± 0.8 19.7 ± 0.8 19.8 ± 0.2 17.7 ± 0.4 18.4 ± 0.8 19.1 ± 1.2 19.3 ± 0.2

821 ± 38 788 ± 72 1696 ± 36 1672 ± 28 879 ± 28 1587 ± 20 1613 ± 22 998 ± 66 954 ± 24 984 ± 50 850 ± 8 949 ± 24 2087 ± 52 863 ± 20 502 ± 134 902 ± 64 810 ± 36 787 ± 80 877 ± 50 817 ± 44 976 ± 34 1589 ± 22

Core Core Core Core Core Core Core Core Core Core Core Core Core Core Core Core Core Core Core Core Core Core Inner Inner Inner Inner Inner Inner Inner

Th/ U

206

Common Pb (%)

206

62 176 231 227 175 1238 661 266 237 321 98 1391 172 2 2

0.75 0.35 0.52 0.51 0.35 0.41 0.31 1.00 0.43 0.85 0.19 0.50 0.50 0.01 0.01

– 0.22 – 0.16 0.03 2.22 0.57 0.40 3.53 0.24 0.99 1.05 – 7.99 0.25

6.92 6.28 2.25 2.22 1.92 5.62 5.96 6.94 4.07 7.00 6.48 5.46 2.16 245.12 301.62

4.4 2.6 2.6 2.6 2.6 2.2 2.2 3.0 2.6 2.6 2.6 2.2 2.6 4.6 4.8

Ky-Sil migmatitic gneiss (H2101) 7.1 1110 251 8.1 922 21 1.3 292 204 3b.1 1815 184 1.1 1301 8 6.1 1927 12 9.1 1296 15 13.1 2228 15 14.1 1976 15 16.1 1230 9 19.1 1449 9 23.2 2680 32 3b.2 1854 13 4b.1 2275 29 5b.1 1679 11 6b.1 2065 24 4.1 496 9 5.1 575 7 7.2 850 11 10.1 678 9 11.1 445 7 12.1 684 11 17.2 579 6 1b.2 1070 17

0.23 0.02 0.70 0.10 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.02 0.01 0.01 0.01 0.01 0.02 0.01 0.02

0.12 0.46 – 0.03 2.22 0.40 2.14 0.66 1.16 3.53 0.99 1.05 1.47 – 0.36 0.67 5.60 4.69 5.97 – – 4.20 7.54 2.83

10.90 73.45 6.55 7.27 316.81 303.50 338.75 310.10 317.03 312.91 323.24 329.08 303.96 324.80 308.36 340.84 384.77 361.62 360.05 353.09 400.16 388.66 387.23 370.92

Sil migmatitic gneiss (H2908) 1.1 412 405 2.1 178 169 3.1 263 143 3.2 486 151 6.2 1080 268 10.1 1552 60 10.2 257 152 11.2 258 117 13.1 1193 509 14.2 173 128 15.1 3512 1205 16.1 3119 42 18.1 1403 93 22.1 817 188 23.1 1274 25 23.2 1423 647 24.2 505 278 25.1 783 66 25.2 301 180 26.1 664 397 27.1 776 316 28.1 441 270 1.2 660 3 1.3 841 5 2.3 1040 9 5.1 1261 12 6.1 1063 9 11.1 757 5 12.1 1366 12

0.98 0.95 0.54 0.31 0.25 0.04 0.59 0.45 0.43 0.74 0.34 0.01 0.07 0.23 0.02 0.45 0.55 0.08 0.60 0.60 0.41 0.61 0.01 0.01 0.01 0.01 0.01 0.01 0.01

0.04 0.57 0.48 1.10 0.12 1.05 0.85 0.32 0.26 0.39 0.21 0.69 3.46 0.77 0.56 – 1.09 0.45 0.18 0.85 0.53 0.81 4.05 2.28 1.28 1.80 3.44 1.73 1.44

6.9 7.5 3.5 3.8 6.9 4.0 3.8 6.2 6.6 6.0 7.5 7.6 3.4 8.8 16.0 6.6 10.5 8.6 6.9 9.9 6.9 3.9 307.3 320.0 320.4 356.5 338.6 331.5 329.4

Grt-mica gneiss (H1205) 1.1 82 2.1 495 3.1 445 3.2 442 4.1 506 5.1 2994 5.2 2145 6.1 267 8.1 549 6.2 377 9.1 508 11.1 2762 4.1 341 1.4 364 1.2 432

Th (ppm)

2σ (%)

207 206

Pb/ Pb

rim rim rim rim rim rim rim

T. Imayama et al. / Lithos 134-135 (2012) 1–22

9

Table 1 (continued) U (ppm)b

Th (ppm)

238

207 Pb*/206Pb* age (Ma, 2σ)d

Common Pb (%)

206

15 9

0.01 0.01

1.74 2.67

335.0 327.2

4.4 3.4

0.0601 0.0676

5.4 3.4

18.9 ± 0.8 19.1 ± 0.6

Sil migmatitic gneiss (H2710) 1.1 223 99 4.1 300 315 6.2 1878 231 8.1 238 155 1b.2 68 103 2b.2 390 190 8b.2 427 165 14b.1 2618 661 1–1.1 1479 32 2.1 1112 16 4.2 888 12 5.2 973 11 6.1 1302 15 10.1 827 6 11.1 2232 34 14.1 776 9 2b.1 924 5 4b.1 682 7 7b.1 797 8 12b.1 812 10 13b.1 837 10 1m.1 448 6 2m.2 475 2 3m.1 853 8 5.1 739 15 9.2 594 14 12.2 837 16 16.1 758 28 18.1 703 12 1b.1 672 21 1b.3 651 17 3b.1 461 14 5b.1 873 22 6b.1 1402 27

0.44 1.05 0.12 0.65 1.52 0.49 0.39 0.25 0.02 0.01 0.01 0.01 0.01 0.01 0.02 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.02 0.02 0.02 0.04 0.02 0.03 0.03 0.03 0.03 0.02

0.53 – 0.85 0.00 0.55 0.08 0.01 0.01 0.28 1.73 – – 1.11 – 0.75 – 1.11 0.00 0.55 3.16 1.94 – 8.43 5.78 2.31 5.50 0.92 3.88 2.16 0.00 – 3.94 – –

14.13 6.48 22.29 6.13 7.50 4.01 6.26 3.95 9.16 214.25 209.77 221.75 208.94 215.63 206.37 218.27 193.21 225.87 209.06 219.12 218.69 209.83 196.35 213.60 270.49 243.62 241.37 282.58 243.34 264.00 251.89 248.88 285.16 272.96

2.6 2.2 2.0 2.2 4.0 7.2 3.2 2.8 2.4 2.2 2.6 2.8 2.2 2.6 2.2 2.6 3.4 3.6 3.4 3.4 3.4 6.0 4.0 3.2 2.8 3.2 3.0 2.8 2.8 4.0 3.6 4.0 3.6 3.2

0.0697 0.0746 0.0683 0.0806 0.0785 0.1179 0.0816 0.0900 0.0835 0.0451 0.0512 0.0383 0.0488 0.0482 0.0496 0.0541 0.0492 0.0519 0.0508 0.0540 0.0499 0.0661 0.0894 0.0605 0.0532 0.0715 0.0498 0.0492 0.0502 0.0521 0.0456 0.0550 0.0492 0.0476

4.8 4.0 5.2 2.2 5.2 4.8 1.4 0.4 1.8 10.0 7.4 21.2 5.4 7.0 6.0 7.0 6.2 7.6 8.2 8.2 6.8 19.4 19.2 24.6 8.2 8.2 7.4 8.6 12.0 8.4 8.8 9.2 7.8 6.2

439 ± 10 925 ± 20 281 ± 6 974 ± 20 803 ± 30 1434 ± 96 956 ± 28 1454 ± 36 666 ± 16 30.1 ± 0.6 30.5 ± 0.8 29.3 ± 0.8 30.7 ± 0.6 29.8 ± 0.8 31.1 ± 0.8 29.2 ± 0.8 33.2 ± 1.2 28.3 ± 1.0 30.6 ± 1.0 29.1 ± 1.0 29.3 ± 1.0 29.9 ± 1.8 31.0 ± 1.4 29.6 ± 1.2 23.6 ± 0.6 25.7 ± 0.8 26.6 ± 0.8 22.7 ± 0.6 26.3 ± 0.8 24.2 ± 1.0 25.6 ± 1.0 25.6 ± 1.0 22.5 ± 0.8 23.5 ± 0.8

780 ± 168 1073 ± 80 640 ± 152 1211 ± 44 1038 ± 170 1914 ± 88 1234 ± 28 1424 ± 10 1230 ± 48

Core Core Core Core Core Core Core Core Core Inner rim Inner rim Inner rim Inner rim Inner rim Inner rim Inner rim Inner rim Inner rim Inner rim Inner rim Inner rim Inner rim Inner rim Inner rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim

Bt–Ms leucogranite (H2708) 1.1 86 2.1 157 4.1 747 5.1 319 6.1 190 1.2 2841 10.1 912 15.2 995 3.1 4095 3.3 6626 4.2 2181 5.2 570 3.2 1129 8.1 1234 8.2 794 1.3 959 9.1 954 10.2 981 11.1 848 11.2 1126 12.1 1026 13.1 893 13.2 559 14.1 2440 14.2 1660 6b.2 1466 11.3 460 15.1 1503

0.75 1.91 0.41 0.02 0.18 0.00 0.05 0.09 0.03 0.03 0.03 0.07 0.08 0.08 0.03 0.11 0.10 0.06 0.07 0.05 0.09 0.06 0.07 0.07 0.10 0.10 0.08 0.05

0.51 0.65 0.37 0.31 0.31 0.07 0.20 0.12 2.41 0.98 2.76 6.57 3.99 0.00 8.13 0.00 – 3.67 11.36 5.95 7.19 11.03 16.92 4.91 1.23 3.05 9.08 6.97

2.00 11.23 11.89 3.07 8.95 11.82 47.69 17.16 352.91 346.92 351.62 393.53 396.40 396.65 387.78 382.76 393.98 398.32 378.23 383.10 397.50 383.73 381.48 382.76 380.11 399.10 348.61 374.70

4.2 3.8 2.6 4.4 4.0 2.4 2.6 2.4 3.2 2.8 2.8 8.2 4.0 3.6 4.3 3.8 3.8 3.8 4.0 3.6 3.6 3.8 5.2 4.2 3.2 3.6 5.4 3.2

0.2049 0.0635 0.0704 0.1113 0.0952 0.0586 0.0589 0.0548 0.0494 0.0483 0.0562 0.0746 0.0670 0.0589 0.0706 0.0839 0.0693 0.0703 0.0662 0.0644 0.0626 0.0631 0.0821 0.0470 0.0636 0.0509 0.1013 0.0475

2.6 7.2 3.2 2.6 5.0 1.6 3.8 4.2 11.4 4.2 11.4 17.4 11.4 19.2 11.2 14.8 13.4 13.8 10.2 10.8 9.2 10.6 25.4 37.8 8.8 8.8 17.2 8.6

2603 ± 88 547 ± 20 519 ± 14 1813 ± 70 681 ± 26 523 ± 12 134 ± 4 365 ± 8 18.2 ± 0.6 18.5 ± 0.6 18.1 ± 0.6 15.8 ± 1.4 15.8 ± 0.6 16.0 ± 0.6 16.1 ± 0.6 16.0 ± 0.6 15.9 ± 0.6 15.7 ± 0.6 16.6 ± 0.6 16.4 ± 0.6 15.9 ± 0.6 16.4 ± 0.6 16.1 ± 1.0 16.8 ± 0.8 16.6 ± 0.6 16.0 ± 0.6 17.2 ± 1.0 17.2 ± 0.6

2840 ± 48 545 ± 264 853 ± 94 1784 ± 54 1487 ± 122 534 ± 40 504 ± 116 367 ± 118

Core Core Core Core Core Core Core Core Inner rim Inner rim Inner rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim Outer rim

a b c d

64 301 307 7 34 10 49 94 122 170 56 38 93 104 25 102 91 55 63 61 88 50 38 167 163 147 37 68

2σ (%)

206

Pb/ Pb

206 Pb*/238U age (Ma, 2σ)c

206

Sil migmatitic gneiss (H2908) 19.1 1528 24.1 1074

U/ Pb

207

Th/ U

Grain. spota

2σ (%)

The first and second numerals divided by a period (.) denote grain and spot numbers, respectively. The SL13 srandard zircon, which contains 238 ppm of U, has been used to determine the U content of target. Common Pb for zircons yielding the Tertiary and pre-Tertiary ages were corrected using measured 207Pb/206Pb and Common Pb were corrected using measured 204Pb/206Pb ratio.

204

Domain

Inner rim Inner rim

Pb/206Pb ratios, respectively.

10

T. Imayama et al. / Lithos 134-135 (2012) 1–22

Table 2 REE data (concentrations in ppm) of zircon. Spot

Eu

Gd

Tb

Kyanite–sillimanite migmatitic gneiss (H2101) 3 0.04 36.3 0.41 0.95 1.74 4 0.04 41.4 0.18 0.36 1.04 1 0.01 4.61 0.03 0.07 0.26 6 0.09 3.14 0.01 0.04 0.28 7 0.02 2.28 0.00 0.03 0.22 2 0.61 7.14 0.24 0.21 0.62 5 0.29 5.59 0.03 0.08 0.41 8 0.06 4.08 0.02 0.06 0.26

0.54 0.54 0.70 0.54 0.43 1.05 0.79 0.81

11.6 8.55 3.83 2.74 3.15 4.50 4.44 4.13

30.6 25.6 8.36 5.22 7.77 13.3 13.5 13.5

Sillimanite migmatitic gneiss (H2710) 1 0.34 42.8 2.51 4.80 8 0.10 61.7 0.29 0.67 4 0.02 1.96 0.06 0.12 7 0.01 5.13 0.08 0.07 3 0.06 1.85 0.08 0.09 5 0.00 2.10 0.02 0.14 6 0.05 1.58 0.05 0.11

1.30 0.36 0.98 0.50 0.28 0.40 0.35

33.8 9.43 5.02 3.77 4.52 4.88 5.18

76.7 24.3 13.3 9.00 13.3 13.4 15.4

a

La

Ce

Pr

Nd

Sm

6.87 1.72 0.44 0.42 0.46 0.54 0.51

Dy 79.9 76.1 8.73 4.39 6.70 25.8 28.2 33.9

184 56.5 19.5 14.6 31.3 23.5 30.0

Ho

Er

Tm

Yd

Lu

LuN/GdN

Eu/Eu*a

Domain

137 151 7.66 2.54 3.73 33.1 36.1 48.0

160 207 4.56 1.97 1.87 27.7 33.7 49.2

127 177 3.50 1.28 1.23 19.1 23.8 34.3

216 321 5.45 2.05 1.51 25.0 34.8 54.9

238 391 5.39 2.22 1.26 28.4 38.1 59.9

26 57 2 1 1 8 11 18

0.08 0.12 0.36 0.37 0.27 0.42 0.34 0.38

Core Core Inner rim Inner rim Inner rim Outer rim Outer rim Outer rim

295 89.6 14.2 14.2 39.8 22.3 33.6

323 97.3 8.15 8.78 35.4 15.7 26.9

232 77.3 4.61 6.96 24.6 9.66 15.3

364 128 6.46 12.1 38.0 14.0 22.5

382 135 5.76 12.2 39.8 12.8 20.2

14 18 1 4 11 3 5

0.06 0.06 0.37 0.24 0.11 0.15 0.13

Core Core Inner rim Inner rim Outer rim Outer rim Outer rim

Eu/Eu* = 2EuN/(SmN + GdN).

with dark-CL at the rims is interpreted as remnant growth zoning, indicating recrystallization of pre-existing zircon (Fig. 5a). The zircon cores and recrystallized domain yielded 207Pb*/ 206Pb* ages of c. 1860 to 1780 Ma and c. 1800 to 1620 Ma, respectively (Fig. 6a), whereas their 206Pb*/ 238U ages are c. 1560 to 350 Ma (Table 3). These discordant ages are also the case for the other samples (see Figs. 6, 7), suggesting younger thermal events, but the discordia is not well defined. Hence, only 207Pb*/ 206Pb* ages are reported for the zircon cores in this study. 5.2.2. Garnet-mica gneiss (H1205) Most grains from sample H1205 at the base of the HHCS are characterized by cores with bright-CL domains (Fig. 5c). In some cases, the primary oscillatory core was slightly modified by the patchy and dull

CL domains. The cores show high Th/U ratio of c. 0.31–1.00 (Table 1). Their 207Pb*/ 206Pb* ages yielded a wide range of c. 3090–810 Ma, including the clusters at nearly concordant ages of c. 980 to 860 Ma, which are possibly interpreted to be ages of detrital zircon (Fig. 7a). Most crystals have a thin overgrown rim with dark-CL domains, which were rarely large enough for analyses. Two spot analyses in a single grain were successfully conducted for the overgrown rim with moderate U content of 360–430 ppm and low Th/U ratio b 0.01 (Table 1), showing the isolated 206Pb*/ 238U ages of 19.9 ± 1.2 Ma and 25.2 ± 1.2 Ma (Fig. 5c). 5.2.3. Kyanite–sillimanite migmatitic gneiss (H2101) Zircon grains from sample H2101 are subhedral with rounded edges, and have a distinct core-rim structure with two rims successively

Table 3 Nano-SIMS U–Pb analytical data of zircon. Grain.spota

238

U/206Pb

Mylonitic augen gneiss (M1806) 16.17 3.33 18.18 3.86 23.20 3.88 33.28 4.13 18.19 4.09 27.23 7.58 30.24 3.66 30.26 18.28 33.30 4.22 Sillimanite migmatitic gneiss (H2710) 5.46 12.61 5.47 17.11 17.54a 12.23 19.57 14.60 17.54b 21.73 17.54c 177.37 Granitic orthogneiss (H90-30) 16.31 11.70 27.38 43.25 25.34 15.15 25.33 14.23 25.36 17.53 27.37 15.20 32.40 15.29 a b c

2σ (abs.)

207

Pb/206Pb

2σ (abs.)

206 Pb*/238U age (Ma, 2σ)b,c

207 Pb*/206Pb* age (Ma, 2σ)b,c

Domain

0.42 0.23 0.26 0.26 0.25 0.51 0.19 1.70 0.28

0.1142 0.1095 0.1129 0.1104 0.1103 0.1101 0.1103 0.1046 0.1098

0.0033 0.0005 0.0004 0.0006 0.0004 0.0003 0.0010 0.0012 0.0003

1693 ± 187 1485 ± 78 1476 ± 90 1396 ± 78 1410 ± 79 798 ± 50 1557 ± 73 342 ± 30 1371 ± 83

1862 ± 56 1781 ± 8 1835 ± 7 1796 ± 11 1798 ± 7 1787 ± 5 1795 ± 18 1624 ± 30 1788 ± 5

Core Core Core Core Recrystallization Recrystallization Recrystallization Recrystallization Recrystallization

1.07 2.57 0.83 0.93 1.98 26.64

0.0708 0.0706 0.0719 0.0774 0.0709 0.0568

0.0012 0.0016 0.0006 0.0021 0.0008 0.0009

491 ± 40 365 ± 53 506 ± 33 427 ± 26 289 ± 26 35.8 ± 5.4

902 ± 40 830 ± 58 938 ± 22 1093 ± 57 896 ± 28 93 ± 100

Core Core Core Core Core Inner rim

0.50 4.34 0.78 0.77 1.04 0.97 0.88

0.0570 0.0564 0.0575 0.0573 0.0593 0.0560 0.0574

0.0002 0.0004 0.0006 0.0005 0.0013 0.0010 0.0003

412 ± 21 437 ± 23 528 ± 22 147 ± 15 410 ± 25 408 ± 23 356 ± 20

469 ± 30 442 ± 25 475 ± 10 430 ± 19 417 ± 45 475 ± 14 457 ± 69

The first and second numerals divided by a period (.) denote grain and spot numbers, respectively. In order to estimate a 206Pb*/238U and 207Pb*/206Pb* ages, the measured 207Pb/206Pb ratio was used for common Pb correction. Uncertainty of the 206Pb*/238U age was estimated by error propagation of the uncertainty of the 206Pb/238U and 204Pb/206Pb ratios.

Core Core Core Core Recrystallization Recrystallization Recrystallization

T. Imayama et al. / Lithos 134-135 (2012) 1–22

(a) M1806

11

(b) M1806

(c) H1205 25.2±1.2

Rim

(947±196) Re-c Core

Core

(1788±5)

(1796±11)

Rim

19.9±1.2 100 µm

100 µm

(d) H2101

Or

(e) H2101

Or

20 µm

(f) H2101

18.9±0.6 Ir

(1007±64)

Core

Core Core

(618±44)

Ir

16.8±0.6

50 µm

20 µm

20 µm

(i) H2710

23.6±0.6 Or

29.3±0.8

Core

Ir

Qtz

Ir

(h) H2710

(821±38)

Kfs

20.8±0.6

(g) H2908 20.1±0.8

Core

20.3±0.4 Or

Or

24.2±1.0 Ir

Ir

19.7±0.8

Core

Ir

(1038±170)

25.6±1.0

50 µm

(j) H2710 Or

20 µm

(k) H2710

20 µm

(l) H2708

Bt

Or

30.7±0.6

15.8±1.4 Core

Core

Ir

Kfs

(1784±54) Ir

Kfs

(m) H2708

(n) H90-30 15.8±0.6 Ir

Ir

Or

18.2±0.6

30 µm

20 µm

20 µm

18.5±0.6

Or

(o) H90-30

(475±13) Re-c

(430±19)

Core

(442±25)

Core

(475±10) Core

Re-c

(417±44) 50 µm

Pb*/ U age ( Pb*/206Pb* age 207

50 µm

238

(

206

20 µm

Fig. 5. Representative cathodoluminescence (CL) and back-scattered electron (BSE) images of dated zircon crystals. (a) CL- and (b) BSE-images of zircon grains in augen gneiss M1806, showing an oscillatory zoned core surrounded by a dark recrystallized rim. (c) CL-image of zircon grains from garnet-mica gneiss H1205, showing a core and rim with bright-CL and dark-CL domains, respectively. (d, e) CL- and (f) BSE-images of zircon grains in kyanite–sillimanite migmatitic gneiss H2101, showing a core–rim structure with two distinct, successively overgrown rims over an inherited core. See the euhedral inner overgrown rim with dark-CL, and quartz and K-feldspar inclusions surrounding the highly corroded cores. (g) CL-image of zircon grains from sillimanite migmatitic gneiss H2908, showing a core and inner rim with bright-CL and dark-CL domains, respectively. Zircons have a very thin outer overgrown rim with a little brighter-CL domains than those of an inner rim. (h, i, j) CL- and (k) BSE-images of zircon grains in sillimanite migmatitic gneiss H2710, showing an inherited core with bright-CL image surrounded by two distinct parts of successively overgrown rim with different brightness of CL. Also see an inclusions-rich band including K-feldspar and biotite situated at the boundary between the core and the inner rim. (l, m) CL-images of zircon grains in biotite–muscovite leucogranite H2708, showing two successively overgrown rims with or without xenocrystic cores. (n, o) CL-images of zircon grains in granitic orthogneiss H90-30, showing well-developed internal oscillatory zoning. Dashed and solid circles indicate the locations of spot analyses with (a, n–o) Nano-SIMS and (c–e, g–j, l–m) SHRIMP, respectively. Numerals show U–Pb ages with ±2 σ errors. Both 206Pb*/238U ages for rims and 207Pb*/206Pb* ages for cores (enclosed by parentheses) are shown. Re-c: recrystallization; Ir: inner rim; Or: outer rim.

overgrown on an inherited core (Fig. 5d, e). Euhedral inner rim with dark-CL is dominant in the grain, whereas the core is highly corroded. Micro-sized mineral inclusions are mostly found at the boundary between the inherited cores and the inner overgrown rims (Fig. 5f). The

inclusions consist of quartz, K-feldspar, plagioclase, apatite, and monazite. The Th/U ratios of the cores vary in a wide range of c. 0.01–0.70, but are generally higher than those of overgrown rims (Table 1). Chondrite-normalized REE patterns of the cores are characterized

12

T. Imayama et al. / Lithos 134-135 (2012) 1–22

(a) M1806 0.12

core recrystallization

1900 1800 1700

0.10

1600

207

Pb/206Pb

0.11

1500

0.09

1400 1300

0.08 0

4

8

12 238

16

20

24

U/206Pb

(b) H2710 0.08 core inner rim

1040 880 720

0.06

560

0.05

240

207

Pb/206Pb

0.07

400

80

0.04 0

40

80

120 238

160

200

240

U/206Pb

(c) H90-30 0.08 1120

0.06

207

Pb/206Pb

0.07

core recrystallization

1040 960 880 800 720 640 560 480

400 320

0.05

240 160

0.04 0

10

20

30 238

40

50

U/206Pb

Fig. 6. Tera–Wasserburg plots for U–Pb isotopic ratios of dated zircon measured using Nano-SIMS from samples (a) M1806, (b) H2710, and (c) H90-30 along the Tamor–Ghunsa section, far-eastern Nepal. Gray, dark-gray and white circles are data-points for the core, inner rim and the recrystallized domain, respectively. All error ellipses are plotted at 2σ level.

by abundance of REE contents with a steep positive slope (LuN/ GdN = 26–57) of MREE to HRRE (Figs. 8a, 9). The 207Pb*/206Pb* ages for cores are scattered between c. 1010–620 Ma (Fig. 7b). The inner overgrown rims with dark-CL show high U content of c. 1300–2700 ppm and low Th/U ratio less than 0.01 (Table 1). MREE and HREE contents of inner rims are significantly lower than those of cores, and the slope (LuN/GdN = 1–2) of the MREE to HREE are considerably flat (Figs. 8a, 9). Twelve analyses on inner rims yielded the 206 Pb*/ 238U ages ranging from 21.2 ± 0.4 to 18.9 ± 0.6 Ma (n = 12, Fig. 7c). The inner rim is further surrounded by a thin outer rim with relatively bright-CL layer, showing moderate U content of c. 450–1070 ppm and low Th/U ratio of c. 0.01–0.02 (Table 1). The outer overgrown rims have relatively enriched HREE and similar MREE, compared to those of inner rims, and have moderate slope (LuN/GdN = 8–18) of the MREE to HREE (Figs. 8a, 9). Eight analyses of the outer overgrown rim yielded the 206Pb*/ 238U ages ranging from 18.1 ± 0.6 to 16.1 ± 0.6 Ma (n = 8, Fig. 7c). 5.2.4. Sillimanite migmatitic gneiss (H2908) Zircon grains from sample H2908 exhibit subhedral to euhedral shape with slightly rounded edges. In most grains, an oscillatory zoned core with bright-CL domains is surrounded by one or two

overgrown rims with dark-CL domains (Fig. 5g). The cores show high Th/U ratio of c. 0.01–0.98 (Table 1), and their 207Pb*/ 206Pb* ages yielded a wide range of c. 2090–500 Ma, including the cluster at nearly concordant ages of c. 980 to 820 Ma (Fig. 7d). Inner overgrown rims with dark-CL image show high U content of c. 660–1530 ppm and low Th/U ratio of c. 0.01 (Table 1), yielding the 206 Pb*/ 238U ages ranging from 20.1 ± 0.8 to 17.7 ± 0.4 Ma (n = 9, Fig. 7e). Most crystals have outer overgrown rims with relatively brighter-CL domains than those of inner rims, but they are too thin for SHRIMP and REE analyses (Fig. 5g). 5.2.5. Sillimanite migmatitic gneiss (H2710) Zircon grains from sample H2710 exhibit subhedral to euhedral shape with slightly rounded edges. The internal structure consists of an inherited core with bright-CL, which is surrounded by at least two different overgrown rims with different brightness of CL (Fig. 5h, i, j). An inclusion-rich band mostly occurs in the xenocrystic core and at the boundary between the core and the inner overgrown rim (Fig. 5k), where K-feldspar, plagioclase, quartz, biotite, apatite, and monazite were found. Most cores show the high Th/U ratio of c. 0.02–1.52 with high Th content (Table 1), and abundant REE contents with a steep slope (LuN/GdN = 14–18) of MREE to HREE (Figs. 8b, 9).

T. Imayama et al. / Lithos 134-135 (2012) 1–22

The 207Pb*/ 206Pb* ages for cores using SHRIMP and Nano-SIMS are c. 1900–640 Ma (Fig. 7f) and c. 1090–830 Ma (Fig. 6b), respectively. Inner overgrown rims with dark-CL show high U content of c. 700–2200 ppm and low Th/U ratio less than 0.02 (Table 1). They have significantly low concentrations of MREE and HREE with a flat slope (LuN/GdN = 1–4) of MREE to HREE (Figs. 8b, 9). The 206Pb*/ 238 U ages of inner rims range from 33.2 ± 1.2 to 28.3 ± 1.0 Ma (n = 15, Fig. 7g). The inner overgrown rim with dark-CL is further surrounded by an outer rim with a brighter-CL domain (Fig. 5h, i, j), which has moderate U content of c. 500–1400 ppm and moderate Th/U ratio of c. 0.02–0.05 (Table 1). The REE patterns show relatively

13

higher concentrations of HREE than those of inner rims, and MREE to HREE patterns with a moderate slope (LuN/GdN = 3–11) (Figs. 8b, 9). The 206Pb*/ 238U ages of the outer rims range from 26.6 ± 0.8 to 22.5 ± 0.8 Ma (n = 10, Fig. 7g). 5.2.6. Muscovite–biotite leucogranite (H2708) Zircon grains from sample H2708 are idiomorphic with large grain size up to 300 μm. They are divided into a xenocrystic core including minute minerals with bright-CL image, and two successively overgrown rims surrounding the core (Fig. 5l). An inner overgrown rim shows dark-CL, whereas an outer rim often has the brighter-CL

(a) H1205 core 0.28

core 3200

0.24 3000 2800 2600

0.16

2400

207

Pb/206Pb

0.20

2200

0.12

2000 1800 1600 1400

0.08

1200

1000 800

0.04 0

2

4

6

8

10

238

U/206Pb

(b) H2101 core

(c) H2101 rims

0.10

0.14

core

inner rim outer rim

1500

0.09

0.12

0.08

700

0.06

500

0.05

300

20

40

60

0.02

80

100

238

U/206Pb

30

25

200

15

20

300

400

238

U/206Pb

(d) H2908 core 0.14

5045 40 35

0.04

100

0.04 0

b

0.06

0.08

207

207

0.07 900

0.10

mon P

Pb/206Pb

1100

To com

Pb/206Pb

1300

(e) H2908 rims

2200

0.14

core

inner rim

0.12

0.12

Pb/206Pb

1400

0.06

1000

Pb

0.06

0.08

207

0.08

0.10

n mmo

207

0.10

To co

Pb/206Pb

1800

600

0.04

60 50

40

20

30

200

0.04 0

10

20

30

238

U/

206

Pb

40

50

0.02 100

140

180

220

260

300

340

380

238

U/206Pb

Fig. 7. Tera–Wasserburg plots for U–Pb isotopic ratios of cores and rims of dated zircon measured using SHRIMP from (a) sample H1205, (b, c) sample H2101, (d, e) sample H2908, (f, g) sample H2710, and (h, i) H2708 along the Tamor–Ghunsa section, far-eastern Nepal. Gray, dark-gray and white ellipses are data-points for the core, inner rim and outer rim domains, respectively. All error ellipses are plotted at 2σ level.

14

T. Imayama et al. / Lithos 134-135 (2012) 1–22

(f) H2710 core

(g) H2710 rims

0.14

0.14

core 2000

0.12

inner rim outer rim

0.12 Pb/206Pb

1600

0.08

207

1400 1200

0.10 omm

207

0.10

To c

Pb/206Pb

1800

0.08

800

0.06

600 400

0.04 0

10

200

20 238 206 U/ Pb

0.04

30

0.02 100

b

0.06

on P

1000

50 45 40

140

35

180

30

20

25

220

260

300

340

380

238

U/206Pb

(h) H2708 core

(i) H2708 rims

0.24

0.16

core 3000

0.14

inner rim outer rim

0.20 0.12 Pb/206Pb

b

0.04 0

0.08

nP

0

0.06 60

0.08

1800 00 14 00 10

0.10

mo

0.12

207

2200

com

207

0.16

To

Pb/206Pb

2600

0.04

28

24

20

16

200

10

20

30

40

50

60

238

U/206Pb

0.02 200

240

280

320

360

400

440

480

238

U/206Pb

Fig. 7 (continued).

(Fig. 5l) with the oscillatory zoning (Fig. 5m). The Th/U ratio of the cores is highly variable in the range of c. 0.01–1.91, but is mostly higher than those of the rims (Table 1). The 207Pb*/ 206Pb* ages for cores are highly variable ranging from 2840 to 370 Ma (Fig. 7h), representing the pre-Himalayan inherited ages from different sources. Inner overgrown rims with dark-CL image show high U content of c. 2200–6600 ppm and low Th/U ratio of c. 0.03 (Table 1). Three 206Pb*/ 238 U ages of the inner rims are consistent within the error, yielding a weighted mean age at 18.3 ± 0.3 Ma (2σ, n = 3; MSWD 0.73, Fig. 7i). On the other hand, the outer rims have moderate U content of 460–2440 ppm and moderate Th/U ratio of c. 0.03–0.11 (Table 1). The outer rims yielded a weighted mean 206Pb*/238U ages of 16.3 ± 0.2 Ma (2σ, n = 17; MSWD 1.8, Fig. 7i). 5.2.7. Granitic orthogneiss (H90-30) Zircon grains from sample H90-30 are subhedral to euhedral with large grain size up to 300 μm in diameter. The CL images of zircon grains show well-developed prismatic facies and internal oscillatory zoning (Fig. 5n, o), which are also recognized in the BSE images. A few zircon grains have xenocrystic cores showing irregular, concentric, and complex zoning. The outer parts in zircons often have darkCL domains, exhibiting weakly oscillatory zoning parallel to the unmodified one in the interior of the crystal (Fig. 5n, o). Such zoning is the relict of primary oscillatory zoning, implying that their domains were formed by recrystallization of pre-existing zircon rather than overgrowth (Hoskin and Black, 2000). The cores and recrystallized

domains yielded the 207Pb*/ 206Pb* ages of c. 475 to 430 Ma and c. 475 to 420 Ma, respectively (Fig. 6c), indicating no clear discrimination in the ages between them. Tertiary ages from sample H90-30 have not been obtained in this study. 6. K–Ar biotite age K–Ar biotite ages were analyzed for six samples along the Tamor– Ghunsa section (Fig. 2); the results are shown in Table 4. Samples were crushed and sieved, and then their fractions were washed in distilled water and dried at 60 °C. Pure biotite fractions were collected by handpicking after electromagnetic separation. Potassium and argon were analyzed by flame photometry using a 2000 ppm Cs buffer and by mass spectrometer with a single collector system using an isotopic dilution method with 38Ar spike, respectively, at Okayama University of Science, following methods described by Itaya et al. (1991). Multiple runs of standard (JG-1 biotite, c. 91 Ma) indicate that analytical error of argon analysis is about 1% at the 2σ confidence level (Itaya et al., 1991). The K–Ar biotite age from sample H2101 yielded the oldest age of 26.8 ± 0.6 Ma, which is older than the U–Pb zircon rim ages (c. 21–16 Ma) from the same rock. Similarly, the K–Ar biotite age of 20.2 ± 0.6 Ma from sample H2908 is similar to within the error, or slightly older than the U–Pb ages of inner rims in zircon (c. 21–18 Ma) from the same rock. Therefore, most K–Ar ages including the ages of c. 20–16 Ma from two orthogneisses (M1806, H90-30) and cordierite-bearing gneiss (H2308) are likely to be apparent,

T. Imayama et al. / Lithos 134-135 (2012) 1–22

15

Fig. 9. Plots of total REE contents against LuN/GdN for dated zircon crystals from samples H2101 and H2710. REE compositions of different domains of zircons are plotted as shown in legends.

Fig. 8. REE chondrite normalized patterns for the different zircon growth domains in samples (a) H2101 and (b) H2710. Gray, dark-gray, and white colors are data-points for the core, inner rim, and outer rim domains, respectively. Data are normalized using the values of chondrite from Sun and McDonough (1989). Data are listed in Table 2.

because the closure temperatures for each system predict the opposite age relationship, which was perhaps caused by the existence of excess argon (e.g. Vannay and Hodges, 1996). On the other hand, sample H2710 collected from the point near the HHT proposed by Goscombe et al. (2006) yielded K–Ar biotite ages of 9.0 ± 0.3 Ma, which is much younger than the U–Pb zircon rim ages (c. 33–23 Ma) from the same rock. We interpret that this age indicates the timing of resetting the K–Ar biotite system in relation to the deformation along the HHT, which occurred at the temperature conditions above or close to the closure temperatures c. 300 °C of K–Ar biotite system (Harrison et al., 1985). 7. Fission track zircon ages In order to estimate the cooling age, zircons were collected from two migmatites (H2101, H2710) for fission track (FT) analyses (Fig. 2). Separated zircons were mounted in a PFA Teflon sheet, polished with diamond pastes, and etched using NaOH–KOH eutectic melts at a constant temperature of 225 °C. The external detector method, described by Danhara et al. (1991), was used for the

fission-track dating. A diallyl phthalate plastic detector was used as the external detector, and placed on each sample mount. Induced tracks were produced by thermal neutron irradiation in a research reactor at the JRR-4 reactor of the Japan Atomic Energy Research Institute, and samples were irradiated together with dosimeter glasses NIST-SRM612. Fission tracks of 30 grains per each sample were counted under a microscope, and age calibration was carried out by the zeta approach (Hurford, 1990) using internal surface of zircon, with a zeta value of 416 ± 3 Ma (Danhara and Iwano, 2009). The age standard Fish Canyon tuff (FCT) sample was used for monitoring the system calibration of the irradiation run. Analytical results are listed in Table 5 and shown in Fig. 10. Zircon grains from kyanite–sillimanite migmatite (H2101) yielded a weighted mean FT age of 6.3 ± 0.2 Ma with the uranium content of zircon grains ranging between 570 and 1680 ppm (Table 5, Fig. 10a). Those from sillimanite migmatite (H2710) yielded a weighted mean FT age of 6.3 ± 0.2 Ma (Fig. 10b) that is same as the age of sample H2101, with the uranium content ranging between 150 and 930 ppm (Table 5). Fission-track data for the zircons failed the χ 2 test (Galbraith, 1981), which can be most probably attributed to the heterogeneous uranium distribution within a crystal (Danhara et al., 1991). On the other hand, the FT ages from 24 grains (80%, sample H2101) and 28 grains (93%, sample H2710) vary within 2σ error (Fig. 10). Hence, these facts probably indicate that the heterogeneous age data do not consist of age groups of more than one different generation, but represent a distinct age group almost following the normal distribution.

Table 4 K–Ar ages of biotites from the Higher Himalayan Crystalline Sequence in far-eastern Nepal. Sample namea

Tectonic unit

K (wt.%, 2σ)

M1806 H2101 H2908 H2710 H2308 H90-30

MCTZ HHCS HHCS HHCS HHCS HHCS

7.009 ± 0.140 7.281 ± 0.146 7.410 ± 0.148 7.314 ± 0.146 7.296 ± 0.146 7.030 ± 0.141

a

Rad. 40Ar (10− 8ccSTP/g, 2σ) 457 ± 12 763 ± 10 585 ± 13 257 ± 6 573 ± 9 445 ± 8

40

Non rad. Ar (%)

K–Ar age (Ma, 2σ)b

55.6 26.8 34.6 53.4 34.8 40.9

16.73 ± 0.56 26.81 ± 0.64 20.23 ± 0.61 9.02 ± 0.28 20.12 ± 0.52 16.23 ± 0.44

Locations for analyzed samples are shown in Fig. 2. The following decay constants were used for age calculation: λe = 0.581 × 10− 10/yr, λβ = 4.962 × 10− 10/yr, and 40K/K = 0.0001167 (Steiger and Jäger, 1977). b

16

T. Imayama et al. / Lithos 134-135 (2012) 1–22

Table 5 Fission-track ages of zircons separated from the migmatitic gneiss, far-eastern Nepal. Sample namea

Spontaneous trackb (Ns) ρs (per cm2)

Induced trackb (Ni) ρi (per cm2)

Dosimeter (Nd) ρd (per cm2)

r

Pr (χ2)c (%)

U (ppm)

Fission-track agesd (Ma ± 1σ)

H2101 H2710

(2158) 3.17 × 106 (2331) 1.53 × 106

(9631) 1.41 × 107 (10,274) 6.74 × 106

(4839) 13.44 × 104 (4824) 13.40 × 104

0.876 0.942

0 0

880 420

6.3 ± 0.2 6.3 ± 0.2

a b c d

Locations for analyzed samples are shown in Fig. 2. ρ and N are fission track densities and total numbers of counted tracks, respectively. Pr (χ2) is probability of χ2 for (number of crystals − 1) degrees of freedom (Galbraith, 1981). Ages were calculated using dosimeter glass NIST-SRM612 and ζED1 = 416 ± 3 (Danhara and Iwano, 2009).

8. Discussion 8.1. Interpretation of U–Pb zircon ages The 207Pb*/ 206Pb* ages of c. 1860 to 1780 Ma (Fig. 6a) for zircon cores from mylonitic augen orthogneiss M1806 are interpreted as inherited ages of magmatic protoliths. Similar protolith ages have been reported by U–Pb zircon ages from the Ulleri augen gneiss (c. 1831 Ma) in the MCT zone (DeCelles et al., 2000) and the granitic orthogneisses (c. 1830 to 1780 Ma) in the stratigraphically lower LHS (Kohn et al., 2010), suggesting that the augen gneiss M1806 along the MCT zone in far-eastern Nepal belongs to the LHS (Imayama and Arita, 2008). In contrast, the 207Pb*/ 206Pb* ages of c. 475–430 Ma (Fig. 6c) for zircon cores in the granitic orthogneiss H90-30 from the upper HHCS are correlated with the U–Pb zircon and monazite ages of c. 470–465 Ma from the Namche Migmatitic Orthogneiss in the HHCS in the Everest area (Viskupic and Hodges,

2001). Hence, although on petrographic basis the granitic orthogneiss in the HHCS is similar to the augen gneiss in the MCT zone, they are clearly distinguished based on geochronology. In eastern Nepal, on the basis of U (–Th)–Pb ages of zircon, monazite, xenotime, and thorium oxide, the two orthogneisses from the Ama Drima Massif in the MCT zone and the overlying HHCS yielded protolith ages of c. 1799 ± 12 Ma with an anatectic age of c.13 Ma, and c. 473 ± 16 Ma with an anatectic age of c. 16 Ma, respectively (Cottle et al., 2009b). Also, Daniel et al. (2003) have reported inherited zircon ages of c. 1840–1760 Ma and monazite ages of c. 1760 ± 7 Ma with metamorphic monazite ages of c. 20–18 Ma from a quartzofeldspathic gneiss from the LHS in eastern Bhutan. These facts indicate that both orthogneisses from the LHS and HHCS with the Paleoproterozoic and Early Ordovician protolith ages, respectively, recorded the Himalayan thermal event in the Tertiary, as also suggested by discordant ages of zircons from sample M1806 and H90-30 (Fig. 6a, c). However, the 207Pb*/ 206Pb* ages of c.

(a) H2101

10

0.30

20 Age histogram

Radial Plot

ages: 6.3 ± 0.2 Ma

0.20

10 0.10

8

+2 0

6

-2

Age (Ma)

Frequency

15

Relative frequency

weighted mean

5

5

Relative standard error (%) 100 50 0 0

2.5

5

7.5

0.00 10

25 20

10

4

0 1 2 3 4 5 6 7 8 9 10 11

Age (Ma)

Precision

(b) H2710 20

0.40

Age histogram

Radial Plot 10

ages:

0.30

6.3 ± 0.2 Ma 10

0.20

5

8

+2 0 6

-2

Age (Ma)

Frequency

15

Relative frequency

weighted mean

5

0.10

Relative standard error (%) 100 50 0

0

2.5

5

Age (Ma)

7.5

0.00 10

2520

10

4

0 1 2 3 4 5 6 7 8 9 101112

Precision

Fig. 10. Age frequency histogram with age spectra (left) and the radial plot (right) for fission track zircon data obtained from samples (a) H2101 and (b) H2710 along the Tamor– Ghunsa section, far-eastern Nepal. Age spectra are presented as curves in the histogram. Weighted mean ages are quoted at 95% confidence level. Data are listed in Table 5.

T. Imayama et al. / Lithos 134-135 (2012) 1–22

1800 to 1620 Ma (M1806) and c. 475 to 420 Ma (H90-30) for recrystallized domains that have dark-CL images are indistinguishable from those in the primary oscillatory-zoned core within the error (Fig. 6a, c), indicating that the U–Pb ages in the zircon were retained during the recrystallization, probably due to the high closure temperature of zircons (>c. 900–1000 °C; Cherniak and Watson, 2000). Garnet-mica gneiss (H1205) and three migmatites (H2101, H2908, and H2710) contain the pre-Himalayan detrital zircon cores, showing high Th/U ratios of c. 0.02–1.52 (Table 1) and high REE concentrations with steep MREE to HREE slopes (Figs. 8, 9), which are generally found in zircon of igneous origin (Hoskin and Schaltegger, 2003). Detrital zircon cores show the variable ages, but a relative probability histogram of the 207Pb*/ 206Pb* ages provides the main cluster at c. 1000 to 800 Ma (Fig. 11a), consistent with the previously reported detrital zircon age data for provenance of the HHCS (e.g. DeCelles et al., 2000; Martin et al., 2005; McQuarrie et al., 2008). Also, inherited cores in zircon from leucogranite H2708 show similar variations in U–Pb ages and Th/U values (Table 1), which is interpreted as the entrainment of older grains from source rocks of the magma. Petrological P–T grids for garnet-mica gneiss H1205 predict that no melt phase forms at peak conditions of c. 640 °C at c. 11 kbar (Imayama et al., 2010). Therefore, the two rims of the dark-CL domain

(a) Core 14 13 12

M1806

11

H1205

10

H2101 H2908

Frequency

9

H2710

8

H2708

7

H90-30

6 5 4 3 2

3400

3200

3000

2800

2600

2400

2200

2000

1800

1600

1400

1200

800

1000

600

400

0

200

1

207

Pb*/206Pb* age (Ma)

(b) Rims 9 8

Frequency

7

H1205 H2101 H2908 H2710 H2708 inner rim outer rim

6 5 4 3 2 1 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33 34 35 36

0

206

Pb*/238U age (Ma)

Fig. 11. Cumulative probability density plots for (a) 207Pb*/206Pb* ages of cores and (b) 206Pb*/238U ages of rims. Ages of different samples and domains of zircons are plotted as shown in legends.

17

with low Th/U ratio in sample H1205 could have been formed at subsolidus conditions. The low Th/U ratio in zircons is often used as indicator of metamorphic or anatectic zircons during high-grade metamorphic conditions (e.g. Hoskin and Schaltegger, 2003). Assuming that the new growth of zircon rims occurred during prograde or peak metamorphism, the garnet-mica gneiss had experienced the increasing temperature paths (c. 590 °C to 640 °C) during both loading and decompression, as inferred from garnet growth zoning (Imayama et al., 2010), at c. 25–20 Ma. This is in agreement with the reverse-sense movements on the MCT at c. 24–17 Ma under temperature conditions of c. 500–600 °C in eastern Nepal (Catlos et al., 2002; Hubbard and Harrison, 1989). Kyanite–sillimanite migmatitic gneiss (H2101) and two sillimanite migmatitic gneisses (H2908, H2710) recorded the two metamorphic overgrowth stages. In addition, low Th/U ratios (b0.02) and MREE contents (Figs. 8, 9) of zircon rims compared with detrital cores (Th/U: 0.02–1.52) are most likely related to partial melting processes of the host rocks (Rubatto, 2002). This is because the MREE relative to HREE and Th relative to U are mainly partitioned into magmatic zircons when compared to the metamorphic zircons (Hoskin and Schaltegger, 2003; Orejana et al., 2011). Alternatively, the coexistence of zircon and monazite could explain low Th/U ratio and low MREE contents detected in zircon, because monazite preferentially hosts Th and L- to MREE (e.g. Harley et al., 2007). Presence of monazite in matrix of samples H2101 and H2710 could be a possible reason for low Th/U ratio and MREE content in metamorphic zircon rims. However, monazite also occurs in the zircon cores of samples H2101 and H2710, showing the co-existence of these two phases during pre-Himalayan magmatic activities as well as Tertiary metamorphism. Thus, the former interpretation is more likely to explain the differences in Th/U ratio and MREE content between the cores and metamorphic rims. The U–Pb analyses of inner rims from samples H2101 and H2710 yielded the 206Pb*/ 238U ages at c. 21–18 Ma and c. 33–28 Ma, respectively. Low HREE content with a nearly flat MREE to HREE slope (LuN/ GdN = 1–2 for H2101, 1–4 for H2710) of these inner rims (Figs. 8, 9) is attributed to the presence of garnet (e.g. Orejana et al., 2011; Whitehouse and Platt, 2003), which hosts substantial MREE to HREE. Xenotime also hosts a large amount of HREE, but it is not observed in the two migmatites. Hence, the inner zircon rims with low HREE content probably grew in equilibrium with garnet: their formation can be explained either by subsolidus reaction with metamorphic fluid (e.g. Dempster et al., 2008) or by reaction with anatectic partial melts (e.g. Rubatto, 2002; Vavra et al., 1996). Homogenous compositions of garnets in migmatites H2101 and H2710 indicate that these garnets in their migmatites grew at c. 750 °C (Imayama et al., 2010). Since pseudosection P–T analyses for samples H2101 and H2710 indicate that melt phase could have been extensively present at temperatures above c. 650 °C (Fig. 4), the garnets in equilibrium with the inner rims of zircons must have grown in the presence of melt, not at subsolidus conditions. In fact, the peritectic garnet such as garnet–quartz intergrowths is present in sample H2710 (Fig. 3e), implying the presence of melt phase during garnet growth (e.g. Waters, 2001). The cores with irregular shape observed in zircons from migmatite samples (Fig. 5e, i, j) indicate that the inherited zircon was melted during partial melting of the rocks, and behaved as the source of Zr for inner zircon rims. As a result, the inner zircon rims were overgrown from partial melt when zirconium was simply saturated in the melt. Based on the pseudosection P–T analysis, the two migmatites H2101 and H2710 are interpreted to record muscovite dehydration melting, rather than biotite dehydration melting (Fig. 4). The 206Pb*/ 238 U ages (c. 21–18 Ma) of inner rim in zircon from muscovite-free sillimanite migmatite (H2908) directly above muscovite-out isograd are similar to those (c. 21–18 Ma) of muscovite-bearing kyanite–sillimanite migmatite (H2101) directly below the isograd within the error. Perhaps, melting could have continued until muscovite was completely

T. Imayama et al. / Lithos 134-135 (2012) 1–22

Temperature-time path for the HHCS leucogranite emplacement (H2708)

800

H2710

H2908

approx. grt breakdown

H2710

C/ M y

700

H2101

c. 4

0

o

600 500

H2101

y /M y o C M o C/ 0 3 c. . 25 y c o C/M 5 1 . c

H1205

400 H90-30

300

H2710

M1806

100 0 0

excess Ar H2101

H2308

y

H2710

200

H2908

H2101

M

Anatexis during Early Oligocene has been determined for kyanite– migmatite exposed in the middle HHCS in eastern Nepal (Groppo et al., 2010). These authors reported partial melting at c. 31 Ma shown by monazite crystallization from the melt, when the garnet rims were growing at peak P–T conditions of c. 820 °C at c. 13 kbar. These ages can be correlated with those related with growth of the inner zircon rims (c. 33–28 Ma) at peak P–T conditions of c. 730–780 °C at c. 7–10 kbar for sillimanite migmatitic gneiss H2710. Subsequent growth of the outer zircon rims took place at c. 27–23 Ma and P–T conditions of c. 725 °C and bc. 7 kbar, in the presence of melt and in absence of stable garnet. These data indicate Early Oligocene anatectic melting in the middle HHCS in eastern to far-eastern Nepal was extensive and was possibly coeval with the Eohimalayan, M1 event. However, our data suggest that the Eohimalayan, M1 event during the Early Oligocene is not always characterized by the kyanite-grade metamorphism (i.e. intermediate P/T-type metamorphism), rather we show that another anatectic event that was recorded in the middle HHCS. Both leucogranite dyke (H2708) and kyanite–sillimanite and sillimanite migmatites (H2101, H2908) formed during partial melting at c. 21–18 Ma, based on the similarity of ages between them. The decompression path of sample H2101 (Fig. 4a), with steep dP/dT slope from the peak-conditions, implies that melting during decompression could have been a continuous process, resulting in the regional exposure of Early Miocene leucogranite plutons (e.g. Hodges, 2000, and references therein). Subsequently, the migmatites cooled at the middle HHCS, leading to the crystallization of outer rim zircons at c. 18–16 Ma and P–T conditions of c. 650–700 °C and bc. 4–6 kbar. On the other hand, the leucogranites intruded at c. 16.3 ± 0.2 Ma into the Early Oligocene sillimanite migmatitic gneiss H2710 at the relatively high structural level of the HHCS. These facts indicate that the leucogranite melt was segregated and migrated upwards for a long distance from the location of melt production. Consequently, the present study provides for the first time reliable evidence for two

45 o C/

8.2. Two-stage partial melting and the different cooling histories within the HHCS

discontinuous multiple melting events occurred in the time interval from 33 to 16 Ma (>17 Myr) within the HHCS in far-eastern Nepal. Different cooling histories for two migmatites are constrained by three points of the FT zircon age, and U–Pb zircon ages of inner and outer rims. The closure temperature of FT zircon ages is considered as c. 280 °C (Tagami et al., 1998) for geologically reasonable cooling time scales of 1–10 Ma. Inner rim zircons in samples H2101 and H2710 grew at c. 21–18 Ma under the peak-T of c. 720–770 °C and at c. 33–28 Ma under the peak-T of c. 730–780 °C, respectively, as discussed above. We take the ranges of the ages and peak temperatures to calculate the maximum and minimum values (i.e. possible ranges) of cooling rates. The average cooling rates from peak temperatures, through those for the garnet breakdown, to the closure temperature of FT zircon ages are calculated as c. 30–40 °C/My for H2101 and c. 15–25 °C/My for H2710, respectively (Fig. 12). Here, the estimated ranges of cooling rates conform to the ranges of the ages and temperature conditions for the garnet breakdown. However, since the cooling rates are perhaps not constant but vary with time, the cooling ages at temperatures between 300 and 600 °C need to be estimated by other methods of radiometric dating to better constrain the cooling curves. Assuming a geothermal gradient of c. 30 °C/km during exhumation, inferred from the P–T paths shown in Fig. 4 (density of metamorphic rocks assumed to be 2.8 g/cm 3), the exhumation rates are calculated to be c. 1.0–1.3 mm/yr and c. 0.5–0.8 mm/yr for the Early Miocene (H2101) and Early Oligocene (H2710) migmatites, respectively, indicating a much higher rate in the former than in the latter. The faster exhumation rate of the Early Miocene than Early Oligocene migmatites is consistent with the higher dP/dT slope of exhumation P–T path of the former than the latter migmatites (Fig. 4). Same zircon FT ages (c. 6 Ma) of both migmatites indicate that they had been juxtaposed by c. 6 Ma, and then have rapidly cooled since the Late Miocene with the same cooling rates of c. 45 °C/My (Fig. 12). The results indicate that these migmatites were exhumed at rates of c. 1.5 mm/yr, i.e. higher than those before juxtaposition.

Zr U-Pb (peak) Zr U-Pb (retrograde) Bt K-Ar Zr FT

c.

consumed, resulting in the formation of sillimanite–garnet–biotite migmatites that do not include primary muscovite (e.g. Searle et al., 2010). The 206Pb*/ 238U ages of outer rims in zircons from samples H2101 and H2710 have been dated at c. 18–16 Ma and c. 27–23 Ma, respectively. Relatively enriched HREE with moderate values of LuN/GdN (8–18 for H2101, 3–11 for H2710) of outer rims in zircons, compared to those of inner rims (Figs. 8, 9), could be explained by the disappearance of garnet from the equilibrium mineral assemblages with zircon. This is because based on the experiments of zircon/garnet trace element partitioning, MREE equally partitions between the two phases, whereas with increasing atomic number the HREE prefers zircon to garnet (Rubatto and Hermann, 2007). Garnet breakdown occurred by the retrograde reaction with residue melt (Grt + Kfs + melt = Bt + Pl + Sil + Qtz) on cooling at c. 675–725 °C below c. 6–7 kbar (Fig. 4). Zircon is abundant in the reaction domain of garnet rims in sample H2710 (Fig. 3g), and the microstructural setting of zircon implies that the garnet breakdown could have promoted the growth of the outer rim zircons. Hence, ages of the outer zircon rims date zircon growth during melt crystallization at retrograde stage outside the stability field of garnet. The mean 206Pb*/ 238U age of 16.3 ± 0.2 Ma for outer rim zircon in leucogranite H2708 represents the timing of new zircon growth by melt crystallization on cooling, based on the presence of oscillatory zoning (Fig. 5m). In contrast, the Th/U ratios in inner overgrown rims, showing the mean 206Pb*/ 238U age of 18.3 ± 0.3 Ma, are significantly low (c. 0.03), compared to the values of the inherited core. Such low Th/U ratio in zircon in granites could be associated with the anatectic melts (e.g. Williams et al., 1996).

Temperature (oC)

18

5

10

15

20

25

30

Time (Ma) Fig. 12. The integrated U–Pb zircon, K–Ar biotite and fission track zircon age data for the HHCS along the Tamor–Ghunsa section, far-eastern Nepal. Yellow and purple shades bounded by thick, dashed lines indicate the cooling curves with the possible range of the inferred cooling rates for samples H2101 and H2710, respectively. The ranges of the U–Pb zircon ages and peak temperatures are shown by bars. The cooling curve with the inferred cooling rate for H2101 and H2710 since c. 6 Ma is shown by thick solid line. See Fig. 2 for locations of samples. Ages of different geochronometries are plotted as shown in legends. See text for discussion.

T. Imayama et al. / Lithos 134-135 (2012) 1–22

19

8.3. Tectonic implications The potential role played by channel flow in exhumation of the Himalaya orogenic belt (i.e. extrusion process) is a highly debated topic (Beaumont et al., 2001, 2004; Grujic et al., 2002; Jamieson et al., 2004). Channel flow is inferred to occur within a weak midcrustal layer between rigid upper- and lower-crustal layers, where continuous rapid heat advection within the channel enables the rocks to continue to be hot and ductile (e.g. Beaumont et al., 2004). Jamieson et al. (2004) predicted in their channel flow models that the Oligocene partial melting of c. 30 Ma occurs at the maximum burial depth (Pmax) related to Eohimalayan metamorphism (M1), followed by lateral return flow related to Neohimalayan metamorphism (M2), which occurs at the maximum temperature (Tmax). Although the formation of Early Oligocene migmatites (H2710) at P = c. 7–10 kbar in far-eastern Nepal is consistent with the prediction within the errors of the P–T estimations, the single continuous episode of channel flow model cannot reproduce the further increasing of pressures after the M1 for the formation of Early Miocene migmatites (H2101) at P = c. 8–14 kbar. In addition, note that the cooling rates for the HHCS in far-eastern Nepal were not as high as the single channel flow models predicted, by which a very rapid cooling rate of c. 75–100 °C/My results from rapid exhumation (2–3 km Myr − 1) since c. 10 Ma (e.g. Jamieson et al., 2004). This discrepancy may be attributed to the fact that the presence of mechanically weak brittle–ductile shear zones such as the HHT within HHCS is not taken into account in the original channel flow models. The strain localization by softening (i.e. formation of shear zones) within the HHCS allows the relatively slow exhumation under relatively low temperatures and hence slow cooling rates of the HHCS, compared to those forced by rapid heat advection and high denudation rates at the surface assumed in the original channel flow models (Beaumont et al., 2004). The coeval thrusting on the MCT and the extension on the STD at the Early to Middle Miocene in eastern Nepal are generally believed to be responsible for the southward extrusion of the entire HHCS in the convergent orogen. However, the different P–T–t paths from two migmatites in far-eastern Nepal suggest two exhumation episodes within the HHCS rather than the single exhumation episode for the entire HHCS only bounded by the MCT and STD. The first exhumation of the HHCS is shown by the cooling P–T path of Early Oligocene migmatites: during exhumation of these migmatites, Early Miocene migmatites were buried. The discontinuity of the P–T–t paths in the HHCS well above the MCT and near the HHT, is probably due to the thrust displacement along the HHT, rather than movement on the MCT. In order to interpret the two stage partial melting and the relation to the HHT, here we present the following scenario. After partial melting at c. 33–28 Ma during Early Oligocene recorded by U–Pb ages of the inner rim zircon from the migmatites (H2710) (Fig. 13a), a kinematic change from progressive burial to return flow of the partially molten rocks (Fig. 13b) could have been induced by the rheological weakening resulting from initial melting reaction (e.g. Hollister and Crawford, 1986). Under such circumstances, the thrust (i.e. the HHT) is easily formed within the HHCS, which brings about exhumation of the Early Oligocene migmatites (H2710) and contemporaneous burial of the Early Miocene migmatites (H2101) (Fig. 13b). This may be similar to the relationship between the tectonic loading and the return flow of the previously buried rocks occurring as corner flow within the orogenic wedges (e.g. Harris and Massey, 1994). Since the HHT is a wide shear zone at depth, not a single narrow thrust, it is likely that the Early Oligocene migmatite (H2710) mapped directly below the HHT constituted a part of the brittle–ductile shear zone, which can be evidenced by the resetting of K–Ar ages (i.e. 9.0 ± 0.3 Ma). During the Late Oligocene, the Early Oligocene migmatite (H2710) were exhumed by ductile extrusion along the

LHS rocks sense of movement along faults movement of rocks

(a) c. 33-28Ma

0 20

? 40 (km)

MHT

(b) c. 27-23Ma

HH

T

0

ST

D?

20 40 (km)

MHT

(c) c. 21-19Ma

HH

0

T

MC

ST

T

D

20

MHT

40 (km)

(d) c. 18-16Ma

0

MC

ST

T

HH T

D

20

MHT

40 (km)

(e) c.16-6Ma MC

T

HH T

STD

0 20

MHT

40 (km)

Fig. 13. Conceptual model of tectono-magmatic process associated with two-stage partial melting in far-eastern Nepal Himalaya, based on the different P–T–t paths of two migmatites. White, black, and gray circles represent the Early Oligocene migmatites, Early Miocene migmatites, and LHS rocks, respectively. Partially melted rocks are shown by gray color. (a) c. 33–28 Ma, showing the Early Oligocene partial melting. (b) c. 27–23 Ma, showing exhumation of the Early Oligocene migmatites by ductile extrusion along the HHT. This caused in situ crystallization of the residue melt in the migmatites on cooling, whereas the melt separated from the migmatites intruded as leucogranite dykes. (c) c. 22–19 Ma, showing the development of normal faults along the STD and thrusting along the MCT and/or the HHT, and the production of the Early Miocene leucogranites. The rocks buried by the overthrusting of the Early Oligocene migmatites were partially melted at depth. (d) c. 18–16 Ma, showing the extensional reactivation of the HHT, leading to the rapid exhumation of the Early Miocene migmatites and the intense magmatic intrusion. (e) c. 16–6 Ma, showing the continued extensional reactivation of the HHT, leading to the juxtaposition of the two migmatites formed at different ages. The LHS rocks were deeply buried beneath the HHCS along the MCT.

HHT, which caused in situ crystallization (i.e. growth of outer rims in zircons at c. 27–23 Ma shown by U–Pb ages) of the residue melt in the migmatites on cooling (Fig. 13b). In fact, Late Oligocene hightemperature shear zone in the structurally higher level of the HHCS has been recently identified in western Nepal (Carosi et al., 2010). Although the Late Oligocene leucogranite plutons are rarely exposed in eastern to far-eastern Nepal, the Late Oligocene leucogranite plutons of c. 28–27 Ma have been reported at the North Himalayan antiform in southern Tibet (Zhang et al., 2004). A large amount of leucogranites was produced during the Early Miocene, as shown by the leucogranite plutons with ages ranging

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between c. 24–17 Ma in eastern Nepal (e.g. Harrison et al., 1999; Hodges et al., 1998) and c. 23–16 Ma in southern Tibet (Lee and Whitehouse, 2007; Lee et al., 2006). This melting is presumably related to muscovite dehydration reaction in low, deep-seated, structural level represented by samples H2101 and H2908 (c. 8–11 kbar; c. 25–35 km): this occurred at c. 21–18 Ma, as shown by U–Pb ages of the inner rim zircon (Fig. 13c). Subsequently, the extensional reactivation on the HHT could have thinned the Himalayan crust, probably accompanied by further melting at shallower depths (c. 4–6 kbar; c. 15–20 km) due to decompression, which led to the rapid exhumation of the lower structural level of the HHCS (Fig. 13d). The exhumation caused in situ crystallization (i.e. growth of outer rims in zircons at c. 18–16 Ma shown by U–Pb ages) of the residual melt in the migmatites (H2101) on cooling (Fig. 13d). A few data indicate that movement along the Kakhtang thrust in Bhutan, which can be correlated with the HHT, began at c. 18–16 Ma and lasted until c. 13 Ma (Daniel et al., 2003). The Early Miocene migmatites are inferred to have been exhumed more rapidly than the Early Oligocene migmatites, as suggested by the dP/dT slope of uplift P–T paths and by the cooling rates of respective migmatites. Juxtaposition of the two migmatites was probably accompanied further extensional movements along the HHT. Concurrently, the LHS rocks were buried beneath the HHCS along the MCT (Fig. 13e). Although timing of juxtaposition of two migmatites can be as early as c. 16 Ma, the same zircon FT age from both migmatites at least indicates that the relative tectonic movement between the two migmatites with the different P–T–t paths terminated by c. 6 Ma (Fig. 13e). 9. Conclusions U–Pb ages and trace element compositions of zircons, together with P–T pseudosection from the HHCS migmatites and associated leucogranite reveal two discontinuous multiple melting events during the time interval from c. 33 to 16 Ma in far-eastern Nepal. Zircons from two migmatites consist of the inherited magmatic cores with high REE contents and high Th/U ratio, and two successively overgrown metamorphic rims with low MREE content and low Th/U ratio. Inner zircon rims with flat slope of the MREE to HREE were in equilibrium with garnet in the presence of melt, and the U–Pb ages represent those at the muscovite dehydration melting during Early Oligocene (c. 33–28 Ma: subaluminous sillimanite migmatite) at P = c. 7–10 kbar and T = c. 730–780 °C, and during Early Miocene (c. 21–18 Ma: aluminous kyanite–sillimanite migmatite) at P = c. 8–14 kbar and T = c. 720–770 °C. Outer rims of zircons have relatively enriched HREE, compared to inner rims, and the U–Pb ages from the two migmatites are c. 27–23 Ma (Late Oligocene) and c. 18–16 Ma (Early Miocene), respectively. These ages are related to melt crystallization accompanying garnet breakdown on cooling at P = c. 4–7 kbar and T = c. 650–725 °C. Zircons from muscovite–biotite leucogranite crosscutting the sillimanite migmatitic gneiss also have two successively overgrown rims with ages of 18.3 ± 0.3 Ma associated with the anatectic melting, and 16.3 ± 0.2 Ma with melt crystallization on cooling, respectively. The formation of leucogranite was genetically related to that of Early Miocene migmatites, because both of them were formed at similar ages. Nearly isothermal P–T paths of Early Miocene migmatites, in contrast to the relatively isobaric P–T paths of Early Oligocene migmatites during exhumation, suggest the faster exhumation rates of the former than the later migmatites. The inferences are consistent with higher average cooling rates for the Early Miocene (c. 30–40 °C/My) than the Early Oligocene (c. 15–25 °C/My) migmatites, inferred from peak-T conditions and FT (c. 6 Ma for both migmatites) and U–Pb zircon ages. Based on the P–T–t paths of the two migmatites, it is likely that the Early Miocene migmatites were buried when the Early Oligocene migmatites were exhumed. The fact suggests that the return flow of the previously buried rocks at the upper structural level

caused tectonic loading on rocks at the lower structural level by overthrusting. The two-stage partial melting and the different cooling history for the two migmatites cannot have been caused only by the single channel flows of the entire HHCS as proposed by Beaumont et al. (2004), suggesting that the formation of large-scale thrust (i.e. HHT) within the HHCS is essential to generate the two migmatites. Acknowledgment We wish to thank K. Suzuki, T. Kato and their colleagues (Nagoya University) for useful discussion on this study. We also greatly thank T. Danhara and H. Iwano (Kyoto Fission Track Corporation) for fission track zircon analyses and H. G. Kim (Korea Basic Science Institute) for U–Pb zircon analyses using the SHRIMP instrument. We appreciate the reviews by three anonymous reviewers, which have greatly improved the quality of manuscript. The authors also thank M. Scambelluri for editorial handling of the manuscript and correction of the English text. 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