Thermal evolution of the Lesser Himalaya, central Nepal: Insights from K-white micas compositional variation

Gondwana Research 9 (2006) 409 – 425 www.elsevier.com/locate/gr

Thermal evolution of the Lesser Himalaya, central Nepal: Insights from K-white micas compositional variation Lalu Prasad Paudel a,⁎, Kazunori Arita b a

Research Institute of Natural Sciences, Okayama University of Science, 1-1 Ridai-cho, Okayama 700-0005, Japan b Hokkaido University Museum, Kita 10, Nishi 8, Sapporo 060-0810, Japan Received 27 May 2005; accepted 6 January 2006 Available online 23 March 2006

Abstract The Lesser Himalayan low- to medium-grade metamorphic rocks in central Nepal are rich in K-white micas occurring as porphyroclasts and in matrix defining S1 and S2. Porphyroclasts are usually zoned with celadonite-poor cores and celadonite-rich rims. The cores are the relics of igneous or high grade metamorphic muscovites, and the rims were re-equilibrated or overgrown under lower T metamorphic conditions. The matrix K-white micas defining S1, pre-dating the Main Central Thrust activity, are generally celadonite-rich. They show heterogeneous compositional zoning with celadonite-rich cores and celadonite-poor rims. They were recrystallized at lower T condition prior to the Main Central Thrust activity, most probably prior to the India–Asia collision (pre-Himalayan metamorphism). The matrix K-white micas along S2, synchronous to the Main Central Thrust activity (Neohimalayan metamorphism), are relatively celadonite-poor and were recrystallized under relatively higher T condition. K-white micas defining S1 also were partially re-equilibrated during the Neohimalayan metamorphism. The average compositions of recrystallized K-white micas defining both S1 and S2 become gradually poor in (Fe + Mg)- and Si-contents and rich in Al- and Ti-contents from south to north showing an increase of metamorphic grade from structurally lower to higher parts in the Lesser Himalaya. This shows that the metamorphism is inverted throughout the inner Lesser Himalaya. The tectono-metamorphic significance of the published K–Ar and 40Ar / 39Ar Kwhite micas ages from the Lesser Himalaya need re-evaluation in the context of observed intrasample compositional variation and zoning, and possible higher closure temperature (∼500°C) for K–Ar system. © 2006 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved. Keywords: K-white micas; Low-grade metamorphism; Inverted metamorphism; Lesser Himalaya; Nepal

1. Introduction The Lesser Himalaya (LH) in central Nepal consists of low- to medium-grade Late Proterozoic–Early Paleozoic metasediments (Stöcklin, 1980; Parrish and Hodges, 1996) which were accreted onto the Himalaya in the Cenozoic (DeCelles et al., 2001; Robinson et al., 2003). The wedge of the LH metasediments occupies the position between high-grade (amphibolite to granulite facies) Proterozoic rocks of the Higher Himalaya in the north and the Tertiary molassic sediments of the SubHimalaya (Siwaliks) in the south (Fig. 1). Two major intracontinental thrusts, the Main Central Thrust (MCT) and the Main Boundary Thrust (MBT), are the major tectonic boundaries ⁎ Corresponding author. Permanent address: Central Department of Geology, Tribhuvan University, Kirtipur, Kathmandu, Nepal. Tel./fax: +81 86 2569669. E-mail address: [email protected] (L.P. Paudel).

limiting the LH to the north and south, respectively (Gansser, 1964). At most locations, the footwall rocks of the MCT display an inverted metamorphic zonation, i.e., an increase in metamorphic grade towards structurally higher levels (Le Fort, 1975; Arita, 1983; Pêcher, 1989). In the past three decades, quite a number of works have been carried out on the metamorphism of the relatively higher grade rocks of the Higher Himalaya and the MCT zone in central Nepal (e.g., Le Fort, 1975; Pêcher, 1975; Arita, 1983; Kaneko, 1995; Vanny and Hodges, 1996; Hodges et al., 1996; Catlos et al., 2001; Kohn et al., 2001). Those studies provided numerous structural, petrological and geochronological data from the area and enabled the researchers to propose a number of thermotectonic models for the Himalayan evolution (e.g., Sorkhabi and Arita, 1997; Harrison et al., 1998; Guillot and Allemand, 2002; Bollinger et al., 2004 and references there in). In general, there are two categories of models to explain the Himalayan inverted

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Fig. 1. Generalized geological map of the Nepal Himalaya showing the location of the present study area (modified after Upreti, 1999). MCT: Main Central Thrust, MBT: Main Boundary Thrust.

metamorphism. One set of models consider inversion of isotherm during MCT movement (Le Fort, 1975; Arita, 1983; Molnar and England, 1990; Guillot and Allemand, 2002) and the other set of models suggest that the inverted metamorphism is apparent due to the inversion of isograds by post-metamorphic kinematics (Searle and Rex, 1989; Jain and Manikavasagam, 1993; Hubbard, 1996; Harrison et al., 1998; Catlos et al., 2001; Bollinger et al., 2004). The LH occupies the footwall of a major intracontinental thrust of the Himalaya, i.e., the MCT, and comprises a sequence of low- to medium-grade metasediments extending for about 90 km N–S section, which are believed to have been metamorphosed and sheared during the MCT activity. Therefore, the metamorphic information from those rocks is critical to test the models of thermo-tectonic evolution of the Himalaya. However, conventional methods of petrographic characterization and thermodynamic calculations are not possible in those rocks because of their fine-grained nature and lack of appropriate mineral assemblages. Paudel and Arita (2000, in press) presented a qualitative thermal profile for the whole section of the LH in central Nepal using crystallinity and b-spacing values of white micas. They suggested that the low-grade sediments record at least two thermal events (pre- and syn-MCT events), and the thermal structure is inverted throughout the LH. Beyssac et al. (2004) and Bollinger et al. (2004) estimated peak metamorphic temperature in some sections of central and western Nepal using Raman spectroscopy of carbonaceous materials (RSCM) and showed that apparent thermal gradient is very steep and inverted. They also suggest that the metamorphism occurred mainly during and after the MCT activity and the inverted thermal structure is due to post-metamorphic deformation of isograds. In the context of contrasting observations, further investigation of those metasediments from different approach may help in better understanding of the low-grade metamorphic evolution of the LH.

K-white micas (KMs) are among the most abundant phyllosilicates in metasedimentary rocks of almost all the bulk compositions. They occur in a wide range of metamorphic conditions, and exhibit extensive compositional variations strongly controlled by the degree of metamorphism (Guidotti and Sassi, 1998). They display a number of chemical substitutions, with complete solid-solution between idealized muscovite (Al-rich) and celadonite (Fe-rich) end members, partial solid-solution between muscovite and paragonite, and substitution by Ti in the octahedral sites (Guidotti, 1984). Among them, the Tschermak substitution [(Mg, Fe2+)vi, Siiv = Alvi, Aliv], which controls the chemistry of phengites, is thought to be particularly sensitive to metamorphic conditions (Guidotti, 1984). Generally, celadonite-rich KMs are stable at lower temperatures and higher pressures (Ernst, 1963; Velde, 1965; Butler, 1967; Powell and Evans, 1983). With increasing T, KMs become celadonite-poor, and Al- and Ti-rich and serve as a metamorphic indicator in low-grade metamorphic rocks, provided they satisfy certain petrological conditions (Guidotti and Sassi, 2002). The present work aims at using KMs compositions to investigate low- to medium-grade metamorphism in the Nepal LH. Metasediments and granitic augen gneisses were collected from Tansen–Pokhara section of central Nepal and KMs occurring in various textural forms were analysed and used to interpret the thermal evolution and paleothermal structure. Mineral abbreviations are used after Kretz (1983). 2. General geology and metamorphism The LH in the Tansen–Pokhara area of central Nepal is divided into several tectonic packages by a series of northdipping thrusts and faults (Figs. 2–4). Basically, it may be divided into the inner (north) and outer (south) belts by the Bari Gad-Kali Gandaki Fault (BKF) (Arita et al., 1982; Sakai, 1985). The outer belt in the southernmost part of the

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Fig. 2. Tectonic map of the Tansen–Pokhara area in central Nepal (modified after Paudel and Arita, 2000). Bt isograd is shown. Grt and Ky isograds coincide with the Lower and Upper MCTs, respectively. STDS: South Tibetan Detachment System, MCT: Main Central Thrust. X–Y: line of cross-section in Fig. 4.

LH is a Parautochthon (Fuchs and Frank, 1970) overlain by the Palpa Klippe. The Palpa Klippe is a tectonic outlier of the LH rocks on the Tansen Group (Sakai, 1983). However, its root thrust in the north is unclear. The inner belt consists of two thrust sheets (Thrust Sheets I and II) and the ductile shear zone of the MCT (MCT zone). The Thrust Sheet I (TS I) and the Thrust Sheet II (TS II) are separated by the Phalebas Thrust (Upreti et al., 1980). The MCT zone is bounded by the Upper MCT (MCT of Le Fort, 1975) in the north and Lower MCT in the south (Paudel and Arita, 2000). The MCT zone consists of the sheared and recrystallized rocks of the LH affinity (Nawakot Complex) forming the footwall of the Upper MCT and lying below the Ky isograd. The LH rocks are overthrust by the Higher Himalayan rocks along the Upper MCT. The LH rocks are basically composed of the Late Proterozoic–Early Paleozoic Nawakot Complex (Stöcklin, 1980; Parrish and Hodges, 1996) and the Late Carboniferous–Early Miocene Tansen Group (Sakai, 1983). The Tansen Group comprises marine and fluvial sedimentary rocks exposed only in the Parautochthon (Fig. 2). The Nawakot Complex is distributed throughout the LH and comprises low- to medium-grade metasediments such as slates, phyllites, schists, metasandstones, quartzites, dolomites and marbles. Bands of granitic augen

gneisses (Ulleri augen gneiss of Le Fort, 1975) are found in the MCT zone. The LH is a complex fold-and-thrust belt with a polyphase deformational history (Paudel and Arita, 2000). Although there are some lines of evidence for pre-Himalayan deformational structures, main deformational features are of the Himalayan stage (Pêcher, 1977). Mainly two prominent foliations can be observed in the inner belt of the LH (Fig. 3). The older one is a bedding-parallel foliation (S1 = S0) predating the MCT activity. The younger one (S2) is either a crenulation cleavage or a shear foliation. S1 is dominant in the TS I and II but it can be observed only as a relic fabric in the MCT zone. S2 is prominent in the MCT zone and cross-cuts S1 but becomes weak or vanishes as going southward from the MCT zone. Several features such as the parallelism of S2 with the Upper MCT (Fig. 3), confinement of stretching and mineral lineations and snowball garnet along S2 indicate that S2 was formed as a result of southward movement of the Higher Himalayan thrust sheet over the LH along the Upper MCT (Bouchez and Pêcher, 1981; Brunel, 1986). The study area is one of the best sections of the LH showing inverted metamorphic zonation, with metamorphic grade increasing from Chl zone at structurally deepest part of the section to Bt, Grt and Ky zones at successively higher structural levels (Le Fort, 1975; Pêcher, 1977, 1989; Arita, 1983; Kaneko, 1995;

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Fig. 3. (A) Detailed geological and structural map of the Syangja–Pokhara area (modified after Paudel and Arita, 2000) with locations of the samples. Foliations and lineations are projected on the Schmidt's lower hemisphere. S1 and S2 (=S3 in Paudel and Arita, 2000) represent older and younger foliations, respectively, but do not necessarily denote the first and second deformation events. Note that S2 is homoclinally dipping to the NNE parallel to the Upper MCT while S1 is folded. Enlarged route maps with sample locations along the Seti River north of Pokhara (B) and Syangja–Haripala motor road (C) are shown in the insets. MCT: Main Central Thrust.

Hodges et al., 1996) (Figs. 2 and 4). The peak-metamorphic temperatures have been estimated to be ca. 325 °C for the Parautochthon and TS I and ca. 400 °C for the TS II on the basis of illite crystallinity and quartz microstructures (Paudel, 2002). Peak temperatures estimated using RSCM method range from b 330 to 340 °C for the Parautochthon and TS II, and 500 to 550 °C for the MCT zone (Beyssac et al., 2004). Garnet–biotite thermometry also yields similar range of temperatures (500 to 600 °C) for the MCT zone (Le Fort et al., 1986a; Kaneko, 1995;

Vanny and Hodges, 1996). Pressure has been estimated to be about 4–5kbar for most parts of the LH on the basis of white micas b-spacing values (Paudel and Arita, in press). 3. Petrography of studied samples Twenty six phyllosilicate–rich metapelites (phyllites and schists) and one metasandstone samples were collected along the Syangja–Pokhara motor road and the Seti Khola valley

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Fig. 4. Geological cross-section across the Tansen–Pokhara area. Approximate positions of the samples are shown. Line of cross-section, patterns and abbreviations as in Fig. 2.

(Fig. 3). Two samples of the Ulleri augen gneiss were collected from the Modi Khola valley. The metapelitic samples from each tectonic unit have similar petrographic characteristics and mineral assemblages. Therefore, the petrography of same lithological group of samples is described together. 3.1. TS I 3.1.1. Phyllites (Nos. 10, 376, 372, 369, 366, 363, 358, 357 and 351) They are from the Chl zone. In thin-section, they are finegrained (b0.1 mm in size) and show well-developed beddingparallel foliation (S1) defined by alternating quartzo-feldspathic and micaceous domains (Fig. 5A). The S1 represents a preMCT (probably pre-Himalayan) deformation (Paudel and Arita, 2000). The S1 is weakly deformed in some samples to form kink folds. All the samples contain the mineral assemblage KMs + Chl + Ab + Qtz ± Gr ± Ilm along with detrital KMs, Tur, Zr and Qtz. Recrystallized Qtz grains occur as polygonal aggregate whereas detrital Qtz shows fluxion texture and strong pressure solution features. Samples close to the Phalebas Thrust (Nos. 351 and 357) are relatively more sheared and deformed compared to the other samples of TS I. 3.2. TS II 3.2.1. Phyllites (Nos. 3, 4 and 5) These samples also belong to the Chl zone but are more coarse-grained than the samples from the TS I. However, they contain two foliations; S1 and S2 defined by KMs and elongated Qtz grains (Fig. 5B). The samples contain the mineral assemblage KMs + Chl + Ab + Qtz ± Gr ± Ilm (KMs, Qtz, Tur and Zr as detrital). Quartz clasts show fluxion texture with relics surrounded by polygonal sub-grains. 3.2.2. Metasandstone (No. 109) The metasandstone contains bimodal grain-size with larger clasts of Qtz and KMs (0.5–1 mm) embedded in fine-grained (b0.1mm) matrix (Fig. 5C). The recrystallized matrix contains assemblage similar to that of phyllites (KMs + Chl + Ab + Qtz ± Gr ± Ilm). The Qtz clasts are ellipsoidal with long axes oriented parallel to the foliation. Detrital KMs grains are deformed and

sheared exhibiting wavy extinction. They are usually kinked and have corroded rims and, thus, can be easily distinguished from the recrystallized ones. The S2 is represented by S–C structure in metasandstones. 3.2.3. Biotite schists (Nos. 344, 341 and 331) The Bt zone samples are more sheared and recrystallized, and S2 makes the dominant foliation. They contain the mineral assemblage Bt + KMs + Chl + Ab + Qtz ± Gr ± Ilm (Qtz and Tur as detrital). Qtz occurs as granoblastic polygonal aggregate elongated along S2. Biotite is fine-grained (b0.1 mm) and shows pale yellow to brown pleochroism. Chlorite is green to pale-green. 3.3. MCT zone 3.3.1. Garnetiferous schists (Nos. 336, 137, 139, 140, 141, 145, 148, 154, 156, and 157) Samples from the MCT zone are strongly sheared and recrystallized. S2 represented by crenulation cleavage and S–C structure is predominant foliation in those samples. However, S1 is preserved as relic fabric in some samples (Fig. 5D). Main mineral assemblage is Grt + Bt + KMs + Chl + Ab + Qtz ± Gr ± Ilm. Bt, KMs and Chl define foliation. Biotite shows pale yellow to brown pleochroism. Qtz is strongly elongated parallel to the foliation showing ribbon texture. Syn-tectonic poikiloblastic Grt in the schists occurs in different shapes (skeletal, elongated, s-shaped and eqidimensional) and sizes (0.1–5mm). Spiral Grt shows up to 360° rotations of the inclusions (Fig. 5E). The rotated inclusion pattern in the Grt porphyroblasts shows a topto-the-south sense of shearing in the MCT zone. 3.3.2. Augen gneiss (Nos. 127 and 624) The augen gneisses are mylonitic to protomylonitic with Kfs (Or and Mc) and Ab augens up to 1 cm in diameter. The matrix contains the assemblage Or + Ab + KMs + Bt + Qtz ± Grt with Tur, Zr, Ap and opaques as accessories. The feldspar augens have fractured and granulated boundaries. Myrmekitic intergrowths of Qtz and Or are also observed along the rims of the augens. KMs occur both as coarse-grained porphyroclasts and fine-grained matrix (Fig. 5F). The KMs porphyroclasts are sheared parallel to the foliation and make mica fishes. They are kinked and bent. Medium-grained (0.1–0.5 mm) Bt is an

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Fig. 5. Photomicrographs of representative samples showing two major foliations and three textural types of KMs (K-white micas). (A) Phyllite from TS I (No. 376) showing only one bedding parallel foliation (S1 = S0) defined by recrystallized KMs. (B) Phyllite from TS II (No. 3) showing two foliations (S1 and S2) defined by recrystallized KMs. (C) Detrital KMs and S–C fabric in metasandstone from the TS II (No. 109). (D) Crenulation cleavages in biotite schist from TS II (No. 139). (E) Well-developed S2 foliation and snowball garnet in the garnet schist from the MCT zone (No. 154). (F) KMs fishes showing right lateral slip in granitic augen gneiss from the MCT zone (No. 127). Scale: one smallest division = 0.01mm.

abundant recrystallized phase defining the foliation. It shows pale yellow to dark-brown pleochroism. 4. Analytical techniques Whole rock compositional analyses were made on five representative samples with Philips PW1404 X-ray Fluorescence Spectrometer (XRF). Samples for XRF analyses were prepared following the procedures described in Murata et al. (1994), and the data were calibrated with GSJ reference samples (Imai et al., 1995). Compositional point analyses of KMs were performed in about 60 porphyroclasts and more than 400 recrystallized KMs. Recrystallized KMs along both S1 and S2 were analyzed from one sample of the TS II (No. 3) and two samples of the MCT zone (Nos. 336 and 139). Analyses in other samples from the TS I and II are on the recrystallized KMs along S1 while those in

remaining samples from the MCT zone are along S2. Analyses were carried out using JEOL Superprobe 733 Electron Probe Microanalyzer with 0.02 mA specimen current, 15kV accelerating voltage. Natural and synthetic silicates and oxides were used as standards. All the raw data were corrected using conventional ZAF correction procedures. X-ray element mapping was performed with JEOL superprobe JXA-8900M. Almost all the analyses were performed at the Hokkaido University. X-ray mapping of one sample was carried out at the Okayama University of Science. 5. Host rock chemical compositions Eleven metasediment samples with diverse lithology collected from all the tectonic units were analysed by XRF to check the bulk compositional variation in the studied samples (Table 1).

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Table 1 Whole rock chemical analyses of representative samples Unit

TS I

Sp. no.

372

366

350

5

TS II 3

336

MCT zone 141

145

154

156

157

Rock

Phyl

Phyl

Phyl

Phyl

Phyl

Schist

Schist

Schist

Schist

Schist

Schist

SiO2 TiO2 Al2O3 Fe2O3* MnO MgO CaO Na2O K2O P2O5 LOI Total

57.19 1.14 18.54 11.83 0.03 2.03 0.15 0.07 6.46 0.11 2.24 99.78

66.79 0.77 13.85 9.65 0.02 1.85 0.08 0.04 5.33 0.03 1.11 99.51

53.91 0.55 28.41 7.90 0.13 2.40 0.18 0.35 6.49 0.12 0.12 100.54

54.30 0.58 27.42 7.26 0.03 2.47 0.20 0.20 7.32 0.15 0.81 100.75

59.01 0.50 24.82 6.19 0.13 2.04 0.18 0.21 6.46 0.12 1.12 100.78

60.53 0.64 22.33 6.75 0.06 2.71 0.24 1.14 4.85 0.11 0.71 100.06

67.37 0.60 19.37 3.38 0.01 2.09 0.72 0.78 4.76 0.12 0.56 99.75

69.23 0.56 17.97 2.63 0.00 2.94 0.46 0.17 4.55 0.09 0.81 99.40

72.48 0.69 12.19 6.12 0.06 1.61 0.35 1.10 3.96 0.09 0.67 99.31

70.95 0.71 14.53 5.27 0.08 1.17 0.72 1.22 4.01 0.08 0.57 99.31

68.00 0.77 13.91 7.31 0.12 2.83 0.96 2.52 2.61 0.14 0.68 99.84

Fe2O3* means total Fe as Fe2O3.

Major oxides contents vary slightly among samples although there is not any systematic trend. SiO2 varies from 53.91% to 72.48%, Al2O3 varies from 12.19% to 28.41%, Fe2O3 varies from 2.63% to 11.83%, MgO varies from 1.17% to 2.85% and K2O varies from 2.61% to 7.32%. Loss on ignition (LOI) is very low (between 0.12% and 2.24%) for all the samples. Comparison of major oxide contents of KMs with host rock shows no systematic relationship between them. 6. Results of KMs analyses The variation in KMs compositions can be demonstrated in a simple plot of Al2O3 against FeO* (total Fe measured as FeO) and the variation can be related with the metamorphic grade (Miyashiro, 1973). Generally, Al2O3 increases and FeO* decreases with increasing metamorphic grade and KMs compositions approach to ideal muscovite (Al2O3 = 38.4%, FeO* = 0%) indicating high-grade metamorphic rocks and granites have KMs compositions close to ideal muscovite. Therefore, KMs compositions are discussed mainly using Al2O3–FeO* diagrams in the present work. However, cation-based histograms and variation diagrams are also used to access the spatial variation in KMs compositions. Representative KMs analyses are given in Table 2. 6.1. Compositional variation and zoning in porphyroclastic and detrital KMs Compositions of porphyroclastic KMs in augen gneisses (Nos. 127 and 624) and detrital KMs in metasandstone (No. 109) are shown in the Al2O3–FeO* diagram along with the compositions of recrystallized KMs in the same samples for comparison (Fig. 6). Porphyroclastic KMs in No. 127 plot near the ideal muscovite and have narrow range of compositions (Al2O3: 34.9–36.5%, FeO*: 1.0–1.6%) whereas those in the Sp. No. 624 are relatively poor in Al2O3 (33.5–34.3%) and rich in FeO* (1.8–2.2%) contents. The recrystallized KMs in Sp. No. 624 are comparably celadonitc (Al2O3: 29.8–31.8%, FeO*: 2.9–4.1%). The difference of compositions among porphyroclastic and re-

crystallized KMs is clearly displayed also in the X-ray element maps (Fig. 7). Further, some of the porphyroclasts show heterogeneous compositional zoning (Figs. 8 and 9). The detrital KMs grains in metasandstone (No. 109) have variable compositions and plot in two different fields on the Miyashiro diagram. Some are relatively celadonite-rich (Al2O3: 34.3–35.0%, FeO*: 1.8–2.3%). Some are celadonite-poor (Al2O3: 35.1–36.1%, FeO*: 0.8–1.0%) and plot near the porphyroclasts from augen gneiss. The compositional variation among the detrital grains suggests their different provenance. The compositions of recrystallized KMs in the same sample are more celadonitic (Al2O3: 32.2–33.7%, FeO*: 2.0–4.0%) compared to those of the detrital grains. 6.2. Temporal variation in composition of recrystallized KMs Compositions of recrystallized KMs along S1 (older KMs) and S2 (younger KMs) in three samples (one from the TS II and two from the MCT zone) are plotted on Al2O3–FeO* diagram (Fig. 10) and histograms of Al- and (Fe + Mg)contents (Fig. 11A–C). For all the samples, KMs defining S1 and S2 plot in separate fields with some overlapping. KMs defining S1 are relatively celadonite-rich and those defining S2 are relatively celadonite-poor. For all the samples, Al-contents of KMs defining S1 lie in the lower and those defining S2 lie in the higher value sides in the histogram plots (Fig. 11). Similarly, (Fe + Mg)-contents of KMs defining S1 lie in the higher and those defining S2 lie in the lower value sides. Such compositional differences between KMs defining S1 and S2 are also observed in the compositional X-ray element mapping of sample No. 139 (the results are discussed in Paudel and Arita, in press). In addition, the compositional X-ray element mapping shows that the KMs porphyroblasts defining S1 are compositionally zoned; the cores are celadonite-rich and the rims are celadonite-poor. Despite of difference in average compositional variation between KMs defining S1 and S2, the chemical compositions of KMs from within a single foliation are remarkably dispersed. The dispersion is wider in KMs defining S1 compared to those

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Table 2 Representative K-white micas analyses from the Lesser Himalaya Unit

TS I

Sp. no.

10

376

372

369

366

363

358

351

109

TS II 109

5

4

3

3

1

344

341

331

Rock

Phyl.

Phyl.

Phyl.

Phyl.

Phyl.

Phyl.

Phyl.

Phyl.

Sst.

Sst.

Phyl.

Phyl.

Phyl.

Phyllite

Phyl.

Phyl.

Phyl.

Phyl.

Mica type

S1

S1

S1

S1

S1

S1

S1

S1

Detrital

S1

S1

S1

S1

S2

S1

S1

S1

S1

SiO2 TiO2 Al2O3 FeO⁎ MnO MgO CaO Na2O K2O Total

49.50 0.59 25.92 6.86 0.01 2.80 0.02 0.08 9.96 95.74

51.41 1.55 24.34 6.78 0.01 3.30 0.02 0.10 7.54 95.04

50.79 0.44 29.50 4.70 0.02 2.13 0.05 0.10 8.53 96.27

50.48 0.46 26.60 3.98 0.01 3.27 0.13 0.14 9.78 94.84

51.32 0.40 27.14 5.19 0.01 2.45 0.00 0.12 7.97 94.60

51.23 0.34 28.32 2.37 0.03 3.07 0.00 0.12 10.12 95.61

49.75 2.79 25.97 6.84 0.03 2.34 0.03 0.05 7.81 95.59

47.59 4.86 26.92 5.03 1.96 0.00 0.08 8.93 0.00 95.37

47.32 0.57 35.85 0.89 0.00 0.40 0.01 0.56 10.08 95.67

50.54 0.31 31.58 2.62 0.03 1.90 0.08 0.14 8.20 95.37

49.30 0.36 30.88 2.37 0.01 1.68 0.03 0.31 9.53 94.47

48.76 0.33 32.88 1.93 0.00 1.31 0.02 0.21 9.76 95.18

48.34 0.36 31.87 2.54 0.03 1.41 0.04 0.27 9.32 94.18

49.87 0.40 32.32 2.60 0.03 1.44 0.02 0.28 8.58 95.53

49.40 0.30 31.57 3.03 0.00 2.15 0.03 0.18 8.63 95.28

49.91 0.37 32.42 2.64 0.04 1.40 0.06 0.31 8.06 95.20

50.05 0.48 31.64 3.91 0.00 2.13 0.06 0.29 7.63 96.19

49.48 0.41 31.93 3.33 0.00 1.28 0.00 0.11 8.28 94.82

Cation for 22 oxygens Si 6.72 6.91 6.69 6.80 6.87 6.78 6.68 6.36 6.23 6.62 6.59 6.46 6.48 6.55 6.53 6.55 6.54 Ti 0.06 0.16 0.04 0.05 0.04 0.03 0.28 0.49 0.06 0.03 0.04 0.03 0.04 0.04 0.03 0.04 0.05 Al 4.15 3.86 4.58 4.22 4.28 4.42 4.11 4.24 5.56 4.88 4.86 5.13 5.04 5.00 4.92 5.02 4.87 Fe⁎ 0.78 0.76 0.52 0.45 0.58 0.26 0.77 0.56 0.10 0.29 0.26 0.21 0.29 0.29 0.33 0.29 0.43 Mn 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.22 0.00 0 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Mg 0.57 0.66 0.42 0.66 0.49 0.61 0.47 0.00 0.08 0.37 0.33 0.26 0.28 0.28 0.42 0.27 0.41 Ca 0.00 0.00 0.01 0.02 0.00 0.00 0.00 0.01 0.00 0.01 0.00 0.00 0.01 0.00 0.00 0.01 0.01 Na 0.02 0.02 0.02 0.04 0.03 0.03 0.01 2.31 0.14 0.03 0.08 0.05 0.07 0.07 0.05 0.08 0.07 K 1.73 1.29 1.43 1.68 1.36 1.71 1.34 0.00 1.69 1.37 1.62 1.65 1.59 1.44 1.45 1.35 1.27 Total 14.02 13.67 13.71 13.91 13.65 13.84 13.66 14.19 13.85 13.61 13.80 13.80 13.80 13.67 13.74 13.61 13.65 FeO means total Fe as FeO. Phyl.: Phyllite; Ppclast: porphyroclast; S1: K-white micas defining first foliation; S2: K-white micas defining second foliation. MCT⁎ zone

defining S2 (Fig. 11). The variation among KMs defining S1 is higher in the samples from the TS I (Fig. 11E–G) and southern part of the TS II (Fig. 11D) than in the northern part of the TS II (Fig. 11B–C) and the MCT zone (Fig. 11A). Implications of these compositional variations are discussed in Section 7.1. 6.3. Spatial variation in composition of recrystallized KMs The compositional histograms in Fig. 11 are arranged from south to north according to their positions in the cross-section in Fig. 4. The diagrams show that Al-content of KMs defining both S1 and S2 gradually increases and (Fe + Mg)-content decreases from the bottom (south) to the top (north). It shows that the KMs become gradually celadonite-poor from south to north. The compositions of recrystallized KMs in all the samples from the present area are plotted in Al2O3–FeO* diagrams in Fig. 12A and B. The diagrams indicate that KMs defining both S1 and S2 become celadonite-poor when going north from the TS I to TS II and to the MCT zone. The KMs defining S1 in the TS I are relatively celadonite-rich and show very wide range of composition (Al2O3: 22.0% to 30.0%, FeO*: 2.5–8.0%). Those from the TS II are relatively celadonite-poor and plot in relatively narrow field (Al2O3: 30.0% to 33.0%, FeO*: 1.5– 4.0%). The KMs defining S1 from the MCT zone also show high variation in compositions (Al2O3: 30.0% to 37.0%, FeO*: 1.0– 4.0%), but most of the grains are celadonite-poor and plot close to the ideal muscovite.

6.55 0.04 4.98 0.37 0.00 0.25 0.00 0.03 1.40 13.63

The KMs defining S2 show less compositional variations compared to those of S1 (Fig. 12B). The Al2O3 content of KMs defining S2 in the TS II varies from 31.5% to 33.5% and FeO*content from 2.0% to 4.5%. In the case of MCT zone, Al2O3content varies from 32.5–38.4% and FeO*-content from 0.05– 2.5%. Miyashiro (1973) pointed out that KMs from the St and Sil zones and those from Chl, Bt and Grt zones plot in distinctly different fields on the Al2O3–FeO* diagram. In the present case, KMs along S1 from the TS I plot within and outside the fields of Chl, Bt and Grt zones (Fig. 12A). In the TS II (Chl and Bt zones), KMs along both S1 and S2 plot almost entirely in the fields of Chl, Bt and Grt zones. KMs along S2 from the MCT zone (Grt zone), partly plot in the fields of Chl, Bt and Grt zones, but part of them plot in the fields of St and Sil zones suggesting higher metamorphic temperature conditions in the MCT zone. 7. Discussions The importance of KMs compositions for metamorphic studies has been highlighted by many researchers (e.g., Ernst, 1963; Velde, 1965, Butler, 1967; Miyashiro, 1973; Sassi and Scolari, 1974; Powell and Evans, 1983; Guidotti, 1984; Guidotti and Sassi, 1998). Guidotti and Sassi (2002) have suggested that, with careful selection of samples, the effects of the petrological parameters such as host rock chemistry, associated mineral phases, and oxygen fugacity (fO2) can be

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417

MCT zone 336

336

137

127

127

624

624

139

139

140

141

145

148

154

156

156

157

Schist

Schist

Schist

Gneiss

Gneiss

Gneiss

Gneiss

Schist

Schist

Schist

Schist

Schist

Schist

Schist

Schist

Schist

Schist

S1

S2

S2

Ppclast

S2

Ppclast

S2

S1

S2

S2

S2

S2

S2

S2

S2

S2

S2

50.30 0.32 32.67 2.12 0.01 1.41 0.00 0.50 9.05 96.39

49.72 0.38 33.28 1.89 0.03 1.23 0.03 0.55 9.20 96.29

49.80 0.29 34.32 1.53 0.01 1.08 0.08 0.47 8.36 95.93

47.78 0.07 36.10 1.05 0.00 0.15 0.00 0.60 9.98 95.74

48.27 0.06 35.21 1.51 0.00 0.28 0.01 0.83 9.48 95.63

47.44 0.15 34.34 1.83 0.00 0.85 0.00 0.41 8.52 2

49.01 0.41 31.12 3.85 0.00 1.46 0.00 0.10 7.94 2

49.79 0.23 32.71 2.43 0.00 0.75 0.01 0.71 9.20 95.84

48.48 0.28 34.07 2.47 0.00 0.68 0.07 0.64 8.62 95.30

48.39 0.35 33.46 1.66 0.02 1.06 0.03 1.08 9.34 95.38

49.02 0.35 34.45 1.54 0.01 0.81 0.06 0.38 8.60 95.23

48.40 0.43 33.66 0.68 0.01 1.39 0.02 0.67 10.18 95.45

48.88 0.76 32.39 1.13 0.00 1.65 0.02 0.40 9.85 95.06

48.41 0.57 33.34 1.87 0.00 1.11 0.01 0.56 10.06 95.94

48.22 0.21 37.58 1.13 0.00 0.21 0.20 5.04 2.05 94.65

47.77 0.07 38.24 1.01 0.01 0.11 0.41 5.98 0.98 94.59

47.56 0.06 38.42 0.83 0.01 0.07 0.42 6.52 1.27 95.16

Cation for 22 oxygens 6.55 6.48 6.46 0.03 0.04 0.03 5.01 5.12 5.25 0.23 0.21 0.17 0.00 0.00 0.00 0.27 0.24 0.21 0.00 0.00 0.01 0.13 0.14 0.12 1.50 1.53 1.38 13.73 13.76 13.63

6.27 0.01 5.59 0.12 0.00 0.03 0.00 0.15 1.67 13.84

6.34 0.01 5.45 0.17 0.00 0.05 0.00 0.21 1.59 13.82

6.35 0.02 5.41 0.20 0.00 0.17 0.00 0.11 1.45 13.71

6.57 0.04 4.91 0.43 0.00 0.29 0.00 0.03 1.36 13.63

6.54 0.02 5.06 0.27 0.00 0.15 0.00 0.18 1.54 13.77

6.39 0.03 5.29 0.27 0.00 0.13 0.01 0.16 1.45 13.74

6.40 0.03 5.21 0.18 0.00 0.21 0.00 0.28 1.57 13.89

6.42 0.03 5.32 0.17 0.00 0.16 0.01 0.10 1.44 13.65

6.38 0.04 5.23 0.08 0.00 0.27 0.00 0.17 1.71 13.90

6.47 0.08 5.05 0.12 0.00 0.32 0.00 0.10 1.66 13.81

6.38 0.06 5.18 0.21 0.00 0.22 0.00 0.14 1.69 13.89

6.21 0.02 5.70 0.12 0.00 0.04 0.03 1.26 0.34 13.72

6.14 0.01 5.79 0.11 0.00 0.02 0.06 1.49 0.16 13.78

6.10 0.01 5.81 0.09 0.00 0.01 0.06 1.62 0.21 13.91

neglected and KMs compositional variations can be related with metamorphic grade. In the present study, samples with limited mineral assemblages were selected so as to minimize the effects of petrological factors other than P and T. Absence of Hm and presence of Gr, Ilm and Mag in the studied samples indicate reducing conditions for the terrain. Plot of whole rock chemistry against the KMs

composition in the sample shows very poor relation between the two. On the other hand, all the samples contain KMs with variable compositions within a single thin-section. It indicates that host rock composition has a little control on the KMs compositional variation. These backgrounds allow us to argue that the compositional variations in the KMs are mainly due to the difference in P–T conditions, and, therefore, can be used as a metamorphic indicator in the study area. The implications of the KMs compositions on metamorphism of the LH are discussed below. 7.1. Two thermal events in the LH

Fig. 6. Al2O3–FeO* plot of porphyroclasts and recrystallized (S2) K-white micas in augen gneisses (No. 127 and 624) from the MCT zone, detrital and recrystallized (S1) K-white micas in metasandstone (No. 109) from the TS II (109). FeO* means total Fe as FeO.

In contrast to the well-established and less-debated twostage prograde metamorphic evolution of the Higher Himalayan crystallines (Le Fort, 1975; Arita, 1983; Pêcher, 1989; Inger and Harris, 1992; Lombardo et al., 1993; Hodges et al., 1996), the thermal evolution of the LH is often not clear. Some researchers think that the inverted metamorphism of the LH is a primary pattern acquired during one-stage prograde metamorphism due to the MCT activity in the Miocene (Le Fort, 1975; Pêcher, 1989; Guillot, 1999). Others argue that it is due to the transposition of right-way-up metamorphic sequences by postmetamorphic kinematics (Catlos et al., 2001; Bollinger et al., 2004). Metamorphism in the LH is supposed to occur during and after the Miocene (∼20Ma) after the initiation of the MCT. On the other hand, Johnson and Oliver (1990), Oliver et al. (1995) and Paudel and Arita (2000) suggested that the LH experienced at least one pre-MCT (probably pre-Himalayan)

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Fig. 7. Photomicrograph (A) and compositional X-ray maps (B, C and D) of porphyroclastic and matrix K-white micas in the Ulleri-type augen gneiss (No. 624). The maps clearly show that the porphyroclast is celadonite-poor and the matrix is celadonite-rich. Matrix grains showing extremely high Fe-contents in (D) are biotites. KMs: K-white micas; Kfs: K-feldspar; Qtz: Quartz.

thermal event based on crystallinity, b-spacings and K–Ar ages. The present results of KMs compositions are in line with the later view.

The S1, which pre-dates the movement along the MCT, contains relatively celadonite-rich KMs. In contrast, the S2, which was formed during the MCT movement, contains

Fig. 8. Photomicrograph (A) and compositional X-ray maps (B, C and D) of a porphyroclastic K-white mica in Ulleri augen gneiss (No. 127). The maps show

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Fig. 9. Compositional profile of K-white mica porphyroclast along X–Y in Fig. 7(A).

relatively celadonite-poor KMs. Provided that P and T are the only controlling factors during the recrystallization of KMs defining S1 and S2 for the observed samples as discussed above, this is a clear evidence for distinct P–T conditions between S1 and S2 events. Generally, celadonite-rich KMs recrystallize at higher P at constant T (Ernst, 1963; Velde, 1965, 1967, 1980). With the increase of T, (Fe + Mg)- and Sicontents decreases and Al-content increases as fine-grained illite changes to coarse-grained phengite and muscovite (McDowell and Elders, 1980; Padan et al., 1982). At higher T conditions with the appearance of Bt and Grt, partitioning of (Fe + Mg) into the ferromagnesian phases results in the recrystallization of celadonite-poor KMs. In the present case, most probably, the celadonite-rich KMs defining S1 were recrystallized under lower T conditions and the celadonitepoor KMs defining S2 were recrystallized under higher T conditions. Higher T condition for S2 is also supported by the presence of Bt and Grt along S2 in the northern part of the TS II and the MCT zone.

419

Despite of difference in average compositions of KMs defining S1 and S2, there are wide compositional variations among grains defining same foliation (Figs. 9 and 10). The compositional variations and zoning in KMs defining S1 may be due to partial re-equilibration of celadonite-rich KMs during the second higher T thermal event (S2 event). The inverted metamorphic zonation from the TS I to the MCT zone shown by KMs defining S1 is likely due to the thermal effect of the S2 event. Parallelism of S2 with shear planes of the MCT suggests that the second (higher T) event is the Neohimalayan (Miocene) event as described by Le Fort (1975), Arita (1983), Pêcher (1989) and Guillot (1999). As S1 is oblique to the MCT shear planes, it is clear that the first event (lower T) is pre-MCT. However, whether it is an Eohimalayan (Eocene) or pre-Himalayan (early Paleozoic) events is debatable. Some lines of evidence such as (i) an abrupt break in illite crystallinity values between Gondwana and pre-Gondwana sediments (Johnson and Oliver, 1990; Paudel and Arita, 2000), (ii) pre-Cenozoic mineral and whole rock ages from the LH (Khan and Tater, 1970; Ashgirei et al., 1977; Oliver et al., 1995; Paudel, 2000) and (iii) finding of slate and phyllite clasts from the Upper Carboniferous sequence of LH in India (Thakur and Pande, 1973) suggest for a pre-Himalayan thermal event in the LH. Bhargava (1980) and Williams et al. (1988) have presented evidence showing pre-Tertiary orogenies in the Himalaya. Presence of early Paleozoic granites (Le Fort et al., 1986b), Carboniferous to Cretaceous monazites in metapelites and ∼270 Ma allanite inclusion in garnet cores (Catlos et al., 2001) also indicate for a preHimalayan thermal event in the LH. 7.2. Inverted thermal structure in the inner LH In the studied area, inverted metamorphic zonation is obviously observed in relatively higher grade rocks of the MCT zone and the Higher Himalaya (Le Fort, 1975; Pêcher, 1975; Arita, 1983, Kaneko, 1995; Hodges et al., 1996). However, its continuity to the low-grade rocks to the south of the MCT zone has been a subject of discussion. Morrison and Oliver (1993) suggest that the inverted metamorphism does not extend to the

Fig. 10. Compositions of recrystallized K-white micas defining S1 and S2 in sample Nos. 3 (A), 336 (B) and 139 (C) plotted on Al2O3–FeO* diagram. FeO* means total Fe as FeO.

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Fig. 11. Compositional histograms of K-white micas defining S1 and S2. The histograms are arranged from north (top) to south (bottom) according to their positions in cross-section (Fig. 3). (A) and (B): MCT zone; (C), (D) and (E): TS II; (F) and (G): TS I.

low-grade metamorphic rocks. On the other hand, Sakai et al. (1999) and Paudel and Arita (2000) revealed inverted thermal structure to continue throughout the Lesser Himalaya including the Gondwana and post-Gondwana sediments. Recently,

Beyssac et al. (2004) and Bollinger et al. (2004) showed that the thermal structure is inverted up to 5-8 km structurally downward from the Upper MCT and that the gradient is very steep (20–50 °C/km).

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421

Fig. 12. Al2O3–FeO* diagram showing K-white micas compositions defining S1 (A) and S2 (B) compositions for each tectonic unit. The fields for chlorite, biotite and garnet and for staurolite and sillimanite zones are after Miyashiro (1973). FeO* means total Fe as FeO.

The KMs compositional data in the present study also show an inverted thermal structure in the inner LH. The average contents of Ti and Al increase and those of (Fe + Mg) and Si in KMs defining both S1 and S2 decrease from south to north (Fig. 11) indicating northward increase in temperature. Fig. 12 also shows that the temperature condition increases from TS I to TS II and to the MCT zone. To evaluate the qualitative thermal profile across the study area, the average contents of Si, Ti, Al and (Fe + Mg) in recrystallized KMs in all the metapelitic samples are plot against their spatial positions in the N–S geological cross-section (Fig. 13). The diagram shows that, despite of some intrasample variation (2σ variance), the average Si- and (Fe + Mg)-contents decrease and Al- and Ticontents increase from south to north towards the upper MCT. Gradual decrease in intrasample compositional variation of KMs from south to north indicates from poor to more complete re-equilibration of celadonite-rich KMs defining S1 and neocrystallization of celadonite-poor KMs defining S2 as a result of northward increasing T. Although the higher metamorphic condition shown by KMs compositions in the TS II may be partly due to the presence of stratigraphically deeper level rocks (Kunchha Formation), gradually northwards increasing trend of metamorphic grade from the TS II to the level of the Upper MCT where rocks become stratigraphically younger (Benighath Slate) should be due to the primary inversion of thermal gradient. Local irregularity in the compositional trend may be due to local-scale shearing or folding. There are some disagreements between the thermal structures revealed from different methodologies in the MCT zone of the present study area. Kaneko (1995) showed a steep inverted thermal gradient at the MCT zone of the present section with peak metamorphic temperature increasing from 400– 450 °C at the Lower MCT level to 600–650°C near the Upper

MCT using Grt–Bt thermometry. Hbl–Pl thermometry also shows peak-temperatures increasing from 540 °C at the base to 610–640 °C at the top of the MCT zone (Paudel et al., 2005). A decrease of celadonite content in KMs from the base to the top of the MCT zone is in agreement with the northward increase in temperature shown by the above thermometric calculations. In contrast, Beyssac et al. (2004) and Bollinger et al. (2004) found relatively lower and fairly uniform temperatures (500–570 °C, mean: 541 °C) with no systematic trend within the MCT zone. Given the higher error of calibration (± 50 °C) of RSCM method (Beyssac et al., 2004) and possible intrasample variation in degree of graphitization of carbonaceous materials due to polyphase metamorphism and shearing, the measured temperature may not reflect the actual thermal profile in the MCT zone. Further, systematically lower peak temperatures shown by this method compared to the temperatures determined by other methods in all the sections of the LH indicates that the RSCM method may not be sensitive at high temperature conditions due to saturation. In the low-grade terrain (especially in the Chl zone) of the LH, the Neohimalayan metamorphism was relatively weaker than the pre-MCT metamorphism (Paudel and Arita, 2000). Therefore, it is uncertain whether the temperatures and geothermal gradient measured with RSCM method actually represent only the Neohimalayan metamorphism. 7.3. Implications for the KMs geochronology in the LH A number of researchers including the present authors have measured K–Ar and 40Ar / 39Ar ages to constrain the metamorphic and cooling history of the central Nepal Lesser Himalaya. Paudel (2000) carried out an reconnaissance K–Ar dating on b2 μm powder fractions of seven phyllite samples in the Nawakot Complex (Chl zone only) in the present section and

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Fig. 13. Average compositional variation of recrystallized K-white micas (KMs) along the N–S section in the study area. KMs compositions in the Thrust Sheet I and II belong to the grains along S1 and those from the MCT zone belong to S2. Note the northward increasing trends of Al- and Ti-contents and decreasing trends of average (Fe + Mg)- and Si-contents showing inverted thermal profile throughout the inner Lesser Himalaya. PT: Phalebas Thrust, LMCT: Lower Main Central Thrust, UMCT: Upper Main Central Thrust.

found that the ages are highly dispersed (Precambrian to Paleocene) but become successively younger from south to north. The ages range from 1119 ± 3 to 315 ± 7 Ma in the Parautochthon, 356 ± 7 to 91 ± 2 Ma in the TS I and 76 ± 2 to 68 ± 2 Ma in the TS II. Radiometric ages presented by other researchers also range from Precambrian to Pliocene. The detrital KMs in the Kunchha Formation and the Tansen Group yield K–Ar ages of 872 and 1280 Ma, respectively (Krummenacher, 1966). K–Ar whole rock ages on phyllite and slates presented by Khan and Tater (1970) and Oliver et al. (1995) range from 560 to 8 Ma. 40 Ar / 39Ar muscovite ages from the MCT zone range from 33 to 2.3 Ma (Copeland et al., 1991; Macfarlane, 1993). Catlos et al. (2001) have presented 40Ar / 39Ar muscovite ages of 116.8 ± 0.2 and 257 ± 1Ma in the Darondi Khola section of central Nepal. 40Ar / 39Ar muscovite ages measured by Bollinger et al. (2004) for the Damauli and Marsyangdi sections range from 31.3 to 15.6 and 29.1 to 2.4 Ma, respectively. If all the available 40 Ar / 39Ar muscovite ages from the LH in central Nepal are plot in a north–south section, they show approximately a linear

trend; ages getting younger northwards (Bollinger et al., 2004). The older muscovite ages (N 20 Ma) from the LH have been interpreted either to represent pre-Himalayan metamorphic event (Krummenacher, 1966; Oliver et al., 1995) or due to extraneous argon (Vanny and Hodges, 1996; Coleman and Hodges, 1998; Catlos et al., 2001). On the other hand, the younger ages (b 20 Ma) are thought to represent the Neohimalayan metamorphism caused by the MCT activity. The youngest ages (about 2–8 Ma) have been interpreted as either due to the reactivation of the MCT (Harrison et al., 1997) or late-stage brittle thrusting (Macfarlane, 1993) or thermal resetting from retrograde hot fluid (Copeland et al., 1991). However, it is to be noted that intrasample and intra-grain compositional variability in KMs that has significant effect on the calculated ages has been neglected in the age interpretations. Several studies have shown that deficiency of total recrystallization by major element diffusion in KMs also entails a lack of total isotopic re-equilibration (Hammerschmidt and Frank, 1991; Villa et al., 1997). It has also been proved by the work of Mulch et al. (2005) who carried out UV-laser 40Ar / 39Ar dating

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of a chemically zoned mica fish and obtained different ages for the cores and rims. Therefore, the presence of chemically zoned KMs and inter-fabric compositional variation in the LH also points to the inequilibration of K–Ar isotopic system. As the MCT zone has experienced temperature of more than 500°C during the Neohimalayan thermal event, the chemical inequilibration at this temperature suggest the closure temperature for the K–Ar system is higher (∼500 °C; De Sigoyer et al., 2000; Villa, 1998; Gouzu, 2004) than previously thought (∼350 °C; Jäger, 1979). It is also consistent with the work of Dempster (1992) who suggested that phengites remain effectively closed system under metamorphic conditions of greenschist and greenschist–amphibolite transitional facies. Above facts allows us to argue that the older K–Ar and 40 Ar / 39Ar ages from the LH are not only due to extrageneous argon but also due to the presence of older KMs (detrital and S1). Northward younging of the cooling ages may be due to the gradual resetting of the older aged KMs during the latest heating event (Neohimalayan metamorphism) similar to that observed in the Sanbagawa metamorphic belt (Itaya and Fukui, 1994). Therefore, we suggest that the tectonometamorphic significance of the published K–Ar and 40Ar / 39 A in the LH should be re-evaluated on the ground of observed compositional variability and possible higher closure temperature for K–Ar system in KMs. 8. Conclusions The compositions of K-white micas (KMs) in the Lesser Himalaya (LH) vary both in time and space. The porphyroclasts and detrital grains are generally celadonite-poor and show heterogeneous zoning with celadonite-rich rims. The cores are the relics of igneous or higher grade metamorphic muscovites, and the rims were re-equilibrated or overgrown at lower T metamorphic conditions. The recrystallized KMs along S1 (preMCT foliation) are generally celadonite-rich and show compositional zoning with celadonite-rich cores and celadonite-poor rims. KMs along S2 (syn-MCT foliation) are generally celadonite-poor. The KMs compositions show at least two prograde thermal events in the LH. The first is a lower T event prior to the MCT activity that caused recrystallization of the celadonite-rich KMs defining S1. The second is a relatively higher T event during the MCT activity that cause recrystallization of celadonite-poor KMs defining S2 and partial re-equilibration of older KMs (porphyroclastic, detrital and S1). Average compositions of recrystallized KMs defining and S2 and partially re-equilibrated KMs defining S1 become gradually poor in (Fe + Mg)- and Si-contents and rich in Al- and Ticontents from south to north (structurally lower to higher levels) in the LH showing an inverted thermal structure throughout the LH. Intrasample compositional variation and chemical zoning, and possible higher closure temperature (∼500 °C) for the K–Ar system in KMs warrants for re-evaluation of the tectonometamorphic significance of the published K–Ar and 40Ar / 39A ages in the LH.

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