Petrology of a non-classical Barrovian inverted metamorphic sequence from the western Arunachal...

Journal of Asian Earth Sciences 36 (2009) 390–406

Contents lists available at ScienceDirect

Journal of Asian Earth Sciences

journal homepage: www.elsevier .com/locate / jseaes

Petrology of a non-classical Barrovian inverted metamorphic sequencefrom the western Arunachal Himalaya, India

S. Goswami a, S.K. Bhowmik a,*, S. Dasgupta b

a Department of Geology and Geophysics, Indian Institute of Technology, Kharagpur 721 302, Indiab Indian Institute of Science Education and Research, HC Block, Sector III, Salt Lake, Kolkata 700 106, India

a r t i c l e i n f o

Article history:Received 3 October 2008Received in revised form 2 July 2009Accepted 2 July 2009

Keywords:BarrovianInverted metamorphismArunachal Himalaya

1367-9120/$ - see front matter � 2009 Elsevier Ltd. Adoi:10.1016/j.jseaes.2009.07.006

* Corresponding author.E-mail address: [email protected] (S.K. Bh

a b s t r a c t

In this study, we reconstruct the inverted metamorphic sequence in the western Arunachal Himalayausing combined structural and metamorphic analyses of rocks of the Lesser and Greater HimalayanSequences. Four thrust-bounded stratigraphic units, which from the lower to higher structural heightsare (a) the Gondwana rocks and relatively weakly deformed metasediments of the Bomdila Group, (b)the tectonically interleaved sequence of Bomdila gneiss and Bomdila Group, (c) the Dirang Formationand (d) the Se La Group are exposed along the transect, Jira–Rupa–Bomdila–Dirang–Se La Pass. The MainCentral thrust, which coincides with intense strain localization and the first appearance of kyanite-gradepartial melt is placed at the base of the Se La Group.

Five metamorphic zones from garnet through kyanite, kyanite migmatite, kyanite-sillimanite migma-tite to K-feldspar-kyanite-sillimanite migmatites are sequentially developed in the metamorphosedlow-alumina pelites of Dirang and Se La Group, with increasing structural heights. Three phases of defor-mation, D1–D2–D3 and two groups of planar structures, S1 and S2 are recognized, and S2 is the mostpervasive one. Mineral growths in all these zones are dominantly late-to post-D2, excepting in some gar-net-zone rocks, where syn-D1 garnet growths are documented. Metamorphic isograds, which are alignedparallel to S2 were subsequently folded during D3. The deformation produced plane-non-cylindrical foldalong NW–SE axis.

In the garnet-zone, peak metamorphism is marked by garnet growth through the reaction biotite + pla-gioclase ? garnet + muscovite. An even earlier phase of syn-D1 garnet growth occurred in the chloritestability field with or without epidote. In the kyanite-zone metapelites, kyanite appeared via the pres-sure-sensitive reaction, garnet + muscovite ? kyanite + biotite + quartz. Staurolite was produced in thesame rock by retrograde replacement of kyanite following the reaction, garnet + kyanite + H2O ? stauro-lite + quartz. These reactions depart from the classical kyanite- and staurolite-isograd reactions in low-alumina pelites, encountered in other segments of eastern Himalaya. In the metapelites, just above thekyanite-zone, melting begins in the kyanite field, through water-saturated and water-undersaturatedmelting of paragonite component in white mica. Leucosomes formed through these reactions are charac-teristically free of K-feldspar, with sodic plagioclase and quartz as the dominant constituents. Withincreasing structural height, the melting shifts to water-undersaturated melting of muscovite componentof white mica, producing an early K-feldspar + kyanite and later K-feldspar + sillimanite assemblages andgranitic leucosomes.

Applications of conventional geothermobarometry and average P–T method reveal near isobaric (atP � 8 kbar) increase in peak metamorphic temperatures from 550 �C in the garnet-zone to >700 �C forK-feldspar-kyanite-sillimanite-zone rocks. The findings of near isobaric metamorphic field gradientand by the reconstruction of the reaction history, reveal that the described inverted metamorphicsequence in the western Arunachal Himalaya, deviates from the classical Barrovian-type metamorphism.The tectonic implication of such a metamorphic evolution is discussed.

� 2009 Elsevier Ltd. All rights reserved.

ll rights reserved.

owmik).

1. Introduction

The Himalayan orogen shows a classic example of invertedmetamorphic sequence (IMS), where higher-grade rocks, appear-ing at progressively higher structural levels have been traced along

http://dx.doi.org/10.1016/j.jseaes.2009.07.006

mailto:[email protected]

http://www.sciencedirect.com/science/journal/13679120

http://www.elsevier.com/locate/jseaes

S. Goswami et al. / Journal of Asian Earth Sciences 36 (2009) 390–406 391

the length scale of the orogen from Nanga Parbat in the west toArunachal in the east. Although a general Barrovian style of meta-morphism has been uniformly reported from these sequences,there is still considerable debate on the status and the location ofthe Main Central Thrust (MCT) in the IMS (see Searle et al.(2008) for a review), the nature of metamorphic P–T–t paths andthe metamorphic field gradient across the Lesser and GreaterHimalayan Sequences and the origin of the inverted thermal gradi-ent in the IMS.

Contrasting signatures of metamorphic field gradient havebeen noted by various workers in the different segments of theIMS. Kohn et al. (2001) documented a discontinuity of metamor-phic P–T–t paths near the base of the Main Central Thrust Zone(MCTZ) in central Nepal based on contrasting garnet growth his-tory. More in detail, the structurally lowest garnets grew withincreasing P and T, while the structurally higher garnets grewwith increasing T, but decreasing P. This implies a metamorphicfield gradient (MFG) with negative slope in the P–T space. In theSutlej section of the western Himalaya, the metamorphic fieldgradient across the High Himalayan crystallines is nearly flat,with nearly isobaric (at P � 8 kbar) increase in metamorphictemperature from 600 to 750 �C higher up in the metamorphicsequence (Vannay and Grasemann, 2001). Lack of inverted meta-morphic pressure gradient across the MCT has also been ob-served in the Bhutan Himalaya (Daniel et al., 2003) and inLangtang and Darondi regions, central Nepal (Kohn, 2008). Onthe other hand, Dasgupta et al. (2004, 2009) established a nearlycontinuous metamorphic field gradient across the IMS from theSikkim Himalaya. The positive slope of the MFG is indicated byprogressive increase of both pressures and temperatures at pro-gressively higher structural levels.

Numerous thermo-tectonic models have been proposed duringthe last three decades to explain the origin of the IMS (reviewed inHubbard (1996), Hodges (2000) and Yin et al. (2006)). In recentyears, significant new information on the origin of the IMS andthe exhumation of the deep crustal rocks has emerged from twosegments of the eastern Himalaya, namely the Sikkim–DarjeelingHimalaya and the Bhutan Himalaya. Daniel et al. (2003) presenteda ductile extrusion model by channel flow for the origin of the IMSin the Bhutan Himalaya. Faccenda et al. (2008) attributed this duc-tile extrusion to the formation of a mid-crustal, partially meltedchannel due to an anomalous concentration of radiogenic heat-producing elements in the protolith. According to these latterauthors, the melting-triggered pro-foreland propagation of thechannel was responsible for exhumation of metamorphic rocks(cf. Greater Himalayan Sequence) from different depths (15–23 kbar, �700 �C) on top of a coherent Lesser Himalayan Sequence,thereby producing a coherent and continuous IMS in the SikkimHimalaya. A two stage exhumation of the rocks of the GreaterHimalayan Sequence has been previously modelled by Gangulyet al. (2000) with an initial rapid decompression (exhumation rate,15 mm/yr) being followed by slow cooling and exhumation(�2 mm/yr).

In contrast, there is no information from the crucial Aruna-chal Himalaya, at the eastern part of the Himalayan Metamor-phic front, in relation to (a) the zonal distribution of the indexmetamorphic minerals and the disposition of the metamorphicisograds, (b) the metamorphic reaction history that was respon-sible for the formation of the different mineral zones, (c) the ex-act location of the MCT separating the Lesser (LHS) and theGreater Himalayan Sequences (GHS) and (d) the peak P–T condi-tions of metamorphism recorded in these metamorphic zones. Inthe absence of this information, it is neither possible to recon-struct the basic framework of the IMS in Arunachal Himalayanor to make a correlation with other segments of the HimalayanMetamorphic front. In this paper, we address these issues using

rocks from the western Arunachal Himalaya (WAH), which areexposed along the transect Jira–Rupa–Bomdila–Dirang–Se LaPass. The results provide the first documentation on the develop-ment of the IMS from the WAH, which has important implica-tions in building a coherent model of inverted metamorphicgradient in the eastern Himalaya. Abbreviations of minerals areafter Kretz (1983).

2. Regional geological setting

The present study area in the WAH is located in between theBhutan Himalaya to the west and the eastern Himalayan syntaxisto the east (Fig. 1). This area has been geologically mapped byDas et al. (1975), Thakur (1986), Singh and Chowdhary (1990),Acharyya (1987), Kumar (1997) and Bhattacharjee and Nandy(2008) and others. A plethora of stratigraphic names have beenproposed by these workers, resulting in considerable confusionon the litho-tectonic framework of the WAH. Three major, thrustbounded tectonic units, namely the sub-Himalaya (Siwaliks), theLesser Himalaya (including the Gondwana Group) and the GreaterHimalaya were recognized in these studies, which are describedbelow.

2.1. Sub-Himalaya (SH)

The southern Sub-Himalayan domain of molasse-type Siwalikdeposits of Lower Miocene to Lower Pleistocene age is separatedfrom the structurally overlying Lesser Himalayan Sequence (LHS)by the Main Boundary Thrust (MBT) (Fig. 1) (Kumar, 1997).

2.2. Lesser Himalayan Sequence (LHS)

2.2.1. Gondwana rocks and Bomdila GroupThe basal part of the LHS is marked by a thin strip of marine and

plant fossils-bearing Gondwana rocks of Permian age (Kumar,1997). These are overlain by weakly metamorphosed sequence ofinterlayered dark grey phyllite–sericite quartzite and metavolca-nics of the Bomdila Group, locally referred to as the Tenga Forma-tion (Das et al., 1975; Kumar, 1997). The phyllites are extremelyfine-grained, with a dominant north dipping foliation defined byoriented muscovite.

2.2.2. Bomdila gneiss and Bomdila GroupThese low-grade metamorphites of the basal LHS are structur-

ally overlain by a thick unit of megacrystic granite gneiss (Bomdilagneiss) (Fig. 1). These orthogneisses, which have yielded Palaeo- toEarly Mesoproterozoic Rb-Sr whole rock isochron ages (Dikshituluet al., 1995) resemble the megacrystic Lingtse gneiss of the SikkimHimalayas (Sinha-Roy, 1982), the Ulleri augen gneiss in the Annap-urna region (LeFort, 1975a,b; Pecher and LeFort, 1977), the Numorthogneiss of Arun valley (Brunel, 1983; Lombardo et al., 1993)and the Phaplu augen gneiss in the Everest region of central Nepal(Maruo and Kizaki, 1983).

Previous studies noted a penetrative tectonic foliation in thegranites in the form of an augen gneiss foliation (Verma and Tan-don, 1976; Bhattacharjee and Nandy, 2008). The localization ofsuch penetrative ductile strain in the Bomdila gneiss is also in-ferred from the map pattern in Fig. 1, where the granites are shownto be interleaved with an interlayered sequence of phyllite–quartz-ite (Tenga Formation) and dolostones (Dedza/Chilliepam Forma-tion). Preliminary kinematic analysis of the Bomdila gneiss ledYin et al. (2006) to correlate the deformation to regional top-southmovement during the Cenozoic thrusting. These workers also doc-umented Late Mesoproterozoic (�950 Ma) depositional ages in themetasediments of the Bomdila Group and the Dirang Formation(see later). This raises the possibility that the Bomdila gneiss could

Fig. 1. Geological map of the western Arunachal and adjoining Bhutan Himalayas (modified after Kumar (1997) and Daniel et al. (2003)), showing the distribution of differenttectono-stratigraphic units and the location of the Main Central Thrust (MCT). Other abbreviations used: SH, Sub Himalaya; LHS: Lesser Himalayan Sequence; GHS: GreaterHimalayan Sequence; BG: Bomdila Group. Inset shows the location of the present study in the eastern Himalayan syntaxis.

392 S. Goswami et al. / Journal of Asian Earth Sciences 36 (2009) 390–406

represent tectonic slivers of Palaeo- to Mesoproterozoic basementrocks. The occurrence of tectonic slivers of basement granitic rockswithin the overlying Shumar Formation has also been reportedfrom the LHS of the adjoining Sikkim (Sinha-Roy, 1982) and Bhutan(Ray et al., 1989; Ray, 1995; Dasgupta, 1995) Himalayas.

2.2.3. Dirang FormationThe Dirang Formation, representing the upper part of the LHS,

consists of thick metasedimentary sequence of garnetiferous Ms-Bt schist (Fig. 1), often calcareous, phyllite, sericite quartzite,calc-silicate and Tr-Act marble, with pods and lenses of amphibo-lites. There is a progressive increase in the garnet size with struc-tural height. The pelitic schist in the upper part of the DirangFormation contains sporadic occurrence of kyanite and staurolite,along with garnet and micas (Bhattacharjee and Nandy, 2008).These low-medium grade schists are exposed in the adjoining Bhu-tan Himalaya as the Jaishidanda Formation (Daniel et al., 2003),immediately below the Greater Himalayan Sequence (GHS)(Fig. 1). In the Lumla, Thimbu and Womingla areas in the northwestern extremity of the studied transect, identical low-mediumgrade metamorphites consisting of Qtz-Chl-Bt-Grt schist, quartzite,marble and calc-silicate rocks, and belonging to the Lumla Forma-tion are exposed as tectonic windows within higher-grade rocks ofthe GHS (Fig. 1). Although previously correlated with the Tethyanmetasediments, recent studies infer that the Lumla Formationhas Late Mesoproterozoic depositional age (Yin et al., 2006) andis equivalent to the Dirang Formation, described above (Bhatta-charjee and Nandy, 2008). Yin et al. (2006) dated monazite inclu-sions within garnet from Dirang area using U–Th ion-microprobedating technique, which yielded an age of 10 ± 1.4 Ma (1r). Thiswas taken as evidence that movement along the MCT was as youngas 10 Ma.

2.3. Greater Himalayan Sequence (GHS)

2.3.1. Se La GroupStructurally overlying the Dirang Formation is the Se La Group,

which represents the exposed GHS in the WAH. It is composed ofmedium-high grade crystalline rocks of pelitic to psammo-peliticbulk rock compositions with subordinate calcareous, arenaceous,felsic and mafic components. The stratigraphic unit derives itsname from the Se La pass in the West Kameng district (Bakliwaland Das, 1971). Dhoundial et al. (1989) classified the Se La Groupinto Lower Taliha Formation and Upper Galensiniak Formationon the basis of differences in lithological association and metamor-phic grade. The former, metamorphosed at relatively lower gradeconsists of graphitic schist, calc-silicate gneiss, marble, amphibo-lite and quartzite. The higher grade Galensiniak Formation consistsof schists and gneisses, intruded by tourmaline granite and pegma-tite. In the Se La–Tawang–Bumla section, the rocks are a complexassociation of migmatite, garnetiferous Bt-Pl gneiss, calc-gneiss/marble (diopside and scapolite-bearing), staurolite-bearing schist,tourmaline-bearing leucogranite, quartzite and pegmatite (Bhatta-charjee and Nandy, 2008). These authors noted the presence ofKy + Bt + St-bearing assemblages in the lower part of the Se LaGroup at places like Lish and north of Changla, while sillimanite-bearing assemblages with garnet and biotite are present in themigmatitic upper part. Scapolite-bearing calc-silicate rocks arepresent as enclaves within migmatite near YJN on the Tawang–Bumla road. Previous workers noted increase in the proportionsof leucogranite with structural height, becoming extremely abun-dant in places like Senge, Se La Pass, Jaswantgarh, Mago and Pang-ila. In agreement with mineralogy from other GHSs, theleucogranites in the Se La Group are characteristically two mica-bearing (Ms-Bt) with accessory tourmaline and garnet.

Fig. 2. Deformation fabrics from the basal part of the Main Central Thrust Zone (MCTZ). (a and b) Well foliated (a) and lineated (b) quartzite of the Bomdila Group in contactwith the mylonitic Bomdila gneiss. The lineation is marked by stretched Qtz ribbon. (c) Protolith of the mylonitic Bomdila gneiss. Megacrystic granite with maficmicrogranular enclave showing the development of a protomylonite foliation. (d) Mylonitic foliation in the megacrystic granite, being defined by shape-preferred alignmentof the Kfs porphyroclasts and by the formation of a gneissic banding, consisting of alternate layers of coarse and fine-grained Qtz-feldspar. (e) Formation of myloniticbanding in the megacrystic granite in domains of the highest ductile shear strain. (f) A penetrative shear foliation transposing an early augen gneiss foliation to the right ofthe photograph. Note alignment of the foliated mafic microgranular enclave parallel to the shear foliation. (g) A back-scattered electron image map of a Grt-zone rock,showing two generations of foliation development (see text for details). Boxes mark the locations of textural features, shown in details in Fig. 5.


Fig. 3. Structural (a), lithological (b) and mineral isograd map (c) along the traverse from Jira to Se La. (d) An interpretative cross section showing the development of invertedmetamorphic sequence in the western Arunachal Himalaya.


2.4. Definition of the Main Central Thrust in WAH

Since its first definition as the thrust fault that places high-grade metamorphic rocks of the GHS over lower-grade rocks ofthe LHS (Heim and Gansser, 1939; Gansser, 1964), the MCT, overthe years, has been variously defined as a discontinuity surface,(a) akin to a Himalayan unconformity that separates the structur-ally underlying Palaeo- to Mesoproterozoic Lesser Himalayan Se-

Fig. 4. AFM plot of pelitic schists from the different metamorphic zones of thewestern Arunachal Himalaya. Also shown are the fields of low- and high-aluminapelites after Spear (1993).

quence from the overlying Greater Himalayan Sequence ofNeoproterozoic to Cambrian age (Searle and Rex, 1989; Hubbard,1996; Kohn et al., 2001), although this has been contradicted byGoscombe et al. (2006), (b) which is coincident with the kyanite-isograd in the IMS (Bordet, 1961; LeFort, 1975a; Colchen et al.,1986), (c) marking a contrast in lithological (e.g., Gansser, 1983;Daniel et al., 2003) and Nd isotope composition (e.g., Parrish andHodges, 1996; Ahmad et al., 2000; Robinson et al., 2001; Martinet al., 2005; Richards et al., 2005, 2006) and, (d) which locatesyoung U–Pb and Th–Pb monazite metamorphic ages down thestructural sequence (e.g., Harrison et al., 1997; Catlos et al., 2001,2002). Because of these differential interpretations, the structuralposition of the MCT is found to be ambiguous, varying both alongand across the orogen, often occupying multiple structural heightsin the IMS (e.g., MCT I and MCT II, Arita, 1983). Emphasizing thestructural criteria, Searle et al. (2008) recently defined the MCTas the base of a domain of high ductile strain, along which Tertiarymetamorphic rocks of the GHS are thrusted over unmetamor-phosed and low-grade rocks of the LHS. According to this defini-tion, the MCT coincides with the base of the zone of invertedmetamorphic isograds.

In the WAH, previous workers have defined the Main CentralThrust (MCT) as a metamorphic discontinuity, which separatesthe low-medium grade Dirang Formation from the medium-highgrade Se La Group (Kumar, 1997; Yin et al., 2006; Bhattacharjeeand Nandy, 2008). As in the adjoining Bhutan Himalaya, theMCT, defined in this way, coincides with the first appearance ofkyanite in the migmatites of pelitic bulk rock composition inthe Se La Group (see later). Intense strain localization at the baseof the Se La Group is also inferred by the development of strongmigmatite banding in these rocks (Figs. 3 and 4, Bhattacharjeeand Nandy, 2008), as processes of melt generation and melt seg-regation in planar bands in the migmatites are aided by tectonic

Table 1Tectono-stratigraphic subdivisions of the western Arunachal Himalaya (modified after Das et al., 1975).

Tectonic domains Stratigraphic units Rock units

Greater HimalayanSequence

Se La Group Grt-Sil gneiss/migmatites, Bt–Hbl-Pl gneiss/migmatite, often garnetiferous, garnetiferous calc-silicate gneiss, Tur-(Grt)-bearing two mica granites

MCTLesser Himalayan

SequenceDirang Formation Garnetiferous mica schist, often St- and Ky-bearing, quartzite, calc schist and marble, amphibolitesShear zoneBomdila gneiss Megacrystic granite with mafic microgranular enclaves, augen gneisses and amphibolitesBomdila Group

Tenga Formation Quartzite, quartz schist etc.Chilliepam Formation/Dedza Formation

Dolomites (white and dark grey) and carbonaceous phyllite

Bomdila gneissShear zoneBomdila Group

Tenga Formation Quartzite, Qtz-Chl-sericite schist, talcose schist, thin intercalation of quartzite bandsGondwana rocks Dirty green sandstone and carbonaceous shale with occasional coal bands, grey micaceous sandstone with slaty

partings, white gritty quartzitic sandstone. Plant fossils recordedMBT

Sub-Himalaya Siwalik Grey to green sandstoneFaultSiltstoneFaultPebble to boulder bed

Abbreviations used: MCT = Main Central Thrust; MBT = Main Boundary Thrust.


deformation (Brown, 1995; Sawyer, 1996). Following these crite-rion, we place MCT at the base of the Se La Group. We have addi-tionally noted considerable strain gradient in the exposed LHS, ata structural height, below the defined MCT in the WAH (see be-low). In particular, intense ductile shear strain has been noticedall along the exposed Bomdila gneiss. Following the criterion pro-posed by Pecher (1989) in Nepal, Grujic et al. (1996) in Bhutanand Stephenson et al. (2000) in Zanskar, the entire domain acrossthe Bomdila gneiss, the Dirang Formation and the Se La Groupconstitutes the Main Central Thrust Zone (MCTZ). In Table 1, wehave summarised the different litho-tectonic domains that areencountered in the studied transect Jira–Tenga–Rupa–Bomdila–Dirang–Se La Pass, and its correlations with the adjoining BhutanHimalaya.

3. Fabric development in the MCTZ

In this section, we describe and interpret the different stages offabric development in the rocks of the LHS and GHS. At the struc-turally lower part of the LHS, the signatures of fabric developmentduring tectonic interleaving of metasediments and the basementgneisses are best preserved along the traverse from Jira to Rupa(Fig. 1). Rocks of both these associations are strongly foliated andlineated (Fig. 2a–d). The protolith of the Bomdila gneiss is a por-phyritic granite, with large K-feldspar megacrysts (Fig. 2c), oftenwith rectangular outline being embedded in a coarse-grained ma-trix of quartz, K-feldspar, plagioclase, muscovite, biotite, epidoteand titanite. Mafic microgranular enclaves of various shapes anddimensions are present in the granites (Fig. 2c and f). Towardsthe contacts with both underlying (Bomdila Group) and overlying(Dirang Formation) metasediments, there is progressive develop-ment of metamorphic foliation in the megacrystic granites. Thisis marked by dynamic recrystallisation of K-feldspar megacrysts,producing lensoidal or augen-shaped porphyroclasts, shape-pre-ferred alignment of the porphyroclasts (Fig. 2d) and ribbon quartz,parallel to the foliation plane and the development of a gneissicbanding consisting of alternate layers of coarse and fine-grainedquartz-feldspar. In domains characterized by the highest ductilestrain, a mylonitic banding is developed (Fig. 2e). A penetrativeshear foliation, which transposed an early mylonitic foliation, isalso present in the granite gneiss, just below the Dirang schists

(Fig. 2f). In many cases, this strain localization was associated withfluid infiltration, producing phyllonites.

The structural data of this exposed section are presented inFig. 3a. In the basal part of the LHS, the mylonitic foliation strikesNE-SW and dips moderately (40–50�) towards NW. The stretchinglineation plunges 5–20� towards north to north-northwest. Thesedata in combination with structural features presented in Figs.2b–f suggest that the contact between the Bomdila gneiss andthe metasediments of the Bomdila Group is tectonic, consistentwith previous interpretations (Verma and Tandon, 1976; Yinet al., 2006). This tectonic contact of the Bomdila gneiss with thestructurally overlying metasediments is folded (Fig. 3a), the axisof which plunges 42� ? 209�.

Metasediments of the Dirang Formation record evidence forthree phases of overprinting deformations. Two groups of majorplanar fabrics (S1–S2) are recognized, the earliest of which is re-corded in the hinges of F2 crenulations and as spiral-shaped inclu-sion trails in syn-D1 garnets (Fig. 2g). S1 foliation strikes NE–SWand dips 58�NW. The axial planar schistosity, S2 is the most pene-trative planar fabric in the rock (Fig. 2g). The intersection lineation(S1/S2), which is parallel to the F2 axis plunges 42–54�W. Progres-sively higher up in the structural sequence, the F2 crenulations aresignificantly tightened. In the kyanite-bearing rock, these crenula-tions lie as detached isoclinal fold trains in the thin inter-folial do-mains. In Fig. 3a, we have plotted the trace of the S2 foliation in theDirang Formation, which shows a general NW–SE to NNW–SSEstrike with moderate to steep (35–60�) southwesterly dip.

In the rocks of the Se La Group, traces of S1 are completely oblit-erated, except in rare preserved inclusion trails in garnet and thedominant foliation S2 is a stromatic migmatite banding (Figs. 3and 4, Bhattacharjee and Nandy, 2008). Subsequent folding of theS2 foliation along NW–SE axis has produced an antiformal struc-ture with axial culminations, being marked by plunge reversalsof F3 axis (Fig. 3a).

4. Evolution of mineral assemblages

We have deduced the relationship between mineral growth andfabric development across the MCTZ, based on the petrographicstudy of 50 samples. Fig. 3b shows that, along the studied transect,the following sequences of metamorphic rocks are exposed: (1)


Granite mylonite (Bomdila gneiss), (2) garnetiferous schist andGrt-Ky ± St schist, both intercalated with quartzite (Dirang Forma-tion), (3) Ky-Grt migmatites, (4) Ky-Sil migmatites with interca-lated calc-silicate gneiss and (5) Bt-Kfs-Ky-Sil-Grt migmatiteswith intercalated Bt-Pl gneiss-migmatites and sills and dykes ofleucogranites (Se La Group). We adopt a metamorphic zonal termi-nology, based on the first appearance of an index mineral or min-eral assemblage in the metapelites. The following mineral zoneshave been identified in this study: (1) garnet-zone, (2) kyanite-zone, (3) kyanite-migmatite zone, (4) kyanite-sillimanite-migma-tite zone and (5) K-feldspar-kyanite-sillimanite-migmatite zone.The distribution of the metamorphic isograds is shown in Fig. 3cand also in cross section in the Fig. 3d. Given that mineral growthsin the different metamorphic zones are controlled by bulk rockcompositions, we present bulk chemical analyses for major ele-ments of representative metapelite samples from the differentmetamorphic zones in the next section.

4.1. Bulk rock compositions

Bulk chemical analysis was carried out by X-Ray fluorescencespectroscopy (XRF) at the Ruhr Universität of Bochum on a PhillipsPW2404 instrument. Bulk rock compositions of these samples arepresented in Table 2 and they are compared with average pelites(after Shaw (1956) and Ague (1991)) and sub-aluminous pelite(after Atherton and Brotherton (1982)) are presented in Table 2.These compositions are also graphically shown in AFM diagram,projected from muscovite, plagioclase, quartz, ilmenite and H2O(Fig. 4). Because of these projections, Al2O3 corrections are madeby extracting the amounts of Al2O3 bound by K2O in muscoviteand by those of Na2O and CaO in plagioclase. FeO correction is alsodone by removing an amount of FeO bound by TiO2 in ilmenite. Thecorrected alumina and FeO are shown in the ternary A0F0M diagram(Fig. 4) by the terms A0 and F0. It should be noted that the projec-tions do not consider the alumina content bound by paragonitecomponent in muscovite solid solution, and thus should be consid-ered as qualitative.

The studied rocks, as well as the average amphibolite facies pel-ite of Ague (1991) and the average sub-aluminous pelite of Ather-ton and Brotherton (1982) broadly fall in the field of low-aluminapelites of Spear (1993). In contrast, the average pelite of Shaw(1956) lies in the field of high-alumina pelites of Spear (1993). Ofthe three garnet-zone samples, AR20 is the most depleted in A0

Table 2Representative bulk chemical analyses of samples from the different metamorphic zones

Metamorphic zone Grt Grt Grt KySample no. AR18 AR19 AR20 AR65

SiO2 65.47 67.91 65.21 67.2TiO2 0.59 0.72 0.82 0.82Al2O3 17.55 15.76 15.54 15.83Fe2O3 5.85a 5.76a 6.28a 6.13a

FeO – – – –P2O5 0.07 0.08 0.11 0.16MnO 0.05 0.05 0.05 0.05MgO 1.84 1.62 2.41 2.1CaO 0.62 0.5 0.96 0.68Na2O 1.28 1.41 1.38 1.06K2O 4.87 3.56 4.38 3.8CO2 0.54 0.74 0.52 0.07L.O.I. 2.23 2.45 2.12 2.03Total 100.42 99.82 99.26 99.86Mg# 0.38 0.36 0.43 0.40Ca# 0.31 0.24 0.40 0.33

Abbreviations used, Mig: migmatite; PS: average pelite composition after Shaw (1956); PA

composition after Atherton and Brotherton (1982); n.d.: not determined.a Total Fe as Fe2O3 basis.

component. Substantial CaO and Na2O contents in this bulk rockcomposition are stored in plagioclase (cf. Na2O) and in muscovite,causing a reduction in A0 value in the A0F0M plane (Table 2). Sam-ples AR20 and AR18 are moderately enriched in Ca# (=CaO/(CaO + Na2O), in the range 0.40–0.31) relative to sample AR19(Ca# = 0.24). Compared to the low-alumina pelite of Ague (1991)and the sub-aluminous pelite of Atherton and Brotherton (1982),all the garnet-zone samples of this study are deficient in MnO(�0.05 wt% vs. 0.18–0.10 wt%) (Table 2). Sample AR65 from thekyanite-zone is relatively aluminous, sodic, MnO-poor (MnO =0.05 wt%) and distinctly magnesian (Mg# = 0.40). Migmatitic sam-ples, AR29 (from the kyanite-migmatite zone) and AR37 (K-feld-spar-kyanite-sillimanite-migmatite zone), in contrast, are themost MnO- (=0.18–0.21 wt%) and FeO-rich (Mg# = 0.33–0.34).

Summarising, the rocks from the different metamorphic zonesshow variations in bulk rock compositions, in relation to Mg#,Ca# and Al2O3 contents, which may cause mineral assemblage var-iability at similar metamorphic grade. The migmatites with rela-tively higher Al2O3 and FeO + MgO and lower H2O, SiO2, Na2Oand CaO contents compared to the non-migmatitic rocks appearto reflect their residual character.

4.2. Petrography, mineral chemistry and metamorphic reaction history

A summary of representative mineral chemical compositions isgiven in Table 3. Electron Microprobe analyses were carried outwith a CAMECA SX-100 electron microprobe at the laboratory ofthe Geological Survey of India, Kolkata, using natural and syntheticmineral standards and PAP (CAMECA) correction procedures.

4.2.1. Garnet-zoneThe first garnet growth has been noted at the base of the Dirang

Formation, in the form of small, idioblastic garnets (diameter�0.01–0.03 mm) in fine-grained, thinly banded Qtz-white mica-Bt-Pl-Grt schist with minor chlorite, epidote/zoisite and accessoryzircon, monazite and allanite (sample AR16). An early stability ofMs-Bt-Chl-Ep-Pl-Qtz assemblage is inferred by the occurrence of(a) preserved inclusions of chlorite and epidote within garnet, (b)relics of biotite and muscovite-defined S1 foliation, which is trans-posed by the pervasive S2 foliation, (c) porphyroblasts of musco-vite, being warped by S2 (Fig. 5a) and (d) plagioclase in thepressure shadow of muscovite (Fig. 5a). Garnet growth in sampleAR16 was late-to post-D2 (Fig. 5b). These garnets are Fe-rich

in the western Arunachal Himalaya.

Ky Mig Kfs-Ky-Sil Mig PS PA PAB

AR29 AR37

63.35 65.19 60.26 56.25 64.20.77 0.8 1.05 1.05 0.8917.88 18.22 20.64 20.18 16.697.83a 7.23a 1.41 9.31 n.d.– – 5.49 n.d. 6.260.09 0.06 n.d. n.d. n.d.0.18 0.21 n.d. 0.18 0.11.92 1.87 1.93 3.23 2.130.3 0.32 0.52 1.54 0.840.69 0.8 1.38 1.8 2.114.63 4.35 3.72 4.02 3.380.07 0.25 n.d. n.d. n.d.2.01 0.7 n.d. n.d. n.d.99.65 99.12 96.4 97.56 96.60.33 0.34 0.34 0.41 0.380.23 0.25 0.29 0.49 0.31

: amphibolite facies pelite composition after Ague (1991); PAS: sub-aluminous pelite

Table 3Representative mineral chemical compositions from different metamorphic zones.

Metamorphiczones

Sample no. Textural site Grt Textural site Bt Textural site Ms Textural site Pl Ep St

XPrp XAlm XSps XGrs XMg Ti p.f.u. XPa XCela XMs XAn XPs XMg

(11O) I^Grt M

Grt AR16 C 0.031 0.697 0.072 0.2 M 0.3 0.128 M-C 0.096 0.257 0.642 M-C 0.27 – –R 0.042 0.721 0.01 0.179 – – – – – – – M-R 0.21 – –– – – – – – – – – – – – M-R 0.24 – –

Grt AR18 Grt1-C 0.034 0.595 0.123 0.247 I^Grt 0.45 0.125 I^Grt 0.132 0.178 0.607 M-C 0.15 0.18 –Grt1-IR 0.066 0.752 0.032 0.15 M-C 0.43 0.111 – – – – M-R 0.27 0.16 –Grt1-OR 0.087 0.774 0.008 0.132 – – – – – – – – – – –

Grt AR19 Grt1-C 0.044 0.696 0.106 0.154 M-C 0.49 0.078 I^Bt 0.259 0.093 0.581 M-C 0.13 – –Grt1-R 0.101 0.802 0.007 0.09 M-C 0.48 0.084 M 0.204 0.146 0.543 M-R 0.17 – –Grt1-OR 0.11 0.808 0.005 0.077 – – – – – – – – – – –Grt2-C 0.094 0.774 0.009 0.123 – – – – – – – – – – –Grt2-R 0.116 0.797 0.007 0.08 – – – – – – – – – – –Grt2-C 0.088 0.775 0.007 0.13 – – – – – – – – – – –Grt2-R 0.113 0.807 0.006 0.073 – – – – – – – – – – –

Grt AR20 Grt2-C 0.089 0.71 0.082 0.119 M 0.49 0.156 M 0.067 0.209 0.631 M-C 0.23 - –Grt2-IR 0.09 0.699 0.055 0.156 M-C 0.51 0.134 – – – – I^Grt 0.26 – –Grt2-OR 0.111 0.726 0.038 0.124 – – – – – – – – – – –

Ky AR65 Grt2A-C 0.074 0.711 0.089 0.126 M 0.53 0.147 I^Grt2A 0.213 0.154 0.553 M-C 0.22 – 0.21Grt2A-IR 0.108 0.744 0.008 0.14 M 0.53 0.111 I^Bt 0.143 0.22 0.598 M-R 0.18 – –Grt2A-OR 0.142 0.779 0.006 0.076 – – – M 0.121 0.19 0.57 – – – –Grt2B-C 0.104 0.745 0.007 0.144 – – – – – – – – – – –Grt2B-R 0.148 0.775 0.007 0.07 – – – – – – – – – – –

Ky Mig AR29 C 0.15 0.761 0.056 0.033 M-C 0.45 0.16 M 0.126 0.16 0.623 M 0.14 – –IR 0.154 0.761 0.046 0.039 – – – M 0.126 0.171 0.628 M-R 0.1 – –OR 0.13 0.806 0.035 0.028 – – – – – – – – – – –

Ky-Sil Mig AR31 C 0.147 0.776 0.042 0.04 M 0.41 0.216 M 0.102 0.176 0.657 M 0.17 – –OR 0.14 0.768 0.055 0.036 – – – – – – – – – – –

Abbreviations used, C = core; IR = Inner Rim; OR = Outer Rim; M = Matrix; I = Inclusion; Mig = Migmatite.a XCel = Combined mole fractions of Mg- and Fe-celadonite end members.

S.Gosw

ami

etal./Journal

ofA

sianEarth

Sciences36

(2009)390–

406397

Fig. 5. Back-scattered electron images of Grt-zone rocks. (a) An early generation of Ms occurs as porphyroblast (Ms(P)), which is warped by oriented Ms (Ms(FOL), where FOLrepresents folial) + Qtz + Bt-bearing S2 foliation. (b) Late syn- to Post-D2 Grt porphyroblasts, overgrowing S2 foliation. (c and d) Snowball Grt (Grt1) of early generation,showing spiral-shaped inclusion trails of Qtz, Ms, Ep (c) and Pl (d). (e) Second generation of Grt (Grt2), with idioblastic habit overgrowing both S2 foliation and F2 crenulations.


(XPrp = 0.03–0.04, XPrp = Mg/(Mg + Fe + Mn + Ca)) and show a slightgrowth zoning with rimward enrichment in XPrp (Prp03?04, where? refers to a change in composition from core to rim) and deple-tion in spessartine (Sps07?01) and grossular contents (Grs20?18)(Table 3). Matrix plagioclase exhibits a calcic core (An27) and a so-dic rim (An21–24) (Table 3). Matrix biotite is moderately titanifer-ous (Ti = 0.13 p.f.u. on the basis of 11O) and ferroan (XMg = 0.30,where XMg = Mg/(Mg + Fe)), consistent with Fe-rich bulk composi-tion of the rock. Matrix muscovite in the S2 foliation domain hasthe composition of Pg10Cel26Ms64 (where Cel refers to combinedMg- and Fe-celadonite components) (Table 3).

A progressive increase in the grain size of garnet is observedwith increase in structural heights. At a height of �100 m abovethe garnet-isograd, Grt (�1.9–2.1 mm in diameter) has a snowballshape, exhibiting a large core with spiral Ms-Bt-Ep-Chl-Ilm inclu-sions trail and a thin inclusion-free rim (Fig. 5c) (Sample AR18).The inclusion trails are consistent with syn-D1 growth. Both epi-dote and chlorite are characteristically absent in the matrix S2 foli-

ation, defined by Bt + Ms + Pl + Qtz. Compared to the basal partof the garnet-zone, these snowball garnets in AR18 are more mag-nesian (XPrp = 0.03–0.09 relative to �0.03–0.04 in AR16) and showgrowth zonation, with rimward depletion in spessartine (Sps12?01)and grossular (Grs25?13) contents and enrichment in pyrope con-tents (Prp03?09). Matrix plagioclase is reversely zoned with sodiccores (An15) and calcic rims (An27). Matrix biotite is magnesian(XMg = 0.43) with low Ti content (Ti = 0.11 p.f.u.). Biotite includedin garnet is slightly more magnesian (XMg = 0.45). Muscovites areenriched in paragonite and depleted in celadonite contents(Pg13Cel18Ms61). Epidote included in garnet contains pistacite con-tents in the range �0.16–0.18 (Table 3).

At a height of 135 m above the garnet-isograd, syn-D1 garnetgrowth (Fig. 5d) is still observed in the pelitic schists, but withsome differences (Sample AR19). Prograde epidote is characteristi-cally absent and protracted garnet persisted beyond D2. Post-D2

garnet (Grt2) occurs as small (�0.32–0.84 mm in diameter), idio-blastic grains (Figs. 2g and 5e), overgrowing the matrix S2 foliation


consisting of muscovite, biotite, plagioclase and quartz. Rarely, gar-net2 also occurs as inclusion-free metamorphic overgrowth on gar-net1. In sample AR19, cores of garnet1 are spessartine-(Sps11) andgrossular-rich (Grs15) and poor in pyrope (Prp04) content. In con-trast, outer rims of garnet1 are characteristically magnesium-rich(Prp11) and grossular- (Grs08) and spessartine-poor (Sps601). Gar-net2 is also zoned with a core to rim enrichment in pyrope(Prp09?11–12) and a decrease in grossular (Grs12-13?07-08) contents.Spessartine in garnet2 remains uniformly low (Sps01). Muscovite inthis sample is uniformly sodic (Pg20–25) relative to that from othergarnet-zone rocks. Matrix biotite is magnesian (XMg = 0.48–0.49).Matrix plagioclase shows reverse zoning (An13 ? An17), with calcicrims being in contact with or in the microdomains of garnet2.

In sample AR20, collected from the same structural level asAR19, prograde chlorite and epidote are characteristically absent,and late- to post-D2 garnets occur in two textural modes: (a) Idio-blastic garnets (�1.84 mm in diameter) enclosing a trapped S2

schistosity or F2 crenulations, and (b) elongated garnets (�3.16mm long and 0.81 mm wide) aligned parallel to the S2 schistosity.Equant garnets are zoned with core to rim enrichment in pyrope(Prp09?11) and depletion in spessartine (Sps08 ? Sps04) contents.Plagioclases included in garnet2 and cores of matrix plagioclaseare An23–26. Cores of matrix biotites are magnesian (XMg = 0.49–0.51) and titaniferous (Ti = 0.13–0.16 p.f.u.).

At all structural levels of the garnet-zone rocks, garnet showsminor replacement by chlorite.

Summarising, textural and compositional features attest to twostages of prograde garnet growth in the garnet-zone rocks. Garnet1

appears to have grown in equilibrium with the Ms + Bt + Chl +Pl + Qtz assemblage with or without epidote, according to thetwo model reactions,

ChlþMsþ Cz component in Epþ Qtz! Grtþ Btþ PlþH2OðR1Þ

Chlþ Plþ Btþ Qtz! GrtþMsþH2O ðR2ÞAfter chlorite and epidote were completely consumed, textural andcompositional features suggest that the next stage of garnet growth(garnet2) occurred at a higher metamorphic grade in equilibriumwith the Ms + Bt + Pl + Qtz assemblage. The following model reac-tion may be suggested for the growth of garnet2,

Btþ Pl! GrtþMs ðR3Þ

Systematic increase in XMg in the rims of growth zoned garnets withstructural height (Table 3) additionally reveals that successivelyhigher level garnet-zone rocks were equilibrated at higher meta-morphic temperatures, consistent with the predicted invertedmetamorphism in the WAH.

4.2.2. Kyanite-zoneThe first appearance of kyanite has been noted in the garnetif-

erous schists, at �210 m above the garnet-isograd. Garnet growthin the kyanite-zone rock (sample AR65) is late-to post-D2. Thereare two varieties of garnet2. Paragenetically, early garnets (gar-net2A) are large, ovoid porphyroblasts (�4 mm in diameter) havingtrapped F2 crenulations (Fig. 6a), while elongated garnet2 (garnet2B,�3 mm long and �0.76 mm wide) (Fig. 6b) overgrows the S2 schis-tosity, which consists of Ms + Bt + Pl + Qtz. Kyanite generally oc-curs as long, elongated grains in close spatial association withmuscovite in the S2. Kyanite partially rims layered garnet2B

(Fig. 6c) and muscovite. Flakes of biotite occur both in the S1 andS2 foliation domains. A third variety of biotite occurs as porphyro-blasts overgrowing S2 (Fig. 6d). A few sub-idioblastic to idioblasticgrains of staurolite, intergrown with quartz have been observed.Staurolite cuts across kyanite (Fig. 6e) or partially rims garnet. Sim-ilar to garnet-zone garnets, garnet2A also shows rimward depletionin spessartine (Sps09?01) and grossular (Grs13?08) and enrichment

in pyrope (Prp07?14) contents. Growth zoning in garnet2B is weak,with cores being characteristically spessartine-poor (Sps01). Themost magnesian compositions are noted in the rims of garnet2B

(Prp15). Muscovite inclusions in garnet2A are characteristically so-dic with moderate phengite contents (Pg21Cel15Ms55), similar tothat in garnet-zone sample AR19. In contrast, matrix muscovitesare depleted in paragonite, but enriched in phengite components(Pg12–14Cel19–22Ms57–60). Matrix biotites are magnesian(XMg = 0.53) and moderately titaniferous (Ti = 0.11–0.15 p.f.u.).Rims of matrix plagioclase are sodic (An18) relative to cores(An22). Staurolite is ferroan (XMg = 0.21)

Sample AR65 from the kyanite-zone records identical metamor-phic reaction history as in the adjoining garnet-zone rocks, up tothe growth of garnet2. In low-alumina pelites such as the studiedsample, kyanite could appear following the KFMASH modelreactions,

StþMsþ Qtz! Grtþ Kyþ BtþH2O ðR4ÞMsþ Stþ Qtz! Kyþ BtþH2O ðR5Þ

However, textural features such as (a) the absence of stauroliteinclusions, either in kyanite or in garnet2B, (b) restricted occur-rence of kyanite in the aluminous layer containing muscovite andgarnet2B, with kyanite partially rimming garnet (Fig. 6c), and (c)late biotite porphyroblast overgrowing the matrix muscovite folia-tion (Fig. 6d) are consistent with kyanite growth following the so-lid–solid reaction,

GrtþMs! Kyþ Btþ Qtz ðR6Þ

The progress of the pressure-sensitive reaction, (R6) is stronglydependent on the bulk rock composition, and operates at higherpressures in magnesian and aluminous bulk rock compositions. Inthe studied sample AR65, with a relatively high magnesian compo-sition (Mg# = 0.40), the reaction appears to have progressed in re-sponse to decompression.

The late appearance of garnet after kyanite and garnet suggeststhat staurolite formed according to the model FeO–Al2O3–SiO2–H2O (FASH) end-member reaction,

Almþ KyþH2O! Fe-Stþ Qtz ðR7Þ

The reaction has steep negative slope, and progresses to theright during near isobaric cooling or cooling accompanyingdecompression.

This aspect of kyanite and staurolite growth will be further dis-cussed in a later section.

4.2.3. Kyanite-migmatite zoneThe initiation of melting, just above the kyanite-zone, rock has

been noticed in the metapelites, at a height of �370 m above thegarnet-isograd. The rocks are Ms-Qtz-Bt-Grt schists, with Pl-Qtzleucosomes (without K-feldspar) occurring as thin discontinuousstringers, pods (Fig. 7a) or veins within the dominantly musco-vite-rich matrix. The leucosomes constitute less than 5–10% ofthe rock volume. Garnet porphyroblasts (diameter �0.4–1.8 mm,sample AR29) show an inclusion-rich core, often with trapped S1

foliation at a high angle with respect to the matrix S2. Garnet rimsare characteristically inclusion-free. Rare chlorite is present as pro-grade inclusions within garnet cores, whereas it lacks in the ma-trix. The leucosome consists of coarse aggregates of sodicplagioclase (An10–14) and quartz (Fig. 7b), with quartz occurringin the interstices of plagioclase, or shows delicate intergrowthswith plagioclase. These features are consistent with crystallizedpartial melt. Volumetrically minor kyanite (<1–5 vol%) occurs assmall-medium-sized inclusions within plagioclase (Fig. 7b) ormuscovite (Fig. 7c). Coarse garnet porphyroblasts (grain diameter�1.75 mm) show a slight core to rim depletion in spessartine

Fig. 6. Microphotographs (a, c) and back-scattered electron images (b, d, e) of Ky-zone rocks. (a) Part of a large, ovoid, porphyroblastic Grt (Grt2A) containing trapped F2

crenulations. (b) Elongated Grt porphyroblast (Grt2B) oriented parallel to the matrix S2 foliation. (c) Grt2B is partially rimmed by Ky. Bt rims folial Ms. Box shows the locationof Fig. 6e. (d) Bt porphyroblast (Bt(P)) overgrowing Ms-defined S2 foliation. (e) St with Qtz intergrowth overgrowing Ky-Ilm assemblage. Note the continuity of the Ilminclusion trail within St.


(Sps06?05), but are broadly homogeneous with respect to pyrope(Prp15) and grossular (Grs03-04) contents. Garnet outer rim in con-tact with biotite is enriched in Fe (Prp13). Leucosome plagioclase isAn10–14. Compared to the garnet- and kyanite-zone rocks, matrixmuscovites are depleted in paragonite content (Pg13 vs. Pg21–26).Matrix biotites are ferroan (XMg = 0.45).

Textural features reveal a possible continuity of metamorphicevolution across the LHS and GHS. Record of an early progrademetamorphic stage is inferred by the presence of preserved pro-grade chlorite within garnet cores. The beginning of melting inthe pelites, producing Pl-Qtz leucosomes takes place at higher tem-peratures. Chatterjee (1972) demonstrated that Ky-Ab assemblagecould be produced by paragonite break down reaction of the type,

NaAl2Si3AlO10ðOHÞ2Pg

þ SiO2Qtz! Al2SiO5

KyþNaAlSi3O8

AbþMelt=H2O

ðR8Þ

Dasgupta et al. (2009) attributed this melting in the Sikkim Himala-yas to water-saturated melting at 580–600 �C, primarily involvingparagonite component of muscovite,

Pgþ Abþ QtzþH2O!Melt ðR9Þ

The reaction (R9) also predicts growth of garnet during the melt-ing due to the consumption of celadonite/Fe-celadonite compo-nent of the white mica (Reaction (1) of Chakraborty et al.,2003). These authors related the presence of kyanite in themigmatites due to the operation of two consecutive melting reac-tions, an early hydrous, aluminosilicate-free reaction (R9) and alater paragonite breakdown reaction (R8) (Dasgupta et al.,2009). Spear et al. (1999) predicts that water-saturated meltingof this type in the metapelites produces low-degrees of partialmelting, in the range of 1–2 vol%. This results in substantial dropin paragonite content of mica in residual rock from about 20% to5%.

Fig. 7. Field- (a, d) and microphotographs (b and c, g and h) and back-scattered electron images (d and e) of Ky- (a–c), Ky-Sil- (d and e) and Kfs-Ky-Sil- (f–h) migmatite zonerocks. (a) Incipient leucosome formation in Ky migmatite. (b) Leucosome showing coarse Pl aggregates of subhedral to euhedral crystal outline and interstitial Qtz. Note Kyinclusions within Pl. (c). Ky inclusion within Ms. (d) Leucosome development in Ky-Sil migmatite. Note localized fibrolite-defined S2 foliation within the leucosome domain.(e) The occurrence of fibrolite as thin stringers in Ms. (f) Kfs-Pl-Qtz assemblage in the leucosome domain. Ms occurs as inclusion within Kfs. (g) Ky-bearing Kfs-Pl-Qtzleucosome assemblage. (h) Sil foliation warping around Grt porphyroblast.



In the studied rocks, textural and compositional features suchas: (a) the presence of kyanite in the sodic plagioclase leucosomes,(b) sharp drop in paragonite content in white micas from 26 to13 mol% from garnet- to kyanite-migmatite zone rocks of compa-rable bulk rock composition, (c) relatively magnesian compositionof garnet cores compared to those of garnet- and kyanite-zonerocks, suggest that reactions (R9) and (R8) were responsible forpartial melting in the kyanite-migmatites. In the adjoining BhutanHimalayas, the P–T conditions for the formation of kyanite-migma-tites has been variously constrained at >�700 �C, 8 kbar (Davidsonet al., 1997) and �800 �C, 12–13 kbar (Daniel et al., 2003).

4.2.4. Kyanite-sillimanite-migmatite zoneAlthough having a general resemblance with the kyanite

migmatites by having diagnostic sodic Pl-Qtz leucosomes, thesemetapelitic rocks have been distinguished basing on the firstappearance of fibrolitic sillimanite and rutile, the latter as rareinclusions within garnet and by the presence of volumetricallysignificant proportion of leucosome (�20–30 vol%). These leuco-somes are segregated to produce stromatic migmatite banding(Fig. 7d). Volumetrically minor kyanite is encased either as smallaggregates within leucosome plagioclase or within muscovite.Fibrolite-quartz assemblage occurs as thin stringers or bands,within matrix muscovite (Fig. 7e) and plagioclase. In most ofthe microdomains, a penetrative fibrolite-defined foliation (=S2),which warps around the leucosomes has been observed(Fig. 7d). As in sample AR29, garnet is compositionally homoge-neous with low grossular (Grs04) and moderate pyrope (Prp15)contents. Garnet rims against matrix biotites are slightly en-riched in Fe (Prp14). Similar to kyanite migmatites, muscovitesare impoverished in paragonite contents (Pg10). Matrix biotitesare ferroan (XMg = 0.41) and titaniferous (Ti = 0.22 p.f.u.). Plagioc-lases in the leucosomes are slightly more calcic (An17) comparedto those in the kyanite migmatites.

The occurrence of fibrolite in the Pl-Qtz leucosomes (Fig. 7d) inexclusion of Kfs, and often replacing plagioclase and muscovite(Fig. 7e) can be best explained by the operation of various baseleaching reactions (cf. Vernon, 1979), postdating the kyanite-faciesmelting described before.

4.2.5. K-feldspar-kyanite-sillimanite-migmatite zoneThe first appearance of K-feldspar in the migmatites, at a struc-

tural height of �560 m above the garnet-isograd coincides withdecreasing muscovite and increasing sillimanite contents in therock and distinct grain coarsening of the fibrolite.

At �670 m above the garnet-isograd, the pelitic migmatitesshow abrupt increase in the modal abundance of biotite overmuscovite, high content of K-feldspar and well-developedmigmatite banding, defined by alternate layers of oriented, bladedsillimanite and biotite, porphyroblastic garnet and Kfs + Pl + Qtzlayers. K-feldspar contains inclusions of muscovite (Fig. 7f). Inthe leucosomes, K-feldspar has lobate to amoeboid grain contactswith plagioclase and quartz (Fig. 7f). Imprints of an even earlierstage of kyanite-facies partial melting are, however, preserved inthese rocks by the occurrence of kyanite-bearing Pl-Kfs-Qtzmineral assemblage (Fig. 7g), as small pods and lenses within per-vasive Bt + Kfs + Qtz + Pl ± Sil leucosomes. The Bt-Sil foliationwarps around porphyroblastic garnets (Fig. 7h).

Textural features suggest progression of muscovite meltingfrom an early kyanite to later sillimanite field, following the modelreaction,

Msþ Plþ Qtz! Kfsss þ Ky=SilþMelt ðR10Þ

Reaction (R10) led to near complete removal of muscovite from therock.

5. Geothermobarometry

P–T conditions of peak metamorphism from the different meta-morphic zones have been constrained using thermobarometryapproach. We have chosen both individually calibrated thermoba-rometers and the THERMOCALC average P–T method. For the latter,we have used the internally consistent thermodynamic dataset 5.5of Holland and Powell (1998: Nov. 2003 upgrade) and the THER-MOCALC program (v.3.26). It has been shown in this study thatgarnet is compositionally zoned. The nature of mineral zonationis a complex function of the thermo-baric history of the rocks,the mineral grain size and their modal abundance and also of thetypes of the involved net transfer reactions (Chakraborty and Gan-guly, 1991; Spear, 1993; Dasgupta et al., 2004, 2009). Inferringmeaningful equilibrium composition is, therefore, central to thereconstruction of the thermal history of the IMS from the WAH.Two critical studies from the adjoining Sikkim Himalayas havedemonstrated that the choice of equilibrium mineral compositionto constrain peak metamorphic temperatures across the IMS is dic-tated by the grade of the concerned metamorphism (Dasguptaet al., 2004, 2009). In our study, we have also noted that contrast-ing garnet compositional zoning across the MCTZ. Up to the kya-nite metamorphic zone, growth zoning in garnet is unaffected bydiffusional modification. This suggests that the garnet outer rim(Mg-rich and Mn-poor) can be paired with matrix biotite to con-strain TMax. From the kyanite-migmatite grade upwards, garnetgrowth zoning is partially modified by diffusion zoning. For rocksof these settings, garnet inner rim, with maximum XMg and mini-mum spessartine contents, and beyond the zone affected by diffu-sion, is assumed to mark the closest approximation of peakmetamorphic composition. Garnets of this composition are thuscombined with matrix biotites to estimate peak thermal condition.Temperatures have been estimated using the Fe–Mg exchangeequilibrium between garnet and biotite at a reference pressure.Pressure has been calculated at a reference temperature, applyingthe Grt-Ms-Bt-Pl-silica (GMBPS) and Grt-aluminosilicate-silica-Pl(GASP) equilibria. For Grt-Bt thermometry, the calibrations ofHoldaway (2000) and Ganguly et al. (1996) have been used. Theformer employs the garnet activity models of Berman and Arano-vich (1996), Ganguly et al. (1996) and Mukhopadhyay et al.(1997). For the GMBPS and GASP barometers, we have used the for-mulations of Hoisch (1990) and Holdaway (2001), respectively.Plagioclase activity model of Newton et al. (1980) has been usedfor this purpose. The results of P–T calculations are presented inTable 4.

For garnet-zone rocks, Grt-Bt thermometry and THERMOCALCaverage T estimates show consistent increase in peak metamorphictemperatures from 535–585 �C (average 560 ± 40 �C, 2r error) insample AR16 at the basal part of the garnet-zone to 575–680 �C(average 620 ± 70 �C) in sample AR20, at the uppermost part ofthe same metamorphic zone. For the kyanite-zone rock (sampleAR65), the peak temperature estimates vary in the range of 570–650 �C with an average of 610 ± 60 �C, which is close to that re-trieved from the uppermost garnet-zone sample, AR20. The differ-ential appearance of kyanite in these two samples under the samethermal and baric conditions (see later) and despite their close spa-tial association is related to the variation in the bulk rock compo-sitions. Kyanite formation in AR20 is possibly prevented by its sub-aluminous and calcic bulk rock composition. In the kyanite- andkyanite-sillimanite-zone migmatites, the estimated peak tempera-tures are broadly the same (690 ± 30 �C and 695 ± 55 �C, respec-tively) and exceed by 80–85 �C from that of the kyanite-zonerocks. Presence of Fe3+ in garnet and biotite reduces temperatureestimates by an average of �15–40 �C in all these metamorphiczones. For this calculation, we have considered 3 mol% of total

Tabl

e4

Resu

lts

ofge

othe

rmob

arom

etry

for

peak

met

amor

phis

min

the

diff

eren

tm

etam

orph

iczo

nes.

Met

amor

phic

zon

eSa

mpl

en

o.G

ASP

GM

BPS

T Ref

GB

P Ref

Ave

rage

P–T

P H(G

)P H

(B)

P H(M

)P H

OI

P(A

v)H

PT H

(G)

T H(B

)T H

(M)

T Gan

T(A

v)H

P

Grt

AR

16A

––

–8.

28

550

584(

563* )

534(

506* )

563(

551* )

563(

554* )

538

88.

0±

0.3;

552

±45

AR

16B

––

–7.

87.

955

058

4(56

3* )53

4(50

6* )56

3(55

1* )56

3(55

4* )58

58

Grt

AR

18A

––

–8

8.7

575

598(

571* )

562(

531* )

594(

572* )

555(

535* )

600

88.

1±

0.9;

574

±50

AR

18B

––

–8

7.5

575

614(

587* )

580(

549* )

607(

586* )

582(

561* )

597

8A

R18

C–

––

88.

657

559

8(57

1* )56

2(53

1* )59

4(57

2* )55

5(53

5* )59

78

Grt

AR

19E

––

–8.

38

575

613(

583* )

594(

562* )

612(

586* )

568(

537* )

640

98.

1±

0.5;

581

±65

AR

19F

––

–8.

27.

757

560

2(57

2* )58

6(55

4* )60

2(57

7* )55

3(52

4* )63

49

AR

19G

––

–8.

37.

857

560

0(57

1* )58

3(55

1* )60

1(57

6* )54

9(52

0* )64

39

Grt

AR

20A

––

–9.

28.

660

063

0(60

0* )60

1(56

9* )62

6(60

1* )59

7(56

6* )67

89

8.9

±1.

0;60

3±

70A

R20

B–

––

9.4

8.4

600

618(

596* )

590(

565* )

614(

597* )

576(

543* )

680

9

Ky

AR

65C

8.03

88.

28.

78.

460

061

9(58

6* )60

0(56

5* )62

0(59

1* )57

1(54

4* )65

19

8.3

±0.

6;59

4±

65

Ky

Mig

AR

29A

8.1

88.

88.

58.

467

569

4(66

3* )68

8(65

4* )68

8(66

0* )70

6(67

3* )66

58

8.4

±0.

6;67

7±

35

Ky-

Sil

Mig

AR

31A

76.

86.

67.

77.

770

069

4(66

5* )69

0(65

9* )68

7(66

2* )73

5(69

7* )68

68

7.4

±1.

0;68

5±

50A

R31

B7

7.1

7.8

7.8

870

068

8(67

5* )68

4(67

0* )68

2(63

0* )72

2(72

0* )68

48

Abb

revi

atio

ns

use

dfo

rP

(kba

r)an

dT

(�C

)es

tim

ates

:P R

ef/T

Ref

:re

fere

nce

Pan

dT;

GA

SP/G

MB

PS:

Grt

-Alu

min

osil

icat

e-Q

tz-P

l/G

rt-M

s-B

t-Pl

-Qtz

geob

arom

eter

s;G

B:

Grt

-Bt

geot

her

mom

eter

;P H

(G)–

T H(G

)/P H

(B)–

T H(B

)/P H

(M)–

T H(M

):G

ASP

baro

met

ryan

dG

rt-B

tth

erm

omet

ryaf

ter

Hol

daw

ay(2

001)

and

Hol

daw

ay(2

000)

,res

pect

ivel

y,u

sin

gac

tivi

tym

odel

sof

Gan

guly

etal

.(19

96),

Ber

man

and

Ara

nov

ich

(199

6)an

dM

ukh

opad

hya

yet

al.(

1997

);P H

OI:

GM

BPS

baro

met

ryaf

ter

Hoi

sch

(199

0);

T GA

N:

Grt

-Bt

ther

mom

etry

afte

rG

angu

lyet

al.(

1996

);P(

Av)

HP/T

(Av)

HP:

aver

age

Pan

dT

esti

mat

esu

sin

gTh

erm

ocal

cpr

ogra

m(v

.3.2

6);

Tes

tim

ate

wit

has

teri

sk:

Tes

tim

ate

assu

min

g3%

ofto

talF

ein

Grt

and

11.6

%of

tota

lFe

inB

tto

bein

Fe3

+st

ate;

Erro

rin

Pan

dT

esti

mat

esas

2rva

lues

.


Fe in garnet and 11.6 mol% of that in biotite to be in the Fe3+

state following Holdaway (2000). A statistical average of peakmetamorphic temperatures in the different zones is presented inTable 4.

Irrespective of calibrations, applications of the GMBPS and GASPgeobarometers yield near identical pressure ranges for the differ-ent metamorphic zones (garnet-zone: 8.0 ± 0.34–8.9 ± 1 kbar, kya-nite-zone: 8.3 ± 0.6 kbar, kyanite-migmatite zone: 8.4 ± 0.6 kbar)(Table 4). Somewhat lower pressure of 7.4 ± 1 kbar has been ob-tained for the kyanite-sillimanite-migmatite zone, consistent withthe stability of sillimanite in the rock.

6. Discussion

We now summarise the results of the reconstructed metamor-phic reaction history and geothermobarometry from the WAH.The results show consistent pattern of distribution of the meta-morphic zones in the MCTZ, with progressive increase in metamor-phic grades with structural height. The following five metamorphiczones have been identified in the studied transect, which withincreasing structural height are: garnet-, kyanite-, kyanite-migma-tite-, kyanite-sillimanite-migmatite, and Kfs-Ky-Sil-migmatitezones. Our observation on the temporal relation between mineralgrowth and fabric development suggests that zonal boundariesand metamorphic isograds are broadly aligned parallel to the S2

foliation domains, except for some garnet-zone rocks. In theserocks, an even earlier phase of syn-D1 growth of garnet1 in chlorite-and/or epidote-bearing mineral assemblage fields records animportant phase of prograde metamorphism. In Fig. 8a, we havepresented the P–T estimates of peak metamorphism from the dif-ferent metamorphic zones, which clearly demonstrate progres-sively higher metamorphic temperatures with increasingstructural height. In agreement with the characteristic mineralassemblage stabilities in the different metamorphic zones, thesecalculated P–T estimates convincingly establish inverted metamor-phism in the WAH. The P–T results shown in Fig. 8a demonstratekey differences with those from the adjoining Sikkim and BhutanHimalayas, in relation to (a) the P–T stability of the garnet-zonerocks, (b) the nature of the dP/dT gradient of the prograde segmentof the P–T path and (c) the metamorphic field gradient. These arediscussed below.

For the Sikkim Himalayas, Dasgupta et al. (2004) estimated theP–T conditions of peak garnet-zone metamorphism at �4.8 kbar,480 �C. Significantly higher P–T conditions at P � 11–12 kbar,�650 �C have been calculated for the Bhutan Himalayas (Danielet al., 2003). There may be two reasons for this apparent discrep-ancy: (1) The Bhutan rocks are not classical garnet-zone rocks.The garnet growth at this P–T condition, which is well within themiddle-upper amphibolite facies metamorphic condition, is con-trolled by bulk rock composition. (2) The P–T gradients at thesetwo locations are significantly different. In the Bhutan Himalaya,the estimated apparent P–T gradient of 15–17 �C/km is steeperthan that in the Sikkim Himalaya (29 �C/km) implying differentburial mechanisms of rocks in these two sectors. For the Bhutanrocks, the burial was along a steep dP/dT gradient, which is akinto continental subduction.

The studied garnet-zone rocks in the WAH were metamor-phosed at much deeper crustal level compared to those of the Sik-kim rocks (Fig. 8a). While the apparent steep P–T gradient of20 �C/km of these rocks is akin to that recorded in the BhutanHimalaya, the estimated peak metamorphic pressures are lowerby 3–5 kbar. This suggests that the two stages of garnet growth(garnet1 and garnet2) recorded here occurred along a progradeP–T path, which is steeper in the dP/dT gradient than the classicalBarrovian-type metamorphism. Significantly, our study alsoshows that classical Barrovian-style staurolite-, kyanite- and

Fig. 8. (a) P–T diagram showing the P–T estimates (with 2r error bar) of peak metamorphism from the different metamorphic zones of western Arunachal Himalaya. Theshaded field formed by the intersection of NKASH Ms melting reaction (after Spear et al., 1999) and Ky = Sil reaction constrains the P–T stability of early Ky-bearingassemblage in Kfs-Ky-Sil migmatites. Also shown are the peak P–T conditions of various metamorphic zones from the adjoining Sikkim (after Dasgupta et al. (2004)) andBhutan Himalayas (after Daniel et al. (2003)). (b) Metamorphic field gradients from the Arunachal (this study), Sikkim and Bhutan Himalayas.


Grt-Ky mineral zones are not exposed in the Bomdila–Se La passtransect, and that the staurolite in the kyanite-zone rock is of ret-rograde origin. This raises the possibility that WAH witnessed anon-classical Barrovian inverted metamorphism.

The metamorphic field gradient established here is one ofincreasing peak temperatures at nearly uniform metamorphicpressures (Fig. 8b). The lack of an inverted pressure gradient acrossthe MCT is similar to that obtained for the Sutlej Valley region, NWHimalaya (Vannay and Grasemann, 1998; Vannay et al., 1999) andthe Bhutan Himalayas (Daniel et al., 2003), but is different fromthat of the Sikkim Himalaya (Dasgupta et al., 2004) (Fig. 8b). Thisfinding provides some constraints on the origin of the IMS in theWAH. Numerous models have been proposed over the years to ex-plain the IMS in the Himalayas, which can be classified into fourbroad types (summarised in Dasgupta et al., 2004, 2009). Theseare: (a) Hot iron model of Frank et al. (1973) and LeFort (1975a),which invokes thermal inversion of the metamorphic sequencedue to the emplacement of a hot thrust sheet on the top of colderfootwall rocks, along the MCT. Since in this model, heating takesplace from the top to the bottom, a metamorphic field gradientof increasing T in the direction of decreasing P is expected, and thatthe footwall rock should show a near isobaric heating P–T trajec-tory. (b) Tectonic inversion models that explain inversion of ‘‘rightside down metamorphic isograds” by recumbent folding (Searleet al., 1999), thrusting (Jain and Manickavasagam, 1993; Hubbard,1996; Grujic et al., 1996) and by other mechanical processes. (c) Amodified model of tectonic imbrication by thrusting in a fold-thrust system, whereby foreland propagating thrust systems pro-duce progressively younger and lower grade metamorphic assem-blages at the base of the structural sequence (Robinson et al.,2003). The model predicts gradients both in metamorphic age aswell as in P and T across the LHS such that peak P–T conditionsare expected to decline at successive stages of metamorphism.From central Nepal, Kohn (2008) recently obtained key metamor-phic and geochronological data in support of foreland propagatingthrust system. (d) Combined thermo-mechanical models, whichinvoke inversion of the right side metamorphic sequence due tosyn- to late-metamorphic ductile extrusion of the partially meltedGHS (Jamieson et al., 1996, 2002; Beaumont et al., 2001; Vannayand Grasemann, 2001; Daniel et al., 2003; Faccenda et al., 2008;Dasgupta et al., 2009).

For the studied IMS in the WAH, the lack of critical geochrono-logical and P–T path information from rocks of the LHS and GHSdo not allow formulation of a unique thermo-mechanical model.

However, our preliminary petrological and microstructural find-ings, in particular the documentation of near isobaric metamor-phic field gradient allow comparison with the end membermodels presented above. From the Sutlej section, Vannay andGrasemann (2001) suggested that during the subduction of theIndian plate, the isotherms were folded with a dip towards theforeland. Subsequent thermal relaxation caused progressive nearisobaric shifting of the peak metamorphic temperatures towardsthe foreland. Despite cooling and tectonically-driven exhumationof these rocks, the record of near isobaric decrease in TMax to-wards the foreland was retained in the rock record. For the Bhu-tan Himalayas, Daniel et al. (2003) proposed near co-evalunderthrusting of the LHS and the GHS to lower crustal depthsalong a tectonic zone, which was precursor to the modern MCTZ.According to these authors, the underthrusting and metamor-phism of the two units in Bhutan Himalaya underwent at c.22 Ma. Because of the likely flattening of the shear zone at depth,there was no inverted pressure gradient. By 16–12 Ma, ductileback flow of the low viscosity, partially melted rocks of theGHS was initiated as part of a south-directed channel flow in be-tween the MCT and the STDS, producing an inversion of the meta-morphic isograds across the MCT. From the Langtang and Darondiregions, central Nepal, Kohn (2008) recently proposed a criticaltaper model for the origin of IMS. The model invokes the forma-tion of a wedge of deforming rock, which is at a condition of crit-ical failure everywhere. The wedge maintains its size andequilibrium by combined processes of erosion at top and accre-tion or underplating of new material at its base. The proposedmodel adequately explains several key features of the IMS fromdifferent sectors of the Himalayan metamorphic front, namely,(1) near isobaric metamorphic field gradient across the LHS andthe GHS, (2) hairpin P–T loops of the rocks of the LHS and (3) pro-gressively younging peak metamorphic ages down the structuralsequence. Seen in this context, the metamorphic field gradientdocumented here appears to suggest some resemblance to themodels described above. However, these models need to be fur-ther tested with detailed geochronological and metamorphic P–T path data.

Acknowledgements

We acknowledge research grant from the Department of Sci-ence and Technology, Government of India (Grant No. ESS/16/242/2005/Kameng/01). This contribution constitutes part of the


PhD dissertation work of the first author, for which she acknowl-edges the support from the Department of Geology and Geophys-ics, IIT Kharagpur. Part of the revision of the manuscript wascarried out in the Department of Applied Geology, Curtin Univer-sity of Technology for which S.K.B. acknowledges TiGER VisitingSenior Research Fellowship. S.D. also acknowledges support re-ceived through J.C. Bose Fellowship. The authors thank Sumit Cha-kraborty for XRF analysis and his helpful suggestions on an earlierdraft. Exhaustive review comments by C. Guilmette and C. Groppoand competent editorial handling by J. Liou have helped to improvethe quality of the manuscript.

References

Acharyya, S.K., 1987. Cenozoic plate motions creating the Eastern Himalaya andIndo-Burmese range around the northeast corner of India. In: Ghosh, M.C.,Varadarajan, S. (Eds.), Ophiolites and Indian Plate Margins. Patna University,Patna, pp. 143–160.

Ague, J.J., 1991. Evidence for major mass transfer and volume strain during regionalmetamorphism of pelites. Geology 19, 855–858.

Ahmad, T., Harris, N., Bickle, M., Chapman, H., Bunbury, J., Prince, C., 2000. Isotopicconstraints on the structural relationships between the Lesser Himalayan Seriesand the High Himalayan Crystalline Series, Garhwal Himalaya. GeologicalSociety of American Bulletin 112, 467–477.

Arita, K., 1983. Origin of the inverted metamorphism of the lower Himalayas,Central Nepal. Tectonophysics 95, 43–60.

Atherton, M.P., Brotherton, M.S., 1982. Major element composition of the pelites ofthe Scottish Dalradian. Geological Journal 17, 185–221.

Bakliwal, P.C., Das, A.K., 1971. Geology of parts of Kameng districts, NEFA.Geological Survey of India Unpublished Progress Report.

Beaumont, C., Jamieson, R.A., Nguyen, M.H., Lee, B., 2001. Himalayan tectonicsexplained by extrusion of a low-viscosity crustal channel coupled to focusedsurface denudation. Nature 414, 738–742.

Berman, R.G., Aranovich, L.Y., 1996. Optimized standard state and solutionproperties of minerals. I. Model calibration for olivine, orthopyroxene,cordierite, garnet, and ilmenite in the system FeO–MgO–CaO–Al2O3–TiO2–SiO2. Contributions to Mineralogy and Petrology 126, 1–24.

Bhattacharjee, S., Nandy, S., 2008. Geology of the Western Arunachal Himalaya inparts of Tawang and West Kameng districts, Arunachal Pradesh. Journal of theGeological Society of India 72, 199–207.

Bordet, P., 1961. Recherches geologiques dans l’Himalaya du Nepal, region duMakalu. Editions du Centre National de la Recherche Scientifique, Paris. 275 pp..

Brown, M., 1995. The late Precambrian geodynamic evolution of the Armoricansegment of the Cadomian belt (France): distortion of an active continentalmargin during south-west directed convergence and subduction of abathymetric high. Gelogie de la France 3, 3–22.

Brunel, M., 1983. Etude petro-structurale des chevauchements ductile en Himalaya(Nepal oriental et Himalaya du Nord-Ouest). These Doct Sciences, Universite’deParis VII, p. 395.

Catlos, E.J., Harrison, T.M., Kohn, M.J., Grove, M., Ryerson, F.J., Manning, C.E., Upreti,B.N., 2001. Geochronologic and thermobarometric constraints on the evolutionof the Main Central Thrust, Central Nepal Himalaya. Journal of GeophysicalResearch 106, 16177–16204.

Catlos, E., Harrison, T.M., Dubey, C.S., Edwards, M.A., 2002. P–T–t constraints on theevolution of the Sikkim Himalaya, Abstract. Journal of Asian Earth Sciences 20,6–7.

Chakraborty, S., Ganguly, J., 1991. Compositional zoning and cation diffusion inaluminosilicate garnets. In: Ganguly, J. (Ed.), Diffusion, Atomic Ordering andMass Transfer, Advances in Physical Geochemistry, vol. 8. Springer-Verlag,Berlin, Heidelberg, NewYork, Toronto, pp. 120–175.

Chakraborty, S., Dasgupta, S., Neogi, S., 2003. Generation of migmatites and thenature of partial melting in a continental collision zone setting: an examplefrom the Sikkim Himalaya. Indian Journal of Geology 75, 38–53.

Chatterjee, N.D., 1972. The upper stability limit of the assemblageparagonite + quartz and its natural occurrences. Contributions to Mineralogyand Petrology 34, 288–303.

Colchen, M., LeFort, P., Pecher, A., 1986. Geological Researches in the Nepal’sHimalayan: Annapurna-Manaslu-Ganesh Himal. Editions du Centre National dela Recherche Scientifique, Paris, France.

Daniel, C.G., Hollister, L.S., Parrish, R.R., Grujic, D., 2003. Exhumation of the MainCentral Thrust from lower crustal depths, Eastern Bhutan Himalaya. Journal ofMetamorphic Geology 21, 317–334.

Das, D.P., Bakliwal, P.C., Dhoundial, D.P., 1975. A brief outline of geology of parts ofKameng district, NEFA. Geological Survey of India Miscellaneous Publication 24,115–127.

Dasgupta, S., 1995. Jaishidanda formation. In: Bhargava, O.N. (Ed.), The BhutanHimalaya: A Geological Account. Geological Survey of India Special Publication,vol. 39, pp. 79–88.

Dasgupta, S., Ganguly, J., Neogi, S., 2004. Inverted metamorphic sequence in theSikkim Himalayas: crystallization history, P–T gradient and implications.Journal of Metamorphic Geology 22, 395–412.

Dasgupta, S., Chakraborty, S., Neogi, S., 2009. Petrology of an Inverted Barroviansequence of metapelites in Sikkim Himalaya, India: constraints on the tectonicsof inversion. American Journal of Science 309, 43–84.

Davidson, C., Grujic, D.E., Hollister, L.S., Schmid, S.M., 1997. Metamorphic reactionsrelated to decompression and synkinematic intrusion of leucogranite, HighHimalayan Crystallines, Bhutan. Journal of Metamorphic Geology 15,593–612.

Dhoundial, D.P., Kumar, G., Singh, S., Reddy, K.V.S., 1989. Geology of ArunachalHimalaya. Geological Survey of India Unpublished Progress Report.

Dikshitulu, G.R., Pandey, B.K., Krishna, V., Dhana, R., 1995. Rb-Sr systematic ofgranitoids of the Central Gneissic Complex, Arunachal Himalaya: implicationson tectonics, stratigraphy, and source. Journal of the Geological Society of India45, 51–56.

Faccenda, M., Gerya, T.V., Chakraborty, S., 2008. Styles of post subduction collisionalorogeny: influence of convergence velocity, crustal rheology and radiogenicheat production. Lithos 103, 257–287.

Frank, W., Hoinkes, G., Miller, C., Purtscheller, F., Richter, W., Thoeni, M., 1973.Relations between metamorphism and orogeny in a typical section of the IndianHimalayas; NW Himalaya; S-Lahul, Kulu; Himachal Pradesh; firstcomprehensive report. TMPM Tschermaks Mineralogische PetrographischeMitteilungen 20, 303–332.

Ganguly, J., Cheng, W., Tirone, M., 1996. Thermodynamics of aluminosilicate garnetsolid solution: new experimental data, an optimized model and thermometricapplication. Contributions to Mineralogy and Petrology 126, 137–151.

Ganguly, J., Dasgupta, S., Cheng, W., Neogi, S., 2000. Exhumation history of a sectionof the Sikkim Himalayas, India: records in the metamorphic mineral equilibriaand compositional zoning of garnet. Earth and Planetary Science Letters 183,471–486.

Gansser, A., 1964. Geology of the Himalaya. Wiley-Interscience, London. p. 284.Gansser, A., 1983. Geology of the Bhutan Himalaya, Basle. Denkenschrift der

SchweizerischenNaturforschendenGesellschaft. Band. 96. Basel, Birkhauser, p.181.

Goscombe, B., Gray, D., Hand, M., 2006. Crustal architecture of the Himalayanmetamorphic front in eastern Nepal. Gondwana Research 10, 232–255.

Grujic, D., Casey, M., Davidson, C., Hollister, L.S., Kündig, R., Pavlis, T., Schmid, S.,1996. Ductile extrusion of the Higher Himalayan Crystalline in Bhutan:evidence from quartz microfabrics. Tectonophysics 260, 21–43.

Harrison, T.M., Ryerson, F.J., LeFort, P., Yin, A., Lovera, O.M., Catlos, E.J., 1997. A LateMiocene–Pliocene origin for the Central Himalayan inverted metamorphism.Earth and Planetary Science Letters 146, E1–E7.

Heim, A., Gansser, A., 1939. Central Himalaya: geological observations of the Swissexpedition 1936. Denkschriften der Schwieizerishen NaturforschendenGasellschaft, Abhandlung 1, 245.

Hodges, K.V., 2000. Tectonics of the Himalaya and Southern Tibet from twoperspectives. Bulletin of the Geological Society of America 112, 324–350.

Hoisch, T.D., 1990. Empirical calibration of six geobarometers for the mineralassemblage quartz + muscovite + biotite + plagioclase + garnet. Contributions toMineralogy and Petrology 104, 225–264.

Holdaway, M.J., 2000. Application of new experimental and garnet margules data tothe garnet-biotite geothermometer. American Mineralogist 85, 881–892.

Holdaway, M.J., 2001. Recalibration of the GASP geobarometer in light of recentgarnet and plagioclase activity models and versions of the garnet-biotitegeothermometer. American Mineralogist 86, 1117–1129.

Holland, T.J.B., Powell, R., 1998. An internally consistent thermodynamic dataset forphases of petrological interest. Journal of Metamorphic Geology 16, 309–343.

Hubbard, M.S., 1996. Ductile shear as a cause of inverted metamorphism: examplefrom the Nepal Himalaya. Journal of Geology 104, 493–499.

Jain, A.K., Manickavasagam, R.M., 1993. Inverted metamorphism in theintracontinental ductile shear zone during Himalayan collision tectonics.Geology 21, 407–410.

Jamieson, R.A., Beaumont, C., Hamilton, J., Fullsack, P., 1996. Tectonic assembly ofinverted metamorphic sequences. Geology 24, 839–842.

Jamieson, R.A., Beaumont, C., Nguyen, M.H., Lee, B., 2002. Interaction ofmetamorphism, deformation and exhumation in large convergent orogens.Journal of Metamorphic Geology 20, 9–24.

Kohn, M.J., 2008. P–T–t data from central Nepal support critical taper and repudiatelarge-scale channel flow of the Greater Himalayan Sequence. Bulletin of theGeological Society of America 120, 259–273.

Kohn, M.J., Catlos, E.J., Ryerson, F.J., Harrison, T.M., 2001. Pressure–temperature–time path discontinuity in the Main Central Thrust Zone, Central Nepal. Geology29, 571–574.

Kumar, G., 1997. Geology of Arunachal Pradesh. Journal of the Geological Society ofIndia, Bangalore. p. 217.

Kretz, R., 1983. Symbols for rock-forming minerals. American Mineralogist 68, 277–279.

LeFort, P., 1975a. Himalayas: the collided range. Present knowledge of thecontinental arc. American Journal of Science 275, 1–44.

LeFort, P., 1975b. Les formations cristallophyliennes de la ‘‘Dalle du Tibet” enMarsyandi. In: Bodet, P. (Ed.), Recherches geologiques dans l’Himalaya du Nepalregion du Nyi-Shang. Paris Editions du Centre National de la RechercheScientifique, Paris, pp. 21–47.

Lombardo, B., Pertusati, P., Borghi, S., 1993. Geology and tectonomagmatic evolutionof the eastern Himalaya along the Chomolungma-Makalu transect. Himalayantectonics 74, 341–355.

Martin, A.J., DeCelles, P.G., Gehrels, G.E., Patchett, P.J., Isachsen, C., 2005. Isotopicand structural constraints on the location of the Main Central Thrust in the


Annapurna Range, central Nepal Himalaya. Bulletin of the Geological Society ofAmerica 117, 926–944.

Maruo, Y., Kizaki, K., 1983. Thermal structure in the nappes of the eastern NepalHimalaya. In: Sharns, F.A. (Ed.), Granites of Himalaya Karakorum and Hindu-Kush. Punjab University, Lahore, Pakistan, pp. 271–286.

Mukhopadhyay, B., Holdaway, M.J., Koziol, A.M., 1997. A statistical model ofthermodynamic mixing properties of Ca–Mg–Fe2+ garnets. AmericanMineralogist 82, 165–181.

Newton, R.C., Charlu, T.V., Kleppa, O.J., 1980. Thermochemistry of the highstructural state plagioclases. Geochimica et Cosmochimica Acta 44, 933–941.

Parrish, R.R., Hodges, K.V., 1996. Isotopic constraints on the age and provenance ofthe Lesser and Greater Himalayan sequences, Nepalese Himalaya. GeologicalSociety of America Bulletin 108, 904–911.

Pecher, A., 1989. The metamorphism in the central Himalaya. Journal ofMetamorphic Geology 7, 31–41.

Pecher, A., LeFort, P., 1977. Origin and significance of the Lesser Himalayan augengneisses. In: Jest, C. (Ed.), Ecologie et geologie de l’Himalaya, vol. 268. C.N.R.S.Paris coll.int, Sci. de la Terre Ed. Cent. Natl. Rech. Sci. Paris, pp. 319–329.

Ray, S.R., 1995. Lateral variation in geometry of thrust planes and its significance, asstudied in the Shumar allochthon, Lesser Himalayas, eastern Bhutan.Tectonophysics 249, 125–139.

Ray, S.R., Bandyopadhyay, B.K., Razdan, R.K., 1989. Tectonics of a part of the ShumarAllochthon in eastern Bhutan. Tectonophysics 169, 51–58.

Richards, A., Argles, T., Harris, N., Parrish, R., Ahmad, T., Darbyshire, F., Draganits, E.,2005. Himalayan architecture constrained by isotopic tracers from clasticsediments. Earth and Planetary Science Letters 236, 773–796.

Richards, A., Parrish, R., Harris, N.B.W., Argles, T., Zhang, L., 2006. Correlation oflithotectonic units across the eastern Himalaya, Bhutan. Geology 34, 341–344.

Robinson, D.M., DeCelles, P.G., Patchett, J., Garzione, C.N., 2001. The kinematicevolution of the Nepalese Himalaya interpreted from Nd isotopes. Earth andPlanetary Science Letters 192, 507–521.

Robinson, D.M., DeCelles, P.G., Garzione, C.N., Harrison, T.M., Catlos, E.G., 2003.Kinematic model for the Main Central Thrust in Nepal. Geology 31, 359–362.

Sawyer, E.W., 1996. Melt Segregation magma flow in migmatites: implications forthe generation of granite magmas. Transactions of the Royal Society ofEdinburgh, Earth Sciences 87, 85–94.

Searle, M.P., Rex, A.J., 1989. Thermal model for the Zanskar Himalaya. Journal ofMetamorphic Geology 7, 127–134.

Searle, M.P., Waters, D.J., Dransfield, M.W., Stephenson, B.J., Walker, C., Walker, J.D.,Rex, D.C., 1999. Thermal and mechanical models for the structural evolution of

Zanskar High Himalaya. In: MacNiocaill, C., Ryan, P.D. (Eds.), ContinentalTectonics, Special Publication, vol. 164. Geological Society, London, pp. 139–156.

Searle, M.P., Law, R.D., Godin, L., Larson, K.P., Streule, M.J., Cottle, J.M., Jessup, M.J.,2008. Defining the Himalayan Main Central Thrust in Nepal. Journal of theGeological Society of London 165, 523–534.

Shaw, D.M., 1956. Geochemistry of pelitic rocks. Part 3: Major elements and generalgeochemistry. Geological Society of American Bulletin 67, 919–934.

Singh, S., Chowdhary, P.K., 1990. An outline of the geological framework of theArunachal Himalaya. Journal of Himalayan Geology 1, 189–197.

Sinha-Roy, S., 1982. Himalayan main central thrust and its implications forHimalayan inverted metamorphism. Tectonophysics 84, 197–224.

Spear, F.S., 1993. Metamorphic Phase Equilibria and Pressure–Temperature–TimePath. Mineralogical Society of America Monograph, Washington, DC.

Spear, F.S., Kohn, M.J., Cheney, J.T., 1999. P–T paths from anatectic pelites.Contributions to Mineralogy and Petrology 134, 17–32.

Stephenson, B.J., Waters, D.J., Searle, M.P., 2000. Inverted metamorphism and theMain Central Thrust: field relations and thermobarometric constraints from theKishtwar window, NW Indian Himalaya. Journal of Metamorphic Geology 18,571–590.

Thakur, V.C., 1986. Tectonic zonation and tectonic framework of Eastern Himalaya.Science de la Terre, Memoir 47, 347–366.

Vannay, J.-C., Grasemann, B., 1998. Inverted metamorphism in the high Himalaya ofHimachal Pradesh (NW India): phase equilibria versus thermobarometry.Schweizerische Mineralogische und Petrographische Mitteilungen 78,107–132.

Vannay, J.-C., Grasemann, B., 2001. Himalayan inverted metamorphism and syn-convergence extension as a consequence of a general shear extrusion.Geological Magazine 138, 253–276.

Vannay, J.C., Sharp, Z.D., Grasemann, B., 1999. Himalayan inverted metamorphismconstrained by oxygen isotope thermometry. Contributions to Mineralogy andPetrology 137, 90–101.

Verma, P.K., Tandon, S.K., 1976. Geological observations in parts of Kameng district,Arunachal Pradesh (NEFA). Himalayan Geology 6, 259–286.

Vernon, R.H., 1979. Formation of late sillimanite by hydrogen metasomatism (base-leaching) in some high grade gneisses. Lithos 12, 143–152.

Yin, A., Dubey, C.S., Kelty, T.K., Gehrels, G.E., Chou, C.Y., Grove, M., Lovera, O., 2006.Structural evolution of the Arunachal Himalaya and implications forasymmetric development of the Himalayan orogen. Current Science 90, 195–206.

Petrology of a non-classical Barrovian inverted metamorphic sequence from the western Arunachal...

Documents

Transcript of Petrology of a non-classical Barrovian inverted metamorphic sequence from the western Arunachal...