Burial of thermally perturbed Lesser Himalayan mid-crust: Evidence from petrochemistry and P–T estimation of the western Arunachal Himalaya, India

Lithos 208–209 (2014) 298–311

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Burial of thermally perturbed Lesser Himalayan mid-crust: Evidence from petrochemistry and P–T estimation of the western Arunachal Himalaya, India Sriparna Goswami-Banerjee a,b,⁎, Santanu Kumar Bhowmik a, Somnath Dasgupta c, Naresh Chandra Pant d a

Department of Geology & Geophysics, Indian Institute of Technology, Kharagpur 721302, India Institute of Geological Sciences, University of Bern, Baltzerstrasse 1 + 3, Ch-3012 Bern, Switzerland c Indian Institute of Science Education & Research, DC35/1, Sector 1 Saltlake, Kolkata 700064, India d Department of Geology, University of Delhi, Delhi 110007, India b

a r t i c l e

i n f o

Article history: Received 4 June 2014 Accepted 16 September 2014 Available online 28 September 2014 Keywords: Dual prograde P–T path Isopleth thermobarometry Arunachal Himalaya Inverted metamorphism

a b s t r a c t In this work, we establish a dual prograde P–T path of the Lesser Himalayan Sequence (LHS) rocks from the western Arunachal Himalaya (WAH). The investigated metagranites, garnet- and kyanite-zone metapelites of the LHS are part of an inverted metamorphic sequence (IMS) that is exposed on the footwall side of the Main Central Thrust (MCT). Integrated petrographic, mineral chemistry, geothermobarometric (conventional and isopleth intersection methods) and P–T pseudosection modeling studies reveal a near isobaric (at P ~ 8–9 kbar) peak Barrovian metamorphism with increase in TMax from ~560 °C in the metagranite through ~590–600 °C in the lower and middle garnet-zone to ~600–630 °C in the upper garnet- and kyanite-zone rocks. The metamorphic sequence of the LHS additionally records a pre-Barrovian near isobaric thermal gradient in the mid crust (at ~6 kbar) from ~515 °C (in the middle garnet zone) to ~560–580 °C (in the upper garnet- and kyanite zone, adjoining the Main Central Thrust). Further burial (along steep dP/dT gradient) to a uniform depth corresponding to ~ 8–9 kbar and prograde heating of the differentially heated LHS rocks led to the formation of near isobaric metamorphic field gradient in the Barrovian metamorphic zones of the WAH. A combined critical taper and channel flow model is presented to explain the inverted metamorphic zonation of the rocks of the WAH. © 2014 Published by Elsevier B.V.

1. Introduction One of the most puzzling features of the Himalayan Metamorphic Front (Goscombe et al., 2006) is the development of an inverted Barrovian metamorphic sequence. This sequence runs along and across the length of the orogen, from Nanga Parbat in the west to Arunachal in the east (Fig. 1a). The most important discontinuity surface of this sequence, the Main Central Thrust (MCT), separates the Lesser Himalayan Sequence (LHS, non-fossiliferous, low-grade metamorphic rocks of Proterozoic age) from the Greater Himalayan Sequence (GHS, Crystalline complex consisting of gneisses, migmatites and granites of Proterozoic to Ordovician age) (Gansser, 1964; Heim and Gansser, 1939; LeFort, 1975a). The origin of the inverted thermal gradient, manifested in the disposition of progressively higher-grade metamorphic zones with higher structural levels in the rocks of the Lesser and Greater Himalayan Sequences is debatable. Dasgupta et al. (2004, 2009) showed that the Barrovian type Metamorphic Field ⁎ Corresponding author at: Institute of Geological Sciences, University of Bern, Baltzerstrasse 1+3, Ch-3012 Bern, Switzerland. E-mail address: [email protected] (S. Goswami-Banerjee).

http://dx.doi.org/10.1016/j.lithos.2014.09.015 0024-4937/© 2014 Published by Elsevier B.V.

Gradient (MFG) in the Sikkim Himalaya has a positive slope, whereas in the adjoining Bhutan Himalaya (Daniel et al., 2003) and the distant Sutlej valley of western Himalaya (Vannay and Grasemann, 2001), MFG is isobaric. Although, numerous thermo-mechanical models have been proposed on the inverted metamorphic zonations (reviewed by Dasgupta et al., 2009; Hodges, 2000; Kohn, 2014), but there is no agreement on the metamorphic processes in the LHS vis-a-vis in the GHS. None of these models considered the possibility of thermal inversion of a pre-Barrovian thermally perturbed mid-crust. In the context of Arunachal Himalaya, there has been very little systematic petrological study, which could relate pre-peak elevated geothermal gradient (ca. 25–30 °C/km) in the LHS rocks to the Barrovian Inverted Metamorphic Zonation. In this study, we address this issue using lower, middle, and upper garnet-zone and kyanitezone rocks from the footwall of the MCT in western Arunachal Himalaya (WAH) (Fig. 1b–c). On the basis of detailed microstructural, textural, mineral compositional, including X-ray element imaging studies, we reconstruct the metamorphic pathways of two stages of garnet growth in these rocks. Applying isopleth and conventional geothermobarometry in combination with calculated P–T pseudosections,

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Fig. 1. (a) Simplified geological map of Arunachal and adjoining Bhutan Himalayas after Yin et al., 2010, showing the location of the present study. (b) The geological map of the western Arunachal Himalayas along the Pinjoli–Rupa–Bomdila–Dirang–Se La Pass transect (this study) and around Tawang–Lumla areas, showing the distribution of different tectono-stratigraphic units and the locations of the Main Central Thrust (MCT) (after Kumar, 1997; Goswami et al., 2009) and previous geochronology (after Yin et al., 2010). Abbreviations used: Mnz/M, U–Th–Pb monazite/40Ar–39Ar muscovite dates; BG, Bomdila Group; SH, Sub Himalaya; A, Arunachal. (c) Mineral isograd map along the traverse from Pinjoli to Se La Pass (after Goswami et al., 2009). Also shown are the locations of the investigated samples in the LHS.

we provide evidence for two stage metamorphic evolution of the studied domain: an early burial and near isobaric heating that produced differentially heated mid crust, and subsequent prograde burial of the perturbed

mid-crust along steep dP/dT gradient, producing peak metamorphism and near isobaric metamorphic field gradient. Based on these new findings, we have modeled the origin of inverted metamorphic zonation.

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2. Regional geological setting The investigated area in the WAH lies in between the Bhutan Himalaya in the west and the eastern Himalayan syntaxis in the east (Fig. 1a). This area has been previously studied by Acharyya (1987), Bhattacharjee and Nandy (2008), Das et al. (1975), Kumar (1997), Mathew et al. (2013), Singh and Chowdhary (1990), Srivastava (2013), Thakur (1986), Warren et al. (2014), Yin et al. (2006, 2010), Goswami et al. (2009) mapped four thrust-bounded stratigraphic units along a transect (Fig. 1b), which, from the lower to higher structural heights are: (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 Formation, and (d) the Se La Group. The MCT, which coincides with intense strain localization and the first appearance of kyanitegrade partial melt, is placed at the base of the Se La Group (Goswami et al., 2009). This separates the rocks of the LHS (stratigraphic units (a) to (c)) from those of the GHS (unit (d)). The different lithotectonic elements and the defined MCT in the WAH can be correlated with those in the adjoining Bhutan Himalaya (Fig. 1 in Goswami et al., 2009). Five metamorphic zones from garnet through kyanite, kyanite migmatite, kyanite–sillimanite migmatite to K-feldspar–kyanite– sillimanite migmatite were developed in the metamorphosed lowalumina pelites of Dirang and Se La Group with increasing structural heights (Fig. 1c). Three phases of deformation, D1–D2–D3, and two groups of planar structures, S1 and S2 are recognized (Goswami et al., 2009), wherein S2 is the most pervasive one. Mineral growths in all these zones are dominantly late to post-D2, excepting in some garnet-zone rocks, where syn-D1 garnet growths have been documented. Yin et al. (2010) concluded that the major contractional structures in the WAH developed at ca. 13–10 Ma, as part of a LHS thrust duplex in the sole of the MCT, ~ 5–10 Ma after the onset of the equivalent structures in the central Himalaya. Warren et al. (2014) recently obtained metamorphic ages from the hanging and footwall rocks of the Zimithung Thrust in the GHS from the western part of the present study, adjoining the Bhutan Himalaya. They showed structural juxtaposition of younger hanging wall (age ~ 16 to 13 Ma) over the older footwall (age ~ 27 to 17 Ma) rocks at ~7 Ma.

indicate a two-stage growth process (designated as: Grt1 and Grt2) (Fig. 2b-1, d-1, d-4). The early garnet is syn- to post-D1 deformation, and the late garnet is late-to post-D2 (Table 1). Mn- and Ca-images in the middle garnet zone sample AR18 show a spiral-shaped geometry of the garnet core (Fig. 2a-1 to a-2). The central spiral band is enriched in both Mn and Ca and depleted in Fe and Mg (Sps12Grs25Prp03Alm60) (Table 1; Fig. 2a-1 to a-4). The inclusion-free garnet outer rim, in contrast, shows progressive enrichment in Mg and Fe and depletion in Ca and Mn contents (Sps01Grs13Prp09Alm77) (Fig. 2a-1 to a-4). These compositional features are consistent with prograde garnet growth. In the upper garnet zone sample AR19, which records two stages of garnet growth (garnet1 and garnet 2 ) (Fig. 2b-1), the maximum spessartine (Sps 09–11 ) and grossular (Grs15–17 ) and lowest pyrope (X Prp = 0.04–0.05) contents are observed in the central axial spine of garnet 1 (Table 1; Fig. 2b-2 to b-4). There is progressive enrichment in X Prp and depletion in Sps and Grs contents along and across the spire of the snowball garnet (Fig. 2b-2 to b-4), implying garnet growth during prograde metamorphism. Some of these large garnet porphyroblasts show the development of an idioblastic garnet2 rim on the garnet1 spiral core (Fig. 2c-1 to c-2). Small idioblastic garnet grain overgrowing S2 is also growth zoned with core to rim enrichment in XPrp (Prp09 → Prp11–12) and fall in grossular (Grs12–13 → Grs07–08) contents. Spessartine remains uniformly low (Sps01) (Table 1; Fig. 2c-3). Muscovites in this sample are uniformly sodic (Pa26–29 to Pa23) relative to those from other garnet-zone rocks. The most sodic muscovites (Pa 46 Ce 10 Ms 44), now occurring as rare paragonite and muscovite intergrowths (Fig. 2c-4) along a spiral-shaped inclusion trail in garnet1 (Fig. 2c-1) stabilized in the S 1 foliation domain. Garnet 1 from the kyanite zone sample AR65 shows progressive rimward depletion in spessartine (Sps09 → 01) (Fig. 2d-1) and enrichment in pyrope (Prp07 → 14) (Fig. 2d-2) contents, consistent with prograde garnet growth as in the middle and upper garnet zone rocks. Mild growth zoning features are also preserved in elongated garnet 2 grains (Fig. 2d-3 to d-4). We now use these textural and compositional constraints to develop the pressure–temperature evolutionary history of the LHS rocks from the WAH. 4. Thermobaric evolutionary history

3. Petrochemistry 4.1. Methods of calculations We collected samples from locations (Table 1 for GPS data) at different structural heights within the garnet-zone (AR2C-lower, AR18-middle, AR19, 20-upper), and kyanite-zone (AR65) metapelitic schists along the traverse as shown in Fig. 1c. An additional sample (AR16) was collected from the upper most part of the structurally underlying Bomdila Gneiss, which is in direct contact with the lower garnet-zone rocks. In Table 1, we have presented a concise summary of mineralogy, bulk compositional characteristics, textural features, mineral chemical data, and mineral reactions (deduced from compositional and textural criteria) of the samples studied in this work. Complete mineral chemical data are summarized in the Supplementary Tables 1–4, and instrumental conditions are enlisted in the Supplementary document 1. Fig. 2 shows a compilation of X-ray images of garnet grains present in different samples. Selected compositional data are plotted on the images, which also document some of the textural features (Fig. 2b-1, c-1, c-3–c-4) mentioned in Table 1. Critical to the reconstruction of the prograde thermal history of the rocks of the LHS is the recognition of early foliation development in the rocks. Even more important is the identification of mineral assemblages, which defines this early foliation. We previously reported the basic outline of progressive fabric developments in the LHS rocks (Goswami et al., 2009). In this study, we adopted similar nomenclature as was used in our previous paper. The nomenclature is summarized in Table 1. Textural and compositional characteristics of garnet in upper garnet-zone and kyanite-zone metapelites

We have adopted two approaches to reconstruct the P–T evolutionary history of the studied sequence: (1) conventional mineralogical thermobarometry and (2) isopleth thermobarometry from the P–T pseudosections computed from the whole rock compositions, and effective bulk composition of the inferred reacting volume. We observed in the studied rocks that up to the kyanite metamorphic zone, the growth zoning in the garnet is unaffected by the diffusional modifications (Table 1, Fig. 2). This suggests that the garnet outer rim, which is compositionally Mg-rich and Mn-poor, could be paired with the matrix biotite compositions to constrain TMax. We used the rim compositions of Grt2 in the upper garnet-zone and kyanite-zone rocks, and the same of Grt1, in the others (Table 1). To estimate the metamorphic pressures at TMax, it is important to know the equilibrium composition of plagioclase. We could show in our analysis that the reverse and normal compositional zoning in the metamorphic plagioclase up to the garnet-zone metamorphism (Table 1) could be linked to metamorphic mineral assemblages, which are associated with the growth of syn-D 1 and syn- to post-D2 garnets. Given that syn-D1 garnets appeared in chlorite + epidote + plagioclase + garnet + biotite + muscovite + quartz-bearing assemblage (e.g. sample AR18), calcic rim of matrix plagioclase is inferred to have equilibrated at the peak metamorphism. For syn- to post-D2 garnets from the plagioclase + garnet + biotite + muscovite + quartz assemblage, sodic plagioclase rims which are formed in the matrix as well as

Table 1 Summary of the petrological history of the studied LHS. Met zone

Deformation

Metamorphic assemblage

Mineral chemistry

Metamorphic reaction

MG: AR16 MnO = 0.02; (Below Ca# = 42; Mg# = 27 Grt isograd) Lat: 27°19'46.4", Lon: 92°21'32" Grt Lower Grt Not available

Bulk composition

D1: S1 assemblage D2: S2 assemblage

Chl + Ms + Pl + Ep + Qtz Chl + Ms + Bt + Pl + Ep + Qtz

Bt(M): 30, 0.13; Ms(P): Pa08Cel21Ms71; Ms(M): Pa09-10Cel26Ms65 Pl(M): An27 → An21-24; Ep(I^Grt): Ps16-24

Formation of Grt: Chl + Ms + Cz component of Ep + Qtz →

Late to post-D2 porphyroblast

Grt2 (growth zoned)

Grt2: Sps07Grs20Prp03Alm70(C ) → Sps01Grs18Prp04Alm77(R)

Grt + Bt + Pl + H2O (R1)

D1: S1 assemblage

zone: AR2C Lat: 27°12'35.6", Lon: 92°24'22.2"

Pre-D2 porphyroblast

Chl + Bt + Ms + Qtz + Ilm + Pl + Gr Bt(M): 51, 0.10-0.11; Ms(I^Grt): Pa17Cel14Ms69; Ms(M): Pa22Cel12Ms65 Grt1 (strongly growth zoned) Grt1: Sps14Grs07Prp06Alm74(C) → Sps05Grs04Prp12Alm79(R)

D2: S2 assemblage

Bt + Ms + Qtz + Ilm + Grt1 + Gr

Middle Grt LAP; Al2O3 = 17.55 zone: MnO = 0.05; AR18 Ca# = 31; Mg# = 38 Lat: 27°20'26.5", Lon: 92°17'49.2"

D1: S1 assemblage Syn-D1 porphyroblast with snowball

I. Grt1-isograd reaction : Bt + Ms + Qtz + Chl + Ilm + Pl + Ep Bt(I^Grt) = 45, 0.13; Bt(M): 43, 0.11; Ms(M): Pa14-15Cel21-22Ms64; Chl + Ms + Cz component of Ep + Qtz → Grt1 (strongly growth zoned) Grt1: Sps12Grs25Prp03Alm60(C) → Sps01Grs13Prp09Alm77(R) Grt + Bt + Pl + H2O (R1)

texture (Figs.2a-1 to 2a-4) S2 assemblage

Figs. (2a-1 to 2a-4) Bt + Ms + Qtz + Pl + Ilm + Grt1

Ep: Ps16-18; Pl (M-C): An15; Pl(M-R): An22.5 Chl(M): 43-47; Bt(M): 48-49, 0.08; Bt(I^Grt) = 51, 0.12

LAP; Al2O3 = 15.76;

D1: S1 assemblage

Bt + Ms + Qtz + Chl + Pl + Ilm

zone: AR19

MnO = 0.05; Ca# = 24; Mg# = 36

Syn-D1 porphyroblast with snowball texture (Figs.2b-2, c-1); D2: S2 assemblage Post-D2 porphyroblast (Fig.2c-3)

Grt1 (strongly growth zoned)

Lat: 27°21'24", Lon: 92°14'32.8"

Chl + Pl + Ms + Qtz → Grt + Bt + H2O (R2)

I. Grt1-isograd reaction : Chl + Pl + Ms + Qtz → Grt + Bt + H2O (R2)

Grt1: Sps11Grs15-17Prp04-05Alm70(C) → Sps01Grs09-08Prp1011Alm80-81(R)

Bt + Ms + Qtz + Pl + Ilm + Grt1 Grt2 (mildly growth zoned)

Ms(I^Grt1): Pa46Cel10Ms44; Ms(M): Pa23Cel16Ms61;

II. Formation of Grt2 :

Grt2: Sps01Grs12-13Prp09Alm77-78(C ) → Sps01Grs07-08Prp11-

Bt + Pl → Grt + Ms (R3)

12Alm80-81(R)

Pl(M-C): An13; Pl(M-R N Grt2): An17: Pl(I^Grt1): An18-32 SAP; Al2O3 = 15.54; MnO = 0.05; Ca# = 40; Mg# = 43 Lat: 27°20'58", Lon: 92°15'11.9"

D1: S1 assemblage D2: S2 assemblage

Bt + Ms + Qtz + Chl + Pl ± Ep(?) Bt + Ms + Qtz + Pl + Ilm

Bt(M): 49-51, 0.13-0.16; Bt(Sym) = 45-46, 0.07-0.10; Ms: Pa07Cel23Ms70; Pl(M/I^Grt): An23-25; Pl(Sym) = An25-30

Early, Syn-D2/Late D2 porphyroblast Grt2 replaced by Bt & Bt + Pl Sym

Grt2 (strongly growth zoned)

Grt2: Sps08Grs12Prp09Alm71(C) → Sps04Grs12Prp11Alm73(R)

LAP; Al2O3 = 15.83;

D1: S1 assemblage

Bt + Ms + Qtz + Chl + Pl

Bt(M): 53, 0.11-0.15; Ms(I^Grt2): Pa23Cel19Ms58

MnO = 0.05; Ca# = 33; Mg# = 40

Early Syn-D1 porphyroblast

Grt1 (strongly growth zoned)

D2: S2 assemblage

Bt + Ms + Qtz + Pl + St + Grt1

Post-D2 porphyroblast;

Grt2 (mildly growth zoned)

St included in Ky

(Figs. 2d-3 to 2d-4) + Ky + Bt

Grt1: Sps09Grs13Prp07Alm71(C) → Sps01Grs07Prp14Alm78(R) (Figs.2d1 to 2d2); St: 21; Ms(M): Pa14-15Cel22-23Ms62-65; Ms(P): Pa1617Cel16Ms68 Grt2: Sps02Grs13Prp11Alm75(C) → Sps01Grs06-07Prp14-15Alm7879(R) Pl(I^Grt1): An24; Pl(M): An22→ 18

Upper Grt zone: AR20

Ky: AR65

Lat: 27°19.845', Lon: 92°15.672'

Formation of Grt2 : Bt + Pl → Grt + Ms (R3)

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Upper Grt

I. Grt1-isograd reaction :

I. Formation of Grt1 : Chl + Pl + Ms + Qtz → Grt + Bt + H2O (R2) II. Formation of St : Chl + Ms + Qtz → St + Bt + H2O (R4) III. Ky + Grt-isograd reaction St + Ms + Qtz → Grt2 + Ky + Bt + H2O (R5)

Grt1 & Grt2 in chronological order of development Abbreviations used: Mineral abbreviations after Kretz (1983); Met: Metamorphic; MG: Metagranite; WRC: Whole rock composition; M: Matrix; I: Inclusion; ^/N: Within/Against; C/R: Core/Rim; Sym: Symplectite;→: Core to rim compositional variation; LAP/SAP: Low alumina/sub-aluminous pelite; Ca# = 100*{Molar CaO/(molar CaO + molar Na2O)}; Mg# = 100*{Molar MgO/(Molar MgO + molar FeO)}; Al2O3 & MnO values in wt.%; Values in Chl, St: 100*XMg; Values in Bt in sequence: 100*XMg, Ti in per formula unit per 11(O); Ps value: 100*(Fe3+/Fe3++Al); Lat: Latitude; Lon: Longitude

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inclusion within garnet, appears to mark peak metamorphic compositions. The syn-D1 and post-D2 garnet (recorded in the sample AR19), the most calcic plagioclase (An32), which is preserved as inclusion within garnet1, appears to be in compositional equilibrium with the most grossular-rich composition of the spiral garnet core. By contrast, sodic plagioclase rim in the matrix is inferred to be in equilibrium with grossular-poor post-D2 garnet rim, and is taken to estimate metamorphic pressures at the thermal peak. Based on these caveats, we calculated the temperatures using the Fe–Mg exchange equilibrium between the garnet and biotite (Holdaway, 2000) at a reference pressure. We calculated the pressure at a reference temperature, applying the equilibria of the garnet–biotite–plagioclase–quartz (GBPQ) (Wu et al., 2004), the garnet–muscovite–plagioclase–quartz (GMPQ) (Wu and Zhao, 2006), and the garnet–aluminosilicate–silica–plagioclase (GASP) (Holdaway, 2001). We used the GASP equilibrium for the kyanitezone rock. The results of P–T calculations are presented in the Supplementary Table 5. The second thermodynamic approach that we adopted in this study was the calculation of the equilibrium metamorphic phase diagrams using Gibb's free energy minimization technique in P–T space (pseudosections) for fixed bulk compositions (Connolly, 1990). Assuming the bulk rock was not significantly fractionated prior to garnet growth, the P–T condition of the first appearance of garnet in the rocks of the LHS can be modeled by knowing the compositions of the earliest stabilized prograde garnet core. Intersection of the three compositional isopleths of garnet, X Prp , X Grs and X Sps that are contoured in the calculated phase diagram, uniquely constrain the P–T condition of the first appearance of garnet in the rock. In this study, we applied this isopleth thermobarometry approach (after Vance and Mahar, 1998) for garnet- and kyanite-zone rocks, where prograde garnet compositional zonation has been preserved (Fig. 2). Due to relatively low modes of garnet growth in the studied metapelites, the garnet mode and compositional isopleths of garnet (XPrp, XGrs and XSps), plagioclase (XAn) and biotite (XMg), in equilibrium at peak metamorphic condition were additionally contoured in the calculated phase diagram to retrieve the P–T conditions of the peak metamorphism. The latter is likely to complement conventional geothermobarometric data. P–T pseudosections can also be constructed with effective bulk rock composition, estimated by computation of mineral modes and mineral compositions. This was adopted in the kyanite-zone sample AR65, where the growth of the peak metamorphic assemblage of garnet2 + kyanite was largely controlled by local compositional domain. P–T pseudosections have been calculated in the simplified model system MnO–Na2 O–CaO–K2O–FeO–MgO–Al2 O3 –SiO 2–H2 O (–CO 2 ) (MnNCKFMASH–(C)) and using the PERPLEX software (Connolly, 1990) and the thermodynamic dataset of Holland and Powell (1998). The following mixing models were applied: garnet: Holland and Powell (1998); plagioclase: Newton et al. (1980); biotite: Powell and Holland (1999); chlorite: Holland and Powell (1998); chloritoid: Holland and Powell (1998); staurolite: Holland and Powell (1998); paragonite: Chatterjee and Froese (1975); phengite: following the description at http://www.esc.cam.ac.uk/astaff/holland/ds5/muscovites/ mu.html. Kyanite, sillimanite, clinozoisite and quartz were considered as pure phases. The metamorphic phase relations for quartz and water saturated conditions for the different metamorphic zones are presented in Fig. 3. 4.2. P–T conditions of peak metamorphism from conventional geothermobarometry The results of the P–T conditions of peak metamorphism from conventional geothermobarometry are presented in the Supplementary Table 5. These results show a consistent pattern of increasing peak metamorphic temperatures with structural height. For metagranite

and garnet–biotite thermometry, temperature ranges from 534° to 584 °C with an average of 560 ± 50 °C (error in 2 sigma). For garnetzone rocks, there is a steady increase in the average peak metamorphic temperature from 590 ± 20/40 °C at the lower (Sample AR2C) and middle (AR18) parts of the garnet-zone through 600 ± 20 °C (sample AR19) to 615 ± 20 °C (AR20) at the uppermost part of the same metamorphic zone. For the kyanite-zone rock (sample AR65), the average peak temperature is constrained as 615 ± 20 °C, which is same as what was retrieved from the upper garnet-zone sample, AR20. The differential appearance of kyanite in these two samples under the same thermal and baric conditions (described later) and despite their close spatial association is related to the possible variation in the bulk rock compositions. Presence of Fe3+ in garnet and biotite (calculated after Holdaway, 2000) reduces temperature estimates by an average of ~20 to 35 °C in all these metamorphic zones (Supplementary Table 5). Irrespective of the calibrations, all geobarometers yielded nearly identical pressure ranges for the peak metamorphism for the different metamorphic zones (metagranite: 8.9 ± 1.8 kbar, error in 2 sigma; garnet zone: 8.4 ± 0.4 kbar to 9.0 ± 1.2 kbar, kyanite-zone: 8.3 ± 0.2 kbar) (Supplementary Table 5). 4.3. P–T conditions of prograde and peak metamorphism from isopleth thermobarometry The computed pseudosections and the results of isopleth thermobarometry for the three garnet-zones (samples: AR18, AR19 and AR20) and one kyanite-zone (sample: AR65) are presented in Fig. 3a–d. Of the three compositional isopleths in garnet, i.e., XPrp , X Grs, and XSps, the first two are strongly T-sensitive. The slope of the XSps isopleth, which parallels with that of the garnet-in curve is variable and moderately pressure sensitive. To make estimations on the prograde P–T conditions in the first two samples of the garnet-zone rocks, the compositions of the cores of garnet1 have been considered. For the subaluminous metapelite sample AR20, the core composition of the earliest formed garnet was taken. The isopleths intersect each other either at a point, or produce a small triangle-shaped space, limiting the P–T conditions of garnet growth. The intersection plots in the field of chlorite-bearing, muscovite + biotite + plagioclase + quartz + garnet + vapor for the two metapelite samples, and in the chlorite-absent, muscovite + biotite + plagioclase + quartz + garnet + vapor field for the subaluminous pelite sample are consistent with the textural observations (Table 1). For the three garnet-zone samples, there is consistent increase in temperature from 515 °C (AR18) through 545 °C (AR19) to 580 °C (AR20) at broadly near constant pressure condition (P ~ 6 to 6.2 kbar). Although samples AR19 and AR20 are from the same structural height, the higher temperatures in the sample AR20 warrant additional comments. We argue that, there could be two possibilities for the higher temperatures in AR20: (1) AR20 represents a thin tectonic slice of the slightly higher grade metamorphic rocks, or, (2) the contrast is due to the differences in bulk rock composition. We consider that relatively low MnO content and the high Mg# of the bulk rock inhibited garnet growth in AR20 till higher metamorphic temperature was attained. For the kyanite zone low-alumina metapelite sample (AR65), the isopleth intersection in the chlorite assemblage field yielded prograde P–T condition of ~ 5.9 kbar, 560 °C. The remarkable constancy in metamorphic pressure estimate of early garnet formation in all the metamorphic zones is noteworthy. Interestingly, the garnet-in isograds in Fig. 3a–d occurred at lower temperatures than those obtained by isopleth thermometry for the early studied garnets. This may imply that the preserved compositions are not those of the garnet isograds. Perfect preservation of growth zoning in the studied garnet grains rules out any relaxation. Further, for grains with 1–3 mm diameter in the studied samples, such diffusional modification of core compositions is likely to be small for peak

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Fig. 2. X-ray element and back-scattered electron (BSE) images and thin-section sketch of representative garnet-zone [AR18, (a-1 to a-4); AR19, (b-1 to b-4 and c-1 to c-4)] and kyanitezone [AR65, (d-1 to d-4)] samples. (a-1 to a-4) Mn-, Ca-, Mg- and Fe-X-ray element images of syn-D1 garnet1 in AR18. Spot values in these and other images mark corresponding mole fractions. (b-1) Schematic thin section sketch in AR19 showing two generations of fabric developments (S1 and S2) and garnet growths (garnet1 and garnet2). Boxes mark the locations of representative syn-D1 garnet1 (Gr-A and Gr-B) and post-D2 garnet2 (Gr-D) grains. (b-2 to b-4) Mn-, Ca- and Mg-X-ray element images of grain B. (c-1 to c-2) BSE and Mn-X-ray element images of grain A, showing sodic-muscovite-defined (boxes marking mica locations) spiral S1 inclusion trail. (c-3) False color image of mildly growth-zoned, garnet2 grain D, overgrowing F2 crenulation. (c-4) Na-X-ray element image of one representative S1 foliation-defined sodic muscovite inclusion within garnet1. The inclusion now occurs as intergrowths of paragonite and muscovite. The spot values mark XPa. (d-1 to d-4) Mn- and Mg-X-ray element images of garnet1 (d-1 to d-2) and garnet2 (d-3 to d-4) in kyanite-zone sample AR65.

temperatures of ≤ 600–625 °C (Chakraborty and Ganguly, 1991). Even if there were some modifications, it would move the intersection of compositional isopleths corresponding to garnet core toward higher temperature and pressure than experienced during initial

garnet growth. Alternatively, the nucleation and growth of garnet could be delayed till significant overstepping of metamorphic temperatures of the garnet-in reactions took place (e.g. Zeh and Holness, 2003). This remains a viable option for the studied samples

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as small overstepping of both pressures and temperatures along a prograde P–T path could have produced the first garnets with a different composition than what was modeled for the garnet-in curve. Nevertheless, the P–T conditions of metamorphism obtained by isopleth thermobarometry provided robust constraints on the prograde segment of the P–T path of evolution in the garnet- and kyanite-metamorphic zones. Compositional parameters of the rims of Grt2 (XPrp, XGrs), biotite XMg (Bt), and the garnet mode isopleths yielded narrow P–T fields of intersection, the central point of which yielded peak P–T estimates in the range of 9 kbar, 615 °C (AR19) and 9.2 kbar, 635 °C (AR20), both belonging to the upper garnet-zone (Fig. 3f–g). For the kyanite-zone sample AR65, the P–T stability of the peak metamorphic assemblage of quartz + biotite + muscovite + garnet2 + kyanite + staurolite + plagioclase + water vapor has been shown by the shaded field in the P–T pseudosection, which was calculated with the effective bulk rock composition (Fig. 3h). The measured Grt2 mode as well as the peak P–T estimate through conventional geothermobarometry lies well within the shaded field. The outermost rim composition of Grt1 in the middle garnet-zone sample AR18 gives peak P–T estimate of ~8.2 kbar, 600 °C (Fig. 3e). These are in perfect agreement with those calculated by conventional thermobarometry (Supplementary Table 5). The deviation of XAn isopleths from the narrow fields of intersections (Fig. 3a–d) may be related to partial bulk rock fractionation in relation to Ca during prograde garnet growth. 4.4. Metamorphic P–T path 4.4.1. Garnet zone In Fig. 3e–g, the different mineral assemblage fields for the three garnet-zone samples (AR18, AR19, AR20) are isoplethed with garnet mode, XGrs, XSps, XPrp (garnet), XAn (plagioclase), XMg–Cel, XPa (muscovite). Although there are variations in the positions of garnet-in, clinozoisite-out, chlorite-out curves in these samples due to the differences in bulk rock compositions (see Table 1), the slopes of the isopleths in P–T space are similar. Contours of garnet mode have moderate to steep negative slopes in assemblages with or without chlorite, implying considerable garnet growths via the R1–R3 reactions (Table 1) with heating or compression, or both. The XSps isopleths mirror the garnet mode isopleths, which can be attributed to the large Kd (Mn) for garnet relative to other metamorphic phases. The pseudosections predict increasing XPrp and decreasing XGrs contents in garnet with progressive heating. XAn (in plagioclase) isopleths have moderate positive slopes at lower temperatures (T b 500 °C). Plagioclase becomes calcic with increasing temperature up to 550 °C. Because of a change in slope of X An isopleths beyond this temperature, plagioclase becomes increasing sodic with pressure. One important prediction of the calculated P–T pseudosection is that for relatively sodic bulk rock compositions (e.g. sample AR19) extremely sodic muscovite in association with chlorite + biotite + plagioclase + quartz may become stable at midcrust and at elevated temperatures of 525–550 °C. Mg–celadonite isopleths, in contrast, are pressure-sensitive, bearing linear relationship with increasing pressure. When evaluated with the calculated P–T pseudosection, the textural features in AR18, such as an early stability of epidote + chlorite + two mica-bearing assemblage pre-dating garnet growth and the final growth of garnet in the epidote + chlorite-absent assemblage field, indicate a prograde P–T segment intersecting garnet-in, clinozoisite-out and chlorite-out curves in sequence (Fig. 3e). In combination with the compositional features of decreasing spessartine and grossular, and increasing pyrope contents in the growth zoned garnet and reverse zoning in matrix plagioclase, the metamorphic reaction history provided a rather tight constraint on the prograde P–T path. Two segments of the prograde P–T path were distinguished: (1) an early stage with a relatively shallow positive dP/dT gradient (Path 1), and (2) a later segment, which has a moderate to steep positive dP/dT gradient (Path 2). An early

mid-crustal heating (Path 1) is also inferred by the stability of sodic muscovite in the S1 foliation pre-dating the appearance of Grt1 in AR19 (Fig. 2c-1, 2c-3). Compositional features of decreasing spessartine and grossular and increasing pyrope contents in the growth zoned garnet, normal zoning in matrix plagioclase, increasing celadonite contents in later stabilized muscovite and garnet mode isopleths in samples AR19 and AR20 also indicated near identical steep dP/dT gradient of the later stage of the prograde P–T path (Path 2) as in AR18.

4.4.2. Kyanite zone Metamorphic phase relations in the kyanite-zone pelite AR65 (Fig. 3h) are broadly similar to that observed in the garnet-zone rocks. The initial phase of garnet growth (Grt1) occurred in the chlorite + muscovite + biotite + plagioclase + garnet assemblage field (Fig. 3d) at P–T conditions of ~ 5.9 kbar, ~ 560 °C. The stability of the relatively sodic muscovite, predating Grt 1 is similar to that observed in AR19, and may be taken as an evidence for possible existence of hotter mid-crust for kyanite-zone rock. As previously mentioned, the peak garnet2 + kyanite assemblage appeared in more aluminous bulk rock composition. Consequently, the peak kyanite-zone metamorphism was evaluated with a P–T pseudosection (Fig. 3h), which was constructed with an effective bulk rock composition. The effective bulk composition was computed with mineral modes and compositions from a microdomain of 1.5 mm × 4.4 mm size and containing staurolite (~ 0.18%), kyanite (~ 9.29%), garnet2 (~ 11.29%), muscovite (~ 19.85%), biotite (~ 14.94%), plagioclase (~ 2%) and quartz (40.44%). Assuming the staurolite + kyanite + garnet2 + muscovite + biotite + plagioclase + quartz as the final stabilized assemblage, the pseudosection predicts successive growths of staurolite and kyanite via classical staurolite and kyanite isograd reactions, respectively (R4 and R5, Table 1). When evaluated with the calculated phase relations, the reconstructed metamorphic sequence of the growth of garnet1 (in the chlorite-assemblage field), the appearance of staurolite, complete disappearance of chlorite and the formation of garnet 2 + kyanite, confirmed the steep dP/dT gradient of the path 2 (Fig. 3h).

5. Discussion 5.1. Thermally perturbed mid-crust pre-dating peak metamorphism The integration of prograde and peak P–T conditions of metamorphism reveals the existence of a pre-peak, isobaric (at ~ 6 kbar) thermal gradient in the studied rocks of the LHS (Path 1, Figs. 3e–g and 4a). The unusual sodic compositions of S1 white micas lend further support to the development of unusually hot upper-to mid-crust beneath the WAH, and predating the peak metamorphism. The pseudosection modeling shows that such white micas containing 25 to 40 mol% paragonite in solid solution (Table 1) would require temperatures in the range of 525–575 °C (Fig. 3f). The thermal conditions of the first-formed garnets from both the garnet- and kyanite-metamorphic zones are found to be hotter (thermal gradient of c. 24–27 °C/km) relative to the normal geothermal gradient (Fig. 4a). This reflects the presence of a thermally perturbed midcrust. The hottest rocks appeared from the upper garnet zone (~ 580 °C, sample AR20) and the kyanite-zone (T ~ 560 °C, Sample AR65), whereas structurally lower garnet-zone rocks were from relatively cooler (T ~ 515 °C, sample AR18) zones. Rocks from all these zones uniformly show subsequent prograde burial along steep dP/dT gradients (Path 2, Fig. 4a). Near identical metamorphic P–T history of prograde burial of thermally perturbed mid-crust was previously established by Vance and Mahar (1998) from the Zanskar Himalayas, but from the rocks of the GHS.

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5.2. Summary of the LHS P–T paths In Fig. 4, the P–T path information from different areas of the LHS is summarized. Stephenson et al. (2000) deduced a hairpin P–T loop from the Kishtwar window (Path 1, Fig. 4b) and argued for a rapid burial and then cooling event. From the staurolite-zone sample of Sutlej section,

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Caddick et al. (2007) inferred an early burial to 7.5 kbar pressure, near-isobaric prograde heating to TMax of ~ 640 °C and finally decompression accompanying cooling P–T path (Path 2, Fig. 4b). Such tight P–T path closure around the TMax in the rocks of the LHS is generally interpreted in terms of footwall heating due to rapid overthrusting of the hot GHS block, followed by tectonically-driven exhumation

Fig. 3. MnNCKFMASH P–T pseudosections of the investigated garnet-zone [AR18 (a, e), AR19 (b, f) and AR20 (c, g)] and kyanite-zone [AR65 (d, h)] samples calculated with whole rock (a–g) and effective bulk rock (h) compositions. The pseudosections, isoplethed with garnet modes and garnet, plagioclase, biotite and muscovite compositional parameters enable estimations of prograde and peak P–T conditions of metamorphism through isopleth intersection method (a–d) and reconstructions of two stage P–T paths (e–h) (see text for detail). Also plotted are peak P–T conditions (error bar 2 sigma), calculated with conventional geothermobarometry.

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Fig. 3 (continued).

(Caddick et al., 2007). Monazite geochronology from the investigated staurolite-zone sample yielded 10–11 Ma for the prograde metamorphism. This was taken as evidence by the authors to distinguish metamorphism of the LHS from that typically predicted for the GHS. From central Nepal, Kohn et al. (2001) calculated different P–T paths from lower to upper part of garnet-zone rocks. The garnets from lower part (Path 3b, Fig. 4b) grew with increasing P and T, whereas the P–T path from mid to upper part of garnet zone demonstrates garnet growth with increasing T but with decreasing P. Monazite ion-microprobe Th–Pb ages yielded 8–9 Ma for structurally lower rocks, while 10–22 Ma for the structurally higher rocks (Kohn et al., 2001). The authors interpreted the metamorphic and geochronologic data in terms of footwall metamorphism in part due to thrust reactivation (ca. 8 Ma) and also thermal relaxation following older (≥20 Ma) thrust movement. From eastern Nepal, Imayama et al. (2010) reported nearly isothermal loading path [dP/dT slope in the range from ~ 3.0 kbar per 10 °C (Gibbs method) to ~2.5 kbar per 75 °C (pseudosection method)] from the staurolite zone rock, in the footwall of the MCT (Paths 4a and 4b respectively in Fig. 4b). This was taken as evidence for very rapid underthrusting of the LHS beneath the GHS. Daniel et al. (2003) examined the rocks from the eastern Bhutan sector, adjoining the present study area in WAH. Although the prograde segments of the P–T paths could not be constructed because of the lack of well-constrained prograde inclusions within garnet, the peak P–T conditions of garnet-zone rocks (11–12 kbar, ~650 °C) as well as the exhumation path of these rocks (Path 5, Fig. 4b after Davidson et al., 1997) were discussed by these authors. For the uppermost LHS rocks, a CW P–T path involving early heating and compression was

inferred based upon the presence of high-Ca rims, syn- to intertectonic cores of growth-zoned garnet (Daniel et al., 2003). A relatively steep decompressive P–T trajectory of the Bhutan garnet-zone rocks indicates rapid exhumation from lower crustal depths. Metamorphic monazites from the rocks of the MCTZ rocks yielded U–Pb ages, which range from 22 to 20–18 Ma (Daniel et al., 2003). The results were interpreted as the timing of prograde metamorphism and deformation within the MCTZ. Summarizing, despite some differences, the LHS metamorphism, in a general way reflects limited heating during prograde burial, relatively cooler peak metamorphism, in the stability field of kyanite and lack of protracted heating following TMax (Kohn, 2014). All these features are recorded by path 2 from the Arunachal Himalaya. The Arunachal LHS rocks are, however, unique in their record of a pre-thickening, relatively hot mid-crust (Fig. 4a). 5.3. Arunachal LHS P–T path and tectonic implication Any model that deals with the origin of the inverted metamorphic sequence (IMS) in the WAH should explain the following key features: 1) The Barrovian metamorphic sequence of the LHS recording two phases of prograde heating. 2) An early episode of heating in the mid-crust (~ 6 kbar), which predates peak metamorphism. 3) Differential heating of the rocks of the LHS. The rocks adjoining the heat source (now kyanite-zone rocks, sample AR65) were heated to a maximum extent, while garnet zone rocks, located furthest

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from the heat source (e.g. sample AR18) were heated to a lesser extent. 4) Steep burial and prograde heating of the differentially heated midcrust of the LHS to a uniform depth of ~8–9 kbar, such that garnetand kyanite-zone rocks record a broadly isobaric MFG. 5) Classical clockwise (CW) P–T path from the deepest section of the Arunachal GHS with the following sequence of metamorphic stages: burial to lower crustal depths, prograde heating leading to kyanite-facies partial melting at T ~ 750–800 °C and steep, near isothermal decompression to mid-crustal depths and subsequent post-decompressional cooling (Fig. 5a after Goswami, 2010 and Goswami-Banerjee et al. in preparation). Near identical GHS metamorphic P–T paths were previously established from the adjoining Bhutan (Daniel et al., 2003) and Sikkim Himalayas (Ganguly et al., 2000; Harris et al., 2004; Neogi et al., 1998; Sorcar et al., 2014). 5.3.1. Inverted metamorphic zonation and comparison of models Numerous models have been proposed over the years to explain the inversion of the thermal gradient in the Himalayas (See Dasgupta et al., 2009 for a summary). Investigation of these models indicates,

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while there are similarities with some aspects of the metamorphism in the WAH, these models are not sufficient to explain all the features, (1) to (5). The discrepancies with the existing models are briefly summarized below. In the Hot Iron Model of Frank et al. (1973) and LeFort (1975a), the thermal inversion of the metamorphic sequence occurs due to the emplacement of a hot thrust sheet on the top of colder footwall rocks, along the MCT. Since in this model, heating takes place from the top to the bottom, a metamorphic field gradient of increasing T in the direction of decreasing P is expected, and that the footwall rock should show a near isobaric heating P–T trajectory. While the reconstructed inverted metamorphic sequence in the WAH does indicate the presence of an initial heat source and a conductive near isobaric heating of the footwall rocks (points 2–3), it is different from the Hot Iron Model by the following two aspects: (1) The MFG in the studied IMS in the WAH is near isobaric (point 4) and not with a negative slope in the P–T space. (2) The rocks of the WAH record a two-stage prograde heating (point 1), with the second heating being synchronous with burial along a steep dP/dT path (point 4). Tectonic inversion models explain inversion of “right side down metamorphic isograds” by recumbent folding (Searle and Rex, 1989),

Fig. 4. (a) The metamorphic P–T condition of metagranite (sample MG) and P–T paths of Grt- (samples AR18, AR19) and Ky-(sample AR65) zone rocks from the western Arunachal Himalaya. SSG1 and SSG2 are limiting steady state geotherms calculated with the following assumptions: 2 layer crustal model of Jamieson et al. (2004) (SSG1) and exponential decrease of radiogenic heat production in the crust (SSG2). (b) Representative metamorphic P–T paths of the LHS rocks from western, central and eastern Himalayas.

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thrusting (Grujic et al., 1996; Hubbard, 1996; Jain and Manickavasagam, 1993) and by other mechanical processes. One of the important consequences of the tectonic inversion model is that the MFG would have a positive slope, which is lacking in the rock record of the WAH (point 4). In a modified model of tectonic imbrication by thrusting in a fold–thrust system, foreland propagating thrust systems produce progressively younger and lower grade metamorphic assemblages at the base of the structural sequence (e.g. Robinson et al., 2003; Webb et al., 2013). The model predicts the sole of the propagating thrust system as the locale of progressively lower grade of metamorphism. Since, in the fold–thrust system, the entire block moves to shallower depth as one unit, the model predicts gradients both in metamorphic age as well as in P and T across the LHS. Since the peak P–T conditions are expected to decline at successive stages of metamorphism, the MFG of the LHS is expected to have a positive slope, but with distinct directions of age progression. Thus, in the Annapurna region of central Nepal, progressively younger ages from 22 Ma to 3.3 Ma were recorded with decreasing structural levels (Catlos et al., 2001, 2004). From Langtang and Darondi regions, central Nepal, Kohn (2008) obtained key metamorphic and geochronological data in support of foreland propagating thrust system. The author proposed a critical taper model for the origin of IMS. The model invokes the formation of a wedge of deforming rock, which is at a condition of critical failure everywhere. The wedge maintains its size and equilibrium by combined processes of erosion at top and accretion or underplating of new material at its base. The proposed model adequately explains several key features of the IMS from different sectors of the Himalayan metamorphic front, namely, (a) continuous variation of ΔT (Fig. 5b) and Δ P (Fig. 5c) and near isobaric metamorphic field gradient across the LHS and the GHS, (b) hairpin P–T loops of the rocks of the LHS (Fig. 5a), (c) progressively younging peak metamorphic ages down the structural sequence (Fig. 5d) and (d) the development of Lesser Himalayan duplex. Recent documentation of a LHS duplex from the WAH (Yin et al., 2010) and records of continuous variation of T and P across the MCT, near isobaric MFG and the steep dP/dT of the prograde P–T path of the peak Barrovian metamorphism (this study) may be consistent with the critical taper model. The model, however, does not predict a pre-Barrovian, near isobaric, mid-crustal heating event (points 2–3). Combined thermo-mechanical models, which invoke inversion of the right side metamorphic sequence due to syn- to late-metamorphic ductile extrusion of the partially melted GHS (Beaumont et al., 2001; Daniel et al., 2003; Dasgupta et al., 2009; Faccenda et al., 2008; Harris, 2007; Jamieson et al., 1996, 2002, 2004; Vannay and Grasemann, 2001) have been widely used to explain the origin of the IMS. The basic points of this model, also referred to as the Channel flow model state that the rocks of the GHS are part of a low-viscosity, crustal channel that was produced in the lower crust. The ductile channel is bounded by two time synchronous tectonic discontinuity surfaces, the STDS above and the MCT below. While the movement along the STDS has a normal sense, that along the MCT has a reverse sense. Because of lateral pressure gradient, the low-viscosity, high-T ductile channel is extruded to the south, producing a zone of inverted, telescoped metamorphic isograds and isotherms along the basal contact of the channel. One of the critical issues of the Channel flow model is the knowledge on the source of heat, which is required to raise the temperatures in excess of 700–750 °C so that metasediments of the GHS can undergo partial melting. Based on numerical calculations, Faccenda et al. (2008) have established that the collective heat budget produced by thrusting, shear heating and heat generation from the burial of highly radiogenic sediments is sufficient enough to trigger melting in the rocks of the GHS. Their model also predicts the partial meltinginduced reduction of density and viscosity of a buried deep crust, which acts as a driving force for uplift and exhumation. Moreover,

~20–30 Ma of timescales of burial, heating/metamorphism and subsequent exhumation, predicted in the different simulations are well within the limit of measured timescales of the Himalayan orogeny (e.g. Rubatto et al., 2013). In a general way, the channel flow models predict (a) very sharp variations in T and P across the MCT and near continuous decrease in T and P with increasing structural heights in the GHS (Fig. 5b–c), (b) a positive metamorphic field gradient of the metamorphic pile across the LHS and GHS, (c) isobaric heating P–T path for the LHS (Fig. 5a) and HT isothermal decompression path for the GHS and (d) similar timescales of the cooling history of the GHS and LHS rocks (Fig. 5e) with limited or no time gap between peak metamorphism in the LHS footwall and melt crystallization in the GHS hanging wall. Seen in this context, the pre-Barrovian near isobaric heating P–T path (Point 2) when evaluated with model LHS P–T paths in Fig. 5b, shows some similarity with channel flow model LHS path 5b (Path L4 in HT1 model of Jamieson et al., 2004). Although not of the same scale the latter path also predicts loading following near isobaric heating (Fig. 5a) as observed in the WAH. The P–T path of the deepest section of the GHS from WAH (Path 3 in Fig. 5a; see Fig. 1c for location of the representative GHS sample AR37) also resembles channel flow model P–T path. 5.3.2. Origin of the IMS in WAH A review of these thermal models indicates that the channel flow and critical taper models partially explain some aspects of tectonic processes that gave rise to the composite LHS P–T path. A hybrid thermomechanical model that combines mechanisms of both channel flow and critical taper models is proposed to explain the changing styles of tectonic processes during the India–Asia collision, implicit in the Arunachal P–T path. We suggest that processes akin to channel flow could have been active at an early stage, when the hanging wall mid-crustal migmatites, immediately after isothermal decompression were still in supra-solidus conditions (T ≥ 750–800 °C). Thermomechanical model of India–Asia collision by Faccenda et al. (2008) and numerical calculations of exhumation velocity of the Sikkim granulites along isothermal decompression and post-decompression cooling paths (Ganguly et al., 2000) both predict dominant horizontal motion of these mid-crustal migmatites. One possible consequences of this motion is the large-scale lateral heat transport. Short-lived conductive heat transfer from the hanging wall hot GHS to the partially buried footwall LHS rocks across a tectonic discontinuity surface, akin to a proto-MCT could have led to near isobaric heating of the LHS rocks (Path 1 in Fig. 5a). Given the near-isobaric thermal gradient, rocks proximal to the heat source (present kyanite- and upper garnet-zone rocks) underwent maximum heating. In an alternate model, midcrustal heating of the LHS crust, in the absence of a channel flow, could still have ensued due to thermal re-equilibration from an earlier thrusting event. At a later stage, as the hanging wall GHS migmatites cooled below the solidus, the heat source was switched off, leading to the deactivation of the channel flow and shifting of the convergent tectonics to the foreland side of the collisional orogen. Continued thrust tectonics along the MCT led to rapid burial of footwall LHS rocks below the cooler hanging wall GHS rocks. In the absence of extraneous heat sources, thrust loading of the LHS rocks, also accentuated by in sequence thrusting within the LHS, as manifest by the LHS duplex in the WAH (Yin et al., 2010) can explain, following the critical taper model of Kohn (2008)(a) much cooler peak metamorphic temperature of the LHS lower crust, (b) prograde burial of the LHS mid-crust along a steep dP/dT gradient (path 2, Fig. 5a) and the inversion of metamorphic isograds in WAH. These conclusions are in good agreement with microstructural information that syn-D1 to post-D2 garnet growth during prograde burial occurred just before and during movement on the MCT. Although meager, available geochronology data (after Yin et al., 2010) from the Bomdila–Dirang–Se La Pass (this study) and Kimin–Geevan traverses from the adjoining eastern Arunachal Himalaya

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Fig. 5. (a) P–T paths of the LHS (paths 1 and 2, this study) and the deepest section of the GHS (path 3, after Goswami, 2010; Goswami-Banerjee et al. in preparation) rocks from the WAH. Paths 1 and 2 are compared with LHS P–T paths predicted by critical taper (Path 4, after Kohn, 2008 and abbreviated as K-08) and channel flow (Paths 5a and 5b, after model HT1 of Jamieson et al., 2004, abbreviated as J-04) mechanisms. Circular points on channel flow P–T paths mark reconstructed geological timescales (in Ma) from model predictions with the assumption that the orogenesis started at 54 Ma. The timing of peak metamorphism in the garnet-zone rock from the WAH is after Yin et al. (2010). (b–c) Variations of TMax (b) and P at TMax (c) with structural distance from the MCT for the rocks of the WAH and critical taper and channel flow predictions. Errors in P–T estimates for peak LHS (this study) and GHS (Goswami et al., 2009; Goswami, 2010; Goswami-Banerjee in preparation) rocks are 2σ. (d–e) Model T–t paths of LHS and GHS rocks are after critical taper (Fig. 5d, Kohn, 2008) and channel flow (Fig. 5e, Jamieson et al., 2004) mechanisms. Also plotted are the published geochronology data of the LHS and GHS rocks (after Yin et al., 2010) and T range of peak metamorphism in the GHS rocks (after Goswami, 2010; Goswami-Banerjee et al. in preparation), all from the WAH.

(Fig. 6, Yin et al., 2010 for location) show a significant time lag between melt crystallization in the hanging wall (U–Pb zircon age of leucogranite at 22 Ma and 40Ar–39Ar muscovite age at 12.2 Ma from the deepest section of the GHS migmatite from the Se La pass area) and peak metamorphism of the footwall LHS rocks (Th–Pb ion microprobe monazite age at 10 Ma) (Fig. 5e), lending further credence to the proposed hybrid model. We must admit that this model of alternating channel flow and critical taper mechanics for the WAH is not a key deduction from our work, but is built upon based on published natural rock data (Chambers et al., 2011; Kellet et al., 2009). However, detailed geochronological investigations of the two metamorphic events, in support of channel flow and critical taper mechanics are required to further strengthen the proposed model. Supplementary data to this article can be found online at http://dx. doi.org/10.1016/j.lithos.2014.09.015.

Acknowledgments This is part of the PhD dissertation work of SGB. SGB and SKB acknowledge research grant from the Department of Science and Technology, Government of India (Grant No. ESS/16/242/2005/ Kameng/01) and support from IIT, Kharagpur. S.D. also acknowledges financial support received through J.C. Bose Fellowship. Constructive reviews by Takeshi Imayama and Alex Webb helped to improve the manuscript. The authors thank Sumit Chakraborty and James A.D. Connolly for their helpful suggestions on an earlier draft.

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