Monazite geochronology unravels the timing of crustal thickening in NW Himalaya

Lithos 210–211 (2014) 111–128

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Monazite geochronology unravels the timing of crustal thickening in NW Himalaya Konstanze Stübner a,⁎, Djordje Grujic b, Randall R. Parrish c,d, Nick M.W. Roberts c, Andreas Kronz a, Joe Wooden e, Talat Ahmad f a

Geowissenschaftliches Zentrum, Universität Göttingen, 37077 Göttingen, Germany Department of Earth Sciences, Dalhousie University, Halifax, NS B3H 4R2, Canada NERC Isotope Geosciences Laboratory, British Geological Survey, Keyworth, Nottingham NG12 5GG, UK d Department of Geology, University of Leicester, LE1 7RH, UK e Department of Geological and Environmental Science, Stanford University, CA 94305, USA f University of Kashmir, Hazratbal, Srinagar, 190 006 Jammu and Kashmir, India b c

a r t i c l e

i n f o

Article history: Received 19 March 2014 Accepted 21 September 2014 Available online 2 October 2014 Keywords: Himachal Pradesh Haimanta Monazite Geochronology Crustal thickening Eohimalayan

a b s t r a c t Greenschist to amphibolite grade Haimanta metasediments of the NW Himalaya preserve much of the prograde metamorphic history of Eohimalayan crustal thickening, which has been erased by Oligo-/Miocene migmatization elsewhere in the Himalaya. Our zircon and monazite U/Th–Pb data unravel a multi-stage prograde metamorphic evolution. The earliest evidence of prograde Barrovian metamorphic monazite growth is ~41 Ma. Peak metamorphic conditions (~8–8.5 kbar, ~600–700 °C) were attained at 37–36 Ma and followed by a prolonged evolution at high temperatures with at least three distinct episodes of monazite growth, which may be related to the formation of the northern Himalayan nappes (e.g., Shikar Beh nappe, Nyimaling nappe). Rapid exhumation of the crystalline started at ~26 Ma and resulted in cooling through the muscovite 40Ar/39Ar closure temperature by 21.8 Ma. Although a local continuation of the South Tibetan detachment is not unambiguously identified in central Himachal Pradesh extrusion was likely facilitated by a system of several minor late Oligocene/early Miocene top-to-the-N to NE shear zones. In contrast to the crystalline of Zanskar and eastern Himachal Pradesh, extrusion was not accompanied by widespread decompression melting. © 2014 Elsevier B.V. All rights reserved.

1. Introduction Exhumation of the metamorphic core of the Himalaya has been the focus of many studies employing petrological, chronological and structural methods, and numerical simulations (e.g., Beaumont et al., 2001; Law et al., 2006; Yin, 2006). It is generally accepted that the Greater Himalayan Crystalline (GHC) exhumed during the early and middle Miocene by extrusion accomplished by contemporaneous southdirected thrusting on the Main Central Thrust (MCT) at its base and top-to-the-N normal shear along the South Tibetan detachment (STD) at its top (e.g., Hodges et al., 1992). The mechanism of extrusion (e.g., Channel Flow, Critical Taper, General Shear) is a matter of ongoing debate (e.g., Kohn, 2008; Vannay and Grasemann, 2001; Webb et al., 2011a). We share the opinion (e.g., Jamieson and Beaumont, 2013 and references therein) that the Himalayan orogeny has evolved by different tectonic modes at different times and different crustal levels according to the rheological properties of the rocks. Our current ⁎ Corresponding author at: Department of Geosciences, University of Tübingen, 72074 Tübingen, Germany. Tel.: +49 7071 29 73151. E-mail address: [email protected] (K. Stübner).

http://dx.doi.org/10.1016/j.lithos.2014.09.024 0024-4937/© 2014 Elsevier B.V. All rights reserved.

understanding of the Himalayan tectonics, in particular of its metamorphic core, the GHC, is based on observations of deformation under peak temperature conditions followed by nearly isothermal decompression during the early and middle Miocene (e.g., Groppo et al., 2010; Mottram et al., 2014; Rubatto et al., 2013). Along most of the Himalaya, the GHC rocks experienced amphibolite to granulite facies metamorphism and intense in-situ partial melting and magmatism, which had fundamental impact of the rheological structure of the crust (Beaumont et al., 2001). The pervasive ductile deformation at high metamorphic conditions has erased the evidence of the prograde metamorphic and tectonic stage (e.g., Groppo et al., 2010; Rubatto et al., 2013). Therefore, the early Himalayan evolution of crustal shortening and thickening following collision at ~ 55 Ma, the Eohimalayan stage, is less well understood (Aikman et al., 2008; Guillot et al., 1999), and numerical and conceptual models of Himalayan tectonics make undocumented assumptions about the prograde metamorphism and associated tectonics of the first 30 Myr of Himalayan orogenesis. The lower-grade metamorphic conditions of the GHC in the central Himachal Himalaya make this an ideal area to study Eohimalayan tectonics. In fact, our current understanding of the Eohimalayan history is largely based on evidence from the western Himalaya (e.g., Guillot

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et al., 1999; Vance and Harris, 1999; Walker et al., 1999; Wiesmayr and Grasemann, 2002). Previous studies indicate that peak pressure conditions were attained by the earliest Oligocene, for example in Zanskar (Walker et al., 1999) and in eastern Himachal Pradesh (Chambers et al., 2009; Langille et al., 2012; Fig. 1). The style of crustal thickening is, however, unknown. This study focuses on the GHC of central Himachal Pradesh, NW India, where Tertiary migmatization is negligible, if at all present, preserving information on the Eohimalayan tectonics. The putative absence of Miocene magmatism and migmatisation offers the opportunity to study the tectonic style of the GHC with different rheological properties than elsewhere along the orogen. Combined with field and microstructural observations, we use zircon and monazite U/Th–Pb geochronology to assess the time of prograde metamorphism and of the related deformation. In addition, we use muscovite 40Ar/39Ar thermochronology to constrain the onset of exhumation of the central Himachal crystalline. Comparison of the metamorphic record and exhumation history of central Himachal Pradesh with other exposures of crystalline rocks reveals a common tectonic evolution of the NW Himalaya crystalline with peak metamorphic conditions attained by 37–36 Ma followed by continuously high-grade metamorphic conditions for ≥ 10 Myr and rapid tectonically controlled exhumation at 26–22 Ma.

2. Geologic background 2.1. Geology of NW Indian Himalaya The GHC is a continuous belt of greenschist to granulite-grade metamorphic rocks and migmatites between the MCT and the STD that occurs along the entire Himalayan range. In the NW Himalaya, the GHC is exposed northeast of the Kullu–Rampur tectonic window and in the Zanskar crystalline (Fig. 1). The exposures are separated from low-grade to unmetamorphosed Tethyan Himalayan Sequence (THS) above by strands of the STD system, locally termed Sangla detachment and Zanskar shear zone (e.g., Dèzes et al., 1999; Steck, 2003; Vannay et al., 2004). Between the Sutlej and Beas valleys, e.g., in the Chamba syncline (Fig. 1), the MCT hanging wall consists of the sub-greenschist to greenschist-grade greywacke sequence of the ‘lower-grade Haimanta’, a metamorphosed equivalent of the Haimanta Fm. (e.g., Frank et al., 1995) usually found in the hanging wall of the STD. The lower-grade Haimanta is overlain by Palaeozoic to Mesozoic THS rocks. Metamorphic grade gradually decreases upsection, and no normal-sense shear zone that correlates with the STD and connects the Zanskar shear zone with the Sangla detachment has been unambiguously identified (Fig. 1).

Fig. 1. Overview geological map of the NW Himalaya (modified from Steck, 2003; Dèzes et al., 1999; Thiede et al., 2006; Leger et al., 2013). Inset shows location within the Himalaya-Tibet orogen. Numbers show locations of geochronological studies addressing the timing of prograde and high-temperature metamorphism and magmatism: 1 (Vance and Harris, 1999); 2, 5 (Walker et al., 1999); 3 (Finch et al., 2014); 4 (Robyr et al., 2006, 2014); 6, 7 (Thöni et al., 2012); 8 (Chambers et al., 2009); 9 (Langille et al., 2012). Location 1 is just north of the map. Mineral abbreviations after Kretz (1983).

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The lower-grade Haimanta of Chamba has been variably attributed to the THS, to the GHC, or mapped as a separate unit. For example, the geologic map of Thakur (1998) shows the crystalline of Zanskar and of the Beas and Sutlej valleys as tectonic window and half-window, respectively. Webb et al. (2007, 2011b) follow this view, but attribute a smaller portion of the Haimanta metasediments in the Beas area to the GHC. Other authors (e.g., Searle et al., 2007; Steck, 2003) recognize neither of the shear zones between crystalline and lower-grade Haimanta as major tectonic boundaries and attribute the Chamba syncline to the GHC. In order to avoid misunderstanding in terminology resulting from ambiguous geology, we use the terms ‘Haimanta metasediments’ and ‘crystalline’ without implication on the structural position within the MCT hanging wall, unless explicitly stated. We use term ‘extrusion’ as syn-convergent exhumation of a crustal slice between two subparallel ductile shear zones with opposite down-dip shear sense, thrusting at the bottom and normal shearing at the top (cf. Godin et al., 2006). 2.2. Tectono-metamorphic evolution of the eastern Himachal (Sutlej) and Zanskar crystalline In Zanskar (e.g., Searle et al., 1992, 1999) and the Sutlej GHC (e.g., Vannay and Grasemann, 1998; Vannay et al., 1999), Eohimalayan crustal thickening accounts for one or several stages of prograde Barrovian metamorphism with peak-metamorphic assemblages in structurally high levels of the crystalline (Fig. 1). Migmatization is widespread, especially in the Gianbul dome (e.g., Dèzes et al., 1999; Robyr et al., 2002) and in the Leo Pargil dome (e.g. Langille et al., 2012; Fig. 1). Late Eocene to Oligocene prograde metamorphism is constrained by garnet and monazite geochronology from Zanskar (~33–28 Ma, Vance and Harris, 1999; Walker et al., 1999; Fig. 1, locations 1–2) and from the Leo Pargil dome and basal sections of the Sutlej THS (~40–30 Ma, Langille et al., 2012; Chambers et al., 2009; Fig. 1, locations 8–9). The crystalline exhumed during the Miocene by SW-ward extrusion between the basal MCT and the top STD (Sangla detachment, Zanskar shear zone; e.g., Vannay and Grasemann, 1998; Searle et al., 1999; Walker et al., 1999; Vannay et al., 1999). Extrusion resulted in rapid cooling constrained by early Miocene K/Ar, Rb/Sr, and 40Ar/39Ar ages (Zanskar: ~ 22–14 Ma and up to 34 Ma along the Zanskar shear zone, Searle et al., 1999 and references therein; Sutlej GHC: 17–15 Ma, Vannay et al., 2004; Leo Pargil dome: 16–14 Ma, Thiede et al., 2006). In the migmatitic gneiss domes, late Oligocene high-temperature decompression was accompanied by widespread leucogranite emplacement (Gianbul dome: ~26–20 Ma, Dèzes et al., 1999; Robyr et al., 2006; Finch et al., 2014; Fig. 1, locations 3–4; Leo Pargil dome: decompression until 23 Ma, leucogranite emplacement 25–18 Ma, Langille et al., 2012; Fig. 1, location 9). 2.3. Central Himachal Pradesh (Beas) In central Himachal Pradesh, Haimanta metasediments are intruded by voluminous early Palaeozoic granitoids (e.g., Hanuman Tibba, Deo Tibba, Mandi granite; Fig. 1); the sequence is overlain by Palaeozoic and Mesozoic greenschist-grade to anchi-metamorphic sediments, for example in the Tandi syncline and northeast of the Chandra valley (Steck et al., 1999). In the Tosh transect, the metamorphic field gradient is inverse reaching 7–8.5 kbar and 650–700 °C at the base of the Deo Tibba intrusion (Wyss, 2000; Fig. 2). In the Beas valley, metamorphic isograds are folded by the km-scale, recumbent, SW-vergent Phojal fold. Peak metamorphic conditions in the core of the fold are N6– 8 kbar and N550–650 °C (Epard et al., 1995; Thöni, 1977). Tertiary migmatization is common and extensive in the central and eastern Himalaya, but in central Himachal Pradesh petrological, textural or geochronologic evidence for it is scarce. Epard et al. (1995) interpret aplitic and pegmatitc dikes as probably related to early Miocene extensional structures. Wyss et al. (1999) distinguish four stages of leucogranite emplacement based on structural relationships, three of

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which are related to shortening structures and are products of in-situ partial melting without melt migration; the latest leucogranitic dikelets are interpreted as decompression melts. The only time constraint is a 40.5 ± 1.3 Ma garnet Sm–Nd age of a pegmatite related to local decompression melting (Thöni et al., 2012; Fig. 1, location 6). The earliest recorded phase of crustal thickening is the NE-directed emplacement of the Shikar Beh antiformal nappe and the related formation of the Tandi syncline (Steck et al., 1999; Vannay and Steck, 1995; Epard et al., 1995; Figs. 1, 2). Emplacement of the Shikar Beh nappe is followed by SW-directed emplacement of the Nyimaling and Mata nappes (Robyr et al., 2002, 2014). Monazite and garnet geochronology indicate the end of crustal thickening at 30–29 Ma (Thöni et al., 2012; Walker et al., 1999; Fig. 1, locations 5 and 7). Cooling of the Himachal crystalline occurred between ~ 26–20 Ma (muscovite Rb–Sr and 40Ar/39Ar ages) and ~18–10 Ma (biotite 40Ar/39Ar and Rb–Sr ages; Schlup et al., 2011, and references therein; Thöni et al., 2012). The existence of a normal sense, top-to-the-NE shear zone at the roof of the crystalline (i.e., the STD) is disputed. Several authors (e.g., Epard et al., 1995; Schlup et al., 2011) point out that despite thorough mapping no STD has been determined in central Himachal Pradesh. Mylonitic shear zones with variable shear sense within and along the margins of intrusions near Manali, Rohtang pass and in the Chandra valley have been interpreted as the STD with a complex (alternating top-to-the-SW and top-to-the-NE) slip history (Jain et al., 1999; Webb et al., 2007, 2011b; Fig. 2). In the Tosh transect, Wyss et al. (1999) place the boundary between GHC and THS above the Deo Tibba intrusion, consistent with the STD trace by Thakur (1998), but point out that displacement is minimal. Top-to-the-ENE shear also occurs below the Deo Tibba intrusion in the Tosh shear zone (Wyss et al., 1999). Webb et al. (2007) propose that in the Tosh transect the STD along with the central Himachal crystalline is folded to a large-scale SWvergent recumbent fold (Figs. 2 and 3). 3. Samples Samples for zircon and monazite U/Th–Pb geochronology and muscovite 40Ar/39Ar thermochronology comprise different lithologies (orthogneiss, leucogranite and Haimanta metasediments) and cover different structural levels (Figs. 2 and 3; coordinates in Table DR1). Major structures, in particular the MCT, are folded around the Kullu– Rampur window and dip to the NW. Therefore the projection of our geochronology samples into the cross-section results in an underestimation of structural elevation for the samples NW of the profile and overestimation of structural elevation for samples SE of the profile. Although the absolute structural positions of our samples within the profile are slightly distorted the relative positions above or below the STD as suggested by different authors are correctly indicated and summarized in Table 1. The following section describes petrography and microstructures of the samples used for monazite age dating; abbreviations are after Kretz (1983). 3.1. Petrography and microstructures on monazite samples The structurally highest sample is 812C3; according to the interpretation of Webb et al. (2007) it belongs to the THS. It is a mica schist with strong metamorphic segregation into mm- to cm-scale bands and lenses of Ms + Bt and Qtz + Kfs with minor Bt. Biotite is partly chloritised. Accessory minerals include Rt and Zrn; large (≤ 100 μm) Mnz is limited to the mica bands. Garnet and Ky are present in the outcrop but thin section 812C3 has no Grt and only one small Ky crystal. In the outcrop, Ky is associated with Qtz segregates and aligned parallel to the WSW-plunging stretching lineation. 016A1 and 016B3 are from the Tosh shear zone (Fig. 2). The samples are from the uppermost GHC (Webb et al., 2007) or from its central structural level (Thakur, 1998; Wyss et al., 1999). 016A1 is a Sil + Bt schist with abundant Ky. Sillimanite forms fibrolite mats parallel to

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Fig. 2. Tectonic map of central Himachal crystalline and the traces of its tectonic boundaries according to interpretations of (1) yellow — Thakur (1998) and (2) red — Webb et al. (2007, 2011b). Thakur (1998) attribute the Deo Tibba granite and a larger part of the Haimanta metasediments along the Beas valley to the GHC. Other authors (e.g., Searle et al., 2007; Steck, 2003) do not differentiate between GHC and Haimanta sediments in western and central Himachal Pradesh. The Tosh shear zone (Wyss et al., 1999) coincides with the STD after Webb et al. (2007). Shades of grey show metamorphic zones with index minerals of the Himachal crystalline (after Steck, 2003; Leger et al., 2013). Locations of geochronology samples and muscovite 40Ar/39Ar ages are indicated.

the foliation, Ky is broken and partially replaced by Ms ± Chl (Fig. 4a, b). Muscovite also occurs as large post-kinematic flakes but does not form foliation (Fig. 4b). A distinct top-to-the-E S–C′ fabric with Sil along shear bands is present. Monazite is large (~100 μm) and very frequent, sometimes associated with Ap and/or Zrn. It occurs in the matrix, in Bt + Sil bands, and as inclusion in Ky (Fig. 4a). Rutile is present as inclusion in Ky and in the matrix.

Garnet mica schist 016B3 has a distinct top-to-the-E S–C′ fabric. Anhedral 2–4 mm Grt shows signs of resorption (embayments). Undeformed Bt occurs as Grt inclusions and in strain shadows together with Qtz and Ab, where it is interpreted to replace Grt (Fig. 4c). Kyanite is rare and partially replaced by Ms. Monazite is most common in the matrix near Grt rims. One grain is located together with Ap in an embayment in Grt. We identified two Mnz inclusions in Ms that replaced Ky and

Fig. 3. Cross section of the study area parallel to the regional NE–SW stretching lineation structural data from Epard et al. (1995) and own observations. In the Chandra valley and in the Tosh shear zone stretching lineations are rotated to a more E–W orientation by a late stage of deformation. Location of the cross section is indicated in Fig. 2. Colours and symbols as in Fig. 2. Major structures, in particular the MCT, are folded around an open, upright antiform (erosion of which to forms the Kullu–Rampur window) but, in general, structures dip to the NW. Therefore, the projection of our geochronology samples into the cross-section results in underestimation of structural elevation for the samples NW of the profile (labels placed above the profile) and overestimation of structural elevation for samples SE of the profile (labels placed below the profile). Specifically, samples 815E1, 815C1, and 8114C1 are located in the upper part of the staurolite-zone Haimanta; 823G1–3 are located closer to the MCT in the lowermost part of the staurolite-zone Haimanta; 814G1 is located several kilometres above the MCT; 806D3 and 807E1 are located above the STD(2).

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Table 1 Structural position and petrography of monazite samples. Sample

Structural position (top to bottom)a

Main mineral phasesb

Textural context of Mnz

812C3

~1 km above STD(2) ~4 km below STD(1) Top of crystalline near STD(2) ~3 km below STD(1) Centre of crystalline, approx. halfway between MCT and STD(2)

Two domains: Ms + Bt ± Ky and Qtz + Kfs ± Bt Bt, Qtz, Sil, Ky, ±Ms, ±Chl Grt, Bt, Qtz, Kfs, Ab, ±Ky Ms, Bt, Grt, Ky, ±St Qtz, Kfs, Pl, Bt, Ms

Limited to the Ms + Bt domains

016A1 016B3 8114C1 815C1

In the matrix; within Sil + Bt bands; inclusions in Ky In the matrix; inclusion in altered Ky (Ms after Ky) and in Grt; rims of Aln ± Ap In the matrix; inclusions in Grt; more frequent in micaceous than in felsic layers In the matrix

a Depending on the interpreted existence and location of the STD, the samples are from the GHC or the THS. Structural position is given with respect to the STD after Thakur (1998) (STD(1)) or Webb et al. (2007) (STD(2); cf. Figs. 2 and 3). b Abbreviations after Kretz (1983).

several small inclusions in Grt. Matrix Mnz is ≤40 μm and typically has a ≤10 μm rim of Aln ± Ap (Fig. 4d). 8114C1 and 815C1 are from the centre of the crystalline, located approximately halfway between the MCT and the STD after Webb et al. (2007). 8114C1 is compositionally layered. Muscovite + Bt rich bands are coarse grained with a strong wavy foliation and weak top-to-theSW S–C′ fabric. Felsic bands are finer grained with gneissic foliation defined by Bt and minor Ms. Mica rich layers have 1–3 mm Grt, rare St, and frequent broken and bent Ky (Fig. 4e). Fig. 4f shows a large Grt with sigmoidal trails of μm-scale inclusions (probably Qtz and Bt) in the core, which is consistent with the S–C′ fabric, and with Ky inclusions parallel to the main foliation in the rim. Where not enclosed by Grt the Ky is partly replaced by Ms. Foliation wraps around Grt, undeformed Bt and Ms occur in the strain shadows, and strain caps are weakly developed. Rutile, Ilm, and Rt with Ilm rims are frequent both in matrix and as Grt inclusions. Monazite is small (b30 μm) and occurs as inclusions in Grt and as matrix grains; it is more frequent in the mica-rich layers. Two-mica gneiss 815C1 is from a dominantly leucogranitic succession. It is rich in Qtz + feldspar (Mc and Pl); Bt and Ms form a slightly wavy foliation around feldspar porphyroclasts (Fig. 4g, h). Mineral lineation trends E–W, a very weak S–C′ fabric indicates top-to-the-SW shear. Rare post-kinematic Ms overgrows the foliation, and Bt is locally chloritised. Mnz is 20–50 μm and associated with Ap, Zrn, and minor xenotime. No textures unambiguously related to in-situ partial melting could be identified in any of the examined samples (ca. 60 thin sections of 53 samples of low-grade to high-grade metamorphic Haimanta metasediments). We also examined several samples of aplitic, leucogranitic and pegmatitc dikes and veins, which are oriented parallel to the bedding and foliation of the Haimanta metasediments, intruded along axial planes of SW-vergent folds, or cross-cut foliation and folds. Of these only one leucogranite (823G1) yielded material suitable for U–Th–Pb geochronology. The dated Zrn yielded Palaeozoic ages consistent with the age data of Cambro-Ordovician granitic intrusions. Although our field observations strongly confirm previous workers' assertion of Tertiary magmatism (e.g., dikes related to Tertiary structures) we were thus not able to identify Tertiary migmatites. 4. Methods 4.1. Zircon geochronology Zircon U/Th–Pb analyses were conducted on the Sensitive High Resolution Ion Microprobe (reverse geometry; SHRIMP-RG) at Stanford University. Minerals, concentrated by standard heavy mineral separation processes and handpicked for final purity, were mounted on double-sided tape on glass slides in 1 × 6 mm rows, cast in epoxy, ground and polished to a 1 μm finish on a 25 mm diameter by 4 mm thick disc. For age depth profiling euhedral zircon grains were pressed into Indium holders of the same size. Polished grains were imaged with transmitted light and reflected light on a petrographic microscope, and with cathodoluminescence on a JEOL 5600 SEM to identify internal structure, inclusions and physical defects. The mounted grains were

washed with 1 N HCl solution and distilled water, dried in a vacuum oven, and coated with Au. Mounts typically sat in a loading chamber at high pressure (1E− 7 Torr) for several hours before being moved into the source chamber of the SHRIMP-RG. Secondary ions were generated from the target spot with an O− 2 primary ion beam varying from 4 to 6 nA. The primary ion beam produced a spot with a diameter of 20–40 μm and a depth of 1–2 μm during six scans through the mass sequence and for an analysis time of 9–12 min. Nine peaks were measured sequentially for zircons (the SHRIMP-RG is limited to a single collector, usually an EDP electron mul204 Pb, Bgd (0.050 mass units above 204Pb), 206Pb, 207Pb, tiplier): 90Zr16 2 O, 208 Pb, 238U, 248Th16O, and 254U16O. Autocentering on selected peaks and guide peaks for low or variable abundance peaks (i.e. 96Zr16 2 O 0.165 mass unit below 204Pb) were used to improve the reliability of locating peak centres. Counting times on each peak were varied according to sample age and U and Th concentrations to improve counting statistics and age precision. Measurements were made at mass resolutions of 6000–8000 (10% peak height), which eliminated all interfering atomic species. Acid washing of the mount and rastering the primary beam for 90–120 s over the area to analysed before data is collected, assures that any counts found at mass 204Pb are actually Pb from the zircon and not surface contamination. In practice greater than 95% of the spots analysed have no common Pb. Concentration data for zircons are standardized against zircon standard R33 (419 Ma, quartz diorite of Braintree complex, Vermont, John Aleinikoff, pers. comm.) which are analysed repeatedly throughout the duration of the analytical session. Data reduction follows the methods described by Williams (1997) and Ireland and Williams (2003) and uses the Squid and Isoplot programs (Ludwig, 2001, 2003). 4.2. Monazite geochronology For monazite U/Th–Pb geochronology we scrutinized thin sections of more than 70 samples from different structural locations and lithologies. Samples for age dating were selected based on the presence of metamorphic index minerals (Grt, St, Ky, Sil) and textural location and size of monazite grains. Samples from higher or lower structural levels than those selected for monazite dating have allanite instead of monazite as the main REE-bearing phase. Analyses were conducted at NERC Isotope Geosciences Laboratory, UK and follow the procedures described in Palin et al. (2013). In-situ monazite U/Th–Pb analyses were conducted on polished thin sections using a Nu Instruments Attom ICP Mass Spectrometer coupled to a New Wave Research 193 nm excimer laser ablation system. Laser parameters include a 15 μm spot size and 30 s dwell time, at 5 Hz with a fluence of 2.7 J/cm2. U–Th–Pb data were normalized to primary reference material ‘Manangotry’ monazite (554 Ma ID-TIMS age; Paquette et al., 1994); in addition, ‘Stern’ and ‘Moacyr’ monazites were used to assess accuracy and precision of the session. For common lead correction, a Pb isotopic composition was measured for each sample on mineral phases that are assumed to be free of radiogenic lead (Ap, Pl, Bt; Table 2). With the exception of sample 812C3, Pb isotopic compositions were similar to values estimated with a Stacey–Kramers model

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Fig. 4. Photo micrographs and backscatter electron (BSE) images of monazite geochronology samples. Where applicable x and z arrows indicate sample orientation. (a) Deformed kyanite with monazite inclusion with two laser spots from in-situ dating in 016A1. Kyanite is partly replaced by sericite (XP). (b) Biotite + sillimanite schist 016A1 with postkinematic muscovite grain growing across the foliation, PP. (c) Garnet mica schist 016B3, XP (45°). Note embayments in garnet with undeformed muscovite and biotite interpreted as partial resorption of garnet. (d) BSE image of monazite in 016B3 with a rim of allanite and apatite interpreted to reflect retrograde metamorphic monazite breakdown. (e) Garnet mica schist layer of 8114C1, XP (45°). Note band of kinked kyanite parallel to the foliation. (f) Idiomorphic garnet in 8114C1. Sigmoidal inclusion trails in the core of garnet indicate early garnet growth during top-to-the-SW shear. The outer 0.5 mm of the garnet grain includes kyanite that is oriented parallel to the external foliation showing that garnet continued to grow post-kinematically. (g) Two-mica gneiss 815C1 with microcline feldspar porphyroclast. Plain polarized light (PP), x points to the W. (h) Same as g, crossed polarized light (XP).

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for Phanerozoic ages (Stacey and Kramers, 1975). For 812C3 measured 208 Pb/206Pb and 207Pb/206Pb were lower than model values; the difference amounts to a difference in corrected Th–Pb or U–Pb ages of ~ 0.03 Ma and is considered insignificant. Data were evaluated using Isoplot (Ludwig, 2003). Ages discussed below are Th–Pb single-spot corrected for common-Pb using measured Pb composition unless otherwise stated. Before LA-ICP-MS measurements, monazite grains were located in thin section by optical microscopy and by backscatter electron (BSE) imaging on a JEOL JXA 8900 microprobe at University Göttingen, Germany. BSE images did not reveal internal structure in any of the monazite grains. Chemical composition (P, Si, Ca, La, Ce, Pr, Nd, Sm, Gd, Dy, Y, U, Th) of monazite was determined by wavelength dispersive spectrometry (WDS) on a subset of those grains used for age dating (acceleration voltage 20 kV, probe current 20 nA, beam diameter 10 μm, counting time of 15–60 s on the X-ray signal and total background respectively). Most grains are 10 to 50 μm in diameter, and only one WDS analysis of the average grain composition was obtained; where possible, 2 to 3 spots were analysed on one grain. Detection limits (LoD) were calculated according to the counting statistic of the background noise. The LoD are 150 μg g−1 for CaO, 340 μg g−1 for UO2 and ThO2 and between 500 and 1000 μg g−1 for the REE. After LA-ICP-MS analysis the element distribution of Ca, Ce, Th, and Y of selected crystals was mapped by WDS to highlight possible zonation. The acceleration voltage was set to 20 kV, a probe current of 60 nA, a focused beam, and a dwell time of 50 ms per counting step were used. We analysed the chemical composition (Si, Al, Fe, Mg, Ca, Mg, Mn, Y, Ti) of garnet grains along profiles from core to rim of the two garnetbearing samples, 016B3 and 8114C1 using WDS analysis (acceleration voltage 15 kV, probe current 20 nA, beam diameter 10 μm, counting time 15–30 s). 4.3. 40Ar/39Ar thermochronology 40

Ar/39Ar thermochronology of bulk coarse-grained (ca. 200 μm average diameter) muscovite separates was carried out at Dalhousie University, Halifax. Pristine muscovite grains were hand-picked. Mica concentrates were individually wrapped in aluminium foil, and then stacked in an aluminium irradiation canister. Interspersed among the samples were five to seven aliquots of the flux monitor, Fish Canyon tuff sanidine, which has an apparent K–Ar age of 28.205 ± 0.046 Ma (Kuiper et al., 2008). The canister was irradiated with fast neutrons in the nuclear reactor at McMaster University in Hamilton, Ontario, Canada. At Dalhousie University, a double-vacuum tantalum resistance furnace was used to carry out step-heating. Isotopic analyses were made in a VG3600 mass spectrometer using both Faraday and electron multiplier collectors to measure the abundance of 39Ar for 40Ar/39Ar and 36 Ar/39Ar ratios, respectively. Errors are reported at the 2σ level and include the uncertainty in the irradiation parameter, J, but do not incorporate uncertainty in the assumed age of the flux monitor. 5. Results The following three sections describe our zircon and monazite U/Th– Pb and muscovite 40Ar/39Ar results (Tables 3–5, Figs. 5–8). The complete data sets including CL images of zircons, garnet chemical composition

Table 2 Lead isotopic compositions used for common lead correction. Sample

Phases

208

Pb/206Pb

816B3 8114C1 815C1 812C3 016A1

Ap, Pl Ap, Pl Ap, Pl Ap, Bt, Pl Ap

2.204 2.211 2.075 1.671 2.114

2σ

208

Pb/206Pb

0.020 0.034 0.055 0.063 0.036

0.849 0.848 0.814 0.630 0.804

2σ 0.015 0.010 0.014 0.019 0.006

117

and zoning, monazite element distribution maps and REE patterns, as well as additional geochronological diagrams can be found in the Supplementary data (Tables DR2–5, Figs. DR1–6). Errors are 2σ; averages are weighted by data point errors unless stated otherwise.

5.1. Zircon U/Th–Pb geochronology All granite and orthogneiss samples show a dominant Early Ordovician age component (Table 3). Granites 827C1 and 804C1 have mean 207 Pb-corrected 206Pb–238U ages of 485.4 ± 3.8 Ma and 500 ± 11 Ma, respectively (Fig. DR2). In both samples Proterozoic cores (oldest spot age 1740 ± 7 Ma) are present. Augengneiss samples 807A1 and 819A2 also yielded dominant Ordovician age clusters (480 ± 13 Ma and 485.0 ± 7.1 Ma, respectively); Proterozoic cores were only detected in 807A1 (oldest spot 2283 ± 10 Ma). In both samples zircons have thin rims that indicate a Tertiary metamorphic overprint (Figs. DR1 and DR2). In 819A2, nine analyses on rims (bright CL signal) were subdivided into two groups, with mixing between a common lead composition and radiogenic compositions equating to ages of 35.8 ± 1.3 Ma and 28.9 ± 0.7 Ma (Fig. 5). Age depth-profiling on prism faces of euhedral grains from sample 819A2 indicates that 10–20% of grains posses end of Oligocene overgrowths on the order of one micron (Table DR3). Their mean 207Pb-corrected 206Pb–238U age is 26.0 ± 1.0 Ma (MSWD = 2.24). Ti-in-zircon thermometry (Ferry and Watson, 2007) on the same spots yielded overgrowth crystallization temperature of ca. 650–700 °C (too few data to be more precise; analytical procedure as in Grujic et al., 2011). All Tertiary ages have low Th (typically b1 ppm, vs. 50 to N 200 for most Palaeozoic ages), U is lower in the ~ 36 Ma group than in the ~ 29 Ma group (40–80 ppm and 250–500 ppm, respectively; Table DR2). Garnet + Bt augengneiss 823G2 yielded two Proterozoic and 10 Ordovician ages with a mean of 480 ± 12 Ma; an Oligocene metamorphic overprint was discernable by age depth-profiling only on two grains and for 1 and 4 scan cycles only (thus b1 μm thickness), and therefore no precise age could be calculated. Leucogranite 823G1 from the same location yielded a range of early Palaeozoic ages. Garnet mica schist 823G3 has dominantly Proterozoic zircons up to 2.6 Ga, some of which have Ordovician (410 to 510 Ma) rims. The average of all early Ordovician signatures (excluding samples 823G1 and 823G3) is 485.8 ± 6.1 Ma (MSWD = 2.1).

Table 3 Zircon U–Pb SHRIMP ages. Sample

Lithology

804C1

2-Mica granite

827C1 807A1

2-Mica granite Augengneiss

819A2

Augengneiss

823G1 823G2

Leucogranite Grt + Bt augengneiss

823G3

Grt mica schist

Interpreted age, 2σ-error, model 500 ± 11 Ma 1740 ± 7 Ma 485.4 ± 3.8 Ma 480 ± 13 Ma 2283 ± 10 Ma 66.6 ± 0.5 M 122 ± 62 Ma 485.0 ± 7.1 Ma 602 ± 0.5 Ma 35.8 ± 1.3 Ma 28.9 ± 0.7 Ma 26.1 ± 1.0 Ma 5 ages 380–500 Ma 480 ± 12 Ma 1825 ± 12 Ma 6 ages 410–510 Ma 2624 ± 11 Ma

Averagea Oldest agea Averagea Concordia upper intercept Oldest agea Youngest agea Concordia lower intercept Concordia upper intercept Oldest agea Lower intercept, cluster I Lower intercept, cluster II Coherent age groupa from depth profiling Averagea Oldest agea On rims Oldest agea

Error-weighted mean of Ordovician components (excluding 823G3 and 823G1) 485.8 ± 6.1 Ma (MSWD = 2.1). a 207-Corrected 206Pb/238U age.

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K. Stübner et al. / Lithos 210–211 (2014) 111–128

812C3

27–23 Ma Youngest ages 22.4 ± 0.5 Ma 36.3 ± 0.6 Ma Youngest age 24.3 Ma Ordovician inclusions in Grt and matrix grains Late Eocene inclusions in Grt and matrix grains 38.1 ± 0.9 Ma Matrix grain age groups at 30.3 ± 0.6 Ma and 26.5 ± 0.5 Ma Youngest inclusions 26.8 ± 1.2 (Grt), 24.9 ± 1.1 Ma (St) Youngest matrix grain ages 22.7 Ma 40.7 ± 0.4 Ma b 36.6 ± 0.4 Ma b 24.7 ± 0.7 Ma youngest age 22.1 ± 1.0 Ma Ordovician Grt inclusions 26.9 ± 0.2 Ma, matrix and in Grt and Ms after Ky

8114C1

016A1

016B3 a

Average of Th–Pb ages corrected for common lead using Pb-compositions measured on main mineral phases. b Average of spot ages from grain #17.

5.2. Monazite U/Th–Pb geochronology Monazite U/Th–Pb ages are Eocene and Oligocene; pre-Cenozoic ages were rarely detected. Most samples have more than one age component (Table 4). Chemical variation of the analysed monazite is generally small. BSE imaging did not reveal any zonation. Element distribution maps reveal zoning in Y only in sample 8114C1; Th zoning is more common (016A1, 815C1, 812C3, Fig. 7; cf. supplementary data). Inclusions in metamorphic phases (e.g., Grt) are too small, 5–10 μm, to permit element mapping. Rare earth element (REE) patterns analysed with WDS on 10 μm spots show enrichment of light REE and depletion in heavy REE + Y with respect to Chondrite with little variation within or between samples. 8114C1 has higher heavy REE + Y in Mnz inclusions in Grt compared to matrix grains (Fig. DR4); the spread in HREE + Y in the matrix grains is probably an effect of mixing between high-Y rims and low-Y cores as revealed by element maps.

5.2.1. 812C3 63 spot ages were obtained on 15 grains. Several grains were large enough for up to 13 spot analyses. Single-spot common Pb-corrected Th–Pb ages range from 27 to 23 Ma (Fig. 6a). Th/U ratios vary between 1 and 35 but do not correlate with spot age. Element maps for three of the largest grains show uniform Y distribution and patchy zoning of Ca, Th, and Ce. The oldest spot ages correspond to high Ca + Th and low Ce zones in the centre of grain #31; two younger ages (22.7 Ma) are recorded in the small adjacent grain (Fig. 7a).

Table 5 Muscovite 40Ar/39Ar ages. Sample

Lithology

Age, 2σ

Model

Initial 40Ar/36Ar

804C1 807E1 806D3 807D1 807C1 803B3

Granite Bt-gneiss Orthogneiss Grt mica schist Paragneiss Leucogranite-gneiss

± ± ± ± ± ±

10 40 7 7 18 6

Bt-gneiss Orthogneiss 2-Mica gneiss Leucogranite-gneiss Orthogneiss Grt mica schist Bt-gneiss

Isochron Plateau Isochron Loss profile Isochron Plateau Isochron –/– Isochron Isochron Isochron Plateau Isochron Isochron

302 262 304 534 369 315

814G1 810B1 815C1 815E1 819A2 823G3 827B1

60.3 ± 4.8 Ma 22.0 ± 0.5 Maa 28.9 ± 3.6 Ma b27.7 ± 0.8 Mab 24.2 ± 1.1 Ma 35.1 ± 1.6 Ma 26.7 ± 2.0 Ma –/– 21.3 ± 1.0 Maa 22.6 ± 0.7 Maa 21.9 ± 0.2 Maa 21.1 ± 0.5 Maa 20.4 ± 2.4 Maa 32.0 ± 3.0 Ma

–/– 317 317 362 312 314 319

± ± ± ± ± ±

5 12 12 6 13 6

a b

Average of early Miocene ages: 21.8 ± 0.4 Ma (MSWD = 3.4). Isochron defined by 6 low-temperature steps.

Model 1 Solution (±95%-conf.) without decay-const. errs on 3 points Lower intercept: 28.89 ± 0.69 Ma Upper intercept: 5088 ± 50 Ma MSWD = 1.12, Probability of fit = 0.29

0.8

Pb/ 206 Pb

Age component, 2σ errora

0.6

207

Sample

815C1

data-point error ellipses are 2σ

1.0

Table 4 Monazite U/Th–Pb LA-ICP-MS ages.

0.4

Model 1 Solution (±95%-conf.) without decay-const. errs on 6 points Lower intercept: 35.8 ± 1.3 Ma Upper intercept: 5101 ± 88 Ma MSWD = 0.99, Probability of fit = 0.41

augengneiss 819A2

0.2

0.0

400 0

40

80

120 238

160

200

240

280

U/206 Pb

Fig. 5. Tera–Wasserburg diagram of zircon U–Pb data from augengneiss. Most ages are Ordovician; Cenozoic ages can be subdivided into two subgroups with lower intercept ages of 35.8 ± 1.3 Ma and 28.9 ± 0.7 Ma.

5.2.2. 815C1 45 spot ages were obtained on 22 grains. Most grains afforded 1–2 spots, and only 2 grains were large enough for 5 spots each. Based on Th/U ratios (Fig. 6b) the data were split into an ‘old’ group with Th/U = ~ 20 and a mean Th–Pb age of 36.3 ± 0.6 Ma (MSWD = 2.3) and a ‘young’ group with Th/U ratios of 7–14 and a range of Th–Pb ages from 36 to 24 Ma. Discordance of U–Pb ages is higher for the older age group (Fig. 6c). The element maps show uniform Y and Ce distributions, weak concentric zoning in Th, and ≤5 μm rims of elevated Ca concentrations; four laser spots at the rims yield younger ages (~27 Ma) than the cores (Fig. 7b). 5.2.3. 8114C1 We analysed 66 spots on 42 matrix grains and 29 spots on Mnz inclusions in Grt, Ilm, and St. The Th–Pb ages of Mnz inclusions range from 592 to 25 Ma and are consistent with mixing between an early Palaeozoic and a Tertiary age component. The 5 youngest ages (28.7 to 25.9 Ma) are indistinguishable from the matrix grain ages; of these, one is an inclusion in Grt, one in St, and three in Ilm (Fig. 6d). Three ages from inclusions in Grt cluster at 38.1 ± 0.9 Ma. Monazite in the intact interior of Grt away from any cracks tends to give the oldest ages (~450–500 Ma), but there is no significant pattern of ages with distance to the Grt rim. The matrix grain ages range from 370 to 22.6 Ma (Fig. 6e). Pre-Cenozoic ages were only obtained from the felsic layer. Tertiary ages were obtained from the felsic and the schistose layer without difference between these lithologies. The Tertiary matrix grain ages can be subdivided into two coherent age groups of 30.3 ± 0.6 Ma and 26.5 ± 0.5 Ma (Fig. 6f). Ca, Ce, and Th are homogenously distributed, but cores are Y-depleted in this sample (Fig. 7d); one age analysis in the centre of a low-Y core yielded 36.4 ± 1.7 Ma. The ratios Dy/Gd and Y/Gd have a high scatter but indicate a tendency to increase towards younger ages (Fig. DR4e–f). This suggests that the earlier Mnz (the ~ 36 Ma, ~ 30 Ma age groups) crystallized while Grt was growing, while the younger Mnz (~26 Ma) crystallized when Grt stopped growing or even started breaking down. In the same sample, Grt grains have uniform chemical compositions (Fig. DR6). The lack of growth zoning profiles indicates complete diffusional re-equilibration of original prograde zoning in these Grt grains during higher-T metamorphic conditions, or of growth at high-T (Yardley, 1977). Similarly, spessartine enrichment at the rims pointing to retrograde effects (Kohn and Spear, 2000; Yardley, 1977) is also lacking.

K. Stübner et al. / Lithos 210–211 (2014) 111–128

5.2.4. 016A1 We analysed 140 spots from 21 grains; several grains were large enough for 10 and more laser spots. Ages range from 43 to 22 Ma with several age clusters (Fig. 6g, h). Element maps reveal patchy zoning of Th and Ca in the cores and chemically more homogenous rims (Fig. 7c). The chemical zoning corresponds to age zones best revealed in grain #17: 18 spots on the core yield 40.7 ± 0.4 Ma and 15 spots on the rim yield 36.6 ± 0.4 Ma (average Th–Pb ages; Fig. 7c). Two ages from a Mnz inclusion in Ky (39.0 ± 1.3 Ma) are consistent with the older of the two age groups. The clusters are less well defined in the overall data set (Fig. 6g, h, j), and we consider the Th–Pb ages of grain #17 as the best example highlighting the different age components. The youngest ages of sample 016A1 are 24.7 ± 0.7 Ma (average Th–Pb ages; Fig. 6h, j, red); the remaining analyses (grey) are considered mixed ages. Fig. 6j highlights a decrease in Th/U with decreasing age, similar to the observations in 815C1 (Fig. 6b). 5.2.5. 016B3 58 spot analyses were performed on 20 matrix grains, six of which have allanite rims; 9 analyses are from inclusions in Grt and Ms after Ky. Four analyses from inclusions in Grt yielded Palaeozoic ages. All the other analyses in inclusions lie on a mixing line between common lead and a lower intercept of 27.1 ± 0.2 Ma (Fig. 6k); the average of Th–Pb ages is 26.9 ± 0.2 Ma (MSWD = 1.3). Two grains are located within Ky that is partly replaced by Ms, and one grain forms an inclusion in Grt together with Bt and albite. Th/U ratios are 10–12; slightly higher Th/U ratios were obtained from the inclusions in Grt. 5.3. Muscovite 40Ar/39Ar thermochronology Of the 13 analysed muscovite samples only three yielded 40Ar/39Ar plateau ages (e.g., Fig. 8b, c), and for 8 samples we report isochron ages (Table 5). For all samples, the initial 40Ar/36Ar is significantly higher than atmospheric composition (typically ~315 ± 10), and results have to be viewed with caution. Our interpretation is therefore based on the regional pattern of 40Ar/39Ar ages rather than the precise age of any individual sample. No age is obtained from 814G1. For 810B1 the plateau age (28.0 ± 1.5 Ma) is significantly older than nearby samples and we preferred the isochron age (21.3 ± 1.0 Ma) in our interpretation. Six samples (807E1, 810B1, 815C1, 815E1, 819A2, 823G3) yielded indistinguishable ages with an average of 21.8 ± 0.4 Ma (MSWD = 3.4). Two structurally and topographically highest samples from the Chandra valley/Rohtang pass are slightly older than average: 807C1 (3647 m a.s.l.) yielded 24.2 ± 1.1 Ma, 807D1 (3813 m a.s.l.; Fig. 8a) yielded a loss profile with high-temperature steps of ~60–100 Ma and 6 low-temperature steps that fit an isochron of 27.7 ± 0.8 Ma. Older 40 Ar/39Ar ages (≤60 Ma) were obtained from structurally high samples north of the Deo Tibba intrusion (803B3, 804C1, 806D3) and from lower-grade Haimanta rocks southwest of the Kullu–Rampur window (827B1). 6. Discussion 6.1. Ordovician magmatism and contact metamorphism Effects of the Cambro-Ordovician Bhimphedian Orogeny (Cawood et al., 2007) can be traced across the Himalaya from Pakistan to the eastern Himalaya. Cawood et al. (2007) interpret that the orogeny was as Andean-type orogenic activity on the northern margin of the Indian continent, following Gondwana assembly. The magmatic arc was associated with andesitic and basaltic volcanism and was active from ca. 530 to 490 Ma. The arc activity overlapped with, and was succeeded by, regional deformation, crustal melting and S-type granite emplacement that extended until 470 Ma. Our zircon U/Th–Pb data demonstrate Early Ordovician crystallization ages of all mapped intrusive rocks in central Himachal Pradesh,

119

and Ordovician intrusive protoliths for Himachal augengneisses (Table 3). We interpret the weighted mean of the zircon U–Pb ages of granites (827C1, 804C1) and augengneisses (807A1, 819A2, 823G2) as the time of magmatism at 485.8 ± 6.1 Ma. This estimate is similar to previously published ages of NW Indian intrusives ranging from 450 Ma to 560 Ma (Miller et al., 2001). Similar ages have been reported to the north for Rupshu (482 ± 1 Ma) and Tso Morari (479 ± 1 Ma) granites (Girard and Bussy, 1999) and in Sutlej valley, Kinnaur Kailas (459 ± 7.7 Ma; Marquer et al., 2000). Our ages are identical (within uncertainty) for intrusions in the frontal part of the lower-grade Haimanta (827C1), in Haimanta crystalline (e.g., 807A1, 823G2), and structurally low and high samples of the Deo Tibba intrusion (819A2, 804C1; Fig. 2). Tourmaline + Garnet leucogranite 823G1 yielded only few zircons but the age signature suggests that this is an Ordovician rather than a Tertiary leucogranite. Emplacement of early Palaeozoic granitoids in central Himachal Pradesh occurred during a short time interval. The large range of previously published crystallization ages of these intrusives (Singh and Jain, 2003) may be at least in part attributed to analytical uncertainties associated with the respective geochronological method (e.g., Rb–Sr, Sm–Nd, U/Th–Pb). Monazite U/Th–Pb ages are Eocene to Oligocene, but two samples yield an Ordovician age component: Ordovician ages were obtained from Mnz included in Grt in 016B3 and 8114C1 (Fig. 6d, k) and from matrix Mnz of 8114C1 (Fig. 6e). The Ordovician Mnz in the metasediments may be of detrital or metamorphic origin. Because of the Al-rich chemical composition of the samples, their close proximity to the Deo Tibba intrusion, and the common observation of contact metamorphism in the vicinity of the Deo Tibba (e.g., pseudomorphs after andalusite) we consider contact metamorphism the most likely cause of Ordovician Mnz growth. In addition, Grt mica schist 823G3 shows that the Ordovician magmatism was sufficient to induce new Zrn growth in the country rock (≤30 μm early Palaeozoic rims on Proterozoic cores of 2.6 Ga to ~600 Ma). 6.2. Himalayan metamorphism The Himalayan metamorphism resulted in overgrowths on zircon and, more prominently, on monazite. Zircon U–Pb data record Tertiary ages only in thin outer rims on Proterozoic detrital or Ordovician magmatic Zrn from three samples, 807A1, 819A2 and 823G2. The Tertiary signature is strongest in 819A2, where two age groups (35.8 ± 1.3 Ma and 28.9 ± 0.7 Ma, Fig. 5) are distinguished, while the latest overgrowth is dated at 26.0 ± 1.0 Ma (Table DR3). Based on our Mnz geochronology and on low Th/U ratios (0.002–0.005) in Zrn we interpret these Zrn overgrowths as results of regional high-grade metamorphism, rather than in-situ melting or magmatism. In 807A1 the timing of the metamorphic overprint is poorly constrained (discordia lower intercept 122 ± 62 Ma). Monazite yields Eocene to Oligocene ages throughout the Himachal crystalline, and pre-Himalayan signatures are preserved only as an exception. A Mnz grain may comprise several age domains, which may correspond to distinct chemical compositions. Tracing chemical variations of grains may help avoid mixing of different age domains within one laser-spot analysis. In our samples, Ca, Th, Y, and Ce distribution maps indicate minimal chemical variation within a sample or within a grain; this observation is confirmed by the remarkably uniform REE patterns analysed by WDS. Nevertheless, grains that are large enough for several spot analyses prove that more than one age domain exist at least in some samples. For example, one grain in sample 016A1 reveals an older core and a younger rim with well-defined respective ages (Fig. 7c). In other samples (e.g., 812C3, 815C1; Fig. 7a, b) age zonation with the youngest ages close to the rim of the grain is evident. Whether a range of ages reflects continuous monazite crystallization, laser pits that overlap distinct age domains resulting in geologically meaningless mixed ages, or Pb-loss by volume diffusion (e.g. Cherniak et al., 2004) or marginal recrystallization of the grains remains elusive.

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K. Stübner et al. / Lithos 210–211 (2014) 111–128

We suggest that our monazite age data set results for the most part from following processes: (1) prograde metamorphic monazite crystallization (e.g., Foster and Parrish, 2003; Kohn and Malloy, 2004; Smith and Barreiro, 1990; Spear and Pyle, 2002; Wing et al., 2003), (2) possibly monazite precipitation triggered by hydrothermal fluids (SeydouxGuillaume et al., 2002), which may, for example, be related to episodes

of migmatization at deeper crustal levels, and (3) mixing of age domains by the size and location of the laser spot. The lack of evidence for in-situ partial melting in the central Himachal Himalaya suggests that the analysed Mnz typically did not crystallize from partial melts (e.g., Kelsey et al., 2008; Rubatto et al., 2013). In most of our samples the large number of individual analyses resulted in one or several

uncertainty is 2σ

a

10

data-point error crosses are 2σ

42

b

812C3 Relative probability

Number

corr. Th-Pb age [Ma]

38 8

6

4

34

high-Th data points: Th/U = 20.0 ± 3.7 Mean = 36.32 ± 0.60 [1.7%] 95% conf. Wtd by data-pt errs only, 0 of 20 rej. MSWD = 2.3, probability = 0.001

30

(error bars are 2σ)

2

26

0

22

815C1 20

22

24

26

28

6

30

10

14

corr. Th-Pb age [Ma] data-point error ellipses are 2σ

0.09

c

d

0.28

Grt+Ky schist 8114C1 inclusions

0.24

Pb/ 206 Pb

815C1

0.20

St: 24.9 ± 1.1 Ma

0.16

Grt: 26.8 ± 1.2 Ma

207

Pb/ 206 Pb 207

0.06

22

data-point error ellipses are 2σ

Model 1 Solution (±95%-conf.) on 20 points Lower intercept: 36.99 ± 0.73 Ma Anchored at 207/206 = 0.814 ± 0.014 MSWD = 5.9, Probability of fit = 0.000

0.08

0.07

18

Th/U

0.12

Grt: 31.1 ± 0.9 Ma

0.05

40

44

0.04 140

36

160

32

180

200 238

28

220

0.08

240

600

0.04

260

Ilm

Grt 0

40

200

40

80

120 238

U/206 Pb

30

160

200

240

U/206 Pb

data-point error ellipses are 2σ

e

f

10

0.09

8114C1 30.3 ± 0.6 Ma

26.5 ± 0.5 Ma

all ages <50 Ma

(n = 18, MSWD = 2.3)

(n= 14, MSWD = 1.4) 8

Relative probability

0.10

Number

0.07

207

Pb/ 206 Pb

0.08

36.4 ± 1.7 Ma on low-Y core of matrix Mnz

4

0.06

0.05

matrix grains 6

8114C1 matrix

400 200

100

50

40

2

inclusions

30

0.04

0 0

40

80

120 238

160

U/206 Pb

200

240

280

20

25

30

35

40

corr. Th-Pb age [Ma]

45

50

55

K. Stübner et al. / Lithos 210–211 (2014) 111–128

18

data-point error ellipses are 2σ

0.16

g

16

36.6 ± 0.4 Ma

40.7 ± 0.4 Ma

(n= 15, MSWD = 0.71)

(n = 18, MSWD = 0.42)

h

Model 1 Solution (±95%-conf.) Anchored at 207/206 = .804 ± .006 55 points: 40.31 ± 0.28 Ma MSWD = 2.1, Prob. of fit = 0.000 43 points: 36.64 ± 0.34 Ma MSWD = 2.8, Prob. of fit 0.000 7 points: 24.94 ± 0.35 Ma MSWD = 0.74, Prob. of fit = 0.62

0.14

8

mixed ages 6

Pb/ 206Pb

grain #17

10

0.12

016A1 0.10

207

12

Relative probability

016A1

14

Number

121

0.08

25 Ma

4

39.0 ± 1.3 Ma inclusion in Ky

2

all other data

0.06 96 80

20

25

30

35

40

45

64

0.04 60

0 50

48

100

32

140

180

k

Model 1 Solution (±95%-conf.) on 62 points Lower intercept: 27.11 ± 0.15 Ma Anchored at 207/206 = .849 ± .015 MSWD = 1.15, Probability of fit = 0.20

0.6 40

30

Grt+Ky schist 016B3

41 Ma Pb/ 206 Pb

35

300

data-point error ellipses are 2σ

j

37 Ma

260

U/206Pb

data-point error crosses are 2σ

45

mixed

0.4

207

corr. Th-Pb age [Ma]

220

238

corr. Th-Pb age [Ma]

0.2 25

Ms after Ky

25 Ma Grt

016A1 0.0

20 5

10

15

20

Th/U

400 200

0

40

Grt 50

100

80

40

120

160

30

200

240

280

238

U/206Pb

Fig. 6. Monazite geochronology. Th–Pb ages are corrected for common-Pb using Pb isotopic compositions measured in plagioclase, apatite and biotite of the same sample. Errors are 2σ; averages are weighted by data-point errors. (a) Probability density distribution of Th–Pb ages of 812C3. (b) and (c) Th–Pb age versus Th/U ratio and Tera–Wasserburg diagram of 815C1. Data are subdivided into a high-Th/U cluster at 36.3 ± 0.6 Ma and a range of ages between 36 and 24 Ma with decreasing Th/U ratios. (d) Monazite inclusions in garnet (black), ilmenite (blue) and staurolite (red) in 8114C1. Grey shading marks three spot ages that cluster at 38 Ma. (e) Matrix monazite in 8114C1. Ages N36 Ma (grey) are obtained from the quartz + feldspar rich layer; most ages ≤36 Ma (black) are obtained from the mica-rich layer. (f) Probability density distribution of all Cenozoic ages has distinct peaks at ~30 Ma and ~27 Ma; youngest ages are 22.6 Ma. (g), (h), (j) Probability density distribution of Th–Pb ages, Tera–Wasserburg diagram and Th–Pb age versus Th/U ratio for sample 016A1. Data are subdivided into clusters at 40.7 ± 0.4 Ma (blue) and 36.6 ± 0.4 Ma (black) based on age distribution in grain #17 (dark grey in Fig. 6 g); the youngest age cluster has a mean Th–Pb age of 24.7 ± 0.7 Ma (red); data shown in grey are interpreted as mixing between two or more age domains. (k) Tera–Wasserburg diagram of 016B3. Except for three inclusions in garnet, all monazites are 26.9 ± 0.2 Ma (mean of Th–Pb ages), including grains within kyanite that are partly replaced by muscovite and within garnet. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

narrow and well-defined peaks in the age distribution. We argue that if diffusive Pb-loss occurred it would have affected individual grains to a variable extent, e.g., as a function of grain size; we take the narrow peaks as evidence against diffusive Pb-loss in the analysed samples, which is consistent with the estimates of peak metamorphic temperatures being lower than the temperature allowing diffusive loss of radiogenic lead (N900 °C; Cherniak et al., 2004).

6.3. Eocene prograde metamorphism The oldest record of Tertiary monazite growth is the 40.7 ± 0.4 Ma age cluster in Bt + Sil ± Ky schist 016A1 (Fig. 6g–j). The lack of Y zonation in this sample is consistent with the absence of Grt and suggests that Y was not depleted by coeval growth of Grt or xenotime. A 39 Ma Mnz inclusion in Ky in this sample indicates middle Eocene prograde metamorphic Mnz growth. In paragneiss 8114C1 some of the Eocene ages may result from mixing with an inherited Ordovician age

component, but several 38 Ma Mnz inclusions in Grt constrain late Eocene prograde metamorphism and Grt crystallization 016A1 (Fig. 6d). A prominent late Eocene age signature is recorded by 016A1 (36.6 ± 0.4 Ma) and 815C1 (36.3 ± 0.6 Ma), and coincides with the age of low-Y cores in Mnz in 8114C1 (36.4 ± 1.7 Ma) and Zrn U–Pb ages in augengneiss 819A2 (35.8 ± 1.3 Ma). In 8114C1, the Y-zoning in Mnz suggests that Mnz cores crystallized while Grt was growing (Fig. 7d). 815C1 may be a strongly sheared Ordovician orthogneiss, but evidence for a magmatic origin of this sample (e.g. relic microcline porphyroclast, Fig. 4g, h) is scarce, and unlike in the nearby sample 8114C1 there is no record of Ordovician Mnz or Zrn. We suggest that 815C1 is a typical Al-poor (no Grt, no aluminosilicates) Haimanta metagreywacke, compositionally similar to the Qtz + feldspar rich layer of 8114C1. High-Th/U Mnz crystallized during a short time interval at ~36.3 Ma probably during prograde metamorphism (Fig. 6b, black). In the central Himachal Haimanta metasediments Mnz is generally rare and small in the Qtz + feldspar rich lithologies (cf. 8114C1), whereas in 815C1 Mnz is frequent, large and idiomorphic with a weak

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a 812C3 (Th)

23.5

b 815C1 (Th)

23.9 23.7 25.2 24.4 25.0 23.4 25.2 24.1 25.9 26.8 23.5 26.0 24.6 26.3 26.1 24.6 24.7 23.9 26.5 26.6 24.7 24.1 23.5 26.4 24.6 22.8 22.7

25.0

36.2

27.9 37.0 * 33.6

26.7 37.0 *

36.3 *

27.3 35.8

22.0

* T h /U > 1 6

23.8

c 016A1

d 8114C1 (Y)

(Th)

37.7 36.8

23.8 39.0

35.2 38.4 37.5 39.2 40.4 33.8 33.0 33.1 35.3 31.6

40.7 ± 0.4 Ma

29.5 26.2

25.0

36.6 ± 0.4 Ma

42.0 32.1 39.5 38.2 36.3

29.8 29.5 27.5

30.5 28.9 36.4 29.2

Fig. 7. Element distribution maps of selected monazite grains and common-Pb corrected Th–Pb ages of single-spot analyses. Warm colours indicate high element content, and cold colours indicate low content. (a) Th in sample 812C3, grains 46, 31, and 37 (left to right); (b) Th in sample 815C1, grains 80 and 81; (c) Th in sample 016A1, grains 36, 17, and 33; (d) Y in sample 8114C1, grains 103 and 51. In (b) analyses with a Th/U ratio N16 are marked with an asterisk. Grain 17 in (c) has a bimodal age distribution and average Th–Pb ages with 2σ error are given instead of individual spot ages. Note that in sample 8114C1 Th is uniformly distributed, but in contrast to the other samples Y is zoned.

concentric Th-zoning (Fig. 7b). This observation together with the occurrence of microcline porphyroclasts that may be relics from aplitic veinlets suggest the alternative interpretation that the ~ 36.3 Ma ages represent the time of melt or pegmatite emplacement likely corresponding to partial melting at deeper crustal levels (c.f. Ayers et al., 1999; Robyr et al., 2014). In this interpretation, top-to-the-SW shearing in this sample postdates vein emplacement and is thus younger than 36.3 Ma. 016A1 from the Tosh shear zone is an Al-rich metapelite (abundant Ky + Sil), but Ms occurs only as rare post-kinematic flakes and as retrograde replacement of Ky. Although there is no textural or petrological

evidence for in-situ partial melting in this sample, the lack of Ms in the main mineral assemblage may indicate muscovite dehydration melting. This interpretation is consistent with petrological data from Wyss (2000), who report peak conditions of 8–8.5 kbar and 650– 700 °C and ‘widespread presence of small quantities of melt’ for this locality. If the sample experienced partial melting it may have occurred in the late Eocene (~36.6 Ma) or the late Oligocene (young age cluster at 24.7 ± 0.7 Ma; Fig. 6g–j). In line with our interpretation of sample 815C1 we propose that the late Eocene marks a time of widespread and intense Mnz crystallization in the central Himachal crystalline during prograde metamorphism, or at near-peak metamorphic conditions,

K. Stübner et al. / Lithos 210–211 (2014) 111–128

120

Plateau steps are red, rejected steps are black

box heights are 2σ

a) 807D1 100

Age (Ma)

80 60 40 20 0 0.0

0.2

0.4

0.6

0.8

1.0

Cumulative 39Ar Fraction 50

Plateau steps are red, rejected steps are black

box heights are 2σ

b) 807E1

Age (Ma)

40

30

20 Plateau age = 21.98 ± 0.51 Ma (2σ, incl. J-error of 1%) MSWD = 0.40, prob. = 0.92 Includes 89.1% of the 39Ar

10

0 0.0

0.2

0.4

0.6

0.8

1.0

Cumulative 39Ar Fraction Plateau steps are red, rejected steps are black

Age (Ma)

40

box heights are 2σ

c) 819A2

30

20

10

0 0.0

Plateau age = 21.09 ± 0.45 Ma (2σ, incl. J-error of 1%) MSWD = 0.59, prob. = 0.82 Includes 96.5% of the 39Ar

0.2

0.4

0.6

0.8

1.0

Cumulative 39Ar Fraction Fig. 8. Muscovite 40Ar/39Ar age spectra of selected samples; see data repository for the complete data set. (a) 807D1 is interpreted as a loss profile resulting from partial resetting of pre-Cenozoic muscovite due to heating and cooling at ≤27 Ma. (b) 807E1, plateau age of 22.0 ± 0.5 Ma includes 89% of the 39Ar. (c) 819A1, plateau age of 21.1 ± 0.5 Ma includes 96.5% of the 39Ar.

possibly associated with localized in-situ partial melting, and likely enhanced by fluid influx from deeper crustal levels, which may have experienced partial melting at this time. 6.4. High-grade conditions throughout the Oligocene Several samples show a range of Mnz ages from the late Eocene to ~24 Ma. Although these data are likely affected by some degree of age domain mixing by the laser spot, we identified several distinct age groups. In paragneiss 8114C1, Mnz of 30.3 ± 0.6 Ma and 26.5 ±

123

0.5 Ma forms rims on late Eocene low-Y cores (Fig. 6f). The decrease of Dy/Gd and Y/Gd with age suggests that late Eocene to early Oligocene Mnz crystallized while Grt was growing, whereas Mnz growth at ~ 26.5 Ma postdates Grt growth and may even be associated with Grt breakdown. Langille et al. (2012) relate a similar transition from Ypoor cores to Y-rich rims in monazite from the Leo Pargil dome to prograde Barrovian garnet breakdown and staurolite growth within the kyanite stability field at ~31 Ma. Although we have no constraints on the relative timing of Grt and St growth in sample 8114C1, the relative ages of Mnz inclusion in Grt (26.8 ± 1.2 Ma) and St (24.9 ± 1.1 Ma; Fig. 6d) are consistent with prograde Barrovian Grt breakdown and St growth. The Grt porphyroblast from 8114C1 depicted in Fig. 4f records at least three stages of its tectono-metamorphic evolution. Sigmoidal inclusion trails in its centre attest to Grt crystallization during crustal thickening by SW-directed nappe stacking. This phase was followed by Ky formation parallel to the main foliation, which was then overgrown by continued or renewed Grt crystallization. Although the crystallization of Ky was followed by further deformation (cf. bent and broken Ky in Fig. 4e), this deformation phase was not associated with significant rotation of Grt porphyroblast, preserving the foliationparallel orientation of Ky inclusion trails in Grt (Fig. 4f). Monazite inclusions were found only in smaller Grt porphyroblast in this sample, which lack textural evidence of this tectono-metamorphic evolution, therefore we have no direct age constraints on the stages of Grt growth. We suggest that crustal thickening by SW-directed thrusting resulted in prograde Barrovian metamorphism of this sample up to the Ky stability field. The timing of SW-directed shear may range from N41 Ma (oldest evidence of prograde metamorphic Mnz growth in central Himachal Pradesh) to at least 37–36 Ma (widespread and intense Mnz crystallization during prograde or near-peak metamorphic conditions). This phase was followed by Ky and Grt growth during minor pervasive deformation and possibly Grt breakdown and St formation at ~ 26 Ma. The prolonged evolution at high-temperature conditions from at least 36 to 26 Ma may account for the chemical homogeneity of Grt in this sample (Fig. DR6). In sample 016B3, matrix Mnz and inclusions consistently yield 26.9 ± 0.2 Ma (Fig. 6k). Because replacement of Ky by Ms is a retrograde, likely greenschist grade reaction (Wyss, 2000) Mnz grains within replaced Ky indicate that Ky was stable after 27 Ma. The preferential occurrence of Mnz in Bt near Grt and two conspicuous Mnz grains located in an embayment of partially resorbed Grt and in a domain of Bt + Pl in the interior of Grt, which we interpret as an embayment in Grt cut by the thin section, suggests that Mnz growth in this sample is related to Grt breakdown during prograde metamorphism (Chambers et al., 2009; Kohn and Malloy, 2004), further substantiating the prolonged high-grade metamorphic evolution inferred from sample 8114C1. Garnet in 016B3 is chemically homogeneous and interpreted to reflect diffusional re-equilibration at high temperatures (Fig. DR6). The age groups detected in zircon from sample 819A2 (35.8 Ma, 28.9 Ma, and the latest overgrowth at 26.0 Ma) correspond to the stages of late Eocene to Oligocene monazite growth and attest to temperatures of ~650–700 °C at 26 Ma (see Section 6.2). 6.5. Onset of exhumation In sample 812C3 Mnz growth occurred at ~27–23 Ma (Fig. 6a). The patchy Th zoning and the arrangement of age domains with oldest ages in the core and younger ages towards the rim (Fig. 7a) suggest continuous crystallization rather than mixing of distinct age domains. These ages postdate the main phases of Mnz crystallization in the other samples, although scattered late Oligocene ages were also recorded in 815C1, 8114C1, and 016A1 (Fig. 6b, f–k). No Zrn age and no Mnz inclusion in metamorphic phases are younger than 26 Ma, and we propose that the end of the Oligocene marks the end of high-grade metamorphic conditions of the Himachal crystalline, and that tectonic exhumation of

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the crystalline began ≤26 Ma. The 22.7 Ma appendage to one grain in 812C3 (Fig. 7a) suggests a distinct episode of renewed Mnz precipitation, which may correspond to the youngest ages in 8114C1 (22.7 Ma) and 016A1 (22.1 Ma) but the tectonic cause of this episode is uncertain. Monazite grains older than 27 Ma are not detected in 812C3 although the sample has reached similar metamorphic conditions as structurally lower ones (presence of Grt and Ky). Either the chemical composition of sample 812C3 has prevented earlier Mnz growth (cf. different age spectra of samples 016A1 and 016B3), or pre-existing Mnz was dissolved or recrystallized in the late Oligocene. Monazite is restricted to Bt bands suggesting that Mnz crystallization is related to metamorphic segregation, but the relative timing of these processes is unclear. We speculate that metamorphic segregation may have erased any earlier Mnz record and resulted in Mnz precipitation or recrystallization at ~27–22 Ma. Timing of exhumation is further constrained by muscovite 40Ar/39Ar cooling ages of 21.8 ± 0.4 Ma. Early decompression was near isothermal from peak temperatures of ~700 °C (Wyss, 2000; our Ti-in zircon thermometry). For the muscovite 40Ar/39Ar system, greatly varying diffusion parameters and closure temperatures Tc are reported in the literature (e.g., Hames and Bowring, 1994; Harrison et al., 2009; Kirschner et al., 1996), but even with a high estimate of Tc ~ 400 °C, the data of this study yield a first order cooling rate of ≥60 °C/Ma from peak temperatures to ~ 400 °C, which was followed by continuously rapid (~50 °C/Ma) cooling to ~ 300 °C (Schlup et al., 2011). Such high early Miocene cooling rates in the central Himachal crystalline can only be explained by tectonic exhumation, i.e. by thrusting at the base and concomitant normal-sense shear at the top of the sequence. The disagreement in the literature on the existence and location of the central Himachal STD, and the assertion of Wyss et al. (1999) that no appreciable offset exists below or above the Deo Tibba intrusion confirmed by our own observations suggest that early Miocene normal shear is not accommodated by one narrow shear zone but distributed over a broad zone of normal shear or partitioned between several small-scale shear zones. We suggest that the central Himachal STD system comprises shear zones at the base of the Deo Tibba intrusion (basal augengneiss exemplified by sample 819A2; Tosh shear zone; STD as suggested by Webb et al., 2007), at the top of the Deo Tibba intrusion (Thakur, 1998; Wyss et al., 1999) and potentially other as yet unidentified shear zones. 6.6. Geodynamic implications Age estimates for Eohimalayan metamorphism recorded throughout the Himalaya typically range from 37 to 30 Ma (e.g., Guillot et al., 1999). Our 40.7 Ma monazite ages are among the earliest for the onset of prograde metamorphism in the GHC, although similar ages are reported from the Gianbul dome in Zanskar (Robyr et al., 2012), from Garhwal (monazite U–Th–Pb 45–25 Ma; Foster et al., 2000; garnet Sm–Nd and Rb–Sr, 40–24 Ma; Prince et al., 2000) and eastern Nepal (monazite Th–Pb, 45 Ma; Catlos et al., 2002). Although the amount of data is still sparse, a pronounced diachroneity of prograde metamorphism along the Himalayan range is not evident; high-grade metamorphic conditions were reached throughout the Himalayan crystalline within ~10 Myr following initial collision at about 55 Ma. The metamorphic evolution of the central Himachal crystalline is comparable to Zanskar, the Gianbul dome and eastern Himachal Pradesh (Sutlej GHC, Sangla detachment hanging wall = ‘Sutlej THS’, Leo Pargil dome; compilation in Fig. 9). Peak pressures in these GHC outcrops are 8–10 kbar, and peak temperatures are 650–800 °C. Peak burial is generally estimated to have occurred at ~30 Ma (e.g., Chambers et al., 2009; Langille et al., 2012; Walker et al., 1999). We propose that the central Himachal crystalline has reached high-grade metamorphic conditions that facilitated crystallization of zircon rims (819A2) and widespread growth of metamorphic monazite (e.g., 016A1, 8114C1) as early as 37–36 Ma. There is no evidence of widespread in-situ partial melting, but our sample 815C1 may indicate deeper crustal migmatization,

which may have resulted in melt and pegmatite emplacement at ~ 36 Ma. This observation is in line with the 40.5 ± 1.3 Ma garnet Sm–Nd age of a pegmatite, which Thöni et al. (2012) relate to localized decompression melting. Although the metamorphic conditions in the central Himachal crystalline in the Eocene are unknown our data suggest high temperatures near the solidus curve. Metamorphic conditions remained high for ≥10 Myr. Our samples record narrow age peaks at 37–36, ~30, and 27–26 Ma. Although in chemically different rocks similar metamorphic conditions lead to different monazite age patterns (compare, for example, 016A1 and 016B3; gneissic and schistose layers of 8114C1) the well-defined age clusters in our data suggest episodic monazite growth controlled by regional changes in metamorphic conditions rather than local variations in rock chemistry. We propose that regional episodes of monazite growth record distinct stages in the tectono-metamorphic evolution of NW Himalaya; these episodes may be correlated with successive emplacement of the Shikar Beh, Nyimaling, and Phojal nappes (Fig. 1). Studies on the Zanskar GHC and the Gianbul and Leo Pargil domes may have missed some of these episodes, either because the areas were less affected by these tectonic episodes or, more likely, because the metamorphic overprint at ~ 30 Ma was the most intense and erased earlier monazite growth. In the THS of the Spiti area, east and structurally above our field area, a SW-directed fold-and-thrust belt developed in the middle Eocene (~ 45–42 Ma, illite 40Ar/39Ar ages; Wiesmayr and Grasemann, 2002; Fig. 1), shortly before the onset of prograde metamorphic monazite growth in central Himachal Pradesh, and ~5–10 Myr before peak temperatures were attained. Crustal thickening thus propagated SW-ward from the fold-and-thrust belt of the THS in Spiti to the central Himachal GHC, i.e. from upper to middle crustal levels. This scenario is similar to Ladakh and Zanskar, where SW-directed folding and thrusting propagated from the Nyimaling nappe down-section into the Zanskar crystalline activating a predecessor of the Zanskar shear zone as SW-directed thrust (Dèzes, 1999; Patel et al., 1993; Robyr et al., 2002). In eastern Himachal Pradesh the Sangla detachment reactivates an early SWdirected thrust likewise indicating SW-ward propagation of crustal thickening (Vannay et al., 2004). In central Himachal Pradesh, minor shear zones with both NE- and SW-directed shear sense are recorded at different structural levels, e.g., in the Chandra valley, near Rohtang pass or at the base of the Deo Tibba intrusion (Jain et al., 1999; Thöni, 1977; Webb et al., 2011b; own observations). We suggest that some of these shear zones may result from Eocene/Oligocene crustal thickening by SW-directed thrusting, which propagated south and down section from the Spiti fold-and-thrust belt to the GHC of the central Himachal Pradesh. Estimates for the exact onset of decompression vary: In the Gianbul dome, decompression started either before ~27 Ma (Robyr et al., 2006) or between 26 and 22 Ma (Finch et al., 2014). The Leo Pargil dome experienced high-temperature decompression at ~ 29–23 Ma followed by decompression and cooling (Langille et al., 2012). In this study, the youngest Mnz inclusions in Grt, Ky and St suggest that high-grade metamorphic conditions prevailed until ~27–26 Ma, and that decompression of the central Himachal crystalline occurred after 26 Ma. We therefore relate the 29 ± 1 Ma garnet Sm–Nd age reported by Thöni et al. (2012) to prograde metamorphic garnet growth rather than continued exhumation of the crystalline. The discrepant interpretations of the exhumation history of the Gianbul dome (e.g., Finch et al., 2014; Robyr et al., 2006, 2014) impede a comprehensive comparison of the onset of exhumation; we suggest that decompression started contemporaneously in the latest Oligocene throughout the NW Himalayan crystalline, but possibly slightly earlier in the Leo Pargil dome. In contrast to the common history of prograde metamorphism and high-temperature decompression, cooling through muscovite 40Ar/39Ar closure temperature was diachronous across the NW Himalaya. Comparison of cooling paths is hampered by the combined uncertainty of the onset of exhumation, the temperature at the onset of exhumation, the thermal evolution during early decompression (heating, isothermal or

K. Stübner et al. / Lithos 210–211 (2014) 111–128

peak burial conditions: 30 Ma 30 Ma 30–29 Ma

melting during heating 29–26 Ma near-peak conditions: 37–36 Ma this study

9

8

dehydr. melting related to heating

prograde metamorphism: 33–28 Ma 42–27 Ma 41–36 Mathis study

? 23 Ma

7

pressure [kbar]

125

21 Ma

high-temperature decompression: 30–21 Ma 26–22 Ma 29–23 Ma 30–23 Ma 26–23 Mathis study

decompression melting & leucogranite emplacement 26–20 Ma

6

ky

decompression melting & leucogranite emplacement 23–18 Ma

5

23 Ma 4

muscovite 40 Ar/ 39Ar

3

sill 15 Ma 21 Ma 21-22 Mathis study 14-16 Ma 13 Ma 15-17 Ma 400

Western Zanskar Gianbul dome Central Himachal crystalline Leo Pargil dome Sutlej THS Sutlej 700 GHC 800

and 500

600

temperature [°C]

temperature [°C] main phases of mnz crystallization in Central Himachal Pradesh

800 600

muscovite 40 Ar/ 39Ar

400 200 45

40

35

30

25

20

15

10

5

geologic time [Ma] Fig. 9. Compilation of metamorphic paths and age constraints from western Zanskar (Searle et al., 1999; Vance and Harris, 1999), the Gianbul dome (Dèzes et al., 1999; Finch et al., 2014; Robyr et al., 2002, 2006, 2014; Walker et al., 1999), the Leo Pargil dome (Langille et al., 2012; Thiede et al., 2006), the footwall and hanging wall of the Sangla detachment (‘Sutlej GHC’ and ‘Sutlej THS’; Chambers et al., 2009; Vannay et al., 1999, 2004), and the central Himachal crystalline (Wyss, 2000; Schlup et al., 2011; Thöni et al., 2012; this study). Aluminosilicate stability fields and dehydration melting curve in top panel from Wyss (2000). Bottom panel shows corresponding time-temperature paths (thin lines) and documented episodes of monazite growth (bold lines) in the NW Himalaya. Black boxes in the upper part of the diagram indicate the main phases of monazite growth documented in this study. Circles show muscovite 40 Ar/39Ar ages; a lower closure temperature (~300 °C) is depicted for the Sutlej THS because of the smaller grain size of samples used in that study (Chambers et al., 2009). See text for further details.

cooling), and the effect of decompression melting and crystallization on the heat budget. For example, whereas most studies suggest onset of decompression and cooling at ~ 23–22 Ma (e.g., Chambers et al., 2009; Walker et al., 1999), Langille et al. (2012) argue that by 23 Ma the Leo Pargil dome had already exhumed to 4 kbar/600 °C. Muscovite 40 Ar/39Ar closure temperatures reported in thermochronologic studies of the NW Himalaya range from 300 to 480 °C; this variation reflects in part variable cooling rates and grain sizes (Dodson, 1973), but also our still incomplete understanding of Ar diffusion in muscovite. Within the eastern Himachal Pradesh 40Ar/39Ar ages correlate with structural position, i.e. older ages are recorded in the crystalline (14–17 Ma, GHC and Leo Pargil) compared to the THS (~13 Ma) reflecting more rapid, structurally controlled exhumation and cooling in the respective footwalls of the Sangla detachment and Leo Pargil shear zone (e.g., Chambers et al., 2009; Thiede et al., 2006; Vannay et al., 2004). The ages are similar to those obtained in the Zanskar crystalline (~15 Ma; Searle et al., 1999 and references therein), and although peak metamorphic conditions are considerably higher in Zanskar than in the Sutlej GHC, the early Miocene exhumation and cooling history is similar. In the Gianbul dome, 21–19 Ma muscovite 40Ar/39Ar ages suggest extremely rapid cooling resulting from rapid exhumation along shear zones north and

south of the dome, probably assisted by buoyant uplift of early Miocene intrusives (Robyr et al., 2006). We suggest that the central Himachal crystalline exhumed as footwall block of top-to-the-ENE shear zones (central Himachal STD system), but in contrast to the Gianbul dome, shear was distributed (e.g., top and base of Deo Tibba intrusion), and shear zones are sub-horizontal and therefore do not account for rapid tectonic denudation and rock uplift in the footwall block. We suggest that the 22–21 Ma 40Ar/39Ar ages do not reflect exceptionally rapid exhumation of the central Himachal crystalline but that early Miocene exhumation rates were similar to the Zanskar and Sutlej GHC. The older 40Ar/39Ar ages result from lower temperatures at the onset of exhumation and probably shallower level of upper crustal exhumation than elsewhere along the GHC. The lack of evidence for migmatization in the central Himachal GHC indicates that its rheology was different (i.e., strength was higher) than elsewhere along the Himalayan orogen. It is possible that partial melting occurred in central Himachal Pradesh in the late Oligocene, but was limited to not-yet exhumed deeper crustal levels. This possibility raises the question why this segment of the GHC underwent much less exhumation, or extrusion, compared to the rest of the belt. Alternatively, there was never partial melting in central Himachal Pradesh at the

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outcrop or at deeper levels of the GHC, except for widespread decompression melting in migmatitic gneiss domes (Leo Pargil, Gianbul dome). If so, this could be either a compositional or temperature effect, or the barrier effect of the Karakoram fault to the southward propagation of partially molten crust from underneath the Tibetan plateau (Leech, 2008). In either case, the structure of central Himachal Pradesh indicates the tectonic style of the GHC in the absence of rheological weakening by partial melting and magmatism, and is characterized by stacking of recumbent fold anticlines and their exhumation along basal thrusts and a broad zone of normal shearing at the top. It may thus represent the style of deformation of the entire GHC belt during the Eocene and Oligocene, i.e. prior to SW-ward extrusion of the GHC and accompanying Miocene partial melting. 7. Conclusions New monazite U/Th–Pb geochronological data in conjunction with zircon U–Pb geochronology and muscovite 40Ar/39Ar thermochronology unravel the history of Eocene–Oligocene crustal thickening in the NW Himalaya and the onset of extrusion of the crystalline. Large igneous bodies (Deo Tibba, Hanuman Tibba, Mandi granite) emplaced at 486 ± 6 Ma are the protoliths for augengneiss within the Himachal crystalline, e.g. in the Chandra valley. Ordovician magmatism was accompanied either by contact metamorphism or by regional metamorphism, which led to monazite growth and crystallization of metamorphic rims around Proterozoic zircon cores in the Haimanta sediment country rock. Contraction the Eohimalayan triggered prograde Barrovian metamorphism as early as ~ 41 Ma. Near peak-metamorphic conditions (~ 8–8.5 kbar, ~ 600–700 °C) were probably attained by ~ 37–36 Ma and may have been accompanied by partial melting in deeper crustal levels. Metamorphic conditions remained high through the Oligocene. Distinct episodes of monazite crystallization at ~36.5, 30 and 27 Ma recognized throughout the crystalline are probably related to stages in the tectonic evolution. If further investigations unravel the regional distribution of relative significance of these episodes it may be possible to correlate monazite growth to the successive emplacement of NE- and SWdirected nappes previously recognized in the NW Himalaya (e.g., Shikar Beh, Nyimaling, Phojal). Exhumation of the central Himachal crystalline started as nearisothermal decompression at ~ 26–23 Ma followed by decompression and cooling to ~ 400 °C by 21.8 Ma. Early Miocene exhumation rates are comparable to those of the GHC in Zanskar and eastern Himachal Pradesh, and older 40Ar/39Ar cooling ages result from differences in the thermal budget, most importantly lower temperatures at the onset of exhumation and the lack of decompression melts. Early Miocene exhumation is attributed to the extrusion of the crystalline between SW-ward thrusting along the base (MCT), and normal shear along the top; the latter was probably distributed over several shear zones including those at the base and at the top of the Deo Tibba intrusion. We consider this system of normal-sense shear zones the local expression of the Sangla detachment and Zanskar shear zone, and suggest that in central Himachal Pradesh this STD system was active in the late Oligocene/early Miocene. In contrast to the Gianbul and Leo Pargil domes and to the majority of the Himalayan orogen, the exhumation of the central Himachal crystalline was not associated with migmatization nor driven by the related rheological weakening. Supplementary data to this article can be found online at http://dx. doi.org/10.1016/j.lithos.2014.09.024. 8. Acknowledgements The study was supported by the German Science Foundation (DFG grant STU 525-1/1) and by the Natural Sciences and Engineering Research Council of Canada (NSERC, grant RGPIN/2274752009). KS was partially funded through a scholarship from the Canadian Bureau

for International Education (CBIE-BCEI). Field work was made possible by Tashi Tsering and the assistance of Borja Antolin and Daria Czaplinska. Matt Horstwood and Andy Smye helped with monazite U– Th/Pb analyses at NERC. The final version of the manuscript has benefitted greatly from constructive and detailed comments by two anonymous reviewers.

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