Protolith lithostratigraphy of the Greater Himalayan Series in Langtang, Nepal: implications for the architecture of the northern Indian margin

Brendan Dyck, Marc St-Onge, Michael P. Searle, Nicole Rayner, David Waters and Owen M. Weller
Geological Society, London, Special Publications, 483, 281-304, 20 September 2018, https://doi.org/10.1144/SP483.9

Abstract

Reconstruction of the protolith lithostratigraphy of amphibolite-facies metasedimentary rocks of the Greater Himalayan Series (GHS) in Nepal documents a single, long-lived passive-margin succession that was deposited along the northern margin of the Indian Craton. In the Langtang area, Paleoproterozoic gneisses are unconformably overlain by a succession of upper Neoproterozoic–Ordovician fluvio-deltaic quartzite, basinal pelite and psammitic beds that grade upsection into micaceous semipelite and pelite. U–Pb zircon geochronology yields maximum depositional ages between c. 815 and 460 Ma for the GHS in Langtang. Regional variations in the composition and thickness of the GHS along the length of the Himalaya are attributed to siliciclastic depocentres centred on Zanskar in northern India, Langtang and Everest in central to western Nepal, which contrast with coeval marine carbonate shelf deposition in the Annapurna region. The protolith lithostratigraphy documented for Langtang provides a coherent framework for interpreting subsequent Cenozoic Himalayan deformation, specifically the homogeneously distributed layer-normal shortening (i.e. flattening) and layer-parallel stretching (i.e. transport-parallel stretching) that characterizes the GHS. Within the context of a single protracted northern Indian marginal sedimentary succession, the distinction between the Lesser, Greater and Tethyan Himalaya is structural rather than lithostratigraphic in origin.

Supplementary material: Concordia plots showing all detrital zircon analyses and representative zircon cathodoluminescence (CL) images are available online at https://doi.org/10.6084/m9.figshare.c.4234313

The Greater Himalayan Series (GHS) is a c. 5–20 km-thick layer of mid-crustal schist, migmatites and leucogranites, bounded by a major south-vergent shear zone, the Main Central Thrust (MCT), below and a major top-north shear zone and low-angle normal fault, the South Tibetan Detachment (STD), above. The Langtang Valley in the Nepal Himalaya cuts through the MCT and GHS zone of inverted metamorphic isograds from biotite-grade upsection through garnet, kyanite, sillimanite + muscovite, sillimanite + K-feldspar to cordierite-grade schists and migmatites (e.g. Reddy et al. 1993; Massey et al. 1994; Kohn et al. 2005; Kohn 2008; Dyck 2016). North of the Nepal–Tibet border, the top of the GHS and the bounding STD are exposed in the Rongbuk Valley north of Everest (Burchfiel et al. 1992; Searle 2003; Cottle et al. 2015), and in the Shisha Pangma region north of Langtang (Searle et al. 1997b). While the Cenozoic metamorphic and structural evolution of the GHS is reasonably well known (e.g. Kohn 2014), details of the protolith and tectonic origin of GHS rocks are still debated. Himalayan protolith studies can be used to constrain several important tectonic issues relating to the GHS, notably: (1) the nature of the passive and rifted margin of the Indian Plate throughout the Proterozoic and Paleozoic; (2) evidence for the Ordovician Bhimpedian (Kurgiakh) Orogeny along the Himalaya (Cawood et al. 2007; Myrow et al. 2010, 2016; Palin et al. 2018); (3) the palaeostratigraphic relationship between the GHS protolith and the basal units of the Tethyan Himalaya; (4) the palaeostratigraphic relationship between the GHS and unmetamorphosed rocks of the Lesser Himalaya; and (5) the nature and location of the MCT (e.g. Searle et al. 2008). This paper documents the protolith stratigraphy of GHS rocks from Langtang, and considers the detrital zircon ages obtained from the study area within a broader regional context along strike.

Geological history of the northern Indian margin

The volcano-sedimentary succession deposited along the northern margin of Peninsular India records a long history of rifting, basin formation and continental margin sedimentation between the Neoproterozoic and Paleogene. Tectonic thickening and uplift of the Indian margin, by folding and thrusting, during the collision between India and Eurasia has exposed this succession across the Himalaya (Searle et al. 1987; Searle 2015). There are three main phases of passive and rifted margin sedimentation preserved in the stratigraphy of the Himalaya, with initiation in the (1) Neoproterozoic, (2) late Carboniferous–early Permian and (3) late Permian–Middle Jurassic (e.g. Kumar & Brookfield 1987; Brookfield 1987; Garzanti et al. 1987; Garzanti 1999). The strata deposited during the Neoproterozoic–Middle Jurassic unconformably overlie Paleoproterozoic and Mesoproterozoic units (DeCelles et al. 2004). Sedimentary rocks deposited on the northern Indian margin during the Neoproterozoic record an early phase of rifting, associated with the opening of the Paleotethys (Kumar & Brookfield 1987; Brookfield 1993), and subsequent passive margin sedimentation (Garzanti 1999; Garzanti et al. 1987; Myrow et al. 2003; Gehrels et al. 2003, 2011). Apart from the growing record of detrital zircon ages in the Himalayan metamorphic core (e.g. Parrish & Hodges 1996; DeCelles et al. 2000; Gehrels et al. 2003, 2011; Spencer et al. 2012, 2018), most of what is known about the architecture of the Indian margin prior to the break-up of Gondwana comes from studying the record of low-grade and non-metamorphosed sedimentary strata that outcrop within the Lesser Himalaya Series (LHS) and the Tethyan Himalayan Series (THS) (Fig. 1). The GHS represents a relatively untapped source of information about the northern Gondwanan margin, which we explore herein.

Fig. 1.
Fig. 1.

Simplified geological map of the Himalaya, and locations of the Everest, Langtang, Annapurna and Zanskar regions. Modified after Searle et al. (2008).

The LHS, GHS and THS are all structural domains bounded by large-scale faults, notably from north to south, the Great Counter Thrust or Indus Suture Zone (ISZ), STD, MCT and Main Boundary Thrust (MBT) (Fig. 1). These faults (and associated ductile shear zones) cut across stratigraphic boundaries, such that a common stratigraphic unit may be present in all three structural domains. Restoration of the Himalaya shows that India was one contiguous plate, and stratigraphic units can be correlated across from the LHS north to the THS (Myrow et al. 2003, 2009; Searle et al. 2006). The LHS comprises a predominately Paleoproterozoic succession that overlies Paleoproterozoic and Archean crystalline basement. Within a series of doubly-plunging synclines (Jiang et al. 2002) and palaeogeographical lows (Godin & Harris 2014), these rocks are overlain by uppermost Mesoproterozoic–Cambrian (Brookfield 1993; McQuarrie et al. 2008; Myrow et al. 2010; Mckenzie et al. 2011) and Upper Cretaceous–Eocene strata (Brookfield 1993), which collectively form the ‘upper LHS’. The low-grade THS lies to the north and sits structurally above the GHS. It comprises the Neoproterozoic–Eocene so-called ‘Tethyan’ Himalayan Series, which preserves a record of sedimentation in the Neoproterozoic, late Carboniferous–early Permian and late Permian–Middle Jurassic (Brookfield 1993; Garzanti 1999; Chatterjee et al. 2013). Younger overlapping strata to the south define the influx of foreland basin sedimentation, which resulted from the crustal thickening and uplift of the Himalayas (DeCelles et al. 2004)

Previous studies of the Himalayan detrital zircon record have shown that there is considerable overlap in the depositional ages of the upper LHS, GHS and lower THS strata (DeCelles et al. 2000, 2004; Gehrels et al. 2003; Myrow et al. 2003, 2009, 2010). Published detrital age spectrums range between c. 3520 and 520 Ma for the upper LHS (Myrow et al. 2003; Colleps et al. 2018), c. 2700 and 500 Ma for the GHS (DeCelles et al. 2000, 2004), and c. 3450 and 500 Ma for the lower Paleozoic strata of the THS (Myrow et al. 2003, 2010).

In the simplest scenario, the LHS, GHS and THS constitute the depositional record of a single continental margin that is both spatially and temporally continuous along the length of the Himalaya (Searle 1986; Brookfield 1993; Myrow et al. 2003, 2015). However, the high metamorphic grade of the GHS and movement along the two major structures that bound it (Fig. 1) has obscured the primary depositional relationships between the LHS, GHS and THS, making it difficult to assess this interpretation.

Regional lithostratigraphic context

Published models that explain the pre-Himalaya distribution of the LHS, GHS and THS strata, based on palinspastic restoration of the imbricate stack of thrust sheets, make unique predictions for the GHS in terms of protolith lithostratigraphy and timing of sediment accumulation. The prevailing model, in which the LHS, GHS and THS cumulatively represent the ancient continuous cover margin succession of the northern Indian Craton, implies a rift/passive-margin setting for accumulation of the GHS during the late Neoproterozoic–upper Cambrian (Searle et al. 1987; Brookfield 1993; Gehrels et al. 2003; Myrow et al. 2003, 2010). Alternative models, such as the crystalline axis model (Saxena 1971; Aharon et al. 1987), interpret the LHS and THS as distinct depositional basins separated by a high-standing crystalline ridge. The crystalline axis model interprets the GHS as a Precambrian basement complex that was unaffected by Cenozoic Himalayan metamorphism, which is wholly inconsistent with the known timing of GHS metamorphism (e.g. Searle et al. 2006; Kohn 2014). A third model, which incorporates the more recent work on upper Cambrian–Lower Ordovician granite suites in the Western Himalaya, suggests that the GHS and THS form the basement and cover units of an accreted terrane juxtaposed against India and its LHS cover (DeCelles et al. 2000). These alternative models notably do not fit either the coeval upper LHS, GHS and lower THS detrital zircon record (Myrow et al. 2003), nor the similar whole-rock εNd isotopic record in the LHS and GHS rocks (Whittington et al. 1999; Ahmad et al. 2000; Clift et al. 2001; Robinson et al. 2001). Moreover, they remain unsupported by the absence of intervening suture zone lithologies (Searle et al. 1999, 2003, 2008). Nevertheless, confidence in the continuous margin interpretation would be greatly improved if the protolith lithostratigraphy of the high-grade GHS was further constrained.

Geology of the Greater Himalayan Series and samples

Having experienced medium–high temperatures (550–750°C) of metamorphism and up to c. 30 km of tectonic burial, the metasedimentary rocks of the GHS preserve a record of crustal orogenic evolution. Metamorphic grade is documented to increase structurally upwards from kyanite–muscovite assemblages in the hanging wall of the MCT zone (Figs 2 & 3), through subsolidus kyanite–sillimanite–muscovite to K-feldspar–sillimanite at the vapour-absent muscovite dehydration reaction isograd (Searle & Rex 1989; Stephenson et al. 2000; Searle & Godin 2003). Leucogranite sills and dykes, originating from the muscovite dehydration melting reaction, intrude the upper half of the GHS and are hosted in K-feldspar–sillimanite-bearing metasedimentary rocks (Dyck 2016). The uppermost leucogranites of the GHS record the lowest pressures of peak-temperature metamorphism (<4 kbar), preserving cordierite–sillimanite assemblages (Streule et al. 2010; Groppo et al. 2013).

Fig. 2.
Fig. 2.

Geological map of Langtang documenting the principal lithostratigraphic units, structures and metamorphic isograds. Labels denote the location of samples used for geochronology.

Fig. 3.
Fig. 3.

Cross-section through the Greater Himalayan Series in Langtang, showing lithostratigraphic units, distribution of metamorphic mineral isograds and the Main Central Thrust zone.

Samples used in this study were collected across the GHS in Langtang (Fig. 1) during field mapping campaigns in 2012 and 2014. Previous work in Langtang (Inger & Harris 1992; Macfarlane 1992, 1993, 1995; Reddy et al. 1993; Parrish & Hodges 1996; Fraser et al. 2000; Pearson 2002; Takagi et al. 2003, 2007; Kohn et al. 2004; Pearson & DeCelles 2005; Kohn 2008) has outlined the general structural, chronological and lithological framework for this area, but in general has focused on the lower MCT zone, which can be readily accessed by trekking routes. In this study, we extended the observations to the uppermost levels of the GHS, beyond the base of the Cenozoic leucoganite bodies along the Langtang glacier valley (Figs 2 & 3). While not directly accessed in the field, the position of the STD zone is constrained in the northern segment of the Langtang transect (Fig. 3) with the use of satellite imagery and field observations in the neighbouring Shisha Pangma region in south Tibet (Searle et al. 1997b).

The degree of deformation is largely consistent along the transect and is characterized by primary lithostratigraphic bedding transposed into a dominant east–west-striking metamorphic foliation, defined by the alignment of biotite and the flattening of quartz and feldspars. The exception is the imbricated base of the GHS, which is punctuated by discrete zones of high strain localized along the base of individual thrust imbricates. Seven imbricate thrust sheets were identified, which repeat the lowermost GHS stratigraphic succession and collectively make up the MCT zone in Langtang (Figs 2 & 3). The thickness of individual imbricates ranges from c. 50 to 700 m. The upper two imbricates comprise granodioritic augen gneiss overlain by quartzite and pelite, whereas the lower five thrust sheets involve only quartzite and overlying pelite. The regional foliation deviates to more northerly orientations near the base of the MCT zone, as a result of late cross-folding and/or the presence of a lateral ramp (Macfarlane 1992), and is accompanied by the development of crenulation folding with a north-plunging fold axis (Takagi et al. 2003). Structurally above the MCT zone, S–C fabrics and both δ- and σ-type feldspar augens record ductile shear, which transitions from top-to-the-south thrust motion in the lower part of the GHS to top-to-the-north normal motion from the centre of the GHS structurally up to the STD zone, similar to the structural transition documented in other transects across the Himalaya (Searle et al. 1997a, 2003). Previous workers have identified thrust faults within the GHS in the Langtang area based on changes in lithology or the presence of metamorphic index minerals (e.g. the Langtang Thrust: Kohn 2008). However, we did not observe any obvious repetitions of lithostratigraphic units, or distinguishable high strain zones, that would suggest structures with major displacement.

Occurrences and previous interpretations of the Greater Himalayan Series augen gneisses

There are two distinct units of augen gneiss within the GHS, which are distinguished based on their lithostratigraphic position (at the base v. the upper half of the GHS) and their age. The lower augen gneiss unit, termed the Ulleri augen gneiss in Annapurna (Pêcher & Le Fort 1977; Colchen et al. 1986; Parsons et al. 2016b) and the Phaplu augen gneiss in Everest (Maruo & Kizaki 1981; Searle et al. 2006), has a granitic to granodioritic composition (Parrish & Hodges 1996), and yields monazite and zircon with ages between c. 1970 and 1400 Ma (Catlos et al. 2002; DeCelles et al. 2004). This lower augen gneiss unit has been interpreted to have intruded into a metasedimentary host, which depending on where the MCT is mapped, may be the uppermost structural unit of the LHS (Le Fort 1975), or the lowermost structural unit of the GHS (Macfarlane 1992).

The upper augen gneiss unit, often referred to as the ‘Formation III augen gneiss’ after Colchen et al. (1986) or ‘Pumori K-feldspar augen gneiss’ (Searle et al. 2003, 2006), is by contrast more varied in both appearance and composition, and yields zircon with ages between c. 490 and 470 Ma (DeCelles et al. 2000; Godin et al. 2001). Unlike the lower augen gneiss unit, which is restricted to the lowermost level of the GHS, the Formation III augen gneiss is widespread, with multiple occurrences within the upper half of the GHS. Le Fort (1975) reported a 300 m-thick layer of coarse augen gneiss in the uppermost levels of the GHS in the Annapurna–Manaslu ranges, which is a notable occurrence because individual K-feldspar crystals are up to 30 cm in diameter. The origin of the Formation III augen gneiss has been attributed to deformation of a belt of Cambrian–Ordovician granites that extends along strike from Zanskar in the west (Pognante et al. 1990) to Bhutan in the east (Gansser 1983). Deformed Formation III augen gneiss layers in the GHS are correlated with undeformed granites exposed in the Kathmandu Klippe (Gehrels et al. 2003), which yield zircon and monazite with an age of c. 470 Ma (Schärer & Allègre 1983). The prevalence of upper Cambrian detrital zircon grains within GHS siliciclastic rocks along the length of the mountain belt (Myrow et al. 2009, 2010, 2015, 2016) implies that formation of the Cambrian–Ordovician granites was nearly contemporaneous with the deposition of some of the GHS rocks. As such, the interpretation of whether or not the Formation III augen gneiss represents deformed granite has a bearing on how the lithostratigraphic history of the GHS is interpreted.

Whereas the change in mineral assemblages as a result of metamorphism for a wide variety of compositions is known (e.g. Bowen 1940; Spear & Cheney 1989; Connolly & Trommsdorff 1991; Frey et al. 1991), and equilibrium phase diagrams can be calculated for most lithologies (cf. Powell et al. 1998; Capitani & Petrakakis 2010), the change in the penetrative rock texture(s) that accompanies mineral reactions can conceal the features and structures required to readily identify a rock's protolith. Of relevance to this work is an understanding of the textural evolution of siliciclastic rocks during vapour-absent muscovite dehydration melting, which was recently documented by analysis of a suite of Langtang and Everest samples (Dyck 2016). Dyck (2016) documented the reaction and mineral growth mechanisms that operated during muscovite dehydration melting, and established the importance of the partial melting process on grain-size coarsening, culminating in the formation of pelitic K-feldspar augen gneiss. The implication is that the Formation III augen gneiss could have been formed by the textural modification of a pelite during muscovite dehydration melting, rather than deformation of granitic bodies. Consequently, a modified lithostratigraphic interpretation of the GHS is required, which does not invoke late Cambrian–Early Ordovician granite magmatism along the entire length of the Himalaya (cf. Gehrels et al. 2003).

In this paper, we identify the first-order protolith lithostratigraphy of the GHS exposed in Langtang (central Nepal: Fig. 1) and contribute new zircon ages from both the augen gneiss units and the surrounding metasedimentary lithologies. We present an alternative lithostratigraphic interpretation for the GHS based on our protolith interpretation of the K-feldspar augen gneiss and metasedimentary units, and discuss how this fits into the continuous margin model for the accumulation of the LHS, GHS and THS.

Langtang lithostratigraphy

The Langtang section exposes a siliciclastic-dominated stratigraphic succession that shares many similarities with parallel sections in the Everest and Zanskar areas (Fig. 1) (Searle et al. 1997a, 2003, 2006). Here we describe the GHS lithostratigraphy as a succession of eight map units, which we have traced along a series of valley sections in the Langtang area. The basal unit of the GHS is a c. 300 m-thick panel of coarse-grained paragneiss and orthogneiss, dominated by biotite augen gneiss of granitic to granodioritic composition, herein called the basement complex (Fig. 4a). The gneiss is strongly deformed and irregularly banded at the centimetre/metre scale and cross-cut by dykes of aplitic monzogranite. This augen gneiss forms a major part of the basement complex, is found across the Nepalese Himalaya, and is correlated with the Ulleri and Phaplu augen gneiss in Annapurna (Pêcher & Le Fort 1977; Colchen et al. 1986; Parsons et al. 2016b, c) and Everest (Maruo & Kizaki 1981).

Fig. 4.
Fig. 4.

Outcrop photographs of the major lithostratigraphic units. The scale card is in centimetres (8 cm). (a) Plagioclase and K-feldspar augen orthogneiss, formed by ductile deformation of a coarse-grained granodiorite. The augen orthogneiss, together with aplitic dykes and layers of paragneiss, comprise the Indian basement, which is unconformably overlain by the Greater Himalayan Sequence metasedimentary units. (b) Medium- to thick-bedded quartzite with heavy mineral banding. (c) Thin- to medium-bedded low-angle quartzite laminasets interbedded with pelitic garnet–biotite schist. (d) K-Feldspar and sillimanite pelitic augen gneiss. (e) Right-way-up interbedded psammite and pelite turbidite bed with reverse metamorphic grading formed as a result of the porphyroblastic intergrowth of sillimanite + quartz in more pelitic layers.

Approximately 400 m of thickly bedded quartzite directly overlies the gneiss, preserving a relationship interpreted as an unconformable basement–cover contact. The thickness of individual beds decreases systematically upwards from the basal contact, from metre- to decimetre-scale. The quartzite is equigranular with 2–3 mm quartz grains that are recrystallized, displaying a sugary texture. Decimetre-scale trough cross-stratification and the accumulation of heavy mineral grains within the centre of concave-up troughs denote a stratigraphic right-way-up orientation (Fig. 4b).

The top of the quartzite unit is interbedded with, and progressively grades into, a c. 6 km-thick pelitic unit that comprises predominantly kyanite–muscovite- to sillimanite–K-feldspar-grade schist. The lower part of the unit is characterized by thin-bedded quartzite intervals (Fig. 4c), whereas the top is interbedded with 2–5 cm-thick layers of arkosic psammite. The texture and grain size of the pelite changes dramatically across the sillimanite–K-feldspar isograd, where it becomes a pelitic K-feldspar augen gneiss with centimetre-sized porphyroblasts of K-feldspar (Fig. 4d). Due to its distinct appearance in the field, the K-feldspar augen gneiss is mapped as a separate unit from the underlying pelitic schist. However, the K-feldspar augen gneiss has a sedimentary protolith as evidenced by the presence of garnet porphyroblasts and abundant schlieren of fibrolitic sillimanite, which comprise 2–10 modal% (modal percentage) of the mineral assemblage (Dyck 2016). A thin (c. 50 m) discontinuous layer of diopside-rich calcareous marl overlies the K-feldspar augen gneiss unit (Fig. 3).

The marl and top of the K-feldspar augen gneiss unit are overlain by a c. 3.7 km-thick unit made up of rhythmically bedded sillimanite–K-feldspar-grade arkosic psammite. Bed thickness decreases upwards from c. 30–45 to 5–10 cm, with overall biotite, K-feldspar and sillimanite content increasing towards the top of the unit. The beds are characterized by reverse metamorphic grading (Fig. 4e) which shows an increase in biotite, K-feldspar and sillimanite content towards the top of each layer, with the contacts between individual beds both discrete and continuous. There is a gradual transition from the rhythmically bedded psammite unit to a micaceous semipelite with no discernible rhythmic bedding. The top of the micaceous semipelite unit grades into an irregularly layered, coarse-grained, pelitic K-feldspar augen gneiss, which forms the uppermost rocks of the GHS in Langtang. While not directly observed in Langtang, correlative successions in Everest and Shisha Pangma expose diopside-grade calc-silicates and marble that overlie the upper pelitic K-feldspar augen gneiss (Searle et al. 1997b, 2003, 2006).

Timing of Greater Himalayan Series deposition

Zircon grains were separated from samples of granodioritic augen gneiss (BD1438), quartzite (BD1441 and BD1437), pelitic K-feldspar augen gneiss (BD1215) and arkosic psammite (BD1435) for analysis with the sensitive high-resolution ion microprobe (SHRIMP) at the Geological Survey of Canada (Fig. 5; Table 1; see also the Supplementary material) in order to constrain the timing of GHS deposition in Langtang.

Fig. 5.
Fig. 5.

U–Pb zircon geochronology results. (a) Concordia diagram showing all analyses of zircon from sample BD1438. Ellipses are shaded to show the interpreted crystallization (light) and inherited (dark) populations. (b) Detrital zircon age spectra (probability density diagrams) from metasedimentary samples (BD1441, BD1437, BD1215 and BD1435) taken across the Greater Himalayan Series in the Langtang region. For ease of comparison of the youngest detrital modes, only results (206Pb/238U ages) less than 1200 Ma are plotted, Neoprotozoic and older detrital results (207Pb/206Pb ages) are listed. Age dates interpreted as metamorphic are shown as dots. Three discordant detrital analyses were excluded from the diagram (11247-68.1, 11247-75.1 and 11245-6.1: Table 1). Grey bands denote the subpopulations of detrital ages from the Arabian–Nubian Shield (ANS) reported by Avigad et al. (2003).

View this table:
Table 1.

U–Pb zircon geochronology data

Analytical methods

Zircon separates were cast in two separate epoxy mounts, along with fragments of primary and secondary standards. The mounts were polished to expose internal grain features prior to cathodoluminescence (CL) and backscattered electron (BSE) imaging of individual grains in order to characterize the internal features of crystals (e.g. zoning, internal microstructure and alteration). SHRIMP analyses followed the methods described in Stern (1997). Analytical details regarding spot size, beam current, data reduction protocol and U–Pb calibration are reported in Table 1. Isoplot (version 4.1: Ludwig 2003) was used to generate Concordia plots and mean ages with related statistics. Uncertainties given for ages of individual analyses are reported at the 2σ level unless otherwise noted (Fig. 5; Table 1). In order to account for the differences in sources of analytical error depending on age, the 206Pb/238U age (precision constrained by the Pb–U calibration) is reported/plotted for analyses younger than 1500 Ma, while the 207Pb/206Pb age is used for analyses older than 1500 Ma (precision constrained by counting statistics). Of these older analyses, those >10% discordant have not been included in the probability density plot (Fig. 5). Concordance was visually evaluated on Concordia diagrams (Supplementary material) for analyses younger than 1500 Ma.

Results

BD1438: granodioritic augen gneiss from the basement complex

The granodioritic augen gneiss contains light pink euhedral zircon with sharp crystal facets. Grains range in size from 100 to 200 µm, and comprise both smaller equant and larger prismatic morphologies. Under CL, most grains are oscillatory zoned; a few grains display a brighter homogeneous overgrowth. The 24 analyses were carried out on 22 individual zircon grains, which yielded ages that fall into two populations (Fig. 5a; Table 1). The older population consists of dominantly 1890–30 Ma zircon (weighted mean 207Pb/206Pb age of 1864 ± 9 Ma, n = 10, MSWD = 1.4), with a single older analysis at 2208 Ma. The younger group of results yielded a weighted mean 207Pb/206Pb age of 1738 ± 10 Ma (n = 13, MSWD = 0.86). Oscillatory zoned zircon grains are present in both age populations; however, analyses of bright overgrowths only fall in the younger grouping. Oscillatory zoned zircon typically represents magmatic growth and, consequently, the younger age is interpreted as the crystallization age of the granodiorite. The older ages, which in two instances (grains 11 and 124) occur as cores with c. 1740 Ma rim, are interpreted to represent inheritance. Field relationships are consistent with an unconformable contact between the rocks comprising the basement complex and the overlying quartzite, and thus this unit is inferred to represent the depositional basement to the GHS.

BD1441: quartzite, lower

A sample of quartzite was collected directly above the basement complex (Fig. 4b). Colourless to light pink zircon present a range of sizes (75–200 µm) and morphology, from small and round equant grains to long tabular crystals with rounded facets. Sixty grains from the lower quartzite were analysed, yielding ages that range from 2913 to 781 Ma (Fig. 5b; Table 1). The majority of analyses cluster between c. 1000 and 800 Ma (n = 49), with smaller modes at c. 2480 Ma (n = 9), c. 2010 Ma (n = 2), c. 1700 Ma (n = 3) and c. 1600 Ma (n = 3). The youngest distinct detrital zircon mode occurs at c. 815 Ma, which is considered the maximum age of deposition of the lower quartzite. Notably, there is only a very minor contribution of 1890–1740 Ma zircon, which is abundant in the underlying basement (Fig. 5b), suggesting non-local derivation for this quartzite.

BD1437: quartzite, upper

A second quartzite sample was collected from the top of the 400 m-thick unit. The zircon grains recovered from the quartzite are very similar in terms of morphology and zoning patterns to those from the lower quartzite. Forty-six analyses of 39 grains yielded ages that range from 2418 to 15.6 Ma. The majority of analyses cluster between c. 1000 and 800 Ma (n = 31), with smaller modes at c. 2450 Ma (n = 2) and c. 1800 Ma (n = 2) (Fig. 5b; Table 1). The youngest distinct, reproducible zircon mode occurs at c. 790 Ma, which is interpreted as a detrital zircon signature and, thus, the maximum age of deposition of the upper basal quartzite. Similarly, there is only a very minor contribution of 1890–1740 Ma zircon, common in the underlying basement (Fig. 5b); again suggesting non-local derivation for this quartzite. Analyses younger than 750 Ma are consistently determined from moderate CL overgrowths with low Th/U ratios (down to 0.002) that increase with progressively older ages. No dates between c. 750 and c. 15 Ma were reproduced. Overgrowths are quite complex and commonly quite thin (<15 µm), and therefore we infer that these younger analyses are the result of mixing due to overlap of the analytical spot across older detrital core domains and Miocene metamorphic rims. This interpretation is consistent with the progressive increase in Th/U that accompanies the increase in ages. As a single, non-reproducible result, age constraint on Miocene metamorphism is speculative (204Pb-corrected 206Pb/238U age = 15.6 ± 0.4 Ma); however, it should be noted that it overlaps with a reported age of metamorphic monazite growth in a nearby sample (16 ± 1 Ma: Kohn et al. 2005).

Both quartzite samples show strong similarities in age distribution and range with previously published spectra from quartzite across the GHS (DeCelles et al. 2000; Myrow et al. 2003, 2010; Spencer et al. 2012). Based on the prominence of 1100–900 Ma zircon, previous studies suggest the East African Orogen and Arabian–Nubian Shield as plausible sources for Himalayan detritus (DeCelles et al. 2000; Myrow et al. 2010; Spencer et al. 2013). Avigad et al. (2003) reported a detrital zircon age from a Cambrian sandstone that was deposited on the northern margin of the Arabian–Nubian Shield, which yielded subpopulations with ages of 2700–2470, 1850–1650, 1100–900 and c. 650–550 Ma (ANS in Fig. 5b). The overlap in the detrital ages presented here and the Arabian–Nubian Shield signature supports the existing source interpretation. Older grains may be attributed to the various Paleoproterozoic orogens and Archean crust interpreted to have been exposed on Gondwana (e.g. Myrow et al. 2010).

BD1215: pelitic K-feldspar augen gneiss

A sample taken from the top of the pelitic K-feldspar augen gneiss unit contains colourless zircon that range in size, from 50 to 150 µm, and morphology, from small and ovoid grains to large prismatic crystals with rounded facets. Forty-one analyses were carried out on 34 individual zircon grains, which yielded ages that range from 1069 ± 14 Ma to 26.4 ± 0.4 Ma. Detrital analyses cluster between c. 500 and 460 Ma (n = 12), with older ages between c. 1100 and 900 Ma (n = 3; pelitic K-feldspar (Kfs) augen gneiss: Fig. 5b; Table 1). Similar to the upper quartzite, younger analyses from CL dark rims exhibit low Th/U ratios (down to 0.008), which increase progressively with older age. None of the younger ages (c. 450–26 Ma) were reproducible. As with the upper quartzite, overgrowths are complex and often thin. We therefore interpret these younger analyses as a result of mixing due to overlap of the analytical spots across older detrital core domains and Miocene metamorphic rims. The youngest distinct detrital zircon mode, where the age of an oscillatory zoned grain was reproduced, is at c. 470 Ma, which is considered the maximum age of deposition of the protolith to the pelitic K-feldspar augen gneiss. In contrast to the underlying quartzite, there is only a very minor contribution of 1100–900 Ma zircon, and no contribution of 1890–1740 Ma zircon, which is present in the basement unit (Fig. 5b).

The modest detrital population is not sufficient for a statistically robust comparison between other published GHS detrital age spectra. However, it is worth noting the prominent upper Cambrian–Lower Ordovician peak matches detrital age spectra from the GHS in Garhwal (Spencer et al. 2012), Sikkim (Mottram et al. 2014) and Bhutan (Spencer et al. 2012). Moreover, these ages overlap with the age of the upper Cambrian–Lower Ordovician magmatic and associated tectonic event recorded elsewhere along the length of the Himalaya (e.g. Gansser 1983; Schärer & Allègre 1983; Gehrels et al. 2003; Pognante et al. 1990), and is known as the Kurgiakh (Srikantia 1981) or Bhimphedian (Cawood et al. 2007; Myrow et al. 2016; Palin et al. 2018) Orogeny. In contrast to Spencer et al. (2012), we interpret the upper Cambrian–Lower Ordovician grains as being detrital rather than igneous or inherited.

BD1435: arkosic psammite

The arkosic psammite unit was sampled immediately above the contact with the underlying pelitic K-feldspar augen gneiss. Only a small number of small (40–100 µm) ovoid, colourless zircon grains were recovered. Eight zircon grains were analysed, yielding ages ranging from 2850 to 240 Ma (arkosic psammite: Fig. 5b; Table 1). Given the limited zircon recovery, these results are not adequate to assess an age profile of the psammite. However, the zircon age spectrum of the arkosic psammite shares many similarities with the sample of pelitic K-feldspar augen gneiss (Fig. 5b). Notably, late Cambrian and Tonian–Stenian age peaks. The maximum age of deposition of the psammite is tentatively interpreted to be c. 500 Ma based on the analysis of a single, oscillatory zoned zircon.

Greater Himalayan Series protolith lithostratigraphy

While there is considerable variation in mineral assemblages and rock textures during prograde metamorphism of pelitic rocks (e.g. Weller et al. 2013), the bulk-rock composition of unmetamorphosed and high-grade pelite differs significantly only in volatile content (Shaw 1956). In the absence of large quantities of fluid, compositional mass transfer during regional metamorphism is restricted to the grain and micro-domain scale (Walther & Orville 1982). Consequently, sedimentary structures that are compositionally defined, such as cross-bedding, graded bedding, heavy-mineral lamination and changes in mudstone content in a turbidite succession, often remain identifiable following regional metamorphism (e.g. Walker 1988; Myers 2001). Likewise, strata that are compositionally homogeneous, such as reworked wave-dominated shoreline deposits, are expected to remain homogeneous on the micro-domain scale (e.g. McClelland et al. 1992). Throughout metamorphism, the most substantive textural evolution and mass transfer occur during anatexis, with the introduction of mobile partial melts (e.g. Waters 2001). The presence of partial melt increases the rate of intergranular diffusion, which enables grain-size coarsening and mass transfer between compositional micro-domains (Dell'Angelo & Tullis 1988; Evans et al. 2001). In cases where a partial melt may have migrated away from the metamorphic source rock, identifying peritectic minerals, such as sillimanite or cordierite, is useful in differentiating melt reactions and protolith compositions in rocks (e.g. Waters 1988).

Another useful criterion for distinguishing primary sedimentary units is the differentiation of event bedding v. deposits of environments with uniform lithofacies. Discrete strata formed by punctuated deposition events, such as storm deposits or turbidite flows (e.g. Immenhauser 2009; Talling et al. 2012), may well survive metamorphism without significant changes to composition or deposit morphology. Thus, despite penetrative fabric development in medium- to high-grade metasedimentary terranes, many aspects of a protolith lithostratigraphy can be determined with observation of rock type and distribution, variation in composition, and preserved primary features.

The amphibolite-grade sequence of GHS metasedimentary units in the Langtang area is interpreted as a right-way-up (at the regional scale) upper Neoproterozoic–Lower Ordovician cratonic margin lithostratigraphic succession (Fig. 6) that accumulated on Paleoproterozoic Indian crystalline basement (BD1438: Fig. 5a). The basement–cover contact is exposed at two structural levels: in-sequence at the base of the GHS, and as thick-skinned thrust imbricates within the MCT zone (e.g. Figs 2 & 3). We interpret the overlying succession as recording two deepening-upwards successions (Fig. 6). The maximum age of basal quartzite deposition (c. 815 Ma: BD1411 and BD1437) indicates that sedimentation along the northern Indian margin and deposition of the GHS began no earlier than the Tonian (early Neoproterozoic). The thinning of quartzite beds upsection and the gradual transition to a finely laminated and fine-grained quartzite- and pelite-dominated siliciclastic unit is interpreted to reflect an increase in water depth during transgression. The thin quartzite strata near the base of the overlying pelite unit are interpreted as the distal part of a transgressing continental margin siliciclastic shoreline succession (cf. McKie 1990). Overall, the pelite unit (BD1215) is devoid of event beds and is interpreted as deposits that accumulated in a distal shelf to upper continental slope basinal setting (cf. Faugères & Stow 2008). The rare calc-silicate horizons reported from Langtang (Reddy et al. 1993), and the presence of the thin unit of calcareous marl at the top of the augen gneiss unit, also support a distal shelf to upper slope depositional environment marked by intervals of carbonate precipitation within a deepening-upwards setting. The upper contact between the discontinuous carbonate unit and the overlying psammite is interpreted as representing the top of a highstand systems tract. Following this interpretation, the overlying interbedded arkosic psammites and pelites (BD1435), the former of which are interpreted as sandy turbidites preserving repeating graded bedding compositional profiles (cf. Mutti 1985), represent a transgressive systems tract. The thinning of the event beds and the overall increase in pelite content upsection marks a second deepening-upwards succession that culminated with an accumulation of carbonate strata recognized elsewhere along the GHS (Searle et al. 2003, 2006; Searle & Godin 2003; Parsons et al. 2016b). In contrast to the siliciclastic-dominated GHS in the Langtang area, the GHS in the Annapurna region is uncharacteristically thin (Parsons et al. 2016a) and comprises a carbonate-dominated succession (Colchen et al. 1986; Hodges et al. 1996; Searle & Godin 2003; Gleeson & Godin 2006). For comparison, the simplified ‘stratigraphy’ of the GHS in the Annapurna region is made up of five units: (1) interlayered quartzite, pelite and psammite (Unit I); (2) calcareous gneisses, marbles and psammitic schist (Unit II); (3) interlayered psammite, pelite and K-feldspar augen gneiss (Unit III); (4) micaceous marble (Unit IV); and (5) interlayered biotite phyllite and biotite–garnet schist (Unit V) (Searle & Godin 2003; Gleeson & Godin 2006; Parsons et al. 2016b).

Fig. 6.
Fig. 6.

(a) Schematic along-strike section from west to east across the Central Himalaya as interpreted for the Silurian, showing lateral variation in protolith lithostratigraphy, detrital zircon ages from the Langtang transect (this study), and the local unit name for the base orthogneiss unit in Annapurna and Everest. Solid lines represent stratigraphic boundaries and wavy lines represent erosional unconformities. The lithostratigraphic position of the Cenozoic South Tibetan Detachment (STD) zone and Main Central Thrust (MCT) zone are denoted by dashed white lines. The protolith lithostratigraphy of Langtang and Everest are interpreted as a siliciclastic depocentre. The thinner, carbonate-dominated stratigraphy in Annapurna is interpreted as a coeval carbonate shelf with units marked 1–5 (see the text for a description of each unit). (b) Alternative interpretation highlighting a potential along-strike correlation of the Everest Yellow Band with the carbonate unit in Langtang and the lower carbonate (Unit 2) in Annapurna. (c) Restored across-strike section for the Langtang transect, showing the in-sequence and upsection movement of the MCT and the relative position of the STD.

Using our lithostratigraphy for Langtang and the lithostratigraphic records reported from Annapurna (Colchen et al. 1986; Searle & Godin 2003; Gleeson & Godin 2006; Parsons et al. 2016b) and Everest (Searle et al. 2003, 2006; Myrow et al. 2009, 2016) allows an interpretation of the relationships (Fig. 6) between the GHS successions in the Annapurna, Langtang and Everest regions. We propose two possible along-strike correlations (Fig. 6) that differ primarily in how the age and depositional setting of the lower carbonate units of Annapurna, Langtang and Everest are interpreted. In our first correlation (Fig. 6a), the age of the Langtang carbonate unit is pinned by the ≤470 Ma age of the underlying K-feldspar augen gneiss. Whereas in our alternate interpretation, which is based largely on our interpretation of a deepening succession culminating in the deposition of the Langtang carbonate (Fig. 6b), we correlate the Yellow Band in Everest straight across to the lower carbonate in Annapurna (Unit 2). The simplicity of this second correlation (Fig. 6b), as well as the potential it offers for correlating a sequence boundary along the length of the orogen, is appealing. However, further detrital studies are warranted before the along-strike correlation presented here is applied to other regions of the Himalaya with confidence. The interlayered psammite, pelite and K-feldspar augen gneiss (Unit 3: Searle & Godin 2003; Gleeson & Godin 2006) are interpreted to correlate with the pelite and K-feldspar augen gneiss in the Langtang area (Fig. 2), whose protolith is suggested to be a distal shelf to basinal mudstone (Fig. 6). The sequence boundary in the Annapurna section is interpreted as the contact between the K-feldspar augen gneiss and the micaceous marble (Unit 3 and Unit 4: Searle & Godin 2003; Gleeson & Godin 2006), above which the depositional environment deepened and the strata are increasingly mudstone-dominated. Furthermore, it is possible that this sequence boundary corresponds to the 22–36 myr unconformity that is associated with the Kurgiakh (Bhimphedian) Orogeny (Cawood et al. 2007; Myrow et al. 2015, 2016).

A basin reconstruction that may partly explain the variation in the thickness of the GHS across the length of the orogen (e.g. Yin 2006; Prasad et al. 2011; Parsons et al. 2016c) involves a series of thick siliciclastic Tonian–Ordovician depocentres localized along the northern India and central (Langtang) Nepal continental margin. In this interpretation, the thinner and carbonate-dominated GHS in the Annapurna area may have formed as a shelf promontory analogous to the present-day Bahamas Bank (e.g. Burke 1988). A first-order comparison of the Langtang section with Annapurna suggests a similar sequence of initial transgression followed by a sequence boundary and a second deepening-upwards succession. Based on the continuity of mapped units (Figs 2 & 3) and the interpreted succession of metasedimentary protoliths, we suggest that the GHS represents a semi-coherent north-facing passive margin succession without any repetition of major map-scale units. Our interpretation of a semi-coherent succession is not incompatible with reported findings of both in-sequence and out-of-sequence thrusts within the GHS map units (e.g. Ambrose et al. 2015; Montomoli et al. 2015; Mukherjee 2015). Rather, we favour the interpretation that strain was, on the first order, largely homogeneous and distributed amongst all map units (Weinberg 2016).

Implications for Himalayan structure

Constraining the nature and contrasting origins of the lower ‘Ulleri’/‘Phaplu’ augen gneiss and the upper pelitic K-feldspar augen gneisses (‘Formation III augen gneiss’ or ‘Pumori K-feldspar augen gneiss’) is important for unravelling the fundamental structure of the Himalaya. The composition and zircon age profile of the lower augen gneiss in Langtang clearly documents an igneous protolith, whereas the presence of peritectic sillimanite and the characteristic pelitic augen gneiss microstructures outlined in Dyck (2016) indicate that the upper K-feldspar augen gneisses has a metasedimentary origin. Although upper Cambrian–Lower Ordovician granites are documented in the Kathmandu Klippe (Schärer & Allègre 1983; Gehrels et al. 2003) and in parts of the GHS in northern India (Foster 2000; Marquer et al. 2000; Catlos et al. 2002) and central Nepal (Stöcklin 1980; Le Fort 1986; Godin et al. 2001), based on the observations and data presented here we suggest that not all of the K-feldspar augen gneiss unit that is commonly mapped as the Formation III augen gneiss is the deformed equivalent of these granites. This interpretation is contrary to previous interpretations of the K-feldspar augen gneiss layer having an igneous origin (e.g. Le Fort 1986; Inger & Harris 1992; Reddy et al. 1993; Fraser et al. 2000; Searle et al. 2003; Mottram et al. 2014). This new interpretation simplifies the tectonic history of the GHS because it does not require the emplacement of a uniformly thick succession of Cambrian–Ordovician granites along the entire length of the northern Indian margin.

The protolith lithostratigraphy documented for the GHS in this paper provides additional support for viewing the stratigraphic evolution of the Himalaya as involving a protracted, north-facing, Indian cratonic margin, with the LHS, GHS and THS Himalaya corresponding to different structural domains rather than different stratigraphic sequences (Myrow et al. 2003, 2009; Searle & Godin 2003; Searle et al. 2003, 2008; Steck 2003). Furthermore, the stratigraphy of the northern Indian margin can be reconstructed without a need to include a suture or plate boundary (DeCelles et al. 2000) along the MCT zone. The imbricate thrust slices of Indian crystalline basement at the base of the GHS, and the exposed basement–cover contact, suggest that the MCT zone roots into the upper part of the Indian basement. Variations in the presence and thickness of Indian basement slices along the length of the orogen (e.g. Searle et al. 1997a, 2003; Searle & Godin 2003) may be influenced by the manner in which evolving thrust faults responded to heterogeneities in the underplating Indian basement (i.e. basement highs or ‘ridges’: Godin & Harris 2014). Repetition of the basal quartzite unit and pelite unit in the MCT zone indicates that thrusting was in-sequence and climbed upsection in the transport direction, placing the GHS above a proximal equivalent of the basinal pelite. The transition from homogeneous, ductile, layer-normal shortening (i.e. flattening) to brittle thrusting, as proposed by Godin et al. (2006), is precisely dated by the in situ monazite chronology of the MCT zone faults in Langtang (Kohn et al. 2005), yielding a Miocene age (16 ± 1 Ma) for thrust initiation. The timing of thrust initiation corresponds to the youngest crystallization ages of the Himalayan cordierite-bearing leucogranites (Searle et al. 2010), suggesting that orogen-wide strengthening of the metamorphic core during high-grade metamorphism allowed the GHS to act as semi-coherent package, which was exhumed by thick-skinned thrusting along the MCT zone (cf. Larson et al. 2015; Parsons et al. 2016a).

The preservation of a semi-coherent right-way-up cratonic margin lithostratigraphy in spite of subsequent Himalayan shortening and metamorphism necessitates that any model constructed to explain the exhumation of the Himalayan metamorphic core must maintain the overall protolith geometry of the GHS at the regional scale. It follows that the extent of mid-crustal extrusion in any crustal model that invokes ductile flow of the Himalayan core must be commensurate with preserving the first-order foreshortened basin geometry. To preserve the passive-margin stratigraphy, deformation across the GHS must have been predominately layer-normal shortening (i.e. flattening: Law et al. 2013; Long et al. 2016; Parsons et al. 2016c). Any significant amount of layer-parallel stretching (i.e. transport-parallel stretching) and out-of-sequence thrusting would have necessarily disrupted the protolith stratigraphic order.

The protolith lithostratigraphy presented here can be tested with further detailed stratigraphic studies of the GHS in other regions along the length of the Himalaya. It is easy to overlook protolith lithostratigraphy in metamorphic terranes, but there is much to be gained by exploring the crossroads of basin stratigraphy, metamorphic petrology and U–Pb geochronology.

Acknowledgements

We are grateful to T. Pestaj and B. Davis (GSC) for SHRIMP support, and P. Hunt (GSC) for SEM support. We are grateful to Paul Myrow, Andy Parsons, Nicholas Christie-Blick and an anonymous reviewer for insightful comments that have greatly improved this contribution.

Funding

This work was generously supported by the Clarendon Fund and the Geological Society of London's Mike Coward Fund (Funding agency ID: http://doi.org/10.13039/100008066) which were awarded to B. Dyck. M. St-Onge's travel and research were funded by the Earth Sciences Sector, Natural Resources Canada. This is Earth Science Sector contribution No. 20160157.

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