2. Geological setting

The TMC is located in the northwestern part of Indian Trans-Himalaya and is characterised by a sub-elliptical outline (Fig. 2a). It occurs as a northwesterly trending antiformal dome, separated from the Indus-Tsangpo suture including the Nidar ophiolite to the north and the Tetraogal Nappe to the south by detachment faults that dip away from the core of the dome. Further to the south, lies the Mata Nappe, which consists of two Early Paleozoic granite intrusions viz. Rupshu (ca. 482 Ma) and Nyimaling (ca. 460 Ma) (Epard & Steck, Reference Epard and Steck2008). The Tso Morari Gneiss forms the core of the TMC and is overlain by a metasedimentary sequence (Fig. 2a, (Buchs & Epard, Reference Buchs and Epard2019). The metabasites including eclogites occur as foliation-parallel boudins (Fig. 2b, c) within the ductilely deformed gneisses (Fig. 2d). Eclogites have been widely retrogressed to amphibolites (Fig. 2e). The tectono-thermal evolution of the TMC can be broadly classified into four stages: (i) prograde metamorphism during Neo-Tethys subduction, (ii) (U)HP metamorphism under eclogite-facies condition (>2 GPa, 500–760°C) at ca. 55 Ma, (iii) near-isothermal decompression under granulite-facies condition at the mid-crustal level at ca. 48–45 Ma, and, finally, (iv) retrograde amphibolite-facies conditions at ca. 31–29 Ma (Guillot et al. Reference Guillot, de Sigoyer, Lardeaux and Mascle1997; St-Onge et al. Reference St-Onge, Rayner, Palin, Searle and Waters2013; O’Brien, Reference O’Brien2019 and references therein). The Early Eocene Indo-Eurasia collision facilitated TMC extrusion along a northeasterly dipping channel (Dutta & Mukherjee, Reference Dutta and Mukherjee2021). At least three deformation phases affected the TMC during extrusion and produced the gently dipping gneissic foliations (de Sigoyer et al. Reference de Sigoyer, Guillot and Dick2004; Epard & Steck, Reference Epard and Steck2008).

Figure 2.

Study area, outcrop photographs and photomicrographs. (a) Geological map of the Tso Morari region (reproduced from Epard and Steck, Reference Epard and Steck2008). Boudins of (b) eclogite and (c) K-rich metabasite in the Tso Morari gneiss parallel to the gneissic foliation. This photo was clicked at the same location as that of Fig. 3 by St-Onge et al. (Reference St-Onge, Rayner, Palin, Searle and Waters2013). The large eclogite boudin is about 4 m thick at its thickest portion. (d) Top-to-the SE ductile shear sense exhibited by a feldspar augen in the quartzofeldspathic gneiss. (e) Retrograded Grt amphibolite. (f) Augen-shaped aggregates of quartz and feldspar within the fine-grained quartzofeldspathic matrix. The mica grains of the gneiss define the foliations in (g) and (h). (i) Prismatic zoisite grain within the coarser phengite grains of the gneiss. (j) Quartz porphyroclasts are common in some gneisses. Photomicrographs of the metabasites (k-m) and Grt amphibolite (n).

Gneisses and metabasites samples were collected for petrographic, geochemical and geochronogical analyses (see Supplementary Materials Table S1 for sample details). Recently, Pan et al. (Reference Pan, Macris and Menold2023) investigated the fluid evolution of the TMC since the Eocene, and they sampled a traverse from the centre of an eclogite boudin out into the host orthogneiss. In this study, we focus on the protoliths of the TMC, and thus samples were not collected from the reaction zone to avoid the effects due to the metasomatism. Quartz and calcite veins are found in the outcrops but fist-sized samples without the veins were selected for whole-rock analyses to avoid the effect of veins. Gneisses consist mainly of Ph + Bt + Pl + Qz ± Mc ± Grt ± Zo (Fig. 2f–j, mineral abbreviations are as per Whitney & Evans, Reference Whitney and Evans2010). They are subdivided into Ph-Bt gneiss, Grt-Ph gneiss, Ph-rich gneiss and quartzofeldspathic gneiss depending on the modal amount of minerals and presence of garnet. Most of the gneisses are orthogneisses whereas the quartzofeldspathic gneisses may be paragneisses in part. The metabasites include (retrograded) eclogites, Grt amphibolites and K-rich metabasites. Eclogite-facies metamorphism is characterised by the mineral assemblage of Grt + Omp + Rt + Coe/Qz ± Ph ± Zo (Fig. 2k–m) with later phases of Amp + Pl ±Di. The eclogites have zoning patterns in garnets associated with prograde and peak metamorphism but compositions and textures related to metasomatism stages are rare (e.g., St-Onge et al. Reference St-Onge, Rayner, Palin, Searle and Waters2013). The Grt amphibolites and K-rich metabasites mainly consist of Grt + Amp + Pl + Qz + Ilm + Ttn (Fig. 2n) and Grt + Bt + Ph + Amp + Pl + Qz + Rt + Zo ± Ttn, respectively.

3. Major and trace element geochemistry

The ten metabasite samples have low SiO₂ (44.09–51.20 wt%) and MgO (5.23–9.40 wt%), high TiO₂ (1.09–3.12 wt%) relative to FeO/MgO ratio and variable alkali contents (1.61–7.21 wt%) (Supplementary Materials Table S2 and Figure S1). Unlike the retrograded eclogites and K-rich metabasites, the Grt amphibolites and eclogite boulder have low alkali contents. The analysed samples plot on the field of subalkaline basalts in the Zr/TiO₂ vs. Nb/Y diagram, and the eclogite boulder has a low Nb/Y ratio (Supplementary Materials Figure S1).

The large ion lithophile elements (e.g., Rb and Ba) concentrations are variable (Fig. 3a, b) due to their mobility during surface weathering and alteration processes (Xia & Li, Reference Xia and Li2019). Thus it is important to check element mobility before applying tectonic discrimination diagrams (Polat & Hofmann, Reference Polat and Hofmann2003; Imayama et al. Reference Imayama, Oh, Jeon and Yi2021). The eclogites and amphibolites lack large Ce anomalies as shown by the Ce* values between 0.96 and 1.04 (Ce* = Ce_N/Sqrt(La_N×Pr_N), which is considered as immobile when the range is 0.9<Ce*<1.1. There is also no significant carbonisation or silicification due to the absence of carbonate and sulfide minerals in samples. These occurrences imply that the major chemical compositions of the sampled eclogites and amphibolites have not been strongly affected by hydrothermal alteration (Polat & Hofmann, Reference Polat and Hofmann2003; Imayama et al. Reference Imayama, Oh, Jeon and Yi2021).

Figure 3.

Spider diagrams of (a) retrograde eclogite and K-rich metabasite and (b) Grt amphibolite and eclogite boulder. Chondrite-normalised rare earth element patterns for (c) retrograde eclogite and K-rich metabasite and (d) Grt amphibolite and eclogite boulder. The compositions of metabasites from the TMC plotted in (e) Nb*2–Zr/4–Y (Meschede, Reference Meschede1986), (f) Zr/Y vs. Zr (Pearce & Norry, Reference Pearce and Norry1979), (g) Nb/La vs. Nb and (h) Y/Nb vs. Zr/Nb diagrams (Wilson, Reference Wilson1989, updated by Xia & Li, Reference Xia and Li2019). Data sources of the N-MORB, E-MORB, mantle plume and depleted asthenosphere compositions are from Sun & McDonough (Reference Sun and McDonough1989) and Salters and Stracke (Reference Salters and Stracke2004) with the PetDB database (http://www.earthchem.org/petdb). A(22): Ahmad et al. (Reference Ahmad, Bhat, Tanaka, Bickle, Asahara, Chapman and Sachan2022), J(19): Jonnalagadda et al. (Reference Jonnalagadda, Karmalkar and Duraiswami2019), R&R(06): Rao and Rai (Reference Rao and Rai2006). WPA: within-plate alkali basalts, WPT: within-plate tholeiites, VAB: volcanic-arc basalts.

All the metabasite samples are enriched in high field strength elements (HFSE) with Nb-Ta negative anomalies, and their chondrite-normalised rare earth element (REE) patterns plot in the fields between oceanic island basalt (OIB) and enriched mid-ocean ridge basalt (E-MORB, Fig. 3a–d). The HFSE and REE contents are lowest in the eclogite boulder along with prominent Nb and Zr–Hf negative anomalies. In the Nb–Zr–Y diagram, the metabasites, except for the eclogite boulder, plot in the fields of within-plate tholeiite or volcanic-arc basalts (Fig. 3e). In the Zr/Y–Zr diagram (Fig. 3f), these metabasites have high Zr/Y representing the within-plate basalts and are thus distinguished from the island-arc origin. In contrast, the eclogite boulder plots in the fields where island-arc basalt and MORB overlap on both diagrams (Fig. 3e, f).

4. Zircon U-Pb geochronology

We separated the zircons from eight gneisses and one Grt amphibolite (Fig. 2a, Supplementary Materials Table S1) to obtain cathodoluminescence (CL) images and U–Pb ages (see Supplementary Materials Text S1). The CL images of the zircon grains from the gneisses show well-developed prismatic faces and internal oscillatory zoning, which is typical igneous-type zircons (Fig. 4). Some grains have modified grey-CL rims showing faint oscillatory zoning or patchy pattern of dark and bright colours in CL-image (Supplementary Materials Figure S2). Thin metamorphic rims showing dark-CL domains are observed in 19-2 and 19-4, but they are too thin to analyse. The U–Pb analyses of two Ph-Bt gneisses (15-3B and 9g) concentrated at a single population of concordant to near-concordant grains with a weighted mean ²⁰⁶Pb/²³⁸U date of 481.2±4.4 Ma (n = 14, mean squared weighted deviation or MSWD = 2.5, Fig. 4a) and 484.3±5.7 Ma (n = 14, MSWD = 2.1, Supplementary Materials Figure S3a), respectively. Some grains were possibly affected by Pb-loss resulting in young ²⁰⁶Pb/²³⁸U ages. The inherited grains yielded scattered dates ranging from 1402 Ma to 606 Ma. The zircons from a Grt-Ph gneiss (18-3B) yielded ²⁰⁶Pb/²³⁸U dates ranging from 500 to 470 Ma, but the weighted mean date was not calculated due to the large MSWD value (Supplementary Materials Figure S3b). The U–Pb analyses of concordant grains from the other two Grt-Ph gneisses (16-2A and 18-4) yielded the weighted mean ²⁰⁶Pb/²³⁸U date of 485.9±3.3 Ma (n = 12, MSWD = 1.7, Fig. 4b) and 493.4±5.2 Ma (n = 13, MSWD = 1.6, Supplementary Materials Figure S3c), respectively. Some inherited grains observed in three Grt-Ph gneisses yield ²⁰⁶Pb/²³⁸U dates of 1841 to 796 Ma. The igneous zircons of 15 concordant grains from Ph-rich gneiss (18-5B) yielded the weighted mean ²⁰⁶Pb/²³⁸U date of 493.3±5.4 Ma (n = 15, MSWD = 2.1, Supplementary Materials Figure S3d). The inherited domains gave ²⁰⁶Pb/²³⁸U dates of 1154 to 766 Ma. Some grains from two quartzofeldspathic gneisses (19-2 and 19-4) show discordant behaviour due to the disturbance of common Pb and Pb loss. The U–Pb analyses of the crystals that yielded concordant dates gave weighted mean ²⁰⁶Pb/²³⁸U dates of 480.4±4.4 Ma (n = 18, MSWD = 1.8, Fig. 4c) and 485.6±4.5 Ma (n = 14, MSWD = 3.0, Supplementary Materials Figure S3e), respectively.

Figure 4.

Concordia diagrams from the SHRIMP zircon U–Pb analyses of (a) Ph-Bt gneiss (15-3B), (b) Grt-Ph gneiss (16-2A), (c) Quartzofeldspathic gneiss (19-2) and (d) Grt amphibolite (20-2) from the TMC. All error ellipses and weighted mean ²⁰⁶Pb/²³⁸U ages are quoted at the 1σ and 2σ levels, respectively. The white bars below the zircons are 20 μm long.

The zircon grains from Grt amphibolite (20-2) show well-developed prismatic faces and show bright-CL oscillatory zoning surrounding the inherited core (Fig. 4d). The igneous zircons of 13 concordant grains yielded the weighted mean ²⁰⁶Pb/²³⁸U date of 475.8±6.9 Ma (MSWD = 2.6, Fig. 4d). The inherited domains gave ²⁰⁶Pb/²³⁸U dates of 1039 to 813 Ma.

5. Discussion

Previous workers suggested that the protoliths of the Tso Morari eclogites are either the Permian Panjal basalts associated with a mantle plume and crustal contamination (Spencer et al. Reference Spencer, Tonarini and Pognante1995; de Sigoyer et al. Reference de Sigoyer, Guillot and Dick2004; Jonnalagadda et al. Reference Jonnalagadda, Karmalkar and Duraiswami2019) or the Late Jurassic to Early Cretaceous Ladakh ophiolitic mafic rocks in the supra-subduction zone (Ahmad et al. Reference Ahmad, Bhat, Tanaka, Bickle, Asahara, Chapman and Sachan2022). The metabasites we analysed have enriched light REE contents (Fig. 3c,d), and the mixing line for the lithospheric contamination in the Y/Nb–Zr/Nb diagram (Fig. 3h) starts from the mantle plume source, rather than the depleted MORB on the mantle array, precluding the possibility of N-MORB components in the mantle source. Although the enriched light REE patterns are comparable to both E-MORB and OIB, the concentrations of incompatible trace elements such as the HFSE are much higher than those of the E-MORB and resemble those of OIB (Fig. 3a, b). The signatures of within-plate basalts are marked by the high Zr/Y content in most metabasites (Fig. 3f). Although the Nb-Ta negative anomalies (Fig. 3a, b) and Nb/La<1 (Fig. 3g) in most of the TMC metabasites imply either continental intraplate basalts that experienced crustal contamination or island-arc basalts as their protoliths, the former is supported by high concentrations of incompatible trace elements (Xia and Li, Reference Xia and Li2019). Crustal contamination leads to low Nb content, increasing Y/Nb and Zr/Nb ratios (Fig. 3h). The trend has a different slope from that of the mantle array line from the plume, through MORB, to the depleted asthenosphere. Based on field occurrences, the contaminant lithology could partially consist of the Early Paleozoic orthogneisses surrounding metabasites. The Zr/Nb ratios from a few orthogneisses reach up to 30 (Ahmad et al. Reference Ahmad, Bhat, Tanaka, Bickle, Asahara, Chapman and Sachan2022), which exceed the average Zr/Nb ratios of around 10 for continental crust compositions (e.g., Wedepohl Reference Wedepohl1995; Rudnick & Gao, Reference Rudnick, Gao, Holland and Turekian2003). The crustal contamination is also supported by the presence of the inherited zircon cores in the garnet amphibolite (sample 20-2). Therefore, the TMC metabasites most likely formed in a continental rift setting and experienced lithospheric contamination. In contrast, the eclogite boulder with prominent Nb and Zr–Hf negative anomalies suggests its island-arc origin, which could correspond to the Nidar ophiolite mafic rocks formed by the subduction initiation in the forearc (Ahmad et al. Reference Ahmad, Bhat, Tanaka, Bickle, Asahara, Chapman and Sachan2022).

The U–Pb SHRIMP data from the eight gneisses suggest an Early Paleozoic (ca. 493–480 Ma) protolith, which is consistent with the previously reported U–Pb zircon date of 479±2 Ma from the Tso Morari Gneiss (Girard & Bussy, Reference Girard and Bussy1999). The youngest zircon date of 475.8±6.9 Ma was obtained from the Grt amphibolite, which represents a retrograde product of the eclogites. Oscillatory zoning is a common feature in zircons from felsic igneous rocks, but some zircon textures in amphibolites from the orogenic belts also include oscillatory zoning (Oh et al. Reference Oh, Imayama, Jeon and Yi2017; Kang et al. Reference Kang, Li, Kang, Dong, Jiang, Liang and Dong2020), representing the timing of mafic magmatism. It is unlikely that contamination of zircons from orthogneisses occurred during intrusion because the garnet amphibolite occurs within the paragneisses, not orthogneisses. The zircons may have crystalised from basaltic magmas if the rocks displayed high Zr contents (Shao et al. Reference Shao, Xia, Ding, Cai and Song2019). Therefore, the ca. 476 Ma zircon date from garnet amphibolite is interpreted as the timing of mafic magmatism and is roughly consistent with those of the gneisses within error, indicating the bimodal magmatism during the Early Paleozoic at least for some time. In terms of trace element geochemistry, continental basalt is similar to the OIB but the presence of coherent granites implies the tectonic setting of continental rift. These Early Paleozoic ages indicate that the TMC eclogites were not derived from the Permian Panjal Traps associated with the opening of the Neo-Tethys Ocean. The crystallisation ages (ca. 493 Ma) from the central part of the TMC near Puga are slightly older than those (ca. 486–476 Ma) in surrounding parts, implying multi-stage magmatism. Felsic rocks volumetrically dominate the TMC, whereas the volume of mafic rocks is quite small. This is also in contrast to what is observed in the NW Himalaya where the Panjal Traps are characterised by abundant mafic magmatism (Shellnutt, Reference Shellnutt2016). We alternatively propose that the protoliths of the UHP TMC formed in the extensional setting at the rifted Indian margin during the Early Paleozoic. During the Pan-African orogeny, the Indian shield was located in the northern East Gondwana (Gray et al. Reference Gray, Foster, Meert, Goscombe, Armstrong, Trouw and Passchier2008, Fig. 5). Epard and Steck (Reference Epard and Steck2008) pointed out that the sequences deposited in the north Indian margin in the Tso Morari region are affected by extensional structures, rather than the compressional structures of the Pan-African orogeny. The Early Paleozoic tectonic events are widely known in the Himalayas (Le Fort & Cronin, Reference Le Fort and Cronin1988; Gehrels et al. Reference Gehrels, DeCelles, Ojha and Upreti2006), but the investigation of the detailed tectonic evolution is beyond the scope of this study as here we only discuss the Early Paleozoic igneous events in the NW Himalaya. The report of A-type alkaline Kaghan metagranite of ca. 470 Ma (Trivedi et al. Reference Trivedi, Sharma and Gopalan1986) also supports the Early Paleozoic rifting at the passive margin of northern India. The Early Paleozoic felsic magmatism is also exposed in Pakistan (483–476 Ma, Ogasawara et al. Reference Ogasawara, Fukuyama, Siddiqui and Zhao2019; ca. 459 Ma, Mughal et al. Reference Mughal, Zhang, Hussain, Rehman, Du and Hameed2022) and in the Garhwal, NW India (495–490 Ma, Imayama et al. Reference Imayama, Bose, Yi, Jeong, Horie, Takehara and Kawabata2023; ca. 512 Ma, Sen et al. Reference Sen, Sen, Chatterjee, Choudhary and Dey2021), which are considered as S-type and A-type, respectively. Thus, the A-type Early Paleozoic granites in the Garhwal are more closely related to the Early Paleozoic rifting events in this study. Coupled with geochemical results in this study, we posit that the Early Paleozoic rifting after the Neoproterozoic Pan-African orogeny produced the protoliths of the TMC. Recently, a close relationship between North India and South China during the Neoproterozoic to Early Paleozoic has been discussed (Qi et al. Reference Qi, Cawood, Xu, Du, Zhang and Zhang2020; Imayama et al. Reference Imayama, Bose, Yi, Jeong, Horie, Takehara and Kawabata2023), and the Early Paleozoic rifting in NW India led to the separation of South China from North India (Fig. 5, Imayama et al. Reference Imayama, Bose, Yi, Jeong, Horie, Takehara and Kawabata2023). However, more study is required to confirm this assumption.

Figure 5.

Map of Gondwana showing the positions of the Indian Shield and the rifting event inferred from this study. Modified from Gray et al. (Reference Gray, Foster, Meert, Goscombe, Armstrong, Trouw and Passchier2008), Meert and Lieberman (Reference Meert and Lieberman2008) and Imayama et al. (Reference Imayama, Bose, Yi, Jeong, Horie, Takehara and Kawabata2023). SF, São Francisco Craton; RP, Río de la Plata Craton.

Combined with the UHP Nanga Parbat eclogites (Rehman et al. Reference Rehman, Lee, Chung, Khan, O’Brien and Yamamoto2016), our results suggest multiple sources and ages for the protolith of the UHP metamorphic rocks in the NW Himalaya. Similarly, the protolith of the HP metamorphic rocks in the central Himalaya also originate from multiple sources and ages (Zhang et al. Reference Zhang, Wang, Webb, Zhang, Liu, Fu, Wu and Wang2022). The HP eclogites of the Ama Drime Massif originated from the Ordovician (∼480–430 Ma) rift-related or E-MORB-like magmatism (Wang et al. Reference Wang, Zhang, Li and Somerville2017; Dong et al. Reference Dong, Zhang, Tian, Niu and Zhang2022) and the Paleoproterozoic (∼1850 Ma) continental flood basalts (Zhang et al. Reference Zhang, Wang, Webb, Zhang, Liu, Fu, Wu and Wang2022). In Bhutan, the HP metabasite was derived from young Paleoproteorozoic rifting (∼1742 Ma, Chakungal et al. Reference Chakungal, Dostal, Grujic, Duchêne and Ghalley2010). These occurrences indicate the sources and ages of the protoliths of the Himalayan HP–UHP metamorphic rocks are diverse in the NW-SE direction along the orogen.