The Geodynamic Significance of Continental UHP Exhumation: New Constraints From the Tso Morari Complex, NW Himalaya

The burial and exhumation of continental crust to and from ultrahigh‐pressure (UHP) is an important orogenic process, often interpreted with respect to the onset and/or subduction dynamics of continent‐continent collision. Here, we investigate the timing and significance of UHP metamorphism and exhumation of the Tso Morari complex, North‐West Himalaya. We present new petrochronological analyses of mafic eclogites and their host‐rock gneisses, combining U‐Pb zircon, rutile and xenotime geochronology (high‐precision CA‐ID‐TIMS and high‐spatial resolution LA‐ICP‐MS), garnet element maps, and petrographic observations. Zircon from mafic eclogite have a CA‐ID‐TIMS age of 46.91 ± 0.07 Ma, with REE profiles indicative of growth at eclogite facies conditions. Those ages overlap with zircon rim ages (48.9 ± 1.2 Ma, LA‐ICP‐MS) and xenotime ages (47.4 ± 1.4 Ma; LA‐ICP‐MS) from the hosting Puga gneiss, which grew during breakdown of UHP garnet rims. We argue that peak zircon growth at 47–46 Ma corresponds to the onset of exhumation from UHP conditions. Subsequent exhumation through the rutile closure temperature, is constrained by new dates of 40.4 ± 1.7 and 36.3 ± 3.8 Ma (LA‐ICP‐MS). Overlapping ages from Kaghan imply a coeval time‐frame for the onset of UHP exhumation across the NW Himalaya. Furthermore, our regional synthesis demonstrates a causative link between changes in the subduction dynamics of the India‐Asia collision zone at 47–46 Ma and the resulting mid‐Eocene plate network reorganization. The onset of UHP exhumation therefore provides a tightly constrained time‐stamp significant geodynamic shifts within the orogen and wider plate network.


Introduction
The exhumation of continental crust from ultra-high pressure (hereafter referred to as "UHP exhumation") is an important yet poorly constrained orogenic process (Guillot et al., 2009;Hacker & Gerya, 2013;Warren, 2013).Many studies relate UHP exhumation to other orogenic processes such as a reduction in the slab-pull force during the subduction of continental crust, slab break-off, and/or onset of various modes of orogenic extension/collapse (e.g., Brun & Faccenna, 2008;Burov et al., 2014;Chen et al., 2022;Guillot et al., 2009;Hacker et al., 2010;Little et al., 2011;Yamato et al., 2008).As such, UHP exhumation in the rock record is often interpreted to signify an important shift in the tectonic and geodynamic regime of an orogen (e.g., Chen et al., 2022;Guillot et al., 2009;Hacker et al., 2010;O'Brien et al., 2001;Soret et al., 2021;Yamato et al., 2008), However, the exact relationships between continental UHP exhumation and the geodynamics of orogenesis and plate tectonics, as well as the mechanisms that facilitate this process, remain poorly constrained due to the difficulty of investigating UHP metamorphism and exhumation from the bedrock record (Hacker & Gerya, 2013;O'Brien, 2019;Warren, 2013).To better understand the processes of continental UHP exhumation and its impacts on orogenesis and plate tectonics, we must combine petrology-based pressure-temperature (P-T) pathways with radioisotope geochronological constraints.This approach allows us to determine the timing and rates of UHP exhumation through P-T space and with respect to the orogeny.However, such attempts to constrain the timing and rates of UHP exhumation are always complicated by the uncertainty with which accessory phase ages are related to P-T paths (e.g., Kohn et al., 2017;O'Brien, 2019a).This problem is clearly demonstrated in the UHP Tso Morari Complex of the NW Himalaya (Epard & Steck, 2008;Guillot et al., 2008;Steck et al., 1998), which forms the central study region of our investigation.
The Tso Morari Complex is one of two UHP terranes in the NW Himalaya, the other being the Kaghan Valley Complex, located ∼450 km to the west of Tso Morari (Buchs & Epard, 2019;Epard & Steck, 2008;Guillot et al., 2008;O'Brien et al., 2001;Parrish et al., 2006;Steck et al., 1998).Over the last three decades, geochronological studies of Tso Morari have yielded a range of estimates for the timing of UHP metamorphism and exhumation, spanning a period of 11 Ma (Donaldson et al., 2013;Leech et al., 2007;St-Onge et al., 2013).Early interpretations were based on a limited number of spot analyses of zircon (e.g., n = 19, Leech et al., 2007), whereas the later publication by Donaldson et al. (2013) provided a more robust record (n = 108) with a zircon age distribution ranging from 53 to 37 Ma, with a peak at 47-43 Ma that provied a substantially narrowed estimate of peak UHP metamorphism.In contrast, estimates from the Kaghan UHP eclogites are tightly constrained to ∼46 Ma (Kaneko et al., 2003;Parrish et al., 2006;Zhang et al., 2022).From these constraints, a diverse but poorly constrained set of models have been proposed for the exhumation of Tso Morari and Kaghan Valley complexes (Boutelier & Cruden, 2018;Kylander-Clark et al., 2008;Möller et al., 2015;Schwartz et al., 2007).
The spread of ages for UHP metamorphism and exhumation in Tso Morari, could, reflect differences in analytical techniques and their inaccuracies, but a likely cause is also the difficulty and uncertainty associated with linking geochronological data with independently constrained metamorphic petrology (Foster et al., 2004;Kohn et al., 2017;O'Brien, 2019a).This problem is compounded further in the Tso Morari complex because UHP rocks exist as meter-scale eclogite facies mafic pods hosted within amphibolite facies felsic gneiss.Accessory phases suitable for geochronology are rare in eclogite pods, so some studies have attempted to relate the P-T evolution of eclogite pods with geochronological constraints from the amphibolite facies gneiss (St-Onge et al., 2013).However, this approach carries additional uncertainties surrounding the structural and metamorphic relationships between the eclogite pods and the amphibolite facies gneiss in which they are hosted (O'Brien, 2019b).
To overcome these problems, estimates of the timing of UHP exhumation are better resolved using a combination of modern petrochronology techniques (Kohn et al., 2017), and where mineral size and zonation allows, high precision techniques (Parrish et al., 2006).The former can utilize a range of approaches, including: (a) the combination of geochronology, trace element geochemistry and metamorphic petrology, to quantitatively relate precise ages to specific stages on a metamorphic P-T path (e.g., Rubatto, 2002;Rubatto & Hermann, 2003); and (b) identifying and selecting a range of accessory phases associated with different metamorphic assemblages and lithologies, which collectively span a wide range of closure temperatures (Lotout et al., 2018;Regis et al., 2016;Tual et al., 2022).Such techniques provide the best opportunity to accurately and precisely constrain the timing of UHP exhumation.
In this study, we employ a range of petrochronological techniques to precisely constrain the timing of UHP metamorphism and exhumation of the Tso Morari Complex.We use detailed petrographic analyses, including major and trace element X-ray maps of garnet, to identify and relate prograde, peak, and retrograde metamorphic assemblages in mafic eclogite samples and amphibolite facies felsic gneiss samples (the Puga gneiss).This allows us to select a variety of accessory phases from different metamorphic assemblages for U-Pb geochronology, to constrain the timing of metamorphism at different points of the P-T path.Using laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS), we analyze zircon and rutile from an eclogite pod sample, and zircon, rutile, and xenotime from a felsic gneiss sample to constrain the timing of prograde, peak, and retrograde metamorphism.In addition, we analyzed zircon from the eclogite using chemical abrasion isotope dilution thermal ion mass spectrometry (CA-ID-TIMS) to more precisely constrain the timing of zircon crystallization.
Our results suggest that the onset of exhumation of Indian continental crust from UHP conditions occurred synchronously across the Tso Morari complex and Kaghan Valley complex at ∼47-46 Ma.

The Geodynamic Significance of UHP Exhumation
After presenting our new geochronology data and interpreting its implications for the metamorphic evolution of the Tso Morari complex, we then consider the geodynamic implications of our new constraints for the Himalayan orogeny and the wider tectonic plate network.The early stages of continental collision involve a shift in the geodynamics of subduction due to the positive buoyancy of continental crust on the down-going plate (Afonso & Zlotnik, 2011;Capitanio et al., 2015;Hacker & Gerya, 2013;Laik et al., 2022;Warren, 2013).The exhumation of continental crust from UHP conditions provides an important marker of this geodynamic shift, due to the association of UHP exhumation with modification(s) to the subduction dynamics of a collision zone such as slab break-off, a reduction in slab dip, and/or an increase in mechanical coupling between the upper and lower plates (e.g., Afonso & Zlotnik, 2011;Chen et al., 2022;Guillot et al., 2008;Warren, 2013).The timing of UHP exhumation is especially important, as this constraint allows scientists to interpret the structural, metamorphic, and magmatic evolution of orogenesis with respect to changes in the subduction dynamics of a collision zone and vice-versa, offering an invaluable opportunity to test geodynamic and plate tectonic models of continental collision against bedrock constraints (e.g., Capitanio et al., 2015;Chen et al., 2022;Guillot et al., 2008;Yamato et al., 2008).
More broadly, the timing of continental UHP exhumation can be used to better understand the impacts of continental collision on tectonic plate networks.Continental collisions change the summative slab pull forces within their encompassing plate network, either by reducing slab pull of the collision zone during subduction of buoyant continental lithosphere or by removing slab pull from the collision zone altogether via slab break off (e.g., Afonso & Zlotnik, 2011;Capitanio et al., 2015;Chen et al., 2022;Laik et al., 2022).In some cases, this can trigger a plate network reorganization event, defined by a short period (∼10-20 Myr) of rapid and widespread modifications to the speed and direction of tectonic plate motions and the topology of plate boundaries (e.g., Gibbons et al., 2015;Gürer et al., 2022;Matthews et al., 2012;Müller et al., 2016).Understanding the causes and occurrences of these reorganizations is important as they can drastically alter the course of Earth's paleogeographic and tectonic evolution in a relatively short period of time.
Our understanding of plate network reorganizations is limited due to the difficulty of robustly demonstrating a causative relationship between temporally overlapping, but spatially disparate tectonic events and coeval changes in plate kinematics.This is particularly problematic when investigating reorganizations associated with continental collision, due to the complexity of their bedrock record and the uncertainties of geochronological constraints that obfuscate the duration and order of localized events and their geodynamic relationships with networkwide modifications of plate kinematics.Faced with this problem, constraints such as the timing of UHP exhumation, which directly correspond to changes in subduction dynamics of a collision zone, can provide crucial geochronological markers for constraining causative relationships between collisions and plate network reorganizations (e.g., Gibbons et al., 2015;Matthews et al., 2012).
A major plate network reorganization for the Indian, Eurasian, African, Australian, and Antarctic plates is recognized by previous studies during the Eocene (Gibbons et al., 2015).This reorganization is arguably associated with the Himalayan orogeny (Gibbons et al., 2015;Zahirovic et al., 2016), but uncertainty in the collision history of the orogeny (e.g., Hu et al., 2016;Kapp & DeCelles, 2019;Parsons et al., 2020;van Hinsbergen et al., 2019) means that a precise and unambiguous timestamp linking the bedrock record of the orogen with the Eocene plate network reorganization has been lacking.It is therefore unclear exactly how the structural and metamorphic evolution of the Himalayan orogeny relates to and/or caused plate kinematic changes during the Eocene reorganization.Our new, precise geochronometric constraints for the UHP exhumation in the NW Himalaya present a valuable opportunity to evaluate the temporal relationships between the Himalayan orogeny and the Eocene plate network reorganization and consider their causative links.With the modern advent of the user-friendly GPlates plate reconstruction application (Müller et al., 2018), similar integrated analyses of collisional metamorphism and plate kinematics could be used to better understand the plate tectonic consequences of other orogenies.

The Himalayan Orogeny and the India-Asia Collision: Definitions
In order to interpret our data with respect to the Himalayan orogeny and India-Asia collision, it is necessary to outline some definitions used hereafter.Working hypotheses for tectonic formation of the Himalayan orogeny and the India-Asia collision can be broadly categorized into "Single Collision" and "Double Collision" models (e.g., Hu et al., 2016;Kapp & DeCelles, 2019;Parsons et al., 2020).Single Collision models propose a single continental collision between India and Eurasia, beginning at ∼61 Ma (An et al., 2021) with convergence and under thrusting continuing to the present day (e.g., Gansser, 1966;Hu et al., 2016;Ingalls et al., 2016;Le Fort, 1975).Such models are not considered tenable as they require extreme volumes of continental subduction and do not adequately explain the significant kinematic and geodynamic changes which occur within the orogen and the surrounding plate network between 50 and 40 Ma (e.g., Parsons et al., 2020, 2021;van Hinsbergen et al., 2019).As such, the results of this study are interpreted in the context of the competing "Double Collision" models, for which two alternative working hypotheses exist, differing with respect to the nature of the first collision.In both hypotheses, "first collision" began at ∼61 Ma (e.g., An et al., 2021;Hu et al., 2015) but corresponds, to either (a) collision of the Indian continent with an equatorial Neotethys intra-oceanic arc (e.g., Bouilhol et al., 2013;Burg & Bouilhol, 2019;Patriat & Achache, 1984;Replumaz et al., 2014;Stampfli & Borel, 2004);or (b) collision between an India-derived microcontinent and the Eurasian active margin (e.g., Sinha Roy, 1976;van Hinsbergen et al., 2019;Zhou & Su, 2019)."Second collision" occurred between the Indian continent and the Eurasian active margin sometime between 50 and 25 Ma (e.g., Burg & Bouilhol, 2019;Parsons et al., 2021;Patriat & Achache, 1984;Replumaz et al., 2014;Searle, 2019;van Hinsbergen et al., 2019).
In the context of these double collision models, the Himalayan orogeny corresponds to the deformation of Indian continental crustwhich initiated during first collision and continued through second collision.In contrast, the "India-Asia collision" sensu-stricto corresponds to the second collision event only (e.g., Parsons et al., 2020Parsons et al., , 2021)).Debate continues to surround the relative validity of the two "double collision" hypotheses and further considerations can be found in recent reviews (e.g., Kapp & DeCelles, 2019;Parsons et al., 2020;Searle, 2019;van Hinsbergen et al., 2019); our study is presented in the context of both double collision models (e.g., Burg & Bouilhol, 2019vs. van Hinsbergen et al., 2019) as it is beyond the scope of the new data presented in this paper to address their relative validity.

Geology of the Tso Morari Complex
The Tso Morari dome is situated on the north-western margin of the Indian plate (Figure 1), and is separated from the Ladakh batholith of the Asian plate located to the north by the Indus suture zone (Fuchs & Linner, 1996).The Tso Morari dome comprises a set of stacked nappes, folded into a northwest-southeast trending periclinal antiform (Buchs & Epard, 2019;Steck et al., 1998).The structurally lowermost nappe is the Tso Morari Complex (also known as the Tso Morari Gneiss, Epard & Steck, 2008), which crops out in the center of the dome.The Tso Morari Complex contains Ordovician granite and disrupted mafic dykes and sills intruded into Cambrian sediments of the Indian continent.Granitic rocks are variably deformed, resulting in an array of undeformed metagranites, augen gneiss and garnet-mica-schists within the Tso Morari Complex.The range of deformation states preserves different parts of the P-T-t history of the Tso Morari Complex.The early subduction-related history is rarely preserved except as thin corona textures in low strain metagranite (Bidgood et al., 2023), whereas the high strain gneisses are often overprinted by later, amphibolite facies metamorphism and exhumation-related deformation.These end-member states are distinguished locally as the Polokongka La granite and the Puga Gneiss, which share the same granitic protolith, and differ only in their state of strain and metamorphic evolution (Girard & Bussy, 1999).
This study describes the record of subduction-related major and accessory mineral growth preserved in a rare garnet-bearing Puga Gneiss which records a part of the P-T-t history not preserved elsewhere.Additionally, new data is presented which tightly constrains zircon growth in the mafic eclogites, which, in conjunction with the observations in the Puga Gneiss, allows us to place the petrographic observations in a geochronologic and geodynamic context.

Geochronology of the Tso Morari Complex
The range of estimates for UHP metamorphism and exhumation in the Tso Morari complex (54-44.5 Ma, Donaldson et al., 2013;Leech et al., 2007;St-Onge et al., 2013) is less tightly resolved than the age estimates for Kaghan (∼46 Ma, Kaneko et al., 2003;Parrish et al., 2006;Zhang et al., 2022).Given the differences in analytical techniques used to date the Tso Morari complex and the difficulty in relating these dates to metamorphic stages, it remains unclear to what degree the spread of ages from Tso Morari can be assigned to geological heterogeneity or to analytical uncertainty and artifacts.For example, the absence of petrographic context in some studies makes it difficult to interpret the age clusters with respect to specific metamorphic conditions (see O'Brien, 2006 for further discussion).Additionally, the link between zircon growth and fluid infiltration (e.g., Chen & Zheng, 2017;Chen et al., 2010;Kohn et al., 2015;Sheng et al., 2012;Skuzovatov et al., 2021) has not been explored in the context of the published data.Importantly, given the types of data yielded from Tso Morari thus far (see following summary), it is not possible to robustly constrain whether zircon growth in the mafic eclogites was prolonged or punctuated, and whether the range of dates correspond to prograde, peak and/or retrograde metamorphic conditions.Given the narrow age range for UHP metamorphism from Kaghan only 450 km away, comparative age estimates from Tso Morari would have significant implications for the processes of continental subduction and UHP exhumation; as such, this is the primary motivation for this study.Here we present new and precise ages of eclogite-facies metamorphism that can be compared with the ages recorded at Kaghan, along with new in situ data which, when combined with the petrographic interpretation, records an additional petrographic age and stage of the evolution of the Tso Morari complex that has not been previously observed.
The Puga Gneiss and Polokongka La granite contain zircon with thin metamorphic rims surrounding igneous cores, with the latter dated by U-Pb at 479 ± 1 Ma (Girard & Bussy, 1999;Leech et al., 2007).Initial estimates of the age of high pressure metamorphism by De Sigoyer et al. (2000) used Lu-Hf isochron ages on garnetclinopyroxene-whole-rock data, Sm-Nd on garnet-glaucophane-whole-rock data, and U-Pb in allanite, giving ages of 45 ± 4.4, 55 ± 12.0, and 55 ± 17.0 Ma, respectively.These estimates attempted to date minerals which can definitively be linked to metamorphic reactions/conditions, but this was hindered by the large uncertainties on these dates.
A spread in zircon dates from 55 to 45 Ma was recorded in the Puga Gneiss by Leech et al. (2007) by U-Pb SHRIMP dating of the metamorphic rims found around igneous zircon grains (n = 19; 4 ages interpreted).Sub-concordant data were split into groups arbitrarily (see O'Brien, 2006), and interpreted to represent ultrahighpressure metamorphism at 53.3 ± 0.7 Ma, followed by further zircon growth in eclogite facies conditions at 50.1 ± 0.6 Ma, amphibolite facies at 47.5 ± 0.6 Ma, and a lower temperature growth at 45.2 0.7 Ma. St-Onge et al. (2013) recorded monazite and allanite U-Pb SHRIMP ages of 45.3 ± 1.1 and 43.3 ± 1.1 Ma respectively, interpreted to represent post-eclogite facies peak temperature metamorphism of 7.0-8.4kbar at 705-755°C , based on pseudosection modeling of the observed garnet breakdown reaction.Cooling through phengite, biotite and muscovite Ar-Ar closure temperatures is recorded at 48.0 ± 2.0, 31.1 ± 0.3, and 29.3 ± 0.3 Ma (De Sigoyer et al., 2000) respectively, with further cooling recorded by apatite and zircon fission track data at ∼23.5-7.5 Ma (Schlup et al., 2003).Dates from within the Puga Gneiss are dominated by post-peak, amphibolite facies ages recording Barrovian metamorphism, cooling and uplift.
The mafic eclogites better preserve the early high-pressure metamorphic history in their major and accessory mineral assemblages.St-Onge et al. (2013) analyzed zircon in situ with SHRIMP U-Pb geochronology, yielding an age of 58.0 ± 2.2 Ma (n = 2) for zircon included in the core of a garnet, and 50.8 ± 1 Ma (n = 4) for zircon included/adjacent to matrix barroisite, phengite and garnet.The older age is interpreted to record zircon crystallization during prograde garnet growth to high pressure, and the younger age is interpreted to represent peak metamorphism in the eclogite facies.Whilst the St. Onge et al. (2013) study provides a textural constraint on the prograde zircon ages which was previously unknown, the data include a limited number of spots and the presence of common lead.Donaldson et al. (2013) used in situ split-stream LA-ICP-MS to measure U-(Th)-Pb combined with REE concentrations of zircon from two mafic eclogites, located ∼10 km apart.Lower intercept dates of each sample overlap at 45.3 ± 1.6 and 44.2 ± 1.2 Ma.These ages broadly overlap with that of Kaghan, as highlighted by Donaldson et al. (2013), but these authors also argue for protracted zircon crystallization in the eclogite facies (ca.47-43 Ma) based on the spread of single spot dates and consistent REE signatures with an absence of an Eu anomaly and flat HREE profiles.The Donaldson et al. (2013) results improve upon the constraints of St. Onge et al. (2013) based on the number of data, but are still hampered by both the presence of common lead and the precision of the in situ method which yields single spot uncertainties of ca.1-10 Ma.
The age of UHP metamorphism in the Kaghan Valley Complex has been estimated using U-Pb SHRIMP and U-Pb ID-TIMS analyses of zircon from eclogite-facies mafic rocks, yielding ages of 46.2 ± 0.7 Ma (Kaneko et al., 2003) and 46.4 ± 0.1 Ma (Parrish et al., 2006), respectively.These zircons were found included in UHP garnet rims with coesite inclusions.Eclogite facies ages of zircon and allanite from Kaghan were also estimated using U-Pb and Th-Pb ID-TIMS analyses at 45.5 ± 6.6 and 46.5 ± 1.0 Ma, respectively (Parrish et al., 2006).An additional age of eclogite facies zircon was estimated using U-Pb SIMS at 46 ± 2 Ma in Naran, 30 km south-west of Kaghan (Zhang et al., 2022).A compilation of geochronology of high pressure metamorphism in the north west Himalaya can be found in Supporting Information S1 and is compared to our new data in Figure 8.

Analytical Methods
Petrographic study of 29 mafic eclogites and 28 Puga Gneiss samples from the Tso Morari Complex was undertaken, with one fresh eclogite (a03-16), one retrogressed eclogite (a03-12) and one gneiss sample (05-02) selected for further petrological analyses.Major element compositions of minerals that exhibit solid solutions were measured using a Cameca SX-5 field emission electron microprobe at the University of Oxford, with a 15 kV acceleration potential, 20 nA beam current, 30 s count time per major element (30 s background) and 60 s count time on Ti (60 s background).A range of natural and synthetic oxide standards were used including albite (Na, Al, Si), Orthoclase (K), MgO (Mg), wollastonite (Ca), andradite (Fe), Mn metal (Mn), and synthetic TiO 2 (Ti) and analyses were verified against secondary mineral standards.Mineral spot analyses and line profiles were taken across garnet to determine the extent of intracrystalline compositional variation (see Supporting Information S1).
Quantitative major element X-ray maps of garnets were collected from polished thin sections using the CAMECA SX-5 field emission electron microprobe at the University of Oxford at a working distance of 10 mm, a 15 k V acceleration potential, 170 nA current, 0.06 s dwell time and a 3 μm step size for the elements P, Ca, Mn, Mg, and Fe.A dwell time of 0.032 s and a 200 nA beam current were used for the elements Al, Si, Ti, Y, and Yb.A range of natural and synthetic oxide standards were used including Durango apatite (P), andradite (Ca, Fe), Mn metal (Mn), MgO (Mg), albite (Al, Si), TiO 2 (Ti), Y metal (Y), Yb metal (Yb).

Puga Gneiss (Sample 05-02)
Puga Gneiss sample 05-02 was collected from the north shore of Kiagar Tso (33.1214°N, 78.2958°E), in the middle of the Tso Morari Complex.Sample 05-02 is a strongly foliated, garnet-bearing gneiss comprising albite, quartz, muscovite, biotite and garnet, with accessory zircon, apatite, rutile, and xenotime.Sample 05-02 shows a schistosity (S2) defined by bands of white mica, albite and quartz (Figure 2a) which wrap around larger garnet porphyroblasts.Quartz occurs in polycrystalline ribbons or lenses separated by narrow bands of white mica, indicative of a high-strain fabric (Figures 2a-2d).The quartz ribbons are cut by discontinuous shear bands forming an S-C′ fabric.Quartz grains are >200 μm and equant with ameboid shapes along grain boundaries indicative of grain boundary migration (GBM) dynamic recrystallization (e.g., Stipp et al., 2002).Albite bands and lenses are largely composed of fine-grained aggregates.Among these, larger feldspar grains show subgrains of similar size to the dominant population, suggesting that dynamic recrystallization occurred in the subgrain rotation (SGR) regime for feldspar (Passchier & Trouw, 2005).Zircon and rutile occur in the matrix.
Garnet porphyroblasts of varying size are scattered through the rock.They are abundantly fractured, and partly replaced by biotite, white mica and chlorite.The studied section contains one large (6 × 5 mm) ovoid garnet porphyroblast (Figure 3) that contains significant textural complexity (see below).Fine-grained inclusion trails of quartz and rutile define a primary foliation (S1), which is oblique to the matrix fabric (S2) and folded at the core-rim boundary (Figure 3a).Kyanite occurs within the garnet core, adjacent to staurolite, surrounded by white mica (Figure 3a).The garnet is subhedral with embayments containing biotite, white mica and xenotime (Figure 3a).
Garnet exhibits concentric zonation, comprising three distinct zones (Figure 3).The garnet core (zone P1, Figure 3c) has an approximately uniform composition of Alm 79.0 Grs 12.5 Pyr 5.5 Spss 4.0-2.3 .Inclusions of quartz, rutile and white mica are found throughout zone 1, whereas inclusions of kyanite are restricted to the inner portion of zone P1 (Figures 3a-3f).Zone P2 (Figures 3b, 3d, and 3f) is marked by a decrease in Ca and increase in Mn.It contains a narrow annulus of high Y content, and its outer boundary is faceted.The narrow rim zone P3 (Figures 3d and 3f) shows a marked increase in Mg (to 16% pyrope) and Ca (to 13% grossular), with lower Mn.It is discontinuous around the garnet margin.Zones P2 and P3 contain quartz, white mica, and rutile inclusions, but to a lesser extent than in zone 1. Kyanite is absent from zones P2 and P3.In zones P1 and P2, magnesium-enriched haloes are developed along the internal fracture network, surrounding many larger inclusions of white mica and quartz, and connecting with the outer zone of the garnet.
Garnet and omphacite (Figure 4a) are in textural equilibrium, forming planar contacts, with coarse grained homogeneous phengite and talc.In some places, symplectites after omphacite are observed, comprised of amphibole, plagioclase and occasional diopside.Dolomite is also present in the matrix as large poikiloblasts containing inclusions of phengite, omphacite and rutile.Amphibole occurs in the matrix and as inclusions in garnet (Figure 4b).In common with other sampled outcrops, coarse-grained matrix amphibole typically has pale green winchite cores, zoned toward a pargasitic composition at rims.Thin mantles of green pargasitic amphibole surround garnet, and amphibole inclusions in garnet are darker blue-green pargasite and ferropargasite.Fine-grained intergrowths of biotite and plagioclase surround phengite.Matrix quartz and omphacite contain clusters of zircon inclusions, whereas garnet, omphacite, quartz, phengite, and talc contain rutile inclusions (Figure 4c).Garnet occurs as subhedral porphyroblasts that are 3-4 mm in diameter (Figure 4b) and contain inclusions of zoned blue-green amphibole, and also quartz and rutile.Rutile commonly forms angular clusters of grains, both in the matrix and within garnets, which we interpret as replacing a former igneous Fe-Ti oxide.
Major and trace element maps of a single garnet in sample a03-16 show three zones (Figure 5).Zone E1 (Figure 5d) defines the core region, which has a composition of Alm48-42 Grs24-29 Pyr20-30 Sps6-1, with increasing almandine and pyrope components from core to rim (Figures 5b and 5c).The rim of the zone E1 core is outlined by a Mn-rich annulus.Zone E2 surrounds the zone E1 core and is defined by a calcium trough with a faceted outline (Figures 5d and 5e).Zone E3 (Figure 5b) forms along the garnet rim defined by a Mg-rich, Mn-  poor rim with shows an increase in pyrope (to 30%-35%) and grossular (to 30%, decreasing to 16% at the outer rim).Mg-rich E3-like domains also extend in some places from the outer rim toward the core of the garnet, cutting across concentric garnet zones E1 and E2, and surrounding inclusions, often forming channelized features (Figures 5b and 5c).

Puga Gneiss (Sample 05-02)
The Puga Gneiss (sample 05-02) displays evidence of two metamorphic assemblages.The first is defined by the garnet compositional zoning, and the inclusion suite of quartz, kyanite, rutile and zircon.Quartz and rutile inclusions define a crenulated primary foliation (S1).Kyanite is found exclusively within garnet.The garnet compositional zoning has features that match the zoning in eclogite garnet (cf.St-Onge et al., 2013; Figure 5), that is, a Ca-rich core (P1 = E1), Ca-poor trough (P2 = E2) and Mg-and Ca-rich rim (P3 = E3).This is consistent with prograde growth culminating in eclogite-facies conditions (e.g., O'Brien, 2019b;St-Onge et al., 2013;Wilke et al., 2015).The Mg distribution in the garnet interior implies that some fracturing occurred at near-peak conditions.The second assemblage is defined by the rock matrix grains outside the garnet porphyroblasts, which form a segregated quartz-feldspar-mica mineral banding modified by an S-C′ fabric.Accessory minerals in this matrix assemblage include rutile and zircon.Dynamic recrystallization of quartz by grain boundary migration suggests a deformation temperature in excess of 530°C (Stipp et al., 2002), and evidence for subgrain rotation in albite is consistent with about 600°C (Passchier & Trouw, 2005).No lower-temperature dynamic recrystallization microstructures are observed, and minor chlorite, generally associated with garnet, is undeformed.Rutile is stable in the rock matrix, commonly enclosed in white mica.Based on these observations, we interpret the matrix assemblage to reflect metamorphism and deformation on the retrograde path, at amphibolite-facies conditions.
A distinct mineral association surrounds the large garnet in sample 05-02, where parts of the Mg-rich rim zone (P3) of the garnet are missing, due to partial resorption of garnet, and micaceous aggregates in embayments and fractures host xenotime grains (Figure 3a).

Mafic Eclogite (Sample a03-16)
Sample a03-16 contains three identifiable metamorphic assemblages.Garnet and omphacite (Figure 4a) in textural equilibrium with coarse grained homogeneous phengite and talc define a peak pressure eclogite-facies mineral assemblage (e.g., M2 of St-Onge et al., 2013).Accessory phases within this peak assemblage include zircon, present as clusters of grains included within matrix quartz and omphacite, and rutile, present as inclusions, commonly clustered, within garnet, omphacite, quartz, phengite, and talc (Figure 4c).Inclusions of zircon and rutile in garnet cores suggest that initial growth of these grains may have begun prior to eclogite facies metamorphism.Within garnet, the increase in pyrope component from core to rim, corresponds to garnet growth during prograde to peak eclogite conditions.We interpret zones E1-E3 as a prograde to peak assemblage, where zone E3 correlates with the omphacite, phengite and talc peak assemblage described above.The second metamorphic assemblage is defined by the early breakdown products of the peak assemblage phases.The dominant example is the zoned blue-green matrix amphibole, which may represent the product of talc dehydration as well as the influx of external fluids under eclogite-facies conditions (cf.Palin, Searle, et al., 2014;Palin, St-Onge, et al., 2014).The third assemblage includes feldspar-bearing symplectites after omphacite, and secondary finegrained white mica aggregates forming in the irregular rims of coarse-grained phengite.These represent an amphibolite-facies overprint consistent with the matrix assemblage of Puga gneiss sample 05-02.

Retrogressed Mafic Eclogite (Sample a03-12)
Sample a03-12 is dominated by a post-peak metamorphic assemblage, whereas the peak assemblage displayed by sample a03-16, is identifiable in a03-12 as relict garnet grains and pseudomorphs after garnet and omphacite.Blue-green amphiboles are interpreted as an early post-peak phase, which grew within the eclogite facies field prior to the appearance of stable sodic feldspar during decompression (e.g., Palin, Searle, et al., 2014;Palin, St-Onge, et al., 2014).The rest of the post-peak assemblage is typical of an amphibolite facies retrograde assemblage characterized by fine-grained amphibole and plagioclase intergrowths.Aggregates of chlorite and biotite characterize a lower amphibolite facies overprint.Rutile is found within prograde relict garnets as well as the matrix and has no indication of internal zonation.We therefore interpret rutile as a relict grains from a prograde or peak assemblage, rather than a new phase that nucleated during latest-stage lower amphibolite facies metamorphism.Zircon is present throughout this sample, predominantly as inclusions in peak omphacite and quartz, suggesting that it crystallized at eclogite-facies conditions.

Correlation of Metamorphic Assemblages (M1, M2, M3) in the Puga Gneiss and Mafic Eclogites
Comparison and correlation of the petrography and petrology of samples 05-02, a03-16, and a03-12, allows for the definition of three distinct metamorphic assemblages that reflect distinct portions of the same P-T path of the Tso Morari complex, experienced and recorded by both lithologies.These assemblages are correlated with the stages defined by St-Onge et al. ( 2013) (M1, M2, M3).Crucially, these assemblages provide a robust means for linking accessory phase geochronology to metamorphic evolution of the Tso Morari complex.These metamorphic assemblages and their constituent accessory phases are summarized as follows.

M3a/M3b/M3c: Post-Peak Assemblage
The M3 post-peak assemblage reflects continuing metamorphism from eclogite to lower amphibolite facies conditions, and is subdivided to reflect this.In the mafic eclogite samples (a03-16, a03-12), M3a is recorded by growth of coarse-grained blue-green amphiboles in the eclogite facies, which elsewhere in the Tso Morari Complex, has been linked to the breakdown of talc and the influx of fluid at 23 and 19 kbar respectively (Palin, Searle, et al., 2014;Palin, St-Onge, et al., 2014).Post-peak assemblage M3b corresponds to upper amphibolite facies retrograde metamorphism.In the mafic eclogite samples (a03-16, a03-12) M3b is defined by symplectites of fine-grained amphibole and plagioclase intergrowths after omphacite.The lower-temperature association of chlorite with biotite, largely as a replacement of garnet cores in a03-12, is assigned to M3c.
In Puga gneiss sample 05-02, initial garnet breakdown and the associated nucleation of xenotime, occurred at eclogite to upper amphibolite facies, and therefore correlates with either the M3a or M3b post-peak assemblages observed in the mafic eclogites.M3b is defined by the matrix assemblage of quartz + albite + muscovite, which also contains zircon and rutile, and displays quartz microstructures indicating post-peak deformation temperatures of >530°C.M3c corresponds to lower amphibolite facies retrograde metamorphism and is represented by overprinting aggregates of chlorite and biotite after garnet.

Analytical Methods
Zircon grains from the heavy, non-magnetic fraction of sample a03-12 were imaged via cathodoluminescence (CL) using an FEI Quanta 650 environmental scanning electron microscope (E-SEM) at the University of Oxford, using a 10 kV electron beam, 16 mm working distance and a beam current of 0.49 nA.Zircon grains from sample 05-02 were also mounted on sticky tape in order to analyze the <10 μm thick rims.Rutile grains chosen for analysis were picked from the non-magnetic fraction and were imaged via backscatter electron imaging (using the same E-SEM) to determine the homogeneity of the grains chosen for analysis.None of the rutile grains showed any evidence of zoning.Xenotime was measured in a polished thin section in order to preserve the petrographic relationships observed.
All geochronology and mineral separation were conducted at the Geochronology and Tracers Facility, British Geological Survey, Nottingham, UK.Laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) was conducted using a Nu Instruments AttoM sector-field single-collector ICP-MS, coupled to an Elemental Scientific Lasers 193 nm UC Excimer laser ablation system fitted with a TV2 cell.The method follows that described in Roberts et al. (2016), with uncertainty propagation following recommendations of Horstwood et al. (2016), and age calculation and plotting using IsoplotR (Vermeesch, 2018).Common lead corrected ages, where quoted, use a 207 Pb-based method (Chew et al., 2014) and assume a Stacey and Kramers (1975) initial lead composition, and concordance of the final age.All uncertainties are quoted and plotted at 2σ. LA-ICP-MS dates are quoted with propagation of systematic uncertainties, including reference material and decay constant uncertainties; age uncertainties are also expanded to account for the MSWD where the data exhibit overdispersion (Vermeesch, 2018).Accuracy based on zircon and rutile validation materials was better than 1% for zircon and xenotime and 5% for rutile (Plešovice = 339.1 ± 1.9 Ma; GJ-1 = 596.8± 2.9 Ma; Manangotry = 555.48± 4.23; R19 = 465.9± 3.9 Na; Sugluk = 1,714 ± 30 Ma; and PCA = 1,859 ± 25 Ma).Trace elements were measured using the same instrumentation as for U-Th-Pb, with the Attom measuring in linkscan mode (see Supporting Information S1 for full analytical protocol), with normalization to GJ-1 zircon (Piazolo et al., 2017) except Ti which used Szymanowski et al. (2018).
Zircon in one sample (a03-12) was further analyzed by CA-ID-TIMS, following analytical and data reduction methods described by Tapster et al. (2016), and utilizing the ET535 EARTHTIME mixed tracer (Condon et al., 2015).The CA-ID-TIMS uncertainties are quoted as analytical only/analytical + tracer-calibration/ analytical + tracer-calibration + decay-constant uncertainty, as per community protocols.Our U-Pb dates can be compared using the intermediate uncertainty, but for the sake of simplicity, we quote the final uncertainty within the discussion such that our dates can be compared with constraints from other absolute and relative chronometry.

Zircon
Zircon grains in the Puga gneiss sample 05-02 show oscillatory zoning and a euhedral shape, indicative of igneous zircon, with a Th/U of 0.05-0.26.Thin, bright rims (<20 μm) are present surrounding the dark cores.From the epoxy mounted zircon, 33 LA-ICP-MS analyses of zircon cores were obtained from 21 grains.Five ages ranging from 2,481 to 1,020 Ma ( 207 Pb/ 206 Pb age), indicate a population of xenocrystic zircon inherited during emplacement of the Puga gneiss igneous protolith.Nine analyses had ages that were discordant by >10%, and were discarded from age calculation.Of the remaining 19 concordant analyses, 17 spots provide a concordia age of 482 ± 5.9 Ma with a mean square of weighted deviates (MSWD) value of 1.5 (Figure 6a).Using spot analysis of tape-mounted zircon, whereby the outer rim can be targeted more confidently, 64 grains yielded a spread of ages from ca. 569 to 48 Ma (Figure 6b).A broad plateau of ages overlaps the mounted zircon at ca. 480 Ma, consistent with our weighted mean concordia age.The youngest three analyses provide a weighted mean, using common lead corrected 206 Pb/ 238 U ages, of 48.9 ± 1.2 Ma (MSWD = 1.3; Figure 6b).

Xenotime
Rare 30-100 μm xenotime grains were observed in micaceous aggregates at corroded margins of garnet in sample 05-02, as described above (Figure 3a).The grains show homogenous brightness in BSE indicating the absence of internal zonation.Nine spots were analyzed across 9 grains, and yield a mixing line between radiogenic and common lead components.Using a free regression, the lower intercept age is calculated at 47.4 ± 1.4 Ma (MSWD = 1.9; Figure 6c).

Rutile
Rutile grains in sample 05-02 measure 50-100 μm in size and show homogenous brightness in BSE, indicating the absence of internal zonation.Fourteen spots were analyzed across 14 rutile grains, and yield a mixing line between radiogenic and common lead components.Using a free regression yields a lower intercept age of 40.4 ± 1.7 Ma (MSWD = 2; Figure 6d).

Zircon
Zircon grains are rarely found in mafic eclogites of the Tso Morari Complex.In thin section, zircons are often found as inclusions within eclogite-facies phases such as quartz and omphacite.Separated zircon grains from sample a03-12 are translucent and colorless with rounded and euhedral grain shapes and a grain size of <70 μm (Figure 7f).They have zoning patterns dominated by broad oscillatory or sector zoning, that in some grains, truncates a darker core region; the latter were not analyzed.
Thirteen laser spot analyses were performed on the cores of 13 zircon grains separated from sample a03-12 for U-Pb, with 3 analyses rejected due to Pb counts below detection.The U concentrations are variable (13.5-1,528 ppm), and the degree of discordance varies between 0.1% and 12%.Of the remaining 10 analyses, the lower intercept of 9 spots is calculated using a free regression at 47.5 ± 1.7 Ma (MSWD = 0.63; Figure 7a).These zircon analyses have low Th/U ratios (<0.01).The result implies that the data conform to a single population.The Tectonics 10.1029/2023TC007976 BIDGOOD ET AL. omitted analysis has a much higher Th/U ratio (0.25), suggesting an igneous core region was clipped during the ablation.
Trace elements were measured on 20 zircon grains, including adjacent spots on the same 13 grains analyzed for U-Pb; 3 analyses were omitted due to inclusions.Ti-in-zircon temperatures are calculated using Si and Ti activities of 1.0, and the equation of Ferry and Watson (2007).The temperatures range from 603 to 724°C, forming a normal distribution around a peak at ca. 630°C (Figure 7c).The REE data are plotted as chondrite-normalized values (Figure 7d).The REE patterns of zircon are broadly consistent across multiple grains, with no Eu anomaly, flat HREE patterns and depleted LREEs.
Several zircon grains were extracted from the resin mounts and prepared as single grain aliquots for CA-ID-TIMS.The resulting data are six reproducible fractions yielding a weighted mean (Th-corrected) 206 Pb/ 238 U age of 46.912 ± 0.068 Ma with an MSWD of 0.2 (Figure 7b).

Rutile
Rutile grains in sample a03-12 show homogenous brightness in BSE with no indication of internal zonation.The majority of analyses had Pb counts below detection, the remaining data comprise six spots analyzed across 6 grains.The analyses are distributed between the radiogenic and common lead components, and using a free regression, yield a lower intercept age of 36.3 ± 3.8 (MSWD = 1.4; Figure 7e).2013).(b) Time chart for Tso Morari and Kaghan, after (Palin et al., 2012), comprising data from multiple sources, including this study.

Puga Gneiss (Sample 05-02)
Zircon core ages of 482.0 ± 5.9 Ma in the Puga Gneiss are interpreted to represent the igneous crystallization age and are comparable with U-Pb zircon ages of 479 ± 1 Ma in both the Polokongka La granite and the Puga Gneiss (Girard & Bussy, 1999).A monazite age of 473 ± 9 Ma in the Polokongka La granite, interpreted to have formed during crystallization of the granite (Bidgood et al., 2023), also overlaps with the age of zircon crystallization.Rim analyses of Ordovician zircon from the Puga Gneiss yielded an age of 48.9 ± 1.2 Ma (LA-ICP-MS), which overlap with high precision CA-ID-TIMS zircon age from mafic eclogite a03-12 (hosted within the gneiss), suggesting that zircon rims in the Puga gneiss also grew at eclogite-facies conditions.
Xenotime is found exclusively in Puga Gneiss sample 05-02, in mica aggregates adjacent to the corroded outer rim of a large garnet porphyroblast (Figures 2 and 3).Breakdown of parts of the Mg-enriched garnet rims, formed under peak eclogite-facies conditions (M2), liberated Y, which is concentrated in garnet zones P2 and P3 (Figure 3a), for xenotime growth.We therefore interpret the U-Pb xenotime age in sample 05-02 (47.4 ± 1.4 Ma) as the age of initial garnet breakdown during decompression.The xenotime age overlaps with the eclogite facies zircon rim age within the same sample 05-02 (48.9 ± 1.2 Ma), as well as the precise eclogite-facies zircon CA-ID-TIMS age of mafic eclogite sample a03-12 (46.912 ± 0.068 Ma), located within the same structural unit.This allows us to assign the xenotime crystallization and garnet breakdown in the Puga Gneiss sample to the M3a metamorphic assemblage, reflecting the earliest phase of decompression at eclogite facies conditions.
Rutile is present as inclusions within prograde (M1) to peak pressure (M2) garnets and the matrix of the Puga Gneiss.Given that the peak temperature of the Tso Morari Complex (705-755°C, St-Onge et al., 2013) is greater than the closure temperature of rutile at ∼400-630°C (Cherniak, 2000;Kooijman et al., 2010;Mezger et al., 1989;Vry & Baker, 2006), which is dependent on cooling rate and grain size (Oriolo et al., 2018;Zack & Kooijman, 2017), the U-Pb age of rutile can therefore be attributed to cooling of the Puga Gneiss through the rutile closure temperature at 40.4 ± 1.1 Ma, after peak temperature conditions.

Retrogressed Mafic Eclogite (Sample a03-12)
LA-ICP-MS analyses of zircon from mafic eclogite sample a03-12 are entirely <50 Ma, and provide a weighted mean age of 47.5 ± 1.7 Ma.Although a limited data set (9 analyses), our result implies a single population of metamorphic zircon, rather than preserving a protracted history of zircon growth.Flat HREE profiles and the lack of an Eu anomaly in the core and mantle of zircon grains are indicative of crystallization in garnet-present, plagioclase-absent conditions, consistent with crystallization in the eclogite facies (Rubatto, 2002;Rubatto & Hermann, 2003;Schaltegger et al., 1999).The measured zircon date is therefore interpreted to correspond to growth within the eclogite-facies.Ti in zircon temperatures of 603-724°C lie below the closer temperature of Ti in zircon (∼900°C, Cherniak et al., 2007;Watson et al., 2006) and is therefore interpreted as the zircon crystallization temperature.
The zircon CA-ID-TIMS data from mafic eclogite sample a03-12 are tightly clustered yielding a date of 46.91 ± 0.068 Ma, which we interpret as the age of zircon growth at eclogite facies conditions (M1-3a).We note that this interpretation is heavily weighted to only the two analyses of the most U-rich zircon, excluding these analyses yields a weighted mean of 46.71 ± 0.33/0.38/0.38Ma (MSWD = 1.04; n = 4) and is therefore within uncertainty of the original interpretation.Regardless of interpretation, the uncertainty associated with the CA-ID-TIMS data is 1-2 orders of magnitude smaller than LA-ICPMS U-Pb data and provides the most precise estimate of the timing of metamorphic zircon growth from Tso Morari thus far.Additionally, even with the improved precision, the analyses and the lack of dispersion indicated by their MSWDs suggests no measurable crystal to crystal variation at around the ∼1 Myr resolution.
Rutile is the peak titanium-bearing phase in mafic rocks during eclogite-facies metamorphism and is present as inclusions in garnet (M1), as well as in the matrix (M2-3b).Peak temperature in the Tso Morari Complex is estimated at 600-755°C (De Sigoyer et al., 2000;St-Onge et al., 2013), above the predicted closure temperature of rutile at ∼400-630°C (Cherniak, 2000;Kooijman et al., 2010;Mezger et al., 1989;Vry & Baker, 2006) which is dependent on cooling rate and grain size (Oriolo et al., 2018;Zack & Kooijman, 2017).The U-Pb age of rutile can therefore be attributed to cooling through this closure temperature range.

Metamorphism, Deformation and Geochronology of the Tso Morari Complex
Previous studies indicate that subduction, exhumation and emplacement of the Tso Morari Complex took place between c. 60 Ma and c. 7.5 Ma (Figure 6), with exhumation to lower crustal conditions by 45.3 ± 1.1 Ma at an average rate of ∼12 mm a 1 (St-Onge et al., 2013).The texture and composition of the mafic eclogites and Puga gneiss samples in this study collectively provide a record of initial exhumation from UHP eclogite-facies conditions, followed by exhumation through crustal conditions.Integrating this information with our new highprecision geochronology from a range of accessory phases which crystallized at different stages of metamorphism allows us to constrain the timing of mineral growth and fabric development with respect to the burial and exhumation of the Tso Morari Complex.
Evidence of prograde metamorphism to peak pressures (M1-2) is preserved within garnets from mafic eclogite and the Puga Gneiss, comparable to that observed in the eclogite studied by St-Onge et al. (2013).The garnetbearing Puga Gneiss, as represented by sample 05-02, is not typical of the host Puga Gneiss, which generally preserves only amphibolite-facies assemblages.The formation and preservation of garnet in this sample is therefore particularly valuable given that it records the same overall prograde compositional trends as seen in the Tso Morari eclogites.The analyzed garnet (Figures 3a and 3b) also contains complex deformation fabrics and inclusion suites, indicating that deformation and transformation of the original granite was already underway prior to garnet growth (M1) in at least part of the complex.Early fabrics are rarely preserved in the Tso Morari Complex, where the earliest stage of macroscopic deformation has been previously identified as the dominant topto-the-east exhumation fabric within the Puga Gneiss, and is attributed to initial M3a exhumation from eclogitefacies conditions (Epard & Steck, 2008).
Our eclogite-facies zircon dates of 46.912 ± 0.068 Ma (CA-ID-TIMS) and 47.5 ± 1.7 Ma (LA-ICPMS) from a mafic eclogite overlap with our zircon rim and xenotime dates in the Puga Gneiss.We interpret this overlap as a record of the earliest phase of decompression at 630°C and eclogite facies conditions, as recorded by the Ti in zircon temperatures and the partial breakdown of yttrium-bearing garnet rims.The crucial breakdown of the garnet rim zones and fracture-based alteration, with xenotime among its products, provides insight into the early exhumation history of these rocks that is not recorded elsewhere, but predates the Barrovian overprint of 7.0-8.4kbar, 705-755°C recorded at 45.3 ± 1.1 and 43.3 ± 1.1 Ma by St-Onge et al. (2013).
The formation of quartz microstructures during high-temperature (>530°C) dynamic recrystallization indicates that deformation took place during exhumation to crustal conditions (M3b).There is no significant later overprinting of quartz deformation fabrics in Puga Gneiss samples from the core of the dome (Bidgood, 2020), implying a lack of pervasive deformation below 530°C.Deformation therefore occurred prior to cooling through the rutile closure temperature of ∼630-400°C (Cherniak, 2000;Li et al., 2013;Mezger et al., 1989;Vry & Baker, 2006) at 40.4 ± 1.1 Ma and was ductile and pervasive.This interpretation is confirmed by Dutta and Mukherjee (2021) who predict deformation temperatures >600°C based on misorientation analysis and micro-textural evidences of incipient partial melting in the Puga gneiss.Subsequent deformation during exhumation relating to doming and emplacement (M3c) was not pervasive, with foliations and lineations developing at the margins of the dome, adjacent to the normal sense shear zones (e.g., Zildat-Ribil fault, Bidgood, 2020;Dutta & Mukherjee, 2021;Epard & Steck, 2008).The age of exhumation of the Tso Morari dome is recorded by the Ar-Ar muscovite and biotite and apatite fission track dates of <32.4 Ma (De Sigoyer et al., 2000;Schlup & Carter, 2003) which record cooling through the Ar closure temperatures of muscovite (∼425°C, Harrison et al., 2009), biotite (<345°C, Harrison et al., 1985) and apatite (<135°C, Warnock et al., 1997) respectively.

Continental Subduction and Exhumation in the NW Himalaya
Our petrographic correlation of zircon and xenotime ages with the M3a assemblage indicates that zircon growth at ∼47-46 Ma took place at subsolidus conditions during the earliest stages of decompression from UHP conditions.These ages overlap with the zircon age distribution peak of 47-43 Ma recorded by Donaldson et al. (2013).Considering their analytical scatter (i.e., MSWDs of 2.4 and 3.4) the Donaldson et  metamorphism, starting at ∼47 Ma.It is therefore necessary to reassess the Donaldson et al. (2013) data in light of our new data and observations, as follows.
Subsolidus zircon growth requires a fluid phase to mediate the liberation of Zr from Zr-bearing phases (ilm, ru, cpx, grt) (Chen & Zheng, 2017;Chen et al., 2010;Kohn et al., 2015;Skuzovatov et al., 2021).The host rock in the Tso Morari complex is a dry granite with intruded mafic dykes (see Bidgood et al., 2023) resulting in restricted opportunity for prograde/peak zircon growth.Palin, Searle, et al. (2014) and Palin, St-Onge, et al. (2014) determined that the first pulse of post-peak fluid in the eclogite facies relates to the destabilization of talc and growth of the coarse-grained amphiboles at ∼23 kbar, followed by fluid infiltration from an external source at ∼19 kbar.Coarse-grained, zoned amphiboles are abundant in mafic eclogites from across the Tso Morari Complex suggesting that post-peak eclogite facies hydration was a common and widespread occurrence, aided further by exhumation-related deformation.These influxes of fluid during exhumation, subsequent to a potentially dry spell at prograde to peak conditions, would have provided conditions favorable for high concentrations of zircon growth.Consequently, we argue that the breakdown of UHP garnet rims recorded by xenotime at 47.4 ± 1.4 Ma indicates that the zircon age peak at 47-43 Ma of Donaldson et al. (2013), along with our own CA-ID-TIMS zircon age of 46.91 ± 0.068 Ma, reflects a pulse of zircon growth during the onset of exhumation from UHP conditions, aided by the exhumation-driven liberation of fluids.
Based on the above conclusion, older dates in previous studies, including a 58 ± 2.2 Ma (n = 2) zircon age from St-Onge et al. ( 2013) and apparent older zircon dates (55-48 Ma) from Donaldson et al. (2013), may record zircon growth during prograde to peak metamorphism.However, it is difficult to place them in a robust context given: (a) the proportion of common lead for most of the analyses; (b) the potential ablation of older zircon domains reflecting the protolith age; and (c) the lack of relationship between textural location and REE chemistry (Figure 8a).The nature of apparent younger common-lead corrected dates from the Donaldson et al. (2013) data set also remain uncertain for the same reasons; younger dates could plausibly record continued zircon growth during decompression, but dates younger than the ca.40 Ma rutile age imply these may relate to analytical inaccuracies.Given that we consider our xenotime date as a marker for the onset of UHP exhumation, we suggest that the Donaldson et al. (2013) data reflect zircon growth due to fluid infiltration at and after 47 Ma.
Our new CA-ID-TIMS zircon age of 46.912 ± 0.068 Ma more closely and precisely correlates with ages from zircon (46.4 ± 0.1 Ma-ID-TIMS; 46.2 ± 0.7-SHRIMP, 46 ± 2 Ma-SIMS) and allanite (46.5 ± 1.0 Ma-ID-TIMS) in UHP assemblages from Kaghan, located 450 km to the west of Tso Morari (Kaneko et al., 2003;Parrish et al., 2006;Zhang et al., 2022) (see Figure 8).We do not think this is coincidental; metamorphic P-T data from both of these units record similar prograde (up to UHP metamorphism) and retrograde (including UHP exhumation and exhumation through the crust) P-T paths.In Kaghan, coesite is found in thin metamorphic zircon rims in the felsic gneiss, indicating that zircon crystallization occurred at UHP conditions (Kaneko et al., 2003).In the Tso Morari Complex, coesite is observed in the outermost rims of prograde garnets but has not yet been observed as inclusions in zircon.Quartz microstructures after the presence of former coesite are also reported from the Polokongka La Granite of the Tso Morari Complex (Bidgood et al., 2021).Based on these similarities, we argue that the overlap in ages between Tso Morari and Kaghan indicates that UHP exhumation and associated fluid flux at 47-46 Ma was responsible for a ubiquitous pulse of zircon growth across the NW Himalaya, occurring at the same time as or shortly after the achievement of maximum P and T under UHP conditions.We suggest that the regional synchroneity between Tso Morari and Kaghan, across a distance of ∼450 km, reflects the scale at which slab dynamics control metamorphism and exhumation with a subduction zone setting.
The mode of UHP exhumation is unclear from our data.Isothermal, triclinic deformation during exhumation of the Tso Morari complex, as reported by Long et al. (2020) and Dutta and Mukherjee (2021), is most compatible with the recirculation model and plunger model of Warren et al. (2008aWarren et al. ( , 2008bWarren et al. ( , 2008c)).These models invoke the transport of crustal slices of the lower plate from UHP pressures along the subduction interface, rather than wholesale exhumation of the subducting slab, and can therefore occur independently from, and without need for slab break-off.The timing of slab break-off cannot be resolved from our data and therefore remains debated, with estimates ranging widely between ∼50 and 10 Ma (see reviewed compilation of Garzanti et al., 2018).
Lastly, cooling through the rutile closure temperature took place ∼3.7 Ma later in Tso Morari than Kaghan (see Figure 8).At face value, this implies a longer period of time between zircon growth and exhumation through the rutile closure temperature for the Tso Morari Complex, relative to Kaghan.However, uncertainties in the exact conditions of zircon growth and rutile closure temperature prevents us from making any meaningful interpretation for the cause of this difference.

Geodynamic Significance of UHP Exhumation During the Himalayan Orogeny
Our data suggest that zircon crystallization recorded in the Tso Morari complex and Kaghan at 47-46 Ma, corresponds to the onset of exhumation of Indian continental crust from UHP conditions.In the following sections, we consider the geodynamic significance of this event with respect to the wider metamorphic, magmatic and plate kinematic evolution of the Himalayan orogeny and the India-Australia-Eurasia plate network (Figures 9 and 10).

UHP Exhumation During the Himalayan Orogeny
When placed in context with the bedrock record of collision, our constraints for the onset of UHP exhumation in the NW Himalayan reveal how changes in subduction dynamics affected the structural, magmatic, and metamorphic evolution of the Himalayan orogeny (Figure 9).
Collectively the structural, metamorphic, and magmatic changes recorded in both the lower and upper plates of the Himalayan orogeny between ∼50 and 40 Ma (Figure 9a) reflect (a) a warming metamorphic thermal gradient (e.g., Chen et al., 2022;Ji et al., 2024;Soret et al., 2021); (b) greater mechanical coupling of the upper and lower plates resulting in increased crustal shortening and thickening; and (c) a greater contribution of Indian continental crustal material to upper plate magmatism.These phenomena are best explained by a reduction in the dip of the subducting Indian plate beneath Eurasia during that time (Figures 10a and 10b) (e.g., Chen et al., 2022;Ji et al., 2024;Soret et al., 2021).We attribute this to the early stages of the India-Asia collision sensu stricto (i.e., second collision of the Himalayan orogeny), whereby the positive buoyancy of the Indian continental crust stalled subduction of the Indian plate beneath the Eurasian margin (Figure 9b).

A Shift in Orogenic Regimes
UHP exhumation at 47-46 Ma marks a significant geodynamic shift in the orogenic regime of the collision zone (Figures 9b,10a,and 10b).Prior to this time, India-Asia convergence was accommodated by subduction of Indian and/or Neotethys oceanic lithosphere beneath Eurasian.Then, following the India-Asia continent-continent collision sensu stricto by 47 Ma, the reduced dip of the Indian slab increased the mechanical coupling between upper and lower plates and the component of convergence accommodated by crustal shortening.This switch in mode of convergence can be viewed as a shift in the orogenic regime of the collision zone (cf.Cawood   Cawood et al., 2009).(c) Reorganization of the India-Australia-Eurasia plate network at 47-36 Ma with plate velocity profiles for the Indian and Australian plates (Doubrovine et al., 2012;Müller et al., 2019;Torsvik et al., 2008) plotted using GPlates (Müller et al., 2018).Data sources for events in   et al., 2009) from (a) an accretionary orogen, in which India-Asia convergence was driven and accommodated by subduction of Indian and/or Neotethys oceanic lithosphere (Figure 10a); to (b) a collisional orogen, in which India-Asia convergence was accommodated by crustal shortening (e.g., Chen et al., 2022;Parsons et al., 2021;Replumaz et al., 2014), and driven by northeastward subduction of Indian-Australian oceanic lithosphere beneath SE Asia, further to the east (Figure 10b) (e.g., Bose et al., 2023;Capitanio et al., 2015;Li et al., 2008;Parsons et al., 2021).The onset of UHP exhumation at 47-46 Ma therefore timestamps the change in orogenic regime of the India-Asia collision zone from an accretionary orogen to a collision orogen (cf.Cawood et al., 2009) (Figure 9b).

Implications for Models of the Himalayan Orogeny and the India-Asia Continent-Continent Collision (Sensu Stricto)
We interpret the onset of the India-Asia collision (sensu stricto) to occur at ∼50-47 Ma, based on (a) our constraint for the onset of UHP exhumation at 47-46 Ma; (b) a reduction in Indian plate velocity starting at ∼50 Ma (Figure 9c); (c) plate reconstructions of the restored Indian and Eurasian margins at 50-45 Ma (Figure 10b) (Parsons et al., 2021;Replumaz et al., 2014); and (d) changes in the magmatic and metamorphic evolution of the Eurasian margin (upper plate) between 50 and 40 Ma (listed above) (Figure 9a).
In the context of our data, single collision models initiating at ∼60 Ma (e.g., Hu et al., 2016;Ingalls et al., 2016) would require ∼14 Myrs of uninterrupted continental subduction prior to UHP exhumation of Tso Morari at 47-46 Ma.This equates to wholesale subduction of a 2,500 × 2,000 km area of Indian continental lithosphere.This scenario remains highly disputed and is not discussed further as it is yet to be shown that such extreme volumes of continental subduction are feasible or sustainable (Afonso & Zlotnik, 2011;Ingalls et al., 2016;Parsons et al., 2020;van Hinsbergen et al., 2019).
In the context of double-collision models for the Himalayan orogeny, prograde to peak metamorphism of the Tso Morari complex between 60 and 50 Ma (Donaldson et al., 2013;Leech et al., 2007;St-Onge et al., 2013), as well as Barrovian metamorphism of the North Himalayan gneiss domes between 56 and 54 Ma (Ding et al., 2016;Smit et al., 2014), may be explained by initial burial of the NW Himalaya during the first collision event of the Himalayan orogeny (Figure 9a) (first collision began at ∼61 Ma, see introduction for definitions).Our new data do not permit us to constrain the validity of the opposing hypotheses for first collision beyond that which is already discussed in recent reviews (e.g., Kapp & DeCelles, 2019;Parsons et al., 2020;van Hinsbergen et al., 2019).However, we note that models proposing first collision of the Himalayan orogeny between India and an equatorial intra-oceanic arc at ∼61 Ma require two collisions at two separate subduction zones (an intra-oceanic subduction zone and a continental margin subduction zone) and therefore require two slab break-off events.We are unaware of any studies (including our own), which present bedrock evidence for more than one break-off event during the Himalayan orogeny.Cawood et al., 2009) in which convergence was driven and accommodated by subduction of the Indian and/or Neotethys slab.(b) UHP exhumation and slab shallowing starting at 47-46 Ma, in response to the positive buoyancy of Indian continental crust following the onset of continental collision.The impacts of the reduced slab dip are annotated on (b), showing an increase in crustal shortening, Barrovian metamorphism, crustal contamination of arc magmatism and lower crustal melting of the Gangdese arc to produce adakite.The collision zone represents a Collisional Orogen (cf.Cawood et al., 2009), whereby convergence is accommodated by crustal shortening ± continental subduction and driven by subduction of the Australian slab to the east (see Section 6).Two alternative slab break-off scenarios are presented.Slab Scenario 1 (presented in (a)): UHP exhumation and slab dip shallowing occurred in response to slab break-off at 47-46 Ma.Slab Scenario 2 (presented in (b)): Alternatively, UHP exhumation occurred via detachment of the Tso Morari complex from the Indian slab, whilst the buoyancy of the Indian continental crust resulted in lateral slab dragging, overturning of the trailing oceanic slab and reduction in slab dip (see Parsons et al., 2021 for further discussion of lateral slab migrations during the India-Asia collision), with slab break-off occurring at a later time (see Section 6).(c) 46 Ma plate reconstruction showing tectonic events of the Eocene plate network reorganization, triggered by changes in the subduction dynamics of the India-Asia collision.UHP exhumation at 47-46 Ma provide a precise timestamp linking changes in the subduction dynamics of the India-Asia collision zone to plate kinematic changes of the wider plate network.Reconstruction drawn from Müller et al. (2019) using GPlates (Müller et al., 2018), with Greater India and Eurasian margin subduction zone outlines constrained by Parsons et al. (2020Parsons et al. ( , 2021)).Thick dashed black lines show present day locations of the Indus Suture Zone (ISZ) and Main Frontal Thrust (MFT) of the Himalayan orogen, requiring 1,000-2,000 km of crustal shortening since ∼47-46 Ma (Parsons et al., 2021).Abbreviations: AFR-African plate; ANT-Antarctic plate; AUS-Australian plate; CIR-Central Indian ridge; GI-Greater India; IND-Indian plate; ISZ-Indus Suture Zone; ja-Java; khl-Khoistan-Ladakh block; lh-Lhasa block; MFT-Main Frontal Thrust; mk-Mogok metamorphic belt; wb-West Burma block; WR-Wharton ridge.
Plate reconstructions and mantle tomographic analyses propose that this reorganization occurred in response to the India-Asia collision (Gibbons et al., 2015;Parsons et al., 2021;Replumaz et al., 2014); however, these studies lacked an unambiguous and precise marker linking the bedrock record of the collision to plate network reorganization.Crucially, our precise age constraint for the onset of UHP exhumation at 47-46 Ma corroborates those previous studies by demonstrating a viable, causative link between changes in the subduction dynamics of the India-Asia collision zone starting at 47-46 Ma (Figure 9) and the resulting mid-Eocene plate network reorganization that immediately followed (Figure 10c).Our synthesis of the bedrock record of metamorphism and magmatism during the Himalayan orogeny demonstrate that UHP exhumation was rapidly, and in some places immediately, followed by widespread, hotter, amphibolite facies metamorphism (Figure 9a), which we attribute to shallowing and underthrusting of the Indian plate beneath Eurasia.Consequently, we argue that our new geochronological constraints not only date the onset of UHP exhumation in the NW Himalaya, but also date the changes in subduction dynamics of the collision zone that triggered the Eocene plate network reorganization.

The Geodynamic Significance of UHP Exhumation for Orogenesis and Plate Tectonics
Our study demonstrates how our understanding of the bedrock record of continental collisions may be enriched by placing it within the context of plate kinematics.At the same time, the bedrock record provides a crucial means of ground truthing lithospheric-to planetary-scale models of plate tectonics and mantle geodynamics.Further understanding may therefore be gained through new applications of similar integrated approaches to other collision zones.
Multiple examples of HP and UHP metamorphism are recorded across much of SE Asia during the Early to Late Cretaceous, but their timing and tectonic significance is poorly constrained.This metamorphism broadly overlaps with multiple terrane accretion events from the region (e.g., Woyla, Southwest Borneo, and Argoland blocks Advokaat & van Hinsbergen, 2024;Hall, 2012;Metcalfe, 2021;Morley, 2012;van de Lagemaat & van Hinsbergen, 2024), as well as a previously recognized plate network reorganization at ∼110-100 Ma (Matthews et al., 2012;Müller et al., 2016).Detailed petrochronology studies of these occurrences of HP/UHP metamorphism are yet to be conducted but may provide important constraints for plate reconstructions and orogenic models of SE Asia which continue to be debated and revised (e.g., Advokaat & van Hinsbergen, 2024;Hall, 2012;Metcalfe, 2021;Morley, 2012;van de Lagemaat & van Hinsbergen, 2024).
Further back in time, there are many examples HP and UHP metamorphic rocks associated with the Paleozoic tectonic evolution of the Rheic and Iapetus Oceans and related collisional events including the Caledonian, Scandian, and Grampian orogenies (see review of Domeier (2016)).The tectonic and geodynamic significance of these HP/UHP rocks is often unclear.Integrated petrochronology-plate kinematic analyses of these rocks may better constrain both the tectonometamorphic evolution of these HP/UHP metamorphic events and the plate kinematic evolution of the surrounding plate network (e.g., Bottrill et al., 2014;Domeier, 2016;Gilotti et al., 2008;Hacker et al., 2010).Whilst plate reconstructions of pre-Jurassic events are less well constrained than those of younger ages, integration of plate kinematics and bedrock data set still has potential to reveal new information for older tectonic events that would remain otherwise hidden when conducting either discipline (bedrock analysis or plate reconstruction modeling) in isolation.

Summary
By associating accessory phase ages with distinct metamorphic assemblages, and combining both high precision and high spatial resolution techniques, we demonstrate that the phase of peak zircon crystallization recorded in the Tso Morari complex at 47-46 Ma, corresponds to the onset of exhumation from UHP conditions.This study also describes the record of subduction-related major and accessory mineral growth preserved in a rare garnet-bearing Puga Gneiss, which records a part of the P-T-t history not preserved elsewhere.Zircon from a mafic eclogite have a U-Pb CA-ID-TIMS age of 46.912 ± 0.068 Ma (2σ) and an LA-ICPMS age of 47.5 ± 1.7 Ma, with REE profiles indicative of zircon crystallization at eclogite facies conditions.Those ages overlap with zircon rim ages (48.9 ± 1.2 Ma, LA-ICP-MS) and xenotime ages (47.4 ± 1.4 Ma; LA-ICP-MS) from the hosting Puga gneiss, which grew during breakdown of UHP garnet rims, as indicated by garnet element maps.Subsequent exhumation through the rutile closure temperature to crustal conditions is constrained by new dates of 40.4 ± 1.7 .
The overlap between our mafic eclogite zircon ages and our xenotime-UHP garnet break down ages, indicate that the pulse of zircon growth recorded in the Tso Morari complex at 47-46 Ma (e.g., Donaldson et al., 2013) took place as a result of fluid infiltration at the onset of exhumation from UHP conditions, rather than as a result of peak UHP metamorphism.These ages from Tso Morari overlap with U-Pb ID-TIMS, SHRIMP, and SIMS analyses of zircon from eclogite-facies mafic rocks in Kaghan and Naran, ∼450-480 km to west of Tso Morari, which yielded ages of 46.4 ± 0.1, 46.2 ± 0.7, and 46 ± 2 Ma, respectively (Kaneko et al., 2003;Parrish et al., 2006;Zhang et al., 2022).We interpret this overlap as an indication that exhumation from UHP conditions occurred synchronously at 46-47 Ma across the whole of the NW Himalaya.
Integration of our new ages plus previously published ages from the NW Himalaya with existing metamorphic, magmatic, and plate kinematic constraints demonstrates that UHP exhumation at 47-46 Ma occurred in response to changes in subduction dynamics following the onset of the India-Asia collision (sensu stricto) between 50 and 47 Ma.Our data suggest that UHP exhumation was rapidly and in some places immediately followed by widespread, hotter, amphibolite facies metamorphism which we attribute to buoyancy-driven shallowing and under thrusting of the continental Indian slab beneath Eurasia beginning at 47-46 Ma.At a broader perspective, these changes in subduction dynamics and convergence mechanisms of the Himalayan orogeny (Figures 9a and 9b) also impacted the geodynamics and kinematics of the encompassing tectonic plate network.
Our new data from Tso Morari demonstrably tie the bedrock record of the Himalayan orogeny to slab dynamics and plate kinematics of the collision zone.Continent-continent collision of India and Asia at 50-47 Ma not only provided the trigger for UHP exhumation, but also resulted in, (a) significant changes in the metamorphic and magmatic evolution of the Himalayan orogen (Figure 9a); (b) the transition of the Himalaya from an accretionary orogen to a collisional orogen (Figure 9b); and (c) a significant reorganization of the wider India-Eurasia-Australia plate network (Figure 9c).Therefore, our new constraints not only date the onset of UHP exhumation in the NW Himalaya, but also date changes in subduction dynamics of the collision zone that triggered the Eocene plate network reorganization.As a proof-of-concept, our study demonstrates the value of integrated bedrock-plate kinematic approaches, which ultimately lead to a more holistic and accurate understanding how bedrock geology and plate tectonics are linked across subgrain to planetary scales.

Figure 2 .
Figure 2. (a) Puga Gneiss Sample 05-02.PPL and XPL image showing gneissose texture consisting of quartz-rich, feldsparrich and mica-rich layers.Large garnet wrapped by fabric (S2) shown in Figure 3. (b and c) Puga Gneiss Sample 05-01, adjacent to 05-02.(b) Polycrystalline quartz-rich band in between mica sheets.Recrystallized quartz shows undulose extinction.Quartz grains in the lower part of the image span the width of the mica ribbons.(c) Stacked white mica sheets curve into high strain fabric showing top to the right sense of shear.Annealed quartz band and dusty albite and apatite present.(d) Puga Gneiss Sample 05-02.Fine-grained albite lens between mica sheets.

Figure 3 .
Figure 3. Puga Gneiss Sample 05-02 garnet.(a) The brightest grains from the Ti map (yellow), Ca map (red), Fe map (green), Y map (cyan) and Al map (purple) are overlain on an Y map adjusted for Y-zonation in garnet.The location of xenotime grains can be identified based on the brightest Y locations to correspond with the micaceous rim after garnet.The map also highlights the orientation of rutile inclusion trails and the location of rectangular rutile clusters.The approximate outline of the original garnet is highlighted as a red dashed line.(b-e) Yttrium, manganese, magnesium and calcium EPMA element maps.Location of the core-rim boundary defined by the yttrium high.(f) Garnet zoning profiles from core (left) to rim (right).(g) Ternary diagram of transect across garnets.Approximate location of garnet transect shown by white dashed line on panel (e).

Figure 5 .
Figure 5. Garnet maps from mafic eclogite sample a03-16.(a) Ti map showing clusters of rutile inclusions forming distinct shapes after Fe-Ti oxides.b-e.Mg, Fe, Mn, and Ca maps showing concentric zoning patterns cut by Mg-Fe embayments emanating from the outside rim of the garnet.(f) Ternary diagram of transect across garnet showing Fe-Mg alteration and garnet rim.

Figure 6 .
Figure 6.(a) U-Pb Wetherill plot showing U-Pb zircon core analyses of Puga Gneiss sample 05-02.(b) Rank-plot of all 206 Pb/ 238 U zircon date s.An older age plateau represents the zircon cores at ∼480 Ma.The youngest dates represent thin rims, and converge on a Himalayan age with a weighted average 206 Pb/ 238 U age of the three zircon dates.(c) Tera-Wasserburg plot of xenotime U-Pb analyses.(d) Tera-Wasserburg plot of rutile U-Pb analyses.Box heights and uncertainty ellipses are 2σ uncertainties.

Figure 8 .
Figure 8.(a) Zircon data for individual samples from this study, St-Onge et al. (2013) and Donaldson et al. (2013).Blue-yellow color scale represents Dy/Yb.Solid colors do not contain individual REE spot data.Sample histograms shown in gray for Donaldson et al. (2013).(b) Time chart for Tso Morari and Kaghan, after(Palin et al., 2012), comprising data from multiple sources, including this study.

Figure 9 .
Figure 9. UHP exhumation at 47-46 Ma and its temporal relationship with the metamorphic, magmatic, and plate kinematic evolution of the Himalayan orogeny and India-Australia-Eurasia plate network.(a) Metamorphic and magmatic events in the Himalayan orogeny.(b) Orogenic regime of the Himalayan orogeny: onset of UHP exhumation marks the transition from an accretionary orogen to a collisional orogen (cf.,Cawood et al., 2009).(c) Reorganization of the India-Australia-Eurasia plate network at 47-36 Ma with plate velocity profiles for the Indian and Australian plates(Doubrovine et al., 2012;Müller et al., 2019;Torsvik et al., 2008) plotted using GPlates(Müller et al., 2018).Data sources for events in (a) and (c) are cited in the main text.(c) Is modified afterParsons et al. (2021).EOB-Eastern Ophiolite Belt; HMC-Himalayan Metamorphic core; MMB-Mogok Metamorphic Belt.
Figure 9. UHP exhumation at 47-46 Ma and its temporal relationship with the metamorphic, magmatic, and plate kinematic evolution of the Himalayan orogeny and India-Australia-Eurasia plate network.(a) Metamorphic and magmatic events in the Himalayan orogeny.(b) Orogenic regime of the Himalayan orogeny: onset of UHP exhumation marks the transition from an accretionary orogen to a collisional orogen (cf.,Cawood et al., 2009).(c) Reorganization of the India-Australia-Eurasia plate network at 47-36 Ma with plate velocity profiles for the Indian and Australian plates(Doubrovine et al., 2012;Müller et al., 2019;Torsvik et al., 2008) plotted using GPlates(Müller et al., 2018).Data sources for events in (a) and (c) are cited in the main text.(c) Is modified afterParsons et al. (2021).EOB-Eastern Ophiolite Belt; HMC-Himalayan Metamorphic core; MMB-Mogok Metamorphic Belt.

Figure 10 .
Figure10.The geodynamic significance of UHP exhumation for orogenesis and plate tectonics.(a and b) Schematic cartoon section of the India-Asia collision showing the postulated shallowing of Indian slab dip (e.g.,Chen et al., 2022), marked by the onset of UHP exhumation at 47-46 Ma (not to scale).(a) Steep subduction and UHP metamorphism during onset of the India-Asia collision (∼50-47 Ma).Prior to collision, the Asian active continental margin represented an Accretionary Orogen (cf.Cawood et al., 2009) in which convergence was driven and accommodated by subduction of the Indian and/or Neotethys slab.(b) UHP exhumation and slab shallowing starting at 47-46 Ma, in response to the positive buoyancy of Indian continental crust following the onset of continental collision.The impacts of the reduced slab dip are annotated on (b), showing an increase in crustal shortening, Barrovian metamorphism, crustal contamination of arc magmatism and lower crustal melting of the Gangdese arc to produce adakite.The collision zone represents a Collisional Orogen (cf.Cawood et al., 2009), whereby convergence is accommodated by crustal shortening ± continental subduction and driven by subduction of the Australian slab to the east (see Section 6).Two alternative slab break-off scenarios are presented.Slab Scenario 1 (presented in (a)): UHP exhumation and slab dip shallowing occurred in response to slab break-off at 47-46 Ma.Slab Scenario 2 (presented in (b)): Alternatively, UHP exhumation occurred via detachment of the Tso Morari complex from the Indian slab, whilst the buoyancy of the Indian continental crust resulted in lateral slab dragging, overturning of the trailing oceanic slab and reduction in slab dip (seeParsons et al., 2021 for further discussion of lateral slab migrations during the India-Asia collision), with slab break-off occurring at a later time (see Section 6).(c) 46 Ma plate reconstruction showing tectonic events of the Eocene plate network reorganization, triggered by changes in the subduction dynamics of the India-Asia collision.UHP exhumation at 47-46 Ma provide a precise timestamp linking changes in the subduction dynamics of the India-Asia collision zone to plate kinematic changes of the wider plate network.Reconstruction drawn fromMüller et al. (2019) using GPlates(Müller et al., 2018), with Greater India and Eurasian margin subduction zone outlines constrained byParsons et al. (2020Parsons et al. ( , 2021)).Thick dashed black lines show present day locations of the Indus Suture Zone (ISZ) and Main Frontal Thrust (MFT) of the Himalayan orogen, requiring 1,000-2,000 km of crustal shortening since ∼47-46 Ma(Parsons et al., 2021).Abbreviations: AFR-African plate; ANT-Antarctic plate; AUS-Australian plate; CIR-Central Indian ridge; GI-Greater India; IND-Indian plate; ISZ-Indus Suture Zone; ja-Java; khl-Khoistan-Ladakh block; lh-Lhasa block; MFT-Main Frontal Thrust; mk-Mogok metamorphic belt; wb-West Burma block; WR-Wharton ridge.
al. dates have reasonable agreement with our 47-46 Ma age, however, those data were previously interpreted as a record of UHP BIDGOOD ET AL.