Structural variations in basal decollement and internal deformation of the Lesser 2 Himalayan Duplex trigger landscape morphology in NW Himalayan interiors ’ 3

Saptarshi Dey, Rasmus Thiede, Arindam Biswas, Pritha Chakravarti and Vikrant Jain 4 Earth Science Discipline, IIT Gandhinagar, Gandhinagar-382355, India. 5 Institute of Geosciences, Christian Albrechts University of Kiel, Kiel-24118, Germany. 6 3 Department of Applied Geology, IIT-ISM Dhanbad, Jharkhand-826004, India. 7 Corresponding author: Saptarshi Dey; saptarshi.dey@iitgn.ac.in 8 Declaration: All photographs provided in the supplement is ether taken by Saptarshi Dey or 9 Rasmus Thiede in the year of 2019 and 2020. 10 11

1. Throughout the paper the authors are unclear or ambiguous how they characterize and interpret patterns associated with "faulting", "growth", or "active surface faulting" across the Kishtwar Window. For instance their cross section and map figures clearly indicate their interpretation of an out-of-sequence surface breaking fault, however, in the text their interpretation is unclear going back and forth between implying active faulting on MHT crustal ramps (no surface faulting), to active duplex growth, to active out-of-sequence deformation (surface faulting) that links to a crustal ramps. The authors need to clarify and revise many of the structural terms and interpretations being used in the text to streamline their interpretations.
2. The authors need to provide more justification in their interpretation that morphometric indices provide evidence for or against active out-of-sequence faulting. In many places, the authors jump to their preferred model but failed to recognize that these morphometric indices or other structural data are non-unique! In other words, the authors do not provide enough justification why or why not the pattern observed can be attributed to a specific deformation pattern. The authors have a tendency to have model-driven interpretations and do not justify why other structural models or non-tectonic controls are not permissible. For instance, many of the arguments used by the authors as "evidence" of an active out-of-sequence fault within the KW are not justified and highly speculative. Many

Introduction
Protracted convergence between the Indian and the Eurasian plate resulted into the growth and evolution of the Himalayan orogen and temporally in-sequence formation of the Southern Tibetan Detachment System (STDS), the Main Central Thrust (MCT), the Main Boundary Thrust (MBT) and the Main Frontal Thrust (MFT) towards the south (Yin and Harrison, 2000;Yin, 2006;Mukherjee, 2013). All these fault-zones emerge from a low-angle basal decollement, viz. the Main Himalayan Thrust (MHT) which forms the base of the Himalayan orogenic wedge (Ni and Barazangi, 1984;Nabelek et al., 2009;Avouac et al., 2016). The MHT was probably established in the late Miocene (Vannay et al., 2004).
The majority of scientists have favored that the late Pleistocene-Holocene shortening is mostly accommodated within the southern fringe of the Himalayan wedge, i.e., the Sub-Himalaya (morphotectonic segment in between the MBT and the MFT) (Wesnousky et al., 1999;Lave and Avouac, 2000;Burgess et al., 2012;Thakur et al., 2014;Mukherjee, 2015;Vassalo et al., 2015;Dey et al., 2016;Dey et al., 2019). This implies that the northerly thrusts, i.e., the MBT and the brittle faults exposed in the vicinity of the southern margin of the Higher Himalaya, are considered inactive over millennial timescales. However, in recent years, several studies that focused on the low-Temperature thermochronologic data and thermal modeling of the interiors of the NW Himalaya have raised questions on this. The recent studies suggested that 10-15% of the total Quaternary shortening has been accommodated within the interiors of the Himalaya as out-of-sequence deformation (Thiede et al., 2004;Deeken et al., 2011;Thiede et al., 2017;Gavillot et al., 2018) ( Supplementary   Fig. B1). Earlier, out-of-sequence deformation of the Himalayan wedge has been explained by two end-member models-(a)the reactivation of the MCT (Wobus et al., 2003), or, (b) enhanced rock uplift over a major ramp on the MHT (Bollinger et al., 2006;Herman et al., 2010;Robert et al., 2009). Landscape evolution models, structural analysis and thermochronologic data from the interior of the Himalaya favor that the Lesser Himalaya has formed a duplex at the base of the southern Himalayan front by sustained internal deformation since late Miocene (Decelles et al., 2001;Mitra et al., 2010;Robinson and Martin, 2014). The growth of the duplex resulted in the uplift of the Higher Himalaya and established the major topographic and orographic barrier as seen today. The Kishtwar Window (KW) in the NW Himalaya represents the north-western termination of the Lesser Himalayan Duplex (LHD). While most of the published cross-sections of the Himalayan orogen today recognize the duplex (Webb et al., 2011;Mitra et al., 2010;DeCelles et al., 2001), usually very little or no data is available whether the duplex is active over millennial timescales, and potentially a source of major Holocene earthquakes.
The pioneering low-temperature thermochron study by Kumar et al. (1995) portrayed the first orogen-perpendicular sampling traverse extending from the south-western margin of the Kishtwar tectonic Window crossing the Zanskar Range. More recent studies link the evolution of the KW to the growth of the LHD . Along its margin, it is surrounded by the Miocene MCT shear zone along the base of the High Himalayan Crystalline Sequence (HHCS). Locally, the bounding fault zone is named the Kishtwar Thrust (KT). Thermochronological constraints suggest higher rates of exhumation within the KW (3.2-3.6 mm/y) . Their findings corroborate well with similar thermochron-based Quaternary exhumation rates published from the of the Kullu-Rampur window in eastern Himachal Pradesh ((Jain et al., 2000;Vannay et al., 2004;Thiede et al., 2004;Stübner et al., 2018). In contrast, geodetic shortening rates lack spatial resolution and only capture inter-seismic deformation (Banerjee and Burgmann, 2002;Kundu et al., 2014), and there exists no chronological data to provide information on ongoing tectonic activity in the interiors of the Himalaya over intermediate timescales. Therefore, to understand the 10 3 -10 4 -year timescale neotectonic evolution, either we have to have geological field evidence, chronologically-constrained geomorphic markers or at least have a rigorous morphometric analysis of potential study areas, such as the KW.
In this study, we will focus on a few long-standing questions on Himalayan neotectonic evolution, which are- To address these questions, we adopted a combination of methods such as morphometric analysis using high-resolution digital elevation models, field observation on rock type, structural variations as well as rock strength data and analysis of satellite images to determine channel width and assessing the spatial distribution and relative differences in the late Quaternary deformation of the KW and surroundings (Fig.1). We aimed to test if the landscape morphology can be explained by changes in the geometry of the basal decollement.
We used basinwide steepness indices and specific stream power calculation (derived from channel gradient and channel width) as a proxy of the fluvial incision. And, lastly but most importantly, we calculated the fluvial bedrock incision rates by using depositional ages of aggraded sediments along the Chenab River. In this study, we show that the regional distribution of faulting is concentrated in the core and along the western margin of the window. We propose that active faulting within the LH Duplex is controlling the ongoing deformation in the Himalayan interior and driving the uplift of Higher Himalaya in its hanging wall. Our new estimates on the bedrock incision rate agree with Quaternary exhumation rates from the KW, which mean consistent active growth of the duplex over million-year to millennial timescales.

Geological background and field observations
The orogenic growth of the Himalaya is defined by overall in-sequence development of the orogen-scale fault systems which broadly define the morphotectonic sectors of the orogen ( Fig. 1b). Notable among those sectors, the Higher Himalaya is bordered by the MCT in the south and is comprised of high-grade metasediments (Haimanta Formation), Higher Himalayan Crystalline Sequence (HHCS), and Ordovician granite intrusives (Yin and Harrison, 2000). The Low-grade metasediments (quartzites, phyllites, schists, slates) of the Proterozoic Lesser Himalayan sequence are exposed between the MCT in the north and MBT in the south. The Lesser Himalayan domain is narrow (4-15 km) in the NW Himalaya except where it is exposed within tectonic windows (Kishtwar window, Kullu-Rampur window etc.) (Steck, 2003). The Sub-Himalayan fold-and-thrust belt lying to the south of the MBT is tectonically the most active sector since the late Quaternary (Thakur et al., 2014;Vignon et al., 2016).
Near the southwest corner of our study area, Proterozoic low-grade Lesser Himalayan metasediments are thrust over the Tertiary Sub-Himalayan sediments along the MBT (Wadia, 1934;Thakur, 1992). Near the Chenab region, Apatite U-Th/He ages suggest that cooling and exhumation related to faulting along the MBT thrust sheet initiated before ~5 ± 3 Myr (Kumar et al., 1995). Geomorphic data obtained across the MBT in Kashmir Himalaya suggest that MBT has not been reactivated for the last 14-17 ky (Vassallo et al., 2015). In the NW Himalaya, the Lesser Himalayan sequence (LHS) exposed between the MBT and the MCT is characterized by a < 10 km-wide zone of sheared schists, slates, quartzites, phyllites and Proterozoic intrusive granite bodies (Bhatia and Bhatia, 1973;Thakur, 1992;Steck, 2003). The LHS is bounded by the MCT shear zone in the hanging wall. The MCT hanging wall forms highly deformed nappe exposing lower and higher Haimantas, which are related to the Higher Himalayan Crystalline Sequence (HHCS) (Fig.2a) (Bhatia and Bhatia, 1973;Thakur, 1992;Yin and Harrison, 2000;Searle et al., 2007). Nearly 40 km NE of the frontal MCT shear zone, the MCT fault zone is re-exposed in the vicinity of KW and is called the Kishtwar Thrust (KT) (Ul Haq et al., 2019) (Fig. 1a, 2b). Within the KW, Lesser Himalayan Rampur quartzites (Fig.2c), low-grade mica schists and phyllites along with the granite intrusives are exposed (Steck, 2003;DiPietro and Pogue, 2004;Yin, 2006;Gavillot et al., 2018).KW exposes a stack of LHS nappes in the footwall of the MCT (in this case, KT) which is related to the Lesser Himalayan Duplex (LHD), characteristic of the central Himalaya (Decelles et al., 2001). Regionally balanced cross-sections (DiPietro and Pogue, 2004;Searle et al., 2007;Gavillot et al., 2018) suggest that the Himalayan wedge is bounded at the base by a low-angle décollement, namely as the MHT. Sub-surface structural formations beneath the KW are not well-constrained. A recent study by Gavillot et al. (2018) propose the existence of two mid-crustal ramp segments beneath the KW, viz., MCR-1 and MCR-2 (Fig. B2). Based on thermochronological constraints, Gavillot et al. (2018) and Kumar et al. (1995) proposed that the core as well as the western margin of the window exhumed with rates ~ 3.2-3.6 mm/y during the Quaternary, at a higher rate when compared to the surroundings.
The Higher Himalayan sequence dips steeply away from the duplex (~65° towards west) (Fig.2a). The frontal horses of the LH duplex expose internally-folded greenschist facies rocks. Although at the western margin of the duplex, the quartzites stand sub-vertically ( Fig.2b), the general dip amount reduces as we move from west to east for the next ~10-15 km up to the core of the KW. Near the core of the KW, we observed highly-deformed (folded and multiply-fractured) quartzite and granites at the core of the KW (Fig.2d, 2e). We also observed deformed quartz veins of at least two generations, as well as macroscopic white mica. Here, the River is also very steep and narrow; the rock units are also steeply-dipping towards the east (~55-65°) and are nearly isoclinal and strongly deformed at places (Fig.2f).
Towards the eastern edge of the window, however, the quartzites dip much gently towards the east (~20-30°), and much lesser folding and faulting have been recognized in the field ( Fig.2g).
The broad, 'U-shaped' valley profile near the town of Padder at the eastern margin of the KW is in contrast with the interior of the window (Fig.3a). At the core of the KW, the Chenab River maintains a narrow channel width and a steep gradient (Fig.3b). The E-W traverse of the Chenab River through the KW is devoid of any significant sediment storage. However, along the N-S traverse parallel to the western margin of the KW, beneath the Kishtwar surface, ~150-170m thick sedimentary deposits are transiently-stored over the steeplydipping Higher Himalayan bedrock (Fig.3c). The height of the Kishtwar surface from the Chenab River is ~450m, which means ~280m of bedrock incision by the River since the formation of the Kishtwar surface. Along the N-S traverse of the River, epigenetic gorges are formed as a result of the damming of paleo-channel by the hillslope debris flow, followed by the establishment of a newer channel path (Ouimet et al., 2008;Kothyari and Juyal, 2013).
One example of such epigenetic gorge formation near the town of Drabshalla is shown in Fig.3d. Downstream from the town of Drabshalla, the River maintains narrow channel width (< 25 m) and flows through a gorge having sub-vertical valley-walls (Fig.3e). The tributaries originating from the Higher Himalayan domain form one major knickpoint close to the confluence with the trunk stream (Fig.3f). We have identified at least three strath surface levels above the present-day river channel, viz., T1 (280±5 m), T2 (170-175 m) and T3 (~120±5 m), respectively (Fig.3g). The first study on sediment aggradation in the middle Chenab valley (transect from Kishtwar to Doda town) was published by Norin (1926). He

3.1.Morphometry
For conducting the morphometric analysis, we have used 12.5m ALOS-PALSAR DEM data (high resolution terrain-corrected) (Fig.5a). This DEM data has lesser issues with artifacts and noises than 30m SRTM data, which fails to capture the drainage network properly in areas populated by narrow channel gorges. We compiled the topographic relief over a circular moving window of 4 km diameter (Fig.5b) and the rainfall distribution of the Chenab region ( Fig.5c). The rainfall distribution is adapted from 12-year-averaged annual rainfall data from TRMM database (Bookhagen and Burbank, 2006).

Basinwide normalized steepness indices
Global observations across a broad spectrum of tectonic and climatic regimes have revealed a power-law scaling between the local river gradient and upstream contributing area: where S is the stream gradient (m/m), k s is the steepness index (m 2θ ), A is the upstream drainage area (m 2 ), and θ is the concavity index (Flint, 1974;Whipple and Tucker, 1999).
Normalized steepness-index values (k sn ) are steepness indices calculated using a reference concavity value (θ ref ), which is useful to compare steepness-indices of different river systems (Wobus et al., 2006). We extracted the k sn values in the study area using the ArcGIS and MATLAB-supported Topographic Analysis Toolkit (Forte and Whipple, 2019) following the procedure of Wobus et al. (2006). We performed an automated k sn extraction using a critical area of 10 6 m 2 for assigning the channel head, a smoothing window of 500 m, a θ ref of 0.45, and an auto-k sn window of 250 m for calculating k sn values. Stream-specific ksn values in and around the KW are drafted in Supplementary Fig.B3. The catchments were delineated by using a maximum threshold of 200 sq. km, so that the basins we pick are smaller in size. The stream-specific k sn values were rasterized in ArcGIS and were extrapolated to the respective catchments using the zonal statistics toolbox. Basinwide mean k sn values for the delineated watersheds are portrayed in Fig.5d.Basinwide mean k sn values are plotted using a 500 km 2 threshold catchment area (Fig. 2d).
A 50-km-wide swath profile along line AB (cf. Fig.5a) show variation in elevation, mean annual rainfall and mean k sn values in the area (Fig.2e).

Drainage network extraction
The drainage network and the longitudinal stream profiles were extracted using the Topographic Analysis Kit toolbox (Forte and Whipple, 2019). An equivalent of 20-pixel smoothing of the raw DEM (250 m smoothing window) data has been applied to remove noises from the DEM (Fig.6). The longitudinal stream profile of the Chenab trunk stream was processed with the Topotoolbox 'Knickpointfinder' tool (Schwanghart and Scherler, 2014).
Several jumps/ kinks in the longitudinal profile are seen and those are marked as knickpoints ( Fig.6). A 30m tolerance threshold was applied to extract only the major knickpoints. Results from knickpointfinder tool were rechecked with chi vs. elevation distribution ( Supplementary   Fig.B4). Identification of the knickpoints/ knickzones and their relationship with the rocktypes as well as with existing structures are necessary to understand the causal mechanism of the respective knickpoints/ knickzones. Knickpoints/(zones) can be generated by lithological, tectonic and structural control. Lithological knickpoints are stationary and anchored at the transition from the soft-to-hard substrate. The tectonic knickpoints originate at the active tectonic boundary and migrate upstream with time. Structural variations, such as ramp-flat geometry of any emerging thrust may cause a quasistatic knickpoint at the transition of the flat-to-ramp of the fault. In such cases, the ramp segment is characterized by higher steepness than the flat segment and often the ramp is characterized by a sequence of rapids, forming a wide knickzone, instead of a single knickpoint.
Longitudinal profile of the entire Himalayan traverse of the Chenab River show oversteepening across the KW (Fig.6), therefore, we focused on that segment (marked by red rectangle, cf. Fig.6) for further analysis. Longitudinal profile of the selected segment is shown in Fig.7a.

Channel Width
Channel width is a parameter of assessment of lateral erosion/incision through bedrocks of equivalent strength (Finnegan et al., 2005;Turowski, 2009). The channel width of the Chenab trunk stream within the elevation range of 600 to 2200 m above the MSL was derived by manual selection and digitization of the channel banks using the Google Earth Digital Globe imagery (http://www.digitalglobe.com/) of minimum 3.2 m spatial resolution. We used the shortest distance between the two banks as the channel width and rejected areas having largely unparallel channel-banks as that would bias the result. We used a 100 m step between two consecutive points for channel width determination. Ten point-averaged channel width data along with elevation of the riverbed is shown in Fig.7b. Variations in channel gradient and k sn values along the longitudinal profile of the selected stretch are shown in Fig.7c and 7d, respectively.

Specific stream power (SSP) calculation
Specific stream power has often been used as a proxy of fluvial incision or differential uplift along the channel (Royden and Perron, 2013;Whipple and Tucker, 1999). Areas of higher uplift/incision are characterized by a transient increase in the specific stream power. Channel slope and channel width data were used to analyze the corresponding changes in the specific stream power (SSP) from upstream of the gorge area to the gorge reaches (Bagnold, 1966).
The SSP (ω) was estimated using the following equation - Where, γ -unit weight of water, Q -water discharge, s -energy slope considered equivalent to the channel slope; w -channel width. With the available TRMM data, we argue that the rainfall distribution in the study area is almost uniformly low (<1.5 m/y) ( Fig.5c and 5e) and therefore, we assumed a uniform discharge (Q) for SSP calculation. SSP data from selected stretches (stretch 1 and stretch 2, cf. Fig.8a) are shown in Table 1.

Structural data
We measured the strike and dip of the foliations and bedding planes of the Lesser and Higher Himalayan rocks using the Freiberg clinometer compass. We took at least five measurements at every location, and the average has been reported in Fig. 8a. Field photos in Fig.2 document the observed variations in the structural styles.

Rock strength data
Recording rock strength data in the field is essential to understand the role of variable rocktype and rock-strength in changes in morphology. It provides us vital insights on the genesis of knickpoints, whether they are lithologically-controlled or not. It also helps to understand the variations in channel steepness across rocks of similar lithological strength. We systematically measured the rock strength of the major geologic units using a hand-held rebound hammer. Repeated measurements (8-10 measurements at each of the 75 locations throughout the study area) were conducted to measure the variability of rock-strength within the major lithologic units. All the measurements were taken perpendicular to the bedding/ foliation plane, and no measurements are from wet surfaces or surfaces showing fractures.
Each reading was taken at least 0.5m apart from the previous one. Average rock strength data collected from each of the test locations are plotted against the longitudinal river profile and channel width data in Fig.7e. Our data from individual sites are smaller in number than what is preferred for checking the statistical robustness of Schmidt hammer data (Niedzielski et al., 2009). Therefore, we combined the data from all sites representing similar lithology and portrayed the mean ±standard deviation for the same. Field data provided in Supplementary   Table C1.

Luminescence dating of transiently-stored sediments in and around Kishtwar
Luminescence dating of Quaternary fluvial sediments is a globally accepted method for constraining the timing of deposition of sediments in a drainage system (Aitken, 1992;Olley et al., 1998;Wallinga et al., 2001;Cunningham and Wallinga, 2012). Although there exists a few persistent problems in luminescence dating of the Himalayan sediments (including low sensitivity of quartz and numerous cases of heterogeneous bleaching of the luminescence signal), studies over the past couple of decades have also provided an adequate control on Himalayan sedimentary chronology by using luminescence dating with quartz (Optically stimulated luminescence, OSL) and feldspar (Infra-red stimulated luminescence, IRSL).
Earlier studies have reported sediment aggradation over the Higher Himalayan bedrocks in the Kishtwar valley (Norin, 1926;Ul Haq et al., 2019).
The samples for luminescence dating were collected in galvanized iron pipes. The pipes were opened in subdued red light (wavelength ~650 nm). The outer ~3 cm of sediment from both the ends of the pipe were removed to omit the possibility of exposure of the sample to daylight during collection. The removed portion was used for moisture content estimation and determination of Uranium (U), Thorium (Th), and Potassium (K) concentrations. The unexposed interior portion of the sample was further processed to obtain quartz and feldspar using standard procedures (e.g., Aitken, 1998). The portion was treated with a sufficient quantity of 1N HCl and 30% H 2 O 2 to remove carbonates and organic materials, respectively.
The sediments were then oven-dried at 45°C and sieved to obtain a size fraction of 90-150 μm. The quartz and feldspar were separated using Frantz isodynamic separator at a magnetic field of ~10,000 gauss and collected separately. Obtained quartz grains were etched with 40% HF for 80 minutes to remove alpha irradiated outer layer (~10 μm), followed by 37% HCl treatment for 20 minutes to dissolve fluorides formed during the previous step. The isodynamic separation procedure was repeated to remove any broken feldspar grain.
However, even after repeating the last step twice, we were unable to eliminate the feldspar contamination from most of the samples thoroughly. Those samples are not suitable for OSL SAR protocol.
Samples K02 and K11 procured from the fine-grain layers of ~1-1.5m thickness, trapped within coarse, angular and poorly-sorted thick layers of clasts (identified as hillslope debris) were used for OSL (Optically stimulated luminescence) dating using Double SAR (Single Aliquot Regenerative) protocol (IRSL wash before OSL measurement) for equivalent dose estimation (Roberts, 2007). The test doses were set for 75 Gy, 225 Gy, and 450 Gy, respectively (Fig.5). The aliquots were considered for ED estimation only if: (i) recycling ratio was within 1±0.1, (ii) ED error was less than 20%, (iii) test dose error was less than 10%, and (iv) recuperation was below 5% of the natural. For samples K16 and K17 (fluvial sand trapped above the T3 strath surface), the feldspar contamination was negligible.
Therefore, the OSL SAR protocol was tried with test doses of 50 Gy, 100 Gy, and 150 Gy, respectively. Samples K16 and K17 returned highly scattered equivalent dose (De) estimates (over-dispersion > 30%) (cf. Table 2), and thus, both of them have been interpreted by the minimum age model (Bailey and Arnold, 2006). Sample K18 (from the silty clay layer found above the T1 strath surface in the wind gap of Maru River) (cf. Fig.9b) was saturated, and hence, we also provided the minimum age estimate for the same. The sample was exhausted after we performed OSL measurements. Therefore, we couldn't proceed towards feldspar dating with sample K18.
OSL dating for the three samples procured from the fluvio-glacial sediments showed saturation; therefore, we tried for IRSL (Infra-red stimulated luminescence) dating of feldspar for those three samples (K07-K09) using standard post infrared (pIR-IR) protocol (Buylaert et al., 2013), in which, the preheat temperature was 320°C for 60s. The samples were first stimulated at 50°C with IR diodes for 100s followed by IR stimulation at 290°C, and a violetblue luminescence emission (395 ± 50 nm) was detected by PMT through the combination of optical filters, Corning 7-59 (4 mm) and BG-39 (2 mm Supplementary   Fig.B5.

4.1.Field observations and measurements
The Chenab River has deeply incised the KW (Fig. 3a). The LH metasediments exposed within the KW are mainly composed of Rampur Quartzites (Fig.2b,2d) and phyllites with occasional schists in between. (Steck, 2003;Gavillot et al., 2018). The LHD has been suggested to be an asymmetric antiformal stack with a steeper western flank (dip: 70°/west) (Fig.8a). The KW is surrounded by rock units related to the Higher Himalayan high-grade metasedimentary sequence (HHCS), mainly garnet-bearing mica schists and gneisses ( Fig.2a). Higher Himalayan rocks close to the western edge of the KW form a syncline with a southwest-verging MCT at its' base. The KT, southern structural boundary of the window margin, accommodating the differential exhumation between the window and the surroundings -and it is expressed as highly deformed sub-vertical shear bands (Fig.2b).
Along the traverse of the Chenab River through the window and further downstream, two prominent stretches of ~20 and ~25-30 km length have been identified where the channel gradients are high (Fig.7c), and we observed a sequence of rapids (FFig.3a,3e). These steep segments are also characterized by a very narrow channel width (< 30m) (Fig. 7b). These two steepened segments define knickzone rather than a single knickpoint. We refer to the knickzone at the core of the window as K1 and the one downstream from the KW as K2 (Fig.6, 7). The knickzones are hosted over bedrock gorges, and field evidence confirms that none of them (downstream from the eastern edge of the KW) are related to damming by landslides or other mass movements. The east margin of the KW is characterized by a broad 'U-shaped' valley filled with thick sand layers and coarser fluvioglacial sediments. The River incises through this Late Pleistocene fill at present (FFig.3a).
The rock strength data taken along the Chenab River shows large variations across different morphotectonic segments (Fig.7e). Within the KW, Lesser Himalayan phyllites and schists have low R values (30-35); however, the low-strength schists and phyllites are sparsely present and therefore, they were ignored while plotting the regional rock strength values in

Steep stream segments and associated knickpoints
The longitudinal stream profile along the Chenab River does not portray a typical adjusted concave-up profile across the Himalaya (Fig. 6,7a). We observed breaks in slope and concavity at least at six occasions within a ~170 km traverse upstream from the MBT (point A, cf. Fig.1a) (Fig.7a). These breaks are defined as knickpoints or knickzones, depending on their type characteristics. The slope breaks represent the upstream reaches of the steep stream segments. The basinwide steepness indices span from ~30->550 m 0.9 across the study area ( Fig. 5d). We assigned a threshold value of k sn >550 for the steepest watersheds/ stream segments. Along the traverse, the major knickpoints are L1 (~1770m), K1 (~1700m), K2 (~1150m), K3 (~950m), L2 (~750m) and D1 (~700m) respectively (Fig.7a).
Already Nennewitz et al. (2018) had proposed a high basin-averaged k sn value of > 300 in the KW. Here in this study, we worked with a much-detailed DEM for stream-specific k sn allocation (FFig.5d), as well as a basinwide steepness calculation with smaller watersheds.
Our results corroborate with the earlier findings, but predicts the zone of interest in greater detail. It is important to note that by setting a higher tolerance level in the 'knickpointfinder' tool in Topotoolbox, we have managed to remove most of the DEM artifacts from consideration (Schwanghart and Scherler, 2014).

Channel width and valley morphology
The channel width of the Chenab river is on average low (30-60m) within the core of the KW (Fig. 7b), and the low channel width continues till the Chenab River flows N-S along the western margin of the KW. However, there are a few exceptions; upstream from the knickpoint L1 in the Padder valley (in which the town of Padder is located), the channel widens (width ~80-100m), and the channel gradient is low (Fig. 3a,7b,7c). The second instance of a broader channel is seen upstream from knickpoint K2, where there is a reservoir for the Dul-Hasti dam (Fig.7b). Downstream from K2 within the Higher Himalaya, the channel width ranges from 50-70 m. However, towards the lower stretches of the N-S traverse, the width is even lower (16-52m). The river width increases to 100-200m as Chenab River takes a westward path after that. The river width increases beyond 300m until it leaves the crystalline rocks in the hanging wall of the MCT and enters the Lesser Himalaya in the hanging wall of the MBT across the Baglihar dam. Within the frontal LH, the channel width is again lowered (50-80 m).

RRChanges in specific stream power (SSP)
Discharge-normalized SSP data calculated from the upstream stretches and the knickzones, K1 and K2, show a significant increase in SSP within the steep knickzones. The rise in SSP from upstream to the knickzones K1 and K2 is 4.44 and 5.02 times, respectively (Table 1).
Such a high increase in SSP is aided by the steepening of channel gradient and narrowing of the channel.

Luminescence chronology
The results for the luminescence chronology experiment are listed in Table 2 (Fig.4c). The initial IRSL ages (before fading-correction), therefore, may be regarded as a minimum age estimate for the fluvioglacial sediment sequence. The finer fraction of the hillslope debris overlying the fluvioglacial deposits yields OSL ages of 81.1±4.6 ky (K02) and 85±5 ky (K11) (Fig.4d, 4e).
OSL samples taken from sparsely-preserved sediment layers above the T3 strath surface show heterogeneous bleaching, and hence we provide a minimum age of 22.8±2.1 ky (sample K16) and 20.5±1.0 ky (sample K17). One sample taken above the T1 strath level is saturated and shows a minimum age of 52.1±2.8 ky (sample K18) ( Table 2).

Discussions
Morphometric parameters are widely used as indicators of active tectonics and transient topography (Hack, 1973;Kirby and Whipple, 2012;Seeber and Gornitz,1983). Many studies have used morphometry as a proxy for understanding the spatial distribution of active deformation across specific segments of the Himalayan front ( Young AFT cooling ages have been interpreted as the result of rapid exhumation of the LH duplex over million-year timescale (Kumar et al., 1995;Gavillot et al., 2018). However, to date, we lack any estimate of deformation on the 10 3 -10 5 -year timescale. Thus, we have come up with a detailed morphometric analysis of the terrain and structural data to decipher the spatial distribution of faulting and fault patterns. With additional chronological constraints from late Quaternary sedimentary deposits, we predict rapid fluvial bedrock incision in the Himalayan interiors.

Knickpoints and their genesis
Already Seeber and Gornitz (1983) showed that the Chenab River is characterized by a zone The longitudinal profile of the lower Chenab traverse (below ~2000 m above MSL) is punctuated by two prominent stretches of knickpoint zones (Fig.7a). Below we will discuss the potential cause of formation of those major knickpoints in the context of detailed field observation, of existing field-collected structural and lithological data, geomorphic features, rock strength, and channel width information (Fig.3b).

Lithologically-controlled knickpoints
The Himalayan traverse of the Chenab River is characterized by large variations in substrate lithology and rock strength (Fig.7a). These variations have inflicted their 'marks' on the river profile. An instance of soft-to-hard substrate transition happens across the knickpoint L1, lying downstream from the Padder valley, at the eastern edge of the KW (Fig.3a, 7). Across L1, the River enters the LH bedrock gorge (R value> 55) after exiting the Padder valley filled with unconsolidated fluvioglacial sediments (Fig.3a). A similar soft-to-hard substrate transition is observed upstream from the MCT shear zone. The corresponding knickpoint L2 represents a change in lithological formation from the sheared and deformed Higher Himalayan crystalline (R value~35-40) to deep-seated Haimantas (R value~40-50) (Fig.7a).
There is no field evidence, such as fault splays or ramps, in support of L2 to be a structurallycontrolled one.

Tectonically-controlled knickpoints
Compiling previously-published data on regional tectonogeomorphic attributes (Gavillot et al., 2018) with detailed field documentation of structural styles and tectonic features, we have deciphered the role of rock-uplift and variable structural styles in the interiors of the NW Himalaya. We have found at least two instances where knickpoints are not related to change in substrate, nor are they artificially altered.
The knickzone K1 (~1700 m above MSL) represents the upstream reach of a steepened stream segment of run-length ~18-20 km. The upstream and downstream side of K1 is characterized by a change in the orientation (dip angle) of the foliation of the LH bedrock ( Fig.2f, 2g). Across K1, the dip amount of the foliation planes change from ~25-30° to ~60-65° (both cases dip towards east). K1 also reflects a narrowing of the channel width (Fig. 7b) and an increase in channel gradient (Fig.7c) and ksn value (Fig.7d). Near the end of the steep segment, we observed intensely-deformed (folded and fractured) LH rocks (Fig.2d, 2e). We explain this as evidence of faulting within the LH duplex and the steep stream segment represents the ramp of the fault or fault zone between two duplex nappes (Fig.8d). K1, therefore reflects the transition from flat to ramp of the existing structure soled to the basal decollement. The steep segment represents a drop of ~420m of the Chenab River across a run-length of ~20 km (Fig.8c). In addition to this, we may comment that the schists and phyllites within the Lesser Himalayan sequence probably act as the basal planes of the thrust nappes.
On the other hand, the other knickpoint K2 nearly coincides with the exposure of the KT (Fig.7a). K2 cannot be a lithologically-controlled knickpoint as it reflects no significant change in substrate hardness, at least not a soft-to-hard substrate transition. LH quartzites (R value: 51±4) and HH migmatites (R value: 49±5) have similar rock hardness (cf. Fig.7e).
However, in the longitudinal profile, K2 does not represent a sharp slope break because the downstream segment runs parallel for ~25-30 km and not perpendicular to the orientation of all major structures of the orogen, including the KT. Therefore, we performed an orthogonal projection of the E-W trending traverses of the Chenab River and tried to estimate an orogenperpendicular drop of the Chenab across K2 (Fig. 8d). The truncated profile across K2 shows a drop of ~230m of the channel across an orogen-perpendicular run-length of ~5 km. The orogen-parallel stretch of the River exhibits narrow channel width (<30-35m) through a moderately hard HH bedrock (R-value: 35-45). The tributaries within this stretch form significant knickpoint at the confluence with the trunk stream (Fig. 3f). These pieces of evidence hint towards a rapid uplift of the HH rocks along the western margin of the KT and are possibly related to the presence of another crustal ramp emerging from the MHT (Fig.8b).
Both the knickzones, K1 and K2 portray transiently-high specific stream power values (Table   1), which we relate to the fact that the knickzones are undergoing much rapid fluvial incision than the rest of the study area. If we consider the fluvial incision as a proxy of relative uplift (assuming a steady-state), we may well say that the knickzones define the spatial extent of the areas undergoing differential uplift caused by movement on the fault ramps.

Knickpoint marking epigenetic gorge
Epigenetic gorges are common geomorphic features in the high-mountain landscape (Ouimet et al., 2008). Epigenetic gorges form when channels of a drainage system are buried by sediment aggradation and during subsequent re-incision, a new river channel is incised. The N-S traverse of the Chenab River is largely affected by hillslope sediment flux from the steep eastern flank. The knickpoint K3 situated near the village of Janwas, mark one such instance of epigenetic gorge where the paleo-valley has been filled initially by fluvioglacial sediments and the channel abandonment was caused by hillslope debris flow ~80 ky (Fig.4b, 4c).

Sediment aggradation in the Chenab valley
The luminescence chronology of the transiently-stored sediments in the Chenab valley point towards protracted sediment aggradation since the onset of the last glacial-interglacial cycle till ~80 ky. Fluvioglacial outwash sediments range from ~110-130 ky, whereas the hillslope debris range from ~90 to ~80 ky (cf. Fig.4). The obtained chronology of the sediment deposits match well with the relative stratigraphic order of the sedimentary units, and therefore, we believe the obtained ages are reliable. The deeply-incised fluvial network as we observe today require net fluvial incision and formation of bedrock strath surfaces sometime after ~80 ky before present.

Drainage re-organization and strath terrace formation along Chenab River
Hillslope debris flow characterized by white quartzite blocks of different sizes and shapes can only originate from the high-relief frontal horses of the Lesser Himalayan Duplex exposed in the eastern flanks of the valley. These hillslope debris overlies the fluvioglacial sediments stored beneath the Kishtwar surface (Fig.9a, 9c). We argue that the hillslope debris flow intervened in the paleo-drainage of the Chenab River, which might have been flowing through an easterly path than now (Fig.9). The Maru River, coming from the north-western corner of our study area, was also joining the Chenab River at a different location (Fig.9).
Our argument is supported by field observation of the thick silt-clay layer in the proposed paleo-valley of the Maru River (Fig.9a, 9c). OSL sample (K18) from the silt-clay layer is saturated and hence only provides the minimum age of 52±3 ky. We suggest that the hillslope sediment flux has ceased the flow of the Chenab River and also propagated through the wind-gap of the Maru River. The decline in depositional energy has resulted in a reduction of grain-size. Post-hillslope debris flow episode, the Chenab River also diverted to a new path. The new course of the Chenab River upstream from the confluence with the Maru River is defined by a very narrow channel flowing through the Higher Himalayan bedrock gorge (Fig.9). Downstream from the confluence, we identified at least three levels of strath terraces lying at heights of ~280-290m (T1), ~170m (T2), and ~120m (T3), respectively (Fig.4g, 10a). Our field observation suggests that the formation of the straths is at least ~52 ky-old. The luminescence chronology samples in this study belong to the ~150-170m-thick soft sediments that are stored stratigraphically-up from the T1 strath level. Our field observations and chronological estimates suggest that the renewed path of the Chenab River must have been formed post the hillslope debris flow ~80-90 ky but before 52 ky.

4. Rapid bedrock incision along Chenab River
Considering the rate of excavation of softer sediments to be at least an order of magnitude higher than the rate of bedrock incision (Ouimet et al., 2008;Kothyari and Juyal, 2013), we calculated the minimum bedrock incision rate using the height of the T1 strath (~280±5 m) and the average age of the sediments from the Hillslope debris flow deposit. It yields a minimum bedrock incision rate of ~3.1-3.5 mm/y over the last 80-90 ky. Considering the saturated OSL sample from the paleo-valley, we estimated the maximum bedrock incision since 52 ky to be 5.1-5.5 mm/y. Similarly, using the minimum age estimate of the T3 terrace abandonment, we deduce a maximum bedrock incision rate of ~5.7-6.1 mm/y since ~21 ky (Fig.10b).
Assuming the channel hypsometry to be constant during the incision period, we may infer the minimum fluvial incision reflects regional rock uplift caused by movement of rocks over the ramp of the MHT at a rate of ~3.1-3.5 mm/y. This minimum uplift rate estimate is in agreement with long-term exhumation rates of 3.2-3.6 mm/y deduced from the KW . The inferred uplift rate can be translated to a shortening rate by using a simple trigonometric function. Our field findings suggest that the larger ramp on the MHT (MCR-1) have an average near-surface dip of ~60˚. Considering a similar geometry for MCR-2, we obtained a minimum shortening rate of 1.8-2.0 mm/y. On the other hand, considering the minimum ages of T3 terrace abandonment, we obtained maximum uplift rates ~5.5-6.0 mm/y, which would translate into a shortening rate of ~3.2-3.5 mm/y since ~21 ky.

Our findings in context with the previously-published data
The young AFT-cooling ages by the pioneering work of Kumar et al., (1995) showcased a rapid exhumation of the KW (AFT ages: ~1-3 My) compared to the surroundings (AFT age: 6-12 My). The exhumation rates proposed by Gavillot et al. (2018) is based on using a geothermal gradient of 35-40˚C/km in Dodson's equation assuming a 1-D model (Dodson, 1973). Unfortunately, a proper thermal modeling of the region is lacking, therefore, regional correlation with the proposed cooling ages and exhumation rates has large uncertainties.
However, lateral similarities of the regional topography and similar thermochronologic age patterns obtained along the Sutlej area, Beas and Dhauladhar Range (Thiede et al., 2017;Thiede et al., 2009;Stübner et al., 2018) have yielded exhumation rates in the range of 2-3 mm/y. Long-term exhumation rates from the NW Himalaya agree well with the findings of Nennewitz et al. (2018). Their study recognized a strong correlation between the young thermochron ages with high basinwide k sn values suggesting high uplift rates over intermediate to longer timescales. Therefore, the proposed range of long-term exhumation rates of 3.2-3.6 mm/y determined by Gavillot et al. (2018) agree with the regional data pattern. Although the geomorphic implications on landscape evolution are valid for shorter timescales than the low-T thermochron studies, we must comment that our field observations and analysis support a protracted growth of the LH duplex exposed within the boundaries of the KW. Unless there has been a recent growth of the duplex, the geomorphic signatures would have been subdued. Young low-T thermochron ages (Kumar et al., 1995) had been sampled from the steepened stream reaches, where the SSP is high. Interestingly, exhumation rates obtained from the steepened stretches is ~ten times more than that of the Higher Himalaya in the hanging wall of the duplex. Our estimates of SSP also reflect an increase by ~five times within the steepened stretches.
Deeply-incised channel morphology, steep channel gradients marked by knickpoints at the upstream reaches in and around the KW could be explained by the presence of at least two orogen-parallel mid-crustal ramps on the MHT (Fig.8b). The existence of two mid-crustal ramps have already been suggested by Gavillot et al., (2018) (Fig.S2). However, the internal structural orientation of the LH duplex published by Gavillot et al., (2018) (Fig. S2) differ considerably from our field observations (Fig.2, Fig.8b). Our morphometric analysis and field observation indicate pronounced deformation at the core of the KW suggesting that this is related to active faulting or internal folding at the base of the steepened stretch of K1F. The ramp of the fault-zone mentioned above triggers the rapid exhumation of the hanging wall. It causes high relief, steep channel gradients and higher basinwide steepness indices over the ramp (Fig.7). Similar ramps have been proposed on the MBT beneath the Dhauladhar Range (Thiede et al., 2017) and in the east of the NW Himalaya (Caldwell et al., 2013;Mahesh et al., 2015;Stübner et al., 2018;Yadav et al., 2019). We don't have any direct field evidence of surface-breaking faults, which could be linked to the MCR-2. However, a rapid fluvial incision along the western margin of KW and an increase in morphometric parameter values (k sn , channel gradient, topographic relief, etc.) probably justify the existence of MCR-2.
Our findings from the Kishtwar region of the NW Himalaya establish the importance of morphometric parameters in the assessment of intermediate timescales of 10 4 -10 6 years. We can resolve variations in the tectonic imprint on landscape evolution by analyzing the topography with high-resolution DEM. Earlier studies used to process larger areas, but the resolution of those data and findings is coarse (Nennewitz et al., 2018).
Models explaining the spatial distribution of the high uplift zone in the interiors of the Himalaya favor the existence of a mid-crustal ramp, which has variable dimension, geometry, and distance from the mountain front along-strike of the Himalayan orogeny (Robert et al., 2009). Our data support the idea of mid-crustal ramps beneath the Higher Himalayan domain (Nennewitz et al., 2018) and we predict that the seismic hypocenters are clustered in the vicinity of the ramp of MHT and within the LHD and are linked to the ongoing growth of the duplex. Our results verify the previously-suggested models that there exist two orogenparallel small ramps beneath the Kishtwar Window instead of one .
However, we must also comment that the previous model, as well as the balanced crosssection, lack detailing and the thermochron data (Kumar et al., 1995) is sparse. Therefore, field observation and the detailed morphometric analysis using high-resolution DEM help to measure the spatial extent of deformation. We can resolve the high-relief Kishtwar Window and the surroundings into two major steep orogen-parallel belts/ zones (Fig. 4a). While the larger one is an active high-angle fault-ramp emerging from the MHT and causing sustained uplift in the core of the duplex, the smaller one lies along the western margin of the KW. We suggest that this has two major implications. One, we have evidence for ongoing internal deformation of duplex, and that entire window is still tectonically-active -and therefore this could be a potential source future seismic activity. Our finding contradicts with the existence of a single major ramp in the interiors of the Himalaya, as described from other sectors of the Himalaya (Gahalaut and Kalpna, 2001;Elliot et al., 2016;Thiede et al., 2017). Nennewitz et al., (2018) have proposed that the million-year-timescale shortening achieved in the interior of the Himalaya near the Sutlej-Beas area in the eastern Himachal Pradesh is caused by accentuated rock uplift over a ramp at a mid-crustal depth of ~ 8-25 km on the MHT. In contrast, studies from the Dhauladhar Range in the north-western Himalaya hints the presence of deep-seated crustal ramp on the MBT and yielded a shortening rate of 3±0.5 mm/y across the MBT over the last 8 My and absence of mid-crustal ramp Thiede et al., 2017). The work by Gavillot et al. (2018) favors the existence of at least two mid-crustal ramps beneath the KW (Supplementary Fig.S2). Their suggestion is in agreement with very young AFT cooling ages (1-3 Ma) (Kumar et al., 1995) in the window (Fig.1a) and the findings of this study. These studies altogether point out the along-strike variation in the location of the rapidly-uplifting crustal ramp with respect to the southern Himalayan front. The crustal ramp in the nearby Kangra recess is located beneath the Dhauladhar Range at the main Himalayan front, whereas, in the Himalayan transects situated towards the east and west of Kangra recess, the ramps are located ~100km inside from the MBT. Topographic relief and basinwide mean k sn distribution (Fig.5) hint towards the existence of a lateral ramp in between the Kangra and the Jammu-Kashmir Himalayan transects. However, at this moment, we have no conclusive data in support of this claim.

Conclusions
Our

2.
Our field observations, morphometric analysis, and rock strength measurements document that at least two of these major knickzones on the trunk stream are non-lithologic and preferably can be related to differential uplift of the rock units. The incision potential in the steepened stretches ~4-5 times higher than the surroundings.

3.
The differential uplift is most-likely related to variations in the geometry of the basal decollement and out-of-sequence surface-breaking fault/s. Our results favor the presence of at least two mid-crustal ramps beneath the Kishtwar Window and the surroundings, as compared to a single crustal ramp proposed from interiors of the nearby sectors of the NW Himalaya.

4.
The larger of the proposed crustal ramps emerge as an active high-angle ramp at the core of the Lesser Himalayan Duplex and cause sustained faulting and uplift of the hanging wall.

5.
Luminescence chronology of the transiently-stored sediments along the Chenab River suggests that the valley had been overfilled with deposits of fluvioglacial origin as well as with hillslope debris flow. Massive sediment aggradation during ~130-80 ky led to drainage re-organization and bedrock incision, leaving behind a set of strath surfaces.

6.
The late Quaternary bedrock incision rates on the mid-crustal ramp beneath the western margin of the KW are high 3.1-3.6 mm/y. We argue that this rapid fluvial incision can potentially be linked to the accommodation of crustal shortening along a mid-crustal ramp of the basal decollement of the Himalaya.
Our results indicate a minimum shortening rate of ~1.8-2.1 mm/y being accommodated along the N-S traverse of the Chenab River.
Our study refutes the long-standing hypothesis of nearly 100% accommodation of crustal shortening within the Sub-Himalaya since late Pleistocene-Holocene time and provides new insights on the structural styles and ongoing out-of-sequence deformation in the Himalayan interiors.

Appendix
Additional maps, figures on morphometric analysis and luminescence dating are listed in Appendix A. Data of rock strength measurements provided in Table C1.

Code availability
Authors used open-source codes of Topotoolbox and Topographic Analysis Kit Toolbox for this study.

Data availability
Field data are already provided in Appendix 1. Additional data on luminescence dating can be provided on request.

Sample availability
Samples used for luminescence dating are already mostly-destroyed, therefore it is beyond sharing.

Author contribution
S.Dey, the first author , this work and completed the fieldwork, sample processing, measurements and writing of this manuscript. R. Thiede helped in fieldwork, discussion and writing of this manuscript. A. Biswas performed the initial morphometric analysis.
N.Chauhan helped in measurement of luminescence signal and assessment of the data.
P.Chakravarti performed the channel width calculations and compiled the rock strength measurements. V. Jain helped in discussion and writing of the manuscript.   Swath profiles (swath window: 50 km) along the line AB (cf. Fig.5a) demonstrate the orogenperpendicular variations in elevation, rainfall and ksn value. KW is characterized by high elevation, high relief and high steepness, but low rainfall. This suggests that tectonics control uplift and shape of topography not climate. River (For present-day glacial extent cf. Fig.1a). Further in this study, we focused on the area marked by red rectangle.   1a). The mean R-value ± σ for each rock type has been plotted against their spatial extent.
We identified two distinct zones (K1 and K2) of high channel gradient and steepness index, which maintain low channel width despite the variable rock strength of the substrate.
Knickpoint K3 may have been generated by the formation of the epigenetic gorge along the N-S traverse of the Chenab River (cf. Fig.3c). Knickpoints L1 and L2 mark the transition of a soft-to-hard bedrock substrate.  .8d). The orthogonal profile projection method has been used in the case of K2 (cf. fig.7) to identify the width of the steep segment.   Fig. 4a. We propose a minimum and a maximum bedrock incision rate of 3.1-3.5 mm/y and 5.2-5.6 mm/y, respectively. Rapid bedrock incision along the N-S traverse of the Chenab River may imply sustained exhumation over the proposed mid-crustal ramp of the MHT.