Spatiotemporal variability of neotectonic activity along the Southern Himalayan front: A geomorphic perspective

Accepted Manuscript Title: Spatiotemporal Variability of Neotectonic Activity Along the Southern Himalayan Front: A Geomorphic Perspective Authors: Saptarshi Dey, Rahul Kumar Kaushal, Sonam, Vikrant Jain PII: DOI: Reference:

S0264-3707(17)30185-0 https://doi.org/10.1016/j.jog.2018.09.003 GEOD 1593

To appear in:

Journal of Geodynamics

Received date: Revised date: Accepted date:

17-8-2017 11-7-2018 2-9-2018

Please cite this article as: Dey S, Kaushal RK, Sonam, Jain V, Spatiotemporal Variability of Neotectonic Activity Along the Southern Himalayan Front: A Geomorphic Perspective, Journal of Geodynamics (2018), https://doi.org/10.1016/j.jog.2018.09.003 This is a PDF file of an unedited manuscript that has been accepted for publication. As a service to our customers we are providing this early version of the manuscript. The manuscript will undergo copyediting, typesetting, and review of the resulting proof before it is published in its final form. Please note that during the production process errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain.

Spatiotemporal Variability of Neotectonic Activity Along the Southern Himalayan Front: A Geomorphic Perspective Saptarshi Dey, Rahul Kumar Kaushal, Sonam, Vikrant Jain Discipline of Earth Sciences, IIT Gandhinagar, Gandhinagar – 382355, India

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Corresponding author: Vikrant Jain ([email protected])

Abstract

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The interplay of tectonics, climate and erosion has been proposed as the driving factor

behind the growth and evolution of the Himalaya. In this review paper, we focus on the neotectonic deformation history within the southernmost morphotectonic sector of the

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Himalaya (the Sub-Himalaya) through synthesis of geomorphic data.

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The Sub-Himalaya is arguably consuming ~100% of the total Himalayan shortening since early Quaternary. We compiled geodetic shortening rates, paleoseismic events

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(historical earthquakes), shortening rates deduced from uplifted strath/fill terraces and

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shortening rates from balanced cross-sections from the north-western, central and eastern Himalayan compartments to obtain an orogen-wide perspective of Quaternary deformation.

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We supported the compiled data with topographic swath, longitudinal river-profile analysis and ksn plots of the existing drainage in those compartments. Review of the existing data shows a mismatch of the trend of the geodetic shortening rates with those of the millennial or

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longer timescales; however, Holocene and modern day-shortening rates are of same range (~14-21 mm.a-1). Quaternary shortening rates are much lower, probably due to a longer time-

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averaging. Except central Nepal, the other sectors show significant out-of-sequence thrusting (~50% of the total) within the Sub-Himalaya since the Holocene. Paleoseismic data show variable recurrence intervals of large earthquakes along-strike (~100-600 years) and large

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seismic gaps or slip-deficit sectors, which could potentially cause surface-rupture earthquakes in the future.

Keywords: Neotectonics, crustal shortening, geomorphic markers, steepness indices, Quaternary, fluvial terraces; Himalayan Front.

1. Introduction and geological background The Cenozoic Himalaya, formed due to collision between the Eurasian plate and the northward-migrating Indian plate, is one of the most active orogens of the Earth. Approximately, 35-45% of the total of 40-50 mm.a-1 Indo-Eurasian convergence is accommodated by the growth and uplift of the Himalayan fold-and-thrust belt (Yin and Harrison, 2000; Yin, 2006). With continued convergence, the Himalayan orogen grew

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southwards with time (Gansser, 1964). During the early Quaternary, the older Himalayan foreland basin sediments were incorporated into the orogenic wedge as new thrusts emerged from the low-angle basal decollement, namely the Main Himalayan Thrust (MHT) (Zhao et

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al., 1993; Bilham et al., 2001; Lave and Avouac, 2001; Mukul, 2010). This newly-formed frontal fold-and-thrust belt is known as the Sub-Himalaya. Sub-Himalaya is bordered by the Main Frontal Thrust (MFT) and the Main Boundary Thrust (MBT) in the south and north, respectively (Fig. 1). The MFT, which is the southernmost of all the Himalayan thrusts, has

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been proposed as the surface/near surface expression of the MHT. It is argued that since the

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Quaternary, ~100% of the Himalayan shortening is accommodated within the Sub-Himalaya

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(Powers et al., 1998; Mukul, 2000; Valdiya, 2003; Jayangondaperumal et al., 2011; Thakur et

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al, 2014; Hirschmiller et al., 2014; Dey et al., 2016a; Srivastava et al., 2017). This manuscript is focussed only on the Sub-Himalayan deformation and its along-

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strike variations over different time-windows, as most of the shortening since Quaternary (~2-3 Ma) has been accommodated by the Sub-Himalayan Fold and Thrust belt (SHFTB) (Valdiya, 2003; Hirschmiller et al., 2014). The Quaternary tectonic history of the Himalaya

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can be envisaged over different time-averaging windows – geodetic (100-101 years; GPSderived shortening), historical (102-103 years; paleoseismic evidence), millennial (103-104

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years; late Pleistocene-Holocene timescales from fluvial landforms) to million years (105-106 years;

from

magnetostratigraphically-dated

retro-deformed/balanced

cross-sections)

timescales. The major challenges remain mapping of the spatial variability in the deformation pattern, identification of active faults and estimation of uplift rate along active faults over

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different timescales (Kirby and Whipple, 2012). The present contribution reviews the application of fluvial markers in demonstrating

the spatial pattern of deformation in the Himalaya. In general, it has been suggested that most of the total Himalayan convergence is consumed by thrusting along the Main Frontal Thrust (MFT) at least since the Holocene (Bilham et al., 1997, 2001; Wesnousky et al., 1999; Lavé and Avouac, 2000). However, out-of-sequence thrusting is reported from several locations

within the Sub-Himalaya (Mukul et al., 2010; Burgess et al., 2012; Thakur et al., 2014; Mukherjee, 2015; Dey et al., 2016a; Kaushal et al., 2017). Here, we aim to address the neotectonic history of the Sub-Himalayan belt all along the Himalayan arc, as deciphered with the usage of tools of fluvial and quantitative geomorphology. The structural style and width of the Sub-Himalaya is different along-strike. Therefore, we prefer to analyse and discuss the Sub-Himalayan deformation in three different

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windows along the mountain front from west to east. 1.1. Western Himalaya

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Despite accommodating lesser amount of arc-perpendicular shortening (Kundu et al.,

2014; Stevens and Avouac, 2015), the north-western Himalaya has developed a relatively wide (max. width ~90 km) and pronounced SHFTB within the stretch of the Chenab River in the west to the Ganga River exit in the east (Site A, in Fig. 1b and Fig. 2a). SHFTB is

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characterized by several sub-parallel thrusts such as the Palampur Thrust (Medlicott-Wadia

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Thrust), Jwalamukhi Thrust, Barsar back-thrust, Soan Thrust and the southernmost MFT (Powers et al., 1998; Steck, 2003) (Fig. 2a). Due to sinuous exposure of the MBT (in

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plan/map view), several structural ‘re-entrants’ and ‘salients’ were formed within this sector.

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The frontal thrust system has developed a frontal anticline (see Fig. 2a) and a related backthrust accommodating high shortening (≥6-7 mm.a-1) since at least the Holocene (Thakur et

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al., 2014; Vassalo et al., 2015). The frontal anticline has different names along-strike: Janauri Anticline in the Kangra re-entrant, Chandigarh Anticline near Nahan Salient-Pinjaur Dun

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area and the Mohand Anticline in Dehradun re-entrant. 1.2. Central/ Nepal Himalaya

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The central Himalayan sector spans across Nepal and has been the recent focus of attention since the 2015 Gorkha earthquake (Avouac et al., 2015) (site B in Fig. 1b; Fig. 2b). The SHFTB is narrower compared to the western Himalaya and only consists of two major thrusts- the Main Dun Thrust (MDT) and the MFT (Lave and Avouac, 2000; Mugnier et al.,

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2004). Some parts of the frontal thrust are blind in nature. 1.3. Eastern Himalaya The Sub-Himalayan belt in the eastern segment of the Himalayan orogen is even narrower and often less than 5-10 km in width and more importantly, in many places such as in the Darjeeling foothills (Teesta-Jaldhaka exit), the frontal thrust is either blind or expressed as E-W trending scarps on the low-relief topography (site C in Fig. 1b; Fig. 2c). The Nameri

Thrust (Srivastava and Misra, 2008) and Chalsa scarp/Thrust (Kar et al., 2014) is probably an extension of the MFT from the Nepal Himalaya (Kumar et al., 2010). Only in the Arunachal Himalaya, the frontal FTB has significant structural variations, and those structures are Balipara Anticline,

Bhalukpong Thrust

and the MFT (Burgess et al.,

2012).

Paleoseismological evidences from the eastern Himalaya (Kumar et al., 2010; Jayangondaperumal et al., 2011) suggest that the eastern Himalaya is susceptive to recurrent high magnitude earthquakes, as documented from the historical events, with the recent one in

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1950 (The Assam Earthquake- Bilham, 2004). Long et al. (2012) and McQuarrie et al. (2014)

provided thermochronological data derived from Bhutanese Himalaya suggesting a slower

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shortening rate (~7-10 mm.a-1) since ~10 Ma.

The variability of the topographic relief among these sectors are aptly portrayed by representative topographic swath profiles (4 km swath width) from these three focus sites – A, B and C (Fig. 3). It showcases the width of the Sub-Himalaya and the locations of active

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deformation, which is causing change in elevation and relief.

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Slope changes can be directly observed in long profile, hence shape of long profiles

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and long-profile based indices provide the first marker of tectonic deformation. Further,

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changes in planform morphology and associated indices have also been used as proxies to understand tectonic deformation pattern in a region. Recently, river terraces have emerged as

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an important marker to measure time-averaged uplift rate along the active faults (Lave and Avouac, 2000; Burgess et al., 2012; Thakur et al., 2014; Vassalo et al., 2015; Dey et al., 2016a; Cortes Aranda et al., 2018). All these fluvial markers act as important tools to unravel

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subsurface tectonic processes and also to measure the rate of tectonic processes (Holbrook and Schumm, 1999; Schumm et al., 2002).

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2. Methods and calibration 2.1. Stream power law

Rivers are excellent markers to record sub-surface deformation, as river processes and

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morphology are sensitive to even minor changes in channel slope or gradient. An uplift rate ranging between 2-3 mm.a-1 may cause significant changes in river processes and morphology (Schumm et al., 2002), which may result into variations in planform and crosssectional fluvial metrics. River processes and morphology are governed by equilibrium between stream power and sediment supply (Stock and Montgomery, 1999). Stream power

(), is defined by liberation rate of kinetic energy from potential energy (Bagnold, 1966). It is expressed as –

= .Q.s

… (1)

where  - is the unit weight of water, Q- is the discharge, s- represents energy slope which is generally considered equivalent to bed slope. Tectonic processes govern channel slope, which

parameters and river processes are sensitive to tectonic deformations.

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2.2. SL index

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finally changes the stream power in a river channel. Hence, channel morphological

Stream length gradient index (SL) index is defined as the product of reach-scale channel slope and distance of that reach from river head (Hack, 1973). It is expressed asΔℎ Δ𝑙

.𝐿

… (2)

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SL =

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where, Δh is the change in elevation along the reach. Δl is the reach length and L is the total

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distance between mid-point of the reach of interest and the channel head. The SL Index is

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sensitive to changes in channel slope caused by tectonic uplift because even minor changes in slope (Δh/Δl) along a reach is enhanced through its product with channel length (L). SL can

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also be used to characterize stream power law. Channel length (L) can be used as a proxy for drainage area or river discharge if climatic conditions within a river basin are assumed to be uniform.

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2.3. Normalized steepness indices

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Normalised steepness index (𝑘𝑠𝑛 ) is another longitudinal profile-based index to analyse subsurface tectonic deformation (Whipple and Tucker, 1999; Kirby and Whipple, 2001). It is a measure of steepness of a river reach normalized for the downstream increase in basin area (A). The function for normalized steepness index could be derived through

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application of stream power incision law on long profile evolution in a steady state condition (see Appendix-A). It is expressed as𝑘𝑠𝑛 = 𝑆. 𝐴−𝜃

… (3)

Here, S is channel slope and 𝜃 is a reference concavity index value normally taken to be equal to 0.45. Normalised channel steepness index facilitates the comparison of river

channels within different study areas and rivers of different drainage basin sizes. This normalized steepness index is only applicable at steady state condition for detachmentlimited river system. 2.4. Calculation of activity rates Horizontal shortening rates can be deduced from slip rates or uplift rates on any active

fault plane (𝜃 ramp) is known from either field evidence or seismic profiles. Shortening = Fault slip/cos 𝜃 = Uplift/ tan 𝜃 Results and discussion

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3.

… (4)

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fault plane, if the geometry (dip amount and direction and its’ variation with depth) of the

3.1. Spatial and temporal variability in shortening

The Quaternary deformation in the Sub-Himalaya varies across different time-

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averaging windows and along different compartments of the mountain-front (Fig. 2).

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3.1.1. Geodetic shortening

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The GPS-based estimates revealed that the rate of arc-perpendicular shortening

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accommodated within the Himalayavaries from 14-21 mm.a-1 (Fig. 4a) (DeMets et al., 1994; Banerjee and Burgmann,2002; Schiffman et al., 2013; Kundu et al., 2014), with the eastern Himalaya consuming the highest convergence (~21 mm.a-1) and it gradually reduces towards

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the west (~13-14 mm.a-1), as a consequence of oblique convergence of the plates. Several other studies all along the orogen conforms to this result (Banerjee and Burgmann, 2002;

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Jade et al., 2007; Mukul et al., 2010; Kundu et al., 2014). GPS-derived shortening rates in the western Himalaya ranges from 13.3±1.7 mm.a-1

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to 18.5±1.8 mm.a-1 (Stevens and Avouac, 2015). Several other studies, such as Banerjee and Burgmann (2002) and Kundu et al. (2014) from the Kangra re-entrant (slip rate on MHT: 14±1-2 mm.a-1) complement this result. GPS motion vectors within the Kangra re-entrant (Kundu et al., 2014) hints strain partitioning within the Sub-Himalaya. Modern-day GPS

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measurements in Nepal provide shortening rate estimates as 19.4±1.4 to 20.1±1.1 mm.a-1 (Stevens and Avouac, 2015; Ader et al., 2012). The GPS-derived shortening rates in the eastern Himalaya varies from 17.6±0.9 to 21.2±2.0 mm.a-1 (Stevens and Avouac, 2015). Another study by Mukul et al. (2010) shows the GPS shortening rates in the eastern Himalaya is 15-20 mm.a-1.

GPS convergence rates provide higher resolution data but within a narrow range of temporal scale (~10 years). As the recurrence intervals of large earthquakes in the Himalayas are larger than 80-100 years (Bilham et al., 2001), some active faults may be underestimated by the GPS data. Therefore, it is essential to integrate GPS data with geomorphological data on tectonic deformation at millennial timescale or beyond to get an overview of the tectonic setting over a larger time window.

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3.1.2. Paleoseismological evidences Paleoseismological data from the Himalaya portrays the significant surface rupture earthquakes over the last two millennia. Trenching studies in and around the MFT of the NW

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Himalaya document the neotectonic activity along the orogenic front. OSL ages from the

Hajipur Fault trench (along the Janauri Anticline- a growing fault-propagation fold) suggest at least two big earthquakes along the MFT over 0.5-1.7 ka (Malik et al., 2010). Another

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study near Chandigarh on the MFT showcases earthquake ruptures along the MFT ~13001400 A.D (Malik et al., 2008) (Fig. 2). Kumahara and Jayangondaperumal (2013) proposed

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that the recent surface ruptures in Janauri Anticline occurred ~1400-1460 A.D.

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Lave et al. (2005) proposed that the MFT in Nepal ruptured ~1100 A.D. along-strike

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for >240 km and Kumar et al. (2010) suggested that this rupture had continued far east for 600 km in the eastern Himalaya. Paleoseismological evidences from the eastern Himalaya

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(Kumar et al., 2010; Jayangondaperumal et al., 2011) suggest that the eastern Himalaya is susceptive to recurrent high magnitude earthquakes, as documented from the historical

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events, with the recent one in 1950 (The Assam Earthquake- Bilham, 2004). Despite showing high convergence rates and potential for large surface-rupture earthquakes, the eastern

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Himalaya is relatively less investigated. Orogen-wide paleoseismic records show a broad recurrence interval for the high-Mw

earthquakes (~100-400 years). This hints that a minimum of ~1.5-6 m of slip on the MHT will be caused by a single seismic event. The compiled data also focuses that in spite of strain

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accumulation, there has been large seismic gaps in some stretches of the orogen (Kumar et al., 2010, Stevens and Avouac, 2015). Significant seismic gaps were found in the areas of Nahan salient, eastern Nepal, Bhutan foothills and the western Arunachal Pradesh (Bilham, 2004). With continued convergence, these sectors may experience new surface ruptures in future, or, extension of major surface-ruptures along-strike.

3.1.3. Late Pleistocene shortening Last two decades witnessed a significant advancement in the area of tectonic geomorphology, where fluvial markers were used to unravel tectonic deformation at millennial (103-104) timescale. Fault activity rates are mainly deduced from OSL (optically stimulated

luminescence)/radiocarbon

(14C)/

TCN

(terrestrial

cosmogenic

nuclide)

chronology of the deformed geomorphic markers such as uplifted fluvial strath terraces,

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offset fluvial fill terraces and folding of fluvial terraces for landforms as old as 50 ka (Fig. 4b).

Detailed chronological study of the transiently-stored intermontane valley sediments

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within the northwestern Himalaya provides excellent insight to the tectonic and climatic impact on the Earth surface processes. Regarding the tectonic scenario, Thakur et al. (2014)

proposed that the total shortening of 14±2 mm.a-1 in the Kangra re-entrant was partitioned

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among several Sub-Himalayan thrusts over the last 30-40 ka. While the Main Frontal Thrust (MFT) accommodates the majority of the shortening (6.9±0.6 mm.a-1), the orogen-interior

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Jwalamukhi Thrust (JMT), Soan Thrust (ST) and the MFT-back thrust accommodates the

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rest. With recently published Terrestrial Cosmogenic Nuclide-based exposure ages from the

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offset fluvial terraces in the Kangra Basin, Dey et al. (2016a) proposed that since the Holocene, the focus of tectonic activity in the re-entrant probably shifted within the Sub-

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Himalayan interior to the JMT. 5.6±0.8 – 7.5±1.1 mm.a-1 shortening rate accommodated on the JMT signifies that nearly 50% of the total Himalayan shortening in this sector is achieved by out-of-sequence thrusting along the JMT. Activity along the JMT will cause reduction in

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activity on the MFT, if the total shortening is consistent within this timescale. Towards the west, in the Jammu sector, Vassalo et al. (2015) proposed shortening rates of 11.2±3.8 mm.a-1

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on Medlicott-Wadia Thrust (MWT, contemporary to JMT in the east, averaged over ~14 ka) and 9.0±3.2 mm.a-1 on the MFT (over ~24 ka). Hence, total shortening on the MHT is between 13.2 and 27.2 mm.a-1 over this time-range. The lower bound of the shortening estimate correlates with the total Himalayan shortening, however it seems that the upper

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bound is probably an over-estimation. In the Dehradun re-entrant, Wesnousky et al. (1999) proposed a minimum Holocene slip rate of 11.9±3.1 mm.a-1 on the MFT over the last 2-3 ka, using 14C dates from the Dehradun basin. It is lower than the expected shortening rate, if we assume gradual decrease in shortening from central Nepal to the NW Himalaya. Mugnier et al., (2004) proposed the Holocene shortening rate of 19±5 mm.a-1 across the Nepalese Sub-Himalaya. It is the sum of shortening accommodated on the MFT and the

MDT (Main Dun Thrust, again invoking the idea of Holocene out-of-sequence thrusting). Deformed fluvial terraces from western Nepal provid a Holocene slip rate on the MFT ranging from 12-17 mm.a-1 (Mugnier et al., 2004). On the contrary, folded terraces in the Bagmati-Bakeya section in east-central Nepal, provide 21±1.5 mm.a-1 arc-perpendicular shortening rate in the Holocene (Lave and Avouac, 2000). It also suggests that MFT is the sole achiever of the Himalayan shortening in this sector.

the help of OSL and

14

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Holocene slip rates measured in western Arunachal Pradesh (Kameng River exit) with C dated uplifted strath terraces, provide a higher shortening rate of

23±6 mm.a-1 within the ~10 km-wide MFT zone (Burgess et al., 2012). Burgess et al., (2012)

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also proposed strain partitioning during the Holocene among the northerly Bhalukpong

Thrust (~8.4 mm.a-1), Balipara anticline in the middle (~10 mm.a-1) and southerly Nameri Thrust (~5 mm.a-1); Nameri Thrust is the local name or an extension of the MFT. Another study by Srivastava and Mishra (2008) in the same region, proposed possible variations in the

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average Holocene uplift or slip rates on the MFT-MHT. Luminescence ages from the uplifted

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terrace-tops provide a hint that during ~14-6 ka, the uplift rate on the MFT-MHT was higher

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(~11.9 mm.a-1) compared to the 14 ka average of 7.5 mm.a-1. Study by Mukul et al. (2007) in the Darjeeling Himalayan foothills suggested an initiation of emplacement of the MFT at ~40

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ka and onset of out-of-sequence activity to the north of MFT at ~20 ka. OSL-dated strath terraces along the Teesta River across the out-of-sequence South Kalijhora Thrust hint an

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uplift rate of ~4.4 mm.a-1 averaged over ~12-1.5 ka (Mukul et al., 2007). Overall, we can comment that the Holocene slip measured on the MHT has probably

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not activated only the MFT/HFF but may have partitioned within the Sub-Himalaya. Studies from the NW Himalaya invoke the idea of out-of-sequence thrusting in the western Sub-

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Himalaya (Mukul et al., 2007; Burgess et al., 2012; Thakur et al., 2014; Vassalo et al., 2015; Mukherjee, 2015; Dey et al., 2016a). This idea of strain partitioning within the Sub-Himalaya was supported by GPS data across the Kangra re-entrant. The GPS vectors in this area showed significant partitioning in convergence within the interior of the Sub-Himalayan FTB

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and not only across the MFT. Out-of-sequence activity within the Himalaya is also justified by the Critical taper theory (Davis et al, 1983) as it could imply the subcritical state of the frontal wedge taper (Prasad et al., 2011). No significant proof of the Holocene activity along the MBT and beyond have been recorded from dated fluvial markers till date.

3.1.4. Quaternary shortening rates Long-term (106 years) shortening rates estimated from balanced/ retro-deformed cross-sections

all

along

the

Himalayan

orogen

were

recalculated

using

new

chronostratigraphic age-bounds by Hirschmiller et al., (2014). The recalculated shortening rate estimates portray high variability over the last 2-3 Ma (Fig. 4c). While the value in the western Himalaya varies from 3.93 ± 1.08 mm.a-1 (Dehradun re-entrant) to 8.52 ± 1.6 mm.a-1

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(Kangra re-entrant) (re-calculated from Powers et al., 1998). In the central Himalaya the value ranges from 5.32 ± 1.19 mm.a-1 to 9.66 ± 1.35 mm.a-1 (re-calculated from Mugnier et al., 1999). In the eastern Himalaya, the shortening rate estimates are surprisingly low (3.20 ±

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0.67- 6.35 ± 1.04 mm.a-1, re-calculated from Chirouze et al., 2013). While the shorter

timescale estimates show an average shortening around ~20 mm.a-1 in the eastern Himalaya, low Quaternary shortening rate estimates may imply either an underestimation of the actual

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shortening or a re-activation of the FTB in this sector.

Hirschmiller et al. (2014) also pointed out that erosion might have played a significant

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role on the growth of the Himalayan FTB. As the eastern Himalaya receives higher rainfall, it

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triggers higher erosion from the orogenic wedge and to maintain the critical taper, the wedge

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has to deform internally (Dahlen, 1990) i.e. it would trigger out-of-sequence activity as we observe it in the far eastern sector over millennial timescale (Burgess et al., 2012). However,

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our review data portrays that even in the western Himalaya, there is significant amount of out-of-sequence activity going on since the Holocene (Dey et al, 2016; Vassalo et al., 2016). Therefore, we cannot clearly predict if only higher rainfall is causing the changes in the

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regional tectonic setting over shorter timescales. To summarize the temporal variability of shortening rate estimates within the Sub-Himalaya, we may say that the trends of balanced

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cross-section derived data and uplifted terrace-derived data are fairly similar, however the values of longer time-averaged balanced cross-section-derived data are significantly lower (Fig. 4). The trend of geodetic shortening is not quite well-matched with the Holocene timescale, but the range of values are similar. There are significant data gaps and to get a

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proper assessment of the spatiotemporal trends, more studies need to be initiated in the underworked sectors, such as the Darjeeling-Bhutan foothills, Pinjore Dun and Nahan salient etc. The concept of time-averaging in geological processes / geochronology, has its’ own merits and demerits. In one hand, that helps us to quantify deformations rates or temporally constrain any phenomenon, but on the other hand, it comes with certain approximations and

‘smoothing’ of any geological signal. As an example, the movement across any fault is a stochastic process. However, while calculating fault-offsets using dated geomorphic markers we actually take into account multiple such slip-events. Therefore, longer time-averaging addresses only the cumulative impact of any recurring geological event and cannot focus on any single significant event. On the other hand, geodetic timescale encompasses a history of ~10 years and usually record micro-seismic events. But, if we consider that an active structure has a large recurrence interval (>500 years) and almost micro-seismicity within the

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stipulated time-range, even the shorter timescales cannot record or can undermine the specific tectonic activity. Along with this, several locations along the orogenic front do not show a

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surface-breaking MFT, which are assumed to be blind in these locations (Kumar et al., 2010).

And, some areas show significant seismic data gaps along the arc. However, it is known that convergence is somehow accommodated within these sectors. Therefore, it can be predicted that the threshold limit for the seismic ‘lock-point’ on the MHT has yet not been breached in

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these segments, therefore, it results no big seismic events.

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As we review the published shortening rates over different time-scales along the

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Himalayan arc, we must point out that, the rates based on balanced/ retro-deformed cross sections, largely varies depending on how the authors chose to cap the timing of onset of

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deformation. While, the newest available rates from Hirschmiller et al., (2014) were calculated using the Upper vs. Middle Siwalik boundary as the marker horizon, previous

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authors (Powers et al., 1998; Mugnier et al., 1999; Chirouze et al., 2003) used the cessation of Upper Siwalik sedimentation as the marker horizon. Therefore, the older age estimates of

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Hirschmiller et al., (2014) has significantly reduced the Quaternary shortening rate. In comparison to the million-year-scale deformation rates, the geodetic (101 years) and

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Holocene-late Pleistocene (104-105 years) time-scale deformation rates are more precise and fairly correlate with each other, and therefore can be useful seismotectonic studies. However, we must comment that the longer timescale deformation rates are more useful for landscape evolution modelling.

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3.2. Morphometric indices River longitudinal profile convexities and knickpoints were used to identify the

locations of active faults and folds at the southernmost deformation zone of the Himalaya, such as the Janauri Anticline (Delcaillau et al., 2006) and the Pinjaur Dun (Singh and Tandon, 2008). Neo-tectonic activity along the Intra-Foreland Thrust (IFT) also known as Nahan Thrust in the Nahan Salient (Singh and Awasthi, 2010) and the Bhimgoda Back

Thrust (BBT) in the Mohand region of the Dehradun re-entrant (Singh and Jain, 2009) were identified on the basis of spatial variability in SL indices. SL index and knickpoints in river long profile were also used to observe neo-tectonic activity at north of MFT in the Nahan Salient, which suggests out-of-sequence tectonic activity in the region (Kaushal et al., 2017). Similar data were used to map neo-tectonic deformation along the Pawalgarh Thrust and Pawalgarh Anticline (a splay of MFT) in the Kumaon Himalayas (Luirei et al., 2015). Thakur et al. (2014) showcased knickpoints on the river profiles within the FTB which hints

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neotectonic activity along the Jwalamukhi Thrust. Recent studies also highlighted the application of channel steepness index and relief data to map tectonic activity and subsurface

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thrust geometry along the Jwalamukhi Thrust (Dey et al., 2016a) and to understand the

spatiotemporal pattern of tectonic deformation. Morphometric data from fold topography were also used to analyse relationship between uplift rate and topographic growth in the Chandigarh and Mohand anticlines (Barnes et al., 2011).

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Longitudinal profile convexity and SL index suggest some out-of-sequence activity

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to north of MFT in the Chitwan Dun (Divyadarshini and Singh, 2017). Low river profile

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concavity (0.3 – 0.6) was estimated for rivers flowing parallel to the strike of anticline and high concavity (0.7 to 2.1) was estimated for river flowing over the flanks of the anticline

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(Kirby and Whipple, 2001) in Baghmati and Bakeya basin. ksn indices revealed spatial variability in tectonic uplift rate in the hanging wall of the MFT in the Nepal Himalaya

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(Lague and Davy, 2003; Wobus et al. 2006; Robl et al., 2008 and Lague, 2014). The ksn estimates for the Bakeya River show a nearly linear relationship with incision / uplift rate

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(Lague, 2014). Further downstream, neotectonic deformation in the alluvial plains of Baghamti River basin due to sub-surface transverse faults were revealed through long profile

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shape and river morphological variability (Jain and Sinha, 2005). Spatial distribution of rock uplift rates along the eastern margin of the Tibetan Plateau

was also interpreted based on the estimates of river long profile steepness by Kirby et al. (2003) and Goode and Burbank (2009). They suggested that rock uplift rates based on

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channel gradient analyses were consistent with independently estimated rock uplift rates derived from low-T thermochronology in this region. SL index values and knickpoints in longitudinalprofiles were used to identify tectonic activity along the MFT of the Arunachal Himalayas (Srivastava and Misra, 2008; De Sarkar et al., 2014 and Devi et al., 2016). Reactivation of Mishmi Thrust was inferred from analysis of long profile and steepness index in the east Himalayan syntaxis (Bhakuni et al., 2017). Srivastava and Misra (2008) and De

Sarkar et al. (2014) reported uplift activity along out-of-sequence thrusts of the SubHimalaya along Kameng River in Arunachal Pradesh. 3.3. Applications and limitations of morphological markers in the Himalaya Geomorphic markers have played an important role in understanding of deformation pattern across the Himalaya at millennial (103-104 years) timescale. Planform and relief-based geomorphic markers provided spatial variability in deformation pattern in the Himalaya.

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Knickpoints on the longitudinal stream profiles have helped to identify the active structures, while presence of terraces within the hanging wall of different thrusts have been used to

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derive uplift (or slip) rate along the active faults. However, applications of various geomorphic markers are also being governed by its limitations, which cause uncertainties and

error in deriving information about neotectonic deformation. The limitations are governed by nonlinear and complex relationship between landform parameter(s) and tectonic forcing.

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Fluvial terraces are the most important geomorphic markers to derive quantitative

N

assessment of tectonic deformation. However, fluvial terraces also form a good example of

A

divergent landform (Schumm et al., 2000). Cause-effect relationship in divergent landforms will be complex as such landforms can be formed by more than one forcing. Fluvial terraces

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are formed in response to tectonic as well as climatic processes. Terraces in different parts of the Himalaya have been formed due to climatic process. It includes terraces in the Sutlej

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River valley (Bookhagen et al., 2006), in the Alakananda River valley (Srivastava et al., 2008; Juyal, 2010) and in the Ganga and Yamuna River valley (Sinha et al., 2010; Dutta et

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al., 2012). Tectonically active sub-Himalaya is also characterised by fluvial terraces in response to climate change due to changes in water and sediment fluxes (Singh et al., 2001;

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Sinha and Sinha, 2016; Densmore et al., 2016; Dey et al., 2016b). Extraction of tectonic information from these landforms will provide erroneous results. However, climaticallygenerated terraces can get tectonically impacted and may be used for quantifying deformation (Dey et al., 2016a). Terraces in the Mohand area (Wesnousky et al., 1999), Kangra

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intermontane valley (Thakur et al., 2014; Dey et al., 2016a), Jammu area (Vassallo et al., 2015) and Nahan salient (Kaushal et al., 2017) in the western Himalaya; Baghmati River basin in the central Himalaya (Lave and Avouac, 2000), and Kameng River (Srivastava and Misra, 2008) and Yarlung Tsangpo/ Siang River in the eastern Himalaya (Burgess et al., 2012) have been used to identify and quantify tectonic imprint on fluvial landforms. Some specific aspects of terraces like its confined distribution within the hanging wall and presence

of deformation features like back-tilting of terraces, vertical offset and broken clasts are useful to define tectonic control on river terraces (Wesnousky et al., 1999; Lave and Avouac, 2000; Burgess et al., 2012; Thakur et al., 2014; Vassallo et al., 2015; Dey et al., 2016a; Kaushal et al., 2017). Hence, study of tectonic signal from geomorphic markers needs careful analysis and should be filtered out from climatic effect. Further, estimation of channel incision through river terraces is not sufficient to

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estimate uplift rate. Relationship between channel incision and uplift rate will be nonlinear at different spatiotemporal scales, as part of the uplift process is also accommodated through planform changes in river system (Lave and Avouac, 2000; Schumm et al., 2002). Increase in

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the vertical drop between two points due to tectonic activity leads to increase in channel planform length to maintain its equilibrium slope. An increase in sinuosity of a channel reach

can accommodate total uplift in a significant way (Appendix-B). The changes in channel sinuosity will be different at different scales, as variations in sinuosity will also be controlled

U

by accommodation space in bedrock river channel. Hence, accommodation of uplift rate

N

through sinuosity change will not be same for all rivers. Further, post-incision channel

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sedimentation can also cause uncertainties in the estimation of bedrock incision rate (Lave and Avouac, 2000). Therefore, incision rate only provides minimum uplift rate along an

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active fault. Impact of planform change and channel sedimentation was included in estimation of uplift rate in the central Himalaya (Lave and Avouac, 2000), whereas terrace

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studies from other parts of the Himalaya provide minimum estimation of uplift rate (Wesnousky et al., 1999; Thakur et al., 2014).

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Some morphological parameters such as SL index are sensitive to scale variation. Tectonic signal area recorded through local slope ((Δh/Δl) variability (Hack, 1973). However,

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SL index enhances this slope variability through its product with channel length (L), which is dependent on basin scale. Hence, even though SL index is an important parameter to identify active structure in a watershed, direct comparison of SL indices of different areas could not be used to interpret spatial variability in tectonic activity. For example, the Himalayan front is

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characterised by extremely high SL index (~6000) for rivers of the far-eastern Himalaya (Bhakuni et al., 2017), moderately high SL index values within the western Himalaya (≤3000) (Luirei et al., 2015, Kaushal et al., 2017) and lower values in the central Himalaya (≤2000) (Divyadarshini and Singh, 2017). However, the aforementioned spatial variability in SL index may not be an indicator for spatial variability in deformation rate. Comparison of scaling parameters of regional variations in SL indices needs to be normalized with basin-

scale parameters (Kaushal et al., 2017), which will help to map the spatial variability in deformation. A big challenge in analysing neotectonic deformation in the Himalayan front is the lack of data owing to the dynamic topographic conditions at the mountain front. With a very low preservation potential and partially due to anthropogenic activities, obtaining dateable geomorphic markers is a rarity.

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4. Conclusions

Geomorphic parameters and indices obtained for the Himalayan terrain provided

valuable information about neotectonic deformation along the mountain front. These are

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summarised as follows –

1. There exists a significant spatial and temporal variability in deformation pattern and 21 mm.a-1 over Quaternary and smaller timescales.

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uplift rate along the Himalayan mountain front. The total shortening rate varies from 13-

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2. The Sub-Himalaya is accommodating ~100% of the Himalayan shortening within

A

millennial (103-104 years) timescale. However, in most of the regions along strike of the mountain front, the convergence is partitioned within the Sub-Himalaya and not only

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concentrated on the MFT. There are also large data gaps along-strike in the areas, such as the Nahan salient, eastern Nepal, Darjeeling-Sikkim Himalaya, Bhutan Himalaya.

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3. The shortening rates derived from balanced cross section and uplifted fluvial terrace match well along the along-strike of the Himalayan arc. However, there is a significant

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mismatch of this trend with that of geodetic shortening rates. 4. Historical earthquakes and paleoseismological investigations along the frontal range

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show significant seismic gaps or blind faults, which could potentially cause surfacerupture earthquakes in future. Geomorphic markers are important tools to analyse sub-surface tectonic deformation

at millennial time scale. However, we would like to conclude that the cause-effect

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relationship in tectonic geomorphology is generally nonlinear and complex. Effective application of these geomorphic markers requires careful analysis with in-depth understanding and quantification of geomorphic processes.

Appendices

Appendix –A: Normalized steepness index The steady-state profile states that river bed incision (E) is balanced by the rock uplift (base level fall, U), and its can be represented as 𝑑𝑧 𝑑𝑡

=𝑈−𝐸 =0

… (A.1)

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Bedrock incision (E) can be expressed stream power law – 𝐸 = 𝐾𝐴𝑚 𝑆 𝑛 … (A.2) where K is an erosion coefficient (rock erodibility), A is contributing upstream drainage area, S is a channel reach slope, m and n are scaling factors which depends on hydrology of the basin and mechanics of erosion processes respectively (Whipple and Tucker, 1999; Wobus et al., 2006).

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Eqs (1) into 2 confirm and solve for the slope S 𝑆 = (𝑈/𝐾)(1/𝑛) (𝐴)(−𝑚/𝑛) or,

... (A.4)

U

is a profile steepness and 𝜃 (= 𝑚⁄𝑛) is profile concavity.

N

where, 𝑘𝑠 (=

𝑈 [𝐾]1/𝑛 )

𝑆 = 𝑘𝑠 𝐴−𝜃

… (A.3)

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Appendix –B: Accommodation of uplift rate through planform (channel sinuosity) change

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Channel sinuosity of a river reach between two points A (upstream) and B (downstream) is defined by –

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𝑆𝑖𝑛𝑢𝑜𝑠𝑖𝑡𝑦 (𝑃) =

𝐿𝑐 ⁄𝐿 𝑣

… (B.1)

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where, 𝐿𝑐 – Channel length, 𝐿𝑣 – valley length, Further, valley and channel slope will be given as … (B.2)

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𝑉𝑎𝑙𝑙𝑒𝑦 𝑠𝑙𝑜𝑝𝑒 (𝑆𝑣 ) = 𝛥𝐻⁄𝐿 𝑣

Cℎ𝑎𝑛𝑛𝑒𝑙 𝑆𝑙𝑜𝑝𝑒 (𝑆𝑣 ) = 𝛥𝐻⁄𝐿 … (B.3) 𝑐 where, 𝛥𝐻 (= 𝐻2 − 𝐻1 ) is vertical difference between points 𝐴 and 𝐵 having height 𝐻2 and 𝐻1 respectively.

A

From on 𝐸𝑞. (2) and (3), 𝐸𝑞. (1) becomes, 𝑃 = 𝑆𝑣 /𝑆𝑐

… (B.4)

This equation represents an important relationship between planform change and vertical change (Schumm, et al., 2002). This relationship further provides an expression for accommodation of uplift rate by planform (channel sinuosity) change. Sinuosity change after uplift of ‘𝛥ℎ′ at upstream point A will be given by the following -

Height of Point A above a datum after uplift (𝐻′2 ) = 𝐻2 + 𝛥ℎ, and channel slope of reach A-B after uplift (𝑆′𝑐 ) = 𝛥𝐻 ′ / 𝐿𝑐 The channel will try to increase its length to maintain the equilibrium channel slope. Hence, 𝑆′𝑐 = 𝛥𝐻 ′ / 𝐿′𝑐 ; and 𝑆′𝑣 = 𝛥𝐻 ′ / 𝐿𝑣

… (B.5)

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The difference between present and past sinuosity will be given as –

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where, 𝐿′𝑐 – Enhanced channel length after its response to uplift. This increase in channel length (sinuosity) will accommodate part of uplift process. The general relationship between uplift and channel length (after Lave and Avouac, 2000) will be defined through the following steps.

𝑃′ − 𝑃 = (𝑆′𝑣 / 𝑆𝑐 ) − (𝑆𝑣 / 𝑆𝑐 )

… (B.6)

Substituting the value of 𝑆′𝑣 in Eq. (6), and after simplification,

U

𝑃′ − 𝑃 = Δh⁄(𝑆𝑐 . 𝐿𝑣 ), or

… (B.7)

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𝛥ℎ = 𝑆𝑐 (𝐿′𝑐 − 𝐿𝑐 )

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Hence, when 𝐿′𝑐 = 2𝐿𝑐 ; 𝛥ℎ = 𝛥𝐻

Acknowledgments

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The equation suggests that doubling of sinuosity between two points will be able to accommodate uplift equivalent to elevation different between two points.

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We thank the reviewer Prof. Dr habil. A.J. (Tom) van Loon, an anonymous reviewer and the Editorin-Chief, Prof. Artemieva for their constructive comments and suggestions which improved the manuscript significantly. RKS received doctoral fellowship from IIT Gandhinagar and Sonam received doctoral fellowship from CSIR-UGC India fellowship no. 061320507-23/06/2013(i)EU-V.

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Figure 1: (a) Overview map of the Himalayan orogen located at the southern margin of the IndoEurasian collisional zone. (b) Simplified geological map of the Himalaya (modified after Parsons et al., 2016) showing major tectonic boundaries and morphotectonic sectors. We focused on the SubHimalaya for this manuscript (fluorescent green colour in the map) and subdivided it into three major segments according to geographic positions. These segments from west to east are as follows –site A. the northwestern Himalaya, site B. the central Himalaya and site C. the eastern Himalaya.

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Figure 2a: Simplified shaded relief map of the northwestern Himalaya showing major thrusts and the locations of major drainage exits. (MFT= Main Frontal Thrust, JMT= Jwalamukhi Thrust, PT= Palampur Thrust/Medlicott Wadia Thrust, ST= Soan Thrust, BrT= Barsar Thrust, MBT= Main Boundary Thrust, HF- Hajipur Fault, ChA- Chandigarh Anticline, JA- Janauri Anticline, MA = Mohand Anticline). Swath window is shown as white box.

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Figure 2b: Simplified shaded relief map of the central (Nepalese) Himalaya showing major tectonic features Swath window is shown as white box. (MDT = Main Dun Thrust, ChD = Chitwan Dun)

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Figure 2c: Simplified shaded relief map of the eastern Himalaya showing major thrust sand folds. (BhT: Bhalukpong Thrust, TT: Tipi Thrust, BA: Balipara Anticline)

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Figure 3: Topographic swath profiles showing the variations in the Sub-Himalaya in the three designated sites – a. northwestern Himalaya, b. central Himalaya and c. eastern Himalaya (see fig.2 for location of the swath profile). The grey colour bars show width of the Sub-Himalayan FTB in different segments. Even though the arc-perpendicular shortening in NW Himalaya is much lesser than the eastern Himalaya, the shortening is accommodated over a wider Sub-Himalayan domain. The reason behind a narrower FTB in the eastern Himalaya can only be owed to either a rapidly-eroding mountain front, which propels out-of-sequence activity in the orogen interior, or a different wedgegeometry than the NW Himalaya. The red curve represents the mean elevation bounded by maximum and minimum elevation within the swath profile. Relative relief can be visualised by the difference between the maximum and the minimum swath-line. If relief can be assumed as a proxy of uplift, these figures illustrate the localised tectonic uplift along the FTB.

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Figure 4: Compilation of spatial variations in Sub-Himalayan shortening rates over three temporal scales: (a) Geodetic/ GPS-derived shortening rates (Stevens and Avouac, 2015; Kundu et al., 2014; Ader et al., 2012; Banerjee and Burgmann, 2002 and the references therein), (b) Holocene-late Pleistocene shortening rate derived from uplifted terraces (Cortes-Aranda et al. 2018; Dey et al., 2016; Vassalo et al., 2015; Thakur et al., 2014; Burgess et al., 2012; Srivastava and Misra, 2008; Mugnier et al., 2004; Lave and Avouac, 2001; Wesnousky et al., 1999) – the difference between shortening on MHT and MFT signifies the amount of out-of-sequence thrusting in the Sub-Himalaya, (c) Quaternary shortening rates derived from magnetostratigraphically-correlated balanced cross-sections (Hirschmiller et al., 2014 and the references therein).