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Miocene initiation and acceleration of extension in the South Lunggar rift, western Tibet: evolution of an active detachment system from structural mapping and (U-Th)/He thermochronology

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Richard H. Styron1*†, Michael H. Taylor1, Kurt E. Sundell1, Daniel F. Stockli1,2, Jeffrey A. G.

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Oalmann1, Andreas Möller1, Andrew T. McCallister1, Deliang Liu3, and Lin Ding3

Authors

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Author Affiliations

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1: Department of Geology, University of Kansas, Lawrence, Kansas, USA.

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2: Department of Geology, Jackson School of Geosciences, University of Texas, Austin, Texas,

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USA.

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3: Institute for Tibetan Plateau Research, Chinese Academy of Sciences, Beijing, China.

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* Now at Department of Earth and Environmental Sciences, University of Michigan, Ann Arbor,

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MI, USA

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Corresponding author: [email protected]

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Abstract

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robust estimates of the rates, timing or magnitude of Neogene deformation within the Tibetan

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plateau. We present a comprehensive study of the seismically active South Lunggar rift in

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southwestern Tibet incorporating mapping, U-Pb geochronology and zircon (U-Th)/He

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thermochronology. The South Lunggar rift is the southern continuation of the North Lunggar

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Rift, and comprises a ~50 km N-S central horst bound by two major normal faults, the west-

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dipping South Lunggar detachment and the east-dipping Palung Co fault. The South Lunggar

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detachment dips at the rangefront ~20° W, and exhumes a well-developed mylonite zone in its

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footwall displaying fabrics indicative of normal-sense shear. The range is composed of felsic

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orthogneiss, mafic amphibolite, and leucogranite intrusions dated at ~16 and 63 Ma. Zircon (U-

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Th)/He cooling ages are Oligocene through late Pliocene, with the youngest ages observed in the

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footwall of the South Lunggar detachment. We tested ~25,000 unique thermokinematic forward

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models in Pecube against the structural and (U-Th)/He data to fully bracket the allowable ranges

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in fault initiations, accelerations and slip rates. We find that normal faulting in the South

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Lunggar rift began in the middle Miocene with horizontal extension rates of ~1 mm a-1, and in

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the north accelerated at 8 Ma to 2.5-3.0 mm a-1 as faulting commenced on the South Lunggar

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detachment. Cumulative horizontal extension across the South Lunggar Rift ranges from <10

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km in the south to 19-21 km in the north.

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[1] Ongoing extension in Tibet may have begun in the middle to late Miocene, but there are few

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1 Introduction

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[2] Tibet is an archetypal example of an orogen undergoing syncontractional extension

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(Figure 1). Many models of Tibetan and Himalayan orogenesis have been proposed that explain

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or incorporate east-directed extension, such as convective removal of lithospheric mantle [e.g.,

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England and Houseman, 1988; Molnar et al., 1993], slab rollback in western Pacific subduction

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zones [Yin, 2000], orogenic collapse and radial spreading [e.g., Dewey et al., 1988; Copley and

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McKenzie, 2007], progressive underthrusting of Indian lithosphere [DeCelles et al., 2002; Copley

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et al., 2011], or other Himalaya-centric models [e.g., Klootwijk et al., 1985; Styron et al., 2011a].

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Most of these models either make predictions or rely on estimates of the age of onset of Tibetan

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and Himalayan extension. Additionally, many of these and other models seek to characterize the

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nature of deformation in the orogen, such as the debate between a continuum-style [e.g., England

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and Houseman, 1988; Bendick and Flesch, 2007] vs. block-style deformation of the orogen [e.g.,

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Avouac and Tapponnier, 1993; Meade, 2007; Thatcher, 2007], or the debate between

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deformation dominantly occurring along the orogen’s boundaries [e.g., Molnar and Tapponnier,

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1975; Lacassin et al., 2004] vs. internal deformation [e.g., Taylor et al., 2003; Searle et al.,

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2011]. These models are similarly reliant upon predictions or estimates of rates and magnitudes

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of deformation on faults in the orogen.

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[3] Despite the great interest in Tibetan rifting, only a small number of published studies

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document the onset of Cenozoic east-west extension within the plateau interior north of the

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Indus-Yarlung Suture Zone (Figures 1, 2), in contrast to the relatively well-studied Himalaya

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(see Lee et al. [2011] for a recent summary). In the eastern plateau, Pan and Kidd [1992] and

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Harrison et al. [1995] documented Pliocene cooling of the Nyainqentanglha detachment footwall

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(Figure 2). Harrison et al. [1995] modeled rifting beginning at 8 ± 1 Ma with a fault slip rate of

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3 mm a-1, a finding that was supported by J. Kapp et al. [2005]. A similar age was inferred by

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Ratschbacher et al., [2011] to the northeast along the same fault system by

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synkinematic micas from the mylonitic shear zone. Blisniuk et al. [2001] studied a fault zone in

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the Shuang Hu graben, a local releasing bend in the Muga Purou fault system in central Tibet

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[Taylor et al., 2003], and dated mineralized fault breccia at ~13.5 Ma through Rb-Sr and

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At face value, these dates suggest that rift inception across the plateau was very diachronous,

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although the sample size is quite small. Furthermore, these studies do not robustly estimate slip

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rates on the faults by rigorously testing many slip histories against the data in order to better

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constrain possible deformation histories.

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Ar/39Ar dating of

Ar/39Ar methods, which they interpreted as the minimum age of rift initiation on the plateau.

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[4] In order to gain a more thorough understanding of the timing, rates and magnitude of

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Cenozoic exension in Tibet, as well as the potential spatial variations in extension, more data are

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needed, especially for western Tibet. This study presents structural and neotectonic mapping,

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zircon (U-Th)/He (zHe) thermochronology, and zircon U-Pb analysis of the little-known South

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Lunggar Rift (SLR) in the western Lhasa block of Tibet. We document a large (>50 km along

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strike) and active north-trending rift containing both high- and low-angle normal faulting. The

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footwall of the low-angle normal fault displays a well-developed mylonitic shear zone and is

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interpreted as a metamorphic core complex. Data collection was combined with extensive 3-D

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thermokinematic modeling (~25,000 forward models) to test possible deformation histories

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against the structural and thermochronometric observations. These results indicate up to 20 km

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E-W extension starting in the early to middle Miocene, at moderate rates (1-3 mm a-1) following

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a late Miocene acceleration. Significant along-strike variability in fault geometry, slip rate and

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net displacement exist as well. In addition to providing new information on extension in SW

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Tibet, our results have implications for the thermal state of the Tibetan crust. Furthermore, the

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observations of low-angle normal ‘detachment’ faulting and large thermochronometric data

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collection allow for testing of various geometric models of detachment faulting.

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1.1 Style and evolution of detachment faulting

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[5] Low-angle (<30˚ dip) normal ‘detachment’ faults, often exhibiting several to tens of

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kilometers of extension, have been mapped throughout the world over the past 30 years [e.g.,

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Wernicke, 1995]. As these structures are not well understood due to the apparent conflict

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between their dip angle and standard Andersonian rock mechanical theory [e.g., Anderson, 1951]

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predicting normal fault dips of ~60˚ and fault locking at low angles, much effort has been put

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into resolving the contradiction of Andersonian fault theory with field [e.g., Lee et al., 1987; Yin

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and Dunn, 1992], geophysical [Abers et al., 2002; Morley, 2009] and geodetic [e.g., Hreinsdottir

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and Bennett, 2009; Niemi et al., 2004] observations. These studies often focus on the geometry

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of the detachment at depth, and whether faulting initiated at low angle or was first high angle and

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later rotated to a low angle [e.g., Spencer, 1984; Wernicke and Axen, 1988]. Prominent models

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include planar, low-angle fault initiation [e.g., Wernicke, 1981]; the ‘rolling hinge’ model, where

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the very shallow and very deep parts of the detachment fault are low angle, but the majority of

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slip within the seismogenic crust occurs at a moderate to high angle [e.g., Axen and Bartley,

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1997]; and antilistric models, where the detachment fault monotonically steepens with depth

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[e.g., Buck, 1988].

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[6] Many of the canonical field studies of detachment faults focused on the Cordillera of

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western North America (indeed, the typically sheared and metamorphosed antiformal footwalls

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of detachment faults were initially known as ‘Cordilleran’ metamorphic core complexes

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[Crittenden et al., 1980], now less parochially ‘metamorphic core complexes’ or more simply

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‘core complexes’), which were generally active in the late Cretaceous through Miocene [e.g.,

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Lister and Davis, 1989]. Therefore, few studies [e.g., Daczko et al., 2011; J. Kapp et al., 2005;

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Kapp et al., 2008] have been done on active structures, where considerably more certainty exists

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on the geometric and geodynamic context, such as the thickness, strain rate and thermal state of

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the crust and upper mantle.

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[7] However, computational simulations of detachment faulting and core complex

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development are numerous.

Though there has been significant variability in the modeling

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approach, the results typically show detachment faulting to form preferentially in areas of hot,

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thick crust capable of ductile flow at depth [e.g., Buck, 1991; Rey et al., 2009]. These studies

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also uniformly show the flexural or isostatic ‘back’ rotation of the footwall away from the

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detachment fault and hanging wall, leading to an up-dip shallowing of the fault dip, i.e. an

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‘antilistric’ fault geometry [e.g., Buck, 1988; Rey et al., 2009; Tirel et al., 2008; Wdowinski and

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Axen, 1992]. In models that do not specify an initial detachment geometry, some material

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heterogeneity is often needed to initially localize deformation; this is typically a magmatic

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intrusion [e.g., Brun et al., 1994; Tirel et al., 2008], which is compelling because of the strong

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association of magmatism and core-complex formation [e.g., Armstrong and Ward, 1991].

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[8] Thermochronologic techniques have proven invaluable in understanding the rate and

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style of deformation in a variety of tectonic settings, especially in extensional regions, where

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progressive down-dip exhumation of a normal fault footwall often leaves a clear thermal

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signature [e.g., Stockli, 2005]. Thermochronologic data in normal fault footwalls are typically

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interpreted in age vs. elevation or age vs. down-dip distance plots, often by the fitting of linear

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regression trend lines [e.g., Fitzgerald et al., 2009; Mahéo et al., 2007]. However, this method

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makes questionable assumptions about the thermal state of the crust, particularly that the

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geothermal gradient is constant with depth and does not change during faulting, and radiogenic

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heating is not significant. These assumptions have been shown to be inaccurate enough to cause

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erroneous interpretations [Ehlers et al., 2001; Ehlers, 2005].

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complications such as progressive rotation of the footwall during extension may distort the

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geometrical relationship of the samples to horizontal geotherms; these complications have to be

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well-constrained [e.g., Stockli et al., 2002]; or ignored. Furthermore, simple regression lines are

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rarely weighted by age uncertainty, thereby failing to take this important age information into

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account.

Additionally, structural

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[9] Advances in 2-D and 3-D thermokinematic modeling [e.g., Harrison et al., 1995;

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Ketcham, 1996; Braun, 2003; Ehlers et al., 2001] have enabled the use of complicated fault

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geometries and dynamic, nonlinear geotherms incorporating radiogenic heating. Furthermore,

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iterative methods [e.g., Campani et al., 2010; Ketcham et al., 2005] allow for model fitting that

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incorporates formal uncertainties in thermochronometer data, producing much more robust

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interpretations than previously possible.

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[10] Many studies of Himalayan and Tibetan rifting have found evidence of active

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detachment faulting [e.g., Harrison et al., 1995; Jessup et al., 2008; J. Kapp et al., 2005; Kapp

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et al., 2008; Murphy et al., 2002; Pan and Kidd, 1992; Robinson et al., 2004], consistent with

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predictions of detachment fault formation in hot, thick crust [e.g., Buck, 1991]. Detachment

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faults have been mapped in both the North and South Lunggar rifts [Kapp et al., 2008; Styron et

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al., 2011b], and are interpreted to be active. The identification of rapidly-exhumed mid-crustal

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rocks in the detachment footwalls suggest that extension is locally very significant, and that

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faults are of significance to deformation of the Tibetan plateau. The structural and

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thermochronological work presented here on the South Lunggar rift give both an understanding

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of the rates and timing of western Tibetan extension and a picture of core-complex activity in a

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hot and thick orogen.

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1.2 Regional Geology

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1.2.1 Pre-extensional geology

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[11] The southern margin of Eurasia has been tectonically active throughout the

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Phanerozoic. This activity mostly consists of the successive accretion of multiple terranes that

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now compose the Tibetan plateau. Accretion of these terranes is generally assumed to young

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southward, with docking of the Qilian and Kunlun terranes in the Paleozoic, the Qiangtang

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terrane in the early-mid-Mesozoic and the Lhasa terrane in the mid-late Mesozoic (forming the

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Bangong-Nujiang Suture Zone, Figure 1) [Yin and Harrison, 2000]. The late Cretaceous to early

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Eocene saw the beginning of India’s ongoing collision with the Lhasa terrane along the Indus-

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Yarlung Suture Zone [Ding et al., 2005], creating much of the crustal shortening observed today.

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[12] Shortening in central Tibet began in the late Jurassic [Murphy et al., 1997] or early

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Cretaceous [P. Kapp et al., 2005] associated with the underthrusting of the Lhasa terrane beneath

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the Qiangtang terrane [e.g. Yin and Harrison, 2000; Kapp et al., 2007]. Shortening, accompanied

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by magmatism, continued throughout the Lhasa terrane until the Paleocene [Murphy et al., 1997;

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P. Kapp et al., 2005; 2007]. Thin-skinned thrust sheets composed of Paleozoic strata were thrust

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over Mesozoic strata (and vice versa) in the south-central Lhasa terrane, and were sporadically

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intruded by granites throughout the Cretaceous [Murphy, et al., 1997]. During the mid to late

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Cretaceous, subduction of Neothethyan lithosphere underneath the southern Lhasa terrane

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produced the Gangdese magmatic arc [e.g., Ding et al., 2003].

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[13] Following the onset of India’s collision, shortening generally ceased in the interior

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of the Lhasa terrane (inferred from the widespread and essentially flat-lying early Tertiary

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Linzizong volcanic rocks) [Murphy et al., 1997], but was still active until ~20 Ma on its northern

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and southern margins [DeCelles et al., 2011; P. Kapp et al., 2005; 2007; Yin et al., 1994] as well

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as in northern Tibet [e.g., Lease et al., 2011]. Several hundred kilometers of shortening were

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accommodated in the Himalaya at this time, as well [DeCelles et al., 2002; Robinson et al.,

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2006; Murphy, 2007]. Synconvergent extension in the direction of plate convergence occurred

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episodically throughout the Oligocene and early Miocene, expressed as activity on the north-

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dipping South Tibetan Detachment system [Burg et al., 1984; Burchfiel et al., 1992] and the

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development of the South Kailas Basin between the Gangdese arc and the thrusts of the Indus-

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Yarlung Suture Zone [DeCelles et al., 2011; Zhang et al., 2011].

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[14] In the middle to late Miocene, a dramatic change in the style of deformation in the

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Himalaya and Tibet occurred. Activity on the Main Central Thrust and South Tibetan

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Detachment, the dominant early Miocene structures in the Himalaya, was significantly reduced if

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not halted altogether [e.g., Murphy et al., 2002; Leloup et al., 2010] while the dominant zone of

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Himalayan shortening propagated south [Meigs et a., 1995; DeCelles et al., 2001]. At this time,

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the High Himalaya began arc-parallel extension along structures cutting the Main Central Thrust

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and South Tibetan Detachment [Thiede et al., 2006; Murphy, et al., 2002; Styron et al., 2011a;

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Garzione et al., 2003; Lee et al., 2011; Jessup et al., 2008; Kali et al., 2010; Leloup et al., 2010].

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Within the central and southern Tibetan plateau, shortening via folding and thrusting essentially

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ceased and an ongoing phase of east-west extension began [Lee et al., 2011; Kapp et al., 2008],

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with ongoing shortening, observed geodetically [e.g., Zhang et al., 2004] ostensibly

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accommodated on NE- and NW-striking V-shaped conjugate strike-slip faults in central Tibet

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[Taylor et al., 2003; Yin and Taylor, 2011].

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1.2.2 Neogene rifts in Tibet

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[15] Neogene extension in Tibet is characterized by roughly north-trending graben in the

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central Lhasa and Qiangtang terranes [e.g. Armijo et al., 1986; Blisniuk et al., 2001] (Figures 1,

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2). These graben are often linked to V-shaped conjugate strike-slip faults emanating from the

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Bangong-Nujiang Suture zone [e.g., Taylor et al., 2003], although in some cases, such as small

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graben in the Tanggula Shan and Gangdese Shan, extension may be isolated to areas of high

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topography (Figure 2). The southern Lhasa terrane contains five major rifts that essentially span

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the north-south length of the Lhasa terrane, and several subordinate rifts. From east to west, the

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major rifts are the Yadong-Gulu rift (this rift cuts from the Himalaya to the Bangong-Nujiang

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suture; the main segment of the rift through the Lhasa block is called the Nyainqentanglha rift

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[Pan and Kidd, 1992; J. Kapp et al., 2005]), the Pumqu-Xainza rift [Armijo et al., 1986; Hager

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et al., 2006], the Tangra Yum Co rift [Dewane et al., 2006], and the Lunggar rift [Kapp et al.,

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2008, this study]. Subordinate rifts include an unnamed and unstudied (to our knowledge), but

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seismically active (Figure 2) rift to the west of the Lunggar Rift, the Lopukangri rift [Murphy et

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al., 2010] to the southeast of the Lunggar rift, the Xiagangjiang rift to the east of the North

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Lunggar rift [Volkmer et al., 2007], and numerous small graben throughout the western

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Gangdese range [e.g., Yin, 2000] (Figure 1, 2).

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[16] Some estimates have been made of net horizontal extension across the plateau.

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Armijo et al. [1986] estimated <3-4 km extension across the Yadong-Gulu and Pumqu-Xainza

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rifts, and extrapolate to suggest roughly 20 km extension across the plateau, assuming these rifts

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are representative of all major Tibetan rifts. More recently, J. Kapp et al. [2005], informed by

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modern ideas of detachment faulting and continental extension, studied the Nyainqentanglha

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segment of the Yadong-Gulu rift. They estimated a minimum of 8 km fault slip based on the

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down-dip length of the detachment fault’s mylonitic shear zone, and combine structural and

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thermobarometric data to suggest 21-26 km of fault slip, assuming the detachment fault slipped

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at 35 degrees dip. Given this fault dip, their corresponding extension estimates would be 17-21

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km based on all results with a minimum of 6.5 km. Taylor et al. [2003] suggest ~48 km of total

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right-lateral slip along the southern, right-slip faults in the conjugate fault zone along the

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Bangong-Nujiang suture that link into south Tibetan graben (i.e. west of the Jiali fault), based on

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mapping and remote sensing interpretation of the Lamu Co and Riganpei Co faults. Slip on

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these faults may be comparable to extension in the linked graben (Figure 1, 2). Preliminary

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mapping in the Tangra Yum Co [M. Taylor, unpublished mapping] and Pum Qu-Xainza rifts [C.

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Hager, personal electronic communication] suggest less than 10 km for each rift.

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[17] The Lopukangri rift to the southeast of the SLR (Figure 2) is a complex fault system

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interpreted either as part of the trailing end an extensional imbricate fan in a fault system

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extending from the Lamu Co fault through the Lunggar rift and southeastward into the Gangdese

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range, or as a more prominent member of a series of crustal tears with the same geographic

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extent [Murphy et al., 2010; Sanchez et al., 2012]. Based on the work of Sanchez et al. [2012]

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and our preliminary field observations, the Lopukangri rift has a long northern segment, with a

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west-dipping, moderate-angle range-bounding normal fault. Although throw on this fault has not

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been enough to exhume basement in its footwall, the fault has extremely large normal fault

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scarps, offsetting Quaternary alluvium by up to 350 m vertically and locally display triangular

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facets several tens of meters high, suggesting that this segment of the rift is reasonably active.

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To the north of the rift proper is a rangefront fault striking NW that is interpreted as an oblique

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slip (dextral-normal) fault that terminates to the NW near the central Lunggar rift and may

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transfer slip from the North Lunggar rift to the Lopukangri rift. The southern segment of the

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Lopukangri rift cuts the southern slopes of the Gangdese range, offsetting contractional

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structures associated with the Indus-Yarlung Suture Zone by ~15 km [Murphy et al., 2010;

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Sanchez et al., 2012].

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rifting began at ~15 Ma [Sanchez et al., 2012].

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Ar/39Ar dating of the footwall of the southern Lopukangri rift suggests

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1.3 Lunggar Rift

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[18] The Lunggar Rift is a major north-trending rift in the western Lhasa terrane [Armijo

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et al., 1986; Kapp et al., 2008; Elliott et al., 2010] (Figure 2, 3). It is kinematically linked in the

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north to the Lamu Co right-lateral strike-slip fault, part of the V-shaped conjugate fault system

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running along the Bangong-Nujiang Suture Zone [Taylor et al., 2003] . The rift is over 150 km

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along strike, and made up of northern and southern segments separated by an accommodation

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zone (Figure 3). The northern segment, or the North Lunggar Rift (called the Lunggar Rift by

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Kapp et al. [2008] and Woodruff et al., 2012), consists of an east-dipping low-angle detachment

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fault separating a narrow (<10 km wide) supradetachment basin from an elevated footwall

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composed of variably-deformed granitoids, orthogneiss, and metamorphosed Paleozoic

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sedimentary rocks. The detachment is inactive at the rangefront, as indicated by unfaulted

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moraine and alluvial material overlying the fault trace. However, both east and west dipping

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normal faults offset Quaternary alluvium in the supradetachment basin and are parallel to the

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range-bounding detachment, suggesting they sole into the detachment at depth [Kapp et al.,

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2008]. Relief in the North Lunggar Rift approaches 2 km, and maximum footwall elevations are

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~6500 m. The accommodation zone between the North and South Lunggar rifts consists of a

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less-elevated (peak elevations generally <6000 m) footwall made up of the Cretaceous thin-

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skinned thrust belt that forms the pre-extensional surface in adjacent regions [Murphy et al.,

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1997].

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[19] The South Lunggar Rift (Figure 4) is made up of a central horst block, the Surla

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Range, which is bounded on both the east and west by normal faults. Well-developed basins are

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found on both sides of the Surla Range. Quaternary cumulative fault scarps on flanks of the Surla

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Range and active seismicity indicate that extension in the SLR is ongoing. The Swedish explorer

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Sven Hedin was likely the first Westerner to describe the geography and geology of the SLR

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during his passage through in June, 1908 [Hedin, 1909]. He noted the extensive glaciation and

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wide distribution of granite boulders in the western rift valley. He also described feeling

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moderate ground shaking due to an earthquake at approximately 9:30 PM (local time) on 28

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June, 1908 while in Sailipu, a short distance to the west. To our knowledge, this is the first field

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geologic study of the SLR since Hedin’s.

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2 Bedrock and surficial geology of the South Lunggar Rift

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2.1 Bedrock Units

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[20] The Surla Range in our map area is dominantly composed of amphibolite-grade

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metamorphic rocks (ma) and greenschist-facies volcanic rocks (mv) intruded by variably-

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deformed leucogranites (gr, myl). Hanging wall rocks on both sides of the rift include

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unmetamorphosed volcanic rocks (v) and sedimentary rocks composing a Cretaceous thin-

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skinned thrust belt (K) (Figure 3, Figure 4). In general, ice and moraine cover and extensive talus

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development limited access to bedrock exposure, inhibiting more extensive sampling,

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measurement of structural data and contact identification, although high walls of glacial valleys

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and sporadic outcrop allowed for confident mapping of rock units.

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[21] The amphibolite-grade metamorphic unit (ma) is a composite unit of different rock

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types, mapped as one unit. The unit is composed of coarse-grained biotite amphibolite, biotite

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granite orthogneiss, and biotite granodiorite orthogneiss. The orthogneiss is locally migmatitic.

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Contacts between the different subunits were not observed, and relationships are uncertain.

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Foliations in the orthogneiss are strongly developed with individual bands mm to 10s of meters

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in thickness. Amphibolites are unfoliated to moderately foliated at the hand sample to meter

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scale, and a well-developed foliation is visible in glacier-polished valley walls (Figure 5c). The

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foliation is generally north-dipping, though significant variability exists (Figure 4), and therefore

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the metamorphic event that deformed these rocks is interpreted to be unrelated to the modern

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phase of extensional deformation. This interpretation is supported by the widespread intrusion of

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undeformed leucogranite (described below).

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were observed in outcrop, although they were occasionally spotted in talus or glacial debris.

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Leucogranite intrusion is widespread, and locally preferentially occurs along foliation planes

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(Figure 5c). Lower-grade (greenschist facies) felsic to intermediate fine-grained metavolcanic

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rocks (mv) are present in two major areas (Figure 4) and in small lenses (below map resolution)

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on the northern margin of the range.

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indicative of greenschist facies metamorphism. These rocks are observed to be intruded by

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granites, though the nature of contact between higher-grade metamorphic rocks was not

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observed. The unit may correlate to the Burial Hill volcanic rocks mapped by Murphy et al.

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[1997] along strike ~150 km to the east.

No ductile stretching fabrics or other lineations

Biotite displays kink banding and alteration to chlorite,

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[22] Biotite leucogranite intrusions (gr) are widespread, from meter to up to 10s of km

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scales (Figure 4, 5c). These intrusions are observed in the metamorphic units (ma, mv) (Figure

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5a, b). U-Pb ages (Section 4) indicate both Gangdese (~65 Ma) and early- to mid-Miocene (22-

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16 Ma) crystallization ages, although there are no significant petrologic differences between

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rocks of different ages. Therefore the large leucogranite bodies may be made up of smaller

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plutons that intruded episodically over 10s of m.y. In the northwestern Surla Range, structurally

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below the South Lunggar Detachment, the leucogranite is heavily sheared into a mylonitic zone

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(myl).

329

pressures are present in the leucogranites.

No garnet or other mineral phases indicative of intrusion or deformation at high

330

[23] Hanging-wall rocks consist of a Cretaceous thrust belt of Paleozoic and Mesozoic

331

supracrustal rocks (K) [Murphy et al., 1997] and unmetamorphosed biotite-and hornblende-

332

bearing felsic volcanic rocks (v). These volcanic rocks have a middle Miocene zHe cooling age

333

(16.8 ± 0.8 Ma; Figure 4, Table 1), interpreted as an eruption age due to the vesicular texture

334

suggesting little to no post-deposition burial and reheating. The presence of hornblende, not

335

found in the leucogranites, suggests an eruptive source external to the Surla Range.

336 337

2.2 Quaternary units

338

[24] Quaternary sedimentary units are dominantly the products of erosion of the Surla

339

Range massif, and include two generations of Quaternary alluvium (Qa and Qo), Quaternary

340

moraine and outwash (Qm) and Quaternary shorelines (Qsh). Qo is the older Quaternary unit,

341

cut off from modern depositional systems by uplift or drainage reorganization. Qa is found in

342

active to recently active depositional environments. The age of Qm is unknown, but it is

343

reworked by the highest shorelines (Qsh), which are dated at the nearby Ngangla Ringco (Figure

344

3) at ~10.4 ka [A.M. Hudson, electronic personal communication, 2012].

345 346 347

3 Structural geology of the South Lunggar Rift

348

[25] The Surla Range is uplifted on its eastern flank by a moderately east dipping normal

349

fault, here named the Palung Co fault, and on its northwestern side by a low-angle west-dipping

350

normal fault, here named the South Lunggar Detachment, which is linked with moderate to high

351

angle west-dipping normal faults on the northwestern and southwestern margins of the Surla

352

Range.

353 354

3.1 Palung Co Fault

355

[26] The Palung Co Fault is a moderate-angle east-dipping normal fault striking 20˚ in

356

the north and 350˚ in the south (Figure 3). The fault is ~80 km along strike, and cuts into the

357

Gangdese (Transhimalaya) Range south of the Surla Range. Where it bounds the Surla Range, it

358

forms 40˚ - 50˚ east-dipping triangular facets up to 1 km high. A lake, Palung Co, occupies much

359

of the ~10 km wide rift basin east of the fault trace (Figure 4). The fault has uplifted

360

leucogranites and amphibolites in the footwall 1.5 km above the sedimentary and volcanic rocks

361

in the hanging wall, giving a minimum amount of throw of 1.5 km; however, the estimated

362

sedimentary and volcanic cover thickness of ~8 km [Murphy et al., 1997; 2010] and young

363

zircon (U-Th)/He cooling ages (Section 5) suggests throw on the fault may be greater.

364

Interpretation of rift morphology [e.g., Friedmann and Burbank, 1995] and thermochronology

365

suggest that the area near Palung Co is the zone of maximum fault displacement. The Palung Co

366

Fault is currently active, as indicated by small fault scarps in ground moraine visible in remote

367

sensing imagery near the fault’s northern tip (Figure 4). Additionally, in August 2008 a series of

368

earthquakes occurred along the fault. The largest was a Mw 6.7 normal faulting event [Elliott et

369

al., 2010; Ryder et al., 2012]. Body-wave seismology and synthetic aperture radar interferometry

370

(InSAR) indicate the rupture occurred on two fault planes, one projecting directly to the

371

rangefront fault in the study area and the other several km to the south [Elliott et al., 2010].

372

Those authors estimated the northern rupture to be striking 20˚ and dipping 43 ± 2˚ E, in very

373

close agreement with our field observations. Their modeling suggests that the top of the northern

374

rupture patch was 2-5 km deep, and the bottom was 14-20 km deep. The shallow termination of

375

seismic slip and InSAR phase continuity across the fault trace [Elliott et al., 2010; Ryder et al.,

376

2012] is consistent with our observations from mapping the area 12 months after the event

377

indicating no obvious surface rupture. The southern rupture is roughly along strike of the

378

northern rupture, but is south of a change in the mapped fault strike and cuts into the high

379

topography in the Surla Range or Gangdese Range (the two ranges merge at this latitude), and

380

has no clear geomorphic expression [Elliot et al., 2010]. This difference in strike between the

381

rangefront and the rupture possibly represents the southward propagation of the northern Palung

382

Co Fault and cessation of activity on the previous southern segment to the east, consistent with

383

models of developing normal fault systems that hypothesize the simplification and organization

384

of rift geometry into one relatively planar master fault through time [e.g., Bosworth, 1985].

385 386

3.2 South Lunggar Detachment

387

[27] The northwestern portion of the Surla Range is uplifted along the South Lunggar

388

Detachment (SLD). The SLD is a shallowly north- to west-dipping normal fault that is

389

interpreted to link at depth with the steeper range-bounding normal faults to the north (Figure 3);

390

however, thick moraine cover obscures the fault linkage at the surface, though possible fault

391

scarps in moraine suggest partitioning of normal and strike-slip motion into two main strands

392

(Figure 4). To the south it is linked with a moderate-angle normal fault (Section 3.3), though we

393

restrict the use of the name ‘South Lunggar Detachment’ to the northern fault. In its footwall, the

394

SLD has exhumed leucogranite and amphibolite. In places on the western rangefront, triangular

395

facets dipping ~20˚ W are preserved (Figure 7), though extensive glaciation has modified the

396

rangefront elsewhere. Except in its southern extent, hanging wall rocks are not observed near the

397

trace of the SLD due to thick moraine cover and the fault is not observed in bedrock.

398

[28] Immediately below the detachment, footwall leucogranites display a mylonitic shear

399

zone >100 m thick. Foliation planes in the mylonites strike parallel to the local trend of the

400

rangefront, and where measured (at lower elevations) dip ~20˚, though the shear zone is

401

observed to flatten out to <10˚ at the crest of the range (Figure 6c). Lineations, defined

402

dominantly by ribbons of quartz, are consistently oriented WNW with much less variance than

403

the strike of the foliations (Figure 4); slip on the northern part of the SLD is highly oblique (see

404

kinematic data on Figure 4). Kinematic indicators such as S-C fabrics, sigma and delta clasts

405

indicate a top down to the west, normal sense of shear (Figure 6a, b). Large feldspar crystals

406

show brittle deformation instead of ductile deformation, indicating that mylonitization was not

407

entirely in the ductile regime as might be expected if the shear zone formed during magma

408

emplacement. The mylonitic shear zone is therefore interpreted to be the exhumed down-dip

409

extension of SLD shear in the brittle-ductile transition zone. The orientation of foliations broadly

410

defines an antiform that we suggest is a single corrugation of the detachment footwall, with the

411

antiformal axis trending in the direction of extension (Figure 4). This configuration is similar to

412

analogous structures well-defined in the Basin and Range extensional province of the western

413

US [e.g., Deubendorfer et al., 2010; John et al., 1987], as well as in metamorphic core

414

complexes world-wide [e.g., Spencer, 2010].

415

[29] Brittle structures in the SLD footwall consist of chatter marks on the west- to north-

416

dipping foliation planes, with steps consistent with top-W (normal sense) displacement. East-

417

dipping low-angle normal faults with mm to cm scale offsets were also observed; these are

418

interpreted to be products of flexural rotation of the footwall through the upper (antilistric) part

419

of a rolling hinge [e.g. Buck, 1988; Axen and Bartley, 1997]. This interpretation is supported by

420

the observed shallowing of the mylonitic shear zone at the crest of the Surla Range, and by (U-

421

Th)/He data and thermokinematic modeling (Section 5).

422

[30] Though no seismic events on the SLD are represented in global catalogs, it is

423

believed to have been recently active due to well-developed fault scarps along its trace (Figures

424

3, 7). The largest fault scarps associated with the SLD are 2 and 3 km basinward of the

425

rangefront (Figure 4). These two west-dipping fault scarps cut a large lateral moraine extending

426

into the basin and have a cumulative ~112 m of down-to-the-west throw, as determined by Jacob

427

staff field measurements. These faults are considered to sole into the SLD at depth; observations

428

of much smaller scarps at the SLD’s trace at the rangefront immediately east implies that the

429

SLD is active and uncut by the faults in its hanging wall. This arrangement of faults is similar to

430

that observed in the North Lunggar Rift [Kapp et al., 2008], although in the North Lunggar Rift

431

the detachment is inactive at the range front; the important point is that the dominant neotectonic

432

expression of faulting has migrated away from the rangefront, but slip is interpreted to occur

433

along the detachment at depth. The dissection of the detachment hanging wall by high-angle

434

normal faults is a very common feature of low-angle normal fault systems in the Basin and

435

Range and may be a consequence of isostatic uplift of the footwall following tectonic

436

exhumation [Kapp et al., 2008], and possibly as part of an evolving rolling-hinge detachment

437

system [e.g. Axen and Bartley, 1997].

438

[31] The structure of the northern Surla Range, where bounded by the SLD, is shown in a

439

cross section across the range (Figure 8a). The observed gently antilistric geometry of the

440

detachment fault and underlying shear zone are continued at depth. The maximum amount of

441

possible extension in the South Lunggar Rift can be estimated by the horizontal distance between

442

pre-extensional strata in the hanging walls. This distance is ~20 km at the latitude of the section,

443

although it increases northward. This estimate is dependent on the depth of the flanking rift

444

basins, which is unconstrained; the deeper the basins, the greater the distance between the pre-rift

445

hanging wall strata. By limiting basin depth to less than 1 km, consistent with observations of

446

other supradetachment basins [Cogan et al., 1998; Friedmann and Burbank, 1995], we obtain a

447

lower-bound estimate for maximum extension (deeper basins would move the hanging-wall

448

pinning points farther away from each other, which would increase the possible maximum

449

extension). It is important to note that the estimate of maximum extension is only influenced by

450

plausible fault geometries in a small way, as can be envisioned by examining the cross section

451

(Figure 8a). If the faults were vertically-dipping, the hanging-wall pinning points would be <5

452

km closer to each other; conversely, if the faults were very shallowly-dipping (both <20°) the

453

maximum extension would be <5 km greater.

454

[32] ZHe ages young westward, suggesting that the SLD has accommodated significantly

455

more exhumation (and therefore extension) than the Palung Co Fault at this latitude. This

456

requires top-east horizontal-axis rotation of the Surla Range away from the detachment fault as

457

has been suggested for detachment footwalls elsewhere in Tibet [e.g., J. Kapp et al., 2005; Kapp

458

et al., 2008] and worldwide [e.g., Buck, 1988]. This is discussed in more detail in Section 5.

459 460

3.3 Moderate angle normal fault

461

[33] To the south of the SLD, uplift of the Surla Range is accommodated by a moderate

462

to high angle normal fault. The fault changes strike from NNE in the north, near the SLD, to E-

463

W in the south as it wraps around west to bound the southern margin of the western South

464

Lunggar rift basin (Figure 3). Though the fault is not exposed, subordinate high-angle (~55˚)

465

west-dipping small-displacement faults in the footwall are likely parallel to the master fault. The

466

rangefront of the Surla Range becomes significantly steeper south of the SLD as well.

467

Quaternary fault scarps on the range-bounding fault were not definitively observed in the field.

468

Cross-section B-B’ (Figure 8b) characterizes the southern part of the Surla Range. Maximum

469

extension across this section is ~16 km, with the same assumptions and caveats as the northern

470

cross-section. As discussed in more detail in Sections 5 and 6, zHe data and thermokinematic

471

modeling suggests that at this latitude, both the west-dipping and east-dipping (Palung Co Fault)

472

structures have accommodated similar amounts of exhumation, and little to no horizontal-axis

473

rotation of the Surla Range has occurred.

474 475

3.4 Interior structures of the Surla Range

476

[34] No major structures were mapped within the interior of the Surla Range; however,

477

fault surfaces were found in almost every outcrop in the southern Surla Range, where older rocks

478

were more exposed and accessible. These dominantly strike roughly E-W, although some high-

479

angle N-striking fault surfaces were observed (Figure 4). No evidence was found for a shallowly

480

west-dipping fabric that could have been reactivated during extension and influenced detachment

481

fault geometry.

482

[35] A well-developed foliation, with planes meters to 10s of meters thick, dips

483

northward moderately to gently (Figure 5c); in places, it is unclear if this fabric is a true

484

metamorphic foliation or if it is an intrusive complex, with younger leucosomes intruding older

485

mafic and felsic rocks (possibly along a pre-existing fabric). In most locations, no foliation was

486

observed in the mafic rocks (mostly amphibolites) at the outcrop scale, although they were

487

generally more pervasively faulted. This may indicate that metamorphism occurred at relatively

488

low temperatures, and the amphibolites deformed brittlely while the felsic orthogneisses

489

deformed ductilely. Alternatively, this could indicate that the mafic rocks are younger than the

490

foliation event; these may be basalt dikes that were subsequently metamorphosed to amphibolite

491

facies under low differential stress.

492 493

3.5 Accommodation zone

494

[36] E-W striking brittle faults bound the northern side of the northwestern Surla Range.

495

Though we found no fault striations on the northern brittle faults that would indicate the rake of

496

slip, the faults are probably oblique slip (normal and left-lateral), given their orientation relative

497

to that of the strain field in the region and throughout Tibet, which is undergoing ~N-S

498

contraction. This would give them similar kinematics to that observed in the mylonites several

499

km to the west along strike. It is likely that these faults serve to relieve stresses related to

500

significant differential exhumation of the Surla Range to the south and less-exhumed rocks to the

501

north. The northern Surla Range decreases in elevation to the north, into the accommodation

502

zone between the North and South Lunggar rifts. The Palung Co Fault and the North Lunggar

503

Detachment tip out on the east side of the range here, and uplift is only accommodated on the

504

western fault, which runs north from the SLD to the southern North Lunggar Shan. ZHe cooling

505

ages are Oligocene for rocks in the footwall of this western fault, indicating limited exhumation

506

(see Section 5). Steep topographic breaks and exhumation of granites and gneisses juxtaposed

507

against lower grade rocks suggest that significant normal faulting exists within this zone in

508

addition to the rangefront faults (Figure 3, Figure 4), although this area was not mapped in detail.

509 510 511

4 Zircon U-Pb geochronology

512

[37] Zircons from a mylonite sample (SLW-NMT-02) from the SLD shear zone and a

513

mildly deformed leucogranite sample (SLW-SFTR-02) were dated by the U-Pb method with

514

laser ablation ICP-MS in order to bracket the timing of magmatism and place age constraints on

515

other map units and geologic events through cross-cutting relationships. Two samples were

516

selected because they both display evidence of ductile deformation and have Pliocene zHe

517

cooling ages, raising the possibility that ductile deformation was syn-kinematic, and that cooling

518

ages may be young because of residual heat from magmatism. The first possibility is relevant

519

because it may indicate that observed mylonitization may be a result of intrusive processes

520

instead of detachment faulting [e.g., Daoudene et al., 2012], challenging our interpretation of the

521

northern Surla Range as a metamorphic core complex. The second possibility is relevant because

522

residual heat from magmatism would invalidate the assumption of cooling due to exhumation,

523

rendering our thermochronologic interpretations inaccurate. Alternately, if these leucogranites

524

are very old, the observed ductile deformation may be due to a previous deformational episode

525

involving ~E-W extension, also challenging our core-complex interpretation.

526 527

4.1 Methods

528

[38] U-Pb ages were determined by laser ablation inductively coupled plasma mass

529

spectrometry (LA-ICP-MS) using a Thermo Scientific Element 2 ICP/MS at the University of

530

Kansas. A Photon Machines 193nm ArF excimer laser was used to ablate 29 µm spots on whole

531

zircon crystals placed on double sided-tape. The laser was set to 3.5 J cm-2 fluency at 10 Hz

532

repetition rate, which produced ablation pits of ~20 µm depth, with the ablated material carried

533

to the ICP/MS in He gas with a flow rate of 0.74 L min-1, tied in with Ar gas at 1.0 L min-1 flow

534

rate with a Y-connector 15 cm down flow from the ablation cell. Elemental fractionation, down-

535

hole fractionation and calibration drift were corrected by bracketing measurements of unknowns

536

with GJ1 zircon reference material [Jackson et al., 2004] and data reduction using the VisualAge

537

data reduction scheme [Petrus et al., 2011] for the IOLITE software package [Paton et al.,

538

2011]. Because the zircon crystals were not polished, multiple growth zones were ablated during

539

some analyses. Ages were calculated only for the outermost growth zones (rims) in these cases.

540

Within run reproducibility of the GJ1 reference material [Jackson et al., 2008] was better than

541

2% on the U-Pb age. Results were corrected for diffusive lead loss and common lead with the

542

methods of Andersen [2002].

543 544

4.2 Results

545

[39] Results are shown in Figure 9 and Data Table S1. Python code to calculate statistics

546

is given in the auxiliary materials (calculate_weighted_means.py and geochron_stats.py).

547

Sample SLW-SFTR-02 (Figure 9a) has a

548

confidence: 2σ/√n, n=19, MSWD=0.51). This suggests it is related to Gangdese magmatism, as

238

U/206U weighted mean age of 63.1 ± 0.78 Ma (95%

549

the sample is ~20 km from the northern margin of the Gangdese range. Zircons from sample

550

SLW-NMT-02 (Figure 9b) display evidence of zoning (major, step-wise changes in isotopic

551

ratios during laser ablation) and yield early Miocene rim ages, with a

552

age of 16.2 ± 0.77 Ma (n=15, MSWD=2.0). The sample also shows two populations of rim ages,

553

a dominant group (n=11) with a

554

and a lesser one (n=4) of 21.0 ± 0.84 (MSWD=0.29). These crystallization ages are ~58 and ~12-

555

17 Ma older than the zHe ages for each sample (5.1 ± 0.5 Ma and 3.4 ± 0.2 Ma, respectively;

556

Section 5, Table 1), confirming that Pliocene cooling age for sample SLW-SFTR-02 is not likely

557

to be the result of residual magmatic heat. It is possible that some residual magmatic heat could

558

influence the zHe age of sample SLW-NMT-02, although this effect is likely small, because the

559

zHe age is only ~1.5 m.y. younger than the zHe age from the Paleocene granites to the south, and

560

is from an area that appears to be exhumed more rapidly based on the much larger Quaternary

561

fault scarps and localization of extension on the SLD at that latitude. Additionally, though SLW-

562

NMT-02 displays evidence of zoning, U and Th concentrations are not systematically higher in

563

the rims than in the cores of the zircons, indicating that the (U-Th)/He ages are not influenced by

564

compositional zoning of parent isotopes.

238

U/206U weighted mean

238

U/206U weighted mean age of 15.9 ± 0.53 Ma (MSWD=2.2)

565

[40] Given that the fabrics in both the Paleocene and early Miocene leucogranites are

566

indicative of ~E-W extension, which has not been documented in southern Tibet before the

567

middle Miocene, these fabrics are likely the result of Neogene extensional processes and

568

unrelated to magmatic processes. This is supported by the observation that feldspars within the

569

mylonitic shear zone show only slight evidence for ductile deformation, and pervasive brittle

570

deformation, suggesting that mylonitization took place at cooler temperatures than would be

571

expected for syn-intrusive deformation [see Daoudenne et al., 2012 for the converse case].

572 573

5 Zircon (U-Th)/He thermochronology

574

[41] In order to understand the history of deformation in the South Lunggar rift in a

575

quantitative fashion, we used zircon (U-Th)/He, or zHe, thermochronology. This is a technique

576

that utilizes the temperature-dependent diffusion of radiogenic 4He out of a mineral grain to

577

understand the cooling history of that grain. More specifically, it quantifies the time since a

578

mineral grain cooled through a temperature range that is a function of the diffusion parameters

579

for that type of mineral and the cooling rate, approximated by a ‘closure temperature’ [Dodson,

580

1973].

581

temperature is ~190-200 ˚C [Reiners, 2005]. The thermal sensitivity window (defined as a

582

temperature range) yields a depth range termed the ‘partial retention zone’ via the geothermal

583

gradient. Below the partial retention zone, radiogenic 4He is diffused out of the grain as fast as it

584

is produced, while above this zone, diffusion is extremely slow.

For rapidly-cooled zircons (e.g., cooling rates of 20-100 ˚C Ma-1), this closure

585

[42] The temperature and depth sensitivity of zHe thermochronometry is ideal for

586

studying rifts with significant amounts of footwall exhumation, because the partial retention zone

587

is deep enough to be less sensitive to surface processes such as erosion and hydrothermal

588

circulation than for lower temperature thermochronometers such as apatite, while still being

589

shallow enough to be responsive to tectonic exhumation [Reiners, 2005].

590

[43] We have collected samples from throughout the Surla range, with emphasis on two

591

transects across the range near the cross-sections A-A’ and B-B’. These transects are analyzed

592

through 3-dimensional thermokinematic modeling in order to place quantitative estimates on the

593

deformational history, and the other samples (spanning a more broad geographical range) are

594

interpreted in a less quantitative fashion.

595 596

5.1 Zircon (U-Th)/He results

597

[44] Zircons from 33 samples (2-6 single-grain aliquots per sample) were run for (U-

598

Th)/He analysis at the University of Kansas Isotope Geochemistry Laboratory following

599

procedures described by Wolfe and Stockli [2010]. Individual aliquot outliers were rejected

600

according to Peirce’s criterion [Ross, 2003] Mean sample results are shown in Table 1, and

601

individual aliquot results are shown in Table DR2. The cooling ages of all samples in the Surla

602

Range are late Miocene to Pliocene (Figure 4, Table 1), indicating that late Miocene to present

603

exhumation for the entire range has been greater than the depth of the pre-extensional zircon He

604

partial retention zone, i.e. >5-10 km for mean geothermal gradients of 40–20 ˚C km-1. In

605

general, ages increase both with elevation and with distance from the SLD. For the northern

606

sampling transect, corresponding to cross section A-A’ (Figures 4, 8a), cooling ages decrease

607

monotonically from 7.3 ± 0.6 (1σ) in the east to 3.4 ± 0.2 in the west, at the SLD trace. This

608

cooling pattern suggests that cooling has been accommodated by progressive exhumation and

609

top-to-the-east rotation of the SLD footwall. For the southern sampling transect (Figures 4, 8b)

610

cooling ages decrease from 7.3 ± 0.6 for the highest sample, in the center of the range, downhill

611

and to the east and west. Age-elevation relationships are generally similar for both sides of the

612

Surla Range, though there are more samples on the west side. This pattern suggests relatively

613

vertical uplift of the Surla Range at this latitude, accommodated equally on both range-bounding

614

faults. The eastern range-bounding fault, the PCF, has a stepover between two fault strands

615

where this sampling transect crosses it. Samples from in between the two fault strands show 10-

616

12 Ma cooling ages (Figure 4), suggesting that the tectonic sliver in between the fault strands

617

was not exhumed as much or as rapidly as the main Surla Range.

618

[45] Cooling ages older than late Miocene are found in two locations: an age of 16.8 ±

619

0.8 Ma for a tuff (unit v) in the western rift basin, which is interpreted to be the depositional age

620

of the tuff, as its brittle, vesicular texture is not indicative of deep burial. A metavolcanic rock

621

(mv) in the footwall of the accommodation zone between the North and South Lunggar Rifts and

622

a leucogranite (gr) that intrudes it yielded cooling ages of 32 ± 7 Ma, 26 ± 6 Ma, These samples

623

are interpreted as being within in or above the mid-Miocene (pre-extensional) zircon He partial

624

retention zone, which limits late Miocene to present exhumation of the accommodation zone to

625

be less than ~5-10 km.

626 627

5.2 Thermal modeling with Pecube

628

[46] Although the data provide a reasonable first-order picture of relative exhumation

629

rates and some information on timing, they do not directly provide the fault slip rates or precisely

630

estimate the timing of rift initiation in the South Lunggar Rift. While some of this information

631

can be obtained through analysis of (U-Th)/He age-elevation or age-fault distance relationships

632

[e.g., Stockli, 2005], complications relating to the dynamic thermal (e.g. unknown or transient

633

geothermal gradient, radiogenic heating) and structural (e.g., footwall rotation) state of the

634

extending crust can introduce inaccuracies to such estimates [Stockli, 2005; Robinson et al.,

635

2010; Ehlers, 2003]. We have therefore chosen to analyze our data with the thermochronological

636

modeling code Pecube [Braun, 2003; Braun et al., 2012]. Pecube uses the finite element method

637

to iteratively solve the 3-D heat transport equation throughout an imposed tectonogeomorphic

638

scenario, and is able to incorporate these aforementioned parameters that affect

639

thermochronometric ages. In order to robustly constrain the Miocene to present deformational

640

history of the SLR, we perform an iterative grid search throughout the parameter space

641

characterizing deformational history, constrained by structural data, and then accept the

642

deformational histories that fit the zHe cooling ages within a 95% confidence limit, to quantify

643

the deformation history and uncertainty of the SLR.

644 645

Model setup

646

[47] Our model setup consists of two cross sections corresponding to our sampling

647

transects and structural cross-sections across the Surla Range. Each transect is ~100 km wide,

648

centered on the rift, and 4-5 km from north-south, just enough to encompass the samples for that

649

transect. The model extends to 80 km depth, the modern thickness of the crust (Nábělek et al.,

650

2011). More details are given in Table 2, with explanation below. Topography is modeled as

651

steady-state, so uplift at the surface is balanced by erosion [e.g., Campani et al., 2010]; though

652

the Surla Range is a regional topographic high, the samples were taken from elevations similar to

653

the surrounding terrain beyond the adjacent rift basins. Each model simulation begins at 20 Ma,

654

although the models are insensitive to the duration of simulation before deformation initiates (at

655

variable times). Configuration of Pecube is done completely through two text configuration

656

files, fault_parameters.txt and topo_parameters.txt, which are well documented and included in

657

the supplementary material.

658

[48] Fault dip and thermal parameters are fixed during the formal iteration (constrained

659

grid search). Sensitivity testing of these parameters is performed on the best-constrained model

660

following the formal inversion. This testing illustrates the influence that particular parameters

661

have on the results. In an ideal situation, a full exploration of these parameters would occur as

662

part of the iteration process; but because of the combinatoric nature of the parameters, including

663

two values for a single parameter such as the radiogenic heat production rate doubles the number

664

of models to be tested. However, unlike the fault slip rate parameters we iterate over, for which

665

we have no prior knowledge save maximum allowable net extension, the thermal parameters

666

produce geotherms comparable to elsewhere in Tibet, and the fault geometry is constrained by

667

structural and seismic observations (see below for discussion).

668 669

Thermal setup

670

[49] Our modeling has assumed a Moho temperature of 1200, radiogenic heating of 20 ˚C

671

Ma-1, and a thermal diffusivity of 20 km2 Ma-1 [Whittington et al., 2009]. The Moho temperature

672

is consistent with estimates of 1069-1248 ˚C from studies of middle Miocene xenoliths from the

673

uppermost mantle (50-65 km depth, likely very close to the Moho before late Miocene crustal

674

thickening) from the Sailipu area ~50 km west of the Lunggar Rift [Liu et al., 2011]. The heat

675

production value converts to heat production of 2.39 μW m-3 for a granite with a density of 2700

676

kg m-3 and heat capacity of 224.607 J mol-1 K-1 (calculated at standard temperature and pressure

677

using the equations of Whittington et al. [2009]). This heat production is low for granite [e.g.,

678

Förster and Förster, 2000], and lower than mean estimates for the Appalachian orogen of ~3 μW

679

m-3 [Jaupart et al., 2007], which may be representative of Phanerozoic collisional orogens.

680

[50] These parameters lead to pre-extensional geothermal gradients of >40 °C km-1 in the

681

upper crust, though the geothermal gradient relaxes rapidly with depth, so that it is <30 °C km-1

682

by 15 km below sea level (Figure 10). This geotherm is quite hot, at the high end of continental

683

geotherms including those from volcanic provinces [e.g., Ehlers, 2005] though it is consistent

684

with values estimated elsewhere in Tibet (Figure 10). Early model testing (part of model

685

refinement before the formal iterative modeling) indicated that temperatures in the upper and

686

middle crust are required to be this high in order to produce thermochronometer ages that match

687

the observed values when exhumed by modeled faults matching our structural data. Because of

688

the consistency with other Tibetan geothermal estimates [Figure 10, J. Kapp et al., 2005; Mechie

689

et al., 2004], and because we consider it unlikely that the crust is much hotter than this hot model

690

geotherm, we fix the thermal parameters during the formal iterative modeling. See section

691

5.3.2.2 for a further discussion of the thermal parameters, including model results where they

692

have been changed.

693 694

Fault setup

695

[51] Our model setup consists of two cross sections corresponding to our sampling

696

transects and structural cross-sections across the Surla Range (Figure 11). The northern transect

697

(corresponding to cross section A-A’) is modeled with an antilistric detachment fault on the west

698

side (the SLD) that has a dip of 8° above the range, 22° at the rangefront, and 45° below 4 km

699

depth, and a planar, moderate-angle (50° dip) fault on the east side (the PCF). The southern

700

transect (corresponding to cross-section B-B’) contains two planar moderate-angle faults, each

701

with a dip of 50°.

702

[52] The velocity field in each model is calculated by Pecube internally from the

703

geometry and slip rate prescribed for each fault in the Pecube configuration file

704

(fault_parameters.txt); i.e. no velocity boundary conditions are applied to the model.

705

example velocity field for each model is shown in Figure 11 though the field varies in each

706

model run due to different fault slip rates. Slip along a fault results in horizontal motions of the

707

hanging walls and horizontal and vertical motions of the footwalls; no hanging wall subsidence

708

is included in our model, as hanging wall subsidence would have a only a minor effect on our

An

709

thermochronometer ages or the overall extension rates (which we are most interested in). We

710

allow Pecube to update a fault’s geometry due to slip on other faults in the model.

711 712

Formal iteration

713

[53] As we seek to constrain the slip histories of all major faults in the study area, we

714

chose to explore a broad range of fault parameters (given in Table 2) that is enough to fully

715

encapsulate the realistic geological possibilities. We allow for a Pliocene change in slip rate

716

(positive or negative), as an acceleration has been suggested for other Tibetan rifts [Dewane et

717

al., 2006; Hager et al., 2006; Lee et al., 2011; Sundell et al., manuscript in review in Teconics].

718

Because the relationship between a thermochronometer age and the input parameters is

719

nonlinear, it is necessary to iteratively model a large number of parameter combinations

720

spanning the parameter space in order to rigorously estimate the probability distribution of the

721

parameters. Our choice of fault parameters yields hundreds of thousands of combinations (each

722

combination represents a unique faulting history), not including any variation in thermal

723

parameters. Though the Pecube code is capable of running in an iterative ‘inversion’ mode

724

designed to seek the combination of parameters that best fits the observed (U-Th)/He cooling

725

ages, our tests with it for the North Transect resulted in convergence towards combinations of

726

parameters that were individually reasonable but yielded magnitudes of net extension that were

727

unacceptably larger than our maximum estimates from geologic mapping; the code moved

728

rapidly toward parameters resulting in 60 km net extension and varied little for hundreds of

729

subsequent iterations (for interested readers, we have posted the results as ‘negative results’:

730

https://www.researchgate.net/publication/235332918_South_Lunggar_Rift_North_Transect_Pec

731

ube_NA_inversion_results). Furthermore, with iterative nonlinear inversion techniques and a

732

large parameter space, it can be difficult to ascertain that the parameter space was fully explored.

733

[54] Therefore, we chose to take all possible fault parameter combinations, calculate net

734

extension for each combination, and only model those that yield magnitudes of extension

735

consistent with our geologic cross-sections. This yielded 10,397 model runs for the north transect

736

and 13,998 model runs for the south transect out of hundreds of thousands of combinations

737

before filtering. This is a large number of possible fault parameter combinations, but is the

738

minimum number necessary to rigorously characterize the history of normal faulting in the South

739

Lunggar rift at a level of precision appropriate for the data. Fortunately, each model is

740

independent of the others, so the problem lends itself well to running models in parallel on many

741

processors; indeed, the number of independent computations qualifies this as an ‘embarrassingly

742

parallel’ computational problem in computer science parlance.

743

[55] In order to run the models in a time-efficient manner, we used PiCloud

744

(www.picloud.com), a Python-based interface to Amazon’s EC2 cloud servers. Identical Linux

745

(Ubuntu 11.04) virtual environments were created on Amazon’s servers, and Pecube v.3 was

746

installed on each. A Python script was executed on a local machine that assembled and filtered

747

the fault parameter combinations, edited the fault_parameters.txt file in the virtual environment,

748

ran Pecube in the cloud via PiCloud for each combination, and concatenated the resulting

749

modeled thermochronometer ages for each sample. For further information on procedure or

750

implementation, see the Python scripts in the supplementary files, which are thoroughly

751

commented. Although official statistics were not provided by PiCloud, the total run time versus

752

individual run time suggests parallelization of 30-50x was achieved.

753

[56] We chose to filter our model results by testing each run to see if all output zHe

754

model ages at the sample locations matched the observed cooling ages at either 1 or 2 standard

755

deviations. Because many ages have very low standard deviations (possibly a consequence of a

756

low aliquot sample size that does not represent the true uncertainty of the cooling age), we

757

obtained no fits at 1σ or 2σ for either model. We re-filtered the data, using the larger of the

758

observed 1σ value or an 8% uncertainty that represents the 2σ standard error for the analytical

759

standard (Fish Canyon Tuff) as the sample error in the modeling. All Pecube input files

760

(configuration files, thermochronometer data, and elevation data), Python code and binary (.npy)

761

modeling results files are in the supplementary materials.

762 763

5.2.1 North Transect (A-A’)

764

[57] The northern transect (Figure 11a) generally corresponds to cross-section A-A’

765

(Figure 8a; see Figure 17 for model location). 58 model runs fit the data at 2σ, and none fit the

766

data at 1σ (Figure 12). Initiation of faulting occurred on the PCF first in all model runs and is

767

distributed fairly equally between 10 and 16 Ma, with a median at 12 Ma and a mode at 10 Ma

768

(Figures 14and 15b). Initiation of the SLD is younger, with the majority of runs (48 out of 58, or

769

82%) showing an initiation at 8 Ma, with the remainder at 9 Ma.

770

[58] Initial extension rates (during initial PCF activity but before SLD initiation) were

771

very low; all runs show the PCF slipping at 0.25 mm a-1. Rapid extension began with the

772

initiation of SLD slip, which has accommodated the large majority of extension across the SLR

773

at this latitude. Results indicate initial slip on the SLD between 1.5 and 3 mm a-1, with a median

774

and strong mode at 2.5 mm a-1.

775

contribution of the PCF to the horizontal extension rate, this is essentially the extension rate

Because of the shallow dip of this fault and the small

776

across the rift at this latitude (Figure 15a). Results do not strongly suggest a change in slip rate

777

in the Pliocene; median and modal values stay the same, although the runs with low initial rates

778

increased to the modal values (Figure 15a).

779

[59] Net extension is well constrained at 19-21 km, with only a few results between 19

780

and 20 km (Figure 15c). The median value is 20.62 km. Exhumation is also significantly greater

781

in the west (due to the SLD) than in the east (due to the PCF). Exhumation of a sample currently

782

at the SLD fault trace is ~10 km given our model geometry and 20 km of slip on the SLD (Figure

783

17). However, exhumation along the PCF is less, between 2 and 4 km for 2σ fits. A vertical

784

difference of 8 km exhumation over the 20 km width of the range yields a differential tilt of ~21˚

785

to the east, indicating significant back rotation of the SLD footwall, consistent with many models

786

of LANF and core complex evolution (e.g., Buck, 1991).

787 788

5.2.2 South Transect (B-B’)

789

[60] The south transect model corresponds to cross section B-B’ (Figure 8b; see Figure

790

17 for model location). The south transect had 786 model fits at 2σ and none at 1σ (Figure 13).

791

Ages of fault initiation of both the PCF and the western fault are fairly similarly distributed

792

between 10 and 18 Ma, with an increasing probability towards the younger ages (Figure 15e).

793

Median initiations for both faults are 12 Ma. Modes for fault initiation are 11 Ma for the western

794

fault and 10 Ma for the PCF (consistent with the northern model).

795

[61] Extension rates across the SLR for this model are more poorly constrained, between

796

0.5 and 3 mm a-1; however, there are significantly more fits between 1 and 1.5 mm a-1 (Figure

797

15d). Modal extension rates are 1.0 mm a-1 for both before and after a possible Pliocene change

798

in fault slip rate, also giving little support to post-Miocene acceleration. However, the later

799

distribution is more skewed to the high end, and the median values change from 1.0 to 1.3 mm

800

yr-1, owing to an increase in median slip rates (total slip, not simply horizontal extension) of 0.5

801

to 1 mm a-1 on the PCF; this acceleration has uniform distribution over the parameter space

802

between 3 and 6 Ma. Median slip rates on the western fault are 1.0 mm a-1 before and after an

803

acceleration. Therefore a subtle change in rate is not ruled out by the modeling, but is unlikely to

804

be of significance.

805

[62] Net extension across this part of the SLR ranges from 10 to 16 km, with a higher

806

probability at the high end (Figure 15f). As the two faults show similar initiation ages and slip

807

rates, footwall tilt is unlikely to be significant, but slight rotation toward the east is possible, as

808

the PCF may have initiated slightly later and slipped slightly more slowly in its early history

809

(Figure 17).

810 811

5.2.3 Sensitivity analysis of fixed model parameters

812

[63] Though the possible fault slip rates and ages of fault initiation and acceleration were

813

robustly tested in the previous section, the fault geometry and thermal parameters (radiogenic

814

heat production and Moho temperature) where held fixed at values that produced good results in

815

trials before the main testing phase and are in accordance with structural data from our mapping

816

and geotherms from elsewhere in Tibet. Here we analyze these parameters to see how their

817

variations affect our results. For these analyses, we use the northern transect with a faulting

818

history that corresponds to the best model fit from the previous testing, and individually vary one

819

parameter at a time. The results (model ages) are compared to the best-fit model and to the

820

observed zHe ages.

821

822

5.3.2.1 Variations in detachment geometry

823

[64] Although the geometry of the mylonitic shear zone is constrained by field

824

observations in the exhumed footwall of the model, the geometry of the detachment at depth is

825

not. As discussed in Section 1.1, several models of detachment fault geometry exist. Here we

826

run the prominent models (antilistric, planar and rolling-hinge) as well as a model where the

827

northern Surla Range is bound on the west by a planar high-angle normal fault (instead of the

828

low-angle SLD), essentially testing our interpretation of the northern Surla Range as a

829

metamorphic core complex. As a point of clarification, references here to ‘antilistric’ geometry

830

refer to the decrease in dip of the detachment fault above the exhumed footwall of the range

831

(leading to folding and flattening of the footwall), as opposed to the fault’s projection upward

832

with the rangefront dip, which we call ‘planar’. We run two ‘antilistric’ models: one with a low-

833

angle geometry at depth (‘low-angle antilistric’) and one with a high-angle geometry at depth

834

(‘high-angle antilistric’); the fault geometry in the previous section has this same antilistric upper

835

detachment and subsurface dip in between these (‘moderate-angle antilistric’).

836

the ‘rolling hinge’ model, with a shallow antilistric geometry and a listric geometry at depth. We

837

also test two ‘planar’ models, a ‘low-angle planar’ and a ‘high-angle planar’ model. Fault

838

geometries are shown in Figure 16a.

Then, we test

839

[65] The results are shown in Figure 16b. All the antilistric models, including the rolling

840

hinge, produce similar age vs. longitude patterns, although only the moderate-angle antilistric

841

model (used in the main model phase) fits all the data at 2σ. The low-angle antilistric model

842

produces ages that are ~1m.y. older than the observed ages near the western rangefront, but good

843

fits to the east. The high-angle antilistric model produces ages that are ~1 m.y. too young in the

844

west and good fits in the east. The rolling-hinge model, with a moderate-angle ramp, produces

845

ages that are in between these two models, similar to the moderate-angle antilistric model. These

846

models all incorporate the same antilistric geometry, which produces older cooling ages into the

847

footwall, as is observed in the data. In contrast, the planar fault models produce ages that are

848

slightly younger into the footwall. The low-angle planar model produces ages that are ~1 m.y.

849

older than observations in the west (and identical to the low-angle antilistric model) but become

850

2-3 m.y. too young in the east. The high-angle model produces ages that are all younger than 2

851

Ma.

852

[66] The increase in model ages into the footwall in all antilistric models and the decrease

853

in ages into the footwall in all planar models is consistent with previous studies [e.g., J. Kapp et

854

al., 2005; Campani et al., 2010; Robinson et al., 2010]. It is easily explained by the recognition

855

that, in antilistric models, pre-extensional sample locations have significant vertical separation,

856

and therefore pass through the PRZ at different times, and are rotated to roughly horizontal

857

above the PRZ. In planar models, the footwall is not internally deformed, and the vertical

858

separation of the samples remains constant; however, isotherms are convex upward in the

859

footwall, as the footwall margins are cooled by the colder hanging wall. In all runs, steeper faults

860

produce younger ages. We interpret this to indicate that a steeper fault exhumes deeper, and

861

therefore hotter, rocks; in other words, a steeper fault advects heat upward more efficiently.

862

[67] The results of this analysis show that the shallow geometry of the detachment fault

863

has a great effect on the cooling ages, and that an antilistric geometry is necessary to reproduce

864

the cooling patterns observed in the northern Surla Range; a planar geometry produces the

865

opposite age-longitude trend. A similar cooling age pattern may be obtained by significant

866

domino-style block rotation, as has been observed in Nevada [e.g. Stockli et al., 2002]; however,

867

this is not a possibility in the Surla Range given the opposing dip directions of the range-

868

bounding normal faults. The model results also indicate that the dip of an antilistric detachment

869

at depth does not have a large control on the thermochronometer ages [e.g., Ketcham, 1996], and

870

therefore precise determination of this dip would require other methods. Although only the

871

moderate-angle antilistric model fits the observations at 1σ, it is quite likely that slight variations

872

in the slip rate and timing parameters with other antilistric models could yield similarly good fits.

873 874

5.2.3.2 Variations in thermal parameters

875

[68] Varying the heat production to half its value, 10 ˚C Ma-1, caused a dramatic change

876

in the modeled ages (Figure 16b). Given the colder resultant geotherm, faulting was insufficient

877

to entirely exhume the footwall from below the pre-extensional zircon He partial retention zone.

878

The samples near the trace of the SLD were exhumed from that depth, but are still several m.y.

879

too old. Lowering the Moho temperature to 900 ˚C lead to ages several m.y. too old in the

880

eastern part of the footwall, but samples near the SLD trace were of acceptable age.

881

[69] The geothermal gradient in our preferred model is over 40˚ km-1 in the upper several

882

km of the crust before extension (Figure 10b). Within the footwall block near the detachment

883

fault trace, rapid uplift and tectonic exhumation lead to vertical advection of heat and a

884

compression of isotherms, giving a geothermal gradient of >70˚ km-1 in the shallowest crust.

885

Though these geothermal gradients decrease rapidly with depth, the geotherm for the crust

886

remains elevated.

887

[70] Because net extension in this preferred model is at the upper limit of what is

888

acceptable given the structural constraints, it is not possible to increase the slip rates on the faults

889

in order to compensate for a colder upper crust. While Tibet is almost uniformly declared to have

890

a hot crust [e.g., Beaumont et al., 2001; Francheteau et al., 1984; Hu et al., 2000; J. Kapp et al.,

891

2005], the extremely high modeled temperatures in the lower crust are almost certainly too high.

892

This may be the weakest result of our study. We suggest that radiogenic heat production in the

893

crust is non-uniform, and is probably greatly concentrated in the upper 10-20 km of the crust,

894

due to pervasive intrusions of leucogranites [e.g., J. Kapp et al., 2005; Kapp et al., 2008, this

895

study] that are highly enriched in radioactive elements. However, it is not possible to implement

896

depth-dependent radiogenic heating in the available version of Pecube.

897 898

6 Discussion

899

6.1 Evolution of the South Lunggar Rift

900

[71] The geology, thermochronology and geochronology of the South Lunggar Rift

901

indicate a rift characterized by a central horst block bounded by east- and west-dipping normal

902

faults. In the northern SLR, extension and exhumation are dominantly accommodated on the

903

west-dipping South Lunggar Detachment. Farther south, the east-dipping Palung Co fault

904

becomes the dominant structure. Horizontal extension across the SLR ranges from 19-21 km at

905

the latitude of the SLD to 10-16 km at the latitude of the southern transect. Extension decreases

906

abruptly to the north and likely to the south as well, although perhaps more gradually. Extension

907

rates also increase from south to north, from ~1 mm a-1 to ~2.5 mm a-1 at the latitude of the SLD.

908

Fault initiation is broadly contemporaneous, though there is some probability of an earlier

909

initiation in the south. The onset of more rapid extension in the north is much better constrained,

910

and is most likely at ~8 Ma, with the initiation of the SLD (Figure 14). Extensional faulting

911

appears to have initiated during or a few million years after episodic magmatism in the rift; it is

912

possible that thermal weakening associated with magmatism allowed for the onset of extension.

913

914

6.2 Comparison with the nearby rifts

915

[72] Observations and modeling results from the SLR are generally similar to the North

916

Lunggar rift [Kapp et al., 2008; Sundell et al., in review]. ZHe ages from the North Lunggar

917

detachment footwall are the same age or up to 1-2 m.y. younger than samples from the same

918

relative position in the SLD footwall, likely indicating more rapid extension, although the

919

detachment could be steeper at depth in the north or the crust could be hotter. Furthermore, the

920

structural and (U-Th)/He age distribution patterns in the North Lunggar rift are much more

921

continuous along strike [Sundell et al., in review].

922

[73] The Lopukangri rift [Sanchez et al., 2012] shows a similar age of fault initiation of

923

~15 Ma for the southernmost rift segment, which cuts the southern Gangdese range and

924

structures of the Indus-Yarlung suture zone. The northern rift segment is undated, but the

925

presence of supracrustal rocks (dominantly volcanic rocks) in the footwall suggests that

926

exhumation is less than in the SLR. However, Quaternary normal fault scarps up to 350 m high

927

suggest that modern rifting is rapid.

928 929

6.3 Thermal state of the Tibetan crust

930

[74] The distribution of late Miocene to Pliocene zHe ages and the upper bounds on net

931

extension across the SLR indicate moderate exhumation rates of very hot upper crust. The

932

inference of hot crust is supported by a variety of observations. Volcanism and magmatism are

933

ubiquitous in southern Tibet, and appear to have continued at least until ~16 Ma in the SLR.

934

Younger (~9 Ma) leucogranites have been dated in the footwall of the North Lunggar Rift [Kapp

935

et al., 2008]. Leucogranites give evidence of magmatism derived from low degrees of partial

936

melting, as might be expected of a high overall geotherm, and the ultrapotassic volcanic rocks

937

containing very hot upper mantle xenoliths [Liu et al., 2011] indicate that the basal temperatures

938

were high as well. Hot springs in the North Lunggar rift also provide independent evidence of

939

elevated modern-day crustal heatflow, although these were not observed in the south. The

940

‘Zhongba’ 2008 earthquakes on the Palung Co fault may give some idea of the local geotherm,

941

as well. InSAR and teleseismic body wave modeling of the events gives a centroid depth of ~8-9

942

km, with slip extending 3-4 km below that [Elliott et al., 2010; Ryder et al., 2012]. If the

943

centroid depth lies just above the brittle-ductile transition, as is commonly inferred [e.g., Sibson,

944

1983; Ellis and Stöckhert, 2004], then temperatures may be ~350 degrees at that depth, which is

945

well in agreement with our model away from the detachment footwall where the geotherm is

946

elevated due to ongoing exhumation (Figure 10).

947

[75] Evidence for high crustal temperatures from outside the Lunggar region is

948

widespread.

We observed geysers near Raka, along the Indus-Yarlung Suture Zone, at

949

approximately 29.60˚ N, 85.75˚ E. Franchetau et al. [1984] estimated high heat flow from

950

elevated temperatures in lake sediment boreholes in south-central Tibet, south of the Indus-

951

Yarlung Suture Zone. Thermobarometry in the Nyainqentanglha Rift [J. Kapp et al., 2005]

952

indicates temperatures within error of our pre-extensional geotherms (Figure 10). Mechie et al.

953

[2004] located the α-β transition in quartz at ~17 km in the Qiangtang block through seismic

954

methods, indicating an elevated geotherm there (mean geothermal gradient 39 ˚C km-1), although

955

in the Lhasa block, they found more typical temperatures (mean geothermal gradient 25 ˚C km-

956

1

957

values (>350 mW m-2) in the Yadong-Gulu rift and moderately high values to the west, although

958

observations are sparse; a similar study by Wang [2001] showed the plateau to have a high mean

959

heatflow of ~80 mW m-2.

). Hu et al. [2000] interpolated heat flow observations over the plateau and found very high

The same argument outlined above for elevated temperatures

960

evidenced by shallow seismicity holds for the entire plateau [e.g., Molnar and Chen, 1987;

961

Molnar and Lyon-Caen, 1989; Priestley et al., 2008; Wei et al., 2010], and is supported by the

962

short wavelength of rift-flank uplifts indicating a long-term effective elastic thickness of only 2-4

963

km [Masek et al., 1994]. Elevated heatflow is a necessary condition for large-scale lower-crustal

964

flow, which has been commonly inferred to explain flat topography and extension within the

965

plateau itself [e.g., Cook and Royden, 2008], ductile injection into eastern Tibet [e.g., Clark and

966

Royden, 2000] and extrusion through the Himalaya [e.g., Beaumont et al., 2001; Nelson, et al.,

967

1996]. Evidence consistent with partial melt in the crust is given by seismic reflections [e.g.,

968

Nelson et al., 1996], low Vp/Vs ratios [e.g., Hirn et al., 1995] and widespread leucogranite

969

magmatism [J. Kapp et al., 2005; Kapp et al., 2008; Sanchez et al., 2012; this study].

970

[76] A hot and mobile middle to lower crust may be a necessary condition for the

971

formation of metamorphic core complexes [e.g., Buck, 1988], and is very likely to be responsible

972

for the lack of a major regional lowering of topography around the Lunggar Rift, despite large

973

amounts of crustal thinning and extension [Block and Royden, 1990]. The mobile crust would be

974

able to flow laterally into the extending region to mitigate the gravitational potential energy

975

contrasts that would be produced by the steep topographic gradients from the crustal thinning.

976

That said, the relatively high number of large lakes, both near the rift and within rift basins

977

bound by moderate to high angle normal faults (but not the supradetachment basins [Kapp et al.,

978

2008]), may be indicative of minor regional subsidence, as nearby crust is drawn into the

979

actively-uplifting core complex footwalls.

980

suggested to explain subsidence of the Zhada basin in the Indian Himalaya between the Leo

981

Pargil and Gurla Mandhata core complexes [Saylor et al., 2010].

982

This phenomenon, on a larger scale, has been

983

6.4 Implications for rift and detachment fault development

984

[77] The SLD is unlike many detachment faults in that core-complex type deformation is

985

a relatively localized phenomenon; the western range-bounding normal fault in the central and

986

southern Lunggar rift is ~70-80 km north-south (not taking into account curves in the fault trace),

987

though the SLD and core complex are only about 15 km N-S. However, despite the relatively

988

restricted areal extent of the SLD, it has accommodated greater and more rapid extension and

989

exhumation than any other fault in the SLR. Additionally, faulting along the western rangefront

990

transitions from a typical moderate to high angle normal fault geometry in the south, to low

991

angle, and then back to high angle in the north. Most of the mapped detachment faults in the

992

western US and elsewhere remain at low angle along strike, and are either buried or truncated by

993

other faults on their ends, so the transition from low to high angle is not observed.

994

analogs exist: the North Lunggar rift [Kapp et al., 2008], the Dixie Valley fault (Nevada)

995

[Caskey et al., 1996], the Cañada David detachment, Baja, Mexico [Axen et al., 2000; Fletcher

996

and Spelz, 2009], possibly the Mount Suckling--Dayman Dome metamorphic core complex

997

[Daczko et al., 2011], and segments of the Kenya rift [Morley, 1999] show an along-strike

998

transition from high to low angle normal faulting, with along-strike widths of low angle faulting

999

similar to the SLD. All of these, including the SLD, are associated with magmatism during or

1000

immediately preceding extension. However, magmatism may not be exclusive to the region of

1001

detachment faulting; certainly in the SLR, two dated samples are insufficient to fully constrain

1002

the extent of Miocene magmatism. These observations show that LANF and core complex

1003

development is not a distinct mode of rifting, nor that it will be the dominant extensional mode in

1004

a certain geodynamic environment (such as rapid extension in hot, thick crust). Instead, it is an

1005

end-member in the spectrum of rifting, but one that is generally associated with high magnitudes

Some

1006

of extension and exhumation [e.g., Abers, 2001; Forsyth, 1992], as well as synkinematic

1007

magmatism [Parsons and Thompson, 1993].

1008

[78] The large along-strike variation in uplift and extension over fairly short distances in

1009

the SLR (Figure 17) is striking, but is well constrained by the structural and thermochronological

1010

observations and modeling. This extension gradient must be accommodated by deformation of

1011

the hanging wall, though no suitable structures were observed. A zone of distributed dextral

1012

shear to the northwest of the SLD may be present, but difficult to observe due to the cover of

1013

water, thick moraine and alluvium.

1014 1015

6.5 Timing and rates of Tibetan extension

1016

[79] Our results in the SLR suggest a minimum age for the onset of extension in

1017

southwestern Tibet of ~16-12 Ma (Figures 14 and 15). Furthermore, we find evidence of a rapid

1018

increase in extension rate at ~8 Ma in the northern part of the rift, as slip on the SLD began.

1019

These results are consistent with, and may reconcile, the few other studies of rifting within the

1020

Tibetan plateau. The modeled age of rift initiation here is similar to the results of Blisniuk et al.

1021

[2001] which uses cross-cutting relationships to provide a minimum age of the onset of rifting in

1022

central Tibet. Our results give a similar regional minimum, in that activity may have begun

1023

earlier on a nearby rift. However, our combination of thermal and structural constraints (limiting

1024

maximum extension) provides both upper and lower bounds on initiation of the SLR itself. Our

1025

suggestions of rapid extension related to slip on the SLD are also consistent with work on the

1026

Nyainqentanghla segment of the Yadong-Gulu rift indicating a phase of rifting beginning at 8

1027

Ma [Harrison et al., 1995; J. Kapp et al., 2005]. The thermal histories of samples in the footwall

1028

of the Nyainqentanghla detachment show a rapid latest Miocene to early Pliocene cooling event

1029

related to exhumation of the range due to slip on a high-angle normal fault [Harrison et al.,

1030

1995] or on the detachment itself [J. Kapp et al., 2005]; the more recent interpretation is

1031

supported by seismic imaging of the rift showing the detachment to continue uncut and at a low

1032

angle below the supradetachment basin, and to project to active fault scarps at the rangefront

1033

[Cogan et al., 1998]. However, evidence of higher-temperature cooling (>300 ˚C) in the middle

1034

Miocene is seen in their thermochronology data as well, and the footwall rocks would have had

1035

to have cooled through 350 ˚C in the middle Miocene if the mylonitic shear zone was formed

1036

during the current extensional phase. A scenario involving slow deformation beginning in the

1037

mid-Miocene followed by acceleration, involving slip on large-magnitude detachment faults, at

1038

~8 Ma is consistent with all datasets. While there is no compelling reason to assume a priori

1039

that extension in the SLR and Nyainqentanglha should be contemporaneous, the larger dataset

1040

and more thorough thermal modeling from the SLR show how an earlier and slower phase of

1041

extension preceding an acceleration could be masked due to sparse sampling and the restricted

1042

thermal modeling limited by older computing technology.

1043

[80] The structural and thermochronological data from the footwalls of the SLD, North

1044

Lunggar detachment [Kapp et al., 2008], and Nyainqentanglha detachment [Harrison et al.,

1045

1995; J. Kapp et al., 2005] involve an earlier phase of ductile deformation with superposed

1046

brittle deformation.

1047

deformation and Pliocene-present brittle deformation from throughout the orogen and suggest

1048

that two distinct deformational events occurred in Tibet, and that the earlier, ductile event is not

1049

necessarily related to crustal extension. However, given that the normal-sense ductile shear

1050

occurs in detachment footwalls of rifts showing evidence of active extensional deformation (e.g.

1051

seismicity, Quaternary fault scarps with at least 10s of meters of throw), we prefer the

Ratschbacher et al. [2011] combine evidence of Miocene ductile

1052

interpretation that the change from ductile to brittle deformation is a consequence of progressive

1053

exhumation and cooling of the footwall. Similar interpretations have been made for the North

1054

Lunggar detachment [Kapp et al., 2008], the Nyainqentanglha detachment [J. Kapp et al., 2005],

1055

Ama Drime detachment [Langille et al., 2010] and for many detachment faults in the Basin and

1056

Range [e.g., Wernicke, 1981; Davis, 1983], the Aegean [e.g., Lee and Lister, 1992], the Alps

1057

[e.g., Campani et al., 2010] and Peru [e.g., McNulty and Farber, 2002].

1058 1059

6.6 Contribution to the Tibetan strain budget

1060

[81] As discussed in Section 1.2.2, some estimates have been made for net extension in

1061

Tibet. Given ~20 km extension for the SLR and about 10 km for the Pum Qu-Xainza and

1062

Tangra Yum Co rifts, and assuming small (~1 km) contribution from the various smaller rifts in

1063

the Lhasa block at the latitude of the SLR, we may broadly estimate net extension at 50-70 km.

1064

However, our results from the SLR show that along-strike variation can be significant, and

1065

therefore that applying a single or narrow range of values for Tibetan extension may be

1066

problematic.

1067

[82] Extension rates across the plateau are better constrained over the decadal scale by

1068

GPS geodesy. Zhang et al. [2004] measured 21.6 ± 2.5 mm a-1 extension between 79˚ and 95˚ E

1069

longitude, or roughly the area showing N-trending rifts. The sites SHIQ and TCOQ are located

1070

~400 and ~150 km to the west and east of the SLR, respectively, and have an 100˚ component of

1071

1.0 ± 1.3 and 4.6 ± 3.5 mm a-1, yielding 3.6 ± 4.8 mm a-1 extension across the western Lhasa

1072

block [Zhang et al., 2004]. Though this figure is very imprecise, it gives a most probable value

1073

that is about 1 mm a-1 higher than extension across the SLR, suggesting that the Lunggar rift is

1074

the dominant extensional structure in the western Lhasa block. Interestingly, it also suggests that

1075

extension rates are considerably higher in the eastern Lhasa block; this is supported by the

1076

analysis of Gan et al. [2007] using GPS data from throughout the orogen. These comparisons

1077

assume that deformation rates may be compared from the 10 year scale to the 106 year scale; our

1078

modeling is not sufficient (or intended) to resolve high-frequency changes in slip rate due to the

1079

earthquake cycle, fault interaction, or other processes.

1080

[83] Support for both block-type [e.g., Meade, 2007; Thatcher, 2007] and continuum

1081

deformation [e.g. England and Houseman, 1989; Cook and Royden, 2008] can be found in the

1082

results from the SLR. Block-type deformation is supported by the results that the SLR has

1083

accommodated ~10-20 km of localized extension, which is of regional significance, and likely

1084

the majority of the extension that has occurred at that latitude in the western Lhasa block;

1085

therefore, the SLR bounds regions that are deforming at a much lower rate. However, the rapid

1086

along-strike variation in rates and magnitudes of extension are possibly accounted for by

1087

extension on neighboring normal faults (known or not), or diffuse deformation within the

1088

adjacent crust; in this case, extensional strain would be penetrative at a regional scale.

1089

Therefore, extensional strain is present throughout the western Lhasa block, but concentrated at

1090

the Lunggar Rift; essentially, strain is localized at rift zones instead of individual faults. This is

1091

in contrast to the preferred model of Loveless and Meade [2011], who consider the western

1092

Lhasa block to be essentially undeforming. However, given the lack of published slip rates

1093

across the Lunggar rift at the time that study was performed, the omission is understandable. The

1094

specific results here are consistent with the more general conclusion of Loveless and Meade

1095

[2011] that deformation type is spatially variable, with different areas occupying different

1096

positions on the continuum-rigid block spectrum. The observed distributed extension from the

1097

SLD through the Lopukangri Rift and to the smaller graben to the east [Murphy et al., 2010] is

1098

consistent with studies predicting or observing wide zones of extension in areas of hot and weak

1099

crust [Buck, 1988; Kogan et al., 2012].

1100 1101

6.7 Causes for Tibetan extension

1102

[84] Extension in Tibet has been attributed to a variety of causes. Thorough reviews of

1103

many of these have been published recently (Lee et al., 2011; Ratschbacher et al., 2011), and we

1104

will not attempt to replicate these efforts, as our results are from a small area and are in broad

1105

agreement with published work. However, we will discuss the implications of our results with

1106

respect to several prominent models that involve timing constraints.

1107

[85] Convective removal of mantle lithosphere is commonly inferred to explain the

1108

elevation and extension of the plateau (e.g., Molnar et al., 1993; England and Houseman, 1988).

1109

This hypothesis is generally supported by geophysical studies indicating a hot upper mantle

1110

under the Qiantang block with low Vp/Vs and Poisson’s ratios [e.g., Owens and Zandt, 1997];

1111

colder mantle lithosphere under southern and far northern Tibet is quite reasonably explained as

1112

post-removal underthrusting of Indian and Tarim lithosphere. The timing of convective removal

1113

was earlier considered to occur at ~8 Ma, largely due to the work of Molnar et al. [1993], Pan

1114

and Kidd [1992] and Harrison et al. [1995]; more recently this date has been allowed to be

1115

pushed back by several million years to explain the Miocene deceleration of Indo-Asian

1116

convergence rate [Molnar and Stock, 2009]. Recent studies of mantle xenoliths in south Tibetan

1117

ultrapotassic rocks [e.g., Liu et al., 2011] show that the upper mantle was very hot and

1118

metasomatized by ~17 Ma, strongly suggesting that removal of mantle lithosphere was underway

1119

by this time. Therefore, an increase in elevation (and excess gravitational potential energy)

1120

shortly after this time may explain the middle Miocene onset of extension in central and western

1121

Tibet [Blisniuk et al., 2001; this study] and in the Nyainqentanglha rift [Harrison et al., 1995]

1122

should an early phase of extension have occurred. However, if convective removal occurred in

1123

the early-middle Miocene and explains extension and Indo-Asian convergence rate reduction

1124

[Molnar and Stock, 2009], then it cannot explain rapid extension beginning at 8 Ma [Harrison et

1125

al., 1995; J. Kapp et al., 2005; this study].

1126

[86] Yin [2000] suggested that the synchronous onset of extension from southern Tibet

1127

and the Himalaya north through Lake Baikal resulted from a sub-continental scale change in

1128

boundary conditions, which he attributed to rollback of the Pacific slab subducting below

1129

Eurasia. This timing was later revised to ~15 Ma, in accordance with the observed cessation of

1130

backarc seafloor spreading in the East China Sea [Yin, 2010]. This hypothesis is consistent with

1131

the onset of Tibetan extension, although the ability of the crust to transmit extensional stresses

1132

over >1000 km (from the Pacific coast across China to Tibet) is questionable, given the low

1133

theoretical tensile strength of the crust [England et al., 1985]; stress transmission may be aided

1134

by east-directed compression on the South and East Chinese cratons due to the Tibetan plateau’s

1135

excess gravitational potential energy [Kong and Bird, 1998]. Furthermore, the change from

1136

tension to compression across the western Pacific subduction zones in the late Miocene indicates

1137

that another mechanism, such as east-directed asthenospheric flow from beneath the Tibetan

1138

plateau is responsible [Yin, 2010; Yin and Taylor, 2011], although estimates for the initiation of

1139

this flow have not yet been made.

1140

[87] In their work considering various changes to Tibetan geodynamics that may induce

1141

extension, England and Houseman [1988] discuss how a reduction in Indo-Asian convergence

1142

rate could lead to extension; essentially, N-S compressional stress is linearly related to

1143

convergence rate, and a reduction in the former would lead to a reduction in the latter. Though

1144

they discount this possibility on the grounds that their models show the decrease would have to

1145

be far more drastic than the contemporaneous data allowed for, we suggest otherwise. The

1146

presence of both normal and strike-slip faulting, both accommodating E-W extension, indicate

1147

that the N-S compressional stress and vertical compressional stress are close to equal, whereas

1148

the E-W stress is the minimum compressive stress [Molnar and Lyon-Caen, 1988]; the change

1149

between normal faulting and strike-slip faulting may be related to modest changes in vertical

1150

stress related to local variations in elevation [Elliott et al., 2010; Styron et al., 2011b]. This

1151

modern near-equilibrium between N-S and vertical stress suggests that a decrease in N-S stress

1152

in the middle Miocene due to convergence deceleration may be sufficient to initiate extension.

1153

Coeval with the mid-Miocene onset of extension on the high plateau is a change from N-S

1154

thrusting to strike-slip faulting along E-W striking faults in northern Tibet [Lease et al., 2011],

1155

also consistent with a decrease in N-S compression (or an increase in vertical stress). However,

1156

middle Miocene Indo-Asian convergence rate decrease does not explain the extensional

1157

acceleration at 8 Ma, either.

1158

[88] It should be noted that none of these models are mutually exclusive, and some of

1159

them may be linked, such as the hypothesis that delamination and uplift caused the Indo-Asian

1160

convergence deceleration [Molnar and Stock, 2009]. Additionally, because the estimates of

1161

timing and rates of Tibetan extension are constrained by sparse data of different types, testing of

1162

these models for Tibetan extension may not be possible with sufficient resolution to falsify any

1163

of them. However, none of these models explain the observed rapid extension at 8 Ma. As this

1164

is based on only two data points, it is unclear whether this is a local signal relating to (for

1165

example) detachment fault evolution, or whether it represents a regionally extensive signal.

1166

1167

7 Conclusions

1168

[89] We provide the first geologic mapping and zircon U-Pb geochronology and zircon

1169

(U-Th)/He thermochronology of the South Lunggar rift in western Tibet. The SLR is a large N-

1170

S trending active rift and is the southern segment of the Lunggar Rift, likely the major

1171

extensional structure in southwestern Tibet. Robust thermokinematic modeling with Pecube

1172

(~25,000 simulations) indicates that extension initiated in the middle Miocene (16-12 Ma) and

1173

accelerated in the late Miocene (~8 Ma). Significant along-strike variation exists in deformation

1174

rates; horizontal extension rates are ~1 mm a-1 in the south and 2.5 mm a-1 in the north, and net

1175

extension is between ~10 and 21 km, respectively. The lower rates and magnitudes of extension

1176

in the southern SLR correlate with higher-angle normal faulting, while the higher rates and

1177

magnitudes correlate with the South Lunggar Detachment, a fairly narrow (~15 km along-strike)

1178

low-angle normal fault that has exhumed a metamorphic core complex. Testing of multiple fault

1179

geometries indicates that an antilistric geometry of the upper SLD (shallowing to sub-horizontal;

1180

i.e. the upper hinge of a rolling-hinge) is necessary to reproduce the zHe cooling age distribution;

1181

although several subsurface geometries are permissible, the best fit was provided by a

1182

detachment geometry that steepens to moderate angles at depth. Our results also show that the

1183

Tibetan crust is very hot; pre-extensional geothermal gradients are ~40 ˚C km-1 in the upper

1184

several km of the crust, and currently higher within the footwall of the SLD. Though several

1185

geodynamic models may explain the timing of rift initiation in the SLR in the early to middle

1186

Miocene, none so far explain the onset of rapid extension at 8 Ma.

1187 1188

1189

[90] Acknowledgments. RS thanks Jhoma Tsering and her associates for help in the field, and

1190

Roman Kislitsyn and the KU IGL group for laboratory assistance. Jean Braun, Frederic Herman

1191

and Dave Whipp provided useful advice about Pecube modeling. Discussions with Tandis

1192

Bidgoli, Paul Kapp and Chris Morley proved enlightening. We thank Doug Walker for his

1193

encouragement in pushing us to test between magma emplacement and faulting as the cause of

1194

mylonitization. This manuscript, especially of modeling description, benefited greatly from

1195

careful reviews and comments by Tectonics editor Todd Ehlers, associate editor Frederic

1196

Herman, reviewers Gweltaz Mahéo, Marion Campani, and an anonymous reviewer. This work

1197

was supported by Tectonics Division of the National Science Foundation (grants EAR-0809408

1198

and EAR-0911652). The conclusions of this work do not necessarily represent those of the NSF.

1199

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1578

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1581

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1582

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1584

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1586 1587 1588 1589 1590 1591 1592 1593

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1595

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1596

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1597

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1598

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1599

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1600

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1601

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1602

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1603

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1604

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1605

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1606

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1609 1610

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1611

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1612

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1613

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1614

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1615

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1616

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1618

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1619

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1621

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1622

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1623

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1624

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1625

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1626

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1627

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1628 1629 1630

1631

Figure and Table Captions

1632 1633 1634 1635 1636 1637 1638

Figure 1: Active tectonic map of the Himalayan-Tibetan orogen. Structures are from HimaTibetMap-1.1 [Styron et al., 2010]. Thick red structures = thrust faults. Thin red lines = fold axes. Blue lines = normal faults. Orange lines = strike-slip faults. Dashed black lines = suture zones. BNS = Bangong-Nujiang Suture Zone. IYS = Indus-Yarlung Suture Zone. Topography is from Shuttle Radar Topographic Mission. Lakes are from GSHHS [Wessel and Smith, 1996]. Black box indicates location of Figure 2.

1639 1640 1641 1642 1643 1644 1645 1646 1647 1648 1649

Figure 2: Active structures of central and southern Tibet. Map symbology is the same as in Figure 1. Earthquake focal mechanisms are from Global CMT (www.globalcmt.org) from 1 Jan 1976 to 20 Mar 2012 above 80 km depth. SH = Shuang Hu graben. LC = Lamu Co fault. NLR = North Lunggar rift. SLR = South Lunggar rift. XGJ = Xianggangjiang rift. LK = Lopukangri rift. TYC = Tangra Yum Co Rift. PX = Pumqu-Xainza rift. NQTL = Nyainqentanglha rift. KF = Karakoram fault. GM = Gurla Mandhata rift. TG = Thakkhola graben. KC = Kung Co rift. AD = Ama Drime rift. Purple circle marks study locations of Pan and Kidd [1992], Harrison et al. [1995] and Kapp et al. [2005]. Green circle marks study location of Edwards and Ratschbacher [2005]. Red circle marks study location of Blisniuk et al. [2001]. Black rectangle indicates location of our mapping in the Lunggar rift (Figure 3).

1650 1651 1652 1653 1654 1655 1656

Figure 3: Bedrock and Quaternary geologic Map of the North and South Lunggar Rifts and Lamu Co fault. Mapping of North Lunggar Rift modified from Kapp et al. [2008] and our field observations. Mapping of Lamu Co fault modified from Taylor et al. [2003]. NLD = North Lunggar Detachment. SLD = South Lunggar Detachment. PCF = Palung Co fault. See Figure 2 for location. Cross-section lines A-A’ and B-B’ are also shown. Black box indicates location of Figure 4.

1657 1658 1659 1660 1661 1662 1663 1664 1665 1666 1667

Figure 4: Geologic map of the South Lunggar Rift. See Figure 3 for location. Note that crosssection lines A-A’ and B-B’ extend off the map to the east; see Figure 3 for full extent. On stereonets, ‘n’ indicates the number of fault planes plotted, and ‘l’ indicates the number of fault striations or stretching lineations measured; red lines indicate average orientation of lineations. Small blue and yellow arrows indicate position and direction field photographs (Figures 5c and 6c) were taken at. Zircon (U-Th)/He mean and 1σ errors are shown as well. Note the age for collocated samples SLE-SCTR-01 and SLE-SCTR-02 (30.94907 °N, 83.52237) is the mean age for all aliquots. Figure 5: Field relationships in the southern Surla Range. (a): Undeformed leucogranites intruding pervasively fractured ampbibolite. Outcrop is approximately 3 m tall. (b)

1668 1669 1670 1671 1672

Leucogranite intruding greenschist-facies metavolcanic rocks. (c): Northwest looking view of moderately north-dipping gneissic foliation above glacier. Photograph taken from location and orientation indicated by yellow arrow in Figure 4; direction of view is perpendicular to arrow with trend shown.

1673 1674 1675 1676 1677 1678 1679 1680 1681 1682

Figure 6: Field photographs of mylonitic shear zone near the detachment fault trace. (a) Closeup of mylonitic fabric showing interpreted normal shear sense. View is to the south. (b) View W (in interpreted direction of slip) of parallel fault striations and stretching lineations, as well as chatters indicating normal-sense brittle slip along the mylonitic foliation planes. (c) Northnortheast looking view of gently west-dipping mylonitic gneiss in foreground, and foliated rocks interpreted to be continuation of shear zone on the ridgeline in the background. Dashed line indicates bottom of mylonitic foliation; leucogranites below are essentially undeformed. Photograph taken from location and orientation indicated by blue arrow in Figure 4; direction of view is perpendicular to arrow with trend shown.

1683 1684 1685 1686 1687

Figure 7: Southeast looking view of Surla Range showing triangular facets and mylonitic shear zone above the approximate trace of the SLD (here buried under Qm), as well as normal fault scarps (with different degrees of weathering) in Quaternary moraine extending past the rangefront. Direction of view is perpendicular to arrow with trend shown.

1688 1689 1690 1691 1692

Figure 8: (a) Structural cross-section through the northern transect of the SLS (A-A’ on Figure 4) showing low-angle South Lunggar Detachment and moderate-angle Palung Co fault. (b) Structural cross-section through the southern transect of the SLS (B-B’ on Figure 4) showing moderately-dipping faults bounding the range.

1693 1694 1695 1696 1697 1698 1699 1700 1701 1702 1703 1704

Figure 9: Concordia plots of zircon U-Pb ICP-MS rim analyses of two samples showing weighted mean 238U/206U ages. (a) Sample zSLW-NMT-02, mylonitized leucogranite from mylonite zone near SLD trace. Best-fit ellipse shown in black, with individual grain analyses shown in blue. Two populations are visible, one at ~15.9 Ma and one at ~21.0 Ma. (b) Sample zSLW-SFTR-02, ductilely deformed leucogranite from SLR footwall. Ages cluster at ~63 Ma, indicating Gangdese-age magmatism. See Figure 4 for sample locations. Figure 10: (a) Pre-extension and modern geotherm of the footwall of the SLD. Data points are other thermobarometric estimates. Black data are from petrologic studies in the Nyainqentanglha footwall [J. Kapp et al., 2005]. Green data point is from the α-β quartz transition in the Qiangtang block, and orange is the α-β quartz transition in the Lhasa block from seismic reflections [Mechie et al., 2004]. (b) Pre-extension and modern geothermal gradient

1705 1706 1707 1708 1709 1710 1711 1712 1713 1714 1715 1716 1717 1718 1719 1720 1721 1722 1723 1724 1725

from the same location as (a). Steps in geothermal gradient are an artefact of the model and indicate the depth resolution of the model. Figure 11: Cross-sections through Pecube thermokinematic models showing the present-day geometries of the faults and example velocity fields relative to the western hanging walls. The velocity of the eastern hanging wall basin reflects the overall extension rate for a given model run. Note that the orientation of the velocity fields in these models is a function of the slip rate on the faults, and is therefore varied in each model simulation. Also note the change in velocity scale between the models. (a) Velocity field for example North Transect model run. (b) Velocity field for example South Transect model run. Figure 12: Pecube modeling results for north transect. (a) Age vs. elevation plot for observed data (black dots with 1σ errorbars) and model results. Blue lines indicate predicted ages at each sample location for runs where all ages fit the data within 2σ. (b) Age vs. longitude plot for same data and predicted ages. Symbology same as (a). Figure 13: zHe data and Pecube model results for south transect. (a) Age vs. elevation plot for observed data (black dots with 1σ errorbars) and model results. Blue lines indicate predicted ages at each sample location for runs where all ages fit the data within 2σ. (b) Age vs. longitude plot for same data and predicted ages. Symbology same as (a).

1726 1727 1728

Figure 14: Extension rate and cumulative extension for north and south transect Pecube modeling results. Blue lines indicate model fits at 2σ.

1729 1730 1731 1732 1733 1734 1735 1736

Figure 15: Histograms for model run results. a: Horizontal extension rate for the northern transect (A-A’) before (blue) and after (red) a possible Pliocene acceleration for the northern transect. Purple in all graphs indicates overlap of red and blue histograms, and vertical lines indicate median values. b: Fault initation ages for the SLD (blue) and PCF (red). c: Net extension across the northern SLR. d: Horizontal exension rate for the southern transect (B-B’) before (blue) and after (red) a possible Pliocene acceleration. e: Initiation ages for the western fault (blue) and PCF (red). f: Net extension across the southern SLR.

1737 1738 1739 1740 1741

Figure 16: (a) SLD geometries tested during sensitivity testing. Orange = planar high-angle. Blue = steep antilistric. Grey = rolling hinge. Black = moderate antilistric (used in main models). Red = low-angle antilistric. Purple = planar low-angle. Red and purple faults have same subsurface geometry. All modeled faults daylight at mapped fault trace. No vertical exaggeration. (b) Results from sensitivity testing, presented as zHe age-longitude plot, with data

1742 1743 1744 1745

as discrete points with 2-sigma errorbars, and model results as lines connecting modeled ages at data point locations. Green = 1/2 modeled crustal heat production. Light blue = lower (900 ˚C) Moho temperature. Other colors same as in (a).

1746 1747 1748 1749 1750 1751 1752 1753 1754 1755

Figure 17: (a) Approximate contours of net uplift estimated from modeling results, zHe ages and structural mapping. Also shown are zHe sample locations (dark grey dots), geologic structures and contacts (see Figure 3 for symbology), and topography. Grey boxes show the width and location of the north and south transect Pecube models. Uncertainty in contour mapping is estimated at 2-5 km in both value and contour position based on variance in model results. (b) Net extension versus north-south distance from southern edge of map in (a); same scale as map. Dark grey band represents the 95% confidence interval, corresponding to the model results in Figures 12 and 13. Blue-grey box labeled ‘cc’ indicates along-strike extent of core complex, as judged by low-angle brittle and mylonitic fault fabrics and domal geometry.

1756 1757 1758 1759 1760

Table 1: Zircon (U-Th)/He sample summary. Individual aliquot analyses shown in the data repository (Table S2). Age error is 8% 2σ laboratory analytical error (see text for discussion).

1761 1762 1763 1764

Table 2: Parameters for models setup, and rates and timing of faulting for Pecube modeling of the north and south zHe sampling transects. See section 5.2 for discussion of parameters as well as references.

1765 1766

Table 1 Sample

Mean (Ma)

St. Dev. (Ma)

Age err. (Ma)

Latitude (˚)

Longitude (˚)

Altitude (m)

SLE-NMT-02

7.3

0.6

0.6

31.05958

83.54151

5823

SLE-NMT-03

6.3

0.2

0.5

31.08004

83.5342

6063

SLE-SCTR-01

9.4

0.8

0.7

30.94907

83.52237

5366

SLE-SCTR-02

11.8

1.5

0.9

30.94907

83.52237

5366

SLE-SCTR-03

12.3

3.2

1.0

30.94915

83.52008

5604

SLE-SCTR-05

6.0

0.6

0.5

30.94966

83.51246

5477

SLE-SCTR-06

7.2

0.4

0.6

30.95814

83.48333

5826

SLE-SCTR-07

7.3

0.6

0.6

30.96490

83.48569

5979

SLW-BSTR-01

10.2

3.4

0.8

30.94404

83.42437

5450

SLW-BSTR-02

8.6

0.9

0.7

30.93364

83.42674

5478

SLW-BSTR-03

8.9

0.8

0.7

30.92383

83.43140

5641

SLW-BSTR-05

8.8

1.8

0.7

30.91829

83.44883

5873

SLW-BSTR-06a

10.2

0.8

0.8

30.91379

83.45148

5874

SLW-CCTR-03

5.3

0.8

0.4

30.96502

83.43724

5622

SLW-CCTR-04

5.3

0.3

0.4

30.96790

83.43554

5490

SLW-CCTR-05

6.0

0.2

0.5

30.97127

83.45555

5663

SLW-CCTR-06

6.3

0.9

0.5

30.97082

83.46420

5719

SLW-CCTR-07

7.2

0.3

0.6

30.97590

83.47744

5848

SLW-HW-01

16.8

0.8

1.3

31.00171

83.30310

4960

SLW-LK-01

25.9

6.3

2.1

31.27406

83.56464

5010

SLW-LK-02

31.5

6.6

2.5

31.27406

83.56464

5010

SLW-NC-02

4.8

0.4

0.4

31.17597

83.46808

5201

SLW-NFT-01

3.8

0.2

0.3

31.13807

83.43247

5811

SLW-NMT-01

3.5

0.2

0.3

31.07366

83.40467

5381

SLW-NMT-02

3.4

0.2

0.3

31.07363

83.40496

5416

SLW-NMT-03

3.7

0.7

0.3

31.06495

83.41171

5538

SLW-NMT-04

4.4

0.4

0.4

31.06623

83.43498

5609

SLW-NMT-05

4.9

0.6

0.4

31.07644

83.4545

5628

SLW-NWC-01

4.0

0.7

0.3

31.13001

83.40368

5701

SLW-SFTR-01

4.8

0.6

0.4

30.99023

83.41145

5676

SLW-SFTR-02

5.1

0.5

0.4

30.99191

83.41448

5724

SLW-SFTR-04

5.5

0.5

0.4

30.99297

83.41912

5810

SLW-STR-01

6.9

0.6

0.6

30.95806

83.41096

5275

1767 1768 1769 1770 1771 1772 1773

Table 2 Thermal and grid parameters

Value

Unit

Model depth North Transect length, width South Transect length, width FEM node spacing (horizontal, vertical) Thermal diffusivity Radiogenic heat production

80 + elev (a.s.l.) 98 E-W, 4.9 N-S 93 E-W, 4.2 N-S 0.785, 3.55 25 20

km km km km km2 Ma-1 °C Ma-1

Moho temperature Surface temperature Atmospheric lapse rate

1200 0 0

°C °C °C km-1

North Transect fault parameter

Range

Step

Unit

SLD initiation SLD initial slip rate SLD acceleration SLD post-acceleration slip rate PCF initiation PCF slip rate

8 - 18 0.25 - 3.0 2 - 6.5 1.5 - 4.5 10 - 18 0.25 - 1.5

1 0.25 - 0.5 0.5 0.5 2 0.25 - 0.5

Ma mm a-1 Ma mm a-1 Ma mm a-1

South Transect fault parameter

Range

Step

Unit

PCF initiation PCF initial slip rate western fault initiation western fault initial slip rate fault acceleration (of both faults) PCF post-acceleration slip rate western fault post-acceleration slip rate

10 - 18 0.5 - 2.0 10 - 18 0.5 - 2.0 3-6 0.5 - 3.0 0.5 - 3.0

2 0.5 - 1 2 0.5 - 1 1 0.5 - 1 0.5 - 1

Ma mm a-1 Ma mm a-1 Ma mm a-1 mm a-1

1774 1775 1776 1777 1778

1779 1780 1781

1782

75˚

80˚

85˚

90˚

105˚

100˚

95˚

40˚

40˚ Tarim basin

China

Figure 2

35˚

35˚

H

Qiangtang block

im

BNS

Tibetan plateau

a

30˚

la

30˚

ya

Lhasa Block

IYS Nepal

India

Range

25˚

75˚

80˚

85˚

90˚

25˚ 95˚

100˚

105˚

80˚

85˚

95˚

90˚

35˚

35˚

>13.5 Ma SH

Tang

gula

Shan

LC Fig. 3 NLR SLR

8 Ma

GM TYC

NQTL

Gang

30˚

dese

Range

80˚

85˚

AD

Ya d

KC

on

g

TG

PX

lu

LK

Gu

30˚

rift

XGJ KF

90˚

95˚

83.00

32.25

82.75

(

Jr

83.25

(

(

83.50

83.75

(

(

( K

Lamu Co fault

Qa

Quaternary alluvium

Qo

Quaternary alluvium (older)

Qsh

Quaternary shorelines

Qm

Quaternary moraine and outwash

N-Q

Neogene - Quaternary strata

myl

mylonites

Pz-Mz

Qsh

K

4

K

ice

Jurassic strata

Qo K

gr ice

Ngangla Ringco

lake

4

4

mv

Qm

Qo

4

amphibolite-facies metamorphic rocks greenschist-facies volcanic rocks

NLD N-Q

4

Qa

Qa

No

4

Qo

Paleozoic-Mesozoic strata

rth

Lu

volcanic rocks Cretaceous strata and Tertiary volcanic rocks

ma

ng

4

v

Pz-Mz

N-Q

Qm

ga

Qa

r ri ft

ma

granite

Jr

31.50

Qa

4 4

31.75

ice and snow

gr

Qa

K

ice

4

32.00

Map Units

Qsh Qa

Pz-Mz

normal fault, active

Qsh

Figure 4

gr

18

ng e

?

4

intrusive contact, inferred or buried v

depositional contact

24 km

Qa

Ra

B

ice

mv

Palung Co Qa

e Range

v

@ A’

PCF

gr

gr

Gangdes

Qm

Su rla

SLD

intrusive contact

12

ice

A

fault, inactive, buried 31.00

4

myl

?

4

4

right-slip fault

4

K

( Pz-Mz

mv

ma

Qm ??

Pz-Mz

4

detachment fault, queried

?

Sou th L ung gar rift

4

31.25

(

4

detachment fault, known

0 3 6

(

Qo

gr

4

(

thrust fault, inferred

?

(

Ringinyubo Co

normal fault, active, inferred

?

ma

@

Contacts

Qsh

B’

5000

5250

Qsh

Pz-Mz

55 0

ma 600 0

0

Pz-Mz

0

0

57 5

0 60

Ringinyubo Co

31.30

50

00

31.5±6.6 n=4 n=7

l = 53

Qa

25.9±6.3 5 57

n=2

0

Qsh

l = 41

5750

5500

mv 0 55

31.20

gr 0

4.7±0.4 0 500

@ 4 @ 4 ͠ @͠

6000

5750

6250

͠

4

@

͠

60 0

͠ ͠

͠

4

͠

ice

( 625(0 ͠

4.9±0.6 ͠

͠ ͠

͠ ͠

4

͠

͠

4

͠ ͠

͠ ͠

4

6250

0

Palung Co Fault

650 0

5750

600 0

j

h h

625 4.4±0.4 0

͠

͠ 6500

62 5

ice

͠

͠

4

4

70

A’

͠

h

37

7.3±0.6

600 0 ͠

4

6000

6000

͠

͠

6000

(

͠ ͠

͠

15

@ 5 62

39

61

6250

mv

Qsh

7.2±0.3 50 62 6250

h

44

5250

5 62

0

12.3±3.2

Palung Co

0

(

30.90

8.9±0.8

( (

8.8±1.8 6250

n=5

62 5

0

Qa

6250

ice 5500

gr 83.30

62

6250

50

n=6

10.3±1.6

7.3±0.6

10.2±0.8625

6250

Palung Co Fault

B’

h

strike and dip of fault metamorphic antiform foliation ~~~ gradational shear contact

4.4±0.4 zHe cooling age (Ma) 23.4±4 U-Pb age (Ma) for other symbols, see Figure 3 Contour interval 50 m

0 1 2

4

6

83.50

8 km

6000

83.40

v

6.0±0.6 ma

7.2±0.4

6250

8.6±0.9

n = 20

0

525 0

h

6.0±0.2

5.3±0.8

10.2±3.4

6250

6.3±0.9

5.3±0.3

6.9±0.6

n = 11



6250

v

5.5±0.5 62.9±2.9 5.1±0.5 4.8±0.6

͠

h

50

00

23

16.8±0.8

j

50

00

͠

31.00

0

͠

15.9±1.6 3.4±0.2 3.7±0.7

n = 20

͠

͠

͠

B

(

͠

4

A

͠

625

͠

͠

Qm

6.3±0.2 ͠

͠

͠ ͠

͠

4

(

31.10

͠

͠

3.5±0.2 Qa

0

͠ ͠͠

myl

South Lunggar Detachment

5250

0

l = 132

4 @

0

4.0±0.7 Qm

0 60

n = 22

5750

3.8±0.2

5 57

ma

0 575

5250

83.60

a

b

g

mv mv

c WNW

foliations

40˚

NNW

~50 m glacier

b

a E

W

W chatter marks fault striation direction

c W

98˚

E

235˚

mylonitic shear zone triangular facets

SLD trace

4 4

4 4

E

fault scarps in Qm

S

elevation (m)

a A 8000

South Lunggar Detachment 4.4±0.4

6000

Qm

4000 2000

N-Q Pz-Mz ~ ~~

3.7±0.2

3.4±0.2 3.5±0.2

~~ ~~ ~~ ~~ ~ ~~ ~~

l ~~

y~~ m ~~ ~~ ~~ ~~

~ ~~~~

4.9±0.6

Palung Co fault 7.2±0.7 6.3±0.2 Qm

v

N-Q

gr

gr

ma

Pz-Mz

0

b 7000

~~~ ~~~ ~~~ ~~~ ~~

A’

~~~ ~~~~~~~ ~~~~~ ~~~~ ~~~~ ~ ~ ~ ~~

1

2

3

4

5 km

7.3±0.6 7.2±0.3 7.2±0.4 Palung Co fault 6.0±0.2 6.3±0.9 10.6±0.3

B

B’

10.3±1.6 5000 3000

N-Q Pz-Mz gr

5.3±0.3

5.6±0.6

6.0±0.6

v

N-Q Pz-Mz

gr gr 0

1

2

3

4

5 km

b

a zSLW-NMT-02

zSLW-SFTR-02 25 Ma

4x10-3

206

Pb /

238

U

21.0 ± 0.84 MSWD = 0.29 n=4

3

10 60 Ma

15 Ma 15.9 ± 0.53 MSWD = 2.2 n = 11

2

10

70 Ma

11x10-3

20

9

30x10 207

-3

Pb /

60 235

U

63.1 ± 0.78 MSWD = 0.51 n = 19

70x10

-3

b

elev (km above sea level)

a

pre-extension modern

temperature (˚C)

geothermal gradient (˚C km-1)

height above model base (km)

90

a

W

E

2 mm a-1

b

W

E

2 mm a-1

80 70 60 50 40 30 20 10 0

0

20

40

60

80

100 0 20 east-west distance (km)

40

60

80

100

a 6100 SLE-NMT-03

elevation (m)

5900 5700

SLW-NMT-03

5500 5300 2.0

zHe age (Ma)

b

SLW-NMT-04

SLE-NMT-02 SLW-NMT-05

SLW-NMT-02 SLW-NMT-01

3.0

4.0

5.0 zHe age (Ma)

6.0

8.0

7.0

8.0

SLE-NMT-02

7.0 6.0 5.0 4.0 3.0 2.0

SLW-NMT-04

SLW-NMT-05 SLE-NMT-03

SLW-NMT-02 SLW-NMT-03 SLW-NMT-01

83.42

83.44

83.46 83.48 longitude (˚)

83.50

83.52

83.54

83.56

a

6000

SLE-SCTR-07

elevation (m)

SLW-CCTR-03

5800 5600 5400

zHe age (Ma)

b

SLW-CCTR-06 SLW-CCTR-05 SLW-CCTR-04 SLW-CCTR-03

6.0 5.0

SLE-SCTR-05

6.0

5.0

8.0 7.0

SLE-SCTR-06

SLW-CCTR-03

zHe age (Ma)

SLE-SCTR-06 SLW-CCTR-07

7.0

8.0

SLE-SCTR-07 SLE-SCTR-05

SLW-CCTR-06

SLW-CCTR-05 SLW-CCTR-04

83.44

83.46

83.48 longitude ( ˚ )

83.50

83.52

rate (mm yr-1)

4.0

north transect extension rate

south transect cumulative extension

north transect cumulative extension

3.0 2.0 1.0 25

extension (km)

south transect extension rate

20 15 10 5 0

20

15

10

5

0 time (Ma)

15

10

5

0

60

north transect

50

a

45

50

35

20

30

30

15

25 20

20

10

15

10

5

10

0

0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0

5

8

extension rate (mm a )

350

180

d

160

12

14

16

18

18

20

19

20

22

21

net extension (km) 200

e

140

300

f

150

120

250

100

200

100

80

150

60

100

50

40

50 0.5

10

fault initiation (Ma)

-1

south transect

c

25

40

40

400

30

b

20 1.0

1.5

2.0

2.5

extension rate (mm a-1)

3.0

8

10

12

14

16

18

fault initiation (Ma)

20

10

11

12

13

14

15

net extension (km)

16

a

topography

12

fault trace

b

10 all antilistric faults

0

-10

zHe age (Ma)

elev (km. above sea level)

10

8 6 4 2

-20 W

30 20 10 0 distance from SLD trace (km)

-10 E

0

83.42

83.46 83.50 longitude (degrees)

83.54

a

b

50 45 40 35 30

north transect

25 20 south transect

15 10 5 0

>10 net uplift, km

7.5-10

5-7.5

2

2.5-5

4

8

12

1-2.5

16 km

<1

0 5 10 15 20 net extension, km

0

distance N-S, km

4

4

cc

Miocene initiation and acceleration of extension in the ...

the north accelerated at 8 Ma to 2.5-3.0 mm a-1 as faulting commenced on the South ...... parallel' computational problem in computer science parlance. 742 ..... Leucogranites give evidence of magmatism derived from low degrees of partial.

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