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ABSTRACT

Five genetic categories of sedimentary basins have been active within the Indus-Yarlung suture zone and in the neighboring High Himalaya since early Cenozoic time. These include: (1) the Xigaze forearc basin (Aptian–early Eocene), (2) the north Himalayan foreland basin (Paleocene–Eocene), (3) the Kailas extensional basin (Oligocene–Miocene), (4) the Liuqu wedge-top basin (early Miocene), and (5) a set of at least six rift and supradetachment basins that formed by arc-parallel extension (late Miocene–Pleistocene). The older basins (categories 1 and 2) were filled with predominantly deep-marine turbiditic deposits, which shoaled through time to subaerial (but very low) elevations. The other basins (categories 3–5) were filled with alluvial-fan, fluvial, and lacustrine sediments, and these formed at progressively higher elevations, culminating in category 5 basins at essentially modern (or slightly higher than modern) elevations (~4000–5000 m). Development of diverse basin types was a response to changing orientations and relative magnitudes of principal stresses in the upper crust of the suture zone and the northern Himalayan thrust belt. Through the Cenozoic, the orientation of maximum compressive principal stress (σ1) changed from approximately horizontal and north-south (Paleocene–Eocene) to approximately vertical with least compressive principal stress (σ3) oriented north-south (Oligocene–Miocene), to horizontal and north-south (early Miocene), to nearly vertical with σ3 oriented approximately east-west (late Miocene–present). Tectonic stresses associated with the degree of coupling between the converging plates were also potentially important, especially during the Oligocene–Miocene, when the subducting Indian slab was rolling backward relative to the upper Eurasian plate, and during middle to late Miocene time, when the Indian slab was subducting nearly flat beneath the High Himalaya and southern Tibet. Preservation of these extensive sedimentary basins in an orogenic system that is generally being eroded rapidly and deeply stems from original basin-forming mechanisms that produced very large-scale basins (the forearc and early foreland basins) and subsequent evolution of the Himalayan thrust belt in a manner that has isolated High Himalayan basins behind an orographic barrier that protects them from erosion. Recent incision by trans-Himalayan and orogen-parallel suture-zone rivers, however, threatens future preservation of these High Himalayan basins (particularly categories 4 and 5).

INTRODUCTION

Intercontinental collision is among the most important tectonic processes on Earth, affecting the size, distribution, and composition of the continents; producing the largest orogenic belts on the planet; and marking sites of nearly complete obliteration of intervening oceanic basins (Burke et al., 1977; Moores, 1981). Although widely acknowledged to be the most significant collision since late Paleozoic time, the Cenozoic event that united India and Eurasia is surprisingly poorly understood from a kinematic point of view. Whereas nearly universal agreement seems to exist that a significant length of continental Indian lithosphere has subducted beneath Eurasia (e.g., Argand, 1924; Powell and Conaghan, 1973; DeCelles et al., 2002; Kind and Yuan, 2010; Replumaz et al., 2010, 2014), its dynamics since initial entry into the subduction zone along southern Eurasia remain a subject of great interest and debate. Another aspect that is intensely debated is the timing of final collision between Eurasia and the Indian subcontinent, with the main arguments focused on apparent differences presented by geological and geophysical data sets, for example, Ding et al. (2005), Najman et al. (2010), DeCelles et al. (2014), and Hu et al. (2015a) versus Aitchison et al. (2002a, 2007) and van Hinsbergen et al. (2012). Broadband seismic experiments shed light on the present architecture of Indian lithosphere beneath Tibet (e.g., Owens and Zandt, 1997; Van der Voo et al., 1999; Hafkenscheid et al., 2006; Nábelek et al., 2009; Kind and Yuan, 2010; Zahirovic et al., 2012), and the Cenozoic magmatic record of the Lhasa terrane provides indications of past Indian plate dynamics (e.g., Chung et al., 2005; Liu et al., 2014; Zhang et al., 2014). Geodynamic models (both analog and numerical) offer a wide range of potential dynamic processes that help to enlighten geological interpretations of the rocks within the suture zone (e.g., Chemenda et al., 2000; Capitanio and Replumaz, 2013; Butler and Beaumont, 2017). Only recently taken into account, however, is the record of diverse sedimentary basins that straddle the suture zone on both the Indian and Eurasian plates. The fact that widespread sedimentary basins have formed within the suture zone and flanking areas almost continuously through Cenozoic time seems surprising, insofar as the region is also the site of Earth’s highest elevation, greatest topographic relief, and greatest rates of regional erosion and uplift (e.g., Thiede and Ehlers, 2013). Sedimentary deposits of these basins provide the history of subsidence, the provenance and paleogeography of sediment dispersal, and the environmental conditions under which the sediments accumulated, including in most cases valuable information about the elevations of the basin surfaces through time. Moreover, the suture-zone basins, which range in age from Cretaceous to the present, formed under widely varying local tectonic conditions, from north-south contraction, to north-south extension, to east-west extension. Considered together, these basins provide an upper-crustal-scale strain gauge for the history of the India-Eurasia collision.

Here, we summarize recent work that documents numerous basins in the broader suture-zone region from the south flank of the Gangdese magmatic arc to the high central axis of the Himalayan fold-and-thrust belt roughly demarcated by the South Tibetan detachment fault (Fig. 1), with the purpose of bringing the disparate data sets into a uniform system of description and interpretation. We rely heavily on interpretations already provided in the original work that we synthesize. What is new here is the temporal and spatial scope of consideration and the insights that result from examining the protracted history of changing upper-crustal strain.

Figure 1.

Simplified geological map of Indus-Yarlung suture zone and Himalayan thrust belt (modified from Yin, 2006), showing locations of discussed sedimentary basins (see legend for abbreviations).

Figure 1.

Simplified geological map of Indus-Yarlung suture zone and Himalayan thrust belt (modified from Yin, 2006), showing locations of discussed sedimentary basins (see legend for abbreviations).

GEODYNAMIC SETTING OF THE INDUS-YARLUNG SUTURE ZONE

The Indus-Yarlung (or Indus-Yalu) suture zone forms the geological boundary at Earth’s surface between the Himalayan thrust belt to the south and the Lhasa terrane to the north (Fig. 1). The suture zone formed in response to ongoing collision between the Indian continental landmass and the rocks that form the southern flank of Eurasia (Burg and Chen, 1984; Dewey et al., 1988; Yin and Harrison, 2000) beginning ca. 60–55 Ma (Garzanti et al., 1987; Beck et al., 1996; Rowley, 1996; Zhu et al., 2005; Green et al., 2008; Najman et al., 2010; Wang et al., 2011; DeCelles et al., 2014; Hu et al., 2015a, 2015b; Zhuang et al., 2015). Cenozoic basins in this region formed within the suture zone proper and on nearby areas of the two colliding continents. The Himalayan thrust belt comprises a south-verging, ~270-km-wide, 2500-km-long orogenic wedge that tapers southward from a thickness of >50 km in the north beneath the surface location of the suture zone (Hauck et al., 1998; Schulte-Pelkum et al., 2005; Nábelek et al., 2009; Gao et al., 2016). Flanking the thrust belt on the south is the actively subsiding Indo-Gangetic foreland basin. From south to north, the thrust belt can be conveniently divided into four major geographic-geologic zones (Fig. 1), consisting of: (1) the frontal Subhimalayan imbricate belt, composed of several major thrust sheets of Miocene–Pliocene foreland-basin deposits of the Siwalik Group; (2) the Lesser Himalayan zone, which is formed of Paleoproterozoic to Lower Miocene sedimentary, low-grade metasedimentary and igneous rocks that are carried in the hanging walls of numerous major thrust faults; (3) the Greater Himalayan zone, a belt of high-grade metasedimentary and meta-igneous rocks that are mostly carried in the hanging wall of the Main Central thrust, as well as locally isolated synformal klippen sitting structurally on top of Lesser Himalayan rocks; and (4) the Tethyan (or Tibetan) Himalayan zone, which is composed of uppermost Proterozoic(?) through Eocene sedimentary and low-grade metasedimentary (in its lower part) rocks. Thrust faults and large folds are ubiquitous, but several faults are particularly important because they are considered to be the structural boundaries between these four zones: The Main Frontal thrust separates the modern foreland basin from the Subhimalayan zone; the Main Boundary thrust divides the Subhimalayan and Lesser Himalayan zones; the Main Central thrust separates the Lesser and Greater Himalayan zones; and the South Tibetan detachment fault forms the boundary between the Greater and Tibetan Himalayan zones (Fig. 1). The first three of these are large-displacement, south-verging thrust faults, whereas the South Tibetan detachment is a top-to-the-north, normal-sense shear zone (e.g., Burg et al., 1984; Burchfiel et al., 1992) that may have experienced southward thrust-sense slip earlier in its history (Hodges, 2000; Robinson et al., 2006; Yin, 2006).

The Indus-Yarlung suture zone consists of two major geological features: the Xigaze accretionary belt and the Xigaze forearc basin (Fig. 2). The Xigaze accretionary belt, which forms the southern part of the suture zone, is dominated by ophiolitic rocks and ophiolitic- and sedimentary-matrix mélanges (Guilmette et al., 2009, 2012; Cai et al., 2012; Hébert et al., 2012; Li et al., 2015b; An et al., 2017; Metcalf and Kapp, 2017). Ophiolitic elements of the accretionary belt crystallized during Early to mid-Cretaceous time (Hébert et al., 2003; Malpas et al., 2003; Bédard et al., 2009; Guilmette et al., 2009, 2012), and they are unconformably overlapped by the Xigaze forearc basin (Orme and Laskowski, 2016; Wang et al., 2017). The northern boundary of the forearc basin is the north-verging Great Counter thrust, which places forearc strata in thrust contact with the Oligocene–Miocene Kailas Formation.

Figure 2.

Geologic map of central Indus-Yarlung suture zone (based on Cai et al., 2011, 2012; Leary et al., 2016a). GCT—Great Counter thrust. See Figure 1 for broader geological context of this map area.

Figure 2.

Geologic map of central Indus-Yarlung suture zone (based on Cai et al., 2011, 2012; Leary et al., 2016a). GCT—Great Counter thrust. See Figure 1 for broader geological context of this map area.

North of the Kailas Formation belt lies the Lhasa terrane, which extends ~300 km northward to the Banggong-Nujiang suture zone. The Lhasa terrane is dominated by the Gangdese magmatic arc, a long-lived cordilleran-type arc with plutonic and volcanic components that formed under variable tectonic conditions above a northward-dipping subduction zone. Onset of magmatism dates back to Triassic time (e.g., Zhu et al., 2011; Wang et al., 2016), and major magma-production events peaked at ca. 111, 90, 52, and 22 Ma (Chapman and Kapp, 2017). Magmatic activity persisted locally until as recently as 8 Ma (Kapp et al., 2005; Ji et al., 2014; Xu et al., 2017). The southern boundary of Kailas basin outcrops is the south-dipping Great Counter thrust, which formed mainly after the basin became inactive (Carrapa et al., 2014).

The timing of collision between India and Eurasia continues to be a subject of active debate. Given recent progress in this debate, we mention only the two most likely times of the collision. Based on first arrival of unequivocally Eurasian detritus and land-mammal faunas on Indian crust, the age of initial contact between the two continental landmasses ranges between 60 and 56 Ma (e.g., Garzanti et al., 1987; Jaeger et al., 1989; Clyde et al., 2003; Ding et al., 2005, 2016; Zhu et al., 2005; Green et al., 2008; Najman et al., 2010; Clementz et al., 2011; DeCelles et al., 2014; Hu et al., 2015a, 2015b; Zhuang et al., 2015). This Paleocene age of collision is in conflict with a younger time of collision between ca. 45 Ma and 25 Ma that is based mainly on plate-circuit reconstructions and paleomagnetic data (e.g., Patriat and Achache, 1984; Aitchison et al., 2007; van Hinsbergen et al., 2012; Huang et al., 2017). Most explanations for a younger collision require a several million square kilometer tract of oceanic lithosphere between Eurasia and India until ca. 45 Ma, but geological evidence for such lithosphere has not been found. Instead, forearc and early foreland-basin records tie the Tethyan Himalaya to the southern Eurasian margin by ca. 60–55 Ma (Garzanti et al., 1987; Najman et al., 2010; DeCelles et al., 2014; Hu et al., 2015a; Orme et al., 2015), and the remainder of what is now the Himalayan thrust belt lay beneath foreland-basin deposits tied by provenance to the Tethyan Himalaya by no later than ca. 45 Ma (DeCelles et al., 1998, 2004, 2014; Najman et al., 2005; Jain et al., 2009; Ravikant et al., 2011). Although younger ages for the collision are consistent with current plate reconstructions and mantle tomographic models (which themselves have large spatial uncertainties), they have yet to find support in multifaceted geological data sets.

Sedimentary basins in the Indus-Yarlung suture can be classified into five genetic types:

XIGAZE FOREARC BASIN

Structural Framework

The Xigaze forearc basin is preserved along ~500 km of the Indus-Yarlung suture zone between ~82°E and 90°E in an ~20-km-wide belt (Einsele et al., 1994; Dürr, 1996; An et al., 2014; Orme et al., 2015). West of ~87.9°E, the southern margin of the Xigaze forearc consists of a south-dipping splay of the Great Counter thrust (Yin and Harrison, 2000) that places ophiolite, ophiolitic mélange, and sedimentary mélange rocks structurally above Xigaze forearc deposits (An et al., 2014; Orme et al., 2015; Hu et al., 2015b; Orme and Laskowski, 2016). East of this point, the southern contact is variable (Fig. 2): North of Lazi (87.9°E), the southern boundary of forearc rocks consists of both a south-dipping thrust and a north-dipping depositional contact with ophiolitic mélange (Orme and Laskowski, 2016); east of 88.2°E, the southern boundary consists of a north-dipping depositional contact with ophiolitic rocks (An et al., 2014). Along its northern flank, Xigaze forearc basin outcrops are juxtaposed with the Kailas Formation by the south-dipping Great Counter thrust. North of Lazi, the forearc shares a boundary with Paleogene intrusive rocks that contain a south-dipping ductile shear zone with top-to-the-north sense of shear (Leary et al., 2016a; Laskowski et al., 2018). Strata of the Xigaze forearc basin are folded into a broad east-west–trending synclinorium (Einsele et al., 1994; Orme and Laskowski, 2016). Deformation at the outcrop scale is highly variable: Some sections are only tilted and moderately folded (Orme et al., 2015), whereas others are densely faulted and isoclinally folded (e.g., ~10 km west-southwest of Xigaze).

Stratigraphy

The Xigaze forearc basin contains ~6000 m of mixed siliciclastic and carbonate strata (Wen, 1974). Various stratigraphic divisions for the Xigaze forearc basin have been proposed (e.g., Wen, 1974; Wu et al., 1977; Einsele et al., 1994; Wu et al., 2010; Wang et al., 2012; An et al., 2014). We use the most recent divisions proposed by Wang et al. (2012) and An et al. (2014).

Forearc strata are divided into the Xigaze and Cuojiangding (or Tso-Jiangding) Groups (Fig. 3). The Xigaze Group contains the Sangzugang, Chongdui (or Chongdoi), and Ngamring Formations. The Sangzugang Formation is the basal part of the Xigaze forearc basin, exposed in tectonic fragments over an along-strike distance of ~40 km west of Xigaze; it is absent from other Xigaze Group exposures. The unit is ~60–230 m thick (Wu et al., 1977) and consists of dark-gray bioclastic limestone containing abundant benthic foraminifera and rudist corals (Wang et al., 2012; An et al., 2014). In fault contact above the Sangzugang Formation, the Chongdui (Chongdoi) Formation makes up the base of the in-place Xigaze Group succession (Cao, 1981). In many locations, the Chongdui Formation is in depositional contact with the Xigaze Ophiolite (Wang et al., 2012; An et al., 2014; Wang et al., 2017); however, the Chongdui Formation is absent from the western part of the basin (An et al., 2014; Orme et al., 2015). This formation is subdivided into a lower member, ~70–100 thick, composed of purplish-red radiolarian chert and siliceous mudrock (Ziabrev et al., 2003), and an upper member, ~200 m thick, consisting of thin-bedded, upward-fining sandstone, green-gray shale, radiolarian mudstone, and tuff (Wang et al., 2012; An et al., 2014; Wang et al., 2017). Conformably above the Chongdui Formation, the highly deformed Ngamring Formation makes up the center of the Xigaze synclinorium. This ~1000–4000-m-thick formation consists of turbiditic volcaniclastic sandstone, lenticular conglomerate, calcareous mudstone, and limestone (Figs. 3 and 4; Einsele et al., 1994; Dürr, 1996; Wang et al., 2012). The Ngamring Formation is the oldest of the Xigaze forearc units preserved along the entire strike length of the basin (Ding et al., 2005; Orme et al., 2015).

Figure 3.

Composite stratigraphic section of Xigaze forearc basin strata (after Wang et al., 2012; An et al., 2014; Orme et al., 2015). Vertical scale is in meters.

Figure 3.

Composite stratigraphic section of Xigaze forearc basin strata (after Wang et al., 2012; An et al., 2014; Orme et al., 2015). Vertical scale is in meters.

Figure 4.

Sandstone modal petrographic data from the Xigaze, Sangdanlin, Kailas, Liuqu, and Zhada basins. All compositions were generated using the Gazzi-Dickinson point-counting method (Ingersoll et al., 1984). Recalculated data are available from the authors. Provenance fields are after Dickinson (1985): UA—undissected arc; TA—transitional arc; DA—dissected arc; LR—lithic recycled; TR—transitional recycled; QR—quartzose recycled; CI—craton interior; TC—transitional continental; BU—basement uplift; M—mixed; RO—recycled orogen. Sources: Xigaze forearc basin: An et al. (2014), Orme et al. (2015), Hu et al. (2016), Orme and Laskowski (2016); Sangdanlin: Zhu et al. (2005), Najman et al. (2010), Wang et al. (2011), DeCelles et al. (2014); Kailas: DeCelles et al. (2011), Leary et al. (2016a), S. Li et al. (2017), Zhang et al. (2017); Liuqu: Leary et al. (2016b); Zhada: Saylor et al. (2010). Abbreviations of grain-types: Qm—monocrystalline quartz; F—total feldspar; Lt—total lithic grains; Qt—total quartzose grains including lithic quartzose grains; L—lithic grains exclusive of quartzose varieties; P—plagioclase; K—K-feldspar; Qp—polycrystalline quartz; Lv—volcanic lithic grains; Ls—sedimentary lithic grains.

Figure 4.

Sandstone modal petrographic data from the Xigaze, Sangdanlin, Kailas, Liuqu, and Zhada basins. All compositions were generated using the Gazzi-Dickinson point-counting method (Ingersoll et al., 1984). Recalculated data are available from the authors. Provenance fields are after Dickinson (1985): UA—undissected arc; TA—transitional arc; DA—dissected arc; LR—lithic recycled; TR—transitional recycled; QR—quartzose recycled; CI—craton interior; TC—transitional continental; BU—basement uplift; M—mixed; RO—recycled orogen. Sources: Xigaze forearc basin: An et al. (2014), Orme et al. (2015), Hu et al. (2016), Orme and Laskowski (2016); Sangdanlin: Zhu et al. (2005), Najman et al. (2010), Wang et al. (2011), DeCelles et al. (2014); Kailas: DeCelles et al. (2011), Leary et al. (2016a), S. Li et al. (2017), Zhang et al. (2017); Liuqu: Leary et al. (2016b); Zhada: Saylor et al. (2010). Abbreviations of grain-types: Qm—monocrystalline quartz; F—total feldspar; Lt—total lithic grains; Qt—total quartzose grains including lithic quartzose grains; L—lithic grains exclusive of quartzose varieties; P—plagioclase; K—K-feldspar; Qp—polycrystalline quartz; Lv—volcanic lithic grains; Ls—sedimentary lithic grains.

The Cuojiangding (or Tso-Jiangding) Group consists of the Padana, Qubeiya, Quxia, and Jialazi Formations (Fig. 3). The Padana and Qubeiya Formations were previously included in the Xigaze Group, but Wang et al. (2012) placed them in the Cuojiangding Group. The Padana Formation, conformable with the Ngamring Formation, is best exposed in the western part of the Xigaze forearc basin. The Padana Formation, ~640–2000 m thick, contains sandstone interbedded with gray to red mudstone; some red mudstone is bioturbated and contains pedogenic carbonate (Wang et al., 2012; Hu et al., 2015b). Facies of the Padana Formation grade into yellow and gray, sandy fossiliferous wackestone and fine-grained sandstone of the ~200–1000-m-thick Qubeiya Formation (Wang et al., 2012; Hu et al., 2015b; Orme et al., 2015). Fossils in the Qubeiya Formation include large benthic foraminifera, bivalves, ammonites, gastropods, and crinoids (Hu et al., 2015b; Orme et al., 2015). The Qubeiya Formation and the overlying Cuojiangding Group are found only in the westernmost part of the basin, west of 85°E. The overlying Quxia Formation is separated from the Qubeiya Formation by an angular, paraconformable or conformable contact (Ding et al., 2005; Wang et al., 2012; Hu et al., 2015b; Orme et al., 2015). This unit consists of ~100 m of sandstone and conglomerate interbedded with shale and thin limestone beds (Wang et al., 2012; Hu et al., 2015b; Orme et al., 2015). The ~145–400-m-thick Jialazi (or Gyalaze) Formation conformably overlies the Quxia Formation and consists of foraminiferal sandy limestone interbedded with sandstone (Ding et al., 2005; Wang et al., 2012; Hu et al., 2015b). The Jialazi Formation represents the youngest marine strata within the Xigaze forearc basin and postdates India-Eurasia collision at ca. 60 Ma (Ding et al., 2005; Hu et al., 2015a, 2015b; Orme et al., 2015).

Age

The Sangzugang Formation was deposited during late Aptian–early Albian (117–110 Ma) time, based on biostratigraphy (Cherchi and Schroeder, 1980; Bassoullet et al., 1980; An et al., 2014). The oldest autochthonous rocks of the Xigaze forearc basin are in the Chongdui Formation, which has yielded Aptian radiolarian biostratigraphic ages (Ziabrev et al., 2003), as well as U-Pb zircon ages between 119 and 113 Ma from interbedded tuffs (Wu et al., 2010; Wang et al., 2017). The overlying Ngamring Formation is ca. 106–83 Ma (late Albian–Santonian) based on foraminifera and U-Pb zircon dating (Wu et al., 2010; Wang et al., 2012; An et al., 2014; Orme et al., 2015).

At the base of the Cuojiangding Group, the Padana Formation is ca. 83–76 Ma, based on U-Pb zircon ages (An et al., 2014). The Qubeiya Formation was deposited from middle Campanian to Maastrichtian time (78–66 Ma), based on its biostratigraphic assemblage and U-Pb zircon maximum depositional ages (Wu et al., 2010; Hu et al., 2015b). Although the overlying Quxia Formation has not been directly dated, its age is constrained by ages from the Qubeiya and Jialazi Formations. Ding et al. (2005) reported two whole-rock 40Ar/39Ar ages of 62 Ma from tuff in the lower Jialazi Formation, as well as a U-Pb zircon maximum depositional age of ca. 60 Ma (weighted mean of all grains analyzed). However, Hu et al. (2015b) reinterpreted the data reported by Ding et al. (2005) to indicate a maximum depositional age of 56 Ma (weighted mean of youngest population). Hu et al. (2015b) also dated two tuff beds at 56 Ma and 55 Ma and interpreted the Jialazi Formation to be 57–54 Ma, based on biostratigraphy. Orme et al. (2015) reported maximum depositional ages of 57–53 Ma.

Deposystems and Paleogeography

The Xigaze and Cuojiangding Groups preserve a record of initial filling of a newly formed forearc basin (Orme and Laskowski, 2016) to the final overfilled basin some 60 m.y. later. Interbedded chert, siliceous mudstone, and tuffaceous layers of the lower Chongdui Formation (119–113 Ma) accumulated on top of ophiolitic basement and are interpreted as abyssal marine sediments deposited on an irregular bathymetric surface (Fig. 5A; Ziabrev et al., 2003; Wang et al., 2017). Based on zircon εHf values, tuff within the Chongdui Formation has been traced to the central Lhasa terrane (Wang et al., 2017), which experienced a magmatic flare-up between ca. 120 and 100 Ma, peaking at 111 Ma (Zhu et al., 2011; Hou et al., 2015). During the same time period, limestone of the Sangzugang Formation accumulated in a carbonate-shelf setting north of the deepest keel of the forearc basin (Wang et al., 2012; An et al., 2014).

Figure 5.

Schematic paleogeographic maps of Xigaze forearc basin (based on Einsele et al., 1994; Wang et al., 2012; Cai et al., 2012; An et al., 2014; DeCelles et al., 2014; Hu et al., 2015a, 2015b; Orme et al., 2015; G. Li et al., 2015b; Orme and Laskowski, 2016; Wang et al., 2017; Metcalf and Kapp, 2017; An et al., 2017). (A) Initial filling of basin began at 119 Ma with deposition of abyssal Chongdui Formation on an irregular basin floor and Sangzugang Formation on a carbonate shelf along southern flank of Gangdese magmatic arc. History of western basin is unknown because rocks of this age are not exposed. (B) Beginning at ca. 113 Ma, basin began to accumulate clastic Ngamring Formation via prograding turbidite fans derived mostly from southern Lhasa terrane. (C) Between 83 and 76 Ma, basin was nearly filled. Although Xigaze forearc rocks of this age are not exposed in eastern basin, provenance data from accretionary mélange suggest that this part of basin was overfilled at this time. (D) Between 76 and 66 Ma, a shallow-marine shelf environment dominated the forearc basin. As India approached from the south, turbidites of the Denggang Formation were shed northward toward the subduction trench. (E) After onset of continental collision at ca. 59 Ma, western Xigaze forearc basin was overfilled and shed sediment into early collisional foreland basin (Sangdanlin and Zheya Formations).

Figure 5.

Schematic paleogeographic maps of Xigaze forearc basin (based on Einsele et al., 1994; Wang et al., 2012; Cai et al., 2012; An et al., 2014; DeCelles et al., 2014; Hu et al., 2015a, 2015b; Orme et al., 2015; G. Li et al., 2015b; Orme and Laskowski, 2016; Wang et al., 2017; Metcalf and Kapp, 2017; An et al., 2017). (A) Initial filling of basin began at 119 Ma with deposition of abyssal Chongdui Formation on an irregular basin floor and Sangzugang Formation on a carbonate shelf along southern flank of Gangdese magmatic arc. History of western basin is unknown because rocks of this age are not exposed. (B) Beginning at ca. 113 Ma, basin began to accumulate clastic Ngamring Formation via prograding turbidite fans derived mostly from southern Lhasa terrane. (C) Between 83 and 76 Ma, basin was nearly filled. Although Xigaze forearc rocks of this age are not exposed in eastern basin, provenance data from accretionary mélange suggest that this part of basin was overfilled at this time. (D) Between 76 and 66 Ma, a shallow-marine shelf environment dominated the forearc basin. As India approached from the south, turbidites of the Denggang Formation were shed northward toward the subduction trench. (E) After onset of continental collision at ca. 59 Ma, western Xigaze forearc basin was overfilled and shed sediment into early collisional foreland basin (Sangdanlin and Zheya Formations).

The upper Chongdui Formation (113–106 Ma; Wang et al., 2012; Wang et al., 2017) or Lower Ngamring Formation, as defined by An et al. (2017), records the earliest influx of siliciclastic sediment into the basin via distal turbidite-fan systems and submarine debris flows (Fig. 5B; Wang et al., 2012; An et al. 2014; Orme and Laskowski, 2016; Wang et al., 2017). Similarly, the Ngamring Formation (106–83 Ma) is interpreted as turbidite-fan deposits that prograded south toward the growing accretionary prism (Einsele et al., 1994).

Whereas no outcrop evidence exists for deposition of the Qubeiya, Quxia, and Jialazi Formations in the eastern part of the forearc basin (An et al., 2014; Orme and Laskowski, 2016), modeling of thermochronologic data suggests significant additional burial by forearc sediment above the preserved section (Einsele et al., 1994; Orme, 2018). This would suggest that the eastern part of the basin continued receiving sediment until perhaps as recently as early Oligocene time (G. Li et al., 2017; Orme, 2018).

The Padana Formation was deposited in a shallower-marine environment than underlying Xigaze forearc strata (Fig. 5C). In the eastern part of the basin, the Padana Formation was deposited in an upward-shoaling succession of shelf, prodelta, delta-distributary, and delta-plain environments (An et al., 2014; Hu et al., 2015b; Orme et al., 2015). The uppermost Padana Formation contains paleosols and fluvial deposits (An et al., 2014; Orme et al., 2015).

The marked shoaling indicated by Padana Formation facies has been interpreted to represent gradual filling of the Xigaze forearc basin (An et al., 2014; Orme et al., 2015). Siliciclastic sediment with Eurasian provenance reached the subduction trench south of the western part of the Xigaze outcrop belt by ca. 60 Ma (DeCelles et al., 2014; Metcalf and Kapp, 2017), suggesting that the western part of the forearc basin was overfilled by that time. However, provenance analysis from sandstone blocks in sedimentary and ophiolitic mélange deposits in the eastern part of the Xigaze forearc basin suggests that sandy siliciclastic sediment was reaching the trench by as early as ca. 95 Ma (An et al., 2017) or ca. 71 Ma (Cai et al., 2012).

The Qubeiya Formation (ca. 78–66 Ma) was deposited in the western part of the forearc basin in a shallow-marine shelf setting that produced bioclastic limestone containing large benthic foraminifera and calcareous mudstone (Fig. 5D; Hu et al., 2015b; Orme et al., 2015). The change from a nearly filled subaerial basin during deposition of the upper Padana Formation to the shallow-marine environment of the Qubeiya Formation likely resulted from a global transgression beginning at 76 Ma (Miller et al., 2005). Deposition of the Qubeiya Formation also roughly corresponded to a lull in magmatic activity in the Gangdese arc at 75–69 Ma (e.g., Zhu et al., 2011; Ji et al., 2014; summarized in Orme et al., 2015), and it is possible that this had some impact on reducing clastic sediment input to the basin.

The contact between the Qubeiya Formation and the overlying Quxia Formation (ca. 66–57 Ma) has been reported as angular (Ding et al., 2005), paraconformable (Wang et al., 2012), and conformable (Orme et al., 2015). Depositional environments of the Quxia Formation varied along strike, including fluvial to fan-delta deposits north of Zhongba (Hu et al., 2016) and cohesive debris-flow deposits shed onto the Qubeiya carbonate platform or reef foreslope northwest of Saga (Orme et al., 2015). Although no direct dates of the Quxia Formation are available, it is plausible that it was deposited in response to increased exhumation associated with the onset of collision ca. 60 Ma.

The Jialazi Formation (57–55 Ma) is interpreted to represent deposition in a shallow-marine, distal fan-delta environment (Fig. 5E; Hu et al., 2015b). The Jialazi Formation temporally overlaps with the oldest foreland-basin strata exposed south of the accretionary complex in the Sangdanlin area (see below). Here, the Sangdanlin and Zheya Formations contain the oldest strata with Eurasian provenance deposited on the Indian continental margin (Ding et al., 2005; Wang et al., 2011; DeCelles et al., 2014; Hu et al., 2015a). Based on similar provenance indicators (Fig. 4), the Jialazi Formation has been interpreted as a source of Sangdanlin and Zheya sediment, which bypassed the overfilled Xigaze forearc basin and was deposited in the nascent peripheral foreland basin (DeCelles et al., 2014; Orme et al., 2015). In this sense, the Quxia and Jialazi Formations may represent deposition in the wedge-top part of the early collisional foreland basin, rather than a forearc-basin setting. Deposition of the Jialazi Formation corresponds with initiation of a magmatic high-flux period in the Gangdese arc that peaked at ca. 52 Ma (e.g., Zhu et al., 2011; Ji et al., 2014), and it is possible that this produced increased clastic flux into the basin and an overfilled basin state. However, increased volcanic input is not recorded by sandstone composition of the Jialazi, Sangdanlin, or Zheya Formations (Fig. 4).

Sandstone Petrography

Modal sandstone petrographic data from the Xigaze forearc basin reflect sediment provenance, tectonic setting, and source-area history (An et al., 2014; Orme et al., 2015; Hu et al., 2016; Orme and Laskowski, 2016). Sandstone from the Ngamring Formation is volcaniclastic, feldspatho-lithic and litho-feldspatho-quartzose (sensu Garzanti, 2016) and plots mostly within undissected-, transitional-, and dissected-arc fields of Dickinson (1985; see Fig. 4 herein). Sandstone of the Ngamring Formation is devoid of K-feldspar, and volcanic grains make up >70% of all lithic grains. Padana Formation and Ngamring Formation sandstone occupies similar provenance fields (Fig. 4), but Padana sandstone is less quartzose than Ngamring sandstone. Padana sandstone is quartzo-lithic and quartzo-litho-feldspathic and plots mostly within transitional-arc, dissected-arc, and recycled-orogen fields of Dickinson (1985). Padana sandstone contains up to 20% K-feldspar (100Fk/Qm + Fk + Fp [Fk—potassium feldspar; Qm—monocrystalline quartz; Fp—plagioclase feldspar). Sandstone from the Quxia Formation is feldspatho-lithic to litho-quartzose and occupies transitional- and dissected-arc provenance fields, with two samples plotting in the recycled-orogen field. Quxia sandstone has up to 56% K-feldspar (100Fk/Qm + Fk + Fp). Sandstone from the Jialazi Formation, which is the most quartzose within the Xigaze forearc basin, is quartzo-litho-feldspathic, quartzo-lithic, and litho-feldspatho-quartzose. Jialazi Formation sandstone plots mostly within dissected-arc and recycled-orogen provenance fields. Jialazi sandstone contains up to 40% K-feldspar (100Fk/Qm + Fk + Fp), with the least plagioclase feldspar within the Xigaze forearc basin.

Overall, compositions of Xigaze forearc sandstone define a trend from arc (undissected, transitional, and dissected) to recycled-orogen provenance fields. This suggests that parts of the Gangdese arc were increasingly dissected as the Xigaze forearc basin filled or that sediment was progressively eroded from more quartzose sources. Gangdese magmatism was active along the northern margin of this basin for nearly the entire life of the basin (DeCelles et al., 2011; Ji et al., 2014; Chapman and Kapp, 2017); the most quartzose sandstone (the Jialazi Formation) was deposited concurrently with the onset of a major magmatic flare-up in the Gangdese arc (e.g., Zhu et al., 2011; Ji et al., 2014). Thus, it seems probable that the shift from undissected-arc to dissected-arc and recycled-orogen provenance was due to expanding drainage systems across the Lhasa terrane that tapped into areas of incised volcanic and/or sedimentary rocks (Dürr, 1996; An et al., 2014), rather than deep incision of the southern arc. However, clear indication of northern sources is not evident in the westernmost part of the basin (Orme et al., 2015).

Paleoelevation

Most strata within the Xigaze forearc basin were deposited in submarine environments. Burial heating of Xigaze basin strata to temperatures between 140 °C and 200 °C (Orme, 2018; G. Li et al., 2017) precludes reconstruction of paleoelevations for subaerial components of the Xigaze forearc basin (Leier et al., 2009; Bera et al., 2010), but it is reasonable to assume that subaerial sediments deposited in delta-plain settings were near sea level (Orme et al., 2015).

NORTH HIMALAYAN FORELAND BASIN

Structural Framework

Outcrops of Paleocene–Lower Eocene strata in the Sangdanlin, Tingri, and Gamba areas constitute the best-documented remnants of the earliest foreland-basin deposits associated with the India-Eurasia collision (Ding et al., 2005; Zhu et al., 2005; Najman et al., 2010). The Sangdanlin outcrops lie within the suture zone proper, in a thrust sheet of the Tethyan Himalayan thrust belt directly south of accretionary mélange (Yang et al., 2003; Ding et al., 2005; DeCelles et al., 2014). Paleocene strata rest conformably upon Upper Cretaceous outer-shelf and slope deposits and, therefore, represent deposition on the northern edge of the Indian continental crust. Approximately 100 km southeast of the Sangdanlin section lies a second area of Paleocene–Eocene outcrop, referred to in the literature as the Zhepure Mountain, Tingri, or Qumiba section(s). Here, uppermost Cretaceous outer-shelf to slope deposits of the Tethyan sequence are overlain by Paleocene–Eocene shallow-marine to nonmarine deposits (Zhu et al., 2005; Najman et al., 2010; Zhang et al., 2012; Hu et al., 2012). In all of these sections, a change from detritus derived exclusively from the Indian continental landmass to detritus with unambiguous Gangdese-arc provenance can be demonstrated within well-dated Paleocene to Eocene strata (e.g., Zhu et al., 2005; Najman et al., 2010; DeCelles et al., 2014; Hu et al., 2015a; Ding et al., 2016).

Stratigraphy

The Sangdanlin section consists of the Denggang (or Zhongzuo), Sangdanlin, and Zheya Formations (Fig. 6). The ~104-m-thick Denggang Formation consists of massive thick-bedded coarse-grained quartzose sandstone beds. A sharp contact separates this unit from the overlying ~100-m-thick Sangdanlin Formation, which contains bedded radiolarite, siliceous mudstone, and turbiditic sandstone beds. Using biostratigraphy, Ding et al. (2005) inferred the presence of a significant unconformity at the base of the Sangdanlin Formation. Turbidite beds include both quartzose and lithic-arkose petrofacies. The Zheya Formation is a >500-m-thick, upward-coarsening succession of lithic-arkose sandstone beds intercalated with siliceous mudrock and local olistostromal beds (DeCelles et al., 2014).

Figure 6.

Composite stratigraphic section at Sangdanlin in Indus-Yarlung suture zone, after DeCelles et al. (2014). Vertical scale is in meters. Rightward slanting lines indicate prominent upward coarsening sequences.

Figure 6.

Composite stratigraphic section at Sangdanlin in Indus-Yarlung suture zone, after DeCelles et al. (2014). Vertical scale is in meters. Rightward slanting lines indicate prominent upward coarsening sequences.

In the Zhepure Mountain area, the critical stratigraphy lies within the Zhepure Shan Formation and the Pengqu Formation, which is divided into the Enba and Zhaguo members (referred to as the Youxia and Shenkezar formations, respectively, by Zhu et al., 2006), each ~100 m thick (Najman et al., 2010; Hu et al., 2012). The Zhepure Shan Formation is a shallow-marine limestone; the Pengqu Formation is dominated by mudstone and fine-grained sandstone of entirely marine (Wang et al., 2002) or marine and nonmarine (Zhu et al., 2005) origin.

Age

The Sangdanlin section has been dated paleontologically and geochronologically. Microfossils from the Denggang Formation indicate an early Maastrichtian age (Ding et al., 2005). The Denggang-Sangdanlin boundary is a disconformity that spans as much as 7 m.y.; the Sangdanlin Formation is middle Paleocene; and the Zheya Formation is Upper Paleocene–Eocene (Ding et al., 2005; DeCelles et al., 2014). A tuff in the Zheya Formation was dated at ca. 58.5 Ma (U-Pb zircon) by DeCelles et al. (2014); sandstone in the Sangdanlin Formation contains significant detrital zircon with U-Pb ages of 60–58.5 Ma (see also Wu et al., 2014; Hu et al., 2015a). Thus, the onset of deposition of sediment with a northerly provenance was between 60 and 58.5 Ma.

Foraminiferal and nannofossil biostratigraphy of the Zhepure Mountain sections indicates an age no younger than early Ypresian (ca. 54.5–52.5 Ma) for the Zhepure Shan Formation, and an age of ca. 52–50 Ma for the Pengqu Formation (Zhu et al., 2005; Najman et al., 2010; Hu et al., 2012). These ages are consistent with maximum ages of deposition determined from detrital-zircon U-Pb ages (Najman et al., 2010; Hu et al., 2012).

Characteristics of the Sangdanlin and Zhepure Shan (Tingri) sections are consistent with interpretations in the Zanskar and Pakistan Himalayas that indicate initial flexural forebulge uplift followed by foredeep subsidence during Paleocene time (ca. 60–56 Ma), as the northern Indian continental margin approached the trench along the southern flank of Asia (Garzanti et al., 1987; Beck et al., 1996; Ding et al., 2016).

Deposystems and Paleogeography

The stratigraphic record in the Sangdanlin section tracks the approach and eventual collision of the northern Indian margin with the arc-trench system that occupied the southern Asian margin during Late Cretaceous–Paleocene time (Figs. 5D and 5E; DeCelles et al., 2014; Hu et al., 2015a). The precollisional Denggang Formation contains a clear detrital-zircon signal of Indian continental provenance, but no evidence of sediment input from the Eurasian side of the system; it was deposited in quartzose, sand-rich turbidite lobes along the base of the Indian continental slope (Fig. 5D). The disconformity at the top of the Denggang Formation (Ding et al., 2005), and ensuing accumulation of pure radiolarian chert in the Sangdanlin Formation record deposition in deep water isolated from detrital input, perhaps on top of the flexural forebulge along the southern flank of the trench. The upper part of the Sangdanlin Formation and all of the Zheya Formation contain voluminous arkosic detritus derived from the Gangdese arc, as well as the overfilled Xigaze forearc basin (Orme et al., 2015), indicating that India and Eurasia were in physical contact by 60–58.5 Ma. What had been an oceanic trench was transformed into a deep-marine peripheral foredeep, filling with sand-rich turbidite fans (Fig. 5E); as India continued to subduct, this foredeep migrated southward into progressively shallower-water environments on the Indian upper slope and shelf (Zhu et al., 2005; Najman et al., 2010). Flexural modeling demonstrates that the north Himalayan foreland-basin record is compatible with the record of foreland-basin sedimentation preserved in the Lesser Himalaya of Nepal and northern India (DeCelles et al., 2014).

Sandstone Petrography

Sandstone from the Sangdanlin and Zhepure Shan sections can be divided into two distinct petrofacies: (1) quartzose and (2) lithic arkose (DeCelles et al., 2014). Quartzose petrofacies samples are >90% monocrystalline quartz and plot within the continental-interior provenance field of Dickinson (1985). These samples lack feldspar, and lithic grains are dominated by sedimentary and/or metamorphic grains (Fig. 4). The lithic-arkose petrofacies ranges from quartzo-litho-feldspathic to litho-quartzose and quartzo-lithic (sensu Garzanti, 2016) and plots mostly within the recycled-orogen, mixed-arc, and dissected-arc provenance fields of Dickinson (1985). These samples contain a mix of K-feldspar and plagioclase feldspars; lithic grains are mostly volcanic. Sandstone in the Denggang (Sangdanlin section) and Jidula (Zhepure Mountain section) Formations consists entirely of the quartzose petrofacies; the Sangdanlin Formation contains sandstone of both petrofacies; and the Zheya (Sangdanlin section), Pengqu, Youxia, and Shenkezar Formations (Zhepure Shan section) contain only lithic-arkose petrofacies (Zhu et al., 2005; Najman et al., 2010; Wang et al., 2011; DeCelles et al., 2014).

Lithic-arkose petrofacies within the north Himalayan foreland basin are similar in composition to sandstone within the Jialazi Formation, supporting the conclusion that sediment shed into the north Himalayan foreland basin bypassed a filled Xigaze forearc basin. Quartzose petrofacies were derived from the Indian craton, a continental-interior provenance (Dickinson, 1985).

Paleoelevation

All Cenozoic strata in the Sangdanlin and Zhepure Shan areas are of marine origin, except, perhaps, the Zhaguo Member of the Pengqu Formation at Zhepure Shan (Zhu et al., 2005) and the upper part of the Zheya Formation at Sangdanlin (DeCelles et al., 2014). Najman et al. (2010) noted that no marine deposits younger than ca. 50.5 Ma have been documented in the northern Himalaya (e.g., Nicora et al., 1987; Fuchs and Willems, 1990; Critelli and Garzanti, 1994; Green et al., 2008; Henderson et al., 2010). Farther south in the Himalayan foreland basin of Nepal and northern India, middle Eocene strata (ca. 45 Ma) contain marine facies (e.g., Sakai, 1983; DeCelles et al., 2004), and nonmarine conditions were established no later than ca. 20 Ma (Burbank et al., 1996; Ojha et al., 2009). The nonmarine foreland basin was at very low elevation, however, as suggested by evidence for marginal marine influence in strata as young as Pliocene in Bhutan (Coutand et al., 2016).

KAILAS BASIN

Structural Framework

The Kailas basin is the basin in which the Kailas Formation (Cheng and Xu, 1986; DeCelles et al., 2011) was deposited in the Indus-Yarlung suture zone. Several local names have been assigned to the Kailas Formation, including the Kailas Conglomerate (Heim and Gansser, 1939; Gansser, 1964), Qiuwu Formation (Wang et al., 2013), Gangdese Conglomerate (S. Li et al., 2017), Dazhuka Formation (Aitchison et al., 2009), and Luobusa Conglomerate (Aitchison et al., 2002b). Aitchison et al. (2002b) grouped these conglomerate units into the Gangrinboche Conglomerate. We use the term Kailas Formation because of its precedence (Gansser, 1964) and because the stratigraphic unit contains diverse lithofacies in addition to conglomerate. The Kailas Formation is exposed intermittently over a distance of ~2000 km along the length of the Indus-Yarlung suture zone. The preserved map-view width of the basin is everywhere <15 km.

The Kailas Formation rests in buttress unconformity on plutonic and volcanic rocks of the Gangdese magmatic arc along its northern margin and is truncated along its southern margin by faults of the north-verging Great Counter thrust system carrying Xigaze Group and suture-zone rocks in the hanging wall (Figs. 2 and 7). Deformation of the Kailas Formation is variable. In western Tibet, the rocks along the southern margin of the Kailas basin are folded into a footwall syncline by northward movement of the Great Counter thrust; however, in most western Tibet exposures, Kailas strata are nearly undeformed other than regional tilting (DeCelles et al., 2011, 2016). In the Lazi area, Kailas strata are truncated along their southern margin by an ~100-m-thick, south-dipping, top-to-the north ductile shear zone, in which granitic clasts of the Kailas Formation are deformed into sigma and delta clasts (Leary et al., 2016a; Laskowski et al., 2017). North of this shear zone, most strata are simply tilted, but small recumbent folds are also present near the southern edge of the exposure. North of Xigaze, Kailas strata are deformed by isoclinal folding and a dense network of small faults (Wang et al., 2015; Leary et al., 2016a). To date, all documented deformation in the Kailas Formation is contractional and postdepositional, with the exception of possible southward-expanding extensional growth strata pictured by Wang et al. (2015).

Figure 7.

Schematic cross section showing geological relationships among synformal Kailas Formation and Gangdese arc rocks (1), Great Counter thrust (GCT), and Xigaze forearc basin and ophiolitic rocks. Other numbered boxes indicate: (2) basal Kailas buttress unconformity; (3) lower conglomerate member; (4) middle shale member; (5) basin-center turbidite lobes; (6) upper sandstone/conglomerate member; and (7) coarse-grained basin-fringing facies, which include Gilbert-type delta deposits, on south side of basin. Buried normal fault beneath Great Counter thrust is hypothetical southern basin-bounding fault. Vertical scale of topography is ~2 km, and vertical exaggeration is ~2. Figure is after DeCelles et al. (2016).

Figure 7.

Schematic cross section showing geological relationships among synformal Kailas Formation and Gangdese arc rocks (1), Great Counter thrust (GCT), and Xigaze forearc basin and ophiolitic rocks. Other numbered boxes indicate: (2) basal Kailas buttress unconformity; (3) lower conglomerate member; (4) middle shale member; (5) basin-center turbidite lobes; (6) upper sandstone/conglomerate member; and (7) coarse-grained basin-fringing facies, which include Gilbert-type delta deposits, on south side of basin. Buried normal fault beneath Great Counter thrust is hypothetical southern basin-bounding fault. Vertical scale of topography is ~2 km, and vertical exaggeration is ~2. Figure is after DeCelles et al. (2016).

Stratigraphy

The Kailas Formation is thickest in western exposures near Mount Kailas, where more than 2.5 km of clastic strata are preserved (DeCelles et al., 2011). Gansser (1964) suggested that the Kailas Formation on and around Mount Kailas is more than 4 km thick. In the central sector of the Indus-Yarlung suture zone near Xigaze, most Kailas sections preserve ~1.5 km of strata, although a thickness of as much as 2.5 km has been reported (e.g., Wang et al., 2013; Leary et al., 2016a). In the eastern Indus-Yarlung suture zone, the exposed Kailas Formation is reported to be no more than 250 m thick (Kong et al., 2015).

Facies and stratigraphy of the Kailas Formation vary over its ~2000 km strike length; however, several first-order characteristics are consistent along the entire exposure (Fig. 7). In every documented exposure, the basal Kailas Formation along the northern limit of its outcrops onlaps an erosional buttress unconformity above Gangdese magmatic-arc rocks (either granite or intermediate volcanic rocks; Gansser, 1964; Searle et al., 1990; Davis et al., 2004; DeCelles et al., 2011, 2016; Leary et al., 2016a). The strata directly above this unconformity are ubiquitously very coarse-grained, arkosic conglomerate derived from the underlying igneous rocks (Aitchison et al., 2002b; Davis et al., 2004; DeCelles et al., 2011, 2016; Leary et al., 2016a).

In western Tibet, the basal Kailas Formation is extremely coarse-grained, with boulders >2 m in long dimension (DeCelles et al., 2011, 2016). Up section, grain size generally diminishes and transitions into predominantly sandstone beds. Near Mount Kailas, these coarse-grained facies are capped by an abrupt lacustrine transgressive surface, which is overlain by several hundred meters of interbedded sandstone and shale (Fig. 7). The sandstone beds are arranged in upward-coarsening and upward-thickening parasequences, and the shale is organic-rich and laminated (DeCelles et al., 2011, 2016). In the eastern Kailas Range, the middle black shale contains abundant coarse-grained sublacustrine conglomerate and sandstone beds, as well as Gilbert-type delta clinoforms (DeCelles et al., 2016). The Kailas Formation is capped by ~100–200 m of interbedded conglomerate, sandstone, and red beds with calcic paleosols. The conglomerate beds in the upper part of the formation contain abundant chert and other sedimentary clasts; paleocurrent indicators document northward sediment transport (Davis et al., 2004; DeCelles et al., 2011, 2016).

In the central sector of the Indus-Yarlung suture zone, the lithofacies of the Kailas Formation follow similar trends to those in western Tibet. The Kailas Formation rests unconformably on Gangdese granite, and the lower ~700 m of strata are poorly organized, matrix- and clast-supported, boulder to pebble conglomerate (Fig. 8; Wang et al., 2013; Leary et al., 2016a; S. Li et al., 2017). Fine-grained lithofacies have been reported in the lower Qiuwu Formation (basal Kailas Formation) near Xigaze (S. Li et al., 2017), but they are absent from Kailas exposures north of Lazi. The middle part of the Kailas Formation near Lazi consists of ~200 m of sandstone and pebble conglomerate interbedded with shale and coal. In thick, probably more stratigraphically complete exposures, Kailas lithofacies typically are pebble and cobble conglomerate in the upper part of the deposit. Where preserved, this upper interval is dominated by lenticular conglomerate interbedded with red, well-developed calcic paleosols (Leary et al., 2016a).

Figure 8.

Composite stratigraphic section of Kailas Formation exposed in central and eastern Indus-Yarlung suture zone, after Leary et al. (2016a). Vertical scale is in meters.

Figure 8.

Composite stratigraphic section of Kailas Formation exposed in central and eastern Indus-Yarlung suture zone, after Leary et al. (2016a). Vertical scale is in meters.

Age

The age of the Kailas Formation is well constrained because it contains abundant volcanic tuff, ash-flow, and basaltic flow deposits (Aitchison et al., 2009; DeCelles et al., 2011). Dates from various sections along the ~2000 km strike length of the Kailas Formation indicate that the onset of Kailas basin subsidence was diachronous. In western Tibet, the Kailas basin began accumulating sediment ca. 26 Ma (DeCelles et al., 2011; Carrapa et al., 2014), and sedimentation continued until ca. 22 Ma (S. Li et al., 2017). In the central Indus-Yarlung suture zone (~90°E), Kailas deposition began at ca. 23 Ma and may have continued as late as 15 Ma (Leary et al., 2016a; S. Li et al., 2017). In the easternmost exposure of the Kailas Formation (~92°E), a tuff yielded a U-Pb zircon age of ca. 19 Ma (Kong et al., 2015). Although granitic rocks beneath this section yielded an age of ca. 30 Ma (Kong et al., 2015), the age of earliest Kailas sedimentation is not directly constrained in this region.

Kailas Deposystems and Paleogeography

In western Tibet, the Kailas Formation was deposited by coarse-grained fluvial and alluvial-fan systems that drained into deep, probably meromictic lakes (DeCelles et al., 2011, 2016; G. Li et al., 2017). The lower coarse-grained part of the formation is dominated by debris-flow and coarse-grained fluvial deposits. Above the lacustrine transgressive surface, progradational shoreface parasequences demonstrate deep (~100 m) lacustrine environments fringed by fluvio-deltaic and swamp environments. In the axial offshore part of the basin, sandy Bouma-style turbidite and coarse-grained, sublacustrine, debris- and grain-flow deposits accumulated (Fig. 7). Floral and faunal assemblages, lake-margin coal deposits, and organic geochemical paleotemperature proxies suggest deposition in warm (>25 °C) freshwater lakes (DeCelles et al., 2016). The lacustrine interval coarsens upward into renewed fluvial and alluvial deposits that cap the Kailas Formation. The vertical facies stacking in this region is characteristic of extensional basins (e.g., Fig. 7). Alluvial-fan conglomerate, overlain by fluvial and lacustrine facies, followed by fluvial sandstone and conglomerate constitute this “lacustrine tripartite succession” or “lacustrine sandwich” (e.g., Schlische, 1992). This facies succession typically develops in rift basins, as early extension produces steep-sided rift flanks, from which coarse detritus is shed into the basin; as subsidence outpaces sedimentation, deep lakes form (Leeder, 1995). Finally, as extension slows, sedimentation begins to outpace subsidence, and the basin is filled by coarse-grained facies.

In the central Indus-Yarlung suture zone, near Lazi and Xigaze, the Kailas Formation is interpreted to represent fluvial and alluvial transport from basin margins into a large axial fluvial system occupying a narrow extensional basin (Fig. 9; Leary et al., 2016a; S. Li et al., 2017).

Figure 9.

Schematic paleogeographic maps of Kailas and Liuqu basins (based on Robinson et al., 2014; Lang and Huntington, 2014; Bracciali et al., 2015; Leary et al., 2016a, 2016b; Orme and Laskowski, 2016). (A) Kailas basin extensional deformation began around 24 Ma in central/eastern suture zone. Because basin opened from west to east, a westward-flowing drainage is inferred. (B) By ca. 22 Ma, extensional deformation ceased, and shortening resumed with movement along Great Counter thrust (GCT). Paleocurrent and provenance data suggest that an eastward-flowing axial drainage was established around this time. (C) As shortening continued, Liuqu Conglomerate was deposited in a wedge-top setting above the Great Counter thrust system. Only a small portion of this deposystem was preserved via structural accommodation in a triangle zone (Leary et al., 2016b).

Figure 9.

Schematic paleogeographic maps of Kailas and Liuqu basins (based on Robinson et al., 2014; Lang and Huntington, 2014; Bracciali et al., 2015; Leary et al., 2016a, 2016b; Orme and Laskowski, 2016). (A) Kailas basin extensional deformation began around 24 Ma in central/eastern suture zone. Because basin opened from west to east, a westward-flowing drainage is inferred. (B) By ca. 22 Ma, extensional deformation ceased, and shortening resumed with movement along Great Counter thrust (GCT). Paleocurrent and provenance data suggest that an eastward-flowing axial drainage was established around this time. (C) As shortening continued, Liuqu Conglomerate was deposited in a wedge-top setting above the Great Counter thrust system. Only a small portion of this deposystem was preserved via structural accommodation in a triangle zone (Leary et al., 2016b).

Kailas Formation paleocurrent data come mostly from pebble and cobble imbrications (DeCelles et al., 2011; Wang et al., 2013; Leary et al., 2016a; S. Li et al., 2017) and turbidite sole marks (DeCelles et al., 2016), with relatively few measurements from trough cross-bedded sandstone (Leary et al., 2016a; S. Li et al., 2017). Lower Kailas strata show predominantly south-directed sediment transport, although sections north of Lazi show dominantly northward paleoflow directions (Leary et al., 2016a; S. Li et al., 2017). In the middle and upper members, paleocurrent indicators show large variation, with overall northeastward trend. Paleocurrent data from north of Xigaze show dominantly northwestward (Wang et al., 2013) or southeastward (Leary et al., 2016a) transport in the upper section, but extreme deformation of these strata makes the accuracy of the restored orientations questionable.

Paleogeographic evolution of the Kailas basin began with alluvial-fan and lacustrine deposition in the western sector (DeCelles et al., 2016), with likely westward axial fluvial drainage in the central and eastern parts of the basin (Fig. 9A; Leary et al., 2016a). The regional drainage direction reversed later during Kailas basin development (Figs. 9B and 9C). The likely existence of an east-flowing paleo–Yarlung River by the early Miocene has been inferred from the presence of suture-zone detritus in the eastern Indian foreland basin (Lang and Huntington, 2014; Bracciali et al., 2015) and in the Indo-Burman ranges (Robinson et al., 2014). However, fluvial deposits at a scale comparable to the modern Yarlung River deposystem have not been documented in the Kailas Formation.

A key aspect of Kailas basin paleogeography is its narrowness. The presence of coarse-grained proximal lithofacies in the lower conglomerate interval, in depositional contact upon, and derived almost exclusively from, underlying Gangdese arc rocks, indicates that the present northern outcrop limit of the Kailas Formation is not far removed from the original location of the northern basin margin. Similarly, coarse-grained Gilbert-type delta and sublacustrine sediment-gravity-flow deposits along the southern flank of preserved basin strata suggest proximity to a topographically steep southern basin margin. Restoration of postdepositional shortening of the basin fill would add no more than ~5 km to the north-south basin width. We conclude that the modern outcrop belt of the Kailas Formation is a nearly complete representation of the original north-south extent of the basin.

Sandstone Petrography

The modal petrographic composition of Kailas Formation sandstone is highly variable (Fig. 4). Nearly all composition and provenance fields of standard quartz-feldspar-lithics (QFL) plots are represented, although quartzo-litho-feldspathic, litho-feldspathic, and quartzo-feldspatho-lithic sandstone is most abundant; these compositions correspond primarily to transitional- and dissected-arc provenance fields of Dickinson (1985). Lithic grains in Kailas sandstone are mostly volcanic, but a few samples show abundant metamorphic lithic grains. At the basin scale, no clear correlation exists between sandstone composition and stratigraphic position. Zhang et al. (2017) interpreted a provenance shift from transitional- and dissected-arc fields in the lower part of the Kailas Formation ~100 km west of Xigaze to dissected-arc and recycled-orogen provenance fields in the upper section; however, this pattern is not evident when combined with data from other sections (Fig. 4).

Up-section provenance changes are clearly resolvable through conglomerate clast counts (DeCelles et al., 2011; Leary et al., 2016a; S. Li et al., 2017; Zhang et al., 2017), with granitic clasts dominating the lower part of the formation, transitioning upward to dominantly volcanic clasts, and including chert, limestone, and quartzite in the upper Kailas Formation. This trend is interpreted to represent erosion of Gangdese granite, on which the Kailas Formation was deposited, transport from more distal volcanic sources within the Gangdese arc, and input from Tethyan sedimentary sources to the south, respectively (DeCelles et al., 2011; Leary et al., 2016a; S. Li et al., 2017; Zhang et al., 2017).

Paleoelevation

Various records suggest that the Kailas basin formed mainly in a low- to moderate-elevation setting. Biomarker analysis of Kailas lacustrine organic-shale samples from western Tibet suggests paleolake temperatures of >25 °C, and fossils indicate an ecosystem capable of supporting large fish and other vertebrates (DeCelles et al., 2016). Paleoenthalpy analysis of leaf fossils collected in the central Indus-Yarlung suture zone suggests that the Kailas basin formed at an elevation of 1.5–3 km (Ding et al., 2017). Very negative δ18O values from the uppermost Kailas Formation in the Kailas Range suggest high elevations in the basin during its final stages of filling (DeCelles et al., 2011). However, much of the Kailas Formation was heated after deposition to temperatures exceeding the closure temperature of the apatite fission-track system (Carrapa et al., 2014), so Kailas paleoelevation estimates based on stable-isotope data alone should be interpreted with caution.

LIUQU BASIN

Structural Framework

The record of the Liuqu basin consists of the Liuqu Conglomerate, which is preserved in tectonic slivers along ~150 km of the central Indus-Yarlung suture zone (Fig. 2). At its northern extent, the unit is in thrust contact structurally beneath ophiolitic rocks, although rare depositional contacts with ophiolitic and radiolarian chert are also present (Davis et al., 2002; Leary et al., 2016b). To the south, the Liuqu Conglomerate is bounded by north-vergent splays of the Great Counter thrust that place Tethyan and ophiolitic rocks structurally above the Liuqu Conglomerate. Contractional growth strata are observed at the Liuqu Conglomerate type locality as well as ~20 km to the west (Leary et al., 2016b).

Stratigraphy

At its type locality, the Liuqu Conglomerate is ~2000 m thick (Fig. 10); other incomplete exposures range from several hundred meters to ~1000 m in thickness. The Liuqu Conglomerate is dominated by clast-supported pebble to cobble conglomerate interbedded with mature, red paleosols (Davis et al., 2002; Leary et al., 2016b). Conglomeratic units commonly have erosive bases and are poorly organized and clast-supported. These intervals are broadly lenticular to tabular and extend laterally for several hundred meters. The upper part of the type section contains thick, matrix-supported boulder conglomerate, which is absent from most other exposures (Leary et al., 2016b). Very little up-section variation in texture and lithofacies is present within most Liuqu Conglomerate exposures (Fig. 10).

Figure 10.

Composite stratigraphic section of Liuqu Conglomerate, after Leary et al. (2016b). See Figure 8 for legend. Vertical scale is in meters.

Figure 10.

Composite stratigraphic section of Liuqu Conglomerate, after Leary et al. (2016b). See Figure 8 for legend. Vertical scale is in meters.

Clay-rich paleosols separate most Liuqu conglomeratic intervals. Individual paleosols are up to ~1 m thick, and some intervals contain multiple stacked paleosols. Liuqu paleosols are typically red, completely lack bedding, and are made up of slickensided, blocky peds. Many paleosols contain red and green mottling. Calcic nodules are rare but present in several sections. Based on clay content and slickensided peds, these paleosols are classified as Paleo-Vertisols (Mack et al., 1993; Leary et al., 2017).

Age

The age of the Liuqu Conglomerate has been much debated, and estimates include Cretaceous–Eocene (Tao, 1988; Fang et al., 2006; Ding et al., 2017), Oligocene (Wei et al., 2011), and early Miocene (G. Li et al., 2015a; Leary et al., 2016b). Although more work is necessary to constrain the age of the Liuqu Conglomerate, when all available data are considered, ca. 20 Ma emerges as the most viable age (Leary et al., 2016b, 2018; Xu et al., 2018).

Deposystems and Paleogeography

The Liuqu Conglomerate is interpreted as high-energy, flash-flood–dominated alluvial-fan deposits. Channelized conglomeratic intervals are interpreted as backfilled medial-fan channels; tabular conglomerate is interpreted as more distal, sheetflood-dominated parts of alluvial fans (e.g., Allen, 1981; DeCelles et al., 1991; Ridgway and DeCelles, 1993; Kumar et al., 2007). Paleosols are interpreted to represent weathering on inactive interfluve regions.

Imbricated clasts indicate that Liuqu sediment was transported generally northwestward, with local variations (Leary et al., 2016b). Clast composition data consistent with this interpretation indicate that most coarse-grained Liuqu sediment was derived from ophiolitic and mélange rocks to the south, with the possibility of some Tethyan source rocks. However, detrital-zircon U-Pb ages and εHf values require a northerly, Gangdese-arc source (Wang et al., 2010; Leary et al., 2016b). Prior to India-Eurasia collision, sand and finer-grained sediment were likely transported across an overfilled Xigaze forearc basin (Orme et al., 2015), deposited within the subduction trench (Figs. 5D and 5E), and incorporated into the accretionary wedge (Cai et al., 2012; An et al., 2017). As these rocks were eroded during Liuqu basin development, sediment derived originally from the Gangdese arc would have been recycled into the Liuqu Conglomerate.

Only a small N-S extent of the Liuqu deposystem is preserved, and the maximum N-S width of exposure is ~5 km. The Liuqu Conglomerate is interpreted to represent the proximal, remnant portion of a larger south-to-north sedimentary transport and depositional system (Fig. 9). The distal part of this system is interpreted to have connected to the east-flowing paleo–Yarlung River (Robinson et al., 2014; Bracciali et al., 2015). Contractional growth strata and facies analysis indicate that the Liuqu Conglomerate was deposited in a wedge-top setting on top of the actively shortening Himalayan-Tibetan orogenic wedge; the currently exposed Liuqu Conglomerate likely accumulated as the result of sediment accommodation adjacent to thrust sheets and in triangle zones (Leary et al., 2016b).

Sandstone Petrography

Composition of sandstone within the Liuqu Conglomerate ranges between litho-quartzose and lithic (sensu Garzanti, 2016); samples plot mostly in the recycled-orogen field of Dickinson (1985). Feldspar is nearly absent from sandstone in the Liuqu Conglomerate, and abundant lithic grains comprise primarily sedimentary, volcanic, and low-grade metasedimentary types (Fig. 4). Composition does not correlate to stratigraphic or geographic position (Leary et al., 2016b).

Sandstone petrography is consistent with conglomerate clast counts and paleocurrent indicators suggesting that Liuqu sandstone was derived from ophiolitic and mélange source rocks, as well as possibly low-grade Tethyan Himalayan metasedimentary rocks to the south (Leary et al., 2016b). The abundance of lithic grains suggests that the Xigaze ophiolite was deeply incised during Liuqu deposition (e.g., Garzanti, 2016).

Paleoelevation

Well-developed paleosols in the Liuqu Conglomerate provide some constraint on the paleoenvironment, and thus paleoelevation, of the Liuqu basin. Average δ13C values of −9.4‰ relative to Vienna Peedee belemnite (VPDB) from Liuqu paleosol carbonate indicate that the Liuqu basin was densely vegetated by C3 plants (Quade and Cerling, 1995; Quade et al., 1995; Leary et al., 2017), and major-element profiles indicate that chemical weathering in the Liuqu basin was comparable to that in the semitropical, pre–11 Ma Himalayan foreland basin; weathering indices indicate that Liuqu soils received rainfall between 780 ± 108 mm/yr and 1575 ± 108 mm/yr (Leary et al., 2017). Paleoenthalpy calculations based on leaf fossils in the Liuqu Conglomerate suggest that paleoelevation of the Liuqu basin was ~0.9 ± 0.9 km (Ding et al., 2017). These results suggest a warm, low-elevation setting for the Liuqu Conglomerate, similar to that of the Kailas Formation (DeCelles et al., 2011, 2016).

BASINS RELATED TO ARC-PARALLEL EXTENSION

Structural Framework

At least six north-south– to northeast-southwest–striking grabens and half grabens of late Miocene to Holocene age cut across the regionally arcuate structural grain of the Indus-Yarlung suture and northern Himalayan thrust belt (e.g., Fort et al., 1981; Armijo et al., 1986; Cogan et al., 1998; Colchen, 1999; Kapp and Guynn, 2004). These grabens include the Pulan, Thakkhola, Gyirong, Yadong, and Cona basins (Armijo et al., 1986; Cogan et al., 1998; Colchen, 1999; Murphy et al., 2002; Garzione et al., 2003; Murphy, 2007; Xu et al., 2012; De Sarkar et al., 2013; Shen et al., 2016). In addition, the Zhada basin is a southeast-northwest–elongate tectonic depression that is isolated from direct fault-controlled basin margins (Murphy et al., 2002; Saylor et al., 2009, 2010). Although the extensional structural setting of most of these basins is well established, only the Zhada, Pulan, Thakkhola, and Gyirong basins have been studied from a sedimentological standpoint.

Zhada basin, by far the largest of the arc-parallel extensional basins in the Indus-Yarlung suture region, is located at elevations of 3500–4500 m between detachment faults that bound the Leo Pargil and Gurla Mandhata structural domes in the northern (Tethyan) part of the Himalayan thrust belt (Fig. 1; Murphy et al., 2002; Saylor et al., 2010). The basin formed by large-magnitude (>60 km), northwest-southeast extension along the Qusum and Gurla Mandhata detachment faults that bound these two domes, and it occupies a releasing bend in the greater Karakoram–Gurla Mandhata–West Nepal strike-slip system (Ratschbacher et al., 1994; Murphy et al., 2002, 2014; Murphy and Copeland, 2005). Zhada basin measures ~150 km by 60 km and, uniquely among the extensional basins of the High Himalaya, is elongated parallel to the strike of the thrust belt. The basin fill is spectacularly exposed in gorges along the headwaters of the Sutlej River (Saylor et al., 2009, 2010).

The Pulan basin is an ~42-km-long, arcuate, north-south–striking half graben lying in the proximal hanging wall of the top-to-west Gurla Mandhata detachment fault (Fig. 1); accordingly, it could be considered the eastern extremity of Zhada basin. Pulan graben is spatially isolated from the bulk of the Zhada basin, however, and seems to have been paleogeographically independent; Murphy et al. (2002) interpreted Pulan graben as a supradetachment basin that formed in the hanging wall of the Gurla Mandhata detachment. The basin rests between elevations of 3800 and 4400 m. Steeply dipping brittle normal faults are present along both eastern and western basin margins, but the eastern margin is also marked by the lower-angle Gurla Mandhata detachment (Murphy et al., 2002). Internal unconformities and local-scale structures attest to the synkinematic character of the Pulan basin fill (Fig. 11; Murphy et al., 2002).

Figure 11.

Representative stratigraphic sections of four arc-parallel extensional basins in High Himalaya. See Figure 1 for basin locations. Units tg1–tg3 in Pulan basin section are defined in Murphy et al. (2002).

Figure 11.

Representative stratigraphic sections of four arc-parallel extensional basins in High Himalaya. See Figure 1 for basin locations. Units tg1–tg3 in Pulan basin section are defined in Murphy et al. (2002).

The next significant extensional basin along strike to the east of the Pulan half graben is the Thakkhola basin, which is perhaps the High Himalayan basin that most classically fits a continental-rift model (e.g., Leeder, 1995). The basin is nearly 75 km long, with an aspect ratio of ~5:1, strikes NNE-SSW, is bounded on its western flank by the high-angle, top-to-the-east Dangarzong normal fault, and rests unconformably upon strongly deformed rocks of the northern Himalayan thrust belt (Colchen, 1999; Hurtado et al., 2001; Garzione et al., 2003). The basin fill is exposed between 3100 and 4200 m elevation and is deeply incised along the Kali Gandaki River.

East of Thakkhola basin lies the Gyirong basin, which is situated between 4100 and 4500 m elevation in the hanging wall of the eastern basin-bounding Gyirong normal fault (top-to-the-west; Xu et al., 2012; Shen et al., 2016). This basin is smaller (~20 km long) than Thakkhola basin but lies in a similar structural setting unconformably on top of deformed northern Himalayan thrust belt rocks just north of the South Tibetan detachment system.

The Pulan, Thakkhola, and Gyirong basins are bounded by normal faults that strike roughly perpendicular to the Himalayan arc (Fig. 1). Zhada basin exhibits a more complex structural setting, with its northern boundary influenced by the Karakoram dextral strike-slip fault, and its western boundary dominated by the Qusum detachment fault (Fig. 1; Saylor et al., 2010). All of these basins are filled by Miocene–Pleistocene strata that were deposited in angular and buttress unconformity upon intensely deformed rocks of the northern Himalayan thrust belt. Each of these basins contains headwaters of major trans-Himalayan rivers: Zhada is drained by the Sutlej River of northern India; Pulan is drained by the Maja River, a tributary of the Karnali River in western Nepal; Thakkhola is the headwaters of the Kali Gandaki in central Nepal; and Gyirong includes the headwaters of the Trisuli River, also in central Nepal. Of the major east-west extensional basins in the Indus-Yarlung suture zone, only the Yadong-Gulu and Cona rifts cut northward across the suture zone. Thakkhola graben penetrates farthest south, well into the northern part of the Greater Himalayan zone.

Stratigraphy

The east-west extensional basins of the Indus-Yarlung suture contain documented thicknesses of 600–1040 m of exclusively nonmarine strata (Fig. 11). Although each basin is stratigraphically unique, alluvial-fan, coarse-grained fluvial, and lacustrine facies dominate. Where documented (in Zhada, Thakkhola, and Gyirong basins), basal stratigraphic contacts of the basin fills invariably are angular and/or buttress unconformities upon previously deformed strata of the northern Himalayan thrust belt. Internal angular unconformities and small-scale syndepositional faults (typically normal faults) are common (e.g., Garzione et al., 2003).

Zhada basin is filled with the >800-m-thick Zhada Formation, a succession of sandstone, siltstone, marl, and conglomerate (Kempf et al., 2009; Saylor et al., 2009). The overall stratigraphic pattern is that of a coarse-fine-coarse sandwich, with siltstone and marl concentrated in the middle ~300 m section (Fig. 11).

Although a detailed sedimentological study of Pulan basin has not been published, Murphy et al. (2002) included a reconnaissance investigation in their broader tectonic synthesis of the Gurla Mandhata dome. These authors reported a >500-m-thick, upward-coarsening succession of sandstone and conglomerate. Lower strata are composed of sandstone, siltstone, and minor conglomerate beds; the upper ~100 m are dominated by boulder conglomerate (Fig. 11).

Thakkhola basin contains the ~240-m-thick Tetang Formation and overlying ~800-m-thick Thakkhola Formation, separated by a low-angle unconformity. The Tetang Formation consists of coarse-grained conglomerate, sandstone, and limestone; the Thakkhola Formation is a mixture of conglomerate, sandstone, and marly siltstone (Garzione et al., 2003).

Gyirong basin includes three mapped formations, the Danzengzhukang (~240 m), Woma (~400 m), and Gongba (32 m) Formations (Xu et al., 2012). The Danzengzhukang Formation is dominated by conglomerate; the Woma Formation is a mixture of sandstone, conglomerate, and fine-grained siliciclastic rocks that contain abundant ostracodes, charophytes, palynomorphs, and shell fragments; and the Gongba Formation forms a coarse-grained conglomeratic cap (Xu et al., 2012).

Age

Strata in Zhada basin have been dated between ca. 9.2 and younger than 1 Ma through a combination of paleomagnetic stratigraphy and paleontology (Hipparion fauna; Saylor et al., 2009, 2010). Although no direct geochronological information is available from the Pulan basin, its synkinematic attributes coupled with thermochronological data from the adjacent Gurla Mandhata core complex suggest deposition during the interval 8.7–6.8 Ma (Murphy et al., 2002).

Both the Thakkhola and Gyirong basins have been dated by paleomagnetic stratigraphy; the Gyirong basin also contains Hipparion fauna, with chronostratigraphic significance. In the Thakkhola basin, the Tetang Formation is estimated to be 11–9.6 Ma, based on paleomagnetic stratigraphy, and the Thakkhola Formation is inferred to be younger than ca. 7 Ma, based on carbon-isotope data that indicate predominance of C4 plants (Garzione et al., 2000, 2003), which generally did not become abundant in southern Eurasia until after that time (Quade et al., 1989). In the Gyirong basin, the Danzengzhukang Formation is dated by magnetostratigraphy at younger than 10.8 Ma to ca. 7.2 Ma, and the Woma Formation is dated at ca. 7.2 to older than 1.7 Ma (interpretation of data from Wang et al. [1996] and Yue et al. [2004] inXu et al., 2012). The Hipparion fauna (Huang and Ji, 1979) occurs in the lower part of the Woma Formation, indicating that these strata must be ca. 7 Ma (Xu et al., 2012). The conglomeratic Gongba Formation is considered to be Pleistocene. Thus, the northern Himalayan basins related to arc-parallel extension generally date to as early as ca. 11 Ma, and some may be considered to have been active through the Pleistocene. Data from other indicators, usually less definitive than the ages of oldest strata (e.g., thermochronology from rocks in the footwall of basin-bounding normal faults), push the onset of extension back to as old as 14–13.5 Ma (Coleman and Hodges, 1995; Liu et al., 2007; McCallister et al., 2014; Shen et al., 2016).

Deposystems and Paleogeography

Zhada basin was filled with diverse nonmarine deposits, including a basal interval of eolian dune deposits, overlain by fluvial, lacustrine, and alluvial-fan facies (Saylor et al., 2010). A through-going large river system flowed northwest and exited the basin through the region now occupied by the Leo Pargil dome. As slip on the Qusum detachment facilitated growth of the Leo Pargil structure, this river (possibly a paleo-Indus tributary) was defeated, and the basin ponded. Initial lacustrine transgression commenced ca. 6 Ma (Saylor et al., 2010, 2016). Fan deltas were eventually replaced with a large lake system, which filled and was replaced by alluvial and fluvial systems (Kempf et al., 2009; Saylor et al., 2010).

Lithofacies descriptions from Murphy et al. (2002) suggest that Pulan basin was filled with fluvial followed by alluvial-fan deposits. The lower, fluvial-dominated section contains siltstone interbeds that could also represent ephemeral lacustrine environments. Although paleocurrent data from Pulan basin are sparse, they generally define a southward-flowing (with both eastward and westward components) paleodrainage pattern, similar to the modern geomorphological pattern of southward axial fluvial drainage fed by marginal alluvial fans (Murphy et al., 2002). Conglomerate clast-counts document sediment supply from both Tethyan sedimentary strata and uplifted metamorphic rocks in the Gurla Mandhata core complex, including mylonitized granitoid clasts that demonstrate syndepositional exhumation of the ductile shear zone along the western flank of the dome (Murphy et al., 2002).

In Thakkhola basin, the Tetang Formation contains lacustrine carbonate and coarse-grained conglomerate suggestive of deposition in marginal alluvial fans and local lacustrine depressions. Garzione et al. (2003) noted that the Tetang Formation contains no direct evidence linking it to extension along the western, basin-bounding Dangarzong fault and suggested that it could be more closely related to the South Tibetan detachment system. A roughly 2 m.y. unconformity separates the Tetang Formation from the overlying Thakkhola Formation; the latter contains clear evidence for deposition during slip on the Dangarzong fault (Garzione et al., 2003). Paleocurrent and lithofacies data indicate that a large gravelly braided trunk river flowed south out of the basin by ca. 7 Ma. Episodes of lacustrine ponding, however, suggest ongoing structural or geomorphological damming of the southern basin outlet.

Gyirong basin experienced early alluvial-fan sedimentation (Danzengzhukang Formation) derived from the western hanging wall, followed by mixed lacustrine and lacustrine fan-delta deposition of sediment derived mainly from the eastern footwall side of the basin (Woma Formation), and finally a return to alluvial-fan deposition derived from the eastern basin margin (Gongba Formation; Xu et al., 2012). A roughly 180° reversal in paleocurrent directions is documented at the boundary between the Danzengzhukang and Woma Formations, from generally eastward to generally southwestward (Xu et al., 2012). Palynological data demonstrate a transition through time from coniferous and broad-leaved forests under warm, humid climate to coniferous forests under colder and drier climate at higher elevations (Xu et al., 2012; Shen et al., 2016).

Sandstone Petrography

Zhada basin is the only arc-parallel extensional basin in which sandstone petrographic analysis has been undertaken (Saylor et al., 2010). Sandstone is typically lithic rich and feldspar poor, ranging from quartzo-lithic to feldspatho-quartzo-lithic and litho-quartzose, with a few feldspathic samples (Fig. 4). Zhada basin samples plot in the recycled-orogen, transitional-arc, and basement-uplift provenance fields of Dickinson (1985). Samples are increasingly quartzose and show an increase in sedimentary lithics and polycrystalline quartz at the expense of volcanic lithics up section.

Sandstone within the Zhada basin is interpreted to reflect mixing of sediment from two primary source areas: the Gangdese arc to the north, and the Tethyan Himalayan and Greater Himalayan belts to the south. Sandstones plotting in transitional-arc and lithic recycled-orogen provenance fields at the base of the section reflect northeasterly Gangdese-arc provenance. As the basin filled due to structural damming to the west, quartzose sand was shed northward from Tethyan Himalayan and Greater Himalayan thrust belts to the south (Saylor et al., 2010).

Paleoelevation

Since the pioneering paleoaltimetry study by Garzione et al. (2000) in Thakkhola basin, arc-parallel extensional basins in the High Himalaya have been primary targets for stable-isotope–based and paleobotanical paleoaltimetry studies because these basins contain well-preserved archives of paleosol and lacustrine carbonate and, locally, plant fossils that are amenable to stable-isotopic and plant-physiognomic methods. Recent studies have taken advantage of more diffuse organic materials that can be analyzed for their organic geochemistry and stable isotopes. In Thakkhola, δ18O values of paleosol and lacustrine carbonates are very negative and suggest that the basin formed at very high elevations (Garzione et al., 2000). Zhada basin fossils and lacustrine carbonates also yield very negative δ18O values, suggesting paleoelevations even higher than modern elevation in this region (Saylor et al., 2009). Clumped-isotope thermometry also indicates high paleoelevation and low temperatures, consistent with up to ~1 km of elevation loss in Zhada basin since Pliocene time (Huntington et al., 2015). The apparent loss of elevation in the Zhada region, which lies in an anomalously low part of the modern Himalaya, is attributed to crustal thinning during paired extension along the Qusum and Gurla Mandhata detachment faults (Saylor et al., 2009).

Studies of Gyirong basin have produced seemingly conflicting stable-isotope and paleobotanical data sets with respect to paleoaltimetry. Leaf fossils and palynomorphs indicate the presence of broad-leaved deciduous trees and plants, consistent with a warm, humid paleoclimate before ca. 10 Ma, which deteriorated somewhat after that time (Xu et al., 2012). Between ca. 7.2 and 3.2 Ma, flora assemblages from mixed broad-leaved and needle-leaved trees indicate both warm humid climate and a source of plant materials derived from higher elevations. On the other hand, δ18O values from lacustrine carbonates throughout the stratigraphic record are between −14‰ and −21‰ (Wang et al., 1996), which suggest very high paleoelevation, more in line with other studies of late Miocene–Pliocene paleoelevation in southern Tibet (e.g., Garzione et al., 2000; Saylor et al., 2009; Quade et al., 2011; Huntington et al., 2015).

SUMMARY AND DISCUSSION

Basins and Regional Stresses

The sedimentary basins considered in this paper contain kinematic and environmental information valuable for reconstructing the general state of stress in the suture zone during Cenozoic time (e.g., Sundell et al., 2013). Kinematic information helps to assess approximate orientations of principal stresses, which, in turn, can be understood in terms of surface elevation and crustal thickness, or potential dynamic stresses associated with movements of the upper mantle. Figure 12 illustrates development of sedimentary basins in the Indus-Yarlung suture and northern Himalaya from ca. 60 Ma to the present in a palinspastic regional context. The modern tectonic framework shown in panel D is progressively restored in panels A–C for shortening in the Himalayan thrust belt and the Xigaze forearc (sensu lato). From Paleocene time onward, the internal part of the Himalayan orogenic wedge and the suture zone experienced localized and transient basin development against a background of dramatically varying principal stress directions. With the understanding that complete characterization of principal stresses is impossible for ancient settings, we propose generalized upper-crustal approximations of principal stress directions based on basin types and orientations.

Figure 12.

Synoptic paleogeological maps illustrating development of sedimentary basins in Indus-Yarlung suture zone and across High Himalaya since ca. 60 Ma. Shortening estimates used to progressively restore maps are from DeCelles et al. (2002), van Hinsbergen et al. (2011a), Long et al. (2012), and Bhattacharyya et al. (2015). In all frames, stippled areas represent active depocenters, gray areas represent old remnant depocenters, white areas indicate regions of uplift and erosion, arrows denote general sediment transport pathways, and principal stresses are highly generalized from orientations of major Cenozoic basins. (A) Late Paleocene–early Eocene time (60–50 Ma), after restoration of 400 km of shortening in Xigaze forearc, accretionary prism, and Tethyan Himalayan thrust belt. Dashed line in southern part of all frames labeled MFT is position of modern Main Frontal thrust, which does not move in restoration. XFB—Xigaze forearc basin. Cross-hachured areas are Xigaze ophiolite. (B) Ca. 26–19 Ma, after 200 km of restoration on Main Central thrust, Kailas basin (KB) is active in hanging wall of a north-dipping normal fault and receiving sediment mainly from incised Gangdese arc (plus signs), while foreland basin on south side of Himalaya is receiving sediment from thrust belt. (C) Ca. 19–16 Ma, after restoration of 250 km of shortening in Lesser Himalaya, Liuqu basin is actively receiving sediment from north and south sides, and early Miocene foreland basin on south side of Himalaya is active. (D) Modern setting, which is representative of basin development since ca. 10 Ma. Lesser Himalayan duplex and Main Boundary thrust are active, as well as arc-parallel extensional basins. Sources: Fuchs (1979), DeCelles et al. (2001), Murphy et al. (2002), Murphy and Copeland (2005), Kapp and Guynn (2004), Yin (2006), and Saylor et al. (2010).

Figure 12.

Synoptic paleogeological maps illustrating development of sedimentary basins in Indus-Yarlung suture zone and across High Himalaya since ca. 60 Ma. Shortening estimates used to progressively restore maps are from DeCelles et al. (2002), van Hinsbergen et al. (2011a), Long et al. (2012), and Bhattacharyya et al. (2015). In all frames, stippled areas represent active depocenters, gray areas represent old remnant depocenters, white areas indicate regions of uplift and erosion, arrows denote general sediment transport pathways, and principal stresses are highly generalized from orientations of major Cenozoic basins. (A) Late Paleocene–early Eocene time (60–50 Ma), after restoration of 400 km of shortening in Xigaze forearc, accretionary prism, and Tethyan Himalayan thrust belt. Dashed line in southern part of all frames labeled MFT is position of modern Main Frontal thrust, which does not move in restoration. XFB—Xigaze forearc basin. Cross-hachured areas are Xigaze ophiolite. (B) Ca. 26–19 Ma, after 200 km of restoration on Main Central thrust, Kailas basin (KB) is active in hanging wall of a north-dipping normal fault and receiving sediment mainly from incised Gangdese arc (plus signs), while foreland basin on south side of Himalaya is receiving sediment from thrust belt. (C) Ca. 19–16 Ma, after restoration of 250 km of shortening in Lesser Himalaya, Liuqu basin is actively receiving sediment from north and south sides, and early Miocene foreland basin on south side of Himalaya is active. (D) Modern setting, which is representative of basin development since ca. 10 Ma. Lesser Himalayan duplex and Main Boundary thrust are active, as well as arc-parallel extensional basins. Sources: Fuchs (1979), DeCelles et al. (2001), Murphy et al. (2002), Murphy and Copeland (2005), Kapp and Guynn (2004), Yin (2006), and Saylor et al. (2010).

During Paleocene–early Eocene time (ca. 60–50 Ma), the Xigaze forearc basin coexisted with the earliest north Himalayan foreland basin; the two depocenters were separated by the accretionary complex, but they were probably connected via gaps in the trench-slope break or overspilling from the forearc into the trench/foredeep (Fig. 12A; Orme et al., 2015). From Paleocene through Oligocene time, the foreland-basin system migrated generally southward relative to India; the oldest foreland-basin deposits in northern Pakistan, northern India, and Nepal are middle Eocene (DeCelles et al., 1998, 2004, 2014; Najman et al., 2005; Jain et al., 2009; Ravikant et al., 2011; Zhuang et al., 2015). From Paleocene to late Oligocene, the northern part of the orogenic wedge grew by horizontal shortening in response to roughly north-south, horizontal maximum compressive principal stress (σ1) associated with India’s northward movement.

By Oligocene–Miocene time (Fig. 12B), a significant topographic barrier existed in the Tethyan Himalaya. Upper-plate north-south extension in the Kailas basin, and perhaps also the South Tibetan detachment system, may have resulted from Indian slab roll-back relative to the overriding Eurasian plate, followed by break-off (DeCelles et al., 2011, 2016; Leary et al., 2016a; Webb et al., 2017). The warm, humid paleoenvironments of the Kailas basin suggest relatively low paleoelevation. At the same time, the Main Central thrust was active along the southern flank of the Himalayan thrust belt (Hodges, 2000; Kohn, 2014), and sediment was shed from the thrust belt into the early Miocene foreland basin system (Fig. 12B). Along the Indus-Yarlung suture and in the northernmost part of the thrust belt, σ1 in the upper crust must have been regionally vertical to drive extension, and least compressive principal stress (σ3) must have been roughly N-S and horizontal. If Kailas basin was indeed at low elevation, the local verticality of σ1 cannot have been related to excess gravitational potential energy in a thick crust. Instead, this state of stress could have resulted from decreased coupling between the Indian and Eurasian plates in response to Indian slab roll-back and “trench” retreat relative to the upper plate (e.g., Burg et al., 1996; Butler and Beaumont, 2017). Farther south in the actively shortening Himalayan thrust belt, σ1 in the upper crust was approximately N-S and horizontal at regional scale.

After Indian slab break-off, which began at ca. 25 Ma in western Tibet and migrated eastward (DeCelles et al., 2011; Leary et al., 2016a; Webb et al., 2017), the shallower part of the Indian plate was able to once again drive northward at a low angle beneath Tibet, and the upper crust of the orogenic system became dominated by horizontal N-S σ1 and approximately vertical σ3. The Liuqu basin formed in this setting (Fig. 12C), trapped in a north-south tectonic vise in the interior part of the orogenic belt. Sedimentological and geochemical indicators of warm humid climate suggest that the Liuqu basin was at low to moderate elevation (see paleoaltimetry interpretation of Ding et al., 2017; with the understanding that their age model differs from that presented here and in Leary et al., 2016b, 2017), consistent with a roughly horizontal σ1.

By middle to late Miocene time, continued crustal shortening in the Himalaya and northward underthrusting of India beneath Tibet produced crust thick enough to isostatically support elevations high enough to drive σ1 into a near-vertical orientation (Fig. 12D; Styron et al., 2015). Kapp and Guynn (2004) suggested that 4500 m is the critical elevation for east-west extension, based on the present distribution of normal faults and topography of southern Tibet. This value is in agreement with available paleoaltimetry data sets in the arc-parallel extensional basins (Garzione et al., 2000; Saylor et al., 2009; Ding et al., 2017). Consistently high paleoaltimetry in these basins from the onset of basin accumulation indicates that east-west extension did not predate attainment of critically high elevation. Unlike the early Miocene episode of vertical σ1, the late Miocene extension was characterized by roughly east-west σ3, resulting in the onset of arc-parallel extension (Fig. 12D). The roughly east-west extension in southern Tibet can be explained as upper-crustal thinning in response to ongoing insertion of Indian lithosphere beneath Tibet and arc-normal collisional boundary stresses along the central part of the Himalayan thrust belt (Kapp and Guynn, 2004; Sundell et al., 2013; Styron et al., 2015). In the frontal thrust belt, where elevation was insufficient to keep σ1 vertical in the upper crust, horizontal shortening in the upper crust prevailed.

Sandstone Provenance

Sandstone composition has long been used as a proxy for provenance and overall tectonic setting of sedimentary basins (e.g., Suttner, 1974; Dickinson and Suczek, 1979; Ingersoll and Suczek, 1979; Schwab, 1981; Ingersoll, 1983, 1990; Dickinson, 1985; Garzanti and Vezzoli, 2003; Garzanti, 2016). In the classic model of Dickinson and Suczek (1979), the Xigaze forearc basin would be classified as having magmatic-arc provenance, and the north Himalayan foreland and Liuqu basins would be assigned recycled-orogen provenance; both interpretations are correct in terms of tectonic settings. However, neither of the extensional basins (Zhada and Kailas) exhibits compositions predicted by the provenance model of Dickinson and Suczek (1979) and Dickinson (1985); instead, Zhada sandstone overlaps completely with the recycled-orogen provenance field, and Kailas sandstone reflects magmatic-arc provenance (Fig. 4). These basins would be broadly classified as successor basins (Ingersoll, 2012) because they were overprinted on, and thus inherited petrographic compositions of, previously established tectonic edifices. As noted by Garzanti (2016), this type of overprinting and compositional smearing are common in rift basins, which may exhibit virtually any modal sand(stone) composition, depending on the lithological composition of the terrane being extended and the depth of erosion. As Ingersoll (1990) outlined, rift basins are dominated by first-order sampling of local source terranes, whereas Dickinson’s (1985) composition model is based on third-order (continental-scale) fluvial and submarine-fan systems. In essence, Zhada and Kailas sandstone compositions are faithful to their source terranes, but the source terranes do not follow the conventional rules linking lithological composition to tectonic setting (Dickinson, 1985). In contrast, sandstone of the regional-scale Xigaze forearc and North Himalayan foreland basins, as well as modern sand from the Brahmaputra River, Bengal/Nicobar and Indus submarine fans, and Miocene sandstone from the Indo-Gangetic foreland basin plot within the expected recycled-orogen provenance field (Ingersoll and Suczek, 1979; Suczek and Ingersoll, 1985; Critelli and Ingersoll, 1994; DeCelles et al., 1998; Najman and Garzanti, 2000; Garzanti et al., 2010).

Larger Questions

Several larger questions pertaining to sedimentary basin development in the Indus-Yarlung suture and northern Himalaya remain:

  • (1) What controlled the drastically changing state of stress in this region? The state of stress in the northern Himalaya and Indus-Yarlung suture cannot be explained simply as a function of plate kinematics because changes in convergence velocity (e.g., Copley et al., 2010; van Hinsbergen et al., 2011b; Zahirovic et al., 2012) generally do not coincide with changing basin characteristics. Instead, the complex history of basin development more likely reflects variations in a delicate near-balance among principal stresses, themselves controlled by crustal thickness, surface elevation, boundary stresses, and possibly stresses associated with behavior of Indian lithosphere and the mantle wedge beneath southern Tibet and the Himalaya (Copley et al., 2011; Styron et al., 2015; Butler and Beaumont, 2017).

  • Strong correlation exists between surface deformation style and elevation across the Himalayan-Tibetan orogen, with normal and thrust faulting dominating at higher (>4500 m) and lower (<3500 m) elevations, respectively (Molnar and Tapponnier, 1978; Mercier et al., 1987). This suggests that the basins discussed in this paper operate as structural “paleoaltimeters,” insofar as lateral variations in gravitational potential energy play a major role in driving extension at higher elevations and outward growth of the Tibetan Plateau (Dewey, 1988; Molnar and Lyon-Caen, 1988). The flatness of Tibet at long wavelengths (Fielding et al., 1994) leaves the impression that a critical surface elevation can be supported by collisional stresses. According to this logic, estimates for the timing of rift initiation have been used as proxies for the timing of attainment of near-modern elevations (Harrison et al., 1992). Employing a similar line of reasoning, the timing of cessation of thrust faulting in a given region may indicate when significant elevation was attained. These structural geologic “paleoaltimeters” are qualitative, however, given temporal-spatial variations in boundary stress conditions and rheology during orogenesis. Rifts in Tibet and the High Himalaya initiated as early as mid-Miocene and show younger episodes of accelerated extension, raising the possibility that the gravitational potential energy (and elevation) of the region may still be increasing (Kapp and Guynn, 2004; Styron et al., 2015). Conversely, midcrustal channel-flow models (e.g., Beaumont et al., 2004; Royden et al., 2008) and some paleoaltimetric studies (Saylor et al., 2009; Huntington et al., 2015; Currie et al., 2016; Ding et al., 2017) suggest that surface elevations in Tibet may have been higher in the past and that rifting is an indicator of true orogenic collapse and net thinning of the crust.

  • (2) The apparent change from sporadically low to consistently high elevation in the Indus-Yarlung suture took place between the end of deposition in the Liuqu basin and the onset of east-west extension, roughly between 18 Ma and 11 Ma (e.g., Garzione et al., 2000; Gébelin et al., 2013; Carrapa et al., 2016; Ding et al., 2017). What caused this elevation change? Some have argued for ~1 km of regional elevation gain during middle to late Miocene time in response to gravitational foundering of thickened mantle lithosphere beneath Tibet (Molnar et al., 1993; Molnar, 2005). In our view, the documented paleoaltimetry record of Tibet remains too sparse to decide this question, and, in any case, 1 km of elevation gain is near the limit of resolution in currently used techniques (Quade et al., 2011). However, several observations can be made: The Gangdese magmatic arc was probably at high elevation during Paleocene time (Ding et al., 2014; Ingalls et al., 2017), and the Nima and Lunpola basins in central Tibet were probably near their modern elevations (~4500 m) by no later than late Oligocene time (Rowley and Currie, 2006; DeCelles et al., 2007a; Polissar et al., 2009; Huntington et al., 2015), concurrent with development of Kailas basin at low to modest elevation along the Indus-Yarlung suture (DeCelles et al., 2016; Ding et al., 2017). Paleoelevation of the Oligocene–Miocene Oiyug basin, which lies within the Gangdese magmatic arc just north of the Indus-Yarlung suture, was 4500–5400 m (Spicer et al., 2003; Khan et al., 2014; Currie et al., 2016; Ding et al., 2017; Ingalls et al., 2017). It appears that Tibet proper (north of the Indus-Yarlung suture) was already at stable high elevation by late Oligocene time, whereas the Himalaya and the Indus-Yarlung suture vacillated among various states of principal stresses (as discussed above) in response to changing Indian plate subduction modes (flat slab vs. rolling back vs. breaking off) and strength of coupling between the two plates (Copley et al., 2011; Butler and Beaumont, 2017). These changing states of stress produced profound changes at the topographic surface, both in terms of elevation and basin dynamics. The “stable” high elevation of central Tibet may be the end result of a similar series of erratic elevation changes within the Lhasa terrane (which was at or below sea level as recently as ca. 95 Ma; Leier et al., 2007) after its collision with Qiangtang.

  • (3) Finally, we consider the question of how so many different manifestations of sediment accumulation are preserved in the Indus-Yarlung suture and the Himalaya, a region where rapid deep erosion (e.g., Thiede and Ehlers, 2013) has exposed Cenozoic upper-amphibolite-grade to eclogite-grade metamorphic rocks and synorogenic plutonic rocks. Suture zones in general may be paradoxically prone to long-lasting sediment accumulation. For example, the Banggong-Nujiang suture in central Tibet, which formed during Late Jurassic–Early Cretaceous collision of the Lhasa and Qiangtang terranes, remains a topographic depression despite nearly 150 m.y. of orogenic activity in and around the suture zone (Taylor et al., 2003). Cenozoic successor basins are also widespread along the Banggong-Nujiang suture (Kapp et al., 2003; Rowley and Currie, 2006; DeCelles et al., 2007b; He et al., 2011). Suture zones are places where records of some of Earth’s largest and deepest sedimentary basins (e.g., trenches and forearc basins) are to be found, albeit extremely deformed (e.g., Lash, 1985). What is perhaps most remarkable are the arc-parallel extensional basins in the High Himalaya north of the South Tibetan detachment. The oldest preserved record of these basins is no older than ca. 11 Ma (and therefore is a minimum age for onset of east-west extension), and all of these basins are currently being incised by headward-eroding trans-Himalayan rivers. It seems unlikely that these basins will long be preserved if India continues its northward movement. Their preservation up to this time stems in part from their isolated, internal locations in the slowly eroding rain shadow of the High Himalaya (Gébelin et al., 2013; Carrapa et al., 2016). On the other hand, if the east-west extensional basins are linked to stretching of upper crust as India underthrusts Tibet (Styron et al., 2015), a younger generation of east-west extensional basins might form in the hinterland of the Himalayan thrust belt, as thrusting continues to propagate toward the foreland.

CONCLUSIONS

  • (1) Five sets of sedimentary basins have existed since the end of Cretaceous time in the Indus-Yarlung suture zone and what is now the northern Himalayan orogenic belt. Paleocene–Eocene basins consisted of the subcontinental scale (a) Xigaze forearc and (b) north Himalayan foreland basins, both of which formed in deep-marine to low-elevation nonmarine settings. No late Eocene to mid-Oligocene sedimentary basins have been documented in this region. (c) Beginning in late Oligocene time, the suture zone experienced rapid transient extension in a north-south direction, opening the low-elevation Kailas rift basin. (d) Soon thereafter (ca. 19–18 Ma), the region was again shortening in a north-south direction, and the Liuqu wedge-top basin formed to the south of the remnant accretionary complex along the south side of the old (by that time, inactive) forearc. (e) Since ca. 11 Ma (and possibly 14 Ma), the northern Himalaya and southern Tibet have been extending in an east-west sense, resulting in the formation of generally north-south–trending rift basins (Pulan, Thakkhola, Gyirong) and a large northwest-southeast–trending supradetachment basin (Zhada) in the High Himalaya.

  • (2) Marine sedimentation ended in this region no later than ca. 50 Ma; subsequent deposition in all sedimentary basins between the High Himalaya and the suture zone involved alluvial, fluvial, and lacustrine depositional systems.

  • (3) Dramatic changes in orientations of principal stresses and resulting basin types in this region reflect changing Indian subduction dynamics and surface elevation. Since ca. 60 Ma, this region has experienced north-south shortening, north-south extension, and east-west extension, all of which produced sedimentary basin records.

  • (4) Preservation of this diverse spectrum of basin types in a region of extreme uplift and erosion is due to the original geodynamic basin-forming mechanisms and subsequent surface uplift history that has orographically isolated the interior of the Himalayan thrust belt and Tibet, thus preserving the region from erosion. Recent (Pliocene–Pleistocene) breaching of the high Himalayan barrier by trans-Himalayan rivers places the preservation of these basins at risk.

ACKNOWLEDGMENTS

We are grateful to the editors of this volume for their efforts to assemble a body of work in honor of the late W.R. Dickinson, whose career achievements remain so inspirational to us. This research was funded by grants from the National Science Foundation Tectonics (EAR-1140068) and Continental Dynamics (EAR-1008527) Programs. Long-time collaborator Ding Lin helped with permits and logistics. We thank Barbara Carrapa, Devon Orme, Andrew Laskowski, Jay Quade, Joel Saylor, Carmala Garzione, Andrew Leier, and Brian Currie for informative discussions about Tibetan basins over many years. We thank Michael Taylor and Ken Ridgway for detailed, constructive reviews, and Ray Ingersoll for careful editorial work that helped us to improve the manuscript.

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Figures & Tables

Figure 1.

Simplified geological map of Indus-Yarlung suture zone and Himalayan thrust belt (modified from Yin, 2006), showing locations of discussed sedimentary basins (see legend for abbreviations).

Figure 1.

Simplified geological map of Indus-Yarlung suture zone and Himalayan thrust belt (modified from Yin, 2006), showing locations of discussed sedimentary basins (see legend for abbreviations).

Figure 2.

Geologic map of central Indus-Yarlung suture zone (based on Cai et al., 2011, 2012; Leary et al., 2016a). GCT—Great Counter thrust. See Figure 1 for broader geological context of this map area.

Figure 2.

Geologic map of central Indus-Yarlung suture zone (based on Cai et al., 2011, 2012; Leary et al., 2016a). GCT—Great Counter thrust. See Figure 1 for broader geological context of this map area.

Figure 3.

Composite stratigraphic section of Xigaze forearc basin strata (after Wang et al., 2012; An et al., 2014; Orme et al., 2015). Vertical scale is in meters.

Figure 3.

Composite stratigraphic section of Xigaze forearc basin strata (after Wang et al., 2012; An et al., 2014; Orme et al., 2015). Vertical scale is in meters.

Figure 4.

Sandstone modal petrographic data from the Xigaze, Sangdanlin, Kailas, Liuqu, and Zhada basins. All compositions were generated using the Gazzi-Dickinson point-counting method (Ingersoll et al., 1984). Recalculated data are available from the authors. Provenance fields are after Dickinson (1985): UA—undissected arc; TA—transitional arc; DA—dissected arc; LR—lithic recycled; TR—transitional recycled; QR—quartzose recycled; CI—craton interior; TC—transitional continental; BU—basement uplift; M—mixed; RO—recycled orogen. Sources: Xigaze forearc basin: An et al. (2014), Orme et al. (2015), Hu et al. (2016), Orme and Laskowski (2016); Sangdanlin: Zhu et al. (2005), Najman et al. (2010), Wang et al. (2011), DeCelles et al. (2014); Kailas: DeCelles et al. (2011), Leary et al. (2016a), S. Li et al. (2017), Zhang et al. (2017); Liuqu: Leary et al. (2016b); Zhada: Saylor et al. (2010). Abbreviations of grain-types: Qm—monocrystalline quartz; F—total feldspar; Lt—total lithic grains; Qt—total quartzose grains including lithic quartzose grains; L—lithic grains exclusive of quartzose varieties; P—plagioclase; K—K-feldspar; Qp—polycrystalline quartz; Lv—volcanic lithic grains; Ls—sedimentary lithic grains.

Figure 4.

Sandstone modal petrographic data from the Xigaze, Sangdanlin, Kailas, Liuqu, and Zhada basins. All compositions were generated using the Gazzi-Dickinson point-counting method (Ingersoll et al., 1984). Recalculated data are available from the authors. Provenance fields are after Dickinson (1985): UA—undissected arc; TA—transitional arc; DA—dissected arc; LR—lithic recycled; TR—transitional recycled; QR—quartzose recycled; CI—craton interior; TC—transitional continental; BU—basement uplift; M—mixed; RO—recycled orogen. Sources: Xigaze forearc basin: An et al. (2014), Orme et al. (2015), Hu et al. (2016), Orme and Laskowski (2016); Sangdanlin: Zhu et al. (2005), Najman et al. (2010), Wang et al. (2011), DeCelles et al. (2014); Kailas: DeCelles et al. (2011), Leary et al. (2016a), S. Li et al. (2017), Zhang et al. (2017); Liuqu: Leary et al. (2016b); Zhada: Saylor et al. (2010). Abbreviations of grain-types: Qm—monocrystalline quartz; F—total feldspar; Lt—total lithic grains; Qt—total quartzose grains including lithic quartzose grains; L—lithic grains exclusive of quartzose varieties; P—plagioclase; K—K-feldspar; Qp—polycrystalline quartz; Lv—volcanic lithic grains; Ls—sedimentary lithic grains.

Figure 5.

Schematic paleogeographic maps of Xigaze forearc basin (based on Einsele et al., 1994; Wang et al., 2012; Cai et al., 2012; An et al., 2014; DeCelles et al., 2014; Hu et al., 2015a, 2015b; Orme et al., 2015; G. Li et al., 2015b; Orme and Laskowski, 2016; Wang et al., 2017; Metcalf and Kapp, 2017; An et al., 2017). (A) Initial filling of basin began at 119 Ma with deposition of abyssal Chongdui Formation on an irregular basin floor and Sangzugang Formation on a carbonate shelf along southern flank of Gangdese magmatic arc. History of western basin is unknown because rocks of this age are not exposed. (B) Beginning at ca. 113 Ma, basin began to accumulate clastic Ngamring Formation via prograding turbidite fans derived mostly from southern Lhasa terrane. (C) Between 83 and 76 Ma, basin was nearly filled. Although Xigaze forearc rocks of this age are not exposed in eastern basin, provenance data from accretionary mélange suggest that this part of basin was overfilled at this time. (D) Between 76 and 66 Ma, a shallow-marine shelf environment dominated the forearc basin. As India approached from the south, turbidites of the Denggang Formation were shed northward toward the subduction trench. (E) After onset of continental collision at ca. 59 Ma, western Xigaze forearc basin was overfilled and shed sediment into early collisional foreland basin (Sangdanlin and Zheya Formations).

Figure 5.

Schematic paleogeographic maps of Xigaze forearc basin (based on Einsele et al., 1994; Wang et al., 2012; Cai et al., 2012; An et al., 2014; DeCelles et al., 2014; Hu et al., 2015a, 2015b; Orme et al., 2015; G. Li et al., 2015b; Orme and Laskowski, 2016; Wang et al., 2017; Metcalf and Kapp, 2017; An et al., 2017). (A) Initial filling of basin began at 119 Ma with deposition of abyssal Chongdui Formation on an irregular basin floor and Sangzugang Formation on a carbonate shelf along southern flank of Gangdese magmatic arc. History of western basin is unknown because rocks of this age are not exposed. (B) Beginning at ca. 113 Ma, basin began to accumulate clastic Ngamring Formation via prograding turbidite fans derived mostly from southern Lhasa terrane. (C) Between 83 and 76 Ma, basin was nearly filled. Although Xigaze forearc rocks of this age are not exposed in eastern basin, provenance data from accretionary mélange suggest that this part of basin was overfilled at this time. (D) Between 76 and 66 Ma, a shallow-marine shelf environment dominated the forearc basin. As India approached from the south, turbidites of the Denggang Formation were shed northward toward the subduction trench. (E) After onset of continental collision at ca. 59 Ma, western Xigaze forearc basin was overfilled and shed sediment into early collisional foreland basin (Sangdanlin and Zheya Formations).

Figure 6.

Composite stratigraphic section at Sangdanlin in Indus-Yarlung suture zone, after DeCelles et al. (2014). Vertical scale is in meters. Rightward slanting lines indicate prominent upward coarsening sequences.

Figure 6.

Composite stratigraphic section at Sangdanlin in Indus-Yarlung suture zone, after DeCelles et al. (2014). Vertical scale is in meters. Rightward slanting lines indicate prominent upward coarsening sequences.

Figure 7.

Schematic cross section showing geological relationships among synformal Kailas Formation and Gangdese arc rocks (1), Great Counter thrust (GCT), and Xigaze forearc basin and ophiolitic rocks. Other numbered boxes indicate: (2) basal Kailas buttress unconformity; (3) lower conglomerate member; (4) middle shale member; (5) basin-center turbidite lobes; (6) upper sandstone/conglomerate member; and (7) coarse-grained basin-fringing facies, which include Gilbert-type delta deposits, on south side of basin. Buried normal fault beneath Great Counter thrust is hypothetical southern basin-bounding fault. Vertical scale of topography is ~2 km, and vertical exaggeration is ~2. Figure is after DeCelles et al. (2016).

Figure 7.

Schematic cross section showing geological relationships among synformal Kailas Formation and Gangdese arc rocks (1), Great Counter thrust (GCT), and Xigaze forearc basin and ophiolitic rocks. Other numbered boxes indicate: (2) basal Kailas buttress unconformity; (3) lower conglomerate member; (4) middle shale member; (5) basin-center turbidite lobes; (6) upper sandstone/conglomerate member; and (7) coarse-grained basin-fringing facies, which include Gilbert-type delta deposits, on south side of basin. Buried normal fault beneath Great Counter thrust is hypothetical southern basin-bounding fault. Vertical scale of topography is ~2 km, and vertical exaggeration is ~2. Figure is after DeCelles et al. (2016).

Figure 8.

Composite stratigraphic section of Kailas Formation exposed in central and eastern Indus-Yarlung suture zone, after Leary et al. (2016a). Vertical scale is in meters.

Figure 8.

Composite stratigraphic section of Kailas Formation exposed in central and eastern Indus-Yarlung suture zone, after Leary et al. (2016a). Vertical scale is in meters.

Figure 9.

Schematic paleogeographic maps of Kailas and Liuqu basins (based on Robinson et al., 2014; Lang and Huntington, 2014; Bracciali et al., 2015; Leary et al., 2016a, 2016b; Orme and Laskowski, 2016). (A) Kailas basin extensional deformation began around 24 Ma in central/eastern suture zone. Because basin opened from west to east, a westward-flowing drainage is inferred. (B) By ca. 22 Ma, extensional deformation ceased, and shortening resumed with movement along Great Counter thrust (GCT). Paleocurrent and provenance data suggest that an eastward-flowing axial drainage was established around this time. (C) As shortening continued, Liuqu Conglomerate was deposited in a wedge-top setting above the Great Counter thrust system. Only a small portion of this deposystem was preserved via structural accommodation in a triangle zone (Leary et al., 2016b).

Figure 9.

Schematic paleogeographic maps of Kailas and Liuqu basins (based on Robinson et al., 2014; Lang and Huntington, 2014; Bracciali et al., 2015; Leary et al., 2016a, 2016b; Orme and Laskowski, 2016). (A) Kailas basin extensional deformation began around 24 Ma in central/eastern suture zone. Because basin opened from west to east, a westward-flowing drainage is inferred. (B) By ca. 22 Ma, extensional deformation ceased, and shortening resumed with movement along Great Counter thrust (GCT). Paleocurrent and provenance data suggest that an eastward-flowing axial drainage was established around this time. (C) As shortening continued, Liuqu Conglomerate was deposited in a wedge-top setting above the Great Counter thrust system. Only a small portion of this deposystem was preserved via structural accommodation in a triangle zone (Leary et al., 2016b).

Figure 10.

Composite stratigraphic section of Liuqu Conglomerate, after Leary et al. (2016b). See Figure 8 for legend. Vertical scale is in meters.

Figure 10.

Composite stratigraphic section of Liuqu Conglomerate, after Leary et al. (2016b). See Figure 8 for legend. Vertical scale is in meters.

Figure 11.

Representative stratigraphic sections of four arc-parallel extensional basins in High Himalaya. See Figure 1 for basin locations. Units tg1–tg3 in Pulan basin section are defined in Murphy et al. (2002).

Figure 11.

Representative stratigraphic sections of four arc-parallel extensional basins in High Himalaya. See Figure 1 for basin locations. Units tg1–tg3 in Pulan basin section are defined in Murphy et al. (2002).

Figure 12.

Synoptic paleogeological maps illustrating development of sedimentary basins in Indus-Yarlung suture zone and across High Himalaya since ca. 60 Ma. Shortening estimates used to progressively restore maps are from DeCelles et al. (2002), van Hinsbergen et al. (2011a), Long et al. (2012), and Bhattacharyya et al. (2015). In all frames, stippled areas represent active depocenters, gray areas represent old remnant depocenters, white areas indicate regions of uplift and erosion, arrows denote general sediment transport pathways, and principal stresses are highly generalized from orientations of major Cenozoic basins. (A) Late Paleocene–early Eocene time (60–50 Ma), after restoration of 400 km of shortening in Xigaze forearc, accretionary prism, and Tethyan Himalayan thrust belt. Dashed line in southern part of all frames labeled MFT is position of modern Main Frontal thrust, which does not move in restoration. XFB—Xigaze forearc basin. Cross-hachured areas are Xigaze ophiolite. (B) Ca. 26–19 Ma, after 200 km of restoration on Main Central thrust, Kailas basin (KB) is active in hanging wall of a north-dipping normal fault and receiving sediment mainly from incised Gangdese arc (plus signs), while foreland basin on south side of Himalaya is receiving sediment from thrust belt. (C) Ca. 19–16 Ma, after restoration of 250 km of shortening in Lesser Himalaya, Liuqu basin is actively receiving sediment from north and south sides, and early Miocene foreland basin on south side of Himalaya is active. (D) Modern setting, which is representative of basin development since ca. 10 Ma. Lesser Himalayan duplex and Main Boundary thrust are active, as well as arc-parallel extensional basins. Sources: Fuchs (1979), DeCelles et al. (2001), Murphy et al. (2002), Murphy and Copeland (2005), Kapp and Guynn (2004), Yin (2006), and Saylor et al. (2010).

Figure 12.

Synoptic paleogeological maps illustrating development of sedimentary basins in Indus-Yarlung suture zone and across High Himalaya since ca. 60 Ma. Shortening estimates used to progressively restore maps are from DeCelles et al. (2002), van Hinsbergen et al. (2011a), Long et al. (2012), and Bhattacharyya et al. (2015). In all frames, stippled areas represent active depocenters, gray areas represent old remnant depocenters, white areas indicate regions of uplift and erosion, arrows denote general sediment transport pathways, and principal stresses are highly generalized from orientations of major Cenozoic basins. (A) Late Paleocene–early Eocene time (60–50 Ma), after restoration of 400 km of shortening in Xigaze forearc, accretionary prism, and Tethyan Himalayan thrust belt. Dashed line in southern part of all frames labeled MFT is position of modern Main Frontal thrust, which does not move in restoration. XFB—Xigaze forearc basin. Cross-hachured areas are Xigaze ophiolite. (B) Ca. 26–19 Ma, after 200 km of restoration on Main Central thrust, Kailas basin (KB) is active in hanging wall of a north-dipping normal fault and receiving sediment mainly from incised Gangdese arc (plus signs), while foreland basin on south side of Himalaya is receiving sediment from thrust belt. (C) Ca. 19–16 Ma, after restoration of 250 km of shortening in Lesser Himalaya, Liuqu basin is actively receiving sediment from north and south sides, and early Miocene foreland basin on south side of Himalaya is active. (D) Modern setting, which is representative of basin development since ca. 10 Ma. Lesser Himalayan duplex and Main Boundary thrust are active, as well as arc-parallel extensional basins. Sources: Fuchs (1979), DeCelles et al. (2001), Murphy et al. (2002), Murphy and Copeland (2005), Kapp and Guynn (2004), Yin (2006), and Saylor et al. (2010).

Contents

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