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Keywords:

  • collisional orogenic belts;
  • Tibetan Plateau;
  • Himalaya;
  • lithosphere dynamics;
  • delamination;
  • thrust belts

Abstract

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Models for Tibetan Plateau Uplift
  5. 3. Geology of the Himalayan Fold-Thrust Belt and Tibetan Plateau
  6. 4. Geophysical Constraints on Sub-Tibetan Lithospheric and Crustal Structure
  7. 5. Shortening in the Himalayan Fold-Thrust Belt
  8. 6. An Integrated Model for Uplift of the Tibetan Plateau
  9. 7. Implications of the Model
  10. 8. Conclusions
  11. Acknowledgments
  12. References

[1] Recent research in the Himalayan fold-thrust belt provides two new sets of observations that are crucial to understanding the evolution of the Himalayan-Tibetan orogenic system. First, U-Pb zircon ages and Sm-Nd isotopic studies demonstrate that (1) Greater Himalayan medium- to high-grade metasedimentary rocks are much younger than true Indian cratonic basement; and (2) these rocks were tectonically mobilized and consolidated with the northern margin of Gondwana during early Paleozoic orogenic activity. These observations require that Greater Himalayan rocks be treated as supracrustal material in restorations of the Himalayan fold-thrust belt, rather than as Indian cratonic basement. In turn, this implies the existence of Greater Himalayan lower crust that is not exposed anywhere in the fold-thrust belt. Second, a regional compilation of shortening estimates along the Himalayan arc from Pakistan to Sikkim reveals that (1) total minimum shortening in the fold-thrust belt is up to ∼670 km; (2) total shortening is greatest in western Nepal and northern India, near the apex of the Himalayan salient; and (3) the amount of Himalayan shortening is equal to the present width of the Tibetan Plateau measured in an arc-normal direction north of the Indus-Yalu suture zone. This information suggests that a slab of Greater Indian lower crust (composed of both Indian cratonic lower crust and Greater Himalayan lower crust) with a north-south length of ∼700 km may have been inserted beneath the Tibetan crust during the Cenozoic orogeny. We present a modified version of the crustal underthrusting model for Himalayan-Tibetan orogenesis that integrates surface geological data, recent results of mantle tomographic studies, and broadband seismic studies of the crust and upper mantle beneath the Tibetan Plateau. Previous studies have shown that incremental Mesozoic and early Cenozoic shortening had probably thickened Tibetan crust to ∼45–55 km before the onset of the main Cenozoic orogenic event. Thus, the insertion of a slab of Greater Indian lower crust with maximum thickness of ∼20 km (tapering northward) could explain the Cenozoic uplift of the Tibetan Plateau. The need for Tibetan crust to stretch laterally as the Greater Indian lower crust was inserted may explain the widespread east-west extension in the southern half of the Plateau. Our reconstruction of the Himalayan fold-thrust belt suggests that Indian cratonic lower crust, of presumed mafic composition and high strength, should extend approximately halfway across the Tibetan Plateau, to the Banggong suture. From there northward, we predict that the Tibetan Plateau is underlain by more felsic, and therefore weaker, lower crust of Greater Himalayan affinity. Two slab break-off events are predicted by the model: the first involved Neotethyan oceanic lithosphere that foundered ∼45–35 Ma, and the second consisted of Greater Indian lithosphere (most likely composed of Greater Himalayan material) that delaminated and foundered ∼20–10 Ma. Asthenospheric upwelling associated with the break-off events may explain patterns of Cenozoic volcanism on the Tibetan Plateau. Although the model predicts a northward migrating topographic front due solely to insertion of Greater Indian lower crust, the actual uplift history of the Plateau was complicated by early-middle Tertiary shortening of Tibetan crust.

1. Introduction

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Models for Tibetan Plateau Uplift
  5. 3. Geology of the Himalayan Fold-Thrust Belt and Tibetan Plateau
  6. 4. Geophysical Constraints on Sub-Tibetan Lithospheric and Crustal Structure
  7. 5. Shortening in the Himalayan Fold-Thrust Belt
  8. 6. An Integrated Model for Uplift of the Tibetan Plateau
  9. 7. Implications of the Model
  10. 8. Conclusions
  11. Acknowledgments
  12. References

[2] The origin of high-elevation orogenic plateaus is a topic of considerable interest, given the potential climatological, geochemical, and environmental side effects of plateau growth and maintenance. Explanations of orogenic plateaus have been best refined for the case of the Tibetan Plateau, which has been studied intensively since the mid-1970's [e.g., Molnar and Tapponnier, 1975]. The Himalayan fold-thrust belt, which forms the southern rim of the Plateau (Figure 1), is obviously a result of shortening of Indian rocks to the south of the suture zone that marks the paleosubduction zone between the Indian and Eurasian plates [Gansser, 1964]. However, no consensus exists on the timing and mechanism(s) of formation of the Tibetan Plateau and the nature of its relationship to the Himalaya [e.g., Dewey et al., 1988; Harrison et al., 1992; Molnar et al., 1993; Matte et al., 1997; Yin and Harrison, 2000; Tapponnier et al., 2001].

image

Figure 1. Regional tectonic map of the Tibetan Plateau, showing major suture zones, terranes, fault systems, Cenozoic magmatic centers, and the Gangdese magmatic arc [after Yin and Harrison, 2000; Hacker et al., 2000; Tapponnier et al., 2001]. Elevations after Fielding et al. [1994].

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[3] Our corporate understanding of the development of the Tibetan Plateau has developed to its current state largely unaided by a thorough consideration of the potential contributions of Greater Indian crust to thickening of the Plateau. Greater India is regarded as the landmass of the Indian subcontinent before the onset of the Indo-Eurasian collision [Veevers et al., 1975]. The north-south length of Greater Indian lower crust available to thicken the Tibetan Plateau should be equivalent to the amount of supracrustal shortening in the Himalayan fold-thrust belt [Klootwijk et al., 1985]. Moreover, the kinematic history of the fold-thrust belt should provide a gauge of the timing of addition of Greater Indian crustal material to the Tibetan Plateau. These simple concepts have been difficult to exploit because estimates of shortening based on geological data from the Himalayan fold-thrust belt are sparse and vary by a factor of roughly two. Instead, the principal constraints on amounts of Himalayan shortening have come from paleomagnetic data, which are inherently subject to large uncertainties and provide only the coarsest kinematic information [Molnar and Tapponnier, 1975; Achache et al., 1984; Patriat and Achache, 1984; Klootwijk et al., 1985; Besse et al., 1984; Besse and Courtillot, 1988, 1991; Patzelt et al., 1996]. In addition, the high-grade metamorphic rocks of the Greater Himalayan sequence that crop out in the medial part of the fold-thrust belt historically have presented a puzzle in terms of how to restore cross sections through the Himalaya: Should these rocks be treated as Indian cratonic basement [e.g., Gansser, 1964; Mattauer, 1986; Srivastava and Mitra, 1994; Hauck et al., 1998], or should they be treated as an exotic tectonostratigraphic terrane that was structurally elevated prior to the Cenozoic orogenic event [e.g., Parrish and Hodges, 1996; DeCelles et al., 2000]? Adding to the dilemma is the uncertainty of the structure in the mantle and lithosphere beneath the Tibetan Plateau. However, recent progress on the Greater Himalayan issue, as well as recent deep crustal seismic reflection profiling [Nelson et al., 1996; Hauck et al., 1998] and broadband seismic experiments across the Tibetan Plateau [Owens and Zandt, 1997; Kosarev et al., 1999; Kind et al., 2002], and mantle tomographic studies [Grand et al., 1997; Van der Voo et al., 1999] provide an opportunity to integrate geological and geophysical data sets at an unprecedented level.

[4] In this paper we summarize the available geological and geophysical data on the timing and magnitude of crustal shortening in the Himalayan fold-thrust belt and the mantle structure beneath the Tibetan Plateau, with the goal of quantitatively assessing potential contributions of Greater Indian lower crust to thickening of the Tibetan Plateau. Although our perspective is informed mainly by geological investigations of the Himalayan fold-thrust belt in Nepal, the work of many previous investigators throughout the Himalaya and Tibetan Plateau helps to regionalize some of the key observations. We propose an integrated model that connects thickening and uplift of the Tibetan Plateau with the development of the Himalayan fold-thrust belt and suggests timing for at least a major increment of Plateau uplift. What is different about our model is the way we treat Indian lower crust. Previous studies of the Himalayan fold-thrust belt have considered the high-grade metamorphic rocks of the Greater Himalaya to be Indian cratonic basement that has been thrust upward and southward, removing the need to dispose of more than ∼7.0 × 106 km3 of Indian lower crust (roughly one-tenth of the Plateau volume) beneath the Tibetan Plateau. Several recent geochronological and geochemical studies, however, show that Greater Himalayan rocks should be considered as supracrustal material, much younger than cratonic India, that was stripped from its lower crustal basement during the Cenozoic orogeny. This new way of looking at the Greater Himalayan rocks revives the need to dispose of a large amount (>1.4 × 107 km3) of lower continental crust and its associated mantle lithosphere beneath the Tibetan Plateau [Ni and Barazangi, 1984]. Exactly how much of this crust must be accounted for, and how much this may have contributed to growth of the Plateau, are the main subjects of this paper.

2. Models for Tibetan Plateau Uplift

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Models for Tibetan Plateau Uplift
  5. 3. Geology of the Himalayan Fold-Thrust Belt and Tibetan Plateau
  6. 4. Geophysical Constraints on Sub-Tibetan Lithospheric and Crustal Structure
  7. 5. Shortening in the Himalayan Fold-Thrust Belt
  8. 6. An Integrated Model for Uplift of the Tibetan Plateau
  9. 7. Implications of the Model
  10. 8. Conclusions
  11. Acknowledgments
  12. References

[5] Models for the origin of the Tibetan Plateau begin with the following boundary conditions: (1) approximately 2500 km of post-collision convergence has taken place between India and Eurasia [Molnar and Tapponnier, 1975; Patriat and Achache, 1984; Besse and Courtillot, 1988, 1991; Patzelt et al., 1996]; (2) the Plateau is underlain by continental crust that is approximately twice as thick as normal crust [Hirn, 1988; Owens and Zandt, 1997; Zhu, 1998; Kind et al., 2002]; and (3) seismic phase velocities in the upper mantle (at 100–300 km depth) are generally fast compared to adjacent regions, and Pn wave velocity and Sn propagation are relatively slow and inefficient, respectively, beneath the northern half of the plateau but relatively fast and efficient beneath the southern half [Ni and Barazangi, 1984; Molnar, 1988; Holt and Wallace, 1990; McNamara et al., 1997; Owens and Zandt, 1997; Tapponnier et al., 2001].

[6] Five general categories of models for the Tibetan Plateau have been proposed: (1) crustal thickening by pure shear during the Mesozoic and Cenozoic [Murphy et al., 1997] or entirely during the Cenozoic [England and Houseman, 1988; Molnar et al., 1993]; (2) crustal underthrusting, or injection, from the south [Argand, 1924; Powell and Conaghan, 1973; Ni and Barazangi, 1984; Coward and Butler, 1985; Mattauer, 1986; Zhao and Morgan, 1987], perhaps accompanied by phase transitions in the lower crust [LePichon et al., 1997; Chemenda et al., 2000]; (3) crustal underthrusting from the north [Willett and Beaumont, 1994; Tapponnier et al., 2001]; (4) uplift owing to dynamic processes in the mantle [Dewey et al., 1988; Molnar et al., 1993; Turner et al., 1993; Platt and England, 1994]; and (5) indentation of a weak Eurasian plate by a rigid, strong Indian plate, possibly accompanied by intraplate subduction and lateral tectonic escape of large Eurasian crustal blocks [Tapponnier et al., 1982; Matte et al., 1996] and local thrust faulting and sediment aggradation [Métivier et al., 1998; Tapponnier et al., 2001]. A number of studies have suggested that the lower crust of Tibet is extremely hot, rheologically weak, and decoupled from the mantle, which helps to explain the remarkably low-relief internal topography of much of the Tibetan Plateau (Figure 2) [Zhao and Morgan, 1987; Bird, 1991; Fielding et al., 1994; Royden et al., 1997; Kirby et al., 2000; Clark and Royden, 2000]. Some studies have proposed combinations of more than one of these mechanisms for thickening and uplift of the Plateau. Most of these models have difficulty accounting for all of the known features of the Tibetan Plateau and the Himalaya (see, for example, discussions by Matte et al. [1997] and Yin and Harrison [2000]).

image

Figure 2. Digital elevation model topography draped onto a cartoon of the crustal and upper mantle structure in the Himalayan-Tibetan region along a north-south swath between 88°E and 93°E. Note the difference between the topographic and subsurface vertical scales. Structure in upper mantle is modified after Jin et al. [1996], Owens and Zandt [1997], Kosarev et al. [1999], and Kind et al. [2002]. Arrows represent possible west-east flow to account for east-west anisotropy in mantle [McNamara et al., 1994]. Abbreviations as follows: MFT, Main Frontal thrust; MBT, Main Boundary thrust; MCT, Main Central thrust; ISZ, Indus-Yalu suture zone; BSZ, Banggong suture zone; QA, Qiangtang anticlinorium; JS, Jinsha suture; KF, Kunlun fault; SGT, Songpan-Ganzi terrane; QB, Qaidam basin; QS, Qilian Shan.

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[7] The crustal thickening by pure shear model implies that the Tibetan Plateau has thickened either incrementally since early Mesozoic time through the accretion of Eurasian crustal blocks or entirely during the Cenozoic collision between the Indian and Eurasian plates. Recent studies indicate that incremental shortening and magmatism may have played important roles in thickening the crust of the Plateau during Mesozoic time [Murphy et al., 1997; Kapp et al., 2000, 2002; Yin et al., 1999; Robinson et al., 2002; Tapponnier et al., 2001]. However, the presence of mid-Cretaceous marine rocks at many localities in the southern part of the Tibetan Plateau indicates that much of the region must have been at or below sea level until late Cretaceous or early Tertiary time, which implies that the crust was not thick. Models that invoke large amounts of pre-Cenozoic crustal thickening are not necessarily at odds with other mechanisms of Plateau development [e.g., Yin and Harrison, 2000]. The crustal underthrusting model, which in its most basic form was first presented by Argand [1924], proposes an increase of crustal thickness by tectonic insertion [Powell and Conaghan, 1973; Ni and Barazangi, 1984; Chemenda et al., 2000] or ductile injection [Zhao and Morgan, 1987] of Indian lower crust beneath the Tibetan crust, with or without Indian lithospheric mantle. Uplift of the Plateau has also been attributed to a phase transition from eclogite to granulite in an intermediate composition lower crust that has been underthrust from the south [LePichon et al., 1997]. The underthrusting model has been challenged because geophysical studies reveal that Indian lithosphere is probably not present under the northern part of the Tibetan Plateau [e.g., Molnar, 1988; McNamara et al., 1997; Owens and Zandt, 1997], and insertion of Indian lithosphere beneath the Plateau would have been impeded by the presence of Eurasian lithosphere. In addition, it is generally thought that no need exists to account for Indian crustal basement beneath the Plateau because it has been thrust (or “extruded”) southward in the form of the Greater Himalayan high-grade metamorphic rocks that constitute much of the higher part of the Himalayan fold-thrust belt [Gansser, 1964; Dewey et al., 1988; Hauck et al., 1998]. Models that require mantle dynamics include delamination-induced isostatic rebound in combination with crustal thickening by pure shear [Molnar et al., 1993] and those that inflate the crust by significant amounts of magmatic underplating [e.g., Powell and Conaghan, 1973; Molnar, 1988; Bird, 1991]. The rigid indentation model [Tapponnier et al., 1982] requires long-distance eastward tectonic escape of large crustal fragments accommodated by strike-slip faulting. Elements of this model have been supported by recent studies of the geology of the Tibetan Plateau and by neotectonic studies of fault slip rates. However, the actual amounts of lateral escape versus distributed shortening and rotation are still debated [e.g., Yin and Harrison, 2000]. It is quite plausible that elements of several of the existing models for the origin of the Tibetan Plateau may have conspired to produce today's Plateau.

3. Geology of the Himalayan Fold-Thrust Belt and Tibetan Plateau

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Models for Tibetan Plateau Uplift
  5. 3. Geology of the Himalayan Fold-Thrust Belt and Tibetan Plateau
  6. 4. Geophysical Constraints on Sub-Tibetan Lithospheric and Crustal Structure
  7. 5. Shortening in the Himalayan Fold-Thrust Belt
  8. 6. An Integrated Model for Uplift of the Tibetan Plateau
  9. 7. Implications of the Model
  10. 8. Conclusions
  11. Acknowledgments
  12. References

3.1. Himalayan Fold-Thrust Belt

[8] The Himalayan fold-thrust belt stretches for an arc-length distance of ∼2400 km between the Hazara (in the west) and Namche Barwa (in the east) syntaxes. Its northern boundary is defined by the Indus-Yalu suture zone and its southern boundary by the Main Frontal thrust system and largely undocumented blind thrusts beneath the northern part of the Indo-Gangetic foreland basin system (Figures 1 and 3). The Indus-Yalu suture zone contains Cretaceous forearc basin deposits and ophiolites that formed above the north dipping subduction zone between the Neotethyan ocean and the southern, Andean-style margin of Eurasia [Gansser, 1964; Searle, 1986; Searle et al., 1997]. South of the suture zone lies the Tibetan Himalayan zone, which comprises Cambrian-Eocene rocks in numerous south-vergent thrust sheets and related folds [Gansser, 1964; Burg and Chen, 1984; Ratschbacher et al., 1994; Searle et al., 1997]. The Tibetan Himalayan rocks are commonly referred to as the Tethyan sequence [Gaetani and Garzanti, 1991; Brookfield, 1993].

image

Figure 3. (a) Cross sectional restoration of the northern margin of Greater India prior to the Cenozoic Himalayan-Tibetan orogenic event. This architecture was disrupted by development of the Himalayan fold-thrust belt during Cenozoic time. Abbreviations as follows: LHT, Lesser Himalayan terrane; GHT, Greater Himalayan terrane; THT, Tibetan Himalayan terrane. (b) Generalized regional cross section of the Himalayan fold-thrust belt and Tibetan Plateau at the longitude of western Nepal-southwestern Tibet, after Yin and Harrison [2000], DeCelles et al. [2001], and Kapp et al. [2002]. Note that the hypothetical Greater Himalayan rocks that form the lower part of the Lhasa terrane in (b) are not structurally contiguous with rocks of the Greater Himalayan zone. Abbreviations as follows: MFT, Main Frontal thrust; LHZ, Lesser Himalayan zone; GHZ, Greater Himalayan zone; STDS, South Tibetan detachment system; THZ, Tibetan Himalayan zone; ISZ, Indus-Yalu suture zone; BSZ, Banggong suture zone; JSZ, Jinsha suture zone; AKMS, Anyimaqin-Kunlun-Muztagh suture; KF, Kunlun fault; QS, Qilian Shan thrust belt; ATF, Altyn Tagh fault.

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[9] The southern boundary of the Tibetan Himalayan zone is marked by the South Tibetan detachment system (Figure 3b), a family of top-to-the-north normal faults that modified the pre-existing, probably depositional, contact between the Tethyan sequence and the underlying Greater Himalayan sequence (GHS) [Burchfiel et al., 1992; Hodges et al., 1996; Godin et al., 1999; Hodges, 2000; Grujic et al., 2002]. The GHS consists of a 5–20 km thick succession of metasedimentary rocks and orthogneiss that were metamorphosed to amphibolite grade during Eocene and Oligocene time [LeFort, 1986; Pêcher, 1989; Vannay and Hodges, 1996; Coleman, 1998; Godin et al., 1999; Catlos et al., 2001]. Early to middle Miocene leucogranites locally intruded the GHS and the overlying Tethyan sequence [LeFort, 1986; Harrison et al., 1999]. Eclogites have been documented locally in structural culminations in the Tibetan Himalaya (the north Himalayan gneiss domes). The rocks that contain the eclogites are probably of Greater Himalayan affinity [Tonarini et al., 1993; de Sigoyer et al., 1997; O'Brien et al., 2001]. These rocks were evidently metamorphosed during Eocene time and subsequently exhumed from depths of up to ∼100 km to mid-crustal levels during late Eocene-Oligocene time. Any model to explain the Himalayan orogenic belt must incorporate a mechanism to exhume these rocks [Chemenda et al., 2000; O'Brien et al., 2001; Kohn and Parkinson, 2002].

[10] The rocks of the GHS have traditionally been regarded as Indian basement that was sliced off of the northern edge of the subcontinent and thrust southward on top of the underlying Lesser Himalayan sequence (LHS) during the Cenozoic orogeny [Gansser, 1964; Dewey et al., 1988]. Recent U-Pb zircon geochronologic studies [Parrish and Hodges, 1996; DeCelles et al., 2000] and Sm-Nd isotopic studies [Parrish and Hodges, 1996; Whittington et al., 1999; Ahmad et al., 2000; Robinson et al., 2001], however, indicate that the GHS is nearly 1 Gyr younger than LHS rocks. Moreover, the GHS protoliths are upper crustal rocks (mainly sediments) that are possibly exotic to India and experienced regional, medium- to high-grade metamorphism and igneous intrusion during Cambrian-Ordovician time along the tectonically active northern margin of Gondwana [Manickavasagam et al., 1999; DeCelles et al., 2000; Marquer et al., 2000]. The southern boundary of the Greater Himalayan zone is the Main Central thrust, which places the high-grade metamorphic rocks of the GHS on top of low-grade metasedimentary rocks of the LHS. Recent studies have proposed that the GHS was extruded southward from beneath the southern Tibetan Plateau, contemporaneous with motion on the Main Central thrust (below) and the South Tibetan detachment system (above) [Beaumont et al., 2001; Hodges et al., 2001; Grujic et al., 2002].

[11] The LHS is composed mainly of a thick (∼10 km) succession of early to middle Proterozoic greenschist-grade metasedimentary rocks, with local mafic and felsic intrusions. Detrital zircons from the lower part of the succession in Nepal are dominated by U-Pb ages of ∼1.86 Ga, and mylonitic augen gneisses that intruded into these rocks have yielded U-Pb zircon ages of ∼1.83 Ga [DeCelles et al., 2000; DePietro and Isachsen, 2001]. On top of the Proterozoic LHS rocks in Nepal are the Permian and Paleocene Gondwanas, which consist of a thin succession of sedimentary and volcanic rocks [Sakai, 1989]. On top of the Gondwanas are Eocene and early Miocene foreland basin deposits [Najman et al., 1997; DeCelles et al., 1998a; Najman and Garzanti, 2000]. In northwestern India and Pakistan the LHS is overlain directly by lower Paleozoic rocks [Pogue et al., 1999].

[12] The structure of the Lesser Himalayan zone throughout much of the central part of the Himalayan fold-thrust belt is dominated by a large antiformal duplex directly south of the Main Central thrust (Figure 3b) [Dhital and Kizaki, 1987; Schelling, 1992; Srivastava and Mitra, 1994; DeCelles et al., 1998b, 2001]. Where best documented in northern India and western Nepal, this duplex consists of a roof thrust sheet (the Ramgarh sheet) and several internal horses of the older LHS rocks. The faults within the duplex have fed slip southward into the frontal part of the fold-thrust belt.

[13] The Main Boundary thrust marks the boundary between the LHS and the Neogene Siwalik Group foreland basin deposits [e.g., Burbank et al., 1996]. The portion of the Himalayan fold-thrust belt between the Main Boundary and Main Frontal thrusts is known as the Subhimalayan zone, and includes several southward verging thrust sheets composed almost exclusively of the Siwalik Group [e.g., Schelling and Arita, 1991; Mugnier et al., 1993; DeCelles et al., 1998b]. The structural front of the fold-thrust belt is generally considered to be the Main Frontal thrust, although an emerging consensus places the true front in the subsurface beneath the northern part of the Indo-Gangetic foreland basin system [Lillie et al., 1987; Powers et al., 1998; Wesnousky et al., 1999; Lavé and Avouac, 2000].

3.2. Pre-Cenozoic Structural Architecture of Greater India

[14] Prior to analyzing the contribution of Himalayan shortening to thickening of crust beneath the Tibetan Plateau, we must outline the pre-Cenozoic structural architecture of the northern part of the Indian subcontinent. As discussed above, the new geochronologic and petrogenetic data from the GHS and LHS suggest that the Greater Himalayan rocks were tectonically mobilized during early Paleozoic orogenic activity along the northern margin of Gondwana. In turn, this requires that the present boundary between the GHS and LHS (i.e., the Main Central thrust) be a tectonic overprint of an older orogenic structure [DeCelles et al., 2000]. The bulk of the Tethyan sequence was deposited unconformably on top of this early Paleozoic orogenic terrane. Thus, the pre-Cenozoic tectonic architecture of the northern margin of Greater India included three tectonostratigraphic terranes (Figure 3a): the Lesser Himalayan terrane, which rested depositionally upon Indian cratonic lower crust; the Greater Himalayan terrane, which was separated from the Lesser Himalayan terrane by a paleostructural zone and consisted of supracrustal (mainly sedimentary) rocks and underlying Greater Himalayan lower crust; and the Tethyan Himalayan terrane, which was deposited on top of the Greater Himalayan terrane (and probably partly overlapped onto the Lesser Himalayan terrane) largely after the latter had been consolidated with India. These three terranes formed the northern part of Greater India as defined by Veevers et al. [1975]. Tectonic stripping of the Greater Indian cover rocks from their lower crust during the Cenozoic orogenic event would have produced a slab of lower crustal rocks that we refer to as Greater Indian lower crust (GILC), which is a composite of Greater Himalayan and Indian cratonic lower crust (Figure 3b).

3.3. Tibetan Plateau

[15] In this section, we briefly highlight the key geological features of the Tibetan Plateau and adjacent areas that must be explained (or at least accommodated) by models for the origin of the Plateau. Our objective is not to provide an exhaustive account of Plateau geology; for this the reader is referred to reviews by Sengor and Natal'in [1996], Yin and Harrison [2000], and Tapponnier et al. [2001].

[16] We adopt the Fielding et al. [1994] definition of the Tibetan Plateau as the contiguous region of high (>4.5 km) elevation that stretches between the longitudes 72°E and 99°E, and from the junction of the Altyn Tagh strike-slip fault and the Kunlun Mountains on the north to the Indus-Yalu suture zone on the south (Figure 1). The Plateau is characterized by low slopes, average relief of ∼1 km at long (100 km) wavelengths, and internal drainage in its central part. In contrast, the edges of the Plateau have extremely steep slopes and relief locally greater than 6 km [Fielding et al., 1994]. In detail, local topographic relief decreases systematically northward across the Plateau (Figure 2), from ∼1–2 km in the southern half, to <1 km in the northern half [Fielding et al., 1994]. The northeastern margin of the Tibetan Plateau is bordered by a 500,000 km2 region in which average elevation is 2–4 km. Tapponnier et al. [2001] referred to this region as “Plio-Quaternary Tibet,” and proposed that it consists of alternating north verging thrust systems and intervening flexural basins. Shortening in Plio-Quaternary Tibet is on the order of 200 km [Métivier et al., 1998; Tapponnier et al., 2001].

[17] The Tibetan Plateau consists of three major crustal blocks: the Lhasa, Qiangtang, and Songpan-Ganzi terranes [Matte et al., 1996; Yin and Harrison, 2000] (Figures 1 and 2). The Songpan-Ganzi terrane is bounded on the north by the Kunlun suture zone and on the south by the Jinsha suture zone. It consists mainly of thick Triassic turbidites (the “Songpan-Ganzi flysch”) that were deposited in a remnant ocean basin to the west of the Qinling-Dabie orogenic belt, which marked the collision zone between the North and South China cratons [Yin and Nie, 1996; Zhou and Graham, 1996; Hacker et al., 1996]. The Songpan-Ganzi basin was tectonically collapsed during late Triassic-early Jurassic time (see summaries by Zhou and Graham [1996]; Sengor and Natal'in [1996]; Yin and Nie [1996]; and Yin and Harrison [2000]). In its eastern part, the Songpan-Ganzi terrane experienced thin-skinned thrusting and shortening of several tens of kilometers during early Tertiary time [Yin and Harrison, 2000].

[18] South of the Songpan-Ganzi terrane lies the Qiangtang terrane, which rifted off of northern Gondwana during late Paleozoic time and collided with Eurasia in late Triassic-early Jurassic time [Yin and Nie, 1996; Yin and Harrison, 2000]. The Qiangtang terrane consists of metamorphic rocks overlain in structural contact by upper Paleozoic and Mesozoic strata. The metamorphic rocks consist of metasedimentary and mafic schists that enclose less deformed blocks of blueschist-bearing metabasites, upper Paleozoic strata, and minor ultramafic rocks [Kapp, 2001; Kapp et al., 2000, 2001]. Kapp et al. [2000] interpreted these rocks as a metamorphosed melange, and the contacts between the melange and the overlying Paleozoic-Mesozoic rocks as low-angle normal faults that were active during Late Triassic-Early Jurassic time. The footwall rocks experienced pressures of >8 kbar and temperatures of 350°–550°C [Kapp et al., 2000]. It is possible that the footwall rocks include Songpan-Ganzi flysch, oceanic melange materials, and Paleozoic passive margin deposits that were subducted beneath the upper part of the Qiangtang terrane during the Triassic and later unroofed by low-angle normal faulting [Kapp et al., 2000]. Hacker et al. [2000] reported xenoliths of metamorphosed sediments in the central part of the Qiangtang terrane that were heated up to ∼1100°C at depths of 30–50 km and then rapidly conveyed (by volcanic eruption) to the surface at approximately 3.5 Ma. This, together with the presence of Panafrican zircons [Kapp et al., 2000] suggests that the lower crust of the Qiangtang terrane has affinities with Greater Himalayan rocks as well as the Songpan-Ganzi flysch. Kapp et al. [2002] documented significant post-mid-Cretaceous shortening in the central Qiangtang terrane, and Horton et al. [2002] showed that ∼55 km of shortening occurred in its eastern part during Paleocene and Eocene time. In addition, the Qiangtang anticlinorium (Figure 1) involves Tertiary rocks and therefore must have formed partly during the Cenozoic [Yin and Harrison, 2000]. However, significant growth of the Qiangtang anticlinorium occurred prior to the Indo-Eurasian collision during Cretaceous northward underthrusting of the Lhasa terrane beneath the southern Qiangtang terrane [Kapp et al., 2002].

[19] The Lhasa terrane consists of medium- to high-grade metasedimentary rocks of late Precambrian-early Paleozoic age overlain by Ordovician and Carboniferous to Cretaceous sedimentary rocks [Dewey et al., 1988]. Panafrican-age zircons in the footwall of the Nyainqentanghla detachment [D'Andrea et al., 1999] and the Amdo gneiss [Xu et al., 1985] suggest the presence of crust with Greater Himalayan affinities beneath the cover rocks. The Lhasa terrane probably rifted off of northern Gondwana during Triassic time and collided with the southern margin of the Qiangtang terrane in Late Jurassic-middle Cretaceous time [Dewey et al, 1988, Gaetani et al., 1993; Matte et al., 1996; Kapp et al., 2002]. The southern fringe of the Lhasa terrane is intruded by the extensive Cretaceous-Eocene Gangdese batholith belt, which formed an Andean-style continental margin arc above the subducting Neotethyan oceanic slab [Allégre et al., 1984]. Murphy et al. [1999] reported that as much as 187 km of shortening may have occurred in the Lhasa terrane during Late Jurassic-Cretaceous time. If this shortening was accommodated entirely within the Lhasa terrane, then the thickness of Lhasa terrane crust prior to the Himalayan orogeny could have been ∼55 km [Murphy et al., 1997], enough to support regional elevations of ∼3 km, assuming Airy isostasy. However, mid-Cretaceous shallow marine limestones are widespread in the Lhasa terrane, indicating that regions of low elevation persisted until the Late Cretaceous [Zhang, 2000]. Kapp et al. [2002] presented evidence for northward underthrusting of Lhasa terrane lower crust beneath the Qiangtang terrane; this process would minimize the amount of Cretaceous crustal thickening in the Lhasa terrane. Cenozoic shortening in the Lhasa terrane is minimal [Dewey et al., 1988; Pan, 1993; Murphy et al., 1997; Yin and Harrison, 2000], being largely confined to displacements on thrust systems that overprinted the Indus-Yalu suture zone [Yin et al., 1999].

[20] From the above summary, we conclude that large portions of the Tibetan Plateau are underlain by rocks that have upper crustal compositions and physical properties [e.g., Owens and Zandt, 1997; Hacker et al., 2000; Kapp et al., 2000]. The isotopic and geochronologic characteristics of these rocks, where known, suggest affinities with Greater Himalayan (Late Proterozoic-early Paleozoic) and Mesozoic metamorphic rocks. The Mesozoic-Cenozoic orogenic events of central Eurasia, therefore, may be viewed as a piecemeal redistribution of exotic tectonostratigraphic assemblages from Gondwana to Eurasia [e.g., Sengor and Natal'in, 1996; Yin and Nie, 1996]. Documented shortening of upper crust in the central and northern Tibetan Plateau during Cenozoic time is <100 km [Horton et al., 2002; Kapp et al., 2002]. To the north of the Kunlun fault in Plio-Quaternary Tibet, an additional ∼200 km of shortening occurred during mid-late Cenozoic time [Métivier et al., 1998; Tapponnier et al., 2001]. In the southern part of the Plateau, Cenozoic shortening was <50 km [Pan, 1993; Yin and Harrison, 2000]. Evidence for Mesozoic shortening in the upper crust of the Plateau (and presumably crustal thickening) is widespread. However, even where the estimates of pre-Cenozoic crustal shortening are greatest (in the Lhasa terrane [Murphy et al., 1999]) the presence of mid-Cretaceous shallow-marine sedimentary rocks suggests that parts of the region lay close to sea level until the late Cretaceous. The oldest widespread nonmarine sedimentary rocks with relevance for the timing of uplift in the Lhasa terrane are Paleocene-Eocene in age, and even these were probably deposited close to sea level [Willems et al., 1996] (summary by Rowley [1996]). On the other hand, the presence of the Gangdese batholith along the southern margin of Eurasia prior to the Cenozoic collision, evidence for up to ∼60% pre-Cenozoic horizontal shortening in the central Lhasa terrane, and evidence for Mesozoic crustal thickening and extensional detachment faulting in other parts of the Tibetan Plateau imply that the crust in the southern and central parts of the Plateau may have been as thick as ∼50–55 km before the Himalayan orogeny [Murphy et al., 1999; Yin et al., 1999; Robinson et al., 2002; Kapp et al., 2000, 2002]. Unless a dense eclogitic root or dynamic mantle processes held this thickened crust at low elevations, it is likely that local elevation may have been on the order of 2–4 km. For the crust of the Tibetan Plateau (excluding Plio-Quaternary Tibet) to reach its present thickness and high regional elevation, it must have thickened by an additional ∼10–20 km during the Himalayan orogenic event, without shortening internally by more than ∼15–20%.

3.4. Extension and Strike-Slip Faulting in the Tibetan Plateau

[21] Two sets of normal faults and several major strike-slip faults dominate the present structure of the Tibetan Plateau [Armijo et al., 1986; Molnar and Lyon-Caen, 1989; Tapponnier et al., 2001]. The normal fault systems include the generally northwest-southeast to east-west striking South Tibetan detachment system and numerous north-south to northeast-southwest striking graben. These basins are largest and most numerous in the southern part of the Plateau [Yin, 2000], where they cut across the Tibetan Himalaya and the Lhasa terrane (Figure 1). Shallow earthquakes in the modern Tibetan Plateau are dominated by strike-slip and extensional focal mechanisms [e.g., Armijo et al., 1986; Molnar and Lyon-Caen, 1989]. The direction of extension is approximately east-west, at nearly a right angle to the direction of compression in the Himalayan fold-thrust belt [Armijo et al., 1986; Molnar and Lyon-Caen, 1989]. Total east-west extension amounts to less than 3%, or ≤40 km [Armijo et al., 1986; Molnar et al., 1993].

[22] Because of a possible linkage between gravitational extension in the upper crust and attainment of a critical thickness of the crust [Molnar and Lyon-Caen, 1989; Stüwe and Barr, 2000], much interest has been focused on the timing of extension in the Tibetan Plateau. North-south extension along the South Tibetan detachment system commenced ∼22 Ma and occurred episodically during early and middle Miocene time along the length of the fault system [Hodges et al., 1996; Edwards and Harrison, 1997; Coleman, 1998; Hodges, 2000; Yin and Harrison, 2000; Grujic et al., 2002]. Neotectonic studies in central Nepal suggest that faults in this system may still be active [Hurtado et al., 2001]. Most of the available evidence suggests that east-west extension in the Plateau began in mid- to late Miocene time (14–8 Ma) [Harrison et al., 1992, 1995; Coleman and Hodges, 1995; Edwards and Harrison, 1997; Garzione et al., 2000a, 2000b; Williams et al., 2001; Blisniuk et al., 2001]. With respect to the issue of the timing of elevation gain, Garzione et al. [2000a, 2000b] reported oxygen isotopic data that indicate the depositional floor of Thakkhola graben (Figure 1) has been at elevations >4 km since at least 11 Ma.

[23] Several major strike-slip faults are present on the Tibetan Plateau, including the Altyn Tagh, Kunlun, Xingshuihe, Jiali, Red River, and Karakoram fault systems. In eastern Tibet left- and right-lateral faults translate Tibetan terranes eastward, although the amounts of translation versus rotation and distributed motion are still debated [e.g., Tapponnier et al., 1982, 1990; Burchfiel et al., 1995; Yin and Harrison, 2000; Wang et al., 2001]. The left-lateral Altyn Tagh fault, which forms the northwestern boundary of the Tibetan Plateau, has accommodated ∼280–550 km [Peltzer and Tapponnier, 1988; Yin and Harrison, 2000] or 375 ± 25 km [Yue et al., 2001] of slip since Oligocene time. The right-lateral Karakoram fault in the western part of the Plateau may transfer slip between the Himalayan fold-thrust belt and the Pamir Range [Ratschbacher et al., 1994; Yin et al., 1999; Murphy et al., 1999].

[24] In addition to the widespread evidence for extensional and strike-slip faulting at shallow levels (<20 km) in Tibetan crust [Molnar and Lyon-Caen, 1989], a small number of moderate-sized earthquakes have been documented at depths of ∼70–113 km beneath the Lhasa terrane and Himalayan fold-thrust belt (Figure 2) [Ekstrom, 1987; Chen and Molnar, 1983; Chen and Kao, 1996; Zhu, 1998]. These earthquakes occurred within the mantle lithosphere, and demonstrate that the upper mantle is cold and strong enough in this region to store elastic strain energy. The handful of these earthquakes for which fault plane solutions are available exhibit horizontal, east-west extension, similar to the faults in the upper crust of Tibet [Molnar and Chen, 1983; Chen and Molnar, 1983; Chen and Kao, 1996].

3.5. Magmatism on the Tibetan Plateau

[25] The southern margin of Eurasia was occupied by the >2000 km long, calc-alkaline Gangdese magmatic arc from mid-Cretaceous through mid-Eocene time (Figure 1) [Chang and Zheng, 1973; Tapponnier et al., 1981; Allégre et al., 1984]. Silicic volcanic rocks, mainly ash flow tuffs of the Upper Cretaceous-Paleocene Linzizong Formation, are widespread in the Lhasa terrane and the southern part of the Qiangtang terrane [Pan, 1993; Kapp et al., 2002; Ding et al., 2002]. High-potassium magmas mixed with a mantle component have erupted on the Tibetan Plateau since ∼45 Ma, with major pulses of activity during late Eocene-Oligocene time in the Qiangtang terrane and during the Miocene-Quaternary in the Songpan-Ganzi and Lhasa terranes [Deng, 1978; Coulon et al., 1986; Harris et al., 1988; Turner et al., 1993; Chung et al., 1998; Miller et al., 1999, 2000; Ding et al., 2002]. Widespread geothermal activity and several independent geophysical and petrological data sets indicate that the crust of the Tibetan Plateau is abnormally hot, although the debate persists whether partial melt is pervasive throughout the plateau or restricted to the extensional grabens [Nelson et al., 1996; Owens and Zandt, 1997; Alsdorf and Nelson, 1999; Hacker et al., 2000; Wei et al., 2001; Kind et al., 2002].

4. Geophysical Constraints on Sub-Tibetan Lithospheric and Crustal Structure

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Models for Tibetan Plateau Uplift
  5. 3. Geology of the Himalayan Fold-Thrust Belt and Tibetan Plateau
  6. 4. Geophysical Constraints on Sub-Tibetan Lithospheric and Crustal Structure
  7. 5. Shortening in the Himalayan Fold-Thrust Belt
  8. 6. An Integrated Model for Uplift of the Tibetan Plateau
  9. 7. Implications of the Model
  10. 8. Conclusions
  11. Acknowledgments
  12. References

[26] Figures 2, 3, and 4 depict an interpretation of the structure of the lithosphere and upper mantle beneath India and central Eurasia based on recent geophysical studies. Results of the 1991–92 Sino-American PASSCAL broadband experiment [McNamara et al., 1994, 1997; Owens and Zandt, 1997], the international and multidisciplinary INDEPTH experiments [Nelson et al., 1996; Kosarev et al., 1999; Zhao et al., 2001], and the series of Sino-French seismic studies [Hirn, 1988; Hirn et al., 1995; Wittlinger et al., 1996] have begun to delineate the lithospheric-scale structure beneath the Tibetan Plateau, principally in the central and eastern Plateau. A comprehensive review of even these selected studies is beyond the scope of this paper; instead a summary is given of the results most relevant to the thesis of this paper. Most studies agree that the Tibetan crust is up to ∼75 km thick in the southern Lhasa terrane and thins to the north [Owens and Zandt, 1997; Zhao et al., 2001]. The crust is ∼65 km thick in the Qiangtang terrane, and thins to ∼60 km in the Songpan-Ganzi terrane [Zhao et al., 2001; Kind et al., 2002]. Some studies have suggested that the crustal thickness variations occur as abrupt steps corresponding to suture boundaries such as the Jinsha suture and major strike-slip faults such as the Kunlun fault [Hirn et al., 1995; Wittlinger et al., 1996]. The seismic evidence is strong for an abrupt step in the Moho beneath the boundary between the North Kunlun Mountains and Qaidam Basin [Zhu and Helmberger, 1998], but the data for the Jinsha suture are more equivocal owing to larger station spacing. The Moho step at the Kunlun-Qaidam border is associated with a change in topography, so it is not particularly surprising. The presence or absence of a step beneath the Jinsha suture zone, where there is no major change in elevation, is potentially more significant but still uncertain. The 10–15 km gradual variation in crustal thickness across the Plateau may owe to the Moho adjusting independently of the surface to provide isostatic compensation for lateral density contrasts in the upper mantle [Molnar, 1988; Bird, 1991; Owens and Zandt, 1997; Zhao et al., 2001].

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Figure 4. (a) Tectonic map of India and central Eurasia, showing major terrane boundaries. North-south line X-Y indicates location of tomographic profile shown in (b). (b) Simplified version of tomographic model of Van der Voo et al. [1999] along profile X-Y. Vertical ruled zones are characterized by slower than average seismic wave propagation; stippled areas are zones of faster than average wave propagation. (c) Geological interpretation of tomographic profile, showing the Greater Indian lithospheric slab (anomaly I) and the Neotethyan oceanic slab (anomaly II). Details of the lithosphere under Tibet are from Jin et al. [1996], Owens and Zandt [1997] and Kosarev et al. [1999]. Sloping lines labeled 45° and 55° represent plausible projections to the surface of the Neotethyan oceanic slab (anomaly II) before it broke off and foundered into the mantle. Abbreviations: MFT, Main Frontal thrust; ISZ, Indus-Yalu suture zone; BSZ, Banggong suture zone; ATF, Altyn Tagh fault.

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[27] Although the details of the plateau crustal structure are still sketchy, a few characteristic traits have been documented. The bulk crustal seismic velocities are generally low, consistent with an average composition more felsic than average continental crust [Owens and Zandt, 1997; Rodgers and Schwartz, 1998; Zhao et al., 2001]. Seismic reflection profiling in the Lhasa terrane observed mid-crustal “bright spots” interpreted as evidence for pervasive partial melting in the crust [Nelson et al., 1996] or a combination of fluids and melt [Makovsky and Klemperer, 1999]. Magnetotelluric studies have found the middle and lower crust of the plateau are anomalously conductive, most likely because of the presence of aqueous fluids and partial melt [Wei et al., 2001]. Combining seismic conversion data from several international experiments, Kind et al. [2002] concluded that over most of the central plateau, the average Vp/Vs is near normal, indicating that despite the apparent high conductivity, the volume of fluid and melt is not greater than a few percent. Available data indicate that much of the Lhasa terrane is underlain by a ∼15 km thick lower crustal layer where the seismic velocity increases up to 7.2 km/s or higher, causing the “doublet” arrival in the seismic conversion data [Owens and Zandt, 1997; Kind et al., 2002]. The correlation of the high velocity lower crustal layer and high velocity upper mantle in southern Tibet suggests the presence of seismically fast lower Indian cratonic crust and lithosphere beneath southern Tibet (Figures 2 and 3). The northern termination of the high velocity layer may represent the end of cratonic Indian lower crust, but the Greater Himalayan lower crust, of differing composition, may continue farther north (Figure 3b).

[28] The interpretation that the Indian mantle lithosphere (including Indian lower crust) underthrusts the Tibetan Plateau nearly horizontally as far north as the Banggong suture zone (32°N) is based on numerous seismological studies that indicate a seismically fast (cold) upper mantle beneath the southern part of the Plateau and a seismically slow (warm) mantle beneath its northern part [e.g., Ni and Barazangi, 1984; McNamara et al., 1997; Owens and Zandt, 1997; Kind et al., 2002]. Even among those who agree on the existence of the Indian lithosphere beneath southern Tibet, the dip of the underthrust Indian lithosphere has been the subject of debate. Owens and Zandt [1997] and others have suggested that the Indian slab is horizontal to the latitude of the Banggong suture. Based on gravity data, Jin et al. [1996] suggested that the Indian plate is subducting under the Lhasa terrane at a moderate angle instead of horizontally. In their model, the Indian-Eurasian mantle suture is located beneath the Indus-Yalu suture zone and the Greater Indian lithosphere continues to subduct into the upper mantle for an indeterminate distance. Kosarev et al. [1999] modeled migrated receiver functions to image a shallow north dipping interface (their Zangbo Conversion Boundary) at a depth of 80 km just north of the Indus-Yalu suture zone to a depth of ∼200 km beneath the Banggong suture. They suggested that the Zangbo Conversion Boundary is evidence for the presence of subducting, cold Indian mantle lithosphere beneath the Plateau, supporting the model of Jin et al. [1996]. Zandt et al. [2000] further supported this interpretation with geoid modeling that suggests a high density, north dipping Indian lithospheric slab (stripped of most of its original crust) underlying much of the Plateau. In the combined data analysis, the Zangbo Conversion Boundary was not enhanced, and the weight of the seismic evidence has swung back to a horizontal Indian lithosphere beneath the southern plateaus [Kind et al., 2002]. The alternative interpretations of the distribution of Greater Indian lithosphere beneath Tibet are depicted in Figure 3b. For the purposes of this paper, the alternatives are not significantly different from each other; the amount of shortening predicted in the Himalaya is essentially the same regardless of which interpretation is adopted.

[29] An important additional piece of evidence on the nature of the plateau mantle is the presence of east-west oriented seismic anisotropy in the upper mantle under the plateau that is especially strong beneath the northern plateau [McNamara et al., 1994]. Two competing interpretations have been offered for these observations. One interpretation is that the upper mantle beneath northern Tibet is composed of a thick keel of north-south shortened Eurasian lithosphere and the seismic anisotropy is reflecting the resulting east-west deformation fabric in the mantle [McNamara et al., 1994]. An alternative interpretation is that the horizontal motion of rigid Tibetan lithospheric blocks shears the asthenosphere below and produces the east-west anisotropy in the asthenospheric mantle [Hirn et al., 1995; Lavé et al., 1996]. Holt [2000] rejected this latter mechanism based on the apparent absence of strong anisotropy beneath the much more rapidly moving Indian subcontinent [Chen and Özalaybey, 1998]. However, recent studies show that asthenospheric-flow induced anisotropy appears less related to absolute plate motions and more related to forced flow around lithospheric slabs and keels [Russo and Silver, 1994; Fouch et al., 2000]. This supports an alternative interpretation that north of the Banggong suture, weak asthenosphere flows eastward in response to compression between converging strong and thick Indian and Eurasian lithospheres (Figures 2 and 3b). In this model, the anisotropy is attributed to a combination of sublithospheric mantle flow [Owens and Zandt, 1997] and lithospheric fabric associated with a south dipping Eurasian lithosphere [Furlong and Owens, 1997]. This idea is consistent with Holt's [2000] conclusion that the vertical coherence of deformation indicators in Tibet “may be influenced more by the velocity boundary conditions imposed on both crust and mantle than by coupling between crust and upper mantle.” Recently produced seismic images showing a southward dipping converter boundary beneath northern Tibet [Kosarev et al., 1999; Tapponnier et al., 2001; Kind et al., 2002] have bearing on this interpretation. A logical interpretation of this seismic feature is that it marks the top of southward subducting Eurasian lithosphere. Magnetotelluric studies have documented a broad conductivity anomaly centered beneath the Qiangtang terrane extending into the upper mantle. The anomaly is interpreted as partial melting due to localized upwelling of asthenosphere, and is bounded on the north and south by relatively resistive zones that may indicate where Eurasian and Indian lithosphere descends [Wei et al., 2001]. Based on this interpretation of multiple data sets there is direct evidence that the Eurasian lithosphere, beneath the crust, is behaving like a plate and not deforming by pure shear.

[30] Recent work on the seismic anisotropy in the Plateau crust using data from the 1991–92 PASSCAL experiment is also relevant to some of these questions (H. Folsom et al., manuscript in preparation, 2002). Middle to lower crustal anisotropy is present at most stations, with a unique fast axis trending north-south to northwest-southeast in the south, nearly east-west in the central Plateau, and north-south to northeast-southwest in the northern Plateau. This pattern appears consistent with recent ductile deformation due to both topographically induced flow and to boundary forces from subducting lithosphere at the northern and southern margins of the Plateau. The orientations of crustal anisotropy are not entirely consistent with shear wave splitting fast polarization directions, potentially implying distinct motions in the crust and mantle (Folsom et al., manuscript in preparation, 2002).

[31] The structure of the deeper mantle beneath India and central Eurasia is revealed by recent tomographic studies. Tomographic models of Grand et al. [1997] and Van der Voo et al. [1999] suggest that relatively cold, seismically fast lithosphere of the Indian plate plunges northward into the asthenosphere to a depth of several hundred kilometers beneath the central part of the Himalaya (Figure 4, anomaly I). The location of this anomaly ∼500 km south of the northern end of Greater Indian lithosphere beneath the Tibetan Plateau (Figure 3b) suggests that the anomaly represents a slab of Greater Indian lithosphere that is no longer attached to the Indian plate. A second, deeper (∼1000–2000 km) region of cold, fast lithosphere that was imaged to the south of the inferred Greater Indian slab was interpreted by Van der Voo et al. [1999] as Neotethyan oceanic lithosphere that broke off of the Indian plate at the onset of collision (Figure 4, anomaly II). This deeper slab of (presumably) oceanic lithosphere can be traced continuously in tomographic images from northern Indonesia to the eastern Mediterranean region [Spakman, 1991; Van der Voo et al., 1999]. Continued northward motion of the Indian plate during early-middle Tertiary time carried the lower end of the inferred Indian slab (anomaly I) northward away from the detached Neotethyan slab (Figure 4). Alternatively, anomaly I (Figure 4) may be interpreted as a slab of Neotethyan oceanic lithosphere. This interpretation, however, conflicts with known rates of Indian plate migration and the offsets between anomalies I and II and the northern end of Greater Indian lithosphere beneath the Tibetan Plateau. The lateral offset between anomaly I and the northern end of Greater Indian lithosphere is only ∼500 km (Figure 5, point 6). At a rate of 50–55 mm/yr of northward migration of the Indian plate, this separation would have developed since 10 Ma, more than 40 Myr after the initial collision. It seems unlikely that Neotethyan lithosphere would have continued to subduct beneath central Eurasia until late Miocene time. Moreover, this reconstruction would predict only a few hundred kilometers of underthrusting of Greater India beneath the Himalaya, which conflicts with the conservative estimates of shortening in the Himalaya discussed in the following section. We therefore concur with Van der Voo et al. [1999] that tomographic anomaly I is composed of Greater Indian lithosphere. In a later section, we further postulate that this fragment of lithosphere is composed of Greater Himalayan material.

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Figure 5. Simplified version of the cross section depicted in Figure 3, annotated with the key geometric constraints that must be explained in a kinematic model of the India-Eurasia collision. Sources of information on Tibetan shortening: Yin and Harrison [2000], Tapponnier et al. [2001], Horton et al. [2002]. See text for discussion.

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[32] Neither the seismic data nor the geoid observations can resolve in detail the presence or absence of GILC beneath the Tibetan Plateau or on the detached slab of inferred Greater Indian lithosphere (anomaly I). However, the broadband seismic results dictate that a slab of GILC must extend at least as far north as 32°N directly beneath the Plateau. From this latitude northward, there is no lithospheric barrier that would have blocked further northward insertion of GILC beneath the Plateau. We therefore suggest that a northward tapering slab of GILC continues northward beneath the Tibetan crust, perhaps as far north as the Kunlun fault (Figure 3b).

5. Shortening in the Himalayan Fold-Thrust Belt

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Models for Tibetan Plateau Uplift
  5. 3. Geology of the Himalayan Fold-Thrust Belt and Tibetan Plateau
  6. 4. Geophysical Constraints on Sub-Tibetan Lithospheric and Crustal Structure
  7. 5. Shortening in the Himalayan Fold-Thrust Belt
  8. 6. An Integrated Model for Uplift of the Tibetan Plateau
  9. 7. Implications of the Model
  10. 8. Conclusions
  11. Acknowledgments
  12. References

[33] The timing of initial contact between Greater India and Eurasia remains a subject of debate. The usual argument employed for initial contact is the timing of disappearance of marine waters between the two continents. However, many modern foreland basin systems are filled partly or wholly by marine waters [DeCelles and Giles, 1996; Sinclair, 1997], and so this is not a meaningful indicator of contact between Greater India and Eurasia. We prefer to use the timing of initial foreland basin development. In Nepal and northern India, it has been shown that shallow marine sediments of early to middle Eocene age were derived from the nascent Himalayan fold-thrust belt and the Indus-Yalu suture zone and deposited in a southward migrating foreland basin system [Pivnik and Wells, 1996; Najman et al., 1997; DeCelles et al., 1998a; Najman and Garzanti, 2000]. We therefore place the time of initial thrusting in the Himalaya as ∼55 Ma, which is consistent with several previous independent estimates [Klootwijk et al., 1985; Besse and Courtillot, 1988; LePichon et al., 1992; Patzelt et al., 1996].

[34] Plate tectonic reconstructions based on paleomagnetic data from the Himalayan orogenic belt suggest that 2600 ± 900 km of post-collision convergence has taken place between Eurasia and Greater India, with 1700 ± 610 km of this total accommodated by north-south shortening in the Tibetan Plateau and lateral tectonic escape [Patriat and Achache, 1984; Achache et al., 1984; Besse and Courtillot, 1988, 1991]. The ∼900 km difference between these average values is available for shortening in the Himalayan fold-thrust belt [LePichon et al., 1992]. Published estimates of shortening in the Himalaya based on paleomagnetic data range between ∼700 km and 1500 km [Patzelt et al., 1996].

[35] An alternative approach to predicting the amount of shortening in the Himalayan fold-thrust belt is given in Figure 6. The present length of Indian crust at the surface is ∼2304 km (measured along 83°E). Paleomagnetic data indicate a long-term convergence rate of ∼55 mm/yr between Greater India and Eurasia since early Tertiary time [e.g., Patriat and Achache, 1984; Dewey et al., 1989; Molnar et al., 1993]. This rate can be used to hindcast the position of the Indian continent at ∼55 Ma, placing its southern tip at latitude ∼25°S at the onset of the collision. Current estimates of the paleolatitude of the Indus-Yalu suture zone place it at 6.5°N ± 2.5° at the onset of collision [Klootwijk et al., 1985; Van der Voo, 1993]. The distance between this paleolatitude and the paleolatitude of the southern tip of the continent provides the length of Indian continental crust at the onset of collision, 3075 ± 192 km. The 771 ± 192 km difference between this length and the present length of Indian continental crust at the surface is the predicted total shortening (Figure 6), which is within error of the estimate by LePichon et al. [1992]. If we assume that the onset of collision was at 50 Ma, then the shortening estimate reduces to 446 ± 192 km. Let us see how these estimates compare with estimates derived from geologic data from the Himalayan fold-thrust belt.

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Figure 6. Plot showing the northward trajectory of Greater India, the Indus-Yalu suture zone, and the southern tip of India since 55 Ma, based on discussion in text. The shaded regions reflect uncertainties in paleomagnetic data.

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[36] Only a handful of regional balanced cross-sections through the Himalayan fold-thrust belt in Nepal, India, and Pakistan have been published (Figure 1). Table 1 lists the available balanced cross sections and estimates of shortening. Coward and Butler [1985] produced the only complete structural transect of the Himalayan fold-thrust belt, in northern Pakistan, which yielded a total shortening estimate of 470 km. Complete sections in the remainder of the Himalaya must be stitched together by combining cross sections by different investigators. The available studies are generally divided into those that cover the Indus-Yalu suture zone, the Tibetan Himalaya, or the portion of the fold-thrust belt south of the South Tibetan detachment system.

Table 1. Shortening Estimates in the Himalayan Fold-Thrust Belt
Reference in Figures 1 and 7ReferenceLocationStructural BoundariesaTectonostratigraphic ZonesbShortening, km
  • a

    MMT, Main Mantle thrust; ISZ, Indus-Yalu suture zone; MFT, Main Frontal thrust; MCT, Main Central thrust; STDS, South Tibetan detachment system.

  • b

    GHZ, Greater Himalayan zone; LHZ, Lesser Himalayan zone; SHZ, Subhimalayan zone; THZ, Tibetan Himalayan zone.

1Coward and Butler [1985]PakistanMMT-MFTGHZ, LHZ, SHZ470 km
2Searle [1986]India, Zanskar and LadakhISZ-MMTTHZ126 km
3Searle et al. [1997]India, Zanskar and LadakhISZ-MCTTHZ150–170
4Srivastava and Mitra [1994]India, Kumaon and GarhwalSTDS-MCTGHZ + Almora thrust sheet193–260 km
5Srivastava and Mitra [1994]India, Kumaon and GarhwalMCT-MFTLHZ, SHZ161 km
6DeCelles et al. [1998a]Western NepalMCT-MFTLHZ, SHZ228 km
7Murphy and Yin [2003]Tibet, Mt. Kailas regionISZ-STDSTHZ + Indus Suture176 km
8DeCelles et al. [2001]Western Nepal, Seti RiverSTDS-MCTGHZ + Dadeldhura thrust sheet131–206 km
9DeCelles et al. [2001]Western Nepal, Seti RiverMCT-MFTLHZ, SHZ287 km
10Ratschbacher et al. [1994]Tibet, north of Arun RiverISZ-MCTTHZ133–139 km
11Schelling [1992]Central NepalSTDS-MCTGHZ140–210 km
12Schelling [1992]Central NepalMCT-MFTLHZ, SHZ70 km
13Schelling and Arita [1991]Far Eastern NepalSTDS-MCTGHZ140–175 km
14Schelling and Arita [1991]Far Eastern NepalMCT-MFTLHZ, SHZ45–70 km

[37] Searle [1986] estimated that ∼126 km of shortening occurred in the Tibetan Himalaya of northern India (near longitude 76°E), and Searle et al. [1997] increased this amount to 150–170 km based on more recent work. These authors argued that it is not feasible to balance cross sections in this part of the Himalaya because the deformation is so intense. Based on detailed mapping, Steck et al. [1998] suggested that hundreds of kilometers of shortening have occurred in the Tibetan Himalayan zone of northern India but they presented no balanced cross sections. Ratschbacher et al. [1994] produced shortening estimates of 133 km and 139 km for balanced cross sections in the Tibetan Himalaya near longitudes 88°E and 90°E.

[38] In the Indus-Yalu suture zone, post-collisional shortening has been documented by Yin et al. [1999]. Several tens of kilometers of shortening occurred during Oligocene and middle Miocene thrusting. These thrusts reactivated and crosscut suture-related features, suggesting that they are related to internal shortening in the Himalayan hinterland.

[39] Aside from Coward and Butler [1985], seven regional balanced cross sections of the Himalaya south of the South Tibetan detachment system have been published. Schelling [1992] constructed two regional sections in eastern Nepal yielding 210–280 km of shortening between the South Tibetan detachment system and the Main Frontal thrust. Schelling and Arita [1991] published a cross section in Sikkim (northeastern India) that yielded shortening estimates of 185–245 km. Srivastava and Mitra [1994] published a pair of cross sections separated by ∼100 km in Kumaon, northern India, which yielded a range of shortening between 353–421 km. DeCelles et al. [1998b] published a balanced cross section from far western Nepal that indicated shortening between the Main Central and Main Frontal thrusts of ∼228 km. DeCelles et al. [2001] constructed a second cross section, ∼50 km to the east of the previous section, that suggests ∼418–493 km of shortening between the South Tibetan detachment system and the Main Frontal thrust.

[40] By cobbling together the shortening estimates derived from nearby cross sections in the Indus-Yalu suture zone, Tibetan Himalaya, and the portion of the Himalaya south of the South Tibetan detachment system, estimates for the entire fold-thrust belt may be obtained in four districts (Figure 7). Hauck et al. [1998] combined the cross sections of Schelling [1992], Schelling and Arita [1991], and Ratschbacher et al. [1994] with the INDEPTH seismic reflection profile to construct a crustal-scale cross section with shortening of ∼323 km from the Indus-Yalu suture zone to the Main Frontal thrust in the eastern Himalaya. Greater shortening estimates are obtained in northern India (480 to 547 km)[Srivastava and Mitra, 1994] and western Nepal (556–623 km [DeCelles et al., 1998b]; and 643–669 km [DeCelles et al. [2001]) (Table 1). All of these estimates are minima because they do not include penetrative strain or small-scale folds and faults, which could significantly increase the total shortening.

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Figure 7. Compilation of shortening estimates along the Himalayan fold-thrust belt from west to east between the Hazara and Namche Barwa syntaxes. Dashed line indicates corresponding width of the Tibetan Plateau as measured in an arc-normal direction. Numbers indicate sources of data as listed in Table 1.

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[41] Figure 7 provides an along-strike comparison of the available estimates of shortening in the Himalayan fold-thrust belt, and the corresponding width of the Tibetan Plateau. Several features of this diagram stand out. First, the eastern third of the fold-thrust belt is terra incognita from the standpoint of crustal shortening estimates. The vastness of this region and its obvious importance for the origin of the Tibetan Plateau beckon geological and geophysical studies. Second, the diagram suggests that the greatest amount of shortening may be accommodated in the central part of the Himalayan arc, in northern India and western Nepal. The increase in displacement toward the apex of the arc is consistent with paleomagnetic data that indicate rotations in a counterclockwise sense in the eastern half of the orogenic belt and a clockwise sense in the western half of the belt [Klootwijk et al., 1985]. Third, the width of the Tibetan Plateau north of the Indus-Yalu suture zone and the amount of shortening in the Himalayan fold-thrust belt are remarkably similar for the western half of the orogenic belt. The mismatch, by a factor of almost two, in the eastern part of the fold-thrust belt and the adjacent Tibetan Plateau we attribute to the following. The estimates of shortening in the Himalaya of eastern Nepal are based on cross sections [Schelling and Arita, 1991; Schelling, 1992] that do not include two major thrust systems that have been mapped in western Nepal and northern India: the Ramgarh thrust and the Almora-Dadeldhura thrust [Valdiya, 1980; Srivastava and Mitra, 1994; DeCelles et al., 1998b, 2001]. Where mapped, these faults accommodated up to ∼250 km of shortening. Reconnaissance mapping in central and eastern Nepal [Pearson, 2002], U-Pb detrital zircon ages [DeCelles et al., 2000], and Sm-Nd isotopic studies [Robinson et al., 2001] indicate that structural and stratigraphic equivalents of both of these thrust sheets are indeed present in eastern Nepal. The state of the mapping is not sufficient to produce balanced cross sections, but if displacements on these two thrusts in eastern Nepal are roughly equivalent to their displacements in western Nepal, then the total minimum shortening in eastern Nepal should increase to ∼650 km (Figure 7).

[42] The long-term average rate of shortening in the Himalayan fold-thrust was ∼20 mm/yr throughout most of Neogene time and decreased to about 13–20 mm/yr during Pliocene to Recent time [Powers et al., 1998; DeCelles et al., 1998b, 2001; Lavé and Avouac, 2000] (Figure 8). The present rate of shortening in the central part of the orogenic belt is ∼17 mm/yr [Bilham et al., 1997; Larson et al., 1999]. Prior to the Neogene, however, the available shortening estimates from the Tibetan Himalaya suggest that shortening was more than three times slower, only ∼6 mm/yr. Based on arguments presented by Searle et al. [1997], it seems quite likely that the existing values of shortening for the Tibetan Himalayan zone are substantial underestimates. If the actual rate of shortening in the Tibetan Himalayan zone were equal to the Neogene shortening rates in the remainder of the Himalaya, then the total shortening would be on the order of 900–1000 km. More work in the Tibetan Himalaya is needed to clarify this issue.

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Figure 8. Plot of cumulative shortening in the Himalayan fold-thrust belt of western Nepal, based on kinematic evidence summarized by Ratschbacher et al. [1994] and DeCelles et al. [1998b, 2001]. Shown along left-hand vertical axis are locations (present coordinates) of major tectonic features in the Tibetan Plateau.

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6. An Integrated Model for Uplift of the Tibetan Plateau

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Models for Tibetan Plateau Uplift
  5. 3. Geology of the Himalayan Fold-Thrust Belt and Tibetan Plateau
  6. 4. Geophysical Constraints on Sub-Tibetan Lithospheric and Crustal Structure
  7. 5. Shortening in the Himalayan Fold-Thrust Belt
  8. 6. An Integrated Model for Uplift of the Tibetan Plateau
  9. 7. Implications of the Model
  10. 8. Conclusions
  11. Acknowledgments
  12. References

[43] In the following, we present a hypothetical, modified version of the crustal underthrusting model (Figure 3) that can accommodate the kinematic constraints and shortening estimates from balanced cross sections in the Himalayan fold-thrust belt (Figures 58), recent revisions in our understanding of the origin of Greater Himalayan rocks, and the available geophysical data from the Tibetan Plateau (Figures 3 and 4). Figure 5 illustrates the geometrical constraints on possible kinematic models for the Tibetan Plateau. Analog modeling provides additional insight into the machinery of the Indo-Eurasian collisional process [Chemenda et al., 2000].

[44] The Greater Indian lithosphere subducts at a low angle beneath the Lhasa terrane and continues subhorizontally at least as far north as the Banggong suture zone (Figure 3b). Eurasian lithosphere subducts at a moderate angle to a depth of ∼300 km beneath the northern part of the Plateau. The intervening region in the upper mantle between approximately 33°N and 35°N is occupied by hot asthenosphere (Figures 2 and 3b) [Owens and Zandt, 1997; Kosarev et al., 1999; Wei et al., 2001; Kind et al., 2002]. The tomographic images of Van der Voo et al. [1999] combined with the broadband seismology results suggest that the total length of subducted Greater Indian lithosphere (beneath the Plateau and detached in the mantle) is on the order of 600–900 km (Figures 4 and 5). The older, Neotethyan oceanic slab is depicted as a vertical planar domain at depths of 1000–2000 km in the mantle at latitude 20°N (Figures 4 and 9f).

image

Figure 9. Kinematic history of the crust, lithosphere, and upper mantle in the India-Eurasia collision zone since 55 Ma, beginning with present-day interpretation shown in Figure 3b. Upward pointing arrow tracks the northern edge of the hypothesized Greater Indian lower crust. Abbreviations not explained in legend are as follows: FTF, frontal thrust fault (equivalent to today's Main Frontal thrust, MFT); ISZ, Indus-Yalu suture zone; BSZ, Banggong suture zone; KF, Kunlun fault; PQT, Plio-Quaternary Tibet [Tapponnier et al., 2001].

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image

Figure 9. (continued)

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[45] The length of the GILC beneath the Plateau is predicted to be at least ∼650–700 km, consistent with the hypothesis that its length should approximately equal the shortening in the Himalayan fold-thrust belt (Figure 5, point 8). This slab of crust is probably a composite of Archean Indian cratonic lower crust (Figure 5, point 9) and late Proterozoic-Cambrian lower crust of Greater Himalayan affinity (Figure 5, point 10). The length of Greater Himalayan lower crust is expected to be on the order of several hundred kilometers (Figure 5, point 10), which is consistent with the length of the detached Greater Indian lithosphere slab (anomaly 1). We therefore suggest that this detached slab mainly consists of Greater Himalayan lithosphere.

[46] Combining some reasonable assumptions about the motion of the Neotethyan slab and kinematic information from the Himalayan fold-thrust belt, we can constrain the tempo of Neotethyan slab break-off, Greater Indian subduction, break-off of the Greater Indian lithosphere slab, and continued insertion of the GILC beneath the Tibetan Plateau (Figure 9). Following Van der Voo et al. [1999], we assume that the mantle has not imparted any significant horizontal velocity to the Neotethyan slab. Thus, the slab is shown plummeting vertically from the upper mantle, with its lower tip fixed since 40 Ma at its present latitude (Figure 9). Recalling that the location of the Indus-Yalu suture zone at the onset of collision was 6.5°N ± 2.5° [Klootwijk et al., 1985; Van der Voo, 1993], we position the Neotethyan slab at 55 Ma in the upper mantle between depths of 350 km and 1200 km with its lower end at 12°N and its upper end projecting to the surface at the Indus-Yalu suture zone. This configuration imparts a 55° northward dip to the slab (Figure 9a).

[47] Subsequent time steps in Figure 9 are calibrated according to the rates of northward migration of the Indus-Yalu suture zone, the southern tip of the Indian continent, and the rate of subduction of Greater Indian lithosphere as shown in Figure 6. These diagrams show the Neotethyan slab breaking off after it had been rotated into a vertical orientation, perhaps in latest Eocene time (Figures 9b and 9c). A minimum age for the break-off of the Neotethyan slab can be calculated by dividing the present offset between the Neotethyan slab and the northern limit of inferred GILC (∼1650 km) by the rate of northward convergence between India and Eurasia (50–55 mm/yr) (Figure 5, point 11), which yields a minimum age range of 33–30 Ma for the break-off event. There is little constraint on the maximum age of Neotethyan slab break-off. It is plausible that break-off occurred immediately after the initial impingement of Greater India against Eurasia, and that the slab rotated freely into a vertical orientation.

[48] Following Davies and von Blanckenburg [1995], several workers have suggested that initial exhumation of Himalayan eclogites was associated with isostatic rebound of partially subducted Indian crust during and immediately after the break-off event [Chemenda et al., 2000; O'Brien et al., 2001; Kohn and Parkinson, 2002]. This would imply a late Eocene-Oligocene maximum age for this event, based on the ages of the eclogites [Tonarini et al., 1993; de Sigoyer et al., 1997]. The minimum age of break-off of the Greater Indian lithosphere slab can be calculated by dividing the offset between it and the northern end of inferred GILC (∼500 km) by the convergence rate between India and Eurasia, yielding an age of 9–10 Ma (Figure 5, point 6). Greater Indian slab break-off probably commenced sometime after ∼20 Ma and was complete by ∼10 Ma (Figures 9d and 9e). Modeling by Chemenda et al. [2000] suggests that break-off of the Greater Indian slab involved progressive delamination followed by complete detachment. Alternatively, Greater Indian lithosphere could have collided with Eurasian lithosphere beneath the central part of the Plateau (near the Banggong suture zone) at ∼20 Ma, causing both to thicken and become gravitationally unstable. Greater Indian lithosphere began to delaminate and eventually broke off, while Eurasian lithosphere began to subduct southward. In any case, ongoing northward migration of the remainder of Greater India has overrun the upper end of the detached Greater Indian lithospheric slab at the longitude of western Nepal (Figures 9e and 9f) [Van der Voo et al., 1999].

[49] Insertion of Greater Indian mantle lithosphere and lower crust beneath the Tibetan Plateau could not have taken place unless this region had been evacuated of Eurasian lithosphere. Removal of Eurasian lithosphere may have been facilitated by Late Cretaceous-Paleocene low-angle northward subduction of the Neotethyan slab prior to initial impingement of Greater India [Ding et al., 2002]. In this model, the widespread Paleocene ignimbrites (Linzizong Formation) in the Lhasa terrane and similar rocks in the southern part of the Qiangtang terrane were produced during a regional “flare-up” in response to southward rollback and steepening of a formerly flat Neotethyan oceanic slab, analogous to models for mid-Cenozoic ignimbrites in the North American Cordillera [e.g., Coney and Reynolds, 1977; Constenius, 1996]. The Ding et al. [2002] model suggests that Eurasian lithosphere was either not present or greatly attenuated beneath the southern part of the Tibetan Plateau at the onset of the Himalayan orogenic event (Figure 9a). One or more segments of Eurasian lithosphere, too small to be detected by present tomographic models, also could have been removed by delamination during early-middle Tertiary shortening of Tibetan upper crust [Tapponnier et al., 2001]. The contrasting behavior of the relatively young (Paleozoic and Mesozoic) Eurasian lithosphere and the subhorizontally subducting Greater Indian lithosphere may have been promoted by density differences. Archean cratonic lithosphere, like that hypothesized to lie beneath the southern part of the Plateau, is thought to be less dense than younger lithosphere owing to long-term melt extraction [Jordan, 1978]. Stripped of its lighter crust, such lithosphere may have near neutral buoyancy with respect to the asthenosphere. Beneath the northern Plateau, younger, denser Eurasian lithosphere (and eventually Greater Himalayan lithosphere as well) may have been more susceptible to gravitational foundering.

[50] Another factor that may influence whether delamination takes place under Tibet is the well-documented difference in mantle temperatures between the northern and southern parts of the Plateau and its effects on phase transitions. Petrologic and thermal modeling [LePichon et al., 1997] suggests that underthrust GILC would have become eclogitized at depths of 60–75 km, but thermal relaxation within ∼20 Myr would have converted GILC back into granulite that was too light to sink into the mantle. On the other hand, thermal modeling by Henry et al. [1997] suggests that GILC has never been in the eclogite stability field. At depths of 80–100 km the gabbro-eclogite phase transition is nearly isothermal at a temperature of ∼500°C. The intermediate-depth earthquakes beneath the Lhasa terrane indicate upper mantle temperatures <400°C, which should inhibit eclogitization. Under northern Tibet, higher mantle temperatures would trigger the phase change at equivalent depths. Any mafic lower crust remaining on top of the southward underthrusting Eurasian lithosphere would eclogitize and increase the negative buoyancy of the slab. We suggest that the petrologic differences between the northern and southern parts of the GILC slab can explain the contrasting behavior of the lithosphere beneath Tibet. The southern part of the GILC slab is underlain by cold cratonic Indian lithosphere that has not been eclogitized, whereas the northern part of the slab was, until mid-Miocene time, underlain by younger, Greater Himalayan, eclogite-prone lithosphere. The documented presence of eclogites in Greater Himalayan upper crustal rocks supports this view. Thus, conversion of Greater Himalayan lithosphere to eclogite could have driven the Miocene delamination event.

7. Implications of the Model

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Models for Tibetan Plateau Uplift
  5. 3. Geology of the Himalayan Fold-Thrust Belt and Tibetan Plateau
  6. 4. Geophysical Constraints on Sub-Tibetan Lithospheric and Crustal Structure
  7. 5. Shortening in the Himalayan Fold-Thrust Belt
  8. 6. An Integrated Model for Uplift of the Tibetan Plateau
  9. 7. Implications of the Model
  10. 8. Conclusions
  11. Acknowledgments
  12. References

[51] The proposed model is similar in various respects to previously published models involving underthrusting of India beneath Tibet [e.g., Argand, 1924; Powell and Conaghan, 1973; Seeber et al., 1981; Ni and Barazangi, 1984; Zhao and Morgan, 1987; Matte et al., 1997; Chemenda et al., 2000]. The model makes several testable predictions about the history of deformation and magmatism in the Tibetan Plateau. Before discussing these, however, we emphasize that we are not suggesting crustal underthrusting operated alone to thicken the Tibetan Plateau. The pre-Cenozoic history of shortening and stresses transmitted ahead of the underthrusting GILC undoubtedly complicate the tectonic history of Tibetan upper crust. Nevertheless, the fact that Himalayan shortening has continued unabated since ∼55 Ma implies that underthrusting of the GILC is one of the few processes in the Himalayan-Tibetan orogenic system that may have operated predictably. In the following, we briefly outline seven predictions that derive from the crustal underthrusting model.

  1. The similarity of the width of the Tibetan Plateau to the amount of shortening in the corresponding portion of the Himalayan fold-thrust belt is circumstantial evidence in favor of the crustal underthrusting model (Figure 5, point 8). One obvious explanation is that the width of the Plateau faithfully reflects the presence of the underthrust slab of GILC. If this is the case, then the cumulative shortening in the Himalaya at any given time should provide an estimate of the location of the northern edge of the underthrusting slab. Thus, the tip of the slab reached the Banggong suture by ∼20 Ma, the Qiangtang anticlinorium by ∼14 Ma, and the Jinsha suture within the last million years or so (Figures 8 and 9). This aspect of the model predicts that a front of upper crustal strain should have advanced across the Tibetan Plateau as the GILC slab propagated northward. The nature of the strain associated with this front might vary according to the strength of the underthrusting GILC and the location relative to the leading edge of the slab. Directly above the leading edge, we would expect a crustal-scale monocline in the overlying Tibetan crust, keeping in mind the possibility that pre-existing structures could render the recognition of such a monocline difficult. A candidate for a structure affected by the proposed underthrusting GILC is the Qiangtang anticlinorium (Figure 1), which experienced broadening and topographic rejuvenation during the mid-Tertiary [Kapp et al., 2002]. This is not to say that all deformation in the Tibetan Plateau followed an orderly northward march. It is plausible that far-field stresses created contractional strain and strike-slip faulting in advance of the underthrusting GILC slab tip [e.g., Yin et al., 1999; Horton et al., 2002; Robinson et al., 2002; Yin et al., 2002]. A key challenge in proving the validity of the crustal underthrusting model would be to distinguish between strain that resulted directly from crustal underthrusting and strain that resulted from the propagation of stresses far in front of the GILC.
  2. The model predicts two slab break-off events. The first involved Neotethyan lithosphere and occurred during late Eocene-Oligocene time. This prediction is consistent with the timing of Neotethyan slab break-off proposed by previous workers based on petrologic considerations [O'Brien et al., 2001; Kohn and Parkinson, 2002], but relies on an independent, geometrically and kinematically reasonable reconstruction. The second predicted break-off event involved mainly Greater Himalayan lithosphere and took place during mid- to Late Miocene time. A similar break-off event was predicted by Chemenda et al. [2000] during late Oligocene-early Miocene time. Crustal tectonic events that might correlate with these break-off events include Eocene-Oligocene eclogite exhumation by crustal-scale duplexing in the North Himalayan domes of the Tibetan Himalayan zone [Steck et al., 1998; Chemenda et al., 2000; Kohn and Parkinson, 2002], and regional extension in the southern Tibetan Plateau and high Himalaya and activation of the Main Central thrust during early and mid-Miocene time [Hodges, 2000]. Insofar as break-off events should cause isostatic rebound of the orogenic belt [Davies and von Blanckenburg, 1995], we would expect that the Himalayan fold-thrust belt was driven into a supercritical state promoting forward propagation [Davis et al., 1983] during the proposed break-off events. Because it must have occurred beneath the Tibetan Plateau, the second break-off event is expected to have isostatically increased regional elevation in the Plateau during mid- to late-Miocene time, consistent with evidence for late Miocene regional environmental and climatic changes in central Eurasia and changes in the regional force distribution surrounding the Plateau [Harrison et al., 1992; Molnar et al., 1993].
  3. The crustal underthrusting model has implications for the distribution of mafic Indian cratonic lower crust beneath the Tibetan Plateau. The amount of shortening of the Lesser Himalayan and Subhimalayan zones (∼300 km), which were deposited on cratonic India, should approximately equal the length of Indian cratonic lower crust that has been inserted beneath the Plateau. This suggests that strong, mafic, Archean lower crust of cratonic India extends ∼300 km to the north of the Indus-Yalu suture zone (Figure 5, point 9), coincident with the distribution of high-velocity lower crust beneath the Lhasa terrane [Owens and Zandt, 1997; Kind et al., 2002]. The remainder of the Plateau is probably underlain by Greater Himalayan lower crust (Figure 3b) composed of weak, felsic metasedimentary rocks that may be close to solidus conditions. This is consistent with the change in the velocity structure in the lower crust north of the Banggong suture zone [Owens and Zandt, 1997; Tapponnier et al., 2001] and crustal anisotropy data that suggest a component of eastward lower crustal flow (Folsom et al., manuscript in preparation, 2002). The presence of strong lower crust beneath the Lhasa terrane may have increased its ability to transmit compressive stresses northward, accounting for the lack of major Cenozoic horizontal shortening in the Lhasa terrane compared to the significant Cenozoic shortening in the Qiangtang and Songpan-Ganzi terranes [Kapp et al., 2000; Yin and Harrison, 2000; Horton et al., 2002]. However, this explanation is only relevant for post-20 Ma deformation in the Plateau, because the insertion of the strong mafic portion of the GILC did not commence until about that time. Restriction of stronger cratonic GILC to the southern part of the Plateau is also consistent with the modern topography of the Plateau, with highest local relief in the south and very low relief in the northern part [Fielding et al., 1994].
  4. The crustal underthrusting model predicts that the elevation of the Tibetan Plateau should have increased in isostatic proportion with the thickness of the GILC slab. If we assume that the crust of the Lhasa terrane was ∼50–55 km thick prior to the Cenozoic orogeny [Murphy et al., 1997], then underthrusting of the GILC slab would have added ∼15–20 km of crust to the southern portion of the Plateau. Addition of this crustal mass without the need for surficial thrusting above the leading edge of the slab would result in the northward propagation of a topographic front without attendant upper crustal shortening and flexural subsidence in a migrating foreland basin system. Noteworthy in this context is the lack of evidence for active, large-scale shortening and flexural subsidence along the northern [Li et al., 1996; Jiang et al., 1999] and eastern [Royden et al., 1997; Kirby et al., 2000] margins of the modern Plateau. Moreover, GPS studies, earthquake seismology, and neotectonic studies along the eastern margin of the Plateau, in the Longmen Shan and Min Shan, indicate that the upper crust is not shortening despite the presence of an immense topographic break along the edge of the Plateau [Chen et al., 2000; Kirby et al., 2000]. Royden et al. [1997] and Clark and Royden [2000] proposed that the ∼3.0 × 106 km2 region of eastward sloping, moderate to high elevation terrain to the east of longitude 96°E is being uplifted from below by eastward flowing lower crust (Figure 10). Similar models involving lower crustal flow have been proposed for other large mountain belts comprising mid- to lower crustal rocks [e.g., Hodges and Walker, 1992; Wernicke and Getty, 1997]. We suggest that lower crustal flow beneath the region to the east of the Plateau is driven by the insertion beneath Tibet of the GILC (Figure 10), which is essentially a geographic modification of the “hydraulic piston” model of Zhao and Morgan [1987]. Our model is consistent with results of recent GPS studies that show eastward and southeastward rotation of surficial velocity vectors [Wang et al., 2001], suggesting that GILC underthrusting and lower crustal lateral extrusion are not completely coupled with deformation in the Tibetan upper crust [Holt, 2000]. Insofar as the present volume of crust beneath the Tibetan Plateau west of 96°E can be accounted for by our model, the volume of excess crustal material that has been extruded eastward and southeastward [Clark and Royden, 2000] implies that our estimates of Greater Indian additions to Tibet, and/or the pre-Cenozoic crustal thickness of Tibet, may be bare minima.
  5. The underthrusting model suggests that east-west extension in the crust of the Tibetan Plateau may be a result of the need for Tibetan crust to stretch in this direction in order to accommodate the insertion of the GILC slab. If the slab is thickest and widest toward the south, then we would expect that east-west extension should be greatest in the southern part of the Plateau. A careful inspection of the digital elevation model of Fielding et al. [1994] strongly suggests that this is the case (Figure 2). The most prominent north-south striking grabens in the Plateau are those that rim its southern edge, just north of the South Tibetan detachment system (e.g., the Thakkhola and Yadong-Gulu grabens; Figure 1). North of the Banggong suture, evidence for east-west extension, though still present, is much more subdued [Yin, 2000]. This mechanism for east-west extension is also consistent with the radial shortening directions in the Himalayan arc [Molnar and Lyon-Caen, 1989].
  6. The reconstruction shown in Figure 9 has bearing on interpretations of the Tertiary magmatism in the Tibetan Plateau. The model satisfies existing geochronologic constraints on the Gangdese arc by predicting that calc-alkaline magmatism should have persisted in the southern Plateau until late Eocene time, when the Neotethyan slab broke off (Figures 9b and 9c). High-potassium magmatism associated with rapid convective flow in the mantle and melting of metasomatized upper mantle is expected in the wake of slab break-off events [Davies and von Blanckenburg, 1995]. Such a process could have pooled magmas at the base of the crust and produced the outbursts of Eocene-Oligocene high-potassium volcanic rocks in the central part of the Plateau (Figure 1). It has also been suggested that the Banggong suture zone was partially reactivated during the mid-Tertiary [Yin and Harrison, 2000], possibly generating Eocene-Oligocene melts in central Tibet [Hacker et al., 2000; Kapp et al., 2000; Tapponnier et al., 2001; Ding et al., 2002]. A second phase of high-potassium magmatism during Miocene-Quaternary time might be explained by the break-off and foundering of Greater Himalayan lithosphere and the onset of southward subduction of Eurasian lithosphere (Figures 9d–9f) [Tapponnier et al., 2001].
  7. Tapponnier et al. [2001] highlighted the potential significance for growth of the Tibetan Plateau of late Miocene-Quaternary crustal shortening in the region to the northeast of the Kunlun fault and associated southward subduction of Eurasian lithosphere beneath the Plateau. These processes are analogous to the crustal shortening and underthrusting that have taken place along the Himalayan margin of the Plateau since Eocene time. If Eurasian lithosphere were to peel back northward, additional space would be created for further northward insertion of GILC. In this sense, perhaps the tectonic processes operating between the Kunlun fault and the northern margin of the Nan Shan thrust belt are analogous to those which operated in the upper mantle and crust of the Tibetan Plateau prior to insertion of the GILC slab.
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Figure 10. Schematic map showing the hypothetical eastward and southeastward flow (arrows) of lower crustal material into the region east of the Tibetan Plateau, with the 1.5 km elevation contour shown for present conditions, after Clark and Royden [2000]. The Tarim and Sichuan basins are underlain by strong rigid lower crust (black), whereas the lower crust north and south of the Sichuan basin is relatively weak. Crustal flow is driven by insertion of the slab of GILC (light gray). Portion of the GILC that is predicted to be composed of strong, cratonic Indian lower crust is shown by diagonal ruling. Present location of the Indus-Yalu suture zone is shown by heavy line labeled ISZ. Present coastline is kept fixed. Approximately 150 km of shortening in the upper crust of Tibet is included.

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8. Conclusions

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Models for Tibetan Plateau Uplift
  5. 3. Geology of the Himalayan Fold-Thrust Belt and Tibetan Plateau
  6. 4. Geophysical Constraints on Sub-Tibetan Lithospheric and Crustal Structure
  7. 5. Shortening in the Himalayan Fold-Thrust Belt
  8. 6. An Integrated Model for Uplift of the Tibetan Plateau
  9. 7. Implications of the Model
  10. 8. Conclusions
  11. Acknowledgments
  12. References

[52] Our main conclusions are as follows:

  1. Reconstruction of pre-Himalayan structural architecture of Greater India indicates that the northern margin of Gondwana constituted the Lesser, Greater, and Tibetan Himalayan terranes. These terranes were underlain by a composite lower crust that included Archean and early Proterozoic cratonic Indian lower crust and late Proterozoic-early Paleozoic Greater Himalayan lower crust. This Greater Indian lower crust has been stripped of its supracrustal sedimentary and metasedimentary cover during the Cenozoic orogenic event, and underthrust beneath the Tibetan Plateau.
  2. Himalayan shortening matches the width of the Tibetan Plateau north of the Indus-Yalu suture zone. The similarity in these two quantities suggests that a northward tapering slab of Greater Indian lower crust has been underthrust beneath the Tibetan crust. Insertion of the slab probably increased the thickness of Tibetan crust by up to ∼20 km beneath the Lhasa terrane. The tempo of Himalayan thrusting provides (1) an estimate of the position of the northern edge of the slab through time, and (2) a chronometer of events in the upper mantle beneath Tibet. Two key predictions of our model are break-off of the Neotethyan oceanic slab at ∼45–35 Ma, and delamination and break-off of a several hundred kilometer long slab of Greater Indian lithosphere between ∼20 and 10 Ma.
  3. The crustal underthrusting model helps to explain the distribution and cause of east-west extension in the Tibetan Plateau. Insertion of a tapering (both vertically and horizontally) slab of Greater Indian lower crust beneath the Plateau forced the Tibetan upper crust to stretch laterally to accommodate the excess crustal mass. This suggests that extension should be greatest in the southern part of the Plateau, and that the age of east-west extension should decrease northward across the Plateau. The region of maximum east-west extension should also correlate with the predicted distribution of strong, relatively thick, cratonic Indian lower crust, which should extend as far north as the Banggong suture zone.
  4. Cenozoic magmatism in the Tibetan Plateau spans the last 45 Myr. We tentatively note that the ages and spatial distributions of these rocks crudely correlate with mantle upwelling events associated with predicted Neotethyan and Greater Indian lithospheric slab break-off events.
  5. Several processes have collaborated to produce the high elevation of the Tibetan Plateau, including Mesozoic and Cenozoic upper crustal shortening in Tibetan terranes, Cenozoic upper crustal shortening in the Himalaya, crustal thickening by northward underthrusting of Greater Indian lower crust, and generally eastward flow of lower crustal material that thickened during the Mesozoic and early Tertiary and was subsequently displaced by the underthrusting Greater Indian lower crust. In addition, since late Miocene time, southward underthrusting of Eurasian lithosphere and associated upper crustal shortening has begun to add a half-million square kilometer region to the northeastern part of the Plateau [Tapponnier et al., 2001]. Although the regional pattern of deformation in the Tibetan Plateau seems chaotic, incorporation of the temporal predictions made by the crustal underthrusting model may help to identify order and process as more kinematic information is obtained from the Tibetan Plateau.
  6. Based in part on the great breadth of continental orogenic systems (relative to orogenic systems developed between converging oceanic plates), a number of models have suggested that continental crust and mantle lithosphere are weak and incapable of maintaining sharp plate boundaries [e.g., England and Houseman, 1988]. If the general aspects of our model are correct, then the Himalayan-Tibetan orogenic system resembles a flat-slab subduction system, rather than the classic, pure-shear-dominated concept of a collisional orogen [e.g., Dewey and Bird, 1970]. Moreover, the Greater Indian and Eurasian lithospheres beneath the Himalaya and Tibet seem to behave like coherent plates. In this sense, the widespread expanse of intracontinental deformation in central Eurasia owes less to the “softness” of Tibetan crust and mantle lithosphere than it does to the bulk of continental material situated above the subduction system.

Acknowledgments

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Models for Tibetan Plateau Uplift
  5. 3. Geology of the Himalayan Fold-Thrust Belt and Tibetan Plateau
  6. 4. Geophysical Constraints on Sub-Tibetan Lithospheric and Crustal Structure
  7. 5. Shortening in the Himalayan Fold-Thrust Belt
  8. 6. An Integrated Model for Uplift of the Tibetan Plateau
  9. 7. Implications of the Model
  10. 8. Conclusions
  11. Acknowledgments
  12. References

[53] Funding for research that contributed to this paper was provided by the U.S. National Science Foundation (grants EAR-9814060, EAR-0105339, EAR-0125121, and EAR-0207179). Additional support was provided by ExxonMobil and Conoco. We are grateful to Gautam Mitra, Kelin Whipple, and Greg Houseman for thoughtful reviews of the original manuscript. Our understanding of the Tibetan-Himalayan system has been improved by discussions with and the generous provision of preprints by Paul Kapp, Clem Chase, Ofori Pearson, Carmala Garzione, Rainer Kind, Doug Nelson, An Yin, Brian Horton, Heather Folsom, Kevin Furlong, Tom Owens, Matt Spurlin, Ding Lin, Mark Harrison, Jessica D'Andrea, Brad Ritts, and Brad Hacker.

References

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Models for Tibetan Plateau Uplift
  5. 3. Geology of the Himalayan Fold-Thrust Belt and Tibetan Plateau
  6. 4. Geophysical Constraints on Sub-Tibetan Lithospheric and Crustal Structure
  7. 5. Shortening in the Himalayan Fold-Thrust Belt
  8. 6. An Integrated Model for Uplift of the Tibetan Plateau
  9. 7. Implications of the Model
  10. 8. Conclusions
  11. Acknowledgments
  12. References
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