TECTONICS, VOL. 21, NO. 6, 1062, doi:10.1029/2001TC001322, 2002
Implications of shortening in the Himalayan fold-thrust belt for
uplift of the Tibetan Plateau
Peter G. DeCelles, Delores M. Robinson, and George Zandt
Department of Geosciences, University of Arizona, Tucson, Arizona, USA
Received 20 August 2001; revised 13 September 2002; accepted 18 September 2002; published 31 December 2002.
[1] Recent research in the Himalayan fold-thrust belt
provides two new sets of observations that are crucial
to understanding the evolution of the HimalayanTibetan 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
Copyright 2002 by the American Geophysical Union.
0278-7407/02/2001TC001322$12.00
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
I NDEX T ERMS: 8102 Tectonophysics:
Tibetan crust.
Continental contractional orogenic belts; 8110 Tectonophysics:
Continental tectonics—general (0905); 8120 Tectonophysics:
Dynamics of lithosphere and mantle—general; 9320 Information
Related to Geographic Region: Asia; 9604 Information Related to
Geologic Time: Cenozoic; KEYWORDS: collisional orogenic belts,
Tibetan Plateau, Himalaya, lithosphere dynamics, delamination,
thrust belts. Citation: DeCelles, P. G., D. M. Robinson, and G.
Zandt, Implications of shortening in the Himalayan fold-thrust belt
for uplift of the Tibetan Plateau, Tectonics, 21(6), 1062,
doi:10.1029/2001TC001322, 2002.
1. Introduction
[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 mid1970’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
12 - 1
12 - 2
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
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].
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].
[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 contri-
butions 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
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
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 onetenth 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
12 - 3
contributed to growth of the Plateau, are the main subjects of
this paper.
2. Models for Tibetan Plateau Uplift
[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]).
[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
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DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
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.
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 midCretaceous 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
3.1. Himalayan Fold-Thrust Belt
[8] The Himalayan fold-thrust belt stretches for an arclength distance of 2400 km between the Hazara (in the
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
12 - 5
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.
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].
[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,
12 - 6
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
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 lowgrade 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
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
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 IndusYalu 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 SongpanGanzi basin was tectonically collapsed during late Triassicearly 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 PaleozoicMesozoic rocks as low-angle normal faults that were active
during Late Triassic-Early Jurassic time. The footwall rocks
12 - 7
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]. Panafricanage 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
12 - 8
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
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 PlioQuaternary 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 midCretaceous 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 PlioQuaternary 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 rightlateral 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
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
12 - 9
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 EoceneOligocene 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
[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
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.
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
12 - 10
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
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].
[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 IndusYalu 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 eastwest 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
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
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
12 - 11
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.
[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
[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]. Pub-
12 - 12
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
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.
lished 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.
[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.
[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
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
12 - 13
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.
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 short-
ening 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.
12 - 14
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
Table 1. Shortening Estimates in the Himalayan Fold-Thrust Belt
Reference in
Figures 1 and 7
Reference
Location
Structural Boundariesa
Tectonostratigraphic Zonesb
Shortening, km
1
2
3
4
5
6
7
8
9
10
11
12
13
14
Coward and Butler [1985]
Searle [1986]
Searle et al. [1997]
Srivastava and Mitra [1994]
Srivastava and Mitra [1994]
DeCelles et al. [1998a]
Murphy and Yin [2003]
DeCelles et al. [2001]
DeCelles et al. [2001]
Ratschbacher et al. [1994]
Schelling [1992]
Schelling [1992]
Schelling and Arita [1991]
Schelling and Arita [1991]
Pakistan
India, Zanskar and Ladakh
India, Zanskar and Ladakh
India, Kumaon and Garhwal
India, Kumaon and Garhwal
Western Nepal
Tibet, Mt. Kailas region
Western Nepal, Seti River
Western Nepal, Seti River
Tibet, north of Arun River
Central Nepal
Central Nepal
Far Eastern Nepal
Far Eastern Nepal
MMT-MFT
ISZ-MMT
ISZ-MCT
STDS-MCT
MCT-MFT
MCT-MFT
ISZ-STDS
STDS-MCT
MCT-MFT
ISZ-MCT
STDS-MCT
MCT-MFT
STDS-MCT
MCT-MFT
GHZ, LHZ, SHZ
THZ
THZ
GHZ + Almora thrust sheet
LHZ, SHZ
LHZ, SHZ
THZ + Indus Suture
GHZ + Dadeldhura thrust sheet
LHZ, SHZ
THZ
GHZ
LHZ, SHZ
GHZ
LHZ, SHZ
470 km
126 km
150 – 170
193 – 260 km
161 km
228 km
176 km
131 – 206 km
287 km
133 – 139 km
140 – 210 km
70 km
140 – 175 km
45 – 70 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.
[41] Figure 7 provides an along-strike comparison of the
available estimates of shortening in the Himalayan foldthrust 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
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.
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
12 - 15
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.
6. An Integrated Model for Uplift
of the Tibetan Plateau
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.
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
[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 5 – 8), 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).
[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,
12 - 16
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
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].
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°
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
12 - 17
Figure 9. (continued)
[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 IndusYalu 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 orienta-
tion, 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.
12 - 18
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
[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 breakoff 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
[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
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
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 EoceneOligocene 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
12 - 19
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
12 - 20
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
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 eastwest 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].
Figure 10. (opposite) 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.
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
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 highpotassium 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 EoceneOligocene 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.
8. Conclusions
[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 Proterozoicearly 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
12 - 21
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-sheardominated 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
12 - 22
DECELLES ET AL.: HIMALAYAN-TIBETAN OROGENY
than it does to the bulk of continental material situated
above the subduction system.
[53] Acknowledgments. Funding for research that contributed to this
paper was provided by the U.S. National Science Foundation (grants EAR9814060, 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.
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