Oligocene–Miocene Kailas basin, southwestern Tibet: Record of
postcollisional upper-plate extension in the Indus-Yarlung suture zone
P.G. DeCelles†, P. Kapp, J. Quade, and G.E. Gehrels
Department of Geosciences, University of Arizona, Tucson, Arizona 85721, USA
“…a sunlit temple of rock and ice. Its remarkable structure, and the peculiar harmony
of its shape, justify my speaking of Kailas as the most sacred mountain in the world.”
—A. Gansser, 1939
ABSTRACT
The Kailas basin developed during late
Oligocene–early Miocene time along the
Indus-Yarlung suture zone in southwestern
Tibet. The >2.5-km-thick basin-filling Kailas
Formation consists of a lower coarse-grained
proximal conglomerate and more distal fluvial sandstone member, a lacustrine shale
and sandstone member, and an upper redbed clastic member. Felsic tuffs and trachyandesite layers are locally present. Detrital
and igneous zircon U-Pb ages indicate deposition of most of the Kailas Formation between
ca. 26 and 24 Ma. The Kailas Formation was
deposited by alluvial-fan, low-sinuosity fluvial, and deep lacustrine depositional systems
in buttress unconformity upon andesitic volcanic (ca. 67 Ma) and granitoid (ca. 55 Ma)
rocks of the Gangdese magmatic arc. Abundant organic material, fish and amphibian
fossils, and sparse palynomorphs suggest
that Kailas lakes developed in a warm tropical climate, quite different from coeval basins in central Tibet, which formed at high
elevation in a dry climate. Provenance and
paleocurrent data indicate that the bulk of
the Kailas Formation was derived from the
northerly Gangdese magmatic arc (Kailas
magmatic complex). Only during the latest
stages of basin filling was abundant sediment
derived from the southerly Tethyan Himalayan thrust belt in the hanging wall of the
Great Counter thrust. Kailas basin stratigraphy resembles a classic lacustrine sandwich
and is most consistent with deposition in an
extensional or transtensional rift that developed along the suture zone some 30 m.y. after
the onset of Indo-Eurasian intercontinental
collision. Correlative coarse-grained syntectonic strata similar to the Kailas Formation crop out along a >1300 km length of the
†
E-mail: decelles@email.arizona.edu
Indus-Yarlung suture zone, suggesting that
the basin-forming mechanism recorded by the
Kailas Formation was of regional significance
and not exclusively related to local kinematics
near the southeastern end of the Karakoram
fault. We propose that extension of the southern edge of the Eurasian plate was caused
by southward rollback of underthrusting
Indian continental lithosphere, followed by
slab break-off. Alternating episodes of hard
and soft collision, associated with regional
contraction and extension, respectively, in
the Tibetan-Himalayan orogenic system may
have been related to changing dynamics of
the subducting/underthrusting Indian plate.
INTRODUCTION
The Indus-Yarlung suture zone formed when
the Indian continental landmass collided with the
southern flank of Eurasia during late Paleocene–
early Eocene time (Besse et al., 1984; Garzanti
et al., 1987; Leech et al., 2005; Green et al.,
2008). Along much of the length of the suture
zone, rocks of the northern (Tethyan) Himalayan
thrust belt and ophiolitic mélange were juxtaposed by the south-dipping Great Counter thrust
against rocks of the Gangdese magmatic arc
complex (Heim and Gansser, 1939; Burg et al.,
1987; Yin et al., 1999; Murphy and Yin, 2003).
Resting unconformably upon the Gangdese arc
rocks in the footwall of the Great Counter thrust
is a several-kilometer-thick middle Cenozoic
alluvial-fluvial-lacustrine deposit (Heim and
Gansser, 1939) referred to by several different
names locally along the suture zone. Aitchison
et al. (2002) proposed the umbrella stratigraphic
term “Gangrinboche conglomerates” to include
all of these sparsely dated units along ~1300 km
of the suture zone. The tectonic significance
of the Gangrinboche conglomerates remains a
fundamental problem in understanding the postcollisional history of the suture zone. Indeed,
based partly on interpretation of these coarse-
grained deposits, some authors (Aitchison
et al., 2002, 2007) have proposed that the IndoEurasian collision did not begin until middle
Cenozoic time.
In this paper, we report the results of an investigation of the Oligocene–Miocene Kailas
Formation (Cheng and Xu, 1986) along the
Indus-Yarlung suture in southwestern Tibet (Fig.
1A); this unit is the type example of the Gangrinboche conglomerates. The Kailas Formation
is of interest because it consists of a complex of
coarse- to fine-grained clastic strata more than
4 km thick (Heim and Gansser, 1939; Gansser,
1964), representing a basin of unknown tectonic
affinity that developed ~30 m.y. after the putative
onset of Indo-Eurasian intercontinental collision
(Garzanti et al., 1987), in a region that otherwise would be expected to have been structurally elevated and deeply eroded as the collision
continued. The enigma of the Kailas Formation
is heightened by the fact that a major portion of
it consists of deep-water lacustrine deposits. We
address the tectonic setting and paleogeography
of the Kailas basin, and its implications for the
postcollisional tectonics of the Indus-Yarlung
suture and southern Tibet. The data we present
indicate that the Kailas Formation accumulated
in an extensional basin, raising the prospect that
upper-plate extension was associated with southward rollback of the underthrusting Indian plate
and/or transtension along the early Karakoram
fault. In either case, these results present new
details about the postcollisional history of this
archetypal suture zone that are not explained by
existing geodynamic models.
GEOLOGICAL SETTING
The Tibetan Plateau and its southward-flanking
Himalayan rampart have developed in the context
of northward subduction of Indian Neotethyan
lithosphere beneath the Eurasian plate, climaxing with the early Cenozoic collision between
the two continental landmasses (Argand, 1924;
GSA Bulletin; July/August 2011; v. 123; no. 7/8; p. 1337–1362; doi: 10.1130/B30258.1; 18 figures; 4 tables; Data Repository item 2011057.
For permission to copy, contact editing@geosociety.org
© 2011 Geological Society of America
1337
DeCelles et al.
76°E
A
80°
84°
0
88°
ag h
Alt y n T
Qa
lt
fau
ida
m
KF
Qian
g
S
IY
76°
terra
North
China
Xi
an
gs
hu
ihe
ne
BSZ
Legend
Folds
Him T
ala
yan
80°
Suture zones
Oi
Strike-slip faults
t hrust
MFT
Low-angle normal faults
Elevation ≥ 4.5 km
South
China
au
lt
28°N
belt
India
96°
84°
88°
92°
Abbreviations
IYS: Indus-Yarlung suture zone BSZ: Bangong suture zone
JSZ: Jinsha suture zone
KF: Karakoram fault
MFT: Main Frontal thrust
Elevation ≥ 3 km
Elevation < 3 km
Gangdese magmatic belt
100°
S
AY
ad
Zh
Fig. 5A
a
b
in
as
ST
D
Jia
li f
Lhasa
32°
fa
ult
IYS
Fig.1B
Tertiary graben
terrane
Lhasa
N
B
104°
36°
Qiangtan
g
tang an
ticlino
r iu m
Z
Thrust faults
100°
Qilian
Sha
n
Ba
sin
Kunlun Sha
n
n-Ganzi ter
n
gpa
So
rane
JSZ
32°
Figure 1. (A) Tectonic map of
the Tibetan Plateau and Himalayan thrust belt, after Yin
and Harrison (2000). Labeled
solid circles indicate locations
of other middle to late Cenozoic basins in which paleoaltimetry studies have been
conducted: Z—Zhada Basin
(Saylor et al., 2009); Oi—Oiyug
Basin (Currie et al., 2005); N—
Nima Basin (DeCelles et al.,
2007b); T—Thakkhola graben (Garzione et al., 2000a).
Rectangle indicates location of
map shown in part B. (B) General geological map of southwestern Tibet and adja cent
portion of Himalayan thrust
belt, from Murphy and Yin
(2003). Major faults in the Himalayan thrust belt include the
Great Counter thrust (GCT),
Main Central thrust (MCT),
Dadeldhura thrust (DT),
Jajarkhot thrust (JT), Main
Boundary thrust (MBT), Main
Frontal thrust (MFT), and the
South Tibetan detachment
(STD). Other abbreviations:
GM—Gurla Mandatha; D—
Daulaghiri. Line across Ayi
Shan (AYS) in northwestern
suture zone is location of cross
section shown in Figure 5A.
96°
100 200 km
Tarim Basin
36°
92°
Tibetan
Karakoram
P
(Geolog
fault
y not lateau
sho
wn
Mt. Kailas
)
(6714 m)
32°N
Te
thy Fig. 2
an
GM
GC
T
MCT
30°N
Him
a
DT
ST
D
JT
Him
ala
yan
0
M
FT
for
elan
d bas
in
120
MB
T
laya
D
MC
T
28°N
240
km
80°E
84°E
Miocene-Pleistocene basin fill
Paleozoic-Eocene Tethyan Sequence
Miocene-Pliocene Siwalik Group
Neoproterozoic-Cambrian Greater Himalayan Sequence
Oligocene-Miocene Kailas Fm.
Paleoproterozoic-Neoproterozoic Lesser Himalayan Sequence
Cretaceous-Eocene Gangdese
arc and magmatic complex
Jurassic-Cretaceous ophiolitic rocks
in Indus Yarlung suture zone
1338
82°E
Strike-slip fault
Detachment fault
Normal fault
Thrust fault
Geological Society of America Bulletin, July/August 2011
Oligocene–Miocene Kailas basin, southwestern Tibet
Powell and Conaghan, 1973; Molnar and Tapponnier, 1975; Tapponnier and Molnar, 1977;
Allégre et al., 1984; Garzanti et al., 1987; Dewey
et al., 1988). The timing of initial collision remains a topic of debate (for a discussion, see
Aitchison et al., 2007), but most workers place
the event between ca. 55 and 50 Ma, at least in the
northwestern syntaxial region (Besse et al., 1984;
Patriat and Achache, 1984; Garzanti et al., 1987;
Leech et al., 2005; Green et al., 2008). As pointed
out by Rowley (1998), Ding et al. (2005), and
Aitchison et al. (2007), however, the timing of
initial collision in the central and eastern parts
of the orogenic system remains weakly constrained, and, in any case, it would be surprising
if the collision happened simultaneously along
the entire Himalayan orogenic belt.
The geology of southwestern Tibet includes
four major tectonic features: the Gangdese magmatic arc, Indus-Yarlung suture, Karakoram fault,
and Tethyan Himalayan thrust belt (Fig. 1). The
Gangdese arc is a complex of calc-alkaline batholiths and related volcanic and volcaniclastic rocks
that formed as a Cordilleran-style magmatic arc
along the southern flank of the Lhasa terrane in
response to subduction of Tethyan oceanic lithosphere from Cretaceous to Eocene time (Allégre
et al., 1984; Pan, 1993; Kapp et al., 2007; Pullen
et al., 2008); igneous activity within the arc
continued during the early stages of the IndoEurasian collision. The feature that separates the
Gangdese arc and related forearc rocks from the
Himalayan thrust belt is the Indus-Yarlung suture. Within the suture zone, ophiolitic slivers and
sedimentary- and serpentinite-matrix mélanges
structurally overlie Tethyan Himalayan strata
in the south and Gangdese forearc strata in the
north (Fig. 1B; Gansser, 1980; Tapponnier et al.,
1981; Burg and Chen, 1984; Girardeau et al.,
1984; Ratschbacher et al., 1994; Yin et al., 1994,
1999; Murphy and Yin, 2003; Ding et al., 2005).
Where dated, ophiolitic fragments are Jurassic–
Cretaceous (Gopel et al., 1984; Zhou et al., 2002;
Malpas et al., 2003; Miller et al., 2003; Ziabrev
et al., 2003). These ophiolitic rocks were obducted onto the northern Indian margin during
Late Cretaceous–Paleocene (Burg and Chen,
1984; Girardeau et al., 1984; Searle et al., 1987;
Malpas et al., 2003; Ding et al., 2005) and possibly Eocene (Davis et al., 2002) time. The Tethyan
Himalaya is composed of generally southwardverging thrust sheets of unmetamorphosed Paleozoic and Mesozoic sedimentary rocks (Burg
et al., 1987; Ratschbacher et al., 1994; Murphy
and Yin, 2003), locally disrupted by large domal
structures involving mylonitic orthogneiss and
paragneiss that were exhumed from midcrustal
depths during middle to late Miocene time (Lee
et al., 2000, 2004, 2006; Murphy et al., 2002; Lee
and Whitehouse, 2007). The northward-verging
Great Counter thrust is the northernmost major
thrust system in the Tethyan Himalaya (Heim
and Gansser, 1939; Burg et al., 1987; Ratschbacher et al., 1994; Yin et al., 1999; Murphy and
Yin, 2003). In the Kailas Range (Fig. 1B), the
Great Counter thrust is referred to as the South
Kailas thrust system (Yin et al., 1999) and consists of several thrusts that carry Tethyan sedimentary and low-grade metasedimentary rocks
as well as rocks that are considered to be part
of the Gangdese forearc basin and accretionary
wedge (Yin et al., 1999; Murphy and Yin, 2003).
These rocks are juxtaposed directly against the
Kailas Formation along the southern fringe of
its outcrop (Fig. 2), and the southernmost part
of the Kailas Formation is locally overturned in
the footwall of the fault system (Gansser, 1964;
Murphy and Yin, 2003).
Yin et al. (1994, 1999) reported a second major
thrust fault, which they referred to as the Gangdese thrust, exposed in the Xigaze and Zedong
areas along the eastern part of the suture zone.
There, the fault dips northward beneath, and
helped to structurally elevate, Gangdese batholith
rocks prior to slip on the Great Counter thrust.
Although the Gangdese thrust is not exposed in
the Mount Kailas region, Yin et al. (1999) and
Murphy and Yin (2003) inferred it to be present
in the subsurface and responsible for uplift of the
Kailas magmatic complex during deposition of
the Kailas Formation. Aitchison et al. (2003) disputed the existence of the Gangdese thrust along
the entire suture zone; our surface observations
in the Kailas region are consistent with their interpretation, but also do not rule out the possible
presence of the Gangdese thrust in the subsurface
(Murphy and Yin, 2003).
The Kailas Formation (Cheng and Xu, 1986)
is named for its type section on Mount Kailas
(6714 m), which is also referred to by its Tibetan
name of Gangrinboche. Based on original work
reported by Heim and Gansser (1939), Gansser
(1964) named this unit the Kailas conglomerate; however, conglomerate forms only a fraction
of the unit, so we use the more inclusive Kailas
Formation. Adherents to the Hindu, Buddhist,
Jain, and Bön faiths regard Mount Kailas to be
sacred and officially off limits. We therefore focused on outcrops in deep canyons 20–40 km
west of the mountain. Searle et al. (1990) correlated sparsely dated lithologically similar parts
of the Indus Group in the northwestern Himalaya
with the Kailas Formation. Aitchison et al. (2002)
correlated the Kailas Formation with Upper
Oligocene–Lower Miocene clastic rocks that crop
out along nearly the entire length of the IndusYarlung suture (e.g., the Qiuwu, Dazhuqu, and
Luobusa Formations; see also Yin et al., 1999).
At Mount Kailas, the Kailas Formation rests
unconformably upon the Gangdese batholith
(Gansser, 1964). In the region of our study,
the Kailas Formation rests in buttress unconformity upon andesitic volcanic rocks dated at
66.6 ± 1.26 Ma by U-Pb zircon (Figs. 3 and 4;
sample 527052, Table DR11). These volcanic
rocks are intruded by granite that yielded
a U-Pb zircon age of 55.0 ± 0.8 Ma (Fig. 4;
sample 61052, Table DR1 [see footnote 1])
and a hornblende 40Ar/39Ar isochron age of ca.
45 Ma (Yin et al., 1999). Together, the volcanic
and granitoid rocks are referred to as the Kailas
magmatic complex (Honegger et al., 1982) and
constitute the local manifestation of the Gangdese magmatic arc.
West of the Mount Kailas region, the IndusYarlung suture zone is offset by the dextral
Karakoram fault. Along the southwestern side
of the fault, the Ayi Shan (Fig. 1B; also referred
to as the Ayilari Shan) consists of ca. 50 Ma
granitic orthogneiss that experienced metamorphic zircon growth at upper-amphibolite
conditions between 41 Ma and 31 Ma, and subsequent partial melting and rapid exhumation
in the footwall of the north-dipping Ayi Shan
detachment fault at 26–18 Ma (Fig. 5A; Zhang
et al., 2010). The Great Counter thrust and a
thin erosional remnant of the Kailas Formation in its footwall are also preserved along the
southwestern flank of the Ayi Shan (Fig. 1B;
Murphy et al., 2000). As in the Kailas Range,
the Kailas Formation here rests unconformably
upon rocks of the Gangdese magmatic arc.
Miocene slip on the Karakoram fault offset the
rocks of the Ayi Shan ~65 km to the northwest
relative to the Kailas Range (Murphy et al.,
2000; Valli et al., 2007). Thus, during late
Oligocene–Miocene time, the rocks exposed
today in the Ayi Shan were probably located
directly beneath the western part of the Kailas
Range and were probably experiencing rapid
exhumation in the footwall of the Ayi Shan detachment fault (Fig. 5B; Zhang et al., 2010).
STRUCTURAL SETTING OF THE
KAILAS FORMATION
The Kailas Range trends parallel to the South
Kailas thrust system and the regional trend of
the Gangdese magmatic arc (Fig. 2). Kailas
strata are mostly only slightly tilted toward
the south, but in the proximal footwall of the
South Kailas thrust system, they are folded into
a northward-verging fold pair that parallels the
fault (Fig. 5B; Gansser, 1964). The northern part
1
GSA Data Repository item 2011057, Table DR1:
U-Pb age data from igneous rocks and Table DR2:
U-Pb age data from detrital zircons, is available at
http://www.geosociety.org/pubs/ft2011.htm or by request to editing@geosociety.org.
Geological Society of America Bulletin, July/August 2011
1339
DeCelles et al.
Pgr
81°E
Kv
81°°15’E
81
Kv
Kv
5KR
18
25
A’
Tuffs
24.6 Ma
24.2 Ma
1015
29
13
8KR
46
Kv
OMk
55 Ma
67 Ma
7KR
41
31°°15’N
31
3KR
8
67
80
26
Pgr
33
48
4KR
Q
1KR
12
47
49 71
Kv
Pgr
2KR
Q
OMk
Kv
A
Pgr
ml
Ice/snow
OMk
Q Quaternary deposits
ml
N-Q Neogene-Quaternary deposits
N-Q
OMk Oligocene-Miocene Kailas Formation
Pgr
Paleogene granite
05'N
31°05'N
31
South Kailas
Thrust (GCT)
OMk
Kv Cretaceous volcanic rocks
ml Mélange
81°°15’E
81
81°E
81
Strike and dip of bedding
26
71
4KR
0
2
4
6
8
Mt. Kailas
(Gangrinboche)
6714 m
10 km
Strike and dip of foliation, with plunge of lineation
Thrust fault, barbs in hanging wall
Trace of syncline axis
Minor strike-slip fault
Trace of anticline axis
55 Ma
Trace of measured section
N
Geochronology sample location and age
Figure 2. Geological map of the Mount Kailas area, showing locations of measured sections (e.g., 1KR, 2KR, etc.) and locations from which
samples were obtained for U-Pb zircon dating. Cross section along line A-A′ is shown in Figure 5B.
of the Kailas Formation onlaps volcanic rocks
of the Kailas magmatic complex, as discussed
previously. These relationships are consistent
with the geology directly east near Mount Kailas (Gansser, 1964; Yin et al., 1999; Murphy and
Yin, 2003) and 65 km to the northwest in the
Ayi Shan (Fig. 5A; Zhang et al., 2010), as well
as elsewhere along the suture zone farther east
(Aitchison et al., 2002).
STRATIGRAPHY AND
SEDIMENTOLOGY
Depositional environments in the Kailas
Formation are reconstructed on the basis
of sedimentological observations archived
1340
in ~2500 m of measured stratigraphic sections at eight localities (Figs. 6–8). Sections
were measured using a Jacob staff and tape
measure at the centimeter scale. Paleocurrent
data were collected by measuring 10 imbricated clasts per station, or, in some cases,
the limbs of trough cross-strata (method I
of DeCelles et al., 1983). Samples for palynology, geochronology, and sedimentary
petrology were collected in the context of the
measured sections.
The Kailas Formation is divisible into four
informal members based on lithofacies assemblages and lithological characteristics (Fig. 9).
The basal unconformity is overlain by the
lower conglomeratic member, which grades
laterally down-depositional dip (southwestward) into the fluvial sandstone member. Both
of these units are capped by a major flooding
surface, which marks the base of the lacustrine
member (Fig. 9). The uppermost part of the
Kailas Formation consists of fluvial deposits of
the red-bed member. In the following sections,
we describe the sedimentological features of
these units, and document their mutual stratigraphic relationships. Because most of the
lithofacies in the Kailas Formation have been
widely documented in the sedimentological
literature, we provide only brief descriptions
and include a summary table in which standard
interpretations are listed alongside abridged
descriptions (Table 1).
Geological Society of America Bulletin, July/August 2011
Oligocene–Miocene Kailas basin, southwestern Tibet
Sample 7KR465
Kailas tuff bed
Relative probability
24.1 ± 0.41 Ma
(1.7%)
A
0
35
33
20
40
60
80 Ma
Sample PD4
Kailas tuff bed
31
Age (Ma)
29
27
25
23
21
19
17
B
40
Age (Ma)
36
28
24
16
68
64
Age (Ma)
Sample PD3
Kailas tuff bed
32
20
Mean = 24.60 ± 0.52 Ma [2.1%]
n = 24, MSWD = 2.3, Error bars are 2σ
Sample 61052
Kailas magmatic complex,
Granite
60
56
52
48
44
86
82
78
Age (Ma)
Figure 3. (A) Mount Kailas, viewed from the south. Subhorizontally
stratified rocks holding up the peak and surrounding buttresses are
the Kailas Formation; dark rocks in right foreground are in the
hanging wall of the South Kailas (or Great Counter) thrust; and
lighter-colored rocks in background at head of the major canyon are
granitoid rocks of the Kailas magmatic complex. (B) View toward
the southwest of the buttress unconformity at the base of section
1KR (see Fig. 2 for location). Ellipse indicates person for scale.
Mean = 24.17 ± 0.51 Ma [2.1%],
n = 25, MSWD = 1.5, Error bars are 2σ
Mean = 54.98 ± 0.77 Ma [1.4%],
n = 28, MSWD = 1.06, Error bars are 2σ
Sample 527052
Kailas magmatic complex,
Andesite
74
70
66
62
58
54
Mean = 66.6 ± 1.2 Ma [1.8%],
n = 21, MSWD = 0.53, Error bars are 2σ
50
Figure 4. U-Pb zircon ages from samples of Kailas Formation tuffs
(7KR465, PD3, PD4) and from the Kailas magmatic complex. See
text for discussion and Table DR 1 for data. Note that all ages include both internal and external errors. MSWD—mean square of
weighted deviates.
Geological Society of America Bulletin, July/August 2011
1341
DeCelles et al.
A
NE-SW cross section in the southeastern
Ayi Shan, from Zhang et al., 2010.
SW
GCT
Elevation (km)
6
4
KF
schist
ss
nei a n d
ll g
twa
o
c fo
niti
Mylo
THS
AY
SD
ca. 50 Ma granitic orthogneiss
rapidly exhumed 26–18 Ma
2
0
NE
AYSD
No vertical exaggeration
NE-SW cross section in the Kailas Range
(see Fig. 2 for location)
A
B
55 Ma
Summit elev. Mt. Kailas
67 Ma
Elevation (km)
6
4
A′
GCT
THS
AY
S
D
2
0
Kailas Fm. lacustrine member
Kailas Fm. lower conglomerate, fluvial sandstone and red-bed members
Tethyan Himalayan sequence
Granitic gneiss (in Ayi Shan), granite and andesite (Kailas Range)
Figure 5. (A) Structural cross section across the southeastern end of the Ayi Shan (see
Fig. 1B for location), showing the Ayi Shan detachment fault (AYSD), Karakoram fault
(KF), Great Counter thrust (GCT), and the Tethyan Himalayan sequence (THS) modified
after Zhang et al. (2010). (B) Structural cross-section A-A′ (see Fig. 2 for location) across
the Kailas Range in the study area. Squares indicate locations approximately on line of
section from which samples of the Kailas magmatic complex were collected and dated by
U-Pb on zircon. The topographic surface is from the line of section, and the dashed topographic surface shows maximum elevations within ~2 km of the plane of the cross section.
Also shown is the projected elevation of the summit of Mt. Kailas. Projection of the Zhang
et al. (2010) cross section onto plane of this cross section is referenced to the location of
the Great Counter thrust on both cross sections. Structure in hanging wall of the Great
Counter thrust is schematic.
Lower Conglomerate Member
The lower conglomerate member is 613 m
thick in section 1KR (Fig. 6) and at least 381 m
thick in section 5KR (Fig. 7), which is located
12 km northwest of section 1KR (Fig. 2). This
member consists of an overall upward-fining
sequence of boulder to pebble conglomerate
with minor sandstone interbeds. Volumetrically
dominant lithofacies include clast-supported,
moderately well-organized boulder-cobble
conglomerate that is massive (Gcm), crudely
horizontally stratified (Gch), and commonly imbricated with long-axes transverse to paleoflow
direction (Gcmi, Gchi) (Figs. 6, 7, 10A, and
1342
10B). These deposits are arranged in 1–6-mthick beds, typically with slightly erosional basal
surfaces and finer-grained upper parts. Maximum clast size averages 40–60 cm in the lower
part of section 1KR, and decreases to 30–40 cm
in its upper part. In section 5KR, several-meterthick beds of clast- and matrix-supported, disorganized boulder conglomerate (Gcm and Gmm;
Fig. 10C) are abundant; some of these beds have
average maximum clast sizes greater than 1 m,
with individual clasts >7 m in long dimension.
Lenticular beds of granular to coarse-grained
sandstone characterized by planar horizontal
lamination and trough cross-stratification are locally intercalated within the otherwise unbroken
succession of conglomerate. In the uppermost
part of the lower conglomerate member, thick
covered intervals are underlain by sandy conglomerate and sandstone.
Paleocurrent data from imbricated clasts
(~270 measurements at 27 stations) in the lower
conglomerate member indicate southward
paleoflow directions. These data are consistent
with provenance data (discussed later herein)
that demonstrate derivation from the Kailas
magmatic complex.
The association of lithofacies in the lower
conglomerate member is consistent with deposition in alluvial fans that built southward off of
high, rugged topography underlain by the volcanic and granitoid rocks of the Kailas magmatic
complex. The relatively scarce occurrence of
sediment-gravity-flow facies (Gmm and disorganized Gcm) and the abundance of imbricated
and stratified beds of conglomerate indicate that
most deposition took place in stream flows, perhaps during flash floods. The lateral transition
into amalgamated fluvial sandstones suggests
that these were stream-dominated alluvial fans
(Ori, 1982; Nemec and Steel, 1984; Ridgway
and DeCelles, 1993; Wilford et al., 2005). In
section 5KR, coarse boulder beds are abundant
and indicate that sediment gravity flows were
more important in this part of the Kailas Formation. It is plausible that sections 5KR and
1KR were deposited on different alluvial fans
that were dominated by debris flows and stream
flows, respectively.
Fluvial Sandstone Member
The fluvial sandstone member is at least
310 m thick; its base is not exposed in the
study area, but its top is marked by an abrupt
transgressive surface. This member consists of
stacked and laterally overlapping, broadly lenticular, 0.5–2-m-thick beds of medium- to very
coarse-grained sandstone. Basal contacts of
individual beds are erosional (Fig. 10D), and
upward-fining sequences are common (Fig. 6,
section 2KR). The uppermost parts of many of
these upward-fining sequences consist of darkgray siltstone, and fragments of woody plant
material and fossil logs are abundant. Sedimentary structures include planar horizontal
lamination, trough cross-stratification, and local
ripples. Many beds appear to be massive, and
exposures form steep, high cliffs.
The broadly lenticular geometry and upwardfining grain size trends of sandstone beds in the
fluvial sandstone unit suggest deposition in shallow, laterally unstable channels. Absence of typical fine-grained levee and overbank facies implies
that the fluvial system was dominated by shallow,
laterally mobile, low-sinuosity stream channels.
Geological Society of America Bulletin, July/August 2011
Oligocene–Miocene Kailas basin, southwestern Tibet
250
500
1KR240
240
250
LEGEND
Gcm
Gcmi
Petrographic sample
Sh
1KR486
Gcmi
St
St,Sh
U-Pb detrital zircon age
Gcm
10
Gchi
Paleocurrent direction
480
Sr
240
St
2KR232
n = number of measurements
220
Fossil shell fragments
St,Sh
Gcmi,Gcm
220
St
Sm
Vertebrate fossils
200
Sm
200
Leaf fossils
Gcmi
Covered interval
Grain size key:
180
Gcmi
160
420
Gchi
10
Gcmi
Sp
Gchi,Gcm
10
Gcmi,Gcm
Gcmi
Sp
Gcmi,Gcm
1KR148
Gcmi
140
1KR630
1KR629
1KR624
DZ
Gcmi
380
Gcmi
Gcmi
120
10
620
1KR615
St
Sr
Sm
2KR
N31°11.201’
E81°01.085’
elev. 5474 m
2KR420
DZ
420
St
Fsl
160
Sm
St
Fsl
Sm
St
400
Fsl
N31°11.731’
E81°01.767’
elev. 5334 m
Gcmi
Sm
Sm
Fsl
Sm
1KR
400
Sp
2KR184
180
si fs cs cg
e
at e
er ton
om s
gl nd
on a e
C se s ton
r
s
oa d
C san
ne e
Fi ton
lts
Si
1KR165
DZ
Sm
Sr
Fossil wood fragments
440
Sm
Sm
Bioturbation
460
St
Sm
140
TS
St
Sm
Sr
St
St
Sm
TS
TS
TS
Sm
Fsl
Sm
St
380
St
Sm
Sm
TS
120
Gcmi
St
Gcmi
Gcmi
100
Gcmi
360
1KR353
DZ
10
Gct
Gcmi
60
10
Fsl
St
320
1KR568
St
St
Gcmi
Fsl
Sp
10
St
Gcmi
Gcmi
Gchi
40
Gcmi
Gcm
280
Gcmi
Gcmi
20
Gcmi
10
1KR267
A,B
Gcm
260
Gcmi
0
si fs cs cg
N31°11.525’
E81°02.171’
elev. 4934 m
10
10
1KR255
250
540
Gcm
2KR40
40
Sm
Gcm
Sm
Gcmi
St
Gcm
St
300
Sm
10
520
Gcm
Sm
Fsm
Sm
Fsl
Sm
Fsl
Gcm
Gct,St
Gct,St
St
Gct,Gcm
St
Gcmi
Gcm
St
Gcmi
si fs cs cg
St
St
St
Gcm
Fsl
Sm
10
St
TS
Sm
Sm
Sm
10
Gcmi
10
320
Sm
Gcm
300
Gcmi
TS
2KR330
St
60
560
St,Sh
TS
Sm
Sr
Gcm,Gch
St
Gcm,Gct
10
Gcmi
340
80
580
Sr
Gcmi
Gcmi
St
Gcmi
Gcm
Sm
Sm
St
Fsl
St
Gcmi
360
St
5
340
Gcmi
1KR73
Sp
Gcmi
Gcmi
St
Gcp
80
600
10
Sm
2KR107
DZ
100
10
Gct,Gch
Gcmi
1KR86
St
Gcm
Sm
20
St
Sm
280
Sm
Sm
Sm
Sm
Sm
Fsm
St
Sm
St
St
Sm
Fsl
Sm
Gcmi
St
Gcmi
10
500
St
Gcmi
2KR1
0
si fs cs cg
St
260
Sm
Sm
si fs cs cg
Sm
St
250
si fs cs cg
Figure 6. Logs of measured sections 1KR and 2KR. See Table 1 for lithofacies codes. TS—transgressive surface; DZ—detrital zircon sample.
Geological Society of America Bulletin, July/August 2011
1343
DeCelles et al.
LEGEND
250
3KR
250
N31°14.011’
E80°56.752’
elev. 5147 m
TS
Petrographic sample
240
U-Pb detrital zircon age
Gcm
n = number of measurements
220
Sm,St
240
Sm
Gcmi
Gcmi
Gcm
Sm
Fossil shell fragments
10
Gcm
St
Sr
Sr
Sr
240
3KR227P
3KR226
N31°11.606’
E81°00.881’
elev. 5816 m
4KR
Fsl
461
Sr TS
460
Sm,Sr
TS
Sm,St
220
220
Gcm
Gcm
Sm
Bioturbation
Leaf fossils
4KR
SrFsl
Sr,Sh
Sm,St
Vertebrate fossils
200
TS
St
Sr
Fsl
Paleocurrent direction
Gcm
5KR226
250
4KR210
200
200
Gcmi
5KR195
Fossil wood fragments
Gcmi
10
Gcm
440
Covered interval
Grain size key:
180
180
180
si fs cs cg
Gcm
e
at e
er ton
om s
gl nd
on a e
C se s ton
r
s
oa d
C san
ne e
Fi ton
lts
Si
Sm
Sm
Gcmi
Gcm
160
160
Fsl
160
Gcm
Gcmi
Gcm
5KR145
TS
3KR147
5KR
10
4KR
408DZ
Sm
Sr
St
Sr
Sm
4KR400
400
Sm
St
N31°17.151’
E80°56.449’
elev. 5203 m
Gcm
10
380
Gcm
140
Sm,St
140
4KR139
Gcm
Gcm
Gcm
Gcm
Sh,St
Sh,St
4KR388
Fsm
120
120
Fsl
Fsl
360
Gcm
Gcm
340
St,Sh
Gcm
Gcm
St,Sh
320
Gcm
80
Sh
Gcm
Gcm
Gcm
3KR72P
20
Sm
60
0
Gcm
Gchi
Gcm
Gcm
Gcmi
Gcmi
Gcm
Gcmi
si fs cs cg
N31°17.110’
E81°56.827’
elev. 5078 m
Sm
Sm
10
20
10
10
Gcm
40
St,Sh
Sr
3KR14
3KR10
si fs cs cg
Sr
Sr,Fsl
Sr
Gcm
Gcmi
Gmm
0
4KR301
300
Sm
Sh
40
3KR39
Gcmi
Gmm
250
60
HCS
Sh
TS
Gcm
Gcm
Gcm
Gcm
Gcm
260
320
St
Sm
Sp
Sm
Sr/Fsl
3KR47P
Gcm
TS
Srw
300
Gcm
TS
St
Gcm
280
80
Fsl
Gcm
10
340
St
Sm,Sr
3KR65
Gcm
Gcm
60
100
Sh
HCS
3KR88
Gcmi
5KR325
Sm
4KR355A
TS
Sm
St,Sh
Gcm
80
St
3KR100
100
Gcm
Tuff
360
Gcm
100
Sr
Basalt
Fsl
4KR363
40
Sr,Sh
380
Sr,Fsl
120
Gcm
Gmm
Gcm
Gmm
Gcm
Gmm
Gcm
Gcm
Gmm
Gcm
Gcm
Gmm
Gcm
Gcm
Gcm
St
Sh,St
Gcmi
140
420
St
Sm
TS
20
St
Sh
Sh
Sm
Sm
TS
Fsl Sr
Sh
Fsl
Sm
Gct
Gcm
Sm,St
4KR25
TS
280
St
Sr
St
Sr
St
260
TS
Sr
Srw,
HCS
Srw
si fs cs cg
N31°14.079’
E80°56.862
elev. 4959 m
Sm
Sm
Sh
St
Sr
0
si fs cs cg
N31°11.183’
E81°00.972’
elev. 5469 m
250
St
si fs cs cg
Figure 7. Logs of measured sections 3KR, 4KR, and 5KR. See Table 1 for lithofacies codes. TS—transgressive surface; DZ—detrital zircon sample.
1344
Geological Society of America Bulletin, July/August 2011
Oligocene–Miocene Kailas basin, southwestern Tibet
250
7KR246
240
St
Sm
Sm
St
St
Sm,Sh,
Sr
Sm,Sh,Sr
500
TS
Sh,Sr,Sm
7KR483
480
220
Sm
Sm
Paleocurrent direction
Sm
n = number of measurements
Sm
Fossil shell fragments
720
Tuff
Fsl
Sr
Sr
Fsm
700
Sm
Sh
Sm,Sh
440
Sm,Sh
Sm
Sm
Sm
Sm
Vertebrate fossils
Sm
Sm
Leaf fossils
St,Sm
Fossil wood fragments
Covered interval
680
420
Fsl
400
M
Sm
Sm
Sr,Sm TS
Sh,Sm
Fsl
Sh
Sm
Sr
St
640
Sm
380
12
Sm
Sm
Sh
Fsl
Sh
Tuff
si fs cs cg
660
e
at e
er ton
om s
gl nd
on a e
C se s ton
r
s
oa d
C san
ne e
Fi ton
lts
St,Sh
Grain size key:
Sm
M
Fsm
Si
7KR408
160
7KR129
DZ
Paleosol carbonate
St
7KR177
140
Bioturbation
Fsm
Sm
Sm
180
U-Pb detrital zircon age
Sm
Fsl
460
200
LEGEND
Petrographic sample
740
Sm
7KR465
DZ
Sm
Sm
Fsr,Sr
Sm,Sh,St
Fsl
Sm
750
8KR
Sm
Sm
Sm
Sm
Fsl
N31°11.857’
E81°54.736’
elev. 4943 m
Gcm
Sm
Sm
120
Gcm
360
Gcm
100
7KR
Sm
M
Sh,Sr
Sh,Sr
Sm
Sm
Gcm
600
Sh,Sm
Sh
Gcmi
120
620
Sh,St
340
Sm
Fsm
Sm
Sm
Sm
Sm
Fsm
Sm
7KR847
840
Sm
M
Sm,Sh
Sm
Sm
Fsm
Sm,St
Sm
7KR337
10
N31°11.603’
E81°55.759’
elev. 4948 m
80
580
Sm
Sm
St
320
Sm
St
Sh,St
20
Fsl
Sm
Sm
Sm
Sm
300
560
40
20
Sm
Gcm
Sm,Sh
Sm
St
Sr
St
Sm
Sm
Gcm
0
Sm
si fs cs cg
N31°11.974’
E81°56.625’
elev. 4877 m
280
Sm
Sm
Sr
Fsm
Sm
St,Sm
TS
TS
St,Sh
St,Sh
St,Sh
Sm,Sh
St
540
Fsl
M
Fsl
M
Sm
Sm
40
St
Sr
520
Fsl
Sh,Sr
Sh,St
Sm
St
si fs cs cg
Fsm
Fsl
780
Fsm
Sm
Fsm
7KR252
250
Sr
Sm
Sm
Sm
760
500
si fs cs cg
20
8KR17
Sm
8KR8
Sm
Fsm
750
20
Fsm
Sr
St
260
M
Fsm PS
Fsl
Sm
Sm
M
M
M
M
Sr
800
Fsl
Fsl
Sr,Sh
Sm,Sh
St
M
St
8KR65
60
Sr
Gcm
M
Fsl
St
Sm
Fsl
Sm
Gcm
St
M
8KR
67DZ
820
Sr
60
Fsl
8KR96
8KR 80
76A,B
St
Sm
Sm
Fsl
8KR104
DZ
100
Sm
M
M Fsl
M
M
Sm
si fs cs cg
0
St
St
Sr
St
St
Sr
Sr
Fsm
Sh
22
Sm,St
22
Sr
Sr,Fsm
si fs cs cg
N31°11.983’
E80°54.726’
elev. 4946 m
Figure 8. Logs of measured sections 7KR and 8KR. See Table 1 for lithofacies codes. TS—transgressive surface; DZ—detrital zircon sample.
Geological Society of America Bulletin, July/August 2011
1345
DeCelles et al.
NNE
SSW
8KR
Red-bed member
Lac
ustr
ine
7KR
4KR
mem
ber
500
Vertical scale in m
6KR
3KR
24.1 Ma
si
fs
cs
cs
cg
cg
25.1 Ma
26.6 Ma
5KR
Lacustrine transgressive surface
25.0 Ma
26.1 Ma
si
fs
cs
agm
atic
UNCONFO
Com
RMITY
Fluvial sandstone member
cg
congLower
lo
memmerate
ber
TT
RE
SS
as M
fs
24.7 Ma
1KR
0
Kail
24.2 Ma
24.6 Ma
Sr
2KR
si
BU
24.4 Ma
U-Pb zircon age from tuff
Minimum age from detrital zircon U-Pb
Generalized paleoslope direction
plex
Figure 9. Tentative correlation diagram for measured sections in the Kailas Formation. See Figure 2 for locations of sections. No horizontal
scale is implied. Correlations were established by tracing key beds in the field and on satellite images. Locations of samples for geochronological analyses and generalized paleoslope directions are also shown.
Unfortunately, the massive, cliffy aspect of
the outcrops precluded measurement of paleocurrent indicators. However, it is possible to
trace beds of the fluvial sandstone member in
section 2KR directly northward into the lower
conglomerate member in section 1KR (Fig. 9),
indicating that the former is simply the more
distal equivalent of the latter, perhaps as a fringing braid-plain down depositional dip from the
alluvial-fan system.
Lacustrine Member
The lacustrine member is composed of
black shale and sandstone beds arranged in
upward-coarsening parasequences that are
separated by sharp, flat transgressive surfaces
(Fig. 7, sections 3KR and 4KR; Figs. 11A
1346
and 11B). Individual parasequences range
in thickness from a few meters to >75 m
(Fig. 7). A typical parasequence consists of:
(1) a basal transgressive surface overlain by
laminated black shale (Figs. 11B and 11D);
(2) a series of beds that become thicker and
coarser up section, from centimeter-scale
layers to >1 m thick; (3) sedimentary structures that increase in scale upward within the
sandy part of the parasequence, from small
ripples (both symmetrical and asymmetrical)
and plane beds in the lowest sandstone beds
to hummocky cross-strata and trough crossstrata in the upper parts of the parasequence;
and (4) uppermost pebbly beds in some
parasequences that contain low-angle planar
cross-strata. Channelized lenticular pebbly
sandstone beds with trough cross-strata, and
large-scale low-angle progradational clinoforms are present in the upper parts of some
parasequences (Fig. 12).
Molluscan debris, fragmented fish fossils, and coaly plant material are common
(Fig. 11E). However, bioturbation is rare and
consists of small, irregular tubular burrows.
Transgressive surfaces are typically marked by
rusty, calcite-cemented horizons, often littered
with pebbles, fish and turtle fossil fragments,
and carbonaceous plant material (Figs. 11E
and 11F). A few beds of biotite-bearing tuff
and phlogopite-bearing trachyandesite/basalt
(388 m level of section 4KR, Fig. 7) are present
within the lacustrine member.
Sets of parasequences are stacked vertically
in upward-thickening packages (Fig. 7, section
3KR). The lacustrine member is at least 700 m
Geological Society of America Bulletin, July/August 2011
Oligocene–Miocene Kailas basin, southwestern Tibet
TABLE 1. LITHOFACIES (AND THEIR CODES) USED IN LOGS OF MEASURED SECTIONS, AND INTERPRETATIONS
IN THIS STUDY, MODIFIED AFTER MIALL (1978) AND DECELLES ET AL. (1991)
Lithofacies
code
Fsl
Fcl
Fsm
Sm
Sr
St
Sp
Sh
Srw
Gcm
Description
Laminated black or gray siltstone
Laminated gray claystone
Massive, bioturbated, mottled siltstone, usually red; carbonate
nodules common
Massive medium- to fine-grained sandstone; bioturbated
Fine- to medium-grained sandstone with small, asymmetric, 2-D and
3-D current ripples
Medium- to very coarse-grained sandstone with trough crossstratification
Medium- to very coarse-grained sandstone with planar crossstratification
Fine- to medium-grained sandstone with plane-parallel lamination
HCS
Fine- to medium-grained sandstone with symmetrical small ripples
Pebble to boulder conglomerate, poorly sorted, clast-supported,
unstratified, poorly organized
Pebble to cobble conglomerate, moderately sorted, clast-supported,
unstratified, imbricated (long-axis transverse to paleoflow)
Pebble to cobble conglomerate, well sorted, clast-supported,
horizontally stratified
Pebble to cobble conglomerate, well sorted, clast-supported,
horizontally stratified, imbricated (long-axis transverse to paleoflow)
Pebble conglomerate, well sorted, clast-supported, trough crossstratified
Pebble to cobble conglomerate, well sorted, clast-supported, planar
cross-stratified
Massive, matrix-supported pebble to boulder conglomerate, poorly
sorted, disorganized, unstratified
Hummocky cross-stratified fine- to medium-grained sandstone
M
Micritic massive gray and yellow marl
Gcmi
Gch
Gchi
Gct
Gcp
Gmm
Process interpretation
Suspension-settling in ponds and lakes
Suspension-settling in ponds and profundal lakes
Paleosols, usually calcic or vertic
Bioturbated sand, penecontemporaneous deformation
Migration of small 2-D and 3-D ripples under weak (~20–40 cm/s),
unidirectional flows in shallow channels
Migration of large 3-D ripples (dunes) under moderately powerful
(40–100 cm/s), unidirectional flows in large channels
Migration of large 2-D ripples under moderately powerful (~40–60 cm/s),
unidirectional channelized flows; migration of sandy transverse bars
Upper plane bed conditions under unidirectional flows, either strong
(>100 cm/s) or very shallow
Deposition of oscillatory current (orbital) ripples in shallow lakes and ponds
Deposition from sheetfloods and clast-rich debris flows
Deposition by traction currents in unsteady fluvial flows
Deposition from shallow traction currents in longitudinal bars and gravel
sheets
Deposition from shallow traction currents in longitudinal bars and gravel
sheets
Deposition by large gravelly ripples under traction flows in relatively deep,
stable fluvial channels
Deposition by large straight-crested gravelly ripples under traction flows in
shallow fluvial channels, gravel bars, and gravelly Gilbert deltas
Deposition by cohesive mud-matrix debris flows
Deposition by combined unidirectional and oscillatory flows during storms
on the lower shoreface
Lacustrine carbonate mud
A
Figure 10. Photographs of
typical lithofacies in the lower
conglomerate and fluvial sandstone members. (A) Bouldercobble conglomerates (mainly
Gcm and Gcmi) in the lower
part of the lower conglomerate member, section 1KR.
(B) Imbricated pebble conglomerate (Gcmi) in middle
part of the lower conglomerate
member, section 1KR. (C) Bed
of matrix-supported boulder
conglomerate (Gmm) in section
5KR. Hammer is 41 cm long.
(D) Lithofacies Sm, St, and Sh
in lenticular sandstone bodies
of the fluvial sandstone member. Note irregular erosional
basal scour surface with ~1 m
of relief. Vertical dimension of
view is ~8 m.
B
C D
Geological Society of America Bulletin, July/August 2011
1347
DeCelles et al.
Figure 11. (A) Measured section
3KR, showing stack of six lacustrine parasequences. Thickness
of visible portion of the section
is 150 m. (B) Upward-coarsening lacustrine parasequence
capped by transgressive surface
(arrow at right). Person highlighted by arrow in lower left
indicates scale. (C) View toward
east from lower part of section
4KR of rhythmically bedded
lacustrine member and underlying massively bedded lower
conglomerate member. (D) Thin
sandstone beds in lacustrine
profundal shale, probably deposited by turbidity currents.
Hammer is 41 cm long. (E) Bedding surface view of transgressive lag composed of fossil fish
vertebrae and mandibles and
small chert pebbles, 497 m level
of section 7KR. (F) Fossil turtle
scute in transgressive lag, section 7KR.
A
B
C
D
E
F
D
thick, and the regular alternation of fine-grained
and sandy lithofacies imparts a rhythmic character to the outcrop (Fig. 11C).
We interpret these parasequences as the records of progradation of sandy nearshore and
deltaic systems into an offshore muddy lacustrine environment. The laminated black shales
are interpreted as offshore profundal deposits.
Their dark color results from abundant finegrained amorphous kerogen (see palynological information in the following). If maximum
parasequence thickness may be taken as a
crude estimate of minimum water depth, then
these shales accumulated in water greater than
80 m deep; this figure would increase substantially (>30%) upon decompaction. Hummocky
cross-stratification and small oscillatory current ripples in the lower parts of the sandy
portions of parasequences indicate lower shoreface storm deposits. The channelized, trough
1348
cross-stratified sandstone units at the tops of
many parasequences represent fluvial/deltaic
distributaries. Large-scale clinoforms (Fig. 12)
in some of these units probably accumulated in
prograding distributary mouth bars.
Red-Bed Member
In the uppermost part of the Kailas Formation, the assemblage of lithofacies changes
abruptly to bright red siltstone with nodular
carbonate zones, pebbly lenticular sandstone
units, and laminated marl layers (Fig. 8, section 8KR; Fig. 13). Sandstone beds contain
small ripples and horizontal laminations.
Lenticular coarse-grained sandstone units are
dominated by trough cross-stratification. The
red siltstones are typically massive, mottled,
and exhibit nodular weathering. Nodular gray
silty carbonate accumulations are common,
and nodular limestone beds become abundant
higher in the section. Unlike the arkosic conglomerates lower in the Kailas Formation,
those in the red-bed member contain clasts
of chert, limestone, and quartzite. Paleocurrent data from trough cross-strata indicate
southward and northward paleoflow directions
(Fig. 8, section 8KR).
We interpret the lenticular sandstone units as
fluvial channel deposits; the mottled and nodular siltstones as calcic paleosols (Mack et al.,
1993); and the nodular and laminated limestone
beds as carbonate lacustrine deposits. This assemblage of lithofacies suggests deposition in a
mixed fluviolacustrine system that experienced
strongly seasonal, semiarid climate. The absence of dark, organic-rich profundal lacustrine
facies indicates that, although lakes persisted
during deposition of the red-bed member, they
were not deep and may have been ephemeral.
Geological Society of America Bulletin, July/August 2011
Oligocene–Miocene Kailas basin, southwestern Tibet
Figure 12. Large-scale (~7 m thick) clinoform bedding in deltaic sandstone at top of a lacustrine parasequence. The sandy unit is capped by
a transgressive surface, which is in turn overlain by profundal lacustrine black shale.
Similarly, the presence of paleosol carbonate, in
marked contrast to the bulk of the Kailas Formation, suggests semiarid climate (Retallack,
1990; Mack et al., 1993).
A
CHRONOSTRATIGRAPHIC CONTROL
The age of the Kailas Formation was determined from U-Pb zircon ages from samples
of tuffs within the stratigraphic sections we
measured, and from maximum age constraints
imposed by the youngest populations of detrital
zircon ages. Detrital grain ages also provide
valuable provenance information.
The U-Pb ages of detrital and first-cycle volcanic zircon grains were determined by multicollector–laser ablation–inductively coupled
plasma–mass spectrometry (MC-LA-ICP-MS)
at the University of Arizona LaserChron Center.
Spots on individual zircon grains were ablated
with a New Wave DUV193 Excimer laser
(operating at a wavelength of 193 nm) using a
spot diameter of 35 μm. The ablated material
is carried in He gas into the plasma source of a
Micromass Isoprobe, which is equipped with
a flight tube of sufficient width that U, Th, and
Pb isotopes are measured simultaneously. All
measurements are made in static mode, using
Faraday detectors for 238U, 232Th, 208–206Pb, and
an ion-counting channel for 204Pb. Ion yields are
~1 mv per ppm. Each analysis consists of one
20 s integration on peaks with the laser off (for
backgrounds), twenty 1 s integrations with the
laser firing, and a 30 s delay to purge the previous sample and prepare for the next analysis.
The ablation pit is ~20 μm deep. Common Pb
correction was made by using the measured 204Pb
and assuming an initial Pb composition from
Stacey and Kramers (1975) (with uncertainties
B
Figure 13. (A) Abrupt change from dark-colored shales and finegrained sandstones in the upper part of the lacustrine member
to the red-bed member in section 8KR. Icy peak in background
is Gurla Mandatha. (B) Nodular paleosol carbonate in red-bed
member, section 8KR.
Geological Society of America Bulletin, July/August 2011
1349
DeCelles et al.
of 1.0 for 206Pb/204Pb and 0.3 for 207Pb/204Pb).
Our measurement of 204Pb is unaffected by the
presence of 204Hg because backgrounds were
measured on peaks (thereby subtracting any
background 204Hg and 204Pb), and because very
little Hg was present in the argon gas.
Interelement fractionation of Pb/U is generally ~15%, whereas fractionation of Pb isotopes
is generally <2%. In-run analysis of fragments of
a large zircon crystal (every sixth measurement)
with known age of 563.5 ± 3.2 Ma (2σ error)
(Gehrels et al., 2008) was used to correct for this
fractionation. Fractionation also increases with
depth into the laser pit by up to 5%. This depthrelated fractionation was accounted for by monitoring the fractionation observed in the standards.
Analyses that displayed >10% change in ratio
during the 20 s measurement are interpreted to
be variable in age (or perhaps compromised by
fractures or inclusions), and were excluded from
further consideration. Also excluded were analyses that yielded >15% uncertainties in 206Pb/238U
ages or were >5% reverse discordant.
The measured isotopic ratios and ages are reported in Tables DR1 and DR2 (see footnote 1).
Errors that propagate from the measurement of
206
Pb/238U, 206Pb/207Pb, and 206Pb/204Pb are reported at the 1σ level. Additional errors that affect all ages include uncertainties from U decay
constants, the composition of common Pb, and
calibration correction; these systematic errors
are 1%–2% for most samples.
An additional factor that complicates analyses
in the 500 to ca. 2000 Ma age range is the change
in precision of 206Pb/238U and 206Pb/207Pb ages—
206
Pb/238U ages are generally more precise for
younger ages, whereas 206Pb/207Pb ages are more
precise for older ages. Therefore, in samples that
contain a cluster of analyses with concordant to
slightly discordant ages, we rely on 206Pb/238U
ages up to 1000 Ma and 206Pb/207Pb ages if the
206
Pb/238U ages are older than 1000 Ma. Additional details about analytical procedures are
described by Gehrels et al. (2008).
Analyses that yielded isotopic data of acceptable discordance, in-run fractionation, and precision are shown in Tables DR1 and DR2 (see
footnote 1). In total, 650 zircon ages are reported
in this paper. Detrital zircon analyses are plotted on relative age-probability diagrams (Fig.
14), which represent the sum of the probability
distributions of all analyses from a sample normalized so that the areas beneath the probability
curves are equal for all samples. Age peaks on
these diagrams are considered robust if defined
by several analyses, whereas less significance is
attributed to peaks defined by single analyses.
Two samples of sandstone from the lower
conglomerate and fluvial sandstone members
yielded detrital zircons with minimum age clus-
1350
ters of 24.7–26.6 Ma (Fig. 14, samples 1KR353,
2KR107), indicating that the age of the sandstones can be no older than that age range. Three
detrital zircon samples from the lower part of
the lacustrine member produced minimum age
clusters of 24.7–26.6 Ma (Fig. 14, samples
1KR624, 2KR420, 7KR129). The abundance
and tight clustering of grains in the 24–26 Ma
range in these sandstone samples suggest that
the zircon ages are good approximations of
the depositional ages. Approximately 500 m
above the base of the lacustrine member, two
tuff beds produced mean zircon ages of 24.6 ±
0.5 and 24.2 ± 0.5 Ma (Fig. 4, samples PD3 and
PD4). A third tuff bed, slightly contaminated
with detrital grains, at approximately the same
stratigraphic level (Fig. 4, sample 7KR465) produced a U-Pb zircon mean age of 24.1 ± 0.4 Ma.
Finally, near the top of the section, a sandstone
sample in the red-bed member (Fig. 14, sample
8KR67) produced a youngest cluster of detrital
zircon ages of 24 Ma, consistent with a depositional age younger than this, but providing no
further constraint on the minimum age of this
interval. Together, the geochronological data indicate that the bulk of the Kailas Formation in
this region was deposited rapidly over a period
of only 2 to 3 m.y. during the latest Oligocene
and early Miocene. Our U-Pb ages are consistent with, but slightly older than, biotite and
plagioclase 40Ar/39Ar intercept ages reported by
Aitchison et al. (2009) between 22.3 ± 0.7 Ma
and 16.9 ± 0.2 Ma from felsic tuffs interbedded
within the Kailas Formation east of Mount Kailas and within the Gangrinboche conglomerate
near Dazhuqu ~700 km to the east.
PROVENANCE
Modal Sandstone Petrology
The modal framework-grain composition
of 41 sandstone samples from the Kailas Formation was documented by identification of
450 grains per slide according to the GazziDickinson method (Ingersoll et al., 1984). In order to aid the identification of feldspars, one half
of each standard petrographic thin section was
stained for potassium- and calcium-feldspars.
Grain types are listed in Table 2, and recalculated modal values are given in Table 3.
Quartzose grains include monocrystalline
(Qm), polycrystalline (Qp), foliated polycrystalline (Qpt), and quartzite/siltstone (Qmss/Qms)
grains. Feldspars include the potassium varieties
(K; orthoclase, perthite, microcline) and plagioclase (P, occasionally with myrmekitic textures).
Lithic grains include limestone, mica schist, serpentine schist, and several varieties of volcanic
grains, including lathwork, vitric, felsic, mafic,
and microlitic. Vitric grains include microcrystalline (devitrified) to glassy types; occasional
bubble-wall shard glasses; K-feldspar–rich
glass; and welded, flow-banded glassy tuffs.
Mafic grains consist of coarse-grained hypabyssal varieties composed of epidote ± pyroxene ±
plagioclase. Felsic grains are typically strongly
altered sericite + quartz ± feldspar aggregates.
Trace minerals include zoisite/clinozoisite,
epidote, chlorite, biotite, muscovite, clino- and
orthopyroxenes, zircon, sphene, monazite, and
opaque grains.
Conglomerate Clast Counts
At least 100 clasts were counted at five stations representing the entire stratigraphic thickness of the Kailas Formation (Fig. 14). The lower
part of the Kailas Formation is dominated by
volcanic, granitoid, and hypabyssal clasts, with
minor amounts of chert. Conglomerates in the
red-bed member contain abundant sedimentary
clasts, including limestone and chert. Conglomerates in section 5KR, which we include in the
lower conglomerate member, contain numerous
clasts of reworked pebbly conglomerate.
Detrital Zircon Ages
Detrital zircon age spectra from all of the
seven sandstone samples from the Kailas Formation are dominated by ages that are younger
than 100 Ma; fewer than 20% of the dated grains
are older than 100 Ma (Fig. 14). Six of the seven
samples exhibit predominant age clusters with
peaks at ca. 24–26 Ma, and lesser peaks in the
80–55 Ma range (Fig. 14). Sample 8KR104,
near the top of the Kailas Formation, produced
a range of ages between ca. 45 Ma and 60 Ma
with a sharp unimodal peak at ca. 49.5 Ma.
One sample (8KR67) also produced several age
clusters with peaks of ca. 30 Ma, 40 Ma, and
50 Ma. Zircons with ca. 75–80 Ma ages are present in several samples (Fig. 14). The older than
100 Ma ages exhibit peaks in the Late Archean,
Mesoproterozoic, Neoproterozoic, and early to
middle Paleozoic. Although some aspects of the
older than 100 Ma grain populations resemble
documented detrital zircon ages of Lhasa terrane
rocks (Leier et al., 2007), statistically rigorous
comparison is not possible because of the small
numbers of grains in our older than 100 Ma data
sets. It is also noteworthy that none of the Kailas detrital zircon age distributions bears much
resemblance to detrital zircon age spectra obtained from the Himalayan thrust belt (Parrish
and Hodges, 1996; DeCelles et al., 2000, 2004;
Gehrels et al., 2003, 2006; Amidon et al., 2005;
Martin et al., 2005) or from Cenozoic strata of
the central Lhasa terrane (DeCelles et al., 2007a).
Geological Society of America Bulletin, July/August 2011
Oligocene–Miocene Kailas basin, southwestern Tibet
Provenance Interpretation
Modal sandstone compositions in terms of
standard ternary diagrams are illustrated in
Figure 15. Overall, Kailas Formation sandstones exhibit compositions typical of sandstones derived from magmatic arcs (Dickinson
et al., 1983; Garzanti et al., 2007), with large
amounts of plagioclase feldspar and felsic to
intermediate volcanic lithic grains. Conglomerate compositions are consistent with this interpretation, as they are dominated by volcanic
and granitoid clasts. Recycled conglomerate
clasts of unknown provenance are common
in the lower conglomerate member in section
5KR; all we can be sure of is that these clasts
were derived from the north, based on robust
paleocurrent data (Fig. 7, section 5KR). We observed no significant changes in composition
throughout most of the Kailas Formation, until
the red-bed member. The red-bed member contains abundant sandstone/quartzite, limestone,
and chert clasts.
The provenance data are consistent with
paleocurrent data that indicate the principal
source terrane was the Kailas magmatic complex in the Gangdese magmatic arc to the north
of the Kailas basin. Additional volcanogenic
material may have been derived from eruptive centers located in the western Lhasa terrane and across the Karakoram fault toward the
west (Miller et al., 1999; Mahéo et al., 2002;
Williams et al., 2004), which were contemporaneously active and possibly produced the airfall tuffs that we sampled. During deposition of
the red-bed member, Tethyan source terranes
became more important, supplying limestone,
sandstone/quartzite, and chert clasts from the
Paleozoic-Mesozoic sedimentary succession
in the hanging wall of the South Kailas thrust
system. The detrital zircon age spectra are
dominated by age clusters that are typical of
the Gangdese arc, with peaks in the ca. 25 Ma,
49–50 Ma, and 70–80 Ma ranges (Fig. 14). Surprisingly, even the uppermost samples in section
Figure 14. Age probability diagrams of detrital zircon ages from sandstones of the
Kailas Formation. Diagrams in left-hand
column show the younger than 100 Ma populations; those in right-hand column show
the older than 100 Ma populations. Letter
“n” indicates number of grains in each
population. See Table DR2 for data (see text
footnote 1). Ages of Kailas magmatic complex rocks (heavy dashed lines; this work)
and intrusions in the Ayi Shan (gray band;
Zhang et al., 2010) are shown for reference.
Age (Ma)
n = 402
80.5%
n = 98
19.5%
Ages >100 Ma
Ages <100 Ma
8KR104
n = 60
8KR67
n = 86
n = 21
7KR129
n = 76
n = 21
2KR420
n = 28
n=5
2KR107
n = 53
n = 12
1KR624
n = 52
n = 20
1KR353
n = 47
n = 19
0
20
Ayi Shan
intrusions
40
60
80
100
0
1000
2000
3000
Gangdese magmatic
complex
Geological Society of America Bulletin, July/August 2011
1351
DeCelles et al.
TABLE 2. MODAL PETROGRAPHIC POINT-COUNTING PARAMETERS
Symbol
Description
Qm
Monocrystalline quartz
Qp
Polycrystalline quartz
Qpt
Foliated polycrystalline quartz
Qms
Monocrystalline quartz in sandstone or quartzite lithic grain
C
S
Qt
Chert
Siltstone
Total quartzose grains (Qm + Qp + Qpt + Qms + C + S)
K
F
Potassium feldspar
(including perthite, myrmekite, microcline)
Plagioclase feldspar
(including Na and Ca varieties)
Total feldspar grains (K + P)
Lvm
Lvf
Lvv
Lvx
Lvl
Lv
Mafic volcanic grains (epidote ± pyx ± plag)
Felsic volcanic grains (sericite + qtz ± feldspar)
Vitric volcanic grains
Microlitic volcanic grains
Lathwork volcanic grains
Total volcanic lithic grains (Lvm + Lvf + Lvv + Lvx + Lvl)
Lsh
Lph
Lsm
Lc
Lm
Ls
Lt
L
Mudstone
Phyllite
Schist (mica schist)
Carbonate lithic grains
Total metamorphic lithic grains (Lph + Lsm + Qpt)
Total sedimentary lithic grains (Lsh + Lc + C + S + Qms)
Total lithic grains (Ls + Lv + Lm + Qp)
Total nonquartzose lithic grains (Lv + Ls + Lph + Lsm + Lc)
P
Accessory minerals
Epidote/zoisite
Chlorite
Muscovite
Biotite
Zircon
Sphene
Clinopyroxene
Orthopyroxene
Monazite
Magnetite
8KR, which contains clasts of clear-cut Tethyan
provenance, do not show the strong pre–100 Ma
age clusters typical of Tethyan strata throughout
the Himalayan thrust belt (DeCelles et al., 2000;
Gehrels et al., 2003, 2006; Amidon et al., 2005;
Martin et al., 2005).
PALYNOLOGY AND VERTEBRATE
PALEONTOLOGY
Palynological analysis of samples from
dark-colored organic-rich shale from the Kailas Formation produced generally poor yields.
The dark colors of Kailas shales result from
high amorphous kerogen content. The few
taxa that were recovered include Milfordia/
Monoporites annulatus (common in circumequatorial early and middle Cenozoic floras),
Ulmoidiepites (Elm), and Cyatheacidites annulatus (an equatorial fern). None of these taxa
is chronologically or environmentally diagnostic. However, it is significant that pollen from
temperate and boreal/high-elevation species
such as oak, alder, maple, basswood, willow,
fir, pine, hemlock, and spruce is absent from
1352
Kailas basin deposits, whereas these taxa are
abundant in coeval high-elevation deposits
of the Nima Basin in central Tibet (DeCelles
et al., 2007a, 2007b).
Vertebrate fossils are abundant in the Kailas
Formation, particularly in lag concentrations
along transgressive surfaces (Figs. 11E and
11F). These include fish fossils (teeth, mandibles, and vertebrae) and turtles (bones and
scutes). Woody plant material is also abundant
throughout the Kailas Formation, except in the
red-bed member in the uppermost part of
the formation.
STABLE ISOTOPE DATA
Oxygen isotopic values from sedimentary
carbonates (expressed as δ18Occ in ‰) and the
waters from which they precipitated (δ18Omw)
are potentially useful paleoaltimeters because
they decrease with increasing elevation (e.g.,
Garzione et al., 2000a, 2000b; Dettman and
Lohmann, 2000; Poage and Chamberlain, 2001;
Mulch et al., 2004; Blisniuk and Stern, 2005;
Currie et al., 2005; Saylor et al., 2009). The sev-
eral key caveats that must be considered in using stable isotopes to reconstruct paleoelevation
are discussed in Quade et al. (2007) and Saylor
et al. (2009).
We collected samples of paleosol carbonate
from the red-bed member of the Kailas Formation for stable isotope paleoaltimetry. Unfortunately, no paleosol carbonate was found
in the main body of the Kailas Formation. We
also analyzed samples of modern and Quaternary soil carbonate from the Kailas region and
from the Nima area in central Tibet (DeCelles
et al., 2007b) in order to calibrate a modern
baseline relationship between elevation and
isotopic values.
All carbonate samples were heated at 150 °C
for 3 h in vacuo and processed using an automated sample preparation device (Kiel III) attached directly to a Finnigan MAT 252 mass
spectrometer at the University of Arizona.
The δ18O and δ13C values were normalized to
NBS-19 based on internal laboratory standards.
Precision of repeated standards was ±0.1‰ for
δ18O and ±0.06‰ for δ13C (1σ).
Paleosol carbonate from the uppermost part
of the Kailas Formation yielded δ18Occ values
ranging mostly between –15‰ and –18‰ (but
with two values around –22‰), and δ13Ccc values between –4.5‰ and –6.0‰ (Table 4; Fig.
16). Modern soil carbonate in the Kailas area
and in the Nima area of central Tibet produced
δ18Occ values ranging between about –10‰ and
–14.5‰, and δ13Ccc values scattered between
+2.3‰ and –5.0‰ (Table 4; Fig. 16). The
modern samples were collected >50 cm below
the ground surface over an elevation range of
~4500–4700 m.
The lower δ18Occ values in the Kailas Formation paleosol carbonates compared to those in
local modern soils could at first glance be taken
to indicate higher paleoelevations in this region
during the early Miocene. However, the most
negative modern values of –13‰ to –14.5‰
are actually comparable to ancient values in
the –15‰ to –18‰ range after accounting
for the modest isotopic effects (~2‰–3‰) of
warmer, and largely ice-free conditions in the
early Miocene (Shackleton and Kennett, 1975;
Zachos et al., 1994; Bowen and Wilkinson,
2002). The higher and more dispersed δ18Occ
values in modern soils probably also reflect the
arid modern climate (greater soil evaporation)
compared to a more humid early Miocene climate, as suggested by the abundance of organic
material in the main body of the Kailas Formation. This difference in climate is firmly supported by the carbon isotope results. The δ13Ccc
values in the Kailas paleosols are much more
negative and less dispersed than those of modern soils in the area, probably reflecting higher
Geological Society of America Bulletin, July/August 2011
Oligocene–Miocene Kailas basin, southwestern Tibet
Sample
Qm%
F%
Lt%
Qt%
1KR73
0.15
0.48
0.38
0.18
1KR86
0.11
0.20
0.69
0.17
1KR148
0.12
0.33
0.55
0.21
1KR165
0.17
0.42
0.41
0.21
1KR240
0.10
0.51
0.39
0.10
1KR255
0.11
0.55
0.34
0.13
1KR486
0.12
0.48
0.40
0.15
1KR568
0.13
0.59
0.28
0.14
1KR630
0.17
0.34
0.49
0.20
2KR1
0.16
0.40
0.44
0.18
2KR40
0.15
0.44
0.41
0.17
2KR184
0.12
0.46
0.42
0.14
2KR232
0.12
0.42
0.45
0.15
2KR330
0.08
0.63
0.29
0.09
2KR420
0.14
0.43
0.43
0.16
3KR14
0.10
0.54
0.36
0.12
3KR39
0.11
0.48
0.41
0.13
3KR65
0.13
0.35
0.51
0.15
3KR100
0.12
0.34
0.54
0.13
3KR147
0.13
0.44
0.43
0.14
4KR25
0.10
0.61
0.29
0.14
4KR139
0.13
0.43
0.44
0.17
4KR210
0.10
0.57
0.33
0.13
4KR301
0.16
0.46
0.38
0.18
4KR400
0.13
0.59
0.28
0.16
4KR461
0.16
0.33
0.51
0.20
5KR18
0.16
0.31
0.53
0.18
5KR195
0.12
0.27
0.61
0.18
5KR226
0.11
0.17
0.72
0.17
5KR325
0.13
0.25
0.62
0.21
7KR177
0.09
0.45
0.46
0.12
7KR246
0.10
0.50
0.41
0.11
7KR252
0.22
0.50
0.28
0.25
7KR337
0.13
0.61
0.26
0.14
7KR483
0.22
0.53
0.26
0.28
7KR542
0.25
0.54
0.20
0.30
7KR847
0.18
0.36
0.46
0.23
8KR8
0.11
0.02
0.87
0.20
8KR17
0.29
0.44
0.27
0.35
8KR65
0.28
0.32
0.40
0.35
8KR96
0.07
0.03
0.90
0.14
Note: Parameter definitions are provided in Table 2.
respiration rates and more plant cover during
the early Miocene.
Thus, the oxygen isotope data suggest that
the paleoelevation of the Kailas Formation
during deposition of the red-bed member was
essentially the same as the modern elevation
(>4500 m). We emphasize, however, that the
data we present are not definitive without a
more rigorous evaluation of the potential effects of diagenesis. We were not able to run an
isotopic “conglomerate test” (DeCelles et al.,
2007b) or other tests (Cyr et al., 2005) to determine whether the nodules have been reset
with respect to δ18O. Barring such tests, several
considerations provide some confidence in the
results: The paleosol carbonate nodules we analyzed are uniformly micritic, burial depth in the
Kailas red-bed member was probably modest
(e.g., less than 2 km), and both our carbon and
oxygen isotope results are quite similar to those
from early Miocene paleosols at Nima in central Tibet, from which we also concluded that
paleoelevations stood at >4500 m at that time
(DeCelles et al., 2007b).
TABLE 3. RECALCULATED MODAL PETROGRAPHIC DATA
F%
L%
Qm%
P%
K%
Lm%
0.47
0.35
0.23
0.59
0.17
0.02
0.20
0.63
0.35
0.65
0.00
0.00
0.33
0.46
0.26
0.58
0.15
0.02
0.42
0.37
0.28
0.49
0.23
0.02
0.51
0.39
0.16
0.58
0.27
0.13
0.55
0.31
0.16
0.61
0.23
0.14
0.48
0.37
0.20
0.58
0.23
0.01
0.59
0.28
0.18
0.51
0.31
0.18
0.34
0.45
0.33
0.56
0.11
0.08
0.40
0.42
0.29
0.51
0.20
0.09
0.44
0.39
0.26
0.54
0.21
0.11
0.46
0.40
0.21
0.55
0.24
0.05
0.42
0.43
0.23
0.66
0.11
0.06
0.63
0.28
0.12
0.49
0.39
0.02
0.43
0.41
0.25
0.50
0.25
0.01
0.54
0.34
0.15
0.49
0.36
0.00
0.48
0.39
0.18
0.54
0.28
0.01
0.35
0.50
0.28
0.47
0.25
0.02
0.34
0.53
0.27
0.53
0.20
0.00
0.44
0.42
0.23
0.52
0.25
0.06
0.61
0.25
0.15
0.60
0.25
0.01
0.43
0.40
0.24
0.54
0.23
0.04
0.57
0.30
0.15
0.59
0.26
0.01
0.46
0.36
0.26
0.56
0.18
0.01
0.59
0.25
0.17
0.54
0.28
0.02
0.33
0.48
0.33
0.56
0.11
0.03
0.31
0.51
0.34
0.60
0.06
0.00
0.27
0.55
0.30
0.62
0.08
0.02
0.17
0.67
0.40
0.59
0.01
0.01
0.25
0.54
0.34
0.62
0.05
0.02
0.45
0.43
0.17
0.67
0.16
0.01
0.50
0.39
0.16
0.54
0.30
0.02
0.50
0.25
0.31
0.43
0.26
0.03
0.61
0.25
0.17
0.73
0.10
0.07
0.52
0.20
0.29
0.50
0.21
0.04
0.54
0.15
0.32
0.37
0.31
0.01
0.36
0.41
0.33
0.38
0.29
0.06
0.02
0.78
0.85
0.15
0.00
0.53
0.44
0.22
0.39
0.31
0.30
0.08
0.32
0.33
0.47
0.28
0.25
0.03
0.03
0.83
0.66
0.32
0.02
0.56
BASIN ARCHITECTURE AND
SUBSIDENCE MECHANISM
The measured sections and additional field
observations demonstrate that the Kailas Formation coarsens northward toward its basal contact
with the Kailas magmatic complex. Figure 9
depicts a tentative, composite lithostratigraphic
correlation based on our measured sections.
This part of the Kailas basin was filled by three
distinct depositional systems: a lower alluvialfluvial system that was derived from the deeply
eroded Gangdese magmatic arc to the north; a
medial lacustrine system into which marginal
deltas and nearshore systems prograded, again
mainly from the north; and an upper fluvial
system that was mainly derived from Tethyan
rocks exposed in the northern Himalayan thrust
belt located directly to the south in the hanging wall of the Great Counter thrust. The basin
fill exhibits a nearly continuous upward-fining
grain-size trend, until the uppermost part of
the section, and, in most respects, it resembles
a classic “lacustrine sandwich,” with a lower
Lv%
0.97
0.97
0.95
0.93
0.87
0.79
0.96
0.78
0.89
0.91
0.87
0.95
0.87
0.98
0.97
0.99
0.99
0.93
1.00
0.90
0.96
0.91
0.98
0.99
0.94
0.85
1.00
0.97
0.98
0.94
0.98
0.93
0.93
0.87
0.70
0.93
0.88
0.15
0.82
0.89
0.30
Ls%
0.01
0.03
0.02
0.04
0.00
0.07
0.02
0.04
0.03
0.01
0.01
0.00
0.07
0.01
0.03
0.01
0.01
0.04
0.00
0.04
0.04
0.05
0.01
0.01
0.04
0.12
0.00
0.01
0.01
0.05
0.02
0.04
0.04
0.05
0.26
0.06
0.06
0.32
0.10
0.07
0.14
Lv/Lt
0.90
0.91
0.82
0.86
0.86
0.73
0.92
0.76
0.84
0.88
0.83
0.90
0.82
0.95
0.94
0.93
0.95
0.91
0.98
0.89
0.85
0.87
0.90
0.92
0.86
0.82
0.97
0.89
0.92
0.84
0.94
0.90
0.85
0.84
0.60
0.75
0.80
0.15
0.68
0.77
0.29
F/F+Qm
76.58
65.47
73.63
71.65
84.13
83.90
80.30
81.96
67.11
70.92
74.14
78.99
77.18
88.05
75.10
85.11
81.82
72.48
73.08
77.17
85.44
76.49
84.56
74.28
82.59
66.52
65.71
69.94
60.00
66.27
82.85
83.86
69.38
82.72
71.00
68.24
66.94
15.25
60.75
53.23
34.09
K:K+P
0.23
0.00
0.21
0.32
0.32
0.28
0.28
0.37
0.17
0.28
0.28
0.31
0.14
0.44
0.34
0.43
0.34
0.35
0.27
0.33
0.29
0.30
0.30
0.24
0.09
0.16
0.09
0.11
0.01
0.07
0.19
0.36
0.37
0.12
0.30
0.46
0.43
0.00
0.49
0.48
0.07
coarse-grained interval overlain by fine-grained
lacustrine deposits, which are in turn overlain
by a return to relatively coarse-grained fluvial
deposits (Lambiase, 1990; Schlische, 1992).
Aitchison et al. (2002) reported that the Kailas
Formation at Mount Kailas is divisible into three
distinct members that are broadly similar to what
we have found, with the exception that all units
seem to be somewhat coarser-grained in outcrops located on Mount Kailas (Gansser, 1964).
A second key feature of the Kailas Formation is that it rests in buttress unconformity directly upon its principal source terrane, without
fault contact or structural disruption (Figs. 3B
and 5B). The contact dips uniformly and gently
southward. This onlapping relationship with
no evidence for growth structures is not typical of thrust-generated wedge-top or foredeep
sediments, which are characteristic of contractional tectonic settings. On the other hand,
the uppermost part of the Kailas Formation is
undoubtedly incorporated into contractional
structures associated with the tip of the northverging South Kailas thrust system (Gansser,
Geological Society of America Bulletin, July/August 2011
1353
DeCelles et al.
8KR96
Qt
Qm
1KR466
TETHYAN
SEDIMENTARY
Dickinson et al. (1983)
magmatic arc
provenance fields
5KR178
F
Lt
F
Lm
L
Qm
1KR82
Outliers are from uppermost
part of section, with Tethyan
component
HYPABYSSAL
Lv
RECYCLED
CONGLOMERATE &
SANDSTONE
Ls
P
1KR2 GRANITOID
CHERT
K
VOLCANIC
Figure 15. Ternary diagrams showing modal sandstone petrographic data (left) and pie charts showing conglomerate clast count data
(right). See Table 2 for definitions of petrographic parameters, and Table 3 for recalculated modal petrographic data.
1964), and the provenance data support derivation of sediment from its hanging wall. Similar structural relationships were reported by
Searle et al. (1990) in the Indus Group and
by Aitchison et al. (2002) in the Gangrinboche
conglomerates.
We propose that the bulk of the Kailas Formation was deposited in a basin produced by
extension or transtension. This would account
for the typical lacustrine stratigraphic sandwich, the overall upward-fining sequence, and
the absence of contractional growth structures.
An overall extensional setting is also consistent with the presence of phlogopitic trachyandesites/basalts in the section. The timing
of extension would be no older than late Oligocene (ca. 26 Ma) and no younger than slip
on the Great Counter thrust, which was completed by ca. 15 Ma (Yin et al., 1994). Late
Oligocene–early Miocene extension in the
Kailas basin is consistent with evidence for
rapid exhumation in the footwall of the Ayi
Shan detachment at that time (Zhang et al.,
2010), which would have been located directly
1354
beneath the Kailas basin during its development (Fig. 5B). The average rate of sediment
accumulation during this relatively brief interval was ~0.5 mm/yr, which is not diagnostic
of any particular tectonic setting, but is high
enough to be typical of rapidly subsiding rift
basins. Transtensional and strike-slip basins
commonly have sediment accumulation rates
that exceed 1.0 mm/yr (Allen and Allen, 2005;
Xie and Heller, 2009). Only during deposition
of the red-bed member was the Kailas Formation influenced directly by shortening; the redbed member signals the onset of shortening
along the South Kailas (or Greater Counter)
thrust. The direction of extension in the Kailas
basin is unknown, but data from the Ayi Shan
suggest oblique dextral transtension (Zhang
et al., 2010). Kailas basin extension predates
by at least 10 m.y. the better-known east-west
extension that formed north-striking graben
in Tibet beginning in late Miocene time (e.g.,
Harrison et al., 1995; Murphy et al., 2002;
Garzione et al., 2003; Kapp and Guynn, 2004;
Saylor et al., 2009).
DISCUSSION
The origin of the Kailas basin has remained
enigmatic ever since Gansser’s observations
were first published (Heim and Gansser, 1939;
Gansser, 1964). Based on recent reconnaissance work, Yin et al. (1999), Aitchison et al.
(2002, 2009), and Davis et al. (2004) concluded
that the Kailas Formation was deposited in response to regional uplift of the southern Lhasa
terrane in an overall contractional tectonic setting associated with the onset of slip on the
Great Counter thrust to the south and local
uplift along the Gangdese thrust system to the
north (Yin et al., 1999). Aitchison et al. (2002,
2009) attributed the Kailas and other units of
the Gangrinboche conglomerates to braided
stream deposition in an axial system that flowed
parallel to the suture zone. Searle et al. (1990)
reported that units in the Indus Group “molasse” (which are likely correlatives with the
Kailas Formation) are composed of lacustrine
and alluvial deposits, derived from both sides
of the suture zone, and deposited on top of the
Geological Society of America Bulletin, July/August 2011
Oligocene–Miocene Kailas basin, southwestern Tibet
TABLE 4. LIGHT STABLE ISOTOPE RESULTS FROM SOILS AND PALEOSOLS
Sample ID
δ13C VPDB
Soil depth (cm)
Elevation (m)
δ18O VPDB
Modern Kailas soils
KAILAS-1 A
–14.42
–4.93
95
4414
KAILAS-1 E
–11.91
–2.92
95
4414
KAILAS-1 B
–12.55
–1.70
95
4414
KAILAS-1 D
–12.57
–1.60
95
4414
KAILAS-24 B
–11.08
–0.15
100
4809
KAILAS-24 A
–12.06
1.45
100
4809
KAILAS-24 C
–13.06
–1.55
100
4809
Modern Nima soils
Nima 1 B
Nima 1 D
Nima 1 F
Nima 1 H
Nima 4 A
Nima 4 D
Nima 4 E
Nima 4 F
Nima 4 G
Nima 4 H
–11.59
–13.38
–11.66
–13.03
–13.83
–10.07
–13.04
–13.47
–12.98
–13.60
–0.47
–3.22
0.09
–2.50
–2.72
2.33
–2.01
–3.39
–1.84
–2.47
50
50
50
50
120
120
120
120
120
120
4495
4495
4495
4495
4500
4500
4500
4500
4500
4500
Kailas Formation paleosols from sections 7KR and 8KR
8KR103PS-C
–5.14
–16.02
8KR76A-A
–5.14
–15.78
8KR76A-B
–4.76
–15.41
8KR76A-C
–6.02
–15.63
8KR76B-A
–5.59
–16.73
8KR76B-B
–5.56
–16.78
8KR76B-C
–5.56
–16.64
8KR52PS-A
–4.53
–15.69
8KR52PS-B
–4.88
–16.07
8KR52PS-C
–4.83
–16.24
8KR51PS-A
–5.37
–16.30
8KR51PS-B
–5.15
–16.26
8KR51PS-C
–4.84
–16.15
7KR408PS-B
–6.91
–21.48
7KR408PS-C
–6.44
–21.44
Note: VPDB—Vienna Peedee belemnite.
Ladakh batholith (the western continuation of
the Gangdese arc). An inherent notion in all
published explanations of the Kailas Formation
(and its equivalents) is that it was deposited in a
contractional tectonic regime.
Although the uppermost part of the Kailas
Formation does contain clasts that were derived from the south (in the hanging wall of the
Great Counter thrust), the bulk of the unit was
derived from the north, as originally indicated
by Gansser (1964). Paleocurrent data from the
Kailas Formation provide no evidence for significant axial paleoflow along the suture zone
(e.g., Searle et al., 1990; Aitchison et al., 2002).
Our results also are not consistent with deposition of the bulk of the Kailas Formation in a contractional setting, but instead suggest deposition
in a rift or transtensional strike-slip basin. The
key observations that do not support deposition in a contractional setting include: (1) The
Kailas Formation sits on local Gangdese arc
“basement,” and was derived directly from this
basement. This is not typical of thrust-generated
proximal facies, which are derived from the
hanging walls of bounding thrust faults and
deposited upon rocks that are structurally below
the associated thrust sheet. The general paucity
of clasts derived from the Tethyan Himalaya
to the south is remarkable because the Tethyan
thrust belt was certainly active during the
Eocene–Oligocene (Ratschbacher et al., 1994;
Hodges, 2000; Murphy and Yin, 2003; DeCelles
et al., 2004; Aikman et al., 2008). (2) We found
no evidence for progressive tilting of bedding
characterizing contractional growth structures
(e.g., Anadòn et al., 1986) in the Kailas Formation, neither in proximity to the South Kailas
thrust system, nor along its northern outcrop
border where it rests unconformably on top of
the Kailas magmatic complex. The outcrop evidence indicates that most shortening related to
the South Kailas thrust system took place after
deposition of the Kailas Formation. (3) The overall textural and lithofacies pattern in the Kailas
Formation is not typical of contractional wedgetop settings or proximal foredeep settings (see,
e.g., DeCelles and Giles, 1996; Lawton et al.,
1999; Heermance et al., 2007), both of which
usually result in upward-coarsening packages
of growth strata. Instead, the Kailas Formation
generally fines upward and consists of a classic
rift-basin lacustrine sandwich, which could have
developed in any narrow, deep extensional environment (including oblique strike-slip). (4) The
presence of phlogopite-bearing basaltic andesites
and adakitic tuffs (Aitchison et al., 2009) in the
Kailas Formation suggests a thermal pulse that
led to melting of Asian lithospheric mantle and
garnet-bearing lower crust, and is consistent with
(though not diagnostic of) eruption in an extensional tectonic environment—perhaps associated with slab break-off (Mahéo et al., 2002;
Chung et al., 2009).
The bulk of the Kailas Formation is characterized by lithofacies assemblages that are consistent with deposition under a warm, tropical
climate. Searle et al. (1990) also reported thick
lacustrine deposits containing abundant plant
remains in the Indus Group, and suggested a
warm temperate climate. The overall pattern of
sedimentation changes abruptly in the red-bed
member of the Kailas Formation, which contains clasts derived from the hanging wall of the
South Kailas thrust system, exhibits northward
paleocurrent directions (Fig. 8), and contrasts
sharply with the rest of the Kailas Formation in
containing abundant paleosol carbonate, marl,
and highly oxidized fine-grained facies (Fig. 13).
Lithofacies in the red-bed member are similar
to those in the roughly contemporaneous Nima
Basin of central Tibet, which is documented to
have been at high paleoelevation by late Oligocene time (DeCelles et al., 2007a, 2007b). The
abrupt change from an essentially noncalcareous to a highly calcareous depositional environment in the red-bed member is unlikely to have
been simply a matter of changing provenance,
because carbonate clasts and calcite cement
are also present in the lower part of the Kailas
Formation. Rather, we suggest that the bulk of
the Kailas Formation accumulated in a rapidly
subsiding, warm tropical basin at relatively
low elevation, whereas the red-bed member
accumulated at relatively high elevation under
an arid or semiarid climate, more analogous
to (but still wetter than, based on the δ13C evidence) the modern setting of the Indus-Yarlung
suture. Although the oxygen and carbon isotope
data that we present are limited in number and
stratigraphic coverage, they are consistent with
the interpretation that the uppermost part of the
Kailas Formation was deposited at high elevation. More thorough paleoaltimetry studies indicate that the Zhada Basin to the west (Saylor
et al., 2009), Thakkhola graben to the southeast
(Garzione et al., 2000a), and the Oiyug Basin
to the east (Currie et al., 2005) were all at high
elevation (>4000 m) by middle to late Miocene
time. We infer that the red-bed member of the
Kailas Formation was deposited when the Great
Counter thrust became active and began to supply sediment from elevated highlands to the
south, and the Kailas basin began to structurally invert and gain elevation, culminating in
its present very high elevation (~4.7–6.7 km).
Unfortunately, the absence of paleosol carbon-
Geological Society of America Bulletin, July/August 2011
1355
DeCelles et al.
3
2
along the length of the suture zone. Definitive
assessment of the regional significance of midCenozoic extension in the Indus-Yarlung suture
zone requires more detailed basin analyses in
the Gangrinboche conglomerates.
modern Kailas and Nima soils
Kailas paleosols
1
Nima paleosols
GEODYNAMIC IMPLICATIONS
0
δ13C VPDB
–1
decreasing soil
respiration
–2
–3
–4
–5
–6
–7
increasing evaporation
decreasing temperature
–8
–25
–20
–15
–10
–5
δ18O VPDB (‰)
Figure 16. δ18O vs. δ13C values of modern soil carbonates and paleosol carbonate from the
upper Kailas Formation. Modern soil samples all come from >50 cm soil depth (see Table 4).
VPDB—Vienna Peedee belemnite.
ate and marl in the bulk of the Kailas Formation precludes documentation of a systematic
increase in paleoelevation through time.
Extension of Asian lithosphere during late
Oligocene–early Miocene time in the Kailas region is supported by recent work by Zhang et al.
(2010), who documented evidence for coeval
ductile extension, anatexis, and exhumation of
high-grade orthogneisses of Gangdese arc affinity from midcrustal depths in the footwall of the
Ayi Shan detachment (Fig. 1B). Restoration of
~65 km of slip on the dextral strike-slip Karakoram fault (Murphy et al., 2000) places the
Ayi Shan directly west along the outcrop belt
that we have studied in the Kailas Range. Thus,
the Ayi Shan detachment fault projects into the
subsurface beneath the Kailas Range and provides a ready explanation for the presence of
extensional basin deposits in the Kailas Formation (Figs. 1B and 5B). Zhang et al. (2010) also
reported kinematic indicators that document
top-southeastward (~120°E) ductile shear on
the Ayi Shan detachment. Coupled with work
1356
by Valli et al. (2007), this would suggest that
exhumation of the Ayi Shan took place in a
transtensional tectonic setting, perhaps feeding
slip eastward from the southeastern end of the
Karakoram fault into either the Indus-Yarlung
suture zone (Valli et al., 2007) or southeastward
across the northern Himalayan thrust belt via
the Gurla Mandatha detachment (Murphy et al.,
2002). Either of these structural-kinematic interpretations would suggest that the Kailas basin
should be localized to the southeastern end of
the Karakoram fault. In fact, the Kailas Formation outcrop belt extends continuously >50 km
east of the end of the Karakoram fault and the
Gurla Mandatha dome, and observations by
Searle et al. (1990), Yin et al. (1999), and Aitchison et al. (2002) indicate that mid-Cenozoic
coarse clastic rocks correlative with the Kailas
Formation crop out along at least 1300 km of
the suture zone. This argues against the localized extension/transtension explanation for the
Kailas Formation, and suggests instead that the
mechanism for Kailas basin formation operated
Independent data from sediment provenance,
depositional facies, paleontology, biostratigraphy, geochronology, and petrology suggest
that the Indo-Eurasia collision occurred during
late Paleocene to early Eocene time (e.g., Garzanti et al., 1987; Searle et al., 1987; Rowley,
1998; di Sigoyer et al., 2000; DeCelles et al.,
2004; Najman et al., 2005; Zhu et al., 2005;
Leech et al., 2005; Chung et al., 2005; Green
et al., 2008) and that crustal shortening in the
Himalayan thrust belt propagated southward
from the suture zone beginning at that time.
Alternatively, Aitchison et al. (2002, 2007) suggested that the true collision did not commence
until ca. 34 Ma, in part based on the interpretation that the Gangrinboche conglomerates were
produced by crustal shortening associated with
the initial Indo-Eurasia collision. Our results
from the type area of the Gangrinboche facies
do not support the notion that these deposits record initial shortening owing to the collision of
India and Eurasia. It is plausible that the Kailas Formation records local tectonic processes
in southwestern Tibet, and that the remainder
of the Gangrinboche conglomerates along the
suture zone were indeed deposited in a contractional setting as suggested by Aitchison et al.
(2002). However, we suggest that a resolution to
this debate may lie in considering mechanisms
for extension in the Indus-Yarlung suture zone
within an overall convergent tectonic setting.
Crustal shortening and extension are not mutually incompatible in intercontinental collision
zones. For example, extension and shortening
operate side-by-side throughout the Mediterranean region, where foundering and retrograde
“rollback” of generally northward-subducting
remnants of Neotethyan oceanic lithosphere,
stranded between African continental promontories, causes upper-plate extension of the
European mainland behind major thrust belts
that propagate from Europe toward Africa
(Malinverno and Ryan, 1986; Doglioni et al.,
1997; Cavinato and DeCelles, 1999; Jolivet and
Faccenna, 2000; Spakman and Wortel, 2004;
Cavazza et al., 2004; Faccenna et al., 2004).
Such coupled extensional-contractional systems
are characteristic of convergent plate boundaries in which the rate of subduction exceeds the
rate of plate convergence (e.g., Royden, 1993;
Jolivet and Faccenna, 2000). Although we view
it as highly unlikely that subduction of oceanic
Geological Society of America Bulletin, July/August 2011
Oligocene–Miocene Kailas basin, southwestern Tibet
lithosphere was occurring along the Indus-Yarlung suture ~30 m.y. after the youngest known
marine sedimentary rocks were deposited (Garzanti et al., 1987; Willems et al., 1996; Rowley,
1998; Zhu et al., 2005; Green et al., 2008),
Mediterranean-style, convergent-margin tectonics controlled by relative rates of convergence
and subduction provide a useful framework for
understanding Oligocene–Miocene extension in
the southern part of the Eurasian plate.
If Indian continental lithosphere, several
hundred kilometers of which could have been
underthrust beneath the Lhasa terrane after the
initial Eocene hard collision and Neotethyan
slab break-off event (e.g., DeCelles et al., 2002;
Kohn and Parkinson, 2002; Guillot et al., 2003;
Ding et al., 2005; Kapp et al., 2007; Lee et al.,
2009), were to begin foundering into the mantle
such that a rolling hinge line migrated southward relative to Indian lithosphere (Ding et al.,
2003), then the subduction rate would equal the
sum of the absolute values of true Indo-Eurasia
convergence and the rate of hinge-line rollback.
Subduction rate would thus exceed the rate of
convergence, and the upper (Eurasian) plate
could be thrown into extension near the plate
boundary (Fig. 17). This process is analogous to
the mechanism driving upper-plate extension in
the Mediterranean, with the important distinction that no oceanic lithosphere was involved in
the Tibetan case.
Available data suggest that such a situation
existed in the Indus-Yarlung suture zone during
the middle Cenozoic. Molnar and Stock (2009)
calculated a middle Cenozoic (ca. 45–20 Ma)
convergence rate of ~59 mm/yr between northwestern India and Eurasia. The rate of southward rollback of the Indian plate hinge line can
be estimated at ~50 mm/yr from a regional,
southward sweep of magmatism over a north to
south distance of ~400 km from the Qiangtang
terrane to the southern Lhasa terrane between
ca. 32 Ma and 25 Ma (Fig. 18) (Ding et al.,
2003; Chung et al., 2005; Kapp, et al., 2007).
For ~8 m.y. during latest Eocene–Oligocene
time, the subduction rate would have been
~84% greater than the rate of convergence.
Capitanio et al. (2010) suggested a likely cause
of continental slab rollback and subduction:
Previously thinned continental lower crust and
lithosphere that have been stripped of uppercrustal rocks by the development of a thrust
belt would be dense enough to founder into the
mantle. In this case, upper-crustal material was
being actively scraped off of Indian lower crust
to form the Tethyan Himalayan thrust belt, leaving behind a dense, possibly eclogitized, lower
crust and mantle lithosphere. Southward rollback of the Indian plate evidently terminated
at ca. 25–20 Ma, when the subducting Indian
slab broke off and a pulse of high-K and adakitic magmatism occurred all along the southern Lhasa terrane (Fig. 17) (Coulon et al., 1986;
Miller et al., 2000; Mahéo et al., 2002; Ding et
al., 2003; Chung et al., 2003; Williams et al.,
2004; Gao et al., 2007; Aitchison et al., 2009;
Chung et al., 2009). The timing of southward
migration of Tibetan magmatism and the hypothetical slab break-off event bracket deposition
of the Kailas Formation and the other Gangrinboche conglomerates. The abundance of tuffs
and detrital zircons with ages of 24–26 Ma reflects this magmatic activity.
After Indian continental slab break-off (ca.
20 Ma), the system would have reverted to a
“hard collisional” mode, with all of the convergence velocity being accommodated by
underthrusting of Indian lithosphere and crustal
shortening (Fig. 17). Early to middle Miocene
activation of several major thrust systems in
the Himalaya, including the Main Central,
Ramgarh, and Great Counter thrusts (which together accommodate at least 300 km of shortening), reflects this new mode of hard collision.
This explanation for the kinematics of the
Kailas basin is advantageous because it does
not require post-Eocene oceanic subduction outboard of an irregular northern Indian continental
margin; it can be reconciled with regional petrologic, metamorphic, structural, and sedimentological data sets; and a drastic revision in the
timing of initial Indo-Eurasia collision (Aitchison et al., 2007) is unnecessary. Moreover, the
proposed mechanism is consistent with recent
tomographic studies of the upper mantle beneath
Tibet, where large, fast P-wave anomalies have
been detected and interpreted as fragments of
foundered Indian lithosphere (van der Voo et al.,
1999; Hafkenscheid et al., 2006; Li et al., 2008;
Replumaz et al., 2010a, 2010b). Replumaz et al.
(2010a) noted that the distribution of fast anomalies beneath western Tibet is consistent with two
slab break-off events, and their reconstruction
suggests that these events were coeval with the
events we postulate (Eocene and late Oligocene–
early Miocene) based on surface data. Similar
timing of slab break-off and/or lithosphere removal events is derived from petrological studies
of Eocene–Miocene igneous rocks in the Lhasa
terrane (e.g., Miller et al., 1999; Chung et al.,
2009; Lee et al., 2009).
Numerous workers have suggested that the
development and subsequent exhumation to
midcrustal depths of eclogites in the northern
Himalayan thrust belt during Eocene time required initial subduction by slab-pull forces,
followed by buoyant rise when the Neotethyan
oceanic lithospheric slab broke off and foundered into the mantle (Chemenda et al., 2000;
O’Brien et al., 2001; Kohn and Parkinson, 2002;
de Sigoyer et al., 2000; Guillot et al., 2003;
Leech et al., 2005). After oceanic slab break-off,
all convergence velocity contributed directly
to crustal shortening (Chemenda et al., 2000;
Guillot et al., 2003). Our explanation of middle
Cenozoic tectonics along the Indus-Yarlung suture suggests that a second significant episode
of slab break-off occurred during early Miocene time. Unlike the Eocene break-off event,
the Miocene event involved continental Indian
lithosphere. It is possible that the Indo-Eurasian
collisional orogeny operated in two different modes, depending on whether subducted
Tethyan/Indian lithosphere was foundering by
rollback (soft collision), or underthrusting at
low angle beneath Tibet (hard collision; Fig.
17). During episodes of rollback and slab foundering (pre–45 Ma and ca. 32–20 Ma), the rate
of subduction exceeded the convergence rate,
resulting in tectonic neutrality or upper-plate
extension. During episodes of underthrusting
(ca. 45–32 Ma and ca. 20–0 Ma), all of the convergence velocity was thrown into shortening of
the Indian and Eurasian plates. This idea could
be tested with a careful assessment of the timing
of north-south crustal shortening and extension
episodes in the Himalaya and Lhasa terrane.
CONCLUSIONS
The Kailas Formation is late Oligocene to
early Miocene in age, and was deposited in
alluvial-fan, low-sinuosity fluvial, and deep
lacustrine depositional systems. Kailas Formation lakes were narrow, deep, and localized
along the Indus-Yarlung suture zone. The bulk
of the Kailas Formation was derived from the
Gangdese arc to the north, upon which it was
deposited; southerly sediment sources in the
hanging wall of the Great Counter thrust and
the Tethyan Himalaya became available only
during the last stages of deposition. Abundant
organic material (plant fragments, disseminated
kerogen, fish and amphibian fossils) and sparse
palynological data suggest that Kailas lakes
were relatively warm and tropical. In contrast,
the uppermost part of the stratigraphic section
contains abundant paleosol carbonate and redbed paleosols, similar to previously documented
Oligocene–Miocene lithofacies in the central
part of the Tibetan Plateau from which we have
obtained carbon and oxygen isotope data that
are consistent with very high paleoelevation by
late Oligocene time. We tentatively suggest that
the Kailas basin developed at comparatively low
elevations until deposition of the uppermost part
of the Kailas Formation.
The Kailas basin developed along the IndusYarlung suture zone contemporaneously with
extension in the Ayi Shan and possibly oblique
Geological Society of America Bulletin, July/August 2011
1357
DeCelles et al.
LT
20–15 Ma
QT
40 mm/yr
Return to northward
underthrusting mode
TH
KB
LT
25–20 Ma
Hard
collision
KB
GCT
MCT
QT
60 mm/yr
KB
~50 mm/yr
TH
32–25 Ma
LT
QT
60 mm/yr
Greater Indian slab rollback
and upper-plate extension
~30 mm/yr
Hard
collision
45–32 Ma
60 mm/yr
Northward underthrusting of Greater
Indian lithosphere beneath Tibet
QT
LT
55–45 Ma
~60 mm/yr
Tethyan slab rollback, followed by
break-off and foundering
BS
IYS
LT
55 Ma
100 mm/yr
Soft
collision
TH
IYS
E
QT
Hard
collision
Figure 17. Schematic northsouth cross sections showing
development of upper-plate
extension in southern Tibet in
response to Tethyan and Indian slab rollback and breakoff events, as discussed in text.
The Kailas basin (KB) formed
during the 32–25 Ma and 25–
20 Ma frames. Abbreviations
as follows: GMA—Gangdese
magmatic arc; MCT—Main
Central thrust; GCT—Great
Counter thrust; IYS—IndusYarlung suture zone; BS—Bangong suture zone; LT—Lhasa
terrane; QT—Qiangtang terrane; TH—Tethyan Himalaya.
Approximate times of hard and
soft collision are indicated by
solid and dashed bars at right,
respectively. Overturned config uration of the subducted
Greater Indian lithosphere in
the top two frames is after
Replumaz et al. (2010a).
Soft
collision
Greater Indian
slab break-off
GMA
60 Ma
1358
100 mm/yr
Geological Society of America Bulletin, July/August 2011
Precollision
Oligocene–Miocene Kailas basin, southwestern Tibet
700
500
Sl a
400
bf
300
100
i ng
200
en
Slab
rollba
ck
la tt
Distance north from IYS (km)
600
0
–100
0
10
20
30
40
50
60
70
Age (Ma)
Figure 18. U-Pb and 40Ar/ 39Ar ages of extrusive rocks in western Tibet (west of ~86.5°E)
plotted against their arc-normal distance from the Indus-Yarlung suture (IYS). Figure is
based on data compiled by P. Kapp and J. Volkmer (sources: Aitchison et al., 2009; Ding
et al., 2003, 2007; Kapp et al., 2003, 2005; Lee et al., 2009; Li et al., 2006; Matte et al., 1996;
Miller et al., 1999, 2000; Nomade et al., 2004; Williams et al., 2001, 2004; Zhao et al., 2009).
slip along the early Karakoram fault. Kailas
Formation stratigraphy is consistent with deposition in an extensional basin elongated parallel
to the suture zone. The basin also formed during
an episode of southward-migrating magmatism,
from the southern Qiangtang terrane to the rejuvenated Gangdese arc. Our results suggest
that extension in the Kailas basin was related
to southward rollback of the hinge line in the
subducting/underthrusting Indian continental
lithosphere. Insofar as chronostratigraphically
equivalent Oligocene–Miocene syntectonic
strata, locally associated with high-K extrusive
rocks, are distributed along nearly the entire
Indus-Yarlung suture zone, mid-Cenozoic extension of the southern margin of the Eurasian
plate may have been a regional phenomenon. If
true, this would suggest that the Indo-Eurasian
collision involved alternating “hard” and “soft”
modes, with hard collision characterized by
shallow underthrusting of India and regional
shortening, and soft collision associated with
foundering and rollback of the underthrusting Indian/Tethyan plate and local upper-plate
extension.
ACKNOWLEDGMENTS
The National Science Foundation (NSF) Tectonics
Program provided funding for this research. We thank
Ding Lin of the Chinese Academy of Sciences for
logistical assistance. Gerald Waanders performed the
palynological and organic matter analyses. We bene-
fited from numerous discussions with M.A. Murphy,
R. Zhang, J. Volkmer, and J. Saylor, who have collaborated with us on the tectonics of Cenozoic basin development in southwestern Tibet. Murphy and Zhang
kindly provided access to a preprint on the tectonics
of the Ayi Shan. We thank Eduardo Garzanti and Laurent Jolivet for thorough and constructive reviews,
and GSA Bulletin Associate Editor Stefano Mazzoli
and Editor Nancy Riggs for additional guidance that
helped us to improve the manuscript.
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