The Cretaceous Evolution of the Lhasa Terrane, Southern Tibet
Item Type text; Electronic Dissertation
Authors Leier, Andrew
Publisher The University of Arizona.
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Link to Item http://hdl.handle.net/10150/193796 THE CRETACEOUS EVOLUTION OF THE LHASA TERRANE,
SOUTHERN TIBET
by
Andrew Lawrence Leier
A Dissertation Submitted to the Faculty of the
DEPARTMENT OF GEOSCIENCES
In Partial Fulfillment of the Requirements For the Degree of
DOCTOR OF PHILOSOPHY
In the Graduate College
THE UNIVERSITY OF ARIZONA
2 0 0 5 2
THE UNIVERSITY OF ARIZONA GRADUATE COLLEGE
As members of the Dissertation Committee, we certify that we have read the dissertation
prepared by: Andrew L. Leier
entitled: The Cretaceous Evolution of the Lhasa Terrane, Southern Tibet
and recommend that it be accepted as fulfilling the dissertation requirement for the
Degree of: Doctor of Philosophy
______Date: 11/07/05 Peter DeCelles
______Date: 11/07/05 Paul Kapp
______Date: 11/07/05 Jay Quade
______Date: 11/07/05 George Gehrels
______Date: 11/07/05 George Zandt
Final approval and acceptance of this dissertation is contingent upon the candidate’s submission of the final copies of the dissertation to the Graduate College.
I hereby certify that I have read this dissertation prepared under my direction and recommend that it be accepted as fulfilling the dissertation requirement.
______Date: 11/07/05 Dissertation Director: Peter G. DeCelles 3
STATEMENT BY THE AUTHOR
This dissertation has been submitted in partial fulfillment of requirements for an advanced degree at The University of Arizona and is deposited in the University Library to be made available to borrowers under rules of the Library.
Brief quotations from this dissertation are allowable without special permission, provided that accurate acknowledgement of source is made. Requests for permission for extended quotation from or reproduction of this manuscript in whole or in part may be granted by the head of the major department of the Dean of the Graduate College when in his or her judgment the proposed use of the material is in the interests of scholarship. In all other instances, however, permission must be obtained from the author.
Signed: Andrew Leier
4
ACKNOWLEDGEMENTS
I owe a great deal to many individuals, more than I can mention in the limited space provided. I would like to thank my advisor, Dr. Pete DeCelles, for his great tutelage and unending patience. I would also like to thank Dr. DeCelles for all of the support he provided, financial and otherwise; I owe him a lot. I am grateful to the members of my committee for all of the help they provided: Dr. Paul Kapp, who led me through Tibet and was always available for questions and discussion; Dr. Jay Quade, who introduced me to stable isotopes and many other aspects of geology; Dr. George Gehrels, who patiently showed me the methods of detrital zircon analysis; and Dr. George Zandt, who introduced me to the processes acting within and below the lithosphere. Many people have assisted me during my time at The University of Arizona. Special thanks to Dave Barbeau for his assistance and instruction and D Robinson for all of her help; I also thank my officemate Aaron Martin for helpful discussions. I want to say ‘thank you’ to Shundong He and Dan Eisenberg for their assistance with field- work. I would like to thank many more people individually and at length, but given the spatial constraints I will simply list their names here: John Chesley, David Dettman, Matt Fabijanic, Majie Fan, Jen Fox, Facundo Fuentes, Jerome Guynn, Dick Hay, Brian Horton, Jessica Kapp, Ding Lin, Alex Pullen, Joel Saylor, Bob Scott, Victor Valencia, John Volkmer, Ross Waldrip, and the Tibetans who helped me while I was in their territory. In addition, there are others to whom I am indebted: Stacie Gibbins, Alex Bump, Nadine McQuarrie, Ofori Pearson, J. Mike Boyles, Ron Steel, Edward Cotter, Mary-Beth Gray, Jim Steidtmann, Nathan English, and Andrea Fildani. I sincerely thank all of the people listed in this and the preceding paragraph. If there are other names that have slipped my mind, I apologize, but please believe me when I say that I am grateful for the help you provided. My family has been very supportive during this time. I thank my brother Joe and my sister Rachel for their encouragement, interest, and support. I would like to thank my parents, Carl and Jolene Leier, for their love, patience, interest, encouragement, and financial assistance. I would also like to thank my parents for all of the support they have provided. 5
TABLE OF CONTENTS
LIST OF FIGURES ...... 9
LIST OF TABLES ...... 11
ABSTRACT...... 12
CHAPTER 1: INTRODUCTION...... 13
CHAPTER 2: FACIES, COMPOSITION, DETRITAL ZIRCON GEOCHRONOLOGY AND TECTONIC IMPLICATIONS OF LOWER CRETACEOUS STRATA IN THE LHASA TERRANE OF SOUTHERN TIBET ...... 18 ABSTRACT...... 18 INTRODUCTION ...... 19 SETTING...... 21 FACIES...... 23 LITHOFACIES ASSOCIATIONS – NORTHERN AREA ...... 23 Offshore Marine Lithofacies Association...... 24 Description...... 24 Interpretation ...... 25 Northern Fluvial Lithofacies Association...... 25 Description...... 25 Interpretation ...... 27 Northern Limestone Lithofacies Association ...... 28 Description...... 28 Interpretation ...... 28 LITHOFACIES ASSOCIATIONS – SOUTHERN AREA...... 29 Lagoonal Lithofacies Association...... 29 Description...... 29 Interpretation ...... 30 Shoreface Lithofacies Association...... 30 Description...... 30 Interpretation ...... 31 Southern Fluvial Lithofacies Association...... 32 Description...... 32 Interpretation ...... 33 Southern Limestone Lithofacies Association ...... 33 Description...... 33 Interpretation ...... 34 PALEOCURRENT MEASUREMENTS ...... 35 PETROGRAPHY ...... 36 6
TABLE OF CONTENTS - Continued
Northern Study Area...... 36 Southern Study Area...... 37 DETRITAL ZIRCON GEOCHRONOLOGY...... 38 Methods...... 38 Results...... 40 REGIONAL INTERPRETATIONS...... 41 Collision to Mid Neocomian...... 41 Northern Study Area ...... 41 Southern Study Area ...... 42 Mid Neocomian to Early Aptian (~130-120 Ma) ...... 43 Northern Study Area ...... 43 Southern Study Area ...... 44 Relationship Between Lower Cretaceous Conglomerate Units...... 44 Aptian-Albian (~120-100 Ma)...... 45 Northern Study Area ...... 45 Southern Study Area ...... 45 DISCUSSION...... 46 Initial Stage of Basin Development ...... 48 Second Stage of Basin Development...... 50 CONCLUSIONS...... 53
CHAPTER 3: THE TAKENA FORMATION OF THE LHASA TERRANE, SOUTHERN TIBET: THE RECORD OF A RETRO-ARC FORELAND BASIN ...... 85 ABSTRACT...... 85 INTRODUCTION ...... 86 REGIONAL GEOLOGY...... 89 Stratigraphy and Age Control ...... 89 Tectonic History...... 90 SEDIMENTOLOGY OF THE TAKENA FORMATION ...... 91 Lithofacies...... 91 Lithofacies Associations and Formation-scale Architecture ...... 91 Lower Marine Limestone Lithofacies Association – Description...... 92 Lower Marine Limestone Lithofacies Association – Interpretation ...... 93 Marginal Marine Siltstone Lithofacies Association – Description...... 94 Marginal Marine Siltstone Lithofacies Association – Interpretation...... 94 Lower Fluvial Lithofacies Association – Description ...... 95 Lower Fluvial Lithofacies Association – Interpretation...... 96 Middle Limestone Lithofacies Association – Description ...... 98 Middle Limestone Lithofacies Association – Interpretation...... 98 Middle Siltstone Lithofacies Association – Description ...... 98 Middle Siltstone Lithofacies Association – Interpretation...... 99 7
TABLE OF CONTENTS – Continued
Upper Fluvial Lithofacies Association – Description ...... 100 Upper Fluvial Lithofacies Association – Interpretation...... 101 PALEOCURRENT DATA...... 101 PETROGRAPHY ...... 102 Description...... 102 Interpretation...... 103 Maqu Provenance...... 104 SUBSIDENCE...... 105 PALEOGEOGRAPHY AND TECTONIC SETTING...... 107 Lower Strata of the Takena Formation ...... 108 Middle Strata of the Takena Formation...... 109 Upper Strata of the Takena Formation...... 110 Deformation of the Takena Formation ...... 110 The Migration of the Deformation Front...... 111 DISCUSSION...... 112 Where is the Cretaceous Fold-Thrust Belt?...... 112 Implications for Plateau Development...... 113 Previous Paleogeographic Interpretations...... 113 CONCLUSIONS...... 114
CHAPTER 4: DETRITAL ZIRCON GEOCHRONOLOGY OF PHANEROZOIC SEDIMENTARY STRATA EXPOSED WITHIN THE LHASA TERRANE OF SOUTHERN TIBET...... 137 ABSTRACT...... 137 INTRODUCTION ...... 138 REGIONAL SETTING ...... 140 SAMPLES...... 142 Carboniferous Strata ...... 143 Jurassic...... 143 Lower Cretaceous-north-Duba Formation...... 144 Lower Cretaceous-south-Chumulong Formation ...... 144 Lower Cretaceous at Nam Co...... 145 Upper Cretaceous-Takena Formation...... 145 METHODS ...... 146 RESULTS ...... 147 Carboniferous...... 147 Jurassic...... 148 Lower Cretaceous at Nam Co...... 148 Lower Cretaceous-south-Chumulong Formation ...... 149 Lower Cretaceous-north-Duba Formation...... 149 Upper Cretaceous-Takena Formation...... 150 8
TABLE OF CONTENTS - Continued DISCUSSION...... 150 General Characteristics of the Detrital Zircon Ages...... 150 “Lower Cretaceous” strata at Nam Co...... 151 Jurassic Strata...... 152 Upper Cretaceous Strata ...... 154 Carboniferous Strata and Gondwanaland ...... 156 A Northern and Southern Provenance Signature...... 156 Carboniferous Strata as a Sediment Source...... 158 CONCLUSIONS...... 159
CHAPTER 5: SEAWATER 187OS/188OS: THE CENTRAL ANDEAN DISCONNECT...... 192 ABSTRACT...... 192 INTRODUCTION ...... 192 OSMIUM AND SEAWATER...... 193 THE PROBLEM...... 195 THE CENTRAL ANDES...... 196 SAMPLES AND RESULTS...... 197 DISCUSSION...... 199 CONCLUSION...... 200
CHAPTER 6: MOUNTAINS, MONSOONS AND MEGAFANS...... 206 ABSTRACT...... 206 INTRODUCTION ...... 206 METHODS ...... 208 RESULTS ...... 209 INTERPRETATION...... 211 CLIMATE PATTERNS AND FLUVIAL MEGAFANS...... 212 APPLICATIONS TO THE STRATIGRAPHIC RECORD ...... 213 DISCUSSION AND CONCLUSION...... 214
REFERENCES...... 218
9
LIST OF FIGURES
FIGURE 2.1: Study Area...... 55 FIGURE 2.2: Hypothesized basins...... 56 FIGURE 2.3: Regional stratigraphy ...... 57 FIGURE 2.4: Stratigraphic symbols...... 58 FIGURE 2.5: Duba Formation measured section...... 59 FIGURE 2.6: Marine strata...... 61 FIGURE 2.7: Duba Formation...... 62 FIGURE 2.8: Langshan Formation...... 63 FIGURE 2.9: Lower Cretaceous in the Penbo area ...... 64 FIGURE 2.10: Lagoonal strata ...... 66 FIGURE 2.11: Chumulong Formation ...... 67 FIGURE 2.12: Penbo Member ...... 68 FIGURE 2.13: Sandstone composition...... 69 FIGURE 2.14: Petrography ...... 70 FIGURE 2.15: Detrital zircon data ...... 71 FIGURE 2.16: Paleogeography and tectonic setting ...... 72 FIGURE 2.17: Volcanism...... 74 FIGURE 2.18: Subsidence...... 75
FIGURE 3.1: Location map and geology of the study area...... 116 FIGURE 3.2: Proposed tectonic models...... 117 FIGURE 3.3: Generalized Cretaceous stratigraphy...... 118 FIGURE 3.4: Stratigraphic symbols...... 119 FIGURE 3.5: The lower part of the Takena Formation in the Penbo area ...... 120 FIGURE 3.6: The upper part of the Takena Formation in the Penbo area ...... 122 FIGURE 3.7: Photos of the Penbo Member of the Takena Formation...... 124 FIGURE 3.8: Photograph of marginal marine clastic deposits...... 125 FIGURE 3.9: Lower part of the Takena Formation...... 126 FIGURE 3.10: Upper part of the Takena Formation ...... 127 FIGURE 3.11: Petrography ...... 128 FIGURE 3.12: Sandstone composition...... 129 FIGURE 3.13: Subsidence...... 130 FIGURE 3.14: Paleogeography ...... 131 FIGURE 3.15: Takena and Linzizong Formations...... 132
FIGURE 4.1: Map of Area...... 162 FIGURE 4.2: Regional Setting ...... 163 FIGURE 4.3: Takena Formation concordia diagram...... 164 FIGURE 4.4: Concordia diagram of non-Takena Formation samples ...... 165 FIGURE 4.5: Distribution of detrital zircon U-Pb ages ...... 166 FIGURE 4.6: Takena Formation detrital zircon U-Pb ages...... 168 FIGURE 4.7: Regional detrital zircon comparison...... 169 10
LIST OF FIGURES - Continued
FIGURE 4.8: U-Pb ages of igneous rocks...... 170 FIGURE 4.9: Northern and southern Tibet...... 171 FIGURE 4.10: Carboniferous and Takena strata...... 173
FIGURE 5.1: Cenozoic osmium seawater values and the central Andes...... 201 FIGURE 5.2: Osmium isotopic compositions and concentrations...... 202 FIGURE 5.3: Total organic carbon and rhenium concentrations ...... 203 FIGURE 5.4: Osmium isotopic composition of seawater ...... 204
FIGURE 6.1: Images of modern fluvial megafans ...... 215 FIGURE 6.2: Studied rivers and climate ...... 216 FIGURE 6.3: Comparison of rivers...... 217
11
LIST OF TABLES
TABLE 2.1: Lithofacies...... 76 TABLE 2.2: Point-counting parameters ...... 77 TABLE 2.3: Basins...... 78 TABLE 2. DR1: Point-counting data...... 79 TABLE 2. DR2: Detrital zircon data ...... 80
TABLE 3.1: Lithofacies...... 133 TABLE 3.2: Point-counting parameters ...... 134 TABLE 3. DR1: Point-counting data...... 135
TABLE 4. DR1: Samples and locations ...... 174 TABLE 4. DR2: Detrital zircon data ...... 175
TABLE 5.1: Sample Composition...... 205
12
ABSTRACT
The Tibetan plateau is arguably the most important geological feature on Earth, yet its formation and evolution are poorly understood. This investigation utilizes
Cretaceous sedimentary strata exposed in the Lhasa terrane of southern Tibet in order to constrain the paleogeography and tectonic setting of the region prior to the Indo-Asian collision. Lower Cretaceous strata consist of clastic sedimentary units that were deposited in shallow marine and fluvial environments. In northern Lhasa these sediments were deposited in a peripheral foreland basin that formed in response to the Lhasa-Qiangtang collision. The lower Cretaceous sediments in southern Lhasa are quartzose and were derived from cover strata exposed by local uplifts. A marine limestone of Aptian-Albian age overlies the lower Cretaceous clastic strata and was deposited in a shallow continental seaway. The paleogeography of the Lhasa terrane during deposition of the carbonate units was dominated by the effects of the Lhasa-Qiangtang collision, although other conditions, such as a high eustatic sea-level, influenced sedimentation as well. The
Upper Cretaceous Takena Formation is composed of a basal member of marine limestone
and an overlying member of fluvial red beds. The arkosic strata of the Takena Formation
were deposited in a retro-arc foreland basin that formed to the north of the Gangdese
magmatic arc. Collectively, the Cretaceous sedimentary strata indicate significant
tectonic activity occurred in southern Tibet prior to the Indo-Asian collision. Moreover,
the data suggest the crust of southern Tibet was thickened and possibly at high elevations
before the Cenozoic. 13
CHAPTER 1: INTRODUCTION
Tectonic processes play a fundamental role in shaping the geography and
physiographic features of the Earth. In no other region is this more evident than in the
Tibetan plateau, which currently stands >5 km above sea-level (Fielding, 1996). This
region is the archetype of continent-continent collisions, and its uplift and weathering are
thought to have influenced both local and global systems (e.g., Raymo et al., 1988,
Raymo and Ruddiman, 1992; Molnar et al., 1993). Despite its importance, much about
the geological history of the Tibetan plateau remains unknown. Traditionally, the high
topography and thickened crust of the plateau are believed to be the result of the
Cenozoic Indo-Asian collision exclusively. Almost all models pertaining to the growth of
the Tibetan plateau implicitly assert that the region was at or near sea-level and relatively
undeformed prior to the Indo-Asian collision. New data from the region, however, have
raised serious questions about the validity of this assumption.
The Mesozoic history of southern Tibet involved significant tectonic activity (Yin and Harrison, 2000). Several recent studies have suggested that a large proportion of
Tibet’s modern topography is actually inherited from these pre-Cenozoic tectonic events
(e.g., Murphy et al., 1997). The Lhasa terrane of southern Tibet is the southernmost terrane of Tibet and was the last of several continental fragments to accrete onto southern
Asia before the Indo-Asian collision, doing so sometime in Late Jurassic-Early
Cretaceous time (e.g., Allégre et al., 1984; Dewey et al., 1988). At the same time that the
Lhasa terrane was accreting onto southern Asia, subduction of Neotethyan oceanic lithosphere commenced along its southern margin; northward subduction of oceanic crust 14
would continue throughout the Cretaceous period (Yin and Harrison, 2000). Before we
can understand how the Tibetan plateau has formed, we need to understand what impact
these pre-Cenozoic events and processes had on southern Tibet.
Several hypotheses have been proposed that describe the tectonic setting and
paleogeography of the Lhasa terrane immediately before the Indo-Asian collision. The
first hypothesis posits that the Lhasa terrane was most similar to an Andean-style
subduction zone, with an arc along the southern margin of the terrane and compressional faulting in the retro-arc region (e.g., England and Searle, 1986). The second hypothesis implies that the Lhasa terrane experienced extension during the Cretaceous, and therefore, most closely resembled a back-arc basin (e.g., Zhang et al., 2004). And the final hypothesis postulates that the accretion of the Lhasa terrane onto southern Asia resulted in crustal shortening, uplift, and the formation of a peripheral foreland basin
(e.g., Leeder et al., 1988; Murphy et al., 1997). Because these hypotheses involve different styles of deformation, determining which of these models is correct is critical for understanding the condition of the crust in southern Tibet immediately before the
Indo-Asian collision. Although these models make specific predictions about the nature of the Cretaceous sedimentary strata in the Lhasa terrane, previous studies of these rocks have not yielded enough information to adequately test these hypotheses.
This investigation focuses on the Cretaceous sedimentary record of the Lhasa terrane of southern Tibet. The purpose is to reconstruct the Cretaceous paleogeography and tectonic setting of the region and then use this information to test the tectonic models that have been proposed for this area. The data within this dissertation include, but are 15
not limited to, >8 km of measured stratigraphic sections; >900 paleocurrent measurements; >80 point-counted thin sections; and >700 detrital zircon U-Pb ages.
The following chapters consist of five separate manuscripts with minor overlap in content. The first manuscript, entitled Facies, composition, detrital zircon geochronology and tectonic implications of lower Cretaceous strata in the Lhasa terrane of southern
Tibet, presents information on the sedimentological, stratigraphic and geochronological characteristics of lower Cretaceous deposits within the Lhasa terrane. The paleogeography of southern Tibet during the Early Cretaceous was dominated by the Late
Jurassic-Early Cretaceous collision between the Lhasa terrane and the Qiangtang terrane.
This manuscript will be submitted to Journal of Sedimentary Research.
The second manuscript, The Takena Formation of the Lhasa terrane, southern
Tibet: the record of a Late Cretaceous retro-arc foreland basin, describes the sedimentological and stratigraphic characteristics of the upper Cretaceous strata within southern Tibet. These strata provide perhaps the most complete record of the conditions within southern Asia immediately before the Indo-Asian collision. The sedimentological evidence indicates they were deposited in a retro-arc foreland basin that was located to the north of a fold-thrust belt. This reconstruction implies that parts of southern Tibet had thickened crust before the Cenozoic collision with the Indian subcontinent. This manuscript will be submitted to Geological Society of America Bulletin.
The third manuscript is entitled, Detrital zircon geochronology of Phanerozoic
sedimentary strata exposed within the Lhasa terrane of southern Tibet. The >700 detrital
zircon U-Pb ages presented in this manuscript help to address several enigmatic aspects 16
of the pre-Cenozoic tectono-sedimentary history of the Lhasa terrane. The data constrain
the maximum depositional ages for several important stratigraphic units in the Lhasa
terrane and suggest that the ages of some rock units in the area need to be reevaluated.
From a regional perspective, particular populations of detrital zircon U-Pb ages within
sandstone units may provide a means by which a northern or southern provenance can be
determined. This manuscript will be submitted to Sedimentology.
The fourth manuscript, entitled Seawater 187Os/188Os: the central Andean
disconnect, examines the role tectonic uplift plays in controlling the chemical
composition of seawater. The uplift and weathering of the Tibetan plateau and the
Himalaya were once thought to control the 187Os/188Os composition of seawater;
however, recent studies indicate this is not the case. The central Andes of Bolivia are an
ideal candidate for a major 187Os source, and therefore the uplift and weathering of this region may have controlled the 187Os/188Os of sweater during Cretaceous-Tertiary time.
Our analyses of the Re and Os content of rocks exposed in the Central Andes suggest this area did not act as a source of 187Os. This greatly limits the list of possible suspects and
suggests that the seawater 187Os/188Os record and continental weathering may be
decoupled. This manuscript will be submitted to Journal of South American Earth
Sciences.
The fifth manuscript, Mountains, Monsoons and Megafans, examines the effects
of orogenic uplift and climate on fluvial systems. In some instances, rivers draining
regions like the Himalaya-Tibetan orogenic system create unusually large fan-shaped
bodies of sediment that are referred to as fluvial megafans. Information from satellite 17
imagery, monthly discharge and precipitation records, digital elevation models, and other
sources, indicates that the formation of fluvial megafans requires particular climatic conditions. Specifically, modern fluvial megafans in actively aggrading basins are produced by rivers that undergo moderate to extreme seasonal fluctuations in discharge that result from highly seasonal precipitation patterns. This implies that ancient fluvial megafan deposits may be useful for reconstruction paleoclimatic conditions. This manuscript has been published in Geology.
18
CHAPTER 2: FACIES, COMPOSITION, DETRITAL ZIRCON
GEOCHRONOLOGY AND TECTONIC IMPLICATIONS OF LOWER
CRETACEOUS STRATA IN THE LHASA TERRANE OF SOUTHERN TIBET
ABSTRACT
Southern Tibet was tectonically active before the Indo-Asian collision, although the extent and style of the tectonism remains controversial. We examined the lower
Cretaceous sedimentary strata exposed within the Lhasa terrane of southern Tibet in order to determine the depositional history of these units and from this, reconstruct the
Early Cretaceous paleogeography of the region and evaluate the proposed models of
Early Cretaceous tectonism. Lower Cretaceous strata in the northern portion of the
Lhasa terrane are coarse-grained and lithic-rich and were deposited in shallow marine and meandering-stream fluvial environments. These strata were derived from local volcanic units and sedimentary cover strata located in the northern Lhasa and southern
Qiangtang terranes. Sandstone exposed in the northern Lhasa terrane is composed of abundant volcanic material and the majority of the detrital zircons have U-Pb ages between 120 and 140 Ma. Lower Cretaceous sedimentary strata in the southern portion of the Lhasa terrane consist of mudstone, quartzose sandstone and local quartzite pebble-conglomerate that were deposited in marginal marine and fluvial environments.
Sandstone in the southern portion of the Lhasa terrane have populations of detrital zircons with ages between 140-160 Ma, 500-580 Ma, 620-710 Ma, 850-950 Ma and 19
1000-1300 Ma. Sandstone composition and U-Pb age data of detrital zircons suggest the
lower Cretaceous conglomerate exposed within the northern and southern portion of the
Lhasa terrane were deposited at different times and possibly share little genetic
relationship. All of the lower Cretaceous clastic strata are overlain by a widespread
Aptian-Albian limestone that was deposited in a shallow carbonate sea that contained
rudist patch reefs and muddy inter-reef environments. The data indicate that the Lhasa
terrane is best characterized as having been a peripheral foreland basin during the latest
Jurassic-earliest Cretaceous, which formed in response to the collision between the
Lhasa and Qiangtang terranes. Overprinted on the peripheral foreland basin setting is arc volcanism associated with northward subduction along the southern margin of the
Lhasa terrane. High eustatic sea-level and crustal loading along the Bangong suture, and possibly some dynamic subsidence, provided the accommodation for deposition of the
Aptian-Albian limestone. The lower Cretaceous stratigraphic record of the Lhasa terrane of southern Tibet indicates significant tectonic activity occurred during this time and is consistent with recent hypotheses suggesting parts of southern-central Tibet were deformed and above sea-level before the Indo-Asian collision.
INTRODUCTION
Although the Tibetan plateau is often invoked as the type example of deformation resulting from a continent-continent collision, the pre-collisional tectonic history of this area is, at best, unclear. This is not a trivial issue given the importance inherited tectonic features have in controlling orogenic evolution, and the fact that 20 southern Asia is thought to have been the site of significant tectonic activity prior to the
Indo-Asian collision (e.g., Sengör and Natal’in, 1996; Yin and Harrison, 2000). Along the southern margin of Asia, numerous collisional events preceded the Cenozoic Indo-
Asian collision as the constituent continental fragments of the Tibetan plateau accreted onto the continent (Figure 2.1; Allégre et al., 1984; Dewey et al., 1988; Sengör and
Natal’in, 1996). The southernmost terrane of Tibet, the Lhasa terrane, was the last of these continental fragments to accrete onto southern Asia before the collision with
India, doing so sometime time during the Late Jurassic - Early Cretaceous (Allégre et al., 1984; Dewey et al., 1988; Yin and Harrison, 2000). Knowledge of the tectonics and paleogeography of the Lhasa terrane during this general time period is critical for determining the initial conditions that existed in southern Asia before the collision, particularly with regard to whether or not parts of Tibet had already attained high elevations by the end of the Cretaceous (e.g., Murphy et al., 1997). However, the scarcity of data pertaining to the Early Cretaceous history of the Lhasa terrane make it difficult to test even the most general paleogeographic reconstructions.
The compositions, thicknesses, provenance, depositional environments, and sedimentary histories of lower Cretaceous strata exposed in the Lhasa terrane are largely unknown. For example, conglomeratic units mapped as lower Cretaceous are exposed in multiple locations throughout the Lhasa terrane, but the data needed to test whether these deposits are genetically related, or even of Early Cretaceous age, have not been reported. From the reconnaissance work that has been done in the region, two very different hypotheses have been proposed to explain the depositional setting of the 21
region during the Early Cretaceous. One postulates that the accretion of the Lhasa
terrane onto the southern margin of Asia produced a collisional orogen; thus the lower
Cretaceous sedimentary strata were deposited in a peripheral foreland basin (Figure 2.2;
Leeder et al., 1988; Murphy et al., 1997). Acknowledging that a small peripheral foreland basin may have developed, the other model proposes that the majority of the lower Cretaceous sediments were deposited within an extensional back-arc basin, located behind a north-directed subduction zone (Figure 2.2; Zhang, 2000; Zhang et al.,
2004). The disparate topography and crustal thicknesses implied by these opposing
ideas have direct implications for determining the mechanisms involved in the uplift of
the Tibetan plateau.
In this paper we describe and interpret the lithologies, facies, regional
distribution, and provenance of lower Cretaceous strata within the eastern Lhasa terrane, and provide new and detailed paleogeographic reconstructions. Utilizing this
new information, we reexamine existing tectonic hypotheses of this region and propose
a variation of the peripheral foreland basin model for the Early Cretaceous evolution of
the area.
SETTING
The study area in the eastern Lhasa terrane is broadly divided into northern and
southern portions, with lake Nam Co acting as a general border between the two regions
(Figure 2.1). Lower Cretaceous strata in both the north and south contain a basal clastic succession and an upper limestone unit (Figure 2.3). Simplifying the stratigraphic nomenclature, the lower Cretaceous clastic deposits in the north are referred to as the 22
Duba Formation and the overlying limestone unit is called the Langshan Formation
(Yin et al., 1988). We divide the Duba Formation into two informal divisions called the
lower Duba Formation and the upper Duba Formation. In the south, the clastic
succession at the bottom of the lower Cretaceous strata is called the Chumulong
Formation and the upper limestone unit is represented by the Penbo Member of the
Takena Formation (Yin et al., 1988). The clastic sediments at the base of the lower
Cretaceous succession in the south overlie Jurassic limestone, shale, and local beds of volcanic material, whereas to the north, the underlying Jurassic units consist of marine shale and sandstone (Leeder et al., 1988; Yin et al., 1988). The clastic rocks of the Duba and Chumulong Formations are generally considered to be lower Cretaceous although some deposition may have taken place during the latest Jurassic (Yin et al., 1988). From north to south, the limestones at the top of the lower Cretaceous succession are slightly time transgressive (Figure 2.3; Yin et al., 1988; this study). The Penbo Member in the south was deposited during Aptian-Albian time and the Langshan Formation in the north was deposited from the late Barremian to the early Cenomanian (Figure 2.3;
Smith and Xu, 1988; Zhang et al., 2004; Leier, 2005). The Penbo Member is roughly
300 m thick, whereas the thickness of the Langshan Formation is poorly constrained and vary between hundreds of meters to >6 km (c.f., Leeder et al., 1988; Zhang, 2000).
Leeder et al. (1988) reported a ~1100 m thickness for the lower Aptain to upper Albian
Langshan Formation in the Duba area, which we use for the total thickness of the limestone. The lower Cretaceous limestone units are overlain by arkosic red beds of the 23
Takena Formation in the southern portion of the study area. In the north, the contact
between the limestone and the upper Cretaceous strata is not exposed.
The Lhasa terrane rifted from Gondwana during Late Permian-Early Triassic
time and migrated northward until Late Jurassic-Early Cretaceous time when it collided
with the southern margin of the Qiangtang terrane (Allégre et al., 1984; Dewey et al.,
1988; Yin and Harrison, 2000). The remnant of this event is the Bangong suture, an
east-west trending zone of ophiolitic fragments and deepwater deposits that mark the border between the Lhasa and Qiangtang terranes. Numerous Early Cretaceous calc- alkaline grantitoids are exposed in the northern half of the Lhasa terrane but their geodynamic significance is still controversial (c.f., Coulon et al., 1986; Pearce and Mei,
1988; Harris et al 1990). Northward subduction of Neotethyan oceanic lithosphere produced the Cretaceous – Tertiary Gangdese magmatic arc along the southern margin of the Lhasa terrane. During deposition of the lower Cretaceous strata, the Lhasa terrane was located at roughly 10 degrees north latitude (Achache et al., 1984; Chen et al.,
1993).
FACIES
Lithofacies in lower Cretaceous rocks of the Lhasa terrane are typical of carbonate and coarse-grained clastic rocks, and are well understood in terms of depositional processes. Individual lithofacies are described and interpreted in Table 2.1.
The following paragraphs focus on the genetic associations of lithofacies that can be more broadly interpreted in terms of depositional systems.
LITHOFACIES ASSOCIATIONS – NORTHERN AREA 24
Offshore Marine Lithofacies Association
Description
The offshore marine lithofacies association comprises the lower Duba
Formation (Figures 2.4, 2.5) and consists of black laminated and massive mudstone with very thin to thick beds of gray tabular sandstone and minor conglomerate
(lithofacies G1, S5, S7, F1; Table 2.1). This lithofacies association is present in the
Duba region, south of Langshan Mountain and also in the Lunpola area (Figure 2.1).
The mudstone contains siltstone laminae and is commonly bioturbated with horizontal to sub-horizontal burrows. Sandstone beds are typically 0.15 meters thick but vary from
1 cm to over 1 m, and commonly have sharp to erosional bases and sharp contacts with the overlying mudstone. Bivalve fragments are present within a few of these sandstone beds. Thinner sandstone beds tend to be fine- to very fine-grained and contain oscillatory current ripples and rare unidirectional current ripples (lithofacies S5, Table
2.1; Figure 2.6A). Sandstone beds greater than 5 cm thick are typically fine- to medium- grained and have plane-parallel laminae and oscillatory current ripples. Locally, the thicker beds are structureless and contain granules and pebbles. Individual sandstone beds lack internal gradations in grains size, but commonly have fine-grained caps.
Vertical, sub-vertical and sub-horizontal unlined burrows are common in all sandstone beds. A few sandstone beds are moderately eroded by scour surfaces associated with overlying clast-supported pebble conglomerate beds (Figure 2.6B). These conglomerate beds are well organized and contain rounded and well-rounded clasts, some with poorly developed imbrication. 25
Interpretation
The offshore marine lithofacies association is interpreted as deposits of a
shallow marine-shelf setting (e.g, Walker and Plint, 1992; Orton and Reading, 1993;
Johnson and Baldwin, 1996). The laminated mudstone is interpreted to have been
deposited in a relatively quiet environment below fair-weather wave base, where fine-
grained sediment was able to settle from suspension. Sand and pebble-sized sediments
were likely introduced into the offshore area during high-energy events by storm
induced currents and/or hyperpycnal flows associated with terrestrial rivers (Duke,
1990; Orton and Reading, 1993; Nemec and Steel, 1984). Sandstone beds with
abundant oscillatory current ripples indicate that wave base was lowered contemporaneously with deposition of the coarser-grained material, which supports a storm-related interpretation (e.g., Johnson and Baldwin, 1996). Following the high- energy events, when low-energy conditions returned, mud-sized particles were deposited from suspension and organisms burrowed into the deposits. The coarse grain- size of some of the sediment suggests a proximity to topographic relief (Orton and
Reading, 1993).
Northern Fluvial Lithofacies Association
Description
The northern fluvial lithofacies association consists of upward-fining sandstone
and conglomerate units interstratified with massive and laminated mudstone. This
association is best exposed on the north side of Langshan Mountain in the Duba area
where it is at least 1300 m thick (Figure 2.5). Sandstone and conglomerate sequences 26 are generally between 6-10 m thick, have lenticular shapes and overlie basal scour surfaces (Figure 2.7A). Laterally along the outcrop, it is not uncommon for these sequences to divide into individual sandstone beds that are separated from one another by thin mudstone intervals (Figure 2.7B). The lower part of a typical sequence consists of a well-organized pebble-conglomerate bed (~2 m thick) with imbricated clasts, crude horizontal stratification and weakly-developed trough cross-stratification (lithofacies
G1, G2). The conglomerate beds occasionally fine-upward into very coarse-grained pebbly sandstone with trough cross-stratification (Figures 2.5, 2.7C). The upper part of the sequence is composed of coarse- to medium-grained trough cross-stratified and horizontally laminated sandstone overlain by medium- to fine-grained sandstone with plane parallel laminations and unidirectional current ripples (lithofacies S1, S2, S4). In some instances, the cross-stratification is planar and has amplitudes of roughly 1 meter
(lithofacies S3). Internally, sandstone sequences locally contain large-scale cross- bedding (amplitudes >2m) consisting of sandstone and siltstone beds that dip orthogonally to the paleocurrent direction (as derived from trough cross-stratification; see below). The uppermost parts of the sandstone and conglomerate sequences are commonly bioturbated or massive and grade into red siltstone and claystone (lithofacies
S7, F2). These mudstone intervals are commonly mottled, contain CaCO3 nodules, root traces, vertical burrows and thin tabular beds of fine-grained sandstone with unidirectional current ripples and climbing ripples (Figure 2.7D). Within the uppermost
100 meters of this lithofacies association, just below the overlying limestones, the 27
proportion of siltstone increases, conglomerate beds are absent, and the sandstone
contains a few bivalve fragments (Figure 2.5).
Interpretation
This lithofacies association exhibits abundance evidence of having been
deposited by gravelly-sandy meandering rivers (e.g., Levey, 1978; Nijman and
Puigdefabregas, 1978; Miall, 1996). The sandstone and conglomerate sequences are
interpreted as fluvial channel deposits and the mudstone with CaCO3 nodules and thin tabular sandstone beds are interpreted as sediments deposited in floodplains adjacent to the channels. Regular, repetitive, upward-fining sequences like those within this lithofacies association are most commonly produced by meandering rivers as are the large-scale cross-beds, which are interpreted as bar accretion sets (Allen, 1964;
Puigdefabregas and Van Vilet, 1978; Smith, D., 1987; Smith, R., 1987; Miall, 1996;
Allen, 2001). The fact that many of these accretion sets dip orthogonally to the paleocurrent direction provides evidence that the channels were migrating laterally. A few of the accretion sets consist of alternating beds of sandstone and siltstone, similar to
Inclined-Heterolithic-Strata (IHS), which have been documented in both modern and ancient lower coastal plain fluvial deposits (Smith, D., 1987; Choi et al., 2004). Crudely stratified pebble-conglomerate was deposited by gravel sheets and low-relief bars along channel thalwegs (Hein and Walker, 1977), and the overlying trough cross-stratified sandstone were deposited by migrating 3-D dunes as the paleochannels filled and migrated laterally (Jackson, 1975). Intervals of massive mudstone that contains CaCO3 nodules are interpreted as paleosols that developed in a relatively well drained 28
floodplain environment (Mack et al., 1993; Kraus, 1999). Thin, tabular, rippled
sandstone within the mudstone intervals are interpreted as crevasse splay deposits.
Northern Limestone Lithofacies Association
Description
Limestone exposed in the northern portion of the Lhasa terrane is considered a
single lithofacies association in order to interpret the depositional environment at a regional scale, although locally the lithofacies tend to vary. The Langshan Formation in the Duba region (Figures 2.1, 2.5) consists primarily of laterally continuous, dark-gray to black orbitolinid wackestone and packstone (lithofacies C1; Figures 2.5, 2.8). Fossils of Mesorbitolina aperta and Mesorbitolina subconcava represent the bulk of the bioclasts, but in some beds there are also abundant gastropods (Tylostoma sp. and
Acteonella sp.) and bivalve and echinoid fragments (R.W. Scott, personal communication). North of Duba, in the Lunpola area (Figure 2.1), are thickly bedded bioclast wackestone and packstone, consisting primarily of rudists, but also containing indeterminate bivalve and cnidaria fragments and the calcareous algae Pycnoporidium sp. and Cayeuxia sp. (lithofacies C2; photo; R. W. Scott, personal communication).
Locally, the rudist limestone is interbedded with packstone and grainstone beds composed of ooids, composite ooids and oncoids (lithofacies C3). Massive pellodial and orbitolinid mudstone beds are the primary lithofacies in the Nam Co area
(lithofacies C4), although this mud-rich lithofacies is present to some extent at all of the locations.
Interpretation 29
Limestone of this lithofacies association is interpreted as having been deposited within a shallow marine epeiric seaway that was dominated by low-energy lagoonal environments and localized patch reefs. These results are similar to the conclusions of
Leeder et al. (1988). The common occurrence of pelloidal and orbitolind mudstone and wackestone indicate much of the region was occupied by muddy, low-energy, lagoonal environments (Leeder et al., 1988; Tucker and Wright, 1990; Jones and Desrochers,
1992). However, localized exposures of rudist-dominated limestone and ooidal packstone and grainstone suggest the former existence of patch reefs where energy levels were relatively higher (Scott, 1979; James and Bourque, 1992; Wright and
Burchette, 1996).
LITHOFACIES ASSOCIATIONS – SOUTHERN AREA
Lagoonal Lithofacies Association
Description
The lagoonal lithofacies association occurs within lower Cretaceous Chumulong
Formation in the southern half of the study area and is composed of brown, organic-rich claystone and siltstone with subordinate amounts of sandstone (Figure 2.9; lithofacies
S4, S5, F1). This lithofacies association is common in many parts of the Chumulong
Formation, including the lower half of the succession and in the strata just below the
Aptian-Albian limestone (Figure 2.9). The mudstone is typically laminated and contains very fine-grained sandstone laminae with oscillatory current ripples; the laminations are often disturbed or destroyed by subhorizontal burrows and vertical to subvertical skolithos burrows. Disarticulated fragments of oysters and unidentifiable bivalves are 30 common within the mudstone intervals. Interbedded with the mudstone are thin- to thickly-bedded, very fine- to fine-grained bioturbated sandstone with oscillatory and unidirectional current ripples, flaser bedding, oyster fragments, and small amounts of fossil wood and fossilized plant debris. Some intervals within this lithofacies association contain 2-4 m thick repetitive packages composed of mudstone in their lower parts and one or more ~0.5 m thick beds of fine-grained, bioturbated sandstone in their upper parts (Figure 2.10).
Interpretation
Sediments of the lagoonal lithofacies association are interpreted as having been deposited in a relatively low-energy lagoonal environment within a clastic marginal marine setting (e.g., Kirschbaum, 1989). Low-diversity fossil assemblages that are dominated by oysters, small mudstone-sandstone packages, flaser bedding, and fossil plant and wood debris are all characteristic of lagoonal deposits (Elliot, 1974; Ward and
Ashley, 1989, Kirschbaum, 1989).
Shoreface Lithofacies Association
Description
The shoreface lithofacies association is characterized by multiple 5-15 m thick upward-coarsening sequences of mudstone and sandstone (Figure 2.9). The lowest part of a typical sequence contains mudstone and thinly bedded very fine-grained sandstone with oscillatory current ripples, occasional plane-parallel laminations, hummocky cross- stratification (HCS), and skolithos burrows (lithofacies S2, S5, S6). Upward within individual sequences, these heterolithic deposits become progressively more sandstone- 31
rich, eventually giving way to white to buff-colored upward-coarsening sandstone units
5-10 meters thick. The lower beds in these sandstone units are very fine- to fine-grained and contain plane-parallel laminations. Overlying these beds are fine- to medium- grained sandstone beds with abundant trough cross-stratification and occasional plane- parallel laminations. Subvertical to subhorizontal burrows are common in the lowermost sandstone, but rare within the upper, trough-cross-stratified beds. The upper surfaces of the sequences are sharp, commonly bioturbated, and overlain by interbedded mudstone, siltstone and sandstone of the overlying upward-coarsening sequence.
Interpretation
Strata of the shoreface lithofacies association were deposited in a wave- dominated shoreface environment that experienced repeated transgressions and regressions. The upward-coarsening sequences are interpreted as shallow marine parasequences that were deposited during shoreline progradation (Van Wagoner et al.,
1988; Van Wagoner et al., 1990; Posamentier and Allen, 1999). Mudstone in the lower part of the parasequence is interpreted to have been deposited in an offshore transition
zone, the overlying mudstone and sandstone with HCS in a lower shoreface
environment and the trough-cross-stratified sandstone within an upper shoreface
environment (Van Wagoner et al., 1990; Battacharya and Walker, 1991; Walker and
Plint, 1992; Johnson and Baldwin, 1996). The sharp, bioturbated upper surface of the
sequences is interpreted as a transgressive surface that records a depositional hiatus as
relative sea-level rose, the shoreline transgressed, and the depositional environment 32 returned to deeper marine conditions (e.g., Van Wagoner et al., 1990; Posamentier and
Allen, 1999).
Southern Fluvial Lithofacies Association
Description
The southern fluvial lithofacies association consists of interbedded siltstone, sandstone and conglomerate (Figures 2.9, 2.11A). Siltstone intervals are brown, massive and often contain thin beds of very fine- and fine-grained sandstone with unidirectional current ripples and burrows. Organic-rich mudstone is present in the lowermost portion of the lower Cretaceous succession where coal layers also have been reported (Yin et al., 1988), although no coal horizons were observed during this study.
Interstratified with the siltstone and mudstone are thick (> 60 m) sequences of fine- to very coarse-grained sandstone and clast-supported pebble conglomerate (Figure 2.11b).
These sandstone-conglomerate sequences overlie basal scour surfaces and are composed of multistory, 2-8 m thick sandstone-conglomerate packages. Individual packages generally contain both conglomerate and sandstone, but packages composed entirely of pebble conglomerate are not uncommon. Typical packages overlie a scour surface, contain lowermost beds of clast-supported, imbricated to crudely-stratified, quartzite-pebble conglomerate, which are overlain by sandstone with trough cross- stratification and plane-parallel laminations (lithofacies G1, S1, S2). The sandstone fines-upward within each package and some contain relatively large (1.5 m) sets of planar, to slightly-trough, cross-stratification (lithofacies S3). Siltstone and very fine- grained sandstone sometimes cap the upper part of individual packages. 33
Interpretation
The sediment, lithofacies and architecture of the southern fluvial lithofacies
association are most similar to those of low-sinuosity, sandy-gravelly braided rivers
(Miall, 1978; Cant and Walker, 1978; Willis, 1993a; Miall, 1996). The imbricated, well- organized pebble-conglomerate beds are interpreted as deposits of gravel sheets and
low-relief bars that migrated along channel thalwegs (e.g., Hein and Walker, 1977).
Trough-cross stratified sandstone was deposited by subaqueous 3-dimensional dunes
that migrated within the paleochannels. The sandstone with relatively larger and more
planar cross-stratification is interpreted as having been deposited migrating 2- dimensional dunes or transverse bars. The plane-parallel laminated sandstone near the top of individual packages are interpreted as having been deposited within either infilled
channels or on bar-tops where the depth of flow is commonly shallow; local siltstone
beds deposited atop the plane-parallel laminated sandstone are interpreted to have been
deposited during the waning stages of flow (e.g., Langford and Bracken, 1987; Bristow,
1993). The mudstone intervals interstratified between the sandstone sequences are
interpreted to have been deposited in interfluvial floodplain environments (e.g., Cant
and Walker, 1978). The massive nature of the mudstone suggests these may be
paleosols but no definitive evidence of this is present.
Southern Limestone Lithofacies Association
Description
In the southern half of the study area, lower Cretaceous clastic strata of the
lagoonal lithofacies association are overlain by bioclast wackestone and packstone of 34 the southern limestone lithofacies association (the Penbo Member of the Takena
Formation; figure). The limestone contain orbitolinids (Mesorbitolina sp.), echinoids
(Macraster sp., Selina sp.), and rare oysters (Ceratosreon (?)) and ammonites
(Kazanskyella sp.) (R.W. Scott, personal communication). A succession of siltstone and local sandstone overlie the wackestone, and are in turn overlain by a series of 3-5 m thick upward-coarsening limestone cycles (Figures 2.9, 2.12). These cycles are composed, from bottom to top, of marly-siltstone, orbitolinid-wackestone
(Mesorbitolina texana (Roemer); R.W. Scott, personal communication) and relatively coarse-grained and fossiliferous orbitolinid-oyster packstone, which also contain a subordinate amount of echinoderm debris and ostracodes. The packstone beds contain numerous fossils, have sharp upper surfaces, and are overlain by marly-siltstone of the overlying cycle. The limestone within the southern portion of the study area eventually grade into siltstone and fine-grained sandstone with marine fauna (Leier, 2005).
Interpretation
Collectively, the southern limestone lithofacies association was deposited in a low energy, carbonate-dominated shallow-water marine environment. The carbonate cycles are interpreted as carbonate parasequences that record repeated shoaling and flooding of a shallow marine environment (e.g., Inden and Moore, 1983; Jones and
Desrochers, 1992). The presence of siltstone and local sandstone within this lithofacies association suggests that although this location was dominated by carbonate sedimentation, it was proximal to limited clastic input, possibly associated with a nearby shoreline. 35
PALEOCURRENT MEASUREMENTS
In order to reconstruct paleocurrent directions, the limbs of trough cross-
stratification sets within fluvial sandstone were measured (following method I of
DeCelles et al., 1983), as were the strike and dip of imbricated pebble and cobble clasts
within fluvial conglomerate beds.
Trough cross-strata of lower Cretaceous fluvial sandstone in the northern half of
the study area record southwest-directed sediment transport (Figure 2.5). Near the
exposed base of the Duba Formation paleocurrents are generally south-southwest
directed, whereas trough cross-stratification orientations near the top of the Duba
Formation indicate a more westerly-directed flow (Figure 2.5). Accretion sets within some of the fluvial sandstone sequences in the Duba area dip toward the northwest, orthogonal to the mean paleoflow direction, and indicate lateral channel migration.
Imbricated clasts within a lower Cretaceous marine conglomerate exposed in the
Lunpola area record south-southeastward sediment transport.
Paleocurrent data were more difficult to obtain in the southern half of the study area owing to poorer exposures and the greater amount of tectonic deformation. Trough cross-stratification and imbricated clasts indicate southward-directed flow within three fluvial sequences, but northward-directed transport is indicated by trough cross- stratification within two fluvial units (Figure 2.9). Paleocurrent data obtained from a fluvial sandstone interstratified with Aptian-Albian limestone indicate northwestward- directed sediment transport, which is similar to the paleocurrent data derived from the 36 upper Cretaceous fluvial sandstones that overlie the lower Cretaceous strata (Leier,
2005).
PETROGRAPHY
Petrographic thin sections were made from 26 medium- and coarse-grained sandstone samples collected while measuring stratigraphic sections. The thin sections were stained for potassium and calcium feldspar and point-counted (450 counts per slide) using a modified Gazzi-Dickinson method (Ingersoll et al., 1984); the modification involves the identification of monocrystalline quartz grains that are part of sedimentary lithic and plutonic fragments (Table 2.2). In addition, 7 clast counts (100 clasts/count) were performed in the field. The petrographic counting parameters are shown in Table 2.2, and raw data are available in Table 2. DR1.
Northern Study Area
Lower Cretaceous sandstones consist of calcite-cemented litharenites and feldspathic litharenites with Qm:F:L modal compositions of 42:16:42 and Qt:F:L of
57:16:27 (Figures 2.13, 2.14A, 2.14B). Subangular to subrounded monocrystalline quartz grains are the dominant form of quartz. Feldspar grains compose 10-20% of the modal composition, the bulk of which is plagioclase. Potassium feldspar is typically minor or absent, although a few samples have 2-5% potassium feldspar relative to the total number of grains. In a few samples, up to 30% of the quartz and feldspar occur not as individual grains, but as constituents of felsic plutonic fragments (these are still counted as quartz or feldspar, following the Gazzi-Dickinson method). Lithic fragments are commonly the most abundant grain-type within the lower Cretaceous sandstones. 37
Particularly abundant are andesitic-dacitic volcanic grains containing microlaths of
plagioclase feldspar (Figures 2.14A, B). Sedimentary and metamorphic lithic grains such as mudstone/shale, phyllite, schist and limestone grains constitute an average of
20% of the modal composition. Some of the sandstone and siltstone lithic grains have minerals and characteristics (e.g., abundant plagioclase and chloritic matrix) that are similar to that of Jurassic sandstone and siltstone sedimentary strata exposed 50 km to the north in the Lunpola area. Data from clast counts are consistent with the point- counting data in that the majority of the clasts consist of meta-sedimentary rocks, felsic volcanic rocks and granite. Numerous accessory minerals are present within the sandstones, particularly zircon and tourmaline, and also include pyroxene (omphacite), and minor amounts of chlorite and serpentine.
Southern Study Area
Unlike the feldspar-rich sandstone exposed in the northern portion of the Lhasa terrane, lower Cretaceous sandstone within the southern portion of the terrane contain no feldspar and typically have compositions of quartz arenites and sublitharenites
(Qm:F:L 73:0:27 and Qt:F:L 92:0:8; Figure 2.13). Of the total quartz within the sandstones, the largest component is monocrysyalline quartz, although in some samples polycrystalline quartz constitutes up to 15% of the total quartz (Table 2. DR1).
Consistent with the petrographic results, the clasts within pebble-conglomerate beds are almost entirely quartzite and vein quartz. Contrasting with lower Cretaceous sandstone in the northern portion of the Lhasa terrane, no feldspar grains are present within the lower Cretaceous sandstones in the south, and lithic fragments are rare (Figures 2.14C, 38
2.14D). Those lithic fragments that are present are generally fragments of phyllite and
shale. Although sandstones just below the Aptian-Albian limestone beds are similarly
quartzose, they contain a small amount of volcanic material and a slightly higher
percentage of sedimentary and metamorphic lithic fragments relative to sandstone lower
in the lower Cretaceous succession. Rare zircon and muscovite are present within a few
sandstone samples but accessory minerals are largely absent.
DETRITAL ZIRCON GEOCHRONOLOGY
Uranium-lead age data of detrital zircons can provide valuable insight into the provenance and depositional history of sedimentary successions. We collected two
samples of lower Cretaceous sandstone, one from an exposure in the Penbo area and the
other from the Duba area, in order to determine maximum depositional ages, constrain
provenance sources and better characterize the genetic relationship between the lower
Cretaceous strata in northern and southern Lhasa. Raw data are presented in Table 2.
DR2.
Methods
The two ~15 kg samples collected from the Penbo (sample PNB) and Duba
areas (sample DBA) were processed for detrital zircon analysis using standard
procedures described in Gehrels (2000). Uranium-lead ages of detrital zircons were
obtained using a laser ablation mutlicollector inductively coupled plasma mass
spectrometer (LA-MC-ICPMS) at The University of Arizona following the methods of
Dickinson and Gehrels (2003). The zircons were ablated with a New Wave DUV193
Excimer laser operating at a wavelength of 193 nm and using spot diameters of 35 or 50 39
µm, depending on grain size; depth of the ablation pits is ~20 µm. The ablated material
was transported via argon gas to the plasma source of a Micromass Isoprobe. Analyses were made in static mode using Faraday collectors for 238U, 232Th, 208Pb, 207Pb, 206Pb, and an ion-counting channel for 204Pb. The analyses consisted of one 20-second integration on peaks without the laser firing, checking for background levels, and twenty 1-second integrations with the laser firing on the zircon grain. Thirty-second delays occurred between analyses to purge the previous sample and allow peak intensities to return to background levels. Common lead corrections were made using
the measured 204Pb of the sample and assuming initial Pb compositions from Stacey and
Kramers (1975). Lead-204 measurements were corrected for contributions from 204Hg by subtracting background values measured during the 20-second integration when the laser was not firing. Every fifth measurement, an analysis of a fragment of a zircon crystal of known age (564 ± 4 Ma, 2-sigma error; G.E. Gehrels unpublished data) was performed to correct for inter-element and Pb isotope fractionation. Fractionation of Pb isotopes is generally ~15%, whereas inter-element fractionation is ~1-2%. The raw data are available in Table 2. DR2. The two samples were run during the same session and on the same mount. The ages presented are 206Pb*/238U ages for grains less than ~800
Ma and 207Pb*/206Pb* ages for grains whose ages are greater than 800 Ma. Those analyses with greater than 10% uncertainty or more than 20% discordance or 5% reverse discordance are omitted from further consideration. The data from each sample are displayed on concordia diagrams and age-probability plots using the programs of
Ludwig (2001). Age-probability plots depict each age and its uncertainty as a normal 40
distribution, summing all ages from the analyzed zircons of a sample and displaying it as one curve.
Results
Sample DBA, from the lower Cretaceous Duba Formation yielded 94 ages of sufficient concordance and precision to be used for provenance analysis (Figure 2.15).
Many of the zircons are euhedral and have U/Th values indicative of a plutonic origin, which is consistent with the composition of the sandstones and conglomerates (see above). The youngest ages cluster at 125 ±1.4 Ma, which provides a maximum depositional age of the sedimentary succession. Grains with ages between 120-150 Ma are the most numerous (peak at 141 Ma) and represent over 65% of the total population
(Figure 2.15). The remaining grains are in clusters between 250-340 Ma, 610-670 Ma,
700-900 Ma (peak at 850 Ma), 1000-1200 Ma (peak at 1040 Ma) and a few of Early
Proterozoic age.
Sample PNB from the lower Cretaceous Chumulong Formation from the southern half of the study area yielded 93 ages that pass the requisite standards for provenance analysis (Figure 2.15). The zircons are typically subhedral to rounded.
Unlike the DBA sample, the PNB sample is composed primarily of zircons of Paleozoic age and older (66 of the 93 grains are >250 Ma). Twenty-eight zircons are Mesozoic in age with the youngest prominent cluster at 143 ±1.3 Ma. Zircons in this sample have age-groupings of 140-160 Ma (the most numerous with 25 zircons; peak at 140 Ma),
500-580 Ma (the second most numerous with 16 grains; peak at 530 Ma), 620-710 Ma, 41
850-950 Ma (peak at 900 Ma), 1000-1300 Ma (age-probability peak at 1130 Ma), and
near 2500 Ma (Figure 2.15).
Implications of the U-Pb age data of the detrital zircons for the relationship
between conglomerates in the northern and southern portion of the Lhasa terrane are
discussed in the paleogeography section below.
REGIONAL INTERPRETATIONS
Incorporating all of the data presented thus far, we attempt to reconstruct the
paleogeography of the Lhasa terrane during Early Cretaceous time. The evolution of the
region is divided into three stages: 1) the period beginning with the collision between
the Lhasa and Qiangtang terranes (latest Jurassic-earliest Cretaceous) up to mid
Neocomian time (~145 to ~130 Ma); 2) the period from mid Neocomian to early Aptian
time (~130 to ~120 Ma); and 3) the period between Aptian and earliest Cenomanian
time (~120 to ~100 Ma). These three stages are based on prominent changes in
lithologies and do not necessarily represent changes in tectonic conditions. The factors
controlling deposition during these time periods are discussed in the following section.
Collision to Mid Neocomian
Northern Study Area
The Early Cretaceous paleogeography in the northern portion of the Lhasa
terrane is characterized by coarse-grained shallow marine and coastal environments south of an elevated region (Figure 2.16). Interbedded mudstone and sandstone with bivalve fragments and oscillatory current ripples suggest shallow marine environments influenced by wave-processes. Southward-directed paleocurrents and lithic-rich 42
sandstones with chlorite, pyroxenes, and siltstone-fragments similar to Jurassic
stratigraphic units exposed near the Bangong suture, indicate that the Bangong suture and southern Qiangtang terrane served as the primary sediment sources. Volcanic grains are abundant within the sandstone indicating that in addition to areas around the
Bangong suture, volcanism within the region contributed sediment to the strata.
Conglomerate beds suggest the region was proximal to uplifted areas and topographic relief. The paleoshoreline trend is inferred to have been roughly parallel to the Bangong suture (east-west trending) based on the distribution of facies, the lack of lower
Cretaceous marine rocks within the Qiangtang terrane (Liu, 1988; Kapp, P. et al., 2005) and assuming that the shoreline trend was roughly perpendicular to the southward paleocurrent direction.
Southern Study Area
During the same period in the southern portion of the Lhasa terrane, interbedded
lagoonal, shoreface, and fluvial deposits suggest marginal marine and fluvial
environments prevailed (Figure 2.16). The quartzose composition of the sandstones and
the textural maturity of the individual grains suggest a stable tectonic setting. However,
also within the succession are beds of pebble conglomerate and sandstone containing
sedimentary lithic fragments, which is more consistent with deposition in an area close
to topographic relief. The most plausible explanation for this apparent contradiction is
that much of the lower Cretaceous sediments in the southern portion of the Lhasa
terrane are recycled and were derived from proximal uplifts that exposed sedimentary
cover rocks (Figure 2.16). Carboniferous strata of the Lhasa terrane likely supplied 43 much of the sediment to the lower Cretaceous sandstones based on: 1) strata of
Carboniferous age are thick and widespread throughout southern Tibet (Yin et al.,
1988); 2) Carboniferous strata are composed of quartzose sandstones and quartzites, making them a logical source for the quartz grains and quartzite clasts within lower
Cretaceous sandstone and conglomerate; 3) Carboniferous sandstones are interbedded with phyllite, and phyllite fragments are the largest component of lithic grains within the lower Cretaceous strata; 4) the relative probability spectra of detrital zircons older than 400 Ma from the lower Cretaceous and Carboniferous rocks display many similarities (Leier et al., 2004); and 5) in some localities within the Lhasa terrane
Cretaceous rocks unconformably lie atop Carboniferous strata (Liu, 1988), indicating
Carboniferous strata were exposed at the surface during this time.
Mid Neocomian to Early Aptian (~130-120 Ma)
Northern Study Area
Southwest-flowing, gravelly, meandering rivers occupied northern Lhasa during the time period between mid Neocomian and Aptian time (Figure 2.16). Sandstone composition and conglomerate clasts indicate the sediment was derived primarily from felsic volcanic and igneous rocks as well as sedimentary cover strata exposed near the
Bangong suture. The numerous pebble and cobble conglomerate beds indicate topographic relief existed in the immediate or surrounding areas. The fluvial sediments deposited during this time are stratigraphically between two marine sedimentary units
(Figure 2.3), and within some fluvial sequences are IHS, suggesting that although these 44
rocks record terrestrial environments, marine conditions were never far from the area
(e.g., Smith, D., 1987).
Southern Study Area
Sedimentary strata in the southern portion of the Lhasa terrane indicate the area
continued to be occupied by marginal marine, lagoonal and lower coastal plain environments during the mid Neocomian to the Aptian (Figure 2.16). Unlike the underlying sedimentary units, conglomerate beds are absent within the strata deposited
during the mid Neocomian to the Aptian, suggesting any local relief that may have once
existed was by this time subdued.
Relationship Between Lower Cretaceous Conglomerate Units
The lithologic evidence suggests that the lower Cretaceous conglomerates of the
Chumulong Formation in the south and those of the Duba Formation in the north are not entirely correlative, although the two units may have shared some sediment source areas. The youngest clusters of detrital zircon ages from samples of the conglomerate units in the north and in the south differ by roughly 16 Ma (Figure 2.15). Although detrital zircon ages can only definitively be used to constrain maximum depositional ages, the abundance of Early Cretaceous volcanism that post-dates 140 Ma, including that in locations close to the Penbo area like the Nyantenaglha (Kapp, J. et al., 2005) suggests that the conglomerates of the Chumulong Formation predate the volcanism and are therefore older than those of the Duba Formation to the north. It is also difficult, though not impossible, to reconcile the disparate compositions of the quartzose
Chumulong Formation and the volcanic-rich Duba Formation. Furthermore, if Late 45
Cretaceous and Cenozoic north-south shortening of 50% is removed (Berg et al., 1983;
Pan, 1993), the distance between exposures of the pebble conglomerate units of the
Chumulong Formation and those of the Duba Formation is roughly 400 km, which is
much greater than fluvial gravel fronts are typically thought to extend (e.g., Robinson and Slingerland, 1998). However, both the Chumulong and the Duba Formation contain significant detrital zircon populations with ages of roughly 140 Ma (Figure 2.15). This suggests some of the sediment within the two stratigraphic units may have been derived from the same source, although locating the exact or possible sources requires further investigation.
Aptian-Albian (~120-100 Ma)
Northern Study Area
Biostratigraphic evidence indicates that deposition of limestone within the
northern portion of the Lhasa terrane commenced in the Barremian, before carbonate
deposition began in the south (Yin et al., 1988; Zhang, 2000; Leier, 2005). The fossil
assemblages and geographic distribution of carbonate facies indicate the northern Lhasa
terrane was covered by a shallow marine sea with rudist patch reefs and muddy inter-
reef areas (Figure 2.16; e.g., Scott, 1979; Leeder et al., 1988). These environments
persisted in this region through the Albian and into the earliest Cenomanian (Yin et al.,
1988). The limestone deposited in northern Lhasa is thicker than coeval rocks exposed
in the southern portion of the Lhasa terrane (~1100 m in the north versus 300 m in the
south), indicating greater subsidence in northern Lhasa.
Southern Study Area 46
Clastic marginal marine environments were present in the southern portion of
the Lhasa terrane until Aptian time, when carbonate deposition became dominant
(Figure 2.16). The limestone was deposited primarily in muddy, low-energy shallow
marine environments that occasionally received influxes of fine-grained clastic material.
Paleocurrent directions derived from a fluvial sandstone interstratified within the
carbonate succession indicate the clastic material was derived from the south. An east-
west trending shoreline is inferred to have existed to the south of the Penbo area, mimicking the southern margin of the Lhasa terrane and perpendicular to the sediment
transport direction. By the end of the Albian, the marine carbonate environment gave
way to a clastic marginal marine shoreline whose sediment was derived from the
Gangdese magmatic arc to the south. Eventually the marine accommodation was infilled and north-northwest flowing streams traversed the region (Leier, 2005).
DISCUSSION
Lower Cretaceous sedimentary rocks within the eastern Lhasa terrane provide
information and constraints that can elucidate the tectonic processes active in southern
Tibet during the Early Cretaceous. In addition to the data presented thus far, there are
several other constraints that have a direct or indirect bearing on reconstructions of the
regional tectonic setting. The oldest forearc strata within the Xigaze Forearc basin were
deposited during the Aptian-Albian (Dürr, 1996); meaning that by this time, northward
subduction of Neothethyan oceanic lithosphere along the southern margin of the Lhasa
terrane had commenced. By roughly 130-125 Ma, calc-alkaline volcanism was
occurring within the Lhasa terrane, particularly in its northern half (Figure 2.17; e.g., 47
Volkmer et al., 2005). The exact thickness of the Langshan Formation is difficult to
determine because of the variable published thicknesses. We use an approximate
thickness of ~1100 m for the limestone succession following Leeder et al. (1988).
Widespread Aptian-Albian limestone deposits are present within strata throughout
Europe and Asia (Vilas, et al., 1995; Zhang, 2000; Pittet et al., 2002). Cross-cutting
relationships between thrust faults and igneous intrusions indicate thrust faults were
active within the northern portion of the Lhasa terrane during the Early Cretaceous
(Murphy et al., 1997).
The Lhasa terrane during the Early Cretaceous is best characterized as having been a peripheral foreland basin; however, it is somewhat unique in that sedimentation
within this basin was strongly influenced by concurrent arc volcanism caused by
northward subduction of Neotethyan crust. The pre-Aptian stage of basin development
provides ample evidence of having been a peripheral foreland basin that formed in response to the Lhasa-Qiangtang collision (e.g., Leeder, et al., 1988; Murphy et al.,
1997; Zhang, 2004). The paleotectonic setting of the Lhasa terrane during Aptian-
Albian time is more ambiguous, but is best categorized as an inundated peripheral
foreland basin. Sedimentation during this time is controlled by high eustatic sea-levels
and possibly some dynamic subsidence; in the northern portion of the Lhasa terrane
subsidence associated with crustal loading created additional accommodation and
allowed for thicker accumulations of marine limestone. Because of the ambiguity
associated with the limestone record, we evaluate several alternative possibilities for the 48
paleotectonic setting of the region and list the supporting and contradictory evidence for
each in Table 2.3.
Initial Stage of Basin Development
The collision between the Lhasa and Qiangtang terranes during the latest
Jurassic-earliest Cretaceous resulted in deformation and uplift along the Bangong suture
and in the southern Qiangtang terrane (Murphy et al., 1997; Kapp et al., 2005). The
crustal load associated with this deformation resulted in the formation of a peripheral
foreland basin in the northern half of the Lhasa terrane (Figure 2.16; Leeder, et al.,
1988; Zhang et al., 2004; this study). This collisional orogen was the dominant control
on regional sedimentation in the Lhasa terrane from the latest Jurassic to the end of the
Barremian. This interpretation is supported by the overall upward-coarsening trend in
the upper Jurassic-Early Cretaceous strata in the northern portion of the Lhasa terrane.
The lowermost strata consist of deep-marine deposits (Lake-Area Flysch; Figure 2.3;
Yin et al., 1988), and are overlain by fine- and coarse-grained shallow marine deposits
(lower Duba Formation; Figure; Yin et al., 1988; this study), which are in turn overlain
by coarse-grained fluvial deposits (upper Duba Formation; Leeder et al, 1988; this
study). This sequence is typical of peripheral foreland basins. Paleocurrent data and
sandstone composition indicate the sediment deposited in the northern Lhasa terrane
during this time was derived from northern source areas (Zhang et al., 2004; this study),
and is consistent with the absence of lower Cretaceous strata in the southern portion of the Qiangtang terrane (Liu et al., 1988). Stratigraphic relationships within the Qiangtang terrane indicate areas north of the Lhasa terrane were deformed, above sea-level, and 49 eroded by the middle of the Cretaceous (Kapp et al., 2005; Figure 2.16). By the end of the Barremian, volcanism, which was by then widespread in the northern half of the
Lhasa terrane (Figure 2.17), began to strongly influence the composition of sediment deposited in the area.
The paleogeographic setting of the southern margin of the Lhasa terrane during this time period is more difficult to reconstruct. This area may have been similar to a passive margin during this period (Figure 2.16), although it is also possible subduction of Neotethyan lithosphere had commenced along the southern margin. The paucity of volcanic grains in the Chumulong Formation suggests that if subduction had begun by this time, the resulting volcanic activity must have been either north (~shallow subduction) or south (~steep subduction) of the area (Figure 2.16). The lower
Cretaceous strata in the southern portion of the Lhasa terrane may have been derived in part from northern sources associated with topography created during the Lhasa-
Qiangtang collision, but compositional evidence indicates that local sources were also important. The pebble-conglomerate within the Chumulong Formation suggests that topographic relief existed in the region, and was probably exposing pre-Cretaceous
(particularly Carboniferous) strata to weathering and erosion. With the available data, it is impossible to determine what was responsible for the local uplifts, although it is likely the uplifts involved the inversion of normal faults along the southern margin of the Lhasa terrane that originally formed as Lhasa rifted from Gondwanaland (Dewey et al., 1988; Leeder et al., 1988). The forces necessary to invert these structures could have 50
been associated with compression along the southern margin of the terrane or the
Lhasa-Qiangtang collision (Figure 2.16).
Second Stage of Basin Development
The primary issues to be addressed with the Aptian-Albian history of the Lhasa
terrane are: 1) how and why did the region become submerged beneath sea-level? and
2) what tectonic or geodynamic mechanism could have accomplished this without
leaving a record of faulting or producing a major influx of clastic material into the
region? No single existing model can adequately explain all aspects of the Aptian-
Albian strata of the Lhasa terrane (Table 2.3). Subsidence associated with crustal
loading near the Bangong suture should have produced local subsidence, but marine
deposition occurred 400-600 km away from the site of crustal loading, which is too
great a distance to be attributed to the flexural response of the crust (e.g., Allen and
Allen, 1990). Tectonic activity related to extensional forces is an appealing mechanism
to explain deposition during this time but this postulation also has notable problems
(Table 2.3). To accommodate the thickness of Aptian-Albian limestone, the area must
have had 25-50% extension (Beta factor=1.3 to 2, depending on reported thickness of
Langshan Formation; Figure 2.18; e.g., McKenzie, 1978), yet no Early Cretaceous
normal faults have been found in the area. Not only have extensional faults not been
documented within the Lhasa terrane, but structural evidence indicates the area was
experiencing compression during this time period (Murphy et al., 1997). The
geographic distribution of volcanic and plutonic rocks with Early Cretaceous
crystallization ages suggest the subduction angle of the Neotethyan slab was relatively 51
low during the Aptian-Albian (Figure 2.17; Coulon et al., 1986) and was not steepening
as implied by extensional models (e.g., Zhang et al., 2004). It is possible that northward
subduction of the Neotethyan slab may have induced dynamic effects within the
overriding Lhasa terrane (e.g., Gurnis, 1992; Burgess and Moresi, 1999), although no
evidence definitely supports this hypothesis. By the Late Cretaceous, a magmatic arc
and a fold-thrust belt had developed along the southern margin of the Lhasa terrane
(Leier, 2005), therefore some of the subsidence in the southern portion of the Lhasa
terrane may have been caused by mid Cretaceous crustal loading in this area. As will be
discussed below, we postulate that dynamic subsidence and crustal loading in southern
Lhasa were secondary in controlling deposition.
The most plausible explanation for the depositional history during Aptian-
Albian time involves the interaction of eustatic sea-level high and continued, but
perhaps diminished, subsidence caused by crustal loading along the Bangong suture
(Figure 2.16). Given the tectonic setting, it is not unreasonable to suspect that dynamic
effects associated with subduction of the Neotethyan slab may have also played a role in
creating accommodation. Aptian-Albian limestone strata do not occur in the Lhasa
terrane exclusively, but are also present within many additional areas of Eurasia such as the Tarim Basin, northern Oman, and Spain (Vilas, et al., 1995; Zhang, 2000; Pittet et al., 2002). Moreover, evidence of a “superplume,” a sharp increase in ocean-crust production rates, widespread flooding of continental crust, and oceanic anoxic events, all suggest global-scale processes were influencing sea-level and sedimentation during the middle Cretaceous (Larson, 1991). Thus, it is highly probable that eustatic sea-level 52 changes at least assisted in allowing the Aptian-Albian incursion into the Lhasa terrane.
A sea-level rise can also help to explain the creation of accommodation coeval with the paucity of Early Cretaceous faulting and the extensive area that was submerged.
Assuming Airy isostacy, accommodation for the ~200 m thick sequence of carbonates and mixed clastic-carbonate deposits in the southern portion of the Lhasa terrane could have been created by an increase in sea-level of ~75 m (e.g., Allen and Allen, 1990;
Angevine et al., 1990). As mentioned earlier, it is possible that dynamic topography associated with the down-going Neotethyan slab may have also caused additional subsidence (e.g., Davila et al., 2005; Figure 2.16).
The thickness of the limestone in the northern portion of the Lhasa terrane cannot be explained solely by a eustatic rise in sea-level and thus requires an additional mechanism to create accommodation (Figure 2.18). We postulate that additional accommodation was created for Aptian-Albian limestone deposition within the northern portion of the Lhasa terrane by peripheral foreland basin subsidence (Figure 2.16).
Thrust faults were active within the northern Lhasa terrane during the Early Cretaceous, including the Aptian-Albian (Murphy et al., 1997), suggesting the presence of a crustal load. Although carbonate deposits are generally not associated with foreland basin settings, they do occur (e.g., Dorobeck, 1995). An analogue to our interpretation is the
Oligocene and Miocene carbonate strata that were deposited within a foreland basin south of Papua New Guinea. In the Papua New Guinea example, limestone units 1-2 km thick were deposited in the foredeep of a foreland basin (Pigram et al., 1989). Similar to our interpretations of the Langshan Formation, the Oligocene-Miocene limestone 53
deposits thicken toward the foredeep and contain facies indicative of shallow water
environments and patch reefs (Pigram et al., 1989). In such a setting, clastic material
produced by the overriding plate is trapped near the thrust front, allowing carbonate
deposition within the foredeep.
CONCLUSIONS
Lower Cretaceous sedimentary strata of the eastern Lhasa terrane were
deposited in a variety of environments, the majority of which are associated with
marginal marine and coastal plain settings. In the southern portion of the Lhasa terrane,
lower Cretaceous strata consist primarily of shoreface and lagoonal deposits with lesser
amounts of braided-stream deposits, some of which contain pebble-sized clasts. Coeval deposition in northern Lhasa occurred in shallow marine environments, which were later replaced by coarse-grained, southwestward-flowing meandering rivers. All of the lower Cretaceous clastic strata are overlain by a widespread Aptian-Albian limestone that was deposited in a shallow carbonate sea that contained rudist patch reefs and muddy inter-reef zones.
Sandstone compositions, U-Pb ages of detrital zircons, and facies suggest the
lower Cretaceous clastic strata exposed in the northern and southern portion of the
Lhasa terrane have only little genetic relationship. Quartzose sandstones and
conglomerates in the southern portion of the Lhasa terrane were derived from proximal sources as Paleozoic cover strata were exposed and eroded. Detrital zircon U-Pb age data from sandstones within the southern portion of the Lhasa terrane contain multiple age populations, including many with Paleozoic ages. The youngest cluster of ages from 54
these sandstones is roughly 140 Ma. In contrast, the lithic-rich lower Cretaceous strata
in northern Lhasa consist of material derived from the southern Qiangtang terrane, the
Bangong suture, and from local volcanic material. Detrital zircon U-Pb ages of
sandstone within the northern Lhasa terrane are almost entirely between 120 and 140
Ma and indicate an igneous source.
During the Early Cretaceous the Lhasa terrane was located south of a continent-
continent suture zone. A peripheral foreland basin formed on the northern half of the
Lhasa terrane in response to the collision between the Lhasa terrane and the southern
margin of the Qiangtang terrane during the latest Jurassic-earliest Cretaceous. This
peripheral foreland basin was the dominant basin-type throughout most of the Early
Cretaceous. However, by the Aptian, the entire Lhasa terrane was submerged beneath a shallow seaway which most likely formed in response to a rise of eustatic sea-level.
Continued peripheral foreland basin subsidence within the northern portion of the Lhasa
terrane allowed for thicker limestone units to be deposited relative to those in the south.
The lower Cretaceous strata of the Lhasa terrane of southern Tibet clearly indicate
this region had a complicated tectonic history. Further investigation of these rocks
should provide a more detailed history of the area and hold the potential to elucidate how much influence pre-collisional events have had on the development of the Tibetan
plateau. 55
Figure 2.1: Study Area. Location and geology of the study area and location of measured sections. Map modified from Liu (1988). 56
Figure 2.2: Hypothesized Basins. Proposed basin and tectonic models for the lower Cretaceous sedimentary strata of the Lhasa terrane, southern Tibet. (A) An extensional back-arc basin (e.g., Zhang et al., 2004), and (B) A peripheral foreland basin (e.g., Leeder et al., 1988). 57
Figure 2.3: Region Stratigraphy. Lithostratigraphy and chronostratigraphy of upper Jurassic through upper Cretaceous sedimentary strata in the Lhasa terrane of southern Tibet. 58
Figure 2.4: Stratigraphic Symbols. Key for symbols used in the stratigraphic sections. 59
Figure 2.5: Duba Formation Measured Section. Measured sections of the lower Cretaceous Duba and Langshan Formations in the northern part of the Lhasa terrane. Scale is in meters. 60
61
Figure 2.6: Marine Strata. Photos of marine strata of the Duba Formation. (A) Thinly bedded very fine-grained sandstone with oscillatory ripples. Distance between black markers on Jake staff is 0.5 m. (B) Clast-supported pebble conglomerate overlying a scour surface that has partially eroded underlying sandstone bed. Beds are dipping steeply to the right. Hammer for scale. 62
Figure 2.7: Duba Formation. Photos of fluvial rocks of the Duba Formation. (A) Lenticular fluvial channel surrounded by floodplain mudstones. (B) Fluvial architecture of the Duba Formation. Note that many of the sandstone beds are laterally discontinuous. The middle Cretaceous Langshan Formation (limestone) forms the ridge in the background. (C) Typical lithologies within sandstone sequences; conglomerate at the bottom of the unit is overlain by cross-stratified sandstone. Beds are dipping to the left. The Jake staff within the photo is approximately 1 meter. (D) Floodplain mudstones within the Duba Formation. Note the burrows and root traces near the top of the photo, beneath the sandstone bed. 63
Figure 2.8: Langshan Formation. Photos of the Langshan Formation. (A) Orbitolinid packstone, Brunton for scale. (B) Thin section of the Langshan Formation limestone with an orbitolinid test. White line is 0.5 mm. 64
Figure 2.9: Lower Cretaceous in the Penbo area. Measured sections of lower Cretaceous strata exposed in southern portion of the Lhasa terrane near the Penbo area. Scale is in meters. 65
66
Figure 2.10: Lagoonal strata. Photograph of a typical example of the lagoonal lithofacies association. The lithofacies association is characterized by heterolithic strata, oyster fossils, and thorough bioturbation. 67
Figure 2.11: Chumulong Formation. Photos of the Chumulong Formation near the Penbo area. (A) Overall succession with the large cliffs near the top of the outcrop consisting of fluvial sandstone. (B) Clast-supported quartzite pebble conglomerate of the Chumulong Formation. Clasts are imbricated, beds are dipping to the left. Distance between the black marks on the Jake staff is 0.5 meter. 68
Figure 2.12: Penbo Member. Photos of the Penbo Member of the Takena Formation within the Penbo area. (A) Carbonate parasequences. Beds are dipping to the left. Cycles begin with marly siltstone, and coarsen-upward into orbitolinid packstones. Jake staff is 1.5 m. (B) Thin section of the Penbo Member limestone with orbitolinid tests. White line is 0.5 mm.
69
Figure 2.13: Sandstone composition. QFL diagrams (Quartz, Feldspar, Lithic grains) of lower Cretaceous sandstone in the Lhasa terrane. Circles represent lower Cretaceous strata in the southern portion of the Lhasa terrane, triangles represent sandstone from the northern portion. See table 2.2 for parameters. 70
Figure 2.14: Petrography. Photos of petrographic thin section taken in cross-polarized light. All images are at the same scale and the white line is 0.5 mm (A) and (B) Lithic- rich sandstone of the Duba Formation. (C) and (D) quartzose sandstone of the Chumulong Formation exposed within the Penbo area. 71
Figure 2.15: Detrital zircon data. Detrital zircon data from sandstone samples collected from the Duba Formation (DBA) and the Chumulong Formation in the Penbo area (PNB). (A) Concordia diagram. (B) Relative probability plots of all of the detrital zircons, and only those of mesozoic age (inset). Histogram bins in the probability plots of all of the zircons are 50 Ma; 5 Ma for the plots of Mesozoic zircons. See text for discussion. 72
Figure 2.16: Paleogeography and Tectonic Setting. Diagrams illustrating the paleogeography of the Lhasa terrane during different parts of the Early Cretaceous and the tectonic setting of the area as inferred from the sedimentary and stratigraphic data. Specific features are constrained by either this study or from previously published material; numbers refer to source of the data.
73
74
Figure 2.17: Volcanism. Diagram displaying the locations of plutonic and igneous rocks that have been data as Early Cretaceous. Most are confined to the northern portion of the Lhasa terrane, suggesting northward subduction of the Neotethyan slab was occurring at a relatively shallow angle. Based on unpublished compilation of Paul Kapp. 75
Figure 2.18: Subsidence. Subsidence history derived from lower Cretaceous strata in the northern portion of the Lhasa terrane. Subsidence caused by crustal loading that resulted from the Lhasa-Qiangtang collision is interpreted to have been the primary control on deposition until the end of the Barremian. During Aptian and Albian time, tectonic subsidence is reduced and sea-level plays an important role in controlling deposition.
76
TABLE 2.1: Lithofacies Description Interpretation G1.Imbricated conglomerate Clast-supported pebble conglomerate; rare cobbles; imbricated clasts; Deposited by traction transport in sandy matrix; crude horizontal stratification; erosional base. high competency water flows G2.Trough cross-stratified conglomerate Clast-supported pebble conglomerate; weakly developed trough cross- Deposited by migrating, sub- stratification (~1m thick); erosional base; internal scour surfaces. aqueous, three-dimensional gravel dunes. S1.Cross-stratified sandstone Fine- to very coarse-grained trough cross-stratified sandstone (sets Deposited by migrating, sub- typically <1m); occasionally with pebbles and granules. aqueous, three-dimensional sand dunes S2.Plane parallel laminated sandstone Horizontal beds of fine- to coarse-grained, well-sorted sandstone; plane Sub-aqueous deposits associated parallel laminations with upper-flow regime conditions S3.Large, planar cross-stratified sandstone Fine- to medium-grained; large (>1 m thick) relatively low-angle (~20 Deposited by large, migrating, sub- degrees) planar stratification; bases are slightly erosional aqueous, two- to three-dimensional transverse dunes or bars S4.Sandstone with current ripples Very fine- to medium-grained; asymmetric ripples and ripple cross- Deposited by migrating subaqueous lamination; beds < 5 cm thick; occasional sub- and super-critical ripples under a unidirectional current; climbing ripples climbing ripples deposited under high sedimentation rates S5.Sandstone with symmetric ripples Very fine- to fine-grained; symmetrical ripple laminae; occasionally Deposited by ripples formed in interbedded with muddy laminae oscillating currents; muddy laminae suggest alternating energy levels S6.HCS sandstone Very fine- to fine-grained; hummocky-cross-stratification; plane parallel Deposited under combined laminae unidirectional and oscillatory currents; typically lower shoreface S7.Massive sandstone Very fine- to coarse-grained; structureless; occasional faint laminae and Deposited by either sandy gravity burrows flows or bioturbation has destroyed sedimentary structures F1.Laminated dark mudstone Laminated black and brown claystone and siltstone with sandstone Deposited in low-energy sub- laminae and occasional bivalve fossil aqueous conditions F2.Massive red mudstone No sedimentary structures; commonly red to purple and mottled; CaCO3 Paleosols formed in a relatively well- nodules; occasional root traces and vertical burrows drained environment C1.Orbitolina packstone and wackestone Black-gray, laterally continuous beds with orbitolinid tests; rare bivalve Deposited in a low-energy shallow and echinoid fragments marine environment C2.Rudist wackestones and packstones Wackestone, packstone and occasional boundstone composed Shallow marine deposits, possibly a primarily of fossilized rudist shells reef environment C3.Coated-grain packstone and grainstone Ooids, oncoids and aggragate grains; rare indeterminate bioclast Shallow-water marine deposits; fragments moderate energy C4.Pelloidal mudstone Massive; micritic; rare laminations Deposited in a low-energy shallow marine environment
77
TABLE 2.2: Point-counting parameters Symbol Definition Qm Monocrystalline quartz Qpq Polycrystalline quartz Qpt Tectonized polycrystalline quartz Qss Recylced sandstone grain - quartz Slt Siltstone Cht Chert P Plagioclase feldspar K Potassium feldspar Lml Low-grade metasedimentary grain Lmh High-grade metamorphic grain (e.g., schist) Lms Metamorphic grain with serpentine or chlorite Lvl Volcanic grain with laths Lvm Volcanic grain with microlitic texture Lvf Volcanic grain of felsic composition Lvv Volcanic grain with vitric texture Lvc Volcanic grain of mafic composition Lsc Carbonate grain Lsm Mudstone or shale grain
Qm Monocrystalline quartz (Qm) Qt Total quartz (Qm+Qpq+Qpt+Qss+Slt+Cht) F Total feldspar (P+K) L Total lithics (All – (Qm+ Qpq+Qpt+Qss+Slt+Cht+P+K)) Lmet Total metamorphic grains (Qpt+Lml+Lmh+Lms) Lv Total volcanic grains (Lvl+ Lvm+ Lvf+ Lvv+ Lvc) Lsed Total sedimentary grains (Slt+Cht+Lsc+Lsm+Qss) Qp Total polycrystalline quartz (Qpq+Cht)
Acc Accessory minerals
78
TABLE 2.3 - Basins Support Contrary Extensional Stratigraphic succession of coarse fluvial ‘Basal’ fluvial strata have compositions identical to deposits overlain by marine limestone is underlying marine strata; i.e., fluvial deposits are not the basin typical of extensional basins. base of a genetic succession. Avaliable geochronology suggests fluvial No Early Cretaceous normal faults in the region; structural units were deposited rapidly, as expected in data indicate the area was under compression. syn-faulting phase of extension. Reported thickness changes in age- Thermal subsidence required to accommodate limestone equivalent deposits to the west - consistent deposition necessitates up to 50% extension (Beta factor with a basin segmented by normal faults. >2) – no evidence. Differences between conglomerates suggest Modal sandstone composition is equivocal and thickness multiple depocenters - consistent with an changes in age-equivalent strata can be produced in extended terrane. many tectonic settings. Volcanic grains within lower Cretaceous Extensional model invokes slab “rollback” during the Early sandstones to the west have compositions Cretaceous, which is inconsistent with convergence suggesting a bimodal volcanic source. rates. No contemporaneous basalts or other volcanic deposits in the area; little reported elsewhere. Evidence of bimodal volcanism is tenuous, based on a single altered tuff.
Foreland Thrust faults were active in the northern No Early Cretaceous thrust faults documented in this basin Lhasa terrane during the Early Cretaceous. specific area. Strata thicken northward, consistent with a History of subsidence rate lacks the convex-up shape peripheral foreland basin model. typical of foreland basins. Majority of paleocurrents indicate a sediment Reconciling a thick, widespread limestone with a source to the north, suggesting uplift in that peripheral foreland basin is difficult; as is explaining locale. subsidence in southern Lhasa. Earliest Cretaceous strata were deposited in Modal sandstone compositions are equivocal and in the a peripheral foreland basin. north are more volcanic-rich than expected. Including turbidite sandstones in underlying upper Jurassic strata, the overall succession is similar to that in other foreland basins (e.g., flysch-molasse).
Dynamic Creates widespread subsidence like that An ephemeral and poorly understood process that can subsidence recorded by the mid Cretaceous limestones leave an ambiguous sedimentary record. Subsidence rates are relatively constant Stratigraphic thickening to the north is incongruent with through time what would be expected. Widespread calc-alkaline volcanism in the Volcanic rocks suggest a flat slab beneath the Lhasa later stages of the Early Cretaceous is terrane – no modern analogue includes an area of flat consistent with a subducting slab beneath slab subduction and widespread subsidence in the the Lhasa terrane. overriding plate. Creates accommodation without major faulting – enables widespread limestone deposition without a significant clastic input.
Others Eustatic sea- Explains widespread nature of limestones; Cannot account for the thickness of the limestone strata level rise accommodation is created without faults. in northern Lhasa Strike-Slip No evidence of contemporaneous strike-slip faulting; Faulting subsidence rates were much lower and more persistent than those of strike-slip basins; strike-slip basins fail to explain widespread limestone deposition; strike slip basins are generally small. Crustal Paleocurrents indicate southward transportation; loading-south inconsistent with stratigraphic thickness patterns.
79
Table 2. DR1 – Point-Counting Data
Sample Qm Qpq Qpt Qss Silt Chert K P Lml Lmh Lms Lvl Lvm Lvf Lvv Lvc Lsc Lsm
Chumulong Fm - south
PB2-56 368 22 4 3 5 15 0 0 2 0 2 0 2 2 0 23 PB2- 104 364 29 5 5 3 17 0 0 3 2 0 0 0 0 5 0 0 6 PB2- 166 381 14 10 1 1 19 0 0 10 4 0 0 0 0 0 0 9 PB2- 329 394 12 1 6 6 6 0 0 11 2 0 0 0 0 0 0 11 PB2- 387 280 68 17 12 4 14 11 16 0 0 1 1 1 16 PB2- 413 277 60 16 13 10 44 0 0 8 4 4 14 PB2- 452 294 70 9 5 8 35 0 0 4 3 1 0 5 4 1 10 PB2- 542 277 73 16 9 3 37 11 15 1 1 7
MQ4-47 331 28 2 6 4 3 11 1 3 0 4 4 1 38
Mq4-82 358 24 3 2 2 11 16 2 1 1 30
AB-1 314 50 8 13 8 21 0 0 6 3 0 0 2 1 10 12
Duba Formation - north
DB1-0 114 24 7 2 51 7 0 65 33 8 2 40 11 19 9 0 29 DB1- 252 101 18 12 11 40 9 0 66 44 11 0 56 13 11 11 0 32 DB1- 321 97 43 6 11 29 2 0 88 39 6 2 47 10 1 15 1 31 DB1- 417 102 24 8 8 36 7 0 79 53 5 4 38 18 18 18 2 21 DB1- 588 139 30 5 5 34 5 0 57 33 11 0 25 3 27 24 1 5 29 DB1- 620 162 37 4 8 17 5 0 49 39 16 0 24 4 20 29 4 1 21
DB2-0 244 23 17 4 20 0 3 55 4 6 37 0 18 4 2 6
DB3-2 262 14 3 2 40 1 5 4 24 3 3 23 0 13 7 1 9
DB3-58 239 14 5 1 6 0 15 93 8 4 9 2 8 8 8
DB4-2 257 18 7 3 2 1 67 14 3 18 4 21 15 1 1 4
DB6-15 236 21 7 5 7 1 15 78 12 1 25 1 16 7 1 5 80
Table 2.DR2 - Detrital zircon data
10%uncertainty 30%disc 6/8 7/5 6/7 BEST Sample age ±(Ma) age ±(Ma) age ±(Ma) AGE ± 5% reverse disc
DBA Duba Formation - from Duba area DB1252-33 123 2 155 22 687 332 123 2 123 2 DB1252-41 123 4 90 33 -706 1090 123 4 123 4 DB1252-21 123 4 175 35 957 449 123 4 123 4 DB1252-90 124 2 132 14 287 261 124 2 124 2 DB1252-76 124 1 137 9 367 156 124 1 124 1 DB1252-4 124 3 134 22 317 393 124 3 124 3 DB1252-71 125 4 144 82 470 1479 125 4 125 4 DB1252-75 126 4 150 29 558 447 126 4 126 4 DB1252-65 126 3 160 26 701 366 126 3 126 3 DB1252-80 126 2 111 11 -205 250 126 2 126 2 DB1252-84 127 3 161 31 706 452 127 3 127 3 DB1252-9 127 3 138 31 323 551 127 3 127 3 DB1252-31 130 6 128 34 89 680 130 6 130 6 DB1252-47 130 2 136 14 242 248 130 2 130 2 DB1252-86 130 4 116 26 -179 594 130 4 130 4 DB1252-92 131 5 144 22 374 359 131 5 131 5 DB1252-39 131 3 130 53 119 1069 131 3 131 3 DB1252-59 131 2 178 27 861 341 131 2 131 2 DB1252-48 131 2 136 4 209 58 131 2 131 2 DB1252-50 132 7 164 44 657 619 132 7 132 7 DB1252-56 132 4 156 22 534 328 132 4 132 4 DB1252-58 132 2 166 42 681 588 132 2 132 2 DB1252-34 137 2 131 19 31 363 137 2 137 2 DB1252-63 137 2 155 18 441 276 137 2 137 2 DB1252-69 137 1 146 11 294 187 137 1 137 1 DB1252-44 137 2 172 30 688 412 137 2 137 2 DB1252-100 138 3 143 16 227 280 138 3 138 3 DB1252-19 138 2 128 18 -52 362 138 2 138 2 DB1252-30 139 2 143 13 209 232 139 2 139 2 DB1252-57 139 1 134 8 42 159 139 1 139 1 DB1252-67 139 1 146 6 264 104 139 1 139 1 DB1252-83 139 2 140 11 153 203 139 2 139 2 DB1252-77 139 2 138 10 120 182 139 2 139 2 DB1252-23 139 3 134 11 35 212 139 3 139 3 DB1252-16 140 4 146 22 254 370 140 4 140 4 DB1252-94 140 4 189 39 859 466 140 4 140 4 DB1252-78 140 1 142 12 175 202 140 1 140 1 DB1252-29 141 2 146 14 239 239 141 2 141 2 DB1252-37 141 4 143 16 189 266 141 4 141 4 81
DB1252-45 141 2 147 14 253 238 141 2 141 2 DB1252-42 141 3 154 10 350 158 141 3 141 3 DB1252-61 141 1 157 22 394 336 141 1 141 1 DB1252-25 142 3 158 9 414 130 142 3 142 3 DB1252-3 142 2 137 37 59 702 142 2 142 2 DB1252-6 142 2 140 25 114 456 142 2 142 2 DB1252-98 142 3 163 14 474 193 142 3 142 3 DB1252-91 143 2 123 25 -252 549 143 2 143 2 DB1252-52 143 2 139 11 74 205 143 2 143 2 DB1252-99 143 2 166 53 512 783 143 2 143 2 DB1252-7 143 3 137 21 38 391 143 3 143 3 DB1252-22 143 1 136 12 18 222 143 1 143 1 DB1252-73 144 2 142 8 101 136 144 2 144 2 DB1252-88 144 2 160 17 405 251 144 2 144 2 DB1252-24 144 3 154 17 309 273 144 3 144 3 DB1252-26 144 2 159 19 382 290 144 2 144 2 DB1252-20 145 2 141 8 70 132 145 2 145 2 DB1252-62 145 6 200 30 911 326 145 6 145 6 DB1252-74 145 3 150 9 230 137 145 3 145 3 DB1252-75 145 4 207 31 987 334 145 4 145 4 DB1252-54 146 2 153 15 264 246 146 2 146 2 DB1252-28 146 2 147 16 156 275 146 2 146 2 DB1252-15 146 2 144 13 105 236 146 2 146 2 DB1252-14 147 2 171 20 512 285 147 2 147 2 DB1252-60 148 2 144 13 81 230 148 2 148 2 DB1252-79 149 2 149 13 146 224 149 2 149 2 DB1252-43 149 3 159 12 308 180 149 3 149 3 DB1252-8 152 8 167 18 380 231 152 8 152 8 DB1252-64 153 3 167 20 360 291 153 3 153 3 DB1252-68 158 11 171 14 355 132 158 11 158 11 DB1252-11 165 6 156 15 23 226 165 6 165 6 DB1252-36 209 16 210 19 218 150 209 16 209 16 DB1252-2 228 4 275 8 704 63 228 4 228 4 DB1252-13 229 2 227 7 206 78 229 2 229 2 DB1252-53 264 5 331 73 836 544 264 5 264 5 DB1252-66 277 3 292 18 418 157 277 3 277 3 DB1252-96 287 3 309 27 476 220 287 3 287 3 DB1252-10 327 5 330 42 347 334 327 5 327 5 DB1252-11 597 11 668 31 914 126 597 11 597 11 DB1252-46 604 6 571 29 443 147 604 6 604 6 DB1252-38 619 12 623 18 638 69 619 12 619 12 DB1252-27 631 6 627 18 612 80 631 6 631 6 DB1252-40 722 8 702 27 638 111 722 8 722 8 DB1252-93 768 11 770 22 775 79 768 11 768 11 DB1252-55 822 21 845 18 907 31 822 21 822 21 DB1252-87 844 13 883 11 980 21 844 13 844 13 DB1252-72 860 36 861 30 864 50 860 36 860 36 DB1252-95 1051 42 1045 29 1031 22 1031 22 1031 22 82
DB1252-5 1105 14 1091 21 1062 58 1062 58 1062 58 DB1252-51 1075 10 1071 29 1062 87 1062 87 1062 87 DB1252-89 1158 11 1171 10 1196 20 1196 20 1196 20 DB1252-85 1115 86 1243 68 1474 81 1474 81 1474 81 DB1252-97 1854 16 1822 38 1785 80 1785 80 1785 80 DB1252-32 2301 29 2325 17 2346 19 2346 19 2346 19 DB1252-1 2432 20 2438 13 2443 17 2443 17 2443 17 DB1252-12 2938 40 2878 19 2836 16 2836 16 2836 16
PNB J-K strata from the Penbo area PB2452-8 138 2 157 27 448 418 138 2 138 2 PB2452-90 141 3 143 10 174 159 141 3 141 3 PB2452-53 141 2 170 19 588 267 141 2 141 2 PB2452-73 142 2 151 24 295 392 142 2 142 2 PB2452-89 142 2 156 19 379 296 142 2 142 2 PB2452-100 142 3 177 18 675 231 142 3 142 3 PB2452-71 142 4 140 30 94 552 142 4 142 4 PB2452-91 143 2 144 18 169 312 143 2 143 2 PB2452-38 144 2 144 10 149 167 144 2 144 2 PB2452-78 144 7 188 46 780 560 144 7 144 7 PB2452-65 144 6 146 72 173 1311 144 6 144 6 PB2452-4 145 3 138 19 15 349 145 3 145 3 PB2452-45 145 3 153 16 284 257 145 3 145 3 PB2452-33 145 2 180 22 666 286 145 2 145 2 PB2452-35 145 5 162 27 414 400 145 5 145 5 PB2452-36 145 2 158 16 358 244 145 2 145 2 PB2452-86 147 5 108 28 -687 758 147 5 147 5 PB2452-37 148 3 140 17 2 305 148 3 148 3 PB2452-13 149 3 146 6 108 100 149 3 149 3 PB2452-60 149 3 162 15 354 227 149 3 149 3 PB2452-85 150 4 175 31 533 427 150 4 150 4 PB2452-3 151 2 146 17 70 289 151 2 151 2 PB2452-97 152 8 187 48 651 605 152 8 152 8 PB2452-39 157 2 190 41 615 519 157 2 157 2 PB2452-12 165 5 185 16 448 194 165 5 165 5 PB2452-87 213 2 223 43 334 495 213 2 213 2 PB2452-25 226 4 186 24 -301 363 226 4 226 4 PB2452-77 285 3 291 13 343 109 285 3 285 3 PB2452-51 412 9 465 70 735 399 412 9 412 9 PB2452-21 505 7 479 24 357 141 505 7 505 7 PB2452-22 512 14 558 108 749 533 512 14 512 14 PB2452-70 518 18 482 26 318 129 518 18 518 18 PB2452-52 519 7 555 48 701 240 519 7 519 7 PB2452-14 521 5 519 8 511 36 521 5 521 5 PB2452-46 523 10 508 15 441 74 523 10 523 10 PB2452-88 524 6 529 18 549 93 524 6 524 6 PB2452-64 526 6 537 15 587 72 526 6 526 6 PB2452-18 528 5 515 11 454 56 528 5 528 5 83
PB2452-11 536 6 526 27 480 142 536 6 536 6 PB2452-29 538 11 515 28 413 149 538 11 538 11 PB2452-47 543 8 544 31 550 156 543 8 543 8 PB2452-30 554 8 561 16 593 74 554 8 554 8 PB2452-81 556 8 556 9 556 35 556 8 556 8 PB2452-63 564 8 563 15 560 68 564 8 564 8 PB2452-84 570 10 554 66 489 342 570 10 570 10 PB2452-26 627 6 632 15 650 64 627 6 627 6 PB2452-10 631 9 628 16 615 67 631 9 631 9 PB2452-34 656 11 674 22 734 85 656 11 656 11 PB2452-96 665 8 664 9 659 29 665 8 665 8 PB2452-27 676 10 689 30 732 124 676 10 676 10 PB2452-54 678 10 714 40 831 161 678 10 678 10 PB2452-58 697 8 700 19 712 75 697 8 697 8 PB2452-28 828 33 852 26 917 29 828 33 828 33 PB2452-93 856 11 870 16 904 48 856 11 856 11 PB2452-31 877 16 930 52 1059 169 877 16 877 16 PB2452-75 880 8 896 10 935 27 880 8 880 8 PB2452-74 889 9 882 18 865 61 889 9 889 9 PB2452-50 893 11 910 20 954 64 893 11 893 11 PB2452-61 902 12 861 29 756 102 902 12 902 12 PB2452-49 902 26 905 25 914 58 902 26 902 26 PB2452-6 904 9 961 9 1096 20 904 9 904 9 PB2452-15 907 10 920 9 952 20 907 10 907 10 PB2452-98 939 9 947 9 966 20 939 9 939 9 PB2452-56 955 10 956 9 956 21 955 10 955 10 PB2452-7 981 11 987 14 1000 37 981 11 981 11 PB2452-94 1032 17 1040 16 1058 34 1058 34 1058 34 PB2452-20 1110 10 1098 22 1075 64 1075 64 1075 64 PB2452-19 930 23 974 35 1075 99 1075 99 1075 23 PB2452-1 922 37 970 30 1081 42 1081 42 1081 37 PB2452-48 1032 19 1049 19 1085 44 1085 44 1085 44 PB2452-43 1077 29 1094 22 1128 29 1128 29 1128 29 PB2452-80 1174 17 1163 25 1144 64 1144 64 1144 64 PB2452-67 1161 21 1161 20 1160 42 1160 42 1160 42 PB2452-17 1109 11 1129 34 1169 97 1169 97 1169 97 PB2452-95 1161 11 1174 13 1196 30 1196 30 1196 30 PB2452-42 1137 27 1160 35 1205 85 1205 85 1205 85 PB2452-59 1178 15 1194 18 1224 41 1224 41 1224 41 PB2452-99 1167 11 1198 10 1255 20 1255 20 1255 20 PB2452-44 1238 15 1251 24 1276 59 1276 59 1276 59 PB2452-79 1283 12 1314 11 1363 20 1363 20 1363 20 PB2452-62 1457 22 1471 60 1491 144 1491 144 1491 144 PB2452-55 1638 35 1648 30 1661 52 1661 52 1661 52 PB2452-23 1705 25 1743 16 1789 18 1789 18 1789 18 PB2452-72 2376 38 2423 20 2462 17 2462 17 2462 17 PB2452-41 2218 30 2364 17 2492 17 2492 17 2492 17 PB2452-83 2650 22 2652 13 2653 17 2653 17 2653 17 84
PB2452-16 2610 21 2644 13 2670 17 2670 17 2670 17 PB2452-76 3114 67 3248 28 3332 16 3332 16 3332 16 PB2452-5 3343 65 3353 26 3358 16 3358 16 3358 16
85
CHAPTER 3: THE TAKENA FORMATION OF THE LHASA TERRANE,
SOUTHERN TIBET: THE RECORD OF A LATE CRETACEOUS RETRO-ARC
FORELAND BASIN
ABSTRACT Before we can understand the development of the Tibetan plateau, we first need to determine the tectonic setting and crustal conditions of southern Tibet prior to the
Indo-Asian collision. Several different hypotheses have been proposed that describe the
Late Cretaceous tectonic and paleogeographic conditions of southern Tibet, but the data needed to test these models are largely lacking. We examined the upper Cretaceous
Takena Formation of the Lhasa terrane in southern Tibet in order to reconstruct the mid-
Late Cretaceous tectono-sedimentary history of this area. The Takena Formation consists of >2 km of sedimentary strata that include a basal limestone member (Penbo
Member) and an upper member of clastic red beds (Lhunzhub Member). The Penbo
Member consists of ~250 meters of orbitolinid-limestone beds that were deposited in a shallow marine setting during Aptian-Albian time. Decompacted stratigraphic thicknesses indicate relatively minor subsidence during limestone deposition. The lower part of the overlying Lhunzhub Member contains fluvial deposits with abundant paleosol horizons. The paleosol-rich section is overlain by the remainder of the
Lhunzhub Member, which consists of >1,500 meters of fluvial sandstone interstratified with floodplain mudstone. The fluvial sandstone units are lithic-rich and contain abundant plagioclase and volcanic grains, suggesting a Gangdese magmatic arc
86
provenance. Paleocurrent directions record northwest-directed transport and
stratigraphic data indicate the subsidence rate increased during deposition of these
fluvial units. In terms of paleogeography, the Lhasa terrane was initially flooded by a
shallow-marine sea during Aptian-Albian time, wherein carbonate sedimentation
dominated. Clastic sediment derived from the south infilled existing accommodation
and the marine environment was replaced by northwest-flowing braided and
meandering river systems. The strata of the Takena Formation were then folded and
partially eroded by 69 ±2.4 Ma. The sedimentary and stratigraphic characteristics of the
Takena Formation are most consistent with deposition in a retro-arc foreland basin
setting. The low subsidence rate during deposition of the basal limestone member and
the overlying paleosol-rich interval record the passage of a foreland basin forebulge.
The overlying southerly-derived fluvial strata record deposition in the foredeep
depozone and Late Cretaceous folding of the Takena Formation indicates the foreland
basin strata were eventually incorporated into a northward-advancing fold-thrust belt.
Thus, the sedimentary record requires that a fold-thrust belt existed along the southern margin of the Lhasa terrane during the Late Cretaceous and prior to the Indo-Asian collision. Collectively, the evidence indicates that the southern margin of the Lhasa terrane had thickened crust and was therefore likely at high elevations immediately prior to the Indo-Asian collision.
INTRODUCTION
Tectonic processes play a fundamental role in shaping the geography and physiographic features of the Earth. In no other region is this more evident than in the
87
Tibetan plateau, which currently stands >5 km above sea-level (Figure 3.1; Fielding,
1996). This region serves as the archetype of continent-continent collisions, and its uplift and weathering are thought to have influenced local and global systems (e.g.,
Raymo et al., 1988, Raymo and Ruddiman, 1992; Molnar et al., 1993). Despite its importance, there is much about the geological history of the area that remains poorly understood.
Little is known of the crustal and topographic conditions of the Lhasa terrane of southern Tibet immediately before the Indo-Asian collision. Current hypotheses run the gamut from thickened crust and high topography (e.g., England and Searle, 1986;
Murphy et al., 1997), to thinned crust with surface elevations below sea-level (e.g.,
Zhang, 2000; Figure 3.2). In other words, as India approached the southern margin of
Asia it is not known whether southern Tibet more closely resembled the modern
Okinawa Trough or the modern Bolivian Andes.
Three models have been proposed to explain the mid-Late Cretaceous tectonic evolution of the Lhasa terrane of southern Tibet (Figure 3.2). Beginning in the Early
Cretaceous, Neotethyan oceanic crust was consumed in a north-verging subduction zone along the southern margin of Eurasia, which ultimately led to the production of the
Gangdese magmatic arc (e.g., Tapponnier et al., 1981; Burg et al., 1984; Allégre et al.,
1984). Thus, this area is often postulated as having been an “Andean-style” convergent margin during the mid-Late Cretaceous (e.g., Allégre et al, 1984; England and Searle,
1986; Ratschbacher et al., 1993; Chen et al., 1993; Fielding, 1996). But how “Andean” was this margin? The retro-arc region in the Andes has a fold-thrust belt, surface
88 elevations greater than 6 km, and crustal thicknesses in excess of 70 km (e.g., Beck and
Zandt, 2002). With the exception of locally folded Cretaceous strata, evidence supporting this hypothesis is relatively sparse. Existing paleocurrent data from upper
Cretaceous rocks indicate sediment was derived from northern Lhasa (Leeder et al.,
1988), contrary to what is expected in the retro-arc foreland basin model. Furthermore, not all retro-arc regions are compressional. The second proposed model suggests that the Lhasa terrane was occupied by a back-arc extensional basin located atop thinned continental crust (Zhang, 2000). Alternatively, the third tectonic model hypothesizes that the defining tectonic event within the Lhasa terrane during this time may have had nothing to do with subduction of Neotethyan oceanic crust. It is possible that tectonism in southern Tibet was dominated by the Late Jurassic-Cretaceous collision of the Lhasa terrane with what was then the southern margin of Asia. The resulting uplift and crustal loading along the accretionary suture should have produced a peripheral foreland basin in the Lhasa terrane (Leeder et al., 1988; Dürr, 1996; Murphy et al., 1997).
The sedimentary record of the ~2 km thick mid-upper Cretaceous Takena
Formation can be used to test the proposed tectonic models. In this paper we present:
(1) detailed descriptions of the lithologies, petrography, facies and regional architecture of the Takena Formation in southern Tibet; (2) reconstructions of sediment sources, transport directions, and depositional histories of the upper Cretaceous sediment; (3) interpretations of Cretaceous paleogeographies and basin development; and (4) an interpretation of the Late Cretaecous tectonic setting of the southern Lhasa terrane that is consistent with the sedimentological and structural data. Our conclusions are based
89 on >5500 meters of detailed measured stratigraphic sections, 593 paleocurrent measurements gathered at 36 locations, new biostratigraphic constraints, and petrographic data from 42 thin sections, augmented with clast counts. Our results indicate the upper Cretaceous Takena Formation was deposited in a retro-arc foreland basin, north of a Cretaceous fold-thrust belt.
REGIONAL GEOLOGY
Stratigraphy and Age Control
The Takena Formation is exposed in the Lhasa terrane of southern Tibet (Figure
3.1). It overlies lower Cretaceous marginal marine strata and is in turn unconformably overlain by upper Cretaceous-lower Tertiary volcanic strata of the Linzizong Formation
(Figure 3.3; Yin et al., 1988). In the southern half of the study area (Penbo and Maqu areas; Figure 3.1) the Takena Formation consists of a basal limestone member (Penbo
Member) and an overlying succession of fluvial red beds (Lhunzhub Member; Figure
3.3; Yin et al., 1988). Fossil ages indicate these limestones were deposited during
Aptian through late Albian time. The fossil assemblage includes: echinoids Macraster and Salenia (common in Albian), amminoids of the family Parahoplitidae, most similar to Kazanskyella fosteri or K. cuchillensis (late Aptian), and foraminifera Orbitolina
(Mesorbitolina) texana Roemer (late Aptian to early Albian) (R.W. Scott, personal communication). In the northern half of the study area, the Cretaceous limestone thickens and biostratigraphic evidence indicates it was deposited over a longer period of time (Barremian-Cenomanian; Smith and Xu, 1988; Zhang, 2004; this study). This thicker limestone unit is called the Langshan Formation (Yin et al., 1988; Leeder et al.,
90
1988). Therefore, the Takena Formation in the northern portion of the study area consists only of fluvial red beds; however, the general lithologic progression from carbonate to clastic red beds remains the same. The depositional age of the terrestrial clastic section overlying the basal limestone is poorly constrained, although Smith and
Xu (1988) reported a Cenomanian bivalve from a carbonate bed within the succession.
The youngest U-Pb ages of detrital zircons collected from fluvial sandstone cluster at
105 ± 2 Ma, providing a maximum depositional age for the upper part of the Takena
Formation (Leier et al., 2004). The minimum depositional age is constrained by U-Pb zircon ages of 69 ± 2.4 Ma from volcanic strata within the overlying Linzizong
Formation (He, 2005). A 90 Ma igneous dyke cuts across unspecified beds of the
Takena Formation (Coulon et al., 1986), so that although unambiguous constraints restrict deposition of the Takena Formation to between ~105-70 Ma, it is likely that deposition began in roughly late Albian time (~105-100 Ma) and ended in Coniacian-
Campanian time (~90-80 Ma).
Tectonic History
The Lhasa terrane was the last of a series of continental fragments to accrete onto southern Asia during the Phanerozoic before the collision with India (Allégre et al., 1984; Dewey et al., 1988; Yin and Harrison, 2000). Rifting from Gondwanaland during the Late Permian-Early Triassic, the Lhasa terrane moved northward as Meso-
Tethyan oceanic crust north of the Lhasa terrane was subducted northward beneath the
Qiangtang terrane (e.g., Allégre et al., 1984; Dewey et al., 1988; Sengör and Natal’in,
1996). The northern boundary of the Lhasa terrane is the Bangong suture, which formed
91 as a result of the collision between the Lhasa and Qiangtang terranes during the Late
Jurassic-Early Cretaceous (Figure 3.1; Dewey et al., 1988; Yin and Harrison, 2000).
Along the southern margin of the Lhasa terrane, Cretaceous-Paleogene subduction of
Neotethyan oceanic crust produced the Cretaceous-Tertiary Gangdese magmatic arc
(e.g., Allégre et al., 1984; Coulon, et al., 1986). The Indus-Yarlung suture is the southern boundary of the Lhasa terrane and formed as a result of the Indo-Asian collision, which occurred sometime between 70-38 Ma (see Butler, 1995; Yin and
Harrison, 2000). During deposition of the Takena Formation, paleomagnetic data place the Lhasa terrane at roughly 10 degrees north latitude (Achache et al., 1984; Besse and
Courtillot, 1988; Chen et al., 1993).
SEDIMENTOLOGY OF THE TAKENA FORMATION
Lithofacies
Lithofacies within the Takena Formation are typical of carbonate and clastic rocks, and are well understood in terms of depositional processes. Individual lithofacies are described and interpreted in Table 3.1. The following paragraphs focus on genetic associations of lithofacies that can be more broadly interpreted in terms of depositional systems.
Lithofacies Associations and Formation-scale Architecture
Depositional environments of the Takena Formation are interpreted from 18 detailed stratigraphic sections measured in the Penbo, Maqu and Nam Co areas (Figure
3.1). A composite stratigraphic section of the Takena Formation is presented in Figures
92
3.4-3.6. The individual lithofacies are grouped into 6 lithofacies associations: 1) a basal marine limestone (between ~50 and 250 meters in Figure 3.5); 2) a marginal marine siltstone (between ~250 and 300 meters in Figure 3.5); 3) a lower fluvial lithofacies association (between ~300 and 1200 meters in Figure 3.5); 4) a middle limestone lithofacies association (between ~0 and 50 meters in Figure 3.6); 5) a middle siltstone lithofacies association (between ~50 and 800 meters in Figure 3.6); and 6) an upper fluvial lithofacies association (between ~800 and 1100 meters in Figure 3.6). These lithofacies associations are presented in the order of their stratigraphic position within the Takena Formation, beginning with the lowermost unit (Figures 3.5, 3.6).
Lower Marine Limestone Lithofacies Association - Description
In the southern portion of the study area, the lowest part of the Takena
Formation consists of orbitolinid-bearing limestone and minor amounts of clastic siltstone (Lithofacies C1, F3). These lithologies conformably overlie lower Cretaceous marginal marine siltstone and sandstone (between ~0-50 meters in Figure 3.5). The lowermost limestone of the Takena Formation consists of wackestone that contains fossils of orbitolinids (Mesorbitolina sp.), echinoids (Macraster sp., Selina sp.), and rare oysters (Ceratosreon (?)) and ammonites (Kazanskyella sp.) (R.W. Scott, personal communication). Overlying these wackestone beds are a series of ~3 m thick limestone cycles composed, from base to top, of marly siltstone, orbitolinid-wackestone
(Mesorbitolina texana (Roemer); R.W. Scott, personal communication), and coarse, fossiliferous orbitolinid packstone, which also contain echinoderm debris, rare ostracods, and disarticulated bivalves (Figure 3.7). The packstone beds at the top of
93 each cycle have sharp upper surfaces, occasionally with fossiliferous lags, and are overlain by marly siltstone of the overlying cycle. Individual cycles can be correlated between sections > 10 km apart. In the Maqu area (Figure 3.1), the lower orbitolinid- bearing limestone lacks obvious cycles but is otherwise similar to the limestone in the
Penbo region. Limestone beds in the Nam Co region are poorly exposed but consist of laminated to massive pelloidal mudstone (Lithofacies C3). Rudist boundstones of
Aptian-Albian age have been described elsewhere in the Lhasa terrane (Leeder et al.,
1988; Leier et al., in prep). Laminated green siltstone and local very fine-grained sandstone with oscillatory current ripples are present in minor amounts within this lithofacies association in the Maqu and Penbo locations (Lithofacies F3).
Lower Marine Limestone Lithofacies Association - Interpretation
The basal marine limestone lithofacies association records deposition in a low- energy, carbonate-dominated shallow-water marine environment. The lowermost wackestone in the Penbo area, the pelloidal mudstone in the Nam Co area, and the wackestone in the Maqu area, are interpreted to have been deposited in quiet, muddy, lagoonal environments (e.g., Leeder et al., 1988; Tucker and Wright, 1990; Jones and
Desrochers, 1992). Based on the presence of rudist boundstone elsewhere in the region, these strata were likely deposited in shallow-marine lagoons that were situated between rudist patch reefs (e.g., Leeder et al., 1988; Leier et al., ). The carbonate cycles are interpreted as carbonate parasequences that record repeated shoaling and flooding in a shallow marine environment (e.g., Inden and Moore, 1983; Jones and Desrochers,
1992). The presence of siltstone within this lithofacies association in the Penbo and
94
Maqu areas suggests that although these areas were characterized by carbonate sedimentation, they also received occasional influxes of fine-grained clastic material.
Marginal Marine Siltstone Lithofacies Association - Description
The marginal marine siltstone lithofacies association conformably overlies the lower marine limestone lithofacies association (Figure 3.5). This lithofacies association consists of green siltstone and interbedded very fine-grained sandstone (Figure 3.8;
Lithofacies F3). Both the siltstone and sandstone contain marine fossils, including disarticulated bivalves and rare ammonites (Kazanskyella - fosteri or cuchillensis; R.W.
Scott, personal communication). Siltstone intervals commonly grade into sandstone.
Beds of sandstone typically contain oscillatory and unidirectional current ripples, wavey-bedding, mud-draped oscillatory current ripples, and subvertical and subhorizontal burrows (Lithofacies S5). Occasionally, stacked beds of sandstone coarsen-upward from very fine-grained to fine-grained.
Marginal Marine Siltstone Lithofacies Association - Interpretation
Strata of the marginal marine siltstone lithofacies association were deposited in a low-energy shallow marine environment. The upward-coarsening siltstone and sandstone sequences containing oscillatory current ripples and mud-draped ripples are interpreted as deposits of prograding, low-energy shorelines or tidal flats (e.g., Walker and Harms, 1971; Elliot, 1974; Reinson, 1992). Deposits with similar characteristics can be associated with deeper-water offshore environments; however, a shallow-water interpretation is preferred because this lithofacies association grades conformably into overlying terrestrial deposits (Figure 3.5; see below).
95
Lower Fluvial Lithofacies Association - Description
The marine siltstone lithofacies association grades into red mudstone, sandstone
and local conglomerate belonging to the lower fluvial lithofacies association (Figure
3.5). Strata of the lower fluvial lithofacies association are characterized by ~10 m thick
sandstone units interstratified with mottled, red, CaCO3 nodule-bearing mudstone
(Figure 3.9). Two types of sandstone units occur within the lower fluvial succession,
which are designated type A and type B sandstone units. The lower fluvial lithofacies
association in the Penbo area consist largely of the type A sandstone units, whereas the
Maqu and Nam Co areas have both type A and type B sandstone units.
The type A sandstone units overlie a basal scour surface, are generally <10 m
thick, and in their lowest portions have medium-grained sandstone beds with abundant mudstone intraclasts. Sandstone in the lower two-thirds of the units is trough cross- stratified, whereas in the upper third the sandstone contains plane parallel laminations and unidirectional current ripples (Lithofacies S1, S2, S4). Large-scale epsilon cross-
bedding (>2 m amplitude) that dips in directions perpendicular to the paleocurrent
direction (as measured by trough cross-stratification; see below) occurs in a few
sandstone units (Figure 3.5). As a whole, the sandstone units fine-upward continuously
and lack prominent internal scour surfaces or abrupt changes in grain size (Figure 3.5).
Sandstone beds at the top of a sequence are typically massive or bioturbated and grade
into overlying red siltstone (Lithofacies S6).
Type-B sandstone units (~10 m thick) overlie basal scour surfaces and are composed of 2 or more ~2-4 meter thick sandstone packages separated from one
96
another by scour surfaces. Individual packages tend to be laterally discontinuous,
typically fine-upward and contain trough-cross stratification as well as occasional large,
planar cross-stratification (~1 m amplitude; Lithofacies S1, S3). Plane parallel
laminations and rare current ripples are present in sandstone beds in the upper ~0.5 m of
many packages, but these beds are often partially eroded by the basal scour of the
overlying sandstone package. Although the constituent sandstone packages may fine-
upward slightly, the grain-size of the sandstone sequence as a whole remains relatively
constant from base to top. In some locations in the Maqu area these units also contain
beds of pebbley-conglomerate (Lithofacies G1). Normally, the contact between the
uppermost sandstone and the overlying siltstone is sharp.
The mudstone and siltstone between both type A and B sandstone units are
typically mottled and contain CaCO3 nodules and rare root traces (Figure 3.9;
Lithofacies F1). Interbedded with the mudstone is thinly-bedded, fine-grained sandstone that contains unidirectional current ripples, climbing ripples, and sub-vertical burrows.
Lower Fluvial Lithofacies Association -Interpretation
The grain-size trend, facies, and architecture of type A sandstone units are typical of sandy meandering river deposits (e.g., Allen, 1964; Miall, 1996). The epsilon
cross-bedding is interpreted as accretion sets (Allen, 1964); their orientation relative to
paleocurrent direction indicates lateral migration of the paleochannel. The upward- fining trend within the sandstone units and the corresponding change from sub-aqueous dune trough cross-stratification to ripple cross stratification (Lithofacies S1, S4) reflect
97
a progressive decrease in flow strength as the river channel was infilled and migrated
laterally across the floodplain (e.g., Allen, 2001).
The architectural elements, grain size trend, and internal facies of the type B
sandstone units are most consistent with deposition by low-sinuosity, sandy braided
rivers (Cant and Walker, 1978; Miall, 1978; Willis, 1993a; Miall, 1996). Multilateral
and multistory sandstone units are most often associated with unstable channels like
those in sandy braided streams (Miall, 1996). Trough-cross stratified sandstone was
deposited by subaqueous 3-dimensional dunes that migrated within the paleochannels.
The sandstone beds with relatively larger and more planar cross-stratification are
interpreted as deposits of migrating 2-dimensional dunes or tranverse bars. The plane-
parallel laminated sandstone near the top of individual packages is interpreted as having been deposited within either infilled channels or on bar-tops where the depth of flow is commonly shallow; local siltstone beds deposited atop the plane-parallel laminated sandstone are interpreted to have been deposited during the waning stages of flow (e.g.,
Langford and Bracken, 1987; Bristow, 1993). Multiple internal scour surfaces within sandstone units indicate periods of channel reoccupation or subsequent increases in
flow strength.
Mudstone intervals within the lower fluvial lithofacies association are
interpreted as overbank sediments deposited in floodplain environments that were
located adjacent to fluvial channels (Allen, 1964). The rippled, fine-grained sandstone
records periodic crevasse splay events. Mottled mudstones that lack primary
sedimentary structures and contain CaCO3 nodules are interpreted as paleosols (calcisol;
98
Mack et al., 1993) that formed on relatively well-drained portions of the floodplain.
Paleosols are abundant within this lithofacies association; in some instances 4 or more paleosols are stacked atop one another (Figure 3.9), suggesting periods with relatively low sedimentation rates (e.g., Kraus, 1999).
Middle Limestone Lithofacies Association - Description
The middle limestone lithofacies association consists of ~50 m of limestone that overlie the lower fluvial lithofacies association in the Penbo and Maqu localities (Figure
3.6); this portion of the Takena Formation is not exposed in the Nam Co area. The limestone consists of oyster-bearing packstone and wackestone beds (<0.5m thick) interbedded with carbonate mudstone (Lithofacies C2). Unlike limestone at the base of the Takena Formation, these deposits do not contain organized cycles and lack orbitolinid tests and macrofossils other than oysters.
Middle Limestone Lithofacies Association -Interpretation
The middle limestone lithofacies association is interpreted as having been deposited in a low-energy marginal marine environment, most likely one with estuarine or lagoonal characteristics. Low-diversity fossil assemblages that are dominated by oysters and fine-grained sediment are common in estuarine/lagoonal environments (e.g.,
Tucker and Wright, 1990).
Middle Siltstone Lithofacies Association - Description
Conformably overlying the middle limestone lithofacies association is a thick succession of monotonous red siltstone with minor amounts of sandstone and local limestone beds, which comprise the middle siltstone lithofacies association (Figure 3.6,
99
10). The red siltstone does not contain CaCO3 nodules or obvious bedding (Lithofacies
F2). The majority of the siltstone has a pervasive cleavage, making the identification of sedimentary structures difficult. Sandstone beds are fine-grained and contain unidirectional current ripples, climbing ripples, occasional convoluted bedding, and small trough cross-stratification (set thickness < 0.5m). Also within this lithofacies association are 2-6 m thick sandstone units. The sandstone units overlie minor scour surfaces, fine-upward, and contain trough-cross-stratification and large amounts of unidirectional current ripples and climbing ripples. All of the sandstone within the lithofacies association contain little to no evidence of bioturbation. The sandstone/mudstone ratio within this lithofacies association is relatively low (e.g.,
Figures 3.6, 3.10). Locally, the middle siltstone lithofacies association contains thin and laterally continuous, ostracod-bearing limestone beds and laminated marly-siltstone.
Middle Siltstone Lithofacies Association - Interpretation
The features within this succession, and its stratigraphic proximity to underlying oyster-bearing limestone beds, suggest a lower-coastal plain depositional environment.
The principle depositional components are interpreted to have been non-migrating streams, large interfluvial areas and small freshwater lakes. The thin- and thickly- bedded sandstone containing unidirectional current ripples and climbing ripples is interpreted as crevasse splay and channel levee deposits. The 2-6 m thick sandstone units encased within the siltstone are interpreted as deposits of relatively fixed channel rivers, most similar to anastomosing streams (e.g., Smith and Smith, 1980). Similar fluvial channels with well developed levees are common in modern and ancient lower
100
coastal plain environments (e.g., Törnqvist, 1993). The siltstone was deposited in
floodplains within low-lying coastal plain areas that were inundated by crevasse splay
events (e.g., Elliot, 1974). The ostracode limestone and laminated siltstone are
interpreted as deposits of lower coastal plain lacustrine environments.
Upper Fluvial Lithofacies Association - Description
The upper fluvial lithofacies association consists of thick sandstone units
interstratified with red mudstone (Figures 3.6, 3.10). The sandstone units of this
lithofacies association are in many ways identical to the type B sandstone units within
the lower fluvial lithofacies association, and will therefore be described only briefly.
The sandstone units consist of multiple, stacked, 2-4 meter thick, medium-grained
trough-cross stratified sandstone packages (Lithofacies S1). Whereas these sandstone
units are less common in the lower fluvial lithofacies association and generally ~10 m
thick, within the upper fluvial lithofacies association they are ubiquitous and tend to be
coarser-grained and much thicker than the equivalents in the lower portion of the
Takena Formation (Figures 3.5, 3.6). The successions within the Maqu region are
composed of the same sandstone successions as those described above, but they also
contain occasional clast-supported pebble-conglomerate beds.
Mudstone beds within the upper fluvial succession are massive and similar to
mudstone in the lower fluvial lithofacies association; however, the mudstone intervals
in the upper fluvial lithofacies association are typically devoid of CaCO3 nodules. The relative proportion of sandstone to mudstone is noticeably higher within the upper fluvial lithofacies association compared to the lower fluvial lithofacies association.
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Upper Fluvial Lithofacies Association - Interpretation
Similar in almost all respects to the type B sandstone units of the lower fluvial lithofacies association, the sandstone units of the upper fluvial succession are interpreted as deposits of sandy braided streams (more detailed description and references presented above). The mudstone is interpreted as having been deposited in floodplain environments adjacent to the fluvial channels.
PALEOCURRENT DATA
Paleocurrent data consist of measurements made from 572 limbs of trough cross-stratification sets and 21 primary current lineations preserved within upper
Cretaceous fluvial sandstone units. The orientations of cross-stratification sets were measured following method I of DeCelles et al. (1983). The paleocurrent data record north-northwest directed sediment transport during the Late Cretaceous (Figures 3.5,
3.6). In the Penbo area, paleocurrent directions at base of the Takena Formation range from southwest- to northeast-directed, but predominately indicate north-northwest directed transport. Sandstone units at the base of the Takena Formation in the Maqu area contain more north- and northeast-directed indicators than the lowermost sandstone units of the Penbo area, and as a whole, record north-directed sediment transport.
Paleocurrent directions in the upper portion of the Takena Formation indicate transport was almost exclusively to the northwest and varied little (Figure 3.6). Based on the paleocurrent data, sediment transport in the Nam Co area was also northwest-directed.
Thus, the paleocurrent data strongly suggest sediment of the Takena Formation was derived from an area ~south-southeast of the Penbo, Maqu and Nam Co locales.
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PETROGRAPHY
Description
Standard petrographic thin sections were made from 42 medium-grained sandstone samples that were collected while measuring stratigraphic sections. The thin sections were stained for potassium and calcium feldspar and point-counted (450 counts per slide) using a modified Gazzi-Dickinson method (Ingersoll, 1984); the modification involves the identification of monocrystalline quartz grains that are part of sedimentary lithic fragments. The petrographic counting parameters are shown in Table 3.2, and raw point-count data can be accessed in Table 3. DR.1.
Sandstone within the Takena Formation consists of calcite-cemented feldspathic litharenite and litharenite beds, which collectively have average modal compositions of
Qm:F:L = 41:18:41, and Qt:F:L = 55:19:26 (Figures 3.11, 3.12). Subangular to well- rounded monocrystalline quartz is the primary constituent of total quartz within the samples. Sandstone units in the Penbo area have the highest percentage of total quartz, and those in the Maqu area have the lowest. Feldspar content is uniform between the different locations and sections (~19% of modal composition), and with the exception of one unit in the Maqu area, consists entirely of plagioclase feldspar (Figure 3.11).
Volcanic fragments are the predominant lithic grain within the Takena Formation, and constitute an average of 17% of total framework grains (Figure 3.12). Volcanic grains are typically andesitic and often contain microlaths of plagioclase feldspar (Figure
3.11). In addition to volcanic grains, the sandstone contains lithic grains derived from sedimentary and metamorphosed strata, including phyllite, mudstone, sandstone, and
103 quartzite (Figure 3.11). Lithic fragments of calcium carbonate are typically micritic, and occasionally contain unidentifiable bioclasts. Tourmaline, epidote, and zircon are the most common accessory minerals.
Interpretation
The composition of sandstone within the Takena Formation is most similar to sandstone derived from arc-related recycled orogens (Figure 3.12). Potential source rocks in the southern Lhasa terrane include: slightly metamorphosed Carboniferous mudstone and quartzose sandstone, with local Permian limestone; Triassic limestone, volcaniclastic strata, and basalt; Jurassic limestone and mudstone; lower Cretaceous quartzose sandstone, mudstone and limestone; and granites and granodiorites of Early-
Late Cretaceous age (Kidd et al., 1988; Yin et al., 1988; Liu, 1988; Leeder et al., 1988).
The significant proportion of volcanic and plagioclase grains within sandstone of the Takena Formation indicates that the Cretaceous Gangdese volcanic arc located along the southern margin of the Lhasa terrane supplied a significant portion of the sediment to the Takena Formation. It is likely some of the quartz grains in the Takena
Formation were also derived from the siliceous Cretaceous arc rocks as well. In addition to volcanic rocks, Paleozoic and Mesozoic sedimentary strata probably supplied sediment to the upper Cretaceous strata in the Lhasa terrane. Lithic grains of sandstone, limestone, mudstone and phyllite all must have been derived from sedimentary cover strata exposed along the southern portion of the Lhasa terrane (Figure 3.11). The presence of well-rounded quartz grains in compositionally immature sandstone suggests at least some of the quartz is recycled from quartzitic sedimentary units (Figure 3.11).
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This is supported by the fact that Cretaceous sedimentary strata locally unconformably overlie Carboniferous units; and the distribution of U-Pb ages of detrital zircons from the Takena Formation is similar to that of Carboniferous sandstone exposed in the
Lhasa terrane (Leier et al., 2004; Leier et al., in prep). Not all of the plagioclase grains within the Takena Formation were necessarily derived from arc-volcanic rocks.
Plagioclase-rich Triassic volcaniclastic strata are exposed in the southern Lhasa terrane
(Figure 1) and could have served as a sediment source to the upper Cretaceous strata.
Maqu Provenance
Several aspects of the Takena Formation sandstone exposed in the Maqu area suggest the sediment deposited in this region had a slightly different provenance than equivalent strata exposed in the Penbo and Nam Co locations. Relative to sandstone compositions at all of the other locations, sandstone in the Maqu area has the highest percentage of volcanic fragments, the largest amount of tectonized polycrystalline grains and recycled sandstone fragments, and Maqu is the only location where there is sandstone containing grains of potassium feldspar. Grain-size, paleocurrent directions and fluvial facies within the Takena Formation in the Maqu region also differ slightly from those features in the Takena Formation in the Penbo and Nam Co locales. The relatively high percentage of tectonized quartz and the presence of potassium feldspar grains imply the fluvial systems that traversed the Maqu area may have been draining a more dissected region of the source area (e.g., Dickinson and Suczek, 1979).
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SUBSIDENCE
The Late Jurassic – Eocene subsidence history of the southern Lhasa terrane is
reconstructed using stratigraphic data from this study and from the literature (Leeder et
al., 1988; Yin et al., 1988; Leier et al., in prep). Late Jurassic-Eocene strata were
decompacted following the methods outlined in Allen and Allen (1990), and tectonic
subsidence was calculated using the techniques described in Allen and Allen (1990) and
Angevine et al. (1990).
The subsidence history associated with the Cretaceous-Eocene deposits of the southern Lhasa terrane is displayed in Figure 3.13. Relatively poor age control exists for the fluvial deposits of the Takena Formation (Lhunzhub Member), so the decompacted stratigraphic thickness curve was calculated using various depositional durations. The general character of the subsidence curve remains the same regardless of the depositional duration; the shorter the period of deposition the steeper the curve.
Subsidence rates were constant during latest Jurassic and earliest Cretaceous time, and then decreased during Aptian and Albian time (Figure 3.13). Beginning in the Late
Cretaceous, subsidence rates increased and continued at high levels through Paleocene time. The subsidence curve calculated for southern Lhasa exhibits an overall convex- upward shape typical of subsidence histories in foreland basins (e.g., Angevine et al.,
1990). In foreland basin settings this trend is ascribed to initial deposition within a backbulge or forebulge depozone (low subsidence rate) followed by deposition within a foredeep depozone (increased subsidence rate; e.g., DeCelles and Giles, 1996).
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Several aspects of the subsidence history warrant further consideration. No
constraints on surface paleoelevation exist for the time period encompassing deposition
of the terrestrial red beds of the Takena Formation. Estuarine/lagoonal limestone beds
within the middle portion of the Takena Formation imply that the succession was
deposited near sea level, at least during deposition of the lower half of the formation.
Any increase in surface elevation that occurred during deposition of the Takena
Formation, or the Linzizong Formation, would reduce the slope of the subsidence curve during the latter stages of deposition. Eustatic sea-level changes were also not taken into account, which again, could reduce the amount of calculated tectonic subsidence.
However, we suspect the total subsidence and subsidence rates calculated for the
Cretaceous-Eocene strata are conservative estimates. A thicker stratigraphic succession or a briefer depositional period would have the effect of increasing the total amount of subsidence and the subsidence rates. The upper surface of the Takena Formation is a regional unconformity; hence, the total thickness of the Takena Formation is unknown and the ~2 km of strata exposed today represent only what has escaped erosion.
Similarly, the 3 km of volcaniclastic sediment of the Linzizong Formation may have once been as thick as 6 km or more (e.g., Copeland et al., 1995), although whether these strata were deposited in a foredeep depozone is debatable. Folded beds of the Takena
Formation that underlie flat-lying beds of the Linzizong Formation (69 ± 2.4 Ma; He,
2005) obviously necessitate that deposition of the Takena Formation must have ceased before ~69 Ma. The presence of a 90 Ma dyke that cross-cuts an unspecified section of the Takena Formation (Coulon et al., 1986) suggests the cessation of deposition was
107 closer to 90 Ma than 69 Ma, and therefore, the decompacted thickness curve calculated using a 15 Myr duration (~105-90 Ma) is probably more realistic than the curve calculated for a 35 Myr duration (~105-70 Ma; Figure 3.13).
PALEOGEOGRAPHY AND TECTONIC SETTING
Utilizing all of the data presented thus far, we present paleogeographic reconstructions and interpretations of the tectonic setting of the Lhasa terrane during
Early-Late Cretaceous time (Figure 3.14). The temporal evolution is constrained by the relative stratigraphic position of particular lithofacies associations, and absolute ages are delimited by fossil data from limestone within the Takena Formation and zircon crystallization ages from samples of the overlying Linzizong Formation (e.g., He,
2005). The major paleogeographic features within the reconstructions (e.g., fluvial systems) are directly supported by evidence presented within this paper, or are otherwise noted as being inferred from other data (e.g., paleoshoreline trends).
The characteristics of the Takena Formation indicate deposition occurred in a retro-arc foreland basin. The modal composition of sandstone within the Takena
Formation is most similar to retro-arc foreland basin deposits (Figure 3.12; e.g.,
Dickinson and Suczek, 1979; DeCelles and Hertel, 1989; Jordan, 1995). The overall upward-coarsening nature of the formation and the upward increase in sandstone thickness and sandstone/mudstone ratio are also common features of foreland basin strata (e.g., Willis, 1993a, 1993b). The temporal and spatial pattern (south-to-north) of clastic deposition is consistent with a foreland basin model (Figures 3.3, 3.14), as are the convex-upward subsidence curve (Figure 3.13), and paleocurrent data (Figures 3.5,
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3.6). With respect to the paleostress field, folded beds of the Takena Formation that underlie undeformed strata of the Linzizong Formation unequivocally show that the retro-arc region of the Lhasa terrane was subjected to compressional forces during the
Late Cretaceous (Figure 3.15), which is consistent with the northward migration of a
Gangdese retro-arc fold-thrust belt into the foreland basin.
Lower Strata of the Takena Formation
Deposition of the Takena Formation initiated in Aptian time as southern Tibet was submerged beneath a shallow epeiric sea that contained rudist patch reefs and low- energy, muddy inter-reef areas (Figure 3.14A). Northward flowing streams delivered minor amounts of clastic material to the area but this sediment was relatively insignificant compared to the volume of carbonate deposition. An east-west trending paleoshoreline located south of the study area is inferred from the relative abundance of interbedded clastic deposits in more southerly exposures and assuming paleoshoreline trend was roughly perpendicular to the fluvial paleocurrent direction (northward- directed; see Figure 3.5). A nascent Gangdese magmatic arc located to south of the study area is interpreted to have existed by this time based on Aptian-Albian clastic strata with volcanic detritus and north-directed paleocurrent indicators, and Xigaze
Forearc sedimentary strata of Aptian-Albian age (Dürr, 1996).
By the end of the Albian, clastic material from the south began to progressively infill the region in a south-to-north manner (Figure 3.14B). Clastic, shallow marine environments replaced carbonate environments in the south, although in northern portions of the Lhasa terrane carbonate deposition continued into the Cenomanian (Yin
109 et al., 1988). Eventually the existing marine accommodation was filled with clastic sediment and the entire area, with the possible exception of the northernmost parts of the Lhasa terrane, was occupied by terrestrial environments. North-northwest-flowing meandering river systems were present in the Penbo area, whereas contemporaneous rivers in the Maqu area were more similar to braided streams and most likely originated in a different watershed than the Penbo fluvial system. Relatively intense soil forming processes were active in the floodplain areas adjacent to the fluvial channels.
The lithology of the Penbo Member of the Takena Formation and the low subsidence rate during its deposition are most consistent with deposition in a backbulge or forebulge depozone (Figures 3.14A, F; e.g., Dorobeck, 1995; DeCelles and Giles,
1996). Evidence of extensive paleosol development within the terrestrial deposits overlying the limestone of the Penbo Member is also suggestive of a forebulge, where very low rates of sediment accumulation result from flexural uplift and the consequent loss of accommodation (Figures 3.14B, G; e.g., DeCelles and Giles, 1996).
Middle Strata of the Takena Formation
Following subaerial exposure and extensive palesol formation, marine to brackish-water conditions returned to the area and deposition occurred in lower coastal plain settings with local estuarine/lagoonal environments (Figure 3.14C). The region at this time was primarily characterized by anastomosing fluvial channels with interfluve regions containing small lakes.
These facies suggest an increase in the accommodation-sedimentation ratio relative to the underlying paleosol deposits (e.g., Törnqvist, 1993). The fine-grained
110 anastomosing stream deposits and estuarine strata are most similar to distal foredeep deposits (Figure 3.14H), which are characterized by an increased subsidence rate
(Figure 3.13) and a relatively low influx of sediment (e.g., Shuster and Steidtmann,
1987).
Upper Strata of the Takena Formation
During the latter stages of deposition of the Takena Formation, northwest- flowing sandy braided stream systems occupied the southern half of the Lhasa terrane
(Figure 3.14D). Whether this area was occupied by a single river system or multiple systems with similar characteristics and flow directions is not discernible at present.
The different sandstone composition in the Maqu area suggests the fluvial system in this location was still draining a slightly different source area.
The upward-coarsening fluvial succession within the upper part of the Takena
Formation is interpreted as recording deposition in more proximal parts of the foredeep depozone (Figure 3.14I). The overall fluvial architecture of the Takena Formation, from isolated meandering stream deposits with abundant paleosols at the base, to multistory and multilateral braided stream deposits at the top, is consistent with what is observed in other foreland basin successions (e.g., Willis 1993a, 1993b; Horton and DeCelles,
2000).
Deformation of the Takena Formation
At some time following the deposition of the upper fluvial succession and before
69 Ma (He, 2005), the beds of the Takena Formation were folded, eroded and then
111 buried by the volcanic flows and tuffs of the Linzizong Formation (Figure 3.14E; Figure
3.15).
The evidence of Late Cretaceous folding of the Takena Formation (Figure 3.15) is interpreted as representing the incorporation of the foreland basin strata into the advancing fold-thrust belt (Figure 3.14J), which occurs in fold-thrust belts and foreland basin systems around the world (e.g., DeCelles and Giles, 1996).
The Migration of the Deformation Front
A simple test of the foreland basin model can be performed by estimating the rate at which a foreland basin and the associated deformation front must have moved in order to explain the stratigraphy and deformation of the Takena Formation (Figure
3.14F-J). These rates then can be compared with similar calculations derived from retro- arc foreland basin settings. We use the following parameters to make an initial estimate of the migration rate: the distance between the deformation front and the forebulge crest is 200-300 km; deposition on the forebulge crest occurred at roughly 105 Ma (end of the Albian; Figure 3.14F, G); and the deformation front reached the Penbo and Maqu locations by 69 Ma (oldest age of the Linzizong Formation overlying folded Takena
Formation beds; Figure 3.14 j). The migration rate of the flexural wave (i.e., rate of the advancing fold-thrust belt) is 6-9 mm/yr given the above assumptions. Using a more restricted duration between forebulge deposition and deformation of the Takena
Formation of 20 Myr yields rates of 10-15 mm/yr. It should be emphasized that we are not implying that these calculations are the actual migration rates, which, with the available data and the myriad unknowns, are impossible to reconstruct at this time.
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However, these values are meaningful in the sense that they are all consistent with rates of frontal thrust belt migration as derived from modern GPS measurements and from ancient foreland basin strata (e.g., Norabuena et al., 1998; DeCelles and Horton, 2003).
DISCUSSION
Where is the Cretaceous Fold-Thrust Belt?
An obvious and mandatory corollary to a Late Cretaceous retro-arc foreland basin in southern Tibet is the existence of a Late Cretaceous retro-arc fold-thrust belt in southern Tibet. Currently, no such fold-thrust belt has been identified. However, evidence of a Cretaceous-aged “Gangdese retro-arc fold-thrust belt” is steadily accumulating (England and Searle, 1986; Kapp et al., 2004). Interformational relationships provide ample evidence that the Takena Formation was folded during the
Late Cretaceous, with deformation recording at least 26-50% Late Cretaceous shortening (Figure 3.15; Pan, 1993; Burg et al., 1983; Kapp et al., 2004). The intensity of this deformation decreases from south to north (Burg et al., 1983; Coward et al.,
1988), which is consistent with a north-verging retro-arc fold thrust belt. Structural relationships, in conjunction with recently discovered kinematic indicators, suggest north-verging thrust sheets were active in the southern portion of the Lhasa terrane during the Late Cretaceous (Kapp et al., 2004). Furthermore, magmatic evidence indicates that by 80 Ma parts of the crust in the southern Lhasa terrane were abnormally thick (Wen et al., 2004). Clearly, more work is needed before this issue can be resolved
113 but existing data suggest that a Late Cretaceous fold-thrust belt was present in southern
Tibet.
Implications for Plateau Development
The marked difference between the extent of deformation in the Cretaceous
Takena Formation and the overlying Late Cretaceous – Paleocene Linzizong Formation
(Figure 3.15; e.g., Pan, 1993) has led to the hypothesis that prior to the collision with
India, the crust of southern Asia had obtained topographic elevations and buoyancy forces of sufficient magnitude such that horizontal forces produced during the Indo-
Asian collision were transmitted inboard (to the north) and not absorbed via internal strain (England and Searle, 1986). The evidence of a retro-arc foreland basin in southern Tibet is consistent with Late Cretaceous thickening of the region prior to the
Indo-Asian collision. Cretaceous marine limestones are commonly invoked as evidence that southern Asia was at or near sea-level prior to the Indo-Asian collision. However, within southern Tibet, the youngest limestone is Cenomanian in age (Smith and Xu,
1988), which leaves a significant period of time (~40 Myr) before the collision during which considerable surface and rock uplift may have occurred. Moreover, the presence of marine limestone can support the hypothesis of low elevations at specific areas but it does not necessarily support low regional elevations, as is evinced in many modern compressional orogens where high topography is flanked by areas below sea-level.
Previous Paleogeographic Interpretations
The Xigaze Group strata were deposited coevally with the Takena Formation in a forearc basin located to the south of the Gangdese magmatic arc (Einsele et al., 1994).
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Based on sandstone composition and previous regional studies (e.g., Leeder et al.,
1988), Dürr (1996) proposed a Late Cretaceous paleogeography wherein sediments derived from the Bangong suture and the Qiangtang terrane (Figure 3.1) were transported southward across the Lhasa terrane, through the Gangdese arc, and deposited in the Xigaze forearc basin. Our data, particularly the paleocurrent directions, indicate the Xigaze forearc basin and the retro-arc basin in the southern Lhasa terrane were not in communication with respect to sediment transport, at least in this area of the
Lhasa terrane. The sediment of the Xigaze Formation must have been derived from the southern side of the Gangdese arc. Sandstone beds within the Xigaze Formation that contain anomalously high percentages of quartz (with respect to other forearc basins) are perplexing, but it is possible the quartz was recycled from quartzose units within
Carboniferous and Early Cretaceous strata.
CONCLUSIONS
The Takena Formation of the Lhasa terrane in southern Tibet is composed of a lowermost marine limestone member and an overlying clastic member that together are approximately 2 km thick. The limestone beds were deposited within a shallow-marine, lagoonal environment that was eventually inundated by clastic sediment derived from southern source areas. The overlying clastic strata consist of fluvial sandstone units and floodplain mudstone. Paleocurrent data indicate the fluvial systems that traversed this portion of the Lhasa terrane flowed toward the northwest. Petrographic data indicate the strata of the Takena Formation were derived primarily from a Late Cretaceous volcanic arc that was located along the southern margin of the Lhasa terrane, although pre-Late
115
Cretaceous sedimentary cover strata also served as a sediment source. The shallow marine seaway that occupied the region during the Aptian-Albian was eventually infilled by clastic material in a south to north progression. Subsidence rates were initially low during deposition of the limestone but increased noticeably during deposition of the terrestrial clastic deposits. Combined, all of the data indicate the
Takena Formation was deposited in retro-arc foreland basin. This requires the existence of a Cretaceous retro-arc fold-thrust belt within the southern portion of the Lhasa terrane and also implies the region may have had thickened crust before the Indo-Asian collision.
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Figure 3.1: Location map and geology of the study area. The Lhasa terrane is the southernmost of several terranes that compose the Tibetan plateau. The study area includes the territory between the city of Lhasa and the lake Nam Co. Sections were measured in three principle locations: Penbo, Maqu and Nam Co. Geology is modified from Liu (1988).
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Figure 3.2: Proposed tectonic models. A. The Takena Formation may have been deposited in a retro-arc foreland basin that developed on the northern side of the Gangdese magmatic arc (e.g., England and Searle, 1986). B. Accommodation for the Takena Formation may have been caused by back arc extension, possibly associated with southward roll-back of Neothethyan oceanic crust (e.g., Zhang, 2000). C. The Cretaceous tectonic setting may have been largely controlled by deformation that resulted from the Late Jurassic-Early Cretaceous collision between the Lhasa terrane and the Qiangtang terrane to its north (e.g., Leeder et al., 1988). Not to scale.
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Figure 3.3: Generalized Cretaceous stratigraphy. The Takena Formation consists of a basal limestone member, but is primarily composed of clastic red beds belonging to the Lhunzhub Member. From south to north, Aptian-Albian limestone increases in thickness and was deposited over a greater time span. Modified from Yin et al., (1988).
119
Figure 3.4: Stratigraphic symbols. Key to symbols used in the stratigraphic sections in figures 3.5 and 3.6.
120
Figure 3.5: The lower part of the Takena Formation in the Penbo area. The Takena Formation contains a basal succession of limestone (~50-250 m), which is conformably overlain by clastic strata of the Lhunzhub Member. The base of the Lhunzhub Member contains numerous paleosols. See also, Figures 3.7, 3.8, and 3.9.
121
122
Figure 3.6: The upper part of the Takena Formation in the Penbo area. A thin oyster- bearing limestone unit is overlain by a thick succession of relatively fine-grained strata, which coarsen-upward and contain thick sandstone sequences. Note, the contact between the lower (Figure 3.5) and upper (this figure) part of the Takena Formation is not exposed. See also, Figure 3.10.
123
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Figure 3.7: Photos of the Penbo Member of the Takena Formation. A. Carbonate parasequence (beds are dipping to the left, person for scale). The Penbo Member contains a series of stacked parasequences that consist of, from base to top, marly siltstone and carbonate mudstone, orbitolinid-wackestone, and an orbitolinid and bivalve packstone. B. Close up of an orbitolinid-packstone. Not the abundance of orbitolinid fossils. C. Orbitolinid fossil in a petrographic thin section of the limestone within the Penbo Member of the Takena Formation (plane-light).
125
Figure 3.8: Photograph of marginal marine clastic deposits. Note the fine-grain size of the sediment, and the alternation between very fine-grained sandstone laminea and mudstone.
126
Figure 3.9: Lower part of the Takena Formation. The Penbo Member of the Takena Formation and the overlying terrestrial red beds of the Lhunzhub Member in the Penbo area (see also Figure 3.5). A. An overview of the transition from the shallow marine limestone beds to the Lhunzhub Member (beds are dipping to the left). B. Photo of the terrestrial red beds overlying the Penbo Member (beds are dipping into the page and to the right). Note the numerous stacked horizons of red-purple siltstone, which are interpreted as paleosols horizons. C. Close-up of paleosol deposits (beds are dipping into the page and to the right). Note the massive nature of the siltstone and the CaCO3 nodules.
127
Figure 3.10: Upper part of the Takena Formation. Overview of the upper part of the Lhunzhub member of the Takena Formation (beds are dipping into the page and to the left; see also Figure 3.6). Thick fluvial sandstone sequences overly a succession of siltstone and fine-grained sandstone.
128
Figure 3.11: Petrography. Photos of petrographic thin sections of sandstone from the Takena Formation (all photos in cross-polarized light). A. Example of a typical sandstone within the Takena Formation. B. Well- to very well-rounded quartz grains surrounded by unstable lithic volcanic grains and plagioclase, suggesting the quartz has been recycled. C. Volcanic grain within the Takena Formation. Note the small laths of plagioclase feldspar within the volcanic grain. D. Sedimentary lithic grains are common within the Takena Formation. E. Two tectonized polycrystalline quartz grains from the Takena Formation.
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Figure 3.12: Sandstone composition. QFL diagrams constructed from point-counting 42 thin sections of Takena Formation sandstone. See Table 3.2 for explanation of modes. A, B, Sandstone composition of the Takena Formation generally plots within or near the recycled orogen field (fields from Dickinson and Suczek, 1979). C. Sandstone of the Takena Formation contains abundant lithic-volcanic grains, suggesting an arc-related source. The average composition of the underlying lower Cretaceous sandstone (and 1 standard deviation) is also shown to highlight the contrast in composition between the two stratigraphic successions. D. Relative amounts of monocrystalline quartz (Qm), plagioclase (P) and potassium feldspar (K) within the Takena Formation.
130
Figure 3.13: Subsidence. Decompacted thickness and tectonic subsidence curve for Cretaceous strata in the southern portion of the study area (Penbo and Maqu areas). Because of poor age-control the decompacted thickness of the Takena Formation is displayed with alternative durations of deposition. Overall, the curve is most similar to those from foreland basin settings. See text for further discussion.
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Figure 3.14: Paleogeography. Paleogeographic reconstructions of the study area based on the presented data (A.-E.), and the inferred tectonic setting of the region based on the data presented here and in the literature. Note that A.-E. are reasonably well-constrained whereas F.-J. are much less so, but consistent with our results. See text for discussion.
132
Figure 3.15: Takena and Linzizong Formations. Folded beds of the Takena Formation overlain by relatively undeformed volcaniclastic beds of the Linzizong Formation in the Maqu area. Ages of the Linzizong Formation indicate that deformation of the Takena Formation occurred prior to 69 Ma.
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TABLE 3.1: Lithofacies Description Interpretation G1. Imbricated conglomerate Clast-supported pebble conglomerate; rare cobbles; imbricated clasts; Deposits associated with traction sandy matrix; crude horizontal stratification; erosional base. transport by high competency water flows S1. Trough cross-stratified sandstone Fine- to very coarse-grained trough cross-stratified sandstone (sets Deposits of migrating, sub-aqueous, typically <1m); occasionally with pebbles and granules. three-dimensional sand dunes S2. Plane parallel laminated sandstone Horizontal beds of fine- to coarse-grained, well-sorted sandstone; plane Sub-aqueous deposits associated with parallel laminations upper-flow regime conditions S3. Large, planar cross-stratified sandstone Fine- to medium-grained; large (>1 m thick) relatively low-angle (~20 Deposits of large, migrating, sub- degrees) planar to slightly trough-shaped cross-stratification; bases are aqueous, two- to three-dimensional slightly erosional transverse dunes or bars S4. Sandstone with asymmetric ripples Very fine- to medium-grained; asymmetric ripples and ripple cross- Deposits of migrating subaqueous lamination; beds < 5 cm thick; occasional sub- and super-critical ripples under a unidirectional current; climbing ripples climbing ripples deposited under high sedimentation rates S5.Sandstone with symmetric ripples Very fine- to fine-grained; symmetrical ripple laminae; muddy laminae Deposits of ripples formed in oscillating occasionally interbedded currents; muddy laminae suggest alternating energy levels S6. Massive sandstone Very fine- to coarse-grained; structureless; occasional faint laminae and Deposits of sandy gravity flows or burrows original sedimentary structures destroyed by bioturbation F1. Massive red mudstone No sedimentary structures; commonly red to purple and mottled; CaCO3 Paleosols formed in a relatively well- nodules; occasional root traces and vertical burrows (~3 cm diameter) drained environment F2. Red siltstone Siltstone, primarily red, with rare laminations and asymmetric ripples; Deposits of a low-energy environment, does not contain CaCO3 nodules; interbedded with very fine- and fine- either terrestrial or shallow marine; red grained sandstone; locally interbedded with thin ostracode-bearing beds color suggests oxidized nonmarine setting F3. Green laminated siltstone Laminate green siltstone, commonly with marine fossils; occasionally Deposits a low-energy, clastic, shallow marly; local thin beds of very-fine grained sandstone marine environment C1. Orbitolinid packstone and wackestone Laterally continuous beds with orbitolinid tests; rare bivalve and Deposits associated with a low-energy echinoid fragments shallow marine environment C2. Oyster wackestone and packstone Wackestone and packstone composed almost entirely of disarticulated Shallow-marine deposits, possibly oyster shells estuarine/lagoonal setting C3. Pelloidal mudstone Massive; micritic; rare laminations Deposits of a low-energy shallow marine environment
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TABLE 3.2: Point-counting parameters Symbol Definition Qm Monocrystalline quartz Qpq Polycrystalline quartz Qpt Tectonized polycrystalline quartz Qss Recylced sandstone grain - quartz Slt Siltstone Cht Chert P Plagioclase feldspar K Potassium feldspar Lml Low-grade metasedimentary grain Lmh High-grade metamorphic grain (e.g., schist) Lms Metamorphic grain with serpentine or chlorite Lvl Volcanic grain with laths Lvm Volcanic grain with microlitic texture Lvf Volcanic grain of felsic composition Lvv Volcanic grain with vitric texture Lvc Volcanic grain of mafic composition Lsc Carbonate grain Lsm Mudstone or shale grain
Qm Monocrystalline quartz (Qm) Qt Total quartz (Qm+Qpq+Qpt+Qss+Slt+Cht) F Total feldspar (P+K) L Total lithics (All – (Qm+ Qpq+Qpt+Qss+Slt+Cht+P+K)) Lmet Total metamorphic grains (Qpt+Lml+Lmh+Lms) Lv Total volcanic grains (Lvl+ Lvm+ Lvf+ Lvv+ Lvc) Lsed Total sedimentary grains (Slt+Cht+Lsc+Lsm+Qss) Qp Total polycrystalline quartz (Qpq+Cht)
Acc Accessory minerals
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TABLE 3. DR1 - Sandstone Composition
Sample Qm Qpq Qpt Qss Silt Chert K P Lml Lmh Lvl Lvm Lvf Lvv Lvc Lsc Lsm
PB1-32 195 26 10 10 7 12 0 82 12 3 14 6 6 20 5 5 31
PB1-42 160 18 8 13 3 7 0 124 10 2 37 9 8 24 6 14 4
PB1-55 161 20 6 10 3 18 0 96 6 1 26 11 9 20 6 44 10
PB1-95 130 23 8 12 4 10 0 91 7 4 52 9 14 43 11 29 2
PB1-105 150 19 7 12 2 7 0 132 2 9 26 1 17 31 8 21 3
PB1-130 124 18 12 19 7 13 1 106 8 3 35 19 15 23 6 38 2
PB1-170 260 26 12 12 5 18 0 48 15 3 14 5 7 4 2 3 16
PB1-188 294 23 11 8 1 8 0 46 13 5 5 1 10 16 5 2 2
PB1-254 131 20 9 9 7 25 1 121 6 3 37 17 13 21 3 11 11
PB1-266 122 30 13 14 6 13 0 80 9 6 32 8 14 26 4 64 2
PB1-275 161 50 5 13 3 25 0 68 8 5 9 7 34 36 7 8 8
PB1-297 116 21 4 12 3 11 0 95 7 2 45 16 16 17 7 38 21
PB3-118 289 47 6 2 1 13 0 2 19 8 1 14 11 0 1 35
PB3-161 233 14 7 13 2 10 103 8 2 19 16 3 16 0 4
PB3-284 249 24 17 13 1 13 0 84 4 11 8 8 2 9 0 4
PB6-8 340 31 17 5 3 11 0 3 11 3 1 4 8 13
PB6-148 240 21 18 18 4 6 0 52 7 8 34 12 3 17 1 0 9
NM2-12 164 17 6 7 3 6 0 97 13 1 49 15 6 31 4 8 22
NM2-40 120 19 12 4 3 20 96 10 3 63 25 21 39 6 6 2
NM2-146 144 34 19 13 6 19 82 35 8 16 8 11 25 1 11 31
NM2-686 203 15 7 6 6 14 0 54 31 9 9 6 6 30 2 7 40
NM2-694 223 21 12 5 7 15 42 28 10 7 6 8 17 0 0 39
MQ2-5 104 19 5 8 3 14 126 5 1 42 29 16 20 5 29 10
MQ2-104 183 26 24 19 4 34 33 5 7 54 32 7 13 5 0 4
MQ2-167 134 20 19 16 6 18 27 7 3 77 52 4 19 5 27 16
MQ3-92 131 22 20 27 12 35 11 10 5 19 22 8 23 4 89 11
MQ5-1 127 15 3 4 3 8 148 7 10 23 40 10 24 5 7 13
MQ5-110 155 29 20 25 16 21 2 64 12 4 19 20 1 15 4 14 29
MQ5-208 60 12 7 10 3 14 3 113 0 0 67 61 11 25 28 32 3
MQ5-514 87 27 7 11 5 7 39 100 5 0 65 33 4 40 10 6 2
02_10 179 32 3 2 9 12 101 27 7 4 49 8 5 11
02_18 179 26 4 15 1 91 30 2 49 3 3 3 15
02_22B 168 23 4 1 5 11 78 24 1 6 45 19 2 1 53
02_67 259 47 5 3 1 15 0 41 10 4 1 13 15 1 20 8
02_69 245 38 5 2 30 44 1 1 3 45 19 10 1
02_78 204 29 6 1 9 116 8 2 8 24 12 2 6 9
02_82 233 32 1 2 28 102 6 2 20 15 1 5
02_83 194 24 7 3 2 23 138 21 8 1 23 14 1 1
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02_84 141 44 4 6 2 74 118 23 1 11 4 4 27 2 14
02_85 195 21 2 21 162 5 2 2 1 8 9 6 12 1
02_89 296 33 2 4 5 1 50 6 1 2 18 9 6 13 3
02_90 166 26 3 1 19 91 2 17 10 44 24 18 9 4
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CHAPTER 4: DETRITAL ZIRCON GEOCHRONOLGY OF PHANEROZOIC
SEDIMENTARY STRATA EXPOSED WITHIN THE LHASA TERRANE OF
SOUTHERN TIBET
ABSTRACT
The many uncertainties surrounding the evolution of the Tibetan plateau are
related in large part to the fact that the much of the geology in this region has not been studied. Presented here is a comprehensive U-Pb detrital zircon data set (743 new U-Pb ages) from Phanerozoic sedimentary rocks exposed in the Lhasa terrane of southern
Tibet. These ages provide new constraints on the depositional and tectonic history of the
Lhasa terrane. The sampled stratigraphy includes Carboniferous, Jurassic, and
Cretaceous units. Strata near Nam Co in central Lhasa are mapped as lower Cretaceous but are devoid of detrital zircons with ages younger than 400 Ma. The detrital zircon age distribution as well as additional evidence suggests these strata are Paleozoic and not lower Cretaceous, which, if correct, will require new mapping and structural interpretations. Lower Jurassic strata exposed near the Bangong suture between the
Lhasa and Qiangtang terranes contain populations of detrital zircons with ages between
200 and 500 Ma and 1700 and 2000 Ma. These populations are noticeably different from the detrital zircon ages of samples collected elsewhere in the Lhasa terrane. The upper Cretaceous Takena Formation contains zircon populations with ages between 100 and 160 Ma, 500 and 600 Ma, and 1000-1400 Ma. Detrital zircon ages from the upper
Cretaceous strata suggest that several slightly different fluvial systems occupied the 138
southern portion of the Lhasa terrane during the Late Cretaceous and that deposition
probably ceased long before 70 Ma. Carboniferous strata exposed within the Lhasa
terrane likely served as a sediment source for Cretaceous strata deposited in the Lhasa
terrane. Similarities between the age-probability plots of Carboniferous rocks in the
Lhasa terrane with those of rocks in the Qiangtang terrane and with Tethyan strata in
the Himalaya support plate tectonic reconstructions of Paleozoic Gondwanaland.
Differences between the geographic distribution of igneous rocks within the Tibetan plateau suggest it might be possible to discern, at least to a first-order, a southern versus northern provenance signature using age populations of detrital zircons.
INTRODUCTION
In creating models depicting the Cenozoic uplift of the Tibetan plateau, we assume the tectonic and crustal conditions of southern Eurasia prior to the Indo-Asian collision are well understood. This is rather dubious considering how little Jurassic-
Cretaceous data from the region exist, and the data that have been collected suggest
Tibet, particularly southern-central Tibet, was tectonically active throughout the
Mesozoic (Burg et al., 1983, Allégre et al., 1984; Dewey et al., 1988; Sengor and
Nata’lin, 1996; Murphy et al., 1997; Yin and Harrison, 2000; Kapp et al., 2005). In terms of lithosphere-scale structure, the Tibetan plateau is not a single entity, but rather, an amalgam of continental fragments that were accreted onto southern Asia during the
Paleozoic and Mesozoic (Allégre et al., 1984; Dewey et al., 1988; Sengor and Nata’lin,
1996; Yin and Harrison, 2000). The Lhasa terrane is the southernmost of these fragments and was the last to accrete onto southern Asia, doing so sometime during 139
Late Jurassic-Early Cretaceous time (Dewey et al., 1988; Yin and Harrison, 2000). The
tectonism associated with the assembly of Tibet shaped the paleogeography and crustal
structure of the region, but exactly how, or to what extent, is the subject of debate (c.f.,
Murphy et al., 1997; Zhang et al., 2004), as is the impact of this deformation on the
subsequent development of the plateau (e.g., England and Houseman, 1986).
The relative abundance and durability of zircon and the resistance of the U-Pb
system to resetting make detrital zircon U-Pb geochronology particularly well suited for
studying an area like southern Tibet, which has experienced multiple phases of
deformation. By comparing the U-Pb ages of detrital zircons between multiple
sedimentary units and with igneous rocks, the ancient sediment sources and pathways
can be reconstructed, from which the timing and location of tectonic uplift can be
inferred (e.g., Fedo et al., 2003). To investigate the Jurassic-Cretaceous tectonic history
of the Lhasa terrane of southern Tibet, and to better characterize the sedimentary strata
of the region, we sampled 9 sandstone units exposed in the Lhasa terrane (Figure 4.1)
and determined the U-Pb ages of their detrital zircons using laser-ablation multi-
collector inductively coupled plasma mass spectrometry. Sampling focused on strata
that were deposited during Jurassic and Cretaceous time, but also included Paleozoic
rocks. The 743 new U-Pb ages reported here constitute the first comprehensive U-Pb
detrital zircon data set from the Lhasa terrane of the Tibetan plateau.
The geology of Tibet is still being explored and categorized, and the majority of the geologic studies from this area are “frontier” in nature. This investigation was undertaken with the purpose of accomplishing four primary goals, each of which has the 140
potential to contribute to our understanding of the pre-collisional geologic history of
Tibet. The first goal is to obtain maximum depositional ages for stratigraphic units in
the Lhasa terrane. The majority of the stratigraphy in the Lhasa terrane has not been
studied in detail and has been mapped only by remote sensing techniques. Depositional
ages of particular rock units are often speculative, especially when they lack
biostratigraphically useful fossils. Correctly determining the ages of stratigraphic units
is necessary for paleogeographic reconstructions and for identifying faults and crustal
deformation in the region. The second goal is to determine the provenance of sedimentary strata and to reconstruct regional paleogeographies. The ability to reconstruct the paleogeography of the Lhasa terrane is a crucial step in constraining the timing, location, and style of tectonism in the region. The third goal relates to developing a better understanding of the Paleozoic history of the Lhasa terrane.
Analyses of pre-Mesozoic detrital zircons within the samples allow for the examination of the Paleozoic relationship between the Lhasa terrane and surrounding tectonostratigraphic regions of Eurasia. And the fourth goal is to provide data that can be used as the impetus for future investigations and which can serve as a point of departure for subsequent studies.
REGIONAL SETTING
The Lhasa terrane was a constituent of Gondwanaland and located in the
southern hemisphere throughout most of the Paleozoic (Dewey et al., 1988; Sengor and
Nata’lin, 1996; Yin and Harrison, 2000). During Triassic time, the Lhasa terrane rifted
from Gondwanaland and migrated northward (Figure 4.2; Dewey et al., 1988; Yin and 141
Harrison 2000). Following rifting, the Lhasa terrane moved northward and approached the southern margin of the Qiangtang terrane, which itself had accreted onto southern
Asia during the latest Triassic (Figure 4.2; Dewey et al., 1988; Sengor and Nata’lin,
1996; Yin and Harrison, 2000). The northward motion of the Lhasa terrane was accommodated by subduction of Mesotethyan oceanic crust beneath the Qiangtang terrane (Figure 4.2). The collision between the Lhasa and Qiangtang terranes occurred during latest Jurassic-earliest Cretaceous time and resulted in a collisional orogen and the formation of a peripheral foreland basin within the northern portion of the Lhasa terrane (Dewey et al., 1988; Leeder et al., 1988; Murphy et al., 1997; Kapp et al., 2005).
The Bangong suture is the modern manifestation of the Lhasa-Qiangtang collision and demarcates the boundary between the Lhasa and Qiangtang terranes (Figures 4.1, 4.2;
Kidd et al., 1988; Liu, 1988). Crustal shortening related to the Lhasa-Qiangtang collision continued into the Early Cretaceous as did the peripheral foreland basin system in the northern half of the terrane (Leeder et al., 1988; Murphy et al., 1997; Zhang et al.,
2004; Leier, 2005). During roughly the same time, northward subduction of Neotethyan oceanic crust beneath the Lhasa terrane produced regional volcanism throughout southern Tibet (e.g., Allégre et al., 1984; Coulon et al., 1986; Kapp et al., 2005).
Widespread exposures of Aptian-Albian limestone indicate the Lhasa terrane was covered by a shallow epeiric seaway during middle Cretaceous time (Leeder et al.,
1998; Liu, 1988; Yin et al., 1988; Zhang et al., 2004; Leier, 2005). Continued subduction of Neotethyan crust along the southern margin of the Lhasa terrane ultimately resulted in the formation of the Gangdese volcanic arc (Figure 4.2; Allégre et 142
al., 1984; Dewey et al., 1988). Available structural and evidence suggests southern
Tibet had a substantially thickened crust by the end of the Cretaceous (Burg et al., 1983;
England and Searle, 1986; Pan, 1993; Ratschbacher et al., 1993; Murphy et al., 1997;
Kapp et al., 2003a, 2005).
Paleozoic sedimentary strata in the Lhasa terrane consist primarily of
Carboniferous sandstone, metasandstone, shale and phyllite, and lesser Ordovician,
Silurian and Permian limestone (Figure 4.2; Leeder et al., 1988; Liu, 1988; Yin et al,
1988). Triassic strata include interbedded limestone and basaltic volcanic units, which
are particularly prevalent along the southern margin of the terrane (Leeder et al., 1988).
Jurassic strata in the northern portion of the Lhasa terrane consist of deepwater sandstone and shale, often associated with ophiolitic assemblages (Leeder et al., 1988;
Yin et al., 1988). Age-equivalent deposits in the southern half of the terrane consist of
marine limestone and mudstone (Yin et al., 1988). Strata deposited during the Early
Cretaceous consist of clastic mudstone, sandstone and local conglomerate (Yin et al.,
1988). The uppermost strata of the lower Cretaceous succession consist of a widely
exposed marine limestone unit that was deposited during Aptian-Albian time (Yin et al.,
1988; Liu, 1988; Leier, 2005). Upper Cretaceous strata in the Lhasa terrane consist of
arkosic fluvial sandstone and mudstone sequences that were deposited in a retro-arc
foreland basin north of the Gangdese volcanic arc (Figure 4.2; Leier, 2005).
SAMPLES
Sandstone samples were collected over the past several field seasons from
Phanerozoic strata exposed within the Lhasa terrane (Table 4.DR1). In addition to 143
detrital zircon geochronology, standard petrographic thin sections from many of the
samples were point-counted (450 counts per slide) using a modified Gazzi-Dickinson
method (Ingersoll et al., 1984). Descriptions of the sampled strata are presented below.
Carboniferous Strata
Sample DMXNG is from a Carboniferous quartzose sandstone exposed
southeast of the Nyainqentanglha range (Figures 4.1, 4.2). These sedimentary units
were deposited under marine conditions and locally contain glaciomarine deposits
(Leeder et al., 1988; Yin et al., 1988). Because they are thick (>1 km) and widely
exposed, Carboniferous strata are an important stratigraphic unit in the Lhasa terrane
(Figure 4.1; Yin et al, 1988; Liu, 1988).
Jurassic
Sample LNPLA is a medium- to coarse-grained sandstone collected from a succession of Jurassic(?) sedimentary rocks exposed along the southern margin of the
Lunpola basin, near the Bangong suture (Figures 4.1, 4.2). The sandstone beds were deposited in relatively deepwater by submarine turbidity currents in the Mesotethyan
Ocean south of the Qiangtang terrane and north of the Lhasa terrane (Figure 4.2; Leeder et al., 1988). The sandstone sample is lithic-rich and has a modal composition of
Qm:F:L = 49:16:35, and Qt:F:L = 62:16:22; volcanic grains are the most abundant type of lithic grains (Qm= monocrystalline quartz, Qt=total quartz, F=feldspar, L=lithic grains). Chlorite, serpentine and pyroxene are common accessory minerals.
144
Lower Cretaceous-north–Duba Formation
Sample DUBA was collected from a medium-grained fluvial sandstone within the Duba Formation, approximately 50 km south of the LNPLA sample location
(Figures 4.1, 4.2). The lower Cretaceous Duba Formation is >1300 m thick and composed primarily of fluvial sandstone and pebble-conglomerate sequences that are interstratified with floodplain mudstone (Leeder et al., 1988; Leier, 2005). Modal composition of the sandstone in the Duba Formation is Qm:F:L = 42:16:42 and Qt:F:L
= 57:16:27 (Leier, 2005). The tectonic setting and provenance of the Duba Formation is debatable (c.f., Leeder et al., 1988; Murphy et al., 1997; Zhang et al., 2004), but paleocurrent data clearly record southwest-directed sediment transport (Leeder et al.,
1988; Leier, 2005). A detailed account of the sedimentology of the Duba Formation is available in Leier (2005).
Lower Cretaceous-south-Chumulong Formation
Sample CHMLN was collected from the lower Cretaceous Chumulong
Formation in the southern portion of the Lhasa terrane (Figures 4.1, 4.2). The
Chumulong Formation is characterized by interbedded sandstone, mudstone, and local quartzite pebble-conglomerate that were deposited in marginal marine and fluvial environments (Leier, 2005). Sandstone units within the formation are quartzose and have modal compositions of Qm:F:L = 73:0:27 and Qt:F:L = 92:0:8; zircon is the primary accessory mineral (Leier, 2005). Conglomerate beds within the lower
Cretaceous succession in the southern Lhasa terrane suggest these strata were deposited 145
relatively close to paleotopographic relief. A more detailed description of the
Chumulong Formation is available in Leier (2005).
Lower Cretaceous at Nam Co
Sample NAMCO was collected from buff-colored, lower Cretaceous(?)
sandstone exposed along the northeastern shore of Nam Co (Figures 4.1, 4.2).
Sandstone within the succession is quartzose, thin to thickly bedded, contains
symmetrical ripples, and is commonly interbedded with mudstone. These strata are
extensively folded and faulted. The sandstone has modal compositions of Qm:F:L =
95:0:5, and Qt:F:L = 96:0:4
Upper Cretaceous-Takena Formation
Four samples were collected from medium-grained fluvial sandstone of the
Upper Cretaceous Takena Formation. The samples were collected from exposures of the
Takena Formation in the Maqu area northeast of the city of Lhasa (MAQU), the Penbo
area northwest of the city of Lhasa (2 samples, PENBO1 and PENBO2), and from the
Nyainqentanglha Range south of Nam Co (NYANT; Figures 4.1, 4.2). The fluvial sandstone sequences that were sampled are part of a >1.5 km thick succession of fluvial sandstone and mudstone that was deposited by northwest-flowing rivers within a retro- arc foreland basin that developed to the north of the Gangdese magmatic arc (Leier,
2005). Sandstone sequences within the Takena Formation contain abundant plagioclase and volcanic grains and have modal compositions of Qm:F:L = 41:18:41 and Qt:F:L =
56:18:26 (Leier, 2005). A more detailed description of the sedimentology of the Takena
Formation is presented in Leier (2005). 146
METHODS
Detrital zircons were extracted from the 9 samples (~15 kg/sample) at the
University of Arizona following procedures outlined in Gehrels (2000). Once separated,
the detrital zircons were encased in epoxy within 1” ring mounts, which were then
sanded and polished to produce a smooth flat surface that exposed the interiors of the
zircon grains. One hundred individual zircon grains were analyzed from each sample.
These were selected randomly from all sizes and shapes, although grains with obvious cracks or inclusions were avoided.
Uranium-lead ages of detrital zircons were obtained using a laser-ablation, multi-collector, inductively coupled plasma mass spectrometer (LA-MC-ICP-MS). The interior of the zircon grains were ablated using a New Wave DUV193 Excimer laser operating at a wavelength of 193 nm and using a spot diameter of 35-50 microns; laser ablation pits are ~20 microns deep. With the LA-MC-ICP-MS, the ablated material is carried via argon gas to the plasma source of a Micromass Isoprobe, which is configured such that U and Pb can be measured simultaneously. Measurements are made in static mode using Faraday collectors for 238U, 232Th, 208Pb, 207Pb, 206Pb, and an ion-counting channel for 204Pb. Analyses consist of one 20-second integration with the peak centered but no laser firing (to check background levels), and twenty 1-second integrations with the laser firing on the zircon grain. At the end of each analysis a 30 second delay occurs during which time the previous sample is purged from the system and peak signal intensity returns to background levels. The contribution of Hg to the
204Pb is accounted for by subtracting the background values. Common lead corrections 147
are made using the measured 204Pb of the sample and assuming initial Pb compositions
from Stacey and Kramers (1975). A fragment of a zircon crystal of known age (564 ± 4
Ma, 2-sigma error; G.E. Gehrels unpublished data) is analyzed after every fifth zircon
analysis to correct for inter-element and Pb isotope fractionation. The ages presented
are 206Pb*/238U ages for grains less than ~1000 Ma and 207Pb*/206Pb* ages for grains whose ages are greater than 1000 Ma. Those analyses with greater than 10% uncertainty or more than 20% discordance or 5% reverse discordance are omitted from further consideration. The data from each sample are displayed on concordia diagrams and age- probability plots/histograms using the programs of Ludwig (2001). Age-probability plots depict each age and its uncertainty as a normal distribution, summing all ages from the analyzed zircons of a sample into one curve, which is then divided by the total number of analyzed zircons.
RESULTS
The data from each sample are plotted in Figures 4.3 through 4.6 (see also Table
4. DR2). The following paragraphs discuss some of the more salient features of the
data. Interpretations and regional implications of the data are discussed in the following section. All period and epoch-to-age correlations are based on the time scale compiled by Palmer and Geissmann (1999).
Carboniferous
The DMXNG sample yielded 85 ages that are of sufficient concordance and
precision to provide provenance information (Figures 4.4, 4.5). The youngest cluster of
ages has a mean of 548 ± 5, although there are a few individual zircon grains with 148
younger ages (~450-455 Ma). The largest population of detrital zircons is between 500-
600 Ma (peak at 540 Ma), with additional clusters between 1050-1300 Ma (peak at
1120 Ma), and 1700-1950 Ma (peak at 1850). Similar to many of the other samples, a
population of detrital zircons has ages around 2500 Ma.
Jurassic
Sample LNPLA yielded 95 ages that are of sufficient concordance and precision
to provide provenance information (Figures 4.4, 4.5). Based on the youngest ages
within the first significant cluster, the most robust maximum depositional age for these
strata is 270 ± 6 Ma. A small cluster (n=2) has a mean age of 179 ± 6 Ma, but the low
number of ages cast doubts on the validity of this age. A large percentage of the detrital zircons have ages between 200-500 Ma, whereas detrital zircons with ages between
500-650 Ma and 1000-1500 Ma are relatively rare. The largest cluster of detrital zircons is between 1850-2050 Ma, with a peak at ~1900 Ma.
Lower Cretaceous at Nam Co
The NAMCO sample yielded 86 ages that are of sufficient concordance and precision to provide provenance information (Figures 4.4, 4.5). The youngest cluster of ages constrains the maximum depositional age at 495 ± 8 Ma; no Mesozoic ages were measured. Similar to many of the other samples, a significant proportion of the detrital zircons have ages between 500-600 Ma. Zircons with ages of 750-1000 Ma are relatively evenly distributed across this interval, although there are peaks at ~760, 890 and 970 Ma. Also within the sample, detrital zircons with ages between 1000-1400 Ma 149
(peak at 1150 Ma) are present, as are small populations between 1700-1800 Ma and
2500-3000 Ma.
Lower Cretaceous-south-Chumulong Formation
The CHMLN sample collected from the Chumulong Formation yielded 89 ages
that are of sufficient concordance and precision to provide provenance information
(Figures 4.4, 4.5). The zircons are euhedral to rounded and the youngest ages indicate a maximum depositional age corresponding to Early Cretaceous time (143 ±1.3 Ma). The largest cluster of zircons in this sample is between 140 and 150 Ma, the second-largest is between 500-600 Ma. Similar to the overlying Takena Formation, the lower
Cretaceous deposits in the Penbo area have a significant number of Middle Proterozoic zircons with ages between 1000-1400 Ma. Very few of the zircons analyzed have ages greater than 1500 Ma.
Lower Cretaceous-north-Duba Formation
The DUBA sample yielded 95 ages that are of sufficient concordance and
precision to provide provenance information (Figures 4.4, 4.5). Almost all of the zircons
are euhedral. The youngest cluster of ages has a mean of 125 ± 1.4 Ma, providing a
maximum depositional age for the Duba Formation. Zircon grains with ages between
125-150 Ma are by far the most numerous (peak at 140 Ma), and represent over 65% of
the total population. The remaining grains are in groups between 250-340 Ma, 610-670
Ma, 700-900 Ma, and 1000-1200 Ma.
150
Upper Cretaceous-Takena Formation
Together, the 4 samples from the Takena Formation yielded 293 ages that are of
sufficient concordance and precision to provide provenance information (Figures 4.3,
4.5, 6). The analyzed zircons vary from euhedral to rounded. Cretaceous ages comprise
the largest population, with the ages generally between 105-140 Ma (peak at 120 Ma);
Early Jurassic zircons with ages between 180-200 Ma (peak at 190 Ma) form a
subordinate group within the Mesozoic population. The youngest cluster of ages is
between 100-105 Ma, which is from a fluvial sandstone within the uppermost portion of
the Takena Formation (sample PENBO2; Figure 4.6). The second-largest population of zircons has ages between 500-600 Ma and a peak at roughly 550 Ma. Each of the samples contains at least a few zircons with ages between 650-1000 Ma, although these ages are evenly distributed across this interval. A significant proportion of zircons within the Takena Formation samples have ages between 1000-1400 Ma (peak at ~1150
Ma). Relatively few zircons have ages between 1500 and 2000 Ma, but there is commonly a scattered group centered at ~2500 Ma.
DISCUSSION
General Characteristics of the Detrital Zircon Ages
Many of the sandstone samples collected from the Lhasa terrane contain zircon populations of similar ages (Figure 4.5). All Cretaceous sandstone samples (CHMLNG,
MAQU, PENBO1, PENBO2, NYANT) contain numerous zircons with ages between
110-160 Ma, and with particular subclusters at ~120 and ~140 Ma. The Cretaceous(?) 151
NYANT sample collected near Nam Co is noticeably lacking zircons younger than 400
Ma (Figure 4.5). Except for the Jurassic sandstone exposed near the Bangong suture
(sample LNPLA), sandstone sequences within the Lhasa terrane are largely devoid of
zircons with ages between 200 and 500 Ma (Figure 4.5). Detrital zircons with 500-600
Ma ages typically represent a substantial proportion of the total population within many of the samples. All samples contain at least a few zircons with ages between 650-950
Ma, although it is difficult to discern a common peak. Many of the samples also contain a significant population of zircons with ages of 1000-1400 Ma (peak at ~1150 Ma).
Whereas most of the samples have few zircons with ages between 1500 and 2000 Ma, the Jurassic LNPLA sample contains a conspicuously large cluster at ~1900 Ma (Figure
4.5). Within almost all of the samples is a population of zircons with ages scattered around 2500 Ma.
“Lower Cretaceous” strata at Nam Co
Sandstone exposed along the northeastern shore of Nam Co, and from which sample NAMCO was collected (Figure1), is mapped as lower Cretaceous (Kidd et al.,
1988), but the detrital zircon evidence suggests this designation may be incorrect. Not a single grain from ample NAMCO yielded an age younger than 400 Ma and the age distribution has characteristics similar to the Carboniferous DMXNG sample (Figure
4.4, 5). Paleozoic strata within the Lhasa terrane were likely an important sediment source (see below; Leier, 2005); it is therefore possible that the lack of <400 Ma detrital zircons in the NAMCO sample is due to the fact that these strata were derived from the erosion of Paleozoic units. However, several lines of evidence suggest the strata 152
northeast of Nam Co were not deposited during the Early Cretaceous. Lower
Cretaceous strata to the south and north of Nam Co are dominated by Early Cretaceous zircons (e.g., samples CHMLNG and DUBA; Figure 4.5) and ages of nearby plutonic and volcanic rocks suggest there was abundant Early Cretaceous volcanism throughout the area (e.g., Coulon et al., 1986; Kapp et al., 2005), which raises the question as to why not a single zircon analyzed from the NAMCO sample had an Early Cretaceous age. Similarly, zircons with ages between 150 and 400 Ma are present within other
Cretaceous stratigraphic units in the Lhasa terrane, but once again, these ages are absent in the NAMCO sample (Figure 4.5). The sandstone composition and the degree of deformation of the strata northeast of Nam Co are also anomalous in comparison to lower Cretaceous strata exposed to the north and west of the area (e.g., sample DUBA;
Figure 4.1).
The most plausible explanation is that the strata northeast of Nam Co are
Paleozoic (Carboniferous?) and not lower Cretaceous, although additional work is needed to confirm this hypothesis. Assuming these strata are indeed Carboniferous, this discovery could have important implications for understanding the crustal structure of the Lhasa terrane. A previously unrecognized fault or unconformity would have to exist in the Nam Co area in order to account for the juxtaposition of Carboniferous strata with Cretaceous rocks (e.g., Figure 4.1). Regardless of whether this relationship is due to a fault or an unconformity, a new structural model of the crust in the Nam Co area will be needed to explain how and why Carboniferous strata are exposed at the surface in this location. The detrital zircon results from the NAMCO sample may therefore 153
provide constraints that can be used to test hypotheses of the style and kinematic history
of crustal thickening in the Tibetan plateau (c.f., Zhao and Morgan, 1987; Dewey et al.,
1988; Clark and Royden, 2000; Tapponnier et al., 2001).
Jurassic Strata
Recent studies suggest the Lhasa-Qiangtang collision may have been the most
important tectonic event in southern Asia prior to the Indo-Asian collision (Figure 4.2;
e.g., Leeder et al., 1988; Murphy et al., 1997; Arnaud et al., 2003; Kapp et al., 2005;
Guynn et al., in review); however, the processes and products of this event remain
poorly understood. The Jurassic strata exposed near the Bangong suture are thought to have been deposited in a basin that was located between the Lhasa and Qiangtang terranes coeval with, or just before, this collision (Yin et al., 1988; Leeder et al., 1988)
and therefore provide a potentially valuable record of this event (Figures 4.1, 4.2).
As the Lhasa terrane moved northward during the Jurassic, reconstructions
depict northward subduction of Mesotethyan oceanic crust beneath the Qiangtang
terrane, which should have ultimately produced a magmatic arc along the southern
margin of the terrane (Figure 4.2; e.g., Guynn et al., in review). Incongruent with this
tectonic reconstruction is the scarcity of Jurassic igneous rocks within the southern
portion of the Qiangtang terrane (Liu, 1988). Provided that the strata associated with the
LNPLA sample were deposited between the southern margin of the Qiangtang terrane
and the northern margin of the Lhasa terrane (e.g., Leeder et al., 1988), it is difficult to
reconcile the few Jurassic zircons within the sample with the former existence of a
robust magmatic arc in the Qiangtang terrane. However, detrital zircon ages within the 154
LNPLA sample do not provide sufficient constraints to refute the hypothesis of a
Jurassic magmatic arc. The ages of the detrital zircons in the LNPLA sample indicate a maximum depositional age of either 270 ± 6 Ma, or a more tentative 179 ± 6 Ma (see above). The primary period of igneous activity in the southern Qiangtang terrane is thought to have been between 180-170 Ma (Guynn et al., in review). Therefore, deposition of these ‘Jurassic’ strata may predate the formation of a Qiangtang magmatic arc. If one uses the more robust 270 Ma age as a reasonable proxy for depositional age, it is possible these strata were deposited during Permo-Triassic time and would therefore be associated with the rifting of the Qiangtang terrane from northern
Gondwanaland (i.e., the Lhasa terrane) and the formation of the Mesotethyan ocean rather than the closing of the ocean.
Upper Cretaceous Strata
Determining the depositional ages of fluvial strata of the upper Cretaceous
Takena Formation is critically important for reconstructing regional paleogeography,
Late Cretaceous subsidence rates, and the timing of regional shortening (e.g., Leier,
2005). Detrital zircon ages can provide only a maximum depositional age but they do shed light on the history of Late Cretaceous sedimentation within the Lhasa terrane.
Currently, estimates of the depositional age of the fluvial strata within the Takena
Formation can be restricted with only minimal resolution; deposition occurred sometime between ~100 Ma (based on fossils in underlying limestone; Leeder et al.,
1988; Leier, 2005) and ~70 Ma (zircon crystallization age of volcanic rocks overlying the Takena Formation; He, 2005). Because the Takena Formation was deposited 155
proximal to and coeval with abundant Late Cretaceous volcanism and contains volcanic
grains and clasts (Leier, 2005), detrital zircon U-Pb geochronology may provide one of
the few approaches that can be used to better constrain depositional age. For example,
sample PENBO1 was collected from a fluvial sandstone located ~200 stratigraphic meters above limestone beds with middle Albian fossils (~105 Ma, based on the time
scale of Palmer and Geissman, 1999; Leier, 2005). The youngest detrital zircon U-Pb
ages from PENBO1 are between 110 and 115 Ma, which suggests the youngest detrital
zircon ages from a sample may be reasonable proxies for depositional ages (± 10 Ma).
Significantly, none of the detrital zircons analyzed from the Takena Formation were
younger than 100 Ma, including those from sample PENBO2, which was collected from the uppermost portion of the succession. Within the known constraints of the depositional age of the Takena Formation (70-100 Ma; Leier, 2005), the youngest ages of detrital zircons (100-105 Ma) suggest deposition probably did not continue from 100
Ma all the way until 70 Ma.
The Late Cretaceous paleogeography of the southern portion of the Lhasa
terrane was characterized by north-northwest flowing rivers and adjacent interfluvial
regions (Leier, 2005). It is unclear whether the fluvial deposits in the area are associated
with one or multiple fluvial systems. Used in concert with petrographic data, the detrital
zircon ages from the different measured sections can be used to refine the Late
Cretaceous paleogeography. A larger percentage of >500 Ma ages are present within the
MAQU sample compared to the other samples from the Takena Formation (NYANT,
PENBO1, PENBO2), which suggests the Late Cretaceous fluvial system in the Maqu 156
area was distinct from the fluvial systems to the east and north (Figure 4.6). This
evidence corroborates previous ideas based on petrographic and paleocurrent data that
suggest the Late Cretaceous river system in the Maqu area had a slightly different
source area compared to coeval fluvial systems in the southern portion of the Lhasa terrane (Leier, 2005).
Carboniferous Strata and Gondwanaland
Reconstructions of Gondwanaland during the early to middle Paleozoic place
the Lhasa terrane to the north of Greater India and south of the Qiangtang terrane (e.g.,
Dewey et al, 1988; Scotese, 2002). The Qiangtang terrane rifted from Gondwanaland
during Permian time and was later followed by the Lhasa terrane during Triassic time
(Dewey et al., 1988; Sengor and Nata’lin, 1996; Yin and Harrison, 2000). Therefore, if
plate tectonic reconstructions are correct, pre-Permian strata within the Lhasa and
Qiangtang terranes should share detrital zircon age-populations. The distribution of
detrital zircon ages within Carboniferous strata (sample DMXNG) in the Lhasa and
Qiangtang terranes do indeed have some zircon populations with similar ages (Figure
4.7), which is consistent with the two terranes having been in close proximity during
that time. Before these two terranes rifted from Gondwanaland, it is thought that both
were either connected to or near the northern margin of Greater India (Dewey et al.,
1988; Sengor and Nata’lin, 1996). Paleozoic Tethyan strata deposited along the
northern margin of Greater India and now exposed in the Himalaya of Tibet and Nepal
have detrital zircon age populations similar to Carboniferous strata in the Lhasa terrane,
which once again supports recent plate tectonic reconstructions (Scotese, 2002). 157
A Northern and Southern Provenance Signature
For detrital zircon geochronology to be useful in provenance studies, a given
source area must have zircon ages that can be distinguished from those of other potential source terranes. Analysis of zircon age data from this study and from the literature suggests it may be possible to delineate a northern versus southern detrital zircon provenance signature in particular stratigraphic units within the Tibetan plateau.
Published zircon crystallization ages of pre-Cenozoic igneous rocks exposed on the
Tibetan plateau are shown in Figure 4.8. It is apparent in this figure that northern and
southern Tibet contain igneous rocks with different zircon crystallization ages. Northern
Tibet is dominated by igneous rocks with zircon crystallization ages ranging between
200 and 500 Ma, as well as lesser amounts with ages of roughly 900 Ma and 1870 Ma
(Figure 4.8; Gehrels et al., 2003a). Although the ages of detrital zircons from
sedimentary and metasedimentary rocks in northern Tibet vary, they generally have
clusters of ages around ~480 Ma, ~900 Ma, ~1400 Ma and 1800 Ma (Figure 4.9;
Gehrels et al., 2003b). In contrast, zircon crystallization ages of igneous rocks in
southern Tibet are largely <200 Ma (Figure 4.8). Sedimentary and metasedimentary
rocks, including those within the northern Himalaya, tend to have clusters of detrital
zircons with ages of 500-600 Ma and 1000-1400 Ma (Figure 4.9; DeCelles et al., 2004;
this study).
Utilizing the spatial differences in zircon ages on the Tibetan plateau, we can
reexamine the detrital zircon age data of strata from the Lhasa terrane and look for
northern or southern provenance signatures (Figure 4.9). The strata from which the 158
samples PENBO1, PENBO2, NYANT, MAQU, and CHMLNG were collected contain
sedimentary characteristics indicating they were derived from rocks exposed in the
southern part of Tibet (Leier, 2005). Each of the samples contains zircon populations
with ages between 110-160, 500-600 and 1000-1400 Ma, which is consistent with a
southern Tibetan provenance (Figures 4.5, 4.9; see Figure 4.6 for the age spectra of
individual samples of upper Cretaceous strata). Sample LNPLA contains very few
detrital zircons with ages of 500-600 or 1000-1400 Ma. Instead, the majority of the zircons have ages between 200-500 Ma and 1700-2000 Ma, which is more consistent
with a northern Tibetan provenance (Figure 4.9). Little information is available on the provenance of the strata associated with the LNPLA sample; however, paleogeographic reconstructions of the area (Leeder et al., 1988; Yin et al., 1988) and a limited amount
of paleocurrent data from Jurassic strata near the LNPLA sampling site (n=28; Leier,
unpublished data), indicate these strata have a northern provenance and were therefore
derived from rocks exposed in the Qiangtang terrane or areas north thereof. This
suggests that, at least to a first-order, it may be possible to use detrital zircon
geochronology to distinguish the generalized provenance of particular sedimentary units
within the Tibetan plateau. Many factors can complicate the provenance history of a
stratigraphic unit and this hypothesis will need to be tested more thoroughly before it
can be applied with confidence.
Carboniferous Strata as a Sediment Source
Cretaceous strata within the Lhasa terrane contain lithic grains that were derived
from the erosion of shale, phyllite, sandstone, and quartzite successions, indicating that 159 they were derived, at least in part, from pre-Cretaceous sedimentary and metasedimentary rocks (Leier, 2005). Carboniferous strata are the thickest pre-
Cretaceous strata in the Lhasa terrane and are widely exposed at the surface, and in some locations they are unconformably overlain by Cretaceous strata (Figure 4.1;Yin et al., 1988; Liu, 1988). Detrital zircon geochronology provides a means of testing whether Carboniferous strata served as a source rock for Cretaceous sedimentary units.
Figure 4.10 displays the distribution of >400 Ma detrital zircons from all of the samples of the upper Cretaceous Takena Formation. Many of the detrital zircon populations from samples of the Takena Formation are similar to those within Carboniferous strata
(Figure 4.10). Both stratigraphic units contain zircons with ages between 500-600 Ma,
1000-1300 Ma, and ~2500 Ma, which supports the hypothesis that Carboniferous strata within the Lhasa terrane were exposed at the surface and being eroded during
Cretaceous time.
CONCLUSIONS
The scarcity of geologic data from the Tibetan plateau presents a major limitation to understanding the Cenozoic and pre-Cenozoic evolution of this region. The 743 new U-
Pb detrital zircon ages presented here help to elucidate several important aspects of the geology of the Tibetan plateau and its tectono-sedimentary history. Moreover, our results provide encouragement for future detrital zircon studies of strata in southern
Tibet.
Ages of detrital zircons from strata exposed near Nam Co, which are mapped as lower Cretaceous, are all older than 400 Ma. Although detrital zircon data cannot be 160 used to determine depositional ages, this information suggests that the strata exposed along the northeastern shore of Nam Co are Paleozoic in age and not lower Cretaceous.
Further investigation is required to test this hypothesis, but if it turns out to be correct, the crustal structure in this area will need to be reinterpreted.
Strata exposed near the Bangong suture contain zircons indicating they were most likely deposited in the Early Jurassic. The small number of detrital zircons with Jurassic ages is difficult to reconcile with the hypothesized existence of a Jurassic arc along on the southern margin of the Qiangtang terrane, although these strata may predate the construction of such a feature. The Jurassic strata contain populations of detrital zircons with ages between 200-500 and 1800-2000 Ma, which is noticeably different from the ages of detrital zircons within sedimentary strata in the southern portion of the Lhasa terrane. Differences between the age-probability plot of Jurassic strata and the geographic distribution of igneous rocks within the Tibetan plateau suggest it might be possible to discern, at least to a first-order, whether sediments were derived primarily from northern or southern Tibet.
The distribution of detrital zircons within the upper Cretaceous Takena Formation is relatively similar between sampling locations although some differences are apparent.
Detrital zircons in fluvial sandstone exposed northwest of Lhasa (sample MAQU) suggest the fluvial systems responsible for sedimentation in this area were distinct from those that traversed the area northeast of the city of Lhasa. Constraining the timing of deposition of the Takena Formation is important for understanding the Late Cretaceous history of the Lhasa terrane, but is difficult given the nature of the strata. Although 161
detrital zircon ages cannot be used to determine the age of deposition, the ages do
suggest that deposition had probably ceased before 70 Ma.
Carboniferous strata exposed within the Lhasa terrane likely served as a sediment
source for much of the Cretaceous strata deposited in the region. Additionally, similarities between the age probability plots of Carboniferous rocks in the Lhasa terrane with those of rocks in the Qiangtang terrane and with Tethyan strata in the
Himalaya support existing plate tectonic models of Paleozoic Gondwanaland.
The Tibetan plateau is commonly used as an archetype for continent-continent
collisions, yet much of the geology in the region is unstudied. A detailed account of the
tectonic evolution of this area during the time period prior to the Indo-Asian collision is crucial for understanding how and when it achieved its current topography. The new U-
Pb ages presented here help to shed light on some facets of this history, but many questions remain. 162
Figure 4.1: Map of area. Nine samples were collected from sandstone units exposed in the Lhasa terrane of southern Tibet. The samples were primarily from Mesozoic strata, although a Carboniferous unit was sampled also. Geologic map modified from Kidd et al. (1988).
163
Figure 4.2: Regional Setting. A. Simplified diagram depicting the assembly of the terranes of the Tibetan plateau. The Lhasa terrane rifted from Greater India (Gondwanaland) during the Triassic and migrated northward. By Late Jurassic-Early Cretaceous time, the Lhasa terrane began to collide with the Qiangtang terrane. As India moved northward, Neotethyan oceanic crust was subducted northward beneath the Lhasa terrane. Not to scale. Modified from Yin and Harrison, (2000). B. Regional stratigraphy, including stratigraphic position of collected samples. Stratigraphy modified from Yin et al. (1988).
164
Figure 4.3: Takena Formation concordia diagram. Ellipses show 2σ errors. The name of the individual sample and the number of zircons that were sufficiently concordant to be used in the study are noted next to the concordia diagram. See text for further details. 165
Figure 4.4: Concordia diagram of non-Takena Formation samples. Ellipses show 2σ errors. The name of the individual sample and the number of zircons that were of sufficient concordance to be used in the study are noted next to the concordia diagram. See text for further details. 166
Figure 4.5: Distribution of detrital zircon U-Pb ages. All 4 of the samples from the Takena Formation have been combined into one curve. Displayed are the relative probability curve and the histogram of each sample. The number of zircons is given on the vertical axis; the age is on the horizontal axis. Bin size for the histogram is 50 Myr. Also shown for samples of Cretaceous strata are the distribution of zircons with ages <245 (e.g., Takena Fm., MZ); bin size for these Mesozoic populations is 5 Myr. Although sample NAMCO is mapped as lower Cretaceous, it did not contain detrital zircons with ages <400 Ma. See text for discussion. 167
168
Figure 4.6: Takena Formation detrital zircon U-Pb ages. Displayed are the relative probability curve and the histogram of each sample. The number of zircons is given on the vertical axis; the age is on the horizontal axis. Bin size for the histogram is 50 Myr. 169
Figure 4.7: Regional detrital zircon comparison. A comparison of detrital zircon ages from Paleozoic strata of 3 different terranes in the Himalayan-Tibetan orogenic system. Figures are arranged from bottom to top as the terranes are currently located south to north. The lowermost relative age-probability curve is from Tethyan Himalayan strata exposed in northern Nepal and southern Tibet. The middle relative age-probability curve is from Carboniferous strata exposed in the Lhasa terrane; sample DMXNG. The uppermost relative age-probability curve is from Carboniferous metasandstones exposed in the Qiangtang terrane. Plate tectonic reconstructions normally depict the Lhasa and Qiangtang terranes as having been directly north of Greater India during most of Paleozoic time. The similarity between the distributions of detrital zircon ages within each sample supports recent plate tectonic models, although there are some differences between the terranes. 170
Figure 4.8: U-Pb age of igneous rocks. The ages are given in Ma and the corresponding reference is denoted by a number in parentheses. The sampling sites as shown on this map (black dots) are not meant to represent the exact sample site but the approximate sampling location. The shaded regions are added to highlight general trends in the ages. The zircon crystallization ages of igneous rocks in northern Tibet are primarily between 200-500 Ma, those in southern Tibet are generally <200 Ma. The geographic distribution suggests it may be possible to determine the first-order provenance of sedimentary rocks using detrital zircon geochronology. See text for full discussion. 171
Figure 4.9: Northern and southern Tibet. Relative age-probability plots for various lithologic units exposed in the Himalayan-Tibetan orogenic system; the age-probability plots are ordered bottom to top as they are geographically located south to north (i.e., Himalaya at the bottom of the figure, the Altyn Tagh fault at the top of the figure). Clusters that appear to be characteristic of either the northern or southern portion of the region are highlighted. The lighter highlights with questions marks represent age populations that may be present in both the north and the south. The northern portion of the Tibetan plateau contains plutonic rocks that record specific periods of emplacement and cooling. The sedimentary and metasedimentary rocks from the northern portion of the plateau tend to have large zircon populations with ages between ~200-500 Ma and ~1700-2000 Ma. Detrital zircons with these ages are generally absent in Mesozoic sedimentary rocks in the Lhasa terrane, with the exception of the strata associated with the LNPLA sample, which may have been derived from the Qiangtang terrane. Detrital zircon populations in the south have clusters between ~50-200 Ma, 500-600 Ma and 1000-1400 Ma. Less confidence should be placed in the latter two clusters because detrital zircons of these ages may be ubiquitous on the Tibetan plateau. 172
173
Figure 4.10: Carboniferous and Takena Strata. Comparing the distribution of detrital zircon U-Pb ages in samples from the Takena Formation with that from a sample of Carboniferous strata exposed within the Lhasa terrane. Enlarged are the detrital zircons from the Takena Formation that have ages >400 Ma; below that are all of the detrital zircon ages from the sample of Carboniferous strata. The purpose of this figure is to highlight the fact that the older zircons within the Takena Formation have ages that are very similar to the ages of zircons in the Carboniferous strata. This information, combined with petrographic data and regional map-relationships, suggest that Carboniferous strata were exposed at the surface, eroded, and redeposited during the Late Cretaceous. See text for further details.
174
TABLE 4. DR1 - Samples and locations
Latitude Longitude Sample Formation Age Location (N) (E) DMXNG Carbonifreous E of Damxung LNPLA Jurassic? Lunpola Basin 31 51.834 89 36.760 DUBA Early Cretaceous Duba 31 24.459 89 42.541 CHMLN Chumulong Early Cretaceous Penbo area 30 00.267 91 16.899 NAMCO Early Cretaceous? North of Nam Co 30 53.149 91 02.683 MAQU Takena Late Cretaceous Maqu 29 57.357 90 45.105 PENBO1 Takena Late Cretaceous Penbo area 29 54.226 91 20.764 PENBO2 Takena Late Cretaceous Penbo area 29 55.318 91 15.116 NYANT Takena Late Cretaceous South of Nam Co 30 36.133 90 48.311 175
Table 4.DR2 - Detrital zircon data
10%uncertainty 30%disc 6/8 7/5 6/7 BEST Sample age ±(Ma) age ±(Ma) age ±(Ma) AGE ± 5% reverse disc
DUBA Duba Formation - from Duba area DB1252-33 123 2 155 22 687 332 123 2 123 2 DB1252-41 123 4 90 33 -706 1090 123 4 123 4 DB1252-21 123 4 175 35 957 449 123 4 123 4 DB1252-90 124 2 132 14 287 261 124 2 124 2 DB1252-76 124 1 137 9 367 156 124 1 124 1 DB1252-4 124 3 134 22 317 393 124 3 124 3 DB1252-71 125 4 144 82 470 1479 125 4 125 4 DB1252-75 126 4 150 29 558 447 126 4 126 4 DB1252-65 126 3 160 26 701 366 126 3 126 3 DB1252-80 126 2 111 11 -205 250 126 2 126 2 DB1252-84 127 3 161 31 706 452 127 3 127 3 DB1252-9 127 3 138 31 323 551 127 3 127 3 DB1252-31 130 6 128 34 89 680 130 6 130 6 DB1252-47 130 2 136 14 242 248 130 2 130 2 DB1252-86 130 4 116 26 -179 594 130 4 130 4 DB1252-92 131 5 144 22 374 359 131 5 131 5 DB1252-39 131 3 130 53 119 1069 131 3 131 3 DB1252-59 131 2 178 27 861 341 131 2 131 2 DB1252-48 131 2 136 4 209 58 131 2 131 2 DB1252-50 132 7 164 44 657 619 132 7 132 7 DB1252-56 132 4 156 22 534 328 132 4 132 4 DB1252-58 132 2 166 42 681 588 132 2 132 2 DB1252-34 137 2 131 19 31 363 137 2 137 2 DB1252-63 137 2 155 18 441 276 137 2 137 2 DB1252-69 137 1 146 11 294 187 137 1 137 1 DB1252-44 137 2 172 30 688 412 137 2 137 2 DB1252-100 138 3 143 16 227 280 138 3 138 3 DB1252-19 138 2 128 18 -52 362 138 2 138 2 DB1252-30 139 2 143 13 209 232 139 2 139 2 DB1252-57 139 1 134 8 42 159 139 1 139 1 DB1252-67 139 1 146 6 264 104 139 1 139 1 DB1252-83 139 2 140 11 153 203 139 2 139 2 DB1252-77 139 2 138 10 120 182 139 2 139 2 DB1252-23 139 3 134 11 35 212 139 3 139 3 DB1252-16 140 4 146 22 254 370 140 4 140 4 DB1252-94 140 4 189 39 859 466 140 4 140 4 DB1252-78 140 1 142 12 175 202 140 1 140 1 DB1252-29 141 2 146 14 239 239 141 2 141 2 DB1252-37 141 4 143 16 189 266 141 4 141 4 176
DB1252-45 141 2 147 14 253 238 141 2 141 2 DB1252-42 141 3 154 10 350 158 141 3 141 3 DB1252-61 141 1 157 22 394 336 141 1 141 1 DB1252-25 142 3 158 9 414 130 142 3 142 3 DB1252-3 142 2 137 37 59 702 142 2 142 2 DB1252-6 142 2 140 25 114 456 142 2 142 2 DB1252-98 142 3 163 14 474 193 142 3 142 3 DB1252-91 143 2 123 25 -252 549 143 2 143 2 DB1252-52 143 2 139 11 74 205 143 2 143 2 DB1252-99 143 2 166 53 512 783 143 2 143 2 DB1252-7 143 3 137 21 38 391 143 3 143 3 DB1252-22 143 1 136 12 18 222 143 1 143 1 DB1252-73 144 2 142 8 101 136 144 2 144 2 DB1252-88 144 2 160 17 405 251 144 2 144 2 DB1252-24 144 3 154 17 309 273 144 3 144 3 DB1252-26 144 2 159 19 382 290 144 2 144 2 DB1252-20 145 2 141 8 70 132 145 2 145 2 DB1252-62 145 6 200 30 911 326 145 6 145 6 DB1252-74 145 3 150 9 230 137 145 3 145 3 DB1252-75 145 4 207 31 987 334 145 4 145 4 DB1252-54 146 2 153 15 264 246 146 2 146 2 DB1252-28 146 2 147 16 156 275 146 2 146 2 DB1252-15 146 2 144 13 105 236 146 2 146 2 DB1252-14 147 2 171 20 512 285 147 2 147 2 DB1252-60 148 2 144 13 81 230 148 2 148 2 DB1252-79 149 2 149 13 146 224 149 2 149 2 DB1252-43 149 3 159 12 308 180 149 3 149 3 DB1252-8 152 8 167 18 380 231 152 8 152 8 DB1252-64 153 3 167 20 360 291 153 3 153 3 DB1252-68 158 11 171 14 355 132 158 11 158 11 DB1252-11 165 6 156 15 23 226 165 6 165 6 DB1252-36 209 16 210 19 218 150 209 16 209 16 DB1252-2 228 4 275 8 704 63 228 4 228 4 DB1252-13 229 2 227 7 206 78 229 2 229 2 DB1252-53 264 5 331 73 836 544 264 5 264 5 DB1252-66 277 3 292 18 418 157 277 3 277 3 DB1252-96 287 3 309 27 476 220 287 3 287 3 DB1252-10 327 5 330 42 347 334 327 5 327 5 DB1252-11 597 11 668 31 914 126 597 11 597 11 DB1252-46 604 6 571 29 443 147 604 6 604 6 DB1252-38 619 12 623 18 638 69 619 12 619 12 DB1252-27 631 6 627 18 612 80 631 6 631 6 DB1252-40 722 8 702 27 638 111 722 8 722 8 DB1252-93 768 11 770 22 775 79 768 11 768 11 DB1252-55 822 21 845 18 907 31 822 21 822 21 DB1252-87 844 13 883 11 980 21 844 13 844 13 DB1252-72 860 36 861 30 864 50 860 36 860 36 DB1252-95 1051 42 1045 29 1031 22 1031 22 1031 22 177
DB1252-5 1105 14 1091 21 1062 58 1062 58 1062 58 DB1252-51 1075 10 1071 29 1062 87 1062 87 1062 87 DB1252-89 1158 11 1171 10 1196 20 1196 20 1196 20 DB1252-85 1115 86 1243 68 1474 81 1474 81 1474 81 DB1252-97 1854 16 1822 38 1785 80 1785 80 1785 80 DB1252-32 2301 29 2325 17 2346 19 2346 19 2346 19 DB1252-1 2432 20 2438 13 2443 17 2443 17 2443 17 DB1252-12 2938 40 2878 19 2836 16 2836 16 2836 16
CHMLN J-K strata from the Penbo area PB2452-8 138 2 157 27 448 418 138 2 138 2 PB2452-90 141 3 143 10 174 159 141 3 141 3 PB2452-53 141 2 170 19 588 267 141 2 141 2 PB2452-73 142 2 151 24 295 392 142 2 142 2 PB2452-89 142 2 156 19 379 296 142 2 142 2 PB2452-100 142 3 177 18 675 231 142 3 142 3 PB2452-71 142 4 140 30 94 552 142 4 142 4 PB2452-91 143 2 144 18 169 312 143 2 143 2 PB2452-38 144 2 144 10 149 167 144 2 144 2 PB2452-78 144 7 188 46 780 560 144 7 144 7 PB2452-65 144 6 146 72 173 1311 144 6 144 6 PB2452-4 145 3 138 19 15 349 145 3 145 3 PB2452-45 145 3 153 16 284 257 145 3 145 3 PB2452-33 145 2 180 22 666 286 145 2 145 2 PB2452-35 145 5 162 27 414 400 145 5 145 5 PB2452-36 145 2 158 16 358 244 145 2 145 2 PB2452-86 147 5 108 28 -687 758 147 5 147 5 PB2452-37 148 3 140 17 2 305 148 3 148 3 PB2452-13 149 3 146 6 108 100 149 3 149 3 PB2452-60 149 3 162 15 354 227 149 3 149 3 PB2452-85 150 4 175 31 533 427 150 4 150 4 PB2452-3 151 2 146 17 70 289 151 2 151 2 PB2452-97 152 8 187 48 651 605 152 8 152 8 PB2452-39 157 2 190 41 615 519 157 2 157 2 PB2452-12 165 5 185 16 448 194 165 5 165 5 PB2452-87 213 2 223 43 334 495 213 2 213 2 PB2452-25 226 4 186 24 -301 363 226 4 226 4 PB2452-77 285 3 291 13 343 109 285 3 285 3 PB2452-51 412 9 465 70 735 399 412 9 412 9 PB2452-21 505 7 479 24 357 141 505 7 505 7 PB2452-22 512 14 558 108 749 533 512 14 512 14 PB2452-70 518 18 482 26 318 129 518 18 518 18 PB2452-52 519 7 555 48 701 240 519 7 519 7 PB2452-14 521 5 519 8 511 36 521 5 521 5 PB2452-46 523 10 508 15 441 74 523 10 523 10 PB2452-88 524 6 529 18 549 93 524 6 524 6 PB2452-64 526 6 537 15 587 72 526 6 526 6 PB2452-18 528 5 515 11 454 56 528 5 528 5 178
PB2452-11 536 6 526 27 480 142 536 6 536 6 PB2452-29 538 11 515 28 413 149 538 11 538 11 PB2452-47 543 8 544 31 550 156 543 8 543 8 PB2452-30 554 8 561 16 593 74 554 8 554 8 PB2452-81 556 8 556 9 556 35 556 8 556 8 PB2452-63 564 8 563 15 560 68 564 8 564 8 PB2452-84 570 10 554 66 489 342 570 10 570 10 PB2452-26 627 6 632 15 650 64 627 6 627 6 PB2452-10 631 9 628 16 615 67 631 9 631 9 PB2452-34 656 11 674 22 734 85 656 11 656 11 PB2452-96 665 8 664 9 659 29 665 8 665 8 PB2452-27 676 10 689 30 732 124 676 10 676 10 PB2452-54 678 10 714 40 831 161 678 10 678 10 PB2452-58 697 8 700 19 712 75 697 8 697 8 PB2452-28 828 33 852 26 917 29 828 33 828 33 PB2452-93 856 11 870 16 904 48 856 11 856 11 PB2452-31 877 16 930 52 1059 169 877 16 877 16 PB2452-75 880 8 896 10 935 27 880 8 880 8 PB2452-74 889 9 882 18 865 61 889 9 889 9 PB2452-50 893 11 910 20 954 64 893 11 893 11 PB2452-61 902 12 861 29 756 102 902 12 902 12 PB2452-49 902 26 905 25 914 58 902 26 902 26 PB2452-6 904 9 961 9 1096 20 904 9 904 9 PB2452-15 907 10 920 9 952 20 907 10 907 10 PB2452-98 939 9 947 9 966 20 939 9 939 9 PB2452-56 955 10 956 9 956 21 955 10 955 10 PB2452-7 981 11 987 14 1000 37 981 11 981 11 PB2452-94 1032 17 1040 16 1058 34 1058 34 1058 34 PB2452-20 1110 10 1098 22 1075 64 1075 64 1075 64 PB2452-19 930 23 974 35 1075 99 1075 99 1075 23 PB2452-1 922 37 970 30 1081 42 1081 42 1081 37 PB2452-48 1032 19 1049 19 1085 44 1085 44 1085 44 PB2452-43 1077 29 1094 22 1128 29 1128 29 1128 29 PB2452-80 1174 17 1163 25 1144 64 1144 64 1144 64 PB2452-67 1161 21 1161 20 1160 42 1160 42 1160 42 PB2452-17 1109 11 1129 34 1169 97 1169 97 1169 97 PB2452-95 1161 11 1174 13 1196 30 1196 30 1196 30 PB2452-42 1137 27 1160 35 1205 85 1205 85 1205 85 PB2452-59 1178 15 1194 18 1224 41 1224 41 1224 41 PB2452-99 1167 11 1198 10 1255 20 1255 20 1255 20 PB2452-44 1238 15 1251 24 1276 59 1276 59 1276 59 PB2452-79 1283 12 1314 11 1363 20 1363 20 1363 20 PB2452-62 1457 22 1471 60 1491 144 1491 144 1491 144 PB2452-55 1638 35 1648 30 1661 52 1661 52 1661 52 PB2452-23 1705 25 1743 16 1789 18 1789 18 1789 18 PB2452-72 2376 38 2423 20 2462 17 2462 17 2462 17 PB2452-41 2218 30 2364 17 2492 17 2492 17 2492 17 PB2452-83 2650 22 2652 13 2653 17 2653 17 2653 17 179
PB2452-16 2610 21 2644 13 2670 17 2670 17 2670 17 PB2452-76 3114 67 3248 28 3332 16 3332 16 3332 16 PB2452-5 3343 65 3353 26 3358 16 3358 16 3358 16
LNPLA Jurassic? Strata from the Lunpola area - it may be Triassic or older LP234-95 179 3 181 16 206 222 179 3 179 3 LP234-31 181 6 228 65 741 688 181 6 181 6 LP234-37 204 3 217 42 356 485 204 3 204 3 LP234-5 209 5 194 13 22 165 209 5 209 5 LP234-61 218 2 216 15 186 174 218 2 218 2 LP234-6 225 5 180 12 -370 184 225 5 225 5 LP234-70 228 5 256 20 517 187 228 5 228 5 LP234-67 237 5 273 45 597 409 237 5 237 5 LP234-33 258 7 254 44 214 448 258 7 258 7 LP234-30 266 4 237 19 -44 211 266 4 266 4 LP234-90 270 7 271 10 287 71 270 7 270 7 LP234-98 274 9 283 18 356 147 274 9 274 9 LP234-10 277 7 268 19 185 177 277 7 277 7 LP234-3 287 10 283 11 253 61 287 10 287 10 LP234-96 329 7 295 26 36 240 329 7 329 7 LP234-53 330 10 328 12 316 64 330 10 330 10 LP234-62 352 6 328 33 163 272 352 6 352 6 LP234-25 378 4 385 11 426 74 378 4 378 4 LP234-73 397 9 456 143 766 842 397 9 397 9 LP234-49 422 5 412 15 354 98 422 5 422 5 LP234-12 425 7 408 13 313 83 425 7 425 7 LP234-76 443 5 433 6 384 27 443 5 443 5 LP234-1 451 10 454 15 469 76 451 10 451 10 LP234-82 458 6 442 15 358 92 458 6 458 6 LP234-42 460 8 456 8 438 34 460 8 460 8 LP234-84 468 8 487 27 581 145 468 8 468 8 LP234-7 477 7 450 13 317 76 477 7 477 7 LP234-91 482 6 483 11 488 58 482 6 482 6 LP234-17 497 6 508 7 558 27 497 6 497 6 LP234-52 522 7 522 16 520 81 522 7 522 7 LP234-57 538 5 540 34 547 178 538 5 538 5 LP234-51 559 11 547 15 499 65 559 11 559 11 LP234-23 669 12 723 37 895 141 669 12 669 12 LP234-80 721 11 744 18 814 63 721 11 721 11 LP234-69 731 11 728 16 720 54 731 11 731 11 LP234-86 759 7 761 14 765 51 759 7 759 7 LP234-44 802 8 783 38 728 146 802 8 802 8 LP234-81 815 11 776 54 665 209 815 11 815 11 LP234-64 850 21 839 23 810 64 850 21 850 21 LP234-46 861 15 862 12 865 22 861 15 861 15 LP234-92 897 13 921 12 979 24 897 13 897 13 LP234-34 950 9 914 28 828 93 950 9 950 9 LP234-87 1020 16 990 16 923 36 1020 16 1020 36 180
LP234-14 1097 10 1094 13 1089 34 1089 34 1089 34 LP234-47 1179 21 1183 24 1190 57 1190 57 1190 57 LP234-89 1159 36 1171 37 1191 83 1191 83 1191 83 LP234-21 1369 12 1349 17 1317 40 1317 40 1317 40 LP234-20 1431 23 1449 16 1476 19 1476 19 1476 19 LP234-94 1591 19 1602 14 1616 19 1616 19 1616 19 LP234-41 1448 25 1538 17 1664 19 1664 19 1664 19 LP234-35 1648 21 1664 29 1683 61 1683 61 1683 61 LP234-24 1749 15 1725 34 1695 73 1695 73 1695 73 LP234-29 1700 39 1700 26 1700 31 1700 31 1700 31 LP234-78 1237 20 1424 38 1715 84 1715 84 1715 84 LP234-68 1386 16 1530 13 1736 20 1736 20 1736 20 LP234-83 1854 26 1849 19 1842 28 1842 28 1842 28 LP234-66 1893 49 1874 30 1852 33 1852 33 1852 33 LP234-56 1858 16 1859 23 1859 44 1859 44 1859 44 LP234-32 1647 31 1746 20 1866 18 1866 18 1866 18 LP234-9 1793 37 1830 22 1873 18 1873 18 1873 18 LP234-11 1840 23 1858 18 1877 27 1877 27 1877 27 LP234-38 1491 61 1661 39 1883 18 1883 18 1883 18 LP234-88 1705 46 1787 27 1883 18 1883 18 1883 18 LP234-79 1956 31 1923 18 1889 18 1889 18 1889 18 LP234-48 1875 26 1882 16 1890 18 1890 18 1890 18 LP234-27 1777 26 1831 17 1894 18 1894 18 1894 18 LP234-26 1896 47 1898 26 1901 19 1901 19 1901 19 LP234-72 1899 16 1902 12 1905 18 1905 18 1905 18 LP234-65 1828 16 1864 12 1905 18 1905 18 1905 18 LP234-43 1782 25 1840 16 1907 18 1907 18 1907 18 LP234-85 1933 38 1923 21 1912 18 1912 18 1912 18 LP234-36 1945 32 1929 19 1913 18 1913 18 1913 18 LP234-60 1761 99 1832 55 1914 18 1914 18 1914 18 LP234-8 1854 29 1883 18 1914 18 1914 18 1914 18 LP234-75 1829 49 1874 28 1925 18 1925 18 1925 18 LP234-39 1906 27 1920 17 1935 18 1935 18 1935 18 LP234-59 1927 26 1949 16 1973 18 1973 18 1973 18 LP234-54 1983 36 1982 20 1980 18 1980 18 1980 18 LP234-71 1831 45 1910 26 1996 18 1996 18 1996 18 LP234-19 1980 22 1991 14 2003 18 2003 18 2003 18 LP234-28 2007 44 2006 24 2005 18 2005 18 2005 18 LP234-22 1911 26 1964 23 2021 37 2021 37 2021 37 LP234-13 1974 17 2002 12 2030 18 2030 18 2030 18 LP234-40 2047 71 2104 37 2161 17 2161 17 2161 17 LP234-58 2195 48 2187 25 2180 17 2180 17 2180 17 LP234-93 2036 19 2114 13 2191 18 2191 18 2191 18 LP234-16 1971 17 2100 13 2229 17 2229 17 2229 17 LP234-100 2286 19 2268 13 2252 17 2252 17 2252 17 LP234-18 2199 66 2231 34 2261 23 2261 23 2261 23 LP234-50 2407 64 2429 31 2447 17 2447 17 2447 17 LP234-97 2052 69 2260 39 2453 29 2453 29 2453 29 181
LP234-2 2046 43 2259 24 2458 17 2458 17 2458 17 LP234-55 2465 23 2499 14 2527 17 2527 17 2527 17 LP234-63 2661 26 2673 29 2683 47 2683 47 2683 47 LP234-77 2979 24 3087 14 3158 16 3158 16 3158 16
NAMCO Supposed to be lower Cretaceous strata - northeast shore of Nam Co - probably Paleozoic NM15-57 422 20 459 24 645 94 422 20 422 20 NM15-87 486 11 534 23 745 105 486 11 486 11 NM15-94 494 10 525 23 663 110 494 10 494 10 NM15-64 495 24 544 158 755 805 495 24 495 24 NM15-22 497 5 534 30 695 152 497 5 497 5 NM15-6 503 15 533 25 664 110 503 15 503 15 NM15-44 511 14 496 38 428 206 511 14 511 14 NM15-56 518 13 570 43 782 201 518 13 518 13 NM15-80 524 7 542 9 619 33 524 7 524 7 NM15-20 529 14 489 90 310 527 529 14 529 14 NM15-54 534 10 524 22 483 110 534 10 534 10 NM15-60 548 9 482 71 180 431 548 9 548 9 NM15-33 568 8 582 46 636 222 568 8 568 8 NM15-89 576 12 586 18 622 77 576 12 576 12 NM15-74 600 9 596 14 582 56 600 9 600 9 NM15-8 602 11 622 14 697 45 602 11 602 11 NM15-24 660 10 692 11 795 33 660 10 660 10 NM15-5 685 11 682 46 671 193 685 11 685 11 NM15-15 751 16 797 19 928 54 751 16 751 16 NM15-55 756 12 788 85 883 319 756 12 756 12 NM15-69 760 7 781 8 840 23 760 7 760 7 NM15-84 762 13 804 46 921 166 762 13 762 13 NM15-23 763 7 824 35 993 125 763 7 763 7 NM15-12 766 11 811 22 936 74 766 11 766 11 NM15-61 778 9 806 18 886 60 778 9 778 9 NM15-3 794 7 779 29 737 109 794 7 794 7 NM15-65 803 14 897 50 1136 163 803 14 803 14 NM15-85 832 23 847 52 888 178 832 23 832 23 NM15-70 860 18 873 43 907 145 860 18 860 18 NM15-26 867 14 897 22 973 69 867 14 867 14 NM15-92 868 28 906 24 997 43 868 28 868 28 NM15-2 870 9 882 21 913 68 870 9 870 9 NM15-4 882 8 890 30 908 103 882 8 882 8 NM15-40 890 8 911 10 962 25 890 8 890 8 NM15-27 891 18 949 22 1088 54 891 18 891 18 NM15-46 896 9 874 30 818 104 896 9 896 9 NM15-62 928 56 384 228 -2323 488 928 56 928 56 NM15-53 949 23 980 19 1050 29 949 23 949 23 NM15-34 950 21 965 18 999 31 950 21 950 21 NM15-88 966 15 966 24 966 71 966 15 966 15 NM15-58 966 9 950 23 914 75 966 9 966 9 NM15-38 974 9 988 15 1020 45 974 9 974 9 182
NM15-18 984 9 976 40 958 129 984 9 984 9 NM15-63 985 14 985 14 985 33 985 14 985 14 NM15-43 988 9 991 9 998 23 988 9 988 9 NM15-31 1003 19 997 22 984 57 1003 19 1003 57 NM15-36 1012 20 1018 17 1030 31 1030 31 1030 31 NM15-75 968 13 993 16 1048 40 1048 40 1048 13 NM15-73 1080 14 1069 18 1048 47 1048 47 1048 47 NM15-90 1056 10 1067 12 1091 30 1091 30 1091 30 NM15-11 936 30 992 34 1118 83 1118 83 1118 30 NM15-10 1078 10 1092 11 1120 26 1120 26 1120 26 NM15-9 1064 10 1088 21 1136 60 1136 60 1136 60 NM15-49 1141 37 1142 28 1145 44 1145 44 1145 44 NM15-32 1015 9 1061 15 1156 40 1156 40 1156 40 NM15-35 1055 10 1102 12 1195 30 1195 30 1195 30 NM15-13 1142 38 1167 29 1214 42 1214 42 1214 42 NM15-48 1179 11 1196 21 1226 56 1226 56 1226 56 NM15-45 1279 12 1265 24 1241 63 1241 63 1241 63 NM15-39 1200 11 1246 40 1325 106 1325 106 1325 106 NM15-68 1355 28 1343 21 1325 29 1325 29 1325 29 NM15-100 1184 97 1274 68 1431 45 1431 45 1431 45 NM15-50 1449 13 1445 15 1439 33 1439 33 1439 33 NM15-59 1422 19 1435 17 1453 30 1453 30 1453 30 NM15-17 1519 27 1503 23 1480 42 1480 42 1480 42 NM15-77 1495 71 1504 46 1516 45 1516 45 1516 45 NM15-21 1661 37 1684 22 1712 19 1712 19 1712 19 NM15-52 1666 23 1699 18 1740 26 1740 26 1740 26 NM15-91 1786 33 1771 23 1754 31 1754 31 1754 31 NM15-42 1647 15 1698 12 1762 19 1762 19 1762 19 NM15-79 1704 46 1732 27 1767 19 1767 19 1767 19 NM15-7 1628 23 1691 16 1771 18 1771 18 1771 18 NM15-97 1685 18 1733 20 1791 38 1791 38 1791 38 NM15-14 1411 27 1606 22 1871 31 1871 31 1871 31 NM15-72 2389 32 2436 17 2476 17 2476 17 2476 17 NM15-51 2258 19 2403 13 2528 17 2528 17 2528 17 NM15-83 2344 26 2445 15 2531 17 2531 17 2531 17 NM15-28 2367 38 2475 20 2565 17 2565 17 2565 17 NM15-71 2135 49 2379 32 2596 37 2596 37 2596 37 NM15-30 2158 81 2427 42 2660 22 2660 22 2660 22 NM15-98 2571 53 2649 26 2710 18 2710 18 2710 18 NM15-25 2517 71 2631 33 2720 17 2720 17 2720 17 NM15-16 2069 104 2466 59 2812 39 2812 39 2812 39 NM15-78 2853 65 2863 28 2869 16 2869 16 2869 16 NM15-95 2976 33 2984 17 2990 16 2990 16 2990 16 NM15-81 2994 125 3238 60 3393 50 3393 50 3393 50
DMXNG Carboniferous - east of Damxung 36 454 4 492 6 669 24 454 4 454 4 19 501 6 251 33 -1738 525 501 6 501 6 183
35 519 5 528 6 567 22 519 5 519 5 89 526 10 554 97 669 490 526 10 526 10 98 529 5 498 22 362 126 529 5 529 5 38 535 5 491 116 289 692 535 5 535 5 49 543 7 578 52 719 250 543 7 543 7 100 544 5 550 31 574 157 544 5 544 5 4 546 5 540 16 516 82 546 5 546 5 45 549 5 581 24 709 113 549 5 549 5 41 552 7 552 30 549 150 552 7 552 7 70 555 5 555 50 555 256 555 5 555 5 2 564 5 613 23 801 102 564 5 564 5 44 584 6 549 18 406 95 584 6 584 6 93 587 9 533 53 307 288 587 9 587 9 53 628 13 553 51 257 268 628 13 628 13 21 645 6 649 12 664 47 645 6 645 6 37 650 6 597 26 404 125 650 6 650 6 63 651 6 708 14 895 53 651 6 651 6 55 666 6 597 21 341 103 666 6 666 6 31 689 7 688 10 682 35 689 7 689 7 10 793 7 853 9 1013 22 793 7 793 7 51 955 9 942 17 912 54 955 9 955 9 8 956 9 942 14 909 44 956 9 956 9 18 978 9 990 14 1015 41 978 9 978 9 25 979 9 983 20 993 61 979 9 979 9 7 1053 13 1032 19 987 52 987 52 987 52 75 1011 10 1012 30 1014 92 1014 92 1014 92 42 1089 10 1075 22 1047 64 1047 64 1047 64 23 897 8 965 11 1124 29 1124 29 1124 8 60 1012 9 1054 12 1142 31 1142 31 1142 31 90 1197 11 1180 25 1147 70 1147 70 1147 70 30 1130 10 1147 18 1180 49 1180 49 1180 49 11 1200 11 1196 10 1188 20 1188 20 1188 20 16 1188 19 982 114 544 394 1188 19 1188 394 33 1185 11 1188 26 1192 72 1192 72 1192 72 22 1126 12 1151 16 1198 40 1198 40 1198 40 54 1103 17 1140 33 1213 89 1213 89 1213 89 1 1121 10 1160 11 1233 24 1233 24 1233 24 73 1191 11 1223 10 1282 20 1282 20 1282 20 99 1203 14 1251 31 1335 79 1335 79 1335 79 83 1279 16 1314 38 1371 94 1371 94 1371 94 20 1336 12 1383 11 1456 19 1456 19 1456 19 87 1104 13 1238 19 1480 43 1480 43 1480 43 9 1567 15 1560 16 1551 31 1551 31 1551 31 59 1441 13 1490 20 1561 44 1561 44 1561 44 95 1439 13 1500 12 1587 20 1587 20 1587 20 52 1489 13 1530 12 1588 22 1588 22 1588 22 5 1404 13 1480 11 1590 19 1590 19 1590 19 69 1350 12 1452 11 1605 19 1605 19 1605 19 184
6 1379 12 1473 11 1610 19 1610 19 1610 19 85 1380 12 1475 11 1615 19 1615 19 1615 19 15 1447 13 1524 11 1634 19 1634 19 1634 19 67 1638 32 1662 24 1691 34 1691 34 1691 34 57 1489 13 1584 17 1713 33 1713 33 1713 33 82 1257 37 1462 51 1773 106 1773 106 1773 106 17 1674 15 1739 12 1817 19 1817 19 1817 19 72 1634 14 1721 13 1827 22 1827 22 1827 22 65 1793 16 1809 31 1827 64 1827 64 1827 64 29 1763 15 1798 15 1838 27 1838 27 1838 27 78 1639 14 1735 12 1852 18 1852 18 1852 18 58 1755 15 1800 13 1852 22 1852 22 1852 22 76 1615 14 1722 12 1854 19 1854 19 1854 19 97 1793 16 1822 14 1855 24 1855 24 1855 24 68 1782 16 1818 12 1859 18 1859 18 1859 18 24 1715 15 1792 12 1882 18 1882 18 1882 18 48 1478 13 1664 12 1907 18 1907 18 1907 18 92 1419 13 1641 12 1939 18 1939 18 1939 18 84 1394 18 1626 14 1941 18 1941 18 1941 18 66 1868 16 1905 17 1946 30 1946 30 1946 30 47 1941 17 1945 12 1949 18 1949 18 1949 18 26 1466 22 1690 16 1980 18 1980 18 1980 18 43 1593 14 1775 12 1996 18 1996 18 1996 18 28 1578 20 1793 15 2054 18 2054 18 2054 18 40 2116 22 2188 16 2255 22 2255 22 2255 22 74 2135 18 2203 13 2267 17 2267 17 2267 17 61 2132 45 2362 25 2567 17 2567 17 2567 17 32 2253 21 2429 14 2580 17 2580 17 2580 17 50 2265 20 2446 13 2600 17 2600 17 2600 17 62 2454 21 2592 13 2703 17 2703 17 2703 17 79 2738 22 2761 13 2777 16 2777 16 2777 16 71 2651 22 2750 13 2823 16 2823 16 2823 16 39 2703 22 2790 13 2853 16 2853 16 2853 16 81 2943 50 3100 28 3203 31 3203 31 3203 31 56 2968 24 3134 15 3242 18 3242 18 3242 18
MAQU Takena Formation from the Maqu area MQ2167-78 120 3 131 29 340 541 120 3 120 3 MQ2167-23 122 2 121 17 103 360 122 2 122 2 MQ2167-31 137 2 156 16 461 250 137 2 137 2 MQ2167-75 138 3 172 27 680 359 138 3 138 3 MQ2167-52 139 2 133 16 19 309 139 2 139 2 MQ2167-82 142 2 166 17 527 238 142 2 142 2 MQ2167-17 143 2 134 25 -17 483 143 2 143 2 MQ2167-57 144 3 139 12 46 215 144 3 144 3 MQ2167-10 162 6 165 8 207 82 162 6 162 6 MQ2167-18 189 6 142 45 -584 933 189 6 189 6 MQ2167-9 192 10 203 55 326 685 192 10 192 10 185
MQ2167-5 197 9 213 30 391 340 197 9 197 9 MQ2167-39 239 4 240 32 253 338 239 4 239 4 MQ2167-64 391 5 372 22 255 160 391 5 391 5 MQ2167-14 506 5 506 22 510 119 506 5 506 5 MQ2167-13 514 5 521 6 549 24 514 5 514 5 MQ2167-70 528 7 558 50 681 249 528 7 528 7 MQ2167-41 532 6 526 25 500 130 532 6 532 6 MQ2167-79 534 8 541 10 570 39 534 8 534 8 MQ2167-49 540 9 549 24 587 116 540 9 540 9 MQ2167-32 547 5 562 56 623 280 547 5 547 5 MQ2167-66 549 6 556 6 586 22 549 6 549 6 MQ2167-42 556 6 534 40 443 212 556 6 556 6 MQ2167-6 559 5 550 22 516 110 559 5 559 5 MQ2167-91 559 6 561 8 567 33 559 6 559 6 MQ2167-83 561 5 553 9 520 39 561 5 561 5 MQ2167-33 564 6 545 37 468 191 564 6 564 6 MQ2167-40 565 5 540 32 432 169 565 5 565 5 MQ2167-16 576 7 587 7 632 22 576 7 576 7 MQ2167-53 589 22 589 113 587 547 589 22 589 22 MQ2167-4 590 7 579 24 539 114 590 7 590 7 MQ2167-55 595 10 588 20 560 90 595 10 595 10 MQ2167-54 602 6 590 10 541 46 602 6 602 6 MQ2167-51 606 8 609 21 620 96 606 8 606 8 MQ2167-67 615 9 618 10 627 28 615 9 615 9 MQ2167-56 625 8 637 12 682 43 625 8 625 8 MQ2167-37 641 10 603 78 460 382 641 10 641 10 MQ2167-84 644 6 640 22 626 97 644 6 644 6 MQ2167-2 669 17 674 32 690 127 669 17 669 17 MQ2167-25 697 11 688 27 657 109 697 11 697 11 MQ2167-71 709 12 679 48 579 207 709 12 709 12 MQ2167-19 729 8 720 19 690 76 729 8 729 8 MQ2167-27 739 16 771 43 863 158 739 16 739 16 MQ2167-8 762 7 784 9 846 25 762 7 762 7 MQ2167-46 788 10 802 62 843 233 788 10 788 10 MQ2167-95 800 8 773 18 698 67 800 8 800 8 MQ2167-44 805 8 813 10 835 30 805 8 805 8 MQ2167-74 850 8 867 28 909 96 850 8 850 8 MQ2167-35 870 8 892 8 947 21 870 8 870 8 MQ2167-47 873 11 877 14 886 41 873 11 873 11 MQ2167-76 882 13 927 18 1038 51 882 13 882 13 MQ2167-81 885 8 894 22 918 74 885 8 885 8 MQ2167-92 898 15 923 63 982 210 898 15 898 15 MQ2167-59 900 13 917 20 958 60 900 13 900 13 MQ2167-90 904 8 922 9 963 20 904 8 904 8 MQ2167-63 927 11 925 28 920 92 927 11 927 11 MQ2167-45 953 13 970 12 1008 24 953 13 953 13 MQ2167-29 978 18 982 14 990 20 978 18 978 18 MQ2167-80 979 9 987 13 1006 37 979 9 979 9 186
MQ2167-89 1008 9 994 47 964 149 1008 9 1008 149 MQ2167-73 994 20 1008 16 1039 22 1039 22 1039 22 MQ2167-15 1085 12 1078 22 1063 63 1063 63 1063 63 MQ2167-77 1090 21 1097 16 1111 21 1111 21 1111 21 MQ2167-43 1082 11 1105 10 1150 22 1150 22 1150 22 MQ2167-87 1117 25 1128 18 1150 23 1150 23 1150 23 MQ2167-1 1138 15 1142 12 1150 20 1150 20 1150 20 MQ2167-69 1161 15 1161 28 1161 76 1161 76 1161 76 MQ2167-11 1157 11 1167 12 1186 28 1186 28 1186 28 MQ2167-30 1177 13 1182 30 1192 82 1192 82 1192 82 MQ2167-20 1176 19 1194 16 1226 30 1226 30 1226 30 MQ2167-72 1160 11 1183 22 1227 57 1227 57 1227 57 MQ2167-38 1198 11 1232 10 1290 20 1290 20 1290 20 MQ2167-86 1492 13 1511 11 1538 19 1538 19 1538 19 MQ2167-12 1573 14 1583 16 1597 32 1597 32 1597 32 MQ2167-22 1674 21 1711 14 1756 18 1756 18 1756 18 MQ2167-58 1839 32 1827 19 1813 18 1813 18 1813 18 MQ2167-48 1838 24 1834 15 1828 18 1828 18 1828 18 MQ2167-21 1759 17 1802 14 1852 22 1852 22 1852 22 MQ2167-24 1889 16 1899 12 1910 18 1910 18 1910 18 MQ2167-50 1925 30 1999 18 2077 18 2077 18 2077 18 MQ2167-93 2262 48 2322 25 2376 17 2376 17 2376 17 MQ2167-3 2528 29 2514 16 2502 17 2502 17 2502 17 MQ2167-61 2500 46 2502 22 2503 17 2503 17 2503 17 MQ2167-34 2348 22 2495 14 2617 17 2617 17 2617 17 MQ2167-68 2885 44 2934 21 2968 16 2968 16 2968 16 MQ2167-96 3311 26 3404 14 3459 15 3459 15 3459 15
PENBO2 Takena Formation from the Penbo area PB1275-14 100 2 97 5 13 133 100 2 100 2 PB1275-61 103 4 127 40 599 734 103 4 103 4 PB1275-24 104 3 103 6 79 125 104 3 104 3 PB1275-13 104 4 128 39 604 702 104 4 104 4 PB1275-71 105 2 134 37 685 634 105 2 105 2 PB1275-10 106 3 138 19 731 316 106 3 106 3 PB1275-32 106 3 124 20 478 375 106 3 106 3 PB1275-48 107 2 122 38 433 747 107 2 107 2 PB1275-42 108 5 104 51 17 1308 108 5 108 5 PB1275-7 110 3 157 34 950 474 110 3 110 3 PB1275-11 112 1 120 11 280 219 112 1 112 1 PB1275-25 112 3 163 42 994 571 112 3 112 3 PB1275-36 113 3 91 54 -461 1816 113 3 113 3 PB1275-6 114 3 113 25 98 552 114 3 114 3 PB1275-60 116 3 110 26 -7 615 116 3 116 3 PB1275-94 117 1 118 6 129 114 117 1 117 1 PB1275-96 120 5 123 23 184 463 120 5 120 5 PB1275-40 120 4 155 31 721 460 120 4 120 4 PB1275-35 123 2 147 23 548 374 123 2 123 2 187
PB1275-41 124 4 153 94 629 1622 124 4 124 4 PB1275-72 127 4 133 48 243 911 127 4 127 4 PB1275-80 131 2 155 27 538 407 131 2 131 2 PB1275-87 132 7 163 68 646 1003 132 7 132 7 PB1275-100 132 2 160 15 599 210 132 2 132 2 PB1275-29 134 3 164 20 623 284 134 3 134 3 PB1275-9 134 4 110 28 -396 707 134 4 134 4 PB1275-12 136 5 107 64 -487 1830 136 5 136 5 PB1275-17 136 2 174 23 726 307 136 2 136 2 PB1275-74 140 3 153 19 369 301 140 3 140 3 PB1275-91 141 3 172 18 624 239 141 3 141 3 PB1275-50 154 2 160 8 254 117 154 2 154 2 PB1275-66 154 4 145 40 -12 733 154 4 154 4 PB1275-31 158 3 162 10 232 150 158 3 158 3 PB1275-93 175 2 184 9 299 122 175 2 175 2 PB1275-88 183 5 179 15 119 213 183 5 183 5 PB1275-4 184 3 221 38 645 412 184 3 184 3 PB1275-5 184 8 162 48 -148 801 184 8 184 8 PB1275-49 186 4 184 38 166 525 186 4 186 4 PB1275-46 188 8 203 87 382 1111 188 8 188 8 PB1275-81 190 7 208 40 412 475 190 7 190 7 PB1275-55 191 3 244 29 792 283 191 3 191 3 PB1275-69 195 8 193 56 172 759 195 8 195 8 PB1275-89 200 5 213 33 363 383 200 5 200 5 PB1275-58 204 6 256 38 759 353 204 6 204 6 PB1275-78 208 2 218 13 333 150 208 2 208 2 PB1275-68 216 7 211 12 160 118 216 7 216 7 PB1275-33 241 5 253 32 374 315 241 5 241 5 PB1275-44 309 3 352 33 647 241 309 3 309 3 PB1275-57 363 6 378 35 476 246 363 6 363 6 PB1275-27 364 5 423 39 764 237 364 5 364 5 PB1275-2 416 8 434 24 532 139 416 8 416 8 PB1275-59 536 8 534 20 523 101 536 8 536 8 PB1275-65 557 9 556 16 552 74 557 9 557 9 PB1275-18 560 6 548 20 494 102 560 6 560 6 PB1275-76 567 8 579 24 629 116 567 8 567 8 PB1275-22 592 8 591 9 589 31 592 8 592 8 PB1275-21 592 6 597 44 615 210 592 6 592 6 PB1275-39 646 9 647 10 649 33 646 9 646 9 PB1275-73 688 7 738 85 893 337 688 7 688 7 PB1275-15 700 13 700 21 699 75 700 13 700 13 PB1275-54 750 7 777 16 853 57 750 7 750 7 PB1275-20 752 7 760 21 783 80 752 7 752 7 PB1275-86 798 12 838 11 947 21 798 12 798 12 PB1275-26 854 17 896 27 1002 80 854 17 854 17 PB1275-34 937 12 948 14 973 36 937 12 937 12 PB1275-85 1051 17 1047 21 1039 55 1039 55 1039 55 PB1275-67 1055 22 1015 50 930 153 1055 22 1055 153 188
PB1275-19 1044 10 1065 9 1108 20 1108 20 1108 20 PB1275-23 1009 32 1057 23 1156 20 1156 20 1156 20 PB1275-47 1122 13 1145 22 1188 59 1188 59 1188 59 PB1275-30 1160 23 1171 18 1191 28 1191 28 1191 28 PB1275-79 1284 21 1266 16 1234 24 1234 24 1234 24 PB1275-3 1238 24 1237 46 1236 120 1236 120 1236 120 PB1275-95 1105 11 1178 10 1317 20 1317 20 1317 20 PB1275-8 1462 13 1507 11 1571 19 1571 19 1571 19 PB1275-51 1586 22 1601 15 1621 19 1621 19 1621 19 PB1275-99 1621 26 1631 18 1644 22 1644 22 1644 22 PB1275-64 1791 16 1793 34 1796 71 1796 71 1796 71 PB1275-92 1632 33 1948 22 2303 17 2303 17 2303 17 PB1275-37 2452 41 2586 21 2693 17 2693 17 2693 17 PB1275-98 2509 64 2641 31 2744 16 2744 16 2744 16 PB1275-38 2670 35 2719 18 2756 17 2756 17 2756 17 PB1275-45 3075 29 3194 15 3269 16 3269 16 3269 16
NYANT Takena Formation from the Nam Co area 23ALDZ-67 95 4 105 52 341 602 95 4 95 4 23ALDZ-39 110 10 126 29 439 257 110 10 110 10 23ALDZ-4 113 3 134 17 530 150 113 3 113 3 23ALDZ-56 113 4 139 46 604 394 113 4 113 4 23ALDZ-8 114 5 130 47 439 431 114 5 114 5 23ALDZ-58 114 6 109 40 -10 476 114 6 114 6 23ALDZ-74 115 6 121 41 235 423 115 6 115 6 23ALDZ-70 116 3 127 29 326 276 116 3 116 3 23ALDZ-36 116 5 141 42 572 355 116 5 116 5 23ALDZ-46 120 7 166 46 877 315 120 7 120 7 23ALDZ-37 122 11 129 49 249 469 122 11 122 11 23ALDZ-14 123 4 146 26 536 206 123 4 123 4 23ALDZ-51 131 9 189 51 984 299 131 9 131 9 23ALDZ-40 135 8 138 44 200 398 135 8 135 8 23ALDZ-12 135 9 144 53 300 454 135 9 135 9 23ALDZ-47 141 7 202 50 981 279 141 7 141 7 23ALDZ-18 490 35 582 119 960 280 490 35 490 35 23ALDZ-53 507 11 577 82 865 199 507 11 507 11 23ALDZ-23 514 14 560 59 753 146 514 14 514 14 23ALDZ-17 529 13 556 60 669 151 529 13 529 13 23ALDZ-19 530 37 533 85 544 221 530 37 530 37 23ALDZ-59 544 29 492 140 258 438 544 29 544 29 23ALDZ-62 563 10 587 66 681 163 563 10 563 10 23ALDZ-41 565 12 558 31 529 78 565 12 565 12 23ALDZ-52 577 10 571 54 547 137 577 10 577 10 23ALDZ-10 581 6 581 51 582 128 581 6 581 6 23ALDZ-87 595 24 674 111 948 240 595 24 595 24 23ALDZ-27 604 13 628 87 716 204 604 13 604 13 23ALDZ-90 620 14 660 65 801 143 620 14 620 14 23ALDZ-57 648 51 705 66 891 109 648 51 648 51 189
23ALDZ-94 765 34 846 97 1064 173 765 34 765 34 23ALDZ-29 798 6 833 19 927 33 798 6 798 6 23ALDZ-38 888 13 911 37 969 63 888 13 888 13 23ALDZ-65 926 9 924 25 918 41 926 9 926 9 23ALDZ-13 937 20 956 42 998 65 937 20 937 20 23ALDZ-21 948 14 957 16 976 20 948 14 948 14 23ALDZ-72 1048 54 1059 39 1084 22 1084 22 1084 22 23ALDZ-98 1139 26 1145 26 1155 28 1155 28 1155 28 23ALDZ-7 1002 19 1058 54 1175 81 1175 81 1175 81 23ALDZ-24 1022 54 1072 53 1176 59 1176 59 1176 59 23ALDZ-26 1101 26 1132 55 1193 78 1193 78 1193 78 23ALDZ-5 954 10 1031 19 1198 27 1198 27 1198 10 23ALDZ-93 1181 39 1200 53 1233 67 1233 67 1233 67 23ALDZ-68 1030 32 1106 47 1259 62 1259 62 1259 62 23ALDZ-97 1121 19 1212 69 1380 95 1380 95 1380 95 23ALDZ-82 1184 30 1301 92 1500 121 1500 121 1500 121 23ALDZ-35 1410 12 1473 65 1565 80 1565 80 1565 80 23ALDZ-22 1627 39 1605 112 1575 135 1575 135 1575 135 23ALDZ-32 1450 19 1512 54 1600 64 1600 64 1600 64 23ALDZ-100 1679 61 1717 49 1763 40 1763 40 1763 40 23ALDZ-66 1639 30 1711 81 1801 91 1801 91 1801 91 23ALDZ-48 1744 33 1791 35 1846 33 1846 33 1846 33 23ALDZ-6 2106 30 2208 23 2305 17 2305 17 2305 17 23ALDZ-75 1941 22 2126 56 2309 54 2309 54 2309 54 23ALDZ-81 2132 29 2263 31 2384 27 2384 27 2384 27 23ALDZ-20 2481 32 2489 18 2495 10 2495 10 2495 10 23ALDZ-25 2386 61 2451 75 2505 66 2505 66 2505 66 23ALDZ-31 2216 62 2399 34 2557 14 2557 14 2557 14 23ALDZ-60 2487 36 2570 22 2636 13 2636 13 2636 13
PENBO1 Takena Formation from the Penbo area 84-27 110 2 144 41 748 327 110 2 110 2 84-51 112 7 25 21 -7730 6397 112 7 112 7 84-78 113 3 145 17 719 128 113 3 113 3 84-26 113 2 136 27 550 231 113 2 113 2 84-3 114 4 77 38 -961 765 114 4 114 4 84-40 118 3 170 38 969 253 118 3 118 3 84-79 118 6 139 52 500 447 118 6 118 6 84-93 119 3 156 34 767 247 119 3 119 3 84-76 119 4 144 24 575 196 119 4 119 4 84-83 119 5 122 26 171 261 119 5 119 5 84-99 121 5 168 83 898 577 121 5 121 5 84-60 122 3 153 34 670 261 122 3 122 3 84-35 122 6 98 41 -450 592 122 6 122 6 84-96 122 4 110 34 -151 406 122 4 122 4 84-73 124 8 138 76 372 683 124 8 124 8 84-53 125 3 173 23 901 149 125 3 125 3 190
84-46 126 4 83 34 -1016 633 126 4 126 4 84-63 134 6 146 40 353 332 134 6 134 6 84-16 135 7 187 49 914 301 135 7 135 7 84-59 142 9 121 74 -271 853 142 9 142 9 84-80 156 9 181 35 529 227 156 9 156 9 84-57 210 10 284 49 948 201 210 10 210 10 84-17 500 33 593 167 966 406 500 33 500 33 84-13 533 8 585 57 790 139 533 8 533 8 84-28 541 13 545 63 559 165 541 13 541 13 84-25 552 20 527 113 418 321 552 20 552 20 84-30 562 10 603 26 763 58 562 10 562 10 84-64 602 10 589 71 539 180 602 10 602 10 84-81 653 28 644 67 611 149 653 28 653 28 84-18 696 49 733 52 850 72 696 49 696 49 84-5 730 15 719 45 685 92 730 15 730 15 84-82 735 18 744 39 771 76 735 18 735 18 84-20 772 46 803 40 890 38 772 46 772 46 84-48 820 20 830 98 856 186 820 20 820 20 84-14 845 15 908 41 1063 67 845 15 845 15 84-22 903 13 853 75 723 141 903 13 903 13 84-52 910 8 928 39 970 66 910 8 910 8 84-39 984 26 974 129 950 223 984 26 984 26 84-72 1093 18 1058 31 985 46 985 46 985 46 84-43 1080 70 728 226 -249 611 1080 70 1080 70 84-36 1070 18 1093 26 1141 34 1141 34 1141 34 84-75 1147 42 897 178 321 368 1147 42 1147 368 84-89 964 34 1032 33 1180 35 1180 35 1180 34 84-62 1193 14 1226 39 1285 52 1285 52 1285 52 84-38 1064 40 1140 63 1289 83 1289 83 1289 83 84-12 1397 45 1372 95 1334 123 1334 123 1334 123 84-91 1236 31 1289 85 1379 114 1379 114 1379 114 84-92 1181 23 1255 37 1385 45 1385 45 1385 45 84-33 1097 26 1202 76 1396 105 1396 105 1396 105 84-15 1298 54 1336 39 1399 26 1399 26 1399 26 84-23 1594 30 1596 67 1598 78 1598 78 1598 78 84-74 1438 15 1510 48 1613 57 1613 57 1613 57 84-97 1544 20 1589 55 1649 64 1649 64 1649 64 84-29 1599 35 1644 53 1703 57 1703 57 1703 57 84-19 1655 33 1704 25 1764 19 1764 19 1764 19 84-85 1621 68 1718 107 1839 115 1839 115 1839 115 84-58 1647 74 1755 44 1885 15 1885 15 1885 15 84-84 1644 27 1779 22 1941 17 1941 17 1941 17 84-11 1921 16 2083 15 2247 11 2247 11 2247 11 84-65 2204 94 2273 48 2336 16 2336 16 2336 16 84-10 2463 43 2432 30 2405 22 2405 22 2405 22 84-70 2310 74 2389 42 2458 23 2458 23 2458 23 84-1 2245 20 2370 14 2479 10 2479 10 2479 10 84-49 2850 62 2875 43 2892 30 2892 30 2892 30 191
84-32 2818 38 2908 21 2972 12 2972 12 2972 12
192
CHAPTER 5: SEAWATER 187OS/188OS: THE CENTRAL ANDEAN
DISCONNECT
ABSTRACT
The Cenozoic rise in seawater 187Os/188Os is commonly attributed to an increased input of radiogenic osmium from continental sources via mountain building and weathering. The lithologies, aerial extent, and uplift history of the central Andes suggest it was a principal driver behind Cenozoic changes in seawater 187Os/188Os.
Samples of Paleozoic shale collected from the central Andes have osmium concentrations of 0.0448 to 0.1170 ng/g and 187Os/188Os values between 0.3201 and
0.9485; below the 187Os/188Os of modern seawater. The paucity of radiogenic osmium in central Andean shale correlates with low concentrations of rhenium, which in turn, correspond to relatively low amounts of organic carbon. Such low 187Os/188Os ratios and osmium concentrations make it highly unlikely that the central Andean orogeny played any role in elevating 187Os/188Os of seawater over the last 40 million years. These findings eliminate one of the most likely candidates for a major source of radiogenic osmium and significantly narrow the list of suspects responsible for the Cenozoic rise in seawater 187Os/188Os.
INTRODUCTION
The seawater record of temporal changes in the relative abundance of particular
geochemical isotopes provides an indirect way of analyzing past phenomena that may 193 otherwise leave no record. Intervals from the seawater record that indicate increased influxes of isotopes associated with continental lithologies are used to infer major
periods of continental weathering and erosion, which are commonly attributed to
tectonic episodes (e.g., Raymo et al., 1988; Richter et al., 1992). Cenozoic orogenies
and the resultant drawdown of atmospheric CO2 through the weathering of silicate minerals are thought to be responsible for the deterioration of the Earth’s climate from its Cretaceous “greenhouse” state to the current “icehouse” state (e.g., Raymo et al.,
1988, Raymo and Ruddiman, 1992). Consequently, delineating the causal mechanisms behind past variations in the relative abundances of continental-affine isotopes in seawater is crucial for reconstructing past episodes of tectonic uplift and for understanding the processes that regulate the Earth’s climate.
OSMIUM AND SEAWATER
Like the more familiar Rb-Sr system, the Re-Os isotopic system can potentially
be used to determine the relative contribution of continental weathering to seawater
chemistry (Peucker-Ehrenbrink and Ravizza, 2000). Rhenium-187 is partitioned into the
crust and decays to 187Os, resulting in elevated 187Os/188Os values of ~1.3 in the
continental crust and low values of ~0.13 in the mantle (Snow and Reisberg, 1995;
Peucker-Ehrenbrink and Ravizza, 2000). Seawater 187Os/188Os has increased ~250% over the past ~65 Ma to its current value of ~1.1 (Pegram et al., 1992; Ravizza, 1993;
Peucker-Ehrenbrink et al., 1995; Fig. 1A); an increase thought to be largely due to the 194 weathering of silicate rocks exposed during continental orogenic episodes (Turekian and Pegram, 1997).
However, many facets of the 187Os/188Os seawater record remain unclear and hamper attempts to correlate seawater 187Os/188Os with continental tectonism and weathering. Both 87Sr/86Sr and 187Os/188Os of seawater have increased during the
Cenozoic, yet significant and unexplained periods exist during which the two systems are decoupled (Peucker-Ehrenbrink et al., 1995). For example, from ~30 to 15 Ma, there is little change in seawater 187Os/188Os, which may reflect a steady but decreased input
from continental sources during that time or an increased influx of low 187Os/188Os from mantle-affine sources. Hydrothermal activity along mid-ocean ridges and their axial flanks (Sharma et al., 2000), extraterrestrial material (Sharma et al., 1997; Ravizza and
Peucker-Ehrenbrink, 2003), and the subaerial weathering of ophiolitic rocks (Turekian
and Pegram, 1997) all provide osmium with low 187Os/188Os, but the relative importance of these inputs to the 187Os/188Os ratio of seawater is poorly constrained. Additionally,
studies of submarine volcanic arcs indicate interactions between these features and
seawater may be more important to oceanic chemistry than previously thought (e.g., de
Ronde et al., 2003). From the continents, organic-rich shales are an important source of osmium with high 187Os/188Os values (e.g., Ravizza and Esser, 1993; Peucker-
Ehrenbrink and Hannigan, 1996; McDaniel et al., 2004), along with old crystalline
basement (Peucker-Ehrenbrink and Blum, 1998). It is unclear whether seawater
187Os/188Os is most influenced by influxes from the volumetrically small, but important terranes dominated by organic-rich shales, or from the much more abundant areas that 195 expose crystalline basement and rocks with compositions similar to that of average upper continental crust (Fig. 2). Because weathering of the latter two lithologies provides a sink for atmospheric CO2, whereas weathering of organic-rich shales releases
CO2 to the atmosphere, determining which sources are responsible for fluctuations in
the seawater 187Os/188Os record is critical for reconstructing potential sources and sinks
of CO2 during the Cenozoic.
THE PROBLEM
The abstruseness of the osmium seawater record is due in no small part to the many continental areas and lithologies that have yet to be analyzed for Re-Os. As investigators searched for culprits behind the Cenozoic rise in seawater 187Os/188Os, the
Himalayan orogen drew first attention (e.g., Turekian and Pegram, 1997; Singh et al.,
1999). However, recent work demonstrates that the increases in seawater 187Os/188Os cannot be attributed to weathering in the Himalaya (Sharma et al., 1999; Chesley et al.,
2000), begging the question: if not the Himalayas, then what has driven the Cenozoic increases in seawater 187Os/188Os? The Andean orogenic belt of South America contains enormous sections of Paleozoic black and grey shales and has been active throughout the Cenozoic, making it an ideal candidate for a major 187Os source. We collected and analyzed samples from Paleozoic shales in the central Andes of Bolivia to test the hypothesis that the weathering of shales within the central Andes has contributed to the
Cenozoic rise in seawater 187Os/188Os.
196
THE CENTRAL ANDES
The Andean thrust belt in Bolivia and Peru comprises a thick succession of
Ordovician, Silurian, and Devonian siliciclastic sedimentary rocks that includes an aggregate thickness of approximately 10 km of organic-rich marine shale and slate
(Sempere, 1995). Black and grey shales typically have abundant, and labile, osmium
with high 187Os/188Os ratios (Peucker-Ehrenbrink and Ravizza, 2000; Peucker-
Ehrenbrink and Hannigan, 1996; Fig 2, herein). Over 260,000 km2 of Cambrian through
Devonian-aged black and grey shales and slates in the central Andes are currently
exposed to weathering, an area roughly twice the size of Nepal (Brazil Divisão de
Geologia e Mineralogia, 1964; Fig 1B, herein). Current runoff from this area (>31,000
m3/s) is equal to that of the major Himalayan rivers (Guyot et al., 1999; Global River
Discharge Database, 2003), and forms key tributaries to the Amazon River, which dominates the current global riverine influx of osmium to the ocean (Levasseur et al.,
1999). Uplift and exposure of Ordovician-Devonian black and grey shales in the
Eastern Cordillera of the central Andes commenced during Eocene time (McQuarrie,
2002), and the foreland basin stratigraphic record suggests that these rocks were widely exposed and weathering by the Oligocene (DeCelles and Horton, 2003). The timing of uplift in the region, therefore, coincides with a large upswing in the relative abundance of radiogenic osmium in seawater (Fig. 1A).
197
SAMPLES AND RESULTS
To test the hypothesis that central Andean uplift and weathering contributed to the rise in seawater 187Os/188Os, we sampled fresh exposures of Ordovician, Silurian, and Devonian shales and slates along a ~200 km, east-west transect from La Paz,
Bolivia to the Andean foreland (Fig. 1B, Table 5.1). Samples were prepared following procedures described in Chesley et al., (2000) and analyzed by NTIMS. Duplicates of
samples show good reproducibility.
Our results indicate the uplift and weathering of shales exposed in the central
Andes played no part in elevating seawater 187Os/188Os during the Cenozoic. Osmium concentrations of the samples are between 0.0448 and 0.1170 ng/g, and 187Os/188Os values range from 0.3201 to 0.9485 (Table 5.1). The 187Os/188Os values are relatively
low; importantly, they are below that of the modern ocean, meaning that current runoff
from these areas would, if anything, tend to lower seawater 187Os/188Os values. Osmium concentrations are generally similar or slightly lower than those of rocks from other localities, whereas Re concentrations (0.035 to 0.352 ng/g) are significantly lower than similar lithologies exposed in other locations around the world. For comparison,
187Os/188Os values of shales in the Lesser Himalaya average ~3 and include values as high as 12, with mean Os and Re concentrations of 0.477 ng/g and 4.24 ng/g, respectively (Singh et al., 1999; Chesley et al., 2000). The small amounts of Re in the
Andean samples correspond to relatively low TOC (Table 5.1) and are consistent with expected values based on the TOC-Re relationships of other shales and organic-rich marine sediments (Fig. 3). 198
The overall conclusion that the Andes played no substantial part in the Cenozoic
rise in seawater 187Os/188Os remains unchanged even when taking into account possible effects of weathering on our samples. Weathering-related losses of Re and Os are accompanied by the removal of other components, particularly organic carbon
(Peucker-Ehrenbrink and Hannigan, 2000; Jaffe et al., 2002). However, the TOC values of the shales we analyzed match those of Paleozoic shales throughout the central Andes, including subsurface, unweathered shales (e.g., Moretti et al., 1996); REEs and trace elements are also in line with other Paleozoic shales in the region (e.g., Egenhoff and
Lucassen, 2003; data repository). Moreover, extensive mobilization of Os requires water with high fO2 (Xiong and Woods, 2000), which should be uncommon in the widespread and ~10 km thick organic-rich strata of the central Andes.
Although the geographic range represented by our samples precludes the categorical elimination of the Andes from the list of possible suspects responsible for the Cenozoic rise in seawater 187Os/188Os, several aspects suggest our results can be
extrapolated to a large portion of the central Andes. The lithostratigraphic units studied
either belong, or are similar, to many of the rock units exposed in the central Andes, and
as mentioned, have similar geochemical compositions. Our samples are from a
continuous tectonostratigraphic unit exposed throughout Bolivia and Peru (Fig. 1B)
wherein much of the sediment was deposited in related marine environments along the
former western margin of South America (Williams, 1995). It is important to note that
shales exposed in the northernmost Andes contain very high TOC (Bralower and
Lorente, 2003) and river water in this region has some of the highest known Re 199 concentrations (Colodner et al., 1993), suggesting this region may be an important location for future study.
DISCUSSION
We contend the low 187Os/188Os and moderate osmium concentrations measured
in the central Andean shales makes it highly unlikely that uplift and weathering of the
central Andes could have caused the rise in seawater 187Os/188Os over the last ~40 Ma.
When combined with studies that discredit a link between Himalayan weathering and
rising seawater 187Os/188Os values (Sharma et al., 1999; Chesley et al., 2000), it becomes clear that alternative components of the global osmium system require further investigation.
Some aspects of the Cenozoic seawater 187Os/188Os record can be explained by
varying influxes from mantle-affine sources with low 187Os/188Os. If changes in seawater 187Os/188Os values are solely a function of continental weathering, the plateaus in the 187Os/188Os seawater curve (Fig. 4) would imply a cessation of tectonic uplift or continental weathering during long periods of the Cenozoic, which is inconsistent with the geologic record. Periods during which 187Os/188Os of seawater does not increase may be attributable, at least in part, by an increased influx of low 187Os/188Os from the
alteration of oceanic crust via hydrothermal activity along mid-oceanic-ridges (MOR)
and axial flanks (e.g., Sharma et al., 2000). Oceanic crust production rates and seawater
187Os/188Os values are inversely related over the past ~30 Ma (Fig. 4), which, assuming oceanic crust production rates are proportional to seawater-oceanic crust interaction 200
(e.g., Richter et al., 1992), is consistent with some degree of causality. Production of oceanic crust increased noticeably at ~30 Ma, coeval with the cessation in the rise of seawater 187Os/188Os values, and remained high over a period during which seawater
187Os/188Os failed to increase (~30-15 Ma). However, from 40 to 30 Ma a positive correlation exists between seawater 187Os/188Os and oceanic crust production rates, and before this, no relationship between these systems is apparent.
CONCLUSION
In this study, we have tried to make a step towards a more thorough understanding the osmium system and the Cenozoic evolution of seawater 187Os/188Os by testing whether uplift and weathering of the central Andes, a likely candidate for a source of radiogenic osmium, was a principal driver behind the rise in seawater
187Os/188Os over the last ~40 Ma. Osmium concentrations (0.0448 to 0.1170 ng/g) and
187Os/188Os values (0.3201 to 0.9485) of Paleozoic shales exposed in the central Andes are surprisingly low, making it very unlikely that this orogeny had any role in elevating seawater 187Os/188Os during the Cenozoic. Although some relationship appears to exist between MOR activity and the seawater 187Os/188Os record over the past 30 Ma, no obvious correlation exists during other periods. Our results eliminate a likely candidate for a major source of radiogenic osmium and highlight the fact that many aspects of the seawater 187Os/188Os record need to be explored further before it can be confidently
interpreted. 201
Figure 5.1. Cenozoic osmium seawater values and the central Andes. A. Cenozoic record of seawater 187Os/188Os and major periods of central Andean tectonic uplift (McQuarrie, 2002; Lamb and Davis, 2003). B. Shaded areas are Paleozoic strata, dominantly shale and slate; Ordovician shales and slates, where divided from other Paleozoic units, are shown by hatchers. Sampling transect from La Paz (star) to the Andean foreland is shown by the dashed line. 202
Figure 5.2. Osmium isotopic compositions and concentrations. River data from Levasseur et al., (1999), seawater from Peucker-Ehrenbrink and Ravizza, (2000). 203
Figure 5.3. Total organic carbon and rhenium concentrations. Central Andean samples from this study are shown with solid diamonds and highlighted by the dashed oval. Other data are from Ravizza et al., (1991), Singh et al., (1999), Peucker-Ehrenbrink and Hannigan (2000), and Jaffe et al., (2002). 204
Figure 5.4. Osmium isotopic composition of seawater. (A) from Peucker-Ehrenbrink et al., (1995), and oceanic crust production rates (B) from Rowley (2002). A. Seawater 187Os/188Os increase is punctuated by a plateau between 30 and 15 Ma. B. Oceanic crust production rate is shown with the solid line, changes in those rates are shown with the dashed line. The seawater 187Os/188Os increase ceases at roughly 30 Ma, coincident with an increase in oceanic crust production rate, and remains level over an interval during which oceanic crust production rates are high. The inverse relationship between oceanic crust production rates and seawater 187Os/188Os suggests a relationship, but little correlation between the two is evident prior to 30 Ma. 205
TABLE 5.1. Sample Compositions Sample Age [Re] [Os] 187Re/188Os 187Os/188Os 2σ TOC *CR-3 Cretaceous 0.028 0.0127 11.798 0.9963 0.0042 0.06 *#CR-3 0.029 0.0077 20.029 1.0194 0.0027 *CR-3 0.027 CR-11 Devonian 0.352 0.1170 15.714 0.7438 0.0116 0.69 CR-11 0.348 CR-14 Silurian 0.053 0.1584 1.647 0.3201 0.0063 0.40 CR-14 0.041 CR-18 Ordovician 0.295 0.23 CR-20 Ordovician 0.035 0.0448 4.191 0.8741 0.0031 0.28 CR-20 0.027 CR-21 Ordovician 0.082 0.1035 4.291 0.9485 0.0430 0.65 CR-21 0.090 2-sigma error includes a variation in the blank of 50%; Re and Os concentrations are ng/g; TOC - total organic carbon (weight %); *sandstone; # - same sample, different split
Table 5.1: Sample Compositions. The samples and their Re, Os and TOC values.
206
CHAPTER 6: MOUNTAINS, MONSOONS AND MEGAFANS
ABSTRACT
In certain cases, the rivers draining mountain ranges create unusually large fan- shaped bodies of sediment that are referred to as fluvial megafans. We combine information from satellite imagery, monthly discharge and precipitation records, digital elevation models, and other sources to show that the formation of fluvial megafans requires particular climatic conditions. Specifically, modern fluvial megafans in actively aggrading basins are produced by rivers that undergo moderate to extreme seasonal fluctuations in discharge that result from highly seasonal precipitation patterns.
The global distribution of modern megafans is primarily restricted to 15°–35° latitude in the Northern and Southern Hemispheres, corresponding to climatic belts that fringe the tropical climatic zone. No relationship exists between megafan occurrence and drainage-basin relief or area. The tendency of rivers with large fluctuations in discharge to construct megafans is related to the instability of channels subject to such conditions.
Because of the correlation between seasonal precipitation and megafan occurrence, the recognition of fluvial megafan deposits in ancient stratigraphic successions may provide critical information for paleoclimate reconstructions.
INTRODUCTION
Fluvial megafans form as rivers exit the topographic front of a mountain belt, migrate laterally in the adjacent basin, and deposit large fan-shaped bodies of sediment 207
(DeCelles and Cavazza, 1999; Fig. 1 herein). Although they share some characteristics with alluvial fans, fluvial megafans are distinct geomorphic features, distinguishable from stream-dominated alluvial fans by their unusually large area (areas of 103–105 km2
for fluvial megafans vs. generally <100 km2 for alluvial fans), low gradient (fluvial megafans, generally ~0.1° to 0.01°; alluvial fans, ~1° to 4°), sedimentary texture
(sediments in fluvial megafans vary from boulders at the apex to predominantly silt and mud at their toes), and depositional processes (fluvial megafans are devoid of sediment- gravity flows) (DeCelles and Cavazza, 1999; Horton and DeCelles, 2001; and references therein). Fluvial megafans play an integral role in the dispersal and deposition of sediment in tectonically active areas. The deposits of these features, therefore, serve as primary repositories for information on climatic conditions and rates of tectonic uplift and erosion in both young and ancient mountain belts. Fluvial megafan deposits have been recognized in stratigraphic successions in the Cordillera of the western United States and Canada (Eisbacher et al., 1974; Lawton et al., 1994; DeCelles and Cavazza, 1999), the Andes (Horton and DeCelles, 2001), the Pyrenees (Hirst and
Nichols, 1986), and the Himalaya (Willis, 1993; DeCelles et al., 1998). Characteristics of these deposits have been used to support a number of contentious hypotheses pertaining to basin dynamics and the structural evolution of mountain belts (cf. Love,
1973; Schmitt and Steidtmann, 1990; Lawton et al., 1994; Janecke et al., 2000).
One of the most important unresolved issues related to fluvial megafans, and one with ramifications for tectonic, paleoclimatic, geomorphic, and sedimentary studies, centers on the fact that only a limited number of fluvial megafans exist today despite the 208 multitude of sizeable rivers around the world that cross faults, exit topographic highlands, and enter basins. Is there something unique about rivers that create fluvial megafans? Studies have focused on fluvial megafan facies, morphologies, and the relationship of megafans to drainage-basin development (Wells and Dorr, 1987a,
1987b; Iriondo, 1993; Stanistreet and McCarthy, 1993; Singh et al., 1993; Sinha and
Friend, 1994; Gupta, 1997; DeCelles and Cavazza, 1999; Horton and DeCelles, 2001;
Shukla et al., 2001), but the underlying factors governing megafan formation and distribution along individual mountain systems and around the globe remain unknown.
We studied the characteristics of rivers that form, and do not form, fluvial megafans in order to address the fundamental questions that surround the construction of these features. Our results show that fluvial megafans are produced by rivers that undergo large seasonal fluctuations in discharge, which result from seasonal or monsoonal precipitation. The same association between seasonal precipitation and fluvial megafans is present in ancient stratigraphic successions, suggesting that fluvial megafan deposits may be useful paleoclimate indicators.
METHODS
We compared the data of all 202 rivers, looking for qualities that discriminate megafan-forming rivers from rivers that fail to create megafans (e.g., Fig. 3A). Because fluvial megafans are depositional features, and therefore necessarily limited to actively aggrading basins, we focused on regions where rivers are obviously entering aggrading basins (Figs. 2, 3B). Much of northern Asia is excluded, primarily because of 209 incomplete data (e.g., few Landsat-5 images); however, those regions in northern Asia that were examined lack fluvial megafans. The areas that received more detailed study are the Andes, the Cordillera of Central America, the Himalaya, the Indonesian orogenic system, parts of sub-Saharan Africa, the Atlas and adjacent Mediterranean mountain systems, the Russo-Sino and southern Mongolian regions north of the Tibetan
Plateau, and the Middle East. This omits some of the rivers that can be used in the analyses (53 of the 202 rivers), but is important for removing the noise that would otherwise be introduced by river systems that occupy nonaggradational areas where fluvial megafans cannot develop. Some of the areas deemed nonaggradational can be disputed, but our findings remain unchanged even with these data points included in the analysis. Furthermore, those basins and river systems that were closely scrutinized represent many different geological and geographical settings.
RESULTS
We identified 15 new fluvial megafans in addition to the 13 megafans documented in the literature (Fig. 2B). Of the 202 rivers in active depositional basins that we surveyed, 115 have no associated megafan. The majority of river systems can be easily classified as either having megafans or not, but a few cases are ambiguous. For example, the upper Indus River displays some signs of radial dispersal, but only over a limited angle. In these cases, the rivers’ deposits are classified as “possible fluvial megafans.” 210
Figure 3 illustrates the occurrence of fluvial megafans in terms of key hydrologic, geomorphologic, and climatic data. Figures 3A and 3B display discharge
peakedness (the average discharge during the month with the greatest discharge, divided
by the average annual discharge) plotted against the average annual discharge
(reflecting river size). The lack of fluvial megafan rivers on the left side of Figure 3A
indicates that a minimum river discharge (~20 m3/s) is required to create a fluvial megafan. As a result, fluvial megafan rivers typically have moderate to large drainage basins with moderate to high relief (Figs. 3C, 3D). However, many rivers with large, high-relief drainage basins exit mountain belts but do not produce fluvial megafans
(Figs. 3C, 3D). The disconnect between megafan formation and drainage basin characteristics is most clearly shown along the eastern front of the Andes, where fluvial megafans are prevalent along parts of the mountain front but absent along other stretches, despite the fact that these rivers have nearly identical mean discharges, drainage-basin areas, and relief (Fig. 2). Thus, although aggradation and large discharge
and drainage area facilitate megafan formation, by themselves they are insufficient. The
data indicate fluvial megafan construction requires rivers that experience seasonal fluctuations in discharge.
Given sufficient aggradation rates and discharges, the single characteristic
shared by all rivers that produce fluvial megafans is significant seasonal variation in
discharge, as measured by discharge peakedness (Fig. 3B). Those rivers that produce
fluvial megafans cluster in the upper half of Figure 3B, reflecting more acute
peakedness. Rivers that fail to construct fluvial megafans have relatively constant 211 discharges throughout the year (Fig. 3B), regardless of average annual discharge, and drainage basin size and relief. A few rivers in aggrading basins have discharges with
relatively high peakedness but fail to form fluvial megafans for reasons that are discussed below. The seasonal fluctuations in discharges observed in megafan-forming rivers correspond to seasonal precipitation patterns within the rivers’ drainage basins
(59 precipitation records; Fig. 3E). Significantly, almost all fluvial megafans are symmetrically disposed between 15° and 35° latitude in both the Northern and Southern
Hemispheres (Fig. 2).
INTERPRETATION
Understanding why river systems with seasonal fluctuations in discharge are
more likely to produce fluvial megafans requires an understanding of how fluvial
megafans differ from typical river-channel belts. The distinctive fan-shaped sediment
lobes associated with megafan rivers indicate lateral instability that promotes rapid
channel migration and frequent avulsion. This tendency has been noted in previous
studies of megafans (e.g., Geddes, 1960; Sinha and Friend, 1994; Horton and DeCelles,
2001) and is exemplified by the Kosi River, which drains a large part of the Himalaya in northeast India and Nepal. The Kosi River has migrated westward >113 km in just
228 years (averaging 0.5 km/yr) (Wells and Dorr, 1987b). Additionally the overbank areas of the megafans observed in satellite images are replete with abandoned channels
(Horton and DeCelles, 2001; this study). 212
The processes and conditions controlling river avulsion are complex (e.g., Jones
and Schumm, 1999; Mohrig et al., 2000) and beyond the scope of this study. However,
evidence suggests that large fluctuations in discharge may promote channel instability
and avulsion. Floods often serve as avulsion-triggering events (Jones and Schumm,
1999), and in the case of megafan rivers, the annual flooding associated with the wet
season serves as an effective, frequently recurring catalyst for avulsion. For example,
the major channel shifts and avulsions of the Kosi River have occurred during the
annual monsoonal floods (Wells and Dorr, 1987a). Periods of rapid channel migration
also may be associated with peak annual discharges, as river banks are eroded by
increased stream power (Ritter et al., 2002). The relatively high sediment yields from
basins that alternate between wet and dry seasons (Wilson, 1973) may at times overtax
a river’s transport capacity. The high sediment yields can lead to channel aggradation and result in avulsion or rapid channel migration (e.g., Wells and Dorr, 1987a; Bryant et al., 1995). Seasonal precipitation also may influence the type and density of vegetation along the river banks, affecting bank stability and, therefore, migration rates and
avulsion frequency.
CLIMATE PATTERNS AND FLUVIAL MEGAFANS
The relationship of fluvial megafans to seasonal discharges, combined with their
latitudinal distributions, suggests that megafans may be primarily controlled by global climatic patterns. Fluvial megafans are absent in the tropical climatic zone (Fig. 2).
Steady month-to-month rainfall totals in this area are reflected in consistent monthly 213 discharges and a lack of fluvial megafans (e.g., Indonesia and northern South America;
Fig. 2). In contrast, areas fringing the tropical climatic zone are characterized by seasonal precipitation (Fig. 3D) and have many fluvial megafans (Fig. 2). Although a small number of fluvial megafans fall outside of these latitudinal belts, the outlying megafan rivers nonetheless undergo significant seasonal fluctuations in discharge. The current climatic pattern in areas of the Andes where modern megafans occur was established by early Miocene time (ca. 23–18 Ma) (Iriondo, 1993), and the South Asian monsoon has existed since at least ca. 10 Ma (Dettman et al., 2000), making it unlikely that these features are relicts of drastically different climates.
APPLICATIONS TO THE STRATIGRAPHIC RECORD
A correlation between seasonal precipitation patterns and fluvial megafan formation can be documented in ancient stratigraphic successions where paleoclimate and fluvial megafan deposits have been well studied. The Asian monsoons began, or at least intensified, between ca. 10 and 8 Ma (Dettman et al., 2000). Coevally, the deposits of the Himalayan foreland basin reflect a change from small sinuous fluvial channels to fluvial megafans (DeCelles et al., 1998). Fluvial megafans also formed in the central
Andean foreland basin while seasonal precipitation patterns prevailed, similar to the present-day climate (Iriondo, 1993). Although less definitive, seasonal precipitation may have occurred in the Cretaceous Cordillera of the western United States (Glancy et al., 1993; J.T. Parrish, 1998, personal commun.); during this period, several large fluvial 214 megafans traversed the western margin of the Cordilleran foreland basin (Lawton et al.,
1994; DeCelles and Cavazza, 1999).
DISCUSSION AND CONCLUSION
All rivers that produce fluvial megafans undergo large fluctuations in discharge, but not all rivers that undergo large fluctuations in discharge form fluvial megafans.
Peripheral conditions exist that can hinder the construction of fluvial megafans. The
majority of factors prohibiting fluvial megafan formation are site-specific. In some
cases, rivers entering narrow or small basins cannot migrate laterally and therefore
cannot construct fan-shaped sediment lobes. The spacing between channel outlets can
also be an important factor in megafan formation. For example, the outlets of the Ravi
and Chenab Rivers in the western Himalayan foreland basin are closely spaced; thus
channel migration is limited by the adjacent river’s deposits (Geddes, 1960) and fluvial
megafan formation is inhibited.
Fluvial megafans are volumetrically important distributary systems that form
adjacent to both extreme and subdued topography. Provided a sufficient aggradation
rate and discharge, a river will form a fluvial megafan if it experiences large seasonal
fluctuations in discharge. Because of the correlation between seasonal precipitation and
modern megafan occurrence, the presence of fluvial megafan deposits in the
stratigraphic record can provide important information for paleoclimate reconstructions. 215
Figure 6.1. Images of modern fluvial megafans. A: Kosi fluvial megafan forms as the
Kosi River exits the Himalaya. B: Pilcomayo fluvial megafan (outlined by dashes) forms as the Pilcomayo River exits the Andes. C: DEM of Pilcomayo megafan (100 times vertical exaggeration). Field of view matches that in B. 216
Figure 6.2. Studied rivers and climate. The tropical climatic zone is based on Köppen’s classification (de Blij and Muller, 1996). Megafan rivers are denoted by black triangles, possible megafan rivers by gray triangles, nonmegafan rivers used in comparisons by open circles, and nonmegafan rivers examined, but omitted from comparisons are shown with open squares. Fluvial megafans are most prevalent in latitudinal belts that fringe the tropical climate zone, corresponding to regions with seasonal precipitation. 217
Figure 6.3. Comparison of rivers. Symbols described in Figure 2. A: Discharge (Q) peakedness (see text for details) vs. average annual discharge for all rivers examined. B:
Discharge peakedness vs. average annual discharge for rivers used in comparisons.
Megafan rivers have greater seasonal fluctuations than nonmegafan rivers. C: Drainage- basin area derived from DEMs and published data. There is no discernible break between drainage-basin area of megafan rivers and nonmegafan rivers. D: Drainage- basin relief (river elevation at mountain front subtracted from highest point in basin) of rivers in the central Andes.
218
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