ABSTRACT

The Min Shan region, located along the east-ern margin of the Tibetan Plateau north of theSichuan Basin, provides an important naturallaboratory in which to study the rates and pat-terns of deformation and their relationship tomountain building at the margin of the plateau.The topographic margin of the plateau is coin-cident with a north-trending mountain range,the Min Shan, that stands nearly 2 km abovethe mean elevation of the plateau (~3500 m inthis region). We exploit the preservation of a se-ries of variably deformed Quaternary sedi-ments along the western flank of the range toinvestigate the Pleistocene-Holocene deforma-tion field within the Min Shan region. Mappingand field observations of remnant alluvial fansof late Pleistocene age indicate that deforma-tion within the Min Shan involved substantial(~10°), rapid, down-to-the-northwest tilting.The geometry of the deposits and the partialpreservation of an erosion surface beneath thebasin suggest that much of the modern relief ofthe Min Shan relative to the Tibetan Plateau isa consequence of this late Pleistocene tilting.Rates of tilting inferred from luminescence dat-ing of interbedded loess have been remarkablyrapid (~10–8 rad/yr). Similarly rapid rates ofHolocene differential rock uplift are inferredfrom tilted lacustrine sediments in the south-western part of the range. The range isbounded on the west by the Min Jiang faultzone, an east-vergent reverse fault. However,Holocene alluvial terraces in headwaters of theMin River are preserved across the fault in sev-eral places, indicating that displacement rates

on the Min Jiang fault are <1 mm/yr. Activefaulting only occurs along the eastern foot ofthe range (Huya fault) for a short distance(~60 km), despite 3 km of relief on the easternrange front. The relationship between thesestructures and the tilting observed in the MinJiang basin is enigmatic; the faults do not ap-pear to exert a strong control on the rates andpattern of deformation within the basin. A sim-ple flexural model demonstrates that rates oftilting on the western flank of the Min Shan aretoo high to be simply attributed to an isostaticresponse to surficial loading and unloading ofthe lithosphere. Present-day horizontal short-ening across the Min Shan is geodetically de-termined to be less than 2–3 mm/yr, suggestingthat only a small part of the observed tilting canbe attributed to horizontal shortening. Thus,tilting and concomitant differential rock upliftin the Min Shan appear to require an addi-tional driving component. We suggest thatQuaternary deformation along the westernMin Shan may reflect the surface response tothickening of a weak lower crust at the marginof the Tibetan Plateau.

Keywords: neotectonics, Tibetan Plateau,mountain building, flexure, Sichuan Basin,lower crust.

INTRODUCTION

Since the recognition that widespread intracon-tinental deformation within central Asia is the re-sult of the Cenozoic collision of India and Eurasia(Argand, 1924; Molnar and Tapponnier, 1975),there has been a long-standing debate centered onthe degree to which crustal thickening within theTibetan Plateau accommodates convergence be-

tween the two continents (cf. England and House-man, 1986; Tapponnier et al., 1982). Althoughmost workers agree that the present crustal thick-ness of nearly 70 km and the great elevation(Fielding et al., 1994) of the plateau are the resultof the collision (Dewey and Burke, 1973; Harri-son et al., 1992; Molnar, 1988), the amount, rates,and style of shortening within Tibet as well as itsspatial and temporal distribution remain largelyunknown (Burchfiel and Royden, 1991; Cowardet al., 1988). Many of the models for Asian defor-mation predict markedly different kinematics inthe region east and northeast of the Indian inden-ter (Avouac and Tapponnier, 1993; Cobbold andDavy, 1988; England and Molnar, 1990b); thus,one means of distinguishing between competingmodels is to determine the kinematics of defor-mation within this region of the plateau. Our firstgoal therefore is to document the distribution andrates of late Cenozoic deformation in the north-trending mountain range termed the Min Shan—a little-known but important segment of the east-ern margin of the Tibetan Plateau north of theSichuan Basin, in Sichuan Province, China.

Our second goal is to evaluate the processesdriving deformation at the margin of the plateau.Along much of the eastern margin there is a lackof correlation between the surface expression ofstructures and topography (Burchfiel et al., 1995;Royden et al., 1997). Faults and folds are oftenoblique to, and commonly older than, the topo-graphic margin. These observations—coupledwith satellite geodesy (King et al., 1997; Chenet al., 2000) and increasing evidence for a weaklower crust beneath Tibet (Jin et al., 1994; Nelsonet al., 1996; Masek et al., 1994a)—led Roydenet al. (1997) to suggest that crustal thickening ineastern Tibet may be focused in the lower crust.In this model, rock and surface uplift may be due

375

Neotectonics of the Min Shan, China: Implications for mechanisms drivingQuaternary deformation along the eastern margin of the Tibetan Plateau

Eric Kirby*Kelin X. WhippleB. Clark Burchfiel

Dpartment of Earth, Atmospheric, and Planetary Sciences, Massachusetts Institute of Technology,Cambridge, Massachusetts 02139

Wenqing Tang Chengdu Institute of Geology and Mineral Resources, Chengdu, Sichuan Province, ChinaGlenn Berger Desert Research Institute, 2215 Raggio Parkway, Reno, Nevada 89512Zhiming SunZhiliang Chen

Chengdu Institute of Geology and Mineral Resources, Chengdu, Sichuan Province, China

GSA Bulletin;March 2000; v. 112; no. 3; p. 375–393; 11 figures; 3 tables.

}

*E-mail: ekirby@mit.edu.

}

to isostatic and dynamic responses to lower-crustal thickening and not directly related to hor-izontal shortening of the upper crust. Our map-ping and field observations along the Min Shansuggest that it is difficult to account for the pat-tern of surface deformation inferred from tiltedQuaternary sediments by displacement on recog-nized structures. Furthermore, rates of late Pleis-tocene tilting along the western flank of the rangefar exceed geodetically measured rates of hori-zontal shortening (King et al., 1997; Chen et al.,2000). Thus, our results suggest the presence ofan additional driving force and are permissive ofthickening driven by lower-crustal flow.

Our third goal, intimately linked to the othertwo goals, is to evaluate the morphologic evolu-tion of the Min Shan and, in particular, to ascer-tain to what degree high mountain ranges alongthe eastern margin of the plateau may reflect anisostatic response to erosion at the margin. Topo-graphically high mountain ranges commonly oc-cur at the edges of continental plateaus (Fieldinget al., 1994) and may be (1) primarily tectonicfeatures related to the dynamics of shortening atthe margin (Lyon-Caen and Molnar, 1983; Roy-den et al., 1997), (2) primarily isostatic featuresrelated to mass removal along the topographicfront (Masek et al., 1994b; Molnar and England,1990; Montgomery, 1994) or, (3) some combina-tion of both mechanisms (Burbank, 1992). Thelack of large-magnitude shortening across theeastern margin (Burchfiel et al., 1995) makes thisan important region in which to test the hypoth-esis of isostatically generated topography. In thispaper, we argue from a simple flexural model thatthe rates of observed surface deformation in theMin Shan require unreasonably high rates ofmass removal by erosion to produce the topogra-phy isostatically and thus point to the importanceof additional driving mechanisms.

BACKGROUND

In this section we summarize the salient partsof the late Cenozoic tectonics along the easternmargin of the Tibetan Plateau to set the back-ground for our work in the Min Shan region. Wefollow Burchfiel et al. (1995) in using the term“Longmen Shan region” to refer to the easternmargin of the Tibetan Plateau along the edge ofthe Sichuan Basin between the Xianshuihe faultand the west Qinling Shan, and we use the term“Min Shan region” to refer to the margin of theplateau north of the Sichuan Basin (Fig. 1).

Active Tectonics of the Longmen Shan

The topographic margin of the Tibetan Plateaualong the Longmen Shan is one of the most im-pressive continental escarpments in the world;

from the Sichuan Basin at 500 to 700 m elevation,the land rises westward to peak elevations ex-ceeding 6000 m over horizontal distances of40–60 km. This topographic front led a number ofworkers to suggest that it was the locus of majoractive shortening (up to 20 mm/yr) between Tibetand south China (Avouac and Tapponnier, 1993;Ratschbacher et al., 1996). Recent results of GPS(Global Positioning System) space geodetic sur-veys indicate, however, that modern shorteningrates across this margin must be less than2–3 mm/yr and are within uncertainty of zero(King et al., 1997; Chen et al., 2000). Seismicity

along this margin is generally restricted to smallevents (M < 5) that occur in a band paralleling thetopographic margin (Editorial Board, State Seis-mological Bureau, 1989). The historic record ofearthquakes in this region extends back over1500 yr and documents a remarkable lack of sig-nificant (M > 4–5) events (Editorial Board, StateSeismological Bureau, 1989), consistent with theslow shortening rates inferred from GPS.

Geologic evidence for active faulting in thenorthern Longmen Shan is equally scant and ap-pears to be restricted to the Min Shan region(Burchfiel et al., 1995; Chen et al., 1994). Thrust

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fold axial trace

Legend

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Precambrian massifs

Metamorphic rocks ofuncertain age

0 100 km

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AT

80° 90° 100°

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30°

India

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Figure 1. Simplified geologic map of the Longmen Shan region of the eastern Tibetan Plateau.Geology modified after Burchfiel et al. (1995), 1:200000 geologic maps (Ministry of Geology andMineral Resources, People’s Republic of China, 1991), and our mapping. Inset shows location ofthe region relative to major tectonic elements of the Indo-Asian collision zone. Abbreviations:AT—Altyn Tagh fault, K—Kunlun fault, L—Longquan anticline, T—Tangwangzhai nappe,TB—Tarim basin, SB—Sichuan Basin,W—Wenchuan-Maowen fault zone, X—Xianshuihe fault.

faults along the topographic front in the northernLongmen Shan (northeast of An Xian along theTangwangzhai nappe; Fig. 1) are overlapped byUpper Triassic terrestrial clastic rocks, restrictingthe latest displacement along these faults toMesozoic time (Burchfiel et al., 1995).

The absence of significant late Cenozoic short-ening across the Longmen Shan is corroboratedby a distinct lack of foreland sedimentation.Cenozoic deposits are restricted to the south-western corner of the Sichuan Basin (Fig. 1) andare generally <500 m in thickness. Quaternarysediments in particular consist of a thin (<100 m)veneer in this region and appear to be ponded be-hind the Longquan anticline (Fig. 1). Althoughthis anticline may be active, definitive evidence islacking (Burchfiel et al., 1995). Elsewherethroughout the Sichuan Basin, modern sedimentsare routed to the East China Sea via tributaries ofthe Yangtze River. Thus, the Sichuan Basin doesnot appear to have been flexurally loaded duringthe Cenozoic (Royden et al., 1997). Indeed, in thenorthwestern corner of the basin (northeast of AnXian, Fig. 1), hogbacks of Mesozoic red bedsramp up toward the plateau, suggesting that thebasin may have been flexurally upwarped duringthe late Cenozoic.

The slow modern velocity field, the lack of aCenozoic foredeep, and the limited Cenozoicshortening across the Longmen Shan led Roydenet al. (1997) to hypothesize that much of the de-velopment of this region of the plateau may bethe result of eastward flow of weak lower crustfrom beneath the central plateau. Where this ma-terial impinges on the strong crust of the SichuanBasin, it may focus crustal thickening (and con-sequent surface uplift) in the absence of signifi-cant upper-crustal shortening.

Tectonic Setting of the Min Shan

The Min Shan is a topographically high-standing mountain range that delineates the edgeof the Tibetan Plateau for nearly 200 km north ofthe Sichuan Basin (Figs. 1 and 2A). The range is40–50 km wide and contains peaks that reach el-evations of 5600 m. The mean elevation of theplateau in this region is ~3500 m, and relief onthe western flank of the range approaches 2 km(Fig. 2B). Although regional topographic gradi-ents across this margin of the plateau are not aslarge as along the Sichuan Basin, the easternflank of the Min Shan drops 3–4 km to the low-lands of the west Qinling orogen (<1000 m) overa lateral distance of 60–80 km.

Despite the impressive mountain front alongthe eastern Min Shan, active faults occur onlyalong the central 60 km of the range front. Herethe Huya fault juxtaposes Devonian–Triassic sili-ciclastic and carbonate rocks on the west with

Proterozoic crystalline basement and cover se-quences of the west Qinling orogen to the east(Fig. 3). The fault strikes just west of north anddips steeply (70°–85°) to the west. A series ofearthquakes along the fault in 1976 demonstratethat active displacement across the fault isoblique reverse motion with a left-lateral compo-nent (Han and Xia, 1980; Jones et al., 1984).Hypocentral locations for these shocks are notwell determined, but appear to be between 10 and15 km deep (Jones et al., 1984), suggesting thatthe fault remains steep throughout much of theupper crust (Fig. 3).

The Min Shan is bounded along much of itswestern flank by another reverse fault, the MinJiang fault zone. The Min Jiang fault zone con-sists of several strands of high-angle faults(Burchfiel et al., 1995; Chen et al., 1994), theeasternmost of which is hereafter referred to asthe Min Jiang fault. The Min Jiang fault cutsQuaternary deposits along the headwaterreaches of the Min Jiang (jiang = river) and isclearly a recently active structure (Burchfielet al., 1995; Chen et al., 1994; Tang et al., 1993;Zhao et al., 1994). However, geologic estimatesof the magnitude, rates, and sense of displace-ment vary widely in the literature (cf. Chenet al., 1994; Zhao et al., 1994). Consequently,the significance of the Min Jiang fault in the ac-tive deformation field at the edge of the plateauis unknown.

Despite the presence of active faults on eitherflank of the Min Shan, recent geodetic surveys inthe northern Longmen Shan region demonstratethat, within uncertainties of ~2–3 mm/yr, there isno active shortening across the range (King et al.,1997; Chen et al., 2000) (Fig. 2). However, asboth the Min Jiang and Huya faults have steep(>60°) westward dips, horizontal shortening as-sociated with active reverse faulting may be mi-nor. It is important to note, however, that the geo-detic data restrict any vertical component ofdeformation associated with shortening acrossthe Min Shan to less than a few millimeters peryear. With the exception of the 1976 earthquakesequence on the Huya fault, historic seismicityalong the eastern Min Shan appears to be limitedto one large (approximate magnitude of 7–8)event in 1879 (Editorial Board, State Seismolog-ical Bureau, 1989). Thus, the geodetic measure-ments are likely representative of the regional ve-locity field over the past 1000–2000 yr and areconsistent with slow horizontal strain accumula-tion across the range.

Neogene–Quaternary deposits are preservedalong the western flank of the Min Shan in theheadwaters of the Min Jiang (hereafter referred toas the Min Jiang basin—Figs. 1 and 3). They rep-resent one of the only significant exposures of up-per Cenozoic sediment preserved in this region of

the Tibetan Plateau. We mapped these depositsand their structural relationships at a scale of1:100 000 over the course of three field seasonsbetween 1996 and 1998 (Fig. 4). Because theolder sediments are progressively more deformed,we are able to place constraints on the displace-ment across the Min Jiang fault, as well as to eval-uate the late Cenozoic deformation field withinthe Min Shan range itself.

STRATIGRAPHY OF THE MIN JIANGBASIN

Pre-Cenozoic Rocks of the Min Shan

Paleozoic–Mesozoic rocks in the Min Shanregion comprise a thick (>4000 m) section ofplatformal, shallow-water siliciclastic rocks andreef-bearing limestones. We follow Burchfielet al. (1995) in referring to this region as the XueShan platform. In the northern Min Shan, theplatformal rocks are continuous to the east withlate Proterozoic–Silurian rocks that rest uncon-formably on metamorphic rocks shown as Pre-cambrian on published Chinese geologic maps(Fig. 3). However, to the west of the Min Jiangfault zone, Carboniferous and Permian rocksthin to nearly half their thickness on the plat-form. The overlying Triassic rocks change fromshallow-water limestones to thick deposits ofdeep-water flysch characteristic of the Songpan-Garze basin. Thus, the Min Jiang fault zonemarked an important paleogeographic boundaryin the Paleozoic and early Mesozoic and mayhave been a structural boundary at this time(Burchfiel et al., 1995).

Pliocene Deposits

The oldest Cenozoic deposits in the studyarea are a series of small, discontinuous lensesof Tertiary conglomerates. The deposits are re-stricted to elongate exposures preserved alongstrands of the Min Jiang fault system (Tp;Figs. 2, 4) and are mapped as late Neogene(Pliocene) on Chinese geologic maps (Ministryof Geology and Mineral Resources, People’sRepublic of China, 1991). Near Kakagou(Fig. 4), well-indurated, coarse conglomerate isfaulted against Triassic limestone. The con-glomerate is weathered to a deep red color thatis suggestive of its relative antiquity. However,the absolute age of these deposits is unknown,owing to the lack of datable material and/or fos-sils. Likewise, the original extent of these de-posits cannot be confidently reconstructed.Their present geometry along the Min Jiangfault system, however, implies that they mayhave been restricted basins ponded during initi-ation of the fault system (Chen et al., 1994).

NEOTECTONICS OF THE MIN SHAN, CHINA

Geological Society of America Bulletin, March 2000 377

Quaternary Deposits

Quaternary deposits in the Min Jiang basin canbe subdivided into two broad categories: middlePleistocene–upper Pleistocene alluvial conglom-erates and upper Pleistocene–Holocene alluvium

related to the current Min Jiang drainage (Figs. 4and 5). We designate relative ages with subscriptsranging from 1 to 4 (oldest to youngest). We cau-tion, however, that these relative-age designa-tions are restricted to this basin and should not beconfused with similarly designated deposits else-

where in China (similar labels are commonlyused to represent absolute age on Chinese geo-logic maps) (Ministry of Geology and MineralResources, People’s Republic of China, 1991).

Middle Pleistocene conglomerates (Q1 inFig. 4) form most of the deposits within the Min

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378 Geological Society of America Bulletin, March 2000

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Figure 2. (A) Regional topography (GTOPO30; nominal resolution 1 km) and GPS velocities of the northern Longmen Shan region (Chen et al.,2000). Station names and coordinates can be found in King et al. (1997). Uncertainty ellipses are plotted at 95% confidence intervals. Major struc-tures in the Min Shan are shown for reference to Figure 1 (box). SB—Sichuan Basin. Dashed rectangle shows region of topography used in flex-ural model (Fig. 10). (B) Schematic cross section depicting position of the Min Shan at the margin of the Tibetan Plateau and its relationship tothe Min Jiang and Huya faults. Topography taken from a single profile near the northwest-southeast centerline of the dashed rectangle in A.

Jiang basin (Guanyin Shan Formation) (Chenet al., 1994). Q1 deposits are bounded on the westby the Min Jiang fault. They are composed of aseries of west-dipping coarse-gravel and boulderconglomerates, interbedded near their base with

10–20 cm beds of fine sand, silt, and reworkedloess. Q1 deposits are locally well indurated andare often exposed in 10–30 m high cliffs. Clastswithin the Q1 conglomerates range from 5 to30 cm in diameter and are well rounded and well

sorted. Clast composition is primarily limestone(>90%), with minor sandstone and siltstone ma-terial. The conglomerates are clast supported andcontain a weak to moderate crystalline calcare-ous cement.

NEOTECTONICS OF THE MIN SHAN, CHINA

Geological Society of America Bulletin, March 2000 379

N

33° 20' N

32° 00'104° 00' E

Qingchuan fault

Min

Jia

ng F

ault

Sys

tem

Xue Shan fault

Huya fault

Songpan

30 km150

Aug. 16, 1976M = 7.2

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Tazang fault

Figure 4

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Aug. 21, 1976M = 6.7

Aug. 23, 1976M = 7.2

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ExplanationCenozoic

Holocene alluvium

Pleistocene alluvium

Pliocene conglomerate

Pre-CenozoicMesozoic intrusive

Triassic Songpan - GarzeflyschSinian - Permian passivemarginPrecambrian metamorphicrocks

fold axis - upright

fold axis - overturned

fault - teeth on overthrust plate

range crest

Jiaocheng

Figure 3. Geologic map of the Min Shan and surrounding region depicting major structures bounding the Min Shan and their relationship toQuaternary sediments in the Min Jiang basin. Dash-dot line represents the approximate trace of the high topography along the range crest. Epi-centers and focal mechanisms of the 1976 Songpan earthquake swarm adapted from Jones et al. (1984). Map was compiled from 1:200000 geo-logic maps (Ministry of Geology and Mineral Resources, People’s Republic of China, 1991) and our work. Location of Figure 4 shown in box.

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380 Geological Society of America Bulletin, March 2000

Q1 conglomerates were deposited on a bedrocksurface of low relief in westward-thickeningwedges. Maximum preserved thickness recon-structed from the mapped geometries is nearly200 m at the western margin of the deposits. Clastcompositions are similar to Mississippian–Permianlimestones exposed in the Min Shan and distinctfrom the turbiditic flysch of the Triassic to the west(Chen et al., 1994). Clast imbrication within the de-posits indicates a component of west-directedtransport during deposition. The alluvial characterof the deposits, their geometry, clast composition,and transport directions all suggest that Q1 con-glomerates were derived from the Min Shan range,and we interpret them as remnant alluvial fansdraped along the west flank of the range.

These alluvial fans are now dipping west be-tween 12° and 16° (Figs. 4 and 6A) and havebeen deeply eroded by the modern streams drain-ing the Min Shan. In the northern end of thebasin, near the town of Gonganin, Q1 alluvialfans are deeply dissected into high remnant hillsand mesas (Fig. 6B). Streams draining this part ofthe Min Shan have incised through the Q1 and40–80 m into the underlying Paleozoic carbon-ates. Thus, total incision along tributaries to theMin Jiang exceeds 200 m since abandonment ofthe alluvial fans. The conglomerates were de-posited on a bedrock surface that currently dipswest at a slightly steeper angle than the Q1 beds(15°–18°) and is preserved along ridge crests eastof the mapped extent of Q1 deposits (Figs. 4 and6C). The age of this unconformity is unknown.

Q1 conglomerates are overlain at one localityeast of Zhangla (Fig. 4) by 10 m of slightly less in-durated gravel. The gravels are separated by a1–2-m-thick loess bed and have a slight (2°–5°)angular discordance across this bed. The loess hasa well-developed reddish weathering profile thatextends into the upper 1 m of the Q1 conglomer-ates beneath it. Small 0.5–1 cm clay nodules oc-cur at the contact of the loess and the Q1 con-glomerates. We interpret the gravels above thishorizon to be a younger, but related, deposit thatwe have mapped as Q1–2. The angular unconfor-mity and paleosol indicate that this horizon marksan important local hiatus in deposition. Beddingin the Q1–2gravels above the unconformity locallydips moderately to the east (Fig. 4), similar to theQ1 gravels below. Both units are folded on thewestern limb of an open syncline (Fig. 4) and thuspredate much of the folding.

Other deposits of probable Q1–2 age are pre-served throughout the southern and western partsof the Min Jiang basin. The deposits are generally

moderately well indurated gravels, which areslightly reddish in color with a clay-rich matrix.We correlate them with the Q1–2east of Zhanglaprimarily on the character of the deposit and thedegree of induration. We recognize the uncertain-ties in such a correlation, but rely on it in the ab-sence of material available for absolute dating. Itis possible that deposits mapped as Q1–2 span arelatively large time between the deposition of Q1and alluvial terraces of Q2 age.

Q1–2 deposits are also exposed immediatelybeneath the Min Jiang fault along the southwest-ern edge of the Min Jiang basin, west of the townof Zhangla (Fig. 4). They consist of two types ofintercalated deposits: rounded alluvial gravelssimilar to those described above and angularbreccias set in a red clay matrix. The gravelsstrike to the northwest and dip northeast. Thebreccias typically consist of clasts similar to theTriassic limestone west of the Min Jiang fault.They are interpreted as colluvial material shed offthe hanging wall of the fault. The presence of lo-cally derived talus within these deposits suggeststhat the former western margin of the basin wasnear its present-day position. A thin veneer ofloess and colluvium (Fig. 4) covers Q1–2depositsthroughout most of the southern Min Jiang basin.

Similar alluvial deposits crop out over a wideregion in the northwest corner of the Min Jiangbasin (west of Gonganin, Fig. 4). Gravels hereare moderately well indurated and consist of10–20 cm cobbles set in a reddish clay matrix.They are interfingered with breccia deposits sim-ilar to those exposed in the southwest. Exposurein this part of the basin is generally poor, butwhere we find exposures of the gravels, they dipmoderately to steeply west and are locally foldedbeneath the Min Jiang fault.

Late Quaternary Alluvial Terraces

Fluvial terraces preserved along the Min Jiangfrom the headwater region near Gonganin to thetown of Songpan (Fig. 2) consist of deposits oflate Pleistocene–Holocene age. Between Gon-ganin and Chuanjiusu, two levels of fill terracesare inset into the older Pleistocene deposits (Chenet al., 1994). The higher terrace (Q2) is preservedin discontinuous exposures only within the MinJiang basin, whereas the lower terrace (Q3) is con-tinuous between the headwater regions and thetown of Songpan.

The Q2 terrace consists of ~40 m of gravel andis best exposed in the northern part of the basin. Inthis region, Q2 gravels have only a moderately

well developed weathering horizon with a clay-rich matrix confined to the upper 1 m of the de-posit. Terrace surfaces preserved along the easternbank of the Min Jiang dip 2°–3°W and appear torepresent tributary fans draining into the MinJiang. Elsewhere in the northern basin, terracesurfaces are not preserved, but discontinuous ex-posures of similar alluvium suggest that they wereonce present along much of the river (Fig. 4).

The Q2 terrace is also exposed along the west-ern side of the Min Jiang in the southern part ofthe basin (Fig. 4). Here it consists of ~30 m ofalluvial gravel and has a colluvial wedge com-posed primarily of reworked loess draped atop it.The colluvial wedge contains rare boulders ofTriassic limestone and of older Pleistocene con-glomerate. The terrace is cut by the Min Jiangfault west of the town of Zhangla and providesan important relative timing constraint on theyoungest displacement along the fault (discussedsubsequently).

The Q3 terrace along the Min Jiang is continu-ous between Gonganin and Songpan (Fig. 4).Within the Min Jiang basin, the terrace consistsof 15–35 m of coarse-gravel fill inset into theolder deposits (Fig. 6A). The terrace has no ap-preciable soil development and is overlain byalluvial fans whose sources were the main tribu-taries in the basin. In the center of the basin, thisterrace is locally a strath bench inset into a largelandslide that fills much of the valley near theGamet monastery (Fig. 4). The Q3 terrace locallyonlaps the Min Jiang fault and thus places an up-per bound on the timing of displacement acrossthe fault (discussed subsequently).

Despite the continuity of the Q3 terrace acrossthe southern margin of the Min Jiang basin, thenature of the terrace changes south of Chuan-jiusu. Terraces within the Min Jiang basin are en-tirely alluvial fill, whereas south of the basin mar-gin, the Q3 terrace is a strath terrace whose treadranges from 5 to 10 m above the modern channel.The Q3 terrace along this reach generally has acap of alluvial gravels that ranges from 5 to 15 min thickness and is overlain by tributary alluvialfans. South of Songpan, the terrace is not present,but tributary alluvial fans are ubiquitous, andtheir distal parts are generally truncated by recentincision of the Min Jiang. These fans are proba-bly contemporaneous with alluvial fans pre-served atop the Q3 terrace.

Modern alluvium is confined to the bed of theMin Jiang and its tributaries, consists of gravel- toboulder-sized clasts of carbonate, and is mappedas Q4 (Fig. 4).

Figure 4. Geologic map of the Min Jiang basin depicting the relationships between the Min Jiang fault and Quaternary sedimentary depositsalong the west flank of the Min Shan. Note the lack of faults east of the basin proper. Topographic base was originally 1:100000. Inset is an equal-area, lower-hemisphere projection of poles to bedding in Q1 conglomerate showing its uniform west-northwest dip.

Geological Society of America Bulletin, March 2000 381

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Min JiangFault

Chuanjiusu

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Explanation

Q4 alluvium

Q4 colluvium

Q3 alluvial fan

Q3 alluvium

Q2 alluvium

Q1-2 alluvium

Q1 conglomerate

Tp conglomerate

Loess - undivided

Landslide

M - Tr bedrock

Maur Ji Hong Xin Yan Shan (~5100m)

10

thrust fault - dashedwhere approximate

fold axis

bedding orientation

town/village

preserved extent oflow-relief surface

Symbols

Q1

Q1-2 Q1

Q1

Q3

Q2

Q3af

Q3

Q3

Q2

Tp

Q3af

contour interval = 200 m

N

Poles to bedding - Q1 conglomerates

n = 49

mean strike and dip:040, 14W

KIRBY ET AL.

382 Geological Society of America Bulletin, March 2000

CHRONOLOGY OF THE MIN JIANGBASIN DEPOSITS

In order to assess the absolute age of the de-posits within the Min Jiang basin we collectedsamples for 14C and photon-stimulated lumines-cence (PSL) dating. The latter technique can be asensitive means for dating the burial time of de-trital grains of sediment (e.g.,Aitken, 1998). Theparticular approach employed here (IRSL, or in-frared-stimulated luminescence) produces a sig-nal from fine, silt-sized feldspars. The techniquerelies on an assumption that the natural lumines-cence of detrital grains was zeroed by exposureto sunlight prior to final deposition. Although theinsolation histories are unknown, we selectedsamples from facies most likely to representdepositional settings favorable to luminescenceresetting (e.g., eolian silt, laminated lacustrinesilt). Furthermore, the IRSL method is extremelysensitive to daylight; as little as a few minutes in-solation can remove >90% of the IRSL signature(Aitken, 1998). IRSL sampling and analyticaltechniques are presented in the Appendix.

14C Chronology: Q2 and Q3 Terraces

Detrital charcoal from fine-sand beds near thetop of the Q3 terrace north of Zhangla (Fig. 4) pro-vides age constraints for the timing of displace-ment on the Min Jiang fault. Analytical data andages are reported in Table 1. Sample EK 97-9(angular fragment of charcoal 1–2 cm in diam-eter) was collected from a fine-sand layer in-terbedded with gravel ~1 m below the surface ofthe Q3 terrace and gave a calibrated age of8129–8420 yr B.P. (2σ). Sample EK 97–8 (multi-ple detrital charcoal fragments <1 mm in diam-eter) was collected from a similar sandy lenswithin the terrace gravels (~0.5 km south of 97–9)and yielded an age of 6670–6944 yr B.P. (2σ). Wealso collected a single sample of charcoal-rich soil(EK 97–10) from a fine-grained alluvial fan atopthe terrace at Zhangla (Fig. 4). Conventional 14Canalysis on the bulk organic fraction gave an aver-age age of 2742–2993 yr B.P. (2σ) that places aminimum bound on the age of the terrace. We ob-tained a fourth sample (EK 97–26) of wood fromwithin a colluvial and landslide deposit drapedatop the Q3 terrace just west of the town ofChuanjiusu (Fig. 4). This sample yielded a con-ventional 14C age of 14 118–15 720 yr B.P. (2σ),suggesting that some of the Q3 terrace gravelswere deposited in the latest Pleistocene. Althoughit is possible that this wood may be recycled tosome degree, we find it unlikely that woody ma-terial would remain intact for many thousands ofyears at the surface, and we interpret this age asthe approximate time of burial of wood within theslide deposit. Taken as a whole,14C chronology

suggests that the Q3 terrace gravels range in agefrom >14000 yr to 6600 yr B.P.

Although we were unable to obtain datable ma-terial from the Q2terrace proper, relationships withalluvial sediments of late Pleistocene age nearNingpo (Fig. 4) suggest that the terraces probablydate to the late Pleistocene. We obtained detritalcharcoal fragments from east-dipping sands andsilts north of the town of Ningpo (Fig. 4). The sed-iments interfinger with coarse gravels near themodern Min Jiang, and we interpret them as fine-grained alluvial-fan deposits transitional with Q2fluvial terraces. Although this correlation is some-what tentative, the unconsolidated nature of thesediments and the presence of a thin (<5 m) loesscap are consistent with relationships observed onwell-preserved Q2 terraces along the western sideof the Min Jiang. Sample EK 96-1 yielded a 14Cage of 22 780 ± 340 yr B.P. (1σ), which impliesthat Q2 deposits date to the late Pleistocene.

IRSL Chronology: Q1 Deposits

We also collected three samples for lumines-cence dating (IRSL method). Two of these sam-ples were collected from loess and silt inter-bedded with the Q1 conglomerates and placeimportant bounds on the age of these deposits.A third sample, collected from young lacustrinesediments along the Min Jiang (north of Jiao-cheng; Fig. 3), allowed us to test the IRSL multi-aliquot zeroing assumption. This assumptionstates that all grains were exposed to daylight formore than tens of minutes prior to final deposi-tion. Analytical techniques for all three samplesare presented in the Appendix, dosimetry data arepresented in Table 2, and luminescence data arepresented in Table 3.

All IRSL experiments employed a two-tem-perature preheating protocol (Table 3). Separatebatches of aliquots (discs, each having ~0.5 mgof grains on an Al substrate) were given differentheat treatments as reported in Table 3. The ration-

ale for this protocol is that statistical agreementof the paired equivalent dose (DE) values is ex-pected to indicate that a thermally stable IRSLsignature has been isolated. In this case, theweighted mean DE value is used in the age calcu-lation (Table 3).

Examples of the IRSL data are shown in Fig-ure 7. Dose-response curves (center) illustrate thereproducibility of individual data points (one foreach aliquot). The IRSL signatures (after preheat-ing) for all samples were not large and evincedrelatively large scatter. For sample TIBT 98-2(EK 98-5) and the 170 °C experiment of sampleTIBT 98-4 (EK 98-9), short-shine normalization(Ollerhead et al., 1994) did not reduce the scattersignificantly, but for the other samples, normal-ization improved the precision of the extrapolatedDE values. A weighted mean of the plateau seg-ments in Figure 7A was used in age calculation,though the apparent (at 1σ) decline in DE valuewith higher preheating is not understood.

Sample EK98-5 (TIBT 98-2) was collectedfrom laminated lacustrine sediments along theMin Jiang, ~50 km south of the Min Jiang basin.Lacustrine sediments are interbedded with tribu-tary alluvial fans along the river, which are cor-relative with fans of Q3 age in the headwaterreaches (Fig. 4). The resulting apparent age of14.2 ± 1.7 ka (Table 3) broadly agrees with 14Cages up to ~15 ka for Q3 terrace gravels in theheadwaters of the Min Jiang and lends a measureof confidence to our IRSL results. Silts within thelacustrine deposits appear to have been effec-tively zeroed by sunlight during transport andresidence in the water column.

Two samples were collected from the Q1 de-posits (Fig. 5). Sample EK 98–9 (TIBT 98-4)was collected from fine silts and reworked loessnear the base of Q1 (~15 m above the basal con-tact with bedrock). The sample yielded an IRSLage of 254 ± 35 ka. A second sample (EK 98-6 /TIBT 98-3) collected from the 1-m-thick loesspaleosol atop Q1 conglomerates near Zhangla

TABLE 1. RADIOCARBON ANALYSES OF CHARCOAL SAMPLES

Sample Stratigraphic δ13C 14C age 14C age (corrected)* Calendric age rangenumber unit (%) (yr B.P. ± 1σ) (yr B.P. ± 1σ) (yr B.P. ± 2σ)EK 97-10† Q4 fan –25.7 2730 ± 75 2810 ± 77 2742–2993EK 97-8§ Q3 terrace –23.4 5945 ± 55 6120 ± 57 6670–6944EK 97-9§ Q3 terrace –12.5 7525 ± 55 7750 ± 57 8129–8420EK 97-26† Q3 colluvium –22.1 12560 ± 180 12940 ± 185 14118–15720EK 96-1§ Q2 alluvium –25.9 22780 ± 340 23460 ± 350 N.D.

Note: Samples are listed in stratigraphic order (see Fig. 5 for relative stratigraphic sample locations). Reported14C ages use Libby’s half life (5568 yr) and are referenced to the year A.D. 1950. Analytical uncertainties are 1σand reflect the total uncertainty in the measurement. Calendric ages calibrated using the University of WashingtonCALIB program (Stuiver and Reimer, 1993). N.D.—not determined.

*The corrected age is the due to the 3% difference between the Libby half life of 14C (5568 yr) and the actualhalf life (5735 yr).

†Conventional 14C measurement performed at Geochron Laboratories, Cambridge, Massachusetts.Sample 97-10 consists of a charcoal-rich organic soil horizon. The clay/organic fraction was isolated by ultrasounddispersion, filtered to remove rootlets, and treated with dilute HCl to remove carbonate. Sample 97-26 consists ofwood. It was washed, split into small pieces, treated with HCl to remove carbonate, and with NaOH to removehumic acids.

§AMS measurement performed by G. Burr at the University of Arizona AMS facility.

NEOTECTONICS OF THE MIN SHAN, CHINA

Geological Society of America Bulletin, March 2000 383

(Fig. 4) yielded a luminescence age of 157 ±28 ka. Without independent age determinations,the IRSL results must be taken as a best approx-imation of the depositional age. However, wenote that both of our age determinations fallwithin known periods of loess deposition innorth-central China (Kukla, 1987). Together,these samples indicate that most of the sedimen-tation in the Min Jiang basin took place in a shorttime period between ca. 250 ka and 160 ka.

Chen et al. (1994) reported two thermolumi-nescence (TL) ages of 158 ± 12 ka and 830 ±69 ka for samples from the Guyian Shan Forma-tion (mapped as Q1; Fig. 4). The first sample wastaken “from the base of the loess” covering theformation and therefore may be correlative withthe paleosol between Q1 and Q1–2(157 ± 28 ka).Their second sample was taken “50 m down inthe Guyian Shan Formation.” Chen et al. (1994)provided no information about the material in

the horizon sampled, except that formation “iscomposed of thick beds (1–3 m) of conglomer-ate interbedded with thin (10–30 cm) beds ofsandstone.” Considering the likelihood of rapid(poor TL-zeroing) deposition in these types ofdeposits and the general question of the upperlimit of TL dating techniques (Berger, 1995;Prescott and Robertson, 1997), we cannot assessthe accuracy of the older TL age, except that it islikely an overestimate. Our ages are also signifi-cantly younger than ages typically assigned to“Q1” deposits throughout southwestern China(ca. 1.6–0.6 Ma) (Ministry of Geology and Min-eral Resources, People’s Republic of China,1991), which suggests caution in interpreting ab-solute age from regional correlation of terrestrialQuaternary deposits in western China.

Our results further suggest that periods ofdeposition were tied to glacial-interglacialcycles. The onset of Q1 deposition occurred

shortly prior to ca. 250 ka and appears to coin-cide with oxygen isotope stage 8 (Imbrie et al.,1984). Loess deposition between Q1 and Q1–2(ca. 160 ka) coincides with the stage 6 termina-tion (Imbrie et al., 1984), and this loess may becorrelative with the L2 loess of the Upper LishiFormation in well-studied sections on the LoessPlateau (Kukla, 1987). Aggradation of Q2 flu-vial terraces is not well dated in the Min Jiangbasin, but may date to ca. 22 ka (e.g.,14C fromNingpo sediments), near the time of the lastglacial maximum (Imbrie et al., 1984). Aggra-dation of the Q3 terrace occurred betweenca. 15 ka and ca. 6–7 ka, coincident with isotopestage 2. Thus, fill-terrace development along theheadwater reaches of the Min Jiang was appar-ently coincident with the last glacial period. In-cision of the Min Jiang into the Q3 terrace grav-els appears to postdate 6 ka. These results arestriking in their correlation to late Pleistoceneclimatic oscillations (Imbrie et al., 1984) andsuggest that late Quaternary terrestrial deposi-tion in the Min Jiang basin was closely coupledto changing sediment and water flux duringglacial-interglacial cycles.

QUATERNARY FAULTING IN THE MIN SHAN

Quaternary deposits within the Min Jiang basinare variably deformed (Fig. 4) and place impor-tant constraints on the kinematics and rates ofQuaternary deformation in the Min Shan. We firstconsider the recent and active displacement alongthe Min Jiang fault and then examine several im-portant structures within and east of the range.

Min Jiang Fault

The Min Jiang fault zone consists of severalstrands of high-angle reverse faults that strikeapproximately due north and parallel the western

TABLE 2. DOSIMETRY DATA FOR IRSL SAMPLES FROM THE MIN JIANG BASIN

Sample Water* K2O Ct§ Cth

§ Th U b value# Dose rate**

(TIBT) (wt%) (ks–1·cm–2) (ks–1·cm–2) (ppm) (ppm) pGy·m2 (Gy/k.y.)(±0.05)

98-2 0.10 ± 0.05 2.13 0.712 ± 0.010 0.403 ± 0.037 10.84 ± 0.98 2.46 ± 0.30 0.923 ± 0.091 3.82 ± 0.232.18† 0.674 ± 0.010 0.285 ± 0.020 7.68 ± 0.70 3.03 ± 0.20

98-3 0.18 ± 0.03 2.43 0.8540 ± 0.0098 0.495 ± 0.036 13.32 ± 0.97 2.81 ± 0.29 1.0 ± 0.1 4.28 ± 0.162.85 0.836 ± 0.010 0.469 ± 0.033 12.6 ± 1.0 2.87 ± 0.20

98-4 0.10 ± 0.05 1.05 0.4858 ± 0.0079 0.164 ± 0.022 4.41 ± 0.59 2.51 ± 0.18 1.0 ± 0.1 2.83 ± 0.192.26† 0.710 ± 0.048 0.323 ± 0.042 8.7 ± 1.2 3.03 ± 0.20

*Estimated average historic ratio of weight of water/weight of dry sample, based on measurements of moist collected samples and laboratory saturatedsamples (3 samples per IRSL sample). Uncertainties here and elsewhere are ±1σ.

†From top to bottom, estimated uncertainty in these values is ±0.09 and ±0.25.§Total and thorium count rates from finely powdered samples for thick-source-alpha-particle-counting (TSAC) method (Huntley and Wintle, 1981).

Cu = Ct – Cth. Equations for the use of these count-rate values in the calculation of dose rate are given by Berger (1988). A second row of data corresponds to the average values from sediment surrounding the IRSL sample at a radius of 12–23 cm. These data are used to calculate the γ dose-ratecomponent.

#Alpha effectiveness factor (Huntley et al., 1988; Berger, 1988). Values of 1.0 are estimated (from this and other projects).**Calculated with the conversion factors and equations given by Berger (1988), and includes a cosmic ray component varying from 0.04 to 0.14 with es-

timated average depth, from the data of Prescott and Hutton (1988). Note that the dose rates are typical of most sediments (2–4 Gy/k.y.).

TABLE 3. IRSL DATA AND AGES FOR LUMINESCENCE SAMPLES

Sample Preheat* DE† Time§ Age

(TIBT) (Gy) (s) (ka)98-2 135 °C/2d 61.3 ± 4.1 1–40

155 °C/2d 49.3 ± 3.4 1–40Weighted mean DE = 54.2 ± 5.9 14.2 ± 1.7

98-3 150 °C/2d 710 ± 180 20–60170 °C/2d 640 ± 160 1–35

Weighted mean DE = 670 ± 120 157 ± 2898-4 140 °C/2d 730 ± 130 1–15

160 °C/2d 710 ± 130 1–15Weighted mean DE = 719 ± 92 254 ± 35

Note: The polymineralic 4–11 µm size fraction was used for all IRSL measurements. IRSL was detected at the 420 ± 20 nm spectral region (bandpass 390–470 nm at 1% cut). All bleaching was for 3 hr with >780 nm solar spectrum passed (based on the bleaching protocol introduced by Huntley and Clague, 1996). Applying the heat treatment after bleaching corrects for any possible thermal-transfer effect (Ollerhead et al., 1994), an apparently negligible effect with these samples.

*The chosen prereadout annealing. For all samples, the weighted mean DE isused for age calculation (see text).

†Weighted mean equivalent dose plus average error over time interval in next column. A weighted-saturating-exponential regression and error model (Berger et al.,1987) was employed for all samples. For some samples, interaliquot scatter was minimized by short-shine normalization (to natural signals) (Ollerhead et al., 1994).

§The readout (LED-on) time interval for which DE is calculated (e.g., Fig. 7).

KIRBY ET AL.

384 Geological Society of America Bulletin, March 2000

side of the Min Shan for nearly 100 km. Theeasternmost strand bounds the western margin ofthe Min Jiang basin and juxtaposes Triassiclimestone and flysch with terrace deposits alongthe Min Jiang river (Chen et al., 1994; Tanget al., 1993; Zhao et al., 1994). Although manyof these authors consider the fault to be active,recent GPS results from a station west of thefault zone (station SBP2; Fig. 2) suggest thatshort-term displacement rates are less than sev-eral millimeters per year and within uncertaintyof zero (Chen et al., 2000; King et al., 1997). Inaddition to uncertainty regarding the Holoceneslip rates of the fault, previous workers disagreedon the sense of displacement. All have agreedthat there is a strong component of shorteningacross the fault, but the strike-parallel compo-nent has been interpreted as both dextral andsinistral (cf. Chen et al., 1994; Zhao et al., 1994).We mapped the easternmost segment of the fault(herein referred to as the Min Jiang fault, sensustricto) from north of Gonganin (Fig. 4) to thetown of Songpan (Fig. 3).

The southern part of the Min Jiang fault fol-lows the eastern side of the Min Jiang valley be-tween the towns of Chuanjiusu and Songpan(Figs. 3 and 4). The fault, although not well ex-posed, offsets tributary drainages and gulliesalong the eastern valley wall. Seven of nine trib-utary gullies along this stretch of the fault appearto be deflected in a manner consistent with acomponent of left-lateral displacement alongthis segment of the fault (Fig. 6D). We were un-able to map the fault south beyond Songpan, al-though a series of northwest-striking folds in theTriassic appears to be continuous across the pro-jection of the Min Jiang fault (Fig. 3). Thus, weinfer that the displacement on the fault dimin-ishes along strike to the south or is transferred todistributed shortening.

Along the southern part of the Min Jiangbasin, a north-striking, west-dipping splay ofthe Min Jiang fault (herein named the Ningpofault after a nearby village and monastery) jux-taposes fine-grained sands and gravels in itshanging wall (Q2 in Fig. 4) with the Q1 con-glomerates to the east. Gravels west of this faultare warped into an asymmetric anticline with anorth-trending axial trace (Fig. 4). The Ningpofault can be followed as a discrete break be-tween the Q1 and Q2 deposits for ~1 km to thenorth, whereupon the trace of the fault transfersinto the axial trace of an open syncline (Fig. 4)within the Q1 gravels. The geometry of the fold-ing within both the Q1 and Q2 deposits suggeststhat the fault has a component of west-side-up,reverse displacement, akin to the Min Jiangfault. It is noteworthy that this is the only part ofthe basin where a fault coincides with the east-ern basin margin.

The Ningpo fault is joined by an east-trend-ing fault east of Chuanjiusu (Fig. 4). This faultdips ~60°N and is marked by a 10–20-cm-widegouge zone within the Q2 sediments in the val-ley bottom. The separation on the fault appearsto be north side up, and the fault may have aright-lateral component. The fault appears totransfer displacement between the Min Jiangand Ningpo faults, some of which may be ac-commodated by folding of Q2 sediments west ofthe Ningpo fault.

The Min Jiang fault itself is best exposed in asmall drainage west of the town of Zhangla, inthe southern end of the Min Jiang basin (Fig. 4).Here the fault dips 55°–65°W and places brec-ciated Triassic limestone above Q2 terrace grav-els. A 1–3-cm-wide zone of clay-rich red gougemarks the fault zone. The Q2 terrace is inset alongthis segment of the Min Jiang valley into olderQ1–2 gravels. These deposits are mixed alluvialgravels and breccias of angular limestone clastsderived from the Triassic in the hanging wall ofthe Min Jiang fault.

Approximately 4 km north of Zhangla, theQ1–2gravels pinch out, and the fault places lime-stone directly over the Q1 conglomerate. Thefault itself is not well exposed along this stretch,but its mapped trace suggests that it dips moder-

ately to the west (Fig. 4). From this point north,there is little evidence for recent displacementsalong the Min Jiang fault. We infer that the traceof the fault parallels the linear range front be-tween Gamet and Kakagou (Fig. 4); however, ex-posure is poor along the steep forested slopes. Onthe western side of the ~4000 m ridge betweenGamet and Kakagou, an east-dipping splay of thefault juxtaposes the Triassic with Pliocene con-glomerates (Chen et al., 1994), indicating thatthis strand of the fault zone is relatively young.However, we were unable to find any scarps ormorphologic evidence that this splay is active atpresent. The Triassic limestone between the twostrands is interpreted as a pop-up block within theMin Jiang fault zone (Fig. 4).

North of Kakagou, a second strand of theMin Jiang fault zone bounds the western edgeof the Min Jiang basin (Fig. 4). This fault iswell exposed in the drainage west of Gonganin,dips ~45°W, and places Triassic limestonesabove Q1–2 alluvium. The Q1–2 gravels are ex-posed over ~15 km2 along the western side ofthe Min Jiang valley headwaters. Gravelswithin this subbasin generally dip moderatelywest and are locally folded into an overturnedfootwall syncline beneath the western faultstrand. They appear to be juxtaposed with Q1

(

Triassic limestones and marls

Pliocene conglomerates (50-70 m)

Q1 conglomerates (180-200 m)

Q1-2 gravels (10-100 m)

Q2 terrace gravels (40 m)

Q3 terrace gravels (15-35 m)Q4 fine grained alluvial fans

EK 98 - 6 (157 +/- 28 ka)

EK 97-8 (5945 +/- 55 yr BP)EK 97-10 (2730 +/- 75 yr BP)

EK 97-9 (7525 +/- 55 yr BP)

loess (1-2 m)

EK 97-26 (12560 +/- 180 yr BP)

EK 98-9 (254 ± 35 ka)

Figure 5. Schematic composite stratigraphic column of Tertiary–Quaternary sediments of theMin Jiang basin. Coarse pattern represents cobble to boulder conglomerates along Min Jiangfault (white) and Q1 alluvial-fan conglomerates (gray). Loess interlayer is shown as heavy blackline. Fluvial-terrace gravels of modern Min Jiang drainage (Q2 and Q3; fine-dot pattern) areoverlain by Q4 alluvial fans (ruled pattern). Approximate sample locations are shown along theleft side of the column (radiocarbon—circles, IRSL—triangles).

Geological Society of America Bulletin, March 2000 385

Q1

Q4

Q3

A B

C D

EFigure 6. Field photographs of sediments within the Min Jiang basin

(A–C) and of the Min Jiang fault (D, E). (A) Looking north at the cen-tral Min Jiang basin. West-dipping ridges in the background are com-posed of Q1 conglomerates dipping ~14°. Alluvial sediments (latePleistocene-Holocene) of the Min Jiang river (Q3, Q4) fill the valley inthe near midground. (B) View looking north-northeast at remnant Q1(mesalike hill in midground) in the northern Min Jiang basin. Hill is~170 m high. Ridge in background is composed of west-dipping Trias-sic limestones. (C) Tilted bedrock surface beneath basin extends alongridge tops toward crest of range. Photograph is looking south along theeastern edge of the Q1 conglomerates. Snow-capped peak is Hong XinYan Shan (~ 5100 m). (D) Shutter ridge along the southern Min Jiangfault indicative of a component of left-lateral displacement. Trace offault shown in white. The terraced fields at the base of the ridge are ona young alluvial fan built across the fault trace. View is to the east.(E) View to the southeast of high erosion surface west of the Min Jiangfault (highlighted by arrow). Min Jiang fault runs along base of ridges(dashed white line). The flat, linear ridges in the foreground at left areQ2 terraces preserved in the headwaters of the Min Jiang.

conglomerates exposed on high hilltops east ofthe Min Jiang. However, the contact is ob-scured by modern alluvium. We infer from themapped geometries that the eastern strand ofthe Min Jiang fault forms the eastern boundaryof this subbasin. The exposure in the steepforested hillslopes north of the drainage divideis poor, and consequently the northern extent ofthis structure is uncertain. However, the west-ern strand of the Min Jiang fault clearly contin-ues to the north, where it juxtaposes Triassiclimestones and Pliocene gravels (Fig. 4). Ourinterpretation of these geometries is that muchof the displacement on the eastern Min Jiangfault transfers to the western strand across thesubbasin of Q1–2gravels.

Long-Term Displacement on the Min JiangFault

The morphology of the Tibetan Plateau to thewest of the Min Jiang fault places loose boundson the total displacement across the fault.Preserved remnants of a low-relief erosion-surface cap ridge crests west of the Min Jiangfault (Figs. 4 and 6D). This surface is ~200–300m above the modern valley of the Min Jiang.Although bedrock surfaces are continuouslyevolving elements in an active landscape, thelow relief and partial preservation of the surfacesuggest that erosion rates on the surface properare probably low. We interpret the surface as thewestern continuation of the low-relief surfacethat floors the Min Jiang basin. This correlationmust be considered tentative, but, if correct,implies a net throw on the Min Jiang fault ofbetween ~350 and ~750 m. Thus, the Min Jiangfault does not appear to have a significantamount of displacement across it in the lateCenozoic.

Timing of this displacement is, however, un-certain. The age of the erosion surface in the MinShan is essentially unknown. Our mapping indi-cates that it must be older than the Q1 conglom-erates (ca. 250 ka), but its relationship to thePliocene(?) deposits is unclear. Regionally, muchof the surface of the northeastern Tibetan Plateauis morphologically similar: a low-relief surfacecut on vertically foliated Triassic flysch. How-ever, the few Tertiary sediments that exist are notwell dated. Consequently, all we can determine isthat reverse-sense displacement on the Min Jiangfault appears to be late Cenozoic, and the initia-tion of movement likely predates the depositionof Q1. If we assume that the Pliocene basinsalong the fault zone mark the onset of displace-ment, long-term slip rates must be significantlyless than 1 mm/yr. We conclude that shorteningacross the Min Jiang fault is likely to have beenrelatively minor during the Quaternary.

Terrace Profiles and Recent DisplacementsAlong the Min Jiang Fault

The modern Min Jiang and associated Q3(Holocene) terrace cross the Min Jiang fault inthree places within the study area (Fig. 4). In eachlocality there are no visible breaks or offsets in theterraces. Although this continuity suggests thatthere has been little displacement on the fault inHolocene time, we surveyed the modern channeland terrace profiles to assess whether low rates ofmotion on the fault were manifested in changes inthe gradient of the channel system and/or terraces(e.g., Pazzaglia et al., 1998). Details of the sur-veying procedure are given in the Appendix.

The modern channel and Q3 terrace surfaceonlap the fault just southwest of the town ofMaur Ji (Fig. 4). The terrace gravels are unusu-ally well indurated here, and the channel is fairlynarrow. However, there are several cold springsand associated tufa mounds that mark the trace ofthe fault, and we infer that the competence of thegravels in this locality reflects the local ground-water flux. The terrace surface does not appear tobe deformed within the resolution of our survey(~2 m; see Appendix). The terrace surface isnearer to the modern channel along this reach ofthe Min Jiang (Fig. 8), however, and we interpretthis proximity as reflecting the increased streamgradient as the river flows over resistant lime-stone. Approximately 5 km north, the fault cross-es back to the western side of the Min Jiang and

is buried beneath a large landslide mass just southof the Gamet monastery (Fig. 4). This slide massis composed of angular, shattered blocks of Trias-sic limestone, and onlap relationships betweenQ3 terraces on either side of the slide mass indi-cate that the slide predates the terrace. Both ofthese relationships suggest that the Min Jiangfault has had little displacement along it duringHolocene time.

South of Chuanjiusu, the fault crosses the riverfor a third time and is also obscured by a smalllandslide. There is a small strath terrace beveledinto the toe of the slide, which appears to be con-tinuous with the Q3 terrace. Previous workers sug-gested that the Q3 terraces were unusually highalong this reach of the Min Jiang, a fact they at-tributed to recent displacement on the adjoiningXue Shan fault (Chen et al., 1994; Tang et al.,1993); our survey fails to confirm this conclusion.The Q3 terrace is continuous between Chuanjiusuand Songpan (Fig. 8). Several large alluvial fansonlap the terrace surface near the mouths of largertributary streams. Although there is significant re-lief (up to 80 m) on the fan surfaces, they do notrepresent warping of the main terrace tread.

Mapped relationships between the Min Jiangfault and terraces along the Min Jiang indicate thatthe youngest displacement on the fault is brack-eted between deposition of the Q2 gravels and theQ3 terrace. There does not appear to have beensignificant displacement on the fault during thepast ~7000 yr. Given the uncertainties of our sur-

KIRBY ET AL.

386 Geological Society of America Bulletin, March 2000

Applied β dose (100 Gy)0 2 4 6 8

Pho

ton

coun

ts/5

s (1

000)

0

5

10

15

20

TIBT98-2

IRSL at 0-5s

135°C preheat

DE

4-11 µm

Time (s)

0 20 40 60

DE (

Gy)

0

300

600

900

1200

N+Bleach+135ûC140°C160°C

Time (s)0 20 40 60 80

DE (

Gy)

0

20

40

60

80

A

B

135°C155°C

TIBT98-2

TIBT98-4

Figure 7. Illustrative luminescence (IRSL) dose-response curve (center) and plots of equiva-lent-dose (DE) vs. readout time (inserts A and B). A weighted-saturating-exponential regressionmodel was used for all samples. Error bars on the DE plots are ±1σ. Equivalent-dose data for155 °C (inset A) and 160 °C (inset B) are shifted by +2.5 s for clarity of plotting.

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Geological Society of America Bulletin, March 2000 387

vey (±2 m), there is no appreciable deformation ofthe Q3 terrace, requiring Holocene displacementrates to have been less than ~1 mm/yr. This con-clusion is consistent with the short-term rates ob-served in the geodetic data (King et al., 1997).However, the Min Jiang fault cuts Q2 (late Pleis-tocene) gravels west of Zhangla (Fig. 4). These re-lationships are consistent with displacement oc-curring during moderate to large seismic eventswith long recurrence intervals (>7000 yr). In ad-dition, the transition from alluvial to strath ter-races south of Chuanjiusu coincides with thepoint where the valley of the Min Jiang is on thehanging-wall block and may reflect bedrock inci-sion in response to slightly increased rates of rockuplift. Although this incision may be an indicationof activity on the Min Jiang fault, a quantitativeassessment of this hypothesis is beyond the scopeof this paper.

Huya Fault

The structural position of the Huya fault alongthe eastern flank of the Min Shan (Fig. 3) and thehistoric seismic activity along it (Jones et al.,1984) suggest that this fault may play an impor-tant role in active deformation along the MinShan. However, the mapped extent of the fault(Ministry of Geology and Mineral Resources,

People s Republic of China, 1991) along the cen-tral ~60 km of the range (Fig. 3) suggests that dis-placement diminishes along strike or is trans-ferred to distributed deformation. In order toassess these possibilities, we examined the Huyafault in reconnaissance fashion along much of itslength. The central segment of the fault crops outalong the Fu Jiang, where it dips steeply west(70°–90°) and juxtaposes Paleozoic limestonesof the Xue Shan platform with metamorphicrocks mapped as Precambrian (Fig. 3). These ge-ologic relationships (younger over older) are un-usual for a reverse fault and suggest that the Huyafault may have a pre-Cenozoic ancestry. Alongthe valley of Fu Jiang, the fault cores a steeprugged canyon and is spatially associated withnumerous rockfalls and landslides, the most re-cent of which are presumably the result of the1976 earthquakes (Jones et al., 1984).

The Huya fault terminates to the north againstthe Xue Shan fault and to the south against theQingchuan fault (Fig. 3). These relationships atthe terminations of the Huya fault have importantimplications for the transfer of displacementalong the eastern range front of the Min Shan. Al-though it is possible that the Huya fault is a blindstructure along the range front north of the XueShan fault, the apparent continuity of Paleozoicand Mesozoic units across the range front (Fig. 3)

suggests that the total amount of displacementmust be relatively small. This geometric relation-ship at the north end of the Huya fault requiresthat displacement across it must decrease alongstrike to the north or be accommodated by dis-tributed deformation between the Xue Shan andTazang faults (Fig. 3).

A similar relationship exists at the southern endof the Huya fault where Triassic rocks west of thefault are folded around its southern terminationand are continuous with Triassic rocks along theQingchuan fault (Fig. 3). Although the Qingchuanfault makes a remarkably linear trace on satelliteimages (Fig. 2), geodetic results (Chen et al.,2000; King et al., 1997) suggest that active dis-placement on the Qingchuan fault is restricted toa few millimeters per year of right-lateral dis-placement. Our reconnaissance observations ofundeformed Pleistocene terraces along theFu Jiang (E. Kirby, unpublished data) suggest thatshortening is negligible. Although the Qingchuanfault may represent a component of right-lateralshear northeast of the topographic margin of theplateau, it does not appear to absorb displacementalong the Huya fault. Thus, displacement alongthe southern Huya fault is likely transferred tosmall, discontinuous structures of similar orienta-tion in the southern Min Shan (Fig. 3).

Xue Shan Fault

The Xue Shan fault is an west-trending zone offaults that terminate against the Min Jiang faultnorth of the town of Songpan (Figs. 3 and 4). Thefault extends beyond the modern topographicmargin of the plateau into the foreland, where itcarries Paleozoic rocks in both its hanging walland footwall (Fig. 3). These geometries suggest,but do not require, relatively minor displacementacross the fault. At its eastern end, the Xue Shanfault is intruded by a Mesozoic pluton, indicatingthat much of the displacement predates Mesozoicmagmatism (Ministry of Geology and MineralResources, People’s Republic of China, 1991).We were unable to verify proposed recent activ-ity along the Xue Shan fault (Tang et al., 1993)near the crest of the range. Such displacement, ifpresent, appears to be limited, and we interpretthe Xue Shan fault as a Mesozoic structure thathas played a minor, if any, role in Cenozoic de-formation within the Min Shan.

RATES AND PATTERNS OF SURFACEDEFORMATION IN THE MIN SHAN

In spite of the apparently minor displacementalong the Min Jiang and other faults west of theMin Shan, there is abundant evidence for tiltingof Pleistocene–Holocene markers within therange. In this section we describe tilting recorded

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Figure 8. Longitudinal profile of the Min Jiang (black) and Q3 terrace surface (gray) surveyedbetween Gonganin and Songpan. Gray dots represent discontinuous Q2 terrace surfaces. Ar-rows show the intersection between the trace of the Min Jiang fault and the river channel. Notethe smooth transition of the Q3 terrace across the northern two intersections. Downstream of thefinal intersection, the channel steepens and the terrace undergoes a transition to a bedrockstrath bench ~1–3 m above the river. This change may reflect slight differential rock uplift be-tween the hanging wall and footwall of the Min Jiang fault.

in two places within the western Min Shan: (1) inyoung lacustrine sediments along the Min Jiangand (2) within the Min Jiang basin. In both local-ities, tilting is present over 10–20 km wave-lengths and does not appear to be directly associ-ated with exposed faults.

Recent Tilting in the Southern Min Shan

Recent tilting is recorded in the southwesternMin Shan in lacustrine sediments north of theDiexi landslide. The Diexi landslide is a large(~1 km3) slide mass that dammed the Min Jiangdrainage near the village of Jiaocheng (Fig. 3).This site is located near the epicenter of a moder-ate-sized earthquake in 1933 that initiated a rock-slide obliterating the village of Diexi (Chen et al.,1994). The prehistoric slide, however, is pre-served on both sides of the valley and appears tohave dammed the entire Min Jiang. The surfaceof the slide is ~200 m above the current riverlevel and preserves characteristic hummockytopography, draped with ~1 m of loess. Somelarge boulders are visible atop the slide mass.

Deposits of lacustrine sediments as thick as30 m occur upstream from the landslide mass fornearly 25 km along the Min Jiang. These depos-its are typically preserved beneath abandoned al-luvial fans debauched from side-tributary drain-ages, although drapes of sediment can be seenhigh on a few hillslopes. We correlate these fanswith the Q3 tributary alluvial fans in the MinJiang basin (Fig. 4). The lacustrine sediments areextremely fine-grained silts and muds, have acharacteristic yellow color, and are generallylaminated on a millimeter scale. The originalthickness of many of the deposits does not appearto be preserved, and often overlying fan materialhas cut deep channels into the silts.

We surveyed the relative elevations of the pre-served lake sediments along the valley for 21 kmnorth of the slide (survey techniques are pre-sented in the Appendix). The northernmost rem-nants of demonstrably lacustrine sediment rangefrom 20 to 40 m above the elevation of the slidesurface at Jiaocheng (Fig. 9). We referenced allelevations to the lowest point on the upstreamside of the dam where we could discern the orig-inal surface of the slide. The edge of the remnantslide on the opposite valley wall is present at al-most the same elevation (1–2 m higher), and weare confident that, although the original spillwaycould have been lower than our estimated posi-tion, it could not have been higher. Furthermore,we consistently surveyed the highest-preservedsediments; any subsequent erosion would serveto reduce this estimate. Although our inferenceof tilting relies on a few exposures of sediment atelevations higher than the slide, these elevationsare minimum estimates of the original eleva-

tions, and our estimate of the degree of tilting islikewise a minimum.

We obtained a sample of the lacustrine sedi-ments for IRSL dating (see Appendix for analyt-ical procedures and Tables 2 and 3 for results).Sample EK 98-5 (TIBT 98-2) was collected froma 10 m face of wavy, laminated lacustrine muds~19 km above the landslide. The sample gives anage of 14.2 ± 1.7 ka and is within error of ourestimate for the age of Q3 deposits upstream(6–13 ka). Given 20–40 m of tilting over a 21 kmbase line, this age implies tilting rates of 6.6 ×10–8 to 1.4 × 10–7 rad/yr. At the northern end ofthe former lake, this range translates to differen-tial rock uplift rates of 1.4–2.8 mm/yr. Theserates are fairly rapid given the apparent absenceof geodetically detectable shortening across therange. The linear geometry of the deposit, unfor-tunately, allows us to resolve only the north-southcomponent of the tilting; any east-west compo-nent remains unknown.

Pleistocene Tilting of Q1 and Relief Development in the Min Shan

East-west tilting is recorded in the west-dip-ping Pleistocene Q1 conglomerates east of theMin Jiang fault. Q1 deposits are preserved overnearly 175 km2 of the Min Jiang basin (Fig. 4).Over most of this region, the conglomeratesstrike 010° to 040° and dip ~14°NW (Fig. 4, in-set). The only exception occurs at the southernend of the basin, where the gravels are gentlyfolded. Throughout the rest of the basin, they arepreserved on long west-sloping ridges betweendeeply incised tributary drainages (Fig. 6A). Thepresent-day eastern extent of the basin is defined

by the onlap of these conglomerates onto an~15°–18°W-dipping bedrock surface of very lowrelief (Fig. 6C). This surface extends from be-neath the conglomerates and continues to eleva-tions exceeding 4200 m along the west flank ofthe Min Shan (Fig. 4).

Determining the magnitude of the tilting re-corded in the Q1 conglomerates depends largelyon an estimate of the initial slope. Depositionalslopes of terrestrial deposits can vary widely, de-pending on the nature of the transport process,sediment size, and fine sediment content (e.g.,Blair and McPherson, 1994). However, threelines of evidence lead us to conclude that the dep-ositional slope of the Q1 conglomerates in theMin Jiang basin was fairly shallow. First, the de-posits are alluvial in character. All exposures ofthe Q1 conglomerates are clast-supported gravelsinterbedded with fine sands and silts. The finer-grained layers dip concordantly with the coarsergravels. We found no exposures within the Q1conglomerates of debris-flow deposits or ofglacial till. Compilations of the slopes of modernalluvial fans in a wide range of environmentsindicate that depositional slopes on fans domi-nated by alluvial deposits rarely exceed 3°–4°(Blair and McPherson, 1994). Second, measure-ments of the surface slope on undeformed Q2 ter-races preserved in the tributary valleys east ofGonganin (Fig. 4) range from 1° to 3°. Althoughthese may not be a direct analogue for the earlierdeposits, the Q2 terrace gravels are of similargrain size and composition and were likely de-posited on a similar depositional slope. Third, wewere able to estimate the paleoflow depth frompreserved channel geometries, and we measuredthe median grain size at one locality within the

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NEOTECTONICS OF THE MIN SHAN, CHINA

Geological Society of America Bulletin, March 2000 389

Q1 conglomerates. This approach yielded a pale-oslope estimate of <1° (Paola and Mohrig, 1996).Although a single measurement may not be a re-liable indicator of the regional dip, taken with theother lines of evidence, it suggests that the Q1conglomerates were deposited with a shallow ini-tial slope. In the discussion that follows, we makea conservative estimate that the initial slope was~4°. This estimate is at the high end of the rangeof likely depositional slopes and thus representsa maximum value.

Given the present mean dip of the Q1 conglom-erates (~14°NW; Fig. 4), the analysis above indi-cates that the Q1 has been tilted by at least 10°.The spatial extent over which it occurred indicatesthat down-to-the-northwest tilting represents asignificant component of Quaternary deformationalong western flank of the Min Shan. Our map-ping demonstrates that there are no faults exposedbetween the eastern edge of the basin and the crestof the Min Shan. Young faults cut the Q1 con-glomerates only in the southernmost end of thebasin (Ningpo fault; Fig. 4). North and east of thisregion, the conglomerates dip uniformly to thewest-northwest. Preserved remnants of the sub-basin erosion surface extend nearly to the crest ofthe Min Shan (Fig. 6C) and suggest that the west-ern part of the range behaved as a relatively coher-ent block during late Quaternary deformation.

Simple geometric considerations suggest thattilting of the Q1 conglomerates can account formost of the modern relief between the Min Shanand the plateau to the west. Q1 gravels are pre-served as far as 6 km east of the Min Jiang fault;the erosion surface upon which they were de-posited is preserved as far as 9 km east of the fault.Tilting of 10° over these horizontal base lines(3–6 km for the Q1 and 7–9 km for the subbasinsurface) leads to differential rock uplift of500–1000 m at the eastern edge of the Q1 gravelsand 1200–1500 m at the eastern edge of the sub-basin surface. Projecting this tilting to the crest ofthe Min Shan (some 12–15 km east of the MinJiang fault) would result in 2000–2500 m of dif-ferential rock uplift. Modern relief between therange crest and the Min Jiang valley ranges from1600 to 2200 m, and it is clear that most, if not all,of it could have developed as a result of Pleis-tocene tilting. Paleoflow indicators and clast com-position suggest that Q1 deposits were derivedfrom the Min Shan; clearly, there must have beensome initial topographic gradient east of the basin.However, the lack of observed structures betweenthe basin and the crest of the range makes a com-pelling argument that the much of the moderntopography developed during Pleistocene tiltingalong the western flank of the range.

Inferred rates of tilting within the Min Shan areremarkably rapid. Tilting of 10° in 250 k.y. re-quires rates on the order of 7 × 10–7rad/yr. Differ-

ential rock uplift rates depend upon the base linechosen, but for the three positions we have alreadydiscussed the eastern extent of Q1, the eastern ex-tent of the subbasin surface, and the crest of therange—rates range from 2.0 to 4.1 mm/yr,4.8–6.2 mm/yr, and 8.3–10.3 mm/yr, respectively.We stress that these values only reflect differentialrock uplift. Inferred mean incision rates of tribu-tary streams through the Q1 and the underlyingbedrock approach ~1.5 mm/yr (e.g., ~200 m in~160 k.y.). The lack of information regarding thelowering rate of hillslopes and ridge crests in theQ1deposits and the lack of denudation rates east ofthe preserved basin margin prevent us from ob-taining a quantitative estimate of surface uplift inthe Min Shan (England and Molnar, 1990a). How-ever, the preservation of the bedrock surface thatfloors the basin at elevations approaching the crestof the range leads us to infer that differential sur-face uplift was very likely positive along the westflank of the Min Shan during the late Pleistocene.

FLEXURAL ISOSTATIC CONTRIBUTION TO TILTING IN THEMIN JIANG BASIN

Some component of the down-to-the-north-west tilting observed in the Min Jiang basin couldbe attributed to a flexural isostatic response to de-nudation along the eastern flank of the range.Such a mechanism has been invoked for highmountain ranges flanking both the TibetanPlateau and Andean Altiplano (Masek et al.,1994b; Molnar and England, 1990; Montgomery,1994), as well as for high-standing continentalescarpments (Kooi and Beaumont, 1994; Tuckerand Slingerland, 1994). In order to assess thenecessary conditions under which an isostatic re-sponse to erosional unloading could significantlyinfluence the deformation observed in the MinShan, we utilize a simple flexural model. We taketwo complementary approaches. We first con-sider the modern topographic relief as the residuefrom the removal of a distributed load from thelithosphere so that we can model the flexural re-sponse to erosional dissection of a plateau mar-gin. Second, we model the effects of uniform ero-sional unloading of the eastern flank of the rangeand ask how much mass removal is required toproduce the observed tilting (deflections) in theMin Jiang basin.

In our first model, the load removed from thelithosphere is calculated from the topographic re-lief in the landscape. We utilize a digital elevationmodel of the topography in eastern Tibet (90 mnominal resolution, data set described in Fieldinget al., 1994) and extract the topographic reliefalong a 75-km-wide swath oriented perpendicu-lar to the topographic margin of the plateau (seeFig. 2 for location). The swath width was chosen

to enclose the entire upper drainage basin of theFu Jiang, the major river draining the eastern MinShan, and so provides a relatively complete pic-ture of the relief structure of the landscape east ofthe Min Shan. We binned elevations across theswath and calculated local relief as the differencebetween the maximum elevation and all other el-evations within a bin. This is functionally equiv-alent to fitting a surface to peak elevations andextracting the difference between that surface andthe landscape (Montgomery, 1994). However, bymeasuring relief along the entire channel system,we obtain a volumetric estimate of local valley-bottom to ridge-top relief. Relief is averaged ineach bin to obtain an estimate of the mean reliefalong the chosen swath. This quantity is imposedas a distributed surficial load removed from thelithosphere.

We model the lithospheric response to removalof this topographic load as that of a thin, uni-formly rigid elastic plate overlying a viscous sub-strate (Turcotte and Schubert, 1982; equa-tion 3.125). In the discussion that follows we usethe term “flexural rigidity” interchangeably withthe term “effective elastic thickness,”Te, wherewe assume a Young’s modulus of 7 × 1010Pa andPoisson’s ratio of 0.25 (Turcotte and Schubert,1982). The density contrast between the crust andmantle in our model is 600 kg/m3 (ρcrust =2700 kg/m3, ρmantle= 3300 kg/m3). The effectiveelastic thickness and the lateral boundary condi-tions control the flexural response to the imposedload. End loads in our model are assumed to bezero. We place the load within a model domainthat is long (>500 km) relative to the load andforce the ends of the model domain to have zerodeflection. The modeled deflection is then calcu-lated over the original length of the swath. In thisway, we avoid edge effects at the ends of oursampled topography.

The topography of the Min Shan region(Fig. 10) is characterized by a gradual increase inelevation from the Tibetan Plateau in the west tothe Min Jiang fault and an abrupt rise in elevationbetween the Min Jiang fault and the crest of therange. Elevations then fall rapidly toward the fore-land. Mean relief across this landscape variesfrom 200 to 500 m west of the Min Jiang fault, tonearly 2000 m along the eastern flank of therange. In order to facilitate interpretation of thepotential isostatic response attributable to this re-lief, we impose the flexural response on an ini-tially horizontal reference surface (Fig. 10) at theelevation of the Tibetan Plateau at the western endof our swath (~3500 m). We caution that we arenot assuming an initial condition of the plateauprior to erosion. The horizontal line simply servesas a marker horizon on which to examine the ef-fects of varying parameters. The imposed load isalways the mean relief measured along the swath.

The lack of an independent estimate of therigidity of the lithosphere in eastern Tibet forcesus to model a range of possible values. Resultsare shown in Figure 10C and are strongly de-pendent on the modeled rigidity of the litho-sphere. For elastic thickness ranging from 30 to40 km, the potential isostatic response is distrib-uted over a wide region of the eastern plateau. Al-though these models match the broad rise intopography between the Tibetan Plateau and theMin Jiang fault fairly well, they fail to reproducethe high topography of the Min Shan. Only forelastic thicknesses approaching 4–5 km do themodel results produce a deflection of similarwavelength as the Min Shan. This result impliesthat in order for the deformation observed alongthe west flank of the Min Shan to be an isostaticresponse to erosional unloading, the crust of theeastern plateau would have to have extremelylow rigidity (~4 to 8 × 1020 N·m; Te = 4–5 km).

An additional constraint can be inferred fromthe amplitude of the modeled deflections. Even ifthe lithosphere of the Min Shan region is ex-tremely weak (Te≈ 5 km), the predicted amplitudeof the deflection between the location of the MinJiang fault and the crest of the range is on the orderof 500 m (Fig. 10C). This result implies that evenif all the present-day relief developed in the last250 k.y., the isostatic response would not repro-duce the observed tilting along the western flankof the Min Shan. We can, however, ask how muchadditional denudation would be required to gener-ate a deflection of the observed amplitude (in-ferred to be 2000–2500 m at the crest of the range).Here we assume that all denudation is concen-trated along the eastern flank of the Min Shan. Al-though this is not true in detail, it should allow usto approximate, to a first order, the amount of massremoval necessary to produce a deflection similarto the observed topography of the Min Shan.

The deflection of a point at the uneroded edgeof a plateau is a strong function of load width upto the flexural wavelength (for Te = 5 km, thiswavelength is ~45 km; Fig. 11). Removal of adistributed load (of uniform thickness) from a re-gion greater than or equal in width to this dis-tance produces a deflection at the margin with amagnitude equal to ~0.48 of the removed thick-ness. Consequently, in order to produce a deflec-tion with amplitude of 2–2.5 km, we would haveto remove at least 4–5 km of material east of thecrest of the range. It is important to note that thissimple model also demonstrates that if denuda-tion is focused in a narrow zone (i.e., less than theflexural wavelength) along the range front, theisostatic response is greatly reduced.

Our model results indicate that in order for theobserved deformation and topography along thewestern flank of the Min Shan to be produced asan isostatic response to erosion, two conditions

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Figure 10. (A) Topographic swath profile across the Min Shan margin. Maximum, minimum(dark lines), and mean elevations (gray) are extracted from ~90 m bins (1 pixel), selected per-pendicular to the swath. (B) Mean relief as a function of distance across the Min Shan. This func-tion is used as a distributed line load removed from the lithosphere in our flexural model. (C)Modeled deflections for various Tevalues for a reference surface calculated from removal of loadin B. Reference surface is originally a horizontal surface at 3500 m. Gray line shows envelope ofmaximum topography for reference. Details of the calculation are discussed in the text.

must hold: (1) the crust must be extremely weak(Te≤ 5 km), and (2) there must have been 4–5 kmof denudation east of the range crest. Unfortu-nately, little is known about the strength of thelithosphere in this region of the plateau. Estimatesof the elastic thickness of the lithosphere in south-ern Tibet range from 5 to 15 km (Masek et al.,1994a). However, these measurements are in a re-gion of active extension and high heat flow. Thereis little geologic evidence for such high heat flowin the Min Shan region, and we find it unlikelythat the lithosphere has such low rigidity.

Removal of 4–5 km of material in ~250 k.y.from the eastern range front would require de-nudation rates on the order of 15–20 mm/yr. Ex-isting thermochronologic data from the regionare sparse, but bear importantly on the long-termdenudation rates along the range. Fission-trackmeasurements in apatite taken from a few locali-ties within the Min Shan (Arne et al., 1997) yieldages between 30 and 120 Ma. Kirby et al. (1999)obtained a cooling trajectory from a small Meso-zoic granite at the eastern foot of the range (westof the Huya fault; Fig. 3). Potassium feldspars(40Ar/39Ar) and apatite (U-Th-He) from this sam-ple record rapid cooling from temperatures of~200 °C beginning at ca. 6 Ma. These authors ar-gued that mean denudation rates over this periodwere probably between 1 and 2 mm/yr. If thesedenudation rates are representative of the lateQuaternary (past ~250 k.y.), they suggest thaterosion along the eastern range front was likely

limited to 250–500 m. This is an order of magni-tude smaller than rates required to produce theobserved tilting along the western flank of theMin Shan and suggests that the isostatic compo-nent is probably limited to 100–200 m (assumingTe = 5 km). Thus, the tilting in the Min Jiangbasin (and the consequent topography) appears tobe primarily a tectonic signature.

DISCUSSION: DYNAMICS OF MOUNTAIN BUILDING IN THE MINSHAN

The most striking observation regarding theactive faulting along the Min Shan is that theactive faults bear little relationship to the topo-graphic range front. Although the Huya fault isnear the eastern foot of the range, it is discontin-uous along strike. North and south of the Huyafault, the topographic margin does not coincidewith recognized active faulting. Indeed, in thenorthern Min Shan (north of the Xue Shan fault),Proterozoic–Triassic rocks form a continuoussection that extends from the crest of the range tothe foreland (Fig. 3). A similarly intact sectionoccurs south of the Huya fault (Fig. 3). These re-lationships preclude large-magnitude Cenozoicshortening across the range and are surprisinggiven the profound variation in elevation (andpresumably crustal thickness) from the plateau tothe foreland. The lack of correlation between ac-tive faults and the topographic margin of the

plateau is reminiscent of the relationships in theLongmen Shan region to the south (Burchfielet al., 1995), which have been interpreted as aconsequence of shortening focused in a weakmiddle and lower crust (Royden et al., 1997). Thegeology of the eastern Min Shan thus arguesagainst a conventional mode of shortening at theplateau margin whereby crustal thickening is ab-sorbed at the surface by large-magnitude thrustfaults and is consistent with the hypothesis ofRoyden et al. (1997).

Late Quaternary deformation within the west-ern Min Shan is similarly enigmatic. The MinJiang fault, although active in the late Pleis-tocene, does not appear to have significant Holo-cene displacement along it. In addition, the senseof displacement across the Min Jiang fault(plateau side up) cannot account for the hightopography of the Min Shan relative to theplateau. The slow shortening rates inferred forthe Min Jiang fault are consistent with regionalhorizontal velocities determined by space geod-esy. However, the evidence for rapid tiltingwithin the range since late Pleistocene time indi-cates that significant vertical deformation is lo-calized beneath the Min Shan itself. The absenceof faults between the eastern margin of the MinJiang basin and the crest of the range and the con-tinuity of the subbasin erosion surface suggestthat, if discrete structures exist within the MinShan, they must end at depth beneath the range.Likewise, if the Huya fault extends north alongthe Min Shan, either as a blind structure or as aseries of distributed faults, it could account fortilting along the west flank of the Min Shan.However, projection of 10° of tilt across a baseline of 50 km (approximately the distance to theinferred trace of the Huya fault) implies that ver-tical displacement rates (over the past 250 k.y.)on the fault exceed 35 mm/yr. On a fault dipping60°W, this projection implies horizontal shorten-ing rates across the fault of nearly 20 mm/yr.These rates are at odds with both the record ofhistoric seismicity (Editorial Board, State Seis-mological Bureau, 1989) and recent GPS meas-urements (Chen et al., 2000). Clearly, the Huyafault cannot be directly responsible for the ob-served tilting in the Min Jiang basin. We infer thatlate Quaternary tilting along the western flank ofthe Min Shan reflects deformation localized be-neath the range itself, the nature of which re-mains somewhat enigmatic.

The discrepancy between geodetically ob-served rates of horizontal shortening (Chenet al., 2000; King et al., 1997) and the tiltingrates we infer from our mapping and chronol-ogy is a second remarkable result of our work.Insofar as the geodetic measurements are repre-sentative of the time-averaged velocity field,they appear to restrict the rates of differential

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Figure 11. Deflection of an elastic plate as a function of the width of a distributed surface load.Inset shows a schematic representation of load geometry. Load is of constant thickness and vari-able width. Deflection is calculated at the inner edge of the load (x = 0) and roughly correspondsto the divide of a dissected plateau margin. Note that the greatest relief in the Min Shan occursalong the steep eastern range front, over a width of 50–100 km (see Fig. 10); thus (for Te≈≈ 5 km),we expect deflections at the range crest to be between 0.4 and 0.45 of the load removed.

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392 Geological Society of America Bulletin, March 2000

rock uplift that can be attributed to horizontalshortening to less than 2–3 mm/yr. Although therecurrence interval of seismic events beneaththe Min Shan is not well known, the historicrecord of seismicity contains only four events ofmagnitude >6 (three of these occurred as part ofa single sequence in 1976; Editorial Board,State Seismological Bureau, 1989; Fig. 2). Asthe record of earthquakes in this region of Chinaextends back over 1500 yr, it seems likely thatmean rates of horizontal shortening associatedwith seismic release are slow.

Although optical dating is still an immaturediscipline, there is a growing body of empiricalevidence (Aitken, 1998; Berger, 1995) that sug-gests that the technique can resolve depositionalages within 20%–30% over the past 300–400 k.y.The correspondence between our IRSL age forthe lacustrine sediments in the southwestern partof the range and 14C ages on correlative terraceslends a measure of confidence to the technique.In addition, our age on the Q1–2loess (ca. 160 ka)corresponds with a time period of documentedloess deposition on the Loess Plateau (Kukla,1987) and suggests that the technique is probablyresolving depositional ages. Even if our age esti-mates are inaccurate by a factor of two, the in-ferred rates of tilting exceed those attributable toregional horizontal shortening. Although the res-olution of this problem awaits further refinementof the Quaternary chronology and/or the estab-lishment of a detailed geodetic survey across therange, we conclude that rapid tilting within theMin Shan is a significant component of the lateQuaternary deformation field along this marginof the Tibetan Plateau.

Rapid Holocene tilting observed in lacustrinesediments in the southern Min Shan supports thisinference and further highlights the discrepancybetween vertical deformation rates inferred fromthe geology and the lack of significant shorteningacross the Min Shan. The age of the deposit(ca. 14 ka) indicates that tilting likely took placethroughout the Holocene and may be active to-day. The lacustrine sediments are within 10 km ofone GPS station (MJZ1; Fig. 2), which resolvesno significant eastward motion relative to theforeland. As in the Min Jiang basin, we find itlikely that distributed tilting is an important com-ponent of the recent deformation field in thesouthwestern Min Shan. As vertical velocities aredifficult to measure by geodetic means, our workhighlights the utility of geologic markers in re-constructing differential rock uplift across re-gionally significant length scales. In both locali-ties along the Min Shan where geologic markersare present, we find evidence for substantial tilt-ing. We can only speculate as to its importanceelsewhere in this region of the plateau.

We have argued above that the observed late

Quaternary tilting within the Min Shan cannot bedirectly attributed to either horizontal shorteningor to an isostatic response to denudation at theplateau margin. Although both of these mecha-nisms undoubtedly contribute to the observedtilting in the Min Jiang basin, even in concertthey do not appear to account for the rapid rateswe infer from the geometry and age of the Q1 de-posits. Our results thus suggest the presence of anadditional driving mechanism. The lack of evi-dence for large-magnitude Cenozoic shorteningalong the eastern range front of the Min Shanleads us to consider the possibility that much ofthe deformation may be driven by flow of a weaklower crust. If the hypothesis of Royden et al.(1997) is correct, lower-crustal material drivenfrom beneath the central plateau may impinge oncold, unthickened crust at the plateau and local-ize crustal thickening. This scenario is especiallytrue of regions of significant anisotropy (for ex-ample, the Sichuan Basin). The presence of pre-Cenozoic structures related to the Xue Shan plat-form (Min Jiang and Huya faults) may helplocalize crustal thickening and, to some degree,set the width of the mountain range. We suggestthat the observed deformation (both faulting anddistributed tilting) in the Min Shan may be a cou-pled upper-crustal response to thickening local-ized in the lower crust.

The nature of the coupling between lower-crustal flow and upper-crustal deformation re-mains somewhat enigmatic, as does the mannerin which tilting along the western flank of theMin Shan is accommodated along the easternrange front. Tilting in the Min Jiang basin couldrepresent the western limb of a range-scaleantiform with a wavelength of between 50 and100 km. Alternatively, west-side-up displace-ment on the Huya fault could be distributed onsmall discrete structures along the eastern rangefront between the Xue Shan and Tazang faults. Ineither case, the lack of significant horizontalshortening across the range appears to requirethat Quaternary deformation along this segmentof the eastern margin of the Tibetan Plateau isdriven from below.

ACKNOWLEDGMENTS

This work was supported by the ContinentalDynamics program at the National ScienceFoundation (EAR-961460) and by NASA (NAG5-6062). We thank B. King, L. Royden, C. Ru-bin, K. Karlstrom, and M. Clark for helpful dis-cussions on various aspects of this work. B. Kingkindly helped prepare Figure 2, and L. Roydenprovided the algorithm for a band matrix inversesolution to the flexural equation. An informal re-view by C. Rubin and formal reviews by R. Ar-rowsmith, M. Roy, and F. Pazzaglia led to im-

provements in the manuscript. Finally, we thankthe staff and drivers at the Chengdu Institute ofGeology and Mineral Resources for making ourwork in China possible.

APPENDIX

Survey Techniques

We surveyed the relative elevations of channel pro-files, stream terraces, and lacustrine sediments by usinga hand-held altimeter (±1 m), clinometer (±0.5°), andlaser range finder (±1 m). Uncertainties are quotedfrom manufacturer specifications. Survey locationswere logged by using a hand-held GPS receiver. As thealong-channel distance varies with the curvature of thechannel, we approximated along-channel distance byusing the vehicle odometer (±0.1 km) for the terraceand channel survey. We corrected for daily barometricchanges by tracking barometric pressure with a datalogger at a fixed location near the center of our surveybase line and corrected for these changes over the timeof the survey. We estimate that individual measure-ments are accurate to ~3 m, but that relative differencesin elevation (i.e., terrace height above channel) areprobably accurate to 1.5–2 m.

Sampling and Analytical Techniques—IRSL Dating

IRSL samples (Table 2 and Figure 5) were col-lected following the general guidelines found inAitken (1998). The outer, desiccated face of each out-crop was removed (generally 10–20 cm) taking careto shield the outcrop from direct sunlight. Sampleswere collected in light-tight Gill Style sample tins andsealed with black tape to preserve moisture contentand minimize sample disturbance. In most cases, itwas possible to push the tins directly into the sedi-ment, although in one case (TIBT 98–4), a block wascarved from the outcrop with a knife. In all cases, theouter 2–3 mm of the block was removed under low-energy light conditions in the laboratory. At each out-crop, three small samples were collected for dosime-try measurements (Table 2). These were spaced over a50–60 cm distance that bracketed the sample (the ef-fective range of gamma radiation in typical sedimentis ~30 cm). For the luminescence measurements, thepolymineralic, noncarbonate, detrital, 4–11µm-diam-eter size fraction was prepared. Laboratory prepara-tion procedures are outlined elsewhere (Berger, 1990;Wintle and Huntley, 1980).

A luminescence age t = DE/DR, where DE is the so-called “equivalent dose,” or paleodose (units of Gy[gray], 1 Gy = 100 rads), and DR is the effective doserate (Gy/k.y.). DE is a measure of the postdepositionalabsorbed ambient ionizing radiation and is derivedfrom luminescence measurements. DR is derived fromindependent measurements of U, Th, K, and water con-centration. For U and Th measurements, the techniqueof thick-source alpha-particle counting (TSAC) (Hunt-ley and Wintle, 1981) was applied to dried powders. Kwas determined by commercial atomic absorptionspectrophotometry. Radioactive secular equilibrium isassumed; this assumption is apparently valid for many,although not all, alluvial and fluvial sediments (Olleyet al., 1996). Of the available analytical techniques,TSAC can apparently provide a clue to the occurrenceof such disequilibrium and, compared to neutron-acti-vation techniques, can provide a more accurate dose-rate measurement for fluvial samples older than~10 k.y. (Olley et al., 1996).

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Geological Society of America Bulletin, March 2000 393

For IRSL measurements, the deep-blue lumines-cence near the known 410 nm emission from mostK-rich feldspars was chosen for detection (Aitken,1998). To convert the sample’s fossil light to an ab-sorbed-energy equivalent, calibrated laboratory beta(90Sr–90Y) and alpha (241Am) sources were used. Thesignature was recorded with an automated, high-capac-ity Daybreak Nuclear Model 1150 reader using anEMI 9635Q photomultiplier tube. Dose-responsecurves and other data were processed with softwareproduced by one of the authors (Berger; see Bergeret al., 1987).

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