Structural evolution and vorticity of ﬂow during extrusion and...

Structural evolution and vorticity of flow during extrusion and

exhumation of the Greater Himalayan Slab, Mount

Everest Massif, Tibet/Nepal: implications for

orogen-scale flow partitioning

M. J. JESSUP1, R. D. LAW1, M. P. SEARLE2 &

M. S. HUBBARD3

1Department of Geosciences, Virginia Tech, Blacksburg, Virginia 24061, USA

(e-mail: [email protected])2Department of Earth Sciences, Oxford University, Oxford, OX1 3PR, UK

3Department of Geology, Kansas State University, Manhattan, Kansas 66506, USA

Abstract: The Greater Himalayan Slab (GHS) is composed of a north-dipping anatectic core,bounded above by the South Tibetan detachment system (STDS) and below by the MainCentral thrust zone (MCTZ). Assuming simultaneous movement on the MCTZ and STDS, theGHS can be modelled as a southward-extruding wedge or channel. New insights into extrusion-related flow within the GHS emerge from detailed kinematic and vorticity analyses in theEverest region. At the highest structural levels, mean kinematic vorticity number (Wm) estimatesof 0.74–0.91 (c. 45–28% pure shear) were obtained from sheared Tethyan limestone and marblefrom the Yellow Band on Mount Everest. Underlying amphibolite-facies schists and gneisses,exposed in Rongbuk valley, yield Wm estimates of 0.57–0.85 (c. 62–35% pure shear) and associ-ated microstructures indicate that flow occurred at close to peak metamorphic conditions. Vorti-city analysis becomes progressively more problematic as deformation temperatures increasetowards the anatectic core. Within the MCTZ, rigid elongate garnet grains yield Wm estimatesof 0.63–0.77 (c. 58–44% pure shear). We attribute flow partitioning in the GHS to spatial andtemporal variations that resulted in the juxtaposition of amphibolite-facies rocks, which recordearly stages of extrusion, with greenschist to unmetamorphosed samples that record later stagesof exhumation.

The .2500 km length of the Himalayan orogen iscored by a suite of north-dipping metamorphicrocks (the Greater Himalayan Slab; GHS), that arebounded above and below by the normal-senseSouth Tibetan detachment system (STDS) andreverse-sense Main Central thrust zone (MCTZ),respectively (Figs 1 & 2). Assuming simultaneousmovement along these crustal-scale boundingshear zones (see review by Godin et al. 2006b),the GHS is often modelled as a north-dippingwedge or channel of mid-crustal rocks that wasextruded southward from beneath the Tibetanplateau (Fig. 2) beginning in early Miocene time(e.g. Burchfiel & Royden 1985; Burchfiel et al.1992; Hodges et al. 1992). Although consensus onthis general concept of extrusion during crustalconvergence exists, and a range of orogen-scalekinematic and thermal–mechanical extrusionmodels have been proposed, surprisingly littleresearch has focused on quantifying the kinematics(vorticity) of flow within the slab and its potential

causal relationship with progressive exhumationof the GHS.

The use of vorticity analysis to quantify flowwithin sheared rocks has proven to be a usefultool for quantifying the nature and distribution offlow regimes within a range of tectonic settingsincluding contractional (e.g. Simpson & De Paor1997; Xypolias & Doutsos 2000; Xypolias &Koukouvelas 2001, Xypolias & Kokkalas 2006),extensional (Wells 2001; Bailey & Eyster 2003)and transpressional (Wallis 1995; Klepeis et al.1999; Holcombe & Little 2001; Bailey et al.2004) regimes. Vorticity analysis enables esti-mation of the relative contributions of pure andsimple shear, yielding important constraints forGHS extrusion models. Identification of a pureshear component is critically important becausesuch flow would result in: (1) thinning and trans-port-parallel extension of the slab itself, and (2)an increase in both strain rates and extrusion ratesrelative to strict simple shear. Attempts to quantify

From: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in ContinentalCollision Zones. Geological Society, London, Special Publications, 268, 379–413.0305-8719/06/$15.00 # The Geological Society of London 2006.

flow within the GHS that accommodated thissouthward extrusion are limited to: (1) a singletransect through the lowermost 900 m of theGHS exposed in the Sutlej valley of NW India(Fig. 1; Grasemann et al. 1999; Vannay & Grase-mann 2001); (2) preliminary results from the topof the GHS exposed in the Rongbuk valley on thenorth side of the Everest massif, Tibet (Law et al.2004); and (3) preliminary results from the middleof the GHS in the Bhutan Himalaya (Carosi et al.2006). Quantifying and characterizing flow withinthe GHS is important for development of morerealistic models for evolution of the Himalaya, par-ticularly those that propose a synergistic interplaybetween extrusion, erosion and exhumation (e.g.Beaumont et al. 2001, 2004, 2006; Hodges et al.2001; Grujic et al. 2002; Jamieson et al. 2004,2006; Hodges 2006).

The topographic relief of the Everest massif,Tibet/Nepal (Fig. 1), provides a window intomid-crustal processes responsible for extrusion ofthe GHS, and a particularly appropriate field areato test the various components of extrusionmodels. In this paper, we combine field-based struc-tural analysis with detailed vorticity analyses ofsamples from a north–south transect through theGHS in the Everest region using the rigid graintechnique of Wallis et al. (1993) and Wallis(1995). Samples collected for vorticity analysesare from a variety of structural and lithologicsettings, including high-altitude and summitsamples collected during two pioneering climbingexpeditions (1933 and 1953) on the north andsouth sides of Mount Everest, respectively. Rigidgrain analysis using elongate garnet porphyroclastsquantify flow along the MCTZ and, by integration

KhatmanduDarjeeling

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Fig. 1. Simplified geological map of the Himalaya including the distribution of the main lithotectonic elements of theorogen. The location of the Everest transect and significant geographical locations and areas referred to in the textare indicated. Inset (a) is a digital elevation model displaying the range of elevation for SE Asia (dark grey ¼ low; lightgrey ¼ high; white ¼ ocean). The Himalaya marks the transition from high elevations of the Tibetan plateau to thelowlands of the Indian plate. MBT, Main Boundary thrust; MCTZ, Main Central thrust zone; MFT, Main Frontal thrust;MMT, Main Mantle thrust; STDS, South Tibetan detachment system; ZSZ, Zanskar shear zone.

M. J. JESSUP ET AL.380

with microstructural analysis, constrain the timingof mylonite formation in relation to peak meta-morphism (Hubbard 1988, 1989). Field-basedstructural analysis characterizes deformation inthe core of the GHS where deformation tempera-tures exceed the upper limit for robust rigid grainvorticity analysis. As the first attempt to quantifyflow in a transect across the entire GHS, manynew insights into vorticity of flow emerge, includ-ing an understanding of how flow was partitionedduring extrusion and exhumation of the GHS.

Tectonic setting

The Himalaya–Tibet orogenic system has accom-modated crustal convergence since initiation ofcollision between India and Asia at c. 55–50 Ma

(Searle et al. 1987; Hodges 2000; Yin & Harrison2000; Figs 1 & 2). The Tibetan plateau encom-passes an area of .5 � 106 km2 of subdued topo-graphy (Fig. 1, inset a), with an average elevationof c. 5000 m (Fielding et al. 1994). To the southstretch the flat, low elevations characteristic of theundeformed internal margin of the Indian plate.Between lies the crest of the Himalaya, whichextends for c. 2500 km along-strike, contains thehighest elevations in the world (8850 m), andprovides exposure of mid-crustal rocks belongingto the GHS (Figs 1 & 2). The GHS forms a 5–30 km thick section of metasedimentary rocks thatare intruded by leucogranite dykes and migmatizedto varying degrees (Hodges 2000). The age ofleucogranite crystallization (c. 23–13 Ma) suggeststhey were part of a protracted event that marks theculmination of peak metamorphism (Searle 1996;Hodges 2000).

The metamorphic evolution of the GHS is oftensplit into two tectonothermal events that maymark distinct thermal pulses or a thermal conti-nuum. Kyanite-grade assemblages are interpretedas relicts of an early event (M1) that U–Th–Pbmonazite geochronology suggests occurred at c.35–30 Ma (Walker et al. 1999; Simpson et al.2000). M1 is often overprinted by the pervasivehigh temperature–low pressure tectonothermalevent (M2; c. 23–17 Ma) associated with decom-pression, migmatization, and emplacement ofleucogranite sills (Hodges 2000; Simpson et al.2000). 40Ar/39Ar thermochronology of muscoviteand biotite from leucogranites yields cooling agesthat are usually only slightly younger than theU–Pb crystallization age of the leucogranites,suggesting rapid decompression following theiremplacement (e.g. Hodges et al. (1998) for theEverest transect).

Two shear zones bound the GHS: the STDSabove and the MCTZ below (Figs 1 & 2). TheSTDS juxtaposes Tethyan sedimentary rocks ofthe Tibetan Zone against the GHS, while theMCTZ separates the GHS above from rocks of theLesser Himalaya Zone (LHZ) below (for detailedreview see Godin et al. 2006b).

Extrusion models

Despite on-going controversy regarding evidencefor simultaneous movement on the MCTZ andSTDS (see Godin et al. 2006b), many researcherscontinue to view the tectonic evolution of theGHS in the context of models involving southwardextrusion of mid-crustal rocks from beneath theTibetan plateau towards the topographic surface atthe plateau margin. All of these models assumesimultaneous motion on the upper and lower

STDS

STDS

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GHS

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(a) Model 1. Adapted from Burchfiel & Royden (1985)

(b) Model 2. Adapted from Grasemann et al. (1999)

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Tibet

Indian Shield

Indian Shield

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GHS

Tibetgneissdome

partial meltzone-

LHS MBT

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Indian lower crust40 km

(V:H ~2:1)

(d) Model 4. Adapted from Beaumont et al. (2004)

Indian Shield

STDS

MCTZ

GHS

(c) Model 3. Adapted from Grujic et al. (1996)

Tibet

Indian Shield

Tibet

Tibet

N S

Fig. 2. (a–d) Extrusion and channel flow modelsproposed for the evolution of the Greater HimalayanSlab (GHS). See text for detailed review of each model.GHS, Greater Himalayan Slab; LHS, Lesser HimalayanSequence; MBT, Main Boundary thrust; MCTZ, MainCentral thrust zone; MFT, Main Frontal thrust; STDS,South Tibetan detachment system.

FLOW PARTITIONING IN THE HIMALAYA 381

surfaces of the extruding unit. Two types of modelsmay be distinguished: (1) kinematic models forwedge extrusion (Fig. 2a–c) based on the assump-tion that the STDS and MCTZ join at depth as,for example, suggested by early interpretationsof INDEPTH seismic data (Nelson et al. 1996);(2) more complex coupled thermal–mechanicalfinite-element models involving lateral flow ofrelatively low-viscosity material within a tabularmid-crustal channel in response to a horizontal gra-dient in lithostatic pressure between the Tibetanplateau and Himalayan foreland (Fig. 2d; Beaumontet al. 2001, 2004, 2006).

Kinematic models

Model 1 In the original wedge extrusion model(Burchfiel & Royden 1985; Royden & Burchfiel1987; Kundig 1988, 1989; Burchfiel et al. 1992;Hodges et al. 1993), the extruding wedge developedby gravity-driven collapse in response to theextreme topographic gradient developed alongthe southern margin of the Himalayan orogen. Inthis basic conceptual framework, only one mainphase of north–south extension, triggered by a fun-damental, non-reversible change in the stress stateof the orogenic system as a whole (e.g. England& Molnar 1993), occurred in the Himalaya.Synconvergence extrusion, rather than gravity-driven collapse, has been emphasized in morerecent models. In the original wedge extrusionmodel the nature of deformation/flow within theinterior portions of the wedge was not explicitlyaddressed (Fig. 2a), although a broad zone ofreverse-sense shearing along the base of thewedge was suggested (e.g. Brunel & Kienast1986; Hubbard 1988, 1989, 1996; but see alsoHarrison et al. 1999) as a potential explanation forthe long-known inversion of metamorphic isogradsadjacent to the MCT (Heim & Gansser 1939).Based on the mapping of antiformally foldedisograds in the Zanskar section of the GHS, thiswas expanded upon in an alternative model bySearle & Rex (1989) who proposed that the entiresequence of rocks contained between the MCTand STDS (Zanskar shear zone) was isoclinallyfolded during extrusion. More recently proposedkinematic extrusion models are largely based onlocal evidence for spatial strain path (or vorticity)partitioning within the GHS, and the relativeimportance of pure shear and simple shear flowcomponents.

Model 2 Quantitative evidence for a significantpure shear deformation component associated withflow along the base of the GHS was reported byGrasemann et al. (1999) from the Sutlej Valley

(NW India) section of the MCTZ (Fig. 1). Basedon correlation between: (1) a downward increasein estimated pure shear component, (2) a downwarddecrease in deformation temperatures within thiszone of inverted isograds, and (3) a high pureshear component indicated by late vein sets,Grasemann et al. (1999) proposed that their vorti-city data were most readily interpreted as indicatinga temporal (rather than spatial) change in flowregime associated with a decelerating strain path(Simpson & De Paor 1997; Fossen & Tikoff1997), where simple shear flow at higher tempera-tures is replaced by pure-shear-dominated flow atlower temperatures. Grasemann et al. (1999) pro-posed a model for wedge extrusion in which defor-mation is concentrated towards the boundariesof the wedge and, due to strain compatibility, thecentre of the wedge extrudes mainly by pure shear-ing (Fig. 2b). An important aspect of this model isthat the wedge is detached from the footwall andhanging wall.

Model 3 Microstructural and quartz petrofabricdata were employed by Grujic et al. (1996) toqualitatively investigate deformation/flow withinthe lower-central portion of the GHS exposed inBhutan. These data indicated that at least thelower-central part of the slab had undergone planestrain to weakly constrictional deformation, withflow involving components of both simple shear(reverse or top-to-the-south shear sense) and pureshear. Based on these data, Grujic et al. (1996) pro-posed a wedge extrusion model in which pervasiveshearing occurred throughout the evolving wedge,with opposite shear senses on the top and bottomhalves of the wedge (and highest extrusionvelocities in the centre) leading, as previously pro-posed by Searle & Rex (1989) for the ZanskarHimalaya, to antiformal folding of isograds(Fig. 2c; see Godin et al. 2006b). The pure shearcomponent indicated by petrofabric data was notexplicitly addressed in this channel flow model,although Grujic et al. (1996) did note that pureshear would lead to thinning and transport-parallelstretching of the wedge/channel during extrusion.Subsequent fieldwork in Bhutan led Grujic et al.(2002) to abandon their wedge-shaped model forthe GHS and to regard the GHS as a 10–15 kmthick mid-crustal layer, or channel, extending forat least 200 km northward beneath Tibet. In thisrevised channel flow model (incorporating com-bined Couette and Poiseuille flow; see review byGrujic 2006) the influence of changing thermal con-ditions (viscosity) on flow patterns was considered,and qualitative predictions made on the likelyinfluence of changes in boundary conditions andviscosities on domainal variation in flow vorticities


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Thermal–mechanical models

Model 4 Thermal–mechanical models build onthe concept originally proposed by Nelson et al.(1996) that the GHS represents hot low-viscositymid-crustal material extruded southwards frombeneath Tibet towards the Himalayan front duringcontinental convergence, and the overlapping pro-posal that this extrusion can be modelled using theconcept of channel flow driven by a horizontalgradient in lithostatic pressure between theTibetan plateau and the Himalayan front (Grujicet al. 1996). These concepts, and a broad range ofHimalayan structural, pressure–temperature–time(P–T–t) and geochronologic data, have been suc-cessfully modelled in two dimensions with time-varying, plane strain, coupled thermal–mechanicalfinite-element models in which channel viscositiesare reduced by mantle heat flux and radiogenicheating (Beaumont et al. 2001, 2004, 2006;Jamieson et al. 2002, 2004, 2006). Models beginwith a tectonically thickened crust, which is thenthermally weakened, and flows in a mid-crustalchannel towards the orogenic front. Varying inputparameters and model specifications produce var-iants of the basic model. In these models, channelsare exhumed and exposed by denudation focusedon the high-relief transition between the plateauand orogenic front (Fig. 2d). Implicit in thesemodels is that the structures now exposed at thetopographic front will probably have formedduring the last stages, or cessation of extrusion/exhumation of the channel material, rather thanbeing directly related to processes operating whenthese rocks were flowing at mid-crustal levelsbeneath the plateau.

Everest transect: geological background

In this section we outline the major lithotectonicunits and structures of the Everest transect inorder to provide a foundation for the detaileddiscussions of key areas used for our vorticity ana-lyses. This transect begins with the STDS at the topof the GHS (as exposed in Rongbuk valley, Tibet)extends southward to the summit of Everest, andcontinues southward through the GHS to themore limited exposures in the Nepalese foothillsof rocks belonging to the MCTZ. Detailed struc-tural, metamorphic and geochronological reviewsof the Everest transect (some limited to specificsections) are given by Lombard (1958), Bordet

(1961), Brunel & Kienast (1986), Hubbard (1988,1989), Lombardo et al. (1993), Pognante &Benna (1993), Carosi et al. (1998, 1999a, b),Searle (1999a, b) and Searle et al. (2003, 2006).Geological maps covering different sections ofthe transect have been published by Bordet(1961), Lombardo et al. (1993), Carosi et al.(1998) and Searle (2003).

Two major detachments belonging to the STDShave been mapped in the sidewalls of theRongbuk valley southward to Changtse andMount Everest: the upper brittle Qomolangmadetachment (QD) and lower ductile Lhotse detach-ment (LD) (Fig. 3; Searle 1999a; Searle et al. 2003;see also Lombardo et al. 1993; Carosi et al. 1998,1999b; Sakai et al. 2005). As originally defined,the LD is a distinct ductile high-strain zone thatmarks a metamorphic break between amphibolite-facies rocks below and greenschist-facies (EverestSeries) rocks above (Searle 1999a). More recentdetailed thermobarometric results from samplescollected at the base of the Lhotse wall (Jessupet al. 2004, 2005) and East Rongbuk glacier(Waters et al. 2006), demonstrate temperatures ofc. 6508C immediately above and below the pro-posed detachment, and suggest that the LD maymark the upper limit of leucogranites, but not abreak in metamorphic grade. The two detachmentsmerge into one major ductile–brittle shear zonenear the northern limit of Rongbuk valley (Fig. 3;Carosi et al. 1998, 1999b; Searle 1999a). Becausethe QD dips more steeply than the LD, a north-ward-tapering wedge of Everest Series is mappedbetween the two detachments (Searle et al. 2002,2003; Searle 2003).

Previous detailed kinematic investigations arelimited to the Rongbuk valley located on the northside of the Everest massif (Fig. 3). Cross-girdlequartz c-axis fabrics from GHS rocks exposed inthe Rongbuk valley demonstrate that penetrativedeformation, along at least this local section ofthe STDS, occurred under approximately planestrain conditions, and their asymmetry confirmsthe top-to-the-north shear sense (Law et al. 2004).Vorticity analysis (using three techniques) onreconnaissance samples, collected from the top ofthe GHS in the Rongbuk area, indicate pure shearcomponents representing c. 13–53% of the totalrecorded deformation, depending on rock type andstructural position (Law et al. 2004). Integrationof strain and vorticity data, in the reconnaissancesamples, indicated a shortening of 10–30% perpen-dicular to the upper surface of the GHS and, as pre-viously suggested by Grasemann et al. (1999) forthe MCTZ at the base of the slab in NW India(see below), confirmed that the STDS is a stretchingfault (in the sense of Means 1989) with estimated


down-dip stretches of 10–40% (assuming planestrain deformation as demonstrated by petrofabricresults) parallel to the flow plane-transportdirection.

A c. 100 m thick mylonite zone, capped by abreccia zone of variable thickness, characterizesthe uppermost section of the GHS in Rongbukvalley and projects c. 35 km southwards to thesummit of Mount Everest. Structure contours ofthe detachment (QD) on the summit of Everest,and two peaks to the north (Changtse, 7583 m,and Chang Zheng, 7583 m), suggest the detachmentdips c. 108 NNE on the summit and shallows to c. 58NNE in the northern limits of Rongbuk valley(Fig. 3). On Mount Everest and Changtse, the QDseparates Tethyan limestone of presumed early–middle Ordovician age (Yin & Kuo 1978) abovefrom underlying marble of the Yellow Band(Everest Series) (Burchfiel et al. 1992; Searle1999a, 2003). On the NE ridge of Everest, the QDis marked by a 5–40 cm thick breccia zone in thebasal limestone, which rests on intensely foliatedYellow Band marble containing shear bands anddrag folds (Sakai et al. 2005).

The structurally highest section of the Everest–Lhotse massif is predominantly composed ofgreenschist to lower-amphibolite facies EverestSeries metasedimentary rocks, while the lower ram-parts consist of sillimanite-grade schist that gradesinto migmatitic gneiss (Fig. 3; Lombardo et al.1993; Pognante & Benna 1993; Carosi et al. 1998,1999b; Searle 1999a, b; Searle et al. 2003).A variably deformed leucogranite sill complexthat parallels the pervasive fabric within theserocks is limited to a zone immediately below theEverest Series. Searle (1999a) proposed thatthe LD is present along this transition and alsoproposed that a late-stage thrust (Khumbu thrust)is present along the base of the underlying, most

extensive, leucogranite sill complex (Fig. 3). Vari-ably migmatized, interlayered gneiss, calc-silicate,quartzite, schist and orthogneiss are predominantbeneath the LD (or the composite LD–QD in thenorthern Rongbuk valley), and extend downwardsthrough the middle section of the GHS to theupper section of the MCTZ (Lombardo et al.1993; Searle et al. 2003). Deformation within thecore of the GHS is characterized by severalphases of folding that culminate in a pervasivefoliation that is broadly warped by late-stage NW-and NE-trending hinge lines of recumbent foldsthat create dome structures (Carosi et al. 1999a,b). Mylonite zones, typically found at the marginsof the slab, are absent in the core.

As exposed in the Duhd Kosi drainage southof Everest, the MCTZ consists of sheared quart-zite, calc-silicate, amphibolite, garnet–kyanite–staurolite schist, graphitic schist and augen gneiss(Hubbard 1988, 1989; Catlos et al. 2002). Althoughdebate continues about the exact location of thethrust zone (see Godin et al. 2006b; Searle et al.2006), we choose to relate our vorticity resultsto the original context proposed by Hubbard(1988, 1989). In the Duhd Kosi drainage, a com-bined downward increase in apparent penetrativestrain, and upward increase in metamorphic temp-eratures that exceed the kyanite stability field,marks the top (MCT) of the 5 km thick high-strainzone, while the Okhandunga orthogneiss marksthe base (MCT I; Fig. 4). Because the apparentincrease in strain coincides with a change in rocktype (migmatitic gneiss above and pelitic schistbelow), the downward increase in penetrative foli-ation intensity may be controlled by lithologyrather than structural position. The pervasivenorth-dipping foliation overprints several phasesof folding and foliation development that are onlypreserved in lower- to moderate-strain domains

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within the high-strain zone. Variation in foliationorientation is the result of late-stage folding alsopresent at structurally higher positions in the coreof the GHS.

Vorticity analysis

Introduction to techniques

Mean kinematic vorticity number (Wm) is ameasure of the relative contributions of pure(Wm ¼ 0) and simple (Wm ¼ 1) shear. Severalanalytical methods exist for estimating Wm inhigh-strain rocks; however, only in rare cases areindividual samples suited for vorticity analysisusing multiple techniques (see Law et al. (2004)for detailed discussion). We focus on a suite ofsamples collected from the Everest transect thatare suitable for the rigid grain-based vorticity analy-sis developed by Wallis et al. (1993) and Wallis(1995). This suite of 51 samples (Table 1) includesthe seven reconnaissance samples described by Lawet al. (2004) from the Rongbuk valley–Changtseridge part of the transect, some of which had pre-viously proved suitable for multiple methods ofvorticity analysis.

Using the founding principles of Ghosh &Ramberg (1976) and Passchier (1987), Walliset al. (1993) and Wallis (1995) proposed that, forrigid clasts rotating in a flowing ductile matrix, aunique relationship exists between Wm, clastaspect ratio (R) and the angle (u) between clastlong axes and matrix foliation. For a given Wm,clasts with a specific aspect ratio will reach aunique stable sink position (i.e. angle from thefoliation). The method involves measuring theclast aspect ratio and angle between the clast longaxis and foliation for both back- and forward-rotated clasts (in sections cut perpendicular to thefoliation and parallel to the macroscopic stretchinglineation). The distribution of clasts is displayed ona plot of R versus u (Fig. 5a–d). A transitionbetween clasts that rotate infinitely and those thatreach a stable sink orientation defines the criticalthreshold (Rc). Rc is then used to calculate Wmusing the relationship proposed by Passchier(1987):

Wm ¼ ðR2c � 1Þ=ðR2

c þ 1Þ (1)

In practice, a range of likely Rc values is usuallyindicated for a given sample using the Wallis plot,leading to a range of estimated Wm values (Lawet al. 2004; see also Carosi et al. 2006; Xypolias& Kokkalas 2006). Whether the Wallis methodmay consistently under- or overestimate Wm

values probably depends on individual samplecharacteristics. The method may tend to underesti-mate Wm if clasts of large aspect ratio are notpresent, and in such samples the upper bound ofthe estimated Wm range is probably closest to thetrue value (Law et al. 2004). In contrast, if finitestrains are low, then clasts of high aspect ratiomay not have had time to reach stable sink orien-tations and the observed range of Rc values wouldtend to overestimate Wm (Bailey et al. in press).

The rigid grain vorticity method assumes: (1) thatthe clasts undergo no internal deformation (e.g. bycrystal plasticity or pressure solution); (2) no mech-anical interaction occurs either between adjacentrotating clasts, or between the clasts and theirmatrix; and (3) high enough strain has developedto ensure that all clasts have rotated into theircurrent position. Samples were avoided wheredeformation temperatures exceeded the onset ofinternal plastic deformation within otherwise rigidphases, or where excessive interaction had occurredbetween rotated grains. We tentatively assumeplane strain deformation for these samples, basedon the original petrofabric data of Law et al.(2004, p. 313), which strongly indicated that flowwas monoclinic to orthorhombic (Law et al. 2004,p. 314). Due to the lack of robust strain markers,it was impossible to quantify strain in any of thesamples from the Everest transect aside fromthose published by Law et al. (2004). Rigid graindata plots for all 51 samples used for vorticityanalysis are reproduced in the Appendix to thispaper. Full details of the mineral(s) used as rigidgrain markers are given on each plot. Details ofsample locations, and estimated range of Wmvalues for each sample, are summarized in Table 1.

Representative rigid grain data plots for

different structural levels

Four representative samples are used to describeand discuss the characteristics of the major rocktypes within the Everest transect used for vorticityanalysis: (1) sheared Tethyan limestone abovethe QD system; (2) biotite gneiss/schist within theuppermost 100 m of the footwall to the compositeQD–LD system; (3) high-grade gneiss at a deeperstructural level (,2 km) in the footwall to theLD; and (4) pelitic rocks within the MCTZ(Fig. 5a–d).

Tethyan limestone Sample GB-25/3, collectedfrom just below the summit (8836 m) by EdmundHillary (Harker Collection records, CambridgeUniversity) during the first accent of MountEverest via the South Col in 1953, is an example


Table 1. Mean kinematic vorticity (Wm) data

SampleRocktype�

Elevation(m)

Distance (m)from QD–LD

Method 1(Wm)

Northern transectR03-10 lim 5010 10 above 0.87–0.91R03-12 mar 5000 0 0.88–0.91R03-15 calc 4995 5 0.75–0.78R03-16 calc 4993 7 0.68–0.79R03-17 qtz 4991 9 0.81–0.84R03-18 calc 4988 12 0.73–0.80R03-18 (A) calc 4988 12 0.75–0.80R03-19 calc 4987 13 0.62–0.71R03-20 leu 4982 18 0.69–0.80R03-21 leu 4979 21 0.76–0.80R03-24 leu 4974 26 0.76–0.79R03-25 leu 4965 35 0.75–0.81R03-26 leu 4964 36 0.76–0.82R03-26 (A) bt 4934 46 0.74–0.79

Rongbuk Monastery transectR03-55 lim 5767 67 above 0.82–0.84R03-56 lim 5716 16 above 0.77–0.80R03-58 lim/mar 5663 37 0.76–0.79R03-59 mar 5655 45 0.79–0.82R03-63 calc 5650 50 0.81–0.84TI-05 bt c. 5600 100 0.77–0.79ET-15 bt 5450 250 0.82–0.85R03-67 calc 5380 320 0.75–0.80ET-14 bt 5350 350 0.67–0.73R03-70 bt 5255 445 0.73–0.77ET-13 bt 5250 450 0.75–0.80ET-12 bt 5100 600 0.72–0.77

Hermit’s Gorge transectR03-46 lim 5748 5 above 0.74–0.77R03-44 mar 5739 0 0.74–0.77R03-43 mar 5737 2 0.72–0.75R03-39 calc 5698 41 0.64–0.70R03-38 bt 5688 51 0.75–0.78ET-08þ bt 5650 89 0.79–0.84R03-31 leu 5950 211 0.57–0.64R03-33 bt 5398 341 0.69–0.80

Everest & Kangshung valley transects25/3 Hillary lim 8836 0.87–0.89E-03-01 Hamilton lim 8840 0.87–0.89ME-124 Wager lim 8568 0.84–0.8625/1&2 Evans lim 8689 0.85–0.87ME-125 Wager lim 8260 0.83–0.85ET-10 bt .7000 0.81–0.84ET-11 bt .7000 0.77–0.79K04-03 gn 5340 0.68–0.77K04-04 gn 5320 0.72–0.76

Main Central Thrust zone transectET-41 grt-sill 0.63–0.73ET-44 grt-sill 0.63–0.7085-H-22E grt-sill 0.72–0.7785-H-21J grt-ky 0.70–0.7285-H-21G grt-ky 0.60–0.6487-H-6B grt 0.69–0.7187-H-5A gn 0.69–0.7787-H-1B grt 0.66–0.70

�bt, biotite schist; calc, calc-silicate; gn, gneiss; grt, garnet schist; grt-ky, garnet þ kyanite schist; grt-sill, garnet þ sillimanite schist; leu,leucogranite; lim, limestone; mar, marble; qtz, quartz-rich layer in calc-silicate.


of sheared limestone in the immediate hanging wallto the QD (Fig. 3). Other samples that share thesame microstructural characteristics and spatialproximity to the QD, were collected inthe sidewalls of the Rongbuk valley. Abundantequant–elongate detrital quartz grains are inter-preted as rigid clasts that rotated within a ductilecalcite matrix (Fig. 6a). The distribution of quartzgrains on the Wallis plot defines (Fig. 5a) anabrupt transition from grains that rotate infinitely(R � 3.80) to those that reach a stable sink orien-tation (R � 4.05). Using this range in Rc values(3.80–4.05) yields a Wm estimate of 0.87–0.89(c. 32–28% pure shear).

Immediate footwall to LD and composite QD–LDsystem Sample R03-38 represents the amphibolite-facies rocks (marble, calc-silicate, leucogranite andbiotite schist/gneiss) within the upper 100 m ofthe GHS. These samples often contain severalrigid phases such as feldspar, epidote, zircon,amphibole and tourmaline in a matrix of dynami-cally recrystallized quartz. At the upper limit tothese deformation temperatures (amphibolitefacies), large feldspar porphyroclasts remain rigidwhile smaller grains begin to deform internally.R03-38 is a biotite schist with abundant feldsparand tourmaline suitable for rigid grain analysis(Fig. 6b). The narrow range in Rc (Fig. 5b) yieldsa fairly robust Wm estimate of 0.75–0.78(45–42% pure shear).

Structurally deeper levels of LD footwall SampleK-04-03 was collected from outcrops of high-grade gneiss in the western end of the Kangshungvalley. These gneisses are situated at a structuraldepth of c. 2 km beneath the LD (Fig. 3) andcontinue downward into the underlying anatecticcore of the GHS. Feldspar grains in these gneissesare generally separated from each other by amatrix of biotite laths and dynamically recrystal-lized (Regime 3 of Hirth & Tullis 1992) quartz(Fig. 7a). However, many of the feldspar grainsexhibit at least moderate undulatory extinctionand minor grain flattening, indicating that they didnot behave as perfectly rigid markers. ‘Rigidgrain’ plots using these feldspar grains are charac-terized by a broad transition in potential Rc values(Fig. 5c), and therefore greater uncertainty in defin-ing Wm. We propose that these plots are typicalof samples that contain a semi-rigid phase, andcaution against overinterpretation of Wm estimatesfrom such samples.

MCTZ The fourth example, sample ET-41, is agarnet–mica schist typical of pelite samples col-lected from the MCTZ. These pelite samplescontain elongate garnet porphyroclasts that arewrapped by biotite and muscovite, and surroundedby a matrix of quartz and feldspar (Fig. 7b). The

tourmalinegarnet

Rc = 2.10

Rc = 2.55

90°

60°

30°

0°

3.0 4.0 5.0 6.0

R

1.0

1.0

-30°

-60°

-90°

angl

e be

twee

n cl

ast l

ong

axis

and

folia

tion

n = 227 Wm = 0.63 - 0.73(d) ET-41

tourmalinefeldspar

Rc = 2.65

Rc = 2.85

90°

60°

30°

0°

3.02.0 4.0 5.0 6.0

R

1.0

(b) R-03-38-30°

-60°

-90°

angl

e be

twee

n cl

ast l

ong

axis

and

folia

tion

n = 200 Wm = 0.75 - 0.78

quartz clast in calcite matrix

Rc = 3.80

Rc = 4.05

90°

60°

30°

0°

4.03.0 5.0 6.0

R

1.0

(a) GB-25/3 - Hillary-30°

-60°

-90°

angl

e be

twee

n cl

ast l

ong

axis

and

folia

tion

n = 208 Wm = 0.87 - 0.89

tourmalinefeldspar

Rc = 2.30

Rc = 2.80

90°

60°

30°

0°

3.0 4.0 5.0 6.0

R

(c) K-04-03-30°

-60°

-90°

angl

e be

twee

n cl

ast l

ong

axis

and

folia

tion

n = 200 Wm = 0.68 - 0.77

Fig. 5. Rigid grain plots using the Wallis et al.(1993) and Wallis (1995) technique (see text for details).Four representative plots are used to discuss themain rock types used in this investigation: (a) shearedlimestone in the hanging wall of the Qomolangmadetachment; (b) mylonitic metapelites, calc-silicates andleucogranites from the upper 600 m of the GHS; (c)GHS gneiss sample from Kangshung valley withbroad range of potential Rc values which are typicalof samples where the originally rigid phase(feldspar) begins to deform internally; (d) garnet schistfrom the Main Central thrust zone where rigidelongate garnets were used to estimate Wm.


Fig. 6. (a) Photomicrograph (crossed polars) of sample GB-25/3 (collected by E. Hillary in 1953 at c. 8836 m)showing microstructures typical of sheared limestone collected near the Qomolangma detachment. Abundant detritalquartz grains act as the rigid phase rotating in a calcite (Cal) matrix. Some randomly orientated white mica (M) ispresent. Rigid grain plot of the sample is shown in Figure 5a. (b) Photomicrograph (crossed polars) of sample R03-38;section cut perpendicular to the foliation and parallel to the lineation. Microstructures include rigid feldspar (Fs)rotating in a ductile quartz (Qtz) matrix. Large feldspar porphyroclasts in the centre of the image have an aspect ratioof c. 1.6 with a long axis c. 808 from the foliation as defined by aligned white mica (M). Rigid grain plot of the sample isshown in Figure 5b. Sample cut perpendicular to foliation and parallel to lineation; plunge and trend indicated.


evolution of these elongate garnets is discussedin detail below. The orientation distribution andrange in aspect ratio of these garnets confirmstheir appropriateness for rigid grain vorticityanalysis (Fig. 5d). A limited range in Rc definesWm estimates of 0.63–0.73 (58–48% pure shear).The major drawback to using metamorphic phasesfor rigid grain analysis in these pelitic MCTZsamples is the number of appropriate porphyro-blasts available within a thin section; where poss-ible we have used combined data from parallelsections in individual samples. For example,sample 85-H-21G contained a minimal number ofgarnet porphyroblasts (n ¼ 59) that just begins todefine a minimum Rc, whereas 87-H-22E containedmany garnets (n ¼ 275) that define Rc much better(Appendix, Sheet 5). For several MCTZ samples,such as ET-41, the rigid grain analysis provedhighly successful and provides a unique opportunityto explore the relationship between peak meta-morphism and mylonite formation.

Petrography and results of vorticity

analyses

Rongbuk valley transects

We collected orientated samples for vorticityanalysis along three transects in the eastern side-walls of the Rongbuk valley (Fig. 8). A similarlithotectonic sequence is observed in each transectconsisting (traced structurally downwards) oflimestone, marble, calc-silicate, leucogranite andbiotite–sillimanite schist/gneiss (Figs 8 & 9).

Tethyan limestone forms the structurally highestlithotectonic unit and is truncated along the base bythe underlying QD. A 5–10 m thick section ofmarble marks the upper limit to pervasive ductiledeformation beneath the detachment. Lensesof mylonitic leucogranite are commonly foundwithin the sheared marble, demonstrating thatductile deformation outlasted their emplacement(c. 17 Ma; Murphy & Harrison 1999). Interlayeredand pervasively foliated calc-silicate and quartzo-feldspathic layers, defined in outcrop by alternatingblack/green and white layers, are present belowthe sheared marble. Dark layers contain diopsideand are either amphibole- or tourmaline-rich,while white layers contain abundant quartz andfeldspar. Feldspar is commonly fractured andwithin one thin section a complete gradation fromangular clasts to rounded porphyroclasts rotatingin a quartz matrix is common.

Microstructures in quartz-rich layers include thedevelopment of subgrains and bulging grain bound-aries, which indicate dynamic recrystallizationunder Regime 2–3 conditions as defined by Hirth

& Tullis (1992), and suggest deformation tempera-tures of c. 490–5308C (Stipp et al. 2002). Micro-boudinage of diopside, garnet and tourmalinegrains suggests that some components of fabricdevelopment post-dated their growth. Tensiongashes, nearly perpendicular to the NNE- orSSW-trending stretching lineation, suggest that aprogression in deformation mechanisms fromductile to brittle occurred during exhumation ofthe GHS.

Structurally beneath the calc-silicate layers(at least at the northern end of the Rongbukvalley) is a 10–20 m thick mylonitic leucogranitesill complex. Quartz and feldspar record evidencefor grain-scale processes operating at similardeformation conditions to those indicated in theoverlying calc-silicate-rich unit. S-C fabrics withextensional shear bands dominate the detachment-parallel sills.

The structurally lowest unit exposed in theRongbuk valley is composed of biotite schist(Rongbuk Formation of Carosi et al. 1998,1999a) that is migmatized and injected by foli-ation-parallel, variably deformed, leucogranitelenses and sills; cross-cutting leucogranites areless commonly observed (Searle et al. 2006, figs6 & 7). Based on quartz c-axis fabric openingangles, Law et al. (2004) documented progressiveincreasing deformation temperatures of 525–625 + 508C in the biotite schists at depths of300–650 m beneath the mapped position of theLD in the Rongbuk Monastery and Hermit’sGorge transects (see below). Rotation of the rigidgrains used as vorticity markers in this paper,either pre-dated or (more likely) were synchronouswith plastic flow of the quartz-rich matrix associ-ated with these deformation temperatures. Fibrolitein the biotite schist is drawn into extensional shearbands, but remains pristine, suggesting that shearband development occurred in the sillimanite stab-ility field (Law et al. 2004). At depths greater than100 m beneath the composite QD–LD system,feldspar begins to deform plastically (as indicatedby undulose extinction) and grains tend to becomemore elongate and orientated subparallel to foli-ation. At a given depth, this brittle–plastic tran-sition in feldspar deformation seems to be grain-size controlled (Law et al. 2004, p. 311). Incipientconjugate sets of shear bands, defined by biotite,create a lattice network that dominates the micro-structure in the structurally deeper samples. Poly-gonal quartz grains are common, suggesting acomponent of annealing. Many of these structu-rally deeper samples are unsuited for vorticityanalysis (as discussed above for sample K-04-03). However, even at depths of 600 m beneaththe detachment, samples with limited evidencefor internal deformation of feldspar yield a


Fig. 7. (a) Photomicrograph (crossed polars) of sample K04-03 collected in Kangshung valley, Tibet. Biotite (Bt)defines the foliation that is aligned NW–SE in the image. Irregularly shaped feldspar (Fs) that begins to align with thefoliation suggests high deformation temperatures where feldspar begins to deform internally. Rigid grain plot of thissample is shown in Figure 5c. These microstructures typify samples from the core of the Greater Himalayan Slab thatare unsuited for rigid grain analysis. (b) Photomicrograph (crossed polars) of sample ET-41 from the Main Centralthrust zone (Fig. 4). Garnets (Grt) of variable aspect ratios and angles from the foliation are present. Quartz (Qtz)inclusions are present in several of the garnet cores. Aligned biotite (Bt) and white mica (M) define the foliation(east–west in image). Rigid grain plot using garnet porphyroblasts is shown in Figure 5d. Details of garnet evolutionare discussed in the text. Sample cut perpendicular to foliation and parallel to lineation; plunge and trend indicated.


Fig. 8. Simplified geological map of Rongbuk valley, Tibet. Insets (a–c) are enlargements of detailed sampletransects. Spatial distribution of samples is also shown on the cross-section through each transect. North is obliqueto the long axis of the figure. Image compilation created using original mapping from this investigation and othersources (Burchfield et al. 1992; Murphy & Harrison 1999; Searle et al. 2003; Law et al. 2004).


well-defined Rc threshold, and therefore a mean-ingful Wm estimate.

Below we summarize the results of vorticity ana-lyses in our three transects through the eastern side-walls of the Rongbuk valley (Fig. 8); each transectbegins in the sheared limestone or within the com-posite QD–LD system and progresses downwardsinto the migmatitic biotite schist. Results are pre-sented on plots of Wm versus relative distancebelow the QD to show the spatial distribution ofWm domains in each transect (Figs 10 & 11).Because the location of the QD is more readilydetermined in the field than the LD, and in thenorthern section of Rongbuk valley the LD eithermerges with or is cut out by the QD, we use theQD as a reference structural level in these plots.

Northern transect Vorticity analysis results fromthe northern transect (Fig. 8 inset a & Fig. 9) areshown in Figure 10a. Sample R03-10 (limestone)and sample R03-12 (marble), collected c. 10 mabove and within the QD, respectively, yield Wmestimates of 0.87–0.91 and 0.88–0.91, indicatingthe lowest component of pure shear (30–25%) forthe entire transect. Four out of six calc-silicatesamples from below the sheared footwall marbleyield a range in Wm of 0.68–0.80 (c. 52–40%pure shear). The other two calc-silicate samplesare outliers to this trend and yield slightly higher(R03-17: Wm ¼ 0.81–0.84, 40–35% pure shear)and lower (R03-19: Wm ¼ 0.62–0.71, c. 58–49%pure shear) Wm estimates. The large range in

potential Rc values recorded by calc-silicatesamples R03-16 and 19 (together with leucogranitesample R03-20) suggests they are less suitable forrigid grain analysis than the other samples. Five leu-cogranite samples yield a range in Wm estimates thatis consistent with the majority of the calc-silicatesamples (0.70–0.82). Although large, the range inWm for sample R03-20 overlaps with Wm valuesin the calc-silicate and leucogranite samples. Thesingle biotite schist sample (R03-26A) at the baseof the transect has a narrow range in estimatedWm values (0.74–0.79) that is indistinguishablefrom the calc-silicate and leucogranite samples.The sheared limestone and marble in the immediatehanging wall and footwall to the QD yield thehighest Wm values (0.87–0.91), and thereforehighest percentage simple shear values, recordedin the Rongbuk valley transects. At distances ofc. 10–46 m beneath the detachment, the majorityof samples yield Wm estimates of 0.70–0.80(c. 50–40% pure shear).

Rongbuk Monastery transect The RongbukMonastery transect is located c. 7 km to the southof the northern transect (Fig. 8, inset b). Three lime-stone samples at the top of the transect (R03-55, 56and 58) yield Wm estimates of 0.76–0.84, with thegreatest simple shear component (c. 63%) recordedin the structurally highest sample (Fig. 10b). Theone marble sample (R03-59), located beneath thedetachment, yields a Wm estimate (0.79–0.82)that is indistinguishable from Wm values for the

Fig. 9. Photograph of the composite Qomolangma and Lhotse detachments (black line) where they are proposed tomerge in the northern limits of Rongbuk valley. Location where image was taken is shown as solid black star onFigure 8a. The northern transect is located where the road and detachment are closest. The general rock types fromstructurally highest to lowest are: (1) limestone, (2) marble, (3) calc-silicate, (4) leucogranite, and (5) migmatizedbiotite schist (also known as Rongbuk Formation). View is towards the NE. An arrow points to jeep on two-lane dirtroad for scale.


limestone sample above and the calc-silicate samplebelow (R03-63: Wm ¼ 0.81–0.84). Five of the sevensamples below calc-silicate R03-63, including onecalc-silicate (R03-67), one hornblende–epidoteschist (TI-5), and three biotite schist samples,record a fairly consistent range in estimated Wmvalues (0.72–0.80). The two outliers yield slightlyhigher (leucogranite ET-15: Wm ¼ 0.82–0.85)

and lower (biotite schist ET-14: Wm ¼ 0.67–0.73)Wm estimates. Samples TI-5, ET-14, ET-13 andET-12 also proved appropriate for several othervorticity analysis techniques (Law et al. 2004),referred to in Figure 10b as method II (the PHDmethod of Simpson & De Paor 1997) and methodIII (the combined strain and quartz c-axis fabricmethod of Wallis 1995). For TI-5, the rigid grain

Rigid grain technique Vorticity method II (Law et al. 2004)Vorticity method III (Law et al. 2004)

L: limestoneM: marbleC: calc-silicateG: leucogranite A: amphibole schistB: biotite gneiss/schist

DETACHMENT

R-03-16

R-03-15

R-03-17

R-03-18

R-03-18A

R-03-19

R-03-20

R-03-21

R-03-24

R-03-25

R-03-26

0.90.80.70.6

mean vorticity number - Wm

1.00.5

C

C

C

G

G

G

5060 40 30 20 0

percent pure shear

10

C

R-03-26A

G

G

B

R-03-12

R-03-10

C

C

Altitude (m)

5010

5000

4995

4993

4991

4988

4988

4987

4982

4979

4974

4965

4964

4934

(a) NORTHERN TRANSECT

M

L

R-03-56 5716

5767

R-03-58

R-03-63

R-03-59

R-03-67

R-03-70

TI-5

ET-15

ET-14

ET-13

ET-12

5663

5665

5655

5380

5250

~ 5600 (talus)

5450

5350

5255

5100

0.90.80.70.6


1.00.5

A

L

B

B

B

5060 40 30 20 0

percent pure shear

10

L

L

B

Altitude (m)

R-03-55

C

MDETACHMENT

G

C

(b) RONGBUK MONASTERY

Fig. 10. Bar charts for range of mean kinematic vorticity numbers (Wm) estimated by the rigid grain method forsamples collected in the Northern (a) and Rongbuk Monastery (b) transects, Rongbuk valley, Tibet. Sample locationsshown in Figure 8. The range of Wm values estimated by alternative methods (Law et al. 2004) is also indicated.


technique of Wallis et al. (1993) and method IIyield indistinguishable results. For the other threesamples, method III consistently yields higher Wmestimates than the rigid grain technique (see Lawet al. (2004) for detailed discussion). In summary,rigid grain analyses from the Rongbuk Monasterytransect yield Wm estimates of 0.72–0.84 (c. 48–36% pure shear) and represent deformation con-ditions to a maximum depth of c. 600 m beneaththe composite QD–LD fault system. We regardthe structurally deepest samples as yielding theleast reliable Wm estimates, as all size fractions offeldspar grains display at least limited evidencefor crystal plasticity, and thereby undermine the

fundamental assumptions of the rigid graintechnique.

Hermit’s Gorge transect Our third transect islocated in Hermit’s Gorge (and one of its sidevalleys) which intersects the Rongbuk valley atEverest Base Camp (Fig. 8, inset c). One shearedlimestone sample (R03-46) was collected c. 5 mabove the top of the marble section and presumedlocation of the QD. It yields a narrow range inWm estimates of 0.74–0.77 (c. 45% pure shear)that is indistinguishable from the two marblesamples (R03-43 and 44) below. The single calc-silicate sample (R03-39) yields a Wm estimate

Fig. 11. Bar charts for range of mean kinematic vorticity numbers (Wm) estimated by the rigid grain methodfor samples collected in the Hermit’s Gorge (a) and Mount Everest & Kangshung valley (b) transects, Tibet.Sample locations shown on Figures 8, 12 & 13. The range of Wm values estimated by alternative methods (Law et al.2004) is also indicated.


(0.64–0.70) that is significantly lower than both themarble above and biotite schist below (R03-38:0.75–0.78). Sample ET-8, a biotite-rich psammite,yields the highest Wm estimate of the entire transect(0.79–0.84); in contrast, method III analysis on thissample yields higher estimated Wm values (Lawet al. 2004), as noted for samples from theRongbuk Monastery traverse. R03-31, a piece ofmylonitic leucogranite float collected at an altitudeof c. 5950 m, yields the lowest Wm estimate of0.57–0.64. Although collected at the highest alti-tude of the transect, due to its position on thesouth side of the gorge and the northerly dip ofthe structural units, this sample probably comesfrom a relatively deep structural position. The struc-turally lowest sample (R03-33), collected near themouth of Hermit’s Gorge at c. 340 m beneath theQD, yields the largest range in estimated Wm

values (0.69–0.80) for the transect. We attributethe large range in uncertainty of Rc (and henceWm) for this sample to the onset of plastic defor-mation in the feldspar marker grains. This sampleis probably close to the maximum structural depthfor robust rigid grain vorticity analysis.

Summit of Mount Everest–Kangshung

valley

Our final transect across the top of the GHS iscomposed of Tethyan limestone and Yellow Band(Everest Series) marble samples from near thesummit of Mount Everest, samples of EverestSeries interlayered pelite and calc-mylonite col-lected from talus piles at Advance Base Campbeneath the North Col–Changtse Ridge, and

Fig. 12. Simplified geological map of the Mount Everest massif and Kangshung valley, Tibet. Compilation basedon mapping during this project (Kangshung valley) and Searle et al. (2003). QD, Qomolangma detachment; LD, Lhotsedetachment.


samples of high-grade gneiss from outcrops in thewestern end of the Kangshung valley (Figs 3, 12& 13). Only the Kangshung valley samples,discussed above (Fig. 5), are orientated.

The highest altitude sample (GB-25/3), from theHarker Collection at Cambridge University, is ofTethyan limestone collected by Edmund Hillaryon 29 May 1953 at ‘40 feet beneath the summitof Mount Everest’ (Harker Collection records).This sample is augmented by a second, lithologi-cally identical, summit sample (E-03-01) collectedby Scottish alpinist David Hamilton in 2003. Ourthird, and structurally deepest, Tethyan limestonesample (ME-124), from the Lawrence Wager Col-lection in the Oxford University Museum ofNatural History, was collected by Wager ‘from aband forming the First Step’ (Wager Collectionrecords; see also Wager 1934, 1939) on the NEridge of Everest during the 1933 Everest expedition.The highest altitude Yellow Band marble sample(GB-25/1þ 2), two pieces of intensely foliatedand lineated white calc-mylonite from the HarkerCollection, was collected by Charles Evans on 26May 1953 at ‘approximately 28500 feet’ (HarkerCollection records) on the SE ridge of Everest. Ourstructurally deeper Yellow Band sample (ME-125)was collected by Wager from a ‘typical yellow schis-tose marble forming Yellow Band’ on the NE ridge

at approximately 300 feet beneath the 1933 CampVI (Wager Collection records).

Tethyan limestone Sheared Tethyan limestonesamples (GB-25/3, E03-01, ME-124) contain abun-dant white mica laths and subangular–subroundeddetrital quartz grains set in a calcite matrix(Fig. 6a). The calcite matrix grains are completelyrecrystallized, and no remnants of a sedimentaryfabric have been preserved (J.F. Read, pers.comm. 2006). The calcite grains are equant–slightly elongate in cross-section, and an incipientfoliation is defined by weak preferred orientationof the more elongate matrix grains, together withaligned films of an extremely fine-grained opaquephase. Calcite- and quartz-filled microfaults trun-cate the incipient foliation at moderate to highangles, particularly in sample ME-124 collectedfrom immediately above the QD (Fig. 13).Anastomosing quartz-filled fractures subparallel tofoliation are also present.

The matrix calcite grains range in size from 20 to50 mm. Larger single and polygonal calcite grains(200–250 mm), together with randomly orientatedwhite mica laths (up to 100 mm in length) andequant–elongate detrital quartz grains (generally40–80 mm long), are scattered throughout thematrix (Fig. 6a). Weak undulose extinction within

Fig. 13. Photograph of the summit of Mount Everest (viewed towards the east) taken from Renjo La (5340 m), Nepal,using a 300 mm lens. Qomolangma detachment is highlighted by line and separates Tethyan limestone (1) abovefrom Yellow Band marble (2) and Everest Series (3) below. Approximate locations of samples collected by Wagerin 1933 (ME-124 and 125), Evans in 1953 (GB-25/1 and 2), Hillary in 1953 (GB-25/3), and Hamilton in 2003(E-03-01) are indicated. Photomicrograph of summit sample collected by Hillary is shown in Figure 6a, and thecorresponding rigid grain plot is shown in Figure 5a.


the quartz grains suggests a minor component ofplastic deformation, and a high concentration offluid inclusions gives a dusty appearance to someof these grains. E-twins in the larger calcitegrains are straight and thin (,5 mm), suggestingdeformation temperatures ,170–2008C (Burkard,1993; Ferrill et al. 2004). The presence of slightlywider twins (.5 mm) in some of the smallermatrix grains suggests that deformation tempera-tures may have reached .2008C (Burkhard 1993;Ferrill et al. 2004). Observed microstructures, andwell-defined Rc values in all three of these sub-greenschist-facies Tethyan ‘limestone’ samplesindicate that the detrital quartz grains acted as atleast semi-rigid clasts rotating in a plasticallyflowing and dynamically recrystallizing calcitematrix. Wm estimates (Fig. 11b) in these samplesrange from 0.84 to 0.89 (35–30% pure shear).

The pristine grain boundaries of the white micalaths in the limestone suggest that they may haverecrystallized during deformation, rather thanbeing of detrital origin (G. Oliver pers comm.2006). We attribute the lack of a well-developedgrain-shape foliation within the calcite matrix,together with the lack of any sedimentary struc-tures, to the operation of grain boundary migrationrecrystallization, as indicated by the observedmicrostructures in this sample.

Samples of Everest summit ‘limestone’ havepreviously been described by Gansser (1964,pp. 164–171) and Sakai et al. (2005). The micro-structures described by Gansser (includingsamples originally described by Gysin & Lombard1959, 1960) are very similar to those recorded inour samples, except for the presence of crinoidfragments (see also Odell 1965). In contrast, thesample described by Sakai et al. (2005), andcollected at c. 6 m beneath the summit (8850 m),contains crinoid, brachiopod and trilobite frag-ments, and seems to be much less extensivelysheared and recrystallized.

Everest Series, Yellow Band marble Microstruc-turally, the most obvious difference between thesummit limestone and the underlying YellowBand marble is the change in size of recrystallizedmatrix calcite grains, which abruptly increasesfrom 20–50 mm in the Tethyan limestone abovethe QD to 150–200 mm in the Yellow Bandmarble beneath the detachment. The calcitegrains are equant to slightly elongate and define aweak foliation in thin section that is parallel tothe strong macroscopic foliation. Larger singlecalcite grains (400–800 mm), together with ran-domly orientated white mica laths (up to 100 mmin length) and equant–elongate detrital quartzgrains (generally 25–80 mm long), are scattered

throughout the matrix. Calcite twins are thickerand more closely spaced than in the Tethyan lime-stone, and both multiple twin sets and tightchevron-style buckling of twin lamellae are com-monly developed in the larger calcite grains (par-ticularly in sample GB-25/1þ 2). The presence ofthick twins and microstructural evidence for wide-spread calcite recrystallization involving grainboundary migration indicates deformation tempera-tures .2508C (Ferrill et al. 2004). However, the det-rital quartz grains exhibit very little unduloseextinction, and appear to have acted as semi-rigidporphyroclasts in the flowing calcite matrix, indicat-ing deformation temperatures ,300–3508C (i.e.below generally accepted minimum temperaturesfor onset of plastic deformation in quartz at naturalstrain rates; see Stipp et al. (2002) and referencestherein). Wm values of 0.83–0.87 (36–32% pureshear) are indicated for samples 25/1þ2 andME-125 using the detrital quartz grains as rigidmarkers (Fig. 11b).

Structurally deeper levels of Everest Series Talussamples of biotite-grade interlayered phyllite–psammite and calc-mylonite (ET-10 and ET-11),shed from the structurally deeper sections of theEverest Series exposed on the North Col–ChangtseRidge (Figs 3 & 12), clearly indicate a strongmatrix control on deformation mechanisms operatingin detrital quartz grains. Even at the thin-sectionscale, a strong partitioning of deformation mechan-isms is observed. Detrital quartz grains deformplastically (with minor pressure solution) in thepelite layers when surrounded by phyllosilicates(biotite and white mica), but remain as rigid clastsin the calc-mylonite layers where the calcite grainshave accommodated the penetrative strain. Wmvalues of 0.77–0.84 (42–37% pure shear) areindicated for samples ET-10 and ET-11 using thedetrital quartz grains in the calcite-rich layers asrigid grains (Fig. 11b). The combined strain andquartz c-axis fabric method of Wallis (1995) in thequartz–mica layers yielded Wm estimates of0.91–0.98 (Law et al. 2004) and correspondinglylower pure shear components (Fig. 11b; ET-10,method III).

Main Central Thrust Zone

The base of the GHS is marked by the MCTZ(Fig. 3). In the Everest transect the MCTZ isapproximately 5 km thick, and characterized by ageneral decrease in metamorphic grade towardsdeeper crustal levels (Fig. 4), as constrained bythe appearance of index minerals and geothermo-barometry (Hubbard 1988, 1989; Searle et al.2003). Seven samples of garnet-bearing schist and


Fig. 14. (a) Scanning electron microscope (SEM) image of garnets typical of sample ET-41 from the Main Centralthrust zone. Foliation is defined by aligned muscovite (intermediate grey). Biotite (light grey) forms tails on somegarnet porphyroblasts and also defines the foliation. (b) SEM image of an elongate garnet in sample ET-41. Foliation isoriented east–west in the image. Sigmoidal inclusion trails defined by quartz (dark grey), biotite (intermediate grey),and oxides (white). Inclusion-free rims are preserved on both ends of the garnet. (c) SEM image of another example ofan elongate garnet porphyroblast in sample ET-41. (d) Three-step evolution of elongate garnets used for rigid grainanalysis (see text for details). Rigid grain plot for this sample is shown in Figure 5d and a photomicrograph in Figure 7b.Sample cut perpendicular to foliation and parallel to lineation.


one sample of orthogneiss were selected fromdifferent structural levels of the MCTZ for rigidgrain analysis (Figs 3 & 4). Three basic typesof garnet grains are distinguished in the schist:round garnets, small irregularly shaped garnets,and elongate garnets (Fig. 14). Round garnets pre-serve concentric zoning defined by inclusion-rich(commonly sigmoidal) cores and inclusion-freerims (Fig. 14a). Elongate garnets commonlycontain sigmoidal inclusion trails or more planarinclusion trails at a high angle to the grain longaxis (Fig. 14b, c). Small irregular garnets aredominantly inclusion-free (Hubbard 1988, 1989).

We propose a three-step evolution for thesegarnets (Fig. 14d, steps 1–3): (1) formation ofinclusion-rich cores during initial garnet nucleationand growth (as preserved by round garnets);(2) growth of inclusion-free rims during a secondphase of garnet growth; (3) local removal ofgarnet rim/core material by a combination ofbrittle fracturing and pressure solution during late-stage penetrative shearing and foliation develop-ment within the MCTZ. At least some of theobserved irregular garnets may be fracture frag-ments. These microstructures (Fig. 14) indicatethat deformation associated with both rotation ofthese elongate truncated garnets, and formation ofthe observed enveloping penetrative foliation,must either post-date or have outlasted peak meta-morphic conditions, as previously suggested byHubbard (1988, 1989, 1996); see also Brunel &Kienast (1986) for a similar interpretation along-strike in the Makalu section of the MCTZ. Resultsof our vorticity analyses, based on the dispersionof these garnet porphyroblasts, must also relate topenetrative flow that outlasted or post-dated peakmetamorphism. Many of the garnets in thesesamples are the same ones used by Hubbard(1988, 1989) to define the inverted metamorphicisograds along the Dudh Kosi section of theMCTZ. Therefore, these isograds may haveformed prior to this phase of deformation andshearing along the MCTZ (Hubbard 1996), whichis potentially associated with relative late-stageextrusion of the GHS.

Locations of samples used for rigid grain analy-sis are shown in a schematic cross-section throughthe MCTZ (Fig. 4). The structurally highestsamples (ET-44, ET-41 and 87-H-22E) are withinthe sillimanite stability field of the MCTZ andyield Wm estimates of 0.63–0.77 (c. 45–55%pure shear; Fig. 15). Elongate garnet porphyroclastsin kyanite-bearing samples (87-H-21J and 87-H-21G) yield Wm estimates of 0.60–0.72 (c. 60–48% pure shear). One sample (87-H-6B), thoughtto roughly coincide with the staurolite zone,yields a similar Wm estimate of 0.69–0.71(c. 50% pure shear). Sample 87-H-5A, collected

from within a sheared section of the Okhandungagneiss, yields a Wm estimate of 0.69–0.77 usingfeldspar grains. Sample 87-H-1B was collectedfurther south, in an essentially unmapped sectionof the MCTZ, yet yields a Wm estimate of0.66–0.70 using garnet porphyroblasts. The rangein Wm estimates from these MCTZ samples(0.60–0.77; average minimum and maximum Wmvalues of 0.67 and 0.72) suggests a c. 55–45%pure shear component at the base of the GHSfollowing peak metamorphic conditions.

Core of the Greater Himalayan Slab

Vorticity analyses in the anatectic core of the GHSwere not possible because these high-grade rocks

Rigid grain techniqueS: schist; GN: gneiss

ET-44

ET-41

85-H-22E

85-H-21J

85-H-21G

87-H-1B

87-H-6B

87-H-5A

0.90.80.70.6


1.00.5

S

S

5060 40 30 20 0

percent pure shear

10

S

S

S

S

S

GN

sillimanitekyanite

MAIN CENTRAL THRUST I

kyanite

staurolite

garnet

staurolite

MAIN CENTRAL THRUST

Fig. 15. Bar chart for range of mean kinematic vorticitynumbers (Wm) estimated by the rigid grain method forsamples collected in the Main Central thrust zone, Khumburegion, Nepal. Metamorphic isograd locations areapproximate. Sample locations shown in Figures 3 & 4.


lack mineral phases that remain rigid at high defor-mation temperatures. Deformation in the core of theGHS is markedly different from along its boundingmargins (STDS and MCTZ). Polyphase defor-mation of the metasedimentary rocks produced atleast two phases of folds that are migmatized tovariable degrees and injected by numerous leuco-granite sill complexes. Key overprinting relation-ships exposed throughout the core provide criticalinsight into the structural evolution of the GHS.

One such exposure is located on the lower ram-parts of the Nuptse–Lhotse wall where at leasttwo phases of folding are preserved in a singleoutcrop composed of interlayered quartzite andpelites (Figs 3 & 16). The first phase of deforma-tion (D1) is recorded by isoclinal F1 folds(148! 298W) that fold graded bedding in quart-zite and create a composite S0–S1 foliation(N158E, 358NW). Within the F1 hinges zones,white mica is aligned at a high angle to S0 anddefines an axial planar foliation. Quartz micro-structures within the quartzite layers record alimited degree of annealing. Isoclinal folds wererefolded during a second phase of deformation

(D2), producing both open F2 folds(138! N388W) that broadly warp the compositeS0 and S1 foliation in the quartzite layers, andtighter crenulation folds (148! N398W) in themechanically weaker pelite layers (Fig. 16). Theaxial planes and fold axes of both the open F2folds in the quartzite, and the crenulations in thepelitic layers, are parallel to each other (N508W,648NE) indicating that they are part of the samedeformation phase. Broad NE- and NW-trendingsubhorizontal folds in the Khumbu region havebeen documented by many previous studies(Hubbard 1988; Carosi et al. 1999a, b; Catloset al. 2002; Searle et al. 2003) and are here termedthe Khumbu Dome Complex (Fig. 3). An unde-formed layer-parallel leucogranite sill, which is par-tially exposed above this outcrop on the Nuptse–Lhotse wall, suggests that at least D1 and D2 pre-dated its emplacement. Other leucogranite sills inthe upper section of the GHS core also containlittle evidence for solid-state fabric development,suggesting that much of the anatectic melting andleucogranite injection post-dated the polyphasefolding that characterizes the core of the slab.

Fig. 16. Photograph of key outcrop exposure of overprinting folds/fabrics used to define at least two phases ofdeformation in the core of the Greater Himalayan Slab. Uniform light grey layer is quartzite with bedding defined bybiotite. Surrounding material is pelitic schist. See text for details. Brunton compass for scale. Image is of a vertical wallviewed towards the west. Approximate location of outcrop shown on Figure 3.


Summary of vorticity data for

Everest transect

Our vorticity data are taken from three lithotectonicunits: (1) sheared Tethyan limestone and underlyinggreenschist-facies Everest Series calc-mylonites,including the Yellow Band marble; (2) shearedleucogranite sills and amphibolite-facies schistsand gneisses in the footwall to the LD and compo-site QD–LD system; (3) schists from the zone ofinverted metamorphic isograds within the MCTZin which shearing either outlasted or post-datedpeak metamorphism. Arithmetic averages forminimum and maximum Wm values obtained ineach of these lithotectonic units are summarizedin Figure 17.

The highest average minimum and maximumWm values (0.81 and 0.84) are recorded in theTethyan limestone and Everest Series calc-mylonites at the top of the GHS, with a totalrange in estimated Wm values of 0.74–0.91 (16samples). Lower average minimum and maximumWm values (0.73 and 0.78) are recorded in theunderlying leucogranites and amphibolite-faciescalc-silicates and schists with a total range inestimated Wm values of 0.57–0.85 (25 samplesnot including Kangshung valley samples). The

lowest average minimum and maximum Wmvalues (0.67 and 0.72) are recorded in the MCTZschists with a total range in estimated Wm valuesof 0.63–0.77 (eight samples). These averageminimum and maximum estimated Wm valuescorrespond to c. 38–36, 48–41 and 53–48% pureshear components in the three lithotectonic units(Fig. 17). To what extent this distribution of esti-mated vorticity values might reflect a structural,lithologic or temporal partitioning of flow withinthe GHS is discussed below. Interpretation of ourdata is limited by the absence of vorticity datafrom the anatectic core of the 20–30 km thickslab, and it should be kept in mind that our dataare limited to samples collected either from thetop 2 km of the slab (with most samples comingfrom the top 600 m or less) or from the bottom5 km of the slab.

Discussion

Potential lithological, structural and

temporal controls on flow partitioning

Accurate assessment of the relative importance oflithological, structural and temporal controls onflow path evolution is critical for interpreting thetectonic evolution of the GHS. Data from theEverest transect indicate that the most pronouncedspatial transition in Wm values occurs at thetop of the slab. Here, the increase in Wm valuestowards the structurally highest parts of theslab coincides with an upward transition fromamphibolite-facies schist (mica-rich) and leuco-granite (quartz–feldspar-rich) to low-grade andunmetamorphosed marble and limestone. Thiscould be interpreted in several different end-member, as well as potentially overlapping, waysincluding: (1) a lithologic control on vorticity offlow; (2) a progressive spatial variation in flowtype controlled by structural position within themargin of the extruding slab or channel, in whichindividual sampling positions have not moved sig-nificant distances laterally from each other duringflow/extrusion; (3) large-scale foreland-directedextrusive lateral flow resulting in tectonic emplace-ment of high-grade rocks (originally flowing undergeneral shear conditions at mid-crustal depths)beneath cooler upper-crustal rocks deforming bysub-simple shear.

From a mechanics approach, rheological compe-tency can partition flow if sub-simple shear defor-mation is concentrated in relatively incompetentunits while general shear is concentrated in morecompetent units (Lister & Williams 1983). Assum-ing that limestones and marbles at the top of the

Tethyan limestone& calc-mylonites

Rongbuk schist & leucogranite

Main Central thrust zone

0.90.80.70.6


1.00.5

5060 40 30 20 0

percent pure shear

10

Rigid grain technique

{{

CORE OF GREATER HIMALAYAN SLAB

Fig. 17. Bar chart of average minimum and averagemaximum Wm values for: Tethyan limestone, calc-mylonites and marble of the Everest Series, and calc-silicates in immediate footwall to Qomolangma-Lhotse Detachment system; schists and myloniticleucogranites of the Rongbuk Formation; and garnetschist from the Main Central thrust zone.


GHS are the rheologically weakest units, and thatsub-simple shear flow has been concentratedin these units, a case could be made for this inter-pretation. This interpretation ignores the micro-structural, petrofabric and petrologic data thatindicate flow in the amphibolite-facies schistsand leucogranites occurred at close to peakmetamorphic conditions (Law et al. 2004), whileshearing in the overlying marble and limestoneoccurred at greenschist- to sub-greenschist-faciesconditions.

At the top of the GHS, the apparent upwardincrease in Wm values approaching the compositeQD–LD system coincides with a progressiveapparent decrease in deformation temperatures. Ifthis correlation between Wm values and upwarddecrease in deformation temperatures is real, thenour data could indicate an original rapid upwardincrease in Wm values within one flow regime orthe structural juxtaposition of different flowregimes during extrusion. In the second interpret-ation, the structurally deeper and higher tempera-ture samples provide information on flow thatoccurred at deeper crustal levels during earlierstages of channel flow/extrusion (low Wm values),while the structurally higher, lower temperaturesamples only record information on flow (athigher Wm values) that occurred at much highercrustal levels.

An upward decrease in deformation temperatureswithin the upper 600 m of the GHS and overlyingEverest Series and Tethyan rocks is indicated by:(1) decreasing opening angles of quartz c-axisfabrics from the schists beneath the compositeQD–LD system (Law et al. 2004, fig. 8b); (2) aprogression from quartz recrystallization domi-nated by grain boundary migration (Regime 3of Hirth & Tullis 1992) in the deeper schists tocombined subgrain rotation (Regime 2) and grainboundary migration recrystallization in theimmediate footwall to the detachment system; (3)an upward increase in brittle deformation offeldspar grains; (4) a transition from plastic defor-mation of quartz below the detachment system tobrittle deformation of quartz (albeit within amechanically weaker calcite matrix) in the over-lying Everest Series calc-mylonites and shearedTethyan limestones; (5) an abrupt transition intwinning regime (and dynamically recrystallizedgrain size) between the Everest Series YellowBand marble and overlying Tethyan limestone.

Law et al. (2004) previously argued that thestrains in our samples are too low for the extremeapparent thermal gradient (c. 3308C km21) indi-cated by microstructures and fabric openingangles to be solely explained by strain-induced tele-scoping of isotherms during extrusion. We suggest

that, traced structurally upwards towards the topof the GHS, the inverse relationship between Wmvalues (increasing) and deformation temperatures(decreasing) reflects both a spatial and temporalflow partitioning, with structurally deeper andhigher temperature samples preserving informationon flow that occurred at deeper crustal levelsduring earlier stages of channel flow/extrusion(low Wm values). The structurally higher, lowertemperature samples only record information onflow (higher Wm values) that occurred at muchhigher crustal levels, probably during the laterstages of exhumation. In this interpretation, strainrate may also play a role in controlling the natureof local flow. However, using the available data,we cannot unequivocally choose between our pre-ferred model, in which the low temperaturesamples record deformation during late stages ofextrusion–exhumation, and a model in which bothlow and high temperature samples were deformedsimultaneously at different structural levels in thecrust. As previously discussed by Grasemannet al. (1999) and Williams et al. (2006), extrusionmodels involving a component of pure shearrequire an increase in strain rate at the margins ofthe extruding slab/wedge traced from the hinter-land to the foreland. Thus the progressive evolutionfrom more ductile to more brittle behaviour, whichwe infer to indicate flow at progressively structu-rally shallower levels and in more forelandparts of the orogenic system, may be a compositeeffect of decreasing temperature and increasingstrain rate.

Distribution of flow regimes within the

GHS and tectonic implications

In both basic channel flow models (e.g. Grujic et al.1996, 2002) and extrusive flow models (e.g.Williams et al. 2006), a symmetric distribution offlow paths is predicted at any one instant in time,with lowest flow vorticities in the centre of thechannel and a progressive increase in flow vortici-ties towards the boundaries of the channel (Grujicet al. 2002, fig. 7; see also Grujic 2006, fig. 2c).In coupled thermal–mechanical finite-elementmodels (which assume a reduction in channel visc-osities by partial melting), material in the centralparts of the channel originates at mid-lowercrustal depths and ‘tunnels’ for great distanceslaterally before extruding into upper-crustal rocksas it approaches the topographic surface (Beaumontet al. 2001, 2004, 2006; Jamieson et al. 2002, 2004,2006; Godin et al. 2006b). Our microstructural andvorticity data from the top of the GHS are compati-ble with components of all these models. Due to


the high deformation temperatures, presumablyassociated with flow, no vorticity markers havebeen preserved in the core of the slab. Our fielddata indicate that this penetrative flow in the coreoccurred during the relatively earlier stages ofdecompression (see above) and therefore is prob-ably slightly earlier than flow recorded in theamphibolite–sub-greenschist-facies rocks at themargins of the slab. These interpretations are alsocompatible with the above channel flow/extrusionmodels. Assuming flow associated with our data isessentially synchronous in the upper and lowerparts of the slab, our vorticity data from theMCTZ are incompatible with these models.Our analyses, although limited, indicate that theMCTZ is characterized by the lowest average Wmvalues (i.e. highest pure shear components) in ourtransect across the slab. The MCTZ lacks anyconvincing progressive increase in Wm values atdeeper structural levels that might, as predictedfor example by channel flow models, mirror theincrease in Wm values at the top of the slab.From a mechanics perspective, an increase in litho-static pressure towards the base of the slab, asimplied in gravity-driven collapse/spreadingmodels for thrust belt evolution, provides a poten-tial explanation for the highest pure shear com-ponent being recorded at the deepest structurallevels (see Simpson & De Paor 1997; review byMerle 1998).

Results of our vorticity analyses from the Everesttransect share some similarities with those fromthe Sutlej River section of the MCTZ whereGrasemann et al. (1999) demonstrated a progressivedownward decrease in Wm values towards the baseof the GHS. Grasemann et al. (1999) proposed thattheir vorticity data indicated a temporal (ratherthan spatial) change in flow regime associatedwith a decelerating strain path (Simpson & DePaor 1997) in which progressively more generalshear replaced high temperature simple shearflow during cooling. In the Everest transect, ourmicrostructural data from the MCTZ demonstratethat low Wm values (indicating a general shear)are related to flow and foliation development thatpost-dates peak metamorphic conditions (see alsoHubbard 1988, 1989). Evidence such as flow athigher temperatures in the structurally higher sec-tions of the MCTZ, and at lower temperatures inthe structurally deeper levels (e.g. chlorite wingson garnet porphyroclasts), suggest that penetrativeflow may have progressed to deeper structurallevels over time, as suggested for the Sutlej Riversection by Grasemann et al. (1999).

Timing constraints on the kinematic evolutionof the GHS are provided by a wealth of geochronolo-gical data from the Everest region. Mylonitic

leucogranite sills parallel to the combined QD–LDsystem along the top of the slab suggest that ductileshearing lasted until c. 17 Ma after which brittlemotion on the upper strand of the detachmentsystem occurred until c. 16 Ma (Murphy andHarrison 1999; Searle et al. 2003). Timing ofamphibolite-facies metamorphism and early defor-mation (c. 500–5508C) on the MCTZ is constrainedto c. 21 Ma (Hubbard & Harrison 1989). Monaziteinclusions in garnets from the Everest section ofthe MCTZ, dated at c. 14 Ma by Catlos et al.(2002), provide a maximum age constraint forgarnet growth (assuming they pre-date garnetgrowth) and hence flow associated with theseWm data. Based on available isotopic age data,this phase of flow along the base of the GHS,which is associated with the lowest average Wmvalues in the Everest transect, may be youngerthan the documented flow regimes at the top ofthe slab.

This interpretation does not exclude the possi-bility that earlier reverse-sense shearing on theMCTZ (possibly involving sub-simple shear asrequired for example by channel flow models)was synchronous with normal-sense shearing atthe top of the slab. Instead, this interpretationsuggests that the microstructural evidence for thisearlier shearing may have been overprinted as thecore of the channel locked up at c. 16 Ma andflow was partitioned into the foreland to accom-modate continued crustal shortening (see Godinet al. 2006a).

Conclusions

Structural analyses of rocks collected along theEverest transect provide the first quantitativeinformation on how flow was partitioned withinthe Greater Himalayan Slab during extrusion andexhumation. Results of vorticity analyses alongthe top of the slab indicate that the higher-grade,structurally deeper rocks, record general shear atclose to peak metamorphic conditions, while thelower-grade, structurally higher rocks, recordsub-simple shear. Vorticity measurements in thecore of the slab are problematic due to the highmetamorphic grade they reached; however, the pen-etrative fabrics in the core most likely developed atmid-crustal depths during the early stages ofdecompression. The highest average pure shearcomponents are recorded at the base of the slab,and are associated with deformation that post-datespeak metamorphism. We attribute the distributionof flow regimes to spatial and temporal partitioningof flow in which higher temperature samples recordthe early stages of channel flow/extrusion (general


shear) at mid-crustal depths in the interior of thechannel. The structurally higher, lower temperaturesamples at the top of the slab only record infor-mation on flow (sub-simple shear) that occurredalong the upper margin of the channel at muchhigher crustal levels, probably during the laterstages of exhumation. Although flow paths in theupper and lower parts of the slab may have beensimilar during channel flow and extrusion, asrequired by the different models, microstructuralevidence for earlier shearing at peak metamorphicconditions along the base of the slab was over-printed as the channel flow/extrusion systemlocked up at c. 16 Ma and the locus of deformationmigrated towards the foreland in order to accommo-date continued crustal shortening.

We are grateful to R. Tracy, D. Waters, F. Read,K. Karlstrom, and G. Oliver for discussion and adviceon various aspects of this project. We thank T. Sherpa,D. Sherpa, R. Rai, J. Cottle, D. Newell, J. Ashby andL. Duncan for assistance in the field and many stimulatingdiscussions over tea and dhal bhat. Also, many thanks toR. Schrama and S. Dhakta for help with organizing fieldseasons in Tibet. We also thank S. Laurie (CambridgeUniversity) and D. Waters (Oxford University Museumof Natural History) for generously providing access tohigh altitude Mount Everest samples from the Harkerand Wager collections. C. Bailey, L. Godin andP. Xypolias are thanked for their detailed and helpfulreviews of an earlier version of the manuscript. Thiswork was funded by National Science Foundation grantEAR 0207524 to R.D.L. and M.P.S. M.J.J. was alsofunded by Geological Society of America and Sigma Xigraduate student research grants, a W.D. LowryGeosciences Graduate Research Award from VirginiaTech Department of Geosciences, and a 2010 GraduateFellowship from Virginia Tech College of Science-Graduate School. Fieldwork by M.S.H. was funded byNSF grant EAR-8414417 to K.V. Hodges.

Appendix: Rigid grain data plots for the 51 Samplesused for vorticity analysis.(See Figs A1–A5, pp. 409–413)

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quar

tz in

cal

cite

mat

rix

quar

tz c

last

in c

alci

te m

atrix

feld

spar

epid

ote

tour

mal

ine

zirc

on

feld

spar

feld

spar

epid

ote

feld

spar

garn

etep

idot

eam

phib

ole

feld

spar

feld

spar

feld

spar

feld

spar

feld

spar

feld

spar

Fig

.A

1.

Nort

her

ntr

aver

se.P

osi

tive

and

neg

ativ

ean

gle

sbet

wee

ncl

astlo

ng

axis

and

foli

atio

nin

dic

ated

clas

tsei

ther

incl

ined

tow

ards

the

loca

lsh

ear

sense

(to

p-t

o-n

ort

h)

or

agai

nst

shea

rse

nse

,re

spec

tivel

y.


Rc

= 2

.68

Rc

= 3

.15

90¡

60¡

30¡

0¡

3.0

4.0

5.0

6.0R

1.0

R-0

3-26

-30¡

-60¡

-90¡


n =

197

Wm

= 0

.76

- 0.

82

Rc

= 2

.60

Rc

= 2

.90

90¡

60¡

30¡

0¡

3.0

4.0

5.0

6.0R

1.0

R-0

3-26

A

-30¡

-60¡

-90¡


n =

260

Wm

= 0

.74

- 0.

79

Rc

= 3

.20

Rc

= 3

.45

90¡

60¡

30¡

0¡

4.0

3.0

5.0

6.0R

1.0

R-0

3-55

-30¡

-60¡

-90¡


n =

200

Wm

= 0

.82

- 0.

84

Rc

= 2

.80

Rc

= 3

.0

90¡

60¡

30¡

0¡

3.0

4.0

5.0

6.0R

1.0

R-0

3-56

-30¡

-60¡

-90¡


n =

200

Wm

= 0

.77

- 0.

80

Rc

= 2

.70

Rc

= 2

.90

90¡

60¡

30¡

0¡

3.0

4.0

5.0

6.0R

1.0

R-0

3-58

-30¡

-60¡

-90¡


n =

225

Wm

= 0

.76

- 0.

79

Rc

= 2

.90

Rc

= 3

.20

90¡

60¡

30¡

0¡

4.0

3.0

5.0

6.0R

1.0

R-0

3-59

-30¡

-60¡

-90¡


n =

258

Wm

= 0

.79

- 0.

82

feld

spar

in c

alci

tequ

artz

in c

alci

te

Rc

= 3

.10

Rc

= 3

.40

90¡

60¡

30¡

0¡

4.0

3.0

5.0

6.0R

1.0

R-0

3-63

-30¡

-60¡

-90¡


n =

352

Wm

= 0

.81

- 0.

84

Rc

= 2

.80

Rc

= 2

.95

90¡

60¡

30¡

0¡

3.0

4.0

5.0

6.0R

1.0

TI-

5

-30¡

-60¡

-90¡


n =

200

Wm

= 0

.77

- 0.

79

Rc

= 3

.20

Rc

= 3

.55

90¡

60¡

30¡

0¡

3.0

4.0

5.0

6.0 R

1.0

ET-

15

n =

305

Wm

= 0

.82

- 0.

85-3

0¡

-60¡

-90¡


Rc

= 2

.65

Rc

= 3

.0

90¡

60¡

30¡

0¡

3.0

4.0

5.0

6.0R

1.0

R-0

3-67

-30¡

-60¡

-90¡


n =

200

Wm

= 0

.75

- 0.

80

Rc

= 2

.25

Rc

= 2

.5

90¡

60¡

30¡

0¡

3.0

4.0

5.0

6.0R

1.0

ET-

14

n =

250

Wm

= 0

.67

- 0.

73-3

0¡

-60¡

-90¡


feld

spar

tour

mal

ine

zirc

on

feld

spar

feld

spar

quar

tz in

cal

cite

mat

rixfe

ldsp

ar in

cal

cite

mat

rix

quar

tz in

cal

cite

mat

rixfe

ldsp

ar in

cal

cite

mat

rix

quar

tz in

cal

cite

mat

rix

epid

ote

in c

alci

te m

atrix

quar

tz in

cal

cite

tour

mal

ine

in c

alci

tefe

ldsp

ar in

cal

cite

feld

spar

feld

spar

amph

ibol

eep

idot

e

feld

spar

tour

mal

ine

epid

ote

zirc

onam

phib

ole

pyro

xene

Nor

ther

n Tr

anse

ct

Ron

gbuk

Mon

aste

ry tr

aver

se

Fig

.A

2.

No

rther

nan

dR

on

gb

uk

Mo

nas

try

trav

erse

s.P

osi

tiv

ean

dn

egat

ive

ang

les

bet

wee

ncl

ast

lon

gax

isan

dfo

liat

ion

indic

ated

clas

tsei

ther

incl

ined

tow

ards

the

loca

lsh

ear

sen

se(t

op

-to

-no

rth

)o

rag

ain

stsh

ear

sen

se,

resp

ecti

vel

y.


Rc

= 2

.65

Rc

= 3

.0

90¡

60¡

30¡

0¡

3.0

4.0

5.0

6.0 R

1.0

ET-

13

n =

304

Wm

= 0

.75

- 0.

8-3

0¡

-60¡

-90¡


Rc

= 2

.5

Rc

= 2

.75

90¡

60¡

30¡

0¡

-30¡

-60¡

-90¡

3.0

4.0

5.0

6.0 R

ET-

12

n =

206

1.0

Wm

= 0

.72

- 0.

77


Rc

= 2

.60

Rc

= 2

.75

90¡

60¡

30¡

0¡

3.0

4.0

5.0

6.0R

1.0

R-0

3-46

-30¡

-60¡

-90¡


n =

202

Wm

= 0

.74

- 0.

77

Rc

= 2

.50

Rc

= 2

.60

90

¡

60¡

30¡

0¡

4.0

3.0

5.0

6.0R

1.0

R-0

3-43

-30¡

-60¡

-90¡


n =

140

Wm

= 0

.72

- 0.

75

Rc

= 2

.15

Rc

= 2

.40

90¡

60¡

30¡

0¡

3.0

4.0

5.0

6.0R

1.0

R-0

3-39

-30¡

-60¡

-90¡


n =

202

Wm

= 0

.64

- 0.

70

Rc

= 2

.95 Rc

= 3

.35

90¡

60¡

30¡

0¡

3.0

4.0

5.0

6.0 R

1.0

ET-

8

n =

325

Wm

= 0

.79

- 0.

84-3

0¡

-60¡

-90¡


Rc

= 1

.90

Rc

= 2

.15

90¡

60¡

30¡

0¡

4.0

3.0

5.0

6.0R

1.0

R-0

3-31

-30¡

-60¡

-90¡


n =

138

Wm

= 0

.57

- 0.

64

Rc

= 2

.35

Rc

= 3

.05

90¡

60¡

30¡

0¡

3.0

2.0

4.0

5.0

6.0R

1.0

R-0

3-33

-30¡

-60¡

-90¡


n =

146

Wm

= 0

.69

- 0.

80

Rc

= 2

.65

Rc

= 2

.85

90¡

60¡

30¡

0¡

3.0

2.0

4.0

5.0

6.0R

1.0

R-0

3-38

-30¡

-60¡

-90¡


n =

200

Wm

= 0

.75

- 0.

78

Rc

= 2

.60

Rc

= 2

.80

90¡

60¡

30¡

0¡

3.0

4.0

5.0

6.0R

1.0

R-0

3-44

-30¡

-60¡

-90¡


n =

200

Wm

= 0

.74

- 0.

77

Rc

= 2

.55

Rc

= 2

.75

90¡

60¡

30¡

0¡

3.0

4.0

5.0

6.0R

1.0

R-0

3-70

-30¡

-60¡

-90¡


n =

200

Wm

= 0

.73

- 0.

77

feld

spar

feld

spar

feld

spar

tour

mal

ine

zirc

on

quar

tz in

cal

cite

mat

rixop

aque

s in

cal

cite

mat

rixqu

artz

in c

alci

te m

atrix

quar

tz in

cal

cite

mat

rix

feld

spar

in c

laci

te m

atrix

feld

spar

tour

mal

ine

epid

ote

zirc

onam

phib

ole

feld

spar

feld

spar

tour

mal

ine

feld

spar

feld

spar

Ron

gbuk

Mon

aste

ry tr

aver

se

Her

mit'

s G

orge

trav

erse

Fig

.A

3.

Ro

ng

buk

Mo

nas

tery

and

Her

mit

’sG

org

etr

aver

ses.

Po

siti

ve

and

neg

ativ

ean

gle

sb

etw

een

clas

tlo

ng

axis

and

foli

atio

nin

dic

ated

clas

tsei

ther

incl

ined

tow

ards

the

loca

lsh

ear

sen

se(t

op

-to

-nort

h)

or

agai

nst

shea

rse

nse

,re

spec

tiv

ely

.


Fig

.A

4.

Mo

un

tE

ver

est–

Kan

gsh

un

gv

alle

y.

On

lysa

mple

sK

-04

-03

and

K-0

4-0

4fr

om

the

Kan

gsh

un

gv

alle

yar

eo

rien

ted

.P

osi

tive

and

neg

ativ

ean

gle

sb

etw

een

clas

tlo

ng

axis

and

foli

atio

nin

dic

ated

clas

tsei

ther

incl

ined

tow

ard

sth

elo

cal

shea

rse

nse

(to

p-t

o-n

ort

hfo

rK

-04

-03

and

04

)o

rag

ain

stsh

ear

sen

sein

dic

ated

by

mic

rost

ruct

ure

sin

indiv

idual

sam

ple

s,re

spec

tivel

y.


Rc

= 2

.1

Rc

= 2

.4

90¡

60¡

30¡

0¡

3.0

4.0

5.0

6.0R

1.0

ET.

44

-30¡

-60¡

-90¡


n =

232

Wm

= 0

.63

- 0.

70

Rc

= 2

.10

Rc

= 2

.55

90¡

60¡

30¡

0¡

3.0

4.0

5.0

6.0R

1.0

ET.

41

-30¡

-60¡

-90¡


n =

227

Wm

= 0

.63

- 0.

73

Rc

= 2

.50

Rc

= 2

.75

90¡

60¡

30¡

0¡

4.0

3.0

5.0

6.0R

1.0

87-H

-22E

-30¡

-60¡

-90¡


n =

275

Wm

= 0

.72

- 0.

77

Rc

= 2

.40

Rc

= 2

.55

90¡

60¡

30¡

0¡

4.0

3.0

2.0

5.0

6.0R

1.0

85-H

-21J

-30¡

-60¡

-90¡


n =

43

Wm

= 0

.70

- 0.

72

Rc

= 2

.00

Rc

= 2

.15

90¡

60¡

30¡

0¡

4.0

3.0

2.0

5.0

6.0R

1.0

85-H

-21G

-30¡

-60¡

-90¡


n =

59

Wm

= 0

.60

- 0.

64

Rc

= 2

.35

Rc

= 2

.45

90¡

60¡

30¡

0¡

4.0

3.0

5.0

6.0R

1.0

87-H

-6B

-30¡

-60¡

-90¡


n =

98

Wm

= 0

.69

- 0.

71

Rc

= 2

.35

Rc

= 2

.80

90¡

60¡

30¡

0¡

4.0

3.0

5.0

6.0R

1.0

87-H

-5A

-30¡

-60¡

-90¡


n =

159

Wm

= 0

.69

- 0.

77

Rc

= 2

.20

Rc

= 2

.40

90¡

60¡

30¡

0¡

4.0

3.0

5.0

6.0R

1.0

87-H

-1B

-30¡

-60¡

-90¡


n =

105

Wm

= 0

.66

- 0.

70

garn

etga

rnet

tour

mal

ine

garn

etga

rnet

garn

etga

rnet

tour

mal

ine

feld

spar

feld

spar

tour

mal

ine

zirc

on

garn

et

Fig

.A

5.

Mai

nC

entr

alth

rust

zon

e.P

osi

tive

and

neg

ativ

ean

gle

sb

etw

een

clas

tlo

ng

axis

and

foli

atio

nin

dic

ated

clas

tsei

ther

incl

ined

tow

ard

sth

elo

cal

shea

rse

nse

(to

p-t

o-s

ou

th)

or

agai

nst

shea

rse

nse

,re

spec

tiv

ely

.


Structural evolution and vorticity of ﬂow during extrusion and...

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Transcript of Structural evolution and vorticity of ﬂow during extrusion and...