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Geological Society of America Bulletin

doi: 10.1130/B25514.1 2005;117;764-782Geological Society of America Bulletin

Zeshan Ismat and Gautam Mitra The role of cataclastic flowFold-thrust belt evolution expressed in an internal thrust sheet, Sevier orogen:

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ABSTRACT

Large hinterland thrust sheets are the dominant structures in retroarc fold-thrust belt (FTB) wedges. These internal sheets remain active throughout the history of an FTB and strongly infl uence the kinematics of the external portions of the FTB and the overall behavior of the entire tectonic wedge. Their emplacement and deformation account for a signifi cant portion of the total mechani-cal energy budget in FTB evolution based on estimates using deformation features preserved within such sheets. Moreover, the deformation within dominant internal sheets can be used as a means to evaluate the tec-tonic history of the wedge as a whole.

The principal deformation mechanisms in different parts of an FTB wedge are strongly controlled by variations in the geothermal gradient. Retroarc FTBs typically show a steeper geothermal gradient in the hinterland than in the foreland, resulting in a foreland-ward slope of the “brittle-ductile transition.” Therefore, a change from plastic to elastico-frictional (EF) mechanisms is expected not only with decreasing depth, but also across strike as a dominant internal thrust sheet is transported toward the foreland and under-goes exhumation by synorogenic erosion. However, evidence for cataclasis overprinting early plastic deformation features is often not preserved in ancient FTBs because of exten-sive erosion of the wedge top. Even where evidence for late-stage cataclastic behavior exists in either ancient or modern (i.e., active) orogens, it is often not recognized or studied.

Our work in the Canyon Range, central Utah segment of the Sevier FTB shows that unraveling the EF mechanisms, such as late-stage cataclastic fl ow, is critical to under-standing the total deformation history of an

internal thrust sheet. After initial emplace-ment under plastic deformation conditions, the CR thrust sheet was continuously reac-tivated under EF conditions, thereby helping to maintain critical taper of the FTB wedge during its development. Structures preserved in the CR thrust sheet, therefore, record a sig-nifi cant portion of the deformation history of the FTB as a whole. The Canyon Range (CR) syncline, part of the CR sheet, underwent fold tightening within the EF regime and the cataclasized rocks are well preserved. The geometry and progressive deformation pat-terns of the CR syncline reveal four distinct to partly overlapping deformation stages. The history of the rocks in the CR thrust sheet has been tracked by incrementally retro-deform-ing the FTB and by unraveling details of the deformation within the CR syncline at vari-ous scales for each stage.

We use the CR syncline as a type example to show that unraveling EF mechanisms is critical to (1) understanding the state of a FTB wedge during its evolution, (2) calculat-ing the energy involved in FTB emplacement, and (3) developing accurate restorations. Placing constraints on the depth of the rocks for successive stages of deformation is essen-tial for developing retrodeformable balanced cross sections, and in turn, geometric and kinematic models for wedge evolution.

Keywords: cataclastic fl ow, fold-thrust belt, Sevier, energy, folding.

INTRODUCTION

Fold-thrust belts (FTBs) form the frontal parts of major contractional orogens and hold critical clues to the overall mechanical evolution of these orogenic belts. The mechanics of FTBs can be broadly explained by critical wedge the-ory (Elliott, 1976; Chapple, 1978; Davis et al., 1983; Dahlen, 1990; Willett, 1992), providing a

general framework for evaluating the large-scale evolution of FTBs. Within an individual FTB, hinterland thrust sheets that developed from the miogeoclinal portion of the basin tend to be larger and have greater displacements (Boyer and Elliott, 1982; Geiser and Boyer, 1987). These internal, dominant sheets form early in the deformational history, are commonly active over the entire history of an FTB (DeCelles, 1994; DeCelles et al., 1995), and partly con-trol the overall mechanical evolution of the FTB (Boyer, 1995; DeCelles and Mitra, 1995). Unraveling the kinematic and dynamic histories of these dominant sheets is one of the keys to understanding the overall evolution of an FTB.

These large sheets are generally assumed to deform in a ductile manner, principally by crystal plastic mechanisms and/or pressure solu-tion. The high geothermal gradient typical of the internal portions of mountain belts results in a shallow elastico-frictional/quasi-plastic (EF/QP) transition (Sibson, 1977; Rutter, 1986; Twiss and Moores, 1992), so that only a thin carapace on these large thrust sheets is expected to deform by EF mechanisms that form brittle structures (e.g., fractures). This minor brittle portion of the wedge is generally ignored for the purposes of mechanical models (e.g., Elliott, 1976; Chapple, 1978). However, several lines of evidence sug-gest that in most retroarc FTBs the EF/QP tran-sition dips toward the foreland (Armstrong and Dick, 1974; Kulik and Schmidt, 1988; Smith and Bruhn, 1984). Thus, we should expect to fi nd the structures formed by EF mechanisms, such as fractures, overprint features resulting from QP mechanisms as an internal thrust sheet is transported toward the foreland and undergoes exhumation by synorogenic erosion. Under a depressed geothermal gradient, as is commonly found toward the foreland, a signifi cant portion of the wedge can deform within the EF regime (Davis et al., 1983).

During the evolution of an FTB very large volumes of rock are typically eroded from

Fold-thrust belt evolution expressed in an internal thrust sheet, Sevier orogen: The role of cataclastic fl ow

Zeshan Ismat†

Department of Earth and Environment, Franklin and Marshall College, Lancaster, Pennsylvania 17603, USA

Gautam MitraDepartment of Earth and Environmental Sciences, University of Rochester, Rochester, New York 14627, USA

GSA Bulletin; May/June 2005; v. 117; no. 5/6; p. 764–782; doi: 10.1130/B25514.1; 12 fi gures; Data Repository item 2005082.

†E-mail: [email protected].

764


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FOLD-THRUST BELT EVOLUTION EXPRESSED IN AN INTERNAL SHEET: ROLE OF CATACLASTIC FLOW

Geological Society of America Bulletin, May/June 2005 765

internal sheets (DeCelles, 1994; DeCelles and Mitra, 1995), constituting the dominant source of synorogenic sediments in the adjoining fore-land basin. Because of this erosion, evidence for early brittle deformation within the upper part of the hinterland portion of a wedge is seldom preserved in FTBs. However, if internal sheets continue to be actively deformed and uplifted through the entire period of thrusting in an FTB we should expect late-stage brittle deformation features to be preserved. In addition, these struc-tures can give rise to large-scale ductile defor-mation by block-controlled cataclastic fl ow (Ismat, 2002), which is often not recognized even where it is present.

This study examines the geometry and pro-gressive deformation patterns in part of the Can-yon Range (CR) thrust sheet, an internal thrust sheet in the central Utah segment of the Sevier FTB. This internal thrust sheet was repeatedly reactivated, carrying rocks within the sheet upward and toward the foreland (DeCelles et al., 1995; Mitra, 1997). Because the sheet continued to deform, rather than being passively carried (Mitra, 2001), the CR thrust sheet contains a record of its hinterland to foreland deformation path during the evolution of the Sevier FTB (Boyer 1992; Gray and Mitra, 1993, 1999).

Detailed mapping and analysis of bedding contacts, faults, synorogenic conglomerates, and outcrop- and microscale structures within the CR thrust sheet reveal four distinct but over-lapping deformation stages. The early stages are characterized by plastic deformation while the later stages have dominantly brittle structures. The deformation that occurred at each stage within the CR thrust sheet was strongly infl u-enced by the variation in the geothermal gradient both vertically and across the strike of the Sevier FTB. The structures preserved at multiple scales are used to constrain the depth of the rocks for successive stages of deformation. In turn, these constraints are used to unravel the geometric and kinematic history of the sheet and the FTB wedge. These stages may also record the passage of deformation fronts (Gray and Mitra, 1999) through the fold-thrust belt as a whole.

This detailed study of the progressive defor-mation in the upper part of the hinterland por-tion of an FTB wedge will illustrate the impor-tant role that elastico-frictional deformation (e.g., fracturing and cataclastic fl ow) can play in FTB wedge evolution. We suggest methods for unraveling the deformation recorded in fracture networks so as to develop accurate restorations and to understand the mechanical state of the wedge during its evolution. This analysis also raises the important question of how FTB wedge behavior varies depending on whether dominant internal sheets were deforming mainly in the

quasi-plastic or elastico-frictional regime. We discuss this by comparing both the deforma-tion patterns and energy requirements of the CR sheet with that of plastically deformed internal sheets in other segments of the Sevier FTB.

GEOLOGIC SETTING

Central Utah Segment

The Sevier FTB defi nes the eastern margin of the North American Cordillera and has sev-eral salients or segments separated by recesses that are typically located at transverse zones (Lageson and Schmidt, 1995; Lawton et al., 1997; Mitra, 1997) (Fig. 1A). The Central Utah segment is the type area of the Sevier FTB (Arm-strong, 1968) and is bound at its northern end by the Leamington transverse zone (Fig. 1A). This segment is composed of two juxtaposed wedges: a thicker rear wedge composed of stronger rocks (mainly quartzites), and a thinner frontal wedge of weaker rocks (e.g., limestone, shale, sand-stone) (Fig. 2) (Hintze, 1988; Mitra, 1997). The rear wedge, or hinterland, is composed of rocks deposited outboard of the miogeoclinal hinge, consisting mainly of Upper Proterozoic clastic rocks and Paleozoic-Mesozoic carbonates and clastic rocks; the stratigraphic package was no more than ~20 km thick. The frontal wedge, or foreland, is composed of rocks deposited on the stable cratonic shelf, including Paleozoic-Meso-zoic shelf sedimentary rocks with a maximum thickness of ~5 km (Hintze, 1988; Royse, 1993; Christie-Blick, 1997; Mitra, 1997). During the Cretaceous Sevier orogeny, rocks from the rear wedge were translated eastward past the miogeo-clinal hinge and onto the foreland, whereas much of the original shelf sedimentary rocks were bur-ied under synorogenic sediments (Fig. 2).

The four main thrusts in the central Utah seg-ment are (from west to east) the Canyon Range (CR), Pavant (PVT), Paxton (PXT), and Gunni-son (GUN) (Fig. 1B). Thrusting continued from the early Cretaceous to the Eocene and gener-ally progressed toward the foreland with signifi -cant out-of-sequence reactivation particularly of the CR and PVT thrust sheets (DeCelles et al., 1995). The thrust sheets are cut by Tertiary Basin-and-Range normal faults (Figs. 1 and 2) so that Sevier age structures can only be studied in normal fault bound ranges.

Canyon Range (CR) Thrust Sheet

A fi rst-order anticline-syncline pair of the folded CR thrust sheet is exposed in the Canyon Mountains, west-central Utah (Figs. 1 and 3). A complete section (~3.5 km thick) of Proterozoic and Cambrian quartzites is well exposed in the

CR syncline (Figs. 3 and 4). The major quartzite units are the Neo-Proterozoic Pocatello (PCp), Caddy Canyon (PCc), and Mutual (PCm) For-mations, the Eocambrian Tintic Formation (Ct) (subdivided into lower, middle, and upper), and the Cambrian Pioche Formation (Cp). These quartzites are overlain by Cambrian carbonates.

All the quartzites are relatively homogeneous compositionally. Parts of the PCm, the middle Ct, and parts of the lower Ct and Cp quartzites have prominent bedding fabrics, are often thin bedded, and served as weak horizons for local-ized slip during deformation; these units also usually do not form ridges. The PCc, parts of the PCm, and the upper- and lower-Ct quartzites, on the other hand, are thick bedded, massive cliff-forming units.

The quartzites in the CR sheet preserve defor-mation increments from a >70 m.y. history of Sevier FTB evolution. Prior to deformation, the quartzites were at ≥15 km depth in the mio-geocline sedimentary sequence. FTB develop-ment brought the quartzites to the surface with accompanying erosion and progressive unroofi ng (Fig. 2) (DeCelles et al., 1995; Mitra and Suss-man, 1997; Mitra, 1997; Mitra and Ismat, 2001).

During initial CR thrusting, the Protero-zoic-Cambrian quartzites were transported up a long ramp to the surface (Fig. 2, Stage A). During movement on the PVT thrust, the CR thrust sheet was carried over a Pavant ramp and formed a ramp anticline with a syncline in front of it (Fig. 2, Stage B). This motion and concur-rent erosion carried the CR quartzites just past the miogeoclinal hinge and onto the foreland to a depth of

ISMAT and MITRA

766 Geological Society of America Bulletin, May/June 2005

The number, size, and motion history of the horses varies along strike, as does the geometry of the adjoining syncline (Mitra and Sussman, 1997; Ismat and Mitra, 2001a).

Coarse Cretaceous CR conglomerates were deposited in the synclinal hinge area during the fold-tightening phase and were infolded in the core of the CR syncline (Figs. 2 and 3). These synorogenic conglomerates include Paleozoic carbonate clasts derived from the Sevier cul-mination and Proterozoic quartzite clasts shed off the CR culmination. As the structure was uplifted during this deformation, conglomerates were erosively recycled and deposited again farther out into the foreland basin. The thick-ness of the CR conglomerate was no more than 2 km at any stage of fold tightening. Therefore, the basal CR quartzites in the syncline hinge zone deformed at a depth of ≤4 km during the last phases of contractional deformation (Fig. 2, Stage C).

PROGRESSIVE DEFORMATION OF THE CANYON RANGE SYNCLINE

Introduction

To decipher the progressive deformation of the CR thrust sheet, we focused our detailed investigation on the CR syncline, which exposes a complete section of the Proterozoic-Cambrian quartzites. The CR syncline changes from an open fold at the southern end of the map area to a tight overturned fold with a moderate plunge to the NNW at the northern end of the range (Figs. 4A and 4B) (Mitra and Sussman, 1997; Ismat and Mitra, 2001a). Beyond an interlimb angle of ~100° (i.e., 40° limb dips), fold tightening was accomplished by thinning and stretching of the west limb during further limb rotation and by hinge migration (Ismat and Mitra, 2001a). The changes in geometry of the CR syncline along strike can be used to derive a temporal history of folding by assuming that tighter parts of the fold went through a similar folding history to that preserved in more openly folded areas (Ismat and Mitra, 2001a). In other

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Figure 1. (A) The Sevier FTB of western USA, showing major salients, segments, and recesses, and the transverse zones bounding them. Part (B) indicated by boxed area. (B) Central Utah segment of the Sevier fold-thrust belt of western USA showing the four main thrusts (Canyon Range, Pavant, Pax-ton, and Gunnison) and major Basin-and-Range normal faults within this segment. Boxed area shows location of Figure 3.





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Figure 2. (A) Kinematic history of the cen-tral Utah segment of the Sevier FTB with incrementally restored regional cross sec-tions showing the approximate shapes of the orogenic wedge during emplacement of successive major thrust sheets. Sevier culmination, Canyon Range culmination (i.e., connecting splay duplex), Canyon Range thrust (CRT), Pavant thrust (PVT), Paxton thrust (PAX), Gunnison thrust (GUN), and the Sevier Desert Detachment (SDD), a low angle normal fault, are shown. Black arrow shows approximate location of the shelf/miogeocline hinge and separates the rear wedge to the west from the frontal wedge to the east. Black dots show location of rocks studied through time; letters (A–D) refer to deformation stages those rocks went through and indicate stages of CR thrust sheet evolution (referred to in text). Sea level (S.L.) for each cross section indicated by small horizontal lines. The upper surface shows the erosional top of the wedge dur-ing each stage of deformation. (B) Chart showing stages of CR thrust sheet evolution, conglomerate deposition, and sequential timing of thrust faulting in the central Utah segment. Horizontal line pattern indicates periods of erosion and/or nondeposition (after DeCelles, et al., 1995). Letters refer to deformation stages illustrated in part (A).



ISMAT and MITRA


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e zo

nes

(Foo

l Cre

ek,

Dav

idso

n C

anyo

n, a

nd O

ak C

reek

) an

d di

vide

d in

to t

hree

tra

nsve

rse-

zone

bou

nd s

egm

ents

(no

rthe

rn, c

entr

al, a

nd s

outh

ern)

. (A

) M

ap w

ith

equa

l are

a pl

ots

of p

oles

to

M-

plan

es th

roug

hout

the

CR

syn

clin

e. (B

) Sev

en d

own-

plun

ge p

roje

ctio

ns th

roug

h th

e C

R s

yncl

ine.

The

CR

thru

st is

sho

wn

as a

thic

k, a

rrow

head

ed li

ne. O

ther

fi rs

t-or

der

faul

ts

are

show

n as

thi

n lin

es. B

ed t

hick

ness

for

the

sev

en-f

old

profi

les

are

dire

ctly

com

para

ble.

In

gene

ral,

a te

mpo

ral a

nd s

pati

al r

elat

ions

hip

for

fold

tig

hten

ing

exis

ts a

long

the

st

rike

of

the

fold

. In

mor

e de

tail,

how

ever

, the

fold

ing

hist

ory

is b

est

deve

lope

d w

ithi

n ea

ch t

rans

vers

e-zo

ne b

ound

seg

men

t.



ISMAT and MITRA


words, deformation patterns preserved in more open fold profi les can be used to represent the early stages of folding, while deformation pat-terns in the tighter fold profi les represent more advanced stages of fold tightening (Ismat and Mitra, 2001a).

A variety of small-scale structures are pre-served throughout the syncline and their geo-metric and kinematic characteristics are used to understand initial shortening of the beds and later folding and fold tightening. Some early structures were reactivated during later deformation phases and preserve a progressive history of deformation. During the deformation, outcrop-scale fracture networks formed and continuously evolved; many of the fractures produced in the early stages of deformation were healed during later stages, whereas others were reactivated (Bott, 1959; Price, 1967; Nick-elsen, 1979; Bles and Feuga, 1986; Ismat and Mitra, 2001a, 2001b). Most of the penetrative deformation was taken up by this fracturing and cataclastic fl ow.

The CR syncline is cut by transverse zones that strike at a high angle to the trend of the fold, defi ning three transverse-zone bound segments (Figs. 4A and 4B). The fold geometry and sec-ondary structures are unique to each transverse-zone bound segment. The syncline tightness correlates with the degree of development of the adjoining duplex-cored anticline. Therefore, the variable evolution of the duplex horses along strike may have controlled the location and/or activity of transverse zones and the geometry and deformation history of the adjacent syncli-nal segments.

Fracture Analysis

Because cataclastic structures are the most abundant tectonic features at the meso- and microscales, fracture networks/cataclasite zones were studied in detail and measured systemati-cally for over 70 representative sites throughout the CR syncline (Ismat and Mitra, 2001a) (Fig. 4A). Fracture networks were classifi ed into different sets based on crosscutting relationships and were used to defi ne different deformation stages. For each site, fracture sets measured at the outcrop scale were plotted on Equal Area ste-reonets, and the shortening direction for each set was determined from the conjugate-conjugate fracture patterns (see Ismat and Mitra, 2001a, for a more detailed explanation of this method). Typically, the youngest and reactivated fractures, as will be described, accommodated fold tight-ening, while the oldest fractures are related to pre- or early-folding deformation.

A goal of this paper is to consider the role of cataclastic deformation in the development of

FTB hinterland thrust sheets; therefore, from the perspective of energy usage, the fracture-bound blocks defi ned by the fracture networks/cataclasite zones were analyzed, both at the meso- and microscales. Average Surface Area/Volume (SA/V) ratios for the outcrop-scale blocks were determined for 17 representative sites in the syncline, using scaled clay models (Data Repository text 1).1 Outcrop-scale mea-surements of average cataclasite zone thickness with the SA values were used to determine the volume of mesoscale cataclasite zones.

As microscale cataclasite zones played a role in deformation, the Surface Area/Volume (S

V)

ratio of the microscale blocks and zone volumes were determined using Spektor-Chord analysis (Underwood, 1970; Mitra, 1978; Ismat and Mitra, 2001a) (Data Repository text 2). For these analyses, parameters were measured on 4–6 tran-sects each on 49 thin sections from throughout the CR syncline and adjacent connecting splay duplex (Data Repository text 2; see footnote 1).

Structural Stages

Four regionally signifi cant structural stages (A–D) are identifi ed in the CR syncline, based on structural overprinting at multiple scales (Fig. 5).

Stage A—Early Layer Parallel Shortening and Fault Parallel Shearing

Layer-parallel shortening (LPS) features are the oldest structures preserved in the CR syn-cline (Fig. 5). Typical microstructures include plastic deformation features such as undulose extinction, deformation bands, and deforma-tion lamellae (Fig. 6A). A weak grain-shape cleavage at a high angle to bedding (Fig. 7A) was modifi ed or crosscut by all later structures (Fig. 6B), suggesting that it developed fi rst. This plastic fabric is best preserved in open hinge profi les, where later folding did not modify the fabric. Accompanying bed-perpendicular sty-lolites are preserved at the micro- and outcrop scales and are most abundant and largest in the fi ne-grained PCp quartzites (Fig. 7B).

At the outcrop scale, thin-bedded quartzites such as the lower- and middle-Ct units devel-oped small-scale folds (Fig. 5, Stage A). We interpret the presence of asymmetric z-folds (looking north, Fig. 5, Stage A) in the hinge

region of the open syncline to result from fore-landward shear in the CR thrust sheet prior to syncline formation. The asymmetric folds are further modifi ed in the limbs of the syncline, probably by fl exural slip during fold tightening (Ismat, 2002; Ismat and Mitra, 2001b). Thicker beds in the PCc, PCm, and upper-Ct quartzites contain wedge faults with the acute bisector parallel to bedding indicating LPS, which is consistent with the prefolding deformation his-tory when beds were still subhorizontal. West-dipping wedge faults are dominant, suggesting a component of top-to-the-east (i.e., toward the foreland) shear (Fig. 5, Stage A) during early motion on the CR thrust.

The oldest outcrop-scale penetrative fracture networks are found primarily in the massive, thick-bedded PCc and Ct quartzites throughout the syncline. Representative sites in the PCc and Ct quartzites reveal shortening directions sub-parallel to bedding (Figs. 5 and 8A). This geom-etry suggests that these fracture sets formed as LPS structures when the beds were subhorizon-tal and were rotated as the beds folded (Ismat and Mitra, 2001a). These early fracture net-works in the quartzites are consistent across the transverse zones, suggesting that they developed before formation of the transverse zones.

Stage B—Early Large-Scale FoldingBed thicknesses in open fold profi les are

nearly constant. Well-developed bed-parallel shear fractures and deformation zones at the outcrop scale and hinge-perpendicular slicken-lines on bedding faces and fractures in quartzite units with prominent bedding surfaces, such as the PCm and middle-Ct quartzites, indicate that slip during folding was mainly localized along bedding. Older LPS structures, such as small-scale folds, were reactivated in the limbs of the CR syncline. For example, asymmetric z-folds within the Ct-quartzite layers are more open in the west limb and tighter in the east limb of the CR syncline, which is consistent with progres-sive modifi cation of the folds by fl exural slip during initial CR folding (Fig. 5, Stage B).

The high density of fractures and evidence for mesoscopic cataclastic fl ow within the hinge of the syncline suggest that tangential longitudinal strain was localized in these areas (Fig. 9A). At the mesoscopic scale, cataclastic fl ow is defi ned by the presence of fractures/cataclasite zones (Figs. 9A and 9B) that bound small blocks, which slide past one another. Massive quartz-ite units with cryptic, recrystallized bedding fabrics, such as the PCc, contain a high density of fractures. Stereographic plots of the fracture patterns/cataclasite zones yield shortening direc-tions approximately perpendicular to bedding in the hinge zone, indicating outer-arc extension

1GSA Data Repository item 2005082, methods of analyses used to calculate the SA/V of outcrop-scale fracture bound blocks, the grain size and S

V at the

microscale, and the percent volume of microscale cataclasite zones, is available on the Web at http://www.geosociety.org/pubs/ft2005.htm. Requests may also be sent to [email protected].





A

B

C

D

ReactivatedLPS fractures

Tightz-folds

Openz-folds

Steepnormal fault

Bed parallelflexural slip

ReactivatedLPS fractures

Low-angleextension fault

Openz-folds

Out-of-core thrust

Figure 5. Schematic diagrams representing the four struc-tural stages (A–D) of progres-sive deformation in the CR syncline. Secondary structures produced the following: Stage A: weak cleavage perpendicu-lar to bedding, bed perpen-dicular stylolites, asymmetric z-folds (looking north), con-jugate wedge faults (west-dip-ping faults are more dominant and shown as bold lines), and outcrop-scale fractures; Stage B: bed-parallel shear zones, tangential longitudinal strain accommodated by fracture networks (primarily in the hinge region), small out-of-core thrusts; Stage C: cataclasite and breccia zones parallel to transverse zones, hinge thick-ening and limb thinning by cataclastic fl ow, gently dipping extension faults, reactivation of bed-parallel slip in the east limb, reactivation of transverse zones, out-of-core thrusts; Stage D: steep normal faults.



ISMAT and MITRA


(Figs. 5, Stage B and 8B). Favorably oriented conjugate fractures and wedge faults of Stage A were reactivated as out-of-the-core thrust faults with small displacements in the inner hinge of the syncline, accommodating modest amounts of inner arc shortening (Fig. 5, Stage B). Over-all, this suite of hinge structures is indicative of hinge deformation with some, but not dominant, tangential longitudinal strain.

Less prominent fracture networks are observed in the fold limbs (Fig. 9B). The net-works have similar geometries in the east and west limbs within a host quartzite unit, which we interpret to indicate that these networks developed in a fold with initially similar limb dips (Figs. 5 and 8B). The east limb has a rela-tively constant dip of ~40° everywhere along the strike of the syncline, in contrast to the west limb, so we interpret the west limb to have had a dip of ≤40° when the early fracture networks developed. Because of this, we surmise that folding history along the syncline was similar for both limbs up to an interlimb angle of ~100°. The deformation features in the open fold pro-fi les (interlimb angle ≥100°) are consistent with initial folding by bed-parallel fl exural slip in limbs and small amounts of tangential longitu-dinal strain in hinges (Fig. 5, Stage B).

Stage C—Segmentation and Fold TighteningAlong most of the CR fold trace, as the syn-

cline tightened to an interlimb angle of



tilted toward the plunge direction (Mitra and Sussman, 1997; Lawton et al., 1997). Older CR conglomerates also show evidence for internal deformation during later stages of folding. For example, PCm quartzite clasts in older mixed-clast carbonate matrix conglomerates were later shattered and reworked into the carbonate matrix (Ismat, 2002). The CR conglomerates are also cut by out-of-the-core thrusts (Labels b and c on Figs. 4A and 4B) or folded by blind out-of-the core thrusts (Label d on Figs. 4A and 4B), which suggests that some of the conglom-erates were deposited before fold tightening.

Stage D—Normal FaultingThe youngest structures in the syncline are

normal faults (Fig. 5, Stage D) exposed in the southernmost segment (Figs. 3 and 4). Most of these faults trend E–W with one fault trending N–S. The E-W trending faults are subvertical and form a series of horsts and grabens. The N-S trending normal fault dips 75° to the west. Fracture intensity increases at the outcrop-scale and microscale toward the normal faults (Mitra and Ismat, 2001).

Where normal faults offset the CR conglom-erate, the original gray sandy-carbonate matrix

is altered to a red-sandy matrix and quartzite clasts are fractured. Travertine deposits are also found where the normal faults crosscut the con-glomerates. The normal faults may have tapped fl uid reservoirs leading to formation of the red matrix and travertine precipitation.

Summary of Structural Stages and its Relationship to FTB Development

Stage A structures (Fig. 5) are correlated with initial thrust propagation and emplacement. Initial shortening with basal shearing is typical of FTB thrust sheet development (Elliott, 1976; Marshak and Engelder, 1985; Geiser, 1988; Gray and Mitra, 1993; Smart and Dunne, 1997). Stage B correlates with early large-scale folding of the CR thrust sheet (Fig. 5) during emplacement of the CR-PVT sheet over a ramp in the PVT thrust in Albian-Aptian time (ca. 115–98 Ma) (Mitra, 1997). Continued uplift and erosion generated the oldest CR conglomerates, which were deposited in the core of the syncline after Stage B (Royse, 1993; Lawton et al., 1997). Stage C correlates with fold tightening of the CR syncline during growth of the adja-cent connecting splay duplex between the CR and PVT thrusts (Fig. 2) (Mitra and Sussman, 1997; Mitra, 1997). Duplex formation resulted from reactivation along the PVT at the time of a new major movement on the basal décol-lement during the Cenomanian to Paleocene (ca. 95–55 Ma) (Fig. 2) (Mitra, 1997). Stage D normal faults (Fig. 5) crosscut the synorogenic conglomerates in the syncline core and presum-ably offset all earlier structures in the CR thrust sheet. The EW-trending normal faults are most likely related to extension parallel to the fold hinge during late stages of fold tightening while the NS-trending normal faults are probably a result of postcontractional collapse of the CR culmination.

Although the CR thrust sheet is an internal thrust sheet in the Sevier FTB, elastico-frictional mechanisms dominated the small-scale processes that accommodated the large-scale changes in the thrust sheet geometry during much of FTB evolution. To compare the magnitude of this EF behavior to the deformation of other internal sheets, we estimated deformation magnitudes from EF processes in the CR thrust sheet.

QUANTIFYING THE DEFORMATION

The deformation in the CR sheet was mea-sured both by determining the total strain in the rocks and by estimating the energy involved in emplacing the CR thrust sheet. These values are compared to published results on strain distribution in the Sheeprock thrust sheet of the Provo salient (Mukul and Mitra, 1998) and

So

0.5 mm

S1

0.5 mm

SoStylolite

DB

U

S1

A

B

Figure 7. Various structures that formed during stage A at the microscale, looking north. Photomicrographs are from samples from the hinge region with bedding (S

o) subhorizontal.

(A) Photomicrograph of the Ct quartzites containing a weak LPS foliation (S1) at a high

angle to bedding (So). The foliation dips steeply to the west, indicating dextral shear due to

motion on the CR thrust. (B) Photomicrograph of PCp quartzites containing bed-perpen-dicular stylolites, deformation bands (DB), and undulose extinction (U). The foliation (S

1),

with respect to bedding (So), also indicates dextral shearing parallel to bedding.



ISMAT and MITRA


N = 55

Site 9

Bedding: 81,265

N = 44

Site 192

Bedding: 62,090

N = 62

N = 47

N = 81Bedding: 70,076

Site 42

N = 37

Site 198

Bedding: 50,079

STAGE A

west limb

PCc Ct

FOOL CREEK TRANSVERSE ZONE

OAK CREEK TRANSVERSE ZONE



N = 248Bedding: 39,268

Site 27

N = 169

Site 43

Bedding: 47,074

hinge zone

westlimb

eastlimb

A B

N = 69

Site 19

Ct

N = 139

Site 17

Ct

Bedding: 30,040 Bedding: 34,314

DAVIDSON CANYON TRANSVERSE ZONE DAVIDSON CANYON TRANSVERSE ZONE

Bedding: 70,080

Site 194

Bedding: 80,079

Site 4

STAGE B

Figure 8. Representative stereograms of fracture networks from various sites within the CR syncline. Based on crosscutting relationships, fracture networks are separated into different sets, which can be correlated to different stages of CR thrust sheet evolution. Thick great circles represent bedding. Shortening directions (dots) are determined from the conjugate-conjugate fracture sets. (A) Stage A: Orientation plots from representative sites within the syncline in the Ct and PCc quartzites. The shortening directions are subparallel to bedding and therefore suggest that early LPS conjugate fractures were rotated with the beds during folding. (B) Stage B: Stereographic plots of fracture poles and dominant fracture sets from the hinge region (sites 17 and 19). The shortening direction (black dot) is subperpendicular to bed-ding, which suggests that there was outer-arc extension in the CR syncline. Stereographic plots of fracture poles and dominant fracture sets from the two limbs (sites 43 and 27, both within the PCc quartzite). Site 43 in the west limb from the northern transverse-zone bound segment and site 27 in the east limb from the southernmost transverse-zone bound segment have comparable bedding dips, but in nearly opposite directions. Their shortening directions are also essentially mirror images. These patterns suggest that early folding along the entire syncline continued in a similar manner in both the east and west limbs up to ~40° dips. (Continued on following page.)





to estimates of work in emplacement of the McConnell thrust sheet in the Canadian Rockies (Elliott, 1976). These sheets are both mechani-cally dominant internal thrust sheets in other parts of the Cordilleran FTB that have similar deformation histories to the CR sheet but exhibit mainly plastic deformation.

Plastic Strain

The early LPS and fault-parallel shearing in the CR sheet took place in the quasi-plas-tic regime, mostly by dislocation creep and minor amounts of pressure solution (Fig. 7). The strains were quantifi ed from the different

quartzite units in the CR syncline by using the Fry technique (Fry, 1979; McNaught, 1994). Strain data were collected parallel to the transport plane on two transects across the syncline: one in the tightest part of the syncline in the northernmost segment and the other in an open part of the syncline in the central segment (Figs. 11A and 11B). The measured plastic strains in the transport plane are quite small, with most axial ratios in the range of 1.1–1.2 and a highest value of 1.29. The strain ellipse geometries along the two transects are approxi-mately the same, suggesting that these strains accumulated early in the deformation and were not modifi ed by fold tightening.

We can use the plastic strain to help unravel the folding history; the strain ellipse orientations are very consistent with respect to one another and with respect to bedding at each site, suggest-ing that there was very little or no independent rotation of fracture-bound blocks during the later phases of cataclastic deformation. In both limbs of the fold the long axes of the strain ellipses plunge toward the hinterland when the bedding dip is removed; they form an acute angle with respect to bedding, indicating top to the foreland shear (Ramsay and Huber, 1983; Mitra, 1994). The angle increases up-section, which indicates a greater contribution by LPS than prefolding fault-parallel shear higher up within the sheet.

Cataclastic Strain

Strain estimates for individual cataclastic structures could not be obtained because they lack discernible offset markers that allow the necessary measurement of displacements (Wojtal, 1989). Instead, cataclastic strain was estimated by comparing the total strain in strongly fractured quartzite beds to the measured early plastic strains and determining the differ-ence (Fig. 11) (Ismat and Mitra, 2001a). Given the lack of other structures, we surmise that removing the early plastic strain from the total strain yields the cataclastic strain (Ismat and Mitra, 2001a). The total strain in the intensely fractured and strongly deformed west limb of the CR syncline can be estimated by assuming plane strain in the transport plane and using the thickening or thinning of the beds there com-pared to the same beds in the east limb where there is a small amount of cataclastic deforma-tion resulting in minor amounts of strain. The average total strain in the west limb ranges from 1.5 to 1.75 as a function of fold geometry (Ismat and Mitra, 2001a). Also, because the early (plastic) strain axial orientations are consistent in adjoining blocks, cataclastic deformation is interpreted to have involved negligible net rota-tion between blocks.

N = 146

Site 194

Bedding: 70,080

N = 101

Site 29

Bedding: 65,080

N = 164

Site 42

Bedding: 70,076

west limb

PCc



C

DAVIDSON CANYON TRANSVERSE ZONE

STAGE C

Figure 8 (continued). (C) Stage C: Stereographic plots, one from each transverse-zone bound segment, representing a tight fold geometry. All three sites are from the PCc quartzites in the west limb. Fracture patterns and shortening direction dramatically change across the trans-verse zones, suggesting that the later stages of fold tightening are unique to each segment.



ISMAT and MITRA


The average cataclastic strain is 1.64 (with a steeply plunging long axis) in the tight northernmost segment of the CR syncline, equivalent to ~30% stretching (and thinning) of the beds in the vertical direction by cata-clastic fl ow; this mechanism allows the fold interlimb angle to be reduced to as low as 50°. In the middle segment, where the CR syncline is open, the cataclastic strain is 1.72 (with a gently plunging long axis), equivalent to 30% thickening of beds in the west limb of the fold. Fold tightening in this segment is accommo-dated by out-of-the-core thrusting, and the minimum interlimb angle reached is 87° (Ismat and Mitra 2001a).

Work Done in Cataclastic Deformation During CR Sheet Emplacement

An analysis of the work done for cataclastic deformation during sheet emplacement begins by considering total work, which is given by

Wt = W

p + W

b + W

g + W

i , (1)

where the total work (Wt) is the sum of the

work in propagating the thrust (Wp), work in

basal sliding (Wb), work against gravity (W

g),

and work in internal deformation (Wi) (Elliott,

1976; Mitra and Boyer, 1986). For thrust sheets of comparable size in a typical FTB the fi rst three terms (W

p + W

b + W

g) will be similar in

magnitude to the CR sheet. Here we evaluate only the work in internal deformation (W

i) in

the CR sheet, which deformed to a large extent by cataclastic processes, in order to compare it with other similar sheets that have undergone dominantly plastic deformation.

The CR sheet has a complex deformation history, and hence the internal work involved in its initial emplacement and later deforma-

tion are broken into different components to get the total internal work. Because the CR sheet deformed primarily by cataclastic fl ow, the total estimated internal work (W

i = W

ct)

required in this thrust sheet is

Wi = W

ct = W

k + W

f + W

m + W

s , (2)

where Wk is the work involved in kinking of

the thrust sheet, Wf is the work associated with

outcrop-scale fracturing, Wm is the work in

forming microscale cataclasite zones, and Ws

is the work in shearing within outcrop-scale cataclasite zones.

A representative volume for the work esti-mates is defi ned as the regional cross-sectional area for the sheet multiplied by a width of 1 km perpendicular to the cross section (Fig. 2A). During initial emplacement, the CR sheet was transported from a lower fl at at the Proterozoic cover-basement contact to an upper fl at in Jurassic rocks over a single ramp dipping ~25° (Mitra, 1997). If we represent this geometry with a simple kink model, the rear portion of the sheet went through one kink bend and the fron-tal part went through two kink bends. Following Mitra and Boyer (1986),

Wk = W

k1 + W

k2

= τ12

. tan θ. A1. 1 + 2 τ

12. tan θ. A

2. 1, (3)

where Wk1

is the work for kinking the rear portion of the wedge with a cross-sectional area A

1 and unit width, and W

k2 is the work for

kinking the frontal portion of the wedge with a cross-sectional area A

2 and unit width over two

bends. τ12

is the interbed shear stress and θ is the ramp angle. We use τ

12 = 2 × 107 Pa based

on long-term yield strength (Elliott 1976), θ

= 25°, A1 = 1.12 × 103 km2, and A

2 = 0.25 ×

103 km2, yielding

Wk = 1.51 × 1019 J. (4)

During emplacement and later reactivation, the energy required to deform the CR sheet by large-scale cataclastic fl ow was determined by estimating the energy needed to create the total outcrop-scale fracture surface area within the sheet and the work done by sliding on fracture

B

DZR

A 10 cm

B

DZ

B

Figure 9. Outcrop photos in the PCc quartzite from the (A) hinge (site 198) and (B) limb (site 3 with m2 grid for scale) regions. Refer to Figures 4A and 4B for site locations. Note recrystallized zones (R), deformation zones (DZ), bedding (B).

Figure 10. Z-fold (looking north) with low-angle extension faults in the Ct quartzite from the west limb of the syncline.





networks. The estimated work done by fractur-ing (W

f) only includes work needed in propa-

gating fractures and ignores fracture nucleation work because that quantity is usually negligible (Mitra and Boyer, 1986).

Wf = (S

e) (SA/V) (V

S) J , (5)

where Se is the surface energy/area of the frac-

ture surfaces, SA/V is the surface area/volume of the outcrop-scale fracture-bound blocks, and V

S is the volume of the thrust sheet. We

used a surface energy/area (Se) value of 55

Jm–2 (typical values range from 10 to 100 Jm–2): This empirical value generally includes the energy needed to propagate the plastic zone at the fracture tip (Paterson, 1978; Mitra and Boyer, 1986). The average SA/V ratio, deter-mined from scaled clay block models, is 7.25 × 104 m–1 (Data Repository text 1; see footnote 1) (Ismat, 2002), and the sheet volume (V

S),

which includes the footwall duplex, is 2.3 × 1012 m3 for a 1-km-thick slice along the line of cross section. Substituting into equation (5),

Wf = 9.4 × 1018 J. (6)

The work done by sliding along fracture net-works is accommodated by micro- and meso-scopic cataclasite zones. Considering the work done in microscale cataclasite zones:

Wm = (S

e) (S

V) (v

c) J , (7)

where SV is the surface area/volume ratio of the

clasts within the microscale cataclasite zones and v

c is the volume of the microscale cata-

clasite zones within the sheet. Microcataclasite zones compose ~1.5% of the rock (Data Reposi-tory text 2; see footnote 1). Therefore, the vol-ume of the microcataclasite zones (v

c) within

the entire CR sheet is ~3.5 × 1010 m3. We used an average S

e of 0.75 Jm–2, which is typical of

values for quartz at the grain scale ranging from 0.41–1.34 Jm–2 (Brace and Walsh, 1962; Mitra, 1984; Atkinson, 1989), and a S

V ratio of 1.12 ×

105 m–1 for the microscale clasts (Data Reposi-tory 2). Substituting into equation (7),

Wm = 1.9 × 1015 J. (8)

This term is negligible compared to the work done for larger fractures and can therefore be ignored. This relationship is supported by King’s (1983) analyses, which showed that within a fracture population, larger fractures make a signifi cantly greater contribution to the overall deformation.

The presence of the microscale cataclasite zones, as well as cataclasites along outcrop-scale fractures, indicates that both granulation and sliding along confi ned zones were important deformation processes accommodating block movement during large-scale (i.e., block-con-trolled) cataclastic fl ow (Ismat, 2002, 2003). Although granulation involves creation of new surface area, the work required for this process is small compared to the sliding term (Archard, 1953; Scholz, 1987; Mitra, 1993). So, we only estimate the work in sliding on the network of cataclastic fractures. Cataclasite zones that bound outcrop-scale blocks are thin, 4–5-sided tabular zones ~0.2 m long (L) and 0.2 m wide with an average thickness (t) of ~0.004 m; the average volume of one cataclasite zone is ~1.6 × 10–4 m3 (Data Repository 2). Based on a detailed analysis of the outcrop-scale fracture morphology, we estimate that ~1% of the total CR sheet is made up of outcrop-scale cataclasite zones. Therefore, the total volume of outcrop-scale cataclasite zones in the CR sheet is ~2.3 × 1010 m3. From this, we estimate that ~1.5 × 1014 outcrop-scale zones developed within the CR sheet.

The displacement along each zone is calcu-lated from a standard relationship of the form

1.14

1.24

1.21

1.17

1.18

1.16

1.12 1.06

1.241.12

A

B

( ) ( ) =totalε εεplastic catacl./ = 1.77 (33% bed extension)totalε

εplastic = 1.08 (at 10 to bedding) = 1.64 (28% bed extension)εcatacl.

1.22

1.14

1.201.05

1.241.10

= 1.14 (at 13 to bedding)

( ) ( ) =totalε εplastic εcatacl. = 1.52 (23% bed thickening)totalε

= 1.73 (31% bed thickening)εcatacl.εplastic

Figure 11. The cataclastic strain (assuming plane strain) is determined from the total strain (i.e., bed thickening or thinning in profi le view) and from the measured plastic strains. Plas-tic strains measured in the transport plane are plotted as strain ellipses on two down-plunge projections (see Fig. 4B); average axial ratios are given next to the strain ellipses. Strain is given as the extension ratio (or stretch) of the long axis.



ISMAT and MITRA


u = βtα , (9)

where u is displacement, t is zone thickness, β is the antilog of the intercept, and α is the exponent (slope) (Hull, 1988; Mitra, 1993). β typically ranges from 10 to 1000 (Argon and McClintock, 1966; Reed-Hill, 1973; Hull, 1988; Evans, 1990; Mitra, 1993). Here, we use a β value of 10 (a conservative estimate using the smallest likely displacement on the fractures) and an α value of 0.97 (Hull, 1988; Mitra, 1993). Substituting these values into equation (9)

u = 5.86 × 10–2 m. (10)

The amount of work required to slide on an individual outcrop-scale fracture or cataclasite zone is

ws = (C) (σ

s) (u) (L2) J , (11)

where C is an averaging factor that accounts for the amount of slip on each fracture surface/cataclasite zone, σ

s is the sliding stress, u is

the average displacement on an outcrop-scale fracture/cataclasite zone, and L2 is the aver-age surface area of each outcrop-scale fracture surface/cataclasite zone. C accounts for the gain and/or loss of displacement along these fractures and is a constant ~1 (Elliott, 1976; Mitra, 1993). The sliding stress (σ

s) is estimated by using the

Navier-Coulomb fracture criterion,

σs = c + µ σ

n , (12)

where c is the cohesive strength (0.5 × 108 Pa) (Twiss and Moores, 1992; Mandl, 2000; Scholz, 2002), σ

n is the normal stress (esti-

mated to be 1.1 × 108 Pa at the average depth of deformation), and µ is the coeffi cient of internal friction of 0.7. The coeffi cient value is the average of the frictional sliding crite-rion (µ = 0.81) and the coeffi cient of friction for the Coulomb fracture criterion (µ = 0.6) because cataclastic fl ow occurs by sliding and fracturing. Substituting into equation (12), σ

s

is estimated to be ~1.16 × 108 Pa for cataclasite zones (Mitra, 1984, 1993). Substituting into equation (11),

ws = 2.72 × 105 J , (13)

for one outcrop-scale zone. The total work required by all outcrop-scale cataclasite zones/fractures in the CR sheet is

Ws = Σ w

s = 4.1 × 1019 J. (14)

The total estimated work required in this thrust sheet deforming by cataclastic fl ow is

Wct = W

k + W

f + W

s (15)

= 7 × 1019 J. (16)

Even with conservative estimates about fractur-ing and sliding, cataclastic structures account for a larger portion of the work done than fold-ing (kinking) of the sheet over a ramp. Since the CR thrust sheet is an internal, i.e., dominant sheet (Woodward et al., 1989; Boyer and Elliott, 1982; Geiser and Boyer, 1987), these work estimates clearly suggest that cataclastic fl ow can be a major contributor to the evolution of a FTB wedge.

DISCUSSION

Comparison with “Plastic” Thrust Sheets

The strain in the CR thrust sheet can be com-pared with that in the Sheeprock thrust sheet in the adjoining Provo salient. The two sheets are in equivalent hinterland positions in their respective segments of the Sevier FTB. They each carry essentially the same Proterozoic through Mesozoic sections in their hanging walls and were emplaced at about the same time (Mitra, 1997). The CR sheet was trans-lated ~100 km in ~15–20 m.y. (a time-averaged rate of 2.1 × 10–10–1.59 × 10–10 m/s) while the Sheeprock was translated ~30 km in ~7 m.y. (a time-averaged rate of 1.36 × 10–10 m/s) although it had additional displacement due to promi-nent fault-propagation folding. The average strain axial ratios in the transport plane of the Sheeprock sheet range from 1.3 high up in the sheet to 1.5 lower down in the sheet, with strain ratios reaching 2.2 close to the thrust (Mukul and Mitra, 1998); most of this deformation took place by plastic deformation processes and by pressure solution, with only minor amounts of cataclasis during the late stages of deformation (Sussman, 1995). These strains are comparable to the strain axial ratios in the CR sheet, which range from average values of 1.5 in the open parts of the CR syncline to 1.75 in tighter parts of the fold, which was accomplished mostly by block-controlled cataclastic fl ow.

Is the energy required for this type of defor-mation within a large thrust sheet any different from a sheet emplaced under higher-temperature conditions? To answer this question, we compare our work estimates for cataclastic deformation in the CR sheet with that of the McConnell thrust sheet, Canadian Rockies (Elliott, 1976), which is the only published estimate for work done during internal deformation of a sheet deformed primarily within the quasi-plastic regime. The McConnell and CR thrust sheets are both large internal sheets in the Cordilleran FTB and are

deformed by younger thrusts. For a 1-km-wide slice the average size of the McConnell sheet is 1.1 × 103 km3, which is comparable to the average size of the CR sheet of 2.3 × 103 km3. The velocities of emplacement for the two thrust sheets are also similar, with the McConnell sheet having moved ~40 km in ~8 m.y.(at a time-aver-aged rate of 1.6 × 10–10 m/s). The estimated energy for internal deformation of the McCon-nell thrust sheet is ~6.4 × 1019 J (Elliott, 1976), which is very similar to the estimate of 7 × 1019 J for the CR sheet.

These comparisons show that the CR sheet is similar to other internal thrust sheets in the Cor-dilleran FTB in terms of deformation magnitude and work done for the deformation. Thus, even though it deforms dominantly by low-tempera-ture (EF) mechanisms, its role in the evolution of the FTB is similar to that of other internal thrust sheets that deform at higher temperatures (by QP mechanisms).

The Role of the EF/QP Transition in the Evolution of CR Structures and Implications for FTB Restoration

The microscopic- to outcrop-scale struc-tures and their crosscutting relationships can be used not only to evaluate the deformation history but also the varying conditions during different deformation stages. The early plastic microstructures in the CR sheet suggest that the rocks initially deformed within the quasi-plas-tic regime during Stage A of the deformation. Mesoscale conjugate fractures also formed, indicating that plastic and brittle deformation occurred together at different scales at the same depth (e.g., Schmid and Handy, 1991); conse-quently, the rocks may have been close to the EF/QP transition zone in the crust. Alternatively, deformation at different scales may have been sensitive to different variables, such as changes in grain size, strain rate, and fl uid pressure.

Later deformation within the CR syncline accommodated fold tightening of the CR syn-cline by block-controlled cataclastic fl ow (Ismat and Mitra, 2001a), which can be described as macroscale ductile deformation in the EF regime (Stearns, 1971; Borradaile, 1981; Jamison and Stearns, 1982; Hadizadeh and Rutter, 1983). At an even larger scale, fi rst-order (km scale) out-of-core thrusts and transverse zones also formed throughout the syncline during fold tightening. Thus, the CR thrust sheet shows a progressive change from small-scale QP mechanisms with larger-scale EF mechanisms in the early stages of deformation to pervasive EF mechanisms in the later stages.

Deformation mechanisms in the footwall of the CR thrust show a similar change from QP





to EF deformation mechanisms (Sussman, 1995; Mitra and Sussman, 1997). Microstructures in the quartzites of the footwall duplex show that the older, structurally higher horses in the duplex preserve a larger proportion of plastic deforma-tion structures than the younger, structurally lower horses, which show evidence for defor-mation dominantly by fracturing and cataclastic fl ow. This increase in cataclasis in the younger horses supports an interpretation of the duplex deforming with erosion due to rapid uplift as motion continued on the PVT-CR thrust system (Sussman, 1995; Mitra and Sussman, 1997).

The actual transition from QP to EF pro-cesses in the internal portion of an FTB wedge is strongly controlled by the location and dis-position of the transition zone between the EF

and QP regimes. This transition for quartz-rich rocks, such as those in the CR sheet, typically occurs at a depth where temperatures reach ~300 °C (Brace, 1965; Tullis and Yund, 1977, 1980; Turcotte and Schubert, 1982). The depth is a function of factors such as strain rate, fl uid pressures, and geothermal gradient. Based on the known translation of the CR sheet and plastic strains within it, the time-averaged strain rate during its initial emplacement over ~15 m.y. is ~1.37 × 10–15/s. But thrust sheet emplacement is typically episodic, as has been shown in other areas (Knipe, 1989), and we believe that the actual strain rate during episodes of deformation was most likely somewhat faster (10–13–10–14/s), similar to most other thrust sheets (Knipe, 1989). Periodic excursions of strain rates (Knipe, 1989)

at the faster end of this range may have caused conjugate fractures to form while the quartzites were still undergoing plastic deformation in the early stages of deformation in the CR sheet.

Estimates for the geothermal gradient inboard of the miogeoclinal hinge during the Creta-ceous thrusting are quite low; values range from 18 °C/km (Craddock, 1986) to 25 °C/km (Fletcher, 1984). Based on an average of these gradient estimates and a strain rate of 10–14/s (comparable strain rate to the CR thrust sheet), the depth to the EF/QP transition in the foreland would range from ~11 km to ~17 km (Fig. 12). By contrast, the estimated geothermal gradient for the Willard thrust sheet in the hinterland of the northern Utah segment of the Sevier FTB is 30–35 °C/km, based on fl uid inclusion and illite crystallinity studies and on thermal modeling (Yonkee et al., 1989). With such steep gradients, the EF/QP transition would have been at a much shallower level (~7.5–9 km) in the hinterland than in the foreland. In other words, the EF/QP transition zone would have a forelandward slope in the general vicinity of the miogeocline–shelf transition (Smith and Bruhn, 1984; Kulik and Schmidt, 1988).

First-order balanced restorations of the cen-tral Utah segment (Fig. 2), based on geometric criteria (line length and equal area balancing) and reasonable estimates of erosional depth, suggest that initial emplacement and erosion of the CR thrust sheet brought the CR quartzites to depths of ~10 km in the hinterland, close to the miogeocline shelf hinge (Mitra, 1997). The step-wise restored sequence (Fig. 2) shows that further erosion and displacement along the PVT thrust carried the rocks of the CR thrust sheet foreland-ward, inboard of the miogeoclinal hinge and onto the stable craton. On the craton, the geothermal gradient was much lower and the EF regime would have extended to a depth of ~11–17 km, whereas the CR thrust sheet would have extended from the surface to depths of ~5 km and, there-fore, well within the EF regime (Fig. 12). Thus, deformation in the CR thrust sheet would have initiated in the QP regime; the sheet was then translated into the EF regime and continued to fold tighten in the EF regime. We suggest that placing such constraints on the depths of rocks for successive stages of deformation is important for developing viable retrodeformable regional balanced cross sections, and in turn, geometric and kinematic models for wedge evolution.

Lithologic Controls in Wedge Evolution

Late-stage thickening in the CR culmina-tion was achieved by the growth of a con-necting splay duplex in the footwall and fold tightening in the hanging wall of the CR thrust

A

B C

150 300 450 600 750

5

10

15

20

25

SHEAR RESISTANCE (kbar)

5

10

15

20

25

30

1 2 3 4

10-1010-11

10-1210-13

10-14

SHEAR RESISTANCE (kbar)

5

10

15

20

25

30

1 2 3 4

10-10

10-11

10-1210-13

10-14

25 /km

17 /km

EF

/QP

TR

AN

SIT

ION

Z

ON

E

TEMPERATURE (ºC)

DE

PT

H (

km)

DE

PT

H (

km)

TS

N

EF/QPTRANSITION ZONE

DE

PT

H (

km)

TS

N

EF/QP TRANSITION ZONE

GEOTHERMALGRADIENT25 C/km GEOTHERMAL

GRADIENT17 C/km

Figure 12. (A) Graph of two different geothermal gradients. The EF/QP transitional zone is ~300 °C for dry quartzites. (B) Depth-strength graph for a geothermal gradient of ~25 °C/km. Note that the average depth for the EF/QP transition is ~11 km. (C) Depth-strength graph for a geothermal gradient of ~17 °C/km. Note that the average depth for the EF/QP transition is ~17 km. The EF-strength envelopes for thrust (T), strike-slip (S), and normal (N) faults are shown in parts (B) and (C).



ISMAT and MITRA


(Mitra, 2001). Both the CR thrust sheet and the footwall duplex (PVT sheet) contain abundant orthoquartzites with similar strengths and were deformed by cataclastic fl ow within the EF regime. Thus, the CR and PVT sheets, which constituted much of the rear wedge in this part of the Sevier FTB, behaved as one body that was actively deformed during the later stages of wedge evolution.

The presence of lithologies with similar strengths in successive thrust sheets of the rear wedge allowed fractures to form throughout the units, ultimately forming a stable network to enable deformation by block-controlled catacla-stic fl ow. Strain is accumulated by the jostling motion of outcrop-scale blocks. Smaller-scale cataclasite zones bound these outcrop-scale blocks and accommodated local strain incom-patibilities between the blocks by zone reactiva-tion, hardening, thickening, or thinning (Ismat and Mitra, 2001a).

For an orogenic wedge with approximately uniform material properties, the EF/QP transi-tion is not a distinct boundary (Rutter, 1976; Sibson, 1977); rather, deformation mechanisms change gradually over a broad range of depths. Moreover, the strengths of rocks reach their maximum magnitudes at the EF/QP transition, so that a detachment is least likely to develop there (Mitra, 1997). Because of these two rea-sons, thrust sheets within the wedge from the EF regime to depths up to and including the EF/QP transition will deform by similar mechanisms (Rutter, 1976; Knipe, 1989; Ord and Hobbs, 1989; Plesch and Oncken, 1999); detachments may form below the EF/QP transition, where the rock strength decreases dramatically (Mitra, 1997). This potential major break in the central-Utah segment would have formed at a depth well below the portion of the wedge that deformed inboard of the miogeoclinal hinge. So, this portion of the wedge deformed essentially as a single body with no major internal detachments and was detached from the basement because of lithological differences.

Should Cataclastic Flow Be Preserved in All FTBs?

Models of orogenic wedge behavior based on long-term fl ow strength (Elliott 1976) or an assumed plastic rheology (Chapple, 1978; Stockmal, 1983) have generally ignored the shallower rocks deforming in the EF regime because these generally comprise a minor component of the total deformation. On the other hand, wedge models that do incorporate EF mechanisms assume Coulomb behavior (Davis et al., 1983) of the wedge as a whole, which accounts for the orientation, location,

and frictional sliding along detachments but not the formation of a penetrative fracture network. Therefore, Coulomb wedge analyses cannot be directly applied to the hinterland portion of an orogenic wedge that deforms by ductile fl ow. Our evidence indicates that ductile deforma-tion in hinterland thrust sheets can take place in the EF regime by mesoscopic block-controlled cataclastic fl ow. This type of deformation would be expected in areas where hinterland sheets are transported past the miogeoclinal hinge and onto the craton where the geotherm is depressed. Fracture networks (and hence cataclastic fl ow) similar to the ones we have described here have been previously reported from the Moine thrust belt (Hadizadeh and Rutter, 1983) and from the Ogden duplex of the Sevier FTB (Yonkee, 1990; Yonkee et al., 1989) but have not been widely reported from other FTBs. This may be related to the low preservation potential of these rocks.

Evidence for cataclastic fl ow is seldom likely to be preserved in older orogenic belts due to erosion of the upper part of the wedge. More-over, synorogenic sediments derived from the wedge rarely preserve evidence for cataclastic fl ow because the clasts in the conglomerates are generally smaller than the original frac-ture-bound blocks found within the source thrust sheet. The clasts or blocks are internally undeformed because the deformation scale for block-controlled cataclastic fl ow is much larger than any of the individual clasts (Borg et al.,1960; Stearns, 1971; Borradaile, 1981). In younger, active orogenic belts, however, such as the Andes and the Himalayas, we should expect to fi nd evidence for cataclastic fl ow. In these cases, the orogenic wedges are still growing and propagating; the rocks are continuously being uplifted into the EF regime so that erosion has not yet removed such evidence. We suggest that structural analysis similar to that presented in this paper would be useful in these active belts.

Evidence for cataclastic fl ow is only pre-served in select circumstances. In the central Utah segment of the Sevier FTB, early extensive erosion (Royse, 1993) covered the entire wedge top with synorogenic sediments. Although these were recycled during subsequent deformation, a veneer of conglomerates ~1–2 km thick formed a protective cover that prevented the CR syn-cline from being eroded, so evidence for cata-clastic fl ow was preserved in these rocks.

CONCLUSIONS

This study examines the geometry and progres-sive deformation patterns in part of the Canyon Range (CR) thrust sheet, an internal thrust sheet in the central Utah segment of the Sevier FTB. The CR thrust sheet initially deformed outboard

of the miogeoclinal hinge and was transported upward and onto the stable craton, where much of the late-stage FTB deformation took place. The study leads to the following conclusions.

(1) Four deformation stages (A–D) are preserved within the CR thrust sheet that may record the passage of deformation fronts (Gray and Mitra, 1999) through the fold-thrust belt as a whole. Evidence for each stage has some com-ponent of EF mechanisms, such as cataclastic fl ow. Thus, cataclastic fl ow may have contrib-uted signifi cantly to FTB evolution.

(2) The preserved structures suggest that the type of deformation that took place at each stage was strongly infl uenced by the geothermal gra-dient, and they are consistent with the sugges-tion that the EF/QP transition slopes toward the foreland in retro-arc FTB settings. This varia-tion across strike strongly infl uences wedge evolution. The estimated geothermal gradient during the Cretaceous Sevier orogeny was ~30 °C/km in the hinterland and ~18–25 °C/km in the foreland.

(3) Within a depressed geotherm (inboard of the miogeoclinal hinge), the wedge may have behaved as one massive unit deforming by cataclastic fl ow.

(4) The average strain axial ratio in the CR thrust sheet ranges from 1.5 to 1.75 and is comparable to strains in other hinterland sheets deformed by crystal plastic processes. A minimum estimate for the amount of energy expended in the dominant CR sheet deforming by EF mechanisms (~7 × 1019 J) is comparable to estimates for such sheets deforming by QP mechanisms (e.g., ~6.4 × 1019 J for the McCon-nell sheet).

(5) Unraveling the deformation mechanisms helps in placing constraints on the location of the rocks through their history. This added con-straint enhances the viability of retrodeformed sections and may be helpful in developing new kinematic models of wedge evolution.

(6) We should expect EF mechanisms, such as cataclastic fl ow, to contribute signifi cantly to wedge evolution. Although evidence for cataclastic fl ow is often not preserved and/or not recognized, this study shows that it can play a dominant role in FTB evolution.

ACKNOWLEDGMENTS

Acknowledgment is made to the donors of the Petroleum Research Fund, administered by the Ameri-can Chemical Society, for support of this research under grant #33387-AC2 to G. Mitra. Support was also provided by NSF grant EAR-0001030 to G. Mitra. We thank the reviewers, William Dunne, Brian Horton, and the associate editor, Basil Tikoff, for their detailed and thoughtful comments, which were very useful in improving the paper. We also thank Adolph Yonkee for helpful comments on an earlier version of this paper.





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DeCelles, P.G., and Mitra, G

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