encouragingtheextrusionof...

Encouraging the extrusion ofdeep-crustal rocks in collisional zones

A. HYNES*

Department of Earth and Planetary Sciences, McGill University, 3450 University St., Montreal, Canada H3A 2A7

ABSTRACT

Most unroofing mechanisms invoked for the exhumation of blueschist-plus-eclogite terranes, includingcorner-flow and extensional collapse of the orogenic wedge, predict steep unroofing paths for thedeeply-buried rocks and are applicable only to unroofing from depths within the crust. Many high-Pand ultrahigh-P rocks of continental affinity are derived from greater depths than this. Their lack ofwarming during unroofing, together with indications that they may rest directly on less deeply buriedequivalents, are suggestive of shallow unroofing paths similar to those for the subduction-channelmodel. They are interpreted to have been emplaced by the upward extrusion of coherent slices ofcontinental crust, bounded below by thrust faults and above by normal faults, with unroofing pathsessentially reversing the original burial paths.Where continental crust has been subducted into the mantle, upward extrusion is probably driven

largely by buoyancy forces, although examples of upward extrusion without subduction into the mantleindicate that buoyancy forces may not be essential. Two features in addition to buoyancy may promoteupward extrusion. Slab breakoff may reduce the pull from the descending slab, and subduction-zonegeometry may change as a continental margin is dragged into the subduction zone. Both features maypromote the extrusion of continental crust at precisely the time at which it has been partially subducted.A close spatial relationship between a lateral ramp and a lobate zone of extruded high-P rocks in the

Mesoproterozoic Grenvillian orogen indicates that lateral ramps may be important in localizingextrusion. Lateral ramps disturb the two-dimensional flow, with channelling of material into the regionof the lateral ramp as it is extruded. Many exhumed ultrahigh-P terrains are associated with jogs in thetrends of orogenic fronts that may reflect the presence of lateral ramps at depth. Ultrahigh-P rocks maybe expected to be concentrated at such jogs, and may record the channelling in their deformationhistory.

KEYWORDS: blueschist, collision zones, eclogite, Grenville orogeny, subduction-channel model, unroofing

mechanism, ultra-high pressure terranes.

Introduction

SINCE the first recognition of the conditions under

which they occur, mechanisms for the preserva-

tion of blueschists and eclogites at the surface

have been the subject of much debate. The

preservation of these high-P-low-T metamorphic

rocks requires their burial in regions with

unusually low geotherms, followed by unroofing

without warming to normal geotherms.

Mechanisms for the burial are not particularly

problematic. Any deep-seated thrust system, and

particularly one associated with a subduction

zone, produces the requisite conditions in its

footwall (e.g. Oxburgh and Turcotte, 1974;

England and Richardson, 1977; England and

Thompson, 1984). Unroofing without warming,

however, requires either rapid unroofing, or

unroofing under conditions that prevent

warming, and the search for such mechanisms

has been challenging. The now widespread

recognition of high-P and ultrahigh-P (in the

field of coesite stability; >~2800 MPa) metamor-

phosed rocks with continental affinities (e.g.

Compagnoni et al., 1977; Seidel et al., 1982;

# 2002 The Mineralogical Society

* E-mail: [email protected]: 10.1180/0026461026610013

Mineralogical Magazine, February 2002, Vol. 66(1), pp. 5–24

Chopin, 1984; Griffin et al., 1985; Okay et al.,

1989) has broadened this challenge, since it

indicates that continental crust partially subducted

into the mantle may fairly commonly be returned

to the surface. In this paper, I review evidence for

the possible paths followed by deeply buried

rocks on their way back towards the surface, and

the mechanisms that have been suggested for such

paths. I also show that formerly deeply-buried

rocks appear to occur at very specific localities

along orogenic belts, which may reflect features

of the geometry in the subsurface that were

instrumental in effecting their unroofing.

Paths to the surface

Blueschists, high-P and ultrahigh-P continental

rocks and associated eclogites almost certainly

develop their mineral assemblages in the foot-

walls of reverse faults, where the requisite high P

may be achieved at low temperatures. Such

reverse faults could develop within the orogenic

wedges of crustal-thickening zones, in which case

the hanging wall would be crustal, or could be the

reverse faults associated with subduction zones, in

which case the hanging wall could include mantle

material. In the case of many blueschist-plus-

eclogite terranes, the melange-like character of

the assemblage and the incorporation of frag-

ments of ophiolite complexes provide compelling

evidence for their formation in subduction-zone

settings (cf. Ernst, 1973, 1984). In the case of

high-P metamorphosed continental rocks, meta-

morphic pressures of 2000 MPa and higher are

indicative of burial to depths greater than that of

the thickest continental crust, and therefore also

require a subduction-zone setting. In this paper,

discussion is largely restricted to rocks that are

assumed to have been metamorphosed during

burial in subduction zones. There are essentially

two paths that might be followed by such buried

material in its passage back to shallower depths.

One path is steep, transporting material up

through the accretionary prism or orogenic wedge

towards the surface. A path such as this was

advocated by Cowan and Silling (Cowan and

Silling, 1978) in an analysis of accretionary

prisms. In their model, the upward flow results

from the inhibition of continued flow of low-

viscosity material down the subduction zone, due

to the presence of a rigid buttress at the rear of the

accretionary wedge. This type of flow is generally

referred to as ‘corner flow’, because the low-

viscosity material flows around the corner defined

by the point at which the subduction zone and the

rigid buttress meet (Fig. 1a). The path followed

by material is a function of the geometry of this

corner and of the position of the material in the

wedge with respect to the buttress. Upward flow

occurs throughout the wedge, which is treated as a

viscous fluid with uniform rheology.

A similar path was invoked by Platt (1986) who

showed that it is a predictable consequence of

wedge dynamics if there is any significant amount

of underplating in the rear of the wedge. As the

wedge is underplated, it is thickened beyond its

critical-taper geometry and responds by

extending. For continuous or repeated under-

plating, such a mechanism could achieve a

substantial amount of unroofing of deeper parts

of the wedge. The steep path followed by the

rocks reflects simply the fact that material has

been removed from above them, probably largely

by normal faulting. One such geometry for the

normal faulting is shown on Fig. 1b, although

there are others (Platt, 1993).

A steep path could also be achieved through the

development of sloping upper surfaces to an

orogenic wedge where it is floored by locally

steep ramps, as suggested by Jamieson and

Beaumont (Jamieson and Beaumont, 1988;

Fig. 1d, this paper). In this model, unroofing of

the rocks is due to greater erosion rates at the

steep upper surface of the wedge where it overlies

the ramp at depth. Erosion alone would give rise

to a vertical path; continued inward flow of

material into the wedge concomitant with erosion

would lead to a path inclined in the same sense as

the thrusting. In this scenario, the most deeply

buried rocks would be exposed above and slightly

in front of the underlying ramp. Although erosion

has traditionally been regarded as a slow

unroofing mechanism, some recent estimates

(Burbank et al., 1996, 2002) indicate that rates

of erosion in the Himalayas may be adequate to

satisfy the time constraints required by the

preservation of metamorphic assemblages.

An alternative path that could be followed by

material is essentially to reverse the path of

subduction. Such a path was invoked by Cloos

(1982) and Cloos and Shreve (1988) in a variant

of the corner-flow model, with flow restricted to a

narrow ‘subduction channel’ above the subduc-

tion zone, due to the limited width of the region of

low-viscosity material (Fig. 1c).

For all the models of Fig. 1 except the

subduction-channel model, the maximum depths

from which metamorphic rocks could be unroofed

6

A.HYNES

are limited by the thickness of the crust overlying

the subduction zone. The pressures of meta-

morphism in the ultrahigh-P terranes of the

Western Alps (Chopin et al., 1991), the Western

Gneiss region of Norway (Andersen et al., 1991)

and the Dabie Shan (Okay et al., 1993) are,

however, as high as 3000 MPa, indicating

exhumation from significantly below the base of

the continental crust (cf. Andersen et al., 1991).

Although, in principle, there is no limit to the

depths from which material might be derived in

the subduction-channel model, very low viscos-

ities are required in the channel, and the resulting

exhumed high-P material should be relatively

incoherent, in contrast to the characteristics of

many ultrahigh-P terranes (cf. Platt, 1993).

Although the subduction-channel model per se,

with low viscosities in the channel, may not be

appropriate for the unroofing of coherent, high-

and ultrahigh-P terranes in collisional orogens,

the terranes may have followed similar paths. As

was pointed out by Ernst (1975), high-P terranes

are commonly well organized, typically with

increasing pressures towards the structural top of

the exhumed complex, and faults dipping

uniformly in one direction. Ernst (1975) attributed

FIG. 1. Proposed mechanisms and paths for the return of rocks deeply buried at convergent plate-boundaries. The

paths are shown by the heavy dashed lines, and the material unroofed has the dotted ornament. (a) Corner flow

(adapted from Cowan and Silling, 1978). (b) Wedge spreading due to underplating (adapted from Platt, 1993). (c)

Subduction-channel-flow (adapted from Cloos and Shreve, 1988). (d) Enhanced erosion above a basal ramp (adapted

from Jamieson and Beaumont, 1988). Scales estimated by author; all figures have approximately 26 vertical

exaggeration.

ENCOURAGINGTHE EXTRUSIONOF HIGH-P ROCKS

7

these relationships to the successive underplating

of tectonic slices in a subduction-zone setting,

followed by buoyant return of the earlier

underplated slices, on a pathway that essentially

reversed the path of the original burial process.

The deep-seated rocks travelled upwards with

respect to both the underlying rocks and the

immediately overlying rocks, in the direction of

vergence for the collisional zone. In a process of

this kind, the slices would be bounded by thrust

faults below and normal faults above, both of

which had deep-seated roots. This process is a

form of extrusion, but in the vertical plane rather

than the horizontal plane as in the extrusion

associated with continental escape (e.g. Molnar

and Tapponnier, 1975).

10

0k

m

Sino-Korean cratonYangtze craton Wudan

block

mantle

mantle

oceanic crust

Great Valley sequence

oceanic crust

central belt

eastern belt

50

km

10

0k

m5

0k

m

LAURENTIABALTICA

Mid-Silurian (pre 425 Ma)

Late Silurian - Early Devonian

a

b

c

FIG. 2. Three examples of upward extrusion reversing the path of subduction: (a) From the Franciscan of California

(after Harms et al., 1992); (b) from the Dabie Shan of SE China (after Ernst and Liou, 1995); and (c) from the

Norwegian Caledonides (after Hurich, 1996).

8

A.HYNES

Three schematic representations of upward

extrusion reversing the path of subduction in

specific geographic localities are depicted in

Fig. 2, including one in which the descending

plate was oceanic and two in which extrusion

occurred during continental collision. All involve

initial subduction of the rocks to depths beneath

the mantle of the overlying plate. Two variants of

this model are what Wheeler (1991) called the

‘pip’ model and the ‘continental-sheet’ model,

depending on whether the extruding body was

continuous to the surface (Fig. 3).

Upward extrusion of deeply buried rocks by

reversal of the path of subduction has been

invoked increasingly in the recent literature,

particularly to explain the unroofing of high-P to

ultrahigh-P continental terranes. Examples include

the Dora Maira massif in the European Western

Alps (Hsu, 1991; Wheeler, 1991), the Dabie Shan

in China (Maruyama et al., 1994; Ernst and Liou,

1995), the Western Gneiss region in the

Caledonides of Norway (Hurich, 1996), the

Maksyutov Complex in the southern Urals

(Hetzel et al., 1998), the Galicia Massif in the

Variscan Belt of northwestern Iberia (Matte,

1998), and the Phyllite-Quartzite nappe of the

Cretan Hellenides (Wijbrans et al., 1993; Jolivet et

al., 1996). Extrusion models have also been

invoked for the classic Franciscan blueschists

(Ernst, 1975; Harms et al., 1992; Ernst and Liou,

1995).

There are two principal reasons for appeal to an

extrusion path that reverses the subduction path.

First, there is evidence in many cases that the

a

b

crust

crust

lithospheric mantle

lithospheric mantle

100

km

100

km

FIG. 3. The ‘pip’ (a) and ‘continental sheet’ (b) models for extrusion of high-P rocks, after Wheeler (1991).


9

high- to ultrahigh-P metamorphosed rocks have

experienced the bulk of their decompression

without significant heating, or even with cooling

during decompression (Hacker and Peacock,

1995). Decompression without associated

warming could be achieved only through very

rapid unroofing, or through unroofing while being

continually underthrust by cooler material (cf.

Davy and Gillet, 1986). Cooling during decom-

pression requires associated continual under-

thrusting of cool material regardless of the

unroofing rate, and is strongly suggestive of an

unroofing path that reverses the subduction path

(Hacker and Peacock, 1995). A second line of

evidence for reversal of the subduction path

comes from the observation that high-P conti-

nental rocks are typically in tectonic contact with

lower-P underlying rocks to which they are

possibly related. That is, the high-P rocks have

been transported over rock units that may have

been their equivalents, without the interposition of

exotic tectonic slices. Lithological units of the

Dora Maira massif in the European Western Alps,

for example, are lithologically similar to those of

the lower-grade Brianconnais zone (Borghi et al.,

1984; cited in Wheeler, 1991), which underlies

them structurally. The Western Gneiss Region of

the Norwegian Caledonides has similarities to the

Precambrian basement in the orogenic foreland

(Gorbatschev, 1985), and the Dabie Shan

Complex of southeastern China, which is

thought to have been derived from the Yangtze

block, is in direct tectonic contact with under-

thrust basement of the Yangtze block (Okay et al.,

1993). These kinds of relationship, although not

as compelling as the metamorphic arguments,

given the problems with identifying the prove-

nance of metamorphosed terranes, are difficult to

explain with pathways to the surface steeper than

those along which the rocks were originally

buried.

As discussed above, the preservation of high-P

metamorphic assemblages does not require rapid

unroofing if unroofing occurs along the subduc-

tion path. Indeed, a slow rate of unroofing would

provide further support for an unroofing path that

reversed the subduction path, since it would

increase the need for a mechanism for keeping

the rocks cool. In this context, it is noteworthy

that unroofing rates for the ultrahigh-P rocks of

the Dabie Shan appear to have been slow enough

to have permitted substantial thermal relaxation

(Hacker et al., 2000). The absence of evidence for

such thermal relaxation therefore provides strong

support for the suggested subduction-zone

unroofing path in this particular case. There is,

however, evidence in other cases that extrusion

occurred shortly after burial. In the Aegean, some

rocks were decompressed 400�600 MPa within

4 Ma of their burial (Jolivet et al., 1996), and the

Tso Morari eclogites of the northwest Himalaya,

which are interpreted as partially subducted parts

of India (de Sigoyer et al., 1997), were

decompressed from ~2000 MPa to ~900 MPa in

only ~8 Ma (de Sigoyer et al., 2000). The balance

of this paper is concerned with a review of

features that might enhance such rapid extrusion

up the subduction-zone path.

Factors contributing to extrusion

The net effect of extrusion is that the zone in

which the crust is thicker than normal becomes

wider (and less thick) than it was before. The

geometric redistribution is therefore similar to

that which would occur with gravitational

spreading of thickened and elevated continental

crust, but the redistribution is probably not driven

by relative elevation. A major potential driving

force for extrusion is presumably the buoyancy of

the continental crust relative to the mantle, when

the continental crust has been partially subducted

into the mantle. This buoyancy supplies a

substantial, upward-directed body force. The

potential importance of buoyancy as a drive for

extrusion was discussed by Platt (1987), was

emphasized by Hsu (1991) and Ernst and Liou

(1995), and has received strong support in recent

years from the analogue modelling studies of

Chemenda and co-workers (Chemenda et al.,

1995, 1996). Although it has been suggested that

the buoyancy drive for the uplift of crustal rocks

from the mantle could be effective only until such

rocks reached the base of the crust, it was pointed

out by Wheeler (1991), and amply confirmed by

the analogue studies of Chemenda et al. (1995,

1996), that if there is still crustal material within

the mantle below that which has already reached

the Moho, the drive is still there.

The buoyancy of partially subducted crustal

material undoubtedly provides a drive tending to

restore these rocks to shallower depths. In and of

itself, this drive does not guarantee the path of this

restoration, but the path is most likely governed

by the disposition of weak regions in the

neighbourhood of the buoyant rocks and in the

zones overlying them (cf. Wijbrans et al., 1993).

In Chemenda’s analogue modelling, it is the

10

A.HYNES

lowermost crust that is the weakest zone,

modelled with a material of lower yield strength,

and the crust above this zone travels as a unit

above it. In reality, the rheology of the crust and

surrounding mantle is considerably more complex

than has so far been addressed in analogue

models. The variation of yield stress in the

region of a subduction zone may be assessed by

combining estimates of the temperature distribu-

tion in subduction zones with estimates of the

strength of the rocks (Fig. 4). Figure 4 was

derived from modelled temperature distributions

for the subduction of continental lithosphere (van

den Beukel, 1992), combined with the assumption

that rocks yield by the lower of the stress required

by Byerlee’s law (Byerlee, 1978) and the stress

necessary for creep. The temperature-dependent

flow laws for creep of wet olivine (Chopra and

Paterson, 1984) and diabase (Shelton and Tullis,

1981) were used for the mantle and the crust

respectively, assuming a strain rate of 10�15 s�1.

Strength variations are illustrated for the tempera-

ture distributions in initially relatively cool

(Fig. 4a) and relatively warm (Fig. 4b) conti-

nental lithosphere. Both temperature distributions

result in a relatively weak zone coincident with

the crust in the descending slab, due to the lower

viscosity of diabase compared with olivine,

overlain by a strong lid of upper mantle, due to

the low temperatures in the mantle there.

Buoyancy forces begin to take effect as soon as

continental crust passes beneath the Moho (A, on

Fig. 4) and increase in intensity as more crust is

subducted. From Fig. 4, it is clear that the return

route followed by crust subducted to point A, or

anywhere deeper, would probably follow the

route of original subduction, which is the

channel in which strength is lowest. Its return

towards the surface by a steeper route is prevented

by the strong lid of mantle material overlying the

descending slab. In Fig. 4, weakening of the

mantle lid due to the generation of melts above

the subduction zone is not considered. This melt

generation occurs, however, typically only above

FIG. 4. Isotherms (8C) and contours of yield strength (MPa) in the neighbourhood of a continent/ocean lithospheric

boundary subducting beneath a continental lithospheric plate. Temperature distributions after van den Beukel (van

den Beukel, 1992). (a) For continental lithosphere with an initial surface heat flow of 50 mWm�2. (b) For continental

lithosphere with an initial surface heat flow of 90 mWm�2.


11

regions in which the Benioff zone has reached

depths of 100 km (Dickinson, 1973) in accor-

dance with the depths estimated for volatile

release from the descending slab (Peacock,

1990); i.e. at depths beyond the right hand side

of Fig. 4, and is therefore unlikely to affect the

return route followed by the crust.

Although buoyancy is undoubtedly instru-

mental in the restoration of partially subducted

crustal material to shallow levels, there may be

other changes at subduction zones that produce

geometric adjustment in the convergent zone.

Two such possibilities are slab breakoff and

change in the configuration of the lower litho-

spheric plate. Both such effects could promote the

extrusion of subducted continental material.

Slab breakoff is a process envisioned by Davies

and von Blanckenburg (1995), in which the partial

subduction of a continental margin gives rise to

extensional failure of the descending plate. This

failure allows access of asthenospheric material to

the upper plate behind the subduction zone and,

more importantly in the context of this paper,

significantly reduces the slab pull on the lower

plate. In the scenario envisaged by Davies and

von Blanckenburg (1995), this results in the

development of a crustal-scale duplex, the lower

sheets of which were originally subducted into the

mantle but were subsequently extruded into the

lower parts of the accretionary wedge. The

driving force for the upward extrusion is again

the buoyancy of the partly subducted continental

crust, but before slab breakoff, this buoyancy

force was effectively counteracted by the slab-

pull effect. This slab-breakoff model has been

used by von Blanckenburg and Davies (1995) to

explain extrusion of the Dora Maira massif in the

Western Alps.

Based on Davies and von Blanckenburg’s

(1995) mechanical arguments, slab breakoff

might be a direct result of the attempted

subduction of a continental margin, so that the

associated extrusion might be expected wherever

continental margins are partially subducted (cf.

Wong et al., 1997). The slab-breakoff model is

testable, in that access of asthenospheric mantle to

the base of the upper plate should raise geotherms

and produce mantle-derived partial melts,

although such melts would also be produced by

foundering of the lithospheric mantle of the upper

plate as envisaged by Houseman et al. (1981).

The effect of changes in the geometry of the

descending slab has been suggested as a possible

mechanism for the unroofing of high-P rocks,

both where the descending slab is oceanic (Harms

et al., 1992) and for the attempted subduction of a

continental margin (Hynes et al., 1996). In the

Franciscan of California, Harms et al. (1992)

suggested that the extrusion of high-P rocks

60�50 Ma ago resulted from shallowing of the

underlying subduction zone. In essence, their

suggestion was that emplacement of the

Franciscan rocks resulted from adjustment of the

wedge geometry due to shallowing of the basal

plane. The Franciscan orogenic wedge appears to

have been built up against and partly beneath a

thick ophiolite sequence that underlies the Great

Valley, and was accreted to North America in the

Late Jurassic (e.g. Godfrey and Klemperer, 1998).

This ophiolite sequence may have provided a

relatively rigid cap to the wedge, thereby

explaining why adjustment of the wedge was

concentrated in its basal regions. Thus, a

relatively small shallowing of the deeper part of

the subduction zone may have driven the basal

unit of the overlying accretionary wedge a

considerable distance back up the subduction

zone. There is independent evidence for shal-

lowing of the subduction zone in the Franciscan

region at the appropriate time, from the distribu-

tion of subduction-related volcanic rocks (Coney

and Reynolds, 1977), providing support for the

hypothesis.

Hynes et al. (1996) conducted a theoretical

study of the effect of a continental margin arriving

at a subduction zone. They showed that thermally

mature continental-margin lithosphere would

have greater strength than typical oceanic litho-

sphere, and they investigated the effect on the

overlying orogenic wedge of the increase in

flexural strength of the subducting slab as the

continental margin enters the subduction zone. If

the load on the descending slab is held constant,

the deeper parts of the descending slab tend to

rise. If, on the other hand, the descending slab is

not permitted to rise, but is held in place, it

deepens at shallower depths. Neither scenario is

strictly correct, but the true situation probably lies

somewhere between them, and together they

permit an assessment of the effect on the

subduction zone of the arrival of a stronger plate

at the trench. The net effect of the change in

geometry is for the lower plate to exert an upward

pressure on the subduction zone at depth, and a

downward pressure at shallower depths (Fig. 5).

This produces a suction effect, tending to pull

material up from depth in the subduction zone.

Calculations (Hynes et al., 1996) suggest that the

12

A.HYNES

suction effects achieve their maxima in the depth

range between 30 and 70 km (Fig. 5), where

pressures due to the change in geometry exceed

600 MPa. These values are certainly adequate to

overcome the yield strength of the lower parts of

the continental crust (cf. Fig. 4). They provide, in

effect, a boost to the buoyancy forces operating

on partially subducted continental crust, which

occurs as a direct result of the conditions that

produce partial subduction of the continental crust

0 100 200 300 400

Dep

th(k

m)

Distance (km)

120

0

0

0

0

0

0

0

0

0

20

20

40

10

30

0

20

30

50

70

60

80

100

120

120

100

+

-

-

-

-

-

-

-

+

+

+

+

+

+

continentocean

FIG. 5. Pressures exerted in a subduction zone as a result of the change in geometry of the descending slab as an

ocean-to-continent transition is subducted, after Hynes et al. (1996). The approximate position of the ocean-to-

continent transition is shown by the heavy stippled region. Increased pressures are shown by the shaded regions

marked with plus signs. Decreased pressures are shown by shaded regions marked with negative signs.


13

in the first place; i.e. the arrival of a continental

margin at the subduction zone.

Given all the possible aids to the extrusion of

continental crust outlined above, it is perhaps not

surprising that evidence for the unroofing of

partially subducted continental crust in continental

collision zones is quite common. Once continental

crust has been subducted any significant distance

into the mantle, the buoyant drive for its return,

combined with its progressive weakening as it

warms up with descent, would almost inevitably

lead to its return, although it may not in general

return all the way to Earth’s surface. This

tendency to return could probably be nullified

only if some adjustment to the subduction-zone

geometry worked in a sense opposite to that

addressed by Hynes et al. (1996). For example, if

progressively weaker lithosphere was subducted,

adjustments to the geometry of the subduction

zone would lead to pressure gradients that tended

to trap partly subducted continental material at

depth, despite its buoyancy.

Upward extrusion without a buoyancy drive

Probably the best support for upward extrusion of

metamorphic rocks as coherent slices bounded

above by normal faults and below by thrust faults

comes from the Himalayas, where Hodges and co-

workers (Hodges et al., 1992, 1996) have

documented the interplay of movements on

thrust faults (the Main Central Thrust system)

and overlying normal faults (the South Tibetan

Detachment Zone) in the Early Miocene exhuma-

tion of the Greater Himalayan crystalline terrane.

In this case, however, metamorphism in the

extruded sequences attained pressures of only

1000�1200 MPa (Hodges et al., 1996), consistent

with rooting of the faults at depths within the

orogenic crustal wedge rather than in the mantle.

Although there are metamorphic rocks in the

Himalayas that exhibit pressures consistent with

continental subduction (de Sigoyer et al., 1997)

they are at a higher structural level, and were

metamorphosed and unroofed much earlier (de

Sigoyer et al., 2000) than the rocks beneath the

South Tibetan Detachment.

The existence of upward extrusion within the

Himalayan orogenic wedge, as distinct from

upward extrusion from beneath the Moho,

suggests that buoyancy resulting from the

subduction of crustal rocks to depths beneath the

Moho is not a necessary requirement for the

upward-extrusion movement pattern. In the

Himalayas, Hodges et al. (1996) appealed to

incremental adjustments to the orogenic wedge as

a mechanism for the extrusion, with former thrust

discontinuities supplying the planes of weakness

used as normal faults and thrust faults in the

adjustment. These adjustments could be in

response to episodic underplating (cf. Platt,

1986), but there is little support from deep

seismic soundings (e.g. Hauck et al., 1998) for

the requisite amount of underplating. A more

compelling case can be made, based on the

seismic data, for adjustment of the mantle wedge

due to shallowing of the dip of the base of the

wedge. In the palinspastic reconstructions of

Hauck et al. (1998, Fig. 8), rocks of the Greater

Himalayan belt and the Tethyan Himalayan belt

must mount a substantial crustal ramp during the

early stages of motion on the Main Central

Thrust. Once this ramp had been mounted, the

basal slope of the orogenic wedge should have

declined markedly. Response to this change alone

could have occasioned much of the adjustment of

the wedge evident in motion on the South Tibetan

Detachment. Later more minor adjustments would

also be expected with increasing transport of the

orogenic wedge southwards onto the Indian

foreland, due to a gradual decline in basal slope

dictated by the geometry of the relatively rigid

Indian lithospheric substrate. On the basis of the

southern Himalayan example, therefore, it is

perhaps unwise to consider buoyancy a necessary

condition for upward extrusion, although it is

probably a very important component of the drive

in circumstances in which continental crust has

been thrust beneath the Moho.

Lateral ramps and the localization of high-Procks in collision zones

Upward extrusion, driven by buoyancy or by

adjustments to the geometry of bounding rigid

bodies or some combination, may therefore be an

important feature of the tectonics of convergent

plate boundaries. There is evidence, in addition,

that lateral ramps may localize the extrusion. The

impetus for this suggestion comes from studies in

the Mesoproterozoic Grenvillian continental-

collision zone of eastern North America. There,

it can be shown that high-P rocks, metamor-

phosed at pressures up to 1800 MPa in a

continental-collisional setting, were extruded

towards the tectonic foreland shortly after their

burial (Indares et al., 1998, 2000; Hynes et al.,

2000). The high-P rocks are particularly well

14

A.HYNES

developed at the surface in a broad lobate zone

which is immediately adjacent to and in front of a

steep lateral ramp that has been delineated in the

subsurface based on seismic-reflection profiling

(Hynes and Eaton, 1999; Fig. 6, this paper).

Regional tectonic trends in the Grenvillian

orogen are ENE, but the subsurface ramp adjacent

to the high-P domain trends NNE (Fig. 6).

Lineation directions, parallel to transport direc-

tions in the high-P rocks, show clearly that the

high-P rocks at the surface have flowed out into

their present positions from the region of the

20 km

4

8

12

16

24

28

32 km

20

Gre

nvFr

illeon

t

lineation

Man

icouag

an Res.

52°N

68°W

69°W

69°W70°W

52°N

51°30’N

68°W

Atla

nticO

cean

Appal

achia

nsG

renvill

e

Superior

Yavapai-

Mazatzal

Tra

ns

-Hudson

Triassic impact-related volcanic rocks

High- rocks(MIZ)

P

PalaeoproterozoicLaurentian metasediments

Undifferentiated gneisses

Palaeoproterozoic meta-sediments & Archaean bsmnt.

Reworked LaurentianArchaean

Cratonal LaurentianArchean

FIG. 6. Structure contours, at 2 km intervals, on the surface forming the base of the high-pressure Manicouagan

Imbricate Zone (MIZ) in the Mesoproterozoic Grenvillian orogen of eastern Quebec, together with the directions of

lineations.


15

lateral ramp (Hynes et al., 2000; Fig. 6, this

paper).

Geometrical relationships in the Grenvillian

orogen make it clear that the presence of high-P

rocks at the surface is directly related to the

existence of the lateral ramp. Hynes and Eaton

(1999) interpreted this relationship as evidence

that the lateral ramp, because it gave rise to a

concavity in the lower surface of the orogenic

wedge, had produced a channel into which

material flowing from the deeper to the shallower

parts of the orogenic wedge was preferentially

concentrated (Fig. 7).

The ability of a concavity on the lower

confining surface (or a convex-upward irregu-

larity on the upper surface) of a flowing medium

to channel flow is not easy to demonstrate

rigorously, because it requires solution of the

FIG. 7. Schematic illustration of the concentration of flow in an orogenic wedge into the concavity produced by a

lateral ramp. The small arrows indicate the general direction of flow; the large arrows illustrate the increased flow

within and towards the region of the lateral ramp.

h

h3

h

h

h×1.25

due to and/or P��

due to shear

FIG. 8. General forms of the flow profiles between two planar, parallel plates, when the upper plate is moved parallel

to the lower. The flow velocity at each point consists of a component due to the shear of the overlying plate (to the

right) and a component due to changes in pressure or gravitational potential (to the left). The net flow due to the

combination is stippled. For a 25% increase in the distance, h, between the plates, illustrative of the flow within a

channel, there is a marked increase in the flow due to changes in pressure/gravitational potential.

16

A.HYNES

Navier-Stokes equation in three dimensions, but

indications of its probable effectiveness can be

derived from two-dimensional considerations.

Consider a Newtonian viscous fluid bounded

below by a stationary and rigid surface in which

there is concavity representing a channel, and

bounded above by a rigid surface moving at a

fixed rate relative to the lower surface (Fig. 8).

Flow of the viscous fluid is driven by the shear

stress exerted by the upper surface, pressure

differences along the path of the fluid, and the

buoyancy of the fluid. Flow outside the channel

may be approximated by flow between parallel

surfaces, which is readily determined using the

stream function (e.g. Turcotte and Schubert,

1982). Flow within the channel may be approxi-

mated likewise, but with a greater distance

between the upper and lower surfaces. Flow

between two surfaces due to the shear stress

from motion of the upper surface is proportional

to the distance between the surfaces, but flow due

to pressure differences along the path or buoyancy

is proportional to the cube of the distance between

the surfaces. There is, therefore, a dispropor-

tionate increase in flow within the channel

compared with that outside it (Fig. 8). A

channel should consequently concentrate flow

markedly if the flow is driven by pressure

gradients or buoyancy forces. It is therefore

clear in principle why a lateral ramp in the

subsurface might lead to a concentration of

exhumed high-P rocks in the region above the

ramp, as is observed in the Grenvillian orogen.

Recorded pressures in the high-P rocks of the

Grenvillian orogen do not exceed 1800 MPa,

which places them potentially still within

thickened continental crust. There is, furthermore,

no evidence to suggest that burial of these rocks

resulted from subduction of the leading edge of a

continent. It appears that the structural setting of

the preserved Grenvillian orogen was on the

northwestern edge of a southeast-facing Andean

margin (Rivers, 1997). The thrusting by which

rocks were deeply buried therefore forms part of a

retro-shear system, rather than a pro-shear system,

in the sense of Beaumont and Quinlan (1994), and

100 km

APENNINES

ADRIATIC

Po Plain

Jura

MolasseForedeep

Perialpine basins

Oligocene plutons

Ophiolitic units

Flysch

Austroalpinelow-grade

Austroalpinemedium-grade

Cret. metamorphism,W Alps; blueschist

Cret. metamorphism,W. Alps; eclogitic

DM

SL

MR

GP

A

FIG. 9. Tectonic map of the Alps, after Polino et al. (1990). A: Adula nappe; DM: Dora Maira nappe; GP: Gran

Paradiso nappe; MR: Monte Rosa nappe; SL: Sesia Lanzo nappe.


17

their probable correlatives, metamorphosed to

much lower pressures, structurally overlie them.

They have therefore been extruded within a

region of continental crustal thickening, rather

than in association with continental subduction.

Their setting is in this sense similar to the Early

Miocene extrusion, from mid-crustal depths, of

the Greater Himalayan sequence (Hodges et al.,

1996). The general theoretical considerations

discussed above, however, indicate that lateral

ramps might also be important in controlling the

locations of high-P rocks that had been extruded

from mantle depths. Some of the better-known

localities of high-P metamorphosed rocks world-

wide provide support for this suggestion.

The classic high-P locality of the Western Alps,

the Dora Maira massif, occurs at the western end

of the Alpine chain, in a region in which orogenic

trends undergo a pronounced leftward jog

(Fig. 9). Indeed, the eclogite-rich terranes of the

Monte-Rosa, Sesia Lanzo, Gran Paradiso and

Dora Maira nappes occupy a broad lobate region

around this leftward jog that is markedly similar

in form to that of the high-P rocks of the

Grenvillian orogen.

In the Aegean, the trends of the two zones of

high-P rocks, formed 45 and 25 Ma ago,

respectively (Jolivet et al., 1996), are at a high

angle to equivalent trends for the Hellenides, again

associated with a leftward jog of the trends, and

the stretching lineations that presumably illustrate

their transport directions are oblique to regional

orogenic trends (Fig. 10). Although the present

configuration of the Hellenic arcs is markedly

affected by subsequent extension in the Aegean

region, the Neogene extension resulted in little net

rotation of Crete (Angelier et al., 1982), and

restoration of the clockwise rotation of the main

Hellenic peninsula still leaves a substantial

leftward jog of the arc in the Early Miocene

(Kissel and Laj, 1988). Thus, emplacement of at

least the later of the two sets of high-P rocks

appears to have occurred at an offset along the

tectonic trend and in a direction subparallel to it.

In southern China, the high-P terranes in the

contact zone between the Sino-Korean and

Yangtze cratons, in the Hong’an, Dabie Shan

and Su-Lu regions (Hacker and Wang, 1995) lie

along and near a prominent leftward offset of the

zone associated with the Tan-Lu fault (Fig. 11),

Aegean Sea

Ionian Sea

Helle

nid

es

lineations

Hellenide tectonictransport direction

100 km

41°N

38°N

35°N23°E 28°E

N

HP

45M

a

HP

25M

a Crete

FIG. 10. Distribution of high-P metamorphic belts (heavily dotted regions) in the Aegean region, and stretching

lineations associated with them, after Jolivet et al. (1996).

18

A.HYNES

which is believed to have been active as a left-

lateral fault in the Triassic, when the high-P rocks

were emplaced. Yin and Nie (1993) interpret this

fault as due to irregularities in the nature of the

pre-Triassic form of the northern margin of the

Yangtze block, whereas Okay et al. (1993)

consider the fault to have developed and offset

the margin of the block only in the Triassic. In

either case, it could have been the site of a side

ramp, and direct support for eastward thrusting

associated with it is evident in the eastward-

verging folds in the sedimentary rocks to the east

of the Dabie Shan Complex (Okay et al., 1993). A

recent interpretation of the dynamics of the region

(Hacker et al., 2000) incorporates eastward

extrusion of the ultrahigh-P rocks of the Dabie

region along the plate margin, towards an eastern

reentrant in the margin, with the re-entrant

channelling the extruding lithosphere much in

the manner argued above.

In the Norwegian Caledonides, the high-P

rocks of the Western Gneiss Region are also

associated with a marked change in direction of

the frontal thrusts of the orogen from their

regional NNE trend to a more easterly trend

(Fig. 12). In this case, the orogenic trends appear

to have jogged towards the right.

Transverse structures are not in and of

themselves evidence of the presence of subsurface

lateral ramps. The transverse structures associated

with the western Himalayan syntaxis, for

example, have been interpreted to indicate

simply unusual amounts of forelandward transport

due to the presence of weak layers in the orogenic

wedge (Burbank, 1983). In general, however,

broad changes in the directions of orogenic fronts

are interpreted to indicate irregularities in the pre-

collisional geometries of the colliding margins

(e.g. Thomas, 1977) and such irregularities will

tend to give rise to lateral ramps. The association

of all these high-P localities with large-scale

transverse structures is therefore highly sugges-

tive of a role for lateral ramps in their effective

transport to the surface.

Summary and conclusions

In summary, there is evidence from the Himalayas

that, even in the absence of a large buoyancy

drive, upward extrusion of major rock units from

deep in the orogenic wedge is an important

process. In circumstances in which continental

crust has been partially subducted into the mantle,

the buoyancy drive of the light continental crust

Pac

ific

140°E125°110°

095°

080°

20°N

20°N

50°N

600 km

Seaof

JapanSino-Koreancraton

Yangtzecraton

S

H

D

FIG. 11. Tectonic scheme for southern China in Triassic time, adapted from Yin and Nie (1993). Heavily stippled

regions are localities of high-P rocks extruded in the collisional zone between the Sino-Korean and Yangtze cratons,

after Ernst and Liou (1995), labelled H (Hong’an), D (Dabie Shan) and S (Su Tun).


19

provides a major boost to the extrusion process,

and this boost may be further enhanced by

adjustments in the geometry of the descending

slab due to arrival of the cool continental margin,

or by reduction of slab pull with detachment of

the oceanic part of the descending slab. Direct

evidence for the sensitivity of the wedge

geometry to the dip of the descending slab is

rare, but is supplied in one instance by the

synchronicity of slab shallowing with wedge

FIG. 12. Terrane map of the Scandinavian Caledonides, adapted from Stephens and Gee (1989), with eclogite

localities in the Western Gneiss region after Coleman and Wang (1995).

20

A.HYNES

inversion in the Franciscan. Extrusion may, then,

be the rule rather than the exception in

circumstances in which continental crust arrives

at a subduction zone on the lower plate and is

dragged partially into the mantle.

The close spatial relationship between a lateral

ramp and a broad lobe of exhumed high-P rocks

in the Grenvillian orogen indicates that a lateral

ramp may have played a role in the exhumation

process there. Many other high-P and ultrahigh-P

rocks are associated with jogs in orogenic trends

that may be indicative of lateral ramps. If material

transport is restricted to two dimensions, partially

subducted continental crust may be expected to be

extruded along paths similar to those along which

it was buried. Lateral ramps may, however, cause

substantial deviations from two-dimensional flow,

with concentration of the extrusion in the regions

of the lateral ramps themselves. In these

circumstances, high-P and ultrahigh-P rocks

may be expected to occur preferentially at right-

handed and left-handed jogs in orogenic trends

(Fig. 13). Confirmation of the importance of flow

patterns of the kind envisaged here could be

derived from regional studies of the spatial and

temporal variations in attitude of the structures

associated with extruded high-P complexes.

Acknowledgements

My work on the emplacement of high-P rocks

was sparked originally by stimulating discussions

with Reinhard Greiling and Zvi Garfunkel. It has

been supported throughout by grants from the

Natural Sciences and Engineering Research

Council of Canada, through its Operating Grants

and Lithoprobe Supporting Geoscience

Programmes. I am particularly grateful to the

organizers of and participants in the Metamorphic

Studies Group Meeting in Rennes for the

opportunity to air my thoughts further and for

many exciting discussions. Constructive

comments on the manuscript by Alexandre

Chemenda were much appreciated.

References

Andersen, T.B., Jamtveit, B., Dewey, J.F. and

Swensson, E. (1991) Subduction and eduction of

continental crust: major mechanisms during con-

tinent-continent collision and orogenic extensional

collapse, a model based on the south Norwegian

Caledonides. Terra Nova, 3, 303�310.

Angelier, J., Lyberis, N., Le Pichon, X., Barrier, E. and

Huchon, P. (1982) The tectonic development of the

Hellenic arc and the Sea of Crete: a synthesis.

Tectonophysics, 86, 159�196.

Beaumont, C. and Quinlan, G. (1994) A geodynamic

framework for interpreting crustal-scale seismic-

reflectivity patterns in compressional orogens.

Geophysical Journal International, 116, 754�783.

Borghi, A., Cadoppi, P., Porro, A., Sacchi, R. and

Sandrone, R. (1984) Osservazioni geologiche nella

Val Germanasca et nella media Val Chisone (Alpi

Cozie) [Geological studies in the Germanasca Valley

and the middle Chisone Valley, Cottian Alps].

Museo Regionale di Scienze Naturali di Torino,

Bollettino, 2, 503�530.

Burbank, D.W. (1983) The chronology of intermontane-

basin development in the northwestern Himalaya and

the evolution of the Northwest Syntaxis. Earth

Planetary Science Letters, 64, 77�92.

Burbank, D.W. (2002) Rates of erosion and their

implications for exhumation. Mineralogical

Magazine, 66, 25�52.

Burbank, D.W., Leland, J., Fielding, E., Anderson, R.S.,

regionalvergence

regionalvergence

vergenceat depth

vergenceat depth

vergenceat surface

vergenceat surface

subduction

subductionHigh- rocksP

High- rocksP

FIG. 13. Regions in which high-P and ultrahigh-P rocks

may concentrate, in relation to orogenic trends and the

directions of historical subduction and regional tectonic

vergence.


21

Brozovic, N., Reid, M. and Duncan, C. (1996)

Bedrock incision, rock uplift and threshold hillslopes

in the northwestern Himalaya. Nature, 379,

505�510.

Byerlee, J. (1978) Friction of rocks. Pure and Applied

Geophysics, 116, 615�626.

Chemenda, A.I., Mattauer, M., Malavieille, J. and

Bokum, A.N. (1995) A mechanism of syn-collisional

rock exhumation and associated normal faulting:

results from physical modelling. Earth and


Chemenda, A.I., Mattauer, M. and Bokun, A.N. (1996)

Continental subduction and a mechanism for

exhumation of high-pressure metamorphic rocks:

new modelling and field data from Oman. Earth and


Chopin, C. (1984) Coesite and pure pyrope in high-

grade blueschists of the Western Alps: a first record

and some consequences. Contributions to

Mineralogy and Petrology, 86, 107�118.

Chopin, C., Henry, C. and Michard, A. (1991) Geology

and petrology of the coesite-bearing terrain, Dora

Maira massif, Western Alps. European Journal of

Mineralogy, 3, 263�291.

Chopra, P.N. and Paterson, M.S. (1984) The role of

water in the deformation of dunite. Journal of

Geophysical Research, 89, 7861�7876.

Cloos, M. (1982) Flow melanges: numerical modelling

and geologic constraints on their origin in the

Franciscan subduction complex, California.

Geological Society of America Bulletin, 93,

330�345.

Cloos, M. and Shreve, R.L. (1988) Subduction-channel

model of prism accretion, melange formation,

sediment subduction, and subduction erosion at

convergent plate margins: 1. Background and

description. Pure and Applied Geophysics, 128,

455�500.

Coleman, R.G. and Wang, X. (1995) Overview of the

geology and tectonics of UHPM. Pp. 1�32 in:

Ultrahigh Pressure Metamorphism (R. G. Coleman

and X. Wang, editors). Cambridge University Press,

Cambridge, UK.

Compagnoni, R., Dal Piaz, G.V., Hunziker, J.C., Gosso,

G., Lombardo, B. and Williams, P.F. (1977) The

Sesia-Lanzo zone, a slice of continental crust with

Alpine high-pressure low-temperature assemblages

in the western Italian Alps. Rendiconti della Societa

Italiana di Mineralogia e Petrologia, 33, 281�334.

Coney, P.J. and Reynolds, S.J. (1977) Cordilleran

Benioff zones. Nature, 270, 403�406.

Cowan, D.S. and Silling, R.M. (1978) A dynamic, scaled

model of accretion at trenches and its implications

for the tectonic evolution of subduction complexes.

Journal of Geophysical Research, 83, 5389�5396.

Davies, J.H. and von Blanckenburg, F. (1995) Slab

breakoff: a model of lithosphere detachment and its

test in the magmatism and deformation of collisional

orogens. Earth and Planetary Science Letters, 129,

85�102.

Davy, P. and Gillet, P. (1986) The stacking of thrust

slices in collisional zones and its thermal con-

sequences. Tectonics, 5, 913�929.

de Sigoyer, J., Guillot, S., Lardeaux, J.M. and Mascle,

G. (1997) Glaucophane-bearing eclogites in the Tso

Morari dome (eastern Ladakh, NW Himalaya).

European Journal of Mineralogy, 9, 1073�1083.

de Sigoyer, J., Chavagnac, V., Blichert-Toft, J., Villa,

I.M., Luais, B., Guillot, S., Cosca, M. and Mascle, G.

(2000) Dating the Indian continental subduction and

collisional thickening in the northwest Himalaya:

multichronology of the Tso Morari eclogites.

Geology, 28, 487�490.

Dickinson, W.R. (1973) Reconstruction of past arc-

trench systems from petrotectonic assemblages in the

island arcs of the western Pacific. Pp. 569�601 in:

The Western Pacific, Island Arcs, Marginal Seas,

Geochemistry (P.J. Coleman, editor). University of

Western Australia Press, Perth, Australia.

England, P.C. and Richardson, S.W. (1977) The

influence of erosion upon the mineral facies of rocks

from different metamorphic environments. Journal

of the Geological Society of London, 134, 201�213.

England, P.C. and Thompson, A.B. (1984) Pressure-

temperature-time paths of regional metamorphism.

1. Heat transfer during the evolution of thickened

crust. Journal of Petrology, 25, 894�928.

Ernst, W.G. (1973) Blueschist metamorphism and P-T

regimes in active subduction zones. Tectonophysics,

17, 255�272.

Ernst, W.G. (1975) Systematics of large-scale tectonics

and age progressions in Alpine and circum-Pacific

blueschist belts. Tectonophysics, 26, 229�246.

Ernst, W.G. (1984) Californian blueschists, subduction,

and the significance of tectonostratigraphic terranes.

Geology, 12, 436�440.

Ernst, W.G. and Liou, J.G. (1995) Contrasting plate-

tectonic styles of the Qinling-Dabie-Sulu and

Franciscan metamorphic belts. Geology, 23,

353�356.

Godfrey, N.J. and Klemperer, S.L. (1998) Ophiolitic

basement to a forearc basin and implications for

continental growth: the Coast Range/Great Valley

ophiolite, California. Tectonics, 17, 558�570.

Gorbatschev, R. (1985) Precambrian basement of the

Scandinavian Caledonides. Pp. 197�212 in: The

Caledonide Orogen � Scandinavia and Related

Areas (D.G. Gee and B.A. Sturt, editors). John

Wiley, New York.

Griffin, W.L., Austrheim, H., Brastad, K., Brynhi, I.,

Krill, A.G., Krogh, E.J., Mork, M.B.E., Qvale, H.

and Torudbakken, B. (1985) High-pressure meta-

22

A.HYNES

morphism in the Scandinavian Caledonides. Pp.

783�902 in: The Caledonian Orogen �Scandinavia and Related Areas (D.G. Gee and

B.A. Sturt, editors). John Wiley, Chichester, UK.

Hacker, B.R. and Peacock, S.M. (1995) Creation,

preservation, and exhumation of UHPM rocks. Pp.

159�181 in: Ultrahigh Pressure Metamorphism

(B.R. Hacker and J.G. Liou, editors). Cambridge

University Press, Cambridge, UK.

Hacker, B.R. and Wang, Q. (1995) Ar/Ar geochronology

of ultrahigh-pressure metamorphism in central

China. Tectonics, 14, 994�1006.

Hacker, B.R., Ratschbacher, L., Webb, L., McWilliams,

M.O., Ireland, T., Calvert, A., Dong, S., Wenk, H.R.

and Chateigner, D. (2000) Exhumation of ultrahigh-

pressure continental crust in east central China: Late

Triassic-Early Jurassic tectonic unroofing. Journal of


Harms, T., Jayko, A.S. and Blake, M.C. (1992)

Kinematic evidence for extensional unroofing of

the Franciscan complex along the Coast Range fault,

northern Diablo Range, California. Tectonics, 11,

228�241.

Hauck, M.L., Nelson, K.D., Brown, L.D., Zhao, W. and

Ross, A.R. (1998) Crustal structure of the Himalayan

orogen at ~908 east longitude from Project

INDEPTH deep reflection profiles. Tectonics, 17,

481�500.

Hetzel, R., Echtler, H.P., Seifert, W., Schulte, B.A. and

Ivanov, K.S. (1998) Subduction- and exhumation-

related fabrics in the Paleozoic high-pressure-low-

temperature Maksyutov Complex, Antingan area,

southern Urals, Russia. Geological Society of

America Bulletin, 110, 916�930.

Hodges, K.V., Parrish, R.R., Housch, T., Lux, D.,

Burchfiel, B.C., Royden, L. and Chen, Z. (1992)

Simultaneous Miocene extension and shortening in

the Himalayan Orogen. Science, 258, 1466�1470.

Hodges, K.V., Parrish, R.R. and Searle, M.P. (1996)

Tectonic evolution of the central Annapurna Range,

Nepalese Himalayas. Tectonics, 15, 1264�1291.

Houseman, G.A., McKenzie, D.P. and Molnar, P. (1981)

Convective instability of a thickened boundary layer

and its relevance for the thermal evolution of

continental convergence belts. Journal of


Hsu, K.J. (1991) Exhumation of high-pressure meta-

morphic rocks. Geology, 19, 107�110.

Hurich, C.A. (1996) Kinematic evolution of the lower

plate during intracontinental subduction: an example

from the Scandinavian Caledonides. Tectonics, 15,

1248�1263.

Hynes, A. and Eaton, D. (1999) Lateral ramps as an aid

to the unroofing of deep-crustal rocks: seismic

evidence from the Grenville province. Tectonics,

18, 343�360.

Hynes, A., Arkani-Hamed, J. and Greiling, R. (1996)

Subduction of continental margins and the uplift of

high-pressure metamorphic rocks. Earth and


Hynes, A., Indares, A., Rivers, R. and Gobeil, A. (2000)

Lithoprobe line 55: integration of out-of-plane

seismic results with surface structure, metamorph-

ism, and geochronology, and the tectonic evolution

of the eastern Grenville Province. Canadian Journal

of Earth Sciences, 37, 341�358.

Indares, A., Dunning, G., Cox, R., Gale, D. and

Connelly, J. (1998) High-pressure, high-temperature

rocks from the base of thick continental crust:

geology and age constraints from the Manicouagan

Imbricate Zone, eastern Grenville province.

Tectonics, 17, 426�440.

Indares, A., Dunning, G. and Cox, R. (2000) Tectono-

thermal evolution of deep crust in a Mesoproterozoic

continental collision setting: the Manicouagan

example. Canadian Journal of Earth Sciences, 37,

325�340.

Jamieson, R.A. and Beaumont, C. (1988) Orogeny and

metamorphism: a model for deformation and

pressure-temperature-time paths with applications

to the central and southern Appalachians.

Tectonics, 7, 417�445.

Jolivet, L., Goffe, B., Monie, P., Truffert, C., Patriat, M.

and Bonneau, M. (1996) Miocene detachment in

Crete and exhumation P-T-t paths of high-pressure

metamorphic rocks. Tectonics, 15, 1129�1153.

Kissel, C. and Laj, C. (1988) The Tertiary geodynamical

evolution of the Aegean arc: a paleomagnetic

reconstruction. Tectonophysics, 146, 183�201.

Maruyama, S., Liou, J.G. and Zhang, R. (1994) Tectonic

evolution of the ultrahigh-pressure (UHP) and high-

pressure (HP) metamorphic belts from central China.

Island Arc, 3, 112�121.

Matte, P. (1998) Continental subduction and exhumation

of HP rocks in Paleozoic orogenic belts: Uralides

and Variscides. GFF, 120, 209�222.

Molnar, P. and Tapponnier, P. (1975) Cenozoic

tectonics of Asia: effects of a continental collision.

Science, 189, 419�426.

Okay, A.I., Sengor, A.M.C. and Satir, M. (1993)

Tectonics of an ultrahigh-pressure metamorphic

terrane: the Dabie Shan/Tongbai Shan orogen,

China. Tectonics, 12, 1320�1334.

Okay, A.I., Shutong, X. and Sengor, A.M.C. (1989)

Coesite from the Dabie Shan eclogites, central

China. European Journal of Mineralogy, 1,

595�598.

Oxburgh, E.R. and Turcotte, D.L. (1974) Thermal

gradients and regional metamorphism in overthrust

terrains with special reference to the Eastern Alps.

Schweizerische Mineralogische und Petrographische

Mitteilungen, 54, 641�662.


23

Peacock, S.M. (1990) Numerical simulation of meta-

morphic pressure-temperature-time paths and fluid

production in subducting slabs. Tectonics, 9,

1197�1211.

Platt, J.P. (1986) Dynamics of orogenic wedges and the

uplift of high-pressure metamorphic rocks.

Geological Society of America Bulletin, 97,

1037�1053.

Platt, J.P. (1987) The uplift of high-pressure-low-

temperature metamorphic rocks. Philosophical

Transactions of the Royal Society Section A, 321,

87�103.

Platt, J.P. (1993) Exhumation of high-pressure rocks: a

review of concepts and processes. Terra Nova, 5,

119�133.

Polino, R., Dal Piaz, G.V. and Gosso, G. (1990)

Tectonic erosion at the Adria margin and accre-

tionary processes for the Cretaceous orogeny of the

Alps. Geological Society of France Memoirs, 156,

345�367.

Rivers, T. (1997) Lithotectonic elements of the

Grenville province: review and tectonic implications.

Precambrian Research, 86, 117�154.

Seidel, E., Kreuzer, H. and Harre, W. (1982) The late

Oligocene/early Miocene high pressure in the

external Hellenides. Geologisches Jahrbuch

Hessen, E23, 165�206.

Shelton, G.L. and Tullis, J.A. (1981) Experimental flow

laws for crustal rocks. Transactions of the American

Geophysical Union, 61, 376.

Stephens, M.B. and Gee, D.G. (1989) Terranes and

polyphase accretionary history in the Scandinavian

Caledonides. Geological Society of America Special

Publication, 230, 17�30.

Thomas, W.A. (1977) Evolution of Appalachian-

Ouachita salients and recesses from reentrants and

promontories in the continental margin. American

Journal of Science, 277, 1233�1278.

Turcotte, D. and Schubert, G. (1982) Geodynamics.

Wiley, New York.

van den Beukel, J. (1992) Some thermomechanical

aspects of the subduction of continental lithosphere.

Tectonics, 11, 316�329.

von Blanckenburg, F. and Davies, J.H. (1995) Slab

breakoff: a model for syncollisional magmatism and

tectonics in the Alps. Tectonics, 14, 120�131.

Wheeler, J. (1991) Structural evolution of a subducted

continental sliver: the northern Dora Maira massif,

Italian Alps. Journal of the Geological Society of

London, 148, 1101�1113.

Wijbrans, J.R., van Wees, J.D., Stephens, R.A. and

Cloeting, S.A. (1993) Pressure-temperature-time

evolution of the high-pressure metamorphic complex

of Sifnos, Greece. Geology, 21, 443�446.

Wong, A., Ton, S.Y.M. and Wortel, M.J.R. (1997) Slab

detachment in continental collision zones: an

analysis of controlling parameters. Geophysical

Research Letters, 24, 2095�2098.

Yin, A. and Nie, S. (1993) An indentation model for the

North and South China collision and the develop-

ment of the Tan-Lu and Honam fault systems,

eastern Asia. Tectonics, 12, 801�813.

[Manuscript received 20 December 2000:

revised 28 February 2001]

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