Impact of arc-continent collision on the conditions of ... Page/Paper Oman, 2001.pdf · exhumation of UHP/LT rocks: experimental and numerical modelling A.I. Chemenda*, D. Hurpin,

Impact of arc-continent collision on the conditions of burial and

exhumation of UHP/LT rocks: experimental

and numerical modelling

A.I. Chemenda*, D. Hurpin, J.-C. Tang, J.-F. Stephan, G. Buffet

Geosciences Azur, UMR 6526, Universite de Nice-Sophia Antipolis and CNRS, 250 Rue Albert Einstein-Sophia Antipolis,

06560 Valbonne, France

Received 14 August 2000; accepted 3 December 2000

Abstract

A 2-D physical and finite-element numerical modelling of arc continent collision was performed to study the deformation

and failure of the overriding lithosphere. The experimental technique allowed us to model the whole subduction/collision

process from oceanic subduction to deep subduction of the continental crust. With the numerical approach we have modelled

the deformation of the overriding plate only through initial stages of its failure and studied the influence of different

parameters on this process. The results obtained by both techniques are coherent and mutually complementary. They show

that the failure of the overriding plate is physically quite plausible or even inevitable during subduction. The conditions for

such a failure (the weakening of this plate) are prepared during oceanic subduction. The weakening occurs due to the

interaction between the subducting lithosphere and the asthenosphere in the mantle corner between the two plates, and due to

back-arc spreading. In oceanic subduction zones with a compressional regime (no back-arc opening, thick and strong back-

arc lithosphere), the weakest zone is volcanic arc area. When weakening becomes sufficient during subduction, the

lithosphere fails in this area. The failure occurs along a fault dipping under the arc in either of two possible directions and

results either in subduction reversal or subduction of the fore-arc. Almost half of the presently active subduction zones are

characterised by a tensional subduction regime with back-arc spreading. In such subduction zones, the weakest place is not

the volcanic arc but the back-arc spreading centre. When a subduction regime changes from tensional to compressional,

failure occurs in the vicinity of the extinct spreading centre. This process can occur during oceanic subduction again along

either a trench-vergent or trenchward-dipping fault, but the formation of a trench-verging fault is most likely. In this latter

case, which is a principal subject of our study, the failure is followed by partial subduction of the arc plate. Complete

subduction occurs during arc-continent collision (subduction of the continental margin) when tectonic compression of the

lithosphere increases rapidly and becomes sufficient to push the arc plate into the mantle. The arc itself can be subducted

completely or be partially or completely scraped-off and accreted. A deeply subducted material (including continental

margin) is preserved at relatively low temperatures between the lithospheric mantle and the ‘‘cold’’ subducted arc plate to

about 150-km depth. Subduction of the arc plate is a major phenomenon, which affects all processes associated with

continental subduction from deep burial and HP/LT metamorphism to exhumation of subducted material. Does this process

occur in nature? Future investigations will allow us to answer this question. In this paper, we analyse the conditions of

0040-1951/01/$ - see front matter D 2001 Elsevier Science B.V. All rights reserved.

PII: S0040-1951 (01 )00160 -3

* Corresponding author. Tel.: +33-4-9294-2661; fax: +33-4-9264-2610.

E-mail address: [email protected] (A.I. Chemenda).

www.elsevier.com/locate/tecto

Tectonophysics 342 (2001) 137–161

emplacement of a very young oceanic lithosphere (Samail ophiolite) on the continental crust in Oman in the late Cretaceous

and argue that this lithosphere formed in a back-arc basin. It reached and overthrust the Arabian continent after complete

subduction of the arc plate. D 2001 Elsevier Science B.V. All rights reserved.

Keywords: Arc-continent collision; Exhumation; UHP/LT rocks; Analogue modelling; Numerical modelling; Oman

1. Introduction

Three principal interrelated geodynamic problems

regarding the ultra-high pressure/low temperature

(UHP/LT) terrains concern the mechanisms of deep

burial, preservation at low temperature, and then

exhumation of UHP rocks. Everyone agrees that the

burial is due to subduction; the remaining question is

how the low-density, low-strength (at depth) conti-

nental crust can be dragged (pushed) to more than

100-km depth and be preserved at a relatively low

(� 700�C) temperature. There is a consensus that the

principal driving force for the exhumation of UHP

rocks is the buoyancy of the subducted crust that

keeps its lower (with respect to the mantle) bulk

density even at great depth. Under debate are the

deformation style and the volume of the rising crustal/

sedimentary material; in particular, the question is

whether this process occurs on the crustal scale or

whether only thin slices of continental material can be

delivered to shallow depths. Burial and exhumation

are elements of the same process of continental

subduction and should be integrated into a coherent

model. Modelling of this process encounters major

difficulties for two reasons. First, continental subduc-

tion is a complex process that includes the interaction,

deformation and failure of media with different rheol-

ogies. Second, this interaction strongly depends on

various ill-constrained parameters and processes,

including: the rheology of the crust and its change

during subduction with temperature and pressure,

mineralogical transformations, variations in water

content, etc. Therefore, exhaustive direct modelling

of continental subduction/exhumation is difficult.

Advances in understanding of this process can be

achieved by testing models of increasing complexity

against the data appropriate to this complexity. It is

difficult, for example, to study the details of crustal

deformation at depth without knowing the kinetics

and spatial distribution of mineralogical transforma-

tion of the crust. These transformations depend in

particular on the temperature distribution and on the

kinematics of continental subduction which, we try to

show in this paper, may both considerably differ from

the traditionally assumed schemes. We first discuss

the thermal structure of subduction zones and partic-

ularly of the overriding plate during subduction of the

oceanic lithosphere. Using both experimental and

numerical modelling, we show that during subduction

of the continental margin this plate may fail and its

frontal part (fore-arc block or arc plate) can be

subducted into the mantle along with the continental

margin. Subduction of these lithospheric units repre-

sents a major (kinematic, dynamic, thermal, etc.)

difference of continental subduction from oceanic

subduction. We then apply one of the obtained sub-

duction scenarios to the evolution of Oman and show

that the disappearance (subduction) of the whole arc

plate is a plausible process.

2. Constraints on thermal regime of subduction

The fact that UHP rocks register low peak temper-

atures (600–800 �C at depth 100–140 km) is tradi-

tionally considered as a consequence of subduction.

Thermal models of continental subduction (e.g. Van

den Beukel, 1992; Davies and von Blanckenburg,

1994) do yield such low temperatures at these depths.

These models (as well as all others) are based on a

number of assumptions and simplifications that define

the solutions. They do not consider, for example, the

secondary (induced) convection in the mantle corner

in the subduction zone. There is a large variety of

thermal models of oceanic subduction that include

induced convection, which has been shown to be an

important factor in increasing the temperature in the

mantle corner (e.g. Furukawa, 1993; Peacock, 1996;

Kincaid and Sacks, 1997). According to petrologic

data, the temperature in this corner at about 100 km-

depth under the volcanic arc is more than 1300 �C(Fig. 1). This is consistent with geothermal (Furu-

A.I. Chemenda et al. / Tectonophysics 342 (2001) 137–161138

kawa, 1993; Lewis et al., 1988) and seismic (Fig. 2)

data showing a low-velocity zone and strong thinning

of the overriding lithosphere under the arc. Subduc-

tion of ‘‘cold’’ oceanic lithosphere thus not only does

not reduce the temperature in the mantle corner, but

on the contrary, increases it and causes strong ‘‘ther-

mal erosion’’ (partial melting?) of the overriding

lithosphere. The thickness of this lithosphere near

Fig. 1. Model for the formation of volcanic arc in subduction zones based on petrological data (simplified after Schmidt and Poli, 1995).

Fig. 2. P wave velocity structure of NE Japan subduction zone (from Zhao et al., 1994).

A.I. Chemenda et al. / Tectonophysics 342 (2001) 137–161 139

the interplate zone can hardly exceed several tens of

kilometres and under the arc a few tens of kilometres

(see for example Fig. 2). The oceanic subduction zone

thus appears as a hot zone, although the chemical,

dynamic and thermal interaction of the subducting

lithosphere (including hydrated crust and sediments)

and the surrounding mantle remains unclear. What

happens when continental crust starts to subduct into

this hot zone? The temperature should increase even

more due to the reduction of subduction rate usually

associated with the beginning of continental subduc-

tion and the decay of radioactive elements in the

continental crust. Thus, the means by which deeply

subducted continental crust is preserved at relatively

low temperatures is not obvious.

Subduction of the continental margin has another,

mechanical consequence: an increase in horizontal

tectonic (non-hydrostatic) compression of the over-

riding lithosphere (Shemenda, 1994). Since this

lithosphere is weakened, it can fail in the arc area,

resulting either in subduction reversal or subduction

of the fore-arc lithosphere (Chemenda et al., 1997;

see also Figs. 5 and 6). The latter process should

strongly affect the thermal (hence, mechanical)

regime of the subducting continental crust: it will

be colder and stronger than without the thermal

shield (fore-arc block). If subduction of a fore arc

block actually occurs in nature, it would strongly

affect all processes associated with burial and exhu-

mation of the continental crust. There is evidence of

complete or partial subduction of this block in the

Urals and the Variscan belt (Matte, 1998), the

Kamchatka (Konstantinivskaya, 2000), Taiwan

(Chemenda et al., 1997, 2001; Malavieille, 1999;

Tang and Chemenda, 2000) and the Himalayas

(Harrison et al., 1992; Anczkiewicz et al., 1998).

As was stated above, fore-arc subduction may

occur due to the presence of hot and weak lithosphere

in the volcanic arc. The volcanic arc is the weakest

place in many subduction zones except those in a

tensional subduction regime with an active back-arc

basin. The lithosphere in the back-arc spreading centre

should be still thinner and weaker than in the arc.

When the tensional regime changes to a compres-

sional regime (due to continental margin subduction,

for example), the overriding plate should fail not in

the volcanic arc but near the back-arc spreading

centre. What are the possible modes of this failure?

How will continental subduction continue after the

failure? Below we address these questions by both

experimental and numerical modelling.

3. Experimental modelling

3.1. Set-up

Fig. 3 shows a scheme of the experimental model,

which includes a one-layer overriding oceanic litho-

sphere thinned in the arc area and back-arc spreading

centre. The subducting plate comprises a one-layer

oceanic lithosphere and a three-layer continental litho-

sphere. All the lithospheric layers possess plastic

properties and are made of hydrocarbon composi-

tional systems. The upper continental crust, the con-

tinental lithospheric mantle and the oceanic

Fig. 3. Scheme of the experimental model; 1 = oceanic overriding lithosphere; 2 = oceanic segment of the subducting lithosphere; 3 = plastic

upper continental crust with strong strain weakening; 4 = ductile, very weak lower crust; 5 = plastic continental lithospheric mantle; 6 = piston;

7 = liquid low-viscosity asthenosphere. Lb is the back-arc spreading centre/trench distance; La is the arc/trench distance.


lithosphere have the same yield limit and are charac-

terised by a strong strain weakening (Fig. 4). The

lower continental crust is considerably weaker (see

Table 1) and more ductile. The lithosphere is under-

lain by a low-viscosity asthenosphere, which in the

experiments is pure water. Convergence is driven by a

piston moving at a constant rate throughout the

experiment. The similarity criteria met by this mod-

elling are the same as those of Chemenda et al. (1995)

who presented similar experiments, but without weak-

ening of the upper plate.

3.2. Results

First, we present two experiments (Figs. 5 and 6)

where the arrival of the continental margin at the

subduction zone is preceded by oceanic subduction

in a compressional regime (no lithosphere weakening

in the back-arc). There is only one difference

between these experiments: in experiment 1, the

continental crust has two layers as shown in Fig.

3, and in experiment 2, the whole crust is made of

the material corresponding to the upper crust in

experiment 1.

Experiment 1 (Fig. 5). During oceanic subduction

(Fig. 5a), the overriding plate experiences compres-

sion, but it is not sufficient to cause its failure. During

subduction of the continental margin, the compression

increases and the overriding plate fails in the arc along

a continent-vergent fault (Fig. 5d). The failure is

followed by the complete subduction of the fore arc

block (Fig. 5e–h).

Fig. 4. Stress–strain diagrams for the experimental and numerical lithosphere models. The experimental curve corresponds to the oceanic

lithosphere, continental mantle and upper crust (see Fig. 3 and Table 1). For the numerical model, the yield limit for the normal stress ss is:ss = 1.8� 108 Pa, Young’s modulus E= 2� 1010 Pa; stress drop during the failure Ds= 3.6� 107 Pa; strain softening parameter k= 0.3;Poisson’s ratio n= 0.25.

Table 1

Model parameters

Parameters ssl = ss1(Pa)

ss2(Pa)

rl = ro = ra(g/cm3)

rc1 = rc2=(g/cm3)

Hl

(cm)

Hc1

(cm)

Hc2

(cm)

V

(m/s)

Upper plate

weakening

La(cm)

Lb(cm)

Experiment 1 43 0.8 1 0.86 1.5 0.8 0.2 10� 4 arc-notch 7

Experiment 2 43 1 0.86 1.5 1 0 10� 4 arc-notch 7

Experiment 3 43 0.8 1 0.86 1.5 0.8 0.2 10� 4 arc + ridge-notch 7 8.8

Experiment 4 43 0.8 1 0.86 1.5 0.8 0.2 10� 4 arc + ridge-notch 7 10.2

Experiment 5 43 1 1.5 0 0 10� 4 arc + ridge-notch 7 8.5

ssl, ss1 and ss2 are the yield limits for the mantle, upper and lower crustal layers of the lithosphere under normal load, respectively; rl, ra, ro, rc1and rc2 are the densities of the mantle lithospheric layer, asthenosphere, overriding plate, upper and lower crustal layers, respectively; Hl, Hc1

,

Hc2are the thicknesses of the mantle lithospheric layer, upper and lower crustal layers; V is the rate of the plate convergence. Lb is the back-arc

spreading center/trench distance; La is the arc/trench distance.


Fig. 5. Experiment 1. Successive stages of subduction of the continental margin which was not preceded by back-arc opening (see Table 1 for

the model parameters).


Experiment 2 (Fig. 6). Failure occurs again during

margin subduction, but at an earlier stage and along a

continent-ward-dipping fault (Fig. 6d).

Experiment 3 (Fig. 7). In this and the next

experiments, the overriding lithosphere is thinned

in both the arc area and the back-arc basin according

to Fig. 3; the back-arc lithosphere thinning is stron-

ger. The overriding lithosphere fails in the back-arc

during subduction of the continental margin along a

continent-vergent fault (Fig. 7b). The failure is

followed by arc plate subduction, which occurs

simultaneously with subduction of the continental

margin (Fig. 7c–f).

Experiment 4 (Fig. 8).The only difference with the

previous experiment is an increase in the trench/back-

arc spreading centre distance Lb by 1.4 cm (� 40 km

in nature). This modification caused the overriding

plate to fail in the opposite direction (Fig. 8f). The

failure was followed by subduction reversal (Fig. 8g).

Experiment 5 (Fig. 9). The overriding plate has the

same geometry as in experiment 3 and is pre-cut as

shown in Fig. 9a. The subducting plate is entirely

oceanic. During the initial stages of this experiment,

one can observe simultaneous subduction of the arc-

plate and the oceanic lithosphere.

3.3. Comments on the experimental models

In the first two experiments, which differ only by

the presence or absence of the weak lower crust, the

overriding plate fails in opposite directions. The

reason is the difference in flexural rigidity D of the

continental margin lithosphere, which is proportional

to H3, where H is the thickness of the bending layer.

Fig. 6. Experiment 2. Same as the previous experiment except that there is no weak lower crust in subducting continental margin. The whole

crust has the same properties as the upper crust in the previous experiment and is welded to the mantle (the coupling between the crust and the

lithospheric mantle is strong).


Fig. 7. Experiment 3. Subduction of the continental margin preceded by back-arc opening. The trench/back-arc ridge distance is Lb = 8.8 cm

(equivalent to 265 km in nature) (see Table 1 for other model parameters).


In experiment 1, D1�Hc1

3 +Hl3, while in experiment

2, D2� (Hc1+Hc2

+Hl)3, where Hl, Hc1

, and Hc2are

the thicknesses of the mantle lithosphere, upper crust

and lower crust of the margin, respectively. It is seen

that D1«D2 (see also Burov and Diament, 1995).

Therefore, the non-hydrostatic pressure (normal

Fig. 8. Experiment 4. Same as the previous experiment except the trench/back-arc ridge distance, which is now 10.2 cm (equivalent to 305 km in

nature).


stress) sr exerted by the subducting plate on the

overriding lithosphere is higher in experiment 2. This

pressure has the form shown in Fig. 10a and produces

both horizontal compression of the overriding plate

(force Fh) and counter-clockwise torque (T1) on the

fore-arc block. When there is no other force acting

along the interplate surface (no interplate friction), the

interplate pressure sr causes the overriding plate to

fail along the ocean-vergent fault dipping under the

arc (Fig. 6) consistent with the counter-clockwise

fore-arc block rotation. This case corresponds to

oceanic subduction (Shemenda, 1994; Tang and

Chemenda, 2000). During the subduction of the con-

tinental margin (experiments 1 and 2), another factor

is involved, the buoyancy of a progressively thicker

subducting continental crust. The buoyancy force is

proportional to the thickness of the crust and produces

the interplate normal stress sb along the interplate

Fig. 9. Experiment 5. Subduction of oceanic lithosphere. The overriding plate was pre-cut at the back-arc ridge along the trench-vergent surface

(see Table 1 for the model parameters).

Fig. 10. Interplate normal non-hydrostatic stresses: (a) due to the flexural rigidity of the subducting lithosphere, sr (Shemenda, 1994); (b) due to

the buoyancy of the subducting crust of the continental margin, sb (Tang and Chemenda, 2000). T1 and T2 are the torques exerted on the

overriding plate and caused by sr and sb. Fp1and Fp2

are the resultant non-isostatic (tectonic) pressure forces caused by sr and sb and acting on

the overriding plate. Fh is the horizontal component of Fp.


surface shown in Fig. 10b (Tang and Chemenda,

2000). This stress causes a clockwise torque (T2,

Fig. 10b) to be exerted on the fore-arc block, which

favours failure of the overriding plate along the

continent-vergent fault (as in Fig. 7). The subducting

continental margin possesses both the rigidity (corre-

sponding to sr) and the buoyancy (corresponding to

sb). Combinations of different sn and sb values may

result in different failure directions. When the flexural

rigidity D is small (case of experiment 1, Fig. 5), the

failure direction will be defined by the crustal buoy-

ancy (failure along the continent-vergent fault). If the

rigidity is high, the torque produced by sr will prevailover that caused by sb, and the overriding plate will

fail along the continent-ward-dipping fault (Fig. 6).

The torque due to the buoyancy of the subducted crust

is proportional to the gradient of crustal thickness

increase and is thus inversionally proportional to the

continental margin width. The margin flexural rigidity

D and the corresponding torque depend on the thick-

ness of the lithosphere (on its age) and on the coupling

between its layers. The failure mode depends also on

the distance La between the trench and the arc axis and

on the interplate friction (Chemenda et al., 1997; Tang

and Chemenda, 2000). In experiments 1 and 2, Lacorresponds to � 200 km and the interplate friction is

zero.

Experiments 3 and 4 show that the failure of the

overriding plate in the back-arc and the subsequent

subduction process are similar to those in the previous

experiments. The failure mode depends on the dis-

tance Lb between the failure location (spreading axis)

and the trench, or more exactly, on the ratio V = Lb/l,where l is the wavelength of the flexural bending of

the arc/back-arc lithosphere. If the trench/arc distance

does not change significantly from one region to

another, and in most subduction zones averages close

to 200 km (the value adopted in experiments 1 and 2),

the variation of the trench/back-arc spreading centre

distance Lb is much higher; typically it ranges

between 250 and 350 km. Therefore, one might

expect a considerable variation in back-arc lithosphere

failure direction that is sensible to the Lb value. One

the other hand, we were not able to ‘‘capture’’ in the

experimental models a sensitivity of the failure mode

to the flexural rigidity of the subducting lithosphere as

it was the case in experiments 1 and 2. To study more

precisely the influence of this parameter, as well as the

buoyancy of the continental margin and the interplate

friction on the failure of the back-arc lithosphere, we

have performed numerical modelling.

In experiments 1 and 3 (Figs. 5 and 7), the

continental crust subducted to a depth equivalent in

nature to about 150 km and it did not fail. The failure

and subsequent rise of the crust driven by buoyancy

occur when the crust reaches greater depth, 200–250

km (see experiments in Chemenda et al., 1996).

4. Numerical modelling

4.1. Set-up

We studied the deformation and failure of the

oceanic overriding lithosphere using the finite-ele-

ment code ADELI (Hassani, 1994). The lithosphere

has the same geometry and elasto-plastic rheology

with strain weakening, as in the above experimental

Fig. 11. Setup of numerical model. sn is a non-hydrostatic normal stress which is equal to sr, sb (defined in Fig. 10), or to sn + sb. tn is theinterplate friction stresses. Arc axis/trench distance La in all experiments is 200 km. Extinct back-arc spreading centre/trench distance Lb varies

in different experiments. H = 60 km; ha = 16 km; hb = 12 km.



models (Fig. 4). The lithosphere floats upon a

Winkler (liquid) base and is deformed under the

normal sn and tangential tn stresses along the

interplate surface (Fig. 11). The normal stress corre-

sponds either to sr or sb (in Fig. 10) or sr + sb. Thetangential stress corresponds to the interplate friction.

As in the experimental models, the lithospheric

density is the same as that of the asthenosphere

and is equal to 3.3� 103 kg/m3. The lithosphere is

covered by 4.5 km of water. The kinematic boundary

condition at the right edge allows no displacement in

the horizontal direction. For more details of the set-

up, see Tang and Chemenda (2000).

4.2. Results

We present here four sets of numerical tests. The

setup of the first set (tests 1.1–1.4, Fig. 12) corre-

sponds to experiments 1 and 2 (Figs. 5 and 6), and is

designed to model the lithosphere failure in the arc.

In three other sets (Fig. 13), we vary the distance Lbbetween the trench and the extinct back-arc spread-

ing centre, keeping the interplate boundary condi-

tions unchanged for each set.

Tests 1.1 and 1.2 (Fig. 12a) show that failure of the

overriding lithosphere only due to the flexural rigidity

of the subducting lithosphere (test 1.1) or only due to

the buoyancy of the continental margin (test 1.2)

occur in opposite directions. The sr value used in this

and the following tests corresponds to the flexural

rigidity of an old oceanic lithosphere (Tang and

Chemenda, 2000). If the rigidity is smaller (the plate

is younger), the horizontal compression of the over-

riding plate is smaller as well. The failure of this plate

occurs after greater thinning, but in the same direc-

tion. The sb applied in test 1.2 and the following

numerical experiments was calculated for a subducted

continental crust density of 2.8� 103 kg/m3. The

thickness of the crust at the trench is near 25 km

and decays linearly to zero at the deepest point of the

interplate surface (see Tang and Chemenda, 2000 for

more details).

Tests 1.3 and 1.4 (Fig. 12b). In these tests, we

varied the strength ss of the lithosphere as shown in

Fig. 12b. ss decreases toward the arc axis, which

simulates to a first approximation the dependence of

the effective lithospheric strength on its thickness:

thinner lithosphere is weaker. We also added an ‘‘arc’’

composed of volcanics with a density of 2.8� 103 kg/

m3, and the same rheologic parameters as the litho-

sphere (variation of these parameters does not affect

the lithosphere failure direction). The arc is isostati-

cally compensated and is 19.3-km thick, which yields

an underwater topographic edifice of 4.2 km. Such

topography is representative of real subduction zones,

although this parameter has little influence on the

modelling results. Tests 1.3 and 1.4 are presented here

to demonstrate that the failure direction of the over-

riding plate is not sensitive to plate structure. In all the

following tests, we use a simple homogeneous struc-

ture of the lithosphere that corresponds to the above

analogue experiments.

Tests 2.1 to 2.3 (Fig. 13a). The setup corresponds

to Fig. 11. Failure of the overriding lithosphere is due

to the buoyancy of the subducted crust. In tests 2.1

and 2.2, the spreading centre/trench distances Lb are

the same (considering the scaling factor) as in experi-

ments 3 and 4 (Figs. 7 and 8), respectively. The failure

direction is also the same. This direction changes

again at larger Lb distance (test 2.3), which was not

tested experimentally.

Tests 3.1 to 3.3 (Fig. 13b). Failure due to the

flexural rigidity of the subducting plate occurs in the

same direction (along the continent-vergent fault) at

different Lb values. For Lb = 305 km, the result does

not correspond to the appropriate analogue experi-

ment (experiment 4, Fig. 8) because the interplate

normal stress in the experiment is equal to sr + sb.When in numerical model we superpose the sr and sbvalues, the failure at Lb = 305 km occurs along a

Fig. 12. Numerical tests 1.1 and 1.4 with overriding lithosphere weakened only in the arc area. (a) Tests 1.1 and 1.2: failure occurs under the

non-hydrostatic normal stress applied to the interplate surface: in test 1.1 this stress (sr) is only due to the flexural rigidity of the subducting

lithosphere, and in test 1.2, the stress sb is only due to the buoyancy of the subducted continental margin crust. (b) Tests 1.3 and 1.4; the

boundary conditions are the same but the structure of the lithosphere is modified (see set up in (b)): ss diminishes toward the arc axis. An ‘‘arc’’

composed of volcanics with a density of 2.8� 103 kg/m3, and the same rheologic parameters as the lithosphere is also added. The arc is

isostatically compensated, and is 19.3-km thick, which yields an underwater topography of 4.2 km (see text for more details). CPD is cumulative

plastic deformation.


continent-dipping fault, i.e. in the same direction as in

experiment 4 (Fig. 8).

Tests 4.1 to 4.3 (Fig. 13c). As in the three previous

tests, the failure of the overriding plate under interplate

friction stress tn occurs along continent-vergent fault

at all tested distances Lb. The tn value for subductionzones is estimated to be a few tens of megapascals

(Maekawa et al., 1993; Tichelaar and Ruff, 1993;

Fig. 13. Failure of the overriding plate near an extinct back-arc spreading centre located at distances of 265, 305 and 345 km from the trench:

numerical experiments. The failure occurs under: (a) normal interplate stress sb due to the buoyancy of the subducted margin; (b) normal

interplate stress sr due to the flexural rigidity of the subducting plate; (c) interplate friction stress tn. CPD is cumulative plastic deformation.


Peacock, 1996). We have used a value of tn = 28 MPa

to obtain the plate failure at the same thickness hb (see

Fig. 11) as in the previous tests. When tn is smaller, the

failure occurs at lower hb value, but always in the same

direction.

5. Discussion of the modelling results

The results obtained by the experimental and

numerical modelling show a remarkable coincidence.

The experimental approach allows us to model the

whole subduction/collision process from oceanic sub-

duction and failure of the overriding plate to the entire

subduction of the fore-arc block or arc plate. Using

numerical technique, we were able to model only the

deformation of the overriding lithosphere and only

through the initial stages of its failure. This technique,

however, allowed us to study the influence of different

parameters on the lithosphere failure mode, which

according to the experimental modelling defines the

subsequent subduction process. It was shown in

particular that if the lithospheric failure occurs in the

arc area, the failure direction is sensitive to the

flexural rigidity of the continental margin: high

rigidity (e.g. old continental margin) favours failure

along the continent-ward-dipping fault and a subduc-

tion reversal. A narrow continental margin (high

gradient of continental crust thickness perpendicular

to the margin) favours failure in the opposite direction

and subduction of the fore-arc. The torque caused by

interplate friction has the same sign as that due to the

flexural rigidity and thus ‘‘works’’ for the plate failure

along the continent-ward-dipping fault. The failure

direction depends on the trade-off between the oppo-

site torques and is little sensitive to the rheology and

lithospheric structure in the arc (Tang and Chemenda,

2000).

When the lithospheric failure occurs in the back-arc

basin, far from the trench, the failure direction is less

sensitive to the conditions (torques) along the inter-

plate surface (Saint Venant principle) and is largely

controlled by the trench/back-arc spreading axis dis-

tance Lb. We cannot claim that the critical Lb values

corresponding to the switch in the failure direction in

simple 2-D models are the same in nature. This value

should depend on the poorly known rheological details

of the lithosphere. Besides, the real situations are

always three-dimensional and the lithosphere (crust)

of the back-arc basin contains pre-existing faults

whose orientation may define the dip of the resultant

lithospheric fault regardless of the Lb value. Thus, the

back-arc lithosphere can fail in either direction: in one

case the failure is followed by the subduction reversal,

in another case (which we will analyse further in this

paper) by subduction of the whole arc plate. Prelimi-

nary experimental tests not presented in this paper

have shown that during arc plate subduction, arc crust

can either be subducted completely into the mantle or

be scraped-off from its lithospheric mantle and

accreted to the other half of the back-arc lithosphere.

The result depends mainly on two factors: the thick-

ness of the arc crust, and the coupling between this

crust and the mantle substratum.

Failure of the overriding lithosphere is most likely

during subduction of the continental margin (arc-

continent collision) because of the increase in hori-

zontal compression of this lithosphere during subduc-

tion of the buoyant and progressively thickening

continental crust of the margin. In this case the

subduction regime switches from tensional (with

back-arc opening) to strongly compressional because

of the margin subduction. Failure of the very young

back-arc lithosphere is inevitable during this process.

It is known that change in the stress regime occurs

during oceanic subduction as well. The period of this

change is of the order of 10 Ma (Zonenshain and

Savostin, 1979; Nakamura and Uyeda, 1980; Yamano

and Uyeda, 1985). For example, there were at least

two episodes of back-arc opening in the rear of the

Mariana subduction zone at 23–15 Ma ago (Okino et

al., 1998) and 6 Ma to present (Hussong and Ueda,

1981). It is not clear what the stress regime was and

what happened between the episodes of the back-arc

spreading in this region. A tensional subduction

regime leaves clear traces (the oceanic lithosphere

formed in the back-arc basins). This is not always

true for compressional regimes, but such regimes can

be identified where currently active. For example, the

Kuriles and Japan underwent tensional subduction in

the Miocene (Zonenshain and Savostin, 1979; Naka-

mura and Uyeda, 1980) and now are under strong

compression nicely shown by tectonic, geodetic, and

seismologic data (e.g. Zonenshain and Savostin, 1979;

Baranov and Lobkovsky, 1980; Nakamura and Uyeda,

1980; Hashimoto and Jackson, 1993). The mechanism


of change from a tensional to a compressional regime

(and vice versa) during oceanic subduction is less

obvious than during subduction of a continental

margin. It may be related to the breakoff of the

subducted lithosphere, to reorganisation of the asthe-

nospheric currents or plate kinematics. Even if com-

pression of the overriding lithosphere during oceanic

subduction (compressional regime) normally is less

than during arc-continent collision, it should be large

enough to cause failure of very thin and weak back arc

lithosphere which has been formed a few million

years before. Experiment 5 (Fig. 9) shows that if this

occurs, the arc plate can be subducted into the mantle

in the same way as during arc-continent collision. In

experiment 5, the arc plate subducts completely into

the mantle because it has mantle density (the same as

the asthenosphere). If the average density of this plate

is lower, the subduction can be partial. Complete

subduction of the arc-plate should be harder in reality

than in the experiments where we neglect the astheno-

spheric viscosity.

Underthrusting of the arc plate during oceanic

subduction in the models was somewhat unexpected.

It is explained by the ease of initiation of a low-angle

underthrust of the arc plate near the spreading centre.

This underthrusting makes the principal subduction

steeper, and hence harder. Therefore the arc plate

subduction continues simultaneously with the princi-

pal subduction. It would be interesting to look for

evidences of a complete or partial subduction (dis-

appearance) of the arc plate in the real oceanic

subduction zones. In this paper, however, we will

analyse the plausibility of this process in arc-continent

collision environment using example of the Oman

Mountain belt.

6. Oman orogen

6.1. Geodynamic setting and geologic constraints

The Oman orogen, located at the northeastern

margin of Arabia, is certainly the world’s most famous

obduction-related mountain system. It has been inten-

sively studied for the last 30 years (e.g. Glennie et al.,

1974; Coleman, 1981; Searle, 1985; Lippard et al.,

1986; Nicolas, 1989; Robertson et al., 1990a; Le

Metour et al., 1995) and numerous models have been

put forward to explain evolution of this belt (e.g.

Coleman, 1981; Alabaster et al., 1982; Boudier et al.,

1988; Goffe et al., 1988; Michard et al., 1994; Searle

et al., 1994; Chemenda et al., 1996; Hacker and Gnos,

1997; Gregory et al., 1998; Searle and Cox, 1999).

Schematically, the Oman Mountains comprise four

major superposed mega-units, from bottom to top

(Robertson et al., 1990a) (Fig. 14):

� the autochtonous and parautochtonous cover

and basement of the Arabian platform;� the Sumeini and Hawasina Nappes;� the Haybi Complex sensu lato including the

ophiolite metamorphic sole;� the Samail Ophiolite Nappe.

The post obduction Maastrichtian to Paleogene

shallow marine sediments disconformably overly the

above units which we describe below.

Autochton and parautochton. The autochton com-

prises a thick pile of Permian to Cenomanian shallow

water carbonates belonging to the former Arabian

shelf. The carbonates are interrupted by a major

sedimentary break in lower to mid-Turonian times

(91–92 Ma, according to the new chart by Hardenbol

et al., 1998), which marks the sudden flexure of the

Arabian margin (subsidence of the shelf and bulging

of the shelf-slope area: Robertson, 1987; Le Metour et

al., 1995).

The sedimentary cover is gently folded and thrust

southwestward in the outer part of the chain. In the

hinterland the so-called ‘‘parautochton’’ crops out

below the allochton in two large culminations, the

Jebel Akhdar and Saih Hatat windows. It is composed

of the same Permian to Upper Cretaceous cover

together with the Precambrian to Paleozoic basement.

In both culminations, the dominant deformation is

characterized by northeast-verging ductile shearing

and folding. In the northern part of the Saih Hatat

window, it is superimposed on an older and deeper

south-verging deformation (Le Metour et al., 1990;

Michard et al., 1994; Mattauer and Ritz, 1996; Jolivet

et al., 1998; Gregory et al., 1998). In the Saih Hatat

window HP/LT peak metamorphism increases north-

eastward from blueschist carpholite facies (8–12 kb,

280–320�C: Goffe et al., 1988; Michard et al., 1994)

to eclogite facies at As Sifah (20–23 kb, 540�: Searleet al., 1994; Wendt et al., 1993; Searle and Cox, 1999).


The age of the HP metamorphism is poorly con-

strained: various K/Ar and 40 Ar/39Ar ages range

between 130 and 70 Ma (Montigny et al., 1988; El-

Shazly and Lanphere, 1992). Based on stratigraphic

and paleogeographic constraints, most authors con-

sider that the Saih Hatat parautochton did not ex-

perienced HP metamorphism before late Cretaceous

time (Michard et al., 1994; Le Metour et al., 1990,

1995). In fact, the youngest parautochton sediments

underthrust in the windows and at the front of the

Hawasina nappes are at least upper Coniacian–Santo-

nian (87–83 Ma) up to lower Campanian (83–77

Fig. 14. The Oman mountains (after Michard et al., 1994; Feinberg et al., 1999). (a) Geodynamic setting and main geological units; (b)

Schematic cross-section trough the southeastern Oman Mountains (along AA0 in (a)). 1 = outcrops of Arabian basement and Mesozoic cover;

2 = parautochton units (JA and SH are Jebel Akhdar and Saih Hatat windows, respectively); 3 = Sumeini and Hawasina nappes; 4 = Samail

ophiolite nappe.


Ma)(Robertson, 1987). In any case, rapid exhumation

of the HP units was already in progress between 80 ± 8

and 68 ± 2 Ma (based on zircon fission-track thermo-

chronology: Saddiqi et al., 1995).

Sumeini and Hawasina nappes form a relatively

well preserved allochtonous tectonic prism with

stacked various paleogeographic domains from the

continental slope (Sumeini) to the deepest outer part

of Arabian margin (Hawasina Basin and related

domains) (Searle et al., 1980; Bechennec et al.,

1990; Le Metour et al., 1995). Genesis and evolution

of the Arabian passive margin is well recorded in the

Hawasina series: a Late Permian rifting phase was

followed by Middle to Late Triassic opening of the

Neo-Tethys. Drowning of the outer margin occurred

in late Tithonian–Berriasian times (Bechennec et al.,

1988, 1990). The youngest sediments tectonically

incorporated into the Hawasina and Sumeini nappes

are middle Turonian to Coniacian (91–86 Ma) (Le

Metour et al., 1995). The Hawasina units are unme-

tamorphosed except at the northwestern border of the

Saih Hatat window where they contain HP/LT para-

genesis (lawsonite or Fe carpholite–pyrophyllite:

Michard et al., 1994).

Haybi Complex sensu lato is a set of tectonic slices

above the Hawasina nappes and below the Samail

peridotites (Searle and Malpas, 1980, 1982; Searle et

al., 1990; Searle and Cox, 1999). This complex

includes the 100–500-m thick HT metamorphic sole

of the ophiolite and the Haybi Complex sensu stricto.

The latter is made of various unmetamorphosed rocks

belonging to the outer Arabian rifted margin (Permian

and Triassic ‘‘exotic’’ limestones, Permian up to

Cenomanian alkalic and (minor) tholeiitic basalts).

The metamorphic sole is welded to the ophiolite and

is separated from the underlying unmetamorphosed

sediments by a brittle thrust. It shows an inverted

metamorphic gradient with HT amphibolite facies at

the top and greenschist facies below; a sheared con-

tact lies in between (Hacker et al., 1996; Searle and

Cox, 1999). The greenschist facies rocks have strong

affinities with the above-mentioned Haybi Complex

and with upper Jurassic to Valanginian radiolarites

(Searle and Malpas, 1980; Robertson et al., 1990b).

Protoliths of the amphibolite-facies rocks are MORB

basalts, pelagic limestones and radiolarites that could

belong to the former Jurassic–Lower Cretaceous

tethysian oceanic crust, although evidences are really

poor (Rabu et al., 1993; Searle and Cox, 1999). HT

metamorphism related to the initial intra-oceanic

thrusting occurred around 94–93 Ma (Hacker et al.,

1996), at peak temperatures of 775–875 �C (Searle

and Malpas, 1980; Ghent and Stout, 1981; Hacker

and Gnos, 1997) and under pressure of 5–7 kb or

locally even 11 kb (Hacker and Gnos, 1997; Searle

and Cox, 1999).

Samail ophiolite. The Samail ophiolite is a 15–

20-km thick remnant of oceanic lithosphere (Lip-

pard et al., 1986; Nicolas, 1989). Two major mag-

matic episodes (V1 and V2: Ernewein et al., 1988)

led to the genesis of its crustal section. The first

and main episode of ridge-type accretion (‘‘Geo-

times unit’’ after Alabaster et al., 1982) is upper-

most Albian to Cenomanian (Tilton et al., 1981;

Tippit et al., 1981; Beurrier et al., 1987, 1989). It

was immediately followed by the second magmatic

episode (‘‘Lasail’’ and ‘‘Alley units’’: Alabaster et

al., 1982) in the middle Cenomanian to late Turo-

nian. The geodynamic setting of these two episodes

is still controversial. A back-arc basin setting has

been proposed by Pearce et al. (1981) and Alabaster

et al. (1982) based on geochemical data. According

to these authors, the first episode (‘‘Geotimes’’)

corresponds to a transition from MORB to arc

tholeiites and the second (‘‘Lasail’’) has island arc

affinities. A normal mid-ocean ridge setting has

more advocates (Boudier and Coleman, 1981; Cole-

man, 1981; Juteau et al., 1988; Ernewein et al.,

1988). Following Ernewein et al. (1988), the first

episode of magmatism corresponds to typical mid-

ocean ridge tholeiites, whereas the second is the

result of intraoceanic thrusting preceding obduction

on the Arabian margin.

The youngest deep-sea sediments preserved at the

top of the Samail oceanic crust are Lower Campanian

(83 up to 80 Ma or less) radiolarites (Schaaf and

Thomas, 1986). The first ophiolitic pebbles appear in

the Upper Campanian (76–72 Ma) (Juweiza Fm.

Rabu et al., 1993; Warburton et al., 1990).

6.2. Principal geological constraints on an evolu-

tionary model

We do not aim to explain all various-scale details

of the geological evolution of the Oman Mountains,

which are complex, often poorly constrained and


sometimes controversial. We are interested in the

lithospheric-scale processes. The evolutionary model

of these processes is a forced simplification of reality

that neglects many details, but it necessarily must take

into account and explain well established first order

events that define the evolution of the mountain belt.

For the Oman Mountains and its principal unit, the

Samail ophiolite, these are the following.� The Samail ophiolite was in a deep and quiet

marine environment between the uppermost Albian

(100–99 Ma) and lower Campanian (80 Ma or less).

HT metamorphism of the ophiolite sole shows that

within this period, at 94 Ma (Hacker et al., 1996) there

was an episode of intraoceanic overthrusting of the

ophiolite, first over a still hot, i.e. very young, oceanic

lithosphere (Boudier et al., 1988; Hacker et al., 1996)

and then over old Jurassic/Early Cretaceous oceanic

crust. This episode had to be very short as it did not

leave any trace in the sedimentary cover of the

ophiolite.� While the Samail ophiolite stayed undeformed,

subduction of the outermost Arabian margin was

certainly already in progress by 88 Ma (Coniacian).

This subduction could not have occurred under the

Samail ophiolite, i.e. this ophiolite could not have

been in the position of the frontal part of the over-

riding plate which in presently active subduction

zones undergoes intense deformation. This deforma-

tion becomes still greater during underthrusting of the

continental margin (in Taiwan, for example: e.g.

Lundberg et al., 1997). The deformation (hence,

obduction) of the Samail ophiolite started at � 81

Ma and by 76 Ma it has already overthrust the upper

Arabian margin and reached sea level. Thus, at least 7

Ma separates the onset of the Samail ophiolite obduc-

tion on the margin (81 Ma) from the beginning of

subduction of the outermost Arabian margin (88 Ma).

It means that, depending on the convergence rate, 350

km (for 5 cm/year) or 700 km (for 10 cm/year) of

lithosphere located between the subducting margin

and the ophiolite has totally disappeared. This con-

clusion is consistent with the fact that in many places

(on the southern side of the Saih-Hatat window, for

example) the Samail thrust is clearly an out-of-

sequence thrust (Gregory et al., 1998), which means

that some tectonic units have been pinched out

between the ophiolite and the underlying Hawasina

prism. These units which correspond to an overriding

(with respect to the subducting margin) plate, may

have been both the old (Jurassic/lower Cretaceous)

Tethyan and the Cretaceous back-arc lithosphere.

Following the above modelling results, we propose

that this lithosphere is the arc plate, which was

subducted into the mantle together with the Arabian

margin. The incorporation of the arc-plate subduction

makes the principal difference between the model

proposed by Chemenda et al. (1996) and that pre-

sented below.

6.3. Evolutionary model of continental subduction in

Oman

The ophiolite of Oman started to form within or

behind an immature intra-oceanic volcanic arc �100

Ma ago. The tensional regime of the associated intra-

oceanic subduction changed to a compressional

regime by 94 Ma. The reason for this change is

not known. It may have been the general reorganiza-

tion of the plate kinematics in this region between

110 and 83 Ma when the relative displacement of

Africa with respect to Eurasia changed from E–W to

SW–NE (Patriat and Achache, 1984). Compression

caused the overriding lithosphere to fail along a NE-

dipping fault in the vicinity of the back-arc ridge

(Fig. 15a). As the ridge was very oblique to the

subduction zone (Boudier et al., 1988), this fault was

not necessarily initiated along the spreading centres

and/or transform faults (Hacker et al., 1996). It was

initiated in the axial zone of the ridge with thin and

weak lithosphere (see Fig. 15a). Thus, in some

places the Oman ophiolite overthrust young back-

arc lithosphere and in the others old (Triassic accord-

ing to Searle and Cox, 1999) oceanic crust. Under-

thrusting of the arc plate occurred simultaneously

with the principal subduction (as in the experimental

model in Fig. 9) and was then blocked until arrival

of the Arabian continental margin to the principal

subduction zone at � 88 Ma (Fig. 15b). Subduction

of the margin caused a progressive increase in the

horizontal compression of the lithosphere. Over-

thrusting of the Oman ophiolite was restarted (reac-

tivated) and this time resulted in complete

subduction of the arc plate (including the immature

arc itself) into the mantle (Fig. 15c and d). The

accretionary prism formed at the front of this plate

and its cover were partially subducted under and



partially accreted to the Samail ophiolite ( = unmeta-

morphosed part of the Haybi Complex?). The Ara-

bian margin underlying the arc plate was subducted

along with it and was preserved between this plate

and the lithospheric mantle at a relatively low

temperature. On reaching a depth of � 200 km,

the continental crust failed and its deeply subducted

part started to rapidly rise up in (intrude) the channel

between the lithospheric mantle and the arc plate

(Fig. 15e) by buoyancy (see Chemenda et al., 1996).

During this rapid exhumation, the HP rocks reached

a few tens of kilometres depth. The subsequent

slower exhumation (Fig. 15f) occurred between 80

and 68 Ma (Campanian to Lower Maastrichtian)

(Saddiqi et al., 1995) due to erosion of the ophiolite

shield (Warburton et al., 1990; Rabu et al., 1993).

The break-off of the dense lithospheric mantle (Fig.

15f ) caused reduction of the driving force of sub-

duction (the pull force) and rapid isostatic uplift of

the orogen. This uplift resulted in the almost com-

plete emergence of the overriding wedge above sea

level at 72–65 Ma (Maastrichtian) (Nolan et al.,

1990). The continental subduction was stopped.

6.4. Comments on the model

We do not bring new geological or geochemical

constraints on the oceanic ridge versus back-arc origin

of the Samail ophiolite, but our choice in the model

(Fig. 15) in favour of the back-arc origin is clear. On

the other hand, the modelling results do not militate

against another initial setting of subduction (Fig. 16)

corresponding to stage a in Fig. 15. In Fig. 16, the

oceanic basin is behind the volcanic arc, but it is not

genetically related to the subduction (analogue of the

Fiji basin with respect to the Tonga arc, for example).

The only difference between Figs. 15a and 16 is the

spreading centre/trench distance Lb, which in Fig. 16

can be 500 km or more. The evolution of the situation

in Fig. 16 could be the same as in Fig. 15 with the

only difference that instead of a small arc plate, the

larger (wider) plate will be subducted into the mantle

before the young oceanic lithosphere overthrust the

Arabian crust. In its new configuration, the Samail

ophiolite will not bear traces of arc volcanism; the

volcanic arc located to the NW is supposed to be

completely subducted.

Fig. 16. Ridge origin of the Samail ophiolite: alternative to Fig. 15a initial setting of plate convergence at 95 Ma (notations in Fig. 15).

Fig. 15. Evolutionary model of continental subduction in Oman (modified after Chemenda et al., 1996 by the addition of a stage of arc plate

subduction). (a) Change in the regime of oceanic subduction from tensional to compressional and initiation of new subduction along the former

back-arc spreading zone. (b) Arrival at the trench and flexural buckling of the Arabian margin. (c) Simultaneous subduction of the continental

margin and the arc plate; uplift of Samail ophiolite. (d) Failure of deeply subducted crust at 70–100-km depth and beginning of a rapid

exhumation. (e) The crust intrude the interplate zone and puches up the Arabian margin cover accreted previously at various depths. (f) Break-

off of the lithospheric mantle, isostatic uplift and erosion of Samail ophiolite, slow exhumation of HP rocks previously uplifted to shallow

depths, end of continental subduction. (1) Arabian crust (a—upper strong and brittle layer, b—lower weak and ductile layer); (2) continental

sedimentary cover (a—Permo-Mesozoic, b—Proterozoic and Paleozoic); (3) oceanic lithosphere formed in the back arc basin; (4) older

oceanic lithosphere; (5) Hawasina nappe; (6) supposed very incipient volcanic arc; (7) ductile fault; (8) cleavage and folding; (9) supposed

present geometry of the lithospheric base; (10) subduction zone; (11) thrust (a) and normal (b) faults; (12) marker corresponding to � 80 km

depth (ca. 20 kbar) at stage in (d); (13) erosion; (14) direction of vertical movement.


In either case, a large amount of lithosphere had to

have disappeared without leaving no trace. There is

nothing strange in complete subduction of the oceanic

lithosphere including its sedimentary cover. Such a

process is observed in many active subduction zones

with tectonic erosion of the overriding plate (Von

Huene and Lallemand, 1990; Lallemand, 1999). In

such zones not only the subducting plate (and its

cover) completely subducts itself, but it also scrapes

and brings to depth part of the overriding lithosphere.

Arc (especially immature arc) can be subducted in the

same way, although it would be very encouraging to

have some products of its activity accreted to the

Samail ophiolite. They have not been found so far. On

the other hand, arc remnants were recently described

in a very similar tectonic situation in Ladakh Hima-

layas (Corfield et al., 1999) where the Spontang

ophiolite has been thrust over an accretionary prism

( = Photang thrust sheet, similar in composition and

age to the Hawasina prism). The important point here

is that slices of arc material ( = Spong Arc) are

preserved in some places between the above units.

7. Conclusion

Experimental and numerical modelling presented

in this paper show that the failure of the overriding

plate is physically quite plausible or even inevitable

during subduction. The conditions for such a failure

(the weakening of this plate) are prepared during

oceanic subduction. The weakening occurs due to

the interaction between the subducting lithosphere

and the asthenosphere in the mantle corner between

the two plates, and due to the back-arc spreading. In

oceanic subduction zones with old and strong back-

arc lithosphere the weakest zone is the volcanic arc

area. When weakening becomes sufficient, the litho-

sphere fails in this area. The modes and conditions for

the failure during compressional subduction regime

were studied in detail by Chemenda et al. (1997) and

Tang and Chemenda (2000).

Almost half of the presently active subduction

zones are characterised by a tensional subduction

regime with back-arc spreading. In such subduction

zones, the weakest place is not the volcanic arc but

the back-arc spreading centre. When a subduction

regime rapidly changes from tensional to compres-

sional, failure occurs in the vicinity of the spreading

centre. This process can occur during oceanic sub-

duction along either trench-vergent or trench-dipping

fault, but the formation of a trench-vergent fault is

most likely. In this latter case, which is a subject of

our analysis, the failure is followed by a partial

subduction of the arc plate. Complete subduction

occurs during arc–continent collision after the con-

tinental margin has arrived at the trench. During

continental margin subduction the tectonic compres-

sion of the lithosphere rapidly increases and becomes

sufficient to push the arc plate into the mantle. The

arc itself, as well as the crust and sedimentary units

of the arc plate (accretionary prism including

deformed margin) are partly scraped-off and accreted

at shallow depth and partly subducted to great

depths. The deeply subducted material is preserved

at relatively low temperatures between the litho-

spheric mantle and the ‘‘cold’’ subducted arc plate,

which acts as a thermal shield. Is this shield really

necessary to explain the low peak temperatures

registered by HP and UHP/LT metamorphic rocks

or can it be explained within the framework of a

classical kinematic scheme of subduction? As was

shown in the introduction the answer is unknown.

On the other hand, it is clear that a first-order process

such as arc plate subduction should have other major

consequences that can be used to test this hypothesis.

One could analyse, for example the absence of

evidence of a volcanic arc in collisional belts that

have undergone a stage of oceanic subduction before

collision (e.g. the Alps) to see whether an arc was

not formed or whether it was formed and then

subducted. In this paper we test the model on the

Oman belt. In trying to answer the question of the

possible mechanism of emplacement of a very young

oceanic lithosphere on the continental crust in Oman,

we come to the conclusion that this lithosphere was

formed in a back-arc basin. The Arabian margin was

first underthrust beneath the arc plate. The Oman

ophiolite was emplaced on the Arabian crust only

after subduction of this plate some 10 Ma after

initiation of the continental margin subduction

beneath the arc plate. Subduction of the ‘‘cold’’ arc

plate in this region had certainly to affect the thermal

conditions of metamorphism of HP/LT rocks

exhumed in the Saih-Hatat window. In other words,

the low peak temperature ( < 600�) registered by


these rocks at � 70–80 km-depth (equivalent to ca.

23 kbar) is due to the subducted arc plate. If this

conclusion is true for Oman, it is possible that UHP/

LT rocks in other mountain belts were also meta-

morphosed and exhumed in the presence of a sub-

ducted arc plate or fore arc block or other ‘‘cold’’

subducted lithospheric unit that played the role of a

thermal shield.

Acknowledgements

We thank B. Hacker and L. Jolivet for helpful

reviews. This work has been supported by the INSU-

CNRS program ‘‘Interieur de la Terre’’ (Contribution

No. 283) and is Geosciences Azur contribution No.

392.

References

Alabaster, T., Pearce, J.A., Malpas, J., 1982. The volcanic stratig-

raphy and petrogenesis of the Oman Ophiolite Complex. Con-

trib. Mineral. Petrol. 81, 168–183.

Anczkiewicz, R., Burg, J.-P., Hussain, S.S., Dawood, H., Ghazan-

far, M., Chaudhry, M.N., 1998. Stratigraphy and structure of the

Indus Suture in the Lower Swat, Pakistan, NW Himalaya. J.

Asian Earth Sci. 16, 225–238.

Baranov, B.V., Lobkovsky, L.I., 1980. Shallow seismicity behind

the Kurile Arc and its relation to the Benioff zone. Dokl. Akad.

Nauk USSR 255, 67–71 (in Russian).

Bechennec, F., Le Metour, J., Rabu, D., Villey, M., Beurrier, M.,

1988. The Hawasina Basin: a fragment of a starved passive

continental margin, thrust over the Arabian Platform during

the obduction of the Samail Nappe. Tectonophysics 151,

323–343.

Bechennec, F., Le Metour, J., Rabu, D., Bourdillon-de-Grissac, Ch.,

De Wever, P., Beurrier, M., Villey, M., 1990. The Hawasina

Napp: stratigraphy, paleogeography and structural evolution of

a fragment of the south-Tethyan passive continental margin. In:

Robertson, A.H.F., Searle, M.P., Ries, A. (Eds.), The Geology

and Tectonics of the Oman Region. Geol. Soc. London, Spec.

Publ., vol. 49, pp. 213–223.

Beurrier, M., Bourdillon-Jeudy de Grissac, C., De Wever, P., Les-

cuyer, J.L., 1987. Biostratigraphie des radiolarites associees aux

volcanites ophiolitiques de la nappe de samail (Sultanat

d’Oman): consequences tectogenetiques. C. R. Acad. Sci., Paris

304 (II), 907–910.

Beurrier, M., Ohnenstetter, M., Cabanis, B., Lescuyer, J.L., Te-

gyey, M., Le Metour, J., 1989. Geochimie des filons doleri-

tiques et des roches volcaniques ophiolitiques de la nappe de

Samail: contraintes sur leur origine geotectonique au Cretace

superieur. Bull. Soc. Geol. Fr. 2, 205–219.

Boudier, F., Coleman, R.G., 1981. Cross-section through the Peri-

dotite in the Samail Ophiolite, Southeastern Oman. J. Geophys.

Res. 86, 2573–2594.

Boudier, F., Ceuleneer, G., Nicolas, A., 1988. Shear zones,

thrusts and related metamorphism in the Oman ophiolite:

initiation of thrusting on an oceanic ridge. Tectonophysics

151, 275–296.

Burov, E.B., Diament, M., 1995. The effective elastic thickness of

continental lithosphere: What it does really mean? J. Geophys.

Res. 100, 3905–3927.

Chemenda, A.I., Mattauer, M., Malavieille, J., Bokun, A.N., 1995.

A mechanism for syn-collisional deep rock exhumation and

associated normal faulting: results from physical modeling.

Earth Planet. Sci. Lett. 132, 225–232.

Chemenda, A., Mattauer, M., Bokun, A.N., 1996. Continental sub-

duction and a mechanism for exhumation of high-pressure meta-

morphic rocks: new modelling and field data from Oman. Earth

Planet. Sci. Lett. 143, 173–182.

Chemenda, A.I., Yang, R.K., Hsieh, C.-H., Groholsky, A.L., 1997.

Evolutionary model for the Taiwan collision based on physical

modelling. Tectonophysics 274, 253–274.

Chemenda, A.I., Yang, R.K., Stephan, J.-F., Konstantinivska-

ya, E.A., Ivanov, L.M., 2001. New results from physical model-

ling of arc-continent collision in Taiwan: evolutionary model.

Tectonophysics 333, 159–178.

Coleman, R.G., 1981. Tectonic setting for ophiolite obduction in

Oman. J. Geophys. Res. 86, 2497–2508.

Corfield, R.I., Searle, M.P., Green, O.R., 1999. Photang thrust sheet:

an accretionary complex structurally below the Spontang ophio-

lite constraining timing and tectonic environment of ophiolite

obduction, Ladakh Himalaya, NW India. J. Geol. Soc. London

156, 1031–1044.

Davies, J.H., von Blanckenburg, F., 1994. Slab breakoff: a model of

lithosphere detachement and its test in the magmatism and de-

formation of collisional orogens. Earth Planet. Sci. Lett. 129,

85–102.

El-Shazly, A.K., Lanphere, M.A., 1992. Two high-pressure meta-

morphic events in NE Oman: evidence from 40Ar/39Ar dating

and petrological data. J. Geol. 100, 731–751.

Ernewein, M., Pflumio, C., Whitechurch, H., 1988. The death of an

accretion zone as evidenced by the magmatic history of the

Sumail ophiolite (Oman). Tectonophysics 151, 247–274.

Feinberg, H., Horen, H., Michard, A., Saddiqi, O., 1999. Obduc-

tion-related remagnetization at the base of an ophiolite: paleo-

magnetism of the Samail nappe lower sequence and its

continental substratum, southeast Oman Mountains. J. Geophys.

Res. 104, 17703–17714.

Furukawa, Y., 1993. Magmatic processes under arcs and formation

of the volcanic front. J. Geophys. Res. 98, 8309–8319.

Ghent, E.D., Stout, M.Z., 1981. Metamorphism at the base of the

Samail ophiolite, southeastern Oman Mountains. J. Geophys.

Res. 86, 2557–2571.

Glennie, K.W., Bœuf, M.G.A., Hughes-Clark, M.W., Moody-

Stuart, M., Pilar, W.F.H., Reinhardt, B.M., 1974. Geology of

the Oman Mountains. Koninklijk Nederlands Geologisch Mijn-

bouwkundig Genootschap, Transactions, vol. 31, p. 473.

Goffe, B., Michard, A., Kienast, J.R., Le Mer, O., 1988. A case of


obduction-related high pressure, low temperature metamor-

phism in upper crustal nappes, Arabian continental margin,

Oman. Tectonophysics 151, 363–386.

Gregory, R.T., Gray, D.R., Miller, J.-McL., 1998. Tectonics of the

Arabian margin associated with the formation and exhumation of

high-pressure rocks, Sultanate of Oman. Tectonics 17, 657–670.

Hacker, B.R., Gnos, E., 1997. The conundrum of Samail: explaining

the metamorphic history. Tectonophysics 279, 215–226.

Hacker, B.R., Mosenfelder, J.L., Gnos, E., 1996. Rapid emplace-

ment of the Oman ophiolite: thermal and geochronologic con-

straints. Tectonics 15, 1230–1247.

Hardenbol, J., Thierry, J., Farley, M.B., Jacquin, T., de Gracian-

sky, P.C., Vail, P., 1998. Mesozoic and Cenozoic Sequence

Chronostratigraphic Framework of European Basins. SEPM

(Society of Sedimentary Geology) Special Publication No.

60.

Harrison, T.M., Copeland, P., Kidd, W.S.F., Yin, An., 1992. Raising

Tibet. Science 255, 1663–1670.

Hashimoto, M., Jackson, D., 1993. Plate tectonics and crustal de-

formation around the Japanese island. J. Geophys. Res. 98,

16149–16166.

Hassani, R., 1994. Modelisation numerique de la deformation des

systemes geologiques. PhD Thesis, Universite de Montpellier II,

France.

Hussong, D., Ueda, S., 1981. Tectonic processes and the history of

the Mariana arc: a synthesis of the results of Deep Sea Drilling

Project Leg 60. In: Lee, M., Powell, R. (Eds.), Init. Reports

DSDP 60, vol. 60. NSF, Washington, DC, pp. 909–929.

Jolivet, L., Goffe, B., Bousquet, R., Oberhansli, R., Michard, A.,

1998. Detachments in high-pressure mountain belts, Tethyan

examples. Earth Planet. Sci. Lett. 160, 31–47.

Juteau, Th., Ernewein, M., Reuber, I., Whitechurch, H., Dahl, R.,

1988. Duality of magmatism in the plutonic sequence of the

Sumail Nappe, Oman. Tectonophysics 151, 107–135.

Kincaid, C., Sacks, I.S., 1997. Thermal and dynamic evolution of

the upper mantle in subduction sones. J. Geophys. Res. 102,

12295–12315.

Konstantinivskaya, E., 2000. Geodynamics of an Early Eocene arc-

continent collision reconstracted from the Kamchatka orogenic

belt, NE Russia. Tectonophysics 325, 87–105.

Lallemand, S., 1999. La Subduction Oceanique. Gordon and

Breach, Newark, NJ., France.

Le Metour, J., Rabu, D., Tegyey, M., Bechennec, F., Beurrier, M.,

Villey, M., 1990. Subduction and obduction: two stages in the

Eo-Alpine tectonometamorphic evolution of the Oman Moun-

tains. In: Robertson, A.H.F., Searle, M.P., Ries, A. (Eds.), The

Geology and Tectonics of the Oman Region. Geol. Soc. Lon-

don, Spec. Publ., vol. 49, pp. 327–339.

Le Metour, J., Michel, J.C., Bechennec, F., Platel, J.P., Roger, J.,

1995. Geology and mineral wealth of the Sultanate of Oman.

Report, Minist. Pet. Miner. Sultanate of Oman, Muscat, pp. 285.

Lewis, T.J., Bentkowski, W.H., Davis, E.E., Hyndman, R.D.,

Souther, J.G., Wright, J.A., 1988. Subduction of the Juan de

Fuca Plate: thermal consequence. J. Geophys. Res. 93,

15207–15225.

Lippard, S.J., Shelton, A.W., Gass, I.G., 1986. The ophiolite of

Northern Oman. Mem. Geol. Soc. London 11, 178.

Lundberg, N., Reed, D.L., Liu, C.S., Lieske Jr., J., 1997. Fore-arc

basin closure and arc accretion in the submarine suture zone

south of Taiwan. Tectonophysics 274, 5–23.

Maekawa, H., Shozui, M., Ishii, T., Fryer, P., Pearce, J.A., 1993.

Blueshist metamorphism in an active subduction zone. Nature

364, 520–523.

Malavieille, J., 1999. Evolutionary model for arc-continent colli-

sion. Conference Abstr., Active Subduction and Collision in

South–East Asia, Montpellier, France, May, 231–234.

Mattauer, M., Ritz, J.-F., 1996. Arguments geologiques en faveur

d’un modele de subduction continentale pour l’exhumation du

metamorphisme haute-pression d’Oman. C. R. Acad. Sci. Paris

322 (IIa), 869–876.

Matte, Ph., 1998. Continental subduction and exhumation of HP

rocks in Paleozoic orogenic belts: Uralides and Variscides.

GFF 120, 209–222.

Michard, A., Goffe, B., Saddini, O., Oberhansli, R., Wendt, A.S.,

1994. Late Cretaceous exhumation of the Oman blueschists and

eclogites: a two-stage extensional mechanism. Terra Nova 6,

404–413.

Montigny, R., Le Mer, O., Thuizat, R., Whitechurch, H., 1988. K–

Ar and 40Ar/39Ar study of metamorphic rocks associated with

the Oman ophiolite: tectonic implications. Tectonophysics 151,

345–362.

Nakamura, K., Uyeda, S., 1980. Stress gradient in arc-back arc

regions and plate subduction. J. Geophys. Res. 85, 6419–6428.

Nicolas, A., 1989. Structures of Ophiolites and Dynamics of Ocean-

ic Lithosphere. Kluwer Academic Publishing, Amsterdam.

Nolan, S.C., Skelton, P.W., Clissold, B.P., Smewing, J.D., 1990.

Maastrichtian to early Tertiary stratigraphy and paleogeogra-

phy of the Central and Northern Oman Mountains. In: Rob-

ertson, A.H.F., Searle, M.P., Ries, A. (Eds.), The Geology and

Tectonics of the Oman Region. Geol. Soc. London, Spec.

Publ., vol. 49, pp. 495–519.

Okino, K., Kasuda, S., Ohara, Y., 1998. A new scenario of the

Parece Vela basin genesis. Mar. Geophys. Res. 20, 21–40.

Patriat, Ph., Achache, J., 1984. India–Eurasia collision chronology

has implications for crustal shortening and driving mechanism

of plates. Nature 311, 615–621.

Peacock, S.M., 1996. Thermal and petrologic structure of subduc-

tion zones. In: Bebout, G.E. et al., (Ed.), Subduction: Top to

Bottom. Geophys. Monogr. Ser. 96 AGU, Washington, DC, pp.

119–133.

Pearce, J.A., Alabaster, T., Shelton, A.W., Searle, M.P., 1981.

The Oman ophiolite as a Cretaceous arc-basin complex: evi-

dence and implications. Philos. Trans. R. Soc. London, A

300, 299–317.

Rabu, D., Nehlig, P., Roger, J., 1993. Stratigraphy and structure of

the Oman Mountains. Doc. B.R.G.M. 221, pp. 262.

Robertson, A.H.F., 1987. Upper Cretaceous Muti Formation: tran-

sition of a Mesozoic carbonate platform to a foreland basin in

the Oman Mountains. Sedimentology 34, 1123–1142.

Robertson, A.H.F., Searle, M.P., Ries, A., 1990a. The northern Oman

Tethyan continental margin: stratigraphy, structure, concepts and

controversies. In: Robertson, A.H.F., Searle, M.P., Ries, A.

(Eds.), The Geology and Tectonics of the Oman Region. Geol.

Soc. London, Spec. Publ., vol. 49, pp. 495–519.


Robertson, A.H.F., Blome, C.D., Cooper, D.W.J., Kemp, A.E.S.,

Searle, M.P., 1990b. Evolution of the Arabian continental mar-

gin in the Dibba zone, Northern Oman Mountains. In: Robert-

son, A.H.F., Searle, M.P., Ries, A. (Eds.), The Geology and

Tectonics of the Oman Region. Geol. Soc. London, Spec. Publ.,

vol. 49, pp. 251–284.

Saddiqi, O., Poupeau, G., Michard, A., Goffe, B., Oberhansli, R.,

1995. Exhumation des roches metamorphiques HP-BT d’Oman:

datation par traces de fission sur zircons. C. R. Acad. Sci., Paris

320, 1071–1077.

Schaaf, A., Thomas, V., 1986. Les radiolaires campaniens du Wadi

Ragmi (nappe de Samail, Oman): un nouveau repere chronolo-

gique de l’obduction omanaise. C. R. Acad. Sci., Paris 303

(serie II), 1593–1598.

Schmidt, M.S., Poli, S., 1995. Experimentally based water budgets

for dehydrating slabs and consequences for arc magma gener-

ation. Earth Planet. Sci. Lett. 163, 361–379.

Searle, M.P., 1985. Sequence of thrusting and origin of culminations

in the northern and central Oman Mountains. J. Struct. Geol. 7,

129–143.

Searle, M.P., Cox, J., 1999. Tectonic setting, origin, and obduction

of the Oman ophiolite. Geol. Soc. Am. Bull. 111 (3), 104–122.

Searle, M.P., Malpas, J., 1980. Structure and metamorphism of

rocks beneath the Samail ophiolite of Oman and their signifi-

cance in ophiolite obduction. Trans. R. Soc. Edinburg 71,

247–262.

Searle, M.P., Malpas, J., 1982. Petrochemistry and origin of sub-

ophiolitic metamorphic and related rocks in the Oman Moun-

tains. J. Geol. Soc. London 139, 235–248.

Searle, M.P., Lippard, S.J., Smewing, J.D., Rex, D.C., 1980.

Volcanic rocks beneath the Samail ophiolite nappe in the

northern Oman Mountains and their significance in the Mes-

ozoic evolution of the Tethys. J. Geol. Soc. London 137,

589–604.

Searle, M.P., Cooper, D.J.W., Watts, K.F., 1990. Structure of the

Jebel Sumeini–Jebel Ghawil area, Northern Oman. In: Robert-


Tectonics of the Oman Region. Geol. Soc. London, Spec. Publ.,

vol. 49, pp. 361–374.

Searle, M.P., Waters, D.J., Martin, H.N., Rex, D.C., 1994. Structure

and metamorphism of blueschist-eclogite facies rocks from the

Northeastern Oman Mountains. J. Geol. Soc. London 151,

555–576.

Shemenda, A.I., 1994. Subduction: Insights from Physical Model-

ing. Kluwer Academic Publishing, Amsterdam.

Tang, J.-C., Chemenda, A.I., 2000. Numerical modelling of arc-

continent collision: application to Taiwan. Tectonophysics

325, 23–42.

Tichelaar, B.W., Ruff, L.J., 1993. Depth of seismic coupling along

subduction zones. J. Geophys. Res. 98, 2017–2037.

Tilton, G.R., Hopson, C.A., Wright, J.E., 1981. Uranium– lead

isotopic ages of the Samail Ophiolite, Oman, with applica-

tions to Tehyan ocean ridge tectonics. J. Geophys. Res. 86,

2763–2775.

Tippit, P.R., Pessagno Jr., E.A., Smewing, J.D. 1981. The biostra-

tigraphy of sediments in the Volcanic Unit of the Samail Ophio-

lite. J. Geophys. Res. 86, 2756–2762.

Van den Beukel, J., 1992. Some thermo-mechanical aspects of the

subduction of continental lithosphere. Tectonics 11, 316–329.

Von Huene, R., Lallemand, S., 1990. Tectonic erosion along con-

tinental margins. Geol. Soc. Am. Bull. 102, 704–720.

Warburton, J., Burnhill, T.J., Graham, R.H., Isaac, K.P., 1990. The

evolution of the Oman Mountains foreland basin. In: Robert-


Tectonics of the Oman Region. Geol. Soc. London, Spec.

Publ., vol. 49, pp. 419–427.

Wendt, A.S., D’Argo, P., Goffe, B., Oberhansli, R., 1993. Radial

cracks around a-quartz inclusions in almanide: constraints on

the metamorphic history of the Oman Mountains. Earth Planet.

Sci. Lett. 114, 449–461.

Yamano, M., Uyeda, S., 1985. Possible effects of collision on plate

motions. Tectonophysics 119, 223–244.

Zhao, D., Hasegawa, A., Kanamori, H., 1994. Deep structure of

Japan subduction zone as derived from local, regional, and tele-

seismic events. J. Geophys. Res. 99, 22313–22329.

Zonenshain, L.P., Savostin, L.A., 1979. Introduction into Geody-

namics. Nedra, Moscow (in Russian).


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