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Tectonophysics 342 (2001) 137 – 161
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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
Géosciences Azur, UMR 6526, Université 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 backarc 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
*
Corresponding author. Tel.: +33-4-9294-2661; fax: +33-4-9264-2610.
E-mail address: [email protected] (A.I. Chemenda).
0040-1951/01/$ - see front matter D 2001 Elsevier Science B.V. All rights reserved.
PII: S 0 0 4 0 - 1 9 5 1 ( 0 1 ) 0 0 1 6 0 - 3
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A.I. Chemenda et al. / 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) continental crust can be dragged (pushed) to more than
100-km depth and be preserved at a relatively low
( 700C) 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 subduction is a complex process that includes the interaction,
deformation and failure of media with different rheologies. 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 transformation 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 particularly 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 represents a major (kinematic, dynamic, thermal, etc.)
difference of continental subduction from oceanic
subduction. We then apply one of the obtained subduction 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 temperatures (600 – 800 C at depth 100 –140 km) is traditionally 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 kmdepth 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) 137–161
139
Fig. 1. Model for the formation of volcanic arc in subduction zones based on petrological data (simplified after Schmidt and Poli, 1995).
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. Subduction 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 ‘‘thermal erosion’’ (partial melting?) of the overriding
lithosphere. The thickness of this lithosphere near
Fig. 2. P wave velocity structure of NE Japan subduction zone (from Zhao et al., 1994).
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A.I. Chemenda et al. / Tectonophysics 342 (2001) 137–161
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 subduction 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 overriding 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 exhumation 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 compressional 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 lithosphere thinned in the arc area and back-arc spreading
centre. The subducting plate comprises a one-layer
oceanic lithosphere and a three-layer continental lithosphere. All the lithospheric layers possess plastic
properties and are made of hydrocarbon compositional systems. The upper continental crust, the continental 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.
A.I. Chemenda et al. / Tectonophysics 342 (2001) 137–161
141
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.
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 compression, 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).
lithosphere have the same yield limit and are characterised by a strong strain weakening (Fig. 4). The
lower continental crust is considerably weaker (see
Table 1) and more ductile. The lithosphere is underlain 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 modelling are the same as those of Chemenda et al. (1995)
who presented similar experiments, but without weakening 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
Table 1
Model parameters
Parameters
Experiment
Experiment
Experiment
Experiment
Experiment
1
2
3
4
5
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)
43
43
43
43
43
0.8
1
1
1
1
1
0.86
0.86
0.86
0.86
1.5
1.5
1.5
1.5
1.5
0.8
1
0.8
0.8
0
0.2
0
0.2
0.2
0
10 4
10 4
10 4
10 4
10 4
arc-notch
arc-notch
arc + ridge-notch
arc + ridge-notch
arc + ridge-notch
7
7
7
7
7
8.8
10.2
8.5
0.8
0.8
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, rc1
and rc2 are the densities of the mantle lithospheric layer, asthenosphere, overriding plate, upper and lower crustal layers, respectively; Hl, Hc1,
Hc2 are 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.
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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).
A.I. Chemenda et al. / Tectonophysics 342 (2001) 137–161
143
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).
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 stronger. 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/backarc 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 arcplate 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.
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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).
A.I. Chemenda et al. / Tectonophysics 342 (2001) 137–161
145
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).
In experiment 1, D1 Hc13 + Hl3, while in experiment
2, D2 (Hc1 + Hc2 + Hl)3, where Hl, Hc1, and Hc2 are
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
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A.I. Chemenda et al. / Tectonophysics 342 (2001) 137–161
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).
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 continental 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. 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. Fp1 and 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.
A.I. Chemenda et al. / Tectonophysics 342 (2001) 137–161
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 (corresponding 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 buoyancy (failure along the continent-vergent fault). If the
rigidity is high, the torque produced by sr will prevail
over 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 thickness 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, La
corresponds 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 distance 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
147
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 L b 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-element 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 the
interplate 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.
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A.I. Chemenda et al. / Tectonophysics 342 (2001) 137–161
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 corresponds either to sr or sb (in Fig. 10) or sr + sb. The
tangential 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 setup, 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) corresponds 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 Lb
between the trench and the extinct back-arc spreading centre, keeping the interplate boundary conditions 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 overriding plate is smaller as well. The failure of this plate
occurs after greater thinning, but in the same direction. 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
149
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 lithosphere (variation of these parameters does not affect
the lithosphere failure direction). The arc is isostatically 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 overriding plate is not sensitive to plate structure. In all the
following tests, we use a simple homogeneous structure 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 experiments 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 experiment (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 sb
values, 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.
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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.
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 subduction
zones is estimated to be a few tens of megapascals
(Maekawa et al., 1993; Tichelaar and Ruff, 1993;
A.I. Chemenda et al. / Tectonophysics 342 (2001) 137–161
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 subduction 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 subduction 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 opposite 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 interplate surface (Saint Venant principle) and is largely
controlled by the trench/back-arc spreading axis distance 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
151
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. Preliminary 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 thickness 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 (arccontinent collision) because of the increase in horizontal compression of this lithosphere during subduction 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; Nakamura 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
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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 asthenospheric currents or plate kinematics. Even if compression 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 asthenospheric 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 principal subduction. It would be interesting to look for
evidences of a complete or partial subduction (disappearance) 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 intensively 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
Métour 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; Goffé 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 comprises 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 Métour 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 Métour 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 northeastward from blueschist carpholite facies (8– 12 kb,
280– 320C: Goffé et al., 1988; Michard et al., 1994)
to eclogite facies at As Sifah (20 – 23 kb, 540: Searle
et al., 1994; Wendt et al., 1993; Searle and Cox, 1999).
A.I. Chemenda et al. / Tectonophysics 342 (2001) 137–161
153
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.
The age of the HP metamorphism is poorly constrained: various K/Ar and 40 Ar/39Ar ages range
between 130 and 70 Ma (Montigny et al., 1988; ElShazly and Lanphere, 1992). Based on stratigraphic
and paleogeographic constraints, most authors consider that the Saih Hatat parautochton did not ex-
perienced HP metamorphism before late Cretaceous
time (Michard et al., 1994; Le Métour 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 –Santonian (87 – 83 Ma) up to lower Campanian (83 – 77
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A.I. Chemenda et al. / Tectonophysics 342 (2001) 137–161
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 thermochronology: 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; Béchennec et al.,
1990; Le Métour 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 (Béchennec et al.,
1988, 1990). The youngest sediments tectonically
incorporated into the Hawasina and Sumeini nappes
are middle Turonian to Coniacian (91 – 86 Ma) (Le
Métour et al., 1995). The Hawasina units are unmetamorphosed except at the northwestern border of the
Saih Hatat window where they contain HP/LT paragenesis (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 contact 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 (Lippard et al., 1986; Nicolas, 1989). Two major magmatic 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 (‘‘Geotimes unit’’ after Alabaster et al., 1982) is uppermost 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 Turonian. 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; Coleman, 1981; Juteau et al., 1988; Ernewein et al.,
1988). Following Ernewein et al. (1988), the first
episode of magmatism corresponds to typical midocean 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 evolutionary 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
A.I. Chemenda et al. / Tectonophysics 342 (2001) 137–161
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 overriding plate which in presently active subduction
zones undergoes intense deformation. This deformation 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 obduction 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 conclusion 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-ofsequence 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
155
(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 presented 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 intraoceanic subduction changed to a compressional
regime by 94 Ma. The reason for this change is
not known. It may have been the general reorganization 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 NEdipping 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 backarc lithosphere and in the others old (Triassic according to Searle and Cox, 1999) oceanic crust. Underthrusting 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. Overthrusting of the Oman ophiolite was restarted (reactivated) 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
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A.I. Chemenda et al. / Tectonophysics 342 (2001) 137–161
157
Fig. 16. Ridge origin of the Samail ophiolite: alternative to Fig. 15a initial setting of plate convergence at 95 Ma (notations in Fig. 15).
partially accreted to the Samail ophiolite ( = unmetamorphosed part of the Haybi Complex?). The Arabian 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 subduction (the pull force) and rapid isostatic uplift of
the orogen. This uplift resulted in the almost complete 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. 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) Breakoff 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.
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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 Himalayas (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 backarc lithosphere the weakest zone is the volcanic arc
area. When weakening becomes sufficient, the lithosphere 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 subduction 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 continental margin has arrived at the trench. During
continental margin subduction the tectonic compression 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 lithospheric 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
A.I. Chemenda et al. / Tectonophysics 342 (2001) 137–161
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 metamorphosed and exhumed in the presence of a subducted 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 INSUCNRS program ‘‘Intérieur de la Terre’’ (Contribution
No. 283) and is Geosciences Azur contribution No.
392.
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