dynamics of continental collision: influence of the plate ... · dynamics of continental collision:...

Geophys. J. Int. (2008) doi: 10.1111/j.1365-246X.2008.03857.x

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Dynamics of continental collision: influence of the plate contact

Roberta De Franco, Rob Govers and Rinus WortelFaculty of Geosciences, Utrecht University, Utrecht, the Netherlands. E-mail: [email protected]

Accepted 2008 May 19. Received 2008 May 19; in original form 2007 October 11

S U M M A R YObservations shows that continental collision may evolve in different ways, resulting in a widerange of tectonic responses. In search of the controlling conditions and parameters, we startfrom the results of our previous work, which demonstrated that the properties of the platecontact are important for the overall dynamics of convergent plate margins. Two fundamentaltypes of subduction plate contact can be distinguished: one based on a fault and the other basedon a weak subduction channel. In this study, we investigate how the plate contact affects theinitial stage of continental collision. We use a finite element method to solve the heat and thetime-dependent momentum equations for elastic, (power-law) viscous and plastic rheologies.For the same rheological properties and driving forces, varying the nature of the plate contactleads to three types of responses. The presence of a subduction channel promotes coherentand, when the boundary conditions allow it, plate-like subduction of the continental margin. Inmodels with a subduction fault, coherent subduction of the incoming continental lithosphereoccurs when the colliding passive margin has a gentle slope. The approaching continental sliverstarts to subduct and the subduction is characterized by a non-plate like behaviour—slowersubduction velocity than in channel models and strong slab deformation. If the continentalmargin is steep and the strength of the incoming continental crust is high, fault models resultin locking of the trench, eventually leading to slab break-off. If the crustal strength is relativelylow, shear delamination of part of the crust is expected. In the channel model, this type ofdelamination never occurs. The tectonic settings used in our experiments (prescribed platevelocity of the subducting plate versus fixed subducting plate corresponding to a landlockedbasin setting) do not significantly influence the nature of the model response. We conclude thatinitial stages of continental collision are strongly affected by whether the subduction contact isa fault or a channel. Neither the slab pull magnitude nor the tectonic setting is very importantto the overall geodynamics at this stage. The plate contact type, along with the slope of theincoming passive margin and the rheology of the continent, controls whether the incomingcrust (1) subducts entirely; (2) separates partially or entirely from the lithospheric mantle or(3) blocks the trench, likely leading to slab break-off.

Key words: Numerical solutions; Subduction zone processes; Continental margins:convergent.

1 I N T RO D U C T I O N

Continental collision, following oceanic subduction, plays a keyrole in connecting plate motion and orogenic events. In general,due to its thick and light crust, continental lithosphere is positivelybuoyant and will resist subduction, whereas cold and dense sub-ducted oceanic lithosphere generates a large downwards force (slabpull). Consequently, continental crust is subducted with difficulty, ifat all, whereas most oceanic crust is subducted easily at an oceanictrench. In a subduction process, the collision between a continentalfragment or terrane and a continental overriding plate may result indifferent modes: (1) the incoming continental material arrives at thetrench and is subducted; (2) all or part of the continental buoyantcrust is separated from the lithospheric mantle and is accreted to the

overriding plate, allowing continuation of plate convergence; (3) theapproaching continental margin locks the trench, eventually result-ing in slab break-off (slab detachment). All these processes appearto be possible, and which of these scenarios takes place or has takenplace in a certain region, depends on several factors: large scaletectonic setting of the two plates involved and more local factorssuch as the rheological parameters, the geometry of the continentalmargin and the buoyancy.

Subduction of continental crust, together with the lithosphericmantle, results in a positive buoyancy force opposing the negativebuoyancy force of the dense mantle part of the lithosphere. Nev-ertheless, deep subduction of continental crust is possible if, forexample, the crust becomes denser because of phase transforma-tions (Austrheim 1987; Le Pichon et al. 1992; Dewey et al. 1993). In

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2 R. De Franco, R. Govers and R. Wortel

nature, there is ample evidence that continental crust may have ex-perienced pressure of about 30–35 kbar. The evidence comes fromultra-high pressure metamorphic minerals (e.g. coesite, diamond)revealing subduction of continental crust to mantle depth and subse-quent exhumation (e.g. Smith 1984; Liou et al. 2000; Chopin 2001;Yang et al. 2003). Numerical studies supported the idea that the up-per crust can reach a burial depth varying between 50 and 450 kmif the crust is tightly welded to the subducting mantle by a strongrheology and if slab break-off does not occur (Ranalli et al. 2000;Regard et al. 2003; Toussaint et al. 2004). Crustal subduction isalso affected by other parameters—plate velocity, thermorheologi-cal state of the lithosphere, collision zone geometry, shear heating atthe plate interface, partial melting in the subduction channel, com-position and erosion/sedimentation processes (Gerya & Yuen 2003;Gerya et al. 2004b; Gerya & Stockhert 2006; Burov & Yamato 2007;Gerya & Burg 2007; Faccenda et al. 2007; Massonne et al. 2007).Another mechanism influencing subduction of the continental crustis the interplate pressure between the overriding and the subductingplate. Chemenda et al. (1996) showed that a low interplate pressureallows the continental crust to subduct to a depth of 200 km withoutfailure. They found that the subducting continental crust reaches amaximum depth that is proportional to the strength of the crust andinversely proportional to the interplate pressure.

The whole crust, or part of it, may separate from the lithosphericmantle and remain at the surface (Cloos 1993; Kerr & Tarney 2005;Vos et al. 2007). This implies that when crust and mantle litho-sphere subduct, their speed and direction can be different. Thisprocess, known as delamination, was proposed originally by Bird(1978) to explain some of the tectonic features in the Himalayasand it is, indeed, a suitable mechanism to facilitate continuation ofsubduction. Delamination is only possible if there is a layer of weakmaterial within the continental crust, for example, the lower crust(Ranalli et al. 2000; Toussaint et al. 2004). Chemenda et al. (1996)showed that continental crust delamination occurs in front of thesubduction zone when the interplate pressure between overridingand subducting plate is high. For several regimes, progressive sepa-ration of the lithospheric mantle from the buoyant continental crusthas been proposed, for example, the Colorado Plateau (Bird 1979),the Barbados Ridge (Westbrook et al. 1988).

Another possibility is that the incoming continental material maylock the trench, possibly resulting in break-off of the slab if theslab is not sufficiently strong (Davies & von Blanckenburg 1995;Wong A Tong & Wortel 1997; Wortel & Spakman 2000) or aftera period of thermal relaxation (depending on slab age, subduc-tion velocity and interplate heating) for any initial strength of theslab (Boutelier et al. 2004; Gerya et al. 2004a; Andrews & Billen2007; Faccenda et al. 2007). Alternatively, the continental litho-sphere thickens until it drips into the deeper mantle as a resultof a Rayleigh–Taylor thermal instability (Houseman et al. 1981;Houseman & Molnar 1997; Pysklywec et al. 2000, 2002). Whetherand at which depth slab break-off occurs is a function of tempera-ture and convergence velocity (Davies & von Blanckenburg 1995;Wong A Tong & Wortel 1997; Sobouti & Arkani-Hamed 2002;Toussaint et al. 2004). Strong slabs preserve their integrity againstthe increase of the integrated crustal buoyancy with the amount ofsubducted crust. The feasibility of slab detachment started to be in-vestigated in the early 1970s. Some studies suggested that the gapsin seismic activity as a function of depth might be explained by theexistence of detached slabs (Isacks & Molnar 1971; Barazangi et al.1973). Subsequently, seismic tomography allowed for the identifi-cation of real gaps in slab structure (e.g. Spakman 1990; Blanco& Spakman 1993). Slab break-off has been used to explain the

upper-mantle structure and tectonic deformation in many regions,for example, the Mediterranean–Carpathian region, the New He-brides and Turkey (Chatelain et al. 1992; Wortel & Spakman 1992,2000; Parlak et al. 2006; Lei & Zhao 2007). Understanding the fac-tors controlling the three different modes of continental collision isimportant for unravelling the tectonic evolution of convergent plateboundary regimes.

At convergent plate boundaries, the properties of the plate contactare important for the overall dynamics. Two fundamental physicalstates of the subduction contact can be distinguished: one based ona fault and the other based on a subduction channel. A subductionzone may evolve from one state to another, for example, through avariation in sediment supply or hydration (mainly serpentinization)of the overriding plate (Gerya et al. 2002; Stoeckhert & Gerya2005; Sobolev & Babeyko 2005; Babeyko & Sobolev 2007). Fornormal oceanic lithosphere subduction, De Franco et al. (2007)have illustrated that the type of plate contact (i.e. channel or fault)changes the way in which the stresses are transmitted from oneplate to the other and affects the velocity of subduction and themagnitude of the interplate pressure. The fact that these are alsotaken to be governing factors in the evolution of continental collision(see above) motivates us to assess for the first time the role of theplate contact in this process. Although numerous studies have beendone to understand the mechanisms governing the different modesof continental collision (e.g. Pysklywec et al. 2000, 2002; Burovet al. 2001; Toussaint & Burov 2004; Burg & Gerya 2005; Burov& Toussaint 2007), and fault models with various incoming slabgeometries, including ridges on the subducting plate, have beenstudied numerically (e.g. van Hunen et al. 2000, 2002, 2004), acomparison between the influence of the two different types ofplate contact is missing.

In this study, we show that critical differences develop duringinitial stages of the collision. Through numerical modelling we il-lustrate how, for the same rheological properties and driving forces,a variation of plate contact type results in different modes of conti-nental collision—subduction of the continental material or delam-ination or break-off. Here we focus on two tectonic settings. Inthe first scenario, the subducting plate is pushed into the subduc-tion zone by virtue of the plate tectonic setting (e.g. Himalayas).This is in contrast with the second setting of a land-locked basin,where there is no net convergence between the surface plates so thatsubduction must occur through roll-back (e.g. the Apennine andHellenic arc systems in the Mediterranean region).

2 N U M E R I C A L M E T H O D A N D M O D E LS E T U P

In this study, we do not make predictions for any real continentalsubduction zone, but we are interested in understanding the physicalprocess involved. For this reason, we choose to analyse a genericsubduction zone in which a passive margin arrives at the platecontact after the oceanic lithosphere has been subducted beneatha continental overriding plate. In our models, the subduction zoneis represented by a 2-D cross-section (Fig. 1). Although continen-tal collision processes have important 3-D features, the first-ordereffects of convergence can be appreciated by analysing a charac-teristic cross-section normal to the trench. With this simplification,we assume that the continent extends infinitely in the out-of-planedirection.

Because the numerical method and the model setup used inthis study are very similar to De Franco et al. (2007), we refer

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Dynamics of continental collision: influence of the plate contact 3

Figure 1. Model setup. On the left-hand side the continental plate with the continental crust on top, on the right-hand side the oceanic plate with the embeddedcontinental margin. Boundary conditions: rigid bottom; on the left-hand side: in the lithospheric mantle horizontal and vertical displacement are not allowed,in the sublithospheric mantle only horizontal displacement is allowed; on the right-hand side: in the lithospheric mantle horizontal and vertical displacementare not allowed or, alternatively a horizontal velocity of 5 cm yr−1 is applied. In the sublithospheric mantle only horizontal displacement is allowed. On top:free surface boundary conditions.

Table 1. Thermal (Ponko & Peacock 1995), elastic (Dziewonski & Anderson 1981) and creep model parameters. n is the stress power exponent.

Thermal properties Conductivity (k) Heat production (H) Specific heat (Cp)

Continental crust 2.8 W m−1 K−1 6.81 × 10−7 W m−3 1.24 × 103 J kg−1 K−1

Lithospheric mantle 3.14 W m−1 K−1 0 1.17 × 103 J kg−1 K−1

Sublithospheric mantle 3.14 W m−1 K−1 0 1.17 × 103 J kg−1 K−1

Elastic properties Young’s modulus (Y) Poisson’s modulus (ν)

Continental crust 5 × 1010 Pa 0.3Lithospheric mantle 1 × 1011 Pa 0.25Sublithospheric mantle 1 × 1011 Pa 0.25

Creep properties Pre-exponent (A) Activation energy (Q) Activation volume (V) n

Continental crust 8.3 × 10−26 Pa−n s−1 163 kJ/mole 17 × 10−6 m3 mole−1 3.1Mantle < 350 km 9.18 × 10−16 Pa−n s−1 485 kJ/mole 25 × 10−6 m3 mole−1 3.25Mantle > 350 km 1.54 × 10−11 Pa−n s−1 270 kJ/mole 6 × 10−6 m3 mole−1 1

to their study for more detailed aspects. The dynamics of the litho-sphere and upper mantle is governed by the momentum equation.Using the plane strain approximation, we solve this partial dif-ferential equation for instantaneous velocities and stresses. Weuse the G-TECTON finite element code to solve the momentumequation (http://www.geo.uu.nl/Research/Tectonophysics; Melosh& Raefsky 1983; Govers 1993; Buiter et al. 2001), and the steadystate diffusion-advection equation (Govers & Wortel 1993; DeFranco et al. 2007). Constitutive laws in the model represent elastic,viscous and plastic deformation. Viscosity is taken to be stronglytemperature-dependent, in accordance with rock mechanical experi-ments in both the power-law and diffusion creep regime (e.g. Kohlst-edt & Zimmerman 1996). Density also depends on temperature andis expressed by a linear equation of state in our model. These de-pendences make it necessary to also solve for model temperature(see De Franco et al. 2007). As the mechanical models of this paperfocus on short timescales relative to a thermal diffusion timescale,we ignore the temporal evolution of conduction. We do accountfor advective heat transport through the Lagrangian motion of thenumerical grid.

We adopt two different ages for the surface oceanic lithosphere atthe right-hand model boundary: one of 33 Myr and one of 25 Myr.Here, the initial geotherms are defined using a half-space cool-ing model. Using a spreading rate of the oceanic lithosphere of4 cm yr−1, the oceanic age at the trench is about 70 and 40 Myr,respectively. Temperatures in the continental overriding plate arebased on a representative steady state geotherm with a surface heatflow of 65 mW m−2 (see Table 1 and Ponko & Peacock 1995). In the

sublithospheric mantle, we use an adiabatic gradient of 0.3◦K km−1

(Ponko & Peacock 1995). The imposed surface temperature is 0◦Cfor both models. The initial temperature field for the first model isdisplayed in Fig. 2(a). Temperatures at the right- and left-hand sideof the model correspond to the two geotherms plotted in Fig. 2(b).These geotherms represent the side boundary conditions used tosolve the diffusion–advection equation.

Fig. 1 displays the model domain. It is characterized by a depthof 700 km and a horizontal extension of 2800 km. The subductingoceanic plate with a continental block embedded is on the right-hand side, the overriding plate is on the left-hand side. The incom-ing continental margin is located at the plate contact. We adopt twoinclination angles of 3◦ and 15◦ for the incoming margin continen-tal slope. These values are derived from our analysis of the currentglobal distribution of slope inclination angles of passive marginsrepresenting the most common and the maximum slope inclinationangle, respectively (see Fig. 3). The histogram is obtained fromthe analysis of a global 5-min latitude/longitude topography dataset (ETOPO5). The crustal thickness of the continental fragmentis assumed to be 26 km. The continent is in isostatic equilibrium.The shape of the curved portion of slab is defined by an errorfunction (Govers & Wortel 2005); we adopt a radius of curvatureR of 1.6 L (L = 80 km being the thickness of the oceanic litho-sphere in our models) and final dip angle θ of 45◦. The slab initiallyextends to 350 km depth. In some models, we also use a shorterslab that extends to 175 km depth. The overriding plate is 100 kmthick, which represents the lithospheric mechanical thickness de-fined by the viscosity structure, and has a crustal thickness of 35




Figure 2. (a) Total temperature distribution. (b) Geotherm at the sides of the model domain. Solid curve, temperature boundary condition at the right-handside of the model; dashed curve, temperature boundary condition at the left-hand side of the model.

Figure 3. Global distribution of inclination angles of continental slopes atpassive margins.

km. The finite element mesh is described by a non-uniform trian-gular grid, characterized by a maximum cell size of 500 km2 inthe asthenosphere and the lower mantle, 100 km2 in the mantlelithosphere and 50 km2 in the trench area.

The rheology of our model is elastic, viscous, or plastic and de-pends on composition, temperature, pressure and effective stress(Table 1). Model viscosities follow from either steady state dis-location creep or steady state diffusion creep. We choose modelparameters following Karato & Wu (1993) for the mantle, wherethe parameters are chosen to be intermediate between their wetand dry values since they concluded that the rheology of the uppermantle is between that of dry (water-free) and wet (water-saturated)olivine. For continental crust of both the overriding plate and thefragment, we use rheological parameters from Freed & Burgmann(2004) (see Table 1). Dislocation creep is assumed to prevail at adepth shallower than 350 km, whereas diffusion creep dominateswithin the transition zone (Karato & Wu 1993).

In the region of the mantle lithosphere and asthenosphere abovethe subducting slab (Fig. 1), the low viscosity mantle wedge (LVW)reduces the dynamic downwarping of the overriding plate in thearc/backarc region (Billen & Gurnis 2001). In our model, we as-sume a uniform viscosity of 5 × 1019 Pa s in the mantle wedge. Fol-lowing Billen & Gurnis (2001), this value has been chosen 10 timessmaller than the minimum asthenosphere viscosity. We apply anisotropic power-law Von Mises criterion to limit deviatoric stressesin accordance with Byerlee’s law (Byerlee 1978) using hydrostaticfluid pressure and horizontal compression. The lower limits of the

viscosity are 1e20 Pa s for the asthenosphere and the lower mantle,whereas in the mantle lithosphere we did not impose a lower limit.The upper limit is 1e24 Pa s and 1e21 for the mantle lithosphere andasthenosphere and lower mantle, respectively.

We use two descriptions for the active plate contact—a subduc-tion fault and a subduction channel—with the following character-istics:

(1) A subduction channel separates the subducting and the over-riding plate. The channel width is assumed to be approximately6 km, similar to what have been used in previous model studies(Shreve & Cloos 1986; Beaumont et al. 1999) and in agreementwith seismic studies that have shown the presence of interplate sed-imentary channel-like units of about 1–8 km (Eberhart-Phillips &Martin 1999; Tsuru et al. 2002; Oncken et al. 2003; Abers 2005).Channel viscosity is taken to be Newtonian, with a value of 7 × 1018

Pa s (England & Holland 1979; Shreve & Cloos 1986). We namethis class of models CM (channel models).

(2) A deformable subduction fault is described via updatedslippery nodes, in which the fault slip is locally kept parallel to thefault (Buiter et al. 2001). Fault friction is negligible. We name thisclass of models FM (fault models).

We separate model densities into a 1-D reference density profileand the remaining density anomalies (Govers & Wortel 2005). Ref-erence densities are used to initialize hydrostatic pressures, whichinitially do not contribute to the forcing. The remaining densityanomalies (i.e. slab pull forces) are used to drive model deforma-tion and are instantaneously applied at the beginning of the modelcalculation. Bending stresses would be expected to exist in the ini-tially curved portion of the slab. We do not include these initialbending stresses. However, bending stresses do develop as the slabdeforms during the model evolution. The continental fragment ispositively buoyant and its density anomaly varies between 100 and400 kg m−3 (Ranalli et al. 2000).

During the numerical experiments, we perform a re-meshingprocedure of the deformed grid. The reason is that in our La-grangian approach, the finite element mesh becomes so distortedthat integrals and derivatives needed in the finite element prob-lem cannot be properly approximated. Our re-meshing methodconsists of updating the connectivity between Lagrangian nodepositions. Taking particular care to preserve material boundariesand physical boundaries, we adopt a Delauney triangulation using




Figure 4. Surface expression for models with open subduction boundary conditions. (a) Horizontal velocity at the surface of the model for the entire domainat 0.56 Myr. Velocities in the yellow areas are directed towards the trench, at horizontal coordinate of about 100 km. (b) Ratio obtained by dividing the slabvelocity (V slab) by the velocity of the horizontal part of the oceanic plate (V horiz.) versus time. (c) The coloured circles indicate the region in which thevelocities used to calculate the ratio in (b) have been taken.

‘Triangle’ (Shewchuk 2002) to define new elements. Subsequentto mapping element quantities from the old to the new grid, wecompute (artificial) forces resulting from the inaccuracies of theprocedure and update stresses to reinforce mechanical balance. Thisprocedure results in very limited artificial diffusion of momentum(De Franco 2008). We perform the re-meshing every 5000 timesteps.

First, we concentrate on models in an ‘open’ tectonic settingin which we impose velocity boundary conditions of 5 cm yr−1

on the lithospheric right-hand side of the model, and we fix thelithosphere at the left-hand side of the overriding plate. The Hi-malayan plate boundary is a classical example for this type of set-ting. Second, we study models for a land-locked basin setting, anexample of which is the Mediterranean basin (Le Pichon 1982);a zero or very small relative motion between continents that sur-round an oceanic basin has as a consequence that basin-internalsubduction needs to be accompanied by backarc extension. Thus,our lithospheric boundary conditions do not allow relative motionbetween the far-field plates. We ignore convective motions beneaththe lithosphere that are not driven by the sinking slab or plate mo-tions, since our focus is on the subduction zone. This has oneimmediate consequence for the boundary conditions acting on thesublithospheric part of the side boundaries; these boundaries arefar enough from the central downwelling region to result in hori-zontal in and out-flow only beneath the lithosphere. We assume asignificantly higher viscosity in the lower mantle beneath the lowerdomain boundary (Mitrovica & Forte 1997); in our model, this isrepresented by no slip boundary conditions at the bottom of themodel.

3 R E S U LT S

In the experiments described below, we applied two types of bound-ary conditions. First, the open subduction setting where we imposea velocity of 5 cm yr−1 at the right-hand side of the subductingplate, and we fix the left-hand side of the overriding plate. Second,we use a land-locked basin setting in which there is no net conver-gence between the surface plates. We also investigate two differenttypes of geometries for the incoming continental sliver; one with acontinental slope inclination angle of 3◦ and a steeper one with acontinental slope inclination angle of 15◦.

The results are displayed in Figs 4–7. Figs 4 and 6 show the sur-face expressions of the models in the open subduction setting andland-locked basin setting, respectively. Figs 4(a) and 6(a) display thehorizontal velocity of the free surface for the entire model domainat 0.56 Myr. At this time, model initialization signatures have de-cayed. The trench is visible as a step change at horizontal coordinate100 km. Figs 4(b) and 6(b) show the evolution with time, of the ratioobtained dividing the velocity magnitude of the slab by that of thehorizontal part of the oceanic plate after the initial spin-up periodof the model. The ratio is one if the plate has a plate-like behaviour.A value greater than one indicates that the slab moves faster thanthe horizontal part of the plate, suggesting that plate thinning oc-curs somewhere in between the two points where the velocities aretaken (red and blue circles in Fig. 4c). When the ratio is less thanone, the slab moves slower than the surface plate, resulting in platethickening.

Figs 5(a1)–(a5) display the total strain rate distribution for theopen subduction model at 0.56 Myr, whereas Figs 5(b1)–(b5) show




Figure 5. (a) Effective strain rate distribution at 560 kyr. Only an enlarged part of the total domain is showed (see Fig. 1). (b) average effective strain rateevolution. Solid black curve corresponds to the region in the solid black circle in (a), red curve corresponds to the red circle in (a). (a1) and (b1) show the strainrate for CMO3, (a2) and (b2) show the strain rate for CMO15, (a3) and (b3) show the strain rate for FMO3, (a4) and (b4) for FMO15 and (a5) and (b5) forFMOw15.

the evolution of the strain rate with time in two areas—the solidblack line corresponds to the area in the black circle in Figs 5(a1)–(a5), the red curve corresponds to the strain rate in the crust ofthe incoming continental margin (red circle in Figs 5a1–a5). Thecurves represent the average strain rate value of the two regions.Figs 7(a) and (b) shows the horizontal stress at the surface of themodel domain for the open and land-locked basin setting.

3.1 Channel models (CM)

In this setup the plate contact is described by a low viscosity channelthat extends to a depth of 100 km. The viscosity of the channel isη = 2 × 1018 Pa s.

3.1.1 Channel model with open subduction boundary conditionsand 3◦ continental slope inclination angle (CMO3)

The horizontal velocity of the subducting plate varies from 5 cm yr−1

at the right-hand side to 5.6 cm yr−1 for the incoming continentalfragment near the trench (Fig. 4a black solid curve). This veloc-ity increment is likely due to the elastic response of the plate to

the positive buoyancy of the continental crust. Close to the platecontact, the overriding plate is characterized by a negative velocity,indicating the overriding plate retreats from the trench with a speedof about 5 mm yr−1. In Fig. 4(b), the velocity ratio is represented bythe black solid curve and is equal to a constant value close to one,denoting that the velocity along the subducting plate is uniform andit remains uniform as a function of time.

The strain rate distribution is displayed in Fig. 5(a1) at0.56 Myr. The strain rate in the slab and in the continentalcrust of the incoming fragment has a small magnitude of about1 × 10−15 s−1. In the entire model domain, the highest valueof the strain rate is in the subduction channel and below thehorizontal subducting plate. These two regions represent rela-tive thin layers of rapidly deforming material. In Fig. 5(a1), thedeformed geometry of the model shows that the continent hasbeen dragged down with the subducting plate to a depth of about30 km. In Fig. 5(b1), the average value of the strain rate in the slabshows a nearly stationary behaviour as a function of time. The strainrate is not affected by the subduction of the buoyant continentalcrust. On the other hand, the crustal strain rate slightly increases withtime.




Figure 5. (Continued.)

3.1.2 Channel model with open subduction boundary conditionsand 15◦ continental slope inclination angle (CMO15)

In this model, the continental slope of the margin is inclined atan angle of 15◦. The steeper geometry of the continental fragmentresults in a similar response of the model as CMO3.

The subducting plate moves with a speed of about 5 cm yr−1 and,close to the plate contact, increases to a value of about 5.6 cm yr−1

(Fig. 4a black dashed curve). The overriding plate velocity is nearlyzero, except close to the plate contact where the velocity is negative;this region of the overriding plate moves towards the left-hand sidewith a speed of about 4 mm yr−1.




Figure 6. Surface expression for models with land-locked subduction boundary conditions. Velocities in the yellow areas are directed towards the trench, athorizontal coordinate of about 100 km. (a) Horizontal velocity component (V x) at the surface of the model for the entire domain at 0.56 Myr. (b) Ratio obtainedby dividing the slab velocity magnitude (V slab) by the velocity magnitude of the horizontal part of the oceanic plate (V horiz.) versus time.

Figure 7. Horizontal stress (σ xx) at 0.56 Myr. Compression is negative. The overriding plate is divided in three regions: region C extends from the left-handedge of the plate to the vertically projected left-hand side of the LVW, region B extends from the vertically projected left-hand side of the LVW to the verticallyprojected intersection of the slab with the Moho of the overriding plate and region A extends from there to the plate contact. (a) Open subduction setting. (b)Land-locked basin subduction.

Like in the previous model, the velocity ratio is equal to one. Thismeans that the velocity along the subducting plate is uniform andthat this plate-like behaviour is not perturbed in time (Fig. 4b blackdashed curve).

Like in CMO3, the strain rate within the subducting plate is small(Fig. 5a2). Strain rate is high in the channel and beneath the oceanicsurface plate. The deformed geometry of the model displays that theincoming continental fragment is subducted to a depth of about 30km. In Fig. 5(b2) the average value of the strain rate shows a nearlystationary behaviour. This curve displays a response very similarto the previous model. The continental crust is characterized by aslightly higher strain rate than in the slab.

We performed an experiment in which the lower crust of thecontinental margin is characterized by a low strength of about 30MPa (CMOw15). The results are very similar to those of CMO15;for this reason, we decide not to show them.

3.1.3 Channel model with land-locked basin boundary conditions(CML3, CML15 )

Here we impose land-locked basin boundary conditions. We adopttwo inclination angles, 3◦ and 15◦, of the incoming continentalmargin.

In Fig. 6(a), the black solid line represents the velocity of CML3 at0.56 Myr. The speed of the oceanic plate increases linearly towards




the trench with a maximum value of about 2 cm yr−1. The overridingplate moves towards the right-hand side with a maximum speed of1 cm yr−1. The dashed black line shows the velocity of CML15. Likebefore, the velocity increases towards the plate contact. The velocitymagnitude is slightly smaller than in CML3. The overriding plate ischaracterized by the same behaviour as CML3.

In Fig. 6(b) the black solid curve and the black dashed curverepresent the velocity ratio of CML3 and CML15, respectively. Theyare characterized by a constant value equal to 1.3. Different fromCMO, velocities in CML are not plate-like.

The strain rate distribution for these models is very similar to thatof CMO3 and CMO15. Therefore, we will subsequently show theseresults in the Appendix.

3.2 Fault Models (FM)

In these models the plate contact is described by a frictionless faultthat extends to a depth of 100 km.

3.2.1 Fault model with open subduction boundary conditions and3◦ continental slope inclination angle (FMO3)

In Fig. 4(a), the grey dashed curve shows the horizontal velocity atthe surface at 0.56 Myr. The convergence velocity is 5 cm yr−1 andit decreases towards the plate contact, to increase again in corre-spondence of the continental fragment. The overriding plate, closeto the plate contact moves away from the trench with a maximumvalue of 1.5 cm yr−1. In this region the velocity is higher than in theCM.

The grey dashed curve in Fig. 4(b) shows the velocity ratio asfunction of time. This velocity ratio increases with time. This in-dicates that the horizontal part of the oceanic plate moves slowerthan part of the slab and that the velocity difference between thetwo regions enlarges with time. Downdip extension thus occurssomewhere within the region between the two points indicated inFig. 4(c).

Fig. 5(a3) displays the total strain rate distribution at 0.56 Myr.The strain rates in the slab, in the continental crust and in theoverriding plate are higher than those in the CM. The slab signif-icantly deforms at a depth of about 150 km, where the strain rateis localized. The deformed geometry of the domain shows that theincoming continental crust is subducted to a depth of about 15 km.The depth reached by the subducting continent is shallower than inthe CM. The overriding plate is characterized by high strain rateand marked downwelling. In Fig. 5(b3) the black and the red curvesshow that the average strain rate in the slab and the crust of the in-coming fragment increases with time, reaching higher values thanthe previous models. Strain rate is high beneath the oceanic surfaceplate.

3.2.2 Fault model with open subduction boundary conditions and15◦ continental slope inclination angle (FMO15)

The black dotted curve in Fig. 4(a) shows that the velocity jump atthe trench is strongly reduced with respect to FMO3. The horizontalvelocity shows that the speed of the subducting plate decreases fromthe right-hand side towards the plate contact to less than 4 cm yr−1.The overriding plate moves in the same direction as the subductingplate in the region between the plate contact and the horizontallyprojected tip of the slab end. The rate of retreat of the overriding

plate from the trench is faster than in all previous models. The restof the overriding plate has velocity magnitude equal to zero.

In Fig. 4(b) black dotted line, the velocity ratio increases withtime faster than in all the models so far. Such a behaviour showsthat the horizontal part of the oceanic plate moves slower than theslab.

The strain rate in the slab and in the overriding plate is higher thanin the CM and in FMO3 (Fig. 5a4). From the deformed geometry ofthe model, it is evident that the incoming continent locks the trenchand does not subduct with the oceanic material. The slab deformsmore than in FMO3 at a depth of about 150 km. The strain ratein the slab is higher than in FMO3. Fig. 5(b4) shows the averagestrain rate in more detail—the strain rate in the slab increases withtime more than in FMO3 reaching a maximum value of about 1 ×10−14 s−1. In the crust the strain rate is stationary and lower than inFMO3. Strain rate is high beneath the oceanic surface plate.

3.2.3 Fault model with open subduction boundary conditions, 15◦

continental slope inclination angle and weak crustal layer(FMOw15)

In this model, the strength of the weak crustal layer of the incom-ing continental fragment is 30 MPa, about 50 per cent lower thanFMO15. The subducting plate moves at a speed of 5 cm yr−1 andjumps to 2 cm yr−1 in the continental fragment (Fig. 4a solid greycurve). The part of the overriding plate that is closest to the platecontact moves towards the left-hand side, while the rest of the over-riding plate advances towards the plate contact. The continentalfragment moves with the same velocity and in the same directionas the near-trench part of the overriding plate. The velocity jumpalong the subducting plate is an expression of shear delaminationof the upper crust from the rest of the lithosphere.

The velocity ratio (grey solid curve in Fig. 4b) increases withtime similar to FMO3. However, it increases less than in FMO15,indicating that crustal delamination decreases downdip extension.

The strain rate distribution in the slab and in the overriding plateis very similar to the one of FMO3 (Fig. 5a5). The strain ratemaximizes in the slab at a depth of about 150 km; at that depth theslab is strongly deformed. A high strain rate region is visible in thecrust of the continental margin. This is the expression of the sheardelamination process, in which the upper crust separates from thelower one. The lower crust continues to subduct reaching a depth ofabout 15 km during the modelled period. The detailed representationof the strain rate in Fig. 5(b5) indicates that the average strain ratein the slab increases with time but less than in FMO15. In the crust,the strain rate is one order of magnitude higher than in the slab andhigher than in all the other models.

3.2.4 Fault models with land-locked basin boundary conditions(FML3, FML15, FMLw15)

The following experiments are characterized by land-locked basinboundary conditions. We adopt two inclination angles of the incom-ing continental margin of 3◦ and 15◦, FML3 and FML15 respectively.In FMLw15 (w=weak), we use the same geometry as in FML15, butthe strength at the transition between lower and upper crust of theincoming passive margin is lower, and its value is about 30 MPa.

In FML3, the surface velocity of the subducting plate is lower thanin either of the channel models (CML3 and CML15); it increasestowards the trench with a maximum value of 1 cm yr−1 (Fig. 6agrey dashed curve). In the overriding plate, the velocity increases




towards the trench with a maximum speed of 2 cm yr−1. The blackdotted curve shows the surface horizontal velocity of FML15. Thesubduction rate is a few mm yr−1 and it is smaller than in FML3. Theoverriding plate moves at a maximum speed of about 0.5 cm yr−1,which is slower than in all the previous models. The subductionvelocity of FMLw15 is a bit lower than the velocity of FML3 buthigher than in FML15 (Fig. 4a solid grey curve). The main differencebetween FML15 and FMLw15 is that the upper crust of FML15 hardlymoves, and that FMLw15 shows substantial movement towards theright-hand side. There is no relative motion across the trench inFMLw15, and the upper crust of the continental fragment movesalong with the crust of the overriding plate at a speed of 1 cm yr−1.

In Fig. 6(b) the grey dashed curve shows the velocity ratio ofFML3. The ratio increases with time. This means that the horizontalpart of the oceanic plate moves more slowly than the slab. Such adifference increases with time. The black dotted line, representsthe velocity ratio of FML15, whereas the grey solid curve is theratio of FMLw15. The FML15 velocity ratio increases faster thanin all other subduction models. The FLMw15 velocity ratio is onlyslightly higher than in FML3 ratio. In both cases, the horizontal partof the oceanic plate moves slower than the slab, indicating downdipextension of the subducting plate. The detailed description of thestrain rate distribution for these models is displayed in Appendix.We performed some tests with stronger rheology of the plates; sincethe main characteristics of our results do not change, we decidednot to show these results.

3.3 Short slab models (CMOs15, FMOs15, FMOws15)

In the next models, we want to estimate the importance of the slabpull in the evolution of our experiments. To reduce the slab pull, wediminish the length of the slab, that in the following experimentsreaches an initial depth of 175 km. In these experiments, we onlyuse the continental margin with slope inclination angle of 15◦. Weimpose a velocity of 5 cm yr−1 at the right-hand side lithosphericboundary of the models, and we fix the left-hand side of the over-riding plate.

The results are summarized in Fig. 4. In Fig. 4(a) the red dashedcurve represents the horizontal velocity for CMOs15 at 0.56 Myr.The only difference with respect to CMO15 is that the velocitymagnitude is slightly reduced. The velocity ratio of CMOs15 alsois similar to CMO15 (Fig. 4b). The velocity of the slab is initiallyslower than in the horizontal part of the plate (curve value lower thanone). With time, it increases until the ratio reaches a constant valueequal to one. This suggests that the continental margin is pushedcoherently into the subduction zone with the oceanic lithosphere ina steady plate-like fashion.

For the same type of geometry, we further reduce the densityanomaly of the slab (we do not show these curves in Fig. 4), andeven in this case, the behaviour of the model does not change much.In the extreme case, in which the slab pull due to the negativelybuoyant oceanic lithosphere is zero, the subducting plate continuesto be driven into the mantle and slowly tends to reach the sametectonic state of the previous CM.

In the FM, reducing the slab length has a stronger impact than inthe CM. The blue dotted curve in Fig. 4 shows results of FMOs15;the horizontal velocity of the plate is lower than in FMO15. Thehorizontal velocity of FMOws15 (Fig. 4 blue curve) is higher thanin FMOs15 but slower than in FMOw15. Reducing the slab pull evenfurther does not lead to a further decrease of the surface velocity.The velocity ratio does, however, change in this case; a slab pull

results in a smaller slab velocity, making it more similar to thevelocity of the surface oceanic plate.

3.4 Horizontal surface stress

In this section, we summarize the horizontal stress response at thesurface of the previous models (Figs 7a and b) after 0.56 Myr. Awayfrom the plate boundary zone, there is an overall tendency for thefault models to be more compressive than the channel models. InCMO3 and CMO15, the subducting plate and the terrane are in ten-sion, whereas in FMO3 and FMOw15, subducting plate the stressstate is more variable; between x = 1600 and 350 km, the surfacestress increases from more compressive to somewhat tensile. Thestresses are compressive in the ocean plate near the terrane andtensile in the continental fragment itself. In FMO15, where the de-lamination of the upper crust does not develop, the stresses in theoverriding plate are strongly compressive. In Fig. 7, the overrid-ing plate is divided in three regions. Region C extends from theleft-hand edge of the plate to the vertically projected left-hand sideof the LVW, region B extends from the vertically projected left-hand side of the LVW to the vertically projected intersection of theslab with the Moho of the overriding plate; region A extends fromthere to the plate contact. Most of the channel models are largelyin tension in region C in the open subduction setting. Also, tensilestresses dominate in region C of FMO3 and FMOw15. In the sameregion, the stress of FMO15 is nearly zero, indicating a more com-pressive tectonic system. Region C in the land-locked basin settingis mostly in tension, irrespective of the plate contact type. In regionB, compressional regime prevails for all the models. Fault modelsin the open subduction setting tend to show more compression thanchannel models. In region A, horizontal stresses show a positive ex-cursion towards more tensile stresses in the fault models. Channelmodels in region A do not show such behaviour.

When we reduce the slab pull using the short slab model, the sub-ducting plate experiences less tension than in the long slab models(CMOs15 and CMO15 in Fig. 7a). Horizontal stresses in region Cshow to be rather insensitive to the slab pull for these models. Thestress change in region B of Fig. 7(a) is largely brought by the factthat the slab is shorter in CMOs15, with the consequence that theupward projection of the slab tip occurs closer to the plate contact.A reduction of the slab pull in the fault models results in a morecompressive oceanic subducting plate and region C (FMOws15).Stresses at the continental terrane becomes neutral. A simultaneousdecrease of slab pull and terrane crustal strength (FMOws15) re-sults in a similar stress in the surface oceanic plate as in the FMO15.Within region C, these models show approximately the same sur-face stresses too. Differences between FMO15 and FMOws15 aresubstantial only within region B.

In the land-locked basin setting, the horizontal stress along thesurface oceanic plate becomes more tensile than in the open subduc-tion setting, for both channel and fault models. The same holds forregion C of the overriding plate. The most relevant difference withrespect to the open subduction is in region B, where the compressiveregime is drastically reduced. Nevertheless, FM and in particularFML15 are characterized by the most compressive stress in that area.Table 2 summarizes all the models used in the experiments.

4 M O D E L A NA LY S I S

Our experiments are divided in two main categories: channel models(CM) and fault models (FM). In both groups, we vary some of the




Table 2. Summary of the models used in this study.

Model name Plate contact type Boundary conditions Continental slope angle

CMO3 Channel model Open subduction 3◦CMO15 Channel model Open subduction 15◦CML3 Channel model Land-locked basin 3◦CML15 Channel model Land-locked basin 15◦FMO3 Fault model Open subduction 3◦FMO15 Fault model Open subduction 15◦FMOw15 Fault model Open subduction 15◦, weak crustFML3 Fault model Land-locked basin 3◦FML15 Fault model Land-locked basin 15◦FMLw15 Fault model Land-locked basin 15◦, weak crustCMOs15 Channel model Open subduction, short slab 15◦FMOs15 Fault model Open subduction, short slab 15◦FMOsw15 Fault model Open subduction, short slab 15◦, weak crust

characteristic parameters: the geometry of the incoming continentalsliver, the tectonic setting, the rheology and the length of the sinkingoceanic lithosphere.

In a previous study, we established that the nature of plate con-tact affects the response of the subduction process (De Franco et al.2007). We found that CM are characterized by faster subductionvelocities, less compression in the overriding plate, less dynamicsubsidence of the upper plate, more plate-like behaviour and lowerinterplate pressure than FM. We refer to this paper for a detailedexplanation of the physical process that causes the differences be-tween subduction channel and fault. These results are also valid inthe present study.

In the CM of the present study, the very nature of the subduc-tion channel, with its weak material, reduces the coupling betweenthe plates and interplate pressure, promoting in this way conti-nental subduction, as already shown by Chemenda et al. (1996).As a consequence, the incoming continental fragment is subductedcoherently with the oceanic lithosphere without perturbing the sub-duction process, at least during the initial stage of the collision.This is expressed by the fact that in the CMO, the subductingplate shows a steady state plate-like behaviour (solid and dashedblack curves in Fig. 4). In CML, as a consequence of the bound-ary conditions, the slab moves faster than the horizontal part ofthe plate, but in such a way that the two velocities are constantlyproportional to each others. This results in the fact that in bothsettings, the continental sliver is coherently subducted without lo-calization of high strain rate in the slab (Figs 5a1 and a2 and A1a1and a2).

On the other hand, FM result in a very different response. First ofall, FM do not show plate-like behaviour for any kind of boundaryconditions, and the horizontal part of the subducting plate movesalways slower than part of the sinking slab (solid and dashed greycurves and black dotted curve in Fig. 4b); moreover, these differ-ences increase with time. Such a response implies that the platedeforms with time. In all FM, the strain rate strongly differs fromthe cases of the corresponding CM-series. The magnitude of thestrain rate is higher than in CM and localizes between −100 and−200 km depth. The slab visibly deforms (e.g. Figs 5a3, b3, a4,b4, a5, b5). These features are explained by two factors. First, thehigh interplate pressure and high normal coupling between lowerand upper plate in the FM hampers the subduction of the continent,whereas the gravitational force in the slab pulls the plate down,causing deformation of the slab. Second, the strong flow gener-ated in the wedge below the overriding plate (see De Franco et al.(2007)) contributes to deformation of the slab and promotes lo-

calization of strain rate. FMO15 shows the highest strain rate andthe most pronounced deformation of the slab. This occurs becausethe continental fragment, with the steep geometry and the high-strength crust, completely locks the trench, preventing subductionof the incoming continent. Since in FMO15 the strain rate in the slabincreases with time (Fig. 5b4 black curve) whereas the strain rate inthe crust is nearly constant (Fig. 5b4 red curve), slab break-off willpossibly take place.

A general result, true for all the tectonic settings, is that thegeometry of the passive continental margin of the terrane doesnot strongly affect the response of CM. On the other hand, theslope inclination angle of the incoming passive margin plays a veryimportant part in FM. Whereas in FMO3 and FML3, the continentstarts to subduct (Figs 4a and 6a dashed grey curve), in FMO15 andFML15, the subduction stops (see Fig.4a and 6a black dotted). Thesurface oceanic plate surface decreases with time. When the crustalrheology of the continental fragment permits shear delamination(i.e. FMOw15 and FMLw15), the response of the models becomessimilar to the one of FMO3 or FML3.

The progressive separation of the upper crust from the lowercrust occurs only in the FM when the crustal strength at the contactbetween lower and upper crust is low with a value of about 30 MPa(e.g. Figs 4a and 6a, solid grey curve, and Figs 5a5 and A1a5). Theupper crust moves with a different velocity and sometimes differentdirection than the lithospheric mantle. Through shear deformation,the upper crust remains at the surface, whereas the lithospheric man-tle continues to subduct. The strong overriding plate and the high in-terplate pressure hamper the subduction of the upper crust, whereasthe weak strength of the lower crust allows the separation of the twounits. When delamination occurs, the horizontal state of stress ofthe overriding plate changes—the far-field plate (Fig. 7a region C)as well as the region immediately close to the plate contact (Fig. 7aregion A) become tensile, whereas in FMO15 the state of stress inthe overriding plate is neutral or compressive. In the subductingplate, compressive stresses decrease as soon as delamination takesplace and subduction again starts to accommodate the convergence.In the land-locked basin setting, the state of stress of FML15 andFMLw15 is not much different since the total convergence is lim-ited by the boundary conditions. In CM, delamination of the crustdoes not take place, for any geometry and strength of the incomingcontinental material. This is due to the fact that the channel flowfacilitates subduction of the continental sliver and does not act as achisel on the incoming continental material. Moreover, the lower in-terplate pressure helps making subduction of the continental marginpossible.




0 1 2 3 4 5 60

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Figure 8. Qualitative–quantitative description of the influence of different parameters on the modes of continental collision. CM are represented by redsymbols. FM with 15◦ slope inclination angle are represented by black symbols. FM with 15◦ slope inclination angle and a weak lower crust are representedby green symbols. FM with 3◦ slope inclination angle are represented by blue symbols.

To understand the role of the slab pull during the initial stageof continental collision, we reduce the net pull force through thereduction of the slab length and keep the velocity on the right-handside of the subducting plate equal to 5 cm yr−1. In the CM, slabpull does not significantly affect the response of the model. In CMs,after a spin up period, the subduction velocity reaches the samevalue as in the model with a long slab (Fig. 4a red dashed curve).The velocity along the subducting plate becomes uniform with time(Fig. 4b red dashed curve), meaning that the incoming continentalsliver is subducted with the oceanic lithosphere without perturbingthe process. However, the horizontal stress at the surface domainbecomes more compressive in the subducting plate and in region Cof the overriding plate, whereas in region B it is less compressive(see Fig. 7a). The higher compressive stress in region C of theupper plate and in the subducting plate is caused by the reductionof the slab pull that increases the inter plate pressure. The lowercompressive stress in region B is due to the fact the tip of the slabdoes not lie below that region anymore, causing downwelling andcompression in that part of the overriding plate. The slab velocityis fast, like in the long slab experiments, because the horizontalvelocity applied on the right-hand side of the subducting plate isthe same as in the long slab experiments and is strongly transmittedalong the plate. The low viscosity channel reduces the transmissionof such a velocity to the overriding plate. Moreover, the short lengthof the slab reduces the frictional forces in the mantle that opposesthe sinking of the slab.

In FM with short slabs (FMs), the subduction velocity becomesslower than in the other FM, and the overriding and subducting plateare in a more compressive regime. This is due to the fact that thehorizontal velocity applied at the right-hand side of the subductingplate is strongly transmitted through the fault to the overridingplate and that the interplate pressure increases (interplate pressureis inversely proportional to the slab pull). In FMOs15, where the

delamination of the crust of the continental margin does not develop,the subduction velocity is reduced even more than in the FMO15.As a consequence, even though the net slab pull is less, the slabmoves faster than the horizontal part of the plate (Fig. 4b bluedotted curve). The same holds for the case in which delaminationdevelops (FMOsw15). We can conclude that a reduction of the slabpull does not affect the general response of the models.

In our models, we considered different tectonic settings in whichthe right-hand side of the subducting plate is pushed towards thetrench and the left-hand side of the overriding plate is fixed, or inwhich there is no net convergence between subducting and overrid-ing plate. In this way, we consider some important possible settingspresent in nature. From our results, we conclude that even thoughthe tectonic setting can change some of the features of the models,the imprint due to the different plate contacts remains the same.

The results of our models are plotted as data points in Fig. 8. Theblack curves that divide the panels in the different mode domainsare qualitatively drawn. In summary, on the basis of our numer-ical experiments, we support the initial idea that three modes ofcontinental collision can develop:

(1) Coherent subduction of the continental sliver. This mode issubdivided in plate-like and non-plate-like continental subduction.

(2) Crustal delamination of the upper crust of the incomingcontinental margin. Subsequent continuation of subduction in anon-plate-like fashion.

(3) Locking of the subduction zone and possibly leading to slabbreak-off.

The three modes are illustrated in Fig. 9. In the upper panel,CM3/CM15 are characterized by plate-like subduction of the in-coming continental margin for any type of geometry and strengthof the continental fragment; in FM3 the continental material is co-herently subducted but the velocity along the slab is not plate-like




Figure 9. Illustration of the modes of continental collision: CM always allows coherent subduction. FM3 is characterized by coherent and non-plate-likesubduction, FM15 is characterized by subduction locking, FMw15 results in shear delamination. Arrows indicate velocity. The thicker crust in the subductingplate represents the embedded continental fragment.

and the slab deforms; in FM15 the passive margin locks the trenchand does not subduct, causing even more pronounced deformationof the slab; in FMw15 the upper crust is separated from the lowercrust in a shear mode and it remains at the surface, whereas thelithospheric mantle descends in the asthenosphere in a non-plate-like fashion.

We conclude that the plate contact type is critical in controllingthe different modes of continental collision during the initial stageof the process. The plate contact plays a more relevant role than themagnitude of slab pull and than the tectonic setting.

5 A L I N K B E T W E E N S H O RT A N D L O N GT I M E S C A L E C O N T I N E N TA LC O L L I S I O N M O D E L S

In previous studies, buoyancy (Ranalli et al. 2000; Sobouti &Arkani-Hamed 2002), rheological properties of the continental andoceanic material (strength, temperature; Wong A Tong & Wortel1997; van de Zedde & Wortel 2001; Sobouti & Arkani-Hamed2002; Li & Liao 2002; Toussaint et al. 2004), subduction velocity(Davies & von Blanckenburg 1995; Li & Liao 2002; Toussaint et al.2004; Burov & Yamato 2007) and interplate pressure (Chemendaet al. 1996) have been indicated as the main variables controllingthe evolution of continental collision. In our study, we took a stepin direction of establishing which are the parameters that dominatethe behaviour of continental collision during the initial stage ofthe process, and we do not study the entire evolution of it. Thereare two main differences between our study and the previous ones:(1) the earlier studies did not take in consideration the nature ofthe plate contact, and (2) we do not analyse long timescale pro-cesses. Faccenda et al. (2007) discussed the importance of havinga weak rheological coupling between the plates to obtain one-sidedasymmetrical subduction zone during continental collision. Theychanged the coupling between the plates increasing or decreasingthe shear heating (that is governed by the convergence velocity) atthe plate contact. In our models, however, shear coupling in bothchannel and fault models is very low or negligible. The fault andchannel represent two physical different states of the subductionplate contact and are characterized by different properties (for moredetails see De Franco et al. 2007).

It is likely that some of the parameters that do not play an impor-tant role in our experiments, become relevant during a second stageof the process developing. In the following sections, we make someconnections between our findings and the results of other studies.

5.1 Deep burial of continental material and slab break-off

In our CMO, continental lithosphere starts to subduct in a coherentand plate-like fashion, indicating the possibility of deep (>30 km)subduction of the passive margin. Many numerical studies and ge-ological data suggest that the entire continental lithosphere may betaken to even greater depths.

Ranalli et al. (2000) concluded that subduction stops when thetotal slab buoyancy, calculated from the summation of positive andnegative density anomalies, is zero. The maximum depth reachedby the continental material is a function of dip angle, geometryof sliver and convergence velocity. If the imposed velocity at theedge of the subducting plate continues to drive the plate, continen-tal subduction can also carry on when the system is neutrally orslightly positively buoyant, as shown in our short slab experiments.This is in agreement with Regard et al. (2003), who suggested thatsubduction will terminate when the subducting system reaches asignificant positive buoyancy. Faccenda et al. (2007) also foundthat, if the plates are efficiently rheologically decoupled (analogueto our channel models), the slab continues to subduct even withoutexternal forcing due to the slab pull and the scrapping off of thebuoyant upper crust. These findings indicate that a small piece ofcontinental crust can be taken to considerable depth, and that buoy-ancy is not very important until it reaches a certain threshold value.As a consequence, the amount of positive buoyancy of the conti-nental crust is not a dominant parameter during the initial stage ofcollision. In our model, the width of the incoming continent is notcrucial, whereas on a long timescale the volume of the subductedcontinental fragment controls the total amount of buoyancy. Thedepth to which continental crust can be subducted, coherently de-creases as the crustal thickness of the subducting continental plateincreases (van den Beukel 1992).

On the other hand, collision modes is conditioned by initial con-vergence velocity and the thermorheological state of the lithosphere(Burov & Yamato 2007; Toussaint et al. 2004). Burov & Yamato(2007) showed that continental subduction is possible only by strongand cold mantle lithosphere and fast initial convergence velocity.

Nevertheless, it is plausible that a relevant change in buoyancydue to the subducting crust may cause slab break-off. Eventuallyslab break-off is likely to occur in collision zones in which a pas-sive continental margin is involved, since the increasing tensionalstresses generated by continued continental subduction makes ex-tensional deformation inevitable and slab break-off probable. Toestimate where and when the break-off takes place, Davies & von




Blanckenburg (1995) compared the integrated strength of the slabwith the variation in buoyancy that results in extensional force.When this force overcomes the integrated strength, slab break-offtakes place. The extensional force is proportional to the subductiondepth of the continental crust and is controlled by its composition.They found that the integrated strength is strongly affected by theconvergence velocity. A slow subduction is likely to promote break-off and at shallower depth than for fast subduction; for a velocityof 1 cm yr−1 break-off might occur between 50 and 120 km depth.Slow subduction convergence promotes failure of the slab in mostof the continental collision models (e.g. Wong A Tong & Wortel1997; Sobouti & Arkani-Hamed 2002; Toussaint et al. 2004; Burov& Yamato 2007). Gerya et al. (2004a) showed that slab break-offcan be initiated in form of slab necking after cessation of active sub-duction. This supports our finding in which we found necking of theslab in our fault models, where subduction velocity is very small.The main difference between these two models is the timescale inwhich the phenomena takes place. In Gerya et al. (2004a) the neck-ing starts at least after 6 Myr after a period of thermal weakeningdue to thermal diffusion. In our models the slab starts necking af-ter 0.5–0.6 Myr and thermal diffusion is not active. The timescaledifference may be due to the fact that the forcings in the modelsare different. Slab necking is only present in our fault models wherethe strong flow generated in the wedge below the overriding plate(see De Franco et al. 2007) contributes to deformation of the slab.

In summary, we have learned from the previous numeri-cal studies that the parameters governing slab break-off are:strength/temperature of the slab; convergence velocity and buoy-ancy variation. In the light of our experiments another parametermust be added to this list: the nature of the plate contact.

5.2 Delamination

Another way in which subduction of the continental crust mayevolve is through delamination of the entire crust or part of it. Inour models, we excite delamination of the upper crust in that partof the model parameter space where the inclination angle of thecontinental slope is unusually steep, the strength of the lower crustis low and the plate contact is described by a fault. Delaminationoccurs through the removal of the upper crust in a shear mode; theupper crust remains at the surface and subsequently is accreted tothe upper plate. In nature, this process may take place in essentiallytwo ways: (1) by terrane or block accretion or (2) by thrusting ofthe delaminated upper crust over the adjacent margin of the upperplate. The latter may be accompanied by intense deformation of theover-thrust units (Willet et al. 1993).

We note, however, that there are other ways in which delaminationcan take place. Bird (1978) proposed that the previously subductedoceanic lithosphere with its excess density could result in a largetension in the upper slab if it was prevented from sinking at itsterminal velocity (Forsyth & Uyeda 1975). This force would tendto vertically separate the subcrustal lithosphere from the alreadysubducted crust. This kind of delamination may generate a so-callednappe stack that consists of a series of slices of continental crustleft at the plate contact by the incoming subducting plate.

The last process would require subduction of the continental crustto a certain depth like, for example, in our coherent subductionmodels (CM and FM3). Since from observations and model stud-ies we know that the upper crust may reach a considerable depth(�50 km) before being exhumed, we envisage that in models likeCM and FMO3 delamination may occur at a later stage than what

we show in our models. Chemenda et al. (1996) and Toussaint et al.(2004) showed the feasibility of this process. After a period of sub-duction, the upper crust may separate from the rest of the plate andbecause of its positive buoyancy, it may return to the surface. Thisprocess depends on the rheology of the lower crust, which has to berelatively weak, and of course on the buoyancy. Burov & Yamato(2007) also agreed with the fact that a strong lower crust and afast subduction velocity are necessary to drag the light upper crustto great depth. Ellis et al. (1999) investigated a small continen-tal terrane entering the subduction zone after a period of sedimentsubduction. In their model (their experiment 5), the plate contactis described by a subduction channel with an initial thickness of5 km. The continental terrane is dragged down beneath the upperplate. This finding is in agreement with our results, even though therheological properties and the forces used in their experiment differfrom ours. In their model, they investigate a larger amount of conver-gence than in our experiments: after 400 km convergence, the weakupper crust of the terrane separates from the lithosphere and createsa large fold nappe. Chemenda et al. (1996) introduced interplatepressure as a pertinent parameter in the process of continental crustsubduction. If the interplate pressure is low, the continental crustcan be subducted to a depth of about 200 km, before the delamina-tion process starts, and relatively easily exhumed (see also Boutelieret al. 2004; Gerya et al. 2007). When the interplate pressure is high,failure of the crust occurs at shallower depth and exhumation is notso easy and not so fast. This finding has some similarity with ourexperiments: in our FMOw15, where the interplate pressure is highand crustal strength weak, the continental crust fails, whereas thein CM, where the interplate pressure is low the crust is subducted.In FM, subduction of the upper crust is more difficult than in CMor does not take place at all. In conclusion, even though our mod-els differ from Chemenda et al.’s (1996) analogous experiments,the physical process behind them is very similar. Chemenda et al.(1996) vary the interplate pressure through the variation of the slabpull; in our modelling, this is achieved through the variation of theplate contact type.

The previous studies pointed out, which are the main parametersgoverning the evolution of continental subduction. The range inwhich these parameters vary is wide; continental collision can notbe explained by a unique combination of parameters. With ourmodels, we show that the initial phase of the process is importantand may control what will happen at longer timescales. The platecontact type is one of the variables to take into account. Duringthe initial stage of continental collision, the plate contact is morerelevant than slab strength, subduction velocity and buoyancy. In thelater stage of the process, these last variables become important.

6 D I S C U S S I O N

6.1 Do subduction channel and fault plate contacts exist?

One of the questions arising from our models is whether it is possi-ble to establish where, in nature, the plate contact is characterizedby a fault or a channel. Several seismic observations have shown alow velocity layer at the plate interface, which in general is few kmthick (Eberhart-Phillips & Martin 1999; Tsuru et al. 2002; Onckenet al. 2003; Abers 2005). This is interpreted in term of a subduct-ing sediment channel. However, a subduction channel cannot bea proper description of all subduction zones. First, great subduc-tion earthquakes suggest that the subduction interface cannot beweak everywhere (Davies & Brune 1971; Kanamori 1977; Ruff &Kanamori 1983; Tichelaar & Ruff 1993). Most interplate seismicity




occurs in the depth range of 0–50 km (Tichelaar & Ruff 1993), im-plying that at least the shallow part of the subduction zone has afinite shear strength. Second, sediment supply varies from site tosite, as indicated by accretionary wedge dimensions and erosivemargin observations (Clift & Vannucchi 2004). This might suggestthat when the amount of sediments is small, the subduction channelbecomes so thin that a more appropriate description of the platecontact is by one or several faults. In general, when a continentalmargin approaches the trench, a high quantity of sediments arrivesat the subduction zone, increasing the possibility of a channel-likeplate contact. On the other hand, erosion, which produces sedi-ments, is influenced by tectonics, topography and climate. Thesefactors change in time and as a consequence the amount of sed-iment changes, resulting in the possibility that subduction platecontacts evolve from one type into the other.

We could speculate that some tectonic settings may facilitate thecreation of a subduction channel, for example, in a land-lockedbasin setting in which it is easier to accumulate sediments. On theother hand, an open subduction setting, in which the subductingplate is pushed towards the trench, is likely to be represented bya fault. Unfortunately, we must conclude that for many subductionzones, relevant information on the nature of the plate contact islacking so far.

6.2 Comparison between our models and natural systems

The application of our numerical results to a natural system isrestricted by the simplifications we made in our models. However,these idealized convergent systems yield some expressions that aresimilar to those observed in some continental collision systems.For instance, examples of the type of delamination we produce inour models (block/terrane accretion) may be represented by theChugach terrane delaminated and accreted to the cratonic NorthAmerica (e.g. Coney et al. 1980; Clift et al. 2005) or the accretionof the Hanshan and Xilin Hot microcontinents, part of the Turkestanocean, to the Dongqiyishan arc. According to Kusky et al. (2007),this collision resulted in the formation of the North China Craton.An example of delamination and thrust accretion is the Qinling–Dabie Shan orogenic belt east of the Tanlu fault. The upper crustof the South China block thrust over the crust of the North Chinablock, followed by continuous subduction of the lower crust andupper mantle of the South China block underneath the North Chinablock (Li 1994). A similar process occurs in the Alps: the Europeanplate subducts, while the Adriatic plate acts as a chisel scraping offthe upper from the lower crust (Schmid et al. 1996).

In a general sense, understanding the possible variations in con-vergent plate boundaries evolution, as addressed in this study, helpsinterpreting observational data from such a setting. For example,in the nappe stacking of the continental basement of the Rhodope,there is ample evidence of ultra-high pressure metamorphism, sug-gesting that continental upper crust was dragged down to a depth of70-80 km before being exhumed (Liati et al. 2002) . At least 900 kmof continental lower crust and upper mantle have been subducted,without resulting in break-off (van Hinsbergen et al. 2005). This re-gion is characterized by extension in the backarc region. The CMLseem to be a good setting to generate a Rhodope-type geologicalstructure: the crust can be easily subducted and exhumed because ofthe low interplate pressure. The low interplate pressure helps bothprocesses and promotes continuation of subduction. Moreover, inthe land locked basin setting, the state of stress in the backarc regionis compressive, but the compression is strongly reduced comparedwith the FM. We argue that an increase in channel thickness or a de-

crease in the viscosity value in the channel may result in extensionin the backarc region.

The fact that most of the above cited regions are characterizedby the presence of ultra-high pressure rocks—which generally in-dicates the presence of a subduction channel—may be explained bytaking into consideration that the nature of the plate contact maychange in time, going from a channel to a fault and vice versa.

Continental collision is highly 3-D: the size of the continent in theout-of-plane direction affects the response of the process. However,we can infer 3-D aspects from our 2-D models. For instance, ifin our models we predict subduction of the incoming continentalmargin or delamination of the upper crust, we can then infer thatif the out-of-plane dimension of the incoming continent is smallcompared with that of the subducting plate, the subduction of thebuoyant material or the removal of the upper crust is facilitated.

7 C O N C LU S I O N S

From a consistent comparison of channel and fault models, we find:

(i) Channel models always promote coherent and steady statesubduction of the continental fragment, irrespective of the geome-try and strength of the continental crust of the incoming continent.In fault models, coherent subduction of the incoming continentalmaterial occurs if the continental rise of the colliding terrane is gen-tle. In this case, the approaching continental sliver starts to subduct,and subduction is characterized by a non-plate-like behaviour andstrong deformation of the slab.

(ii) Fault models result in locking of the trench and likely insubsequent slab break-off, if the margin of the terrane is steep andif the strength of its lower crust is high.

(iii) The combination of a strong plate contact (i.e. fault), a steepslope angle and a low crustal strength of the incoming continentalmargin results in efficient shear delamination of the upper crust.The remaining lower crust and lithospheric mantle of the continentalfragment continue to subduct. The subducting plate is characterizedby a non-plate-like behaviour and strong deformation of the slab.

In summary, our results indicate that the following types of re-sponses can occur by solely varying the plate contact type:

(1) Coherent subduction of the continental sliver. This mode issubdivided in two groups: plate-like and non-plate-like continentalsubduction.

(2) Shear crustal delamination of the upper crust of the incom-ing continental margin; subsequent continuation of subduction in anon-plate-like fashion.

(3) Locking of the subduction zone and possible developmentof slab break-off.

We conclude that initial stages of continental collision arestrongly affected by whether the subduction contact is a fault ora channel. Neither the slab pull magnitude nor tectonic setting (inour case prescribed plate velocity of the subducting plate versusfixed subducting plate) is very important to the overall geodynam-ics at this stage. The plate contact type, along with the slope of theincoming passive margin and the rheology of the continent, decideswhether the incoming crust (partially) overthrusts, subducts entirelyor blocks the trench leading to slab break-off.

A C K N OW L E D G M E N T S

Partial support for RDF and computational resources for this workwere provided by the Netherlands Research Center for Integrated




Solid Earth Science (ISES). We would like to thank the anonymousreviewer and Taras Gerya for their many critical and constructivecomments and corrections. We also thank Anna Maria Marotta andSerge Lallemand for the very useful suggestions that helped toimprove the manuscript.

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A P P E N D I X A : M O D E L S F O R A L A N DL O C K E D B A S I N S E T T I N G

In this section, we show the strain rate distribution for the exper-iments in the land-locked basin setting in which there is no netconvergence between the surface plates. We again use two differentgeometries to describe the incoming continental sliver: one with acontinental slope inclination angle of 3◦ and a steeper one with acontinental slope inclination angle of 15◦. Figs A1(a1)–(a5) displaythe total strain rate distribution, whereas Figs A1(b1)–(b5) showthe evolution of the average strain rate with time in two areas: solidblack line corresponds to the solid circle in Figs A1(a1)–(a5), redcurve corresponds to the average crustal strain rate (red circle inFigs A1a1–a5.

A1 Channel models (CM)

In this setup the plate contact is described by a low viscosity channelthat extends to a depth of 100 km. The viscosity of the channel is η

= 2 × 1018 Pa s.

A.1.1 Channel model with land-locked basin subduction boundaryconditions and 3◦ continental slope inclination angle (CML3)

The strain rate distribution is displayed in Fig. A1(a1) at 0.64 Myr.The strain rate in the slab and in the continental crust of the in-coming fragment has a small magnitude of about 1 × 1015 s−1.The horizontal part of the subducting plate is characterized by ahigher strain rate than in CMO (channel models with open subduc-tion setting). In the entire model domain, the highest value of thestrain rate is in the subduction channel and below the horizontalsubducting plate. These two regions represent relative thin layers ofrapidly deforming material. In Fig. A1(a1), the deformed geometryof the model shows that the continent has been dragged down withthe subducting plate to a depth of about 12 km. In Fig. A1(b1), theaverage value of the strain rate shows a nearly stationary behaviouras a function of time with a value of about 1 × 10−15 s−1. The sameholds for the crustal strain rate. However, in this region the strainrate is higher than in the slab. This is a consequence of the fact thatthe crust deforms while entering the trench.

A.1.2 Channel model with land-locked basin subduction boundaryconditions and 15◦ continental slope inclination angle (CML15)

In this model, the continental slope of the margin is inclined at anangle of 15◦. The different geometry of the sliver slightly affects the

behaviour of the subduction process and the results are very similarto the previous model.

Like in CML3, the strain rate within the subducting plate issmall (Fig. A1a2). Strain rate is high in the channel and beneaththe oceanic surface plate. The deformed geometry of the modeldisplays that the incoming continental fragment has been subductedto a depth of about 12 km. In Fig. A1(b2) the average value of thestrain rate shows a nearly stationary behaviour. The black curve inFig. A1(b2) displays a response very similar to the previous model.In the crust, the strain rate magnitude is slightly higher than inthe slab. The fact that the average strain rate in the crust is higherthan in CML3 suggests that the continental crust experiences moredeformation than the in previous model.

We performed an experiment in which the lower crust of thecontinental margin is characterized by low strength of 30 MPa(CMLw15). Because the results are very similar to those the CML15,we decided not to show them.

A2 Fault Models (FM)

In this model the plate contact is described by a frictionless faultthat extends to a depth of 100 km.

A.2.1 Fault model with land-locked basin subduction boundaryconditions and 3◦ continental slope inclination angle (FML3)

Fig. A1(a3) displays the total strain rate distribution at 0.64 Myr.The strain rate that develops in the slab, in the continental crustand in the overriding plate is higher than that in the CML. The slabsignificantly deforms at a depth of about 150 km, where the strainrate is localized, but less than in FML3. The deformed geometry ofthe domain shows that the incoming continental crust is subductedto a depth of about 4 km. The depth reached within the integrationtime by the subducting continent is shallower than in the CML.The overriding plate is characterized by high strain rate and markeddownwelling. In Fig. A1(b3) the black and the red curves show thatthe average strain rate in the slab and in the crust of the incomingfragment increases with time, reaching higher values than in theprevious models. Strain rate is high beneath the oceanic surfaceplate.

A.2.2 Fault model with land-locked basin subduction boundaryconditions and 15◦ continental slope inclination angle (FML15)

The strain rate in the slab and in the overriding plate is higher thanin the CML and in FML3 (Fig. A1a4). From the deformed geometryof the model, we note that the incoming continent locks the trenchand does not subduct with the oceanic material. The slab deformsmore than in FML3 at a depth of about 150 km, but less than inFMO15. The strain rate is higher than in FML3. Fig. A1(b4) showsthe average strain rate in more detail: the strain rate in the slabincreases with time more than in FML3 reaching a maximum valueof about 0.6 × 10−14 s−1. In the crust the strain rate is stationary andlower than in FML3. Like in FML3, the strain rate is high beneaththe oceanic surface plate.

A.2.3 Fault model with land-locked basin subduction boundaryconditions and 15◦ continental slope inclination angle and weakcrustal layer (FMLw15)

The strain rate distribution that develops in the slab and in theoverriding plate is very similar to FML3 (Fig. A1a5). The strain




Figure A1. (a) Effective strain rate at 640 kyr. Only an enlarged part of the total domain is shown (see Fig. 1). (b) Effective strain rate evolution in time. Solidblack curve corresponds to the region in the solid black circle and red curve corresponds to the red circle in (a). (a1) and (b1) show strain rate for CML3, (a2)and (b2) show the strain rate for CML15, (a3) and (b3) show strain rate for FML3, (a4) and (b4) for FML15 and (a5) and (b5) for FMLw15.

rate maximizes in the slab at a depth of about 150 km; at that depththe slab is deformed. A high strain rate region is visible in thecrust of the continental margin. This is the expression of the sheardelamination process in which the upper crust separates from thelower one. The lower crust continues to subduct reaching a depth of

about 4 km, during the modelled period. The detailed representationof the strain rate in Fig. A1(b5) indicates that the average strain ratein the slab increases with time but less than in FML15. In the crust,the strain rate is one order of magnitude higher than in the slab andhigher than in all other land-locked basin models.




Figure A1. (Continued.)



dynamics of continental collision: influence of the plate ... · dynamics of continental collision:...

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