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Abstract

In this paper we examine the mechanical and penological interactions during formation and exhumation of ultrahigh pressure (UHP) terrains with the intention of producing a general three-dimensional dynamic model of continental subduction. We consider aspects of ancient and modern continental subduction to provide boundary conditions, rheological constraints and characteristic scales of time and space for the dynamic model in the generation and evolution of UHP during continental collision. The UHP and HP assemblages of the Western Gneiss of Norway provide rheological, geometric, and geochronological information for the modelling, while the active obliquely convergent plate boundary of central New Zealand serves as a modern analogue of the collision-subduction transition. The geodynamic models of oblique convergence, conditioned by these observations, provide orogen-wide velocity and strain rates and identify the characteristic length scales of strain partitioning within oblique subduction. The petro-structural evolution of individual packets within the oblique orogen are examined within a Lagrangian mixing model currently being developed that allows us to relate the large scale dynamic model to observations made at the level of an outcrop.

Introduction

In this paper we examine the mechanical and penological interactions during formation and exhumation of ultrahigh pressure (UHP) terrains with the intention of producing a general three-dimensional dynamic model of continental subduction. We consider aspects of ancient and modern continental subduction to provide boundary conditions, rheological constraints and characteristic scales of time and space for the dynamic model in the generation and evolution of UHP during continental collision. The UHP and HP assemblages of the Western Gneiss of Norway provide rheological, geometric, and geochronological information for the modelling, while the active obliquely convergent plate boundary of central New Zealand serves as a modern analogue of the collision-subduction transition. The geodynamic models of oblique convergence, conditioned by these observations, provide orogen-wide velocity and strain rates and identify the characteristic length scales of strain partitioning within oblique subduction. The petro-structural evolution of individual packets within the oblique orogen are examined within a Lagrangian mixing model currently being developed that allows us to relate the large scale dynamic model to observations made at the level of an outcrop.

Analytical and analogue models of continental subduction and ultrahigh pressure terrains identify a rather delicate force balance between buoyancy-generated body forces and viscous drag forces within a thickening crust and a downgoing slab. This balance is controlled by the strength of lithosphere materials that glue the heavy slab to the light crust and also produce a strong lid on top of the convergent zone (e.g. Chemenda et al., 1996; Ernst & Peacock, 1996; Davies & von Blackenburg, 1995). Factors affecting buoyancy are composition, pressure, temperature, the degree of metamorphic equilibration (Austrheim, 1998; Ryan, 2001) and the degree of asthenosphere involvement in the convergent zone. Material strength within the orogen depends upon the same factors and, in addition, is a function of local fluid pressure and erosion along the upper surface (Koons et al., 2002). Given a steady plate tectonic setting, i.e. a convergent or obliquely convergent plate boundary with a roughly constant ridge-orogen separation, the mechanical, thermal, structural, and petrological evolution of that margin is dominated by the time-dependent material behaviour of the rocks transferred into the margin. Another source of time-dependent velocity patterns arises from interaction of the lithospheric slab with upper and lower mantle flow (Funiciello et al., 2003a,b). Consequently, any of the questions that arise regarding the geodynamics of UHP terrains in convergent margins, relate to the interaction of metamorphic evolution of orogen material with mantle convection. Information on chemical and mechanical interaction of UHP evolution is therefore contained in the petrological and structural signatures of ancient orogens and the kinematics of modern orogens. In this respect, the inter-relationship of deformation and metamorphism is similar to that identified in subduction/accretionary complexes where links between intermediate depth seismicity and dehydration are well-established (e.g. Abers, 2000; Hacker et al., 2003a,b).

General mechanical considerations

Recognition of the widespread, if infrequent, occurrence of ultrahigh pressure assemblages indicating subduction of crustal material to depths of more than 100km has forced a paradigm shift in geodynamic understanding of convergent zones (e.g.Chopin, 1984; Smith, 1984; Hacker & Peacock, 1994; Carswell & Compagnoni, 2003). In order to model the particle paths of crustal material that exhibited UHP assemblages in dynamic models, it is necessary not only to include a large component of downward velocity associated with the downgoing slab, but also to consider a slab trajectory that is not necessarily parallel to the slab (e.g.Funiciello et al., 2003a,b). These two conditions introduce a dominant vertical velocity component to the orogen as opposed to a dominantly horizontal component of earlier models, and, more importantly, a component of divergence within the velocity field associated with volume increase. The requirement of continuity, or mass balance, in the mechanics of an orogen: 

formula
(in which ρ = mass density and v = the three-dimensional velocity field with components u, v, w in the Cartesian directions x, y, z, respectively) is therefore difficult to satisfy in the absence of synchronous contraction and expansion within the orogen. Consequently, both divergent and contractional velocity fields will coexist in an orogen cross-section with positive and negative dilatation rates: 
formula
in the absence of changes in the plate tectonic setting. In this paper, we consider the implications of this divergence to the petro-structural evolution of orogens, and we concentrate on those rock-related features that either influence or record the geodynamic evolution of continental subduction.

The gross mechanical problems associated with formation and preservation of UHP terrains are relatively well-defined (e.g. Molnar & Gray, 1979) and can be arbitrarily separated into the conditions controlling descent and those controlling exhumation of the buoyant crustal material.

The descent phase is capable of transporting relatively deep crustal material to depths of > 100km material within time frames on the order of 2 Ma. During this descent phase, the crustal material remains cool and strong and appropriate compositions can record the attainment of UHP conditions. This phase is driven by the sinking of the heavy lithosphere, dragging with it the crust that adheres to the slab. The amount and depth attained of crustal material is a function of the plate velocity, buoyancy and crustal strength, all of which are variable (Ryan, 2001).

During the exhumation phase UHP terrains detach from the heavy slab and are displaced upwards to regions within high-pressure (HP) metamorphic conditions. The cause of detachment can be related to crust-mantle delamination, to slab breakoff (Davies & von Blackenburg, 1995), or steady state channel flow (Cloos, 1982). The thermal signals of these three mechanisms are potentially quite different as discussed below, but the mechanical effect is similar. The detached block moves up to the base of the crust with a rate of vertical displacement dependent upon relative buoyancy, strength of the crustal lid and viscous forces applied to the base. Kinetics provides an important, potentially rate-limiting, step in detachment and some observations exist on relative reaction rates, fluid evolution, and deformation in HP and UHP terrains (e.g. Rubie, 1983; Austrheim, 1998; Ernst et al., 1998). Further vertical displacement and tectonic imbrication carries the previously subducted material into the upper crust. This stage of exhumation is generally associated with kinematic evidence for crustal extension.

Above, in the simplified treatment of the mechanical setting, we have arbitrarily separated descent and exhumation, implying a discreet break between the two regimes. However, under certain rheological conditions, these two regimes can be part of a continuous channel flow process (Cloos, 1982).

The balance between the plate tectonic forces is therefore maintained by the metamorphic behaviour of the rocks caught within the convergent zone. The determining variables in convergent dynamics thus include those related to relative, three-dimensional plate velocities as well as to the processes of metamorphic evolution. A measure of the tendency for delamination for very viscous systems is provided by the dimensionless Grashof number relating the ratio of buoyancy forces to viscous forces: 

formula
(L = characteristic length, g = gravitational acceleration, Δρ = density difference between lighter crust and heavier mantle and deep crust, μ = effective viscosity for a given strain rate). For Gr less than a critical value, Grcr, viscous drag dominates and crustal material can be subducted. If Gr > Grcr then buoyant crust will tend to separate from the heavy lithospheric slab.

An approximation of Grcr = 1 may be produced by comparing buoyancy stresses generated in continental subduction with effective viscosity predictions based upon experimentally constrained behaviour of crustal material: 

formula
Where Δσbuoyancy = ρgh; Δσviscous = flow stress at strain rates of 10−14 sec−1 for material exhibiting power law behaviour of the form: 
formula
(Brace & Kohlstedt, 1980; Ranalli, 1995).

Figure 1 illustrates the division between material likely to descend with the slab (Gr < 1) and the material tending to rise under buoyancy (Gr > 1). The specific evolutionary paths are related to thermal equilibration discussed in a later section. We emphasise that this dimensionless number is offered as a guide to the qualitative response of a buoyant/viscous system and does not consider the coupled dynamics of an orogen, which will certainly influence the local material velocities.

Fig. 1.

Diagram of the effective Grashof number representing buoyancy stress versus resisting stress for viscous crustal material. Critical Grashof (Grcr = 1) separates regions that tend to descend (subcritical) from those where buoyancy forces dominate (supercritical). The labelled trajectories indicate the evolution predicted for thermally activated materials (governed, for example by Eqn. 5) undergoing conductive heating after subduction (Eqn. 9; Fig. 7.). The plagio-clase and wet granite rheologies are calculated for strain rates of 10−14. The eclogite trajectory assumes only a buoyancy effect due to relevant phase changes.

Fig. 1.

Diagram of the effective Grashof number representing buoyancy stress versus resisting stress for viscous crustal material. Critical Grashof (Grcr = 1) separates regions that tend to descend (subcritical) from those where buoyancy forces dominate (supercritical). The labelled trajectories indicate the evolution predicted for thermally activated materials (governed, for example by Eqn. 5) undergoing conductive heating after subduction (Eqn. 9; Fig. 7.). The plagio-clase and wet granite rheologies are calculated for strain rates of 10−14. The eclogite trajectory assumes only a buoyancy effect due to relevant phase changes.

Mechanical model: Constraints from natural analogues

In order to reduce some of the variance in our model conditions, we have used specific ancient and modern natural analogues to provide initial velocities, geometries and scales. The Western Gneiss Region is an erosional window through a stack of overlying thrust nappes that were transported hundreds of kilometres onto the former craton of Baltica during the Late Silurian-Early Devonian Scandian continental collision. The Baltica basement and its tectonic cover were then ductilely deformed together in the later phases of this collision. The distinct rock types associated with these nappes make excellent stratigraphic and structural markers in an area where Baltica basement has experienced high and ultrahigh pressures during Scandian continental collision. Following evidence expanded on below, the Scandian orogen appears to have been deformed with distinct vertical partitioning of strain into dominantly contractional at UHP depth, synchronous with dominantly extensional strain in the mid and upper crust (Austrheim et al., 2003).

Although the general shape and conditions of the Laurentia-Baltica plate boundary during the Early Devonian are approximately known, additional kinematic and rheological observations available for modern convergent boundaries are unavailable for ancient orogens. The relevant observations include seismicity patterns, seismic velocities, relative plate vectors, surface strain field and gravity field that all provide potential constraints on a dynamic model. We have, therefore, chosen the accessible and relatively well-characterised Pacific-Australian plate boundary in the transition from oblique subduction to oblique collision of central New Zealand. Associated with the transition of convergent mode is also a transition from a high-temperature convergent zone in the north near Taupo, to a low-temperature, high-pressure continental collision in the south. This tectonic and thermal pattern association is analogous to the coeval Acadian/Baltica convergent margin of the Devonian.

Natural occurrences: Western Gneiss Region, Norway

The ancient analogue is located in the Western Gneiss Region (Fig. 2), which represents a window through tectonic cover nappes and exposes high-pressure (Griffin et al., 1985) and ultrahigh pressure metamorphic rocks (Smith, 1984; Dobrzhinetskaya et al., 1995; Wain, 1997; Wain et al., 2000; Terry et al., 2000b). The Western Gneiss Region is predominantly composed of Proterozoic allochthonous and para-autochthonous granitoid gneisses that contain eclogite (Gee et al., 1985). Temperature estimates of metamorphism from isolated crustal (in situ) eclogites in Proterozoic basement by Krogh & Carswell (1995) show a systematic increase in temperature from ~ 600 °C to ~ 800 °C (Fig. 3). Detailed mapping by Robinson (1995, 1997), and Terry & Robinson, (1998, 2003) indicates that numerous tectonic cover sequences have been tightly folded into the Baltica basement and that the basement is imbricated (Fig. 2).

Fig. 2.

(a)Generalised geological map showing narrow refolded synclines of the tectonic cover in Baltica crust and the location of the study area. (b) Generalised tectonostratigraphic map modified from Gee et al. (1985). Large arrow shows the orientation of a time averaged (450-425 Ma) relative motion vector for Baltica with respect to Laurentia, approximated from reconstructions of Torsvik (1998). Small arrows show movement vectors of allochthons. From Terry et al. (2000b).

Fig. 2.

(a)Generalised geological map showing narrow refolded synclines of the tectonic cover in Baltica crust and the location of the study area. (b) Generalised tectonostratigraphic map modified from Gee et al. (1985). Large arrow shows the orientation of a time averaged (450-425 Ma) relative motion vector for Baltica with respect to Laurentia, approximated from reconstructions of Torsvik (1998). Small arrows show movement vectors of allochthons. From Terry et al. (2000b).

Fig. 3.

Pressure-temperature-time-deformation histories for high- and ultrahigh-pressure rocks of the northern segment (Terry et al., 2000b).

Fig. 3.

Pressure-temperature-time-deformation histories for high- and ultrahigh-pressure rocks of the northern segment (Terry et al., 2000b).

Near the area of the transition to orogen parallel structures there are four regionally extensive thrust nappes overlying Baltica basement (Robinson, 1995; Lutro et al., 1997). The nappes, from bottom to top, include: 1) The Risberget Nappe, dominated by 1190 Ma rapakivi granite and subordinate metamorphosed gabbro, representing a slice of an unknown part of the Baltica basement. 2) The Sætra Nappe, deformed and metamorphosed Late Proterozoic feldspathic quartzite and amphibolite representing sandstone cut by diabase dikes. This represents a sedimentary assemblage with related dikes formed on the margin of Baltica during rifting that formed the Iapetus Ocean. 3) The Blåhø Nappe, metamorphosed, garnet-mica schist and amphibolite representing a volcanic arc assemblage formed and metamorphosed proximal to Baltica during the Early Ordovician or earlier. 4) The Støren Nappe dominated by low amphibolite facies metamorphosed mafic volcanic rocks with subordinate volcanogenic sedimentary rocks. The Støren Nappe and equivalents in the Trondheim region contain early Ordovician fossils of Laurentian affinity and are interpreted as an oceanic and island-arc assemblage produced within Iapetus or on its Laurentian margin, and thrust over Baltica during the Late Silurian. These nappes correlate directly with the Tånnås Augen Gneiss Nappe and the Särv Nappe of the middle allochthon of the Swedish Caledonides, and the Seve Nappe and the Köli Nappe of the upper allochthon of the Swedish Caledonides and the upper allochthon of the Trondheim basin, Norway. The underlying basement is dominated by 1680-1650 Ma granitoid gneisses cut by ~ 1500 Ma rapakivi granite (now augen gneiss) and by a variety of 1450-950 Ma gabbros, now variably eclogitised.

Tectonic framework

The Scandian orogeny, which resulted from collision between Baltica and Laurentia, includes emplacement of thrust nappes followed by and synchronous with subduction of Baltica beneath Laurentia, that resulted in production of HP and UHP eclogites and extensional orogenic collapse. Available paleomagnetic and paleogeographic data indicate that Baltica and Laurentia probably began to collide about 425 Ma with a latitudinal velocity vector of 8-10 cm per year, that resulted in oblique collision (Torsvik, 1998). Production of UHP rocks continued at least until 402 ± 2 Ma (R.D. Tucker in Lutro et al., 1997; Carswell et al., 2003). This was accompanied by rapid extensional exhumation that is divided into two phases that occurred between ca. 400 and 390 Ma. The first phase, associated with deformed ductile detachment faults in the hinterland, is interpreted to have been active during continued convergence of Baltica and Laurentia (Lutro et al., 1997) and to have continued until 395 Ma. The second phase was associated with the development of low-angle detachments that ultimately brought Devonian sedimentary rocks and underlying nappes into contact with the high-pressure Baltica basement (Andersen & Jamtveit, 1990; Andersen et al., 1991; Andersen et al., 1994).

Kinematic framework

Structural analysis by Terry (2000) and Terry & Robinson (2003) integrated with well-constrained P–T–t paths from UHP and HP rocks (Terry et al., 2000a,b) on Nordøyane allow development of a kinematic framework for the Baltica margin during the Scandian orogeny from eclogite to low amphibolite facies. The structural features identified are divided into two categories, 1) those that formed at depths greater than 65 km associated with top-SE shearing (contraction) and 2) those that formed at less than 45 km to less than 20 km associated with top-W or left-lateral shearing (extension).

The HP and UHP structures are seen where partitioning of late strain around metamorphosed gabbro and diorite gneiss in Baltica crust allows direct inferences to be made regarding the geometry of folds, kinematics and original structural orientations related to production and exhumation of HP rocks. In the diorite gneiss, the geometry of folds associated with eclogite facies fabrics is isoclinal to tubular with axes parallel to the trend of a stretching lineation. Results from strain estimates and the presence of L > S or L >> S fabrics indicate that these structures were formed in a constrictional strain field. Eclogite facies mylonite zones that locally have a minimum thickness of 40 m cut Proterozoic gabbro and adjacent gneiss. Removal of late folding by small-circle rotation indicates that these structures were formed during top-SE shearing (140°) during prograde metamorphism in a direction that was consistent with thrusting parallel to plate motion. These structures are interpreted to be segment(s) responsible for juxtaposing HP and UHP rocks at a minimum depth of 65 km.

The transition to late structural features is constrained to occur between 65 and 45 km depth by PT estimates of 780 °C and 1.3 GPa of well equilibrated mafic augen gneiss. The progression of these structures is also well preserved by strain partitioning. The earliest of these were extensional detachments juxtaposing eclogite facies rocks against overlying amphibolite facies rocks that show no evidence for eclogite facies metamorphism. These detachments are strongly overprinted and complexly folded, and they represent a phase of upper crustal extension that was active during continued convergence at deeper levels. Younger, more localised mylonite zones formed synchronously, by tubular, sheath, isoclinal, tight and open folding that shows a progression from WNW to ENE trends. The earliest mylonite zones, interpreted as originally subhorizontal, range in strike through a 20° angle from 110° to 90°. Later steeply dipping mylonite zones formed under lower amphibolite facies conditions, strike 75° and locally truncate earlier structures. The youngest mylonite zones formed at lowest amphibolite conditions, strike 50° and truncate all earlier structures. Folds developed during this progression show the range in orientation from WNW to ENE reflected in the orientations of the mylonite zones that is interpreted to represent progressive evolution during top-west shearing. These changes in orientation of the late structural features are from orogen-normal to orogen-parallel.

There are two major difficulties that must be overcome in order to apply the above interpretation to the scale of the orogen. The first deals with extrapolation of these structual features to the scale of the orogen. The HP structures show a transport direction that is normal to the strike of the metamorphic field gradient (Griffin et al., 1985), identical to the transport direction at the base of the Jotun Nappe (Fossen, 2000), and consistent with plate motion (Torsvik, 1998). The extensional structures dominate the Western Gneiss Region and the extensional fabrics change from WNW to EW in the area of the Jotun nappe and the Devonian sedimentary basins (Fossen, 2000; Fig. 2). Identical changes have been observed in syn-depositional deformational structures in the Devonian basins, which provide a connection to very shallow levels in the crust (Osmundsen et al., 1998). In the north part of the Western Gneiss region narrow synclines of the nappes can be seen in the basement, which show a progressive change from E-W to NE. Together the changes mimic the detailed changes in the extensional structures seen on Nordøyane. The structural trends and changes observed on Nordøyane that rocks experienced during exhumation are in agreement with structure seen on an orogen scale. Thus, we believe it is reasonable to extend our interpretation to provide a kinematic framework for the orogen.

The second problem is time. Obviously, applying the structures that developed along P–T–t paths relates structures that formed at different times. Therefore, it is necessary to show that both compression at depth and extension of the upper crust were operating synchronously. Recent geochronologic and structural data (Williams et al., 1999; Terry, 2000; Terry et al., 2000a,b; Carswell et al., 2003; Austrheim et al., 2003) indicate overlap between the deposition of clastic sediments in extensional basins (403-394 Ma) at high structural levels and prograde eclogite facies metamorphism (407 ± 2-402 ± 2 Ma) at greater depth. Austrheim et al. (2003) recently reported a U-Pb zircon age of 401 ± 2 Ma for pegmatites that are approximately normal to the late nearly orogen-parallel stretching direction. This age is indistinguishable from the age of the UHP metamorphism for the Ulsteinvik eclogite (402 ± 2 Ma) and indicates that syn-collisional exhumation involving upper crustal extension was probably an important mechanism for exhuming HP and UHP rocks.

With these problems considered, it is possible to assemble a generalised kinematic model at the time of UHP metamorphism which involves orogen-normal thrusting at depths below 65 km and orogen-normal extension at 45 km depth that evolve to orogen-parallel at < 20 depth km with an intervening transition zone of 20 km. This transition is well exposed in the area of Hustad but is poorly understood in terms of its detailed P–T–t–d evolution and the nature of strain partitioning. This kinematic framework provides a starting place to test a variety of different exhumation models for UHP rocks.

Observations of the modern analogue: Central New Zealand

Oblique continental subduction is currently occurring between the Pacific and Australian plates as a result of relative motion of ca. 39 mm yr−1 at an azimuth of ca. 255° (Fig. 4) (de Mets et al., 1994). The geometry and plate velocities along this boundary are used to constrain our dynamic model. Oblique convergence and a change in the composition of the Pacific plate from oceanic plateau in the north to continental in the south has led to the subduction of continental material beneath the northern South Island over at least the last 10 Ma (Fig. 4). South of the Hikurangi subduction margin, around Kaikoura, dramatic changes in the mechanics occur. The crust of the Hikurangi Plateau subducting beneath the North Island thickens from ca. 10 km in the north to ca. 15 km adjacent to the Chatham Rise from northeast to southwest (Davy & Wood, 1994). The Chatham Rise itself has a crustal thickness of 23-26 km. The Benioff zone of the Tonga-Kermadec-Hikurangi subduction zone ends abruptly at a line trending northwest through Kaikoura (Fig. 4). The southernmost edge of the subducted plate is identified by an abrupt step in the depth of the deepest earthquakes from 250 km to about 100 km (Fig. 4) (Anderson & Webb, 1994). This step is aligned with the 2000 m bathymetric contour at the transition from oceanic lithosphere of the Pacific plate to continental lithosphere of the Chatham Rise. On the basis of microearthquakes recorded during a detailed survey of the North Canterbury region, Reyners & Cowan (1993) suggest that the subducted slab continues farther to the southwest. They find that the dip of the seismic lower crust changes near 43°S from being subhorizontal in the southwest to a dip of ca. 10°NW in the northeast. It is unclear whether the lower crust near 43°S is flexed or torn (Reyners & Cowan, 1993).

Fig. 4.(a)

Digital Elevation Model (GEOgraphix) of the northern South Island showing the Marlborough Fault System, the Alpine Fault and the Kaikoura Ranges. Dashed grey line shows the edge of the Hikurangi subduction zone as recorded by an abrupt shallowing of the deepest earthquakes from 250 km to 100 km in depth (Anderson & Webb 1994). Inset shows the plate tectonic setting of central New Zealand, the 1000 and 2000 m bathymetric contours defining the transition from the oceanic Hikurangi Plateau to the continental Chatham Rise. Plate vector is from De Mets et al. (1994). (b) Schematic interpretation of the tectonics based on the 3D velocity model at the cross-section shown in (a). Modified from Eberhart-Phillips & Reyners (1998).

Fig. 4.(a)

Digital Elevation Model (GEOgraphix) of the northern South Island showing the Marlborough Fault System, the Alpine Fault and the Kaikoura Ranges. Dashed grey line shows the edge of the Hikurangi subduction zone as recorded by an abrupt shallowing of the deepest earthquakes from 250 km to 100 km in depth (Anderson & Webb 1994). Inset shows the plate tectonic setting of central New Zealand, the 1000 and 2000 m bathymetric contours defining the transition from the oceanic Hikurangi Plateau to the continental Chatham Rise. Plate vector is from De Mets et al. (1994). (b) Schematic interpretation of the tectonics based on the 3D velocity model at the cross-section shown in (a). Modified from Eberhart-Phillips & Reyners (1998).

Determination of the three-dimensional velocity structure to a depth of 100 km images the subducted slab between 40 and 100 km as a relatively low-velocity feature in the upper mantle (Eberhard-Phillips & Reyners, 1998) (Fig. 4). This is interpreted to reflect the continental nature of the subducting crust in this region. The amplitude of the low-velocity feature increases from northeast to southwest as the crust increases in thickness. Using existing calculations of thermal regime in a downgoing slab (Peacock, 1996; Hacker et al., 2003a,b), we calculate the predicted positions of coesite and jadeite stability beneath the northern part of the South Island (Fig. 4).

Above the zone of continental subduction is the Marlborough Region, a broad zone of diffuse deformation in the upper crust extending on the order of 200 km to either side of the Wairau Fault, the terrane boundary between Western province rocks of the Australian plate and the Torlesse terrane of the Pacific Plate. The style of late Cenozoic deformation differs across the Wairau Fault; in the west, Miocene to recent reverse faulting has resulted in crustal thickening and uplift; to the east, the dominantly strike-slip Marlborough Fault System overlies the southernmost edge of the Hikurangi subduction zone (Walcott, 1998).

Mechanical framework I: Orogen-scale dynamics

Our mechanical discussion follows both numerical and analytical approaches. The numerical approach consists of a three-dimensional dynamic model of continental subduction for a region with dimensions of 1000 km normal to the plate boundary, 600 km parallel to the plate boundary, and with a vertical extent of 200 km (Fig. 5). The resultant model provides three-dimensional velocity, strain rate, vorticity and dilatation rates, and stress fields for a variety of boundary and rheological conditions (e.g. Koons et al., 2002). Our solutions also track the thermal evolution of the problem domain, discussed below, to provide information on equilibrium and kinetic states in an externally fixed reference frame. In general, these preliminary numerical models serve as the vehicle for testing the sensitivity of continental subduction to variation in boundary and internal variables. Initial boundary conditions for the example presented here are characterised by a set of oblique far field velocities of lateral velocity: normal velocity (Vy : Vx) of 20 mm yr−1 : 10 mm yr−1 with differing vertical slab velocities imposed upon the base illustrated in Figure 5. These initial models have no along-strike variation in the applied boundary or rheological conditions.

Fig. 5.

Mechanical model of continental subduction. The block consists of three layers: 0–15 km depth consists of a pressure-sensitive Mohr-Coulomb material representing the upper crust; the lower crust, from 15–;40 km, consists of a pressure-insensitive, temperature-sensitive rheology based on wet quartzite; the upper mantle, from 40–;200 km, consists of a pressure-insensitive, temperature-sensitive rheology based on diabase. (a) Contours of velocity parallel to the plate margin (Vy). Dashed line shows position of cross sections below. (b) Contours of rotation. (c) Contours of shear strain rate. (d) Contours of lateral strain rate.

Fig. 5.

Mechanical model of continental subduction. The block consists of three layers: 0–15 km depth consists of a pressure-sensitive Mohr-Coulomb material representing the upper crust; the lower crust, from 15–;40 km, consists of a pressure-insensitive, temperature-sensitive rheology based on wet quartzite; the upper mantle, from 40–;200 km, consists of a pressure-insensitive, temperature-sensitive rheology based on diabase. (a) Contours of velocity parallel to the plate margin (Vy). Dashed line shows position of cross sections below. (b) Contours of rotation. (c) Contours of shear strain rate. (d) Contours of lateral strain rate.

Numerical solution of the motion and stress equations is based upon algorithms from ITASCA (FLAC3D; Cundall & Board, 1988), a three-dimensional finite difference code, which we have modified to accommodate large strains and local erosion. Materials are represented by polyhedral elements within a three-dimensional grid that uses an explicit, time-marching solution scheme and a form of dynamic relaxation. Each element behaves according to a prescribed linear or non-linear stress/strain law in response to applied forces or boundary restraints. The inertial terms in the equations of motion, 

formula
(Cauchy’s equations of motion where σijis the stress tensor, xi, vi are the vector components of position and velocity, respectively, ρ is the density of the material, [b] is the body force) are used as numerical means to reach the equilibrium state of the system under consideration. The resulting system of ordinary differential equations is then solved numerically using an explicit finite difference approach in time. The drawbacks of the explicit formulation (i.e., small timestep limitation and the question of required damping) are overcome by automatic inertia scaling and automatic damping that does not influence the mode of failure. The governing differential equations are solved alternately, with the output for the solutions of the equations of motion used as input to the constitutive equations for a progressive calculation. Solution is achieved by approximating first-order space and time derivatives of a variable using finite differences, assuming linear variations of the variable over finite space and time intervals, respectively. The continuous medium is replaced by a discrete equivalent with all forces involved, concentrated at the nodes of a three-dimensional mesh used in the medium representation.

Initially, we employ standard steady-state rheological models with pressure-dependent upper crustal rheology on top of temperature-dependent lower rheology for various compositions (e.g. Brace & Kohlstedt, 1980; Ranalli, 1995) (Fig. 5).

Numerical results

Application of the conditions discussed above produces the velocity and deformation rate fields of Figure 5. The incoming and sinking slab generates deformation in the orogen along the top of the slab and within the slab itself. The vertical velocity field in the orogen separates into material moving toward the surface and that entrained with the downgoing slab. Maximum downward vertical velocities within the slab are ~ 2.5 mm yr−;1, while maximum upward mass velocities attained at the upper free surface are ~ 1.5 mm yr. The petrological implications of this velocity structure are considered below.

Strain in the numerical models is certainly not homogeneous with clockwise vorticity about the y axis (ω = −0.5(∂Vx/∂z − ∂Vz/∂x), concentrated in the lower and descending crust along the top of the dipping slab down to depths of ~ 200km. Another region of clockwise vorticity exists in the eastern (x ≈ 850 km), upper crust representing the oblique thrust belts of the upper, outboard side of the orogen (Koons, 1990, 1994; Willett et al., 1993). Anti-clockwise vorticity is concentrated further to the west (x ≈ 600 km) in the upper crust where material is rapidly exhumed. The shape of this upper crustal vorticity concentration is strongly influenced by the surface erosion condition (Koons et al., 2003). Material passing from the east to the west through the orogen is subjected to both senses of vorticity and, consequently, should record an opposite sense of shear before it is exposed in the west. It should be noted that this shear sense reversal requires no change in the net plate vectors.

Shear strain rates in the x–z plane (∂ϵxz/∂t = ∂VZ/∂x + ∂Vx/∂z; Fig. 5b) are highest along a mid-crustal detachment and along the slab. Lateral strain rates (∂ϵyz/∂t = ∂Vy/∂x + ∂Vy/∂z; Fig. 5d) provide an indication of the location of strike-slip deformation within the orogen. For the rheological model that we have chosen in this initial model, in addition to a concentration of lateral deformation within the upper crust and along the slab, there is also strike-slip deformation along a vertical zone that slices through the slab at ~ 850 km. Partitioning of lateral velocity along this cross-cutting vertical structure reduces the degree of obliquity in the down going slab.

The dynamic model presented in Figure 5 captures some of the general features of continental subduction with UHP descent and exhumation. However, it fails to incorporate numerous, important features recognised in natural occurrences. The model’s utility lies primarily in being a vehicle for testing its response to perturbations and sensitivity to basic rheological and boundary assumptions including variation in surface erosional regimes. Most of these potential perturbations involve coupling among the processes of reaction and deformation related to both equilibrium and disequilibrium effects of reaction weakening/hardening behaviour, deformation enhanced reaction and the complex interplay of fluid production, migration and deformation.

Role of disequilibrium

Evidence for large departures from this steady-state rheological model, derived from HP and UHP models (e.g. Rubie, 1983; Koons et al., 1987; Austrheim, 1998), suggest a strong degree of strain-rheological coupling. Rubie (1983) suggested that strain-related reaction has led to super-plastic flow in HP terrains, thereby providing a positive feedback for rapid and large strains and is considered in later models.

Rate of metamorphic equilibration influences crustal strength (Rubie, 1983), crustal buoyancy (Ryan & Dewey, 1997), and the record of equilibration (Austrheim, 1998). To provide useful information on the dynamics of convergent margins, it is insufficient for assemblages to travel through PT space, they must also be capable of recording at least some locations on the path. It has long been recognised that the kinetic barriers to equilibration for dehydration reactions are relatively low and that those for relatively anhydrous rocks are high (e.g. Fyfe et al., 1958; Austrheim, 1998; Koons et al., 1987; Rubie, 1983). Seismic investigations of devolatilisation in active subduction zones reinforces this empirical kinetic model of near-equilibrium conditions for hydrated assemblages (e.g. Peacock, 1996; Abers, 2000; Hacker et al., 2003a,b). It is also clear that disequilibrium persists in polymetamorphic assemblages from mid- to deep crustal levels and that metamorphic reaction in these rocks is spatially and causatively linked to deformation and transient fluid pulses (e.g. Austrheim, 1998; Rubie, 1983).

Influence of surface processes

Although the paleoclimate of Baltica in the Devonian is not immediately obvious, the influence of erosion on orogen dynamics, as identified in both numerical models (e.g. Koons, 1990; Willett et al., 1993; Beaumont et al., 1992; Avouac & Burov, 1996; Blythe, 1998) and analogue models (Chemenda et al., 1996), is sufficiently important that we consider several different erosional schemes in our geodynamic models. Processes at the earth’s surface influence the large-scale orogen dynamics in several ways related to load development and rheological modification (Koons et al., 2002; 2003). Local relief similar to that of the Himalaya can produce stress perturbations within the upper crust that strengthen the regions beneath mountain loads relative to the valleys where significant shear stress concentrations occur (Koons et al., 2002).

Focussed exhumation due to either fluvial erosion (Zeitler et al., 2001a,b; Koons et al., 2002) or from an orographically generated erosional pattern significantly alter the integrated strength of an orogen. Over very long time frames, erosion can deplete the upper crust in radiogenic components, lower the geothermal gradient, and consequently strengthen the upper crust (e.g.Dewey et al., 1993). However, on the time scales of most orogenic processes (< 50 Ma) the effect of concentrated exhumation is to significantly weaken rather than strengthen the crust. Exhumation drives the reduction of the integrated strength of the crust (= Fc;Sonder & England, 1986) by thermal thinning due to advection of isotherms within a concentrated region (e.g.Allis et al., 1979; Koons 1987). The amount of advection-driven weakening for a thermally activated lower crust and Mohr-Coulomb upper crust can be approximated by: 

formula
(Koons et al., 2002) in which N is Terzaghi’s (1943) flow value 
formula
for a Mohr-Coulomb material with angle of friction, ϕ.

This approximation emphasises the sensitivity of Fc to the square of the thickness of the frictional upper crust. Upward advection rapidly thins this upper frictional layer and is consequently very efficient at reducing the total strength of the crust. In regions of active convergence, this thermal weakening leads to strain concentration and can lead to the generation of a tectonic aneurysm with an associated high temperature metamorphic signal (Zeitler et al., 2001a,b). In extensional regions, thermal weakening can produce necking behaviour and further concentrate strain within the region of rapid erosion (Fletcher & Hallet, 1983).

Strain partitioning within the orogen

Along oblique plate boundaries, thermal thinning associated with exhumation can result in the concentration of lateral and convergent strain onto a single plate boundary, as along the Alpine Fault in southern New Zealand (Koons et al., 2003). The resulting kinematic evidence recorded in the structural fabric could suggest a change in plate convergence vectors in the opposite sense of the actual vector change.

Further north along the same plate boundary, contraction, strike-slip and extension displacements along the Hikurangi margin coincide with strong lateral gradients in material properties (Fig. 6). A recent modelling study (Upton et al., in press) investigated the influence of material properties on velocity partitioning within oblique subduction zones, using a geometry based on the Hikurangi margin to the north (Fig. 6a). Rheological variation in the oblique models was constrained by seismic velocity and attenuation information available from the Hikurangi margin. Extension and velocity partitioning occur if the subduction interface is weak, but neither develops if the subduction interface is strong (Fig. 6b). The simple mechanical model incorporating rheological variation based on seismic observations produced kinematics that closely match those published from the Hikurangi margin, including: extension within the Taupo Volcanic Zone, uplift of ponded sediments (Eberhart-Phillips & Reyners 1998) and dextral contraction in the south of the North Island.

Fig. 6.

The effect of coupling strength on extension and partitioning behind a subduction zone (Upton et al., in press). (a) Northern New Zealand and the Hikurangi Trough. The interpreted degree of plate coupling along the Hikurangi subduction margin is shown in italics (Reyners, 1998). (b) Model results showing margin normal velocity. In the north, where the interface is weak, margin normal motion is accommodated largely on the plate interface. Extension within the upper plate occurs above the downgoing slab. In the south, where the interface is strong, margin normal motion is accommodated within the upper plate and no extension occurs. (c) Model results showing margin-parallel velocity. Across the whole margin, this component is accommodated within the upper plate. (d) Margin-normal (no dash) and margin-parallel (dashed) components across the weak interface and strong interface (bold lines). Partitioning and extension occur above a weak interface but not a strong interface. Modified from Upton et al. (in press).

Fig. 6.

The effect of coupling strength on extension and partitioning behind a subduction zone (Upton et al., in press). (a) Northern New Zealand and the Hikurangi Trough. The interpreted degree of plate coupling along the Hikurangi subduction margin is shown in italics (Reyners, 1998). (b) Model results showing margin normal velocity. In the north, where the interface is weak, margin normal motion is accommodated largely on the plate interface. Extension within the upper plate occurs above the downgoing slab. In the south, where the interface is strong, margin normal motion is accommodated within the upper plate and no extension occurs. (c) Model results showing margin-parallel velocity. Across the whole margin, this component is accommodated within the upper plate. (d) Margin-normal (no dash) and margin-parallel (dashed) components across the weak interface and strong interface (bold lines). Partitioning and extension occur above a weak interface but not a strong interface. Modified from Upton et al. (in press).

Mechanical framework II: Viscous mixing, local kinematics and nappe formation

One of the primary goals in our mechanical treatment is to relate petro-structural observations of natural UHP terrains to the dynamics of convergent margins. Required for this is an estimate of global flow pattern within an orogen-fixed reference frame in an Eulerian sense as a function of bounding velocities and rheological parameters as produced in our dynamic model. By itself, however, the Eulerian velocity field is insufficient for describing the generation of folds with their associated thermal and compositional mixing, which constitute the petro-structural architecture of the orogen (Escher & Beaumont, 1997). The natural petro-structural architecture contains the kinematic and petrological information necessary to reconstruct position and rates of processes that formed the orogen. This architecture evolves as a function of local stretching and vorticity fields and is best defined within a local Lagrangian reference frame where displacements relative to individual particles can be tracked for at least part of their history. The evolution of the material derivatives of a Lagrangian system can lead quickly to a complex pattern of tendrils and whorls similar to that found along the route to chaos characterised by the “baker transformation” (Ottino, 1989). This Lagrangian complexity in the midst of Eulerian simplicity is responsible for the large and small scale viscous folds and geochemical mixing of geophysical fluids (e.g. Allegre & Turcotte, 1986; Passchier, 1997), and natural and analogue examples are provided in Figure 7. In these examples, non-coaxial, complex refolding is associated with vortex flow as opposed to dominantly shear flow (Ishii, 1992).

Fig. 7.

Large scale folding in natural (Malaspina Glacier) and analogue materials illustrate local complexity in Lagrangian mixing for a simple Eulerian velocity field (vectors on the right image indicate instantaneous surface velocity field). The transition from simple shear flow to the complex folded nappe-type structure occurs where the velocity field is characterised by divergent shear and vortex flow becomes significant. (Image of Malaspina Glacier from http://earthobservatory.nasa.gov/Newsroom/NewImages/images.php3?img_id=10307)

Fig. 7.

Large scale folding in natural (Malaspina Glacier) and analogue materials illustrate local complexity in Lagrangian mixing for a simple Eulerian velocity field (vectors on the right image indicate instantaneous surface velocity field). The transition from simple shear flow to the complex folded nappe-type structure occurs where the velocity field is characterised by divergent shear and vortex flow becomes significant. (Image of Malaspina Glacier from http://earthobservatory.nasa.gov/Newsroom/NewImages/images.php3?img_id=10307)

In the following discussion, we assume that nappe formation results from a Lagrangian mixing process similar to that demonstrated for other geophysical fluids (e.g. Fig. 7), and that this process is facilitated in regions with a large component of vortex flow relative to shear flow, and we examine the thermal implications of nappe formation. High vortex flow is predicted within regions where the Eulerian velocity field contains a component of divergent shear similar to that illustrated in our dynamic model (Fig. 5) and in natural and analogue examples (Fig. 7). We predict that divergent shear will concentrate vortex flow, with associated nappe formation, immediately beneath the upper plate and, further, that the orientation and wavelength is a function of the relative amounts of vortex vs. shear flow and, consequently, that these folds will contain information on relative slab displacement.

Thermal and petrological evolution

Since at least Albarede (1976), there has been a recognition that the mass velocity in an active region strongly influences the thermal state of that region and, consequently, that the thermal state of an orogen can be modelled with the three-dimensional transient partial differential equation for an advecting medium: 

formula
in which k = thermal conductivity as a spatial function, Cp = heat capacity, ρ= density, u, v, w = velocity in an externally-fixed, Cartesian reference frame, A = spatially varying radioactive heat production term discussed below.

In our numerical treatment, two-dimensional solutions employ a standard Crank-Nicholson approximation for the mildly non-linear advective conductive system (Koons, 1987). Three-dimensional solutions for the non-linear advection-conduction equations employ approximations derived from ITASCA (Cundall & Board, 1988). We address questions related to petrological and mechanical evolution with static, kinematic and dynamic models. Many of the difficulties and strategies involved in tracking the thermal history are thoroughly discussed by Roselle et al. (2002).

The standard model geometry used in the static models is 400 km on a side with a depth of 200 km. The block is discretised at variable resolutions depending upon the model, with grid spacing in the standard model set at 20 × 20 × 50 grid elements that are further divided into tetrahedral elements with linear thermal distribution across sub-elements (Fig. 8). In the unperturbed model with 40 km of crust, 110 km of mantle lithosphere, and 50 km of asthenosphere, radioactive heat production is concentrated in the upper 20 km of the crust and reduced to 10% in the lower crust and further reduced to 1% in the mantle (Table 1). The asthenosphere/lithosphere boundary is defined as the 1380 °C isotherm at which the basal boundary is maintained. Model sides are no net-flow boundaries and generally are sufficiently removed from thermal perturbations that they exert no thermal influence on the solutions. The steady state geotherm for these conditions is shown in Figure 8 outside the region of perturbation and is similar to those predicted from older, continental crust. We recognise that natural parameter ranges will result in a spread of temperatures and that our calculations therefore represent reasonable approximations to thermal trends but not absolute thermal values.

Fig. 8.

P–T deformation histories for the UHP Upper Plate (path 1) and HP Lower Plate (path 2) of northern Nordøyane. Grey areas indicate P–T estimates for the Upper Plate based on thermobarometry (Lerry et al., 2000a,b; Ravna, 2000). (b) Lhermal model of an instantaneously thickened crustal root, showing thermal profile at 3 and 6 Ma after emplacement. Position through time on the P–T diagram of the base of the crustal root (shown by x) is indicated by “model 1”. (c) Lhermal model of slab break-off leading to instantaneous emplacement of asthenosphere to within 10 km of the crustal root, showing thermal profile at 3 and 6 Ma after break-off. Position through time on the P–T diagram of the base of the crustal root (shown by x) is indicated by “model 2”.

Fig. 8.

P–T deformation histories for the UHP Upper Plate (path 1) and HP Lower Plate (path 2) of northern Nordøyane. Grey areas indicate P–T estimates for the Upper Plate based on thermobarometry (Lerry et al., 2000a,b; Ravna, 2000). (b) Lhermal model of an instantaneously thickened crustal root, showing thermal profile at 3 and 6 Ma after emplacement. Position through time on the P–T diagram of the base of the crustal root (shown by x) is indicated by “model 1”. (c) Lhermal model of slab break-off leading to instantaneous emplacement of asthenosphere to within 10 km of the crustal root, showing thermal profile at 3 and 6 Ma after break-off. Position through time on the P–T diagram of the base of the crustal root (shown by x) is indicated by “model 2”.

Table 1.

Properties of thermal models

ρAλCPT(°C)
Crust27003.5 × 10−63.01.0 × 103Top fixed at 10 °C
Crustal Root27003.5 × 10−73.01.0 × 103-----------
Lithosphere33001.2 × 10−83.751.0 × 103-----------
Asthenosphere34001.2 × 10−83.751.0 × 1031380 °C
ρAλCPT(°C)
Crust27003.5 × 10−63.01.0 × 103Top fixed at 10 °C
Crustal Root27003.5 × 10−73.01.0 × 103-----------
Lithosphere33001.2 × 10−83.751.0 × 103-----------
Asthenosphere34001.2 × 10−83.751.0 × 1031380 °C

Where p is density (kg m−3), A is radioactive heat production term (W kg−1), λ is the thermal conductivity (W m−1K−1) and Cp is specific heat (J kg−1K−1).

Results of static solutions: Crustal thickening

Perturbation of the lithosphere due to crustal thickening to the 120 km depths indicated by UHP assemblages is represented in the models of Figure 8. Thickening was accomplished in the model by instantaneous thickening of material with the same properties as the lower crust by 80 km over a horizontal distance of 100 km in the x direction and 400 km in the ydirection. Consequently, while the solution is fully three-dimensional, ∂T/∂y is minimised in the static models and the solutions are effectively two-dimensional. The initial temperature of this crustal root was homogenised to 600 °C. This assumption and the assumption of instantaneous emplacement are related to viscous mixing and are discussed below. The thermal solutions behave in a manner predictable from the initial solutions of England & Thompson (1984) and Grasemann et al. (1998). Initial cooling of the lithosphere by a vertical flow of cooler crustal material leads to UHP conditions in the underplated wedge. Thermal response of the wedge is very slow due to insulation from the asthenosphere by the thickened mantle lithosphere and results from radioactive heating with characteristic times of > 40 Ma. Eventually, the increased radioactive source distribution in the overthickened welt warms the welt and the steady-state solution, relevant for t > 100 Ma, contains a thermal hotspot in the overthickened crustal root (Fig. 8b). The time required to approach steady state is so great that it is largely irrelevant for our investigation of the processes at active convergent margins with time frames of < 20 Ma. On these short time frames, underplating has a cooling effect and thermal equilibration is very limited (Liou et al., 1996).

Asthenospheric involvement

The thermal effect on the crustal root of lithospheric slab breakoff (Davies & von Blackenburg, 1995) is represented by instantaneous upward displacement of the asthenosphere to within 10 km of the crustal root (Fig. 8c). For reference in the discussion, the evolution of a particle at 110 km below the surface, 10 km above the Moho, is chosen and plotted together with the position of the 800 °C isotherm (Fig. 8). The 10 km thick remaining lithospheric lid insulates the crustal root for more than 1 Ma. By 2 Ma, the thermal state of the crustal root has been significantly perturbed to a level that should be visible with present geothermometrics, assuming metamorphic reequilibration of the assemblage. The 800 °C isotherm has reached 10 km into the root by 3 Ma and may be representative of the thermal evolution of the Western Gneiss UHP assemblages (Krogh & Carswell, 1995) (Fig. 8a). Removal of more of the insulating lithospheric lid, bringing hot asthenosphere closer to the crust, will reduce the response time in the root in a manner well approximated by the tx = x2/K expression. Evolution of the thermal structure following asthenospheric involvement at longer time scales is similar to that discussed extensively by Ryan & Dewey (1997).

Thermal modelling of actively deforming regions has generally proceeded through solutions of transient non-linear equations for conductive and advective heat transport in two or three spatial dimensions (e.g. Koons, 1987; Peacock, 1995; Grasemann et al., 1998; Lin & Roecker, 1998; Roselle et al., 2002; Hacker et al., 2003a,b). While the thermal field in these solutions is transient, the models are kinematic and assumed velocity fields have been steady and smooth. As indicated by Sleep (1979) the process of internal deformation in nappe formation can have a significant influence upon the thermal and therefore petrological evolution of an orogen. This influence is in part approximated by the characteristic length scale of conduction  

formula
and is consequently sensitive to fold wavelength (Sleep, 1979).

The generation and evolution of nappes and folded isotherms is, as discussed above, spatially linked to regions of high vorticity (Fig. 9). The thermal evolution of these nappes is examined for various dominant wavelengths in the models presented in Figure 9 to see if petrological details of divergent P–T–t paths as a function of folding are likely to persist. For wavelengths of less than 5 km, thermal homogenisation occurs within the folded structures during subduction of crustal material (Fig. 9a). For wavelengths ~ 10 km, the thermal structure remains folded for periods of > 5 Ma (Fig. 9b) and should be visible within the assemblage record.

Fig. 9.

Thermal relaxation of instantaneously placed folds of varying thickness. (a) A series of nappes of thickness 5 km are placed into the upper mantle as a crustal root. After 2 Ma, a pod of colder material exists within the crustal root. (b) Thermal relaxation for nappes of thickness 10 km.

Fig. 9.

Thermal relaxation of instantaneously placed folds of varying thickness. (a) A series of nappes of thickness 5 km are placed into the upper mantle as a crustal root. After 2 Ma, a pod of colder material exists within the crustal root. (b) Thermal relaxation for nappes of thickness 10 km.

Dynamic model with thermal results

In regions of steep velocity gradients, the advective terms of the heat flow equation (Eqn. 9) dominate and the resultant thermal field carries the signature of the velocity field in a manner represented by the dimensionless Peclet number (Pe = VL/ k; where V = velocity normal to the thermal gradient, k = thermal diffusivity, and L = thermal length scale). The relevant length scale (L) for Peclet varies, but is generally on the order of either slab thickness or nappe thickness. As demonstrated in the initial static models, the former leads to thermal insulation represented by high Peclet, while the latter may allow local thermal homogenisation if nappe thickness is less than 5 km (Fig. 9).

The velocity distribution of our dynamic model produces regions of high Peclet near the surface and within the downgoing slab where thermal recovery significantly lags behind pressure changes leading to near isothermal exhumation or burial. High Peclet along the slab results in formation of lenses of coesite and diamond stability within the subducted material (Fig. 10b). Particle paths of UHP must pass through those stability regions at rates that are, in the Western Gneiss region and the Pacific-Australian plate boundary, at least within the time frame of several million years.

Fig. 10.

Mechanical and thermal model of continental subduction. (a) As for Figure 5a. (b) Cross-section through the model illustrating the thermal regime associated with the subducting slab. Stability fields of coesite and diamond are shown within the slab.

Fig. 10.

Mechanical and thermal model of continental subduction. (a) As for Figure 5a. (b) Cross-section through the model illustrating the thermal regime associated with the subducting slab. Stability fields of coesite and diamond are shown within the slab.

Depending upon the surface erosion conditions, a thermal hotspot can form in the near surface where exhumation is rapid, concentrating strain and leading to local high temperature metamorphism (Koons et al., 2002, 2003). In addition to the regions of high Peclet, there are regions in the mid to lower crust of low Peclet where there is little advective influence on the thermal structure. These regions of greenschist to amphibolite grade metamorphism are associated with high strain rates (Fig. 10b).

Thermal-mechanical coupling

In addition to the petrological evolution of crustal assemblages, thermal transience also influences the mechanical behaviour of the orogen for materials with thermally activated rheology (Eqn. 5). The degree of heating related to thermal influence from the asthenosphere reduces the flow resistance of material in the wedge. This thermally induced weakening can shift subducted continental material from subcritical to supercritical Grashof and permit this buoyant material to ascend (Fig. 1). The composition and rate of heating determines the trajectory of subducted material within Grashof space.

Discussion

Our three-dimensional dynamic model is successful in reproducing some of the observations from ancient and modern collisional orogens, but fails in several important aspects. In general, those thermal, petrological and structural features associated with the exhumation phase of UHP formation are at least in part captured in the dynamic model. These include depression of isotherms to produce UHP conditions, length and time scales similar those of the modern Pacific-Australian plate boundary, and significant vertical and horizontal strain partitioning. All of these features result from a standard, time-independent rheological and boundary model in which the crustal material is separated from high temperature asthenosphere by a lithospheric lid.

The exhumation phase is poorly represented with time-independent rheologies and constant boundary conditions. In the absence of the strong non-linear feedbacks that arise through mechanical and rheological coupling, the detachment of deeply subducted blocks of mafic material from the down going slab is unlikely. Rheological time-dependence sufficient to permit exhumation can result from heating of thermally activated crustal material, causing the subducted crust to become supercritical. In addition, more rapid rheological transitions can be produced through metamorphic transitions such as fluid production or reaction-enhanced ductility (e.g. Rubie, 1983). If these processes are widespread, then rapid fluctuations in crustal viscosity are possible and the conduction time scale of crustal heating no longer constrains the rates of UHP separation and exhumation.

In the dynamic model, we have not considered any shift in the far field velocities of the lithospheric slab. Given the recent observations from natural and analogue models of trench rollback (Funiciello et al., 2003b) it is clear that interaction with mantle flows, both in the upper and lower mantle, can profoundly affect the boundary conditions that influence continental subduction. Similarly, the high pressure heating part of the Western Gneiss P–T–t history (at 405 Ma) suggests an influence of high temperature asthenosphere on at least a limited scale prior to exhumation, in a manner not compatible with the steady state boundary models. Incorporation of these boundary variations in the dynamic model would initiate linking of mantle convection models with crustal dynamics and appears to be a rich direction for future investigation.

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Acknowledgements

We would like to thank D.A. Carswell and R. Compagnoni for the opportunity to present this study. Kirsten Jones is thanked for her contribution to the thermal modelling presented here and her assistance with many of the figures. Leigh Sterns and Erich Osterberg are thanked for Figure 7.

Figures & Tables

Fig. 1.

Diagram of the effective Grashof number representing buoyancy stress versus resisting stress for viscous crustal material. Critical Grashof (Grcr = 1) separates regions that tend to descend (subcritical) from those where buoyancy forces dominate (supercritical). The labelled trajectories indicate the evolution predicted for thermally activated materials (governed, for example by Eqn. 5) undergoing conductive heating after subduction (Eqn. 9; Fig. 7.). The plagio-clase and wet granite rheologies are calculated for strain rates of 10−14. The eclogite trajectory assumes only a buoyancy effect due to relevant phase changes.

Fig. 1.

Diagram of the effective Grashof number representing buoyancy stress versus resisting stress for viscous crustal material. Critical Grashof (Grcr = 1) separates regions that tend to descend (subcritical) from those where buoyancy forces dominate (supercritical). The labelled trajectories indicate the evolution predicted for thermally activated materials (governed, for example by Eqn. 5) undergoing conductive heating after subduction (Eqn. 9; Fig. 7.). The plagio-clase and wet granite rheologies are calculated for strain rates of 10−14. The eclogite trajectory assumes only a buoyancy effect due to relevant phase changes.

Fig. 2.

(a)Generalised geological map showing narrow refolded synclines of the tectonic cover in Baltica crust and the location of the study area. (b) Generalised tectonostratigraphic map modified from Gee et al. (1985). Large arrow shows the orientation of a time averaged (450-425 Ma) relative motion vector for Baltica with respect to Laurentia, approximated from reconstructions of Torsvik (1998). Small arrows show movement vectors of allochthons. From Terry et al. (2000b).

Fig. 2.

(a)Generalised geological map showing narrow refolded synclines of the tectonic cover in Baltica crust and the location of the study area. (b) Generalised tectonostratigraphic map modified from Gee et al. (1985). Large arrow shows the orientation of a time averaged (450-425 Ma) relative motion vector for Baltica with respect to Laurentia, approximated from reconstructions of Torsvik (1998). Small arrows show movement vectors of allochthons. From Terry et al. (2000b).

Fig. 3.

Pressure-temperature-time-deformation histories for high- and ultrahigh-pressure rocks of the northern segment (Terry et al., 2000b).

Fig. 3.

Pressure-temperature-time-deformation histories for high- and ultrahigh-pressure rocks of the northern segment (Terry et al., 2000b).

Fig. 4.(a)

Digital Elevation Model (GEOgraphix) of the northern South Island showing the Marlborough Fault System, the Alpine Fault and the Kaikoura Ranges. Dashed grey line shows the edge of the Hikurangi subduction zone as recorded by an abrupt shallowing of the deepest earthquakes from 250 km to 100 km in depth (Anderson & Webb 1994). Inset shows the plate tectonic setting of central New Zealand, the 1000 and 2000 m bathymetric contours defining the transition from the oceanic Hikurangi Plateau to the continental Chatham Rise. Plate vector is from De Mets et al. (1994). (b) Schematic interpretation of the tectonics based on the 3D velocity model at the cross-section shown in (a). Modified from Eberhart-Phillips & Reyners (1998).

Fig. 4.(a)

Digital Elevation Model (GEOgraphix) of the northern South Island showing the Marlborough Fault System, the Alpine Fault and the Kaikoura Ranges. Dashed grey line shows the edge of the Hikurangi subduction zone as recorded by an abrupt shallowing of the deepest earthquakes from 250 km to 100 km in depth (Anderson & Webb 1994). Inset shows the plate tectonic setting of central New Zealand, the 1000 and 2000 m bathymetric contours defining the transition from the oceanic Hikurangi Plateau to the continental Chatham Rise. Plate vector is from De Mets et al. (1994). (b) Schematic interpretation of the tectonics based on the 3D velocity model at the cross-section shown in (a). Modified from Eberhart-Phillips & Reyners (1998).

Fig. 5.

Mechanical model of continental subduction. The block consists of three layers: 0–15 km depth consists of a pressure-sensitive Mohr-Coulomb material representing the upper crust; the lower crust, from 15–;40 km, consists of a pressure-insensitive, temperature-sensitive rheology based on wet quartzite; the upper mantle, from 40–;200 km, consists of a pressure-insensitive, temperature-sensitive rheology based on diabase. (a) Contours of velocity parallel to the plate margin (Vy). Dashed line shows position of cross sections below. (b) Contours of rotation. (c) Contours of shear strain rate. (d) Contours of lateral strain rate.

Fig. 5.

Mechanical model of continental subduction. The block consists of three layers: 0–15 km depth consists of a pressure-sensitive Mohr-Coulomb material representing the upper crust; the lower crust, from 15–;40 km, consists of a pressure-insensitive, temperature-sensitive rheology based on wet quartzite; the upper mantle, from 40–;200 km, consists of a pressure-insensitive, temperature-sensitive rheology based on diabase. (a) Contours of velocity parallel to the plate margin (Vy). Dashed line shows position of cross sections below. (b) Contours of rotation. (c) Contours of shear strain rate. (d) Contours of lateral strain rate.

Fig. 6.

The effect of coupling strength on extension and partitioning behind a subduction zone (Upton et al., in press). (a) Northern New Zealand and the Hikurangi Trough. The interpreted degree of plate coupling along the Hikurangi subduction margin is shown in italics (Reyners, 1998). (b) Model results showing margin normal velocity. In the north, where the interface is weak, margin normal motion is accommodated largely on the plate interface. Extension within the upper plate occurs above the downgoing slab. In the south, where the interface is strong, margin normal motion is accommodated within the upper plate and no extension occurs. (c) Model results showing margin-parallel velocity. Across the whole margin, this component is accommodated within the upper plate. (d) Margin-normal (no dash) and margin-parallel (dashed) components across the weak interface and strong interface (bold lines). Partitioning and extension occur above a weak interface but not a strong interface. Modified from Upton et al. (in press).

Fig. 6.

The effect of coupling strength on extension and partitioning behind a subduction zone (Upton et al., in press). (a) Northern New Zealand and the Hikurangi Trough. The interpreted degree of plate coupling along the Hikurangi subduction margin is shown in italics (Reyners, 1998). (b) Model results showing margin normal velocity. In the north, where the interface is weak, margin normal motion is accommodated largely on the plate interface. Extension within the upper plate occurs above the downgoing slab. In the south, where the interface is strong, margin normal motion is accommodated within the upper plate and no extension occurs. (c) Model results showing margin-parallel velocity. Across the whole margin, this component is accommodated within the upper plate. (d) Margin-normal (no dash) and margin-parallel (dashed) components across the weak interface and strong interface (bold lines). Partitioning and extension occur above a weak interface but not a strong interface. Modified from Upton et al. (in press).

Fig. 7.

Large scale folding in natural (Malaspina Glacier) and analogue materials illustrate local complexity in Lagrangian mixing for a simple Eulerian velocity field (vectors on the right image indicate instantaneous surface velocity field). The transition from simple shear flow to the complex folded nappe-type structure occurs where the velocity field is characterised by divergent shear and vortex flow becomes significant. (Image of Malaspina Glacier from http://earthobservatory.nasa.gov/Newsroom/NewImages/images.php3?img_id=10307)

Fig. 7.

Large scale folding in natural (Malaspina Glacier) and analogue materials illustrate local complexity in Lagrangian mixing for a simple Eulerian velocity field (vectors on the right image indicate instantaneous surface velocity field). The transition from simple shear flow to the complex folded nappe-type structure occurs where the velocity field is characterised by divergent shear and vortex flow becomes significant. (Image of Malaspina Glacier from http://earthobservatory.nasa.gov/Newsroom/NewImages/images.php3?img_id=10307)

Fig. 8.

P–T deformation histories for the UHP Upper Plate (path 1) and HP Lower Plate (path 2) of northern Nordøyane. Grey areas indicate P–T estimates for the Upper Plate based on thermobarometry (Lerry et al., 2000a,b; Ravna, 2000). (b) Lhermal model of an instantaneously thickened crustal root, showing thermal profile at 3 and 6 Ma after emplacement. Position through time on the P–T diagram of the base of the crustal root (shown by x) is indicated by “model 1”. (c) Lhermal model of slab break-off leading to instantaneous emplacement of asthenosphere to within 10 km of the crustal root, showing thermal profile at 3 and 6 Ma after break-off. Position through time on the P–T diagram of the base of the crustal root (shown by x) is indicated by “model 2”.

Fig. 8.

P–T deformation histories for the UHP Upper Plate (path 1) and HP Lower Plate (path 2) of northern Nordøyane. Grey areas indicate P–T estimates for the Upper Plate based on thermobarometry (Lerry et al., 2000a,b; Ravna, 2000). (b) Lhermal model of an instantaneously thickened crustal root, showing thermal profile at 3 and 6 Ma after emplacement. Position through time on the P–T diagram of the base of the crustal root (shown by x) is indicated by “model 1”. (c) Lhermal model of slab break-off leading to instantaneous emplacement of asthenosphere to within 10 km of the crustal root, showing thermal profile at 3 and 6 Ma after break-off. Position through time on the P–T diagram of the base of the crustal root (shown by x) is indicated by “model 2”.

Fig. 9.

Thermal relaxation of instantaneously placed folds of varying thickness. (a) A series of nappes of thickness 5 km are placed into the upper mantle as a crustal root. After 2 Ma, a pod of colder material exists within the crustal root. (b) Thermal relaxation for nappes of thickness 10 km.

Fig. 9.

Thermal relaxation of instantaneously placed folds of varying thickness. (a) A series of nappes of thickness 5 km are placed into the upper mantle as a crustal root. After 2 Ma, a pod of colder material exists within the crustal root. (b) Thermal relaxation for nappes of thickness 10 km.

Fig. 10.

Mechanical and thermal model of continental subduction. (a) As for Figure 5a. (b) Cross-section through the model illustrating the thermal regime associated with the subducting slab. Stability fields of coesite and diamond are shown within the slab.

Fig. 10.

Mechanical and thermal model of continental subduction. (a) As for Figure 5a. (b) Cross-section through the model illustrating the thermal regime associated with the subducting slab. Stability fields of coesite and diamond are shown within the slab.

Table 1.

Properties of thermal models

ρAλCPT(°C)
Crust27003.5 × 10−63.01.0 × 103Top fixed at 10 °C
Crustal Root27003.5 × 10−73.01.0 × 103-----------
Lithosphere33001.2 × 10−83.751.0 × 103-----------
Asthenosphere34001.2 × 10−83.751.0 × 1031380 °C
ρAλCPT(°C)
Crust27003.5 × 10−63.01.0 × 103Top fixed at 10 °C
Crustal Root27003.5 × 10−73.01.0 × 103-----------
Lithosphere33001.2 × 10−83.751.0 × 103-----------
Asthenosphere34001.2 × 10−83.751.0 × 1031380 °C

Where p is density (kg m−3), A is radioactive heat production term (W kg−1), λ is the thermal conductivity (W m−1K−1) and Cp is specific heat (J kg−1K−1).

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