a geodynamic model of subduction evolution and slab

976 www.gsapubs.org | Volume 9 | Number 6 | LITHOSPHERE

A geodynamic model of subduction evolution and slab detachment to explain Australian plate acceleration and deceleration during the latest Cretaceous–early Cenozoic

W.P. Schellart1,*1DEPARTMENT OF EARTH SCIENCES, VRIJE UNIVERSITEIT AMSTERDAM, DE BOELELAAN 1085, 1081 HV AMSTERDAM, NETHERLANDS

ABSTRACT

During the latest Cretaceous–early Cenozoic, the northern margin of the Australian plate was characterized by a large (4000 km wide) north- to northeast-dipping subduction zone (New Guinea–Pocklington subduction zone) consuming a marginal basin. Geological and geophysical data imply that the subduction zone was active ca. 71–50 Ma, and suggest that it was responsible for plate acceleration from ~1.0 to ~7.3 cm/yr ca. 64–59 Ma, and plate deceleration from ~7.3 to ~0.3 cm/yr at 52–49 Ma. This paper presents a numerical model of buoyancy-driven subduction to test if the rates of Australian plate acceleration and deceleration can be ascribed to the progressive evolution of a subducting slab. The geodynamic model reproduces the first-order plate velocity evolution of the Australian plate, with a transient ~5 m.y. period of acceleration from 2 to 8 cm/yr during upper mantle slab lengthening, an ~5 m.y. period of rapid plate motion (~5–8 cm/yr), and a short, 3.9 m.y., period of plate deceleration, starting with a 2 cm/yr velocity drop during 3.1 m.y. of continental subduction and followed by ~0.8 m.y. of rapid deceleration (4 cm/yr velocity drop) during slab detachment. The geodynamic model demonstrates that plate velocity increases or decreases of ~4–6 cm/yr can occur over a period lasting <1 m.y. to a few million years, comparable to what is observed for the latest Cretaceous–early Cenozoic evolution of the Australian plate. Such rates of plate acceleration and deceleration could be tested against plate kinematic data for other subduction settings on Earth.

LITHOSPHERE; v. 9; no. 6; p. 976–986 | Published online 11 October 2017 https: // doi .org /10 .1130 /L675 .1

INTRODUCTION

The Earth’s lithosphere is segmented into numerous larger and smaller plates (Bird, 2003) that move at a wide variety of plate tectonic speeds, to as much as ~10 cm/yr. The velocities of these plates have traditionally been calculated in hotspot reference frames, such as the Pacific (Minster and Jordan, 1978; Gripp and Gordon, 2002; Wessel and Kroenke, 2008), Indo-Atlantic (e.g., O’Neill et al., 2005), or global (e.g., Gordon and Jurdy, 1986; Doubrovine et al., 2012) hotspot reference frames, as well as no-net-rotation reference frames (e.g., Argus and Gordon, 1991; Kreemer et al., 2003; Argus et al., 2011). More recently, other geological and geophysical features have been used to constrain the absolute velocities of the tectonics plates, such as subducted slabs in the upper and lower mantle (e.g., van der Meer et al., 2010; Schellart, 2011; Butterworth et al., 2014), upper mantle seismic anisotropy (e.g., Long and Silver, 2009; Kreemer, 2009; Conrad and Behn, 2010), and spreading ridges at the Earth’s surface (e.g., Becker et al., 2015; Wessel and Müller, 2016). A number of conceptual and geodynamic models have been put forward to explain the variations in plate velocities. Although such models vary as to what might be the main physical mechanism that explains variation in absolute plate velocities, such as subducting plate age (e.g., Carlson et al.,

1983; Goes et al., 2008), slab width (Schellart et al., 2010), plate circum-ference attached to a slab (Forsyth and Uyeda, 1975; Gripp and Gordon, 2002), double subduction (Jagoutz et al., 2015), and the transition from normal (i.e., oceanic) subduction to collision and/or continental subduc-tion (e.g., Patriat and Achache, 1984; Klootwijk et al., 1992; Molnar and Stock, 2009; van Hinsbergen et al., 2011; White and Lister, 2012; Zahi-rovic et al., 2012; Capitanio and Replumaz, 2013), the models generally have the common element of subduction to explain rapid plate motion (e.g., Conrad and Lithgow-Bertelloni, 2002). An important exception is the so-called plume push force, which has been proposed to explain the extreme plate velocity of the Indian plate in the latest Cretaceous–earliest Cenozoic (Cande and Stegman, 2011; van Hinsbergen et al., 2011). The ridge push force is generally thought to be an order of magnitude smaller than the slab pull force (Forsyth and Uyeda, 1975).

The Australian plate is home to the Australian continent, which is cur-rently the fastest moving continent on Earth (Keep and Schellart, 2012) with north- to northeast-directed velocities as high as 6.5 cm/yr (Fig. 1). These high velocities can be ascribed to the wide subduction zones that are present along the northern boundary of the Australian plate, namely the north- to northeast-dipping Sunda subduction zone in the northwest and the north- to northeast-dipping Melanesian subduction zone in the northeast. Over the past ~75 m.y., the Australian plate has generally under-gone rapid plate velocities, except for 2 periods, at 73–63 Ma and 49–40

© 2017 Geological Society of America. For permission to copy, contact [email protected]

*[email protected]

Downloaded from https://pubs.geoscienceworld.org/gsa/lithosphere/article-pdf/9/6/976/3985681/976.pdfby gueston 18 September 2018

https://doi.org/10.1130/L675.1

mailto:w.p.schellart%40vu.nl?subject=

http://www.geosociety.org

LITHOSPHERE | Volume 9 | Number 6 | www.gsapubs.org 977

Australian plate motion change due to subduction evolution and slab detachment | RESEARCH

Ma, when plate velocities reached minima of 0.8–2.9 cm/yr and 0.3–1.6 cm/yr, respectively, in the Indo-Atlantic moving hotspot reference frame (Fig. 2). During the intervening period the velocities were much higher with a maximum of 7.3 cm/yr at 59–52 Ma. During the latest Cretaceous (Maastrichtian) to early Eocene, the Australian plate was bordered to the east, south, and west by spreading ridges in the Tasman Sea, South-ern Ocean, and Indian Ocean (e.g., Royer and Sandwell, 1989; Yan and Kroenke, 1993; Gaina et al., 1998; Hall, 2012; Seton et al., 2012); this cannot explain the period of rapid plate motion and the preceding and following phases of plate acceleration and deceleration. In Schellart and

Spakman (2015), a 4000-km-wide (trench-parallel extent) subduction zone was identified in the New Guinea region active ca. 71–50 Ma, and the phases of plate acceleration, rapid plate motion, and plate deceleration of the Australian plate were ascribed to the geodynamic evolution of this subduction zone. The subduction zone consumed a small oceanic back-arc basin, the Emo backarc basin, with a maximum north-south extent of

~1100 km. The phase of plate acceleration was ascribed (in Schellart and Spakman, 2015) to progressive upper mantle slab lengthening during the initial transient subduction phase and the phase of plate deceleration was ascribed to slab detachment during the final stage of subduction.

I present here a geodynamic numerical model of buoyancy-driven subduction to test if the rates of plate acceleration and deceleration, and the high rate of Australian plate motion during the intervening period, can be ascribed to the progressive evolution of the subduction zone in the New Guinea region, from transient slab lengthening to the final stage of slab detachment. The geodynamic model provides new insight into the rates of plate acceleration and deceleration during subduction of the edge of a plate that contains a large continent.

METHODS

Numerical models are presented that have specifically been designed to investigate the subduction evolution of a small oceanic backarc basin (Emo backarc basin), with a particular focus on the latest stage of subduc-tion when all oceanic lithosphere is consumed and continental passive margin lithosphere approaches the trench. The models use the code Under-world (Stegman et al., 2006; Moresi et al., 2007, 2014), in which plate motion, subduction and mantle flow are modeled in a Cartesian domain using compositional buoyancy contrasts in an incompressible Boussinesq fluid at a very low Reynolds number. Distinct volumes (e.g., continental upper crust, lower crust and lithospheric mantle, sublithospheric upper mantle, lower mantle) are represented by sets of Lagrangian particles that are embedded within a standard Eulerian finite element mesh, which discretizes the problem to solve the governing equations. For additional

40

14012080 160 200

30

20

10

o

ooo o o

o

o

o

5 cm/yr

180o

0o

100o

10o

Collision/

Fossil

Subduction

obduction

subduction

subductionIncipient

Tasman

t

New

-

olin

fos sub

zo

n a

SeaBasin

IndianOcean

Southern Ocean

PacificOcean

Australian plate

Antarctic plate

Sunda platePacific plate

eGui

Pck g on

silduct on ne

i

800

1

2

3

4

5

6

7

8

9

70 60 50 40Time [Ma]Ve

loci

ty o

f poi

nt o

n A

ustr

alia

n pl

ate

[cm

/yr]

30 20 10 0

10

800

1

2

3

4

5

6

7

8

9

70 60 50 40Time [Ma]

Nor

thw

ard

velo

city

com

pone

nt [c

m/y

r]

30 20 10 0

10

IAHS-2005GHS-2012

A BIAHS-2005GHS-2012

Figure 1. Plate tectonic map of the central and eastern part of the Aus-tralian plate, showing the present-day plate boundaries and velocities (in cm/yr), as well as the fossil New Guinea–Pocklington subduction zone. The velocities are based on the relative plate motion model of DeMets et al. (1994) using the Indo-Atlantic moving hotspot reference frame from O’Neill et al. (2005).

Figure 2. Absolute velocity of the Australian plate at 74–0 Ma for a point currently located in the Gulf of Carpentaria (i.e., relatively close to the center of the New Guinea–Pocklington fossil subduction zone, circle with cross in Fig. 1). (A) Total velocity. (B) Northward velocity component (i.e., approximately perpendicular to the strike of the fossil subduction zone). Velocities were calculated in an Indo-Atlantic moving hotspot reference frame (O’Neill et al., 2005; IAHS-2005) and global moving hotspot reference frame (Doubrovine et al., 2012; GHS-2012) using the relative plate motion model of Müller et al. (2008). Note that the Indo-Atlantic hotspot reference frame is the preferred reference frame (see discussion of Plate Velocity Changes in Model and Nature). The velocities in the Indo-Atlantic frame are after Schellart and Spakman (2015).


SCHELLART


information on the numerical technique and the nondimensional equa-tions, see Moresi et al. (2007, 2014) and Stegman et al. (2006). Velocities in the models are scaled following the scaling formulations presented in Schellart and Moresi (2013).

The modeling domain is 1100 km deep and 6200 km long and has free-slip boundary conditions along the top surface, bottom surface, and lateral side walls (Fig. 3). Mesh resolution in the 6200 × 1100 km numeri-cal domain is 1024 (length) by 256 (depth) elements. An adaptive spatial mesh refinement has been implemented to increase resolution in the region of the subduction zone and close to the surface. The smallest cell dimen-sions in this central region (rectangle with dotted outline in Fig. 3), 220 km thick and 1860 km wide, are 2.79 km (length) by 2.15 km (depth). Initial particle distribution is 20 particles per cell (total of 5,242,880 particles).

The reference model A involves a layered mantle volume incorporat-ing a free, four-layer subducting plate, a fixed, neutrally buoyant, non-stratified overriding plate, a low-viscosity sublithospheric upper mantle, and a high-viscosity lower mantle. The top 100 km of the model domain consists of a low-viscosity, low-density air layer to provide a free surface to the tectonic plates (sticky air layer following Schmeling et al., 2008; Crameri et al., 2012). The subduction zone interface is implemented in the same way as in the numerical models in Schellart and Moresi (2013) and the laboratory models of Duarte et al. (2013, 2015), in which a weak zone develops that consists of the weak materials from the top layer of the subducting plate, which allows for progressive one-sided subduction. The subducting plate has a negatively buoyant, 1120-km-long oceanic lithosphere segment representing the Emo oceanic backarc basin and a

trailing, positively buoyant 3600-km-long continental lithosphere segment representing the Australian continent (which includes a 100 km passive margin). The length of the oceanic lithosphere segment is based on the reconstruction from Schellart and Spakman (2015), while the length of the continental part is based on the north-south extent of the Australian continent in the latest Cretaceous–earliest Cenozoic.

The oceanic part of the subducting plate is 80 km thick and con-sists of 4 layers with their own thickness and rheology following earlier modeling approaches (e.g., Schellart et al., 2010; Stegman et al., 2010), except that the models presented here use viscoplastic rheologies with a von Mises yield stress throughout the lithosphere to facilitate possible detachment. The physical properties of the different layers are listed in Table 1. The 80 km thickness of the oceanic lithosphere is based on the age of the Emo backarc basin, which is derived from the age of the pro-toliths of the Emo metamorphics and Owen Stanly metamorphic rocks that are thought to originate from this basin. Early research on the meta-morphics proposed a Mesozoic age (Davies and Warren, 1992) or Late Cretaceous age (Worthing and Crawford, 1996). In more recent work, the plate tectonic reconstruction from Hall (2012) implies that an ocean basin or backarc basin north of the New Guinea passive margin already existed in the Late Jurassic–Early Cretaceous. In other recent work, the age of the protolith of the metamorphics has been described as Cretaceous (Baldwin et al., 2012), while Davies (2012) recorded a mid-Cretaceous age with radiometric dates of 120–107 Ma. This would imply an oceanic backarc basin with an age range of ~36–70 million years at the time of subduction ca. 71–50 Ma. Considering this age range, an 80-km-thick

Sticky air

Sub-lithospheric upper mantle

Lower mantle

Overriding plate Subducting plate

Australian continental lithosphere

6200 km1000 km

oceanicEmo backarc

lithosphere

Forearc

10-3 10-2 10-1 100 101 102 103 104

Non-dimensional effective viscosity

Passivemargin

Passivemargin

Trench

Figure 3. Model set-up of the numerical model of buoyancy-driven progressive subduction in a layered mantle in two-dimen-sional space with free-slip boundaries. In the model, the subducting plate consists of a 1120-km-long oceanic lithosphere segment (which includes the 280-km-long initial slab perturbation) representing the Emo oceanic backarc basin and a 3600-km-long continental lithosphere segment representing the Australian continental lithosphere. The oceanic and conti-nental lithosphere of the subducting plate have the same four-layer rheology, but a different density in which the oceanic lithosphere is negatively buoyant, while the continental lithosphere has a continental crust (top two layers) that is posi-tively buoyant. See Table 1 for details of the physical properties. The initial slab perturbation curves into the subduction zone with a maximum dip angle at the slab tip of 27.3°. Note that the top panel shows the different domains in the model, while the bottom panel shows the nondimensional effective viscosity. The rectangle with the dotted outline shows the region with maximum horizontal and vertical resolution. The lateral extent of the passive margin is 100 km. Note that the continental crustal thickness across the passive margin changes linearly (top panel), while the rheology across the passive margin is constant (bottom panel).




oceanic lithosphere (implying an average age of ~50 million years at the time of subduction) is justified. An average density contrast of 80 kg/m3 between the oceanic slab and ambient sublithospheric mantle is adopted in the models, as it is assumed that there is pervasive eclogitization of the basaltic crust into eclogite facies rocks (Cloos, 1993).

The continental part of the subducting plate, which includes the 100-km-long passive margin, has the same rheological layering with the same thickness as the oceanic part (Fig. 3, bottom panel). The only differ-ence between the oceanic and continental domains is the density structure, with a lighter, positively buoyant, continental lithosphere due to a 30 km continental crust, a denser, negatively buoyant, oceanic lithosphere (Table 1), and a passive margin across which the density changes linearly due to a linear change in continental crustal thickness from 30 km at the continental side to 10 km at the oceanic side. This set-up was specifically chosen to investigate where strain localization and slab detachment would occur in a scenario without any lateral rheological contrasts, but with only lateral density contrasts in the subducting plate.

The model overriding plate is neutrally buoyant, implying that it is oceanic in nature with a density that is comparable to that of the sublitho-spheric mantle. This is most likely if one considers the tectonic setting to the north of Australia in the latest Cretaceous–early Eocene, where reconstructions indicate that there was an ocean to the north (Hall, 2012). Another reason for implementing a neutral buoyancy was to minimize the effect of the overriding plate on the subduction process, such that the change from oceanic to continental subduction and the process of slab detachment could be studied without other complicating factors. For the rheology the simplest set-up was chosen (following Duarte et al., 2013; Schellart and Moresi, 2013) that allows for continuous and progressive subduction, because little is known about the properties and rheology of the overriding plate. The model overriding plate has a forearc region with a higher viscosity than the backarc region (Table 1) (following Schellart and Moresi, 2013), as it is generally thought that the forearc is relatively cool and strong with depressed isotherms due to the cold subducting plate diving below it, and therefore has a higher effective viscosity than the arc and backarc regions, where isotherms are elevated.

Apart from reference model A, one other model is presented, model B, with a higher yield stress compared to the reference model (Table 1). A three-dimensional geometrical set-up would be ideal for the geodynamic models, as presented in other modeling work of subduction and slab

detachment (e.g., van Hunen and Allen, 2011; Capitanio and Replumaz, 2013; Chertova et al., 2014; Capitanio et al., 2015), because such models can take into account three-dimensional (3D) spatial effects, including lateral migration of slab detachment, tearing, and trench-parallel flow of mantle material. However, this was found to be unfeasible at this time, considering the very large size of the subduction zone (~4000 km), the required resolution, and the required number of time steps (several thou-sand). These earlier works generally used lower spatial resolution and modeled smaller plates and subduction zones. Furthermore, considering that the natural prototype is a wide subduction zone, its geodynamic behavior, in particular that of its central portion, can be approximated with a model using a 2D spatial set-up (Schellart and Moresi, 2013).

MODEL RESULTS

Reference model A shows a general style of subduction with progres-sive slab lengthening and an increase of the slab dip angle (averaged along the entire slab length) during the free sinking phase, from 21° at the start (Fig. 3) to 40° at an intermediate stage (Fig. 4A), to 46° at an advanced stage (Fig. 4B). As the slab approaches the 660 discontinuity its tip is slightly deflected to a lower slab dip angle. During the free sinking phase the trench-normal subducting plate velocity v

SP⊥ increases rapidly to a max-imum of ~8 cm/yr, when the slab tip is located at ~578 km depth (82 km above the 660 km discontinuity) (Fig. 5A). The trench-normal trench velocity (v

T⊥) is significantly less than vSP⊥ and increases to a maximum of

~2.5 cm/yr, which is reached at a time before the maximum of vSP⊥. Dur-

ing the slab tip–660 km discontinuity interaction phase (~8–12 m.y.), vSP⊥

decreases from ~8 to 5–6 cm/yr, while vT⊥ remains relatively unchanged.

From ~12.2 to 15.3 m.y. continental subduction takes place (Figs. 6A, 6B), during which v

SP⊥ decreases from ~5 to 3 cm/yr. This is followed by a short phase of slab detachment from 15.3 to 16.1 m.y. (Figs. 6C–6E), during which v

SP⊥ decreases from ~3 to –1 cm/yr and then increases again to ~0 cm/yr (Fig. 5A). During continental subduction and slab detachment, v

T⊥ decreases from ~2 cm/yr to 0 cm/yr. After slab detachment, subduction stops, v

SP⊥ and vT⊥ are close to zero, while the detached slab continues to

sink into the mantle (Figs. 4H, 4I, and 6F).During the phase of continental subduction, the continental lithosphere

is pulled into the mantle by the oceanic slab segment, and the maximum subduction depth of the continental crust (its base, originally at 30 km

TABLE 1. NUMERICAL MODEL PROPERTIES FOR REFERENCE MODEL A

Domain Density (kg/m3)

Rheology Nondimensional viscosity

Yield stress (MPa)

Thickness (km)

Subducting plate (oceanic part)layer 1 (top) 3280 viscoplastic (von Mises) 100 ηUM-Max

5.88 10layer 2 3280 viscoplastic (von Mises) 100 ηUM-Max

29.4 20layer 3 3280 viscoplastic (von Mises) 1000 ηUM-Max

352.8 20layer 4 (bottom) 3280 viscoplastic (von Mises) 50 ηUM-Max

29.4 30Subducting plate (continental part)

top crustbottom crusttop LithMbottom LithM

2800280032003200

viscoplastic (von Mises)viscoplastic (von Mises)viscoplastic (von Mises)viscoplastic (von Mises)

100 ηUM-Max100 ηUM-Max

1000 ηUM-Max50 ηUM-Max

5.8829.4

352.829.4

10202030

Overriding plate forearc 3200 linear viscous 1000 ηUM-Max60

Overriding plate backarc 3200 linear viscous 50 ηUM-Max60

Sublithospheric upper mantle 3200 nonlinear viscous 0.1–1 ηUM-Max580–660

Lower mantle 3200 linear viscous 100 ηUM-Max340

Air 0 linear viscous 0.1 ηUM-Max100

Note: ηUM-Max is the maximum viscosity of the sublithospheric upper mantle, which has a nonlinear stress-dependent viscosity with a stress exponent n = 3.5 (following Mackwell et al., 1990). The scaled viscosity ηUM-Max = 5 × 1020 Pa∙s. The continental part of the subducting plate has two crustal layers and two lithospheric mantle (LithM) layers. Model B is the same as reference model A except that the yield stress of layer 3 and top LithM in the subducting plate is 382.2 MPa.


SCHELLART


depth), just before slab detachment is complete, is 96 km. This indicates a depth increase of 66 km. The slab detachment process starts with the formation of two conjugate shear zones (Fig. 6C) and occurs mostly by symmetrical necking (Figs. 6C–6E). Once slab detachment is complete, the trailing slab segment slightly rebounds.

The other model (model B with a stronger subducting plate than model A) generally shows the same evolution as model A until ~12 m.y.; model B starts continental subduction just after ~12 m.y., and this continues until

~20 m.y. when vSP⊥ has decreased to zero, v

T⊥ is slightly negative (~–0.5

cm/yr), continental subduction stops and the only remaining activity is the very slow progressive steepening of the slab (Figs. 5B and 7). The maximum subduction depth of the continental crust (its base, originally at 30 km depth) at the end of the model run is 144 km.

DISCUSSION

Plate Velocity Changes in Model and Nature

The subduction models A and B show a general style of subduction behavior during the oceanic subduction period that is comparable to previ-ous buoyancy-driven subduction modeling studies, with progressive slab steepening (e.g., Jacoby, 1973; Funiciello et al., 2004; Schellart, 2004a; Guillaume et al., 2009) and an increasing trenchward subducting plate velocity during the free sinking phase to a maximum velocity when the slab tip is just above (<~100 km from) the 660 km viscosity discontinu-ity (e.g., Schellart, 2008; Capitanio et al., 2010; Schellart et al., 2011; Meyer and Schellart, 2013; Chen et al., 2015, 2016) (Fig. 5). In the cur-rent models it takes only ~5 m.y. to increase from v

SP⊥ = ~2 cm/yr to ~8 cm/yr. Such a velocity increase over such a time period is comparable to that implied by the velocity curve for the Australian plate in the Indo-Atlantic moving hotspot reference frame (Fig. 2), showing an increase from ~1 to ~7.4 cm/yr between ca. 64 and ca. 59 Ma and an increase in the northward velocity component, which is approximately perpendicular to the strike of the fossil subduction zone, from ~0.2 to ~5.0 cm/yr. The increase in subducting plate velocity in the models and for the Austra-lian plate can be ascribed to the progressive lengthening of the slab and deepening of the slab tip (Figs. 4A, 4B, and 5), and thereby the increase in slab negative buoyancy and net slab pull (Schellart, 2004b). These high subducting plate velocities can be reached, despite the fact that the plate is relatively long in the latest Cretaceous–early Cenozoic (~4000–5000 km long Australian plate in the north-south direction).

The decrease in subducting plate velocity during the slab tip–660 km discontinuity interaction phase is generally also observed in earlier works (e.g., Schellart, 2008; Capitanio et al., 2010; van Hunen and Allen, 2011; Schellart et al., 2011; Meyer and Schellart, 2013; Chen et al., 2015, 2016). Reference model A further shows a period of decrease in v

SP⊥ from ~4.5 cm/yr to ~3 cm/yr during ~3.1 m.y. of continental subduction, followed by an ~0.8 m.y. phase of slab detachment, during which v

SP⊥ decreases from ~3 cm/yr to –1 cm/yr (Fig. 5A). The slowdown during continental subduction can be explained by the decrease in buoyancy force of the slab (i.e., becoming less negative) due to addition of positively buoyant continental slab material (Figs. 6A–6C, images on left). The rapid slowdown during slab detachment can be explained by the rapid decrease in coupling between the detaching slab segment and the trailing slab seg-ment due to the rapid decrease in effective viscosity in the detachment zone (Figs. 6B–6E, images on right). During the ~3.9 m.y. period of con-tinental subduction and slab detachment the velocity decreases ~5.5 cm/yr, which is comparable to the ~3 m.y. period from ca. 52 Ma to ca. 49 Ma during which the Australian plate velocity decreased ~7 cm/yr (Fig. 2A) and the northward velocity component of the Australian plate decreased

~5 cm/yr (Fig. 2B) in the Indo-Atlantic moving hotspot reference frame. Subduction model A thereby demonstrates the physical viability of the conceptual model in which slowdown of the Australian plate in the latter part of the early Eocene is ascribed to continental subduction and subse-quent slab detachment, as illustrated in Figures 5A and 6.

In the preceding two paragraphs the subducting plate velocities from the numerical models have been compared to the subducting Australian plate velocities in the Indo-Atlantic moving hotspot reference frame from O’Neill et al. (2005). As explained in detail elsewhere (e.g., Schellart

10-3 10-2 10-1 100 101 102 103 104


13.9 m.y.

15.3 m.y.

16.0 m.y.

16.1 m.y.

16.4 m.y.

17.0 m.y.

12.2 m.y.

7.4 m.y.

0 km

500

100010000 km 2000 3000 4000 5000 6000

4.2 m.y.

0 km

500

1000

0 km

500

1000

0 km

500

1000

0 km

500

1000

0 km

500

1000

0 km

500

1000

0 km

500

1000

0 km

500

1000

B

C

D

E

F

G

H

A

I

Figure 4. Model results showing the evolution of the reference numerical subduction model (reference model A) at nine different stages. All images show the nondimensional effective viscosity field. (A) Early stage of the initial transient subduction phase (4.2 m.y.). (B) Late stage of the initial tran-sient subduction phase with maximum subducting plate velocity (7.4 m.y.). (C) Start of continental subduction (12.2 m.y.). (D) Continental subduction phase (13.9 m.y.). (E) Start of slab detachment with necking of lithosphere and formation of conjugate shear zones (15.3 m.y.). (F) Late stage of slab detachment with necking almost complete (16.0 m.y.). (G) Completion of slab detachment (16.1 m.y.). (H) Early stage after slab detachment with sinking of detached slab (16.4 m.y.). (I) Late stage after slab detachment with sinking of detached slab (17.0 m.y.).




10-3 10-2 10-1 100 101 102 103 104


1800 2000 2200 2400 km16001400120010001800 2000 2200 2400 km1600140012001000

10-7 10-6 10-5 10-4 10-3 10-2 10-1

Non-dimensional second invariant of the strain rate

0 km100200300

1800 2000 2200 2400 km1600140012001000

12.2 m.y.

13.9 m.y.

15.3 m.y.

16.0 m.y.

16.1 m.y.

16.4 m.y.

0 km100200300

0 km100200300

0 km100200300

0 km100200300

0 km100200300

Startcontinentalsubduction

Start slabdetachment(conjugateshears form)

Slabdetachmentcomplete

A

B

C

D

E

F

0 2 4 6 8 10 12 14 16 18 20 22Time [m.y.]

0

-1

-2

1

2

3

4

5

6

7

8

9

10Ve

loci

ty [c

m/y

r]


Start slabdetachment(conjugate

shears form)

Slabdetachment

complete

Slab tiptouches 660 km

discontinuity

A

0 2 4 6 8 10 12 14 16 18 20 22Time [m.y.]

0

-1

-2

1

2

3

4

5

6

7

8

9

10

Velo

city

[cm

/yr]


Slab tiptouches660 km

discontinuity

BvSP

T

vT

T

218

578

337

441

610 model BvSP

T

model BvT

T

308

224

449

Reference model A

Figure 5. Diagrams showing the evolution of the trench-normal subducting plate velocity (vSP⊥, trenchward is positive) and trench-normal trench velocity (vT⊥, retreat is positive). (A) Subducting plate velocity (black line) and trench velocity (gray line) for reference model A with a core layer with yield stress C = 352.8 MPa. (B) Subducting plate velocity (black line) and the trench velocity (gray line) for model B with a stronger core layer (C = 382.2 MPa). Note that only the reference model shows slab detachment (see Figures 4 and 6), while model B does not (Fig. 7). The gray arrows with the gray numbers indicate the depth of subduction (in km) of the slab tip. Note that vSP⊥ is measured at the trench; vSP⊥ at the trailing edge of the subducting plate is less due to (unwanted) stretching of the subducting plate, which is a consequence of the great initial length (4720 km) and finite strength of the subducting plate.

Figure 6. Zoom in of model results (reference model A) in the region of the subduction zone showing the evolution of the numerical subduction model at different stages during continental subduction, slab detachment, and after slab detachment. (A) Time at the start of continental subduction (12.2 m.y.). (B) Continental subduction phase (13.9 m.y.). (C) Start of slab detachment with necking of litho-sphere and formation of conjugate shear zones (15.3 m.y.). (D) Late stage of slab detachment with necking almost complete (16.0 m.y.). (E) Completion of slab detachment (16.1 m.y.). (F) Early stage after slab detachment with sinking of detached slab (16.4 m.y.). Images on the left show the different domains in the numerical model, images in the middle show the second invariant of the nondimensional strain rate, and images on the right show the nondimensional effective viscosity field.


SCHELLART


et al., 2008; Schellart, 2011; Schellart and Spakman, 2012, 2015), this Indo-Atlantic hotspot reference frame is preferred over other reference frames for calculating plate velocities and plate boundary velocities, which is particularly evident for the western Pacific domain. The reasons are many, including minimization of global trench migration velocities and rollback-induced toroidal volume fluxes in this reference frame (Schellart et al., 2008), optimal agreement between predicted and observed fossil slab locations in the southwest Pacific (Schellart and Spakman, 2012) and slab structures in the western Pacific (Schellart, 2011), and highest coincidence with fossil slab sinking-induced dynamic topography in the southwest Pacific (Schellart and Spakman, 2015). For comparison, the velocities have also been plotted in the global moving hotspot reference frame from Doubrovine et al. (2012). The velocity profile for the north-ward velocity component is roughly comparable in shape to the Indo-Atlantic velocity profile in the period 65–49 Ma (Fig. 2B), but the total acceleration, peak velocity, and total deceleration are reduced.

In Schellart and Spakman (2015) geological and geophysical evi-dence was presented to propose the existence and the timing of activity (ca. 71–50 Ma) of the fossil New Guinea–Pocklington subduction zone; it was argued that this subduction zone was largely responsible for velocity changes of the Australian plate, with an increase in speed from ~64 Ma to

~59 Ma due to slab lengthening, rapid plate motion ca. 59–52 Ma during mature subduction, and slowing ca. 52–49 Ma due to slab detachment. The reference numerical model A is largely consistent with this conceptual model, but also shows that part of the slowing (~27%) can be ascribed to the phase of continental subduction prior to slab detachment.

The numerical geodynamic model can explain the subducting plate velocity changes and velocity magnitudes in the period 65–45 Ma, but does not explain the significant north-directed plate acceleration that started ca. 44 Ma. This new phase of plate acceleration has been ascribed to the formation of new subduction zones along the northwestern, northern, and northeastern boundaries of the Australian plate between ca. 50 and ca. 40 Ma (Schellart and Spakman, 2015), and this phase of renewed subduction has not been included in the numerical model. Formation of the north- to northeast-dipping Sunda subduction zone ca. 45 Ma (Hall, 2012), North Sulawesi–Halmahera subduction zone ca. 45 Ma (Hall, 2012), and north-east- to east-dipping New Caledonia–Northland subduction zone at 50–40 Ma (Schellart and Spakman, 2012) would have resulted in increasing slab

pull forces, much like the free slab sinking phase and subducting plate acceleration shown in Figure 5, causing fast plate acceleration at 45–40 Ma and fast north- to northeast-directed plate velocities since ca. 40 Ma.

Continental Subduction and Slab Detachment

Velocities and Duration of Continental Subduction and Slab Detachment

Reference model A with slab detachment shows a phase of continental subduction and slab detachment lasting only ~3.9 m.y., with a decrease in v

SP⊥ of ~5.5 cm/yr. Earlier works on slab detachment generally show comparable reductions in subducting plate velocity (e.g., Burkett and Bil-len, 2010; van Hunen and Allen, 2011; Capitanio and Replumaz, 2013; Capitanio, 2014). Normal (oceanic) subduction models from Burkett and Billen (2010) show comparable reductions in v

SP⊥ of ~5–9 cm/yr during slab detachment, as do the continental subduction break-off models of Capitanio (2014) and Capitanio and Replumaz (2013) with a reduction in v

SP⊥ of ~5 cm/yr and a reduction in convergence velocity of 2–4 cm/yr, respectively. The continental subduction break-off model of van Hunen and Allen (2011) shows a higher reduction of ~14 cm/yr, but their model generally shows much faster subducting plate velocities, to ~23 cm/yr.

The duration of slab detachment for model A is relatively short (~0.8–1.0 m.y.), but similar short durations (~0.3–3.0 m.y.) have been reported in earlier works on 2D numerical modeling of subduction and slab detach-ment (Duretz et al., 2012, 2014), as well as 2D and 3D numerical modeling of normal (oceanic) subduction (~1–3 m.y.) by Burkett and Billen (2010). The 2D numerical models of subduction followed by continental subduc-tion of van Hunen and Allen (2011) report a long delay time between the start of continental subduction and slab detachment (in the range ~9–23 m.y.), much longer than reported here (~3.9 m.y.), which is possibly due to the higher yield stress in their models (400 MPa) compared to the one used here for reference model A (352.8 MPa). Model B, with a yield stress of 382.2 MPa, did not show any yielding for the duration of the simulation, which was allowed to continue for ~9.5 m.y. since the start of continental subduction.

The models presented here have a 2D spatial set-up, and thus represent a subduction scenario in which slab detachment occurs contemporane-ously along the entire lateral extent of the subduction zone. Earlier work,

10-3 10-2 10-1 100 101 102 103 104


A

B

C

4.3 m.y.

14.3 m.y.

20.3 m.y.

0 km

500

100010000 km 2000 3000 4000 5000 6000 10000 km 2000 3000 4000 5000 6000

0 km

500

1000

0 km

500

1000

Figure 7. Model results showing three evolutionary stages of numerical subduction model B with a relatively strong core layer (C = 382.2 MPa), which does not develop slab detachment. (A) Early stage of the initial transient subduction phase (4.3 m.y.). (B) Early phase of continental subduction (14.3 m.y.). (C) Very late stage of continental subduction phase (20.3 m.y.), when subduct-ing plate velocity is close to zero. Images on the left show the different domains in the model, while images on the right show the nondimensional effective viscosity field.




however, proposed that slab detachment and slab tears migrate subhori-zontally (e.g., Wortel and Spakman, 2000). Numerical subduction models with a 3D spatial set-up have reproduced this process, demonstrating that lateral tear migration can be very rapid, as much as ~80 cm/yr (van Hunen and Allen, 2011). For a 4000-km-wide subduction zone this would imply that slab detachment can take place in 2.5 m.y. if it initiated at one loca-tion in the center and migrated outward. It is plausible that along such a wide subduction zone, slab detachment starts at multiple locations, and migrates laterally from these places, thereby reducing the time of detach-ment along the entire extent of the subduction zone.

Earlier work has found that temperature effects on slab detachment duration are relatively small (decrease of ~8%–10%) (Gerya et al., 2004; Duretz et al., 2012), indicating that the duration of slab detachment for the isothermal model A presented here is of the right order of magnitude and would only be marginally affected in case thermal gradients were incorporated. Instead, the variability in duration of slab detachment has been ascribed mainly to differences in rheology and yield stress of the slab material (e.g., Andrews and Billen, 2009; Duretz et al., 2012). The yield stress of 352.8 MPa used here for model A with slab detachment is comparable to that reported by Andrews and Billen (2009; 300 MPa) to allow for slab detachment.

Model B without slab detachment shows a decrease in vSP⊥ from ~4.5

cm/yr to ~0 cm/yr during continental subduction over an ~8 m.y. period. Laboratory models of subduction and late-stage continental subduction from Edwards et al. (2015) also lack slab detachment, and show periods of continental subduction that generally last longer, ~10–30 m.y. This can be partly ascribed to the neutral buoyancy of the continental lithospheric mantle in the current models compared to the negative buoyancy for the models from Edwards et al. (2015), giving the continental lithosphere in the current work a larger net positive buoyancy force, thereby reducing continental subduction. Most of the discrepancy, however, can be ascribed to the fact that 14 out of 15 models in Edwards et al. (2015) have a larger passive margin (250–500 km). The one model with the same passive

margin size as here (100 km) has a very comparable continental subduc-tion period of ~10 m.y.

Forces During Continental Subduction and Slab DetachmentIn reference model A, the slab detaches in a late stage of continen-

tal subduction, while there is no detachment in model B. This can be explained by the buoyancy forces that operate during continental subduc-tion and the different yield stress in models A and B. Figure 8A shows the evolution of the buoyancy force of the oceanic slab segment F

Bu-Oc

(negative) (per meter length of the trench) and the buoyancy force of the continental slab segment F

Bu-Co (positive) (per meter length of the trench).

These forces have been calculated with the following equations:

∆F L T gBu-Oc Oc Oc Oc= ρ , (1)and ∆F L T gBu-Co Co Co Co= ρ , (2)

where ΔρOc

is the density contrast between ambient mantle and oceanic slab (Δρ

Oc = –80 kg/m3), L

Oc is the length of the oceanic slab segment that

contributes to the buoyancy force (so excluding the segment subhorizon-tally overlying the 660 km discontinuity; this parameter changes with time), T

Oc is the thickness of the oceanic slab segment (changes with time

and space), g is gravitational acceleration (9.8 m/s2), ΔρCo

is the density contrast between ambient mantle and continental crust (Δρ

Co = 400 kg/m3),

LCo

is the length of the continental slab segment (changes with time), and T

Co is the thickness of the continental crustal slab segment (changes with

time and space; note that the continental lithospheric mantle is neutrally buoyant). As can be observed, subduction of oceanic lithosphere creates a negative (downward) buoyancy force with a minimum of ~–5.2 × 1013 N/m that is reached between the start of continental subduction and the start of slab detachment. Subduction of continental lithosphere creates a posi-tive (upward) buoyancy force with a maximum of ~6.4 × 1012 N/m that is reached during slab detachment. The differential tensional buoyancy force

F Bu-Diff

12% FBu-Diff

10%

0 2 4 6 8 10 12 14 16 18 20 22Time [m.y.]

-4-5-6

-3

-101234567

Forc

e pe

r met

er tr

ench

[x 1

0

N/m

]



shears form)

Slabdetachment

complete

A

0 2 4 6 8 10 12 14 16 18 20 22Time [m.y.]

0

1

2

3

4

5

6

7

8BFBu-Oc

FBu-Co

-2

13

FBu-Diff

Forc

e pe

r met

er tr

ench

[x 1

0

N/m

]12



shears form)

Slabdetachment

complete

FS-MaxFSP-Eff

Figure 8. Model results showing the evolution of the forces in reference model A. (A) Development of the buoyancy force of the oceanic slab segment (FBu-Oc), the buoyancy force of the continental slab segment (FBu-Co), and the differential buoyancy force (FBu-Diff = FBu-Co – FBu-Oc) with progressive time. (B) Development of the effective slab pull force (FSP-Eff) and the yield force in the subducted slab (FS-Max) with pro-gressive time. Note that all the forces are per meter length of the subduction zone trench. Error bars are indicated for data points of FBu-Oc, FBu-Co, FBu-Diff, and FS-Max (based on maximum errors in measuring TOc, LOc, TCo and LCo; see text for definitions). (For discussion see Forces During Continental Subduction and Slab Detachment.)


SCHELLART


(per meter trench length) between the continental slab segment and oceanic slab segment, with F

Bu-Diff = F

Bu-Co – F

Bu-Oc, is focused in the boundary region

between the oceanic and continental slab segments and reaches a peak at the onset of slab detachment with F

Bu-Diff = ~5.6 × 1013 N/m (Fig. 8A).

The strength of the slab is effectively determined by the yield stress of the strong core layer of the subducting plate (layer 3 and Top LithM in Table 1), which is 352.8 MPa. With such a yield stress and a 15–20-km-thick core layer in the vicinity of the ocean-continent transition during the continental subduction phase, this gives a yield force, the maximum in-plane slab pull force (per meter trench length) in the slab, F

S-Max, of

5.3–7.1 × 1012 N/m. The evolution of this yield force for model A is plot-ted in Figure 8B. We can see that F

S-Max is about an order of magnitude

smaller than FBu-Diff

, which might suggest at first that the slab is much too weak to sustain such a large buoyancy force and that slab detachment should have occurred in a much earlier phase of subduction. A significant part of the buoyancy force, however, is sustained elsewhere in the sys-tem, in particular by viscous forces in the ambient mantle (Conrad and Lithgow-Bertelloni, 2002; Schellart, 2004b; Krien and Fleitout, 2008; Leng and Zhong, 2010). Earlier work showed that only ~10%–12% of the buoyancy forces is transmitted in the plane of the subducted slab toward the trailing plate (Schellart, 2004b; Sandiford et al., 2005). This would imply that the effective slab pull force (per meter trench length) F

SP-Eff =

0.10–0.12 × FBu-Diff

. Figure 8B shows the evolution of FSP-Eff

, demonstrat-ing that a maximum effective slab pull force is reached at the start of slab detachment, with F

SP-Eff = 5.6–6.7 × 1012 N/m. Figures 6C and 8 further

show that slab detachment starts at a time when a significant section of continental lithosphere has been subducted (L

Co ≈128 km) and when F

S-Max

≈0.10FBu-Diff

, indicating that at this time the effective slab pull force is large enough to start rapid shearing, necking, and yielding in the slab. Before slab detachment, the amount of subducted continental lithosphere was insufficient to create a F

Bu-Diff of sufficient magnitude, and the slab was

not sufficiently stretched yet to decrease FS-Max

, giving FSP-Eff

< FS-Max

, and so slab detachment was not possible yet.

Applying calculations as above to model B, for a late stage as depicted in Figure 7C, F

Bu-Oc = –3.40 × 1013 N/m (±9.5%), F

Bu-Co = 1.07 × 1013 N/m

(±17.2%), and FBu-Diff

= 4.47 × 1013 N/m (±26.6%). This would imply an effective slab pull force (per meter trench length) F

SP-Eff = 4.47–5.37 ×

1012 N/m. This is less than the estimated yield force (per meter trench length) in the slab of F

S-Max = 7.64 × 1012 N/m (±5%), thereby explaining

why slab detachment did not occur in model B.

Continental Subduction DepthModels A and B show different maximum depths of continental sub-

duction, 96 km and 144 km, respectively, which can simply be explained by the weaker subducting plate in A compared to B. The weaker plate in A resulted in slab detachment prior to the continental slab segment reach-ing its maximum depth that could have been reached with the available negative buoyancy force of the oceanic slab segment. This resulted in a shorter duration of slab pull during continental subduction and thus a shallower depth of continental subduction.

The observed continental subduction depths reported here are compa-rable to, although in the lower range of, depths reported in previous mod-eling studies of buoyancy-driven subduction (e.g., van Hunen and Allen, 2011; Sizova et al., 2012; Edwards et al., 2015). For example, Edwards et al. (2015) reported depths in the range 70–560 km for experiments with different passive margin width, continental crustal thickness, and oceanic lithospheric thickness. Their experiment with a comparable passive margin width (100 km), continental crustal thickness (30 km), density contrast (366 kg/m3), and oceanic lithospheric thickness (100 km) showed a low conti-nental subduction depth (70 km), comparable to the work reported here.

CONCLUSIONS

In this work a numerical geodynamic model of subduction, continental subduction, and slab detachment has been presented to simulate the sub-duction evolution at the northern margin of the Australian plate during the latest Cretaceous and early Cenozoic, and to quantify the velocity of the Australian plate. The numerical model produces a comparable subducting plate velocity evolution as documented for the Australian plate between ca. 66 and 49 Ma, with a first phase of plate acceleration due to progres-sive upper mantle slab lengthening, which increases the slab pull force, a period of relatively rapid plate motion during mature subduction, and a phase of plate deceleration due to continental subduction and subsequent slab detachment. Most (~73%) of the plate deceleration is caused during the short (~0.8 m.y.) phase of slab detachment, while a smaller compo-nent of plate deceleration (~27%) is caused by the preceding phase of continental subduction, which lasted ~3.1 m.y. A geodynamic subduction model that shows continental subduction but lacks a final phase of slab detachment is characterized by a slower and more gradual plate decelera-tion that lasts ~8 m.y., which is in conflict with observations of fast plate deceleration for the Australian plate between ca. 52 and 49 Ma. It is thus concluded that slab detachment is an essential element to explain the rapid plate deceleration of the Australian plate. This conclusion is consistent with the evidence presented in Schellart and Spakman (2015) that the fossil New Guinea–Pocklington slab detached in the early Eocene and is currently located in the upper part of the lower mantle below central and southeastern Australia.

More generally, I propose here that rapid plate deceleration (~3–5 cm/yr) over periods lasting not more than ~1 m.y. can be ascribed to the process of slab detachment following continental subduction. It is thus conceivable that rapid decelerations that have occurred in the geological past for other tectonic plates have a comparable underlying cause, namely slab detachment.

ACKNOWLEDGMENTSI thank Anne Replumaz and two anonymous reviewers for their constructive and helpful com-ments. This research was supported through a Vici Fellowship from the Dutch Science Founda-tion (NWO) and has been supported by computational resources from the NCI National Facility in Australia through the National Computational Merit Allocation Scheme (projects ei8 and qk0).

REFERENCES CITEDAndrews, E.R., and Billen, M.I., 2009, Rheologic controls on the dynamics of slab detachment:

Tectonophysics, v. 464, p. 60–69, doi: 10 .1016 /j .tecto .2007 .09 .004.Argus, D.F., and Gordon, R.G., 1991, No-net-rotation model of current plate velocities incor-

porating plate motion model Nuvel-1: Geophysical Research Letters, v. 18, p. 2039–2042, doi: 10 .1029 /91GL01532.

Argus, D.F., Gordon, R.G., and Demets, C., 2011, Geologically current motion of 56 plates relative to the no-net-rotation reference frame: Geochemistry Geophysics Geosystems, v. 12, Q11001, doi: 10 .1029 /2011GC003751.

Baldwin, S.L., Fitzgerald, P.G., and Webb, L.E., 2012, Tectonics of the New Guinea region: Annual Review of Earth and Planetary Sciences, v. 40, p. 495–520, doi: 10 .1146 /annurev

-earth -040809 -152540.Becker, T.W., Schaeffer, A.J., Lebedev, S., and Conrad, C.P., 2015, Toward a generalized plate

motion reference frame: Geophysical Research Letters, v. 42, p. 3188–3196, doi: 10 .1002 /2015GL063695.

Bird, P., 2003, An updated digital model of plate boundaries: Geochemistry, Geophysics, Geo-systems, v. 4, 1027, doi: 1010 .1029 /2001GC000252.

Burkett, E.R., and Billen, M.I., 2010, Three‐dimensionality of slab detachment due to ridge‐trench collision: Laterally simultaneous boudinage versus tear propagation: Geochemistry Geophysics Geosystems, v. 11, Q11012, doi: 10 .1029 /2010GC003286.

Butterworth, N.P., Talsma, A.S., Müller, R.D., Seton, M., Bunge, H.-P., Schuberth, B.S.A., Sheph-ard, G.E., and Heine, C., 2014, Geological, tomographic, kinematic and geodynamic con-straints on the dynamics of sinking slabs: Journal of Geodynamics, v. 73, p. 1–13, doi: 10 .1016 /j .jog .2013 .10 .006.

Cande, S.C., and Stegman, D.R., 2011, Indian and African plate motions driven by the push force of the Reunion plume head: Nature, v. 475, p. 47–52, doi: 10 .1038 /nature10174.

Capitanio, F.A., 2014, The dynamics of extrusion tectonics: Insights from numerical modeling: Tectonics, v. 33, p. 2361–2381, doi: 10 .1002 /2014TC003688.

Capitanio, F.A., and Replumaz, A., 2013, Subduction and slab breakoff controls on Asian in-dentation tectonics and Himalayan western syntaxis formation: Geochemistry Geophys-ics Geosystems, v. 14, p. 3515–3531, doi: 10 .1002 /ggge .20171.




Capitanio, F.A., Stegman, D.R., Moresi, L.N., and Sharples, W., 2010, Upper plate controls on deep subduction, trench migrations and deformations at convergent margins: Tectono-physics, v. 483, p. 80–92, doi: 10 .1016 /j .tecto .2009 .08 .020.

Capitanio, F.A., Replumaz, A., and Riel, N., 2015, Reconciling subduction dynamics during Tethys closure with large-scale Asian tectonics: Insights from numerical modeling: Geo-chemistry Geophysics Geosystems, v. 16, p. 962–982, doi: 10 .1002 /2014GC005660.

Carlson, R.L., Hilde, T.W.C., and Uyeda, S., 1983, The driving mechanism of plate tecton-ics: Relation to age of the lithosphere at trenches: Geophysical Research Letters, v. 10, p. 297–300, doi: 10 .1029 /GL010i004p00297.

Chen, Z., Schellart, W.P., and Duarte, J.C., 2015, Overriding plate deformation and variability of fore-arc deformation during subduction: Insight from geodynamic models and ap-plication to the Calabria subduction zone: Geochemistry Geophysics Geosystems, v. 16, p. 3697–3715, doi: 10 .1002 /2015GC005958.

Chen, Z., Schellart, W.P., Strak, V., and Duarte, J.C., 2016, Does subduction-induced mantle flow drive backarc extension?: Earth and Planetary Science Letters, v. 441, p. 200–210, doi: 10 .1016 /j .epsl .2016 .02 .027.

Chertova, M.V., Spakman, W., Geenen, T., van den Berg, A.P., and van Hinsbergen, D.J.J., 2014, Underpinning tectonic reconstructions of the western Mediterranean region with dynamic slab evolution from 3-D numerical modeling: Journal of Geophysical Research, v. 119, p. 5876–5902, doi: 10 .1002 /2014JB011150.

Cloos, M., 1993, Lithospheric buoyancy and collisional orogenesis: Subduction of oceanic plateaus, continental margins, island arcs, spreading ridges and seamounts: Geological Society of America Bulletin, v. 105, p. 715–737, doi: 10 .1130 /0016 -7606 (1993) 105 <0715: LBACOS >2 .3 .CO;2.

Conrad, C.P., and Behn, M.D., 2010, Constraints on lithosphere net rotation and asthenospheric viscosity from global mantle flow models and seismic anisotropy: Geochemistry Geo-physics Geosystems, v. 11, Q05W05, doi: 10 .1029 /2009GC002970.

Conrad, C.P., and Lithgow-Bertelloni, C., 2002, How mantle slabs drive plate tectonics: Sci-ence, v. 298, p. 207–209, doi: 10 .1126 /science .1074161.

Crameri, F.H., Schmeling, H., Golabek, G.J., Duretz, T., Orendt, R., Buiter, S.J.H., May, D.A., Kaus, B.J.P., Gerya, T.V., and Tackley, P.J., 2012, A comparison of numerical surface topog-raphy calculations in geodynamic modelling: An evaluation of the “sticky air” method: Geophysical Journal International, v. 189, p. 38–54, doi: 10 .1111 /j .1365 -246X .2012 .05388 .x.

Davies, H.L., 2012, The geology of New Guinea—The cordilleran margin of the Australian continent: Episodes, v. 35, p. 87–102.

Davies, H.L., and Warren, R.G., 1992, Eclogites of the D’Entrecasteaux Islands: Contributions to Mineralogy and Petrology, v. 112, p. 463–474, doi: 10 .1007 /BF00310778.

DeMets, C., Gordon, R.G., Argus, D.F., and Stein, S., 1994, Effect of recent revisions to the geomagnetic reversal time scale on estimates of current plate motions: Geophysical Research Letters, v. 21, p. 2191–2194, doi: 10 .1029 /94GL02118.

Doubrovine, P.V., Steinberger, B., and Torsvik, T.H., 2012, Absolute plate motions in a reference frame defined by moving hot spots in the Pacific, Atlantic, and Indian Oceans: Journal of Geophysical Research, v. 117, p. B09101, doi: 10 .1029 /2011JB009072.

Duarte, J.C., Schellart, W.P., and Cruden, A.R., 2013, Three-dimensional dynamic laboratory models of subduction with an overriding plate and variable interplate rheology: Geo-physical Journal International, v. 195, p. 47–66, doi: 10 .1093 /gji /ggt257.

Duarte, J.C., Schellart, W.P., and Cruden, A.R., 2015, How weak is the subduction zone in-terface?: Geophysical Research Letters, v. 42, p. 2664–2673, doi: 10 .1002 /2014GL062876.

Duretz, T., Schmalholz, S.M., and Gerya, T.V., 2012, Dynamics of slab detachment: Geochem-istry Geophysics Geosystems, v. 13, Q03020, doi: 10 .1029 /2011GC004024.

Duretz, T., Gerya, T.V., and Spakman, W., 2014, Slab detachment in laterally varying subduc-tion zones: 3-D numerical modeling: Geophysical Research Letters, v. 41, p. 1951–1956, doi: 10 .1002 /2014GL059472.

Edwards, S.J., Schellart, W.P., and Duarte, J.C., 2015, Geodynamic models of continental subduction and obduction of overriding plate forearc oceanic lithosphere on top of con-tinental crust: Tectonics, v. 34, p. 1494–1515, doi: 10 .1002 /2015TC003884.

Forsyth, D.W., and Uyeda, S., 1975, On the relative importance of the driving forces of plate motion: Royal Astronomical Society Geophysical Journal, v. 43, p. 163–200, doi: 10 .1111 /j .1365 -246X .1975 .tb00631 .x.

Funiciello, F., Faccenna, C., and Giardini, D., 2004, Role of lateral mantle flow in the evolu-tion of subduction systems: insights from laboratory experiments: Geophysical Journal International, v. 157, p. 1393–1406, doi: 10 .1111 /j .1365 -246X .2004 .02313 .x.

Gaina, C., Müller, D.R., Royer, J.-Y., Stock, J., Hardebeck, J.L., and Symonds, P., 1998, The tectonic history of the Tasman Sea; a puzzle with 13 pieces: Journal of Geophysical Re-search, v. 103, p. 12,413–12,433, doi: 10 .1029 /98JB00386.

Gerya, T.V., Yuen, D.A., and Maresch, W.V., 2004, Thermomechanical modelling of slab detach-ment: Earth and Planetary Science Letters, v. 226, p. 101–116, doi: 10 .1016 /j .epsl .2004 .07 .022.

Goes, S., Capitanio, F.A., and Morra, G., 2008, Evidence of lower-mantle slab penetration phases in plate motions: Nature, v. 451, p. 981–984, doi: 10 .1038 /nature06691.

Gordon, R.G., and Jurdy, D.M., 1986, Cenozoic global plate motions: Journal of Geophysical Research, v. 91, p. 12,389–12,406, doi: 10 .1029 /JB091iB12p12389.

Gripp, A.E., and Gordon, R.G., 2002, Young tracks of hotspots and current plate velocities: Geophysical Journal International, v. 150, p. 321–361, doi: 10 .1046 /j .1365 -246X .2002 .01627 .x.

Guillaume, B., Martinod, J., and Espurt, N., 2009, Variations of slab dip and overriding plate tectonics during subduction: Insights from analogue modelling: Tectonophysics, v. 463, p. 167–174, doi: 10 .1016 /j .tecto .2008 .09 .043.

Hall, R., 2012, Late Jurassic–Cenozoic reconstructions of the Indonesian region and the Indian Ocean: Tectonophysics, v. 570–571, p. 1–41, doi: 10 .1016 /j .tecto .2012 .04 .021.

Jacoby, W.R., 1973, Model experiment of plate movements: Nature. Physical Science (Lon-don), v. 242, p. 130–134, doi: 10 .1038 /physci242130a0.

Jagoutz, O., Royden, L., Holt, A.F., and Becker, T.W., 2015, Anomalously fast convergence of India and Eurasia caused by double subduction: Nature Geoscience, v. 8, p. 475–478, doi: 10 .1038 /ngeo2418.

Keep, M., and Schellart, W.P., 2012, Introduction to the thematic issue on the evolution and dynamics of the Indo-Australian plate: Australian Journal of Earth Sciences, v. 59, p. 807–808, doi: 10 .1080 /08120099 .2012 .708360.

Klootwijk, C.T., Gee, J.S., Peirce, J.W., Smith, G.M., and McFadden, P.L., 1992, An early India-Asia contact: Paleomagnetic constraints from Ninety east Ridge, ODP Leg 121: Geology, v. 20, p. 395–398, doi: 10 .1130 /0091 -7613 (1992)020 <0395: AEIACP>2 .3 .CO;2.

Kreemer, C., 2009, Absolute plate motions constrained by shear wave splitting orientations with implications for hot spot motions and mantle flow: Journal of Geophysical Research, v. 114, B10405, doi: 10 .1029 /2009JB006416.

Kreemer, C., Holt, W.E., and Haines, A.J., 2003, An integrated global model of present-day plate motions and plate boundary deformation: Geophysical Journal International, v. 154, p. 8–34, doi: 10 .1046 /j .1365 -246X .2003 .01917.x.

Krien, Y., and Fleitout, L., 2008, Gravity above subduction zones and forces controlling plate motions: Journal of Geophysical Research, v. 113, B09407, doi: 10 .1029 /2007JB005270.

Leng, W., and Zhong, S., 2010, Constraints on viscous dissipation of plate bending from compressible mantle convection: Earth and Planetary Science Letters, v. 297, p. 154–164, doi: 10 .1016 /j .epsl .2010 .06 .016.

Long, M.D., and Silver, P.G., 2009, Mantle flow in subduction systems: The subslab flow field and implications for mantle dynamics: Journal of Geophysical Research, v. 114, B10312, doi: 10 .1029 /2008JB006200.

Mackwell, S.J., Bai, Q., and Kohlstedt, D.L., 1990, Rheology of olivine and the strength of the lithosphere: Geophysical Research Letters, v. 17, p. 9–12, doi: 10 .1029 /GL017i001p00009.

Meyer, C., and Schellart, W.P., 2013, Three-dimensional dynamic models of subducting plate-overriding plate-upper mantle interaction: Journal of Geophysical Research, v. 118, p. 775–790, doi: 10 .1002 /jgrb .50078.

Minster, J.B., and Jordan, T.H., 1978, Present-day plate motions: Journal of Geophysical Re-search, v. 83, p. 5331–5354, doi: 10 .1029 /JB083iB11p05331.

Molnar, P., and Stock, J.M., 2009, Slowing of India’s convergence with Eurasia since 20 Ma and its implications for Tibetan mantle dynamics: Tectonics, v. 28, TC3001, doi: 10 .1029 /2008TC002271.

Moresi, L., Quenette, S., Lemiale, V., Mériaux, C., Appelbe, B., and Mühlhaus, H.-B., 2007, Computational approaches to studying non-linear dynamics of the crust and mantle: Physics of the Earth and Planetary Interiors, v. 163, p. 69–82, doi: 10 .1016 /j .pepi .2007 .06 .009.

Moresi, L., Betts, P.G., Miller, M.S., and Cayley, R.A., 2014, Dynamics of continental accretion: Nature, v. 508, p. 245–248, doi: 10 .1038 /nature13033.

Müller, R.D., Sdrolias, M., Gaina, C., and Roest, W.R., 2008, Age, spreading rates, and spread-ing asymmetry of the world’s ocean crust: Geochemistry, Geophysics, Geosystems, v. 9, Q04006, doi: 10 .1029 /2007GC001743.

O’Neill, C., Müller, D., and Steinberger, B., 2005, On the uncertainties in hot spot reconstruc-tions and the significance of moving hot spot reference frames: Geochemistry, Geophys-ics, Geosystems, v. 6, Q04003, doi: 10 .1029 /2004GC000784.

Patriat, P., and Achache, J., 1984, India-Eurasia collision chronology has implications for crustal shortening and driving mechanism of plates: Nature, v. 311, p. 615–621, doi: 10 .1038 /311615a0.

Royer, J.-Y., and Sandwell, D.T., 1989, Evolution of the eastern Indian Ocean since the Late Cretaceous; constraints from Geosat altimetry: Journal of Geophysical Research, v. 94, p. 13,755–13,782, doi: 10 .1029 /JB094iB10p13755.

Sandiford, M., Coblentz, D., and Schellart, W.P., 2005, Evaluating slab-plate coupling in the Indo-Australian plate: Geology, v. 33, p. 113–116, doi: 10 .1130 /G20898 .1.

Schellart, W.P., 2004a, Kinematics of subduction and subduction-induced flow in the upper mantle: Journal of Geophysical Research, v. 109, B07401, doi: 10 .1029 /2004JB002970.

Schellart, W.P., 2004b, Quantifying the net slab pull force as a driving mechanism for plate tectonics: Geophysical Research Letters, v. 31, L07611, doi: 10 .1029 /2004GL019528.

Schellart, W.P., 2008, Kinematics and flow patterns in deep mantle and upper mantle subduc-tion models: Influence of the mantle depth and slab to mantle viscosity ratio: Geochem-istry Geophysics Geosystems, v. 9, Q03014, doi: 10 .1029 /2007GC001656.

Schellart, W.P., 2011, A subduction zone reference frame based on slab geometry and sub-duction partitioning of plate motion and trench migration: Geophysical Research Letters, v. 38, L16317, doi: 10 .1029 /2011GL048197.

Schellart, W.P., and Moresi, L., 2013, A new driving mechanism for backarc extension and backarc shortening through slab sinking induced toroidal and poloidal mantle flow: Re-sults from dynamic subduction models with an overriding plate: Journal of Geophysical Research, v. 118, p. 3221–3248, doi: 10 .1002 /jgrb .50173.

Schellart, W.P., and Spakman, W., 2012, Mantle constraints on the plate tectonic evolution of the Tonga-Kermadec-Hikurangi subduction zone and the South Fiji Basin: Australian Journal of Earth Sciences, v. 59, p. 933–952, doi: 10 .1080 /08120099 .2012 .679692.

Schellart, W.P., and Spakman, W., 2015, Australian plate motion and topography linked to fossil New Guinea slab below Lake Eyre: Earth and Planetary Science Letters, v. 421, p. 107–116, doi: 10 .1016 /j .epsl .2015 .03 .036.

Schellart, W.P., Stegman, D.R., and Freeman, J., 2008, Global trench migration velocities and slab migration induced upper mantle volume fluxes: Constraints to find an Earth reference frame based on minimizing viscous dissipation: Earth-Science Reviews, v. 88, p. 118–144, doi: 10 .1016 /j .earscirev .2008 .01 .005.

Schellart, W.P., Stegman, D.R., Farrington, R.J., Freeman, J., and Moresi, L., 2010, Cenozoic tectonics of western North America controlled by evolving width of Farallon slab: Sci-ence, v. 329, p. 316–319, doi: 10 .1126 /science .1190366.

Schellart, W.P., Stegman, D.R., Farrington, R.J., and Moresi, L., 2011, Influence of lateral slab edge distance on plate velocity, trench velocity and subduction partitioning: Journal of Geophysical Research, v. 116, B10408, doi: 10 .1029 /2011JB008535.

Schmeling, H., et al., 2008, A benchmark comparison of spontaneous subduction models—Towards a free surface: Physics of the Earth and Planetary Interiors, v. 171, p. 198–223, doi: 10 .1016 /j .pepi .2008 .06 .028.


SCHELLART


Seton, M., et al., 2012, Global continental and ocean basin reconstructions since 200 Ma: Earth-Science Reviews, v. 113, p. 212–270, doi: 10 .1016 /j .earscirev .2012 .03 .002.

Sizova, E., Gerya, T., and Brown, M., 2012, Exhumation mechanisms of melt-bearing ultrahigh pressure crustal rocks during collision of spontaneously moving plates: Journal of Meta-morphic Geology, v. 30, p. 927–955, doi: 10 .1111 /j .1525 -1314 .2012 .01004 .x.

Stegman, D.R., Freeman, J., Schellart, W.P., Moresi, L., and May, D., 2006, Influence of trench width on subduction hinge retreat rates in 3-D models of slab rollback: Geochemistry Geophysics Geosystems, v. 7, Q03012, doi: 10 .1029 /2005GC001056.

Stegman, D.R., Farrington, R., Capitanio, F.A., and Schellart, W.P., 2010, A regime diagram for subduction styles from 3-D numerical models of free subduction: Tectonophysics, v. 483, p. 29–45, doi: 10 .1016 /j .tecto .2009 .08 .041.

van der Meer, D.G., Spakman, W., van Hinsbergen, D.J.J., Amaru, M.L., and Torsvik, T.H., 2010, Towards absolute plate motions constrained by lower-mantle slab remnants: Nature Geoscience, v. 3, p. 36–40, doi: 10 .1038 /ngeo708.

van Hinsbergen, D.J.J., Steinberger, B., Doubrovine, P.V., and Gassmöller, R., 2011, Accelera-tion and deceleration of India‐Asia convergence since the Cretaceous: Roles of mantle plumes and continental collision: Journal of Geophysical Research, v. 116, B06101, doi: 10 .1029 /2010JB008051.

van Hunen, J., and Allen, M.B., 2011, Continental collision and slab break-off: A comparison of 3-D numerical models with observations: Earth and Planetary Science Letters, v. 302, p. 27–37, doi: 10 .1016 /j .epsl .2010 .11 .035.

Wessel, P., and Kroenke, L.W., 2008, Pacific absolute plate motion since 145 Ma: An assess-ment of the fixed hot spot hypothesis: Journal of Geophysical Research, v. 113, B06101, doi: 10 .1029 /2007JB005499.

Wessel, P., and Müller, R.D., 2016, Ridge-spotting: A new test for Pacific absolute plate mo-tion models: Geochemistry Geophysics Geosystems, v. 17, p. 2408–2420, doi: 10 .1002 /2016GC006404.

White, L.T., and Lister, G.S., 2012, The collision of India with Asia: Journal of Geodynamics, v. 56–57, p. 7–17, doi: 10 .1016 /j .jog .2011 .06 .006.

Wortel, M.J.R., and Spakman, W., 2000, Subduction and slab detachment in the Mediterra-nean-Carpathian region: Science, v. 290, p. 1910–1917, doi: 10 .1126 /science .290 .5498 .1910.

Worthing, M.A., and Crawford, A.J., 1996, The igneous geochemistry and tectonic setting of metabasites from the Emo metamorphics, Papua New Guinea; a record of the evolu-tion and destruction of a backarc basin: Mineralogy and Petrology, v. 58, p. 79–100, doi: 10 .1007 /BF01165765.

Yan, C.Y., and Kroenke, L.W., 1993, A plate tectonic reconstruction of the Southwest Pacific, 0–100 Ma, in Berger, W.H., et al., Proceedings of the Ocean Drilling Program, Scientific Results, Volume 130: College Station, Texas, Ocean Drilling Program, p. 697–709, doi: 10

.2973 /odp .proc .sr .130 .055 .1993.Zahirovic, S., Müller, R.D., Seton, M., Flament, N., Gurnis, M., and Whittaker, J., 2012, Insights

on the kinematics of the India-Eurasia collision from global geodynamic models: Geo-chemistry Geophysics Geosystems, v. 13, Q04W11, doi: 10 .1029 /2011GC003883.

MANUSCRIPT RECEIVED 22 APRIL 2017 REVISED MANUSCRIPT RECEIVED 17 JULY 2017 MANUSCRIPT ACCEPTED 24 AUGUST 2017

Printed in the USA


a geodynamic model of subduction evolution and slab

Documents