Earth and Planetary Science Letters 274 (2008) 392–400

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Earth and Planetary Science Letters

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Intraplate continental deformation: Influence of a heat-producing layer in thelithospheric mantle

Sérgio P. Neves a,⁎, Andréa Tommasi b, Alain Vauchez b, Riad Hassani c

a Departamento de Geologia, Universidade Federal de Pernambuco, 50740-530, Recife, Brazilb Géosciences Montpellier, CNRS & Université Montpellier 2, 34095 Montpellier, Francec Laboratoire de Géophysique Interne et Tectonophysique, CNRS & Université de Savoie, 73376 Le Bourget du Lac, France

⁎ Corresponding author.E-mail address: serpane@hotlink.com.br (S.P. Neves)

0012-821X/$ – see front matter © 2008 Elsevier B.V. Adoi:10.1016/j.epsl.2008.07.040

a b s t r a c t

a r t i c l e i n f o

Article history:

Deformation of the contine Received 1 November 2007Received in revised form 21 July 2008Accepted 25 July 2008Available online 2 September 2008

Keywords:

intracontinental deformationlithospheremantlemetasomatismheat productionstrain localization

ntal lithosphere is traditionally modeled assuming that heat production in thesubcontinental lithospheric mantle is negligible. Although this may be appropriate for highly depletedArchean lithosphere, studies of mantle xenoliths and lithospheric-derived mafic rocks suggest that heatproductions of up to 0.4 μW/m3 may be attained in the mantle of Phanerozoic and Neoproterozoic regionsdue to modal metasomatism. To investigate the effect of a local enrichment in heat-producing elementswithin the lithospheric mantle on the continental deformation, we performed a series of 2D thermo-mechanical models simulating the deformation of a lithospheric plate 1000 km long, which contains in itscentral part a 200 km wide and 20 km-thick mantle layer with a heat production of 0.05–0.25 μW/m3

between 40 and 80 km depth, leading to a local enhancement of the surface heat flow by 1–5 mW/m2.Compression models show that a 20 km-thick layer with a heat production as low as 0.05 μW/m3 within theshallow lithospheric mantle leads to strain localization in both the enriched mantle and the overlying crust.Strain localization depends exponentially on the temperature increase in the lithosphere section, which iscontrolled by the volume of metasomatized mantle and the intensity of the metasomatic enrichment in heat-producing elements. A heat production of 0.25 μW/m3 between 60 and 80 km depth results, for instance, instrains within and above the metasomatized mantle 5 times higher than in the surrounding lithosphere.Strain localization also depends on the location of the heat-producing domain and on the plate rheologicalstructure; deeper heat-producing domains and weaker plates leading to stronger localization. These resultsprovide an explanation for strain localization in intraplate environments, leading to the formation oforogenic belts situated hundreds to thousands of kilometers away from known plate boundaries, as observedin the Neoproterozoic belts of Gondwana and in the present-day deformation of the Asian plate.

© 2008 Elsevier B.V. All rights reserved.

1. Introduction

Large-scale deformation of continental lithosphere may extend faraway from convergent plate boundaries, as shown, for instance, by theactive intraplate deformation of central Asia (Yang and Liu, 2002;Cunningham, 2005). Intraplate settings have been suggested fororogenic belts of various ages since the Archean. Examples includeseveral Paleoproterozoic to Paleozoic orogenies in Australia (Sandifordand Hand, 1998; McLaren et al., 1999; Giles et al., 2002), thePaleoproterozoic Kapuskasing Structural Zone in Canada (Perryet al., 2006), the Neoproterozoic Damara belt in S. Africa (Coward,1976), and the Neoproterozoic Borborema Province in northeasternBrazil (Vauchez et al. 1995; Neves, 2003).

Strain distribution within continental plates is to first ordercontrolled by lateral variations in rheology (rheological heterogene-ities; e.g., Vauchez et al., 1994; Tommasi and Vauchez, 1997). Due to

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the lower thermal gradient and resulting high stiffness of cratonicdomains, deformation is mainly confined to regions of Proterozoic oryounger lithosphere. Weak intraplate rheological heterogeneities, dueto a thicker crust or a higher geotherm, also localize strain (Dunbarand Sawyer, 1988, Tommasi et al., 1995, Tommasi and Vauchez, 1997).An aspect that has been fully neglected in the study of the deformationof the continental lithosphere and that may also influence its responseto applied stresses is the effect of heat production in the lithosphericmantle. Thermo-mechanical models often assume that heat produc-tion in this layer is negligible (0–0.02 μW/m3). However, studiesof xenoliths and mantle-derived magmas suggest this may be anoversimplification.

The neoproterozoic Borborema province in northeast Brazil, forinstance, is characterized by a network of intracontinental transcur-rent shear zones that are spatially associated with magmaticintrusions (Vauchez et al., 1995). Dioritic rocks are widespread, eitherforming small isolated plutons or associated with large graniticbatoliths up to 2000 km2 in area. The geochemical and isotopiccharacteristics of the diorites points to an origin in the lithospheric

393S.P. Neves et al. / Earth and Planetary Science Letters 274 (2008) 392–400

mantle (Neves et al., 2000). Based on estimates of U, Th and K contentsin the source region of the diorites, Neves and Mariano (2004)calculated heat productions of up to 0.25 μW/m3. These results aresimilar to the lithospheric mantle heat production of up to 0.4 μW/m3

estimated from U, Th and K contents of mantle xenoliths fromPhanerozoic volcanic regions in Australia, Europe and Alaska (O’Reillyand Griffin, 2000). In the lithospheric mantle, high concentrations oflarge-ion lithophile elements (LILE), including the heat-producingelements (HPE) K, U and Th, can be stored in the structure of mineralslike apatite (O’Reilly and Griffin, 2000), amphibole, pyroxene andphlogopite (Eggins et al., 1998; Harlow and Davies, 2004; Pearsonet al., 2004), in reaction rims of phlogopite and titanium oxides atthe surface of spinel crystals (Bodinier et al., 1996), and/or at grainboundaries (Hiraga et al., 2004).

Since the lithosphere rheology is strongly dependent on itsthermal state, a local increase in heat production in the lithosphericmantle may lead to strain localization and allow intracontinentaldeformation. The aim of this paper is to investigate the influence ofradioactive heat production in the mantle on the mechanical behaviorof a continental lithospheric plate. The processes allowing to generateand preserve an enrichment in heat-producing elements within thelithospheric mantle are discussed. Then, we perform 2D numericalmodels of deformation of a continental plate of normal thicknesscontaining a HPE-enriched domain in the lithospheric mantle andcompare the resulting deformation field to observations in intracon-tinental orogenic belts.

2. Generation and long-term preservation of HPE-enriched mantle

HPE-enrichment is produced by interaction of the lithosphericmantlewith fluids or H2O- or CO2-richmagmas. Several scenariosmaybe envisioned. Transport of mantle-derived basalts through thecontinental lithosphere may result in refertilization (enrichment inclinopyroxene or plagioclase). This process however does not neces-sarily produce HPE-enrichment (e.g., Le Roux et al. 2007). Semi-continuous infiltration of small melt fractions due to partial melting ofthe asthenosphere (McKenzie, 1989) and metasomatic enrichment byfluids or melts derived either from mantle plumes or from slabdehydration/partial melting in subduction zones are more prone toproduce such enrichment. Studies of mantle xenoliths suggest that thelithosphere above a mantle plume may undergo significant change incomposition (Baker et al., 1998; Witt-Eickschen et al., 2003; Tommasiet al., 2004). Similarly, observations and thermal modeling of modernsubduction zones provide strong evidence for extensive metasoma-tism of the lithosphere above subducting slabs (Peacock, 1993).Analysis of subduction-related peridotites massifs, such as Finero inthe Alps, also show clear evidence for strong modal metasomatism byhydrous or CO2-bearing fluids, characterized by crystallization ofphlogopite, amphibole, and apatite (Morishita et al., 2008).

In normal subduction zones, metasomatismwill preferentially affectthe base of the lithosphere of a laterally limited region (a few hundredkilometres at most). Enrichment in HPE at shallow levels of thelithosphere may occur by underthrusting of arc lithosphere duringcollisionof apassive continentalmarginwith an island arc (Peltonen andBrügmann, 2006), but this process is also of limited lateral extent. Anefficient way to produce extensive HPE-enrichment at shallow mantledepths is flat subduction. Indeed, flat subduction leads to a laterallyexpanded volatiles release domain, resulting in metasomatism of theoverlaying plate up to several hundred kilometers away from the trench.Flat subduction occurs in 10% of modern convergent margins (Gutscheret al., 2000), but it might have been more common in the Precambrianbecause of increased buoyancy of oceanic lithosphere resulting fromyoung age, presence of oceanic plateaus, and greater thickness ofoceanic crust (e.g., Abbott et al., 1994; Gutscher et al., 2000; Kerrich andPolat, 2006). On the other hand, numerical models show that formantletemperatures 75 °C higher than at present, steep subduction is favored

due to lower asthenosphere viscosities, suggesting that shallowsubduction might have been more common in the Proterozoic than inthe Archean (van Hunen et al., 2004).

Themain factor controlling the potential of preservation of a radio-active mantle layer is probably the interaction with the convectingmantle. An enriched layer at the base of the lithosphere, althoughpossibly buoyant due to its higher temperature, is weak and can beentrained by asthenospheric flow (e.g., King, 2005; Michaut andJaupart, 2007). Radioactive elements must therefore be trapped inthe shallow lithospheric mantle in order to escape recycling in theconvective mantle. This probably favors flat subduction as the mainmechanism by which radioactive subcontinental mantle can be estab-lished and preserved. Cessation of flat subduction (through collision orotherwise) or transition from flat to normal subduction will bothresult, subsequently to thermal equilibration, in thickening of themantle lithosphere, placing the metasomatic layer in the middle orupper mantle lithosphere.

The influence of HPE-enriched mantle on continental rheologydepends on how long the thermal anomaly produced by its presencewill persist. Due to the exponential decay in radioactive elements,after an initial period of fast reduction, a protracted time will followwhere decrease in heat production will be much slower. Reducingheat production to a third of its initial value takes about 3 Ga (Jaupartand Mareschal, 1999). Since heat production of up to 0.4 μW/m3

has been estimated for some segments of Phanerozoic continentalmantle lithosphere (O’Reilly and Griffin, 2000), this indicates that oncea HPE-enriched reservoir is formed, the resulting thermal anomalymay subsist for a long time.

3. Modeling approach

We use the 2D thermo-mechanical finite-element (FE) code Adeli2D(Hassani, 1994; Hassani et al., 1997; Chéry et al., 2004; Chéry andHassani, 2005) to investigate the deformation of a lithospheric plateinwhich themantle is locally enriched in radioactive elements. Adeli2Dsolves for the quasi-static mechanical behavior of the lithosphereusing an explicit scheme based on the Dynamic Relaxation Method(Cundall, 1988).

Viscoplastic or viscoelastic deformation, depending on whichbehavior results in the smaller work, is given by

dτij=dt ¼ 2G �etij− �eaij� �

dp=dt ¼ 3K � J1 �eð Þð1Þ

where and ε̇tij,τij are the i–jth component of the total strain rate anddeviatoric stress tensors, respectively, ε̇aij indicates the anelastic(viscous or plastic) strain rate tensor, p is the pressure, G and K arethe shear and bulk modules.

Frictional behavior follows a standard Drucker–Prager yieldrelationship:

f σð Þ ¼ J2 σð Þ−α J1 σð Þ þ ctan

/0

h ib0

α ¼ 6 sin/0

3− sin/0

ð2Þ

where J1(σ) and J2(σ) are the 1st and 2nd invariants of the stresstensor, ϕ0 is the internal angle of friction and c is cohesion, respec-tively set to 30° and 1 MPa in the present numerical experiments.

For deviatoric stresses below the power law breakdown limit(ca. 750 MPa for olivine, Goetze and Evans, 1979), thermally activatedcreep obeys a power law relationship:

�eij ¼ 3=2ð ÞγJ2 τð Þ n−1ð Þτij exp −Q=RTð Þ ð3Þ

where έij,τij are the deviatoric i–jth component of the strain rate andstress tensors, respectively, J2(τ) is the 2nd invariant of the deviatoric

Fig. 1. Geometry, boundary conditions, and initial temperature field of the reference model am01w.

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stress tensor, n, Q, and γ are the experimentally-derived power lawexponent, activation energy, and fluidity in Pa− n s− 1, R is the gasconstant, and T is temperature in Kelvin degrees. Above the power lawbreakdown stress, viscous creep either obeys the power law (Eq. (3))or an exponential relationship depending on which produces lowerstresses. The exponential flow law, which describes deformation bylow-temperature plasticity mechanisms (Peierls creep), is given by:

�eij ¼ 3e0=2J2 τð Þð Þ exp −Q=RTð Þ 1−J2 τð Þ=τ0ð Þ2� �

τij ð4Þ

where ε̇0 and τ0 are experimentally determined reference strain rateand stress (Goetze and Evans, 1979).

The model geometry (Fig. 1) simulates a 1000 km long and 80 kmdeep lithosphere consisting of: (a) a 40 km thick two-layer crust withheat productions of 2 and 0.25 μW/m3 in the upper and lower crust,respectively and (b) a 40 km mantle layer containing a 200 km-wideand 20 km-thick sectionwith a heat production ranging between 0.05and 0.25 μW/m3 (referred in the following as the anomalous, HPE-enriched or radioactive mantle) embedded in mantle section with anull heat production (referred to as the normal mantle).

In all simulations, a symmetry condition is imposed at the center ofthe anomalous mantle domain and calculations are carried out onlyfor the left half of the model. Deformation is driven by imposing aconstant convergent velocity of 1e− 9 m/s (∼ 3 cm/y) to the leftboundary of the model. Gravity-induced body forces are taken intoaccount; the upper surface is free and a hydrostatic pressure thatequilibrates the lithostatic pressure produced by the overlying rockcolumn is applied to the lower boundary. Models are discretized using∼45,000 triangular elements, leading to a resolution on the order ofthe km.

A fixed temperature of 300 K and a constant heat flux of 15mW/m2

are imposed at the upper and lower boundaries, respectively. Heat fluxis null on both lateral boundaries. The initial thermal structure (Fig. 1)corresponds to the stationary temperature field for these boundaryconditions and the imposed heat production distribution. With theseconditions, themodeled 80 km-thick lithospheric section correspondsroughly to the mechanical lithosphere, where long-term storage ofheat-producing elements is thought to be possible (McKenzie, 1989).

Table 1Physical parameters used in the simulations

Wet quartzite Quartz-diorite Felsic

Density (kg m−3) 2.65e3 2.65e3 2.7e3Young modulus (GPa) 70 70 70Poisson ratio 0.25 0.25 0.25Fluidity γ (Pa−n s− 1) 1.63e−26 1.14e−28 4.e−25Activation energy Q (kJ mol− 1 K− 1) 135 123 260Stress exponent n 3.1 3.2 3.4Power law breakdown stress – – 1000Reference stress σ0 (GPa) – – 8.5Reference strain rate e0 (s−1) – – 5.7e11Thermal conductivity k (W m−1) 2.7 2.7 2.7Heat production A (µW m−3) 2 2 0.25

⁎Heat production in the HPE-enriched mantle domain varies between 0.05 and 0.25 µW m

Both crust and mantle lithosphere are assumed to have an elasto-viscoplastic rheologywith constant elastic strength, pressure-dependentyield, and thermally activated viscous creepwith non-linear viscous flowlaws. The choice of a 40 km-thick crust is based on seismic dataand geothermobarometry of lower crustal xenoliths (Christensen andMooney, 1995; Rudnick and Fountain, 1995; Artemieva, 2002) that showthat post-Archean continental crust has a mean thickness closer to thisvalue rather than the 30 km or 35 km frequently used in numericalmodels. In neoproterozoic domains of Central and SE Brazil, for instance,inversion of receiver functions showed crustal thicknesses varying from37 to 42 km (Assumpção et al., 2004; França and Assumpção, 2004).Combination of seismological, petrophysical and crustal xenoliths studiesindicates that the structure of most crustal sections is better modeled asconsisting of three layers: an upper felsic crust, an intermediate middlecrust, and a mafic lower crust (Rudnick and Fountain, 1995). For sim-plicity, only two crustal layers were used in the present simulations. Therheology of the ductile upper crust is based on either awet quartzite or aquartz-diorite flow law (Paterson and Luan, 1990; Ranalli, 1997). For thelower crust, experiments were carried out with a variety of compositionsranging from “strong” (dry diabase or basic granulite; Mackwell et al.,1998, Wilks and Carter, 1990 ) to “soft” (wet diabase or felsic granulite;Caristan, 1982, Wilks and Carter, 1990) rheologies. For the mantle, wetand dry olivine rheologies were used (Chopra and Paterson, 1981, 1984).Material parameters used in all simulations are summarized in Table 1.

Although simple, these 2D thermo-mechanical models allowevaluating the potential of a local enrichment in HPE within thelithospheric mantle to induce intraplate strain localization. Experi-ments were carried out to verify the effects of variations in: (1) theheat production in the radioactive mantle, (2) the location of the heat-producing domain, (3) the rheological structure of the lithosphere,and (4) the background thermal state of the lithosphere.

The Lagrangian formulation of the finite-element code ADELI 2-Ddoes not allow for very large deformations without remeshing. How-ever, after the first thousands of years, there is little variation in bothdeviatoric stress and strain rate fields as a function of time indicatingthat the systemapproached steady state. All experimentswere thereforerun for 1 Ma, corresponding to 6% shortening, i.e., a finite strain of 0.06if the deformation was homogeneous within the model.

granulite Mafic granulite Dry diabase Dry dunite Wet dunite

2.75e3 2.75e3 3.3e3 3.3e370 70 160 1600.25 0.25 0.28 0.288.83e−22 1.2e−26 2.42e−16 3.98e−25445 485 545 4984.2 4.7 3.5 4.51000 1000 500 5008.5 8.5 8.5 8.55.7e11 5.7e11 5.7e11 5.7e112.7 2.7 3.5 3.50.25 0.25 0⁎ 0⁎

−3 in the different models.

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4. Results

4.1. Reference model

The reference model — am01w — is composed by a ‘weak’ uppercrust and a ‘strong’ lower crust, which behave mechanically as a wetquartzite and a dry diabase, respectively, and a weak, hydrated upper

mantle, which has a 20 km thick heat-producing domain with a heatproduction of 0.1 μW/m3 between 60 km and 80 km depth (Figs. 1 and3, Table 1). This heat production in the mantle lithosphere results in asmall enhancement of the thermal gradient, characterized by atemperature increase within and above the anomalous mantledomain, which varies almost linearly from 30 K at 80 km to 0 K atthe surface (Fig. 2a,b). The surface heat flow varies therefore across themodel from 60 mW/m2 in the ‘normal’ lithosphere to 62 mW/m2 atopthe metasomatized mantle.

Analysis of the finite strain distribution after 1 Ma of convergence(Figs. 2c, 3a) highlights that the strain distribution is stronglycontrolled by the enrichment in HPE in the mantle. In contrast, thelater has a minor effect on the deviatoric stress field (Fig. 3c), which isessentially controlled by the compositional structure and verticaltemperature gradient. Strain in the anomalous mantle and in the crustabove it is in average two times higher than in the normal mantle,decreasing gradually to attain a roughly constant value ca. 50 km awayfrom the boundary of themetasomatizedmantle domain (Fig. 2c). Thisresults in crustal and mantle thickening with a maximum atop thecenter of the anomalous mantle domain (Fig. 3a). Strain localization isslightly weaker in the upper crust (Fig. 2c). Although brittle behaviordoes not depend on temperature and, by consequence, is notinfluenced by the lateral variation in heat production, the interactionbetween the elastoplastic deformation and ductile flow resultsnevertheless in localization of brittle deformation. Effective brittleyielding occurs within 100 km of the anomalous mantle and thehighest brittle strain is observed in a few ‘faults’ at the external limits ofthe strain localization domain (Fig. 3a). These ‘faults’ continue intosubvertical high strain domains in the ductile upper crust that end in adetachment horizon at the base of upper crust. This detachmenthorizon aswell as the strain localization level located at the base of thelower crust results from the strong gradient in strength at the upper—lower crust and lower crust — mantle boundaries (Figs. 2d, 3c).

4.2. Effect of the heat production intensity

Increasing the heat production in the metasomatized mantle to0.2 μW/m3 does not change significantly the strain and deviatoricstress patterns (Figs. 2d, 3b,d), but the higher heat production results inmore effective strain localization. We tested the effect on varying theheat production between 0.05 and 0.25 μW/m3. These values result insurface heat flows above the metasomatized domain ranging from 61to 65 mW/m2, i.e. in a 1 to 5 mW/m2 anomaly relative to the ‘normal’mantle surface heat flow of 60 mW/m2. The temperature increase atthe base of the mantle is directly proportional to the heat production,varying from ∼15 to ∼70 K (Fig. 2a,b). Strain localization dependsexponentially on temperature and hence on heat production; averagestrains within and above the metasomatized domain range from 1.2to 5 times higher than in those in the normal mantle (Fig. 4). Thisexponential dependence on temperature is coherent with a mechan-ical behavior controlled by thermally activated creep in the mantlelithosphere and crust (Eqs. (3) and (4)). Based on this behavior, wemaypredict that a mantle heat production of 0.4 μW/m3 as estimated fromU, Th and K contents of mantle xenoliths from Phanerozoic volcanicregions in Australia, Europe and Alaska (O’Reilly and Griffin, 2000) will

Fig. 2. Results formodelswithasimilar compositional structure:awetquartziteuppercrust, adry diabase lower crust, and wet mantle, but an anomalous mantle between 60 and 80 kmdepthwithaheat production rangingbetween0.05 and0.25 μW/m3 (black symbols / lines) aswell as amodel with an anomalousmantlewith a heat production of 0.1 μW/m3 between 40and 60 km depth (red or gray symbols/lines). (a) Initial geotherm at 100 km of the plateboundary (normal lithosphere) andat 450kmof theplateboundary (anomalous lithosphere).(b) Initial temperature anomalyestimated as the difference between the temperature profilesat 100 and 450 km. (c) Strain localization characterized by the ratio between the 2n invariantof the strain tensor along a profile at 450 km of the plate boundary relative to a profile at150 kmof the plate boundary. (d) 2nd invariant of the deviatoric stress tensor as a function ofdepth for a profile at 450 km of the plate boundary.

Fig. 3. 2nd invariant of the finite strain (a,b) and deviatoric stress tensors (c,d) for models with the reference geometry (Fig. 1), but heat productions in the HPE-enriched mantle of0.1 μW/m3 (a,c) and 0.2 μW/m3 (b,d) after 1 Ma of convergence.

396 S.P. Neves et al. / Earth and Planetary Science Letters 274 (2008) 392–400

result in strains more then 1 order of magnitude higher than those inthe ‘normal’ surrounding mantle. We may also conclude that heatproductions b 0.05 μW/m3 will result in negligible strain localization.Finally, analysis of the strain localization profiles for different heatproductions (Fig. 2c) highlights that strain localization at the base ofthe upper and lower crust due to the steep rheological contrast at theinterface with the lower crust and mantle, respectively, is moreeffective in the hotter and, hence, weaker models.

4.3. Effect of the location at depth of the heat-producing mantle layer

Moving the location of the anomalous mantle layer to shallowerlevels, just below the Moho, results in weaker strain localization.Strain in the mantle section containing the HPE-enriched domain andin the crust above it is in average 1.5 times higher than in the normalmantle (Fig. 2c). The weaker strain localization results from the factthat for a shallow HPE-enriched layer, the temperature enhancementdoes not increases regularly with depth up to the base of the plate asin the reference model, but attains a constant value below 50 km(Fig. 2b). The temperatures at the base of the lithosphere are thereforesignificantly lower than the ones in the reference model (Fig. 2a),leading to a weaker variation in lithospheric strength (Fig. 2d).

Models with a shallow HPE-enriched layer also show an exponen-tial dependence of strain localization relative to the heat productionintensity, but with a weaker slope (Fig. 4). Thus for a shallowermetasomatized layer, a higher HPE enrichment is necessary forproducing the same strain localization. Indeed, a model (not shown)with an increase of the lower crust heat production by 0.1 μW/m3 in a

Fig. 4. Strain localization as a function of heat production in the metasomatized mantlefor models with heat productions varying from 0.05 to 0.25 μW/m3 at 40–60 km depth(in red) and 60–80 km depth (in black).

domain 20 km thick and 200 kmwide (the same size as the anomalousmantle domain in the previous models) does not show strain locali-zation in the HPE-enriched domain.

4.4. Role of the lithosphere compositional structure

The effect of the lithosphere compositional and, consequently,rheological structure was tested by varying the viscous flow laws forthe three layers (Fig. 5). The upper crust was modeled using either aweak ‘wet quartzite’ or a strong ‘quartz-diorite’ rheology. The com-position of the lower crust is more controversial. Although worldaverages point to a basic, plagioclase-rich composition (Rudnick andFountain, 1995), in some regions, the continental crust appears toconsist mostly of felsic material (Villaseca et al., 1999; van den Berget al., 2005). We choose therefore to test three different rheologies forthis layer: a strong dry diabase, an intermediate strength maficgranulite, and a rather weak felsic granulite flow law. Finally, a pos-sible variation in strength of the lithospheric mantle due to incor-poration of water in the olivine structure was considered by using wetand dry dunite flow laws.

Comparison of models with the reference geometry, i.e. a 20 kmthick metasomatized mantle with a heat production of 0.1 μW/m3

between 60 and 80 km depth, but different structures highlights thatstrain localization shows an inverse dependence on the plate strengthestimated by integrating the deviatoric stress over the plate thickness(Figs. 5 and 6). Weaker plates show a stronger strain localization.The mantle lithosphere rheology plays a major role on the straindistribution: ‘wet mantle’ models (am01a and am01w; Figs. 3 and 6a)display a significantly stronger strain localization than most modelswith a ‘dry mantle’ rheology (am01c, am01d, and am01f; Fig. 6b,c,d).However, model am01g (Fig. 6e), which has a dry mantle, but a weakfelsic lower crust and hence relatively low lithosphere strength showsa strain localization similar to the one of the wet mantle models. Incontrast, comparison between models characterized by a similarlithospheric strength shows that strain localization in the lower crustis directly proportional to the strength of this layer (Figs. 5 and 6a). Aweaker rheology also favors strain localization at the base of the lowercrust. Finally, models with a strong lithosphere show weaker strainlocalization in the upper crust (Fig. 6). A stronger rheology in theupper crust leads to a deeper brittle–ductile transition and to a moreregular distribution of the yielding. One should note however that theminimum ‘fault’ spacing in the present models is still dependent onthe mesh resolution, which is ∼1 km.

5. Discussion

The numerical models suggest that the presence of a HPE-enrichedlayer in the lithospheric mantle may explain the localization of the

Fig. 5. Finite strain (J2, 2nd invariant, left) and strength profiles (right) for models with the same geometry and initial thermal structure as the reference one but differentcompositional structures. Material parameters for all compositions are shown in Table 1.

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deformation in intraplate settings, far from any plate boundary. Strainlocalization develops in the mantle lithosphere, in the ductile crust,but also in the shallow brittle crust, where it results from perturba-tions in the stress field due to thermally induced strain localization in

Fig. 6. (a) Strain localization as a function of depth, estimated as the ratio between the2nd invariant of the strain tensor along a profile at 450 km of the plate boundaryrelative to a profile at 150 km of the plate boundary, and (b) average strain localizationas a function of the lithosphere strength, estimated by integrating the deviatoricstresses over the lithosphere thickness, for the models displayed in Fig. 5.

the ductile levels. Strain localization at the plate scale depends expo-nentially on the local increase in the geotherm, which is controlled bythe interplay between the local enrichment in HPE and the volume ofthe mantle that has been metasomatized. Heat productions between60 and 80 km depth ranging from 0.05 to 0.4 μW/m3, which arecoherent with geochemical data in mantle peridotites and mantle-derived magmas, result, for instance, in strains within and above themetasomatized mantle 1.2 times to more than 1 order of magnitudehigher than in the surrounding lithosphere. Strain localization alsodepends on the location of the heat-producing domain. Shallow heat-producing domains result in a smaller increase in temperature in thedeep mantle lithosphere and hence in weaker localization. The platerheological structure also affects the response to local increase in heatproduction; weaker plates display stronger localization. However,for plates with similar strength, strain localization in the crust isproportional to the crustal strength.

An effect that has not been considered in the present models, butthat may significantly enhance strain localization, is further thermalweakening by shear heating (Kaus and Podladchikov, 2006; Rege-nauer-Lieb et al., 2006; Hartz and Podladchikov, 2008). Indeed,depending on composition, high stresses (≥1 GPa) are attained inthe lower crust and the uppermost mantle (Figs. 2, 3, and 5).Strain rates within and above the metasomatized domain varybetween 4.5e−15 s− 1 and 1.e−14 s−1. Such an association of highstresses and strain rates may result, if viscous heating is the dominantenergy dissipation mechanism, in heat production by viscousdissipation of up to 10 µW/m3 (Hartz and Podladchikov, 2008), i.e.,up to one order of magnitude higher than the one due tometasomatism, and the resulting increase in temperature will leadto further strain localization. However, it should be noted that therheology of mantle and lower crustal rocks under low-temperatureconditions is poorly constrained (data used in the present models arederived from a single dataset by Goetze and Evans, 1979) and thatdeviatoric stresses and hence shear heating may be overestimated.

5.1. Implications for intracontinental orogenesis

The occurrence of deformation and metamorphism in ancientcontinental domains far from any convergent plate boundary has

398 S.P. Neves et al. / Earth and Planetary Science Letters 274 (2008) 392–400

puzzled geologists for a long time. The best-known examples areArchean granite-greenstone terrains in the Pilbara (Australia) andDharwar (India) cratons (e.g., Bouhallier et al., 1995; Van Kranendonket al., 2004), and intraplate orogenic belts with ages ranging from thePaleoproterozoic to the Paleozoic in the Australian continent(Sandiford and Hand, 1998; McLaren et al., 1999; Giles et al., 2002).Other clear examples of intracontinental deformation includes thePaleoproterozoic Kapuskazing Structural Zone in Canada (Perry et al.,2006), an unnamed contractional Mesoproterozoic event in SW USAduring which ca. 1.4-Ga A-type plutons were emplaced (Nyman et al.,1994; Nyman and Karlstrom, 1997; Ferguson et al., 2004), and thePaleozoic Ancestral Rocky Mountains of western North America(Dickerson, 2003). An intracontinental setting has also been proposedfor the Taltson-Thelon orogenic belt of northwestern Canada (De et al.,2000), for the Neoproterozoic Kaoko and Damara belts of south-western Africa (Dürr and Dingeldey, 1996; Jung et al., 1998), and forthe Pan-African/Brasiliano belts of Northeast Brazil, Cameroon andNigeria (Vauchez et al. 1995; Neves, 2003). Some of these intraplateorogens are still today characterized by high surface heat flow (Jaupartand Mareschal, 1999 and references therein), probably reflecting heatproduction at depth.

Deformation and metamorphism in intraplate settings have beenascribed to a number of causes. In classical Archean dome-and-keelterranes, the most popular hypothesis is partial convective overturndriven by density inversion (Collins et al., 1998). However, thestructural style of Proterozoic and younger belts clearly requiresdeformation to be driven by horizontal forces. Moreover, to localizedeformation in a specific domain of a plate, either rheologicalheterogeneities or mechanical anisotropy is required (e.g., Tommasiand Vauchez,1997, 2001). In the absence of large-scale plate boundaryforces, deformation of continental interiors may result from develop-ment and growth of gravitational (Rayleigh–Taylor) instabilities at thebase of the lithosphere. Numerical experiments (Neil and Houseman,1999; Pysklywec and Beaumont, 2004; Pysklywec and Cruden, 2004)have shown that the continental crust can be thickened and upliftedover a lithospheric downwelling. However, development of aconvective instability requires: (i) higher density of the lithosphererelative to the asthenosphere, (ii) existence of an initial perturbationin the topography of the base of the lithosphere, (iii) fast growinginstabilities (their rate of growthmust exceed the dissipative effects ofthermal diffusion), and (iv) a weak ductile crust. In presence ofhorizontal stresses transmitted through the lithosphere from remoteplate boundaries, intracontinental deformation has been related tolithospheric weakening resulting from the burial of high heat-producing granitic basement beneath a thick sedimentary cover(McLaren et al., 1999; Sandiford and McLaren, 2002). However, toproduce strain localization in an intracontinental setting, exceedinglyhigh heat production values in the buried basement are required. Inaddition, a refractory lower crust is essential so as to avoid wholesalecrustal melting (McLaren et al., 2005). Furthermore, this model ofintraplate orogenesis has been challenged in the type areas in centralAustralia where it was first proposed based on evidence that thebasement was emergent— and not covered during the orogenic event(Camacho et al., 2002).

In contrast with these previous studies, the results described herecan be of more general applicability. Radioactive heating in themantlecan be more efficient in weakening the lithosphere than high heatproduction restricted to the upper crust since this will impart thestrength of the whole lithosphere, nor only of its upper levels. A largeregion characterized by slow seismic shear waves, located at depths of75–100 km, has recently been detected beneath central Australia(Fishwick andReading, inpress). As temperature is the dominant causeof seismic velocity anomalies, the lowvelocities probably relate to highheat production in this zone. This finding is on line with the mainimplication of our numerical simulations: that segments of uppermantle enriched inHPE are very efficient in localizing strain in both the

crust and the mantle during contractional deformation of the con-tinental lithosphere. Strain localization in the crust atop the enrichedmantle is observed in all simulations and it is enhanced inmodels withweak lithosphere or with strong lower crust rheology (Fig. 4). Inmodelswith aweak upperor lower crust, strain localization tends to beless effective and we observe the development of horizontal highstrain zones at the base of the weak layers transferring strain to moreexternal domains (Figs. 3 and 4). Geologically, these could correspondto regions of localized channel flow or detachment zones.

5.2. Application to modern intraplate orogens

Delamination or convective removal of a root of cold mantlematerial 8–15Ma ago has been suggested to explain the rapid increasein mean elevation and the potassic magmatism in the Tibetan plateau(e.g., Platt and England, 1994). However, shoshonitic volcanics displayages from the Eocene onwards and their geochemical and isotopiccharacteristics indicate derivation by small-degree partial melting ofmetasomatically-enriched mantle lithosphere (Turner et al., 1996;Miller et al., 1999; Nomade et al., 2004; Chung et al., 2005). Nd modelages of the shoshonites indicate that the main event of metasomaticenrichment occurred 1.3–1.2 Ga ago (Turner et al., 1996), and thuswas not associated with subduction episodes during the Paleozoic/Mesozoic amalgamation of Asia. Recent seismic results show thatTibet is underlain by thick lithosphere (260 km), implying that large-scale lithospheric removal has not occurred (McKenzie and Priestley,2008). These petrological and geophysical observations might beexplained if substantial portions of the Tibetan plateauwere underlainby HPE-enriched lithospheric mantle at the onset of the India–Asiacollision.

Metasomatized mantle sections produced by addition of incom-patible elements (but without significant alteration on major-elementchemistry) should not only be weaker but also more buoyant thanthe surrounding mantle, because of their higher temperature.During contraction, it is therefore unlikely that lithospheric thicken-ing can lead to delamination. Thus for equilibrium to be reached,shortening has to be compensated by exhumation, progressivelyreturning the anomalous mantle to shallow depths. This in turn canlead to decompression melting, surface uplift, and increased tem-peratures in the lower crust, leading to crustal melting, all of whichare geological features commonly ascribed to delamination or con-vective removal of the lithosphere. A possible placewhere this processmight be currently occurring is the Tien Shan, where a low seismicvelocity zone was detected at ca. 90 km underneath the highestelevations in the region (Vinnik et al., 2004; Kumar et al., 2005). Thetop of this low velocity zone was interpreted as marking thelithosphere/asthenosphere boundary, but it could equally well beinferred to represent the boundary between 'normal' and anoma-lously hot lithospheric mantle.

5.3. Other geological implications

Apart from intraplate orogenesis, enrichment of the mantle litho-sphere in HPE can have other geological consequences. Geologicalevidence suggest existence of large continents/supercontinents in thePaleo- and Mesoproterozoic periods, which formed in consequence ofcollisional orogenesis following convergence of major crustal blocks(Rogers, 1996; Rogers and Santosh, 2002, Zhao et al., 2004). Asdiscussed in Section 2, the Proterozoic eon was a period of the Earthhistory during which shallow subduction was probably more wide-spread than in the previous and latter eons. As a result, HPE-enrichedmantle may have formed within the lithospheric mantle of Proter-ozoic continental masses. It is perhaps significant that the 2.0–1.0 Gatimespan is characterized by abundant anorogenic magmatism andthat most anorogenic plutons postdate the last period of orogenicactivity by several tens of million of years to a few hundred million of

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years (Windley, 1993). Because radioactive decay is slow, these timescales correspond to those required for the temperature in theanomalous mantle to attain its maximum value before declining dueto the dissipative effect of thermal diffusion. In highly metasomatizedregions of the mantle, the temperature increase may have beenenough to surpass that of water-undersaturated peridotite solidus.Indeed, temperatures up to 1080°C may be attained at a depth of80 km for a heat production of 0.25 μW/m3 in the mantle lithosphere(Neves and Mariano, 2004). This is comparable to the experimentalsolidus at 1.5 GPa of phlogopite–pargasite-bearing peridotite, whichranges from 1025 °C to 1125 °C (Conceição and Green, 2004).Therefore, partial melting of mantle rocks and, in turn, of the lowercrust can occur in absence of external tectonic forcing, in accord withthe general lack of evidence, except in a few cases (e.g., Hutton et al.,1990; Nyman et al., 1994), for deformation, either contractional orextensional, coeval with emplacement of anorogenic plutons.

If heat production in the anomalous mantle is not enough topromote partial melting when the climax temperature is reached,melting of the enriched mantle may still happen if extension occursconcomitantly or shortly afterwards. Indeed, experimental results formetasomatized peridotite (Conceição and Green, 2004) show a 100 °Creduction in the solidus temperature between 1.5 and 1.0 GPa,implying that decompression melting of lithospheric mantle, incontrast with that of sublithospheric mantle (McKenzie and Bickle,1988), may start at moderate degrees of extension. Furthermore, thecomposition of the melts is K-rich and thus similar to those of potassiclavas erupted in some Cenozoic extensional regions, such as the Basin-and-Range (e.g., Lopez and Cameron, 1997).

6. Conclusion

Deformation and metamorphism in continents often occur far fromany plate boundary. The results of the present 2D numerical experi-ments suggest that intraplate orogeny may be a consequence of a localincrease in the geotherm due to an enrichment of HPE in thesubcontinental lithospheric mantle. In a compressional setting, aradioactive mantle section 200 km-wide and 20 km-thick with heatproduction of 0.1 μW/m3 between 60 and 80 km depth results in strainsin the anomalous mantle and in the overlying crust twice as high thanthose in the normal lithosphere. Strain localization depends exponen-tially on the increase in the geotherm, it is thus enhanced for higherheat productions, thicker and deeper radioactive layers. Metasomaticintroduction of HPE due to plume activity or subduction produces,therefore, weakness zones within the lithospheric mantle that canpersist for long geological periods and be deformed by stressestransmitted through the lithosphere from remote plate boundaries.

Acknowledgements

This work was funded by the Brazilian agency Conselho Nacionalde Desenvolvimento Científico e Tecnológico (CNPq). Jean Chéry andRiad Hassani are thanked for making the code ADELI freely available tothe geosciences community (http://www.isteem.univ-montp2.fr/PERSO/chery/Adeli_web/). Jean-Claude Mareschal, Taras Gerya, andMaya Kopilova are thanked for the constructive reviews that greatlyhelped to improve the manuscript.

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