Plume head lithosphere interactions near intra-continental...

(2007) 15–38www.elsevier.com/locate/tecto

Tectonophysics 434

Plume head–lithosphere interactions near intra-continentalplate boundaries

Evgene Burov a,⁎, L. Guillou-Frottier b, E. d'Acremont a,L. Le Pourhiet c, S. Cloetingh d

a Laboratoire de Tectonique, Université Pierre and Marie Curie, Paris, Franceb Service des Ressources Minérales, Bureau de Recherches Géologiques et Minières, Orléans, France

c Geological and Planetary Sciences, California Institute of Technology, Pasadena, USAd Department of Earth Sciences, Vrije University, Amsterdam, Netherlands

Received 3 April 2006; received in revised form 5 January 2007; accepted 5 January 2007Available online 24 January 2007

Abstract

Plume–lithosphere interactions (PLI) have important consequences both for tectonic and mineralogical evolution of the lithosphere:for example, Archean metallogenic crises at the boundaries of the West African and Australian cratons coincide with postulated plumeevents. In continents, PLI are often located near boundaries between younger plates (e.g., orogenic) and older stable plates (e.g., cratons),which represent important geometrical, thermal and rheological barriers that interact with the emplacement of the plume head (e.g.,Archean West Africa, East Africa, Pannonian–Carpathian system). The observable PLI signatures are conditioned by plume dynamicsbut also by lithosphere rheology and structure. We address the latter problem by considering a free-surface numerical model of PLI withtwo stratified elasto-viscous–plastic (EVP) lithospheric plates, one of which is older and thicker than another. The results show that:(1) plume head flattening is asymmetric, it is blocked from one side by the cold vertical boundary of the older plate, which leads to themechanical decoupling of the crust from the mantle lithosphere, and to localized faulting at the cratonic margin; (2) the return flow fromthe plume head results in sub-vertical down-thrusting (delamination) of the lithosphere at the margin, producing sharp vertical coldboundary down to the 400 km depth; (3) plume head flattening and migration towards the younger plate results in concurrent surfaceextension above the centre of the plume and in compression (pushing), down-thrusting and magmatic events at the cratonic margin(down-thrusting is also produced at the opposite border of the younger plate); these processes may result in continental growth at the“craton side”; (4) topographic signatures of PLI show basin-scale uplifts and subsidences preferentially located at cratonic margins.Negative Rayleigh–Taylor instabilities in the lithosphere above the plume head provide a mechanism for crustal delamination. Inferredconsequences of PLI near intra-continental plate boundaries, such as faulting at cratonic edges and enhanced magmatic activity, couldexplain plume-related metallogenic crises, as suggested for West Africa and Australia.© 2007 Elsevier B.V. All rights reserved.

Keywords: Plume–lithosphere interaction; Cratons; Metallogenic crises; Rheology

⁎ Corresponding author.E-mail addresses: [email protected] (E. Burov),

[email protected] (L. Guillou-Frottier), [email protected](L. Le Pourhiet), [email protected] (S. Cloetingh).

0040-1951/$ - see front matter © 2007 Elsevier B.V. All rights reserved.doi:10.1016/j.tecto.2007.01.002

1. Introduction

Mantle plumes are regarded as focused mantleupwellings that originate from deep mantle sources (e.g.,Campbell et al., 1989; Campbell and Griffiths, 1990;

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http://dx.doi.org/10.1016/j.tecto.2007.01.002

16 E. Burov et al. / Tectonophysics 434 (2007) 15–38

Olson, 1990). If the link with the source region iscontinuous, such upwellings are considered as mantleplumes senso stricto, otherwise they are referred asmantle “diapirs” senso lato. Yet, there is no differencein physical mechanisms driving “plumes” and “dia-pirs”, and in nature it is difficult to discriminatebetween, continuous, multiple discontinuous or inter-rupted continuous mantle upwellings. For this reason,we call here “plumes” all focused plate-scale mantleupwellings. This is justified by the fact that, in thisstudy, we focus only on the initial response of thelithosphere to plume head emplacement.

The possible consequences of plume–lithosphereinteractions and even the existence of mantle plumes aretopic of vivid discussions (e.g., Anderson, 1982:Farnetani et al., 1996; Sleep, 1997; Anderson, 1998;Ebinger and Sleep, 1998; Kukkonen and Peltonen,1999; Sheth, 1999; Tackley, 2000; Sleep et al., 2002;Foulger, 2002; Sleep, 2003a,b, 2005; Nyblade andSleep, 2003; Ingle and Coffin, 2004; King, 2005). Thesediscussions are not only fuelled by ambiguous obser-vational evidence for the existence of mantle plumes butalso – and at a large extent – by the discrepanciesbetween the geological and geophysical observationsand model-based predictions for consequences ofplume–lithosphere interactions such as the expectedlarge-scale (N1000 km) dynamic topography, thermal,tomographic and gravity anomalies. Indeed, in theirclassical definition, mantle plumes should be associatedwith a large-scale domal uplift at surface and volcanicactivity. However, these features are not always orunambiguously observed in the areas that present othersignatures of plume activity, including such undisputa-ble hotspots as Hawaii (e.g., Davies, 1988, 1992; Wesseland Keating, 1994). For example, the dynamic topog-raphy detected in the “Afar Triangle” — Red Sea zonediffers from a large-scale domal uplift, although there islittle doubt in the presence of large-scale mantleanomaly below this area that bears tomographic (e.g.Pik et al., 2006) but also some geochemical signatures ofa mantle plume (e.g., Moreira et al., 1996).

Despite of the difficulties of their detection, mantleplumes exist and have considerable impact on tectonicevolution of plates, because they can rapidly bring deepmaterial to the surface, such as Fe+Mg rich magmas(e.g. Ebinger and Sleep, 1998; Isley and Abbott, 1999;Sleep et al., 2002; Sleep, 2003a,b; Nyblade and Sleep,2003). The plumes may also trigger heating and meltingof the lower crust and mantle lithosphere thus leading tocrustal or mantle delamination (Smithies et al., 2003;Saleeby et al., 2003) and continental growth (McCul-logh and Benett, 1994; Condie, 2002). A number of

advanced plume–lithosphere interaction models haveconsidered various aspects of these interactions assum-ing self-consistent starting conditions and treatingcomplicated problems such as longevity of xenolithsat the base of the lithosphere and channeling of theplume material (e.g., Ebinger and Sleep, 1998; Sleepet al., 2002; Sleep, 2003a,b; Nyblade and Sleep, 2003).Recent tectonically-oriented models incorporate rheo-logically stratified lithosphere with a free surface(d'Acremont et al., 2003; Burov and Guillou-Frottier,2005; Guillou-Frottier et al., in press). The models showthat plumes can create particular small-scale topograph-ic signatures such as the succession of uplift and sub-sidence phases. Finally, plumes may enable lithosphericfaulting (d'Acremont et al., 2003; Burov and Guillou-Frottier, 2005) as well as crustal growth by accretion ofjuvenile magmas (Boher et al., 1992).

When a tectonically realistic formulation for thelithosphere is accounted for in models of plume–lithosphere interactions, the predicted surface deforma-tion shows new intuitively unexpected features, anddeeper lithosphere may become gravitationally unstable(Burov and Guillou-Frottier, 2005). Since variations inlithospheric thickness and other heterogeneities arecommon in intra-continental settings (e.g., Artemievaand Mooney, 2002), a number of PLI-related geologicalissues cannot be tackled without considering lateralvariations of lithosphere thickness. For example, lateralvariations in thermal structure of the continentallithosphere, as underlined by Perry and Jaupart (2004),Mareschal and Jaupart (2004), should have strongimplications for plume–lithosphere interactions nearcratonic boundaries. To account for these phenomenawe have conducted a series of new numerical models ofPLI, in which we have included lateral heterogeneitiesassociated with intra-continental plate boundaries. Themodels account for tectonically realistic lithosphericstructure, free surface deformation and elastic–viscous–plastic (EVP) rheology of the lithosphere. The main goalof the experiments is to study plume–lithosphere in-teractions in tectonically realistic settings such as thosethat include intra-lithospheric blocks and cratonicborders. This study is also driven by the idea thatsurface signatures resulting from plume–continentallithosphere interaction near cratonic boundaries canexplain a number of phenomena such as the concentra-tion of magmatic activity at these boundaries orsimultaneous extension–compression events. Predictedsurface topography features are then used to interpretuplifts and subsidences in terms of mechanical coupling/uncoupling between different rheological layers withinthe lithosphere. In addition, heat and mass transfer at

17E. Burov et al. / Tectonophysics 434 (2007) 15–38

lithospheric depths, following the plume impact phase,is detailed and discussed in the framework of a numberof well-known geological concepts such as continentalgrowth or crustal delamination.

2. Plume–lithosphere interactions with a realisticformulation for the lithosphere

2.1. Plume physics, lithosphere layering and surfacedeformation

The physical side of the mantle plume conceptincludes upwelling of low density, high temperature andlow viscosity mantle material separated from large-scaleconvective motions. When active, this upwelling ischaracterized by large geoid/topography ratio and canform a «superplume», i.e. a persisting diapir of a verylarge scale coming from D″ boundary (e.g. Condie et al.,2000; Romanowicz and Gung, 2002), or smaller-scalediapirs originating from different depths (Courtillotet al., 2003; Montelli et al., 2004; Ritter, 2005). Whenconvection is not chaotic (high Raleigh and Reynoldsnumbers), the plume tail is connected to its “mushroom”head, but when more appropriate rheologies areconsidered (e.g., Davies, 1995; Trompert and Hansen,1998) the plume can ascent in a chaotic regime with ahigh ascent rate.

The commonly predicted consequences of PH–CL(plume head–continental lithosphere) interactions aremainly derived from conventional fluid dynamics/viscous models of mantle convection (e.g., Farnetaniet al., 1996; Cserepes et al., 2000; Nyblade and Sleep,2003). The predictive power of these models forsurface features is limited because they are notdesigned to adequately account for the mechanicalresponse of the lithosphere. In typical convectivemodels (e.g., Doin et al., 1997; Solomatov and Moresi,2000), the lithosphere is considered as a thin viscousor quasi-viscous layer with undeformable surface. Thisassumption might be partly acceptable for oceanicplates, but the continental lithosphere cannot beregarded as a thin layer because its thickness (150–300 km) constitutes 50%–90% of the thickness of theconvective part of the upper mantle (350–500 km).The continental lithosphere cannot be also regarded asa single layer, because, in contrast to oceaniclithosphere, the continents are highly stratified. Thisstratification counts 2 to 4 layers: a 30–70 km thickhighly buoyant quartz-rich crust, which may besubdivided into 2–3 sub-layers of different rheology,and a 60–250 km thick, negatively or neutrallybuoyant olivine-rich mantle layer. The crustal sub-

layers include a brittle–elastic–ductile upper, middleand lower crust. The ductile and other mechanicalproperties of these layers may be highly contrasting.For example, the upper crust is mainly brittle while thelower crust is mainly ductile. Due to the presence ofquartz-rich aggregates and increasing temperature withdepth, the lower crust may have a viscosity as low as1020–1021 Pa s below 20 km depth, at temperature of250 °C–400 °C (e.g., Kirby and Kronenberg, 1987). Inthe underlying mantle lithosphere, the ductile creep isactivated at considerably higher temperatures (750–800 °C). This infers the presence of a weak separationlayer (lower or middle crust) between the strong upperor intermediate crust and mantle lithosphere (e.g.,Watts and Burov, 2003; Burov and Watts, 2006). Theweak layer between the strong upper crust andlithospheric mantle allows them to deform relativelyindependently (Fig. 1, Burov and Guillou-Frottier,2005). The ductile crust serves not only as a decou-pling layer but also as a damping layer because thedeformation at its bottom may not be directly trans-mitted to its surface, and vise-versa. Indeed, dependingon the rheology and temperature distribution, thethickness of the weak lower crustal layer may reach15–20 km, while the vertical undulations of its bottomdue to the plume impact (=Moho boundary, top of themantle lithosphere) are not expected to exceed 1.5 km(e.g., Ribe and Christensen, 1994). These smallundulations can be accommodated, or absorbed, withinthe lowermost parts of the ductile crust without congruentsurface deformation. However, they also should producesome indirect impact on the surface. This impact mayresult from lateral and vertical shear flow in the ductilecrust cased by the plume-related uplift of the Mohoboundary. This flow exerts basal shear at the base of theupper (strong) crust, that may result in its thinning and inexpression–compression instabilities. The related surfacesignatures would primarily depend on the mechanicalproperties of the crust and less on the characteristics of theplume itself. Similarly, plume head emplacement belowthe strong mantle lithosphere would not only result in itslarge-scale upward deflection but also in its thinning,basal erosion and extensional–compressional instabil-ities. In all cases, the wavelength of plume-related verticalundulations of the lithospheric layers will be proportionalto 5–10 thicknesses of their resistant cores (10–50 km,Burov and Diament, 1995). This evokes wavelengths ofdeformation on the order of 50–400 km, or up to 1000 kmin case of very strong lithosphere. Most of thesewavelengths are smaller than the wavelengths usuallyattributed to the plume-related dynamic topography(N500–1000 km). The thermo-mechanical consequences

Fig. 1. Comparison of expected plate reaction to plume head impact in case of a single viscous layer (A) and a multilayer brittle–elastic–ductile platemodel (B). In case of a tectonically realistic formulation for the continental lithosphere as in (B), surface topography shows periodic undulations atvarious wavelengths, and multiple zones of compression and extension are present. In this case, a part of lithosphere may be eroded and thrusteddown at the edges of the plume head. Internal Rayleigh–Taylor instabilities (grey arrows) may also develop in the thermally weakened lithosphereabove the plume head leading to drips-off of the lithosphere mantle and lower crust inside the plume head.


of PLI for surface geodynamics and geology thus appearto be largely dependent on the structure and rheologicalproperties of the continental lithosphere (Burov andGuillou-Frottier, 2005).

The idea that in continents, mantle deformation due toplume head impingement produces only moderate orstrongly modulated impact on surface deformation, wasconfirmed by recent numerical studies that take intoaccount the stratified elasto-viscous–plastic lithosphere(Fig. 1, Burov and Guillou-Frottier, 2005;). Asmentionedabove, conventional monolayer models of plume–lithosphere interaction predict only long-wavelengthdeformation (N1000 km) with dynamic topography of1–1.5 km (Ribe and Christensen, 1994). In these models,surface deformation is not computed senso stricto but isestimated from the assumption that vertical undulations oflithosphere–mantle boundary are directly translated to thesurface. This assumption is questionable not only becausethe mantle undulations can be accommodated, or dampedwithin the thick ductile lower crust, but also because thecrust exhibits visco-elastic properties that permit relaxa-tion of plume-related stresses over geologically importantperiods of time (up to 10 Myr).

2.2. Previous results: the need for lateralheterogeneities

After exploring some of the possible consequences ofplume–lithosphere interactions, Burov and Guillou-Frottier (2005) have suggested that the continental

lithosphere serves as a frequency modulator of plumeimpact, as it transforms the large-scale deformation ofmantle–lithosphere boundary (N1000 km wavelength)into a short- and mid-wavelength deformation (50–500 km wavelength) observed at the surface. Thisdeformation occurs at “tectonic scales” (i.e., same scalesas sedimentary basins or orogens) and can be directlyobserved. When the structure and rheology of conti-nental lithosphere is taken into account in plume–lithosphere models (Burov and Guillou-Frottier, 2005),quite complex behavior is predicted: small and middle-scale deformation attenuates large-scale topographicfeatures, and a series of uplifts and subsidences form.Some of the predicted behaviors evoke those usuallyattributed to “passive” tectonic events.

The study by Burov and Guillou-Frottier (2005) hasalso shown that as the plume head locally erodes andthins the lithosphere, the buoyant crust may bemechanically decoupled from the mantle lithosphere,either above the center of the plume head or at itsborders. At the borders, the edge of the laterallyspreading plume head intrudes between the positivelybuoyant crust and heavier mantle. As a result, thelithospheric mantle is separated from the buoyant crustand is down-thrusted forming a “slab” that is drivendown with the return flow in the plume head. Theintrusion of the spreading plume head into thelithospheric mantle results in the compression abovethe borders of the plume, whereas the lithosphere aboveits center is extended.


The two main results briefly described above, i.e. thedevelopment of a complex surface topography and thepossible occurrence of lithospheric instabilities, areexpected to be seriously affected by lateral heterogene-ities of the lithosphere (e.g. presence of a cratonic root).As it was shown by Burov and Guillou-Frottier (2005),wavelengths and amplitudes of surface topography arehighly dependent on thickness of crust and mantle litho-sphere. Moreover, the spreading of the plume head thatfollows the initial impact (see examples in Burov andGuillou-Frottier, 2005) should be seriously hampered oreven stopped by deepening of the lithosphere–astheno-sphere boundary. As already suggested by severalstudies, variations in lithosphere thickness (Artemievaand Mooney, 2002) may help to focus strain and riftlocalization (Pascal et al., 2002; Davis and Slack, 2002),or may enhance the plume-induced decompressionmelting rate (Manglik and Christensen, 1997, 2006).Consequently, a new series of experiments including thepresence of a cratonic root is needed, especially becausenumerous plume-related events would have occurrednear margins of cratonic blocks. In particular, continentalflood basalts (e.g. Deccan, Parana), some of which wererecently called “cratonic flood basalts” (Silver et al.,2006), would have been erupted after abrupt stresschanges in the lithosphere. The Archean greenstone beltsthat contain komatiites have been linked to plumes. Inaddition, some carbonatites and kimberlites may origi-nate from plumes that have stalled beneath thicklithosphere (Ernst and Buchan, 2003). Abundance ofbanded iron formations (e.g., Barley et al., 1997) at theArchean–Proterozoic boundary (e.g. Western Australia,Brazil and Canada) as well as similar locations for othermineralized provinces (e.g. the east African gemstonebelt), suggest that some very particular dynamicalprocesses occur between Archean cratons, their adjacentbelts, and the underlying mantle. Margins of cratonicblocks, associated with transition from younger litho-sphere to older lithosphere, are characterized by lateralvariations in thermo-rheological structure and platethickness, which are reflected in variations of platestrength. In these transitional zones, plume head materialand heat may be channeled below plate boundariesproducing a number of important effects such as local-ized plate weakening and basal erosion (e.g., Monnereauet al., 1993; Davies, 1994: Molnar and Jones, 2004)including rifting, localized thermal and magmaticactivity, attenuation or amplification of the materialcontrasts between the neighboring lithospheric blocks.In the present study, numerical modelling of plumehead–continental lithosphere interaction in the presenceof lithospheric boundaries of the cratonic blocks has

been performed for a variety of thermo-mechanicalstructures of the lithosphere. Models and results arepresented in the next section. We then discuss geologicalimplications of these results.

3. Models and experiments

3.1. Model setup

Our experiments are based on the models andtechniques developed in the previous parametric studiesby Burov and Guillou-Frottier (2005) and d'Acremontet al. (2003). These studies have modelled plume–continental lithosphere interactions for various visco-elastic–plastic lithospheric structures, plume head sizesand Rayleigh numbers. We use the same approach basedon the numerical code Para(o)voz derived from FLACalgorithm (Cundall, 1989; Poliakov et al., 1993). Thismethod allows for handling realistic rheologies and uppersurface boundary conditions. To account for all possiblemodes of lithospheric deformation and for consistentaccount of pressure and of free upper surface, this methodsolves Newtonian equations of motion (second law) inLagrangean framework instead of flow equations(Navier–Stokes) in Eulerian framework commonly usedin plume models. The algorithm explicitly takes intoaccount elastic–brittle–ductile properties of the continen-tal lithosphere, and handles strain localization, whichpermits to model formation of brittle and ductile faults andshear zones. The free upper surface boundary conditionallows for adequate account of deformation of lithospher-ic layers. Due to the explicit nature of the code, thetemperature, gravity and pressure dependent body forcesare computed directly without necessity to introduceadditional assumptions like those of the reference thermalgradient. The last versions of the code were described indetail in (Burov and Poliakov, 2001; Burov et al., 2001,2003; Burov and Guillou-Frottier, 2005). Therefore, wereduce the description of the method to a summaryprovided in Appendices A and B.

3.1.1. Model geometriesIn the first preliminary set of experiments, the

lithosphere is vertically stratified but laterally homoge-neous. In the second, main set of experiments weconsider two bordering lithospheric plates of contrastingage and thickness (see next sections).

Following previous plume models (e.g., Ribe andChristensen, 1994), we skip the initial stages of plumerise, as they are of minor importance for near-surfaceevolution. The initial mantle plume is located at the baseof the model box, and has a diameter of 100–200 km.


We chose a spherical shape for the initial plume becausetranslating viscous bodies take spherical shape inlaminar regime at important distance from the surface(Batchelor, 1967). However, since the model plume isdeformable, the choice of its initial geometry is of minorimportance. The vertical size of the model is 650 km(upper mantle), its horizontal size is 1800 km (Fig. 2A).

3.1.2. Density and rheological structureEach element of the numerical grid is assigned its

specific material phase defined as a subset of physicalparameters of the corresponding material: density,

Fig. 2. A. The model setup (see description in the text). Left side of the model:initial positions of the plume were tested: plume centered below the plate bouand 300 km away from the boundary below the craton. Color code: purple—corange—marker layer at the bottom that otherwise do not differ from ththermotectonic age (left), corresponding yield-stress envelops, computed for ainitial geotherm for craton (left). The lithosphere approaches thermal steadyproperties. In the present case, the upper crustal composition is dominated by qolivine controlled (Table 1). (For interpretation of the references to colour in th

thermal and rheology parameters (Tables 1 and 2). In-stead of using viscous rheology that is common forplume and convection models, we use elastic–viscous–plastic (EVP) rheology, which is more realistic for thelithospheric part of the model. The EVP rheology can berepresented as a serial (Maxwell type) elastic, viscous(ductile) and plastic (brittle) body:

eij ¼ eij e þ eij v þ eij p ð1Þ

where εij stands for the components of strain tensor andsubscripts refer to “elastic”, “viscous” and “plastic”

younger continental lithosphere. Right side of the model: craton. Threendary, 300 km away from the boundary below the younger lithosphererust, blue—mantle lithosphere, green—upper mantle, yellow—plume,e upper mantle. B. Initial lithospheric geotherms as a function ofconstant strain rate of 10−15 s−1 (middle), and example of mantle-scalestate starting from 400 Ma age, which also applies to its mechanicaluartz (granite), the lower crust is composed of diabase and the mantle isis figure legend, the reader is referred to the web version of this article.)

Table 1Notations and physical values common for all experiments (Turcotteand Schubert, 2002; Schubert et al., 2001)

Parameter Values and units Definition

σ, τ Pa, MPa StressP Pa, MPa Pressureu m s−1, mm y−1 Velocity vectorμ 1019−1025 Pa s Effective viscosityε˙ s−

1

Strain rateT °C Temperaturehc 40 km Moho depthh| 100–200 km Thickness of lithosphereD 100–200 km Plume diameterρ| 3330 kg m− 3

Density of mantle lithosphereρm 3400 kg m− 3

Reference mantle density(at 200 km depth)

ρp ρm+Δρch+αρmΔT Plume densityΔρch 0–25 kg m− 3

Chemical density contrastg 9.8 m s−

2

Acceleration due to gravityCp 103 J kg−

1

°C− 1

Specific heatΔT 250°C Initial temperature contrast

plume—backgroundα 3×10−

5

°C− 1

Thermal expansivity

Table 2Specific rheology and related thermal parameters

Parameter Value

All rocksλ, G Lamé elastic constants (λ=G) 30 GPaϕ friction angle (Mohr–Coulomb criterion) 30°C0 cohesion (Mohr–Coulomb criterion) 20 MPa

Specific properties of the crust (quartz dominated)ρ (upper crust) 2700 kg m−3

ρ (lower crust) 2900 kg m−3

n 2.4A 6.7×10−6 MPa−n·s−1

Q 1.56×105 kJ·mol−1

Specific properties of strong lower crust (diabase)ρ 2980 kg m−3

n 3.4A 2×10−4 MPa−n·s−1

Q 2.6×105 kJ·mol−1

Specific mantle properties (olivine)ρ (lithosphere) 3330 kg m−3

n 3A 1×104 MPa−n·s−1

Q 5.2×105 kJ·mol−1

Thermal modelSurface temperature, T0 (fixed) 0 °CCharacteristic temperature at the bottom of the

mechanical lithosphere750°–1000 °C

Characteristic temperature length scale forviscosity drop (Sleep, 2003b), Tη

50–60 °C

Characteristic temperature change inrheological boundary layer,

in Sleep (2003b) notation, Trheo ∼150 °CInitial base of thermal lithosphere, in Sleep

(2003b) notation, Tc

1330 °C

Temperature at 660 km depth (fixed),T1 2000 °CThermal conductivity of crust, k 2.5 W m−1 °C−1

Thermal conductivity of mantle, k 3.5 W m−1 °C−1

Thermal diffusivity of mantle, χ 10−6 m2.s−1

Thermal expansion, α 3×10−5 K−1

Surface radiogenic heat production, Hs 9.5×10−10 W kg−1

Radiogenic heat production decay depth, hr 10 kmThermotectonic age of the lithosphere, a 50–200 Myr and

800 Myr (craton)

Compilation by Burov et al. (2001). ρ is density; Q, n, A are material-dependent parameters of ductile flow laws (Kirby and Kronenberg,1987; Kohlstedt et al., 1995). Other parameters from Turcotte andSchubert (2002).


contributions to the total strain (see Appendix forformulations of the elastic, viscous and brittle rheologiesused). Eq. (1) is solved in the incremental form. Eachstrain increment is computed according to its constitutivelaw, or “rheological term”: linear elastic, non-associatedMohr–Coulomb plastic or non-linear power law ductilecreep (Appendix).

The deep mantle part of the model is effectivelyviscous since at temperatures N1330 °C the viscous/ductile term provides largest contribution to the strainincrement. Two notions of the effective viscosity can beassociated with Eq. (1). The first notion relates to the“true” effective viscosity of the non-linear viscous termdetermined as:

leff ¼ sjj=Aejjv=At ð2Þ

where τ|| is the effective shear stress (second invariant),ε||v is the effective incremental viscous strain (secondinvariant). The second notion relates to the apparentviscosity determined as:

la ¼ sjj=Aðee þ ev þ epÞjj=At: ð3Þ

The numerical code uses the “true” viscosity (Eq. (2)),but the apparent viscosity is also evaluated on the output,for comparison with pure viscous models.

We use conventional density and rheology structure forthe mantle and lithosphere (Turcotte and Schubert, 2002,Fig. 2A, B). All models include 40 km thick crust and four

horizontal rheological layers (Tables 1 and 2; Kirby andKronenberg, 1987; Carter and Tsenn, 1987): (1) a 20 kmthick quartz-rich upper crust with density of 2700 kg/m3;(2) a 20 km thick quartz-rich lower crust with density of2900 kg/m3 (in some experiments we used diabase lowercrust with density of 2980 kg/m3); (3) a 60 or 160 kmthick olivine mantle lithosphere with density of 3330 kg/


m3; (4) deeper mantle with reference density ofρm=3330 kg/m3 below the lithosphere (Turcotte andSchubert, 2002). Although the cratonic mantle may bechemically different from normal lithosphere (depletion),the recent data show that this has much smaller effect onits density than it was previously thought (e.g., PoudjomDjomani et al., 2001). For this reason we use the samedensity for the mantle of young and old lithosphere. Thedensity of the plume has the same dependence on pressureas the background, so the absolute values of thebackground density are not important. A uniformnumerical grid provides vertical and horizontal resolutionof 5 km/element (Fig. 2A). The lithosphere has 30 to 50elements in vertical cross-section, which provides 10–15times better resolution than typical plume models (e.g.Ribe and Christensen, 1994).

3.1.3. Mechanical boundary conditions.Themechanical boundary conditions are assigned at the

four sides of themodel (Fig. 2A). The surface is free. At theleft and right sides, the horizontal velocity is ux=0. At thebottom, hydrostatic pressure is applied. All other compo-nents of stress, displacement or velocity are free (e.g.,horizontal boundaries have vertical free slip condition, thebottom has a horizontal free slip condition etc).

3.1.4. Thermal boundary and initial conditions. Keythermal parameters.

Zero outflow is set as lateral thermal condition atboth sides of the model. Fixed temperatures of 0 °C and2000 °C are imposed as surface and bottom boundaryconditions, respectively (Table 2). The initial thermalcondition refers to application of initial backgroundgeotherm. This geotherm is constructed by joining con-tinental geotherm with deep mantle (adiabatic)geotherm. The continental geotherms are age-dependentand computed from the half-space cooling model ofParsons and Sclater (1977) and Burov and Diament(1995).

The thermal and rheological structure of the older platerefers to a thick (200 km) craton with a nearly steadygeotherm (Parsons and Sclater, 1977) that corresponds toa thermotectonic age of 800 Ma. This age is smaller thanthe geological age of cratons. However, the thermalstructure of continental lithosphere becomes steadyalready after 400 Myr (e.g., Burov and Diament, 1995;Burov, in press). Thus there is no significant difference inthe thermal structure of a 800 Ma and 2000 Ma old plate.For these reasons we call “cratons” all thermally stableplates older than 400–800 Ma.

The thermotectonic age of the younger lithosphericplate has been varied from very young, weak and hot

(50 Ma and 100 Ma) to “normal” (150 Ma and 200 Ma)lithosphere. The thermotectonic age of 150–200 Ma is ofspecial importance for geodynamic implications as itcharacterizes the thermo-rheological state of many con-tinental plates (e.g., Burov, in press). For example, theAlpine lithosphere was thermally reset in Jurassic time.

The combined (lithosphere+mantle) initial geotherm(Fig. 2B) becomes adiabatic (approx. 0.3 °C/km, Sleep,2003b) at depths below 1330 °C; the temperature slowlyincreases from 1330 °C to 1400 °C at 400 km depth, to2000 °C at 650 km depth (e.g., Ribe and Christensen,1994). Peculiar features of the initial geotherm are ofminor importance as the thermal field is continuouslyupdated during further computations (Appendix).

The initial thermal model is validated by computingkey stagnant parameters according to (Solomatov, 1995;Solomatov and Moresi, 1995, 2000; Nyblade and Sleep,2003; Sleep, 2003b). For that, basal heat flow, qm , isfirst roughly estimated from:

dckmðTc−TmohoÞ=qm þ kcðTmoho−T0Þ=ðqs−qmÞ¼ kmðTc−TmohoÞ=qm þ hc;

ð4Þ

where δ is the thickness of the “stagnant lid” (stronglithosphere), qm is the basal heat flux, qc is the meancrustal heat flux, Tc is the temperature at the bottom ofthe lid, Tmoho is the temperature at Moho depth (36 km).For qm=15–20 mW t m−2, δ=200–250 km (craton).For qm=30 mW t m− super 2, δ=120–150 km (normallithosphere). In the stagnant mantle lithosphere, theeffective viscosity drops exponentially as the tempera-ture increases with depth (Fig. 2, Appendix). Theviscosity becomes nearly constant in the convectingsub-lithosphere mantle as the temperature gradientbecomes adiabatic. The depth Zrheo to the zone oftransition from fast viscosity drop (pre-dominantthermal conduction, lithosphere) to nearly constantviscosity (predominant thermal advection, mantle)defines the rheological thickness of the lithosphere.This zone is referred as rheology boundary layer. Thislayer separates strong lithosphere (“stagnant lid”) fromweak convective mantle and its thickness characterizesthe mechanical stability of the “lid”. For deep mantlerheology (Table 2), the estimated characteristic temper-ature length scale for viscosity reduction in the rheologyboundary layer is Tη=50–65 K. This length-scale yieldstemperature change, Trheo, across the rheology boundarylayer. In our model, Trheo, is on the order of 120–150 K(Sleep, 2003a):

Trheoc2:4TDcðTc−TmohoÞ=hc1=A; ð5Þ


where θ is Frank–Kamenetskii parameter (Solomatov,1995), θ≈AΔT where A (∼10−2) is the activationparameter in Schubert et al. (2001, p 618) notation. Thethickness of the rheological boundary layer, Zrheo can beestimated as (Nyblade and Sleep, 2003): Zrheo≈Trheoδ /Tc=50 km. Consequently, the characteristic stress scale,τb∼ (ZrheogρmαTrheo)≈5–7 MPa. Using the estimate forτb, we verify if the model craton would be stagnant in theabsence of a plume. For the geotherms and rheology used(Fig. 2B), the strength of the model craton is higher thanτb down to the depth of 200–250 km. Depending on theage (50 to 200 Ma), the strength of the youngerlithosphere is higher than τb down to the depth of 100–150 km (Fig. 2B). For τb∼5–7MPa and typical tectonicstrain rates of 10−14–10−15 s−1, the effective mantleviscosity is on the order of 1020–1021 Pa s, which is areasonable value for the sub-lithosphere mantle.

The horizontally variable lithospheric structure and“stagnant-lid” convection at the base of the lithospheremay modulate plume convection (Sleep, 2003a). Theconventional Rayleigh number is thus not informative,and one needs a specific “plume” Rayleigh number orplume Stokes velocity to characterize the model (e.g.,Burov and Guillou-Frottier, 2005). Nevertheless, ourmodel is automatically self-consistent because thetemperature field is updated at each time step and thebuoyancy forces are re-adjusted according to the actual(and not reference) thermal distribution.

3.2. Results

The results of the experiments are shown in theFigs. 3–7. We have started from the simplest scenariothat considers a single continental plate impinged by arising plume (Fig. 3) and then introduced two plates ofdifferent age (Figs. 4–6). In the following we describe indetail experiments with most representative 150 Macontinental plate bordering a near steady craton. Theresults of experiments with 50, 100 and 200 Ma platesare outlined in the summary of Fig. 7.

3.2.1. Laterally homogeneous lithosphere (benchmarkreference experiment)

Experiments with a single “normal” (150 Mageotherm) lithosphere (Fig. 3) show that such litho-sphere develops RT instabilities at its base, as aconsequence of the plume impact (d'Acremont et al.,2003; Cloetingh et al., 2004; Burov and Guillou-Frottier, 2005). These instabilities are accelerated bythe return flow at the borders of the plume head. Theplume head arrives at sub-lithospheric depth 0.5–1 Maafter its initialization (this time is compatible with the

results of Burov and Guillou-Frottier, 2005). Plumehead emplacement commences at about 2 Ma andachieves an apex at 5–6 Ma after. At this stage theplume destabilizes the colder (and consequently denser)mantle lithosphere in the areas adjacent to its borders.As the plume head arrives at the base of the lithosphericplate, it starts to flatten horizontally, which results in: (1)initial regional domal uplift of the surface (1000–1500 m), (2) mechanical erosion of the lithosphereabove the centre of the plume head, (3) differentialmantle thinning, (4) forced delamination of thelithospheric mantle from the crust, and (5) down-thrusting of this mantle at the plume head borders. At15 Ma this down-thrusting forms a negative mantledownwelling reaching 400–500 km depth. The litho-sphere above the plume head undergoes thermalweakening and important differential thinning, whichis most prominent in the mantle part (Fig. 3); thethickness of the resistant core of the mantle lithosphereis reduced by a factor of 3–5 already 10 Ma after plumeinitialization. As to the initial domal uplift, in less than1–2 Myr it is replaced by a 300 km wide central basinsurrounded by elevated flanks and two secondarybasins. This subsidence occurs when the plume headtouches the base of the lithosphere at 1–2 Ma. At 5 Mathe secondary basins move away as the plume headspreads laterally and the central basin continues towiden (600 km). At the same time, lithospheric mantlein the central zone is destabilized due to a negative RTinstability and is sagged down inside the plume head,obviously because it is colder and thus denser than theplume material. At about 10 Ma since plume heademplacement, a series of new small basins/rifts and500 m elevations form at surface. The undulations arespaced 200–400 km, while there is no more detectableremnants of the initial large-scale dome.

3.2.2. Young lithosphere and a craton, plume rising belowthe lithosphere–craton boundary

The second set of experiments (Fig. 4) explores theconsequences of plume–lithosphere interaction below thetransition betweenweaker and hotter “young” or “normal”lithosphere and an old cold thick “craton” lithosphere of400–800 Ma age (analogous to the Eastern Europeanplate). We vary the thermotectonic age of the youngerlithosphere from 50 Ma to 200 Ma.

3.2.2.1. Experiments with 50, 100 and 150 and 200 Maold lithosphere (Figs. 4 and 7). In these experiments,sharp change in the vertical density and thermalstructure at the boundary between cold craton andyounger lithosphere leads to development of a marginal

Fig. 3. Preliminary test with a single continental lithosphere of representative 150 Ma age and a 200 km-diameter plume. This experiment in based onthe model from Burov and Guillou-Frottier (2005). Color code for the phase field: purple—crust, blue—mantle lithosphere, green—upper mantle,yellow—plume, orange—marker layer at the bottom. Note significant erosion of the mantle lithosphere, drip-like down-sagging of the lithosphericmantle, with some bits of the lower crust, above the middle of the plume head, complex behavior of surface (periodic undulations/basins and uplifts).Surface topography (right) is obtained directly from the simulations. Circular diagrams show zones of plastic yielding (= faulting), following samegraphical conventions as the Schmidt's diagrams for focal depth solutions used in seismology. Note that in nature drip-like instabilities may beenhanced by metamorphic phase changes in the lower crust that basically produce significantly denser material. The small secondary diapirs refer tothe fact that we use a source layer at the bottom and fixed temperature boundary condition, instead of a localized infinite source region and fixedbottom heat flux as in some models. This assumption fits with recent tomography-based models (i.e. Goes et al., 2004) that suggest that plumes havethick tails comparable with their head size when they arrive to the upper–lower mantle boundary and then a smaller-headed plum with much thinnertail rises through the upper mantle. The horizontal axis is distance in km (from 0 to 1500 km).


gravitational instability. Such instability may resulteither in a drip-off of the border material or in smoothingof the border. In the beginning (first 2 Myr), the verticalborder of the craton starts to collapse, that is, it getssmoothed due to the horizontal pressure gradient acrossthe boundary between the lithospheres. Yet, when thelaterally spreading plum head material reaches thecratonic boundary, it pushes aside and drags down the

material of the craton boundary resulting in restorationand even in downward stretching of the initial verticalboundary between the younger and older lithosphere.The difference in the thermo-rheological structurebetween the two lithospheres eases preferential plumeflattening under the younger hotter and weaker litho-sphere. Horizontally spreading plume head materialproducesmultiple instabilities in the overlying lithospheric

Fig. 4. Experiment with two plates (younger 150 Ma old lithosphere and 800 Ma old craton); plume head is located in the centre, below the plateboundary. Same color code as in Fig. 3. Note significant erosion of the mantle lithosphere, “verticalization” of the cratonic boundary, multiple down-sagging of the lithospheric mantle above the middle of the plume head with hot material arriving at Moho depth, asymmetric flattening of the plumehead towards the younger lithosphere, down-thrusting of the mantle lithosphere at both sides of the plume, and asymmetric doming at surface.


mantle, in particular, drip-like down-sagging of thisheavier mantle with some bits of the lower crust insidethe plume head. At the border of the plume head, themantle lithosphere delaminates from the positively buoy-ant part of the crust and starts to sink down.An asymmetricdomal uplift of approx. 1500 m initially forms above theplume head. The amplitude of local shorter-wavelengthdeformations, for example above the plate borders, laterreaches 2000–3000m.Upwellings of hotmaterial arrivingfrom the flattening plume head erode mantle lithosphereand stall at Moho boundary. At this boundary, theupwellings are blocked by higher buoyancy of thecontinental crust (density of 2700–2900 kg/m3). Largeamounts of hot mantle plume material thus can bestocked below the Moho for an important period of time,and only a fraction of light magmas produced by partialmelting of this material may be able to ascend to thesurface through dikes or by pluton emplacement. In suchcontexts, the absence of magmatic events should not beinterpreted as a decisive evidence for the absence of plumeevent. In this study, we do not model partial melting.However, it may be suggested that even if some melts/

magmatic events will be observed at surface, they willnot necessarily have unambiguous deep signatures as thehot sub-Moho reservoir also serves as heat source forcrustal melting and thus may produce induce crustalmelting.

The difference between very young (50 Ma–100 Ma)and older (colder and more resistant) normal lithosphere(150 Ma and 200 Ma, Figs. 4 and 7) is that the domalsurface uplift is rapidly replaced by a series of smallerscale ups and downs. The other features basicallyresemble the case of young lithospheres except that hotmaterial is less penetrative upwards but has a higherability for horizontal flattening. In this case it also has ahigher ability for producing gravity instabilities at thebottom of the overlying lithosphere (because of higherdensity contrast). In all cases, plume head spreadsasymmetrically: it propagates to a much larger distancetowards the external boundary of the normal lithospherethan towards the craton. The plume erodes a part of thenormal lithosphere and provokes down-thrusting of thelithospheric mantle at the front of the laterally pro-pagating plume head material.


3.2.3. Normal (younger) lithosphere and a craton,plume rising below the normal lithosphere

Experiments with 50, 100 ,150 and 200 Ma old plate(Figs. 5 and 7). In this set of experiments, the plumeascends below the normal lithosphere, 300 km away

from the cratonic border (Fig. 5). As in the previousexperiments, the gravitational instability of the cratonicborder ,is prevented by horizontal push from the laterallyspreading plume head. Near the border, this spreadingtransforms in the localized downward return flow in the


plume head. This flow drags the cratonic material downand forms a sharp vertical boundary between the plumeand the craton. This cold high viscosity boundaryprevents craton-ward propagation of the horizontallyspreading plume head material, which is thereforechanneled towards the weaker normal lithosphere. Thisresults in asymmetric flattening of the plume head and inpreferential thinning of the normal lithosphere and inslab-like mantle down-thrusting at its borders. Thedown-thrusting of the borders of weaker and thinnerlithosphere is less pronounced (200 km depth) than of thecratonic border, whichmaterial reaches the 350 km depthin 15 Myr (see temperature and effective viscosity fieldin Fig. 5B). The total width of the normal lithosphereaffected by thinning and weakening due to plume headflattening is about 1000 km, i.e. 5 times wider than theinitial plume head diameter. The surface response to theplume head uprise shows moderate doming above theplume head almost compensated by subsidence due toregional extension and thinning of the lithosphere (rightplot in Fig. 5A).

Plume flattening is strongly asymmetric, with pro-nounced extension of the young lithosphere and withcentre of spreading moving away from the craton–lithosphere boundary. The area above the plume head(extension) is surrounded by peripheral zone of com-pression and down-thrusting that form small compres-sion belts. One of the striking features of the surfacetopography refers to the pronounced subsidence at thecratonic border (Fig. 5A) caused by the subduction-likedown-thrusting of the cratonic mantle. The resultingtopography resembles that of the Pannonian basin–Carpathian transect (Cloetingh et al., 2004, see discus-sion below).

3.2.4. Normal (younger) lithosphere and a craton,plume rising below the craton

3.2.4.1. Experiments with 50, 100, 150, 200 Ma oldplate (Figs. 6 and 7). In this set of experiments, theplume rises below the craton, 300 km away from itsborder (Fig. 6). The hot plume head material arrivingbelow the craton, erodes its central part and spreadshorizontally to its borders (please note that on the right

Fig. 5. A. Experiment like in Fig. 4 but with plume head located below thsignificant erosion of the mantle lithosphere, “verticalization” of the cratonicmiddle of the plume head with hot material arriving at Moho depth, asymmetrthrusting of the mantle lithosphere at both sides of the plume, asymmetric do(ratio of stress to strain rate for all, brittle, elastic and ductile domains) and tempmantle lithosphere at both sides of the plume, drip-like down-sagging above thcrust and is “subducts” down the depth of 300 km (“slab-like” instability, i“subducts” to 200 km depth.

side of the craton, the effect of this flow should beignored due to the vicinity of the right wall of themodel). The material arriving at the transition zonebetween the normal lithosphere heats up this zoneconsiderably. The gravitationally unstable border of thecraton then gets smoothed within 3 Myr. The mantle ofthermally-weakened normal lithosphere delaminatesfrom the crust and is down-thrusted, together with thecratonic mantle, below the craton (see the material phasefield in Fig. 6). It may be suggested that the flow of theplume material to the cratonic borders may producesome volcanic activity focalized at the transition zonebetween the normal lithosphere and the craton. It is alsonoteworthy that surface topography of the craton isstrongly affected by the spreading of the plume: thecentral region of the craton takes plateau-lake shapeswhereas the borders are characterized by verticalundulations (Fig. 7B). We later discuss that thistopography profile resembles the Tanzanian craton.

The plume impact produces a similar effect for allages of the “normal” plate (50–200 Ma). The majordifference concerns the surface amplitude and wave-length that tend to increase with the plate age. Zones ofextension and thermal heating are first localized at theborders of the craton but then start to move progres-sively away inside the normal lithosphere. These zonesare more localized than in the case of hot initiallithosphere and in nature may present narrow deep riftbasins. Importantly, these basins are found far awayfrom the centre of the plume head.

4. Discussion. Implications for interaction betweenthe upper mantle and lithospheric processes

Results of the experiments fit first-order rheologicaland geophysical features of a number of known areaswhere plume–lithosphere interactions are suspected, suchas the present-day active Pannonian–Carpathian system,or cratonic areas such asWest Africa, Tanzania, southwestAustralia, or possible plume-induced delamination of theColumbia River FloodBasalt Province (Hales et al., 2005;V. Camp, personal communication, 2006). A number ofkey results for mantle and lithosphere dynamics aredetailed below with geological implications.

e younger lithosphere, 300 km away from intra-plate boundary. Noteboundary, moderate extension and multiple down-sagging above theic flattening of the plume head towards the younger lithosphere, down-ming at surface and multi-scale undulations. B. The effective viscosityerature for the experiment shown in Fig. 5A. Note down-thrusting of theemiddle of the plume head. Strong cratonic mantle delaminates from then green in the temperature field). The mantle of younger lithosphere

Fig. 6. Experiment like in Fig. 4 but with plume head located below the craton, 300 km away from plate boundary. Same color code as in Fig. 3 Notesignificant erosion of the cratonic mantle lithosphere, asymmetric flattening of the plume head and down-thrusting “subduction” of the younger andcratonic mantle lithospheres below the craton, and permanence of a topographic “plateau” above the craton. The effect of plume head spreading on theright side of the craton should be ignored due to the vicinity of the boundary of the model.


4.1. Lithospheric thinning, extension and compression.Predicted upper mantle structure and constraints fromseismic tomography

The thermal structure of the upper mantle created at12–14 Ma (Fig. 4) after the arrival of the mantle plumedisplays a number of common features with thetomographic images of cratonic margins.

For example, one of the most striking features ofCarpathians dynamics is the formation and evolution of theFocsani Depression, a 13 km thick Miocene–Pliocenebasin (e.g. , Cloetingh et al., 2004), situated in the vicinityof the Vrancea earthquake cluster (Fig. 8A) expressed in a“cold” 300–400 km deep tomographic anomaly (e.g.Wortel and Spakman, 2000).Whether or not this representsthe present position of a Beniof plane is disputed, mostlybecause the peculiar 85° dip of the downgoing lithospheric“slab” is situated almost in the front of the orogen. This hasbeen so far explained as a late Miocene to Pliocene roll-back effect in the SE Carpathians, which could have cul-minated when the subducting slab tore (slab detachmentsensu) along the trend of the orogenic arc. The latter iscommonly thought to have induced a rebound in the

northern, torn, parts of the slab, and subsidence of the lowerplate in the bending area, where the slab still holds itsintegrity: this occurred simultaneouslywith a fastmigrationtowards the foreland, leading to the opening of the thickforedeep (Focsani) basin. One can suggest, however, thatthe Miocene plume that has initiated extension in thePannonian basin could also drive down-thrusting of thelithosphere around its borders (Fig. 8A), which can explainthe simultaneously occurring extension and compression inthe surrounding Carpathians. In this context, particularlystriking are the model-predicted downwellings of the coldmantle to depths of 300–400 km that can be matched withthose observed in the Pannonian–Carpathian system(Cloetingh et al., 2004, Fig. 8A). Inspection of the modelshown in Figs. 4, 5 or 7 shows the presence of cold sub-vertical cratonic lithosphere down to the depth of 300 kmflanked by attenuated thinned stretched lithosphere with adramatic thinning of its mantle component (e.g., see vis-cosity structure in Fig. 5B at 16.1My). This abrupt verticaltransition between the two types of lithosphere is strikinglysimilar to the pronounced near vertical contrast in seismicvelocities in the Vrancea zone (Wortel and Spakman, 2000,Fig. 8A). The predicted down-thrusting of the normal


lithosphere at the left border of themodel (Fig. 4, snapshots5.2 and 11.3My or viscosity structure in Fig. 5B, snapshots11 and 16.1 My) also resembles cold downwellings belowthe Western border of the Pannonian basin observed fromthe seismic tomography (Fig. 8A) Consequently, ourexperiments may provide a simple explanation to: (1)simultaneous extension in the middle of the basin andcompression at its borders, (2) the occurrence of anomalousthinning of the mantle lithosphere below the Pannonianbasin, and (3) the extreme low flexural rigidities recorded inthis basin strikingly contrasting with the presence of therelatively strong lithosphere in the surrounding areas (e.g.,Wortel and Spakman, 2000; Cloetingh et al., 2004).

In the Tanzanian craton area (Fig. 8B, top), two verticalcross-sections of velocity anomalies published in (Weer-aratne et al., 2003) clearly show near vertical cratonicedges (high velocity zones, extending down to 150 kmdepth), some ofwhich (south-west and north-east borders)being surrounded by 100 km-wide low velocity zones. On

Fig. 7. A. Comparison of mature stages (7.5 Myr since initialization) of expeyounger lithosphere (thermal field). Left column: plume rising below the younplume rising below the craton. In all cases, the presence of intraplate bolithosphere and in down-thrusting of the mantle lithosphere at plume borders.thermal ages (topography). Left column: plume rising below the younger lithoplume rising below the craton. In all cases, the presence of intraplate boundaryundulations superimposed onto larger scale features.

the opposite side, the south-west border of the craton doesnot show such a vertical boundary and is not associatedwith a low velocity zone. The regional tomographicmodel of Sebai et al. (2006) supports the possibility of hotmaterial such as a plume head beneath the Tanzaniancraton, as does the two-plume scenario of Lin et al. (2005).

4.2. Lithospheric instabilities and crustal growth

Formation and growth of continental crust has beenassociated with various accretionary processes, includingoceanic plateau or island arc accretion, back-arc basin ornormal oceanic crust accretion (e.g., Kusky and Polat,1999; Condie, 2002). The plume head–craton margininteraction may in fact provide another mechanism forcrustal growth, by lateral spreading of lower crustalmaterial. Indeed, the physical possibility for lower crustalconvective instability has been confirmed by Jull andKelemen (2001), who found that the conditions needed

riments for different initial thermal ages (50, 100, 150, 200 Ma) of theger plate; Middle—plume rising below the cratonic boundary; Right—undary results in asymmetric plume spreading towards the youngerB. Comparison of developed stages of experiments for different initialsphere; Middle—plume rising below the intraplate boundary; Right—results in asymmetric surface features and in small-scale (100–250 km)

Fig. 7 (continued ).


for development of Rayleigh–Taylor (RT) instabilities onthe time scale of 10Myr can be reachedwithin continentalplates underlain by a mantle plume. Recently, alithospheric instability has been proposed as a plume-activated mechanism of formation of flood basalts andtopographic signatures. In particular, Hales et al. (2005)have suggested that the flood basalts of the ColumbiaRiver are associated with mantle dynamics of theYellowstone hotspot. Hales et al. (2005) interpret thesebasalts and the particular topographic signatures of thearea in terms of interactions between the convectivedownwelling and detachment of a dense plutonic root.These authors suggest that convective instabilities mighthave been triggered by propagating hot mantle materialalong the north Precambrian continental margin, where anincrease in lithosphere thickness is suggested (Artemievaand Mooney, 2002). Plume-induced horizontal migrationof deep continental material towards cratons edges maythus represent another mechanism for lateral growth ofcontinental blocks.

Numerical experiments of Figs. 4 and 5 have de-monstrated the possibility for Rayleigh–Taylor insta-bilities to occur few million years after the plumearrival at the base of the lithosphere. Thermal weaken-ing of the lithosphere by plume head involves twotypes of RT instabilities, that we call “drip-like” and“slab-like” instabilities. The drip-like downwellings(easily identifiable in thermal fields of Figs. 4 and 5B)are located above the plume head center, where heatingis most efficient. These dynamic features have beensuggested by several authors to explain convectiveremoval of the deep crust (e.g. Saleeby et al., 2003).Due to subsequent arrival of hot upwelling material,mixing of lithospheric and asthenospheric materialdestroys drip-like instabilities in a few million years.The slab-like instabilities seem to develop more slowlyand to last longer. Fig. 5B (thermal and viscosity fields)illustrate such instabilities, having formed at bothedges of plume head. Spreading of plume head towardsyoung lithosphere would probably have a mixing


effect, similarly to drip-like instabilities. On the op-posite side, the near-craton slab-like instability is en-trained by the return downward flow of the plumematerial, leading to a deep subsidence of lithosphericmaterial, despite its high viscosity. These two kinds oflithospheric instabilities may trigger and enhancemelting processes, since the lithospheric material isbrought in contact with hot mantle, especially near

Fig. 8. A. Top: Crustal-scale cross-sections across the East Carpathians (modifinorthern cross-section in the area affected by the collision with the East Europein the bend area influenced by the collision with highly buoyant Moesian blotrace. Bottom: Seismic tomography (after Wortel and Spakman, 2000) acrosspositive and negative P-wave anomalies corresponding to the presence of coldVrancea earthquakes. B. Topographic profiles across the Tanzanian craton areundulations can be identified, from latitude−2°S to−6°S. Bottom: comparisonwFig. 7A and B, in case of a plume ascending below a lithospheric plate of limitedplume head center. (For interpretation of the references to colour in this figure

cratonic margins where hot plume head stalls. As re-sult, focusing of newly formed crust at cratonic mar-gins may favor continental growth.

In West Africa, the major evolution of Birimian crust(∼2.1 Ga) was suggested to be completed in less than50 Ma (Boher et al., 1992), allowing these authors tosuggest an enhanced plume activity close to Archeannuclei followed by collision between oceanic plateaus and

ed after Radulescu, 1988; Cloetingh et al., 2004). Red lines represent thean/Scythian rigid block. Blue lines represent the southern cross-sectionck. Red dots represent earthquake projection into the southern sectionthe bend zone of the East Carpathians in the MB domain. Blue and red:and hot mantle in the area, respectively. Open circles give the location ofa, where a “topographic plateau” surrounded by short-scale topographicith topographic signature generated by themodels (Fig. 7B). As shown insize, plume edges may have more impact on surface topography than thelegend, the reader is referred to the web version of this article.)

Fig. 8 (continued ).


the Man Archean craton. Geochemical mixing betweenmantle and lithospheric material (as for the “slab-like”lithospheric instability) would also fit the geochronolog-ical and petrological constraints detailed in Abouchamiet al. (1990). InWestern Australia, a similar scenario witha mantle plume event at 2.0 Ga around cratonic areas, hasalso been suggested by Pirajno (2004).

4.3. Lithospheric faulting and near-craton ore deposits

Deep suture zones, identified in Figs. 4 and 5, tend tolocalize at cratonic margins and are associated withhighest thermo-mechanical contrasts. Combined with ahigh melting potential, these sutures may provide path-ways for fluids and granitoids to migrate upwards andform ore deposits (Kusky and Polat, 1999). Anomalousconcentration of metals in the lithosphere has beenassociated with large-scale geodynamic processes (e.g.Neubauer et al., 2005), and the role of deep lithosphericstructures and discontinuities has been outlined by severalauthors (e.g. Kutina et al., 2003). In particular, the role ofplume head is often evoked (Barley et al., 1997, 1998;Pirajno, 2000) to account for pulses of magmatismbetween 2.8 and 2.6 Ga and associated mineralization,such as gold-rich granite–greenstone terrains in theSuperior Province and Yilgarn craton (e.g. Groves andBarley, 1994).

According to Isley and Abbott (1999), mantle plumeactivity in Proterozoic and Archean times correlatesstrongly with the ages of banded iron formations (BIFs).Indeed, Fe-rich (tholeitic) magmas are mainly eruptedduring mantle plume events, and it appears that BIFs arespatially correlated with Archean cratons. Another possibleexample relating to cratonic area, ore deposits andlithospheric discontinuities is the east African gemstonebelt, including the Tanzanian Ruby belt (Le Goff et al.,2004), several hundred of kilometers long nearby theeastern branch of the East African rift. Thermobarometricconditions for genesis of these gemstones would involveparticular isothermal deepening of lower crust, as identifiedwith the "slab-like" lithospheric instability. Relationsbetween archean cratons, diamond deposits and kimberliteevents have also been discussed by several authors, amongwhich O'Neill et al. (2005) focus on abrupt changes inlithospheric thickness to explain coeval lithosphericdiscontinuities and enhanced melt production.

4.4. Surface behavior

Depending on the assumed plate rheology, modelspredict initial doming before the plume emplacement, andlocalized subsidence during the emplacement. The latter is


associated with extension and lithospheric instabilitydowning inside the plume head (b2 Ma, first riftingphase, Burov and Guillou-Frottier, 2005). This phase isfollowed, after 3–5Ma, by a second rifting phase, which isassociated with plume head flattening and is marked byuplifts (∼1000 m) on the sides of the basin. Plume headflattening provokes tensional instabilities in the litho-sphere,which result in formation of series of narrowbasins.At the end, the system of highs and lows of intra-basintopography may develop, with wavelengths between 300and 500 km. It is noteworthy that such wavelengths andperiodic surface undulations are often attributed totectonically induced lithospheric instabilities (e.g., Smith,1975). In case of a plume rising below the normallithosphere or below the cratonic margin, there is animportant subsidence on the side of the craton and strongtopography elevation on the side of the normal lithosphereafter 10–16Myr. This type of behavior is strikingly similarto what is observed in a number of plume–lithosphereinteraction zones such as Pannonian basin (e.g., Cloetinghet al., 2004, Fig. 8A), where East-Carpathian “foredeep”resembles the subsidence predicted in the model (Figs. 4and 5A), whereas the East Carpathians mountains areformed on the weak lithosphere, which refers to thepredicted uplift on the side of the normal lithosphere.

In the presence of highly contrasting plate boundaries(normal lithosphere–craton), when the plume head hasbegun to spread, surface topography shows particularareas where elevation gradients are higher than elsewhere(Figs. 4, 5A, 6, and 7B). In other words, surfacetopography above the cratonic area shows little variationwhereas cratonic borders are characterized by short-scaleundulations. Fig. 8B shows west–east topographicprofiles across the Tanzanian craton area, where a“topographic plateau” surrounded by short-scale topo-graphic undulations can be identified, from latitude −2°Sto −6°S (see also Weeraratne et al., 2003). Model resultsare highly consistent with such signatures (Fig. 7B).Clearly, as shown in Fig. 7A, B, plume edges have muchmore impact on surface topography than the plume headcenter, above which the lithospheric stresses relax withinfirst 1–4 millions of years.

5. Discussion and conclusions

Lateral heterogeneities in the lithospheric struc-ture have four major effects on plume–lithosphereinteractions:

(1) Horizontal plume head flattening is blocked fromone side by cold vertical boundary of the craton,leading to crust–mantle decoupling, stress and

heat concentration and faulting at the cratonicmargin; mantle lithosphere can be thrusted downat both sides of the plum head as a result of lateralspreading of the plume head material.

(2) Horizontal push exerted on the cratonic boundaryby the spreading plume, and the return flow fromthe plume head, may result in delamination andnear vertical down-thrusting of the mantle litho-sphere at the cratonic margin producing sharp sub-vertical “cold” boundaries going down to thedepths of up to 300–400 km. This delaminationresembles the delamination of the “rheologicalboundary layer” at the bottom of the lithospherepredicted by Nyblade and Sleep (2003) for amono-layer lithosphere. However, in case ofstratified rheology, the delamination is likely tooccur at Moho boundary or within the lower crust.A large part of the mantle lithosphere, and somepart of the lower crust can be delaminated fromthe buoyant part of the crust and trapped by thespreading plume head.

(3) The presence of a cratonic boundary promotesasymmetric plume head flattening towards theyounger plate, followed by surface extension andlithospheric down-thrusting both at the cratonicmargin and at the far end of the young plate. Thiscreates a possibility that extension and compres-sion occur simultaneously; same process wouldresult in enhanced melting events at cratonicedges, and favor continental growth at the side ofthe craton.

(4) Topographic signatures of the plume–litho-sphere interaction show basin-scale uplifts andsubsidences due to the interplay between rheo-logical stratification, visco-elastic relaxation andplume-induced mantle instabilities. These sig-natures have “typical tectonic” wavelengths(b500 km) while long-wavelength deformationusually associated with plume impact is almostundetectable.

Rayleigh–Taylor instabilities that develop at the topand the sides of the plume head provide a mechanism forcrustal delamination. Lateral flow of mantle lithosphere,from plume head to the base of the craton, suggestsa new mechanism for crustal growth, where surfacemagmatism is not required. Lithospheric faulting atcratonic edges and enhanced magmatic activity couldexplain the apparent plume-related metallogenic crises,as suggested for West Africa and Australia. Plume–lithosphere interactions can explain a number of otherkey phenomena such as the simultaneous occurrence of


climax of extension of the young plates/segments andclimax of compression in the surrounding belts. Thisscenario is consistent with differential mantle litho-sphere thinning of the younger plate.

This study suggests that the plume impact may havecomplex consequences for surface evolution. In par-ticular, absence of magmatic events should not beinterpreted as evidence for the absence of plumeevents. If some melts/magmatic events are observed atthe surface, they will not necessarily have unambig-uous deep geochemical signatures, as the hot sourceplume material stalls below Moho and forms a long-lasting (10 to 100 Myr) sub-Moho reservoir. Thisshould induce strong crustal melting that mayoverprint deeper signatures since crustal melts aregenerated at much lower temperatures than mantle,and produce light low-viscous rapidly ascending mag-mas. In addition, drip-like down-sagging of the li-thospheric mantle and metamorphic lower crustalmaterial inside the plume head may contaminatethe latter and thus alter the geochemical signatureof plume-related magmas. However, this conclusionneeds further investigation as we did not take intoaccount metamorphic phase changes in the crust andpartial melting.

Acknowledgements

We are gratefully thankful to N. Sleep and theanonymous reviewer for detailed and highly construc-tive comments. We thank for the financial support ofBRGM (Bureau des Recherches Geologiques et Mini-eres). This is a BRGM contribution 4572. The code Para(o)voz v.9d originates from Paravoz v.3 by Poliakov andPodladchikov (Poliakov et al., 1993).

Appendix A. Numerical model

A.1. Basic equations

The Para(o)voz (previously Paravoz) is “2.5 D”FLAC-like code (Cundall, 1989). The term “2.5D”refers to the fact that strain/stress relations are solved in3D formulation but the model geometry is cylindrical2D (out-of plane boundary conditions for stress are setfor plain stress). The code has a mixed finite-difference/finite element numerical scheme, in which the coordi-nate frame is Cartesian 2D, but stress/strain relations arecomputed in 3D formulation. The Lagrangian mesh ofPara(o)voz is composed of quadrilateral elementssubdivided onto 2 couples of triangular sub-elementswith tri-linear shape functions. Para(o)voz is a large

strain fully explicit time-marching algorithm. It locallysolves full Newtonian equations of motion in continuummechanics approximation:

hq �ui−divs−qg ¼ 0 ðA1Þ

coupled with constitutive equations:

DrDt

¼ Fðr; e;ju; N T N Þ ðA2Þ

and with the equations of heat transfer and Boussinesqapproximation:

qCpAT=AT þ ujT−k divðjTÞ−Hr ¼ 0q ¼ q0ð1−aTÞ ðA3Þ

Here u, σ, g, k are the respective terms for velocity,stress, acceleration due to body forces and thermalconductivity. The triangular brackets in Eq. (A1) specifyconditional use of the related term (in quasi-static modeinertial terms are dumped using inertial mass scaling(Cundall, 1989). The terms t, ρ, Cp, T and Hr designaterespectively time, density, specific heat, temperature andinternal heat production. The terms ∂ /∂t, Dσ /Dt, F area time derivative, an objective Jaumann time derivativeof stress and a functional, respectively. In the Lagrang-ian method, incremental displacements are added to thegrid coordinates allowing the mesh to move and deformwith material. This enables solution of large-strainproblems locally using small-strain formulation: on eachtime step the solution is obtained in local coordinates,which are then updated in large strain mode.

Solution of Eq. (A1) provides velocities at meshpoints; these velocities are used for computation ofelement strains and of heat advection u∇T. The strains areused in Eq. (A2) to calculate element stresses and forces;the forces are then used to compute velocities for the nexttime step. Due to the explicit approach, there are noconvergence issues in case of non-linear rheologies. Thealgorithm automatically checks and adopts the internaltime step using Courant’s criterion for propagation ofinformation, which warrants stable solution.

A.2. Explicit EVP rheology

We use serial (Maxwell type) body (Eq. (1)), in whichtotal strain increment in each numeric element is definedby a sum of elastic, viscous and brittle strain increments.In contrast to fluid dynamic approaches, where non-viscous rheological terms are not implemented directly


but simulated using pseudo-plastic and pseudo-elasticviscous terms (e.g. Solomatov and Moresi, 2000;Bercovici et al., 2001), our method explicitly treats allrheological terms. The parameters of elastic–ductile–plastic rheology laws for crust and mantle come fromrock mechanics data (Tables 1 and 2; Kirby andKronenberg, 1987; Kohlstedt et al., 1995).

A.3. Plastic (brittle) behavior

The brittle behavior of rocks is described byByerlee's law (Byerlee, 1978; Ranalli, 1995), whichcorresponds to Mohr–Coulomb material with frictionangle ϕ=30°, small dilatation angle (5°) and cohesion∣C0∣b20 MPa (e.g., Gerbault et al., 1998):

jsj ¼ C0−rntan� ðA4Þwhereσn is normal stressσn=1/3σ| +σ||

dev sinϕ, 1 /3σ| =Pis the effective pressure, σ||

dev is the second invariant ofdeviatoric stress, or effective shear stress. The condition oftransition to brittle deformation (function of rupture f)reads as: f=σ||

dev+P sinϕ−C0 cosϕ=0 and ∂f /∂t=0. Interms of principal stresses, the equivalent of the yieldcriterion (5) reads as:

r1−r3 ¼ −sin�ðr1 þ r3−2C0=tan�Þ ðA5Þ

A.4. Elastic behavior

The elastic behavior is described by linear Hooke'slaw:

rij ¼ keiidij þ 2Geij ðA6Þ

where λ and G are Lame's constants. Repeating indexesmean summation and δ is Kronecker's operator.

A.5. Viscous (ductile) behavior

Mantle convection and a part of lithosphericdeformation is controlled by thermally activated creep(Kirby and Kronenberg, 1987; Ranalli, 1995). Withindeep lithosphere and underlying mantle regions, creep-ing flow is non-Newtonian since the effective viscositycan vary within 10 orders of magnitude as a function ofdifferential stress:

edjj ¼ Aðr1−r3Þnexpð−QR−1T−1Þ ðA7Þ

Where ed|| =ε˙=(Inv|| (eij))1 / 2 is effective shear strain

rate, A is material constant, n is power law exponent, Qis activation enthalpy, R is Boltzman's gaz constant, and

T is temperature in K, σ1 and σ3 are the principalstresses. The effective viscosity μeff for this law isdefined as:

sij ¼ leffedjj

which yields:

leff ¼ edjjð−1=nÞ=nA−1=nexpðQðnRTÞ−1Þ ðA8Þ

For non-uniaxial deformation, the law (A8) isconverted to a triaxial form, using invariant of strainrate and geometrical proportionality factors:

(A9)leff ¼ edjjð1−nÞ=nðA⁎Þ−1=nexpðQðnRTÞ−1Þ

where A⁎ ¼ 0:5Ad 3ðnþ1Þ=2

parameters A, n,Q are experimentally determined materialconstants (Table 2). Using olivine parameters (Table 2),one can verify that the predicted effective viscosity at thebase of the lithosphere is 1019–5×1019 Pa s matching thepost-glacial rebound data (Turcotte and Schubert, 2002). Inthe depth interval of 200 km–0 km the effective viscosityexponentially grows from 1019 to 1025–1027 Pa s withdecreasing temperature. Within the adiabatic temperatureinterval below200 kmdepth, the viscosity slowly increasesfrom 1019–1021 Pa s (e.g., Turcotte and Schubert, 2002).Weinberg and Podladchikov (1994) have also shown thatthe effective viscosity in close vicinity of an ascendingdiapir is influenced by local strain rate field and partly byheat exchanges between the diapir and surrounding rock.When such a diapir translates to the surface, theenvironment above and in the vicinity of the plum headis weakened, which eases its ascent (self-lubrification).

References

Abouchami, W., Boher, M., Michard, A., Albarede, F., 1990. A major2.1 Ga event of mafic magmatism in west Africa: an early stage ofcrustal accretion. J. Geophys. Res. 95, 17605–17629.

Anderson, D.L., 1982. Hotspots, polar wander, Mesozoic convectionand the geoid. Nature 297, 391–393.

Anderson, D.L., 1998. The scales of mantle convection. Tectonophy-sics 284, 1–17.

Artemieva, I.M.,Mooney,W.D., 2002. On the relations between cratoniclithosphere thickness, plate motions, and basal drag. Tectonophysics358, 211–231.

Barley, M.E., Pickard, A.L., Sylvester, P.J., 1997. Emplacement of alarge igneous province as a possible cause of banded ironformation 2.45 billion years ago. Nature 385, 55–58.

Barley, M.E., Krapez, B., Groves, D.I., Kerrich, R., 1998. The LateArchean bonanza: metallogenic and environmental consequencesof ther interaction between mantle plumes, lithospheric tectonicsand global cyclicity. Precambrian Res. 91, 65–90.


Batchelor, G.K., 1967. An Introduction to Fluid Dynamics. CambridgeUniv. Press. 615pp.

Bercovici, D., Ricard, Y., Schubert, G., 2001. A two-phase model forcompaction and damage 3. Applications to shear localization andplate boundary formation. J. Geophys. Res. 106, 8925–8939.

Boher, M., Abouchami, W., Michard, A., Albarède, F., Arndt, N.T.,1992. Crustal growth in west Africa at 2.1 Ga. J. Geophys. Res. 97,345–369.

Burov, E.B., Plate rheology and mechanics, Ed.: G. Schubert, in press,Treatise on Geophysics, Volume 6 — Crust and LithosphereDynamics (Volume Edt. A.B. Watts), AGU-Elsevier.

Burov, E.B., Diament, M., 1995. The effective elastic thickness (Te) ofcontinental lithosphere: What does it really mean? J. Geophys.Res. 100, 3905–3927.

Burov, E.B., Poliakov, A.N.B., 2001. Erosion and rheology controlson synrift and postrift evolution: verifying old and new ideasusing a fully coupled numerical model. J. Geophys. Res. 106,16461–16481.

Burov, E.B., Guillou-Frottier, L., 2005. The plume head–continentallithosphere interaction using a tectonically realistic formulation forthe lithosphere. Geophy. J. Int. 161, 469–490.

Burov, E., Watts, A.B., 2006. The long-term strength of continentallithosphere: “jelly-sandwich” or “crème-brûlé”? GSAToday 16, 1.doi:10.1130/1052-5173(2006)016b4:TLTSOC.

Burov, E., Jolivet, L., Le Pourhiet, L., Poliakov, A.N.B., 2001. Athermomechanical model of exhumation of high pressure (HP) andultra-high pressure (UHP) metamorphic rocks in alpine-typecollision belts. Tectonophysics 342, 113–136.

Burov, E.B., Jaupart, C., Guillou-Frottier, L., 2003. Ascent andemplacement of magma reservoirs in brittle–ductile upper crust.J. Geophys. Res. 108, 2177. doi:10.1029/2002JB001904.

Byerlee, J., 1978. Rock friction and earthquake prediction. Pure Appl.Geophys. 116, 615–626.

Campbell, I.H., Griffiths, R.W., Hill, R.I., 1989. Melting in an Archeanmantle plume: Heads it's basalts, tails it's komatiites. Nature 339,697–699.

Campbell, I.H., Griffiths, R.W., 1990. Implications of mantle plumestructure for the evolution of flood basalts. Earth Planet. Sci. Lett.99, 79–93.

Carter,N.L., Tsenn,M.C., 1987. Flowproperties of continental lithosphere.Tectonophysics 36, 27–63.

Cloetingh, S., Burov, E., Matenco, L., Toussaint, G., Bertotti, G., 2004.Thermo-mechanical constraints for the continental collision modein the SE Carpathians (Romania). Earth Planet. Sci. Lett. 218 (1–2), 57–76.

Condie, K.C., 2002. Continental growth during a 1.9-Ga superplumeevent. J. Geodyn. 34, 249–264.

Condie, K.C., Des Marais, Abbott, 2000. Geologic evidence for amantle superplume event at 1.9 Ga. Geochem. Geophys. Geosyst.1 (Paper number 2000GC000095).

Courtillot, V., Davaille, A., Besse, J., Stock, J., 2003. Three distincttypes of hotspots in the Earth's mantle. Earth Planet. Sci. Lett. 205,295–308.

Cserepes, L., Christensen, U.R., Ribe, N.M., 2000. Geoid heightversus topography for a plume model of the Hawaiian swell. EarthPlanet. Sci. Lett. 178, 29–38.

Cundall, P.A., 1989. Numerical experiments on localization infrictional materials. Ing. Arch. 59, 148–159.

d'Acremont, E., Leroy, S., Burov, E.B., 2003. Numerical modelling ofa mantle plume: the plume head–lithosphere interaction in theformation of an oceanic large igneous province. Earth Planet. Sci.Lett. 206, 379–396.

Davies, G.F., 1988. Ocean bathymetry and mantle convection. 1.Large-scale flow and hotspots. J. Geophys. Res. 93, 10467–10480.

Davies, G.F., 1992. Temporal variation of the Hawaiian plume flux.Earth Planet. Sci. Lett. 113, 277–286.

Davies, G.F., 1994. Thermomechanical erosion of the lithosphere bymantle plumes. J. Geophys. Res. 99, 15709–15722.

Davies, G.F., 1995. Penetration of plates and plumes through themantle transition zone. Earth Planet. Sci. Lett. 133, 507–516.

Davis, P.M., Slack, P.D., 2002. The uppermost mantle beneath theKenya dome and relation to melting, rifting and uplift in EastAfrica. Geophys. Res. Lett. 29. doi:10.1029/2001GL013676.

Doin, M.-P., Fleitout, L., Christensen, U., 1997. Mantle convectionand stability of depleted and undepleted continental lithosphere.J. Geophys. Res. 102, 2771–2787.

Ebinger, C.J., Sleep, N.H., 1998. Cenozoic magmatism throughoutEast Africa resulting from impact of a single plume. Nature 395,1788–1791.

Ernst, R.E., Buchan, K.L., 2003. Recognizing mantle plumes in thegeological record. Annu. Rev. Earth Planet. Sci. 31, 469–523.

Farnetani, C.G., Richards,M.A., Ghiorso,M.S., 1996. Petrologicalmodelsof magma evolution and deep crustal structure beneath hotspots andflood basalt provinces. Earth Planet. Sci. Lett. 143, 81–94.

Foulger, G., 2002. Plumes, or plate tectonic processes ?Astron. Geophys.43, 619–623.

Gerbault, M., Poliakov, A.N.B., Daignières, M., 1998. Prediction offaulting from the theories of elasticity and plasticity; what are thelimits? J. Struct. Geol. 20, 301–320.

Goes, S., Cammarano, F., Hansen, U., 2004. Synthetic seismic signatureof thermal mantle plumes. Earth Planet. Sci. Lett. 218, 403–419.

Groves, D.I., Barley, M.E., 1994. In: Condie, K.C. (Ed.), Archeanmineralization, in: Archean Crustal Evolution. Elsevier, Amster-dam. 528pp.

Guillou-Frottier, L., Burov, E., Nehlig, P., Wyns, R., in press,Deciphering plume–lithosphere interactions beneath Europefrom topographic signatures, Global and Planetary Change,Spec. vol. on Topography of Europe.

Hales, T.C., Abt, D.L., Humphreys, E.D., Roering, J.J., 2005. Alithospheric instability origin for Columbia River flood basalts andWallowamountains uplift in northeast Oregon. Nature 438, 842–845.

Ingle, S., Coffin, M.F., 2004. Impact origin for the greater Ontong Javaplateau? Earth Planet. Sci. Lett. 218, 123–134.

Isley, A.E., Abbott, D.H., 1999. Plume-related mafic volcanism andthe deposition of banded iron formation. J. Geophys. Res. 104,15,461–15,477.

Jull, M., Kelemen, P.B., 2001. On the conditions for lower crustalconvective instability. J. Geophys. Res. 106, 6423–6446.

King, S.D., 2005. Archean cratons and mantle dynamics. Earth Planet.Sci. Lett. 234, 1–14.

Kirby, S.H., Kronenberg, A.K., 1987. Rheology of the lithosphere:selected topics. Rev. Geophys. 25, 1219–1244.

Kohlstedt, D.L., Evans, B., Mackwell, S.J., 1995. Strength of thelithosphere: constraints imposed by laboratory experiments.J. Geophys. Res. 100, 17587–17602.

Kukkonen, I.T., Peltonen, P., 1999. Xenolith-controlled geotherm forthe central Fennoscandian Shield: implications for lithosphere–astenosphere relations. Tectonophysics 304, 301–315.

Kusky, T.M., Polat, A., 1999. Growth of granite–greenstone terranes atconvergent margins, and stabilization of Archean cratons.Tectonophysics 305, 43–73.

Kutina, J., Pei, R., Heyl, A.V., 2003. The role of deep lithospheric structurein the genesis and distribution of giant and supergiant concentrationsof metals in the crust. Global Tecton. Metallog. 8, 9–49.

http://dx.doi.org/10.1130/1052-2006)016<4:TLTSOC

http://dx.doi.org/10.1130/1052-2006)016<4:TLTSOC

http://dx.doi.org/10.1029/2002JB001904

http://dx.doi.org/10.1029/2001GL013676


Le Goff, E., Deschamps, Y., Muhongo, S., Cocherie, A., Milesi, J.P.,Pinna, P., Msechu, M., Mshihili, A., 2004. The Tanzanian “RubyBelt”: structural, petrological and geochronological constraintswithin the pan-African orogeny — examples from the Morogoroand Mahenge districts. Colloquium of African Geology, CAG20,Orléans, France, pp. 2–7. June.

Lin, S-C., Kuo, B.Y., Chiao, L.-.Y., van Keken, P.E., 2005. Thermalplume models and melt generation in East Africa: a dynamicmodeling approach. Earth Planet. Sci. Lett. 237, 175–192.

Manglik, A., Christensen, U.R., 1997. Effect of mantle depletionbuoyancy on 2flow and melting beneath a stationary plate.J. Geophys Res. 102, 5019–5028.

Manglik, A., Christensen, U.R., 2006. Effect of lithospheric root ondecompression melting in plume–lithosphere interaction models.Geophys. J. Int. 164, 259–270.

Mareschal, J.-C., Jaupart, C., 2004. Variations of surface heat flow andlithospheric thermal structure beneath the North American craton.Earth Planet. Sci. Lett. 223, 65–77.

McCullogh, M.T., Benett, V.C., 1994. Progressive growth of theEarth’s continental crust and depleted mantle: geochemicalconstraints. Geochim. Cosmochim. Acta 58, 4717–4738.

Molnar, P., Jones, G.H., 2004. A test of laboratory based rheologicalparameters of olivine from an analysis of late Cenozoic convectiveremoval of mantle lithosphere beneath the Sierra Nevada,California, USA. Geophys. J. Int. 156, 555–564.

Monnereau, M., Rabinowicz, M., Arquis, E., 1993. Mechanicalerosion and reheating of the lithosphere: a numerical model forhotspot swells. J. Geophys. Res. 98, 809–823.

Montelli, R., Nolet, G., Dahlen, F.A., Masters, G., Engdahl, E.R.,Hung, S.H., 2004. Finite-frequency tomography reveals a varietyof plume in the mantle. Science 303, 338–343.

Moreira, M., Valbracht, P.J., Allègre, Staudacher T., 1996. Rare gassystematics in Red Sea ridge basalts. Geophys. Res. Lett. 23,182453–182456.

Neubauer, F., Lips, A., Kouzmanov, K., Lexa, J., Ivascanu, P., 2005.Subduction, slab detachment and mineralization: the Neogene inthe Apuseni mountains and Carpathians. Ore Geol. Rev. 27, 13–44.

Nyblade, A.A., Sleep, N.H., 2003. Long lasting epeirogenic upliftfrom mantle plumes and the origin of the Southern African Plateau.Geochem. Geophys. Geosyst. 4 (2003GC000573).

Olson, P., 1990. Hotspots, swells and mantle plumes. In: Ryan, M.P.(Ed.), Magma transport and storage. Wiley, Chichester, pp. 33–51.

O'Neill, C.J., Moresi, L., Jaques, A.L., 2005. Geodynamic controls ondiamond deposits: implications for Australian occurrences.Tectonophysics 404, 217–236.

Parsons, B., Sclater, J.G., 1977. An analysis of the variation of oceanfloor bathymetry and heat flow with age. J. Geophys. Res. 82,803–827.

Pascal, C., van Wijk, J.W., Cloetingh, S.AP.L., Davies, G.R., 2002.Effect of lithosphere thickness heterogeneities in controlling riftlocalization: numerical modeling of the Oslo graben. Geophys.Res. Lett. 29, 1355.

Perry, H.C., Jaupart, C., 2004. Particular mantle dynamics induced bycontinental roots. Eos Trans. AGU 85 (17) (Jt. Assem. Suppl.,Abstract T41C-04).

Pik, R., Marty, R., Hilton, D.R., 2006. How many mantle plumes inAfrica? The geochemical point of view. Chem. Geol. 226, 100–114.

Pirajno, F., 2000. Ore Deposits and Mantle Plumes. Kluwer Acad,Dordrecht, Neth. 556 pp.

Pirajno, F., 2004. Oceanic plateau accretion onto the northwesternmargin of the Yilgarn craton, western Australia: implications for amantle plume event at ca. 2.0 Ga. J. Geodyn. 37, 205–231.

Poliakov, A.N.B., Cundall, P., Podladchilov, Y., Lyakhovsky, V., 1993.An explicit inertial method for the simulation of visco-elastic flow:an evaluation of elastic effects on diapiric flow in two-or three-layers models. In: Stone, D.B., Runcorn, S.K. (Eds.), Flow andCreep in the Solar System: Observations, Modelling and Theory,Dynamic Modelling and Flow in the Earth and Planet Series,pp. 175–195.

PoudjomDjomani, Y.H., O'Reilly, S.Y., Griffin,W.L., Morgan, P., 2001.The density structure of subcontinental lithosphere: constraints ondelamination models. Earth Planet. Sci. Lett. 184, 605–621.

Ranalli, G., 1995. Rheology of the Earth, 2nd ed. Chapman and Hall,London. 413pp.

Ribe, N.M., Christensen, U.R., 1994. Three-dimensional modeling ofplume–lithosphere interaction. J. Geophys. Res. 99, 669–682.

Ritter, J.R.R., 2005. Small-scale mantle plumes: imaging andgeodynamic aspects, in Perspectives in Modern Geology. In:Wenzel, F. (Ed.), Lectures Notes in Earth Sciences, vol. 105.Springer Verlag, Heidelberg, pp. 69–94.

Romanowicz, B., Gung, Y.C., 2002. Superplumes from the core-mantleboundary to the base of the lithosphere. Science 296, 513–516.

Radulescu, F., 1988. Seismic models of the crustal structure in Romania.Rev. Roum. Géol. Géophys. Géogr., Sér. Géophys. 32, 3–17.

Saleeby, J., Ducea, M., Clemens-Knott, D., 2003. Production and lossof high-density batholitic root, southern Sierra Nevada, California.Tectonics 22, 1064. doi:10.1029/2002TC001374.

Schubert, G., Turcotte, D.L., Olson, P., 2001. Mantle Convection in theEarth and Planets. Cambridge Univ. Press. 956pp.

Sebai, A., Stutzmann, E., Montagner, J.-P., Sicilia, D., Beucler, E., 2006.Anisotropic structure of the African upper mantle fromRayleigh andLove wave tomography. Phys. Earth Planet. Int. 155, 48–62.

Sheth, H.C., 1999. Flood basalts and large igneous provinces from deepmantle plumes: fact, fiction, and fallacy. Tectonophysics 311, 1–29.

Silver, P.G., Behn, M.D., Kelley, K., Schmitz, M., Savage, B., 2006.Understanding cratonic flood basalts. Earth Planet. Sci Lett. 245,190–201.

Sleep, N., 1997. Lateral flow and ponding of starting plume material.J. Geophys. Res. 102, 10001–10012.

Sleep, N.H., Ebinger, C.J., Kendall, J.-M., 2002. Defection of mantleplume material by cratonic keels, in The Early Earth: Physical,Chemical and Biological Development. In: Fowler, C.M.R., Ebinger,C.J., Hawkesworth, C.J. (Eds.), Geological Society London, Spec.Pub., vol. 199, pp. 135–150.

Sleep, N., 2003a. Survival of Archean cratonal lithosphere.J. Geophys. Res. 108, 2302. doi:10.1029/2001JB000169.

Sleep, N.H., 2003b. Geodynamic implications of xenolith geotherms.Geochem. Geophys. Geosyst. 4 (2003GC000511).

Sleep, N.H., 2005. Evolution of continental lithosphere. Annu. Rev.Earth Planet. Sci. 33, 369–393.

Smith, R.B., 1975. Unified Theory of the Onset of Folding, Boudinageand Mullion Structure. Geol. Soc. Amer. Bull. 88, 1601–1609.

Smithies, R.H., Champion, D.C., Cassidy, K.F., 2003. Formation ofEarth's early Archean continental crust. Precambrian Res. 127,89–101.

Solomatov, V.S., 1995. Scaling of temperature- and stress-dependentviscosity convection. Phys. Fluids 7, 266–274.

Solomatov, V.S., Moresi, L.N., 1995. Stagnant lid convection onVenus. J. Geophys. Res. 101, 4737–4753.

Solomatov, V.S., Moresi, L.N., 2000. Scaling of time-dependent stagnantlid convection; application to small-scale convection on Earth andother terrestrial planets. J. Geophys. Res. 105, 21795–21817.

Tackley, P.J., 2000. Mantle convection and plate tectonics; toward anintegrated physical and chemical theory. Science 288, 2002–2007.

http://dx.doi.org/10.1029/2002TC001374

http://dx.doi.org/10.1029/2001JB000169


Trompert, R., Hansen, U., 1998. Mantle convection simulations withrheologies that generate plate-like behaviour. Nature 395, 686–689.

Turcotte, D.L., Schubert, G., 2002. Geodynamics, applications ofContinuum Physics to Geological Problems, Second edition. Cam-bridge Univ. Press. 456 pp.

Watts, A.B., Burov, E., 2003. Lithospheric strength and its relationshipto the elastic and seismogenic layer thickness. Earth Planet. Sci.Lett. 213, 113–131.

Weeraratne, D.S., Forsyth, D.W., Fischer, K.M., Nyblade, A.A., 2003.Evidence for an upper mantle plume beneath the Tanzanian craton

from Rayleigh wave tomography. J. Geophys. Res. 108 (B9), 2427.doi:10.1029/2002JB002273.

Weinberg, R.F., Podladchikov, Y., 1994. Diapiric ascent of magmasthrough power law crust andmantle. J. Geophys. Res. 99, 9543–9559.

Wessel, P., Keating, B.H., 1994. Temporal variations of flexuraldeformation in Hawaii. J. Geophys. Res. 99, 2747–2756.

Wortel, M.J.R., Spakman, W., 2000. Subduction and slab detachmentin theMediterranean–Carpathian region. Science 290, 1910–1917.

http://dx.doi.org/10.1029/2002JB002273

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