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Tectonophysics, 199 1991) 343-314

Elsevier Science publishers B.V., Amsterdam

343

A two level concept of plate tectonics: application to geod~a~cs

L.I. Lobkovsky a and V.I. Kerchman b

y

nstitute of Oceanology The U.S.S.R. Academy

of Scien ces

Moscow ii 7238, USSR

h The Interuniversity Computer Center, Kishinev 279003, USSR

(Received August 15.1989; revised version accepted July 15.1990)

ABSTRACT

Lobkovsky, L.I. and Kerchman, V.I., 1991. A two-level concept of plate tectonics: application to geodynamics. In: L.P.

Zonenshain (Editor), The Achievements of Plate Tectonics in the U.S.S.R. Tectonophysics, 199: 343-374.

A twolevel plate tectonics concept is developed on the basis of data on lithosphere rbeological stratification. This

approach differentiates between crust and subcrust plate ensembles separated by a lower-crust viscoplastic asthenolayer.

Similarly to classical plate tectonics, three types of boundaries are distinguished in the lower layer which do not always

coincide with crust-plate boundaries (especially for continents). Applications of this concept to geodynamics are considered,

and a corresponding quantitative analysis for several important

processes is carried out.

A

quantitative mode1 of mountain

formation and collision-plateau origins is proposed.

Also, a geodynamic model of the evolution of passive margins, taking into

account a lower-crust viscous flow, is considered and its geological consequences are discussed. A mechanism

of

rifting, taking

into consideration rheological lithospheric layering and its vertical movements caused by extension, is developed. Both a

qualitative scheme and quantitative analysis of the slow evolution of intracraton structures of “shield-basin” type, taking into

account erosion and sedimentation processes, are worked through. Also, historical aspects of plate tectonics are discussed from

the point of view of the proposed concept.

Intmduetion

The orthodox theory of plate tectonics seems to

have certain restrictions on its application. Thus,

tectonic processes of a regional scale, that is of

several hundreds of kilometres, cannot be de-

scribed sufficiently well by standard plate-tectonic

models. This is of special concern for the conti-

nents (Molar, 1988; Lobkovsky, 1988a) because,

when one analyses regional processes, those inho-

mogeneities (both horizontal and vertical) and dis-

tributions of intraplate strains which have not

been considered on a global scale (since they have

virtually been averaged) become the main objects

of the research. Since many researchers have been

guided by the idea of scale restrictions imposed on

the plate-tectonic processes, they have put this

above the quantitative study of plate non-rigidity

and patterns of intraplate strains and stresses

(Molnar and Tapponier, 1978; England and Mc-

Kenzie, 1982; Vilotte et al., 1982; Cloetingh et al.,

1984; England et al., 1985; Khain, 1986; Bruhn,

1987; Kirby and Kronenberg, 1987).

As for geomechanics, the analysis of the prob-

lem mentioned is reduced to the description of the

rheological behaviour of lithospheric rocks under

the real P-T conditions and for the various regi-

mes of tectonic strain.

In geodynamics, an “effective strength)* is usu-

ally used as a generalized rheological characteristic

of the lithosphere (Ranalli and Murphy, 1987;

Kirby and Kronenberg, 1987); this may be inter-

preted in different ways, depending upon the

P-T

conditions and mechanisms of rock strain. For the

domain of quasi-elastic strains and brittle fractur-

ing, this characteristic coincides with the defini-

tion of a material strength as used in mechanics

(Byerlie, 1968; Sibson, 1974; Brace and Kohlstedt,

1980). For the ductile (non-linear viscous) flow of

the medium, the notion of “creep strength” has

been introduced (Ranalli and Murphy, 1987;

Kirby, 1983).

~-1951/91/ 03.50 6 1991 - Elsevier Scienw Publishers

B.V. All rights reserved

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344

L..l. LOBKOVSKY AND ‘4.1. KtRtHMAN

a

Temperature

rengh

b Tempera ure

Fig. 1. Model strength profiles (solid lines): (a) for normal

continental lithosphere with a 15 km thick granitic upper crust

and a 25 km thick mafic lower crust; (b) for normal oceanic

lithosphere with a 5 km thick basaltic upper crust and a 2 km

thick serpentinitic lower crust. Rashed lines denote the geo-

therms for the two cases. The dotted line denotes the strength

envelope. The strain fate is lO_” s-‘.

Ahhough each particular profile of an “effec-

tive strength” compiled from the data of labora-

tory tests, depends upon the assumed composition

of the Earth’s crust and upper mantle, and upon

temperature distributions, deformation rate, water

saturation of the media, and so on, all these pro-

files plainly indicate the main feature of the rheo-

logical s~at~~ca~on of the exosphere; th8t is, the

occurrence of an extremely low strength (viscosity)

of the medium in the lower crust in contrast to the

stronger and more brittle layers of the upper crust

and the subcrustal part of the lithosphere.

As an example, in Fig. la we illustrate a typical

profile of the generalized strength of the litho-

sphere, which is composed of an upper “granitic”

crust (Z) - 15 km thick, a lower “mafic” crust

(II) - 25 km thick and an underlying olivine

mantle (III). This plot has been drawn from a

typical continental geotherm that corresponds to

the average heat flow on the surface; that is, about

50 mW/m* {Morgan and Sass, 1984), in accor-

dance with experimental rheological data (Byerlie,

1968; Brace and Kohlstedt, 1980; Kirby, 1983;

Ranalli and Murphy, 1987), the deformation rate

being constant at C? 1O-‘5 s-j. The envelope line

shows the resistance of the medium to brittle

fracturing (Byerlie, 1968), and corresponds to the

generalized strength of almost the entire upper

crustal layer (I), a small part of a “cold” lower

crust (II’) and a quasi-rigid core of the subcrustal

mantle (III ‘). The main part of the lower crust

(II) - that is, the crustal asthenolayer - shows

ductile properties which may be described by a

creep law (Kriby, 1983). The lower quasi-viscous

lithosphere (III) is a tradition to the mantle

asthenosphere.

The profile of the generalized strength for a

normal oceanic lithosphere has a similar character

(Fig. lb). Here, the lower serpentinite layer of the

oceanic crust is analogous to the crustal

asthenolayer (Raleigh and Paterson, 1965;

Lobkovsky et al., 1986). Note that each particular

profile of a generalized strength (limiting shear

stress) varies considerably from region to region as

the heat regime and tectonic strain of the media

vary (for instance, as the rates of various processes

change in different layers).

The present-day concepts described concerning

the rheological stratification of geological media,

distinguishing quasi-rigid and ductile layers in the

crust, are analogous to the traditional geophysical

concept of the existence of the lithosphere and

asthenosphere in the upper mantle, the concept on

which the orthodox theory of plate tectonics rests.

The understanding of the geodynamic essence of

this analogue has resulted in the fo~ula~on by

Lobkovsky (1987a,b; 1988a,b) of the pfincipaily

new concept of two-level plate tectonics. The main

idea of this concept lies in the recognition of two

main levels (crustal and subcrustal) and in the

introduction of different scales in the conventional

theory of plate tectonics. When global horizontal

motions of several thousands of kilometres (a typi-

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A TWO-LEVEL CONCEPT OF PLATE TECTONICS

345

cal scale) are described, it is the lower subcnustal

(upper-mantle) level of the system that works,

whereas the processes in the upper-crustal level

are subordinate. However, to describe adequately

the majority of the regiond te&m.ic processes, the

typical scales of which are only some hundreds of

kilometres and even less, it is necessary to turn to

the upper-crustal level of the geodynamic model

considered. Within this level, the upper brittle

crustal sublayer is dissected into individual micro-

plates and geoblocks, their dimensions being from

several hundreds to several tens of kilometres.

Such crustal blocks are capable of horizontal dis-

placements over the underlying crustal astheno-

layer relative to the mantle part of the lithosphere.

If one ignores these displacements - in the case

of relatively strong coupling of crustal blocks with

the mantle and at lesser differential stresses - it

is possible to hold to the standard plate-tectonic

constructions. If differential stresses are suffi-

ciently great, as in the case of collision and con-

~nent~-~ft~g zones? the upper-crnstal layer may

move and be deformed independently, to a con-

siderable extent, on the subcrustal layer of the

system due to the developed ductile flow of the

lower crust. One such situation is illustrated in

Fig. 2. A general two-level scheme for the continental c&&ion

of India and Eurasia (after Lobkovsky, 19 8a; see text

for

explanation). I - Upper brittle crust; 2 = lower ductile crust;

3 = subcrustal Lithosphere; 4 = mantle asthenosphere; 5 =

thrust-type (convergent) crusti boundaries; 6 = strikaslip-type

(transform) txustal bounties; 7 = she-ax- elo@ity distributio+x

ill a cross-section of lowex crust; 8 = direetioll of sllbcrustal

lithosphere movement; 9 = subduction zone.

Fig. 2: the continental collision of India and

Eurasia (Lobkovsky, 1988a,b). This shows the

mosaic of the upper-crustal microplates, the rela-

tive

displacements of which are determined both

by the indentation of the Hindustan to Eurasia

boundary (at crustal level) and by the dragging of

crustal blocks by the mantle part of the litho-

sphere7 which is slipping beneath the crust. It is

also determined by the flow of the lower crust

caused by the above-mentioned process. As a re-

sult of the viscous flow of the lower crust, its

material may be forced (injected) into the vicinity

of the suture zone, thus causing thickening of the

lower crust and formation of the roots of the

mount~s and isostatic uplifts of the territory.

The present paper not only gives a quantitative

analysis of the collision process (according to the

scheme mentioned above) (Lobkovsky, 1988a,b;

1990; Khain and Lobkovsky, 3990), but also con-

siders how a two-level concept of plate tectonics

can be applied to some geodynarnic problems. It

describes the evolution of passive continental

margins and the formation of rift zones at their

rear (Lobkovsky, 1989; Lobkovsky and Khain,

1989); it gives the general scheme of continental

rifting at two levels and analyses the geodynamic

behaviour of the crust within cratons with regard

to the ‘~erosion-sedimen~tion-rnet~o~~srn-

flow in the lower crust” rnate~~-circula~on cycle

(Lobkovsky, 1989); finally, it considers some his-

torical aspects of two-level evolution of the litho-

sphere.

Geological-geophysical grounds and general ideas

of a two-level, plate-tectonic concept

First, we briefly consider the factual data that

provide evidence for a two-layered (in first ap-

proximation) rheological structure of the Earths

crust. The results of deep seismic studies, by pro-

jects such as COCORP (Allmendinger et al., 1987),

ECDRS (Choukroune and Garridq 1989) and

EUGENO-S (h4eissner et al., 1987), have allowed

their authors to determine a fine-layered structure

of the lower continental crust. We believe that

such a stratification probabIy originates from the

horizontal flow of ductile mater&I in the Iower-

crustal asthenolayer.

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346

L 1. LOBKOVSKY AND V.I. KEKCHMAN

Considerable progress has been achieved in in-

strumental seismology in the past decade. As a

result, the accuracy with earthquake hypocentres

can be determined has increased by one order of

magnitude (Chen and Molnar, 1983). From the

studies that have been carried out, it turns out that

the hypocentres of continental earthquakes are

predominantly concentrated in the uppermost 20

km of the crust. The lower layer of the continental

crust, which is 15-20 km thick, is virtually

aseismic; earthquake foci appear again only in the

subcrustal horizons of the upper mantle in active

tectonic regions of the Earth. Thus, the lower

Continental crust can be considered to be an

aseismic layer (Chen and Molnar, 1983; Meissner,

1985; Jackson, 1987) and this verifies its astheno-

spheric nature.

A comparison of the observed distribution pat-

terns of earthquake hypocentres in the lithosphere

and the temperature field in the crust and upper

mantle, made for various regions of the Earth, has

shown that seismicity within the continental crust

is limited by about the 3SO” C isotherm, whereas

seismicity in the upper mantle is controlled by the

700 o C isotherm (Chen and Molnar, 1983; Wiens

and Stein, 1983). The temperatures mentioned

correspond to the transition from a quasi-brittle to

ductile deformation regime for the geomaterial of

the crust (at

- 350 o C) and mantle (at - 700 0 C)

obtained in laboratory tests (DeRito et al., 1986;

Jackson, 1987).

The observed pattern of the distribution of

seismicity serves as independent evidence for the

rheological stratification of the Earth’s crust into

the lithosphere and asthenosphere. Note that the

correlation between the heat regime of the litho-

sphere and the character of seismicity in one re-

gion or another allows us to draw conclusions on

the lateral variability of the properties of the

crustal asthenosphere; and on the changes in its

thickness, effective viscosity, etc. in particular.

It follows from a proposed pattern of plate

tectonics (Lobkovsky, 1988a) that the process of

isostatic equilibrium should occur at two levels at

least; namely, at the crustal and hthosphe~c levels.

As for the lateral scales of isostatic compensation

of several tens (or perhaps hundreds) of kilo-

metres, the main role is played by the crustal

asthenosphere, this being verified by analysis of

local isostasy on the continents (McAdoo, 1985).

When the scale is increased to several hundreds of

kilometres, the mantle asthenosphere acquires an

important meaning. The given supposition of a

two-level isostatic compensation is verified by

analysing the isostasy of each particular region.

On the basis of such an analysis, it is possible to

show which part of the isostatic compensation is

taken up by the crust (it is usually 70%) and how

much remains in the mantle lithosphere and

asthenosphere (Artemjev and Kaban, 1987).

The rheologieal statification of the crust and

lithosphere is verified by numerous geological data

concerning their tectonic layering in each particu-

lar region &nipper and Ruzhentsev, 1977; Peive

et al., 1983; Ranalli and Murphy, 1987). Extensive

granite-gneiss allochthons with displacement am-

plitudes of hundreds of kilometres, in the Alpine

belt, Appalachian Mountains and other fold belts,

are the direct geological evidence for the occur-

rence of the lower-crustal asthenolayer (Cook et

al., 1979; Hsii, 1979; Peive et al., 1983).

We shall now formulate some general state-

ments initially proposed by Lobkovsky (1987a,

1988a) which complement the plate-tectonic ap-

proach to a two-level pattern. At each level we

distinguish a geodynamic system with its own

ensemble of plates (~croplates) which, generally

speaking, do not coincide. This is the principal

difference in the proposed model from the ortho-

dox plate-tectonic concept. The system of plates

and microplates of the upper-crustal “deck” coin-

cides with the orthodox plate-tectonic concept (Le

Pichon et al., 1973; Zonenshain and Savostin,

1979). The tectonic system of the lower (mantle>

floor of the lithosphere can be explained by the

occurrence of a quasi-rigid subcrustal layer III’

(Fig. la) dissected into large plates, the size of

which is on the scale of thousands of kilometres.

They are dissected by boundaries of three types, in

a similar manner to the orthodox concept, namely

divergent, convergent and transform.

If these plates and the boundaries between

them coincide (the character of motion included),

the plates may be considered as a monolith and

their boundaries as a common one. In this case,

we can use constructions of conventional plate

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A TWO-LEVEL CONCEPT OF PLATE TECTONICS

341

tectonics: this is true for the most part of the

oceanic lithosphere, since in the oceans, the

governing behaviour is that of the lower subcrustal

lithosphere with a quasi-rigid “core” III’ (- 20-

40 km thick), which is considerably thicker and

stronger compared to the brittle overlying basalt

crust (3-5 km).

If the plates (and their boundaries) of the upper

and lower floors do not ocincide, a more complex

two-level model of tectonics originates, where

various motions of crustal plates (blocks) are pos-

sible relative to each other and relative to the

mantle part of the lithosphere (due to the flow of

a ductile lower-crustal layer). The collision scheme

for such motions is shown in Fig. 2.

A two-level tectonic structure is most distinct

in the continental regions where, in contrast to

oceans, the total strength and thickness of the

upper-crustal brittle layer and the quasi-rigid sub-

crustal core of the lithosphere are comparable. In

this connection, a problem arises as to how to

single out individual plates of the lower level

under the continents, the boundaries of which do

not coincide with the traditional ones.

As for the divergent plate boundaries of the

lower level, it seems most logical to attribute them

to the extensive linear zones of basaltic (alkaline

and tholeiitic) magmatism (Milanovsky, 1975),

which usually occurs long before the rifting in the

system at the upper level and which produces a

complete splitting up of the plates in the tradi-

tional sense. So, for instance, it is known

(Razvalyaev, 1984) that during the past 800 Ma

intrusions of alkaline basalts happened many times

(at around 770-450, 290, 185, 120 and 80 Ma)

along the East African rift system. The periods of

basaltic magmatism alternated with the periods of

secondary alkaline-granitic magmatism (570-450,

185 and 50 Ma) (Razvalyaev, 1984). In addition to

the East African rift system, a spatial coincidence

of the linear zones of alkaline magmatism with rift

structures has been revealed in the Gardar prov-

ince (Greenland), the St. Lawrence rift, the Baikal

rift zone and a number of others. In many cases,

such magmatism occurred before not only the

Cenozoic but also the Mesozoic rifting.

In the framework of a two-level plate-tectonic

concept, it is natural to consider such prolonged

linear zones of alkaline magmatism as divergent

plate boundaries in the lower level. It is difficult

to trace these boundaries through the entire area

of the continents, as the magmatism attributed to

them is not always exposed on the Earth’s surface;

for example, due to the compression regime in the

upper level of the system. As an example of a

divergent plate boundary in the lower level, we

can draw a linear belt extending from the North

Sea to the Lower and Upper Rhine grabens, to the

grabens of the Seine and Rh&e, then continuing

through the Mediterranean (Tunisia strait, grabens

in Pantelleria and Malta), further on into Africa

to the Gulf of Guinea, and along the Cameroon

line into the South Atlantic. The East African rift

system and the East Asian system from Anadyr

Bay to Tungting Hu Lake in Southern China may

be attributed to the same category of divergent

plate boundaries (Lobkovsky and Khain, 1989).

As in orthodox plate tectonics, the divergent seg-

ments of plate boundaries in the lower level might

be connected by lower-layer transform segments,

although there is little direct evidence of this.

The convergent plate boundaries of the lower

level usually occur in collision belts where in-

tracontinental subduction processes develop. Ex-

amples are the Alpine-Himalayas zone of sub-

crustal subduction (Fig. 2) and possible in-

tracontinental subduction beneath the Rocky

Mountains during the Laramide orogeny. As a

rule, lower-level subduction is not marked by mid-

dle- and deep-foci earthquakes, although rare ex-

ceptions occur in individual segments with a nar-

row frontal part (Calabria, Hindu Kush, etc.). The

aseismic behaviour of such subduction zones can

be explained first by the fact that cold hydrated

crust is not subducted, as in the case of the

oceanic plates and, second, by the considerable

frictional heating of the subducted lithosphere (this

mechanism will be considered in detail below).

The aseismic regions of convergent plate

boundaries in the lower level can be revealed using

seismic tomography data. The inclined high-veloc-

ity layers and the Q-factor correspond to such

zones (Spakman, 1986). Post-collisional granitic

magmatism generated by frictional heating may

serve as a geological indicator of collisional sub-

duction (Debon et al., 1986; Lobkovsky, 1988a;

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34x

1, I. LOBKOVSKY AND V.I. KEKCHMAN

Khain and Lobkovsky, 1990; Kerchman and

Lobkovsky, 1990b).

The two-level concept of plate tectonics poses a

general problem, how to recognize the plate

boundaries of the lower layer using the mass of

complex geological and geophysical data availa-

ble, including gravity field, heat flow and geo-

chemical information (in particular, the zones of

mantle helium).

The possible structure of divergent plate

boundaries in the lower layer may be predicted by

analogy with the central parts of slowly expanding

mid-oceanic ridges (Lobkovsky, 1989). Thus, they

should show a bilaterally symmetrical inhomo-

geneity of the subcrustal Lithosphere thickness. An

example of a divergent plate boundary in the

lower level may be the asthenospheric bulge under

the Sayan-Baikal arch uplift and its northeast and

southwest extensions (Rogozhina and Kozhevni-

kov, 1979; Logachev and Zorin, 1988). Another

example of a divergent boundary in the subcrustal

layer is an extension of an anomalous mantle

bulge from the Rio-Grande rift into the inner

parts of the North American continent. In con-

trast to the oceanic divergent boundaries, which

are characterized by extension only, crustal struc-

tures of both extension (Baikal rift) and compres-

sion (the Mongolian Altai, Han-hai, the Gobi AI-

tai and Eastern Sayan) can be located over the

divergent boundaries of the lower level. Within the

African continent, the latter may be represented

by a considerably thinner subcrustal lithosphere

(Fairhead and Reeves, 1977; Kazmin, 1987).

As has been noted above, the divergent

boundaries of the lithospheric lower layer are part

of the global system of plate boundaries. In par-

ticular, they continue from the continental to the

oceanic regions and vice versa. The continuation

of the East Pacific Rise under the crustal layer of

North America in the Basin and Range Province

and the ‘“Cameroon line” that continues the is-

lands of the South Atlantic into the lower-level

extension boundary under West Africa are typical

of this feature (Lobkovsky, 1989).

Plate bound~es of the subcrustal layer seem to

be as stable in time as the principal boundaries of

large oceanic plates, i.e. mid-oceanic ridges and

subduction zones. They can change due to changes

in the mantle convection regime, and also due to

large-scale geodynamie processes which occur at

both levels of the lithosphere (collision, for in-

stance).

A two-level plate-tectonic model allows us to

incorporate the principal mechanisms of rift prop-

agation (Co~tillot, 1982; Martin, 1984). In par-

ticular, the propagation of rifts from oceanic basins

to the upper-crustal layer of continents may be

due to a pre-existing divergent boundary in their

lower layer.

This new view of the plate tectonics of the

lithosphe~c lower layer allows us to consider hot

spots within the continents as zones where mag-

matic activity is localized at triple junctions of

divergent and transform lower-level boundaries,

where the quasi-rigid layer of subcrustal litho-

sphere is completely absent. Examples of such

isolated points are Yellowstone Park in the U.S.A.,

and the volcanic areas of Tibesti, Darfur and

Achaggar in Central and Western Africa (Lobkov-

sky, 1989). The tectonic evolution of the Earth in

the earliest historical stages (Archean and Pro-

terozoic) is of great interest. At present, the majot-

ity of researchers believe that plate-tectonic

processes began on the Earth as late as in the

Proterozoic (Khain and Mikhailov, 1985; Khain

and Bozhko, 1988). According to the concept de-

scribed above, the plate-tectonic regime may begin

when a general cooling of the Earth occurs, the

intensity of mantle convection decreases and, as a

result, a continuous subcrustal layer is formed in

the lithosphere. The quasi-rigid core of the lower

lithospheric layer prevents chaotic motion and

decouples the previously formed massifs of sialic

protocrust from the unsteady convective flow of

the mantle. Such “capturing” of individual proto-

crustal blocks by the lithospheric lower layer may

explain the formation of stable cratons and con-

tinental cores. The mosaic pattern of Archean

permobile tectonics may be explained by the lack

of a quasi-rigid frame in the lower level. Similar

conclusions can be reached about the oceanic

proto-crust. This concept of the evolution of the

Earth’s tectonic regime from a hot chaotic state to

an organized plate-tectonic one by the formation

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A TWO-LEVEL CONCEPr OF PLATE TECTONKS

of a continuous lower-level “skeleton” (frame)

may be applied to the evolution of all terrestrial

planets.

We note that in a two-level plate-tectonic model

the main driving forces are applied to the plates of

the lower level, either by coupling of subcrustal

plates with convective mantle flows or by a “pull-

ing” effect that works on the subducted mantle

parts of these plates or, finally, by spreading of

anomalous mantle and “ridge push” (Forsyth and

Uyeda, 1975; Cloetingh and Wortel, 1985; Seki-

guchi, 1985). The forces driving the upper-crustal

layer are determined by their coupling with the

plates of the lower level; they also may be con-

nected with inhomogeneities of crustal thickness

and density. One other special aspect of the two-

level model is that plate interactions at the crustal

level play a greater role in the geodynamics of

crustal blocks and microplates than the corre-

sponding plate interactions at the lower level.

Mechanical aspects of the two-level plate-tectonic

model

To analyse the mechanical problems which arise

in two-level tectonic processes, we first consider

briefly some aspects of the rheological behaviour

of the continental lithosphere under various con-

ditions and strain regimes. The ductile flow law

derived experimentally for crustal and upper-man-

tle rocks at high temperature is:

t=Ar”exp[-Q/R(T+273)]

(1)

where P is the strain rate, 7 = u1 - a3 is twice the

shear stress, T is the temperature in *C; and A, n

and Q are material constants depending upon the

composition, structure and water content of the

rocks (Kirby, 1983).

In Figs. 3a and b are presented (a) the modified

curves of generalized strength (limiting stress) for

different temperature regimes and (b) the non-ho-

mogeneous strain-rate dist~bution in sublayers of

viscous flow that correspond to real tectonic regi-

mes; for instance, to the slip of a lithospheric

subcrusml layer during intr~ntinental subduc-

tion. This difference from the idealized curves

(solid lines, Figs. 1 and 3) should be taken into

account when analysing each particular geody-

namic situation qu~titatively.

Strmgth

I 60

%

4

80

.

”

b

st rain rote

Fig. 3. Model strength profiles of continental lithosphere with

a 15 km thick granitic upper crust and a 25 km thick mafic

lower crust: (a) for different temperature regimes correspond-

ing to g-therms with heat flows of 45 mW/m’ (solid line) and

60 mW/m* (dashed line); (b) for ho~~~us (solid line) and

non-homogeneous (dashed line) strain-rate distribution in the

lithosphere (dotted line).

We now consider some mathemati~l models in

which the main role is played by ductile flow of

the material in the lower crust due to changes in

its thickness.

Consider a two-dimensional mathematical

model in which the x axis is oriented along the

dominantly horizontal motion and the z axis is

oriented vertically upwards (Fig, 4). We assume

that there is a viscous deformable asthenolayer

(II) of the lower crust, with varying thickness,

which is overlain by an elastic-brittle upper crust

(I) and underlain by a horizontally moving

quasi-rigid plate (III’) of the lower level. For

simplicity, we ignore the bending rigidity of the

upper layer (I) and assume it to be non-deforma-

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350

L.I. LOBKOVSKY

AND V.I. KERCHMAN

Fig. 4. Main layers interacting with each other within the

framework of the two-level plate-tectonic concept (see text for

explanation). I = upper brittle crust; 2 = lower ductile crust;

3 = subcrustal quasi-rigid lithosphere.

ble in a horizontal direction. If we assume a

Newtonian viscosity of layer (II) and local iso-

static compensation of the crust (due to a rather

rapid response of the subcrustal lithosphere), the

equation for the evolution of the lower-crustal

thickness h is (Lobkovsky, 1988a):

- ; (Uh)

where K = (pM -

&~cg/lh.m; P,,., ad pc are

the densities of the mantle and crust, respectively,

g is the acceleration due to gravity, q, is an

effective mean viscosity of the crustal lower layer,

and U is the horizontal velocity of the subcrustal

part of the lithosphere relative to the upper-crustal

layer. We consider this in greater detail in the

Appendix.

When the behaviour of the lower crust follows

a more realistic rheological law (1) with index

n = 3, the equation of the evolution of the layer in

an isothermic approximation for a fixed sub-

crustal basement is (see Appendix):

Here:

/3=bB

Pc(Ph4 - Pc)g

3

PM

1

(3)

(4)

where

B = I

exp[ - Q/R(T + 273)] and

b

is a

normalizing factor that depends on the behaviour

of the upper crust: when this layer is horizontally

rigid (flow “under a cap”), b = l/80; when the

upper layer is fractured and blocks of the upper

crust can move freely in a hotintal direction,

b = l/S (Kerchman, 1990).

In the following sections we shall consider

mathematical models of particular geodynamic

processes, using eqns. (2) and (3).

A two-layer collision model at an intrawutiuental

subduction zone

According to orthodox plate-tectonic theory,

the mountain fold belts of the Earth are formed

by the collision of large continental plates and by

the accretion of smaller crustal blocks (terranes) in

zones of lithospheric convergence (Dewey and

Bird, 1970; Ben-Avraham et al., 1981; Zonenshain,

1986). Various mechanisms for crustal thickening

and uplift of mountain areas have been proposed:

namely, frontal compression, shortening and

warping of the crust and lithosphere (Dewey et al.,

1988); partial subhorizontal subduction of one

continental plate under another to cause a

mechanical doubling of the crust (Powell and

Conaghan, 1973); penetration of rigid continental

“indentors” (Adria, Arabia and Hindustan) into

the elastoplastic (Molnar and Tapponnier, 1978)

or viscoplastic (England and McKenzie, 1982;

England and Houseman, 1986) body of an ad-

jacent plate (Eurasian lithcsphere); and various

types of piling up and delamination of the crust

(Oxburgh, 1972; Bird, 1978; Hsti, 1979).

Although the above-mentioned models each

have their own merits and may accurately describe

certain aspects of collision belts, they do not ex-

plain some important special features of their

structure and evolution. For example, they do not

explain the data on thickening of the continental

crust due to accretion of its lower layer (Giese,

1980; Choukroune and Garrido, 1989), the high

geothermal gradient of collision belts (Artyush-

kov, 1979; Morgan and Sass, 1984), post-collision

granitic magmatism (Debon et al., 1986), or the

lower convergence rate of continental blocks in

the collision zone in comparison with that of the

lithosphere plates which carry them (collision of

Hindustan with Eurasia; Trifonov, 1987). As yet,

there is no explanation for phenomena such as the

absence of mid- and deep-focal earthquakes in the

greater part of the Alpine-Himalayas collision

belt, although the entire geological history of the

Tethys closure (Sborschikov, 1988; Kazmin et al.,

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A TWO-LEVEL CONCEFT OF PLATE TECTONICS

351

1986 and the recent data from seismic tomogra-

phy (Spakrnan, 1986) provide evidence for sub-

duction of the lithosphere under Eurasia.

A new model of continental collision was

worked out by Lobkovsky (1988a,b). It arises from

the assumption of two-level plate tectonics, by

which the above aspects of the structure and

evolution of collision belts are explained (Fig 2).

The traditional analysis of the plate-tectonic

evolution of the Alpine-Himalayas belt is based

on the assumption of a mosaic pattern of litho-

spheric microplates (- 100-200 km thick) and on

the standard kinematic models (McKenzie, 1972,

Zonenshain and Savostin, 1979). In contrast to

this approach, the two-level plate-tectonic model

(Fig. 2) supposes that the microplates that we

observe at the surface are of crustal origin and

may move relative to the mantle part of the litho-

sphere, undergoing large rotations and non-elastic

deformations. Palaeomagnetic data actually give

evidence of remarkable rotations of microplates

and blocks in the process of collision (Klootwijk

et al., 1986).

In accordance with this new approach (Bird,

1978; Lobkovsky, 1988a,b), at an early stage in

the collision between continental plates (preceded

by subduction of the oceanic part of the plate,

which carries a “climbing” continent; Hindustan,

for instance), a sharp reduction in the convergence

velocity of the upper brittle layer of the crust

occurs. The mantle part of the lithosphere con-

tinues to move and to subduct under the forces of

convective “dragging” and “pulling” of the sink-

ing edge of the plate (Figs. 2 and 5).

At the same time, an intensive shear flow

evolves in the lower plastic layer of the crust. This

Fig. 5. Two-lewl model of mountain formation in the wmsse of

continental collision (after Lobkovsky, 1988a; see text for

explanation).

flow embraces one area after another (squeezed

between the upper crust and mantle part of the

lithosphere) as the front of deceleration of the

upper crust advances toward the inner parts of the

overriding subcontinent. This propagation of the

deformation front in the crust can be described by

Elsasser’s equation (Lobkovsky, 1988a). Non-elas-

tic compression of the upper crust in the vicinity

of a suture zone is manifested as a system of

thrusts that form the front of the developing

orogen. The principal agent of crustal thickening

and regional uplift is the pumping of the plastic

material of the lower crustal layer into the vicinity

of the subduction zone by the movement of the

underlying mantle part of the lithosphere (Figs. 2

and 5). A similar model of the detachment of

subducted lithosphere from the crust in a colli-

sional orogen, but with a different crustal thicken-

ing mechanism, is considered in papers by Mat-

tauer (1986) and Dewey et al. (1988).

Consider a quantitative analysis of this ap-

proach (Lobkovsky,

1988a; Kerchman and

Lobkovsky, 1990b). Changes in thickness h of the

lower-crustal layer in the approximation of New-

ton quasi-isomers rheology are described by

eqn. (2). The co-ordinate system moves with the

conventional non-deformed upper crust of the col-

liding continent (India, for instance) (Figs. 2 and

5). The slip rate U =

U, of the subcrustal litho-

sphere relative to the upper crust increases from

zero, starting from the moment of detachment

after the first collision phase (which is manifested

by frontal compression thrusts and a slight thick-

ening of the crust), as shown in Fig. 6b.

The following parameter values were assumed

in order to solve eqn. (2) numerically: pc = 2.8-2.9

g/cm3 (for the lower crust), PM = 3.3-3.35 g/cm3

and nc = (0.3-1.5)

X

lo*’ Pa s; therefore K =

(0.5-3) X lo-’ km-’ yr-‘. Time variations in the

rate of slip have been adopted to correspond with

the conditions of the intracontinental subduction

of the lower layer of the Indian plate (Fig. 6b).

In Fig. 6a are shown calculated curves for the

subsequent thickening of the lower-crustal layer

from the moment when the shear flow started.

Thus, the total crustal thickness (its brittle upper

layer is 18-25 km thick) increases to about 70 km

after a time of 40 Ma in the collision zone, which

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352

L I LOBKOVSKY AND V I KERCHMAN

IO

20

30 40

t , Ma

time

Fig. 6. (a) The curves of succe.ssivehickeningof the lower ductile crust in the wurse of continentaJcolIisionobtained by numerical

simulation; values near the curves denote post-collision time in Ma. (b) Time dependence of subduction velocity of subcrustal

litbosphere elative o the upper crust since the beginningof the collision. See text for explanation.)

agrees well with data on the crustal thickness in

the Himalayas. The typical lateral dimension of

such an area with thickened crust (formation of

the Himalayas and Tibet “roots”) is 600-700 km,

corresponding well with the actual area of isostatic

uplift of the territory.

The calculated asymmetry of relief about the

suture zone, which implies a steeper topography of

the frontal erogenic area, a considerably more

gradual slope at its rear, and the existence of a

collisional plateau, is more pronounced in the

calculations when the non-linear viscous rheologi-

cal model of eqn. (1) and the evolutionary equa-

tion (3) are used. In Fig. 7 is shown the evolution

of a collisional plateau due to the northward prop-

agation of the material “pumped” under the orog-

eny.

We should mention that the model described

above is based on the principle of local isostasy

bending rigidity of the lithospheric lower layer

(III’) should be included in estimates of the

isostatic subsidence of the subcrustal part of the

lithosphere under the load of a thick crust (Karner

h.

km

L

L

I

1W 200 300 400 ml

600

X.km

dlstonce from orogen ax,s

Fig. 7. Succe&on of the cakzuiated urves showiiq the evolu-

tion of a collision rdateaa values near the curves denotina

(see eqn. (Al) in the Appendix). In reality, the

post-colliSionime n Ma.

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A TWO-LEVEL CONCEPT OF PL4TE TECTONICS

Fig. 8. Thermal scheme of a three-layered lithosphere witbin an

intracontinental subduction zone, based on the model of con-

tinental collision (see text for explanation). 1 = Granite upper

crust; 2 = “basalt” lower crust, 3 = olivine subcrusti litho-

sphere; 4 = direction of subduction; 5 = horizontal velocity

distribution in the layered lithosphere; 6 = frontal thrusts.

and Watts, 1983; Lyon-Caen and Mohmr, 1985).

Therefore, the above dynamic model of a colli-

sional orogeny should be generalized to include

quasi-elastic bending of the hthospheric lower

layer (III’) under the influence of a thickening

crust. With such a model it would be possible to

study the dynamics of the formation of foredeep

basins, and to have a more realistic view of the

theoretical history of uplift.

We shah now analyse the thermal regime of the

crust and subcrustal lithospheric layer in the zone

of a continental collision, using the above geody-

namic model and considering the additional factor

of dissipative heating of the media. Assuming that

all displacements are horizontal, we consider a

tw~~~ion~ thermal model of a three-layered

lithosphere that includes a “granite” upper layer

of the crust (I), 0 G z B h,(x), a “basalt” lower

crustal layer (II), kl (x)gz G k2(x), and an

olivine subcrustal part of the lithosphere (Fii. 8;

Taylor and McLennan, 1985). Taking into account

the motion of the lower layers (the upper layer of

the crust is assumed to be locked and therefore

stable), dissipation in the ductile layer of the lower

crust and radioactive heat generation in the crust

lead to the following non-stations

heat conduction:

aT

2

a+fi

37 =‘I ax2

(

ar2

1

+Q,=P -z/M

wP)1

O

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354

lL.l. LOBKOVSKY AND V.I. KERCHMAN

km further beneath the Himalayas and Tibet; Fig.

8). The thickness of the “granite” layer was as-

sumed to increase from 12.5 km in the south to 20

km in the north.

The rheological parameters of the rocks com-

posing the deformable “basalt” crustal layer were

taken from experimental data on diabase (Kirby

and Kronenberg, 1987): n = 3.4; A = 3 x lo6

GPa-” s-’ and Q = 260 kJ/mol in the tempera-

ture range from 400 to 800°C and at strain rates

from 10-‘4-10-‘3 s-i.

The temperature distribution in a “normal”

continental lithosphere (Morgan and Sass, 1984;

Taylor and McLennan, 1985) was taken as the

initial state. These temperatures were held con-

stant for the entire period of the modelling process

at the boundary of the study area. Numerical

estimates were made using an explicit finite-dif-

ference scheme with steps of Ax = 6 km; AZ = 2.5

km, At = 0.05 Ma. The numerical analysis de-

scribed here used the following parameter values:

X, = 2.5 W/mK,

(pc,,), = 3 x lo6 J/m3K,

X, = 2 W/mK,

(PC,), = 3 x lo6 J/m3K,

X, = 3.5 W/mK, (PC,), = 4 x lo6 J/m3K,

Q,=2X10e6 W/m3;

qb = 0.5

x

1O-6 W/m3.

In Fig. 9 is shown the evolution of the calcu-

lated geotherms (for 30 Ma) of the plastic flow

developed in the lower crust, with the relative

velocity of the subcrustal lithosphere given by Fig.

6b. It can be seen that the temperature in the

lower-crustal layer increases to 680-700 o C as a

result of dissipative heating. The mantle heat flow

is screened by the anomalously hot lower crust,

leading to temperature increases in the mantle: in

the adjacent 10 km thick layer of the subcrustal

lithosphere, for t = 15-20 Ma, the increase is 80-

150’ C (up to 700-750” C) and, below that, it is

40-80” C (up to 750-800” C). Thus, the litho-

sphere subducted under the Himalayas is already

heated to 700-800” C in its upper part, and so

brittle rupture does not occur in the zone of high

shear strain (Molnar and Chen, 1983; Jackson,

1987). This may explain the aseismic behaviour of

this entire intracontinental subduction zone. Mid-

c

km

Fig. 9. Calculated lithosphere gmtherrns for different times

since the beginning of the collision. I = Initial geotherm; 2 =

geotherm for 10 Ma; 3 = geotherm for 20 Ma; 4 = geotherm

for 35 Ma.

depth and deep-focal seismicity may occur in some

regions of a subducted lithosphere due to some

local underheating (for instance, because of the

sensitivity of the dissipation process to rheological

properties of the media, or because of the ex-

istence of more brittle lithosphere as a result of

water release during deserpentinization of the sub-

siding parts of the oceanic crust (Lobkovsky et al.,

1986).

The estimated dissipative heating of the lower

crust is sufficient to generate collisional and post-

collisional magmatism, since the melt temperature

for damp (hydrous) granites at depths of 15-25

km is 650-700°C (S&mid and Wood, 1976;

Dobretsov, 1980; Reverdatto and Kalinin, 1980;

Taylor and McLerqmn, 1985).

We note that for a period typical of a continen-

tal collision, i.e. 30-50 Ma, the thermal dis-

turbance in the lower crust has only a slight effect

on the surface heat flow. Nevertheless, the high

mean heat flow of Hinchrstan (especiaRy in the

north; &I&, 1985) may he a result of dieraipative

heating of the media; local variations caused by

various types of radiogenic heat release in the

surface crystal layer and by other factors are not

considered in this paper.

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A TWO-LEVEL CONCEPT OF PLATE TECTONICS

Geodynaniics

,of

passive tlwghw taking

into

account te49onic flow of the lower crust and sub-

lithosphere

3nombm mantle

The general features of the genesis of passive

margins have been explained within the frame-

work of the classical theory of plate tectonics (Le

Pichon et al., 1973; Burke and Drake, 1974). How-

ever, certain details of their structure and evolu-

tion cannot, as yet, be explained sufficiently well.

In particular, these phenomena are (1) the forma-

tion of elongated uplifts on the continents, parallel

to passive margins (Ollier, 1985); (2) the existence

and evolution of a system of rift zones at the rear

of passive margins (Milanovsky, 1976; Ziegler,

1982); (3) the process of tearing away of crustal

blocks (such as microcontinents and terranes) from

large continental massifs - a process which, in

fact, is the opposite of accretion tectonics (Sengor,

1984; Vink et al., 1984; Kazmin, 1989); and (4)

the occurrence of a belt of crystalline basement

(100-200 km wide) with an anomalous P-wave

velocity of about 7 km/s between the ~ntinent~

slope and normal oceanic crust (Emery and

Uchupi, 1984).

The two-level geodynamic model of the evolu-

tion of passive continental margins proposed by

Lobkovsky (1989) and Lobkovsky and Khain

(1989) explains the features listed above qualita-

tively. We shall consider this model briefly and

then describe the evolution of passive margins

within its framework quantitatively. To a first

appro~mation, the spreading of the ~thosphere

subsequent to the process of continental rifting

leads to the formation of passive margins. This

spreading is determined by the interaction of the

four main layers of the crust and upper mantle,

namely (I + II’), (II), (111’) and (M), which

were described in the Introduction (Fig. la;

Lobkovsky, 1989).

An accumulation of anomalous mantle first

occurs under the pre-rift uplift during the exten-

sion of the ~thosph~e. This accusation results

in the breaking through of a portion of partially

melted mantle into the crustal level of the geody-

namic system. Thus, a typical continental rift

structure develops that includes upper subcrustal

(IV) and lower sub~~osphe~c (V) lenses of

mf @2 @j-j@ m4 loo;;lS /36 B7

Fig. 10. Successive stages of rifting and spreading processes:

(a) continental rifting; (b) initial stage of ocean spreading; (c)

mature stage of ocean spreading (after Lobkovsky, 1989).

I = Upper brittle crust; 2 = lower ductile crust; 3 = subcrusti

lithosphere; 4 = anomalous mantle; 5 = normal mantle; 6 =

owau crust; 7 = sediments (see text for explanation).

anomalous mantle (Fig. lOa). The size of the

former lens is several tens or hundreds of kilo-

metres, whereas that of the latter is, at first, several

hundreds of kilometres; later, in the stage of ac-

tive spreading, becoming thousands of kilometres.

The appearances of a subcrustal diapir of anoma-

lous mantle and related additional local extension

and heating of the crust lead to thinning of the

lower plastic layer of the crust due to the outflow

of its material off the rift axis, resulting in an

&static near-axial subsidence. The development

of extensional fissures and normal faults in the

brittle upper layer of the crust that forms the

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356

L.I. LOBKOVSKY AND V.I. KERCHMAN

structure of the rift graben can be observed (the

lower crust and anomalous mantle propagating

under the peripheral zones of the rift lift their

“shoulders” simultaneously (Fig. 10a). Thus, the

dynamics of the crustal layer governs this system’s

interaction with the subcrustal mantle lens: the

sublithospheric anomalous mantle causes an uplift

of the terrane and acts as a feeding chamber for

the upper subcrustal lens.

Continental rifting is completed by extreme

thinning of the lower crust and rupturing of the

upper brittle layer, which is also considerably

thinned by that time. A final stage of developed

spreading follows this one. In the final stage, the

separated parts of a continental plate move away

from each other and undergo further evolution

(Figs. lob and c).

The previously cited theoretical papers devoted

to the evolution of the continental passive margins

are mainly concerned with the problem of vertical

subsidence due, first, to the cooling and densifica-

tion of anomalous mantle preserved under the

crust (Artyushkov, 1979; Meissner and Kopnick,

1988) and, second, to accretion of a “heavy”

oceanic lithosphere “brazed” onto the continental

lithosphere along the line of the prime rupture

(Sorokhtin, 1979). Note that the subsidence of the

basement of the passive margin is enhanced by the

additional loading of a rapidly accumulating sedi-

mentary cover (Cloetingh et al., 1984), whereas rift

expansion and subsidence of a continental edge is

accompanied by the development of listric faults

(Le Pichon and Sibuet, 1981, Kazmin, 1987).

In addition to the phenomena described above,

within the passive transitional zones from conti-

nents to oceans, new processes arise that are in-

duced by horizontal motions in various directions

of the media in the upper and lower levels of the

heterogenous system illustrated in Figs. lob and c.

The main process at the lower level is the propa-

gation of a large lens of sublithospheric mantle

away from the axis of a mid-oceanic ridge, which

is usually accompanied by “dragging” through

convection. The formation of wide swells along

passive margins seems to be connected with this

process. It occurs about 30-100 Ma after the

onset of relative spreading (Lobkovsky and

Khain,

1989).

We now consider a quantitative model of this

phenomenon (Figs. lob and c). The equation de-

scribing the evolution of a rheologically homoge-

neous layer of anomalous mantle with variable

thickness h and viscosity na (which is consider-

ably less than that of the over- and underlying

media) reduces to eqn. (2) in view of the observed

isostatic condition (Artyushkov, 1979; Lobkovsky,

1988a).

The coefficient of eqn. (2), K = pM -

Pa)Pag/12Plvrnla, where pM and p, are the densi-

ties of the normal and anomalous sublithospheric

mantle, respectively. Typical values of the parame-

ters are as follows: plLl= 3.4 g/cm3, pM -

Pa =

0.1-0.5 g/cm3; and the viscosity qa of anomalous

mantle is - lo’* Pa s. For the coefficient

K we

thus have an estimate K = (l-5) X lo-’ km-’

yr-‘.

First consider the simplest model problem,

which ignores the convective motion of the sub-

lithospheric mantle (this motion seems to be im-

portant only in a narrow belt, several hundreds of

kilometres wide, along the margin). In this case, if

the axis x is oriented horizontally towards the

motion of the anomalous mantle and the reference

point x = 0 is located under the crest of the coastal

slope of the margin, eqn. (2) is valid only if x >, 0.

Using the equation:

-=KK-

(6)

we solve the problem of the propagation of the

anomalous mantle lens beneath the continent, sub

ject to the boundary condition that its thickness is

constant at x = 0.

hl,_,,=H, hl,=,=O

(7)

We choose an initial condition that represents the

early stage of the under-flow of a localized massif

of anomalous mantle, its width being I, = 100-200

km:

(usually H =

50-80 km,

the dimension of a study

area S being 2500-5000 km; see Fig. 1Oc).

The solution of eqn. (6) with such initial condi-

tions is characterized by a finite rate of propa-

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ATWO-LEVELCONCEI'TOFPLATETECTONICS

I 1

I

/

I

I

Q

I‘? A4

d6 A8 mc

Fig. 11. Graph of the function f(t) for a self-similar asymp-

totic law of anomalous mantle lens propagation.

gation of the frontal boundary of the disturbed

area X(t) (Barenblatt, 1980). The boundary con-

ditions (7) permit self-similar “intermediate”

asymptotics, such as:

h=f fW, E= -

Substituting (8) into eqn. (6) we obtain an

ordinary differential equation for the function

f 5):

the conditions being:

fl

+,=l

and

f

It_co=O

(10)

Here to = lim, _ f0

X(t)/dKH3tl

so that f I ,to

= 0; the function d f 4/dc should be continuous at

the point t = &, to satisfy the flw-continuity con-

dition of anomalous mantle.

A plot of the numerical solution

of this

boundary-value problem is presented in

Fig. 11.

To estimate the time of propagation

for the

351

anomalous mantle front at a distance L, from the

edge of the continental margin, we use the

asymptotic law of propagation of the front:

t=o %Li/~H~

(11)

When

H =

70 km (which agrees with an isostatic

uplift of S = 2 km) and uplift width L, = 1200

km, we have (for K = 0.15 km-’ Ma-‘) I* = 30

Ma.

The general solution for the full eqn. (2) has

been obtained by an explicit finite-difference

scheme. This method permits us to consider a

non-stationary boundary condition at the left-hand

edge of the study area, reflecting, for instance, a

local thinning of an anomalous mantle lens due to

the subsidence of an oceanic plate margin under

the increasing load of accumulating sediments.

The initial form of the protruding massif of

anomalous mantle under the continental margin is

assumed to be:

ho(x)=

i

Hew[-0.2(x/H)2], O~XGI,,

0, l,

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358

L.I LOBKOVSKY AND V.I. KERCHMAN

The evolution of the anomalous asthenospheric

layer beneath the continent is presented in Fig. 12.

The self-similar law (11) is a good description of

the propagation of the lens for a time t > 10 Ma.

Thus, the flow of anomalous mantle in the

lower level of the system uplifts the edges of

continental passive margins, and lags considerably

behind the primary rifting (Lobkovsky and Khain,

1989; Kerchman and Lobkovsky, 1990a). As flow

beneath the continental passive margin reaches

maturity (Fig. lOc), the margin undergoes various

vertical motions: (1) gradual uplift as a wide

marginal swell is formed as a result of the flow of

asthenospheric matter under the continent

(another competing uplift mechanism may be con-

nected with thermal erosion of the lithosphere

basement, its thinning and isostatic uplift); and

(2) very sharp, large-amplitude, subsidence which

occurs in a much narrower belt, caused mainly by

the cooling and compaction of the upper lens of

anomalous mantle, probably accompanied by

phase transitions of the basalt-eclogite type

(Artyushkov, 1979).

The two main mechanisms presented cannot

completely explain the evolution of continental

passive margins within which, in accordance with

a two-level model of plate tectonics (Lobkovsky,

1989; Lobkovsky and Khain, 1989), one more

geodynamic process affecting the development of

its structure should occur. This is the flow in the

lower ductile layer of the crust within a continen-

tal margin at the spreading stage (Bott, 1972)

(Figs. lob and c). Such a flow, from the continent

towards the ocean, starts when the continental

crust splits. By that time, as was mentioned earlier,

both layers of the continental crust in the rift zone

are already considerably thinned. When the conti-

nents begin to move apart, the ductile lower crust

at the edge of the continental margin tends to be

squeezed out towards the ocean, as it is subjected

to an uncompensated horizontal loading. Taking

into account isostasy on the mantle surface, the

outflow of the non-linear viscous material of the

lower crust towards the ocean can be described by

the evolution equation (3) of the “Mechanical

aspects” section, derived in the Appendix.

Let us consider the problem of a self-squeezed,

semi-infinite layer of the lower crust when it is

I

I

I

I

I

I I

-\

,

-42 -60

-48 -46 -44

-42 a 42

44 E

Fig. 13. Oceanward and continent-ward propagation of a lower

crust thickness inhomogeneity on a continental passive margin.

(a) Numerical solution (values near the curves are time in Ma).

(b) Asymptotic self-similar solution: I, for unfractured upper

crust “cap”; 2, for dissected upper-crust cover of the “tongue”.

(See text for explanation.)

able to propagate horizontally after the continents

start to move apart; that is, under conditions such

as:

H, xx> -1, (hl,i,

h(,=,=O, hl,,_,=H

A numerical solution of the problem of the

flow for H = 18 km, /I = 10e9 kmP3 yr-’ and a

small initial wedge or “ tongue” (I, = 5 km, I, = 10

km) is shown in Fig. 13a. For longer times (t 2 20

Ma), the propagation of the “tongue” towards the

ocean and the depression front towards the conti-

nent are well described by the asymptotic self-sim-

ilar solution (Kerchman and Lobkovsky, 1990a):

h=Hf(E),

E= “-“‘:,a

(BH7f)

where function f satisfies the equation:

4;r;i

fS

3 3++-J

[ i

)I

(15)

in the domain &, f 5) = 0 and for t < t2,

f(E) = 1 (corresponding to an undisturbed state of

the media into which the depression wave propa-

gates at a finite rate). Moving towards the con-

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A TWO-LEVEL CONCEPT OF PLATE TECTONICS

59

tinental interior, such a front captures new por-

tions of the lower crust in accordance with the law

’

,=W t)

‘/4, from which it is seen that the

expansion of the flow of the lower crust decel-

erates sharply with time. The function f(t) is

given in Fig. 13b. Its construction is described in

Kerchman (1990).

The above analysis of ductile flow in the lower

layer of the crust predicts the appearance of ten-

sional stresses in the more brittle layers of the

overlying “granite” crust and underlying rigid

lithosphere. These tensional-stress maxima de-

velop in the vicinity of the disturbed wavefront as

it propagates towards the continental interior. The

maximum tensional stress in the upper brittle layer

of the crust is estimated as:

1

I

x2

e=-

d

rdx

x1

I Pch - P&( If2 -q

%.&

,

h1= Wx,) 07)

where d is the thickness of the upper, mechani-

cally strong layer of the brittle crust; X, is the

section separating the non-fractured part of the

brittle crust from the frontal, oceanward “frac-

tured” area, including the zone of listric faults

(Figs. 14a and b); and x2 is the continent-ward

km

b

Fig. 14. Structure and presumable evolution of continental passive margin: (a) North Atlantic passive margin, Flemish Cap bank

(after Emery and Uchupi, 1984); (b) model pattern of passive-margin dynamics (after Lobkovsky, 1989) (see text for explanation).

I = Synrift sediments; 2 = post-rift sediments; 3 - upper “granitic” crust; 4 = lower “mafic” crust; 5 = subcrustal lithosphere;

6 - anomalous mantle fens; 7 = normal mantle (asthenosphere.); 8 = fractured zone in subcrustal lithosphere; 9 = lower-crustal flow;

10 = filtration in partially melted astbenosphere. (Values on the upper figure denote P-wave velocities in different layers.)

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MO

L 1 LOBKOVSKY AND V 1. KERC’HMAN

front of the flow propagation in the lower crust.

For

d =

8-10 km, H = 15-20 km and h, = 5-10

km, the mean tensional stress in the upper crust

reaches u = 20-40 MPa. A long-term tension of

this size seems sufficient to cause brittle rupture of

the upper weaker part of the crust (Sawyer, 1985)

rifting, and the formation of faults and effective

“removal” of the blocks of the upper crust ocean-

ward (Fig. 14).

The proposed rifting mechanism operates

through the lateral inhomogeneity and rheological

stratification of the crust within the transition

zone. It seems natural to attribute to it the ob-

served evolution of rift structures that are sub-

parallel to continental passive margins - near the

continental slope, in the early stages of ocean

opening, and at the rear of the transition zone in

the later evolutionary stages (Fig. 14b). The same

mechanism acts throughout the geological history

of a passive margin, causing the splitting and

subsequent break-off of continental blocks from

the margins of large continents (Vink et al., 1984;

Lobkovsky and Khain, 1989).

An obvious result of the creep-propagation

model of the lower-crustal “tongue” is the ap-

pearance of a layer of anomalous basement rocks

with a P-wave velocity of about 7.0 km/s (typical

of the lowermost continental crust) in the transi-

tion zone between the continental slope and nor-

mal oceanic crust (Lobkovsky and Khain, 1989).

It is clear from seismic cross-sections of the con-

tinental passive margins that such an anomalous

rock layer does, in fact, underly sediments in the

transition zone, occupying a belt about 100-200

km wide (Emery and Uchupi, 1984). Geochemical

data also exist which provide evidence for the

continental origin of the lower crust in the transi-

tion zone of passive margins, even in cases where

the upper-crustal layers are composed of basalts

(Morton and Taylor, 1987). The model presented

also explains the quiet magnetic field in this tran-

sition zone (Boillot, 1983).

Once local fracturing and the formation of a

fault structure have occurred in the upper crust,

the boundary conditions on the top of the ductile

lower layer change. In particular, it shifts to the

regime of a quasi-free horizontal displacement of

the top boundary, leading to changes in the effec-

tive coefficient in equations of the type (2) or (3).

The coefficient ratio becomes /?,//? - 10 for the

fractured, oceanward-moving tongue (see the Ap-

pendix). This, in turn, causes an approximately

double acceleration of the flow under the newly

fractured part of the brittle crust, and the effective

transport of the corresponding block toward the

ocean, as well as substantial thinning of the dis-

sected crust at the rear of the block and the

consequent formation of a sedimentary basin there

(Figs. 14a and b). The scheme described thus

causes a discrete sequence of events: when the

previous oceanward block of the upper continen-

tal crust is faulted off, a new cycle of preparation

for the subsequent rifting stage starts within its

continent-ward part. According to the above anal-

ysis, the first stages of marginal rifting (at an early

stage in the ocean opening) often break off “small”

blocks such as the Blake plateau, the Flemish Cup

bank, the Rockall plateau, the Vijring plateau, the

San Paulo plateau, the Exmouth plateau and

others. When the “squeezing-out” of the lower-

crustal layer decelerates (due to the cooling and

blocking of the front of the tongue), the flow wave

propagating toward the continent behind decel-

erates. As the pressure gradients on the roof of the

lower crust decrease, a strain sufficient for rifting

accumulates at considerably greater distances from

the continental edge and over the considerably

greater time of hundreds of millions of years. This

time can, for instance, even exceed the duration of

one Wilson cycle in the Atlantic. The graben

systems of China and the Rhine, the Labrador rift

zone and others may be due to this process of

rifting at large distances from the ocean. Very

often, such distant rift zones are located in ancient

weak sections of the subcrustal lithosphere (or of

the crust), especially along suture zones (Vink et

al., 1984; Dunbar and Sawyer, 1988). The ad-

ditional tension of a quasi-rigid subcrustal litho-

spheric core, caused by the flow of the lower crust,

might reactivate the lower-level plate boundaries

(in particular, transform and suture boundaries

might change to divergent).

This might cause complete splitting up of the

lithosphere through formation of a divergent litho-

spheric boundary in the lower level. This may

occur long before the rupturing of the crust. The

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A TWO LEVEL CONCEPi OF PLATE TECTONICS

361

possibility of the complete disintegration of the

continental lithosphere and a jump of an oceanic

spreading axis to a new position at the continental

margin arises when the above rifting rn~h~srn is

superimposed on the general (“background”) ten-

sion of the lithosphere. Such a background tension

may be created by horizontal flow in the astheno-

sphere and/or by a pull due to a subduction zone

on the opposite edge of the oceanic part of the

plate (Bott, 1982). This situation may have oc-

curred during the structural evolution (destruc-

tion) of Gondwanaland, when a powerful horizon-

tal mantle flow (from the south to the north) and

the “pulling” of a subducted oceanic part of the

plate under Laurasia, together with the continu-

ously acting mechanism of the marginal continen-

tal rifting, caused the onset of new rift zones at

the northern boundary of the Gondw~a super-

continent. This mechanism may even have caused

the subsequent break-off of blocks of microconti-

nents, lithosphere terranes and corresponding

jumps of the spreading axes to a new southward

position (Sengiir, 1984; Kazmin, 1989; Lobkovsky

and Khain, 1989).

We note that the problem of the structure and

evolution of passive continental margins has been

studied extensively in the past decade. Various

models for passive margins have been developed,

particularly: models of different types of crustal

and lithospheric extension (McKenzie, 1978; Le

Pichon and Sibuet, 1981; Beaumont et al., 1982;

Wernicke, 1985; Lister et al., 1986); models incor-

porating a different rheology of the upper and

lower parts of the continental crust during the

extension (Bott, 1971, 1982; Meissner, 1985);

“ volcanic” models of the continental margins

(Royden et al., 1980; Mutter et al., 1988; Meissner

and Kiipnick, 1988); models of isostasy of passive

margins (Karner and Watts, 1982); thermomecha-

nical models of the evolution of passive margins,

taking sedimentation into account (Cloetingh et

al., 1984) and others. All of these models consider

various aspects of the structure and evolution of

passive margins, many aspects of which may also

be interpreted within the framework of the pro-

posed two-level scheme.

Modei fur ~~ q~-sy~ extension

of

the continental lithosphereand some featuresof

rifting

We now consider several aspects of continental

rifting from the viewpoint of the two-level plate-

tectonic model. A great deal of attention has re-

cently been devoted to the problem of the out-

crops of ultramafic mantle rocks on the Earth’s

surface (oceanic bottom) as a result of the com-

plete tectonic denudation of the continental crust

during rifting. In fact, the data show the existence

of anomalous ophiolitic sequences in which the

basalt layer and dyke complex are completely

missing, where oceanic sediments directly overlie

serpentinized peridotites which are sometimes as-

sociated with gabbro. Such incomplete ophiolitic

sequences are typical of the Ligurian segment of

the Mesozoic Tethys and occur, for instance, in

the Alps, Apennines and Corsica (Lemoine et al.,

1987). A similar anomalous structure of the oc-

eanic crust has been seen in a few places on recent

continental passive margins; particularly, on the

Galicia Bank and in the North Atlantin (Boillot et

al., 1987) on Zabargad island in the Red Sea

(Bonatti et al., 1986), and on the Sardinia passive

margin in the Tyrrhenian Sea (Lemoine et al.,

1987), where serpentinites or peridotites have been

found beneath sediments by deep-sea drilling.

These data suggested to some geologists that, at

least in the early stages of ocean opening, a special

geodynamic regime exposed the ultramafic base-

ment between diverging blocks of the continental

lithosphere (Boillot et al., 1987; Lemoine et al.,

1987). A model of the splitting and asymmetrical

divergence of continental lithosphere along a gen-

tle, through-going fault has been proposed to ex-

plain the formation of rather extended segments

of oceanic bottom (now composed of se~ent~iz~

peridotites) at divergent plate boundaries. This

model was first introduced by Wernicke (1981) in

an analysis of the tectonic situation in the Basin

and Range Province in the western U.S.A., and it

has subsequently been widely used for various

schemes of rifting (Wemicke, 1985; Lister et al.,

1986).

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362

From the standpoint of mechanics, one cannot

ignore the fact that in the Wemicke scheme the

postulated complete fracture of the lithosphere

along the single, gently dipping fault contradicts

the entire sum of our knowledge of the rheological

stratification of the crust and lithosphere. It seems

more realistic to suppose that the fracture and

deformation of the various layers of the crust and

lithosphere in the course of the rifting occur

according to the particular flow laws acting in

each layer.

Another purely mechanical objection to the

Wernicke scheme arises from the fact that the

lithosphere cannot bend sharply without sec-

ondary faulting during the normal slip of the

adjacent walls of the fixed inclined fault. As early

as 1958, Heiskanen and Vening-Meinesz, in their

L..l LOBKDVSKY AND V.I. KkKCHMAN

fundamental work, showed that during crustal ex-

tension, for instance, a second fault appears due

to the elastic bending of the crustal layer and

results in the formation of a graben structure. It

follows from this argument that even if we assume

an initial fracture of the lithosphere in accordance

with the Wemicke model - a single, gently dip-

ping fault - the further development of extension

would follow another course (see below) and would

be determined mainly by the appearance of an

additional fault in the lithosphere and by the

motion of a lithospheric block cut off by it (Us-

sami et al., 1986).

We now describe a new model of the develop-

ment of continental rifting, proposed by Lobkov-

sky (1989). This scheme is guided by the estab-

lished rheological and tectonic stratification of the

b

d

Fig. 15. Two-level model of continental rifting: successive stages (after Lobkovsky, 1989).

1 =

Upper brittle crust; 2 = lower ductile

crust; 3 = subcrustd lithosphere; 4 = asthenosphere; 5 = volcanics. (See text for explanation.)

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A ‘IWO-LEVEL CONCEPT OF PLATE TECTONICS

lithosphere. Derived from general mechanical con-

siderations and from the results of physical mod-

elling of the extension of an elastic-plastic litho-

sphere floating on a liquid basement (Shemenda,

1984) this model postulates that the loss of stabil-

ity and strain localization in the quasi-rigid mantle

core of the lithosphere would cause the formation

of two conjugate inclined shear planes. These shear

planes bound a central subcrustal lithospheric

block of a wedge-like or of a trapeziform geome-

try, which would rise under the influence of the

applied system of forces and squeeze the viscous

material of the lower-crustal layer away the axis

(Lobkovsky, 1989) (Figs. 15a and b). This, in turn,

would cause thinning of the crust (neck-forma-

tion), as well as additional extension and isostatic

subsidence of the upper brittle layer. It is im-

portant to stress that in this scheme the material

flow that leads to the thinning of the lower ductile

layer of the crust is induced not so much by the

external tensional force applied to the crust but by

the squeezing effect of the rising mantle block

(Fig. 15b).

Unlike the majority of the proposed schemes

(McKenzie, 1978; Le Pichon and Sibuet, 1981), in

this theory the crust is thus not subject to uniform

363

extension during continental rifting. This permits

us to explain the observed discrepancy between

the degree of extension of the crust, determined

from the system of faults, and the estimates of the

extension of the crustal layer derived from its

thinning and the isostatic subsidence of the surface

(Artyushkov, 1988). In fact, the analysis of the

structure of recent and ancient large depressions

within continents shows that in the majority of

cases crustal rifting involves extensions of only

several percent, whereas the thinning of a con-

solidated crust determined from seismic data is

50-lOO’%, i.e. approximately one order of magni-

tude greater.

Let us consider the above-described model in

detail, highlighting the principal stages of rifting

(Lobkovsky, 1989). These stages, which are char-

acterized by the occurrence of two inclined fault

planes, dipping from the axis of the future rift,

and formed in the mantle part of the lithosphere

during its extension, are illustrated in Figs. 15a-c.

As extension increases, a trapeziform block of

subcrustal lithosphere, bounded by fault planes,

begins to move upward under the influence of the

non-u~fo~y applied stresses (Fig. 16). This

movement was verified by physical modelling ex-

b

II..*....

. . . . . .

* -   . . . . , . s .

. . . . . . . .

Fig. 16. Rift scheme for cakulation of (a} balanceof forces and (b) uplift amplitudeof the axial sub~~phe~ block {see text for

explanation). I = Weight of the axial block 1rl; 2 = upper brittle crust; 3 = lower ductile crust; 4 = subcrustal lithosphere;

S = asthenosphere.

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364

I..1 LOBKOVSKY AND V.I. KFKCHMAIU

perimems, taking similarity criteria into account

(Shemenda, 1984).

The qualitative explanation of this phenome-

non is as follows. When the ~thosphere blocks Z

and ZZ move away from the central block ZZZ see

Fig. 16a), the average pressure on its inclined

planes AB and CD drops and, as a result of the

buoyancy force applied to the base of block ZZZ, t

becomes unbalanced. A subsequent drop in the

pressure, transmitted from the overhanging parts

of blocks I and ZZ, as they move away, occurs due

to their own elasticity.

Let us consider the uplift of block ZZZquantita-

tively and estimate its ultimate amplitude. We

assume that block ZZZ protrudes into the plastic

layer of the lower crust by an amount Ah. Then

the total hydrostatic pressure on the base of the

block equals:

where ZZ, and H, are the thicknesses of the crust

and of the mechanically strong subcrustal part of

the lithosphere, respectively; p, and p, are their

densities; 1 is the width of the upper surface of the

uplifted block; and a is the inclination of the side

planes of the block (Fig. 16a). The weight of block

ZZZ per unit of length in the plane, perpendicular

to the figure) is p&Z 4 H, cot cu)ZZ,; the pres-

sure on its upper plane is p,g(ZZ, - Ah)Z. The

vertical component of the pressure of the over-

hanging lithospheric blocks Z and ZZ on the side

planes of block ZZZcan be calculated as:

[2p,H,H, - pc(Ahj2 -t P&Z,,, - Ah)‘] g cot a

i-2

J

Hm-Ah(r - Ao cot a) dt

0

where the first term of this expression contains the

pure hydrostatic part of the forces, whereas the

second term reflects the deviation from the hydro-

statics on the side planes of block ZZZ n the lower

quasi-rigid lithospheric layer due to (a) the drop

Au of pressure transmitted from the overhanging

regions; and (b) the friction r along the inclined

faults. From the balance of forces applied to block

ZZZ and its weight, it is seen that at Ah = 0 (the

initial state after the inclined faults are formed;

see (Fig. 16a) the net force affecting block ZZZ is

negative (oriented upward), if the following condi-

tion is satisfied:

Aa > 7 tan cy

It should be noted that the magnitude ha is

proportional to the geomaterial strength up to its

fracture, whereas r has the meaning of a residual

shear strength; then from eqn. (19) it follows that

the uplift of the block may start when the angle a

is not very large, i.e. approximately 0 2 60 ‘.

On the other hand, angle LT annot be too small

(as is assumed in the Wemicke model; Wemicke,

1981); hence, if angles (Y are small, bending of

relatively extended near-fault parts of blocks Z

and ZZof a quasi-rigid lithosphere begins to play a

main role, and this would cause the secondary

faulting of the lithosphere. Such secondary faults

are proposed in the paper by Ussami et al. (1986).

Let us estimate the ultimate amplitude of the

uplift of mantle block ZZZ. First we note that, as

the block uplifts, friction 7 in the faults would

decrease due to the frictional heating of the media

and other factors. Ignoring friction 7 at the ma-

ture uplift stage of the block and assuming the

magnitude of Aa being approximately equal to the

mean strength a, of a quasi-rigid subcrustal litho-

sphere, we would write the condition of the force

balance as:

2u,( ZZ, - Ah) cot IY=

Apg[iAh + (Ah)2 cot CY]

(20)

where Ap = pm - pc.

For typical values of the parameters which are

used in eqn. (20) (Ap 5=0.4 g/cm3, g = lo3 cm/s*,

I = 20-30 km, a* = 100-300 MPa, ZZ,,,= 30-50

km), we have the uplift amplitude of block Ah -

lo-25 km. This means that, in the course of

rifting, a mantle block may be uplifted to the base

of the upper brittle layer of the crust. When a

quasi-rigid mantle block is uplifted, it squeezes

aside plastic matter of the lower crust, the flow of

which - in turn - causes additional extension of

the upper brittle crustal layer (Fig. 15b). Thus,

non-uniform extension of the crust occurs and its

lower ductile layer undergoes almost complete de-

nudation, whereas the upper brittle layer is ex-

tended and becomes thinner, although only by

several percent (Lobkovsky, 1989).

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A TWO-LEVEL CONCEPT OF PLATE TECTONICS

365

If the lithospheric extension ceases at

this stage,

a structure typical of a continental depression of

rift origin is formed, which has a sharply thinned

consolidated crust and a slight extension on its

surface. Some large aulacogens within continents

seem to have been formed according to this model.

Another situation is possible when much more

extension of the upper brittle layer of the crust

leads to its complete splitting. This may lead to an

exposure of mantle rocks directly on the bottom

surface of the opening oceanic basin (Fig. 1%).

Thus the proposed model may explain the ex-

istence of mantle rocks on some parts of the

oceanic bottom which correspond to the initial

stages of the ocean’s formation (in particular, on

passive continental margins). From this stand-

point, the Ligurian ophiolites seem to be attri-

buted to a rather narrow oceanic basin of the type

described.

An important consequence of the described

non-magmatic model of continental rifting is as

follows. Since the uplift amplitude of the mantle

block is proportional to the thickness of the

quasi-rigid core of a subcrustal lithosphere H,,

thinning of this strong core may hinder a rather

considerable uplift of the central block. In such a

case, a further divergence of lithospheric blocks I

and II cannot be compensated by the uplift of

block 111, and their adjacent sides would move

apart. The asthenospheric matter would flow into

the gaps formed as a consequence of this diver-

gence (Fig. 15d). As a result, two large channels

would be formed at the lateral sides of central

block 117, which would provide for the uplift of

the as~enosphe~c matter to the base of the crust

and further onto the Earth’s surface as alkaline

basalt outflows (Lobkovsky, 1989) (Fig. 15d).

It should be noted that the existence of mag-

matic channels along the margins of central block

111 may explain a very curious feature of rift

volcanism: the majority of large volcanoes are

usually attributed not to the central, but to the

peripheral parts of the rift zone. Powerful perip