Two types of rifting: dependence on the condition of extension

Tectonophysics,

85

143 (1987) 85-92

Elsevier Science Publishers B.V., Amsterdam - Printed in The Netherlands

Two types of rifting: dependence of extension

on the condition

V. KAZMIN Shirshov Institute of Oceanology, Academy of Science of the U.S.S.R., (Received August 8,1985;

I I 7218 Moscow (U.S.S.R.)

revised version accepted April 10,1987)

Abstract Kaxmin, V., 1987. Two types of rifting: dependence on the condition of extension. In: I.B. Ramberg, E.E. Milanovsky and G. Qvale (Editors), Continental Rifts-Principal

and Regional Characteristics. Tectonophysics,

143:85-92.

Rifting of continental lithosphere occurs in two ways: (I), through continuous extension and subsequent sinking and (2), in a series of pulses separated by quiet periods, sometimes of considerable duration. In the first case rifting is not accompanied by significant uplifts; in the second, uplifts of various sixes from 100 to 2000 km are typical. Rifting of the first type takes place when continental fragments are freely dispersed, as was the case with Pangea during its Late Triassic-Early

Jurassic break up. In the Cenozoic, rifts developed in the course of continental collision, when

relative plate motion was restricted and occurred in short pulses. Uplift is caused by accumulation of anomalous mantle at the base of the crust during the intervals between phases of extension. Geological evidence is not consistent with the idea of rifting being directly related to a local source of energy in the mantle (mantle plumes and diapirs). It is produced by regional stresses in the lithosphere which, in their turn, may have different origins. The “lithospheric inhomogeneities” observed below the present day rifts are not the cause, but a consequence of extension.

Introduction

ing is regarded as an “active” process (Gass, 1975;

As was pointed out by Talwani and Langseth (1981), the cause of continental rifting remains

Gratchev, 1977; Gass et al., 1978; Artyushkov, 1979; Ramberg, 1981; Logatchev et al., 1983 and many others). However, in this case the litho-

poorly understood. Geophysical studies indicate that large-scale lithospheric inhomogeneities exist below continental rifts. These inhomogeneities are in fact uplifts of the asthenosphere, forming large domes, much wider than individual rifts or grabens and elongated concordantly with the strike of the

sphere must be linked to the energy source for a long period to produce an appreciable effect on the surface. Consequently, rifting is correlated to “stable” intervals in plate motion (Burke and Dewey, 1973; Gass, 1975; Gass et al., 1978; Kazmin, 1980).

rift systems. Usually they are interpreted as active mantle diapirs rising from the deeper mantle. The light material of diapirs (or mantle plumes) accumulates at the base of the crust, resulting in isostatic uplift. Lateral flow of this material due to gravitational instability creates tensional stress in the crust, which finally breaks up to form grabens. According to this point of view, rifts are linked to energy sources immediately below them, and rift-

On the other hand, in many cases there is no visible indication of plume activity (volcanism, uplift etc.) and rifts appear to be “passive” fractures of the lithosphere caused by regional stresses applied to a plate or to an assemblage of plates (Kazmin, 1984 and others). The stress can be created by convective flow causing viscous friction at the bases of the lithosphere (Sorochtin, 1974), by gravity pull of the subducted plate (Bott, 1982),

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0 1987 Elsevier Science Publishers B.V.

86

by

membrane

stresses

(Oxburg

1974), or by the impact and Tapponier,

and

Turcotte,

of plate collision

(Molnar

1975). Pulses of the Earth’s expan-

to be controlled

great rift systems originated

1980). Passive rifting seems

in Europe,

by inherited

zones of weakness

in

pends

evidence

and an attempt

that the difference

on the conditions

is made to

between

in which

for the rifts de-

extension

oc-

Conditions of rifting

were epochs

in geological

history

when

we would

developed Pichakun,

and Antarctica.

In some

areas (South Africa and Antarctica), initial rifting was accompanied by extensive plateaubasalt volcanism which occurred simultaneously with subsidence. Other rifts of this epoch had practically no volcanism although extension was of the same order as in the volcanic rifts. An even more impressive example

Fig. 1. Triassic-Early

of a world-

Jurassic rifts of Pangea. (Reconstruction

to assume

margin

Siberia.

Some of

of kilometers

above

mantle

the unlikely

of a great number

rifts on the southern are of great interest.

of

long. plumes,

simulta-

of plumes. margin

of the

Comparatively

small slices of continental lithosphere (microcontinents of Palagonia, Kirsehir, South Armenia and Iran) split from the Gondwana continent and were carried away across the Tethys. A linear system of

rifting of Gondwana,

in Africa, India, Australia

have

The Triassic

continental Afghanistan

systems formed

and in Western

were thousands

neous development

numerous rifts developed sim~taneousIy over a wide area. A good example is the Late Paleozoic when graben

of

in the North

along the southern

If all these rifts originated

Paleo-Tethys

curs.

There

the Paleo-Tethys, these systems

the geological

break-up

Pangea (Fig. 1). In addition to the Late Paleozoic rifts of Gondwana, of which many continued to develop,

origin of rifts is studied demonstrate

is the Triassic

Atlantic,

as a possible

the plate body. In this paper,

of rifting

cause

sion have also been considered of rifting (Milanovsky,

wide epoch

rifts extending from Greece to was formed behind them and later into the Iiawasina,

deep-sea basins of Pind, and others. This disruption

of the margin can hardly be accounted for by local plumes or diapirs, but instead agrees with the conception of tension created by NE-directed convective flow. Development of some rifts is related to continent-continent collision. As demonstrated by Molnar

and Tapponnier

f 1975) and Tapponnier

et

of Pangea after Zonenshain and Savostin, 1979). I = faults; 2 = major

rift faults; 3 = subduction zones; 4 = mid-oceanic ridge; 5 = areas of diffused extension and volcanism.

al. (1982),

the Baikal rift is part of a structural

pattern resulting from the collision of India and Eurasia. Since the Late Eocene-Oligocene,

South-

east Asia has divided into a system of microplates with independent movements (Zonenshain and Savostin, 1979). In particular, the relative motion of the Siberian and Amur plates produce tension which is responsible for the opening of the Baikal rift. In the same way, movements on the faults of the Rhine rift system are caused by stresses generated in the collision zone of the Alps, where the Apulian microcontinent

collided with Europe (Il-

lies, 1975; Bergerat and Geyssant, 1980; Illies and Baumann, 1982). Stages of Rhine graben development coincide with the stages of collision; changes in the direction of collision resulting in transition from normal faulting to strike-slip displacement on major faults. Both the Baikal and the Rhine rift systems are apparently passive; however, their deep structure is similar to that of the East African rift which is usually considered to be active (Fig. 2). It thus seems that the domal projections of the asthenosphere below all these rifts do not represent active mantle structures (plumes and asthenoliths) and therefore represent the consequence and not a cause of rifting. To

produce

a visible

change

140

/

l.l

in the plate’s

surface (doming, rifting etc.), a mantle plume must be in contact with a specific part of the plate for a considerable time, i.e., the plate must be stable or slowly moving. This condition was applicable to Africa and Eurasia in the Late Cenozoic, when “absolute” motion of these plates was slowed down due to collision (Burke and Dewey, 1973; Gass, 1975; Kazmin, 1980). However, reconstructions of the “absolute” plate motion with reference to hot spots has shown that in the Late Paleozoic all major plates moved rapidly (Zonenshain et al., 1984). Nevertheless, extensive rifting occurred at that time. It is generally believed that domal uplift of the crust precedes rifting. In the mantle plume or diapiric models of rifting this is an essential point: the light material (“anomalous mantle”) accumulates at the base of the crust so that isostatic uplift is generated. Tensional features (grabens and normal faults) appear later when lateral flow

B

Fig. 2. Lithospheric inhomogeneities and rifts. A. In Central Asia (after Zorin et al., 1982) B. In Africa (after Fairhead and Reeves,

1977).

I = normal

faults;

2 = other

faults;

3 =

plateaubasalts; I = volcanic centers; 5 = depth of the asthenosphere-lithosphere

boundary measured from the base of the

crust in A and from the surface in B (in kilometers);

6=

boundary of the lithospheric inhomogeneity.

of the anomalous mantle begins. In fact, many rifts in the past were not accompanied by domal uplift. This applies to the Triassic rifts of Europe and North Atlantic, the Neo-Tethys, Western Siberia, Late Paleozoic rifts of Gondwana and the Mesozoic rifts of the North Sea. Moreover, when rifting is accompanied by uplifts, the succession of events is usually inverse-the uplift follows stages of extension (rifting) and subsidence. A good example is the Red Sea and the Gulf of Aden (Azzaroh, 1968; Kazmin, 1987). In the prerift

88

50-70

Ma

period,

most of the future Red Sea and Aden rifts

were covered by a shallow Late Cretaceous-Paleogene sea (Fig. 3A). The first protorift

depressions

were formed

in the Late Eocene close to sea level.

Immediately

after

sions were partly

sinking, invaded

these shallow

depres-

by sea which advanced

from the east (the Gulf of Aden and the southemmost

part

formation

of the Red Sea, Fig. 3B). Uplift of half-domes

in the Early Miocene, coarse

elastics

bordering

as indicated

to the marginal

and

the rifts began by the influx of parts

of grabens

(Fig. 4). The first stages of rifting in the Baikal rift system

occurred

in the Oligocene

m.y. or more ago), whereas

uplift

or Eocene

(30

and mountain

building is relatively young (3.5 m.y. ago) (Logatchev et al., 1983). The pattern in the Kenya rift is less clear. An uplift of the Kenya dome presumably began in the Middle Miocene, whereas the first indication

“\

35-40

Ma

of subsidence

(extension?)

is

dated as Early Miocene (McCall, 1967). However, some authors relate the start of doming to the Late Cretaceous (Milanovsky, 1976), although the nature of this early uplift is not clear (whether it was the Kenya dome sensu stricto or a larger uplift of the entire continental plate). In summary, geological evidence indicates that rifts are not connected to local sources of energy (mantle plumes and diapirs), but are created by regional stresses. At the same time, two groups of rifts can be recognized; those developing significant uplifts and those accompanied lifts on various scales. The first group includes

rifts developed

without by upduring

the Mesozoic break-up of Pangea. Their evolution is well estimated by models of uniform or of nonuniform extensions described by McKenzie (1978), Le Pichon and Sibuet (1981) and Beaumont et al. (1982). In this group of rifts, rapid extension and attenuation of the lithosphere are followed by Fig

3. Paleogeology of the Red Sea-Gulf

Late Cretaceous-Paleocene B. Late Eocene-Early

of Aden area. A.

(partly after Basahel et al., 1982).

Oligocene. Note the area of diffused

extension in the southern part of the Red Sea rift coinciding with the area of plateau-basalts. To the north this diffused extension is substituted with concentrated extension within the Red Sea graben. I = continental sediments, 2 = shallow marine sediments;

3 = evaporites,

ci = major, b = others.

4 = voIcanies, 5 = normal faults:

isostatic subsidence. Isostatic subsidence is replaced in the time by slow thermal subsidence due to thickening of the cooling lithosphere from below. Extension may lead to the break-up of continental lithosphere and spreading or to formation of rifts with strongly attenuated continental crust. This type of development is typical when continental fragments move freely. in this case volcanism occurs against a background of subsi-

89 2. 51O-

3% O-

S:: O-

Fig. 5. Possible

origin

moved with convective tal collision

block movements

tion created

tension,

causing

ness. I = continental volcanic

in lithospheric

2:::;3*4-

A, Fig. 4. Stages continental

of the development

crust,

4 = direction

2 = anomalous

of the Red Sea rift. mantle,

3 = oceanic

Arabian

fric-

lithosphere,

4 = ancient

flow, 6 = direction plate,

3=

zone of weak-

E-Eurasian

of tension plate,

plate.

I=

of movement.

as is seen in the Mesozoic

A-African

and viscous

along old zones of weak-

2 = oceanic

type,

of convective

plate.

plate

crust,

can rifts which originated and developed simultaneously with the Africa-Eurasia collision. This collision

dence,

of lithosphere rupture

lithosphere,

belt of the Andean

ness, 5 = direction q+++

of the Red Sea rift. The African

flow until the end of Eocene. Continen-

plateau-basalts,

in the Triassic deep-water basins of the Tethys passive margin and in the rifts of Western Siberia. The development of Cenozoic rifts is different. They are accompanied by uplifts varying in size from gigantic (East Africa and Baikal) to moderate (the Rhine rift system). Plateau-basalts related to Cenozoic rifts erupted on the uplifts (the Trap Series of the Ethiopian plateau, plateau-phonolites in Kenya and plateau-basalts of the Baikal

began in the Late Eocene and at the same

time, tensional structures were formed behind the collision zone (Fig. 5). The kinematics were different from that in the Baikal and the Rhine rifts. It seems plausible that when the subduction zone was blocked by continental lithosphere, the convective flow created an excessive tension in the African plate and finally ruptured the latter.

area). As already mentioned, these uplifts are caused by accumulation of light material (anomalous mantle) at the base of the crust. Seismic data show that the anomalous mantle forms 5-30 km thick lenses separated from the asthenosphere by mantle rocks of normal density. Structural interpretations based on seismic and gravity data were presented by Makris et al. (1975), Ginsburg et al. (1981) and Prijdehl (1981) for the Dead Sea and Ethiopian rifts, the Afar region and the Central European rifts. Anomalous mantle has also been demonstrated by geoelectric studies of the Baikal rift (Vanian and Shilovsky, 1983). In contrast to Mesozoic rifts (with free dispersion of Pangean fragments), Cenozoic rifting occurs invariably under the conditions of continental collision. This is true not only for the Baikal and the Rhine rift systems but also for the East Afri-

Fig. 6. Extensional occurs

in the area extension rifts.

structures

in northeastern

not only in Afar (I) and the Ethiopian of the Ethiopian

plateau.

is 700 km wide. 1 = normal

Africa.

Extension

rift (II) but also

The zone of diffused faults,

2 = axial zones of

90

Regardless of the kinematics, all Cenozoic rifts possess an important common feature: Their opening occurred in infrequent pulses separated by long, quiet periods. In the East African rift system, major rifting events are dated at 40, 22, 15, 10 and 5 Ma. These events are reflected in uncomformities, sharp changes in the character of sedimentation, outbursts of volcanism etc., whereas between events, the situation was stable. Similar short pulses of rifting are known in the Baikal rift and the Rhine graben (Illies, 1975; Logatchev et al., 1983). Pulses in rift development (as well as phases of folding in the fold belts) reflect the character of plate motion in the course of collision. This motion, obstructed by collision, apparently occurs periodically at the time of maximum stress release. Another important characteristic of continental rifting should be considered. In early stages, extension is usually diffuse occurring within a zone 300-1000 km wide (Fig. 6), and only later is it concentrated in single grabens, axial zones, and so forth. The pattern of deformation in early stages can be rather complex, with areas of both diffuse and concentrated extension substituting for each other along the strike of the rift system (Fig. 3b). When extension of the continental lithosphere

takes place, it results not only in subsidence of the rift floor but also in passive rise of the asthenosphere. Combination of the two processes leads to nonsymmetrical attenuation of the lithosphere: The rise of the asthenosphere is much greater than surface subsidence {Fig. 7). Physical models of this phenomenon were presented by Shemenda (1984). It may be further assumed that rapid (instantaneous) rise of the asthenosphere at near-solidus temperature causes an increase in partia1 melting. The light material (anomalous mantle) starts to rise to the base of the crust and accumulates there. Conclusion

The differences between Mesozoic and Cenozoic rifts may be explained if one takes all the above points into consideration. When the movement of plates was unrestricted (the Mesozoic rifts), a considerable attenuation of the lithosphere occurred over a short time interval (Fig. 7a). In this case the process proceeds as described by McKenzie (1978) and Le Pichon and Sibuet (1981), i.e., initial isostatic subsidence accompanied the extension and was followed by thermal subsidence of the cooling and thickening lithosphere. The extension may lead directly to the break-up of the

Fig. 7. Two types of continental rifting. A. Extension and subsidence during free dispersion of lithospheric plates (the Mesozoic rifting of Pangea). B. Extension in pulses during collision (the Cenozoic rifts). I = continental crust, 2 = Lithospheric part of the mantle, 3 = top of the asthenosphere, 4 = lens of low-density material (anomalous mantle) at the base of the crust, 5 = volcanism. &I = direction of vertical motion, b = direction of the relative plate motion, 7 = flow direction of light. rising material (anomalous mantle), 8 = boundary of the hthospheric i~omogen~ty.

91

continental lithosphere and onset of spreading, or it may cease before that occurs. In the latter case, large areas of subsidence are formed, e.g., the North Sea and Western Siberia. Anomalous mantle accumulated at the base of the crust, but no uplift occurred since attenuation and subsidence were the dominant factors. If rifting occurs in periodic pulses, and attenuation of the lithosphere during each pulse is small (the Cenozoic rifts), the picture is different (Fig. 7b). Initial attenuation occurs in a zone several hundred kilometers wide and the corresponding rise of the asthenosphere is followed by a long quiet period. During this period, anomalous mantle accumulates in the domal part of the rise and then ascends to higher levels due to the density difference. A lense of anomalous mantle is formed at the base of the crust, causing isostatic uplift. The whole cycle is repeated at the next stage of extension. From geological evidence it seems that large uplifts are formed 15-30 m.y. after the first stage of extension. This time span is probably needed for the anomalous mantle to migrate through the lithosphere mantle.

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