Rift initiation by lithospheric rupture

EPSL Earth and Planetary Science Letters 123 (1994) 281-298

ELSEVIER

Rift initiation by lithospheric rupture A. Nicolas a, U. Achauer b, M. Daignieres c a Laboratoire de Tectonophysique, URA 1370 CNRS, USTL, Montpellier, France b IPG, Strasbourg, France c CGG, UPR 361 CNRS, USTL, MontpeUier, France

(Received March 22, 1994; revision accepted February 23, 1994)

Abstract

Field evidence from lherzolite massifs related to asthenosphere upwelling below rifts and recent tomographic imaging of the East African Rift suggests that the upwelling occurred within a narrow and steep conduit, implying a lithospheric rupture. This evidence prompts a two-stage model of rift development. The first stage is characterized by a lithospheric rupture creating, in the upper crust, narrow troughs, with expansion being mainly caused by intrusion of mantle wedges at depth and expansion of basalt dykes above. The second stage is characterized by homogeneous stretching of the lithosphere. This stage is prepared by thermal relaxation following, 10 Ma later, the fracturing stage and hot injections. The thermomechanical behaviour of a stable continental lithosphere subjected to stretching suggests that the lithosphere has been previously weakened (a small thermal plume or presence of a mechanically weak zone). Rifting of a structurally and thermally homogeneous lithosphere seems to require both plume heating and tensional forces, thus limiting the meaning of the concept of 'active' and 'passive' rifting.

1. Introduction

McKenzie's [1] model of subsidence caused by thermal relaxation implying homogeneous lithospheric stretching has successfully explained many observations made in rifting environments and in passive margins. Evidence [2-4] supports the idea of large stretching occurring in the continental crust before oceanic spreading is initiated. The question discussed here is how the stretching of continental lithosphere initiates. Is it possible that, when a structurally and thermally homogeneous lithosphere is subjected to rifting, it immediately starts stretching in a grossly homo-

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geneous manner, as assumed in the general models mentioned above? This would imply that the most resistant part of this lithosphere, which we would expect to be the upper mantle, can yield plastically. Considering the temperature at Moho depths, which is estimated to be 500°C in 'normal' lithosphere, both experimentally determined rheology [5] and natural evidence [6] suggest that mantle peridotites would not yield to realistic applied stresses. We wish here to elaborate on this and present evidence obtained in mantle peridotite massifs emplaced in relation to rifting together with recent results of seismic tomographic experiments conducted in the Kenya Rift to suggest that the regional stretching stage needs to be preceded by a local lithospheric rupture, at least in the case of an initially structurally and

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thermally homogeneous lithosphere. Conditions for this rupture and subsequent evolution toward oceanic rifting will be also discussed. Lithospheric rupture allowing the intrusion of an asthenospheric wedge has been proposed by a number of authors. Fletcher and Hallet [7], Ricard and Froidevaux [8] and Keen et al. [9] envisage the rupture as a consequence of necking following an episode of lithospheric stretching that coincides with our second stage of rifting. McMullen and McMohraz [10] develop a model of early fracturing induced by the thermoelastic response of the lithosphere to an asthenospheric intrusion; this model is closer to our view. Evidence of an asymmetric fault pattern found in rifts has prompted the idea that the asymmetry could be caused by a detachment fault affecting the entire lithosphere, the so-called 'simple shear' model [11,12]. In keeping with that which is suggested above, we feel that this mechanism can operate in all situations where the lithosphere was locally weak at the onset of rifting. This could be typically obtained if this lithosphere was already crossed by a shear zone resulting from a former episode of lithospheric thrusting. We discuss briefly this problem of an inhomogeneous lithosphere in the case of the Red Sea and Kenya Rift and show that in these situations the available evidence from regional geology, rift-related peridotites and the seismic tomography results does not support the lithospheric detachment model.

2. Lithosphere rupture: geological evidence Evidence for lithosphere rupture during the early stage of rifting comes from a combined structural and phase equilibrium analysis of lherzolite massifs. The best example is the lherzolite massif of Zabargad Island in the Red Sea because this massif was intruded during an early stage of rifting and because it has not been affected by collisional tectonics, in contrast with the other examples cited here (i.e., Lanzo massif in the Western Alps, Alpujata massif in the Betic Cordilleras and Trinity massif in the Klamath Mountains of California). All of these massifs are

entirely or partly composed of plagioclase lherzolites; a detailed comparative study has been already presented [6]. A distinct case is that of spinel lherzolite xenoliths in basalts, as illustrated by the Massif Central rift in France. Lithosphere rupture is suggested there on the basis of both a study of mantle xenoliths carried to the surface by volcanoes and geophysical modelling. These occurrences are briefly described after a summary of the method used in peridotite analysis. Finally, we shall address the problem of how the fracturing in the lithospheric mantle is accommodated within the overlying continental crust. Our analysis is only two-dimensional, meaning that we have little constraint on whether the lithosphere rupture occurred in a dominantly tensional or strike-slip regime. Kinematic analysis of the early rifting in the Red Sea suggests, however, that sinistral strike-slip motion played a role [13]. In contrast, the regional geology in the Massif Central points instead to a purely tensional regime.

2.1. General analysis in mantle peridotites The method used here is more fully explained in [6]. In brief, it relies on (1) the capability to distinguish in peridotites plastic deformation which can be ascribed to asthenospheric flow (i.e., flow operating close to or above the temperature of the peridotite solidus) from deformation ascribed to lithospheric flow operating at temperatures between about 1100 ° and 850°C, and (2) the relationship established between petrological peridotite facies and geodynamic environments. The results of these studies have been published in several papers [e.g., 14-16]. Plastic deformation in a peridotite is ascribed to asthenospheric flow when it can be demonstrated that the deformation took place in the presence of a basaltic melt (i.e., in solidus or hypersolidus P,T conditions). This is shown by the presence of (i) gabbroic lenses parallel to the determined flow plane, (ii) undeformed gabbro dykes perpendicular to the stretching lineation (tension fracture orientation), and (iii) gabbroic impregnation patches devoid of any plastic strain whereas the impregnated peridotite presents a

A. Nicolas et al. / Earth and Planetary Science Letters 123 (1994) 281-298

strong lattice fabric induced by plastic strain. The nature of the lattice fabrics, of the recrystallization and of the dislocation substructure also reflects high-temperature deformation. In contrast, deformation is ascribed to lithospheric conditions when all dyke: are plastically deformed or when there is no relationship between late dykes and plastic flow structure and when the textures indicate low-temperature conditions (low-temperature slip systems, small neoblasts, tight substructure). In the lherzolite occurrences considered, this lower temperature deformation is commonly superimposed on the hypersolidus one, suggesting that the high-temperature intrusion was continued through a shallower and cooler environment. Dealing in all the analysed situations with large strain in a dominantly shear regime, we equate below the foliation-stretching lineation and flow reference systems. Systematic analysis of mantle peridotites has suggested that there is a relationship between their petrological nature and their environment of origin [6,15]. Spinel lherzolites, in which hightemperature flow was frozen at depths greater than 30 kin, are preferentially associated with continental rifts. Such rocks commonly occur as xenoliths in basaltic volcanoes and occassionally as massifs pinched in continental rift sutures, such as in the Pyrenees and in the Gibraltar Arc. Plagioclase lherzolites, in which high-temperature flow operated at depths shallower than 30 kin, seem to characterize oceanic rifts and slowspreading oceanic ridges. Finally, harzburgites, in which high-temperature flow operated down to the Moho, are associated with medium to fastspreading ridges. This segregation reflects an increasing degree of partial melting from spinel lherzolite to harzburgite, which results from the decreasing effect of conductive cooling in ascending mantle diapirs, which is itself related to ascent and spreading rates. In other words, spinel lherzolites would melt while deforming down to 30 km, plagioclase lherzolites down to about 1020 km, and harzburgites down to the oceanic Moho, at 6 km below the seafloor. This simple scheme does not apply to all lherzolites, as it has been shown that, geochemically, a number of them were not derived from ascending astheno-

283

spheric diapirs but from the subcontinental lithosphere. The duality of origin is illustrated by the case of the Lanzo massif [17] and possibly by that of Zabargad Island [18]. In conclusion, the situation of incipient rifting considered here should lead to a situation in which solidus to hypersolidus plastic flow is imprinted in lherzolites that range from plagioclase to spinel lherzolites. This seems to be the case with the massifs discussed above.

2.2. Zabargad Island, Red Sea There is now some agreement in considering that the opening of the Red Sea was characterized by two stages that were not necessarily temporaly discontinuous [see 13,19-22]. The first stage initiated during the late Oligocene-Early Miocene (24-22 Ma) (and perhaps as early as the middle Eocene [13,23]) and consisted in strike-slip extension, which opened narrow rifts into a continental crust (see below). The second stage was one of seafloor spreading. It is not yet clear when this started as the process was dependent on the type of crust believed to be present below the Miocene salt layer. Spreading could have started 10-12 Ma ago [24], although it was certainly operating 4-5 Ma ago, which is the age of exposed oceanic crust [13]. Between stages 1 and 2, the amount and timing of continental stretching is still an open question. Zabargad Island is mainly composed of lherzolites which have intruded and metamorphosed granulites that are partly of deep continental origin [25-28]. N - S basaltic dykes have intruded both the cooling peridotites at the top and the walls of the intrusions and the gneisses which are in contact with them. These dykes have been, to various degrees, recrystallized into amphibolites and rotated into parallel with the gneiss foliation, depending on their proximity to the intruding peridotites as shown by Boudier et al. [27]. These authors also show that, together with metamorphosed gabbroic layers, the amphibolitized diabase lenses constitute a large volume fraction of the gneiss in excess of 0.5 in most outcrops. All this points to a large volume of basaltic material

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continental crust lithosphere asthenosphere Fig. 1. Zabargad intrusion of an asthenospheric wedge through a lithospheric fracture below a narrow trough (modified from [26]).

rift continental crust lithosphere

asthenosphere A

-

Rifting by fracturing and injection oceanic CFII.R[

B

-

Rifting by lithospheric stretching

Fig. 2. Sketches contrasting the two main stages of rifting, with moderate extension by injection of mantle wedge and basaltic dykes during the early stage of lithospheric fracturing and large extension by dominantly tectonic stretching of the lithosphere during the following stage.

A. Nicolas et al. / Earth and Planetary Science Letters 123 (1994) 281-298

injected into the crustal gneisses at the time of the peridotite intrusion. A high-temperature deformation, ascribed to asthenospheric ascent by isotope [29] and trace element [18] geochemistry, occurred in the spinel and plagioclase lherzolites of the southern body in the island, as deduced from the observation of gabbroic veins parallel to the peridotite foliation and gabbro dykes crosscutting this plane. The dykes indicate that melt extraction was taking place during mantle flow. Spinel lherzolites in the central and northern bodies could either represent the skin of the inferred asthenosphere ascent, cooled in this facies by the surrounding lithosphere at depths of >~30 kin, or a fragment of the lithosphere in contact with the asthenospheric intrusion [18]. Intrusion continued below the solidus temperature, as the frozen-in foliation indicates temperatures of ca. 1000°C. Intrusion would have ceased at sea level [30], with motion restricted to numerous shear zones locally inducing a new mylonitic deformation in the peridotites and a granulitic metamorphism in the continental gneisses, which by this time had already been significantly contaminated by basaltic intrusions (Figs. 1 and 2). This important intrusion-related deformation was accompanied by an episode of hydrous metasomatism [18,31], with the development of syntectonic amphiboles. The peridotite intrusion has been dated at 18.4-21 Ma by U-Pb ages on zircons in gneisses [32], at and 20.2 + 0.3 Ma [33] and 26.1 + 1.1 Ma [Maluski, pets. commun.] by 4°Ar/36Ar ages affected by excess argon in diabases and related amphibolites. This is the period of initiation of the Red Sea rifting (stage 1 above). Considering the geophysical indication of a large peridotite root below the island [34] and the absence of any compressive activity which could have displaced the massif, it is believed that the observed hightemperature structure reflects the asthenospheric flow orientation. From the steep attitude of the foliation in the peridotites and kinematic study, it is concluded that an asthenospheric wedge could have penetrated the continental lithosphere along a steep, deep fracture oriented NW-SE, parallel to the Red Sea orientation, with a flow direction dipping 50°NW. A crustal model derived from

285

seismic profiles crossing the Egyptian continental margin, 250 km north of Zabargad Island, reveals a sharp, steep (50-60 °) Moho transition between this margin and the new oceanic crust [35] which could be the wall of a mantle intrusion similar to that of Zabargad. This is also suggested by the occurrence of gabbros outcropping on the Brothers Island located in the vicinity of the profile. Coleman [36] has documented the opening of a number of 'closed morphotectonic basins' during the late Oligocene, including, on the other side of the Red Sea, the Tihama Asir area. There, the continental crust is intruded by diabase dykes parallel to the Red Sea trend. Their orientation is close to that of the dykes intruding the gneisses on Zabargad and they have a similar age (4°Ar/ 36Ar plateau ages of 21-24 Ma) [37]. Locally, these dykes are so abundant that they constitute a sheeted dyke complex associated with a layered gabbro unit, a situation very comparable to that of an ophiolite. Bohannon [20] describes the Asir intrusion as filling a fracture that is 10 km across with steep walls through continental crust. At the top of the intrusion, extensional faulting has been documented. Such rifting-related intrusions within the continental crust were possibly caused by an asthenospheric intrusion through the lithosphere and lower crust, as at Zabargad. This episode of massive basaltic dyke injection parallel to the Red Sea at 20 Ma is recorded as far north as Sinai [38] and may be related to the copious flood basalt volcanism of a similar age in Yemen and Ethiopia [39,40]. An early episode of opening of linked grabens of limited extent ( ~ 35 kin) is also envisaged in the northern Red Sea by Cochran and Martinez [41]. 2.3. O t h e r massifs

The Trinity massif in the Klamath Mountains of California is composed of plagioclase lherzolites overlain by an ophiolitic crust [42]. The limited development of layered gabbros and the pervasive hydrothermal alteration penetrating down to the Moho have been interpreted as indications that the oceanic crust was thin with a magmatic inflow insufficient to maintain a permanent magma chamber. This is in keeping with the

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presence, as underlying mantle, of plagioclase lherzolites which are less depleted than the harzburgites encountered in most ophiolites and which have thus yielded less melt to the overlying crust. As a result, an oceanic rift or a slowspreading oceanic environment has been proposed for the site of origin of the Trinity ophiolite. The point of special interest here is that the high-temperature foliation in the lherzolites, which was active at the time of melt segregation, makes a high angle with the base of the crustal section and is broadly parallel to the sheeted dyke attitude which is assumed to be itself parallel to the rift trend. This can be established in spite of displacements on faults and plutonic intrusions subsequent to the ophiolite generation which render precise reconstructions difficult. The steep foliation attitude at the spreading site of origin suggests again that the flow plane attitude was determined by steep lithospheric walls. Incidentally, the flow direction within the flow plane is nearly horizontal. This longitudinal and horizontal flow has been interpreted as a consequence of the fact that the mantle flow, thought to be diverging from discrete asthenosphere diapirs, is channelled by lithospheric walls parallel to the rift trend [6]. In the Lanzo plagioclase lherzolite massif in the western Alps, the high-temperature foliation is again steep and parallel to the proposed trend of the rift of origin; lineations are moderately plunging. There is a zonation with higher temperature foliations recorded in the core of the massif and lower temperature foliations on the western and northern margins [43]. Geochemical evidence also indicates a higher degree of melting in the centre and ascribes an asthenospheric origin to the main part of the massif. Interestingly, the northern edge would be a piece of the lithospheric wall of the intrusion [17]. Unfortunately, the paleohorizontal reference frame is poorly constrained. The massif stands on the gravity high of the Ivrea zone and above a mantle swell determined by seismic refraction experiments [44]. This would represent the mantle underlying an oceanic rift comparable to the Red Sea, an interpretation now widely accepted [45]. We assume here that the steep foliations observed in this

massif have been little tilted by the Alpine collision and thus reflect the existence of steep lithospheric walls controlling the asthenospheric ascent. Finally, in the Alpujata plagioclase lherzolite massif in the Betic Cordilleras, Tubia and Cuevas [46] have proposed the interpretation of an asthenospheric wedge rising at high temperature within the crust of a continental rift. The shape of this wedge is speculative, however.

2.4. Massif Central rifting and mantle diapirism A systematic structural and petrological study of peridotite xenoliths in 70 volcanic vents of the Massif Central, France has revealed a regional pattern of variously deformed xenoliths that have equilibrated at various temperatures. Their deformation and P - T conditions of equilibration reflect the situation in the mantle at the time they were extracted by the basalt [16]. From this it has been possible to contour an area thought to be the surface projection of a similar area in the mantle where high-temperature strain was important with respect to the surrounding undeformed or little deformed mantle [47]. Considering the high-temperature nature of its deformation, the internal area is assumed to be the site of asthenospheric diapirism. In view of a thermal relaxation model [48] incorporating regional geophysical data and geothermometry results obtained on the peridotite xenoliths it has been assumed that the diapiric ascent ceased 10 m.y. ago, soon after the paroxysm of volcanic and tectonic activity. Between that time and their incorporation in the basalt carrying them to the surface, the deformed peridotite xenoliths have experienced various periods of static cooling. This cooling followed their deformation, which is thought to be related to the diapiric ascent and to have taken place at temperatures close to the dry solidus (1200-1300°C), either within the diapir as asthenospheric material or outside as pieces of the lithospheric walls. Assuming first that the area (60-80 km across) where deformed xenoliths have been collected corresponds to a single diapiric ascent results in relaxation models predicting temperature conditions in the depth interval where the xenoliths

A. Nicolas et al. / Earth and Planetary Science Letters 123 (1994) 281-298

have originated, which are, under all conditions, at least 100°C above the 920-1100°C temperature range estimated for the xenoliths by classical geothermometers. A g r e e m e n t between the two t e m p e r a t u r e estimates is obtained if one considers that the area of mantle deformation results from the ascent of a few smaller diapirs (10-20 km across), the extent of which has been indicated by field mapping [47]. This can be easily understood, because thermal relaxation is much faster within these smaller intrusions. By combining deformation and cooling history in basalt xenoliths, a similar conclusion regarding the ascent of small mantle diapirs has been drawn by Witt and Seck [49] for the Eifel rift in Germany and by Zipfel and Worner [50] for the Ross Sea rifted margin in Antarctica. Such small diapirs are represented in cross section in Fig. 3. The shaping of the diapirs is guided by the consideration developed above that the upper lithospheric mantle cannot accommodate deformation that is significant plastically. A round head would indeed induce large differential flow within the upper lithosphere mantle. It is proposed that the space for the diapir is progressively opened by a lithospheric rupture. In two dimensions, the 10 km space opened within the lithosphere at the base of the diapir should be filled above by an increasing number of basaltic dykes whose cumulative width above the diapir head also needs to be ~ 10 km (Fig. 2A). This

287

interpretation of extension caused by intrusion of new material at crustal levels is supported by the important fraction of basaltic dykes seen on Zabargad Island. Such dykes are not exposed at shallow depth in the Massif Central, where tectonic stretching, like on Zabargad Island, exceeds a few percent. There is however a need for large extension at deep crustal levels, because the crustal thickness is now reduced by one-sixth, without the need to invoke much erosion. This large discrepancy between surficial and deep extension is not specific to the Massif Central [see 51]. It might be explained by a decoupling horizon located between the lower and upper crust above the volcanic area. Thus, in the Massif Central rift, which would not have evolved farther than step 1 (see below), lithospheric expansion would result from intrusion of new material into the lithosphere and not from tectonic stretching. Souriau et al. [52] have estimated that, below the Massif Central, the lithosphere is now only 60 km thick. In the discussion developed below, it is suggested that this thinning might be necessary to make lithosphere rupture possible. 2.5. S u m m a r y o f results on m a n t l e occurrences

The information related to our purpose and which is derived from the various peridotite occurrences is summarized in Table 1.

Fig. 3. Model of asthenospheric intrusions beneath the Massif Central showing also the thermal structure established 10 Ma after the beginning of the emplacement of these intrusions [48].

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3. Mantle diapirism beneath the Kenya rift: Evidence from seismic tomography 3.1. Geology

T h e K e n y a rift forms p a r t o f the E a s t A f r i c a n Rift system, which c o m p r i s e s a series of rift z o n e s s t r e t c h i n g for m o r e t h a n 3000 k m f r o m the A f a r triple j u n c t i o n in t h e n o r t h to the Z a m b e z i R i v e r in s o u t h e r n Africa. T h e rift system b i f u r c a t e s a r o u n d t h e A r c h e a n N y a n z a craton, which coincides with t h e u p l i f t e d E a s t A f r i c a n p l a t e a u . T h e K e n y a rift t r a n s e c t s t h e K e n y a D o m e , which itself is s u p e r i m p o s e d on t h e e a s t e r n m a r g i n o f t h e E a s t A f r i c a n p l a t e a u . T h e K e n y a rift is d i v i d e d a l o n g strike into i n d e p e n d e n t subbasins, which in t h e c e n t r a l p a r t a r e c h a r a c t e r i z e d by w e l l - d e f i n e d 5 0 - 7 0 k m w i d e h a l f - g r a b e n s . T h e rift, w h o s e floor r e a c h e s an e l e v a t i o n o f up to 1800 m in t h e c e n t r a l p a r t b e w e e n N a k u r u a n d N a i v a s h a , shows b r e a k a w a y fault scarps with up to 1600 m differe n c e in e l e v a t i o n b e t w e e n t h e floor a n d t h e shoulders. At the northern and southern extremities t h e e l e v a t i o n o f t h e rift floor d r o p s b e l o w 600

m a n d the fault s t r u c t u r e s splay o u t over widths o f 200 km, t h e rift losing its g r a b e n - l i k e a p p e a r ance. T h e rift is a s s o c i a t e d with 14,400 k m 3 of volcanics e r u p t e d since t h e early M i o c e n e [53] a n d is l o c a t e d close to the b o u n d a r y b e t w e e n the A r c h e a n N y a n z a c r a t o n a n d the P a n - A f r i c a n M o z a m b i q u e s h e a r belt. This b o u n d a r y strikes N W - S E a n d crosses t h e rift at a b o u t t h e l a t i t u d e o f N a i r o b i [54]. 3.2. Tomographic imaging

T h r e e - d i m e n s i o n a l P-wave d e l a y t i m e t o m o g raphy, using a m o d i f i e d v e r s i o n o f the r o b u s t A C H m e t h o d [55,56], has b e e n a p p l i e d to i m a g e t h e s t r u c t u r e of t h e crust a n d u p p e r m a n t l e in t h e vicinity of t h e K e n y a rift. S h o w n in Fig. 4 a r e t h e velocity p e r t u r b a t i o n s p e r layer with r e s p e c t to a s t a r t i n g m o d e [ with c o n s t a n t layer velocities. F o r p r e s e n t a t i o n p u r p o s e s t h e original block-wise inversion results (the ' o f f s e t - a n d - a v e r a g e ' m e t h o d , see [56]) have b e e n s m o o t h e d (quasi-low-pass filter), a c o m m o n p r o c e d u r e in p r e s e n t i n g t o m o -

Table 1 Mantle occurrences presumably derived from asthenospheric ascent related to rift situations Massif

Nature of peridotite

Depth of last Asthenospheric Attitude of equilibration flow structures flow plane in rift framework

Attitude of flow line in ridge framework

Associated formations

Zabargad Plagioclase lherzolite

Subsurface

Trinity

Plagioclase lherzolite

3-6 km

Lanzo

Plagioclase lherzolite

3-6 km

Alpujata

Plagioclase lherzolite

< 30 km

Massif Central

Spinel lherzolite

> 30 km

Presumed geodynamic environment

Yes (partial melting evidence); lower T structures superimposed As above

Steep, parallel to rift trend

Parallel to ridge trend, 50° plunge

Deep continental Small oceanic crust; lithorift spheric mantle?

As above

Parallel to ridge trend, ~ 10° plunge

Ophiolitic cover

As above, with zonation: high T in core and low T along west margin High T in core; low T along margins

Uncertain, Parallel to probably steep presumed and parallel ridge trend, to rift trend 30° plunge

Ophiolitic cover; lithospheric wall

Unknown

Deep to middle continental crust

Continental rift

Unknown (xenoliths)

Unknown (xenoliths)

Various xenoliths

Continental rift

Parallel to presumed ridge trend, 0° plunge Unknown (xenoliths)

Oceanic rift or slowspreading oceanic ridge As above

0

100

5 -100

r--

13 Z

E

:/')

.-u

:1)

/3 £

E

r.3 t-

/3 7 (1)

E

-100 0 100 200 Distance E W [km]

100

3

I00 -

_~00 -

I

I

' ' ' ' •1 0 0 0 100 200 Distance E W [km]

I

ayer 5 195-125km]

.100

0

100

layer 3 [35-65km]

-100 0 100 200 Distance E W [km]

0-

100-

200

-

-200

I

I

L

100 0 100 200 Distance E W [km]

I

layer 6 [125-165km]

..5 - 1 0 0 -

t"

0

/3 Z

E ,~

0

100

5 -100

0 c--

Z

E

layer 4 [65-95km]

Zig. 4. Seismic velocity perturbations per layer with respect to a starting model with constant layer velocities below the Kenya rift [57].

-9 -6 - 3 - 1 1 3 6 9 P velocity perturbations [%]

100 0 100 Distance E W [km]

ayer2 [10-35km]

.,j

oe

I

',o ',o

r,~

r~

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graphic images. A detailed discussion of the method, its implications and limitations is given in Evans and Achauer [56] and Achauer et al. [57]. At all depths an average P-wave velocity contrast of 6 - 1 2 % is observed across the rift, with standard errors of the order of 1-1.5%. The variance improvement is approximately 65%, leaving about 0.28 s of the signal unexplained (which can be allocated to the short-wavelength strong heterogeneities in the crust, which are in the main not resolved by this method). The diagonal elements of the resolution matrix R (see section 3.5 in [57] for a detailed discussion of the resolution obtained) exhibits values between 0.5 and 0.8 in the centre of the layers, with the values decreasing to the edges. A careful examination of the resolution matrix reveals no sign of strong lateral coupling, but some vertical smearing between layers 2 and 3 [57]. Because we are here mainly interested in the lithospheric mantle structure our discussion will focus on the observations for the u p p e r mantle layers (Fig. 4, layers 3 and 4). The lowest velocities are observed beneath the rift and a small zone under the eastern shoulder. About 100 km north of the array centre, the low-velocity zone beneath the rift seems to deviate to the northeast, although this zone does terminate within the rift. The variability in velocities along the rift axis is of an order that is similar to that across the rift, indicating the three-dimensionality of the rift's deeper features. The lowvelocity zone, shown here in light tones, clearly exhibits two directional trends that are best visible in the depth range 65-95 km (layer 4), although these trends may be seen at other levels too. Whereas part of the low-velocity zone follows the direction of the present day rift, i.e. approximately N - S , the second trend of decreased velocity has a N W - S E trend, which coincides with the trend of the old Pan-African Aswa suture zone separating the Archean Nyanza craton to the southwest and the Proterozoic Mozambique shear belt to the northeast [54]. Down to a depth of approximately 100 km the low-velocity zone seems more or less confined to the rift itself, broadening only at greater depths

(layer 6) and giving the impression of a steep walled wedge of asthenosphere. This is supported by synthetic modelling [see 55]. As the low velocity persists vertically through the whole model, the interpretation of asthenospheric material seems to be quite likely. The lithosphereasthenosphere boundary is, then, somewhere between 35 and 65 km, i.e. at the depth where the top of the low-velocity zone begins. The variation in the amplitude of the anomaly along the strike of the rift and the small zones (some of an elliptical shape) of lowest velocities further suggests that these zones represent diapirs containing partial melt which are embedded in the asthenosphere with a velocity that is generally lower than that of the lithospheric material beneath the shoulders. The interpretation of the low-velocity zone as having an asthenospheric character with embedded diapiric structures containing partial melt is supported by studies comparing the Bouguer anomaly and the observed travel-time residuals [58,59], as well as by a study which relates the observed refraction seismic velocities to mineral composition, temperature and pressure [60].

4. Discussion 4.1. R i f t i n g m e c h a n i s m

In view of the forthcoming discussion, it is useful to recall the following results from the geological study of peridotites in rift locations and the seismic study of the Kenya rift: (1) Plagioclase lherzolite massifs display evidence of emplacement at shallow depth as steeply walled intrusions. An earlier episode of hypersolidus deformation is generally followed by progressively lower t e m p e r a t u r e deformations, suggesting a continuous intrusion at progressively shallower levels. Geochemistry indicates an asthenosphcric origin for more central parts of the massifs and possibly a lithospheric origin for some margins, as in the cases of the Zabargad and Lanzo intrusions. (2) A regional study of structures and equilibration conditions of spinel peridotite xenoliths

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in the Massif Central, together with geophysical constraints, suggests that the ascending mantle was divided into small ( ~ 10 kin) diapirs (Fig. 3). This is supported by the preliminary results from the recent joint F r e n c h - G e r m a n Massif Central teleseismic tomography experiment [61]. (3) The tomographic study conducted in the Kenya Rift and the modelling of the seismic structures in terms of diapirs leads to results that are very comparable with those from the Massif Central. In both cases a broadly distributed lowvelocity zone in the upper mantle ( - 3 to - 5 % velocity reduction), which roughly coincides with the Bouguer gravity anomaly, is observed. This zone embeds smaller, oval-shaped zones (of some tens of kilometres in diameter) of even more reduced velocities in the depth range 50-70 kin; these oval-shaped zones could possibly be diapiric structures within an asthenospheric wedge [57]. (4) The well-constrained environment of the Red Sea allows us to place the Zabargad intrusion in the framework of oceanic rifting. Two main episodes of rifting have been proposed [13,64]: An earlier episode of magmatic and tectonic activity occurred between 25 and 20 Ma, opening small basins aligned along a narrow line that pre-figured the future rift [41]. On Zabargad, the mantle was intruded and exposed during this episode. After a period of relative quiescence between 15 and 10 Ma, a change in plate motion during the early Pliocene initiated the oceanic opening of the Red Sea, involving continental stretching and generation of oceanic crust. Interestingly, spatial variation in the activity of the Kenya rift may illustrate a similar two-stage history. The central part of the rift, which is characterized by an approximately 60 km wide rift valley bordered by steep-sided flanks, where the narrow mantle intrusions have been recorded, is essentially a young structure (main activity between 16 and 1.6 Ma [54]). This could correspond to the first episode of the opening of the Red Sea. In contrast, the northern part of the rift, where the rifting activity is more widespread and where the crust seems to be attenuated, is approximately 10 Ma older. This area could correspond to the beginning of the second stage of the Red Sea

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rifling, just preceding the creation of oceanic crust. The Baikal rift may be contrasted with the Kenya rift (less volcanism but more tectonic and seismic activity) and is located in a very distinct geodynamic environment, where lithospheric stresses may be more influential [62]. However, its rifting may also be separated into an earlier and longer stage of minor extension followed ~ 10-20 Ma later by the rapid development of a second stage of faster rifting [62,63]. These data can be now incorporated into the model of two-stage rifting: The first step, of rift initiation, would consist in a few asthenospheric diapirs ascending through the cold brittle upper mantle, and dykes intruding the crust. This hypothesis is supported by the observations presented above in sections 2 and 3. Warming of the lithosphere would result from the ascent of asthenospheric diapirs reaching the bottom of the crust and of tholeiitic magmas being injected into the crust. This would create a localized weakness zone allowing the tectonic stretching of the second step. The second step, of rifting, would thus consist in homogeneous stretching of the lithosphere, which is now made possible by the 10-15 Ma of thermal relaxation following step 1. Although our knowledge of the rheology of the lithosphere is too poor to actually constrain quantitatively the mechanism, numerical models of the extension of a locally weakened lithosphere under distant pull [65-67] do lead to concordant results: Starting with the initial conditions of step 1, stretching begins by rapid subsidence in the rift axis, a strong uplift of the flanks, no doming, and an important necking of crust and subcrustal mantle above the thermally perturbed zone. This mechanism avoids a large uplift over a wide horizontal region which is not recorded in the Red Sea [68,69] or in several other rifts (see review in [39]). The necking of the entire lithosphere evolves rapidly to fracturing of the continental crust and expansion of the oceanic area. Analogue experiments of lithospheric rifting using a four-layer system [70] suggest that, with a chosen viscosity contrast between weak and strong layers, it is possible to obtain an intrusion of lower lithosphere and asthenosphere at very shal-

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low levels owing to rupture and boudinage of the strong layers (upper crust and mantle) and largescale flow of the weak lower crust. Because the strong layers are now scattered in boudins, it is expected that further stretching would directly result in large lithospheric attenuation over a widespread area, without the requirement of the delay imposed, in our case, by thermal relaxation. Although this possibility of purely mechanical rifting cannot be discarded, it should be observed that the use of a Newtonian theology and the absence of thermal effects in the analogue experiments favour a more distributed deformation. In particular, lateral cooling of the asthenospheric wedge within the lithosphere will tend to narrow its intruding head. This effect has been recently well illustrated by numerical modelling of mantle ascent at the slow spreading Mid-Atlantic Ridge [71]. At the same time, partial melting within the wedge will help to propel it to shallower depth owing to the increased difference in density with respect to the surrounding lithosphere, whereas melt extraction by dykes will locally weaken the overlying lithosphere. All these effects should help to localize the early intrusions, thus requiring an episode of regional weakening before the onset of homogeneous stretching.

4.2. R i f t initiation

We discuss here the problem of localizing and initiating a lithospheric fracture corresponding to our rifting step 1. In this respect, it is important to recall that in the Red Sea the 25-20 Ma episode of early rifting and magmatism may have been prepared by a still older magmatic activity [40,72]. This activity may be as old as Eocene (50 Ma) [69], but it is mainly recorded during the Oligocene and is important along the Ethiopian and Yemen margins of the southern Red Sea, where copious flood basalt volcanism has been dated at between 30 and 16 Ma [40]. North of this, in the areas of interest here, the volcanism is more of a fissural type. Several geochemical studies [73,21,79,38,40] have contrasted this early alkaline volcanism with the modern tholeitic vol-

canism of the axial trough and concluded that the source may have been asthenospheric with various degrees of contamination by the subcontinental lithosphere. It has also been suggested that this contribution may be due to thermal erosion and thinning of this lithosphere by a mantle plume located below Ethiopia and Yemen [74,75,38]. Thermal erosion of the lithosphere suggested by the study of early magmatic activity may be one way of localizing rifting. Another way may be the existence of a mechanically weak zone within the continental lithosphere. This has been suggested in the case of the Red Sea [76], whose opening would have been locally controlled by two N - S trending suture zones. A somewhat similar situation seems to exist along the Kenya rift, which locally departs from its dominant N - S trend to follow a series of older, Precambrian, parallel and N W - S E trending shear zones [54]. In both situations, the structural control seems to deflect only locally a direction of rift opening, which would be independently controlled.

4.3. Tensile strength of normal continental lithosphere With the concept of integrated horizontal yield stress of the lithosphere [77], which is dependent on its thermal and mineralogical structure as well as on its strain rate, it is possible to estimate the tensile strength of a 'normal' homogeneous continental lithosphere (i.e., with a 30-40 krn thick crust, and a surface heat flow ranging from 50 to 60 mW m-2). The deviatoric horizontal force per unit length (generated by density variations associated with ocean ridges and subducting slabs) applied to the lithosphere in tension has been estimated in the range 1012-1013 N m -1 [78,80]. If the applied lithospheric force (distant pull) is equal to the tensile strength of the lithosphere, for a geodynamical strain rate ranging from 10-15 to 10 14 s-~ rifting is initiated by stretching of the lithosphere. This is a mechanism of passive rifting [81]. All calculations using experimental rock mechanics results lead to similar conclusions

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[82-85]: This mechanism is impossible in a continental lithosphere with a 30-35 km thick crust, except if the temperature at the Moho level were to be around 700°C. In this latter case the 1100 ° and 1350°C isotherms lie at depths of around 55 and 65 km respectively, a situation of thin lithosphere. It seems thus impossible to initiate rifting by passive stretching in a normal continental lithosphere which is thicker than 55 km. The situation is obviously different for a lithosphere with an overthickened crust, because in that case the lithospheric strength is lowered, even with a normal geotherm. The model proposed here of a small asthenospheric wedge penetrating the lithosphere raises the question of its mechanical contribution to rifting. All models of buoyant mantle diapirism developed for oceanic ridges, even those implying the largest buoyancy [86], result in a pressure exerted against their walls that does not exceed a few megaPascals. This contribution to rifting is negligable. It should be quite clear that in this model of rifting the lithosphere must be in a state of tensile stress. 4.4. Thermomechanical aspects o f initiation and evolution o f rifting

As discussed above, it seems impossible to stretch or to fracture a normal homogeneous lithosphere by tectonic forces: the upper mantle is in that case too cold and the lithosphere strength too high. For stretching to begin or rupture to take place it is first necessary to warm up the mantle beneath the crust by around 150°C. Rift initiation in the studied cases thus requires first heating of the lower part of the lithosphere, resulting in heterogeneous thinning of the lithospheric mantle without stretching. Conductive heat tranfer in the lithospheric mantle above a mantle plume is a possible mechanism. Spohn and Schubert [87] have shown that heat supply by a plume may be high enough (5-10 times the normal mantle heat flow) to reduce by thermal erosion the lithosphere thickness by a factor 2 in less than 30 Ma. Braun and Beaumont [65] propose a thermomechanical mechanism in which

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sublithospheric convection-induced gravity push superimposed on distant pull mechanically erodes the mantle part of the lithosphere, with almost no crustal thinning, in less than 20 Ma. These two mechanisms are good candidates for explaining this step of lithosphere thinning without stretching, which is needed to weaken the lithosphere. But they predict [82] or call for [65] an initial, pre-rift, regional doming of some hundreds to a few thousand metres, which is much disputed in geological situations such as those prevailing in the Red Sea rift [68,69,88]. Contrasting situations may exist depending on the volcanic/non-volcanic nature of the rifted margins, doming being more important in the former, which could also be underlain by a large mantle plume [73]. In this respect, it may be appropriate to distinguish the southern Red Sea, where the copious flood volcanism of Yemen and Ethiopia has been related to the presence of a major mantle plume as early as 30 Ma ago [73], from the central and northern Red Sea, the two areas which are more specifically considered here. A solution may also be conductive thermal erosion by a convective heat supply only twice as large as the normal supply and maintained over a longer time (more than 100 Ma): In this case, Spohn and Schubert [87] calculated lithospheric thinning by a factor 2, with less than 300 m of uplift. The horizontal spread of the doming could also be reduced in the case of a narrow mantle plume, but then heating will be less efficient. During this step, partial melting should occur in the heated upper mantle and magmas should intrude the subcrustal mantle and crust. Taking this into account in the models would have increased the thermal weakening of the lithosphere. Such mechanisms lead to temperatures that are high enough in the mantle to decrease the strength of the lithosphere below the available tectonic forces, and make expansion possible. As the steady-state thermal profile is not reached this expansion must, at least in some cases, be accommodated by brittle fracturing of the lithosphere. The thermoelastic effect of an asthenospheric wedge modelled by McMullen and Mohraz [10] should contribute to the fracturing.

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5. Conclusion 5.1. Lithosphere rupture and stretching

We have presented evidence showing that the first stage of rifting in a continental lithosphere supposed to be structurally and thermally homogeneous was the consequence of rupture at the scale of the entire lithosphere. One piece of evidence deduced from seismic tomography in the Kenya Rift is the existence of a narrow, steep-walled zone of low velocity beneath the rift proper. This zone is interpreted as an asthenospheric wedge [89,53,58] extending down to close to the Moho. Another piece of evidence, this time deduced from field studies in rift-related peridotites (mainly Zabargad, Trinity and Lanzo), is the fact that the high-temperature flow was parallel to the rift trend and along steep surfaces (Fig. 1). This contrasts with the flat-lying frozen flow surfaces observed in most ophiolites [6] and suggests the intrusion of a hot and narrow mantle wedge into the lithosphere below rifts and slowspreading ridges {71]. The size of individual mantle wedges would be about 10-20 km across at their base, as deduced from the Massif Central results (Fig. 3), a size which, interestingly, coincides with that of mantle diapirs mapped in the Oman ophiolite [6]. In this model, and at this early stage, rifting is limited and it does not rely on tectonic stretching but on intrusion of new material, namely asthenospheric mantle at the base of the lithosphere and basaltic dykes above this asthenospheric wedge (Fig. 2A). The very large fraction (around 50%) of new basaltic material in the Zabargad gneisses [22,27] suggests that this addition to the crust may account for underplating but also for expansion, at least in the lower crust. Above, extension faulting may have operated [20], or possibly mechanical decoupling. Massive addition of new basaltic material to the crust before crustal splitting is also predicted by Lachenbruch et al. [90] in the case of the Salton trough, which in many respects may be a presentday equivalent of the Zabargad rift. Thermomechanical modelling also suggests that it is difficult to stretch plastically the top of the mantle until it is warmed 150°C above its

initial temperature. In the modelling of Massif Central rifting (Fig. 3), this is achieved after 10 Ma, by heat transfer from a few asthenosphere diapirs and by basaltic dykes which were injected above the melting diapirs. This model is illustrated by the case of the Red Sea, and possibly by that of the Kenya rift. In the Red Sea, the Zabargad peridotite intrusion may represent part of an asthenosphere diapir (Fig. 1); similar intrusions, still concealed, would be present below narrow basins such as Tihama Asir. These intrusions mark the initiation of rifting 20-25 Ma ago. It is only 10-15 Ma later (i.e., between 12 and 5 Ma) that the second stage of rifting begins, with faster extension now dominated by lithospheric stretching and necking. Once the continental lithosphere has been sufficiently stretched it grades into oceanic lithosphere (Fig. 2B). 5.2. Symmetric or asymmetric rifting

The description of narrow asthenosphere diapirs penetrating the lithosphere below a rift (Fig. 1) does not support, in the situations considered here, a 'simple shear' mechanism of rift opening with a major listric fault cutting through the entire lithosphere. In the simple shear model the exposed mantle is (1) the old infracontinental mantle composed of spinel lherzolites [16] and (2) its foliation should be found to be flat-lying and characterized by low-temperature flow structures. In contrast, the exposed mantle from environments related to rifts show, at least in their core, steep foliation active in a melting mantle equilibrated in the plagioclase lherzolite field, at depths shallower than 30 km. An oceanic rift environment is recorded in the case of Lanzo (Corsica, Western Alps) and Zabargad (Red Sea), and such an environment seems possible in the case of Trinity (California) and Alpujata (Southern Spain). This conclusion can be extended to the passive margin of Galicia where the early penetrative foliation in plagioclase lherzolites is also steep and related to melting [91]. From this field evidence we conclude that, in the environments considered, the opening was symmetric within the mantle. This does not preclude the existence of some asymmetry in the

A. Nicolas et al. / Earth and Planetary Science Letters 123 (1994) 281-298 crustal f e a t u r e s of the rift, i n d u c e d for instance by a d e t a c h m e n t fault levelling down within the crust or along the M o h o surface. 5.3. L o c u s o f rifting A m a j o r p r o b l e m r a i s e d by the R e d Sea a n d K e n y a rifts is t h a t of the locus of rifting and initiation. A h o m o g e n e o u s c o n t i n e n t a l lithos p h e r e s e e m s to b e too r e s i s t a n t to yield by fracturing, c o n s i d e r i n g the e s t i m a t e of forces a p p l i e d to t h e p l a t e s (see above), T h e location of the rifting can b e c o n t r o l l e d by a m e c h a n i c a l or thermal w e a k n e s s z o n e within the l i t h o s p h e r e [92]. In t h e case of an old a n d h o m o g e n e o u s c o n t i n e n t a l l i t h o s p h e r e , it w o u l d s e e m t h a t an e p i s o d e of l i t h o s p h e r i c t h i n n i n g by a hot m a n t l e p l u m e is necessary. This is the c o n c e p t of 'active' versus ' p a s s i v e ' rifting [76] or ' p l u m e ' versus ' i n t e r p l u m e ' rifting [60]. In active rifting, doming, which is the surficial m a n i f e s t a t i o n of l i t h o s p h e r i c thinning, p r e c e d e s the o p e n i n g of the rift, w h e r e a s in passive rifting it is t h e o t h e r way a r o u n d . Early d o m i n g has b e e n e n v i s a g e d for m a n y rifts, such as the R i o G r a n d e [92], t h e R h i n e g r a b e n [93] a n d the K e n y a rift [54,94], a l t h o u g h this question is still b e i n g deb a t e d [see 39]. In the case of t h e n o r t h e r n R e d Sea the r e c e n t l i t e r a t u r e c o n c l u d e s that the d o m ing p o s t d a t e s t h e stage 1 rifting [68,89], l e a d i n g to the conclusion that the R e d Sea was a ' p a s s i v e ' rift [69,95] a n d thus raising the p r o b l e m of the locus of t h e rifting. T h r e e possibilities can be envisaged: (1) T h e A f r i c a n - A r a b i a n p l a t e was not h o m o g e n e o u s a n d it has y i e l d e d along a w e a k n e s s zone, (2) t h e r e is no n e e d for any w e a k e n i n g , a n d t h e h o m o g e n e o u s l i t h o s p h e r e can yield anywhere, a n d (3) a n a r r o w t h e r m a l p l u m e has t h i n n e d the h o m o g e n e o u s l i t h o s p h e r e but has r e s u l t e d in no n o t i c e a b l e uplift. A l t h o u g h possibility (1) can p l a y a role, locally d e f l e c t i n g the rift t r e n d (as seen above), t h e r e is no i n d i c a t i o n of such a w e a k n e s s affecting the e n t i r e pre-rift lithos p h e r e . Possibility (2) has b e e n c o n s i d e r e d above to be unsatisfactory. W e t h e r e f o r e favour possibility (3), a l t h o u g h it m u s t be r e c o g n i z e d t h a t this is so far largely conjectural. In this case, g e n e r a t ing a rift w o u l d r e q u i r e b o t h an initial p l u m e

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h e a t i n g a n d a tensile stress a p p l i e d to the lithos p h e r e , as the p r e s s u r e e x e r t e d by the rising m a n t l e d i a p i r is negligable. If this c o n c e p t of rift g e n e r a t i o n can b e e x t e n d e d it will call into question the c o n c e p t of 'active' versus ' p a s s i v e ' rifting.

Acknowledegments This w o r k was s u p p o r t e d by the Institut National des Sciences de l'Unit,ers ( D B T c o n t r i b u t i o n 192).

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