Geo-Marine Letters (1997) 17 : 70—78 ( Springer-Verlag 1997

M. Guiraud · J. Benkhelil · J. Mascle · C. BasileG. Mascle · J. P. Bouillin · M. Cousin

Synrift to syntransform deformation along the Cote d’Ivoire–Ghanatransform margin: evidence from deep-sea dives

Received: 16 March 1995 / Revision received: 8 August 1995

Abstract Fourteen deep dives were conducted along thesouthern slope of the Cote d’Ivoire—Ghana marginalridge. Detrital rock samples reveal Early Cretaceoussynrift sediments. In situ observations and microtectonicanalysis identify the nature and the timing of three suc-cessive episodes of deformation. Clay layers are affectedby early diagenetic deformational structures such asslump microfolds and normal microfaults. Most of theclastic deposits display a second deformational episode.These structures result from a layer-parallel extension andare related to extensional synrift tectonics. The last stageof deformation is identified by right lateral wrench foldsassociated with brittle joints and slaty cleavage. It isattributed to the syntransform evolution of the Coted’Ivoire—Ghana marginal ridge.

Introduction

The Cote d’Ivoire—Ghana (CIG) transform margin is theproduct of continental break-up due to transform motionand subsequent opening of the south Atlantic Ocean dur-ing Early Cretaceous times (Mascle 1976; Mascle et al.

M. GuiraudUMR-CNRS no.5561 Dijon, Universite de Nancy, BP 239,54506 Vandoeuvre-Les-Nancy Cedex France

J. BenkhelilURA-CNRS no. 718, Universite de Perpignan, 52 avenue deVilleneuve, 66860 Perpignan Cedex France

J. MascleURA-CNRS no. 718, Laboratoire de Geodynamique sous-marine,BP 48, 06230 Villefranche/Mer, France

C. Basile · G. Mascle · J. P. BouillinURA-CNRS no. 69, Universite de Grenoble I, Institut Dolomieu,15 rue Maurice-Gignoux, 38031 Grenoble Cedex France

M. CousinLaboratoire de Geodynamique, Universite de Savoie,Campus technolac, BP 1104 Chambery Cedex, France

1988). One of the most prominent features of this marginis a NE—SW-trending 130-Km-long marginal ridge(Fig. 1). To the north, this ridge is bounded by a riftedbasin, the Deep Ivorian Basin, and to the south it towersover the oceanic abyssal plain by more than over 2500 m.According to Basile et al. (1993) the area has been subjectto three successive stages of deformation: (1) An earlyE—W rifting stage resulted in crustal thinning and asso-ciated block tilting of the Deep Ivorian Basin (Fig. 2A).(2) Before the end of rifting, strike-slip faulting beganalong the southern border of the Ivorian Basin. After therifting, strike-slip faulting and slight folding were re-stricted to an intracontinental transform fault (Fig. 2B).(3) This transform fault became a continent—ocean trans-form fault when oceanic crust was formed to the south ofthe margin. Thermal exchanges between the two litho-spheres resulted in marginal ridge uplift. Finally, tectonicactivity ceased after the oceanic spreading center hadpassed south of the transform margin. After this the pass-ive continental margin subsided by passive thermal sub-sidence.

Aim of the study

As transform margin studies were chiefly based on marinegeophysical data (Arens et al. 1971; Delteil et al. 1974;Scrutton 1976; Mascle and Blarez 1987; Eldholm et al.1987; Keen et al. 1990; Lorenzo et al. 1991; Basile et al.1993), little is known about their tectonic and microtectonicprocesses. These were the target of the Equanaute cruisein June 1992, during which 14 dives were performed alongthe southern slope of the CIG marginal ridge using thesubmersible Nautile. Seventy-five kilometers of geologicalsection were surveyed between water depths 2211 and4905 m, and 165 rock samples were retrieved chiefly fromEarly Cretaceous synrift sediments (Mascle et al. 1993)(Fig. 3). The sedimentary sequence is made of three mainfacies corresponding to dark clayey layers, interbedded

Fig. 1 Simplified Seabeambathymetric map (contours inmeters) of the Coted@Ivoire—Ghana marginal ridge.Locations of the Equanaute divesconducted along the southernsteep slope are shown

Fig. 2 Structural evolution of the Deep Ivorian Basin inferred fromseismic data (location on Fig. 1) (Basile et al. 1993). A: Rift stage andN—S-trending normal faulting: B: Dextral shear, induced folding,block rotation and structural inversion

with slightly tilted beds of massive siltstones or sand-stones, (Guiraud et al. this volume). Few rocks of facies1 were recovered; they consist of black organic-rich shalesand brown clays, and they usually display diffuse bandingthat constitutes a primary fabric (Fig. 4: 1 and 2). Facies2 makes up 60% of the sampled material; it correspondsto alternations a few millimeter to 1 cm thick of wavy-bedded yellowish clays and siltstones interlayered withlenticular-bedded to flasered-bedded greenish fine sand-stones (Fig. 4: 3, 5, 6, and 8). The fine to coarse quartz andcalcareous sandstones of facies 3 are yellowish to greycolored (Fig. 4: 4 and 9). These deposits may display basallayers of microconglomerates only a few millimeter thickthat grade up into diffuse crossbedded and massive sand-stone beds. Both in situ observations and sample analysisprovided a wide range of tectonic microstructures (micro-faults, veins, microfolds, joints, and cleavages). No dataare available on the absolute chronology of the successivedeformations, but several stages were identified through

consideration of sedimentary and advancing diagenesisprocesses.

Prelithification deformation

Many microfolds were observed in clay layers. Folds aredisharmonic, isoclinal (Fig. 4:1 and 2), and with horizon-tal or parallel to bedding as well as curvilinear hinges.They may take the shape of recumbent sheath folds. Sharpthickening, thinning, and boudinage of the layers suggestsfor these features a highly ductile mode of deformation,occurring in unlithified muddy sediments with a highwater content. Some flat-lying microfolds show late,parallel-to-bedding, mineral fabric. This demonstratesthat compaction affected sediments that had alreadyundergone some kind of folding. The style of deformationindicates slumping of soft sediments.

Slump microfolds appear to be cross-cut by slidingalong bedding planes. Lustrous slicken-sided beddingplanes are often glossy, probably because of hydroplasticdeformation in soft sediment (Guiraud and Seguret 1987).The same kind of microfaults resulted from shear testsmade on clays with water content exceeding 50%(Maltman 1987).

Other structures indicating deformation in soft sedi-ment include load casts and associated flames. Mud pipesare almost orthogonal to bedding and are injected intosandstone and siltstone beds. They exhibit a fan-shapedhydroschistosity, indicated by zones of aligned phyllosili-cates, induced by rotation of clay particles during dew-atering and upward expulsion of mud.

Finally, the latest tectonic features can be related toprelithification deformation. These are listric normalmicrofaults, dipping 30° in their upper part, and rootedinto bedding planes (a in Fig. 4:3 and 6). The amount ofdip slip motion ranges from 0.5 to 1 cm. Sedimentarylayers thicken along the hanging wall, where frequentprominent rollovers into the shear planes are observed.Those faults are never mineralized, while later ones are

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Fig. 3 Synthesis of structural and microtectonic observations madeduring Equanaute dive EN6 (location of Fig. 1): stereographic pro-jections of bedding planes (So) and plotting of the joint directions onrose diagrams

syn- to postlithification faults (see below). In thin sections,the fault zones are 0.5 mm thick, without disruption butwith shearing and particle redistribution of soft sedimen-tary layers. Upward, the fault plane ends in a chaotic,entirely liquefied silty and clayey sediment in (Fig. 4:3b).According to Ringrose (1988), similar liquefaction-induced microstructures may be due to the triggering ofa water-rich sediment during earthquake shocks. Bothslumps and the oldest set of microfaults occurred in softsediment. Both can be related either to gravitational de-tachment, to liquefaction of the sediment as a result ofpalaeoseismic shocks, or to overloading.

Syn- and postdiagenesis extension

Numerous types of extensional microstructures have beenrecognized on slabbed hand-samples and thin sections.They include normal faults and two main sets of min-

eralized veins that can be distinguished from the shapes oftheir walls and infilling.

In the oldest set of extensional features, tensile quartzveins are characterized by very angular and indented walls(Fig. 4: 4a and b). In thin sections, the veins exhibit suture-like contacts that closely follow but do not truncate thesedimentary quartz grains (Fig. 4:7a; Fig. 5). Moreover,the vein infillings show euhedral radial quartz crystals anda few angular detrital grains (Fig. 4: 7b). These appear tobe fragments of sediment that mixed together with thevein fluids as the vein formed. These observations suggestthat this kind of vein originated from grain-to-grain separ-ation when the sedimentary matrix was weaker than theframework grains, that is, during diagenesis.

The shape of these synlithification veins changes withthe sediments. Clayey and silty samples exhibit con-jugated sets of fractures associated with en-echelon arraysof quartz-filled veins (Fig. 4:6b), while the sandstones dis-play veins oriented normal to bedding (Fig. 4: 4a and b;Fig. 4:5b). A few normal microfaults are associated withthe synlithification veins. They dip at 50—60° with respectto the layering and show millimeter to centimeter offsets(Fig. 4: 5a, Fig. 4:8). All these extensional features can berelated to extension parallel to bedding, probably when itwas horizontal. Moreover, two samples (EN6-2, EN7-1,

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Fig. 4 1—7 See page 75 for legend

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Fig. 4 8—13 See page 75 for legend

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bFig. 4 1: isoclinal and disharmonic slump microfolds showing signsof disaggregation (Al on Fig. 8), (clays, sample 3, EN11 dive, waterdepth 4879 m). 2: Rounded slump microfolds (Al on Fig. 8). Beneaththe microfolds are decollement surfaces corresponding to beddingsliding planes (clays, sample 7, EN8 dive, water depth 3711 m).3: Slabbed hand-sample showing listric normal microfaults (a) pas-sing upwards into a chaotic entirely liquefied sediment (b) (Fig. 8A3—4, clays and silts, sample 7, EN14 dive, water depth 3169 m).4: Slabbed hand-sample showing syndiagenesis extensional branchquartz veins (b and Fig. 8 B6). The contacts of the veins are irregularand some dark mica-rich layers (bedding So) cross-cut the earlyveins (a) (medium to fine sandstones, sample 6, EN12 dive, waterdepth 3183 m). 5: Pervasive normal syndiagenesis microfaults (a)disrupting unconsolidated laminated silts and clays (Fig. 8 B9). Themicrofault surfaces are associated with early en echelon quartz veins(b and Fig. 8 B8) (laminated silstones, sample 1 EN7 dive, waterdepth 4396 m). 6: Slabbed hand-sample showing listric normalmicrofaults (Fig. 8 A3) and en-echelon quartz veins in normal shearzones (Fig. 8 B8). The extensional microstructures were generatedwhen the bedding was horizontal (greenish fine sandstones andsiltstones with lenticular bedding, sample 2, EN6 dive, water depth3937 m). 7: Photomicrograph of syndiagenesis extensional quartzveins (Fig. 4,6b). The veins (a) do not truncate or deform the singlegrains and include detrital quartz grains (b) and mica-rich layers (c),(fine silstones: sample 2, EN6 dive, water depth 3937 m, natural light,scale: box length approximatley 3.3 mm across). 8: Normal syn-diagenesis microfault surface is associated with early en-echelon andnormal to bedding quartz veins (sample 15, EN5 dive, water depth2601 m). 9: Straight calcite veins affecting a calcitic cement sand-stone. The brittle extensional veins are normal to bedding (a andFig. 8 C11) or strike at 75° to the bedding So (branch curving veinsb), (sample 1, dive EN4, water depth 3401 m). 10: Photomicrographof a normal to bedding, brittle vein (Fig. 8 C11). The contact be-tween the quartz vein and fine sandstone is straight and the veintruncates and deforms the single grains. These observations suggestgrain breaking during postdiagenesis deformation (sample 8, EN11dive, water depth 4685 m, natural light, scale: box length approxi-matley 14 mm across). 11: Sandstone beds fractured by late brittlejoints and folded into an asymmetric 2-m-wide anticline (EN7 dive,water depth 3012 m). 12: Large-scale trough cross-bedded sandybar overlain by beds tilted 60—70° to the south. The layers arefractured by N110°E striking joints (EN13 dive, water depth4513 m). 13: Laminated fine siltstones affected by brittle joints (EN7dive, water depth 4000 m)

Fig. 5 Line drawing ofa syndiagenetic quartz vein(sample 2, dive EN6). The contactbetween the extensional vein andthe fine sandstone is irregular andsuturelike and suggests non-grainbreaking

Fig. 4:5 and 6) were oriented during the dives, and thestrike of the faults and veins appears to be parallel toN—S-trending normal faults seen to the north on seismiclines (Basile et al. 1993).

Extensional veins from the second set are in sharp,straight contacts with the host rock (Fig. 4:9 and 10;Fig. 6). These rather clean veins contain calcite or quartzcrystals orthogonal to the tension-gash walls, indicatinga layer-parallel extension. The vein-filling crystals displaya characteristic fibrous fabric (Ramsay 1980), which couldbe related to the opening of successive microcracks andsubsequent instantaneous sealing. Straight walls, the ab-sence of detrital inclusions, as well as the fibrous texturestrongly suggest that the material was highly cohesiveduring vein generation, that is, after sediment diagenesis.

Most of these late veins are normal to bedding(Fig. 4: 9a), but the sediments are also cut by branchcurving veins (Fig. 4:9b). Moreover, in calcite-cementedsandstones, some stylolitic joints may occur along bed-ding planes; in this case the stylolitic columns are normalto sedimentary layering (Fig. 6). Thus, as for the synlithifi-cation veins, the postlithification veins are related to ex-tension parallel to bedding.

Late wrench deformation

During many of the dives, concentric microfolds havebeen found with subvertical axial planes and N—S toN60E trends (Fig. 7; Fig. 4: 11) (Mascle et al. 1993). More-over, a spaced fracture cleavage commonly occurs in tiltedsandstones (Fig. 4: 12). Those brittle joints are more or lessperpendicular to the bedding (Fig. 4: 13). At least in someareas, tilted beds and probably cleavages can be related to

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Fig. 6 Line drawing of lateextensional calcite-filled veinsand bedding parallel styloliticjoints (sample 1, dive EN4). Thebrittle veins break and displacethe detrital quartz grains,suggesting post diagenetic andpressure solution-relateddeformation

Fig. 7 Stereographic projection of the microfold and fold axis ob-served respectively during Equanaute dives and deduced from bed-ding analysis

folds with the same axial planes and trends as the micro-folds (Fig. 3). Associated with folding, a penetrative slatycleavage was commonly observed in the clayey beds. Inthin section, this cleavage is characterized by recrystal-lized phyllosilicates (probably of low-grade metamorphicgrade) and is seen to cross-cut the earlier extensionalstructures (Benkhelil personal communication 1995). Webelieve that those structures were associated with wrenchfolding, as the N—S to N60E azimuths of fold axes are ingood agreement with a dextral displacement along theN60-trending CIG transform margin (Mascle 1976; Basileet al. 1993). This does not exclude brittle joints from beingat least partly related to strike-slip faulting or later brittledeformation, and that part of the tilting is related to

a subsequent uplift of the CIG marginal ridge (Basile et al.1993).

Conclusions

Microstructural analysis of the numerous rocks sampledduring the Equanaute cruise greatly improve our know-ledge of the deformation history of the Coted@Ivoire—Ghana transform margin (Fig. 8). The whole his-tory has been recorded by clastics sediments coeval withthe rifting of the Ivorian Basin. The successive timing anddiagenetic relationships between the different, well-identi-fied, microstructures reveal a three-stage deformationhistory:

1. Earliest prelithification microfolds and hydroplasticshear fractures are thought to result from gravity-inducedslumping and faulting in a shallow, unstable, deltaic envi-ronment (Fig. 8A) (Guiraud et al. this volume). The pre-lithification phase can be related to the Early Cretaceousrifting of the Ivorian Basin.

2. Synlithification (Fig. 8B) and postlithification(Fig. 8C) quartz veins, as well as normal microfaults, pred-ate a folding episode. Since the strike of these structures isN—S parallel to the main normal faults of the IvorianBasin, this stage is believed to record the synrift deforma-tion during diagenesis.

3. Finally (Fig. 8D), late tilting, brittle joints, slatycleavage, and folding can be genetically related to thegeneral dextral, N60-trending, faulting that was prevailingalong the marginal ridge during the transform motionalong the CIG margin.

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Fig. 8 Summary of thedeformation history of the CIGtransform margin. Thecharacteristic sediment sequenceconsists of laminated silts andshales interbedded withsandstones and rare gravel casts.A: Prelithification liquefaction-induced microstructures: type 1,tight to disharmonic isoclinalslump microfold and associatedsliding bedding planes; type 2,mud pipe with fan-shapedhydroschistosity (S1); type 3,listric normal microfaults; type 4,entirely liquefied sediment.B: Syndiagenesis extensionalmicrostructures: type 5, irregularnormal to bedding quartz vein;type 6, indented branch quartzveins; type 7, mica-rich layer; type8, en-echelon quartz veins innormal shear zone; type 9,pervasive normal hydroplasticmicrofaults; type 10, normal tobedding clastic microdike.C: Postdiagenesis extensionalmicrostructures: type 11, normalto bedding straight vein; type 12,brittle branch curving vein, type13, en-echelon veins in normalshear zone; type 14, parallel tobedding stylolitic joint. D: Latebrittle wrench microstructures:type 15, vertical brittle joint; type16, vertical to highly dippedstylolitic joint; type 17, horizontalquartz vein; type 18, brittlereverse microfault?; type 19, slatycleavage S2; type 20, tiltedextensional microstructures

Acknowledgments The captain and crew of the R/» Nadir and theteam of the submersible Nautile are thanked for their assistanceduring the Equanaute cruise. We are grateful to the editorial refereesE. J. W. Jones, P. Clift, and A. H. Bouma for their comments.

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