continental break-up along strike-slip fault zones ... · strong dextral strike-slip-related...

Continental break-up along strike-slip fault zones; observations from

the Equatorial Atlantic

M. NEMCOK1,2*, A. HENK3, R. ALLEN1, P. J. SIKORA1,4 & C. STUART1

1Energy & Geoscience Institute, University of Utah, 423 Wakara Way, Suite 300,

Salt Lake City, UT 84108, USA2EGI Laboratory at GU SAV, Dubravska cesta 9, 840 05 Bratislava, Slovakia

3Institut fur Geowissenschaften, Universitat Freiburg, Albertstrasse 23b,

D-79104 Freiburg, Germany4Paleontological Subcontracting L.L.C., 10075 Turquoise Circle, Sandy, UT 84094, USA

*Corresponding author (e-mail: [email protected])

Abstract: The study focuses on Equatorial Atlantic margins, and draws from seismic, well, gravi-metric and magnetic data combined with thermo-mechanical numerical modelling.Our data and numerical modelling indicates that early drift along strike-slip-originated margins

is frequently characterized by up to 108–208 spreading vector adjustments. In combination with thewarm, thinned crust of the continental margin, these adjustments control localized transpression.Our observations indicate that early-drift margin slopes are too steep to hold sedimentary cover,

which results in their inability to develop a moderately steep slope undergoing cycles of gravita-tional instability resulting in cyclic gravity gliding. These slopes either never develop such con-ditions or gain them at later development stages.Our modelling suggests that the continental margin undergoing strike-slip-controlled break-up

experiences warming due to thinning along pull-apart basin systems. Pull-apart basins eventuallydevelop sea-floor spreading ridges. Margins bounded by strike-slip faults located among pull-apartbasins with these ridges first undergo cooling. However, spreading ridges leaving the break-uptrace along its strike eventually pass by these cooling margins, warming them again before thefinal cooling proceeds. As a result, the structural highs surrounded by several source rock kitchenswitness a sequential expulsion onset in different kitchens along the trajectory of spreading ridges.

Supplementary material: Discussion of the methods used, chronostratigraphic results and strike-slip margin characteristics are available at http://www.geolsoc.org.uk/SUP18518

Early development stages of the Equatorial Atlantic(Pierce et al. 1996; De Matos 2000) (Fig. 1) indicateinteresting structural and thermal consequences ofrift–drift transition in areas controlled by significantstrike-slip movement components on top of theregional extension, which makes thermo-dynamichistories of these settings very different from thosewe know of from literature on classic passivemargins (see e.g. Royden & Keen 1980; Steckler1981, 1985; Alvarez et al. 1984; Keen 1985;Morgan et al.1985;Buck1986;Buck et al.1988; vanBalen et al. 1995; van der Beek et al. 1995; Burov &Cloetingh, 1997). These consequences will beaddressed by numerical modelling applied to theGhana Ridge (Fig. 2) using the finite-element pro-gram ANSYSw (Ansys Inc., Houston, TX, USA).Subsequently, modelling results are discussed incombination with reflection seismic, well and gravi-metric data. Our approach is to apply models tovarious stages of the Ghana Ridge developmentduring the late-rift and early-drift time periods,

focusing on thermal regime perturbation and verticalmovement history. Models have been generatedfrom published transects based on different typesof geophysical data (e.g. Edwards et al. 1997; Sageet al. 2000). Modelling results provide temperatureand uplift histories during a time interval of 104–68 Ma and a snap-shot situation at 78 Ma. Modelsare designed to provide regional-scale results.

Geological setting

Equatorial Atlantic setting

The initial propagation of intracontinental strike-slip zones in the Equatorial Atlantic region(Fig. 1) started from east to west under progressivelyexpanding transtensional conditions (e.g. Szatmari2000). The onset is of late Barremian age (e.g.Popoff 1988; De Matos 1992; Guiraud & Maurin1992), based on the age of the basal transgressivesediments in the Benue Trough, and the ages of later

From: Mohriak, W. U., Danforth, A., Post, P. J., Brown, D. E., Tari, G. C., Nemcok, M. & Sinha, S. T. (eds) 2012.Conjugate Divergent Margins. Geological Society, London, Special Publications, 369,http://dx.doi.org/10.1144/SP369.8 # The Geological Society of London 2012. Publishing disclaimer:www.geolsoc.org.uk/pub_ethics

10.1144/SP369.8Geological Society, London, Special Publications published online March 27, 2012 as doi:

syn-rift sediments on the Brazilian continent and inthe Ivory Coast coastal area. The westward young-ing of basal sediments can be demonstrated on theBrazilian side of the system. The Potiguar Basindoes not have a fine enough stratigraphic resolu-tion to indicate the exact time within the Aptianstage but it does indicate that Aptian sedimentsrest on Neocomian sediments transgressively (DeAzevedo 1991). The oldest sediments of theMundausub-basin of the Ceara Basin have an Early Aptian

age (Costa et al. 1989). Studies available from theBarreirinhas Basin, Gurupı graben system andPara Maranhao Basin do not have fine resolutionand, therefore, they only document the Aptian ageof basal sediments (Cainelli et al. 1986; DeAzevedo 1991). The westernmost basins, theMarajo and Foz do Amazonas basins, have basalsediments as old as the Aptian–Albian boundaryand the Late Aptian–Albian, respectively (DeAzevedo 1991). Therefore, we can interpret that

Fig. 1. Closing match of Brazilian and African oceanic–continental crustal boundaries restored for different times inthe 105–110 (Albian) interval for different segments of the Equatorial Atlantic. Geological maps are taken fromCGMW (1990, 2000). Although the data, such as the mixed fauna, indicate a full marine connection through the regionsince the middle–late Albian (Koutsoukos 1992), restoration does not indicate a broad and deep-water connection at thetime of the continental break-up. The black and blue interrupted lines show a location for the oceanic–continentalcrustal boundary for the Brazilian and African sides, respectively. Grey and yellow areas along the fit indicate areas ofunderlap and overlap in reconstruction. Note that the strike-slip faults causing the future break-up in the EquatorialAtlantic region with an overall dextral strike-slip displacement have P-shear geometry. This is in contrast to analoguematerial models, which start first with the propagation of R-shears (see e.g. Cloos 1928; Riedel 1929; Hills 1963;Tchalenko 1970). Abbreviations: GP, Guinea Plateau; M, Monrovia Basin; IC, offshore Ivory Coast; ND, NigerDelta; BT, Benue Trough; RM, Rio Muni Basin; G, offshore Gabon; DP, Demerara Plateau; FA, Foz do AmazonasBasin; P-M, Para-Maranhao Basin; B, Barreirinhas Basin; C, Ceara Basin; P, Potiguar Basin; PP, Pernambuco ParaibaBasin; S-A, Sergipe–Alagoas Basin.

M. NEMCOK ET AL.10.1144/SP369.8

Geological Society, London, Special Publications published online March 27, 2012 as doi:

the westward propagation of rifting in the futureEquatorial Atlantic, controlled by overall dextraltranstension, took place over an approximately10–11 Ma time period.

The tectonic stresses created a right-steppingsystem of ENE–WSW dextral strike-slip faultsand pull-apart basins between them (Fig. 1).P-shear orientation of dextral faults instead ofR-shear orientation (sensu Christie-Blick & Biddle1985; Naylor et al. 1986) indicates that eitherthey were not initiated by transtensional stress orthat their positions were affected by prominentcrustal weaknesses of NE–SW strikes caused byProterozoic and Palaeozoic orogenies (e.g. Mascleet al. 1988; Genik 1992; Guiraud & Maurin 1992).The Late Cretaceous reactivation of the NE–SW-striking Sobral–Pedro II Fault in the Borboremaprovince to the south of the Barreirinhas Basin(Miranda et al. 1986) serves as support for the latterargument. Similar dextral reactivation of NNE–SSW-striking Pan-African shear zones was reportedfrom Nigeria (Ball 1980; Caby 1989).

Transition from rifting to drifting took placeduring the Albian, displaying continental–marineenvironments in equatorial pull-apart basins (e.g.Potiguar and Barreirinhas) or small basins in com-pressional quadrants near tips of boundary strike-slip faults (e.g. Piauı and Camocim basins)(De Matos 2000). Mixed fauna confirms the onsetof the oceanic communication through the Equa-torial Atlantic since the middle–late Albian(Koutsoukos 1992).

As extension in pull-apart basins led to theorigin of the oceanic crust in their centres, transformfaults started to form, linking spreading centres, anddefining the break-up between Africa and SouthAmerica during the Albian–Cenomanian period(De Matos 2000) (Fig. 1).

As the break-up zone between Africa and Brazilbroadened, newly formed transform faults weredeveloped.Active transform faultsmigrated towards

new active spreading centres, and some older basinportions became inactive and sealed by youngersediments (De Matos 2000). This phenomenonresulted in the very complex and dissimilar deposi-tional and thermal histories of various local basins.Spreading centres travelling oceanwards alongridges of the continental margin (Fig. 2) causedlocal heating and uplift of these ridges (e.g. Pierceet al. 1996), indicated by the oceanward youngingof the associated unconformity.

Spreading directions during the Albian–Cenomanian, affected by a pre-existing block struc-ture of the African and Brazilian continental crusts,were striking from ENE to WSW, as interpretedfrom the magnetic anomalies related to sea-floorspreading (Fig. 3) (Muller et al. (1997). Later, thisearly NE–SW divergence progressively changedinto an almost east–west divergence. It can berecognized through a change in flowline directionsreflected in fracture-zone geometry in the datafrom Muller et al. (1997) (Fig. 3). The timing of amajor portion of this change is in the Coniacian–Maastrichtian (89–65 Ma) (Tables 1 & 2). Tables1 & 2 show that the change in drifting direc-tion took part in two main events; the Coniacian–Maastrichtian and theDanian–Thanetian, the formerbeing more important.

Seismic imaging constrained by well data, whichwill be discussed later, indicates that the Romanchefracture zone itself experienced two phases of trans-pressionally driven deformation: 108–92 Ma and65–52 Ma. As shown by modelling, discussedlater, the former event could have been controlledby a 98 change in divergence direction, whichwould generate a convergence component of2970 m per 1 Ma (m Ma21) across the RomancheFracture Zone. A close observation of Figure 3shows that a 98 change in trace of the RomancheFracture Zone is adjacent to the Ghana Ridge.This segment of the continental margin will bedescribed in the following section.

Fig. 2. The Ghana Ridge area showing the locations of sea-floor spreading ridges in time (Ma) on their relativemovement away from contact with the broken-up continent. The youngest age shown is 72 Ma. Locations (dark greylines) were derived from the map of the digital ages of oceanic crust produced by Muller et al. (1997). Green lines,oceanic fracture zones, red lines, continental faults. Spreading centres migrating oceanwards along ridges of thecontinental margin caused the local heating and uplift of these ridges (e.g. Pierce et al. 1996), indicated by theoceanward younging of the associated unconformity. The Ghana Ridge itself is in the rectangle labelled as themodel area.

BREAK-UP ALONG STRIKE-SLIP FAULT ZONES10.1144/SP369.8


Fig.3.Map

oftheEquatorialAtlanticoceaniccrustagebased

onmagneticanomalydataandthegeologyofadjacentcontinents(m

apsarefrom

CGMW

1990,2000;Muller

etal.

1997).FortheabbreviationsseeFigure

1.



Ghana Ridge setting

Before we enter the discussion of our data andnumerical models, wewill describe theGhana Ridgedevelopment based on available literature (seeMascle & Blarez 1987; Basile 1990; Pontoiseet al. 1990; de Caprona 1992; Basile et al. 1993,1998; Mascle et al. 1995, 1996, 1998; Benkhelilet al. 1998a, b; Clift&Lorenzo 1999;Bigot-Cormieret al. 2005). Although these models have someminor differences, they agree in their major points.

The first sediments in the region were probablydeposited during the Early Cretaceous, in an intra-continental rift setting. It is not clear when exactlythis rift developed because the oldest sedimentshave not been penetrated by wells, and only athick sequence of later syn-rift continental clasticand lacustrine sediments deposited in the Aptianwas penetrated. Sampling along the steep slope onthe southern margin of the Ghana Ridge documentsAlbian–Cenomanian sediments at least 3 km thick(Mascle et al. 1993, 1996). Benkhelil et al. (1998b)indicated that the samples from the slope are more

proximal than the sediments in the wells to thenorth, possibly suggesting that the sediments werederived, in part, from the southern block of theCentral African Fault Zone.

The Albian–Cenomanian sequence underwentstrong dextral strike-slip-related deformation alongthe southern side of the Ghana Ridge, in a zoneabout 40 km wide (see Basile et al. 1998; Benkhelilet al. 1998b). Intense deformation was also devel-oped over the crest of the ridge. A shear zonecharacterized by a semi-chaotic reflection seismicimage marks the zone with the maximum defor-mation (Benkhelil et al. 1998b). North of thiszone, a series of dextral strike-slip faults were ident-ified. Together, the shear zone and the zone ofdextral strike-slip faults represent the damage zonealong the transform. These structures are restrictedto lie under a widespread unconformity. Sedimentsabove the unconformity, undeformed by shearzone and dextral strike-slip faults, have been datedas Cenomanian–Turonian or younger.

The timing of the end of intracontinentalwrenching varies somewhat in the literature. Clift

Table 2. Timing of the change of divergence trajectories from NE–SW to east–west indicated by the geometryof the oceanic fracture zones

Stage Beginning (Ma) End (Ma) Intermediate rotation Younger rotation

Coniacian 89 86 RomancheChainAscension

Santonian 86 83.5 RomancheChainNo nameAscension

Campanian 83.5 71.5 ChainNo nameAscension

Maastrichtian 71.5 AscensionDanian 65 61 RomancheThanetian 61 52 Romanche

ChainNo name

Table 1. Changes of the divergence during the activity of oceanic fractures zones (measured from a map ofdigital isochrones of the sea floor made by Muller et al. 1997)

Oceanic fracturezone

Initial strike(8)

Intermediate strike(8)

Age of rotation(Ma)

Present-day strike(8)

Age of rotation(Ma)

St Paul 70 70 78Romanche 60 71 88–84 78 64–54Chain 54 76 88–76 79 60–52No name 52 74 84–80 78 60Ascension 54 78 88–68 78



& Lorenzo (1999) interpreted that wrenching endedat 105 Ma, representing middle Albian, in the vicin-ity of sites 959 and 960; being half-way between riftinitiation at about 120 Ma and spreading ridgepassage at about 90 Ma. Tectonized Aptian–Albiansediments deposited during wrenching predate anyuplift recorded by erosional unconformity (Cliftet al. 1998; Clift & Lorenzo 1999). Bigot-Cormieret al. (2005) dated the timing of this phase as span-ning until the Albian–Cenomanian, representing105–95 Ma, when oceanic crust formed to thewest of the present-day Deep Ivorian Basin. TheDeep Ivorian Basin started with lacustrine sedi-ments in its individual pull-apart basins, eventuallyoverlain by northward-prograding marine sedi-ments, coming from the Brazilian side (Blarez &Mascle 1988; Lamarche et al. 1997). These sedi-ments are beneath the unconformity along the north-ern slope and crop out on the southern slope of theGhana Ridge. The youngest ages of this sedimentarysection that have proven biostratigraphically fallinto the late Albian (Bigot-Cormier et al. 2005).Clift et al. (1998) understood the end of the intra-continental wrenching phase to be synchronouswith the inversion that took part between the lateAlbian and early Turonian, dating it on the basisof the ages of the oldest post-tectonic sedimentsand youngest tectonized sediments.

Development models suggest that, betweenthe Cenomanian and the Turonian, the GhanaRidge was marked by an active transform margin;that is, during the period between the onset of orga-nized sea-floor spreading between South Americaand Africa, and the passing of the spreading centre(Basile et al. 1998; Benkhelil et al. 1998b). Asequence of carbonates was formed during theTuronian–Coniacian around the crest of the GhanaRidge, migrating up to the crest of ridge. Basile et al.(1998) suggested that the abrupt end of carbonatedeposition in the Coniacian was followed by rapidsubsidence of the ridge that marked the passing ofthe spreading centre.

The approximate locations of the oceanicspreading centre and the oceanic crust just southof the Romanche and St Paul transform faults canbe seen in Figure 2, based on the oceanic ages fromMuller et al. (1997) (Fig. 3). The age data showthat the first age contour along the margin was at108 Ma or in the Albian, probably just after thefinal rifting between South America and Africa,or at the time of the break-up unconformity. The108 Ma contour is located just west of the transi-tional crust, south of the two transform faults.

The detailed age of the early break-up of the con-tinent cannot be shown based on the magneticanomalies. This is because the magnetic anomaliesdeveloped during the break-up fall into the mid-Cretaceous magnetically quiet zone. South of the

St Paul Fracture Zone, Figure 2 shows that the con-tinental crusts of Africa and South America movedaway from being adjacent some time during 100–104 Ma (late Albian), and that the spreadingcentre passed the African crust at about 96 Ma(Cenomanian). South of the Romanche FractureZone, the continental crusts of Africa and SouthAmerica cleared in about 94 Ma (Cenomanian),and the spreading centre passed the African crustat about 76 Ma (Campanian). These ages for theRomanche Transform Fault are similar to theages mentioned in the ODP Leg 159 project-related publications, based on other types of datasuch as cooling ages derived from the apatitefission tracks and erosional unconformity pene-trated by sites 959 and 960 (see Basile et al. 1998;Bouillin et al. 1998; Clift et al. 1998; Clift &Lorenzo 1999; Bigot-Cormier et al. 2005). The pri-mary transcurrent deformation in the Ghana Ridgewas active in the Albian, around the time betweenthe first drifting along the margin, and the timewhen the African and South American crusts wereno longer adjacent. The crest of the Ghana Ridgestarted to subside rapidly in Coniacian–Santoniantime, just before the spreading centre passed theGhana Ridge.

Methods

A synthetic interpretation approach has been usedto define the crustal architecture, combining well,reflection seismic and gravity data. Well datacame from 94 industrial wells, 11 DSDP and 21ODP sites. They constrained the seismic interpret-ation, which was carried out on: 587 km of TGSNOPEC profiles in offshore Liberia, Togo andBenin; 2289 km of Western Geophysical profilesin offshore Ivory Coast and Ghana; 7239 km ofPROBE profiles in offshore Nigeria, EquatorialGuinea, Cameroon and Gabon; and 20 996 km ofLEPLAC profiles covering the Brazilian offshoreside of the Equatorial Atlantic region. The charac-teristics of individual surveys are given in Sup-plementary publication SUP18518.

The interpretation of gravity data, used forthe extrapolation of interpreted boundaries amongreflection seismic profiles, was performed onfree air data, which had onshore covered by theEGM96 dataset and offshore by 30 minute resol-ution data from Sandwell (1992, version 9.2). TheBouguer gravity anomaly and isostatic residualgravity anomaly maps were calculated from freeair data.

The palaeontological record of ODP Leg 159sites 959, 960, 961 and 962 located at GhanaRidge and undisclosed industrial well in the DeepIvorian Basin was processed by the composite



standard technique discussed in Supplementarypublication SUP18518.

Numerical modelling of the Ghana Ridgedevelopment was made using the finite-elementprogramANSYSw (Ansys Inc., Houston, TX,USA).It focused on a quantitative evaluation of the causesfor the uplift of the ridge. The tested alternativesincluded: (1) thermal uplift due to the lateral heatflow from the passing-by spreading centre; and (2)the uplift due to transpressional movements alongthe Romanche Fracture Zone.

The aim of the modelling included the predictionof the timing and spatial extent of the Ghana Ridgeuplift. The model had to prescribe various geody-namic positions of the Ghana Ridge, including:

† rifting of continental lithosphere with the GhanaRidge located at a dextral strike-slip fault;

† the Ghana Ridge opposite to the normal conti-nental crust;

† the Ghana Ridge opposite to the extended conti-nental crust;

† the Ghana Ridge opposite to the progressivelyyounger – that is, hotter – oceanic crust;

† the Ghana Ridge opposite to the sea-floorspreading centre;

† the Ghana Ridge opposite to the cooling oceaniccrust and inactive transform fault.

Three geodynamic scenarios have been simulatedby modelling. Each of them starts at 104 Ma(Albian).

Scenario 1 prescribes the case of pure strike-slipfaulting along the Romanche Transform Fault. Themigrating spreading centre is situated right next tothe transform fault and Ghana Ridge. It containsthree subscenarios S1a, S1b and S1c, characterizedby a dip of the transform fault of 908, 608 and 458,respectively. The reason for sensitivity analysis onthe dip of the transform fault is, first, a range ofdips from 318 to 678 determined by this and otherstudies of the Ghana Ridge (e.g. Clift & Lorenzo1999), and, second, a general understanding thattransform zones are typically associated with verti-cal faults (e.g. Harding et al. 1985).

Scenario 2 also prescribes the case of pure strike-slip faulting but the spreading centre is situatedfurther away from the transform fault and GhanaRidge. The scenario simulates a geological case ofatypically thin oceanic crust forming a thermalbuffer between the Ghana Ridge and a spreadingcentre (see Sage et al. 2000). It would indicateoceanic crust formation in the setting with lowmagma budget. The 70 km-wide zone representinga thermal buffer should reduce the lateral heat trans-fer into the Ghana Ridge and uplift due to thermalexpansion, and cause a time delay between thespreading centre passage and maximum heat flowin the Ghana Ridge.

The scenario contains nine sub-scenarios, repre-senting combinations between:

† the spreading centre located at a distance of 11,31 and 65 km from the transform fault; and

† a dip of transform fault of 908, 608 and 458,respectively.

Scenario 3 prescribes a transpression along theRomanche Transform Fault. In this case, theGhana Ridge uplift is controlled not primarily bythermal expansion but by thermal weakening ofthe thinned continental lithosphere with respect tothe stronger oceanic crust adjacent to it. The scen-ario contains eight sub-scenarios, representing com-binations between:

† spreading centre located 31 and 65 km from thetransform fault;

† transform fault dipping at 608 and 458; and† fixed boundary conditions or lithostatic pressure

gradient.

The geological justification for this scenario isa transpression generated by a spreading vectorrotation described earlier, in the text on the Equator-ial Atlantic setting (Fig. 3).

While the time interval 104–83 Ma in the vicin-ity of the Ghana Ridge can be characterized by purestrike-slip movements, a progressive rotation of thespreading direction resulted in a contraction com-ponent of 2970 m Ma21 at 83 Ma.

Owing to the 20 sub-scenarios included in themain scenarios, only a limited amount of modellingresults can be described here for spatial reasons.As a consequence, we choose end-member sub-scenarios and discuss general trends in uplift andtemperature distributions in time. Modelling resultsare represented as maps of uplift and temperaturedistribution, and diagrams showing uplift withtime for specific Ghana Ridge checkpoints, togetherwith an indication of the ridge at the distance closestto a chosen check point. As only an approximatedescription of real rheologies is used, we do notaspire to predict exact temperature and upliftvalues but focus on trends in their distribution andtheir controlling factors.

Results

The comparison of modelled uplifts in the threesub-scenarios of Scenario 1, characterized by themigrating spreading centre situated right next tothe transform fault and Ghana Ridge, indicates theuplift decrease with decrease in dip of the transformfault through which the lateral heat transfer fromthe migrating spreading centre occurs (Fig. 4). Dipvalues of 908, 608 and 458 are associated withuplift of the Ghana Ridge of about 340, 270 and220 m, respectively, at the point right at the contact



with the passing-by centre. The surface heat flowsassociated with the first and last sub-scenario areabout 160 and 85 mW m22 (Fig. 5).

Comparison of modelled uplifts in the threesub-scenarios of Scenario 2, characterized by a dipof the transform fault of 608 and the transform fault-spreading centre distances of 11, 31 and 65 km,indicates the uplift decrease with increase in thedistance between transform fault and spreadingcentre (Fig. 6). Distances of 11, 31 and 65 km areassociated with uplifts of 120, 20 and 220 m,respectively. The surface heat flows associated withthese sub-scenarios are 70, 60 and 55 mW m22,respectively (Fig. 7). An important consequence ofthe distance between the passing-by spreadingcentre and the Ghana Ridge is a time delay betweenthe centre passing and the peak surface uplift/heatflow, which increases with distance. The delays are4.8, 9.5 and more than 11.3 Ma for distances of 11,31 and 65 km. Another trend is the decrease inthe peak value with distance, being 70, 60 and54.5 mW m22 for distances of 11, 31 and 65 km.

Discussion of the results simulating Scenario 3,characterized by transpression, can be divided intotwo categories. The first one keeps the front, rearand African sides of the model fixed, which tendsto exaggerate the resulting uplift. The second onekeeps the rear side of the model fixed, while theAfrican and front sides are prescribed with a litho-static pressure gradient condition, which decreasesthe resultant uplift.

The category with exaggerated uplift documentsthat there is a less distinct decreasing trend of upliftwith distance between the passing-by spreadingcentre and the transform fault (Fig. 8) but theuplift peak is rather large, reaching values of theorder of 4.4–5.5 km for distances of 31–65 km.The category without exaggerated uplift also docu-ments the uplift decrease with the spreading centre-fracture zone distance, which is, however, verysubtle (Fig. 9). The uplift peak values vary around2.5 km, which is just 50% of the values frommodels with fixed walls but almost an order of mag-nitude larger than uplifts triggered by thermalexpansion from scenarios 1 and 2.

Apart from thepeakuplift, it is interesting to com-pare the overall volume of uplift associated withdifferent uplift scenarios. We do this roughly bymeans of contrasting the widths of the zone of uplift.

Uplift controlled by the passage of the sea-floorspreading-centre-related thermal anomaly affectsthe zone, which is narrower than the zone of thetranspression-related uplift. For the spreadingcentre passing right next to the Ghana Ridge, thewidth of the zone with uplift greater than 125 m is43 and 15 km for a transform fault dip of 908 and458, respectively. This indicates decreasing heattransfer effectiveness through the contact betweencontinental and oceanic crust with decreasingcontact fault angle.

For the spreading centre passing at a distancefrom the Ghana Ridge itself, being separated fromit by a corridor of atypically thin oceanic crust, thewidth of the zone with uplift greater than 125 m is

Fig. 4. Comparison of surface uplifts on Ghana Ridge,modelled by all three sub-scenarios of Scenario 1,representing the spreading ridge passage adjacent to theGhana Ridge. S1a, transform fault with a dip of 908; S1b,transform fault with a dip of 608; S1c, transform faultwith a dip of 458. Note that the modelled uplift decreaseswith a decrease in the transform fault dip.

Fig. 5. Comparison of surface uplifts and heat flows on Ghana Ridge, for a transform fault dip of (S1a) 908 and (S1c)458 for Scenario 1, representing the spreading ridge passage adjacent to the Ghana Ridge. Note that the modelled upliftand heat flow decrease with a decrease in the transform fault dip.



3, 0 and 0 km for a distance/fault dip combinationof 11 km/608, 31 km/608 and 65/608, respectively.This indicates decreasing heat transfer with increas-ing distance between the passing spreading centreand the Ghana Ridge itself.

For the uplift controlled by the transpressionand transpressional contact being at a distanceof 65 km from the Ghana Ridge, the width of thezone with uplift greater than 750 m is 160 and 173for fault dips of 608 and 458, respectively. Whenthe distance is decreased to 31 km and the dipof the fault is kept at 608, the width increases tothe entire width of the model. This indicates that atranspression is a more important uplift driver thanthe thermal anomaly of the passing spreadingcentre and its effect increases with decreasingdistance of the Ghana Ridge from the primarytranspressional contact.

Another important observation is that there isno clear indication of flexural uplift in the models.Our explanation of the lack of a distinct flexuraluplift or isostatic rebound would be the high temp-eratures. They strongly weaken the lithosphere,which in turn results in a very low flexural rigidity.Thus, the lithosphere is always close to local Airyisostasy, so that flexural and rebound effects areunlikely to play a major role.

Discussion

Regardless of their differences, scenarios 1 and 2basically show how the thermally driven upliftrelated to the passing-by spreading ridge movesfrom east to west along the Ghana Ridge (Fig. 10).As the thermal anomaly decays in time, the upliftbecomes progressively replaced by subsidence.Shifting location in a westerly direction, bothearlier uplift and subsequent subsidence travel aswestwards migrating waves. Such migrating subsi-dence should explain the abrupt end of Turonian–Coniacian carbonate deposition some time in theConiacian, and a more rapid subsidence of theridge noted by Basile et al. (1998). However, inorder to prove modelled migrating thermally con-trolled uplift and subsequent subsidence, one needsto find a system of hiatuses/erosional unconformitiesin ODP Leg 159 sites compatible with this model.

In order to do this rigorously, we can look atsites 959, 960, 961 and 962 of the ODP Leg 159,an unspecified industrial well in the Deep IvorianBasin to the NW of the Ghana Ridge, and discussour own chronostratigraphic and palaeoenviron-mental results provided by the determined faunacorrelation to the composite standard.

A closer look at the chronostratigraphy of sitesin the Ghana Ridge (discussed in SUP 18518)

Fig. 6. Comparison of surface uplifts on Ghana Ridge,modelled by all three sub-scenarios of Scenario 2,representing the spreading ridge passage at a distancefrom the Ghana Ridge. S2b represents a distance of11 km; S2e represents a distance of 31 km; and S2hrepresents a distance of 65 km. All three scenarios have atransform fault dip of 608. Note the decrease in upliftwith the increase in distance between the Ghana Ridgeand the spreading centre.

Fig. 7. Comparison of surface uplifts and heat flows on Ghana Ridge, for a transform fault dip of 608 for spreadingcentre–Ghana Ridge distances of (S2b) 11, (S2e) 31 and (S2h) 65 km for Scenario 2, simulating the spreadingridge passage at a distance from the Ghana Ridge. The dashed vertical line indicates the passing-by sea-floor spreadingridge at the shortest distance from Ghana Ridge. Note the modelled uplift and heat flow decrease with increase indistance between the spreading centre and Ghana Ridge. See the text for further explanation.



does not allow one to interpret any systematic wavesof uplift and the following subsidence along theGhana Ridge that could be associated with thermaleffects of the passing spreading ridge. The mostimportant chronostratigraphy control seems to bethe ridge topography plunging in a westerly direc-tion. Therefore, the numerically modelled scenarioof the thermally driven uplift, characterized by themigrating wave of uplift, does not look like a prob-able interpretation, while the modelled transpres-sionally driven uplift does.

Furthermore, the problem with the modelledGhana Ridge uplift, driven only thermally, is par-tially in its relatively small magnitude (see Figs 4& 6), and mainly in the insignificant width of theGhana Ridge Zone affected by the uplift, whichwas described at the end of Results section. Amaximum uplift of 340 m can be achieved by ascenario with the spreading ridge directly next tothe transform margin and vertical transform fault.The first condition is slightly unrealistic, as thereis a good chance of fault blocks separating themargin from the ridge. Examples, such as the Coro-mondal Transform Fault of the East Indian margin(Nemcok et al. 2012), show that such fault blocksare common. The second condition is an unrealisticdip angle of the transform faults controllingbreak-up, because the dip is typically less than 908(Clift & Lorenzo 1999 on the Romanche TransformFault; Rosendahl 2005 pers. comm. on worldwidetransform faults). Models that have a smaller dipangle of the transform fault indicate a less effectiveheat transfer through it – that is, a smaller magni-tude of uplift (Fig. 4). Placing the heat source –the passing-by spreading ridge – a bit further fromthe Ghana Ridge further reduces the thermally con-trolled uplift. Maintaining the 608 dip of the trans-form fault and changing the spreading ridge

distance from the Ghana ridge from 0 km, throughto 11 and 31 km and up to 65 km, reduces themaximum uplift of the adjacent zone of the GhanaRidge from 275 m, down to 125 and 25 m, andfinally to practically nothing. This is associatedwith a dramatic decrease in the width of the zoneaffected by uplift to reach practically no width foruplifts greater than 125 m and already at a distanceof 31 km. Furthermore, the modelling results indi-cating insignificant thermally driven uplift are inaccordance with the fact that apatites were not sig-nificantly thermally reset by the passing spreadingridge (see Clift et al. 1998).

Fig. 8. Comparison of surface uplifts on Ghana Ridge,for a spreading centre–Ghana Ridge distance of 31 km,for transform fault dips of 608 (S3a) and 458 (S3c) forScenario 3, representing a dextral transpression driven bya change in drift vector. The models have front, rear andAfrican walls fixed. Note a modelled decrease in upliftwith a decrease in transform fault dip.

Fig. 9. Comparison of surface uplifts on Ghana Ridge,for a spreading centre–Ghana Ridge distance of 65 km,for transform fault dips of 608 (S3e) and 458 (S3g) forScenario 3, representing a dextral transpression driven bya change in drift vector. The models have only the rearwall fixed. Note a modelled decrease in uplift with adecrease in transform fault dip.

Fig. 10. Vertical movement distribution at 78 Ma alongthe Ghana Ridge in response to the passing-by spreadingridge on its way out of the Gulf of Guinea. The modelrepresents Scenario S1a, characterized by a spreadingcentre–Ghana Ridge distance of 0 km and a transformfault dip of 908. The black arrow indicates the location ofthe spreading ridge. Note that the westward migratingwave of uplift is followed by a wave of moderatesubsidence.



Another uplift mechanism to discuss is flexuraluplift, which is one of the alternative mechanismsduring the transform motion between continentaland oceanic plates. Flexural unloading of a seaward-dipping transform fault in a manner analogous tofootwall uplift in extensional settings could be aprocess causing permanent uplift. This unloadingwould depend on the magnitude of the transformfault dip and the flexural rigidity of the continentalplate, as discussed by Clift & Lorenzo (1999). How-ever, our numerical simulation provided no clearindication of flexural uplift. Our explanation of thelack of a distinct flexural uplift or isostatic reboundwould be the high temperatures. They stronglyweaken the lithosphere, which in turn results invery low flexural rigidity. Thus, the lithosphere isalways close to local Airy isostasy, such that flexuraland rebound effects are unlikely to play amajor role.Our observation of very low flexural rigidity is inaccordance with findings made by Karner & Watts(1982), Holt & Stern (1991) and Kooi et al. (1992)for the situation shortly after the rifting.

The flexural uplift of the Ghana Ridge wassuggested by Clift & Lorenzo (1999), who used atwo-dimensional (2D) flexural cantilever modelfor its determination. However, the use of a 2Dapproach along a section perpendicular to theGhana Ridge brings important constraints to theinterpretation of flexural uplift. First, the directionof the unloading of the transform fault footwall inreality is not perpendicular to the ridge, as calcu-lated in their model (see Fig. 1 in Clift & Lorenzo1999), but at a very low angle to the ridge. Thismeans that their 2D flexural model should havemost of the forces controlling the uplift out ofplane of the model. Secondly, Clift & Lorenzo(1999) determined about 3 km of crustal thinningover a distance of approximately 18 km in thesouthern Ghana Ridge. However, the thinning isnot homogeneous but occurs in the upper levels ofthe continental crust because the underlying Mohoin their profile is flat. We would suggest that thisupper crustal thickness reduction is due to a lateralretreat of the transform fault scarp. It would be indi-cated by apatite fission track data (see Clift &Lorenzo 1999; Bigot-Cormier et al. 2005), whichhave a landwards-younging trend in their agesalong profiles from the oceanic–continental crustboundary to the Ghana Ridge crest. This trendforms a mirror image of the trend usually observedalong apatite fission track data profiles startingsomewhere deep in the contintent and finishing atthe highest point of the marginal uplift (see e.g.van der Beek et al. 1995). Neither Clift & Lorenzo(1999) nor Bigot-Cormier et al. (2005) could seethat this is a mirror image because they only haddata from the ODP sites at the ridge crest and datafrom submarine sampling performed on the slope

oceanwards of the ridge crest. van der Beek et al.(1999) explained the phenomenon of fission trackages at the margin edge being younger than theage of rifting and the relatively high track lengthsby the younger age apatites having experiencedtemperatures above 120 8C prior to rifting. Theapatite fission track ages usually increasing land-wards from the margin edge, forming the highestpart of the margin, and mean track lengths firstdecreasing and subsequently increasing away fromthe coastline indicate that samples were exhumedfrom subsequently lower temperature ranges –that is, shallower depth levels. In fact, the tracklengths along the Ghana Ridge crest are consistentlylonger than 14 mm, which indicate a rapidly cooledmargin at Ghana Ridge (Bigot-Cormier et al. 2005).

When we correlate the apatite fission track agetrend with the change in dip of the transform faultscarp, represented by the southern Ghana Ridgeslope, we see two dip provinces. The upper slopeportion is not a fault scarp itself but a retreatedone. This means that real transform fault dip is inthe lower portion, which is sealed by sediments. Adip of the retreated fault scarp in the depth-migratedseismic section of Clift & Lorenzo (1999, their fig.6) is 318, while a dip of the fault is 678.

This brings us to the third problem with the 2Dflexural cantilever model. The fact that sensitivityanalysis (Clift & Lorenzo 1999), focused onvarying four parameters in control of the flexuraluplift, which are:

† the amount of extension represented by the bfactor;

† the effective elastic thickness;† the transform fault dip; and† the transform fault heave,

cannot explain the modern Ghana Ridge, which isrepresented by parameters such as (1) the amountof uplift experienced and (2) its geometric shape,with a transform fault that has a dip larger than308 where we already know from the previous dis-cussion that an appropriate dip would be about 678.

Clift & Lorenzo (1999) would need to combinetheir flexural uplift with spreading-ridge-passage-related thermal uplift to reach the uplift valuesthey need, given the inappropriate shallow-diptransform fault, to match the observation of hiatusfrom 105 to 90 Ma. However, the heat needed foruplift at sites 959 and 960 would occur there onlyafter a significant delay of more than 12 Ma afterthe centre passage owing to their distance from thecontinental–oceanic crust boundary, which wouldbe too late. Furthermore, the passage of the centrecould be younger than previously thought, whichwould further enhance the problem.

The passage of the oceanic crust adjacent to theGhana Ridge has been understood as indicated by



Cenomanian bathyal marine conditions at Site 962(Mascle et al. 1988, 1996). Even younger, earlyTuronian, spreading ridge passage was suggested byClift et al. (1998), based on the regional plate tec-tonic reconstructions and unconformity penetratedby the ODP Leg 159 sites. The extrapolated ageof oceanic crust in our Figure 3 indicates that thepassage was probably even younger than 90 Ma;that is, 81–80 Ma. This timing is closer to thatderived by Bigot-Cormier et al. (2005) fromapatite fission track data, which dated the passageat 85–80 Ma, representing the Santonian. Thedifference in Bigot-Cormier et al. (2005) and ourridge passage timing from the timing determinedin the 1990s (Bouillin et al. 1998; Clift et al.1998; Clift & Lorenzo 1999 and references inthese studies) could simply be a result of the pro-gress achieved in one decade in understanding thechronostratigraphy of Late Cretaceous stages.

This brings us to the fourth problemwith flexuraluplift, this time enhanced by suggested synchronousthermal uplift. As concluded by Bigot-Cormier et al.(2005) based on apatite fission track data, the GhanaRidge has appeared as an uplifted area since the lateAlbian (see alsoBasile et al. 1998), a significant timebefore the oceanic accretion against the transformmargin, and the spreading centre passage thus came

too late to be synchronous with the required timingof flexural uplift. The interpretation of uplifted areasince the late Albian is further supported by ourbiostratigraphical work (see Supplementary publi-cation SUP18518); in particular, a set of hiatuses/erosional unconformities separated by subordinatedepositional events in Site 959 indicates a long-lasting high.

After ruling out flexural uplift, in order toenhance the thermally driven uplift but mainly tosignificantly increase its influence zone in theGhana Ridge, we need to search for the existenceof yet another mechanism. We would have tocouple the thermal effect of the passing spreadingcentre with the transpression associated with theearly syn-drift change in the spreading vector,which was mentioned earlier (Fig. 3, Tables 1 & 2).

Drift vector change, described as occurring aslate as theTuronian–earlyCampanian,wasobservedby numerous workers (Haxby 1985; Klitgord &Schouten 1986; Fairhead 1988a, b; Fairhead &Binks 1991; Binks & Fairhead 1992). This event isassociated with the reactivation of many oceanicfracture zones and transform faults, which arelocated in the Equatorial Atlantic (Mascle et al.1997; Dailly 2000). It has been postulated byDailly (2000), based on a broad timing correlation,

Fig. 11. Reflection seismic section through offshore Ghana showing the Albian–Cenomanian (108–92 Ma)transtension followed by Coniacian–Santonian (89–83.5 Ma) and Danian–Thanetian (65–52 Ma) transpressionevents. The scale and location are not given for confidentiality reasons.



that this event caused transpressional eventsalong transform faults that resulted in the shorten-ing of the Aptian–Albian sediments in the RioMuni Basin.

The areal extent of these transpressional featuresis interesting. They can be observed in the Rio MuniBasin (Dailly 2000) and in the Northern GabonBasin (Teisserenc & Villemin 1990) but cannot befound in the Congo Basin (Seiglie & Baker 1984)and basins located further south. Therefore, theevent is characteristic only for the Equatorial Atlan-tic and did not occur to the south of the NorthernGabon Basin (e.g. Dailly 2000). The northern limitfor the event can be determined with help fromreflection seismic imaging. While the seismicimaging in the Ivory Coast–Ghana region indicatesthat the Coniacian–Thanetian inversion affectedseveral former transtensional strike-slip faults tothe north of the Romanche Transform Fault andtransform fault itself (Fig. 11), it did not seem toaffect the St Paul Transform Fault bounding theIvory Coast Basin from the north (Fig. 12).

Similar transpressional folding of rift sedi-ments has been reported from the Foz do AmazonasBasin (Figueiredo 1985; Aguiar et al. 1986). Aguiaret al. (1986) observed that folding also affected theLower Tertiary strata. This event was also reportedin the Benue Trough, onshore Nigeria (Benkhelil1988), where it is diachronous, starting during theSantonian in the south and finishing during the endof the Cretaceous in the NE.

The coupling of bothmechanisms can actually bequiteimportantfortranspressionaldeformationenha-ncement. This can be impliedwith a closer look at thecontinental margin in the Ghana Ridge region (Figs13 & 14). A small segment of the distal margin inthe offshore Ivory Coast shows a clear pull-apartcontrol of the continental break-up and a significantwarming-up of the distal margin around its develop-ment (Fig. 13). The former is documented by theAfrican half of the rhomb-shaped geometry of thepull-apart basin. The latter can be implied from avolcanic ridge, indicating a significant emplacementof magmatic bodies during the break-up (Fig. 13).

Fig. 12. Reflection seismic section through the offshore Ivory Coast, showing the Albian–Cenomanian (108–92 Ma)transtension and no apparent subsequent larger-scale transpression. The scale and location are not given forconfidentiality reasons.



Further east, in offshore Togo and Benin, lies amargin segment that allows the comparison of theproximal and distal margins (Fig. 14). The formerand latter are labelled as continental and transitionalcrusts in the figure. Their structural grains are verydifferent from one another. The two sets of fault pat-terns are divided by the boundary between the prox-imal margin and the distal margin, whose crust is

understood to be highly attenuated continentalcrust in its last stages of stretching preceding conti-nental break-up. Both fault patterns define systemsof relatively small pull-apart basins. Those formedin a thicker continental crust of the proximalmargin have a clear distinction between normaland strike-slip faults controlling their rhomb-shapedgeometries. They are analogous to the analogue

Fig. 13. Tectonic map of a small passive margin segment initiated by a pull-apart basin, offshore Ivory Coast. Note arelatively clear distinction of strike-slip faults and normal faults on the landward side of the pull-apart and theoccurrence of a volcanic ridge near the boundary between the continental and oceanic crusts terminating the pull-apartbasin on its oceanward side.



Fig.14.Tectonicmap

oftheTogo–Beninmarginsegmentwithfailed

pull-apartbasinsdetached

atbothbrittleandductiledeform

ationlevels.Analogoustotheresultsofanalogue

materialmodellingcarriedoutbySim

setal.(1999),therhomb-shaped

pull-apartbasinslocatedclosertotheshoreline,whichhaveaclearidentificationofcontrollingstrike-slipand

norm

alfaults,havebeendetached

insidethebrittle

deform

ationzone.Thenarrowpull-apartbasinslocatedcloserto

theoceanic

crust,whichdonothaveacleardistinctionof

controllingstrike-slip

andnorm

alfaultsbutarecontrolled

byoblique-slip

faults,havebeendetached

insidetheductiledeform

ationzone.



material model results of Sims et al. (1999), whodetached a set of their pull-apart basin modelsinside a brittle deformation zone. Those formed ona thinner crust of the distal margin lack a clear dis-tinction between normal and strike-slip faults. Theyare narrower than those on the thicker continentalcrust and are controlled somewhat by single Riedelshears, which are dextral strike-slip faults with anormal fault component. These pull-aparts are analo-gous to the pull-apart models of Sims et al. (1999)detached inside a ductile deformation zone.

When we compare this map (Fig. 14) with theresults of the analogue material modelling of pull-apart basins detached at brittle v. ductile defor-mation levels described by Sims et al. (1999), itbecomes apparent that the pull-apart basins devel-oped in the continental crust of the proximalmargin were detached along brittle detachments astheir host crust was thicker and cooler during theearly stages of stretching. It is also apparent thatpull-apart basins developed in the continental crustof the distal margin were detached along ductiledetachments because their host crust was thinnerand warmer during the mature stages of stretching.

Because the spreading vector adjustments occur-red relatively soon after the continental break-up,

characterized by its warmed thinned crust, themargin and, mainly, the distal margin did not coolenough to become stronger. It was relatively weakand prone to deformation, which was driven by theextra contraction component driven by the spreadingvector rotation. The thicker and cooler continentalcrust further landwards did not experience any sig-nificant transpressional deformation.

An example of the continentalmargin affected bytranspression post-dating the continental break-upcomes from offshore Benin (Fig. 15). Figure 15shows that while break-up unconformity developedsome time during the Albian in this area, the strike-slip faulting lasted a little longer. Furthermore,strike-slip faulting was accompanied by folding.While the lower portion of the post-Albian UpperCretaceous strata does not show any response to thegrowing anticline, the overlying Upper Cretaceous–lower Palaeocene strata thin over its crest thus indi-cating its growth during that time.

It needs to be emphasized at this point that itis not the boundary between the distal and prox-imal margin that forms a landwards constraint forthe occurrence of transpressional deformation. Thiscan be seen in Figures 13 & 14, which containseveral folds located in thicker crust. Furthermore,

Fig. 15. The early syn-drift deformation of the previously extended continental crust in offshore Benin. Thedextral-transpression-driven fold has chevron geometry. The Upper Cretaceous–lower Palaeogene sediments thinningover its crest indicate its Senonian–early Palaeogene growth. The fold shows hardly any erosion as early Palaeogenesediments seem to be removed by gravity gliding. The scale and location are not given for confidentiality reasons.



transpressional folds occur in shelfal zones of severalEquatorial margin segments. Offshore Benin servesas such an example. Other examples come fromthe Barreirinhas Basin (Rodrigues et al. 1984; DeAzevedo1991)and theCearaBasin (Zalanetal.1985).

Although the neighbouring transform marginsegments may have different histories, as indicatedby data from the Romanche and St Paul continentalmargins (see Figs 11 & 12), transformmargins sharecertain common features with a direct applicationto understanding the maturation histories of individ-ual hydrocarbon source kitchens and the timing ofstructural traps (discussed in SUP 18518).

Conclusions

† The main result of this paper is that there wasa strike-slip movement during the crustalbreak-up in the Equatorial Atlantic. After thebreak-up, there was a progressive change in thespreading direction from ENE–WSW to east–west, which caused transpression in severalregions of the Gulf of Guinea. The transpressioneffect was enhanced owing to continental crustbeing weakened by both the thermal effects ofthe rifting that resulted in break-up and the sub-sequent passage of spreading centres along trans-form faults. The thermal effect of the passingridges was a secondary factor in the control ofuplifts along several margin segments.

† Numerical simulation shows that the thermallydriven uplift related to the ‘passing-by’ spreadingridge after the continental break-up must havemoved from the east to the west along theGhana Ridge. As the induced thermal anomalydecayed in time, the uplift became replaced bysubsidence. However, the stratigraphic recordfrom individual sites of the ODP Leg 159locatedalong the ridge doesnot allowone to inter-pret any distinct east-to-west migrating wave ofuplift and subsequent subsidence. This indicatesthat the thermal control of the uplift must havepaled out in comparison to the transpression.

† Simulation also shows that both positive heatflow anomaly and thermally induced uplift inmodelled scenarios decrease with the dip of thetransform fault dividing the oceanic crust fromthe continental margin. Dips of 908 and 458,for example, are associated with heat flows of160 and 85 mW m22, and uplifts of 340 and220 m, respectively. The smaller dip angles areassociated with less effective heat transfer.

† Both the heat flow anomaly and uplift decreasewith distance between the bounding transformfault and the spreading centre. An importantconsequence of the distance is the delay in ther-mal perturbation of the continental margin with

respect to the timing of the passing spreadingcentre.

† The modelling further shows that transpression-driven uplift also decreases with distance fromthe spreading centre, although not as signifi-cantly as it does in the thermally controlledcase. Compared to thermally induced uplift, itis about one magnitude larger.

† The comparison of modelled transpression-driven and thermally induced uplifts, and zonesof their influence, indicates that thermallydriven uplift would not be enough to explainthe observations. However, the thermal pertur-bation driven by a passing-by spreading centreis significant enough to control laterally varyingthermal maturation histories.

† A combination of the spatial distribution ofoceanic crust with evidence of the early driftvector rotation with spatial distribution of trans-pressional events in the continental margin andwith numerical models indicates that transpres-sional events inside the margin were driven bythe early-drift vector rotations. This mechanismonly characterized the Equatorial Atlantic, withthe southern and northern limits at the SouthGabon–North Gabon basin boundary and theSt Paul Transform Fault, respectively.

† Observations and modelling indicates that theincreasedheatflowregimeof thedistal continentalmargin after continental break-up enhanced thetranspressional effect of the transpression drivenby early-drift spreading vector rotation. Obser-vations include evidence such as the proximalmargin containing pull-apart basins detached atbrittle deformational levels, while the distalmargin has them detached at ductile levels or theevidence for transpressional folding locatedmostly in the distal margin and mostly missingin the proximal margin, apart from exceptions inthe Barreirinhas and Ceara basins, and offshoreBenin. In fact, the modelling indicates that trans-pressional deformation could have penetratedinto the Ghana Ridge quite a distance from itsdriving force. Furthermore, the distal margin issometimes found to contain large magmaticbodies indicating a warmer heat flow regimeduring the break-up, which made it especiallyprone to subsequent transpressional deformation.

† The key behind transpression being so effectivewas the fact that the continental margin, andmainly its distal portion, did not have sufficienttime to cool between the break-up and theonset of transpression driven by the spreadingvector rotation.

† Apart from the unique uplift and heat flow his-tories of the transform margin segments, theirslopegeometry interactingwithdepositionusuallydelays the onset of the cycles of gravity gliding



significantly, which differentiates them frompassive margins initiated by normal faulting.

The work on Equatorial Atlantic passive margins wascarried out within the framework of EGI project01-00059-5000-50500936 (2002–2003 ‘Equatorial Atlan-tic Margins Basins Project’) funded by Apache,BHP-Billiton, EnCana, Kerr-McGee, Nexen, Occidental,Ocean Energy, Phillips, Repsol, Shell and Wintershall.The paper benefited from rigorous reviews made byI. Davison and N. Kumar. We are grateful for help fromI. Nemcokova in editing the text, and help provided byL. Ledvenyiova, M. Gazi, O. Pelech, T. Kluciar,I. Perichtova and P. Ekkertova in drafting the figures.

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continental break-up along strike-slip fault zones ... · strong dextral strike-slip-related...

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