Tectonophysics 607 (2013) 32–50

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Pre-existing oblique transfer zones and transfer/transform relationships in continentalmargins: New insights from the southeastern Gulf of Aden, Socotra Island, Yemen

N. Bellahsen a,b,⁎, S. Leroy a,b, J. Autin c, P. Razin d, E. d'Acremont a,b, H. Sloan e, R. Pik f, A. Ahmed g, K. Khanbari h

a UPMC Univ. Paris 06, UMR 7193, ISTeP, F-75005 Paris, Franceb CNRS, UMR 7193, ISTeP, F-75005 Paris, Francec CGS-EOST, CNRS-Univ. Louis Pasteur, F-67084 Strasbourg, Franced ENSEGID, IPB, Univ. Bordeaux 3, 33607 Pessac Cedex, Francee Lehman College, CUNY, Environmental, Geographical and Geological Sciences, New York, USAf CRPG-CNRS, BP20, 54501 Vandoeuvre-Lès-Nancy Cedex, Franceg Seismological and Volcanological Observatory Center, Dhamar, Yemenh Yemen Remote Sensing and GIS Center, Sana'a University, Sana'a, Yemen

⁎ Corresponding author at: UPMC Univ. Paris 06, UMFrance. Tel.: +33 1 44 27 74 64.

E-mail address: nicolas.bellahsen@upmc.Fr (N. Bellahs

0040-1951/$ – see front matter © 2013 Elsevier B.V. All rihttp://dx.doi.org/10.1016/j.tecto.2013.07.036

a b s t r a c t

a r t i c l e i n f o

Article history:Received 8 October 2012Received in revised form 8 July 2013Accepted 26 July 2013Available online 9 August 2013

Keywords:Transform faultTransfer fault zoneFracture zoneSegmentationStructural inheritanceOblique rifting

Transfer zones are ubiquitous features in continental rifts and margins, as are transform faults in oceanic litho-sphere. Here, we present a structural study of the Hadibo Transfer Zone (HTZ), located in Socotra Island(Yemen) in the southeastern Gulf of Aden. There, we interpret this continental transfer fault zone to representa reactivated pre-existing structure. Its trend is oblique to the direction of divergence and it has been activefrom the early up to the latest stages of rifting. One of the main oceanic fracture zones (FZ), the Hadibo–Sharbithat FZ, is aligned with and appears to be an extension of the HTZ and is probably genetically linked toit. Comparing this settingwith observations from other Afro-Arabian rifts as well as with passive margins world-wide, it appears that many continental transfer zones are reactivated pre-existing structures, oblique to diver-gence. We therefore establish a classification system for oceanic FZ based upon their relationship with syn-riftstructures. Type 1 FZ form at syn-rift structures and are late syn-rift to early syn-OCT. Type 2 FZ form duringthe OCT formation and Type 3 FZ form within the oceanic domain, after the oceanic spreading onset. The latterare controlled by far-field forces, magmatic processes, spreading rates, and oceanic crust rheology.

© 2013 Elsevier B.V. All rights reserved.

1. Introduction

Transfer zones and accommodation zones are ubiquitous features incontinental rifts. These zones accommodate along strike structuralchanges within rifts (amount of extension, dip of tilted half grabens;e.g. Bellahsen and Daniel, 2005; Colletta et al., 1988; Corti et al., 2007;Jarrigue et al., 1990; Lezzar et al., 2002; Young et al., 2001) and havebeen extensively studied (e.g., for conceptual models, Bosworth, 1994;Rosendahl, 1987). In the literature, accommodation zones are definedas areas of diffuse strain transfer between two offset basins, while trans-fer zones are more localized fault systems. Acocella et al. (2005) sug-gested that accommodation zones are found in rifts and transfer zonesin rifted continental margins. However, examples from the Gulf ofSuez (e.g. Jarrigue et al., 1990; Moustafa, 1997) and the East AfricanRift System (e.g. Lezzar et al., 2002; Rosendahl, 1987) show that transferzones were active during early stages of rifting. It is noteworthy that, in

R 7193, ISTeP, F-75005 Paris,

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ghts reserved.

these two examples, the transfer zones are both pre-existing andoblique to the divergence. This may suggest that transfer zones doformduring early rifting, only if they reactivated pre-existing structures.

Geometric consistencies between continental margin transfer zonesandmid-ocean ridge transform faults suggest that transform faults orig-inate at structures inherited from the rifting stage (Behn and Lin, 2000;Cochran and Martinez, 1988; Brekke, 2000; d'Acremont et al., 2005,2006; McClay and Khalil, 1998; Watts and Stewart, 1998; Wilson,1965). However, a significant number of studies suggest that transformfaults are inherent to spreading processes and thus not related to conti-nental margin segmentation (e.g., Taylor et al., 2009) but to thermalstress (Sandwell, 1986), and that theymay evolve rapidly after their for-mation (e.g. d'Acremont et al., 2010). Their origin remains unclear asshown in recent reviews (Choi et al., 2008; Gerya, 2012). Analogmodelsof transform-ridge systems (Freund and Merzer, 1976; O'Bryan et al.,1975; Oldenburg and Brune, 1972, 1975) showed that transform faultscan nucleate spontaneously in accreting and cooling plates of wax.However, surprisingly, later experiments (e.g. Katz et al., 2005;Ragnarsson et al., 1996) were not as successful in reproducingtransform faults (see Gerya, 2012 for more details and the complete

33N. Bellahsen et al. / Tectonophysics 607 (2013) 32–50

review). Numerical models confirmed that transform faults may indeedform spontaneously during oceanic spreading, even without initialridge offset (Choi et al., 2008; Gerya, 2010, 2012 and referencestherein).

So far, few studies have attempted to determine, from onshore fielddata, whether or not transfer zones influence the location of oceanictransform faults. MacDonald et al. (1991) proposed a hierarchy forridge discontinuities. From first order discontinuities (transform faults)to 2nd, 3rd, and 4th order (overlapping or connecting segments), thisclassification does not explicitly incorporate considerations about in-heritance from syn-rift times. As noted above, the fact that transferzones in rifts are often oblique to the divergence directionmay be of im-portance. Indeed, in continental rifted margins, syn-rift faults that areparallel to the future transform faults, i.e. parallel to divergence, are al-most never observed. These results led Taylor et al. (2009) to suggestthat transform faults do not form at transfer zones. It may indeed bethe case in some passive margins, as demonstrated by the numerousabove-cited analog and numerical models. However, the observationthat transform faults are rarely or never parallel to any transfer zones

Fig. 1. Topographic–bathymetric map of the Atlantic and Afro-Arabian oceanic and rifted domatem) and one oceanic domain (Eastern Gulf of Aden) are briefly reviewed or studied here. Exa

in the proximal margin domain does not necessarily imply that thereis no genetic link between them. Two cases are possible. First, sometransfer zones (parallel to divergence) can be imaged in distal marginsand their relations with transform faults can be discussed. Of course,the imaging of such sub-vertical features is challenging and can bequestioned, but several datasets show their existence especially in thedistal margins (e.g. d'Acremont et al., 2005). Second, the fact that trans-fer zones are oblique does not imply that they are unlike precursors oftransform faults as stated by Taylor et al. (2009). The question is, canwe find any structures in the continental margin that determine the lo-cation and geometry of the transform faults? And, if so, can we base aclassification for oceanic FZ on their relationships with continentaltransfer zones?

In this contribution, we present a study of the structural evolution ofan onshore oblique transfer zone, based on a new geologicalmap (Leroyet al., 2012; Razin et al., 2010). This feature is located in Socotra Island(in Yemen, Fig. 1), which constitutes a part of the southeastern marginof the Gulf of Aden. From structural and fault slip data, we detail thestructural evolution of transfer zone, that continues offshore as an

ins (from GeoMapApp, Ryan et al., 2009). Two rifts (Gulf of Suez and East African Rift Sys-mples from the Norwegian margin or South Atlantic domain are indicated.

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oceanic FZ (Leroy et al., 2012) here named the Hadibo–Sharbithat FZ.We show that the transfer zone is composed of reactivated, pre-existing faults that have been activated as oblique-slip normal faultssince early syn-rift times. We compare these results to Afro-Arabianrift systems (Fig. 1) to suggest that most of the transfer zones areactually pre-existing structures that are oblique to the divergence. Fur-thermore, we derive a classification for oceanic FZ based on their rela-tionships with syn-rift structures.

2. Geological setting of Socotra Island, Yemen, Eastern Gulf of Aden

In the Gulf of Aden (Figs. 1 and 2), rifting commenced at around34 Ma (Leroy et al., 2012; Pik et al., 2013-this volume; Robinet et al.,2013-this volume; Roger et al., 1989; Watchorn et al., 1998). At thistime, the subduction of Tethyan slabs beneath the Eurasian plates in-duced extension in the Afro-Arabian plate, mainly because the northernArabian plate entered the collision zonewhile the remainder of the sub-duction zone was still active (Bellahsen et al., 2003). The extension waslocated in the Afro-Arabian rifts (Red Sea, Gulf of Aden and East Africa),which were localized by the activity of the Afar hot spot. The combina-tion of intraplate stresses with the Afar weak zone produced obliquerifts that did not involve significant reactivation of anymajor preexistinglithospheric weaknesses (with trend comparable with that of the futureGulf of Aden, Autin et al., 2010a, 2013-this volume; Bellahsen et al.,

Fig. 2. Structure, topography, bathymetry, and oceanic ages from Leroy et al. (2012) and referencEarlyMiocene times. The spreading started at A5d (17.6 Ma). Twomain transform fault zones (Azones are located between them andmigrated through time. The structural map of the Dhofaris composed of normal faults trending 070°E to 110°E. The main grabens are disposed en éche070°E near the shore, while they trend 110°E near or within the OCT. Two transfer zones segmtransfer zones were mapped further north, close to the shore within the continental margin; th

2003, 2006). The rifting pattern is characterized by Tertiary reactivationand formation of E–W to 140°E grabens (Fig. 2) inherited from a Creta-ceous intraplate extensional event (see Birse et al., 1997; Brannan et al.,1997; Leroy et al., 2012). As a result of the rift obliquity, the reactivatedgrabens aswell as newly formed ones are arranged en échelon along thecontinental margins. Many recent studies have shown that duringrifting there were several successive directions of extension rangingfrom 020°E to 160°E (Bellahsen et al., 2006; Fournier et al., 2004;Huchon and Khanbari, 2003; Lepvrier et al., 2002). These extensionalstresses most probably rotated counter-clockwise from 020°E to 160°E.

In the eastern Gulf of Aden, the Ocean–Continent Transition (OCT)ridge (Fig. 2) may represent exhumed serpentinized mantle, whichmay or may not be intruded by post-rift magmatic material that modi-fied the OCT after its emplacement (Autin et al., 2010b; d'Acremontet al., 2006, 2010; Leroy et al., 2004, 2010a, 2010b; Watremez et al.,2011). The continental margin can be divided into several domains(Autin et al., 2010b; d'Acremont et al., 2005, 2006) with contrastingstyles of rifting (Leroy et al., 2010a). Three major transform-generatedFZ can be identified. They correspond to threemajor offsets of the coast-line: fromwest to east, the Alula Fartak FZ, the Socotra–Hadbeen FZ, andthe Hadibo–Sharbithat FZ. The latter is aligned with transverse struc-tures cropping out on Socotra Island, the Hadibo Transfer Zone (HTZ,Figs. 2, 3, and 4). Between the two former FZ, transfer zones can bemapped out from seismic data (Fig. 2) in the distal continental margin

es therein. The grabens alongbothmargins attest for the rifting that occurredduringOligo–lula Fartak and Socotra–Hadbeen FZ) offset the Aden–Sheba ridge. Smaller transform fault

continental margin is from Bellahsen et al. (2013-this volume). Onshore, the fault networklon along the margin and witness the rift obliquity (50°E). Offshore, the faults are mainlyent the margin within the OCT and trend around 025°E, parallel to the divergence. Twoey trend 160°E and are perpendicular to the 070°E-trending rift parallel faults in this area.

Fig. 3.Geologicalmap of Socotra Islandmodified from Leroy et al. (2012). The Island is primarily divided in twodomains, separated by theHadibo Transfer Zone (HTZ). East of the HTZ, thegeologymainly consists of a basement high, with both steep and low-angle normal faults in the northern part, without any significant syn-rift basins. In thewestern part, threemain basincrop out: the Allan Kadarma basin in the Northwestern part, partly controlled by a low-angle segment (see Figs. 5, 6), the Sherubrub–Balan basin in the Southwest, also controlled by low-angle normal fault segments, the central basin that is partly controlled by the transfer zone.

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and thus probably formed during the final stages of rifting, possibly dur-ing formation of the OCT (d'Acremont et al., 2005). These two FZ can beobserved to be continuous with transfer zones (e.g. Salalah FZ, Fig. 2).

Oceanic spreading along the central part of the Aden–Sheba ridgecommenced at 17.6 Ma (Fig. 2, d'Acremont et al., 2006, 2010; Leroyet al., 2012). In the westernmost part it started at 9 Ma (Audin et al.,2004; Leroy et al., 2012) and in the easternmost part, likely at 20 Ma(Fournier et al., 2010). From west to east, present-day spreading ratesrange from 13 mm/yr (azimuth 035°E) to 18 mm/yr (azimuth 025°E)(Jestin et al., 1994; Vigny et al., 2006). This opening direction demon-strates oblique divergence as the rift axis trend is about 075°E.

The following description of the stratigraphy is based on previousdescriptions from South Arabia in Sultanate of Oman (Leroy et al.,2012; Razin et al., 2010; Robinet et al., 2013-this volume; Figs. 6 and 8to 10). Pre-rift Formations (Fms.) are Cretaceous to Eocene and includeCenomanian limestones overlain by the Umm Er Radhuma platformcarbonates (Thanetian to Ypresian) and the Rus peritidal platform car-bonates (Ypresian). The middle to late Eocene is represented by theDammam (Lutetian to Bartonian) and Aydim (Priabonian) carbonateformations. Syn-rift series are composed of the Ashawq (platform,Rupelian) and Mughsayl (calciturbidites, Chattian) Fms. Locally, post-rift sedimentary rocks can be observed and consist of (late) Burdigalianto Langhian conglomerates (Ayaft Fm.).

3. Structural analysis of Socotra Island

Socotra Island is divided in two parts by the large HTZ (Figs. 3 and 4).The HTZ trends NE–SW and consists of several normal and oblique-slipfaults. In the eastern part of the island, Mount Haggier and surroundingpeaks comprise a large outcrop of basement rocks. Themassif is bound-ed northward by two main north-dipping steep normal faults, onetrending ENE–WSW and one trending NW–SE (Fig. 3). Other normalfaults trending NW–SE, E–W, and 070°E have been mapped in thefield. In the hanging-walls of these steep faults, low-angle normal faults(dip around 10–20°) affect the Mesozoic and Tertiary sedimentarylayers, overlying the basement through tectonic contacts (Figs. 3, 5and 6).

In thewestern part of the Island, the structural pattern is more com-plex. Three syn-rift basins (Allan–Kadarma, Sherubrub–Balan, and

Central basins) formed during Oligo–Miocene times and few smallerones formed along the transfer zone (Fig. 3). The two western basinsare aligned along E–W to 110°E trending normal faults, while the cen-tral one is partly aligned along the southern part of the HTZ. The presentstudy is focused on the northwestern and the central parts of the islandthat are described in detail in the following. We describe the structuralsetting along with fault slip data that provide an idea of the syn-riftstress states. The paleostress tensors have been calculated followingthe method of Angelier (1984). For the fault slip inversions, only smallfaults have been used (mm to dm scale displacements, Fig. 7). Balancedcross-sections are presented and based on the assumptions of constantarea and length of sedimentary layers (Dalstrom, 1969).

3.1. The Allan–Kadarma basin

In its western part, the Allan–Kadarma basin is controlled by a low-angle normal fault segment that trends WNW–ESE, between theQalansiya coast and the Jebel Kadarma (Figs. 5 and 8) and was activeduring deposition of the Mughsayl Fm. (Figs. 3 and 5, upper Oligo-cene–Miocene, Leroy et al., 2012). In the eastern part, south of theJebel Allan, the basin is controlled by a steeper fault (dip around 60–70°) (Figs. 5 and 8). Between these two sub-basins, a NE–SW normalfault occurs, which is considered as a transfer fault (see below).

In the western part, the main normal fault is a low-angle segmenttrending WNW–ESE that was active under an extension around N–Sto 020°E (sites 28, 25, 60, 271, Fig. 5). North of the main low-anglenormal fault, in Jebel Kadarma, two fault trends can be mapped. AWNW–ESE south-dipping fault branches at depth from this low-anglesegment (Fig. 8a). A segment trending 070°Ewas activated subsequent-ly, as indicated by its abutment on theWNW–ESE fault segment. There,most of the fault slip data indicate a 020°E extension (sites 37, 141, 233,234, Fig. 5). However, some fault-slip data suggest that theremight havebeen a counterclockwise rotation of the extension, from N–S/020°E toaround 160°E: the distribution of some fault slip data is not symmetrical(Fig. 4). Some stereonet analyses have a wide and asymmetrical distri-bution of fault slip (sites 37 and 234, but also 28, 265, 274, and 275 inother areas, Fig. 5). They differ from other sites such as sites 25, 326,328, or 60B for example (Fig. 5), where the distribution is symmetricalaround the perpendicular to the fault strike. When the distribution is

Fig. 4.Hadibo Transfer Zone: (a) 3DGoogle Earth view(earth.google.com) and (b) photo. In the fault's hangingwall,Mesozoic strata are sub-horizontal above thePaleozoicmetasedimentsand basement. In the footwall, basement rocks crop out up to about 1500 m high (Mount Haggier). The Hadibo Transfer Zone presents both a right-lateral and a normal displacement(see Fig. 9).

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symmetrical, althoughwide, the local tectonic history can be consideredasmonophase (if no stria overprinting is observed on fault planes). Suchwide distribution is indeed due to local fault interactions and/or faultplane irregularities. However, when the stria distribution is both wideand asymmetrical (sites 37 and 234, Fig. 5), it may be interpreted as in-dicative of stress rotations either counterclockwise (sites 37 and 234) orclockwise (site 274A, as well as site 275 in another area, Fig. 5, see cap-tion for more details). The counterclockwise rotation (020°E then160°E) of the least compressive stress is in accordancewith the abuttingof the 070°E fault against the WSW–ESE one (Jebel Kadarma, Fig. 5).

Finally, above the low-angle fault segment, an E–W extension isrecorded (Fig. 5, sites 242B and 60B). Such extension is also witnessedby N–S faults that branch on the main low-angle fault segment, East ofQalansiya.

In the eastern part of the basin, the faults are steeper (Fig. 8b). Twomain extension directions are observed, 160°E and 020°E, although in-termediates are also found (sites 207, 273, 274, 267, 265, 263, 260,Fig. 5). At site 265 (Fig. 5), chronological relationships between striae

(on the same fault plane), indicating a 020°E extension followed by a160°E extension, have been observed in the field (see more examplesin the next sub-section). Moreover, the analysis of some asymmetricfault slip distribution (see above) suggests that a counterclockwise rota-tion of the least compressive stress occurred (also at site 265, Fig. 5).

Between the two sub-basins, the NE–SW fault can be considered as atransfer fault as it separates two sub-domainswith different geometriesand amounts of extension (Fig. 8). We have little data from within thefault zone, but it might have been active during the entire period ofrifting, particularly from early stages, to accommodate the along-strikevariations in geometry and amount of extension. At site 274 (Fig. 5),two stress tensors are compatible with a N–S extension (one in an ex-tensive regime, one in strike-slip regime). Such a stress permutationmay be due to the proximity to the transfer zone (see similar interpre-tation in Fournier et al., 2007).

All the faults described above, as well as those described in thenext sub-section, are considered to be Oligo–Miocene in age. Allfault population types affect the Oligocene formations (see the

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location of the measurement site on Fig. 5). Faults observed in olderformations (Cretaceous, Paleocene, and Eocene) display similar ge-ometry and kinematics. We therefore consider that these are alsoOligo-Miocene in age.

Fig. 5.Geologicalmap of theNWpart of the Islandwith fault-slip data. Extension directionsmaiFig. 8 are represented by the two thin lines. Few stereonets have a wide and asymmetrical distr326, 328 for example, where the distribution is symmetrical around the perpendicular to the fauclockwise (28, 37, 265) or counterclockwise (274, 275). We have drawn the perpendicular to tatically lying on a line that was rotated either clockwise or counterclockwise. On these sites, thbeen represented to suggest the polyphased history. Numbers 1 and 2 attest for probable first

3.2. The Hadibo Transfer Zone

In the HTZ and surrounding area, normal faults trend NE–SW to070°E and around 110°E (Figs. 3, 4 and 9). In the transfer zone sensu

nly range between 020°E and 160°E, with several exceptions around E–W. Cross sections ofibution of fault slip: sites 28, 37, 265 and 274, 275. They differ from other sites such as 25,lt strike. The asymmetry can be interpreted aswitnessing stress rotations that were eitherhe fault strike (dashed line). One may note that the fault striae are, in these cases, system-e fault slip inversion has been performed, but manually drawn arrows for extension haveand second phases.

Fig. 6. Field photos of low-angle normal fault segments. a) East of Hadibo (Figs. 3 and 9). Eocene UmmEr Radhuma and Cretaceous limestone over low-angle normal faults, with basementrocks in the footwall. b) East of Qalansya (Figs. 3 and 5). Tertiary sedimentary rocks above low-angle normal fault with basement rocks in the footwall. The cliff is about 500 m high.

38 N. Bellahsen et al. / Tectonophysics 607 (2013) 32–50

stricto, the fault kinematics indicate that it was activated at some pointof its history as a normal fault zone trending NE–SW to 070°E, with asub-perpendicular extension, around 160°E (Fig. 9, sites 18A, 196,259B, 367). There are also evidences of strike-slip events: ENE–WSWfaults were active as right-lateral faults and record strike-slip regimeswith a probable 020°E extension (Fig. 9, sites 18B, 258).

Indeed, faults oblique to the HTZ have been observed: there arefaults trending E–W to 110°E with extension directions from N–S to020° (Fig. 9, sites 1, 194, 259A for example). Large faults also have thesame trends: around site 256, large faults are trending around 110°E(Fig. 9); at about 10 to 20 km west of the HTZ, two E–W normal faultsare mapped (sites 15, 249, and 246) with syn-rift sedimentary rocksin their hanging wall.

Relative chronologies between two sets of slickenlines on the samefault plane have been observed at few sites and provide us with a rela-tive chronology between states of stress. At sites 4, 246A, 249B, and259A (Fig. 9), striations on WNW–ESE normal faults display dip slipmovements (extension around 020°E) followed by more strike-slipleft lateralmovement (extension aroundNNW–SSE), indicating a coun-terclockwise rotation of the extension. At sites 194 and 249B (and pos-sibly at sites 246A and 259A, although more data is needed), theasymmetrical fault slip distribution (see above for details) can also beindicative of least compressive stress rotation. Moreover, at sites 138(Fig. 9), we found chronologies between striations that suggest bothclockwise and counterclockwise rotations. Finally, at sites 18 and 133(Fig. 9), we found evidences of the opposite chronology on single faultplanes.

From a structural point of view, at a larger scale, the overall geome-try of the area suggests that the 110°E toNW–SE normal faults abut ontothe NE–SW to ENE–WSW faults of the HTZ (Fig. 5 and 9). Moreover, thetwo domains east and west of the HTZ are very distinct and different.The western one presents many normal faults and associated syn-riftbasins as described above. The eastern domain is almost undeformed,except in the northernmost part. It consists of a large, partly eroded

monocline. In the northernmost part, NE–SW and NW–SE steep normalfaults offset the monocline as well as low-angle faults now observed inthe steep fault's hangingwall (Figs 6, 9 and 10). The steep normal faultsare thus considered to post-date the low-angle fault segment. Similarobservation was made in Pik et al. (2013-this volume).

These observations (abutting relationships and different domains oneach side of the transfer zone) suggest that the HTZ was active duringearly syn-rift times. Moreover, as it is strongly oblique to the global020°E divergence, it may be a reactivated pre-existing structure.

4. Discussion

4.1. Structural evolution of Socotra Island

4.1.1. Pre-existing transfer zonesBased on our structural analysis, we propose that the HTZ is a

reactivated pre-existing structure. The HTZ separates two very differentdomains and thus became active during the earliest stages of extensionprior to other nearby faults that all abut on it. Moreover, the HTZ trendsat angles of about 20°–50° to divergence, strongly suggesting that it waspre-existing and that it has been reactivated. Without tectonic inheri-tance, it is most likely that all extensional faults should trend sub-perpendicular to the divergence or the main directions of extension.

Another NE–SW-striking fault is also observed in the easternmostpart of the island (Ras Momi, Fig. 3). The fault may be Mesozoic in ageas it offsets the Permo–Triassic sedimentary rocks (600 m of throw). Itis similar to faults observed in Huqf (north of Oman) and Sawqrah(southeast of Oman) regions (Leroy et al., 2012). These NE–SW faultsthus have polyphased history and are prone to reactivation. Moreover,Proterozoic acidic and basaltic dykes in the basement at Socotra strikeNE–SW (Denele et al., 2012). These dykes may very well have beenreactivated as faults during the rifting.

Similar geometries and interpretations have been reported for theSuez Rift (Bosworth, 1994; Jarrigue et al., 1990; Moustafa, 1997) and

Fig. 7. Field photos of small normal faults. a) ESE–WNW normal fault at site 141 in Eocene Umm Er Radhuma Fm. (Fig. 5). b) ENE–WSW normal faults in Jebel Allan in Oligo–MioceneMughsayl Fm. (Fig. 5). c) ESE–WNW normal fault at site 275 in Oligocene Ashawq Fm. (Fig. 5).

39N. Bellahsen et al. / Tectonophysics 607 (2013) 32–50

East African Rift System (Corti et al., 2007; Rosendahl, 1987). In the Gulfof Suez, the Tertiary rift trends NW–SE to NNW–SSE with an extensiondirection around 060°E; transverse structures are described as transferzones that separate asymmetrical rift segments (Bosworth andMcClay, 2001; Bosworth et al., 2005; Jarrigue et al., 1990, and referencestherein). Three domains can be defined, the Darag, the October, and theZeit domains (Fig. 12) (Younes and McClay, 2002), separated by theZaafarana (ZAZ, NW–SE) and the Morgan (MAZ, N–S to NNE–SSW) ac-commodation zones, respectively (Fig. 12). These zones are oblique to

Fig. 8. Balanced cross sections through the Allan–Kadarma basin. Constant lines and thick-nesses of sedimentary layerswere respected to balance the sections. The basement area onrestored sections is not represented as equal to the basement area on field-based sections.However, the top basement length is conserved. a) Thewestern cross-section, through theJebel Kadarma, displays the low-angle normal fault segment that probably becomes steep-er at depth. The amount of extension is high (5.6 km). b) The Eastern cross-section,through the Jebel Allan, show steep normal faults that accommodatedmuch less extensionthan in the western part (0.9 km).

the divergence direction and reactivated pre-existing basement linea-ments (e.g. the Rihba shear zone for the ZAZ, Younes and McClay,2002). At a smaller scale, numerous transfer zones trend around NNE–SSW. These faults are widespread within the rift and also oblique tothe divergence.

The East African Rifts in Kenya and Tanzania (Fig. 13) formed duringMiocene times (12–10 Ma, see Ebinger, 1989). Within that rift system,along-strike variability and alternating basin asymmetry are observed,and transfer zones are well documented (e.g. Corti et al., 2007; Morleyet al., 1990; Rosendahl, 1987, and references therein) and trend NW–

SE (Fig. 10). This pattern can be mapped at small and large scales. Forexample, the Rukwa rift (Morley et al., 1992) is a large transfer zonecharacterized by oblique-slip normal faults. At a smaller scale, withinthe Lake Tanganyika rift (Fig. 13), several (small) faults trending NW–

SE are also considered as transfer faults (Lezzar et al., 2002). Similarly,in theN–S Kenya rift (Fig. 13, Gregory rift), NW–SE toNNW–SSE obliquefaults are pre-existing basement shear zones reactivated during rifting(Smith and Mosley, 1993). Similar observations were reported for theMalawi rift (Fig. 13) where pre-existing basement structures controlledthe internal rift geometry (Ring, 1994).

The direction of divergence along the East African Rifts has long beendebated (Fig. 13). Some authors conclude that the main extension isabout NW–SE (Chorowicz, 2005; Sander and Rosendhal, 1989; Scottet al., 1992; Specht and Rosendahl, 1989; Versfelt and Rosendahl,1989) while others conclude that the direction is E–W (Bosworth,1992; Ebinger, 1989; Morley, 1989). Based upon analog and numericalmodels, Corti et al. (2007) suggested that it might be E–W. GPS datashow that the present-day kinematics consists of E–WtoWNW–ESE di-vergence between Nubia and Somalia (Calais et al., 2006; Déprez et al.,2013; Fernandes et al., 2004; Nocquet et al., 2006; Royer et al., 2006;Stamps et al., 2008). Thus, depending on the orientation of the diver-gence, the transfer zones might either be parallel or oblique to diver-gence. However, considering the present-day kinematics and also theprobable Miocene kinematics, the transfer zones are oblique to diver-gence. Moreover, most of the transfer zones are strongly influenced bypre-existing older structures (see Fig. 13 and Rosendahl, 1987 fordiscussion).

To summarize, the Gulf of Aden, the Gulf of Suez and the East AfricanRifts have common characteristics among which is the geometry of thetransfer zones: they trend oblique to both the divergence direction andthe main normal faults (orthogonal to the divergence) and they are

40 N. Bellahsen et al. / Tectonophysics 607 (2013) 32–50

often basement structures reactivated as oblique-slip normal faults. Inthese rifts, it appears that there are no faults (or very few) strictly paral-lel to the divergence.

4.1.2. Early syn-rift depocentresEarly transfer zones usually control the location and geometry of

early depocentres (e.g. Bellahsen and Daniel, 2005; Lezzar et al.,2002). However, looking at the basin and depocentre distribution on So-cotra Island (Fig. 11), it appears that the main ones are located west ofthe HTZ. Moreover, the three main basins are located far away fromthis zone and the smaller ones are located at about 10 kmwestward, ex-cept the Central Basin (Figs. 3 and 11) that is probably partly controlledby the HTZ.Why is there no large syn-rift basin in the HTZ hanging wallas it seems active since early syn-rift times? Two answersmay be envis-aged: either the early syn-rift sediments were deposited and subse-quently eroded or they were not deposited in this area that may havebeen relatively elevated since early rifting times.

There are late (syn-OCT, Burdigalian) detrital sedimentary rocks un-conformably overlying the pre-rift, southwest of site 184 (Fig. 9). Thus,this area was already elevated at the beginning of OCT times. UTh/Hethermochronology suggests that the main vertical offset along a 070°Esteep normal faults (star symbol on Fig. 9, east of Hadibo) is ratherlate (around 18–20 Ma, syn-OCT, Pik et al., 2013-this volume). Finally,the syn-OCT to post-rift Miocene conglomerates (Burdigalian toLanghian) located North of the Mount Haggier, in the hanging wall ofthe normal faults east of Hadibo (star symbol, east of site 138, Fig. 9)contains many basement pebbles. No other detrital formation is ob-served between these conglomerates and the underlying Cretaceousand Tertiary carbonates. This suggests that the main uplift of MountHaggier can be dated Late Oligocene to middle Miocene.

This does not, however, allowus to determine if the area in the hang-ing wall of the HTZ was uplifted during early rifting stages or only later.Thus, it is difficult to determine if the normal faults connected very earlyto the HTZ and controlled significant syn-rift basins or not.

In any case, the amount of extension just west of the HTZ is very low(Fig. 10a),much lower than in the eastern (Fig. 10b) andwestern part ofthe island (Fig. 8). Thus, the extension is accommodated elsewhere,most likely offshore, further north, where basins may also be found.

4.1.3. Polyphase tectonic historyThe second main result is the structural evolution of the island and

especially its early structural pattern (Fig. 11). We divide this evolutioninto three stages: (1) a first stage when the main WNW–ESE faultsformed, along with the oblique reactivation of the NE–SW transferzones during extension parallel to 020°E (Fig. 11). The faults withlow-angle segments were active in both the western part (Kadarmaand Sherubrub basins, Figs. 5, 6, and 8) and the eastern part (east ofHadibo, Fig. 6, 9, and 10) of the island. The transfer zones separatedblocks with different geometries and different amounts of extension.Thus, the island was already divided into three domains: theKadarma–Sherubrub, the Central, and the Haggier (Fig. 11). The differ-ent amounts of extension (Figs. 5 and 8) accommodated in thesethree domains can be explained in two ways. Firstly, part of the exten-sion in the central domain is not expressed at the surface, and some ex-tension may, for example, be hidden beneath the coastal plain in thecentral-northern part of the island (Fig. 3). Small parts of syn-rift basinseast of this plain can be observed, indicating that there may be normalfaults hidden below. Secondly, and most likely, some of the extension

Fig. 9.Detailedmap displaying theHTZwith fault-slip data. Extension directionsmainly range bare represented by the two thin lines. Few stereonets have a wide and asymmetrical distribuinterpretation.

is accommodated offshore, north of the island. In any case, we showhere that tectonics evolve along strike, including the dip angle of themain normal faults and associated amounts of extension. During thisstage, the transfer zones (especially the HTZ) are pre-existing structuresreactivated as oblique-slip faults (right-lateral and normal), most prob-ably with dominant strike-slip motion over dip-slip motion, given theangle between the extension (020°E) and the fault zone (~NE–SW).

(2) During a second stage, with increasing extension, 110°E-trendingnormal faults became more numerous, especially in the western part.During this stage, the extension may have rotated counterclockwisefrom 020°E to around 160°E resulting in the formation and reactivationof faults trending ENE–WSW to NE–SW. The transfer zone, strikingNE–SW, was most probably also active as normal faults during thisstage, as they become well-aligned at an angle of 65° to extension. Rift-parallel normal faults (trending 070°E) formed (for example in JebelKadarma, Fig. 5). The other normal faults (WSW–ESE) were still activeand have an oblique-slip behavior. This stage probably did not last longerthan fewmillion years but is important as the onset of rift localization oc-curred during this time (Bellahsen et al., 2013-this volume). This stressrotation has been documented elsewhere in the Gulf (Bellahsen et al.,2006; Huchon and Khanbari, 2003; Lepvrier et al., 2002). Although, itcannot be unambiguously demonstrated by field data at Socotra Island,this is themost probable scenario at the scale of the entire Gulf as strong-ly suggested by field data and confirmed by analog models (Autin et al.,2010a, 2013-this volume) and numerical models (Autin et al., 2013;Bellahsen et al., 2013-this volume).

(3) During a third stage that led to the present-day fault network,the deformation migrated and localized, probably offshore to thenorth, followed by formation of the ocean-continent transition. Thethird stage cannot be documented at Socotra Island because of its prox-imal position. During this time, the extension was likely 020°E againaccording to Autin et al. (2010a). The field data (clockwise rotation ofextension from 160°E to 020°E) presented above may support such aninterpretation, however much more data is necessary to confirm it.The late syn-rift structure and the OCT emplacement may have beencontrolled by the far-field divergence (Bellahsen et al., 2013-thisvolume). Thermochronological data suggest that theHTZwas still activeat this time (Pik et al., 2013-this volume).

4.1.4. Hadibo Transfer Zone/Hadibo–Sharbithat FZ continuityThe HTZ and the Hadibo–Sharbithat FZ continuity is supported by

several observations and interpretations. (1) The FZ is aligned and con-tinuous with the HTZ (Figs. 2 and 11). This alignment can be seen in theoffset of oceanic magnetic anomalies offshore (Leroy et al., 2012). (2) Inthe distal margin between the HTZ and the Hadibo–Sharbithat FZ, a sig-nificant bathymetric escarpment defines two domains, suggesting thepresence of a major structure (Fig. 3). This escarpment roughly linksthe transfer and the fracture zones, though more detailed bathymetrycoverage and subsurface data are needed for a precise structural map-ping. (3) U–Th/He thermochronology on apatites from basement rockin Socotra Island shows that the tectonic denudation, which started atthe rifting onset, during late Eocene–early Oligocene time, continuesduring early Miocene (Aquitanian–Burdigalian) time (Pik et al., 2013-this volume). Thus, the HTZ was still active during the emplacementof the OCT, just before the formation of the oceanic transform fault.We therefore conclude that the HTZ and the Hadibo–Sharbithat FZ arestructurally connected and that the Hadibo transform fault formed atthe Hadibo Transfer Zone.

etween 020°E and 160°E,with several exceptions around E–W. The cross sections in Fig. 10tion of fault slip: sites 15, 246A, 249B. See caption of Fig. 5 for explanation and possible

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Fig. 10.Balanced cross sections through theHTZ. Constant lines and thicknesses of sedimentary layerswere respected to balance the sections. The basement area on restored sections is notrepresented as equal to the basement area on field-based sections. However, the top basement length is conserved. a) The western cross-section, through the HTZ, displays steep normalfault segments; the amount of extension is low (0.8 km). b) The Eastern cross-section, east of Hadibo, show low-angle normal fault segments that accommodatedmore extension than inthe western part. The movements along the low-angle segments was probably along a 020°E direction as attested by fault slip data from this area (Fig. 7, sites 132b, 133a, 138b, 138c)although there are indications of 160°E extension (Fig. 7, site 138a). Thus, we did not estimated the amount of extension as it may occur in the 010° direction along low-angle segments,while itmayhave occurred before, along a 160°E direction, on the steep faults. In any case, the amount of extension in this cross section ismuch larger that on thewestern one. This justifiesthat the location of the Hadibo Transfer Zone runs north of the outcrops east of Hadibo.

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4.2. Axial segmentation in the Gulf of Aden and classification of oceanicfracture zone

In theGulf of Aden, several transform faults andnon-transformdiscon-tinuities offset the Aden–Sheba oceanic spreading ridge (Figs. 2 and 18).Here, we describe in detail the segmentation of the spreading ridge andpresent a new classification of the oceanic fracture zones. As they repre-sent paleo-transform fault zones, the new classification will apply onlyto paleo-1st order transform fault zones (sensu MacDonald et al., 1991).

First, in the light of our results regarding the structural evolution onSocotra Island, we propose that a classification of oceanic fracture zonescan bemade based on the stage of rifting duringwhich they or their pre-cursors formed, and their relationships to the continentalmargin geom-etry and structures. In particular, we base our classification on whetherthe paleo-transform faults/FZ had offset and deformed either thinnedcontinental crust and OCT, OCT only, or oceanic crust only. We also em-phasize their relationship with pre-existing continental structuresreactivated during rifting.

We name FZ's generated by transform faults that deformed conti-nental crust Type 1 FZ. The Type 1 FZ offsets are usually (but not neces-sarily) large and result from a large offset between the two thinnedcontinental domains that the proto-transform joined. Such geometryimplies that part of the margin is a transform margin. Figs. 15 and 16represent conceptual models of the transform fault zone and fracturezone evolution for orthogonal and oblique rifts, respectively. Three dif-ferent Type 1 FZ are detailed below.

Fracture zones generated by transform faults that formedwithin theOCT during its formation (Figs. 15 and 16) are named Type 2 FZ. In theGulf of Aden, Type 2 FZ are observed between Shukra El Sheik and AlulaFartak FZ (Fig. 18a, b and in d'Acremont et al., 2005). These FZ offset thedistal OCT boundary but do not offset its proximal boundary.Detailed studies of each FZ are required, but they clearly display geom-etries that differ from Type 1 FZ. They aremore numerous in the centralpart of the Gulf because there the orientation of the rift and OCT areoblique to the direction of divergence (Bellahsen et al., 2013-thisvolume). Somequestions arise:What is their trend relative to the diver-gence? Are they reactivated pre-existing structures? Are they linked toany continental transfer zones? More data from distal margin is re-quired to answer these questions, in particular high-resolution 3D seis-mic data.

We name FZ generated by transform faults that formed after theonset of oceanic spreading as Type 3 FZ or simply oceanic FZ. Thistype of FZ is not linked to any syn-rift structures and formed afterrifting and OCT emplacement. Turcotte (1974) and Sandwell(1986) showed, on the basis of their spacing, that the oceanic trans-form faults are thermal contraction cracks, mainly due to thethermomechanical evolution of the oceanic lithosphere (Detricket al., 1995; Gente et al., 1995; Hooft et al., 2000; and see Gerya,2012, for a recent review). Type 3 FZ can be identified in the Gulf ofAden (Fig. 18a). They are indeed not correlated to any structure inthe continental margin or the OCT and formed after the onset of oce-anic spreading (see also d'Acremont et al., 2006, 2010).

Fig. 11. Structuralmapof Socotrawith bathymetry andOCT and oceanic domains offshore. Bathymetricmap is builtwith the compiled data fromall cruises in the database of Autin et al., inpress. During the Oligo–Miocene rifting, the island is divided into three domains: theQalansiya–Sherubrub domain (western domain), the Central domain, and theMountHaggier domain(eastern domain). Twomain transfer zones are defined: the Hadibo Transfer Zone, and the Balan–Kadarma transfer zone. An extra transfer zone is located east of the Island, the RasMomitransfer zone that bounds eastward theMount Haggier domain. The Qalansya–Sherubrub domain ismainly characterized by low-angle normal fault segments; the amount of extension ishigh. The eastern boundary of this domain, the Balan–Kadarma transfer zone is a rather diffuse one. It is clearly observed in the Allan–Kadarma basin. Further south, it may cross theSherubrub–Balan basin. However, the continuity in the basin geometry at the suggested location of the transfer zone shows that there is a structural complexity that needs furtherwork. The central domain is structured by steep normal faults and does not exhibit high amounts of extension. The Mount Haggier domain is weakly deformed except in the northernpart, where low-angle segments are observed, but where probably most of the deformation is located further north offshore. Early depocentres are in light yellow. After the riftingstage, the OCTwas emplaced andwas strongly segmented, aswas the oceanic domain. In the distal margins, there is a clear bathymetric signature of the HTZ. Surprisingly, the distal struc-ture bounds a structural high westwards and a relative low eastwards. Further work is needed to better characterize the structural evolution of this area.

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Three cases of syn-rift Type 1 FZ can be distinguished, we describethem below.

4.2.1. Type 1-C (syn-rift) FZThe first Type 1 FZ is named Type 1-C (C for continental, see below):

such FZ formed as a transform fault that is linked to a blockage in alongstrike ridge propagation. In this case the transform fault is a “continentaltransform fault” that, at some point, separated an oceanic domain froma continental domain. Thus, it was located at the tip of a propagatingspreading center (Figs. 15 and 16). Such a setting necessitates the pres-ence of “continental” transform fault zone to accommodate the differ-ence in the amount of divergence. This type of FZ can be found inseveral locations globally.

In the SouthAtlantic Ocean (Fig. 1), onset of seafloor spreading alongthe ridge between the Ascension and the Rio Grande FZ is younger thanbetween the Rio Grande and the Falkland FZ (see Moulin et al., 2010).Thus, the Rio Grande FZ likely formed as a continental transform faultzone, which eventually corresponded with the northern tip of the oce-anic spreading center. Such setting can last during a significant timeonly if there are continental transverse structures that accommodatethe important differential opening (see reconstructions in Moulinet al., 2010; and references therein). The Rio Grande FZ is, for example,

genetically related to continental transfer zones, namely the Pernambu-co shear zone (Basile et al., 2005; Moulin et al., 2010) and Patos shearzone (Darros de Matos, 1999; Françolin and Cobbold, 1994; Sénantand Popoff, 1991) in South America and the Sanaga and N'Gaoundereshear zones in West Africa (Basile et al., 2005; Moulin et al., 2010). Forthe Ascension and the Romanche FZ, the structural evolution may besimilar although less straight forward (see a discussion in Basile et al.,2005 for Romanche FZ; compare Mohriak and Rosendahl (2003),Blaich et al. (2008), and Meyers et al. (1996) for Ascension FZ).

In the North Atlantic Ocean, similarly, the Senja–Greenland FZ (seeMosar et al., 2002 and references therein) separates two oceanic do-mains with different ages of spreading onset. In central Atlantic, thesame is true for the Charlie-Gibbs, 15°20′N-Guinean, and Jacksonville-Cap Vert FZ (see Labails et al., 2010). The Owen FZ (easternmost Gulfof Aden, Fig. 1) is alsomost probably a Type 1-C FZ but with the notabledifference that it likely formed within an oceanic domain (Mountainand Prell, 1990). These faults are parallel to divergence and their loca-tion and trend are probably (but not necessarily) determined by pre-existing structures. Indeed, as these faults are not particularly well ori-ented to be reactivated in an extensive stress field (because they trendparallel to the least compressive stress), it is more likely that they arereactivated pre-existing faults or shear zones. Finally, these transform

Fig. 12. Structural map of the Gulf of Suez (after Younes andMcClay, 2002). Themain nor-mal faults are trending around 150°E. Accommodation or transfer zones divide the rift intoseveral domains with opposite sense of dip of half graben. The Zaafarana AccommodationZone (ZAZ, NW–SE) separates the Darag and the October basins and probably reactivatesthe Rihba shear zone; the Morgan Accommodation Zone (MAZ, N–S to NNE–SSW) sepa-rates the October and the Zeit basins. Many smaller transfer zones trending parallel tothe MAZ can be observed in the Zeit basin.

Fig. 13. The southern East African Rift System (after Rosendahl, 1987 and Corti et al.,2007). The global trend of the rift is N–S (as in Turkana, Gregory, Malawi rifts, for exam-ple), except in transfer zones where it trend NW–SE (as in the Rukwa rift, for example).The transfer zones reactivated pre-existing structures in the basement (such as theAswa lineament). The direction of extension is still debated. Present-day divergencebased on GPS is around 100°E–110°E, close to the long term E–W extension (e.g.Bosworth, 1992; Corti et al., 2007; Ebinger, 1989; Morley, 1999). However, NW–SE exten-sion has been proposed (e.g. Rosendahl, 1987; Chorowicz, 2005). Our study that showsthat transfer zones are commonly oblique to the divergence direction would suggestthat the syn-rift extension was E–W to WNW–ESE.

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faults clearly form during rifting (at least for the continental domain sit-uated on one side of the transform fault) and thus can be named syn-riftFZ.

In the Gulf of Aden, the Shukra El Sheik FZ is a Type 1-C FZ from theonset of spreading (17.6 Ma) and until 10 Ma, there was no spreadingwest of this fault (Leroy et al., 2012). Thus,motion on this fault activatedbecause the westward ridge propagation was blocked there. This majorfault extended southwestward into Afar where it was probablyconnected to the other active faults (Audin et al., 2004) (Fig. 18).

4.2.2. Type 1-T (syn-rift to syn-OCT) fracture zonesA second Type 1 FZ is named Type 1-T (T for transfer) when they

formed at continental transfer zone, as in the Hadibo–Sharbithat FZcase (Figs. 2 and 11). In this case, a transform fault formedwhere a con-tinental transfer zone was active during rifting (this study) and proba-bly during OCT formation (Pik et al., 2013-this volume), leaving afracture zone behind. The FZ and transfer zone have different trends(Figs. 11 and 17). In Figs. 15 and 16, the Type 1-T FZ are located in azone that was acting as a transfer zone during rifting. The extension isaccommodated by faults oblique to both the extension and the mainrift-related faults. As shown above, such transfer zones are very often

reactivated pre-existing faults. In the East African Rift System, theRukwa rift zone (Figs. 13 and 15a) is an oblique rift zone where slip isoccurring along old reactivated lineaments. If this zone eventually be-comes an ocean basin, it is most likely that a Type 1-T FZ would form.In such case, one could not find any faults parallel to the divergencewithin the continental margin.

These Type 1-T FZ may form either during the latest rifting stages orduring earliest OCT times. If the OCT proximal boundary is clearly offset,it probably means that transform fault formation is slightly older thanthe onset of OCT formation, thus late syn-rift (Fig. 17a, upper panel). Ifthe continental transfer zone and the FZ join at the proximal boundaryof the OCT, the formation of the transform fault zone may be dated tothe age of the transition between syn-rift and syn-OCT (Fig. 17a, middlepanel). If the offset of the OCT proximal boundary is controlled by thecontinental transfer zone (Fig. 17a, lower panel), it is most likely thatthe transform fault formed during OCT times. In the Hadibo–SharbithatFZ, the OCT proximal boundary is clearly offset (Figs. 2 and 13, Leroyet al., 2012) and this offset suggests that the paleo-transform faultformed during syn-rift to syn-OCT times (compare Figs. 11 and 17a).This provisional conclusion requires more data in the distal continentaldomains in order to be confirmed.

Fig. 15. Conceptual model of oceanic FZ in relation with rifting geometry, for orthogonalrifts. Three kinds of type 1 FZ are differentiated: Type 1-C when they derive from a “con-tinental transform zone” that has been separating an oceanic basin from a continentalone (as the Levant fault zone for example), Type 1-T when their location is influencedby an (often pre-existing) oblique continental transfer zone (re)activated at the onset of(or during) the rifting, and Type 1-P when their location it due to a pre-existing continen-tal structure reactivated at the very end of the rifting. Type 2 FZ are defined as deformingonly the OCT lithosphere: it means that the proximal boundary of the OCT is not offset,while the distal one is offset. Type 3 FZ initiate after the oceanic spreading onset and arerelated to evolving stresses within the oceanic crust.

Fig. 14. The Norwegian continental margin. Structural map simplified after Brekke (2000)andOlesen et al. (2007). Themargin aswell as a number of important normal faults trendsNE–SW, except within the JanMayen accommodation zone, that are aligned along the JanMayen fracture zone. In this zone, the normal faults are oblique to both the divergence di-rection (Mosar et al., 2002, and references therein) and the future FZ.

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4.2.3. Type 1-P (syn-rift) fracture zonesOn the Mid-Norwegian continental margin, a final Cretaceous–

Tertiary rifting stage led to continental breakup and opening of thenortheastern Atlantic Ocean during Early Eocene times (for recent stud-ies see for example Brekke, 2000; Skogseid et al., 2000; Mosar et al.,2002). The long Jan Mayen FZ separates the margin into two basins:the Vøring basin in the north (Brekke, 2000) and the Møre basin inthe south (Gabrielsen et al., 1999) (Fig. 14). This FZ is thought to contin-ue onshore (Doré et al., 1997; Olesen et al., 2007) in the form of a Pro-terozoic lineament. The Norwegian continental margin is affected bynumerous normal faults sub-perpendicular to the future transformfault, thus perpendicular to the Cretaceous–Tertiary divergence (seeMosar et al., 2002 for the various directions of divergence). However,in the accommodation zone, i.e. around the supposed trace of the JanMayen lineament, normal faults are strongly oblique to divergence(NW–SE) and strike N–S to NNE–SSW (Fig. 14). During rifting, the de-formation and the depocentre locations across the Jan Mayen transferzone clearly evolved from a diffuse accommodation zone (with someorientations that are not perpendicular to the divergence) to more lo-calized transfer zone (with divergence-perpendicular faults, compareFigs. 7 and 13 in Brekke, 2000, for Cretaceous and Figs. 15 and 16 forearly Tertiary times). This clearly indicates that an accommodationzone existed during rifting and that this zone evolved into a transfer/transform (see below) fault zone during the late rifting phase; i.e.when the amount of extension became large. Note that this is consistentwith the review of Acocella et al. (2005) on transfer zones, where it issuggested that transfer zones form during late syn-rift phases, while ac-commodation zones are active during early syn-rift times. However,here the late syn-rift structures that segment the margin are named

Fig. 16. Conceptualmodel of oceanic FZ in relationwith rifting geometry, in the specific case of oblique rifts. Same legend that is in Fig. 15. Three Type 1 FZ are differentiated: Type 1-C, Type1-T, and Type 1-P. Type 2 and Type 3 FZ are also represented.

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transform fault (Fig. 17b), as it divergence parallel and represent aproto-plate boundary.

We name such faults syn-rift or Type 1-P FZ (P for pre-existing)when the transform fault location is influenced by a pre-existing faultparallel to the divergence (Figs. 15 and 16). This does not necessarilymean that such a transform fault was active in the proximal margin. In-deed, such a pre-existing faultmay be reactivated only in the distalmar-gin, where the continental crust is extremely thinwithin a narrow zone.This probably occurs during or slightly before OCT formation and/ormantle exhumation. The timing of paleo-transform fault zone formationdepends on the structural relationships between the FZ and the proxi-mal boundary of the OCT (Fig. 17b), similarly to the discussion aboutType 1-T FZ (see 4.2.2). In Norway, the JanMayen FZ clearly strongly off-set the proximal boundary of the OCT (Fig. 14) and ismost likely relatedto a paleo-transform fault zone that formed during the latest stages ofrifting.

In theGulf of Aden, Alula Fartak FZ is a large-offset FZ. There is amaincoastal offset and early syn-rift normal faults in Qamar basin are also off-set by the FZ. There is no clear continuity of the FZ in the proximal do-main (Fig. 2), near the Yemen/Oman border; this led Taylor et al.(2009) to propose that it nucleated during the onset of spreading orsoon after. However, Autin et al. (2013-this volume) suggest that thisFZ formed because of the presence of the pre-existing Jiza/Qamar–Gardafui Mesozoic basin that implied a stronger crust. There, the strainlocalization occurred on each side of the pre-existing basin (d'Acremontet al., 2005) as a consequence of rift obliquity; this setting necessitated atransform fault when the stretching localized during the late syn-rift/OCT. In this case, the FZ does not correspond to any continentaltransfer zone but its location is “inherited” from the rifted basin pattern.The FZmay also be a Type 1-P FZ, in our classification, as NNE–SSWpre-existing structures were and are present in the Arabian shield (Al-Husseini, 2000; Denele et al., 2012; Stern and Johnson, 2010).

Fig. 17.Detail of a) Type 1-T and b) Type 1-P FZ in distal margins, structures parallel to the oceanic FZ and transform fault zone are explicitly named FZ (and not transfer zones as in Fig. 2).Continental transfer zones are usually oblique to both the divergence and the normal faults. See text Section 4.2. for a discussion of the timing of transform fault zones/oceanic FZ initiation.

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5. Conclusions

In this geological study of Socotra Island (Yemen), we have de-scribed in detail the structural evolution of an oblique transfer zone ina continental margin. This zone was active as an oblique-slip normalfault zone and is not parallel to the plate divergence, even if during itshistory, some pure strike-slip movements have occurred. We alsoshow that the faults in this zone are pre-existing structures that werereactivated at the onset of rifting. Similar features can be observed inmany extensional settings such as in the Gulf of Suez and the EastAfrican Rifts. On the basis of our findings, we propose a new classifica-tion of oceanic FZ based on their spatial-temporal relationships withsyn-rift structures. We distinguish three distinct Type 1 FZ, i.e. oceanicFZ that are related to paleo-transform fault zones that deformed conti-nental margins: Type 1-C that form at “continental transform faultzones” that are syn-rift in age, Type 1-T that form at continental transfer

zones during late syn-rift to early syn-OCT times, and Type1-P that formduring late syn-rift to early syn-OCT at pre-existing structures that werenot reactivated during early rifting. Type 2 FZ initiated during the OCTemplacement; Type 3 FZ form after the onset of seafloor spreading.Our classification requires the acquisition of more high-resolution dataparticularly in distal oceanic domainswhere transfer zones and fracturezones connect, to better constrain the geometry, the structural evolu-tion, and the timing of transform fault zone formation.

Acknowledgments

The comments by C. Ebinger and an anonymous reviewer and thesignificant work of G. Peron-Pinvidic and P.T. Osmundsen as invited ed-itors greatly improved the initial version of this contribution. This workbenefited from financial support from GDR Marges, Action Marges(CNRS-INSU, TOTAL, BRGM, IFREMER), ANR YOCMAL and Rift2Ridge.

Fig. 18. Segmentation of the Gulf of Aden through time. a) The Gulf of Aden rift at around 20 Ma, at the time of OCT onset (purple line). Light orange color is for the pre-existingMesozoic basins, bright orange represents the OCT. Type 1 FZ (in red) that initiated at that time are represented by: a Type 1-C FZ in the west (S. El S., Shukra El Sheik), a Type 1-T FZin the east (S.H., Socotra Hadbeen and H.S. Hadibo–Sharbithat), and possibly a Type 1-P FZ in the center (A.F., Alula Fartak). b) The Gulf of Aden rift at the end of the OCT (yellowareas) formation at ~18 Ma. Type 2 FZ (syn-OCT, in blue). c) Present-day Gulf of Aden. Type 1 FZ are still active (except Shukra El Sheik). Some Type 2 FZ died out (in theeastern part). Type 3 FZ were initiated after the onset of oceanic spreading.

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49N. Bellahsen et al. / Tectonophysics 607 (2013) 32–50

The authors thank Y. Denele for fruitful discussions. GSMRB is thankedfor technical support.

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