Seismic stratigraphy and structure of the Caribbean igneous province

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Tectonophysics 283 (1997) 61-104

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Seismic stratigraphy and structure of the Caribbean igneous province A . M a u f f r e t , S. L e r o y * Dept. de Geotectonique, CNRS URA 1759, Case 129, 4 Place Jussieu, 75252 Paris Cedex 05, France

Received 15 May 1996; accepted 23 April 1997

Abstract Multichannel seismic reflection profiles reveal the tectonic and stratigraphic complexity of the Caribbean province. A smooth reflector B" forms the top of widespread Cretaceous volcanic flows drilled during DSDP Leg 15 and ODP Leg 165. In addition we also show other seismic evidence of volcanism (sills and dipping horizons). The southern parts of the Venezuela and Colombia basins are underlain by rough basement that shows many characteristics of oceanic crust. Several thick volcanic plateaus separated by deep basins with thin crust comprise the Caribbean volcanic province. In the Venezuela basin only normal oceanic crust is present, but in other basins this crust is underlain by underplal~l volcanic material. The crust of the Venezuela basin is less than 5 km thick and was probably formed in the Pacific during the Jurassic. Refraction data indicate that a thin layer, 2V, is formed by original oceanic crust overlain by a thin volcanic layer. An intra basement reflection (sub-B") marks the top of original oceanic crust that is sandwiched between an upper volcanic layer and lower underplated material. This underplated layer forms a very thick layer, 3V, beneath the Beata and Nicaragua volcanic plateaus. The upper part of layer 3V is gabbroic and outcrops on the Beata ridge. The thickening occurs mainly in the high-velocity lower part of layer 3V and is attributed to the presence of magnesian rocks (picrites or ultramafic cumulates). A highly reflective horizon (R) is located at the top of this layer. We present a model where the Galapagos plume is 2500 km wide but has narrow heads that formed the Caribbean volcanic plateaus separated by deep basins. The Caribbean igneous province is compared to similar volcanic features in the Indian Ocean (Kerguelen plateau) and in the Western Pacific Ocean (Ontong Java plateau, Nauru and Pigafetta basins. The thick plateaus of the Caribbean province resemble the Kergu¢len and Ontong Java plateaus, whereas the deep basins present some similarities with the Nauru and Pigafetta basins. The Caribbean igneous province is younger (90-75 Ma) than the Indian and Pacific plateaus (110-130 Ma); however, several geological evidences from land and geophysical offshore observations indicate that the onset of volcanic activity in the Caribbean province could be as old as 115 Ma. Keywords: Caribbean plate; ocean basins; plateau basalts; Cretaceous; underplating; seismic stratigraphy

1. Introduction The general characteristics and principal topographic features of the Caribbean region are shown in Fig. 1. To simplify the description we propose new names, mainly derived from Indian tribes, and subdi* Corresponding author. E-mail: [email protected]

vide the main basins (Fig. 1). The former Colombia basin is divided into two units: the Colombia basin to the south and the Haiti sub-basin to the north. A complex buried basement unit, which includes the Warao rise, separates these basins. Three features abut obliquely on the Hess escarpment: the Mono rise (Bowland and Rosencrantz, 1988; Bowland, 1993), the Chibchas rise where the ODP Site

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Taino ridge. The former Venezuela basin is divided into three units: the Dominican and Puerto Rico subbasins, and the Venezuela basin that is located in the southeastern part of the study area. Previous studies of seismic stratigraphy and structure of the Caribbean region have consisted of descriptions of seismic units in the Venezuela basin

A. Mauffret, S. Leroy / Tectonophysics 283 (1997) 61-104

(Ladd and Watkins, 1980; Ladd et al., 1990) and the Colombia basin (Lu and McMillen, 1982; Bowland, 1993), in addition to interpretation of refraction data (Officer et al., 1957, 1959; Ewing et al., 1960; Edgar et al., 1971; Ludwig et al., 1975; Houtz and Ludwig, 1977; Diebold et al., 1981; Case et al., 1990). Two prominent reflection horizons (A" and B") were mapped in the Venezuela basin. Although attempts to follow these reflections beneath the thick sediment pile of the Colombia basin have met with varying success, the Carib Beds (Caribbean typical reflectors, A" and B") have been identified in the southwestern part of this basin (Bowland, 1993). Sub-B" reflections were identified with multi-channel data in the Venezuela basin (Ladd and Watkins, 1980; Diebold et al., 1981), the southern Beata ridge (Hopkins, 1973; Stoffa et al., 1981) and in the western part of the Colombia basin (Bowland and Rosencrantz, 1988). DSDP and ODP drilling in the Caribbean region (Bader et al., 1970; Edgar et al., 1973a; (Scientific Party, 1996) sampled much of the section down through reflector B". Horizon A" is composed of Middle Eocene chert and limestone; horizon B" correlates with Coniacian to Santonian basalt flows and sills, except at DSDP Site 152 where the basalt is younger (Campanian). At DSDP Site 1001 the same Campanian basalts were sampled (Scientific Party, 1996). Refraction and reflection results (Edgar et al., 1971; Ludwig et al., 1975; Houtz and Ludwig, 1977; Diebold et al., 1981; Case et al., 1990) indicate the presence of oceanic crust with irregular topography in the southwestern Venezuela and Colombia basins (Biju Duval et al., 1978; Diebold et al., 1981; Bowland and Rosencrantz, 1988). However, the largest part of the Caribbean acoustic basement is formed by the smooth B" horizon which correlates with the Cretaceous basalt flows and sills drilled during Leg 15 and ODP Leg 165. These observations provide evidence that the Caribbean province is a large igneous province (LIP) that covers an area of 800,000 km 2. It probably formed during the Cretaceous, in the Pacific on the Farallon plate, and above the Galapagos hot spot (Duncan and Hargraves, 1984). As this thick and hot volcanic plateau migrated eastward it choked the Pacific subduction zone, causing the arc system to

63

flip (Pindell et al., 1988). Many pieces of this plateau were subsequently accreted to the continent (Donnelly et al., 1990). Since Eocene time, the Caribbean plate has continued to migrate towards the east relative to the American plates (Pindell et al., 1988; Pindell and Barrett, 1990). To clarify some questions about Caribbean crustal structure and to survey the northern Colombia basin and the Beata ridge for the Ocean Drilling Program (ODP), we carried out several multi-channel surveys during a cruise named Casis. In this paper, we present the main results of a regional seismic survey, as well as some portions of two detailed seismic surveys of the Beata ridge.

2. Data collection and processing The multichannel seismic (MCS) data used in this paper were acquired during the 1992 Casis cruise (Fig. 1). The data were collected using a 96-channel, 2.4-kin long hydrophore streamer. The R/V Nadir fired an array of 6 GI airguns in harmonic mode, total volume 450 cubic inches (7.36 1), at a pressure of 140-180 bar every 20 s. The source array was maintained at 18-20 m depth to obtain low frequencies. Although the volume of the guns was low we obtained an excellent penetration. The multichannel seismic profiles were processed with standard techniques: deconvolution with stacked signature of six guns, semblance, 24-fold stack and fk migration.

3. Seismic stratigraphy Our main purpose is to define and characterize the main stratigraphic sequences and to extend the coverage to areas that were poorly explored before the Casis cruise. We focus on the various facies related to the Cretaceous volcanic event. The seismic stratigraphy of the central Venezuela basin was defined by Ladd and Watkins (1980). DSDP results reveal the presence of four main seismic intervals (Fig. 2). An upper seismic interval in Hole 31 (Bader et al., 1970) and Holes 146 and 153 (Edgar et al., 1973b) corresponds to lithic unit 1, a chalk marl ooze and clay of Early Miocene to Recent age (eM, Fig. 2). The second seismic interval corresponds to a Middle Eocene to Lower Miocene radiolarian chalk and radiolarian chalk unit. Hori-

A. Mauffrer S. Leroy /Tectonophysics 283 (1997) 61-104

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zon A" (Fig. 2) forms the base of this interval. The third seismic interval corresponds to lithified chalks, cherts, limestones and black shales which rest upon Santonian to Coniacian basalts. Horizon B" (Fig. 4) correlates with these basalts. In addition, we have identified in the Aruba gap area a lower sedimentary unit between an equivalent of B" and a deeper horizon named V (Fig. 2; Leroy et al., 1996; Leroy and Mauffret, 1996). The compressional velocities of seismic intervals were well determined during DSDP Leg 15 using a combination of laboratory studies on core samples, and correlations between the main reflectors (A" and B") and the corresponding lithic horizons. We were able to perform many semblance analyses (Taner and Koehler, 1969) in the Aruba gap and the eastern Beata ridge areas, where there is a dense grid of seismic profiles, and in Fig. 2B we present the statistical results of analyses performed in the Aruba gap. The compressional velocity varies from 1.8 km/s for the interval sea floor-A" reflector to 4.8 km/s in the volcanic acoustic basement (Fig. 2B). These velocities are similar to those obtained (1.9 km/s to 4.7 km/s) by migration before stack performed in the same area (Leroy et al., 1996). It should be noted that RMS (root mean square) velocities determined by the semblance technique are always higher than the DSDP results (Fig. 2). On the other hand, these velocities are lower than the refraction velocities because ray paths used for the latter are mainly horizontal, whereas those providing the RKIS velocities are almost vertical. 4. Distribution of seismic facies and structural provinces in the Caribbean Sea We have identified three main reflectors, dated (Saunders et al., 1973) Early Miocene (eM), Middle Eocene (A") and Late Cretaceous (B"). The thicknesses of seismic intervals vary within the different structural provinces; reflectors may in places lose their seismic characteristics, along seismic profiles. This is especially true for the Colombia basin, where the A" reflector changes in character to the east and north but has normal characteristics in the western part (Fig. 2; Bowland, 1993).

65

4.1. Lower Nicaragua rise

The lower Nicaragua rise, located between the Pedro and Hess escarpments (Fig. 1), is relatively poorly explored. Several submarine hills probably have a volcanic origin (Case et al., 1990). The Casis 01 seismic line (Fig. 3) shows, over a short distance, the previously defined typical seismic sequence (A" and B"). Several reflectors can be seen below B" (white arrow, Fig. 3). These seismic units outcrop along the Hess escarpment (Fig. 4), as demonstrated by DSDP Site 152 and ODP Site 1001 which sampled Middle Eocene cherts and Campanian volcanic flows (Edgar et al., 1973a; Scientific Party, 1996). Several reflectors are observed below the Campanian basalt flows (white arrows, Fig. 4, enlargement A). A" and B" reflectors are shallower on the lower Nicaragua rise than in the Haiti sub-basin (Fig. 4). A. Droxler, who recently performed a single-channel seismic survey in the ODP Site 1001 area, proposes (in: Scientific Party, 1996) that the Hess escarpment is a recent normal fault. The Casis line has a small vertical exaggeration (about 3.3 at the sea floor) and normal faulting is not evident around shot point 5200. However, we cannot rule out a low-angle listric fault. We observe a fan-shaped wedge (enlargement B) but we do not know if this wedge is a synrift fill or volcanic flows. The continuity of the base of the wedge and a reflector below the small hill (white arrow) favours a volcanic origin. In both hypotheses, the wedge is an old feature. The lower part of the Hess escarpment is characterized by a small hill (between shot points 5500 and 5600) with a magnetic anomaly (Fig. 4). Recent sedimentary fill of the Colombian basin onlaps this probable volcano. This seismic configuration provides good evidence that the Hess escarpment is an old feature which has been inactive since the Late Cretaceous (Case et al., 1990), although the southwestern part of the Hess escarpment was clearly active in recent times (Bowland, 1993). On the other hand, the normal faulting, identified by A. Droxler in the ODP Site 1001 area, looks impressive but it is probable that at least some of the disrupted sediments on the Hess escarpment in this area are the result of mass wasting (L. Abrams, written pers. commun.). In conclusion, although we cannot exclude the possibility that post-Eocene normal faulting offsets the A"-B"

66

A. MauJJ~'et, S. Leroy /Tectonophysics 283 (1997) 61-104

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interval, the evidence for such faulting is weak in the portion of Hess escarpment crossed by the Casis line but other parts of the Hess escarpment may be faulted.

4.2. Haiti sub-basin

The northern part of the Haiti sub-basin is poorly known. One seismic profile (Fig. 5; Bien-Aim6

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Momplaisir, 1986) shows deep acoustic basement (8 s TWTT). Refraction data (Fig. 5, R36 W; Ewing et al., 1960) indicate that the crust is thin beneath the basin. Compressional deformation is evident on the upper part of the margin (lie ~ Vache anticline and adjacent syncline, Fig. 5; Bien-Aim6 Momplaisir, 1986) but onlapping of a small tectonic wedge at the base of the scarp by the sedimentary layers of the basin demonstrates that the compression is presently restricted to the upper part of the margin. The upper Hispaniola margin is deformed by transpression related to the collision of the Beata ridge with the northern part of Hispaniola and the left-lateral motion of the Enriquillo fault (Mercier de Lepinay et al., 1988; Heubeck and Mann, 1991). The southern part of the Haiti sub-basin is characterized by a shallow and rough acoustic basement with several prominent highs and lows (Fig. 6). It is evident that this province is related to the Beata ridge. Beneath acoustic basement we observe several areas in which reflectors dip southwards (1 s TWTT thick, about 2.5 km; Fig. 6). The Haiti sub-basin is separated from the Colombia basin by the Warao rise (Fig. 1). This rise has a 14.4-km-thick crust (refraction line 14 NE; Ewing et al., 1960). The crust is thinner to the south (11.9 km in refraction line 14 SW; Ewing et al., 1960). The 10 s TWTT deep reflection in Fig. 6 clearly correlates with the Moho discontinuity indicated by the refraction results. In the Haiti sub-basin we tentatively identify reflector A", but we do not have enough data to be confident with this identification. The slightly deformed southern edge of the Warao rise (shot point 9250, Fig. 6) is onlapped by post-Lower Miocene sediments. Early Miocene deformation is also recorded at the boundaries of the Beata ridge (Mauffret et al., 1994; Leroy and Mauffret, 1996; Mauffret and Leroy, 1997). 4.3. Colombia basin

The western part of the Colombia basin has been well described by Lu and McMillen (1982) and Bowland and Rosencrantz (1988). We focus our description on the eastern part that is crossed by the Casis survey. Acoustic basement plunges regularly towards the Colombia deformed belt (Fig. 7; Houtz and Ludwig, 1977) and some dipping reflectors are again evident

69

beneath B" reflector (Fig. 7). The dipping reflector package is 2.5 km thick (1 s TWTT with a velocity of 5 km/s). In the deep part of the Colombia basin, sediments of the Magdalena deep-sea fan (Kolla et al., 1984) are affected by prominent growth faults (Figs. 8 and 9). The chaotic uppermost layer contrasts with the well-layered underlying seismic units (Figs. 7-9). These differences are related to the formation of the Magdalena deep-sea fan (Kolla et al., 1984). In the southernmost part of Cas C02, we observe a transition between a smooth B" reflector and rough basement (Fig. 8). This rough basement corresponds to oceanic crust defined by Bowland and Rosencrantz (1988) in the western part of the Colombian basin. However, beneath this rough basement some reflectors dip towards the north (Figs. 8 and 9) and this is taken to indicate that an underplated volcanic unit probably underlies true oceanic basement, as discussed later. Horizontal and dipping reflectors beneath B" reflectors are particularly well displayed in Fig. 8. Deep reflections at 10 s TWTT may correlate with Moho (Figs. 8-10). These reflectors are evident on the stacked seismic sections but they are not well displayed in the migrated sections (Figs. 8-10). The easternmost part of the Colombia basin is characterized by several buried highs. The structure shown in Fig. 10 is clearly the southern extension of the DSDP Site 151 ridge (Fig. 1). 4.4. Beata ridge

The northern part of the Beata ridge impinges on central Hispaniola (Ladd et al., 1981; Mercier de Lepinay et al., 1988; Jany, 1989; Heubeck and Mann, 1991), where part of the Caribbean large igneous province (LIP) has been uplifted and accreted (Maurasse et al., 1979; Sen et al., 1988) to the island. The central part of the ridge has been described by Fox et al. (1970), Roemer et al. (1973), Fox and Heezen (1975), Moore and Fahlquist (1976), and Holcombe et al. (1990). A complete structural interpretation of this ridge will be presented elsewhere. The Aruba gap is a strait between the Colombia and Venezuela basins which separates the southern tip of the Beata ridge from the south Caribbean deformed belt (Fig. 1). A prominent reflector (sub-B") was described in this area beneath the smooth B" reflector

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(Hopkins, 1973; Stoffa et al., 1981). The velocity of the B"-sub-B" interval is high (4.5-5 kin/s; Stoffa et al., 1981). This sub-B" feature, an internal reflection 2.5 km below the top of the oceanic plateau, is evident on the Casis lines (Figs. 11 and 12; Mauffret et al., 1994; Leroy and Mauffret, 1996) and particularly well displayed in the pre-stack migrated seismic lines (Fig. 12; Leroy et al., 1996). The Pecos fault zone (Fig. 11) is placed at the boundary between the Beata ridge and the Colombia basin. Previous studies (Mauffret et al., 1994; Leroy and Mauffret, 1996) showed that this structure is an old structure that was reactivated during Miocene compression. A profile crossing this structure (Fig. 11) clearly reveals Carib beds (A" and B") in the deep Colombia basin. Therefore the smooth B" exists in several parts of the eastern Colombia basin that are not underlain by rough oceanic crust. The A"-B" interval is thin and most of the sedimentary fill of the Colombia basin is post-Eocene in age. Subduction of the Caribbean plateau beneath the south Caribbean deformed belt is illustrated in Fig. 13. The decollement is located along the Early Miocene reflector and the oceanic plateau is undeformed. However, we observe landwards-vergent reflectors at the toe of the prism which contrasts with seaward vergence for the reverse faults of the inner prism. The inner prism pushes and tilts the layer of the outer prism and a triangular zone developed at the leading edge (Price, 1986; Alavarez-Marron et al., 1993). This triangular zone and seaward vergence of the toe are probably related to overpressured fluids along the decollement (MacKay et al., 1992). Moreover, several mud diapirs were observed in the upper part of the deformed belt (Vitali, 1985). In addition to the prominent sub-B" reflector, we note several reflections that extend into volcanic crust. The correlation (Fig. 12) between seismic profile Cas 02 and refraction line 80 of Edgar et al. (1971) indicates that a deep reflection may correspond to the Moho discontinuity. The refraction line is located on the Beata ridge (Fig. 1), where the oceanic plateau is shallower and the crust thicker (10.2 kin) than in the Aruba gap (6.6 km). We note that thickening of the plateau is mainly restricted to the deep part of the high-velocity crust (6.5 km/s), although the thickness of the upper crust is the same in the two areas (3.5 km, Fig. 12). The high-velocity deep layer is 6.7 km

75

thick on the Beata ridge, but only 2.8 km thick in the Aruba gap. North of DSDP Site 153 and below B" we identify a new reflector, named V (Figs. 2, 12 and 14; Leroy et al., 1996) that shows prominent but discontinuous reflections. This type of seismic configuration, particularly evident on the migrated before stack seismic profile (Fig. 12), is characteristic of volcanic sills. The smooth B" that is the top of volcanic flows loses its reflective character and a sedimentary layer appears between B" and V. A similar seismic configuration can be seen in the Venezuela basin (Fig. 15). Several dipping reflectors can be seen on the seismic profile shown in Fig. 14, with apparent dips to either the north or the south. Some reflectors seem to cross one another. Certain diffractions, not removed by migration, cannot be excluded, but the crossing reflectors seem to be real. A possible explanation is the presence of successive volcanic flow units with opposite dips. A buried basement high with a magnetic anomaly (SP 10200; Fig. 14) could be a volcano. We note that dipping reflectors beneath this probable volcano do not dip away of this feature, as would be expected if the flows had erupted from the volcano. A more probable volcanic centre is a depression between shot points 10100 and 10000 rather than the adjacent high. Later, we will compare these dipping reflectors to those of the Kerguelen plateau (Rotstein et al., 1990; Schaming and Rotstein, 1990; Schlich and Wise, 1992; Schlich et al., 1993). The edges of the Beata ridge were reactivated during the Early Miocene and in recent times by prominent transpressional tectonics in the Aruba gap (Mauffret et al., 1994; Leroy and Mauffret, 1996). The eastern flank of the Beata ridge is also affected by this deformation. A 'push-up' (Taino ridge, Fig. 1), which has elevated typical Carib beds, was identified during the Casis cruise. The smooth B" reflector overlies a layered sequence which resembles sub-B" reflections (Figs. 16 and 17). A deep reflection, at 7 to 8 s TWTT, is evident in this area (horizon R, Figs. 16 and 17). The B"-horizon R is 6 km thick and the depth of horizon R varies from 10.5 km on the Beata ridge to 11.5 km in the Dominican subbasin. The velocities above this reflector, determined using the semblance technique (6.2 km/s, Fig. 16 and 6.1 km/s, Fig. 17) indicate that this reflector is located deep in the crust. Similar reflections, 8 s

76

A. Mat~ff?et. S. Leroy / Tectonophysit's 283 (1997) 61 - 104

A. Mauffret, S. Leroy /Tectonophysics 283 (1997) 61-104

TWTT deep, were observed in the Venezuela basin (Ladd and Watkins, 1980). These may correspond to a 10 km deep, 6.2 km/s-7.2 km/s interface identified on refraction line 4 from G. Sutton, reported in Officer et al. (1957). Two basement highs, which correlate with magnetic anomalies (Fig. 17) are probably volcanic; however, recent tectonic reactivation is probable for the eastern structure (shot point 4900, Fig. 17). 4.5. Dominican, Puerto Rico and Venezuela basins The western Venezuela basin was well described by Ladd and Watkins (1980). The Puerto Rico basin is characterized (Figs. 18 and 19) by several volcano-like structures either buffed or protruding from the sea floor (Matthews and Holcombe, 1976; Jany, 1989). An example with a large magnetic anomaly was surveyed but not drilled during DSDP Leg 15 (Raft, 1973). Similar structures are prominent on the Gloria mosaic performed in the EEZ (Exclusive Economic Zone) of Puerto Rico (EEZ-Scan Scientific Staff, 1987). Magnetic anomalies, that trend NWSE, of the Puerto Rico, Dominican and Venezuela basins are clearly related to these volcanoes (Leroy, 1995), and were not generated by sea-floor spreading (Ghosh et al., 1984). In the well-studied Venezuela basin (Biju Duval et al., 1978; Diebold et al., 1981), the boundary between the smooth B" reflector and rough basement (interpreted as oceanic crust) is commonly abrupt. This boundary was named the central Venezuelan basin fault (Biju Duval et al., 1978; Diebold et al., 1981). However, this fault is not evident towards the west, where the smooth B" reflector plunges regularly towards the Venezuela deformed belt (Talwani et al., 1977; Diebold et al., 1981; RC19-3 Line 8, Fig. 18). A N-S-trending scarp limits the rough basement towards the west (Ew 1319, Fig. 18; Vitali, 1985; Mascle and Letouzey, 1990). This scarp is crossed by the ESP 6 (Fig. 18; Vitali, 1985). We

77

completed the previous survey of this zone (Talwani et al., 1977; Diebold et al., 1981) with Institut Fran~ais du Petrole (IFP) multichannel lines. In the western part of the Venezuela basin the smooth B" reflector is affected by normal faults related to the Venezuela accretionary prism (Silver et al., 1975). Towards the east, the smooth B" southern boundary is also affected by this E-W-trending normal faulting that offsets Eocene horizon A" and younger strata. Between the typical smooth B" and deep rough oceanic crust we have defined a transitional zone where the basement is deep but does not yet present the characteristic roughness of oceanic crust. This transitional zone corresponds to the rise of the Moho from 17 km beneath the smooth B" basement to 13 km beneath the oceanic crust (Fig. 18, Line IFP 120 and R.C. 21-03, Line 108; Diebold et al., 1981). Deep oceanic crust is covered by a 0.45-kin-thick sedimentary layer that is restricted to the deep part of the Venezuela basin. A 0.012 km/myr (0.012 mm/year) sediment accumulation rate is observed for the 0.45-km-thick A"-B" interval that overlies smooth basement (Fig. 18), and if a similar rate applies in the southern area, we calculate that ages of transitional zone basement and oceanic crust could be 114 Ma and 152 Ma, respectively (Fig. 18). However, our estimation is rough because we did not take into account the compaction of these deeply buried sedimentary layers, the accumulation rate maybe higher in the deep part of the layer that shows a turbiditic seismic facies (Driscoll and Diebold, 1997) but we cannot exclude a basal thin pelagic layer resting upon the oceanic crust. In the Western Cordillera of Colombia the magmatic complex is interlayered and overlain by Albian-Turonian cherts that are covered by Coniacian-Santonian 12-km-thick flysch and turbiditic formation (Bourgois et al., 1985). A Late Jurassic-Early Cretaceous (dated by radiolarian micropalaeontogical identification) radiolarite sequence was described in the Nicoya Peninsula where a fragment of the Caribbean

Fig. 11. Seismic profile crossing the Pecos fault (Mautfret et al., 1994; Leroy and Mauffret, 1996). Note the recent deformation. The Caribbean sequences (A", B") are identical on both sides of the Pecos fault but the acoustic basement (B") is much deeper in the Colombia basin than on the other side of the fault zone. Note also that the A " - B " interval is not thick and is localized at the base of the sedimentary cover. The sub-B r' reflector (Stoffa et al., 1981) is identified on the right side of the figure but is less clear on the left side. See Fig. 1 and inset for location.

A. Mauffret, S. Leroy /Tectonophysics 283 (1997) 61-104

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large igneous province has been accreted (Mauffret et al., 1997). A prominent reflector, identical to the sub-B" reflector, can be followed beneath smooth B" and it merges with rough basement in the southern part of the Venezuela basin (Diebold et al., 1981). Correlation between the seismic reflection profiles and refraction studies (expanded spread profile, ESP 3) indicates that the sub-B" reflector does not always correspond to a velocity contrast. Reflection coeffi-

cient depends on acoustic impedance, which is the product of velocity and density. The sub-B" reflector may correspond to a contrast of density between what was originally a hot Cretaceous volcanic layer and cold dense Jurassic oceanic crust. We therefore propose to place the sub-B" reflector at the top of original oceanic crust that is buried beneath Cretaceous volcanic layers. The velocity break, seen at l0 km in ESP 3, is deeper and may correspond to layer 2V (V = volcanic) - layer 3V interface (Fig. 18). A

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A. Mauffret, S. Leroy l Tectonophysics 283 (1997) 61-104

recent publication (Diebold et al., 1997) presents a detailed study of this area. 5. Thickness and distribution of the seismic units

The upper seismic interval has a variable thickness, relatively thin in the Dominican basin and on the flank of the Beata ridge, and increasing to 2 km thick in the southern part of the Venezuela and Colombia basins where it fills a trench related to the south Caribbean deformed belt (Talwani et al., 1977; Biju-Duval et al., 1982). In the distal northern part of the Venezuela trench, the upper seismic interval onlaps older sequences that dip towards the south. In the Colombia basin, the thickness of the Early Miocene to Recent age layer is related to the formation of the Magdalena deep-sea fan (Kolla et al., 1984). The DSDP sites and the seismic profiles indicate that this interval consists of turbiditic sediments. The second seismic interval corresponds to a pelagic unit. The upper part, and in some places the entire interval, is chaotic. This layer is current-

85

controlled, as confirmed by the presence of several hiatuses in the DSDP holes that have been related to strong Early Miocene currents which were active as the Caribbean Sea opened towards the Pacific Ocean (Edgar et al., 1973b; Holcombe and Moore, 1977). A layer in the western part of the Colombia basin (ODP Site 999, Fig. 2; Chibehas rise) has the same chaotic facies. However, Bowland (1993) correlated the top of this layer with a Late Miocene undercompacted and overpressured lithic unit identified at DSDP Site 502 (Prell and Gardner, 1982). This interpretation is confirmed by the ODP Site 999 (Scientific Party, 1996), but a second deeper chaotic interval corresponds to the Early Miocene (Fig. 2). The corrugated aspect of the Early Miocene reflector is widespread and the top of this unit has been described as a prominent horizon in the Colombia basin (PR reflector; Houtz and Ludwig, 1977). The Early Miocene to Middle Eocene seismic interval dips toward the south Caribbean di~formed belt and this disposition indicates that c o ~ s i o n a l deformation along South America oectmred since the Early Miocene (Biju-Duval et al., I982). At this time

Fig. 18. Depth sections illustrating the main features of the Venezuela basin and depth to basement in the Venezuela basin. (A) Ew-1319 shows the east-west normal faults that are related to the subduction of the Venezuela basin (Silver et al., 1975) beneath the south Caribbean deformed belt (Curaqao prism). The promingmt fault beneath the prism has a north-south trend (Vitali, 1985). (B) R.C. 19-L 8 (Talwani et al., 1977; Diebold et al., 1981) showing A" and B" reflectors dipping regularly southwards. No fault can be evidenced and the central Venezuelan basin fault (Biju Duval et al., 1978; Diebold et ai., 1981) is not a continuous feature but is formed by several east-west faults en echelon (E; Diebold et al., 1997). (C) Depth section from IFP 120 and R.C. 21-03-L 108 (Diebold et al., 1981). A" and B" reflectors are offset by the central Venezuelan basin fault. This fault zone is located at the southern boundary of the smooth B" characteristic of volcanic flows. This boundary has been reactivated during the subduction beneath the Curaqao prism because the Eocene A" horizon is offset on this depth section, but this reactivation is not evident northeast of this profile (Diebold et al., 1997). After a transitional zone where the acoustic basement is faulted and some reflections resemble the B" smooth reflector (Diebold et al., 1997), the acoustic basement is rough and 8 km deep. The roughness of basement is characteristic of oceanic crust (Diebold et al., 1981). Beneath the smooth B" reflector a prominent reflection (sub-B") has the same depth as the rough basement and seems to merge with it (Diebold et al., 1981). This horizon is supposed to be the oceanic crust buffed beneath 2.5 km basalts flows. This layer thins to 1.5 km to the northwest. The sub-B" reflector does not correspond to any break in the ESP 3 curve (Diebold et al., 1981). Therefore this reflector may result from a contrast of density between the upper volcanic flows and the original oceanic crust, and not to a contrast of velocity. A lower break in the curve may be located between layer 2V (upper volcanic flows + original layer 2 of the original oceanic crust and layer 3V (original layer 3 + underplated material). Note the difference between the ESP-derived velocity-depth functions on the volcanic plateau (ESP 3) and those on the oceanic crust (ESP 2). The Moho is 17 km deep beneath the volcanic plateau and 13 km deep below the oceanic crust. The Moho rises in the transitional zone where the volcanic plateau thickens progressively. Location of the depth section indicated on the depth to basement map, E). (D) The sedimentation rate (0.012 km/myr = 0.012 mm/yr) between Eocene (50 Ma) horizon A" and the Coniacian (88.5 Ma) volcanic B" reflector is extrapolated to the transitional zone and the deep basin. We found ages of 114 Ma and 152 Ma, respectively. (E) Depth to basement. Contour interval 0.2 km. Note the difference (about 1 km) in the deep basins between the depth obtained with a velocity law attributing a velocity of 1.8 km/s for the supra A" and 3.1 krn/s for the A"-B" interval and the greater depth derived from ESP. Volcanoes of the Puerto Rico sub-basin are indicated. Observe the east-west normal faults near the front of the south Caribbean deformed belt. The central Ver~zuelan basin fault is formed by several en-echelon faults that trend east-west. We identify a transitional zone between the thick volcanic plateau and the thin oceanic crust. Position of depth sections and ESP are indicated.

A. Mauffret, S. Leroy /Tectonophysics 283 (1997) 61-104

86

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A. Mauffret, S. Leroy /Tectonophysics 283 (1997) 61-104

the floor of the Caribbean Sea underwent large wavelength deformation, with a general tilting towards the south related to the formation of trenches and uplift along the flanks of the Beam ridge (Mauffret et al., 1994; Leroy and Mauffret, 1996; Mauffret and Leroy, 1997). The A"-B" interval is evident in the Venezuela and Dominican basins (Ladd and Watldns, 1980) and in the western part of the Colombian basin (Bowland, 1993). However, horizon A", which correlates with Middle Eocene chert, is not clearly identified in the eastern Colombia basin and on the Beata ridge. The A"-B" interval is very thick in the Venezuela basin where the 1.2-km-thick layer overlies the oceanic crust. The 0.45-kin lower part (Fig. 18) of this deep layer is not observed in other parts of the Caribbean basin. The A"-B" interval is thick in the western part of the Colombia basin (Fig. 2), but thinner in the Aruba gap area where it was penetrated during drilling at DSDP Site 153 (Fig. 2). In a basin located to the north of DSDP Site 153, the A"-B" interval correlates in thickness and seismic facies with units encountered at the DSDP Site 153 (Fig. 2). However, acoustic basement (V, Figs. 2, 12 and 14) is very different in this area from the typical smooth B" reflector and is overlain by a sedimentary layer with a compressional velocity of 3.94 km/s (Figs. 2 and 12).

6. Depth to basement map The depth-to-basement map, derived from all the existing seismic lines (Fig. 19) shows the different morpho-tectonlc provinces that make up the Caribbean Sea. The steep Pedro escarpment separates the upper Nicaragua rise from the lower Nicaragua rise (Mascle et al., 1985; Case et al., 1990; Mascle and Letouzey, 1990). The Nicaragua rise is bounded to the south by the Hess escarpment. The former Colombia basin can be divided into four zones. A southwestern region (the western Colombia basin) is characterized by moderate (45 km) basement depths, except near the Panamanian accretionary prism where they exceed 6 krn (Vitali, 1985; Bowland and Rosencrantz, 1988; Bowland, 1993). The basement is more than 8 km deep in the eastern Colombia basin facing the Colombia accretionary prism. A shallow basement rise, the 5 km

87

deep Warao rise (Fig. 19), extends from the Hess escarpment to the Beam ridge. The basement is also deep (more than 5 km) in the Haiti sub-basin north of the Warao ridge. This basement is separated from the northern Beam ridge by a steep NE-trending escarpment. A N-S-trending ridge, on which DSDP Site 151 lies (Edgar et al., 1973a), forms the southern central part of the Beam ridge. Several plateaus and seamounts are located between the DSDP Site 151 ridge and the Venezuela basin. The Aruba gap connects the Colombia and the Venezuela basins but the Pecos fault zone crosses this zone obliquely. The former Venezuela basin can be divided into three zones: the Dominican sub-basin in the centre, where the basement is relatively shallow (5 km, Fig. 19); the Puerto Rico sub-basin in the north which is characterized by several seamounts protruding from the sea floor or buried (Jany, 1989; Holcombe et al., 1990); and the Venezuela basin sensu stricto in the southeast, adjacent to the Venezuela accretionary prism where the basement is deep (7 to 9 km; Talwani et al., 1977; Biju Duval et al., 1978; Diebold et al., 1981). Towards the east, the Venezuela basin is flanked by the Aves ridge, an extinct volcanic arc (Holcombe et al., 1990). Three areas are characterized by a deep basement: the eastern Colombia basin, the Haiti sub-basin and the Venezuela basin. However, a rough crust, characteristic of oceanic crust, is defined only in the Venezuela basin. In the two others basins a rough crust can be identified in few places (Bowland and Rosencrantz, 1988) but smooth basement, related to the B" volcanic flows, underlies the sedimentary cover in several places. 7. Crustal structure of the Caribbe~ large igneous province Maps of depth-to-Moho and crustal thickness were presented by Case et al. (1990). However, the crustal thickness map included the sedimentary cover, which is quite variable, and can be very thick, particularly in the southern part of the Caribbean Sea. We recomputed the thickness of the crust without its sedimentary cover (Fig. 20), from the refraction results (Officer et al., 1957, 1959; Ewing et al., 1960; Ludwig et al., 1975; Houtz and Ludwig, 1977; Diebold et al., 1981), augmented by results

A. Mauffret, S. Leroy /Tectonophysics 283 (1997) 61-104

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A. Mauffret, S. Leroy/Tectonophysics 283 (1997) 61-104

from sonobuoy experiments performed by the Lamont Earth Science Observatory before the Ewing 9501 cruise (Diebold et al., 1997). The deep crust is resolved by refraction data, whereas information on the upper crust comes mainly from sonobuoys, although the Moho is sometimes reached with this method (Houtz and Ludwig, 1977; Talwani et al., 1977). Upper crust is not well resolved by long refraction lines, and we adjusted the depth to basement with the help of multichannel profiles whenever possible (for example refraction lines north of Colombia performed by Ewing et al., 1960). The structural provinces defined by the depth-tobasement map (Fig. 19) and seismic character are evident on the crustal thickness map (Fig. 20). In deep parts of the Caribbean Sea, Haiti and Puerto-Rico sub-basins, and the eastern Colombia and Venezuela basins, the crust is less than 10 km thick. In contrast, shallow parts such as the Nicaragua rise, Beata ridge and the western Colombia basin have thick crusts. The same may be true for the Dominican basin, but refraction data are scarce in this region. The crust of the Beata ridge is particularly thick (20 kin). The Puerto Rico basin has a thin crust although the acoustic basement is smooth. Rough oceanic basement, identified by Diebold et al. (1981), corresponds to a crust that is less than 5 km thick. The 5-km thickness in the Venezuela basin is confirmed by refraction work (Officer et al., 1957, 1959), ESP data (Diebold et al., 1981) and seismic profiles that show a reflection corresponding to the Moho (Talwani et al., 1977; Diebold et al., 1981; Diebold et al., 1997). A 5-km thickness is not normal for Jurassic-Cretaceous oceanic crust in the Atlantic region. Normal oceanic crust is about 7 km thick (White et al., 1992), and we observe such a thickness throughout the Atlantic Ocean (Fig. 20). The Jurassic crust of the Pigafetta basin (Pacific Ocean) is also 7 km thick (Abrams et al., 1993) and a comparison between the velocity curves of the Pacific crust and those of the Venezuela basin (ESP 2, Fig. 18) indicates that the upper part of the curves (from 5 km/s to 7 km/s) are similar, but the lower part is thicker in the first case than in the second. Original crust may have been thinned by extension related to the formation of the volcanic plateau or the rough crust was formed as a thin crust (Diebold et al., 1997).

89

In the Haiti sub-basin, however, the 5-km contour is based on only one sonobuoy result which gives an apparent value of 9.2 km/s for the Moho refracted arrival, A dipping Moho may explain this high value. Nevertheless, two-ship refraction data indicate that the crust is less than 10 km thick in this basin. Moreover, the seismic reflection profiles shot during the Ewing 9501 cruise show a 5-km thick crust in this basin (J. Diebold, pers. commun.). In the southeastern Colombia basin, the acoustic basement in some places has the rough character of oceanic crust (Bowland and Rosencrantz, 1988; Fig. 8). However, a smooth B" reflector can be also identified in the deepest Colombia basin (Fig. 9). We discuss this problem in the next section. The crust can be divided into two parts: an upper part with compressional velocities between 4.5 and 6 km/s, and a lower part with velocities from 6 to 8 km/s. Refraction data are old and the results are derived from slope-intercept solution. Modern techniques yielding high-quality data (Ocean Bottom Seismometer, OBS and Expanding Spread Profile, ESP), and utilizing methods including geometry raytracing and synthetic modelling, give velocities of 2.5-6.6 km/s and 6.6-7.6 km/s for layers 2 and 3, respectively, and mean thicknesses of 2.1 km for layer 2 and 4.97 km for layer 3 (White et al., 1992). However, the high velocity (6.6 km 3) of layer 2 in fact represents a velocity gradient, whereas the mean velocities obtained during the older refraction studies are lower. This suggests that if the mean velocity is 6 km/s, the layer must have an interval of high velocity at its base. Moreover, ESP work in the Venezuela basin (Diebold et al., 1981) confirms that the crust can be divided into two parts: layer 2V (V for volcanic plateau) with velocities ranging from 5 to 6 krn/s, and layer 3V with velocities higher than 6 km/s (Fig. 18). A third break, at a depth of 15.5 km, is interpreted as the Moho interface. Although layer 2V has a mean thickness of 2.2 km (Fig. 21), the total thickness varies considerably, from low values in the basins to 4 km on the Beata ridge and Nicaragua rise. Layer 3 has a variable thickness: thin beneath the deep Haiti sub-basin, Colombia and Venezuela basins, and thick beneath the lower Nicaragua rise and Beata ridge.

90

A. Mauffret, S. Leroy /Tectonophysics 283 (1997) 61-104

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1977; Diebold et al., 1981) of total crustal thickness and layer 2V (V for volcanic) thickness. Most of the sonobuoy results do not include the Moho refraction and only the upper crustal structure is documented (black bar). Layer 2V is thin and the thickening of the crust beneath the Nicaragua and Beata thick plateaus is observed in layer 3V. The 2.2-km mean thickness includes the very thin layer 2 of the basins, whereas layer 2V is about 4 km thick beneath the volcanic plateaus. (B) In our model, layer 2 of the original oceanic crust is sandwiched between an upper basaltic layer and layer 3V (layer 3 of the original oceanic crust underplated). B" smooth, sub-B" and R represent the top of the basaltic layer, the top of the original oceanic crust and the top of a high velocity layer, respectively. 8. D i s c u s s i o n

8.1. Nature of reflector V and its age In the Nauru basin, characteristic sills were observed in a sedimentary layer between the smooth top of the Cretaceous volcanic unit and oceanic basement (Shipley et al., 1993). Similar bright spots and abrupt shifts in stratigraphic levels are observed in the Aruba gap (Figs. 2, 12 and 14) for reflector V, which is also interpreted as sills in volcanic basement or the overlying sedimentary layer. Although the sills may have the same age as the Cretaceous B" reflector, the acoustic basement and the sedimentary layer that rest upon this basement are definitely older. In this region, the B - V interval is 0.5 k m thick. A 0.02-km/myr sediment accumulation

rate was observed for the S a n t o n i a n - e a r l y Coniacian black shales drilled at DSDP Site 153 (Saunders et al., t973), and if a similar rate applies here, horizon V may be early Albian (113 Ma) in age. The transparent facies o f the B " - V interval is consistent with Albian black shales.

8.2. Cretaceous black shales restricted circulation During Campanian and Maastrichtian time, deep waters circulated among the Pacific Ocean, Caribbean Sea and Atlantic Ocean. In contrast, the ethology of late T u r o n i a n - C o n i a c i a n Santonian units reflects periods o f current activity interrupted by periods of stagnation (Saunders et al., 1973). There is conflicting evidence on the Cretaceous setting of the Caribbean Sea; an Atlantic affinity is indi-

A. Mauffret, S. Leroy /Tectonophysics 283 (1997) 61-104

cated by euxinic conditions, whereas radiolarians favours Pacific affinity. Sedimentation in the Pacific during the Cretaceous was dominated by abundant biogenic silica; these sediments were deposited under oxygenated bottom water conditions, while anoxic conditions were restricted to structural highs (Schlanger and Moberly, 1985; Ogg et al., 1992). Pacific surface waters penetrated the Caribbean region, as shown by radiolarian-rich sediments and planktonic foraminiferal assemblages (Premoli Silva and Bolli, 1973), but a barrier nonetheless existed between well-oxygenated deep waters of the Pacific Ocean and stagnant waters of the Caribbean Sea (Saunders et al., 1973). The formation of Cretaceous oceanic plateaus in the Pacific may have coincided with an increase in global temperatures and CO2 production and a consequent rise in sea level. Cretaceous anoxic events (from Aptian to Santonian) may be related to these global events (Larson, 1991a). In the Atlantic Ocean about 120-125 myr ago, volcanic units of the J anomaly ridge (Tucholke and Ludwig, 1982) and the Walvis ridge (White and McKenzie, 1989) may have prevented free interchange of bottom water with the Southern Atlantic Ocean, and this is believed to have led to deposition of Aptian-Albian organic-rich sediments. In the same way, the Caribbean igneous province formed a barrier between Pacific and Atlantic waters, and the Proto-Caribbean Sea was also affected by restricted circulation, as evidenced in Venezuela by the oil-rich La Luna Formation (A1bian black shales; Tribovillard et al., 1991).

8.3. Dipping reflectors and volcanic sources We observed several dipping reflector packages that apparently dip southward towards the deep part of the Colombia basin (Figs. 7 and 8). These are similar in their form to the seaward-dipping reflectors described along volcanic passive margins (Mutter et al., 1982). However, in the deepest part of the Colombia basin, the reflectors dip to the north (Figs. 9 and 10). Several areas with dipping reflectors have also been observed in the Venezuela basin (Ladd and Watkins, 1980). Most of these reflectors are dipping to the south but some are dipping to the north. On the Kerguelen plateau, dipping reflectors are

91

associated with subaerial volcanic flows whose apparent source is found updip in the form of extinct and eroded volcanoes (Rotstein et al., 1990; Schaming and Rotstein, 1990; Schlich et al., 1993). In the Caribbean area, although there is some magnetic evidence of volcanic features (Figs. 4, 14 and 17), there is no clear relationship between volcanoes and reflectors. Some reflectors dip towards and beneath a presumed volcano (Fig. 14); others with opposing dips cross one another. It is probable that several sources (volcanoes and fissures) fed volcanic flows at different levels and at different times, with the consequence that many volcanic sources may be buried beneath a thick volcanic pile. Subaerial volcanic flows cannot be excluded for the Aruba gap area (Fig. 14). Nevertheless, DSDP Site 153, located nearby, does not indicate a shallow depth of deposition for the late Turonian--early Santonian sediments deposited on basaltic flow (Saunders et al., 1973). On the other hand, a shallow environment is proposed for the basalts sampled on the Hess escarpment (ODP Site 1001, Leg 165, Scientific Party, 1996). The crust is only 11.9 km thick beneath the 2.5-km-thick dipping reflectors illustrated in Fig. 6 and a subaerial volcanic formation can be excluded above a so thin crust. We concluded that a subaerial volcanic construction is possible on the high plateaus (lower Nicaragua rise and Beata ridge) but unlikely in the deep part of the Caribbean basin. In the Puerto Rico sub-basin several volcanoes are observed (Figs. 18 and 19). Although smooth B" is widespread, crust is thin in this area (Fig. 20). We propose that volcanic flows were too thin in this basin to have buried volcanic sources. As in the Nauru basin (Shipley et al., 1993), there is little evidence of fissures or long-lived erupting sources that might explain the wide area extent of basin-wide volcanic rocks.

8.4. Crustal structure of the Caribbean large igneous province In the Caribbean, the thickness of layer 3V controis the structure of oceanic crust. Layer 2V, with velocities between 4.5 krn/s and 6 kin/s, is normally thin (around 2.2 km) and only in places reaches 4 km. Layer 3V, with velocities ranging from 6 km/s to 8 km/s, is more variable: thin beneath the basins

92

A. Mat~[Jl"et, S. Leroy / Tectonophysi( s 283 (1997) 61-104

(abnormally thin beneath the Venezuela basin, 2 kin) but very thick below the Beata ridge (13 kin) and Nicaragua rise (14 km; Fig. 21A). In the world's oceans, variation in thickness of layer 3 dominates the oceanic crustal structure (Mutter and Mutter, 1993) This velocity structure can be compared with that of the Kerguelen plateau. Throughout the entire region, the upper crust is a relatively uniform 5-kin thick basaltic layer with velocities ranging from 4.8 km/s to 6.3 km/s. Beneath the northern Kerguelen plateau, a Cretaceous oceanic plateau like the Caribbean LIE a similar crustal structure has been observed. Layer 2 (4.5-5.5 km/s) is 3.8 km thick, and layer 3 (6.6 km/s at the top to 7.3 km/s at the base) is 17 km thick (Charvis et al., 1993). In contrast, in the southern Kerguelen plateau, the structure of layer 3 is different. In the upper part of the lower crust, velocities increase from 6.3 km/s to 6.6 km/s; in the lower part the gradient is less, reaching only 6.8 km/s at the base. A 6-kin-thick reflective low-velocity zone (6.7 kin/s) overlies the Moho discontinuity (Operto and Charvis, 1995). Operto and Charvis (1995) concluded that the southern Kerguelen plateau is a fragment of a volcanic passive margin, whereas the northern Kerguelen plateau is a purely oceanic feature. In an igneous province, intrusion and extrusion of volcanic material might be expected to disturb the oceanic crustal structure to the extent that the original structure could no longer be identified. This is probably true for layer 3V, which appears to have been thickened by deep magmatic processes and has a velocity structure very different from normal oceanic crust. However, layer 2V of the igneous provinces is not very different from layer 2 of normal oceanic crust. Magnetic anomalies of original oceanic crust in the Nauru basin (Shipley et al.. 1993) are consistent with formation of a thin basaltic sheet from multiple volcanic sources. We propose

that layer 2V in the Caribbean igneous province is composed by a thin volcanic layer (1 to 2.5 km and maximum 4 km) overlying the preserved layer 2 of the original oceanic crust. Several rock samples collected by dredging along the eastern scarp of Beata ridge (Fox et al., 1970; Fox and Heezen, 1975) were re-studied recently by T. Donnelly (written commun., 1995). He identified fine-grained and cumulate gabbros, rock types characteristic of layer 3. This observation is confirmed by a recent sampling survey by submersible (Mauffret et al., in preparation). Gabbros on the Beata ridge constitute good evidence of the thinness of layer 2V, even on the thick Beata oceanic plateau. The absence of layer 2V on the Beata ridge can be explained by strong tectonics and subsequent erosion (Mauffret et al., in preparation).

8.5. Nature of the reflector sub-B" and horizon R We propose to place the sub-B" reflector at the top of the original oceanic crust. This interpretation is based on the correlation between a prominent intravolcanic basement reflector and the rough crust and the interpretation of the ESP 3 in the Venezuela basin (Fig. 18; Diebold et al., 1981). Nevertheless we have not enough seismic data to correlate the sub-B" reflector of the Venezuela basin with those of the Aruba gap. The original oceanic crust could be sandwiched between an upper volcanic layer and lower underplated volcanic material (Fig. 21B). The prominent horizon R is highly reflective and is located in the deep crust (Figs. 16 and 17). A similar reflector, in the Venezuela basin, has been placed at the top of the 7.2-km/s layer (Ladd et al., 1984). The thickness of interval B"-R is about 6 kin, and the interval R-Moho may be as thick as 7 km underneath the eastern flank of the Beata ridge but neither the reflector R nor Moho reflection can be seen on the seismic profiles crossing the 20-kin-thick (Fig. 20) central part of the Beata ridge because the sea-floor multi-

Fig. 22. (A) Columns of the refraction results from right to left in the Pigaletta basin (Abrams et al.. 1993) in the Venezuela basin (ESP from Diebold et al., 1981; SB 32 R.C. 19 from Talwani et al., 1977) and in the Colombia basin. Positions of these refraction experiments are indicated in Fig. 1 (B-E): age-depth relationship on the curve of Parsons and Sclater (1977); (B) ESP 2 and 4 are off the curve because the crust is thin and depressed by the Curaqao prism (see Fig. 18). SB 32 R.C. 19 shows a 160-Ma-old crust close to the ODP 801 results. ESP 3, SB 24 and SB 32 R.C. 20 results show a shallower basement. (C) Upper volcanic layer and underplated material have been backstripped. The final result is close to the SB 32 R.C. 19 result. (D) The underplated material beneath SB 24 R.C. 20 has been backstripped. (E) The underplated material beneath SB 32 R.C. 20 has been backstripped.

A. Mauffret, S. Leroy/Tectonophysics 283 (1997) 61-104

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pie prevents any deep penetration. The high-velocity layer (7.3-7.45 km/s; sonobuoys 24 R.C. 20 and 32 R.C. 20; Fig. 22) is about 5 km thick in the deep part of the Colombia and Haiti sub-basins. We propose that this layer results from the emplacement of unusually magnesian-rich high-temperature magmas and ultramafic cumulates (Fig. 21B). A thermal anomaly of 150°C in the mantle is estimated to generate a 20-km-thick magmatic crust (White and McKenzie, 1989, 1995). Melt generated from abnormally hot mantle is more magnesianrich than that from mantle of normal temperature. Such magma is also more dense than normal basalt, and is prone to be trapped at the base of oceanic crust where it would evolve through fractional crystallization to produce a thick layer of olivine and pyroxene cumulates. A layer of picritic sills and/or ultramafic cumulates should have a seismic velocity of >7.2 km/s (White and McKenzie, 1989), significantly higher than the 6.8 km/s of normal oceanic crust (Farnetani et al., 1996). Widespread eruption of picrites is evident throughout the Caribbean. The 5-km-thick volcanic section of Curaqao Island (Fig. 23) includes picrites at its base (Klaver, 1987, Kerr et al., 1995). Komatiites of Gorgona Island, off the Pacific coast of Colombia, are associated with tholeiites erupted at 88 Ma, and are also believed to be part of the Caribbean plateau (Kerr et al., 1995, 1995). Other oceanic plateaus have high-velocity layers at the base of the lower crust (Coffin and Eldholm, 1994, Farnetani et al., 1996). A 15-km-thick high-velocity (7.3 km/s) layer has been detected beneath Hatton Bank, a volcanic passive margin (White et al., 1987). Reflective lower crust was also identified in the southern Kerguelen plateau but its velocity is low (Operto and Charvis, 1995). 8.6. Age of the oceanic crust

N. Driscoll and J. Diebold (written commun.) estimated the age of Venezuela oceanic crust by ther-

95

mal subsidence of unloaded basement (Parsons and Sclater, 1977), and they compared their results with the 167 4- 4.5 Ma unloaded basement of the Pigafetta basin drilled at ODP Site 801 (Abrams et al., 1992). They found that corrected ESP 2 and 4 are off the age scale (Fig. 22), but this can be attributed to loading of the Venezuela deformed belt whose thickness may reach 11 km (Edgar et al., 1971). This loading and the formation of the Venezuela trench (Talwani et al., 1977) depress acoustic basement in the area where ESP 2 and 4 are located (Fig. 18). In addition, the anomalous thinness of oceanic crust may explain excess subsidence observed for ESP 2 and 4. Sonobuoy 32 RC 19 (Talwani et al., 1977) lies updip in rough basement of the Venezuela basin (Figs. 18 and 22). A Jurassic age (160 Ma) is determined for the 6.1-km deep unloaded acoustic basement; nevertheless a regional flexure cannot be excluded and oceanic basement can be younger than 160 Ma. Much younger ages were obtained (Fig. 21) for ESP 3 (Venezuela basin): SB 32 RC20 (Colombia basin) and SB 24 RC 20 (Haiti sub-basin). These two sonobuoy results were determined by Houtz and Ludwig (1977) and we used refraction profile 84 (Edgar et al., 1971) to constrain Moho depth in the Haiti sub-basin. However, if we accept that layer 2 is formed by a basaltic body overlying the original crust, and layer 3 has been thickened by underplating of high-velocity and high-density material the basaltic layer as well as the underplated material must be also backstripped. A density of 2.7 was determined for samples of Cretaceous basaltic rocks obtained during the DSDP Leg 15 (Edgar et al., 1973b), and we choose reflector sub-B" as the top of the oceanic crust for ESP 3 (Fig. 22B). For the Haiti sub-basin and Colombia basin (Fig. 22C and D) we assumed that rough oceanic crust has not been covered by the upper volcanic layer but was substantially underplated (SB 24 R.C. 20 and SB 32 R.C. 20, Fig. 22). Corrected depth varies from 6.18 km to 6.37 km. These values are relatively close to the corrected depth (6.1 km) obtained for oceanic crust (SB32 RC19) and

Fig. 23. Plate tectonic framework of the Caribbean volcanic province. The Caribbean is fixed relative to the other plates. The Caribbean volcanic province is formed by thick Cretaceous volcanic plateaus separated by deep basins. The Venezuela basin is underlain by oceanic crust probably Jurassic in age. Several fragments of Caribbean plateaus have been accreted to the continent (western Cordillera of Colombia, Central America, Hispaniola, Curaqao and Aruba, Venezuela). The position of the crustal section shown Fig. 24 is indicated.

96

A. Mauff)'et, S. l.eroy/Tectonophysics 283 (1997) 6 I - 104

we concluded that oceanic crust could be as old as 160 Ma in the three basins. However, the Colombia basin and Haiti sub-basin were closer to the Pacific spreading axis during the Jurassic than the Venezuela basin, and consequently they should be younger, but our method is not enough sensitive to appreciate a small difference in age if the sea-floor spreading was fast in the Pacific during this period (5-7 cm/y; Engebretson et al., 1985; Pindell and Barrett, 1990). We also note that in deep basins, subsidence curves do not show a Cretaceous thermal rise that may be restricted to the Beata oceanic plateau. An estimation based on the sedimentation accumulation rate, suggests that the Venezuela oceanic crust is 152 Ma old (Fig. 18) and this age corresponds to the determination deduced from subsidence curves (160 Ma, Fig. 22). A comparison of the Venezuela basin deep-seated sediments and those of the Pacific fragments (Western Cordillera of Colombia, Bourgois et al., 1985; Nicoya Peninsula, Mauffret et al., 1997) of the Caribbean igneous province suggests that the oceanic crust was covered first by Late Jurassic-Early Cretaceous deep-sea sediments (radiolarites and cherts), and then this thin abyssal layer has been covered by a thick Cretaceous turbiditic formation when the oceanic plateau and the adjacent oceanic crust approached the South American continent (Driscoll and Diebold, 1997). We conclude that oceanic crust of the Caribbean province may be Jurassic in age. Lower Jurassic to Lower Cretaceous ophiolites have been dated using radiolarian flora in cherts of the Antilles Islands. Montgomery et al. (1992) obtained 160-148 myr for the Duarte complex in Hispaniola, and 145 myr for the La Desirade complex in French Antilles. These ophiolites may be fragments of original oceanic crust.

8. 7. Age of the oceanic plateau and related volcanics The Caribbean igneous province is made up of several distinct plateaus: the lower Nicaragua rise, the western Colombia basin, and the Beata ridge, separated by deep basins. In the basins the crust is oceanic (Venezuela basin), in some places (Haiti, Dominican and Puerto Rico sub-basins; Colombia basin) covered by a thin volcanic layer and underplated by volcanic material. Although the basins

were deep during the formation of the igneous province, there is some evidence that the tops of the plateaus were shallow. According to Bien-Aim6 Momplaisir (1986), certain tholeiitic basalts exposed in Haiti could have been deposited in a shallow environment. A pebbly mudstone containing shallow marine fauna is intercalated with basalts on Curaqao (Beets et al., 1984). The vesicular nature of the basalts sampled on the Hess escarpment may indicate a shallow environment (ODP Site 1001, Leg 165, Scientific Party, 1996). Cretaceous bathymetry of the Caribbean resembles that of the Pigafetta basin where Cretaceous shoals, guyots and islands are separated by deep basins (Abrams et al., 1993). The Caribbean igneous province is believed to have formed on the Farallon plate to the east of the Pacific subduction zone. Its arrival may have blocked the subduction zone, causing the arc system to flip its polarity, and allowing the insertion of the plateau between the American plates (Pindell and Barrett, 1990). Many pieces of the plateau were accreted to the continent (Fig. 23; Donnelly et al., 1990). From our data and interpretations, it appears that crust in the Venezuela and Puerto Rico basins was too thin to block the Cretaceous eastern Pacific subduction zone. Consequently we propose that the Aves and Lesser Antilles island arcs rest upon a deep thick portion of the Caribbean plateau. This eastern extent of the igneous province may explain the exposure in Trinidad of the 89-87 myr Sans Souci basalts (Wedge and Macdonald, 1985; Erlich and Barrett, 1990). The distance from the Lesser Antilles to the Nicoya complex in Costa Rica is 2500 km. It is 1300 km from the Nicoya complex to Cretaceous volcanic rocks in Ecuador. The Panama Arc and the North Andean Block, which separate Costa Rica from Ecuador, have been deformed since the Cretaceous (Kellogg and Vega, 1993), and the initial distance between the two areas was probably also 2500 km. The overall diameter of the Caribbean igneous province can therefore be assumed to be larger than the suggested 2000 km diameter of plumegenerated plateaus (White and McKenzie, 1989). In islands of the Netherlands Antilles, there is evidence that oceanic plateaus collided and accreted to the continent only shortly after their formation (Beets et al., 1984; Kerr et al., 1995), and that the plume

A. Mauffret, S. Leroy/Tectonophysics 283 (1997) 61-104

head of the ancestral Galapagos hot spot (Duncan and Hargraves, 1984) was probably about 1000 km from the eastern Pacific subduction zone. The main igneous activity throughout the Caribbean igneous plateau, from Colombia (Walker et al., 1991) to Haiti and Curacao occurred over a short period of time in the late Turonian-early Conacian (88 myr; Sinton and Duncan, 1992). Within the Caribbean basins (Edgar et al., 1973a), however, ages of volcanic rocks range from mid-Campanian (76 Ma, Sites 152 and 1001) to late TuronianConacian (88.5 Ma; Sites 146, 150 153) and we conclude that the smooth B" reflector cannot be an isochron. Moreover, volcanic activity persisted after the main igneous event: Santonian (84 Ma) basaltic ash was found 90 m above a late Turonian basaltic flow in DSDP 153 (Edgar et al., 1973b). The time of initiation of the Caribbean plateau is unknown, but sub-aerial volcanic formations exposed on land seem to be older than Turonian. A 1500 m thick sequence of tholeiitic basalts is exposed in Haiti (Maurasse et al., 1979; Sen et al., 1988). The oldest sedimentary layers overlying basalts and/or interlayered between volcanic sills are dated Albian by micropalaeontologic determinations (Bien-Aim6 Momplaisir, 1986). The youngest volcanic rocks in Haiti are Campanian in age. Basalts of MORB affinities of the same age (Albian to Coniacian) and up to 5000 m thick outcrop on Curaqao and Aruba (Netherlands Antilles; Beets et al., 1984). An Albian age was also determined by micropalaeontological studies for the base of the Diabase Group in the Western Cordillera of Colombia (Kolla et al., 1984; Bourgois et al., 1985). Moreover, the volcanic basement (V reflector) north of DSDP Site 153 is probably early Albian in age (113 Ma) if we extrapolate to the B"-V interval (Figs. 12 and 14) the sedimentation rate of the Santonian-early Coniacian sediments drilled at DSDP Site 153 (Saunders et al., 1973). Prolonged volcanic activity is evident in other oceanic plateaus. In the Nauru basin, the igneous complex formed and erupted rapidly onto Jurassic oceanic crust. The earliest volcanic flow drilled in this area is 130 Ma (late Valanginian), but at least two other episodes of mid-plate igneous activity occurred, one Aptian-Albian (100 Ma), the other Campanian (75 Ma) (Schlanger and Moberly, 1985). A 126 Ma dolerite sill has been found in the Pigafetta

97

basin but the main basaltic sequence that overlies Jurassic crust is Aptian (114 Ma) in age (Abrams et al., 1992). Radiometric ages for basalts dredged on seamounts of the Pigafetta basin indicate a long range of formation (from 120 Ma to 74 Ma; Abrams et al., 1992, 1993). Although the volcanic eruptions responsible for the construction of the Ontong Java plateau ended more or less simultaneously during the early Aptian (122 Ma; Tarduno et al., 1991; Mahoney et al., 1993), some volcanism also occurred in the Turonian (90 Ma; Mahoney et al., 1993; Tejeda et al., 1996). In conclusion, the main (88 Ma) volcanic episode in the Caribbean province probably was geologically brief, but several lines of evidence suggest that activity started earlier, in the early Albian, at about 113 Ma, and lasted until later, in the Campanian (76 Ma). Volcanic plateau formation before the Aptian (124 Ma) is proposed in Hispaniola (Lebron and Perfit, 1994). However, this older volcanic event is probably different from the one that thickened the Caribbean plate.

9. Snmmary and conclusions The main observations and interpretations are summarized in Fig. 24 (modified from Case et al., 1990). Continental crust of the upper Nicaragua rise is separated from the lower Nicaragua rise by the Pedro escarpment. The lower part of the rise is covered by typical Caribbean strata (A" and B"), and we infer that this area, as well as the western Colombian basin, is underlain by a thick oceanic plateau. In the Haiti sub-basin and eastern Colombia basin, thin crust of oceanic character is underplated by volcanic material. Beneath the Beata ridge, the crust is very thick and complex. A thin layer, 2V, is formed by original oceanic crust overlain by a 1- to 4-km-thick volcanic layer (B-sub-B"). This relationship may explain the outcrop of gabbros from layer 3 along the western escarpment of the Beata ridge. We found a window in the smooth Bt' reflector where a sedimentary layer, probably Albian in age, is intruded by sills (V reflector; Leroy et al., 1996). Volcanic basement (V reflector) contains dipping reflectors that may indicate discrete volcanic sources. However, typical volcanic cones are rare and most probably are buried by subsequent basaltic flows. Dipping reflectors of

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A. Mauffret, S. Leroy /Tectonophysics 283 (1997) 61-104

2.5-km thickness are identified on the northern edge of the Colombia basin. They resemble the dipping volcanic layers observed along the volcanic passive margins but the thinness of the crust excludes a shallow origin for the volcanic sources. In contrast with layer 2V, layer 3V has a variable but appreciable thickness, and where it is thick, the lowermost portion has high velocities (7.2-7.4 1500W

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Fig. 25. Conceptual model of formation of the Caribbean igneous province. The plate tectonic situation at 80 Ma (Campanian) is modified from Duncan and Hargraves (1984), Engebretson et al. (1985) and Pindell and Barrett (1990). The Galapagos plume is 2500 km wide but its head is narrow. The volcanic plateaus (Albian, Coniacian and Campanian) moved with the Caribbean plate and the external parts collided and were obducted onto the continent (Ecuador and western Cordillera of Colombia). The Venezuela oceanic crust was formed initially on the Jurassic Farallon plate south of the future Galapagos hot spot and transported toward the east by the Caribbean plate. The Campanian Galapagos hot spot may have been formed upon a l18-133-Ma-old oceanic crust if we assume a 5-7-cm/yr convergence plate (Engebretson et al., 1985; Pindell and Barrett, 1990) during and before this period. The hot spot may also have formed the lower Nicaragua plateau. The rough crust of the Venezuela basin is 770 km eastwards of this plateau and the corresponding age is estimated at 129-148 Ma.

100

A. Mauff~'et, S. Len~y/Tectonophysics 283 (1997) 61-104

underlain by a oceanic crust which is apparently normal, but thinner (less than 5 km) than standard oceanic crust. The Ayes ridge and Lesser Antilles island arcs rest upon and penetrate a probable oceanic plateau. We estimate a Jurassic age (150-160 Ma) for oceanic crust from depth-age relationships and sediment accumulation rate. The Caribbean igneous province began to form during the Albian, but the main, widespread short-lived volcanic event was in the Coniacian (88 Ma). We could not calculate the total volume of volcanic material because the Caribbean igneous province comprises several oceanic plateaus separated by basins in which the crust is thin and dominantly oceanic; moreover, large volumes of the thick Caribbean plateau have been obducted on the adjacent islands or continents and portions of thin crust have been subducted. We propose (Fig. 25) a three-stage model for construction of the Caribbean igneous province. This model is modified from Duncan and Hargraves (1984); Engebretson et al. (1985) and Pindell and Barrett (1990). In this reconstruction, the Galapagos hot spot is located 2650 km eastwards of an 80-Maold spreading centre. The Galapagos hot spot does not belong to a Pacific super-plume (Larson, 1991 b), but was constructed off-axis upon Jurassic crust during or shortly after the formation of several other plumes not only in the Pacific Ocean, but also in the Indian and Atlantic Oceans. With a 57-cm/y spreading rate (Engebretson et al., 1985; Pindell and Barrett, 1990), the Campanian Galapagos hot spot may have initially formed on l18-133-Ma-old oceanic crust. Volcanism due to this hot spot may have created the lower Nicaragua plateau. The rough crust of the Venezuela basin is 770 km eastwards of this plateau and the corresponding age is estimated to be 129-148 Ma. Palaeomagnetic data indicate that the Caribbean volcanic province originated to the south (between 0 ° and 5°N; Macdonald, 1990) which is compatible with formation of the volcanic province above the Galapagos hot spot. In our model, the Galapagos plume generated narrow heads forming oceanic plateaus and extensive volcanic flows. The Albian plateau (Fig. 25) collided rapidly with the eastern Pacific subduction zone and most of it was accreted to the Western Cordillera of Colombia and Ecuador, or obducted onto Curaqao and Haiti. The

largest plateau is the Beata ridge, mainly Coniacian (88 Ma) in age, but volcanic flows of this age may overlie the Albian plateau. Conversely, some AIbian flows may underlie the Coniacian plateau. The small Campanian plateau is represented by the lower Nicaragua rise. The Jurassic Venezuela oceanic crust was placed southeast of the Cretaceous Galapagos hot spot and was carried along the eastwards drift of the Caribbean plate. However, we do not know the southern extent of this oceanic crust and its relationship with the Western Cordillera Colombian plateau because larger pieces of these geologic features have been subducted beneath South America. If plateaus older than Aptian have been formed above the Galapagos hot spot these plateaus have been captured by the Caribbean plate by successive flips of the eastern Pacific subduction zone and the remains of these plateaus are dispersed in the Greater Antilles.

Acknowledgements Supported by grants INSU ATP 733 and 780. We thank the officers and crew of the R/V Nadir for assistance in this project. We are especially indebted to the technical crew of GENAVIR who greatly helped in the acquisition of multichannel data. These data were processed at Ecole et Observatoire de Physique du Globe de Strasbourg, and we thank R. Schlich and M. Schaming who have facilitated access to the processing centre and helped us to use the Geovecteur software. The authors would like to express their thanks to H.G. Ave Lallemant and M.E Coffin for constructive reviews, and J. Diebold, who greatly improved the manuscript.

References Abrams. L.J., Larson, R.L., Shipley, T.H., Lancelot, Y., 1992. The seismic stratigraphy and sedimentary history of the East Mariana and Pigafetta Basins of the western Pacific. Proc. ODP Sci. Rcsults 129, 551-570. Abrams, L.J.. Larson, R.L., Shipley, T.H., Lancelot, Y., 1993. Cretaceous volcanic sequences and Jurassic oceanic crust in the East Mariana and Pigafetta Basins of the Western Pacific. In: Late Jurassic-Early Cretaceous Oceanic Crust and Early Oceanic Volcanic Sequences of the Nauru Basin, Western Pacific. Am. Geophys. Union, Geophys. Monogr. 77. 77-101. Alavarez-Marron, J., McClay, K.R., Harambour, S,, Rojas, L.. Skarmeta. J., 1993. Geometry and evolution of the frontal part of the Magellanes foreland thrust and fold belt (Vicuna area),

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