Pelagic clays, marls, and oozes are present in lithologic Units I and II. At this water depth, the proportion of calcium carbonate varies cyclically, in relation to the variation in bottom-water composition, a well-known phenomenon that occurs in other North Atlantic basins (Weaver and Kuijpers, 1983; Raymo et al., 1987). Pelagic layers vary in thickness between 0.01 and 1.89 m, with the thickest pelagic units being observed at Site 900. The pelagic sediment contains an average of 45% CaC03 and 0.31% organic carbon (Milkert et al., this volume). Pelagic sediments show slight to moderate mottling caused by bioturbation, but no distinct ichnofauna is visible. Bioturbation continues downward from the pelagic sediments into the tops of the underlying turbidites.
Brownish red, pelagic clays are common during the Eocene and Paleocene in lithologic Subunit IIIA at Sites 897 and 899; during this time interval, the regional carbonate compensation depth (CCD) was relatively shallow (Emery and Uchupi, 1984). The clay shows vague color banding, and bioturbation is not apparent. The claystones at Site 899 contain several black, organic-rich layers (Sawyer, Whitmarsh, Klaus, et al., 1994), and brownish black concretions, which are barren.
Pleistocene and Pliocene sedimentation at Sites 897, 898, 899, and 900 is dominated by deposition of terrigenous turbidites, separated by thin pelagic layers (Fig. 6). Turbidite sequences are easily distinguished from pelagic sediments by their graded bases, distinctive darker colors, and usual lack of bioturbation (Sawyer, Whitmarsh, Klaus, et al., 1994; Milkert et al., this volume). The base of each turbidite is clearly defined as a sharp boundary over mottled, lighter colored pelagic sediment. The structureless-to-laminated basal sand and silt layers in the turbidites correspond to the Tc-d division of Bouma (1962); the Ta-b interval is absent. Clay-rich intervals are bioturbated to laminated, and contain significant amounts of reworked nannofossils, and correspond to the turbiditic mud divisions (Te1-e3) of Stow and Piper (1984). Visually distinguishing the Te facies from the pelagic interval is difficult in areas where fine-grained pelagic sediments are mixed downwards by bioturbation and drilling disturbance.
A huge number of turbidites was deposited in the last 2 million years. A maximum of 48 turbidites (4 to 60 cm thick) was recognized in a single 9.5-m core (Core 149-898A-14H). In general, the frequency of turbidites varies from 2 to 7 per meter throughout the different cores. The large number of turbidites (e.g., 600 in Hole 898A) makes it difficult to correlate individual turbidites between sites (Milkert et al., this volume).
The terrigenous turbidites appeared very similar throughout the sedimentary sequence during shipboard description. More detailed post-cruise examination and additional visual core descriptions allowed a separation of these turbidites into four groups on the basis of different color and lithologic composition (Milkert et al., this volume).
Four different types of turbidite can be determined on the basis of color, carbonate, and organic carbon content:
- Type 1. Turbidites with gray terrigenous/siliceous and mica-rich bases.
- Type 2. Gray, calcareous, terrigenous turbidites with sandy, silty bases.
- Type 3. Calcareous turbidites with foraminifer-rich bases.
- Type 4. Lower Pleistocene brownish red sandy turbidites, which are common at all sites.
Despite these visual differences, analysis of major and minor elements by inductively coupled plasma atomic emission spectroscopy (ICP-AES) shows a very similar composition for all turbidites (Milkert et al., this volume). The turbidites originate mainly from continental erosion. In general, the Iberia Abyssal Plain turbidites have low Fe, Mg, Ti, and Zr contents with high proportions of Al and alkali metals. This indicates a continental provenance for the aluminosilicate phases and confirms a continental source region, similar to the element distribution described for organic-rich turbidites on the Madeira Abyssal Plain by DeLange et al. (1987). It probably indicates a uniform composition throughout the continental source area, rather than a single source for the Iberia Abyssal Plain turbidites.
Contourite facies and mixed contourite/pelagic facies are common throughout the entire Miocene and Oligocene at Sites 897, 898, 899, and 900.
The distinction between turbidites and contourites is difficult to make, especially when bottom currents modify deposits containing turbidites (Faugères and Stow, 1993). Despite these problems, evidence of occasional bottom-current activity was found in several sandy layers that have low pelite content and contain high concentrations of coarse foraminiferal tests (Alonso et al., this volume). Comas and Maldonado (1988) described similar foraminiferal sands in the Iberia Abyssal Plain, and Faugères et al. (1984) obtained comparable pure foraminiferal contourite sands in several North Atlantic basins and hypothesize that these deposits seem to be winnowed concentrations formed by normal bottom currents.
In addition to sandy sediments in the Iberia Abyssal Plain, silty sediments of mixed biogenic/terrigenous or only biogenic composition were recognized throughout lithologic Subunit IIB at all Leg 149 sites (Alonso et al., this volume). Silty sediments of this type could be recognized (Fig. 7; Sawyer, Whitmarsh, Klaus, et al., 1994). These deposits could be attributed to the effect of bottom-current processes. Silty deposits are differentiated from the Td division turbidites on the basis of their bimodal or polymodal grain size distribution in the total cumulative curve, the lack of vertical grading, and the poor preservation of nannoplankton.
Therefore, Alonso et al., (this volume) and the Shipboard Scientific Party (Sawyer, Whitmarsh, Klaus, et al., 1994) distinguished two types of contourite facies:
Slumps and slides are present within several single cores in the Pleistocene and Pliocene section. A 1.3-m-thick debris flow occurred in the early Pliocene (Section 149-897C-22R-03, Zone NN18 after Liu et al., this volume; Zone N20/N19 after Gervais, this volume). It reveals a highly folded and scrambled mixture of pelagic oozes and silty sands with a small amount of turbidite clay. An inverted, 83-cm-thick sequence, containing three single turbidites, was obtained from Section 149-898A-12H-01. It is underlain by an unusual, greater than 1-m-thick, homogeneous turbidite mud. This flow is of late Pleistocene age (Zone NN19 after Liu et al., this volume or Zone N23/N22 after Gervais, this volume). This section is interpreted as a single block, which was overturned by a gravity flow. The above-described structures cannot be correlated between the drill sites, suggesting that they originated from different events that did not reach the other locations.
Slump and slide structures are equally common in lithologic Subunit IIB, mainly at Hole 899B (e.g., intervals 149-899B-3R-4, 85-110 cm; 5R-3, 30-80 cm; 7R-4, 90-95 cm; 8R-3, 100-145 cm; Shipboard Scientific Party, 1994c). Small slump folds are formed in nannofossil claystone and silty claystone.
In Holes 897C and 897D, Subunit IIIB consists of gravity-flow deposits of poorly sorted, poorly cemented granule-to-pebble clayey conglomerate that grades upward to granule conglomerate and very coarse-grained, lithic, ferruginous clayey sandstone. This upward-fining sequence occurred over an interval of approximately 20 m in Hole 897C and in excess of 10 m in Hole 897D; the sequence was dark-reddish and variegated in both holes (Sawyer, Whitmarsh, Klaus, et al., 1994). Resedimented clasts included in the conglomerates have diverse lithologies and facies: white limestone, micritic limestone, marlstones, dolomite, and varied turbiditic arkosic-to-lithic sandstones. Minor clasts of shallow-water carbonates included in the conglomerates of Subunit IIIB are interpreted as being deposited from high-concentration sediment flows that were transported by high-density turbidity currents or as fluidized sand-silt-clay debris flows on relatively gentle slopes (Shipboard Scientific Party, 1994a). The basal, poorly sorted, conglomerate/gravel interval can be considered as a deposit from a single debris flow. Clasts and granules of basalt and serpentinized peridotite are present as minor components in the conglomerate and coarse sandstone. Samples at the base of this conglomerate yield an early late Eocene age (de Kaenel and Bergen, this volume).
At Site 899, reddish variegated ferruginous conglomerate and coarse sandstone, including basalt pebbles, were sampled in one interval of about 1 m within Subunit IIIB (Shipboard Scientific Party, 1994c). The Subunit IIIB lithologies at Site 899 are dated as middle to late Eocene by Kuhnt and Collins (this volume), using benthic foraminifers, or as Eocene by Gervais (this volume), using planktonic foraminifers (Fig. 3).
In lithologic Unit IV blocks and clasts of serpentinized peridotite were recovered at Site 897 (Shipboard Scientific Party, 1994a) and Site 899 (Shipboard Scientific Party, 1994c). At Site 899, three serpentinite breccias were recovered that contain peridotite blocks and intercalated Lower Cretaceous claystones and siltstones, clasts of basalt, microgabbro, and mylonite. These unusual breccia deposits from Site 899 did not have an obvious equivalent at Site 897. The mechanism of deposition of these blocks was discussed by Comas et al. (this volume) and Gibson et al. (this volume).
Comas et al. (this volume) examined the origin of the Lower Cretaceous deposits recovered at Sites 897 and 899, which were included in lithologic Unit IV (Sawyer, Whitmarsh, Klaus, et al., 1994), and considered the serpentinized peridotite breccia at Site 899 to be a tectonic cataclasite. The igneous-sedimentary complex at both sites is interpreted as an olistostrome, and breccia intervals are interpreted as olistoliths or blocks from cataclastic breccias. The terms "olistostrome" or "gravitational mélange" are widely used in the Alpine and Tethys literature (Abate, et al. 1970; Einsele, 1992) and describe thick, stratiform, heterogeneous, more or less chaotic, deposits that occur over wide areas and that accumulated as a result of tectonically induced massive gravitational sliding in various tectonic settings. These units can contain large blocks up to several hundred meters in thickness and can travel long distances.
The Leg 149 olistostrome, involving sedimentary, ultramafic and mafic rocks, is thought to be derived from a lower Aptian ridge related to an inferred transform fault boundary between the Galicia Bank margin and the Iberia Abyssal Plain at this time Figueiro Fault Zone (Whitmarsh et al, 1990). Cataclastic breccias and other fault rocks (basalt and gabbro), which form part of the olistostrome, mainly originated from wrench tectonics in the transform fault zone. Activity on the Figueiro Fault Zone occurred only during the development period for transitional crust and early stages of seafloor spreading in the Iberia Abyssal Plain (from approximately late Hauterivian, 129 Ma, after Gradstein et al., 1994) and terminated at the time of break-up on the Galicia Bank margin (latest Aptian-early Aptian, approximately 114 Ma).
Comas et al. (this volume) support the idea that the olistostrome originated at an active major fault far away from the drilling area and outside the ocean/continent transition because of the basin plain character of the early Aptian sedimentary environment at both Sites 897 and 899. The Aptian olistostrome was cored on discrete fault-bounded basement highs at Sites 897 and 899 (Sawyer, Whitmarsh, Klaus, et al., 1994). Comas et al. (this volume) suggest that these gravity slide deposits were originally deposited on a flat basin-plain environment and were uplifted during the postrift stage of the west Iberian Margin. The "long-distance transport" of the olistostrome is expressed by internal deformation, and by the roundness and alteration of the involved clasts. Strong affinities are present between the sedimentary facies involved in the olistostrome, the sediments drilled during Leg 103 at ODP Site 638 (Boillot Winterer, Meyer, et al., 1987), and a similar facies at DSDP Site 398 (Shipboard Scientific Party, 1979). Correlations were made between metabasite and gabbro clasts in the olistostrome and chlorite-bearing schists sampled on top of the peridotite in the Galicia Bank margin (Boillot et al, 1988; Beslier et al., 1993).
The hypothesis of Comas et al. (this volume) emphasizes the role of the Figueiro Fault Zone (Whitmarsh et al, 1990) as a major tectonic lineament during the Aptian. The reality of this fault zone is a matter of discussion because of its weak signature on seismic profiles (Beslier et al, 1993). The present-day seismic signature from the paleotransform fault may be masked either by a later transtensional situation or by post-Aptian faulting of the western Iberian Margin (Comas et al., this volume).
Gibson et al. (this volume) interpret the lower Aptian serpentinite breccia unit at Site 899 as being a giant submarine landslide, generated by slope failure on a large, nearby serpentinite fault scarp. Their interpretation focuses on the breccia unit at Site 899. Giant landslides are a typical feature of regions undergoing rapid extensional deformation (Bonatti et al., 1973; Bonatti et al., 1974; Bernoulli and Weissert, 1985). It is proposed that the basement topography at the time of formation of the serpentinite unit differed significantly from the present topography buried beneath the Iberia Abyssal Plain, and, as a result, the location of the source serpentinite escarpment is unknown. The authors indicate that the source was likely to have been within several kilometers of Site 899.
Gibson et al. (this volume) discuss the following possibilities for the mechanism of formation of the breccia units: in situ brecciation, formation in cataclastic shear zones, diapiric extrusion, and mass- flow deposits. They conclude that the breccia units at Site 899 show many similarities to subaerial landslides and associated rock avalanche deposits (Watson and Wright, 1969; Siebert, 1984; Stoopes and Sheridan, 1992; Yarnold, 1993). These similarities include the presence of angular fragments with jigsaw/crackle texture; a general lack of sorting, compatible with very rapid, essentially instantaneous formation, but with larger boulders occurring near the upper surface of the units; a matrix generated by fragmentation of the clasts and a size continuum between the clasts and the matrix; and internal deformation zones or slip zones. Observations at Site 899 and the general stratigraphic framework for Unit IV (Shipboard Scientific Party, 1994c) show that the serpentinite breccias are interbedded with sedimentary units. Biostratigraphic ages suggest a normal stratigraphic succession (de Kaenel and Bergen, this volume) and, therefore, the breccias are most simply interpreted as normal bedded units. The source for the serpentinite breccias was presumably a nearby fault scarp exposing massive serpentinite, but Gibson et al. (this volume) do not provide any distance relations.
Wilson et al. (this volume) could find no indications from seismic profiles to suggest that basement topography established during rifting was modified by later deformation. Seismostratigraphic unit 6 of Wilson et al. (this volume) is of Aptian to possible Valanginian age, which increases the difficulty of interpreting the setting of the Aptian debris flows and rock-fall breccias encountered on the crest of basement highs at Sites 897 and 899. There is no evidence for the basement ridge at Site 897 being tectonically uplifted or rising diapirically during deposition of the seismostratigraphic units 1 to 5. This conclusion contradicts suggestions made by Comas et al. (this volume) and Gibson et al. (this volume) that the basement topography that existed at the time of formation of the Aptian olistostromes at Sites 897 and 899 differed significantly from that observed today (see also Whitmarsh and Sawyer, this volume). Debris-flow deposits, such as those occurring in lithologic Subunits IVA and Subunit IVB at Site 899 (Shipboard Scientific Party, 1994a, c), require only gentle (<1°) slopes for their formation (Stow, 1994) and thus could have been generated on the basement topography observed today. However, the rock-fall origin (Gibson et al., this volume) for the serpentinite breccias of Subunit IVA at Site 899 implies the existence of steep slopes nearby (Stow 1994). As this site was only surveyed by JOIDES Resolution using low-resolution single-channel seismic survey, its topographic setting is not well constrained.