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LATE QUATERNARY TURBIDITE EMPLACEMENT ON THE HORSESHOE ABYSSAL PLAIN (IBERIAN MARGIN) SUSANA M. LEBREIRO1*, I. NICHOLAS MCCAVE1, AND PHILIP P.E. WEAVER2 1

Instituto de Ciencas del Mar, CSIC, P. Juan de Borbon, s/n, 08039 Barcelona, Spain Southampton Oceanographic Centre, Empress Dock, Southampton SO14 3ZH, U.K. * Present address: Southampton Oceanographic Centre, Empress Dock, Southampton SO14 3ZH, U.K. 2

ABSTRACT: The Horseshoe abyssal plain (HAP) is located on the Iberian margin at approximately 4800 m water depth, and is confined within topographic elevations with a relief of about 3000 m, with the exception of Gorringe Bank, which rises to only 20 m below sealevel. The plain is built up by an alternation of turbidites of the order of meter-thick beds, numbered H2, H3, H8, and H13, contrasting with sets of thinner beds of the order of decimeters thick (H4,5,6,7 and H9,10,11,12), and pelagites centimeters in thickness. An intensive study of eight piston cores, using visual observation (color and thickness), relative stratigraphic position of units, magnetic susceptibility logs, calcium carbonate content, and mineralogy of turbidite bases, concludes in a bed-bybed correlation of all individual turbidites. The major source of terrigenous material feeding the HAP is the Saõ Vicente canyon, which incises the Portuguese shelf, while minor sources are the surrounding seamounts. The elongate geometry of the abyssal plain with its single dominant source of sediments produced laterally continuous deposits from both large and small flows. These covered the entire 356 km length of the plain. Grain-size analysis of the four thickest turbidites (H2, H3, H8, and H9) demonstrates only slight downcurrent fining and vertical grading, expressed best in the ratio of coarse silt to fine silt plus clay. A small amount of sand is carried the length of the plain. The thicker units all show the same pattern of thickness, with a maximum in the middle of the plain around a topographic constriction and bend. This is most plausibly explained as due to reduction in flow speed of an initially supercritical flow causing enhanced deposition. Some of the beds appear to have a double coarse layer in the base, which may indicate partial reflection of flows from the side of the basin. It is suggested by application of equations for flow behavior that both thick (; 3 m) and thin (; 0.3 m) beds are due to supercritical flows a few tens of meters high. However, the thick beds resulted from high-concentration flows (Cv ; 4% by volume) whereas thinner beds require low concentration (Cv , 1%) to run out over the full length of the basin. The stratigraphy is tied into the dated oceanic pelagic record by analysis of foraminifera in the pelagic layers above the turbidites and through recognition of two Heinrich layers (H-1 and H-2, ages 14.3 and 21 ka). The resulting age framework shows higher turbidite frequency in the glacial (2.7/kyr) than interglacial (Holocene) (1.0/kyr). This also gives higher mass flux during the glacial. Emplacement of turbidites cannot be clearly related to sea-level changes but may well be due to seismic activity. However, one of the largest earthquakes in human experience (Lisbon in 1755) triggered only a thin turbidite, invalidating the term ‘‘seismite’’ for thick turbidites. INTRODUCTION

Problems of Abyssal-Plain Sedimentation Pilkey (1987) said that occasional giant flows covering the entire plain were necessary to maintain a basin plain, which otherwise would contain a series of coalesced fans with local relief. Only very thick turbidites have the capacity to smooth the bottom topography, by masking the local marginal relief, individual or coalesced fans, and regional slopes. Individual JOURNAL OF SEDIMENTARY RESEARCH, VOL. 67, NO. 5, SEPTEMBER, 1997, P. 856–870 Copyright q 1997, SEPM (Society for Sedimentary Geology) 1073-130X/97/067-856/$03.00

turbidites deposited by a single flow can contain significant volumes of sediment in excess of 100 km3 (‘‘Black Shell turbidite’’ from the Hatteras AP; Elmore et al. 1979), 120 km3 (‘‘turbidite f ’’, from the Madeira AP; Weaver and Rothwell 1987), 160 km3 (turbidite triggered by the Grand Banks earthquake in 1929, Sohm AP; Heezen and Ewing 1952), or turbidite H13 in the Horseshoe Abyssal Plain (HAP), which reaches a volume in excess of 33 km3 (this paper). On the other hand, the existence of topographic features determines the paths of distribution of sediments, and depressions favor thickening of turbidites. Furthermore, there is often a relationship between thickness of beds and entry points: giant flows frequently arrive via submarine canyons or adjacent continental margins whereas local, thin flows originate on the surrounding elevations. The buildup of the HAP, as with many other abyssal plains, consists of a sequence of alternating turbiditic and pelagic beds. Turbidites include sand, silt, mud, and biogenic types (nomenclature after Stow and Piper 1984), and pelagites/hemipelagites consist of marls (60–20% CaCO3) and clays (, 20% CaCO3). Pelagic/hemipelagic sediments are composed of mainly biogenic calcareous skeletons (foraminifers, nannofossils, and rare echinoderms and ostracods) and a few siliceous skeletons (radiolarians, sponges, and diatoms), wind-blown dust, and ice-rafted debris. Turbiditic units represent most of the volume of the basin fill, but pelagic intervals are of major importance in providing the dating framework. The major factors controlling the pelagic sediment composition are the position of the CCD (calcium carbonate compensation depth), the deep oceanic circulation, and the organic productivity of the surface waters. Outstanding problems in abyssal-plain sedimentation that can be attacked using the data from HAP are the magnitude and frequency of turbidity currents and the dynamics of these flows. In particular, what were the most plausible values of concentration, flow speed, thickness, and Froude number, which indicates supercritical or subcritical conditions? The evidence of the volume of individual units and their grain-size characteristics are important inputs into such an evaluation. The present paper examines these problems from an enclosed-basin setting. Geological Setting of the Horseshoe Abyssal Plain The Horseshoe abyssal plain (HAP) is located southwest of the Iberian margin in the subtropical eastern north Atlantic Ocean (Fig. 1). The Saõ Vicente canyon incises the 20 km wide Portuguese shelf at a latitude of 378N, just north of Cape Saõ Vicente, and debouches onto the plain. The HAP is a narrow and elongate basin, 356 km long by 64 km wide, with a surface area of 21,157 km2 below the 4800 m isobath. In the middle of this plain, a strait 11 km wide divides the west and east subbasins. At the western end of the basin the sill depth is 4150 m, 720 m above the floor of the western basin. The average water depth for both subbasins is 4870 m. The HAP is surrounded by the Gorringe Bank to the north, the Madeira– Tore Rise to the west, and the Coral Patch and Ampere Seamounts to the south. The trends of these topographic highs reflect a complex geological structural control related to the proximity of the African/Eurasian plate boundary and the Early Cretaceous rifting of the North Atlantic. The onset of seafloor spreading in the Tagus Abyssal Plain, to the north, has been dated at ca. 133 Ma (Pinheiro et al. 1992; Whitmarsh and Miles 1995; Pinheiro et al. 1996). Block faulting in the basement of the HAP is therefore likely associated with the early rifting history of the North Atlantic,

QUATERNARY TURBIDITE EMPLACEMENT, HORSESHOE ABYSSAL PLAIN

857

FIG. 1.—Location of the Horseshoe Abyssal Plain (HAP) and the eight cores studied. Coordinates of each core are given in Table 1. The two solid lines indicate the seismic profiles shown in Figure 2. Isobaths are in kilometers. The inner isobath in the HAP is of 4.8 km. Small numbers attached to circles in the HAP indicate the position of Hoyt (1976) cores referred to in the text.

but its precise nature is not yet known. Multichannel reflection seismic profiles obtained during cruise SONNE-75, in 1991 (Fig. 2) show a smoothly diffractive basement with well defined block-like features of the order of 20 km wide and a monotonous record of acoustically layered turbidites along the plain (Roeser et al. 1992). Within the horst-and-graben complex, the graben contain up to 1.8 km thickness of sediment infilling (Fig. 2). Previous Studies on the HAP The only previous work concerning the HAP is that of Hoyt (1976) and Hoyt and Fox (1977), who reported the correlation of six individual turbidites over 300 km, based on nine cores ranging from 4307 to 4709 m depth (now known to be 4870 m). Some of these were located closer (within 10 km) to the margins of the plain than cores used in this paper. METHODS

The eight piston cores used for this study (Table 1) were collected during cruise R.R.S. Discovery 187 (20 October–20 November 1989). The cores are distributed on an E–W transect. Once collected, the cores were split, photographed, and stored in sealed plastic containers at a temperature of 48C in the Institute of Oceanographic Sciences–Deacon Laboratory repository, U.K. Thus, when later described and sampled the cores showed an unadulterated aspect, with the exception of a black to grayish-green color change due to oxidation of organic-rich turbidites in Core D11942P. Descriptions comprised color, texture, structure, type of contacts, degree of bioturbation, and appreciation of disturbance during recovery.

Volume magnetic susceptibility was measured with a Bartington Instruments MS 2 magnetic susceptibility meter. Measurements were taken at a constant interval of 2 cm, doing an air blank between each real measurement. Because magnetic susceptibility is sensitive to temperature variation, the accumulated drift is overcome by subtracting the recorded blank value from the measured sediment value. Data were recorded inside the cool store at a temperature of 48C. The carbonate content of Core D11948P was analyzed at the Instituto Geolo´gico e Mineiro (I.G.M.), Portugal, using the Hulsemann (1966) gasometric method, which gives an accuracy in replicated analyses of 6 2%. To analyze the grain size of individual turbidites, samples taken from variable depth intervals were disaggregated by shaking for 24 hours with sodium hexametaphosphate (0.1 vol %). The sand fraction (. 63 mm) was then separated by wet sieving. The rest of the fine fraction, also dispersed with Calgon (0.1%), was gently shaken by using an automatic wheel shaker for 24 hours plus immersion in an ultrasonic bath for 2 minutes prior to analysis of the samples in a SediGraph 5000 ET (Micromeritics 1984; Jones et al. 1988; Coakley and Syvitski 1991). The mineralogy of the bases of turbidites and of the Portuguese continental shelf were examined using light microscope and heavy-mineral slides (slides provided by the I.G.M.). In both cases, the sand fraction selected was 3–4 f. Heavy minerals were separated by density with acetylene tetrabromide (C2H2Br4, density 5 2.87 g/cm3), and mounted with Arkanson resin (R.I. ø Canada balsam). On the Portuguese shelf, bottom samples from 11 transects, perpendicular to the coast, located to the north/south (western shelf) or west/east (southern shelf) of each canyon, and containing samples from three different depths (shal-

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S.M. LEBREIRO ET AL.

FIG. 2.—Infilling of the Horseshoe Abyssal Plain (HAP) from seismic profiles recorded during cruise SONNE SO-75, 1991. Top, profile SO75-003; bottom, profile SO75-002. See Figure 1 for locations of tracks.

lower than 50 m, 50–100 m, and 100–200 m) were examined (Fig. 3). The sampling includes the three canyons that feed the HAP and the Tagus Abyssal Plain (Fig. 3). Because of abundance of micas and to prevent skewing the results, micas were excluded when calculating the percentage of mineral species. On the plain, samples from all the turbidite sand bases with a sufficient volume of sediment to allow a microscopic mineral analysis were collected from three cores: D11944P, located in the west subbasin; D11941P, in the central part of the plain; and D11948P, in the east subbasin. In both cases, the limited number of grains available in each slide, frequently less than 100, is insufficient for a satisfactory statistical analysis. Rather, a semiquantitative analysis was undertaken. For the X-radiographs, slabs of sediment 1 cm thick were manually sliced from the surface of half core sections, and an X-ray cabinet Faxitron (Vinten Instruments) used with the following sets: 80 kVP, 0.3 min of exposure, 1.5 min developing , 8–10 min fixing, and 20 min washing. TABLE 1.—Coordinates of the eight piston cores used in this study. Core Number

Latitude (N)

Longitude (W)

Depth Corrected (cm)

Core Length (cm)

D11944P D11942P D11941P D11947P D11946P D11948P D11949P D11950P

35844.49 35847.29 35843.79 35836.39 35836.39 35853.39 35800.09 36813.69

13837.79 12852.29 12821.09 11847.39 11814.99 11811.79 10843.99 10829.59

4865 4875 4875 4870 4872 4870 4865 4850

823 878 (164) 908 450 700 730 550 700

Value in parenthesis indicates existence of trigger core necessary in this study to complete the record available from the piston core.

Turbidites are numbered from the surface downwards as H1, H2, etc., and intervening pelagic layers as Pn, where n is the number of the underlying turbidite. Additional lettered suffixes were used for local turbidites (for example, H12a); H stands for the Horseshoe Abyssal Plain. For foraminiferal analysis a minimum of 300 planktic specimens were counted in the grain-size fraction . 150 mm. The benthic species, other biogenic groups (echinoderms, ostracods, gastropods, radiolarians, and sponges) and detrital grains encountered in the same split were counted separately. Partially dissolved or broken specimens were excluded. Foraminifera species were grouped into climatic assemblages following Be´ (1977). These data are available in Lebreiro (1995) or upon request to the first author. RESULTS

Turbidite Sedimentation Provenance of Turbidites.—Examination of heavy minerals from both the Portuguese continental shelf and the bases of turbidites in the HAP shows close similarity in time and space between the two domains, though with a few compositional differences. The Portuguese shelf is compartmented by several canyons, and if different turbidites were sourced from different compartments they should have distinctive heavy-mineral assemblages, as long as the canyons were supplied from one side with distinctive mineral assemblages. The results from this shelf suggested, however, a continuous platform with a single mineral assemblage homogenized by longshore currents. The following few peculiarities (Fig. 3, Table 2) could be outlined: (1) abundance of tourmaline, andalusite, and staurolite along the entire shelf stud-


859

TABLE 2.—Correlation of heavy-mineral species between the source (Portuguese Shelf) and the Horseshoe Abyssal Plain (HAP). HAP

Portuguese Shelf

Biotite Clorite Olivine Garnet Andalusite Sillimanite Kyanite Staurolite Sphene Topaze Zircon Idocrase Epidote Tourmaline Orthopx (Hyp) Clinopx (Mg) Actinolite Hornblende Tremolite Rutile Corundum Monazite Apatite Glauconite Sheelite Cassiterite Anatase Opaques

Western

Southern

. * * . . * * . ,, * , ,, , . . * * *

. * * * .

, ,, * ,, * ,, ,, ,, ..

44P

41P

48P

.

* *

. * * *

* .

no staurolite *

* * , . * * * * ,, ,, . ,, * ,, ..

* * * *

*

*

* * . * * *

* no glauconite

.

.

,, 5 very rare; , 5 rare; * 5 present; . 5 abundant; .. 5 very abundant.

FIG. 3.—Locations of heavy-mineral bottom samples (shown by open circles) from the Portuguese Continental Shelf used to study the provenance of turbidites to the Horseshoe Abyssal Plain. Complete reference of location coordinates are in Lebreiro (1995). Isobaths are in meters. TAP, Tagus Abyssal Plain; HAP, Horseshoe Abyssal Plain; SAP, Seine Abyssal Plain.

ied; (2) exclusivity of sillimanite in the western shelf and predominance of hypersthene and garnet in the western shelf in comparison with the southern one; (3) higher percentage of monazite in the Portimaõ canyon–Faro canyon segment; (4) higher percentage of micas in both the Sado canyon–Saõ Vicente canyon segment and Saõ Vicente canyon–Portimaõ canyon segment; and (5) increased abundance of glauconite with distance from the coast (samples from 100–200 m water depth), as earlier reported by Monteiro et al. (1983) and Alveirinho Dias and Nittrouer (1984). In the HAP, the core closest to the Portuguese margin (D11948P) apparently recorded a higher variety of mineral species (Table 2). The mineral composition of the sands in general closely resembles that of the Portuguese continental shelf. The suites comprise minerals from a wide range of rocks, from low-grade to high-grade metamorphic, sialic and mafic igneous, pegmatites, and authigenics, reflecting the variability of the Iberian eroded outcrops (Teixeira 1972). When turbidites are compared downcore, homogeneity is the main characteristic, with absence of any particular mineralogical marker. Consequently, one can infer a common origin for all the turbidites examined. Absence of certain mineral grains in Cores D11944P and D11941P at the end of the transport path is size dependent. The correspondence between the heavy-mineral species in the shelf and the bases of turbidites undoubtedly confirms the provenance of turbiditic flows from the Portuguese shelf but does not specify a particular area of the shelf. The presence of glauconite only at the shelf break and its absence in the plain suggests the origin of turbiditic flows in the canyon comprising sediment derived from the inner part of the shelf, rather than outer shelf/ upper slope failure.

Bed-by-Bed Correlation of Turbidites.—The criteria of relative stratigraphic position, thickness of visually distinct beds, and magnetic susceptibility signature are used to correlate individual turbidites. Horizontal and vertical variations in grain size and the petrographic character of the bases provide additional, useful information. The results show a reliable correlation of every single turbidite bed in the HAP, reflecting a laterally continuous sedimentation derived almost exclusively from the Saõ Vicente canyon source. A typical turbidite consists of a sandy silt base a few centimeters thick, which may be structureless, parallel-laminated or cross-laminated, the grain size fining upwards and with sharp lower contact. Above it lies dark olive green (Munsell color value 5G 7/1) homogeneous mud, a decimeter to a meter in thickness, which is poorly to extensively bioturbated towards the top and has a gradual upper contact with the commonly overlying pelagic unit (Fig. 4). The brownish color (10YR 5/4) in the upper part of some turbidites indicates postdepositional oxidation in situ of the organic matter contained in originally olive-green units. Four thicker beds of order one meter in thickness (H2, H3, H8, and H13, Table 3) contrast with sets of thinner beds of the order of decimeter thickness (H4,5,6,7 and H9,10,11,12). Moreover, the thick beds (H2, H8, and H13) have diagenetic, colored oxidation fronts explained by downward migration of metals regulated by a redox gradient. These banded fronts are common in the Madeira (Colley et al. 1989; Wilson et al. 1986) and Cape Verde Abyssal Plains (Colley et al. 1984; Wilson et al. 1985; Thomson et al. 1989; Thomson et al. 1993). The fact that most of the turbidites cover the entire plain, irrespective of derivation from giant flows or smaller flows, also facilitates the correlation. The oldest record appears exclusively in the westernmost core, D11944P, with turbidites coded H21,20,19,18,17,16,15,14 (Fig. 4). Each pair of turbidites is separated by a pelagic unit except between H19 and H18. Above this set, the largest turbidite, H13, covers the entire plain with a volume estimated as . 33 km3. Next, a local, 0.12 m thick biogenic turbidite (Hc) containing abundant corals appears in the eastern subbasin (Core D11948P). It originates from Gorringe Bank to the north, which rises up to 20 m below sealevel. This turbidite could possibly be correlated with a foraminifer-rich turbidite emplaced at a similar position in Core D11946P, although corals

FIG. 4.—Correlation of individual turbidites in the Horseshoe Abyssal Plain. The profiles used to show the correlation are of magnetic susceptibility (and CaCO3 only for Core D11948P), but several criteria were considered (see text). The four larger turbidites that were the subject of detailed grain-size analysis are outlined by a shaded pattern. H, turbidite; P, pelagic; Hc is a local calcareous turbidite.

860 S.M. LEBREIRO ET AL.


861

TABLE 3.—Sediment volume of the five largest flows (above) and the smallest flow (below) from the Horseshoe Abyssal Plain. Turbidite

Thickness (m)

Sediment Volume (km3

H2 H3 H8 H9 H13

0.6 0.7 1.0 0.3 .4.9

3.5 4.1 5.8 2.0 .32.8

------------------------------------------------------------------------------------------------------------------------------

H7

0.12

0.7

The surface of the HAP is 21,157 km2 (isobath of 4800 m). Sediment volume includes a porosity of 0.69– 0.79 g/cc (see Table 7) at a standard density of 2500 kg/m3.

are rare in the latter. Then, alternating with hemipelagites, canyon-derived relatively thin turbidites cover the whole eastern subbasin but are limited to the entrance of the western subbasin (Core D11941P); these are H12a,b, H12, H11a, H11, and H10a. The succeeding turbidites, H10 and H9, although having the same character as the previous thin ones, reach both subbasins and do not have an intervening hemipelagite. Overlying H9 and P9, one more large turbidite, H8, was deposited. Following this, a set of four thin turbidites (H7,6,5,4) are present. These were initially emplaced with alternating pelagics, although the lithological record has preserved only two (P5 and P7 have been partially eroded). Two final large turbidites were then emplaced (H3 and H2). Although no record of P3 between them was found, vestiges of bioturbation on the top of H3 with burrows containing pelagic sediment indicate its deposition but complete erosion by the flow depositing H2. The most recent turbidite, H1, probably covers the entire surface of the plain, although this part of the record is missing in the westernmost core, D11944P. Finally, the currently accumulating pelagic layer P1 overlies H1. Core D11942P lacks the record between H5 and H3, and Core D11944P lacks the period from H7 to present, most likely because of loss during coring. The larger turbidites (H2, H3, H8, H9, and H13) show maximum thickness in the cores close to the middle plain constriction (D11946P, D11947P, D11941P; Figs. 5, 6). In Core D11950P, located in the canyon

FIG. 6.—Downcurrent variation of thickness for the four major turbidites penetrated in the Horseshoe Abyssal Plain. Asterisk (*) indicates data from Hoyt (1976).

mouth, the identification of individual turbidites for correlation with the rest of the cores is almost impossible on account of the presence of numerous very thin beds, some of which could have originated from the Gorringe Bank because they show more common but still rare corals. Deposition of this type of sequence is expected in a core situated at the mouth of a canyon, where large turbidity currents pass by leaving just a trace of sediment or cause erosion, and where small turbidites unable to reach the whole plain deposit their load. The compositional biogenic or lithic character of the turbidites in the HAP is shown by the abundance of foraminifers or detritals (mainly quartz and micas), respectively. Most of the turbidites, however, have a mixed composition. High-resolution downcore magnetic susceptibility provides information on the physical properties of the sediment. This parameter measures the capacity of particles to be magnetized by a local and weak magnetic field. Thus, this technique reveals the presence and proportion (biogenic to lithogenic) of magnetizable detritus, such as continent-derived or ice-rafted material. The logs reflect the input of coarser rock-fragment grains derived

FIG. 5.—Isopachs (in centimeters) of the largest turbidites—H2, H3, H8 and H9—recorded everywhere over the Horseshoe Abyssal Plain.

862


from the Iberian margin, which settle preferentially at the base of the turbidites. Additionally, the material just below the oxidation fronts has a characteristic low magnetic susceptibility value, probably because of reduction or removal of iron, which can be used as a correlation feature. The magnetic susceptibility profiles in Figure 4 also bring out (1) the persistently high magnetic susceptibility values at the base of H9 and H10 throughout the plain, due to the presence of heavy minerals, and (2) a pronounced step of magnetic susceptibility values in the core closest to the mouth of Saõ Vicente canyon (D11950P), differentiating two sets of values downcore. The latter contrast is explained by the different climatic conditions in which the turbidites were emplaced: low values correspond with interglacial stage 1 and high values with glacial stage 2. This correlation is consistent with the close relation between magnetic susceptibility measurements and climatic interpretations suggested by Robinson (1986). The calcium carbonate content of the turbidites is determined by the calcareous material available in the source area and the dilution by noncarbonate material. The calcareous composition of some bases clearly indicates a seamount origin, as is the case for Hc (just above H13) in Cores D11948P and D11946P, with 38–80% CaCO3, increasing towards the base of the unit. A calcium carbonate profile was obtained for HAP Core D11948P (Fig. 4), in which the ungraded mud part of the turbidites shows a constant content of carbonate, with the exception of the upper bioturbated mixed layer. Pelagic units gave CaCO3 values of 27–30% during glacials and 30–47% during interglacials, whereas the turbidites showed a lower and constant value of 23–28% for the muddy main body and 30–35% for the bases. P11 contains a Heinrich layer and shows the expected lower value of 22–23.8% due to dilution by detrital material. In summary, the three criteria above, when integrated, yield a bed-bybed correlation of turbidites over the whole HAP, as shown in Figure 4, with a few exceptions for local, calcareous turbidites. Nevertheless, it should be kept in mind that detailed bathymetry shows small fans emplaced on the borders of the plain, probably composed of biogenic-rich sand turbidites. ‘‘Correlatability’’ of individual beds depends principally upon the volume of sediment supplied and the number of entry points. In the Columbus Abyssal Plain (SW North Atlantic), for example, the continuity of sand layers is limited because flows are small and introduced from any one of three sides (Pilkey et al. 1980). Accordingly, lateral continuity is expected in the HAP turbidites because only one main source, Saõ Vicente canyon, provides large flows that have accelerated down from the shelf and have run out over the entire narrow and elongate plain. Lateral and Vertical Grain-Size Distribution.—Abyssal plains are the final destination for much fine terrigenous sediment. Grain-size analysis demonstrates the lateral variation of a single turbidite over hundreds of kilometers and the vertical variation within an individual turbidite. It also allows location of entry points of sediment to the plain, and delineation of paths of transport and preferred sites of deposition. The information obtained from a deposit can be further used to characterize the parent flow. Individual silt turbidites (. 70% silt-size particles) very rarely show the complete Bouma grading sequence; more often they contain only divisions C, D, and E or D and E, both here and on the Madeira Abyssal Plain (Jones et al. 1992). The nomenclature of Stow and Shanmugam (1980) for fine turbidites is therefore applied here. Mud turbidites (50–80% mud) are commonly present with either base-cut-out units T4 to T8 or top-cut-out units T0 to T4 (Fig. 7). Correlation of individual turbidites between the cores presented here and with those of Hoyt (1976) indicate lateral continuity of single turbidites covering the entire HAP. Turbidites H2, H3, H8, and H13 are quite clear in Hoyt’s correlation diagrams, and H13 reaches a thickness in excess of 4.8 m in his Core V27-148. In the present study, the four thickest turbidites that can be identified in most of the cores, namely H2, H3, H8, and H9, were selected for detailed grain-size analysis (Fig. 4, Table 4). Although H13 is the thickest turbidite

deposited, it was not fully penetrated in four of the six cores where it is present (D11941P, D11946P, D11948P, and D11949P) (Fig. 4). The downcurrent grain-size variation of turbidites 2 and 8 shown in Figure 8 is representative of the four turbidites analyzed, clearly indicating a unique and expected direction of transport: from east (canyon mouth) to west. In the HAP, sand appears in the bases of all turbidites but decreases gradually from the source to the most distal point, although in all the cases some sand reaches the very end of the plain. In general, the thickness of the basal sand ranges from 0.25 m in Core D11949P, averages 0.12 m in the middle of the plain (Core D11948P), and decreases to , 0.01 m in the distal cores (D11941P, D11942P). Exceptionally, for turbidites H8 (Fig. 8) and H9, the proportion of sand diminishes at the strait located in the middle of the plain (Cores D11941P and D11947P for H8 and only D11941P for H9), and increases again to the west. With the decrease of sand away from source in the bases of turbidites, silt proportionately increases (Fig. 8). The vertical granulometric distribution of the deposit shows a basal concentration of the coarser material, i.e., all the sand and part of the silt, and a relatively thick main body of silt and clay in a constant ratio. Clay is scarce in the base (, 20%). Vertical grain-size variation for turbidite H8 shows two abrupt steps of major grain-size change with depth (Transitions 1 and 2 in Figure 9). ‘‘Transition 2’’ occurs within the base, and marks the upper limit of sand. ‘‘Transition 1’’ is the boundary between variable (below) and constant (above) proportion of silt in the fine fraction (40%). Sedigraph analysis of the 10–4 f (silt and part of the clay) fraction shows well sorted silt in the base (mode 5–5.5 f), almost always fining upwards to modal size ø 6 f. This mode is fairly constant and can be traced through the entire base. ‘‘Transition 1’’ is a boundary above which polymodal grain-size distributions are present, although most of the modes are not at consistently similar sizes between samples to permit definition of trends. Several samples display a consistent mode at 6.8 6 0.1 f (9.0 mm) (see Fig. 9, especially D11942P and D11941P), which may be related to a particular component, as suggested by McCave et al. (1995). A similar pattern is also found for turbidites H2, H8, and H9 in the HAP (data in Lebreiro 1995), and for three turbidites (D, B1, and B) in the Madeira Abyssal Plain (Jones et al. 1992). Double Bases.—Occasionally some turbidites present a ‘‘double coarse base’’, i.e., two sets of coarser layers, inversely graded, separated by structureless, nonbioturbated mud material as part of a single turbiditic event; of the two coarser layers the lower contains sand- and silt-size particles whereas the upper often has only silt. Examples are in possibly all cores except D11944P for H8 (Figs. 8, 10). This abnormal vertical size distribution within the turbidite facies is interpreted as caused by reflections due to basin topography, as previously invoked by other authors for similar deposits (Kneller 1991). Turbidity currents confined to a basin are subject to reflections due to topography in the center or at the borders of the basin. This may lead to fluctuations in either arrival rate of new material or remobilization of recently deposited material at the bed. Pelagic Sedimentation Although pelagic units are volumetrically insignificant in the overall lithological column, they provide microfossils, which are not only useful paleoclimatic and paleoceanographic indicators but also a chronostratigraphic tool for correlation proposes. Interglacial/Glacial Succession and Correlation of Pelagic Beds.— Identification and counts of planktic foraminifers from each pelagic unit in the record enabled their grouping into bioclimatic assemblages and construction of a biostratigraphy. Climatic foraminiferal assemblages (Be´ 1977), controlled principally by the temperature of the surface waters where the organisms lived, allow one to infer the succession of glacial and interglacial periods. Knowledge of both faunal assemblage succession and oxygen isotope composition of foraminifera in a type pelagic section on the nearby Tore Seamount (Lebreiro et al. 1996) serves as a template for the


863

FIG. 7.—X-radiographs representing typical turbidite facies encountered in the Horseshoe Abyssal Plain. Core D11949P/S1 shows a ‘‘double coarse-sand/silt base’’, likely part of the same turbidite and emplaced during a single event. D11950P/S3 shows a complete turbidite sequence from the bottom sandy base to the top gradational contact with the overlying pelagic unit. Nomenclature is after Stow and Shanmugam (1980).

climatic assemblages of the HAP pelagic units. These are tied into the oxygen isotope stage stratigraphy of the late Quaternary and its temporal scale (Martinson et al. 1987). Because of the reliable correlation between individual turbidites as detailed above, and thus that of the intervening pelagic units, three cores were sufficient to construct the age model: D11948P contains the more recent record (above emplacement of H13), and Core D11944P completes it back in time (earlier than emplacement of H13), to indistinguishable glacial stage 2–3 (Fig. 11). Coccoliths do show agreement (van Niel, personal communication) with limits inferred from foraminiferal climatic assemblages. In the third core, D11950P, which is difficult to correlate with the rest of

the cores in the plain, foraminiferal distribution and bioclimatic assemblages indicate glacial stage 2 as the oldest age (Fig. 11). Correlation of Heinrich Layers between the Tore Seamount and the HAP.—The presence of Heinrich layers in the HAP cores is of crucial significance when developing an age model, because they contribute key time horizons. Heinrich layers are influxes of ice-rafted detritus (IRD) originating from the Laurentian ice cap (and possibly Greenland and Norway) and disseminated by icebergs in the North Atlantic during the last glacial period (Bond et al. 1992). Layers of IRD resembling those identified as Heinrich layers at the Tore Seamount to the north (Fig. 1; Lebreiro et al. 1996) were found within specific pelagic intervals in the HAP. These levels

TABLE 4.—Thickness of turbidites in the HAP, in centimeters. VEMA 27 (Hoyt 1976)

DISCOVERY 187 (this paper) 44P

42P

41P

47P

H2 H3 H8 H9

— — 49 14

29 — 136 42

80 97 134 55

93 81 123 —

H13

419

555

.352

—

46P 57 (TC) 94 126 53 .132

TC, trigger core; addition indicates possibly two episodes of a unique event.

48P

49P

148

150

151

152

153

154

37 1 9 40 73 29

42 38 57 6 12 10

20 70 103

57 65 99

45 52 92

32 48 60

28 28 37

67 82 94

.329.5

.149

.500

.220

.170

.280

.50

.380

864


FIG. 8.—Lateral distribution of grain-size percentage of Turbidites 2 and 8 across the Horseshoe Abyssal Plain. These turbidites are representative of the four largest turbidites analyzed, H2, H3, H8, and H9. Depths used in the diagram are relative to the depth in the main core section from where samples were collected; negative depths refer to the preceding section.

contain abundant detrital minerals, mainly transparent quartz, some reddish feldspar, and minor dark grains. Only two Heinrich layers were detected in Cores D11948P and D11944P, and one in Core D11950P. This is because of the short period recorded in the two cores, back to the last glacial stages 2–3, (Fig. 11), or possibly because some levels were missed either by erosion or missampling. Rates of sedimentation (Table 5) and erosion (Table 6) and frequency of turbidite emplacement (Fig. 11, Table 7) can only be estimated within a time framework. DISCUSSION

Depositional Mechanisms, Lateral Thickness Variation, and Ponded Turbidites Two mechanisms have been proposed for emplacement of fine-grained turbidites: either by low-concentration flows (Piper 1978; Stow 1979; Stow and Bowen 1980; Stow and Piper 1984; Stow and Shanmugam 1980) or by high-concentration flows (Elmore et al. 1979; McCave and Jones 1988). McCave and Jones (1988), assuming a high-concentration flow of 100 kg m23, and a slope gradient of 1:5000, estimated a steady speed of 0.65 m s21 during deposition (which corresponds to the velocity for a subcritical flow, Fr , 1.0). They argue that a low-concentration flow (2 kg m23)

would require a flow lasting about six weeks, to deposit a 5 m thick turbidite, which is unrealistic. More recently, Dade and Huppert (1994, 1995) have developed a model to predict runout distance and flow duration of a turbidity current and the resulting thickness distribution of the deposit, depending on the initial conditions of grain size, concentration, and thickness. The model applies to deep-water suspension gravity flows that are initially channelized and subsequently laterally constrained. When the flow enters the plain, it collapses by self-weight and decelerates on a sea floor of negligible slope. The flow is assumed to be primarily deposit-forming, and so reworking of newly deposited material is not considered. The principal governing equations of Dade and Huppert used here are for runout distance xr, xr 5 3(g9o qo 3 /ws 2 )1/5 and runout time tr , tr 5 2(qo2 /g9o ws 3 )1/5 where qo is the initial volume of suspension per unit width of the flow, g9o is initial reduced gravity (gCvoDrs/r), where g is acceleration due to gravity, Cvo is initial volume concentration, and Drg 5 rs 2 r, where rs and r are densities of sediment and fluid, and ws is average settling velocity of


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FIG. 9.—Vertical variation of silt grain-size (10–4 f) for Turbidite 8, representative of the four largest turbidites analyzed with SediGraph. Large black dots mark the basal mode (5–5.5 f) and small dots, mode 6.8 f. See text for meaning of labels ‘‘Transition 1 and 2’’. Depths of samples are in centimeters.

the particles in suspension. The thickness of the current at time t can be approximated by ht 5 0.68(qo 2 /go )(Cv /Cvo )1/5 t If the two-dimensional model for channelized deposit-forming surges is applied to the large turbidites H2, H3, H8, and H13 of the HAP, and if the plain is assumed to be long and parallel-sided, the run-out distances greatly exceed the 356 km length of the plain. This implies either ponding of the turbidites at the western end of the plain, or escape of the flow over the sill at the end of the basin, or some incorrect assumption. A current containing the amount of material in a turbidite 3 m thick at an initial volume concentration of 4% (about 100 kg m23) and a flow density of 1114 kg m23 in sea water of 1050 kg m23 would have an initial volume of 8 3 106 m3 per meter width and with initial thickness of 500 m (the relief of lower Saõ Vicente Canyon) would be ø 16 km long. (However, if the zone of generation in the lower canyon were one-fifth the width of the abyssal plain, the actual length would be nearer 90 km). With sediment settling velocity of 5 3 1024 m s21 (30 mm silt), Dade and Huppert’s (1995) relationship suggests a runout distance of 785 km in over 122 3 103 s (U ø 6.5 m s21) and characteristic thickness of 0.82 m. This is considerably longer and thinner than the plain and initially assumed thickness, respectively. Thickness data show that the bigger turbidites thin towards the end of the basin (Figs. 5, 6) and so are not ponded. The sill to be surmounted is nearly 720 m high, which makes it most unlikely that flow about 50 m thick could escape. At high volume concentration (Cvo 5 0.04) the thin turbidity currents have a predicted runout distance less than the length of the plain. For a thin turbidity current (0.3 m, one-tenth of the previous example) with Cvo 5 0.01, that is about 25 kg m23 , not very low, qo is 3.2 3 106 m3/m width, g9o is 0.15, and the runout distance is 343 km in more than 112 3 106 s (U ø 3 m s21). A concentration near the lower fluid mud limit of 10 kg m23, Fo 5 0.04, yields a runout distance of 495 km in more than 3.33 3 103 s (U ø 1.5 m s21). All of these cases yield turbidity currents

a few tens of meters thick (Dade and Huppert 1995, eq 5a) flowing supercritically in the vicinity of the narrows. We think it most likely that the assumption of unimpeded flow down a parallel-sided plain is significantly in error. The cross section narrows in the middle of the plain, where a sharp bend is also encountered (Fig. 12). These two features provide increased flow resistance and slow the flow down. Turbidity-current flow is similar to open-channel flow, being described by a Che´zy-like equation with reduced gravity and friction on upper and lower boundaries (Middleton 1967). The internal distortion resistance for channelized flows of width-to-curvature ratio of one is greater than the flat-bed resistance by a factor of two (Leopold et al. 1960). The hydraulics of flow through a constriction are governed by dh 1 h 52 2 db (1 2 g9h/U ) b where h is flow depth and b breadth, g9 is reduced gravity (g9 5 gDr/r), where g is acceleration due to gravity and Dr is the flow density contrast over ambient fluid density r, and U is mean flow speed. For the conditions described, the incoming flow is supercritical (g9h/U2 , 1), dh/db is negative, and h increases going into the throat of the constriction. Because dh/ db 5 0 in the throat, g9h/U2 decreases to unity there, i.e., the flow just becomes critical. This involves a decrease in flow speed as well as an increase in flow depth. This situation is too complex to be amenable to simple sediment transport modeling, but the tendency with a constriction, in even a straight flow, is for the deposition rate to be greater than without the constriction. With the sharp bend in the flow this tendency is magnified. We suggest that this is the primary control on the thickness pattern seen along the plain (Figs. 5, 6). There remains the possibility of reflection from the end of the basin, suggested by upward coarsening to a maximum several centimeters above the base of H8 (Fig. 8), and the visible silt layers on X-radiographs (Fig. 10). These are quite close to the base, however, and do not fit with the notion of reflection of a substantially depositing flow from the far end of

866


FIG. 10.—X-radiographs of turbidite H8 across the HAP, showing ‘‘double coarse bases’’ of sand (lower) and silt (upper) due to reflection of flow on the sides of the plain. Magnifications indicated by A, B, C, and D.

the basin, because one would expect that by the time the flow arrived back at D11947 and D11946 to the east of the narrows, more than 10% of the thickness of the unit would have been emplaced. If these units do indicate reflection it is more likely to come from slight excursions of parts of the current up the side of the basin and flow back into the main flow path, rather than reflection from the far end. Estimation of Erosion and Sedimentation Rates Abyssal plains are dominantly depositional domains, with localized erosion as a secondary process. We can analyze the erosion of individual pelagic layers due to turbidity currents by examining the lateral variation in thickness of correlated pelagic units (Table 6). Pelagic layers of the order of 0.08–0.24 m and 0.07–0.17 m thick (P2 and P8, respectively) survived. Deposition and complete erosion of a pelagic unit over the plain may, however, be detected by the preservation of bioturbated pelagic sediment penetrating the top of a turbidite, as is the case for P3 and P10, on the top of turbidites 3 and 10, respectively. There is no means of estimating the

volume of eroded material in this case. In some other cases evidence for partial or local erosion exists, by absence of a particular pelagic interval in a few of the cores, as occurs for instance with P5 and P7. However, the data do not suggest that large flows have higher erosive capacity than small flows (observe for instance P9 and P6, both potentially partially eroded by turbidity currents emplacing turbidites H8 and H5, a large and a small turbidite, respectively). The data also suggest patchy erosion at any location in the plain. Evaluation of erosion rates is not a simple task, because the possible change of volume of units by compression in coring and differential compaction are difficult to assess, but the evidence suggests that turbidity currents partially eroded some of the pelagic layers deposited in the HAP by up to 0.16 m depending on cases and locations (Table 6). If we assume that the pelagic sedimentation rate of 3.7 cm/kyr at Tore Seamount, located above the CCD, applies to HAP for the stages 2/1 boundary (12,050 yr) to the present, an erosion-free accumulation of #0.45 m of pelagic sediment would be expected. More dissolution is expected at the HAP, located


867

FIG. 11.—Number and frequency of turbidites (N/ka) emplaced between pelagic beds in the Horseshoe Abyssal Plain during interglacial and glacial periods. The biostratigraphy is determined from planktic foraminifera, synthesized into climatic assemblages (Be´ 1977) and correlated with oxygen isotope stages. Pn, Pelagic n; H, turbidite; HL, Heinrich layer. Depth is in centimeters.

well below the Tore Seamount. The total thickness measured on the lithologic record from P8 to P1 is, however, 0.575 m (Table 5), so that approach is unprofitable. It is possible that the plain receives a little more pelagic material via resuspension from the sides, i.e., a slight pelagic focusing. We can therefore conclude from the data available from the HAP only that up to 0.16 m of erosion has occurred in patches. In cases such as the Madeira Abyssal Plain, Weaver and Kuijpers (1983) found no significant erosion due to turbidity currents. However, oxygen isotope stages 2 and 8, which are absent in their Core 82PCS34, were suggested to have been removed by intensification of bottom currents adjacent to the hills, during glacial periods. A synthesis of 13 abyssal plains in the western North Atlantic, Caribbean Sea, and Mediterranean Sea by Pilkey (1987) reports virtual absence of erosion in the majority of cases, and estimates a maximum of 0.20 m of sediment column eroded and usuTABLE 5.—Comparison of thicknesses of pelagic units for particular time intervals and sedimentation rates for the Tore Seamount (D11957P) and the HAP. D11957P Time Control Points

Age* (ky)

Boundary stages 2/1 HL 1 HL 2 Boundary stages 5/4 Boundary stages 6/5

12.05 14.3 21 73.91 129.84

Total Sediment Thickness Rate (cm) (cm/ky) 45 4 34 142 170

3.7 1.8 5.1 2.7 3

D11948P144P

D11950P

Total Sediment Total Sediment Thickness Rate Thickness Rate (cm) (cm/ky) (cm) (cm/ky) 57.5 7 134

4.7 3.1

60 17 49

5 7.5

* Ages for oxygen isotope stage boundaries are after Martinson et al. (1987) and ages for Heinrich Layers (HL) are from Bond et al. (1992).

ally less than 0.10 m in three exceptional examples from the Hispaniola– Caicos Abyssal Plain. Our results are comparable. Sedimentation rates are based on time control points (oxygen isotope stage boundary 2/1 and Heinrich layer 1) and cumulative thickness of pelagic units (Table 5). These are minimum estimates, because no correction has been made for eroded material. Nevertheless, values of 4.7 and 3.1 cm/kyr estimated in Core D11948P and D11944P for the Holocene and for the stage boundary 2/1 to HL1, respectively, are fairly consistent with values of 3.7 and 1.8 cm/kyr from flanks of the Tore Seamount, where no turbidite erosion occurred but from which some loss by resuspension can be expected. Frequency and Magnitude of Flows The age model provides the time control points yielding the necessary framework for a high-resolution evaluation of the frequency of turbidite emplacement. Figure 11 illustrates the emplacement of turbidites between the intervening pelagic units, as well as the frequency of turbidites as number per thousand years. The frequency is consistently higher during glacial periods (27 turbidites/10,000 yr) than interglacials (10 turbidites/10,000 yr). This is also true in the case of the neighboring Tagus Abyssal Plain to the north (Lebreiro 1995). Higher glacial frequency also coincides with higher mass flux during the last glacial (Table 7). A second aspect concerns the moment when turbidites are emplaced. In the HAP, a very constant rhythm occurs with alternation of one or a very few turbidites between every pelagic unit (Fig. 11). Weaver and Kuijpers

868

S.M. LEBREIRO ET AL. TABLE 6.—Analysis of erosion in the HAP, based on the lateral variation of thickness (in centimeters) of pelagic units.

Pelagic Units

D11944P — — — — — — — 8 (29) 5 (25)

P1 P2 P3 P4 P5 P6 P7 P8 P9 P10 P11 P11a P12 P12a P13 Pc P14 P15 P16 P17 P18 P20 P21

11

D11942P

D11941P

6 24 (0) drill gap drill gap drill gap 6 (28) 11 (0) 12 (25) 8 (22) eroded

4 (211) 4 (210) 2 (29) 11 (28) 5 (disturbed) eroded 5 11

7 — — — — — — — —

eroded 5 eroded 3, P16?

12 9 20 9 7 6 3

D11947P

— 11 (213)

— 8 (216) eroded 8 (27) 2 11 (23) 1 (210) 17 (0) disturbed, eroded?

D11946P — 12 (212) eroded (2 pelagics of 4 cm and 2 eroded, indistinguishable P4, P5, P6 or P7) 7 (210) 6 (24) eroded

D11948P

— 20 (24)

15 (0) eroded 14 (0) eroded 10.5 (26.5) 8 (22)

3 (212) eroded 7 (27) 7 (24) 12 (25) 10 (very disturbed)

13 5

— — — — — — — —

eroded eroded? mixed up? — — — — — — —

D11949P

— 23 (21)

13 22 (include turbs?) — — — — — — — —

6 20 11 8 — — — — — — — —

Core D11950P is not considered because of unreliability correlating the units. A few pelagic units are disturbed by drilling and cannot be considered for analysis. Blanks, no geological evidences in the record for pelagic units; 2, no recovery.

(1983) and Weaver et al. (1992) suggested a dependence of the scale and intensity of turbidity flow processes on the magnitude of sea-level rises or falls. In the particular case of the HAP, of the four largest flows, the second largest flow (H8) occurred during a major sea-level rise, at the transition to the Holocene. Because of lack of precise dating it is not known whether the thickest turbidite H13 was emplaced during glacial stage 3, stage 2, or at the transition 3/2, when a large sea-level fall with a magnitude of 70– 120 m occurred (Shackleton 1987). On the other hand, H2 and H3 occurred within the Holocene along with other smaller flows. As suggested above, sea-level fluctuations might be responsible for the emplacement of some large turbidites in this plain, though not all of them. Another triggering mechanism is consequently needed to explain the emplacement of both small flows and some of the large flows. The HAP is located close to the south border of the Eurasian tectonic plate, which provides an active seismic environment, producing a random emplacement of turbidites. In fact, the emplacement of the most recent but thin turbidite in the HAP has been dated by Thomson and Weaver (1994) to coincide with the Lisbon earthquake of 1755, which almost completely destroyed the city and affected all of the south of Iberia (Pereira de Sousa 1928). This confounds the recent fashion for calling large turbidites ‘‘seismites’’ (Mutti et al. 1994). Clearly small ones can be seismites too! CONCLUSIONS

Cores collected in the HAP reveal stratified sediment with alternating brown hemipelagite layers (of the order of centimeters) and olive-green thick mud units (of the order of decimeters, in some cases up to 5 m) TABLE 7.—Mass flux of turbidites for the Horseshoe Abyssal Plain.

Turbidites

Number of Turbidites

Dry Bulk Density (g/cm3)

Total Thickness (cm)

Mass Flux (g/cm2/ky)

H1 to H8

8

0.69

286

16

------------------------------------------------------------------------------------------------------(2/1)

H9 to H10a

3

0.79

64

22

------------------------------------------------------------------------------------------------------(HL1)

H11 to H16

10

0.79

591

70

------------------------------------------------------------------------------------------------------(HL2)

(2/1) 5 boundary stage 2/1 5 12.05 ky (after Martinson et al. 1987); (HL1) 5 Heinrich layer 1 5 14.3 ky; (HL2) 5 Heinrich layer 2 5 21.0 ky (ages after Bond et al. 1992). Mass flux as a function of time 5 (dry bulk density 3 total turbidite thickness)/ky.

deposited by turbidity currents. Most of the turbidites are , 1 m, i.e., of decimeter scale. The turbidites display extreme lateral continuity over the 21,157 km2 plain surface, irrespective of their thickness, enabling the correlation of all thin and thick individual beds. For the total time record, glacial oxygen isotope stage 2–3 to the present, only one 0.12 m thin, shallow-water calcareous turbidite (Hc), originating from the nearby Gorringe Bank, was emplaced. Thus, in practical terms, sediments are derived from a unique canyon source. The similarity of the mineral components found both in the Portuguese shelf and the plain proves the provenance of sediments from the Saõ Vicente canyon. The elongate shape of the plain and this primary feeding point significantly control the distribution of facies and sedimentation. The major fill of this enclosed basin consists of turbidites of visually homogeneous mud, but whose detailed grain-size analysis (interval 10–4 f) and X-radiographs reveal a more complex distribution. A clear 5–6 f mode appears for the basal sand-silt layer, and several modes within the interval 6.5–8.5 f for the overlying mud layer. Although the thickness of the slightly sandy bases of the turbidites is partly conditioned by the presence of a constriction 11 km wide in the middle of the plain (average width of the plain is 64 km), its grain size always decreases with distance from the source and sand-size material often reaches the far end of the plain. Application of Dade and Huppert’s (1994, 1995) model for channelized deep-water suspension gravity flows to our turbidite H9, of 0.3 m thickness, gives a velocity U ø 1.5–3 m s21 and a run-out distance of 343–495 km, which approximates the total length of the HAP. Flows depositing turbidites much thicker than H9, for example a flow containing the material of a deposit 3 m thick, have a run-out distance far larger than the 356 km length available in the plain, even for high concentration. The plain narrows and has a kink in the middle. We suggest that these two features are responsible for slowing down the flow, causing a thicker deposit in the central part of the plain than at either end. Some examples of inversely graded silt facies above the H2, H3, and H8 turbidite bases, shown by grain-size data and X-radiographs, may result from partial reflection of flows from the sides of the basin. Slight erosion is commonplace in the HAP, as proved by the absence of some pelagic units, though traces of their existence were preserved on the tops of underlying turbidites as burrow fills of pelagic material. The magnitude of erosion detected is never more than 16 cm and likely negligible


869

may be closely related to seismic activity in this area, which is tectonically situated near the compressional border of the Eurasian–African plate. However, climatic changes are responsible for increased frequency of flows at lower sea level. Support for the seismic trigger comes from the timing of the most recent turbidite, whose age is coincident with the famous 1755 Lisbon earthquake. This produced a very thin ‘‘seismite’’ or ‘‘seismoturbidite’’, which invalidates those terms for thick turbidites (Mutti et al. 1984). ACKNOWLEDGMENTS

The first author is grateful to all the staff and participants of cruise RRS Discovery-187, during which the cores used in this study were collected. This work was carried out during a Ph.D. research program and benefited from stimulating discussions with Dr. J.H. Monteiro. Dr. W.B. Dade assisted significantly in our evaluation of the consequence of flow constriction for the deposit thickness. Reviews from A.R. Prave and an anonymous referee triggered significant improvements in the manuscript. J.C. Moreno contributed with his experience of foraminiferal stratigraphy on the Portuguese margin. Laboratory technical support is acknowledged to A. Ineˆs, R. Lourenço, and G. Foreman. Financial support was provided by Junta Nacional de Investigaçaõ Cientı´fica e Tecnolo´gica–Programa Cieˆncia-BD/1215/91-IG and the European Commission on Marine Science and Technology (MAST II) STEAM programme (Sediment Transport on European Atlantic Margins) grant MAS2-CT94-0083. This is Dept. of Earth Sciences, University of Cambridge, contribution number 4841. REFERENCES

FIG. 12.—(upper) Detailed bathymetric map of the constriction and kink in the middle of the abyssal plain. (lower) Cross sections showing the degree of narrowing at the throat of the constriction.

in the general volumetric context, as has also been hypothesized for other modern abyssal plains. Ages inferred from foraminiferal bioclimatic assemblages and Heinrich layers 1 and 2 (?) of the thin intervening pelagic intervals dated the recovered lithological column back to glacial stage 2–3. The succession of interglacial/glacial periods is tied into the dated isotope stratigraphy. During the last glacial, icebergs originating from the Labrador Sea and Greenland cooled the surface waters in this area off the Portuguese margin, as demonstrated by the record of Heinrich layers in the HAP, as well as in the nearby Tore Seamount. Sedimentation rates could be estimated precisely only for the Holocene. The overall sedimentation rate, i.e., including pelagics and turbidites, is 20 cm/kyr. If only pelagic sedimentation is considered, the average values decrease to 5 cm/kyr. The frequency of emplacement of turbidites can be estimated on the basis of the above age model. Frequency of turbidites was higher during the low stand than the Holocene high stand of sealevel. The Holocene shows a frequency of 1.0 turbidites/kyr. For the glacial stage 3–2, the frequency is 2.7 turbidites/kyr. Both small and giant flows are associated with glacial and interglacial periods and transitions. The triggering of turbidity currents

ALVEIRINHO DIAS, J., AND NITTROUER, C.A., 1984, Continental shelf sediments of northern Portugal: Continental Shelf Research, v. 3, p. 147–165. BE´, A.W.H., 1977, An ecological, zoogeographic and taxonomic review of recent planktonic foraminifera, in Ramsay, A.T.S., ed., Oceanic Micropaleontology, v. 1, p. 1–100. BOND, G., HEINRICH, H., BROECKER, W., LABEYRIE, L., MCMANUS, J., ANDREWS, J., HUON, S., JANTSCHIK, R., CLASEN, S., SIMET, C., TEDESCO, K., KLAS, M., BONANI, G., AND IVY, S., 1992, Evidence for massive discharges of icebergs into the North Atlantic ocean during the last glacial period: Nature, v. 360, p. 245–249. COAKLEY, J.P., AND SYVITSKI, J.P.M., 1991, Sedigraph technique, in Syvitski, J.P.M., ed., Principles, Methods and Application of Particle Size Analysis: New York, Cambridge University Press, p. 129–142. COLLEY, S., THOMSON, J., WILSON, T.R.S., AND HIGGS, N.C., 1984, Post-depositional migration of elements during diagenesis in brown clay and turbidite sequences in the North East Atlantic: Geochimica et Cosmochimica Acta, v. 48, p. 1223–1235. COLLEY, S., THOMSON, J., AND TOOLE, J., 1989, Uranium relocation and derivation of quasiisochrons for a turbidite/pelagic sequence in the northeast Atlantic: Geochimica et Cosmochimica Acta, v. 53, p. 1223–1234. DADE, B., AND HUPPERT, H.E., 1994, Predicting the geometry of channelized deep-sea turbidites: Geology, v. 22, p. 645–648. DADE, B., AND HUPPERT, H.E., 1995, A box model for non-entraining, suspension-driven gravity surges on horizontal surfaces: Sedimentology, v. 42, p. 453–471. ELMORE, R.D., PILKEY, O.H., CLEARY, W.J., AND CURRAN, H.A., 1979, Black Shell Turbidite, Hatteras Abyssal Plain, Western North Atlantic Ocean Basin: Geological Society of America, Bulletin, v. 90, p. 1165–1176. HEEZEN, B.C., AND EWING, M., 1952, Turbidity currents and submarine slumps and the 1929 Grand Banks earthquake: American Journal of Science, v. 250, p. 849–873. HOYT, W.H., 1976, Long-distance turbidite correlations in the Horseshoe abyssal plain [unpublished MS thesis]: State University of New York at Albany, 136 p. HOYT, W.H., AND FOX, P.J., 1977, Long-distance turbidite correlations in the Horseshoe abyssal plain (abstract): American Association of Petroleum Geologists, Bulletin, v. 61, p. 797. HULSEMANN, J., 1966, On the routine analysis of carbonates in unconsolidated sediments: Journal of Sedimentary Petrology, v. 36, p. 622–625. JONES, K.P.N., MCCAVE, I.N., AND PATEL, P.D., 1988, A computer-interfaced SediGraph for the modal size analysis of fine-grained sediment: Sedimentology, v. 35, p. 163–172. JONES, K.P.N., MCCAVE, I.N., AND WEAVER, P.P.E., 1992, Textural and dispersal patterns of thick mud turbidites from the Madeira Abyssal Plain: Marine Geology, v. 107, p. 149–173. KNELLER, B., 1991, Oblique reflection of turbidity currents: Geology, v. 14, p. 250–252. LEBREIRO, S.M., 1995, Sedimentation history off Iberia: Tore seamount, Tagus and Horseshoe abyssal plains [unpublished Ph.D. thesis]: University of Cambridge, 192 p. LEBREIRO, S.M., MORENO, J.C., MCCAVE, I.N., AND WEAVER, P.P.E., 1996, Evidence for Heinrich layers off Portugal (Tore Seamount: 398N, 128W): Marine Geology, v. 131, p. 47–56. LEOPOLD, L.B., BAGNOLD, R.A., WOLMAN, M.G., AND BRUSH, L.M., 1960, Flow resistance in sinuous or irregular channels: U.S. Geological Survey, Professional Paper 282-D, p. 107– 134. MARTINSON, D.G., PISIAS, N.G., HAYES, J.D., IMBRIE, J., MOORE, T.C., AND SHACKLETON, N.J., 1987, Age dating and the orbital theory of the ice ages: development of a high-resolution 0 to 300,000 year chronostratigraphy: Quaternary Research, v. 27, p. 1–29. MCCAVE, I.N., AND JONES, K.P.N., 1988, Deposition of ungraded muds from high-density nonturbulent turbidity currents: Nature, v. 333, p. 250–252.

870


MCCAVE, I.N., MANIGHETTI, B., AND ROBINSON, S.G., 1995, Sortable silt and fine sediment sizecomposition slicing: parameters for palaeocurrent speed and palaeoceanography: Paleoceanography, v. 10, p. 593–610. MICROMERITICS, 1984, Instruction Manual, Sedigraph 5000 ET particle sizer. MIDDLETON, G.V., 1967, Experiments on density and turbidity currents: III. Deposition of sediment: Canadian Journal of Earth Sciences, v. 4, p. 475–505. MONTEIRO, J.H., ABRANTES, F.G., ALVEIRINHO-DIAS, J.M., AND GASPAR, L.C., 1983, Upwelling records in recent sediments from southern Portugal: a reconnaissance survey, in Thiede, J., and Suess, E., eds., Coastal Upwelling. Its Sediment Record. Part B: Sedimentary Records of Ancient Coastal Upwelling: New York, Plenum Press, p. 145–162. MUTTI, E., RICCI LUCCHI, F., SEGURET, M., AND ZANZICCHI, G., 1984, Seismoturbidites: a new group of resedimented deposits: Marine Geology, v. 55, p. 103–116. PEREIRA DE SOUSA, F.L., 1928, O terremoto do 18 de Novembro de 1755 em Portugal e um estudo demogra´fico, v. 3: Serviços Geolo´gicos de Portugal, Lisbon, p. 516–160. PILKEY, O.H., LOCKER, S.D., AND CLEARY, W.J., 1980, Comparison of sand-layer geometry on flat floors of 10 modern depositional basins: American Association of Petroleum Geologists, Bulletin, v. 64, p. 841–856. PILKEY, O.H., 1987, Sedimentology of basin plains, in Weaver, P.P.E., and Thomson, J., eds., Geology and Geochemistry of the Abyssal Plains: Geological Society of London, Special Publication 31, p. 1–12. PINHEIRO, L.M., WHITMARSH, R.B., AND MILES, P.R., 1992, The ocean–continent boundary off the western margin of Iberia. Part II: crustal structure in the Tagus Abyssal Plain: Geophysical Journal International, v. 109, p. 106–124. PINHEIRO, L.M., WILSON, R.C.L., REIS, R.P., WHITMARSH, R.B., AND RIBEIRO, A., 1996, The Western Iberia Margin: a geophysical and geological overview, in Whitmarsh, R.B., Sawer, D.S., Klaus, A., and Masson, D.G., Proceedings of Ocean Drilling Program, Scientific Results, vol. 149, p. 3–23. PIPER, D.J.W., 1978, Turbidite muds and silts on deep-sea fans and abyssal plains, in Stanley, D.J., and Kelling, G., eds., Sedimentation in Submarine Canyons, Fans, and Trenches: Stroudsburg, Pennsylvania, Dowden, Hutchinson & Ross, p. 163–176. ROBINSON, S.G., 1986, The late Pleistocene palaeoclimatic record of North Atlantic deep-sea sediments revealed by mineral-magnetic measurement: Physics of Earth and Planetary Interiors, v. 42, p. 22–47. ROESER, H.A., BARGELOH, O., EILERS, G., FRITSCH, J., KEPPLER, H., KEWITSCH, P., KLEIN, A., PINHEIRO, L.M., AND SCHRECKENBERGER, B., 1992, Sonne cruise SO-75: geophysical investigations of the crustal structure of the north Atlantic off Portugal: Bundesanstalt fu¨r Geowissenschaften und Rohstoffe, Cruise Report 109.997, Hannover.

SHACKLETON, N.J., 1987, Oxygen isotopes, ice volume and sea level: Quaternary Science Reviews, v. 6, p. 183–190. STOW, D.A.V., 1979, Distinguishing between fine-grained turbidites and contourites on the Nova Scotian deep water margin: Sedimentology, v. 26, p. 371–387. STOW, D.A.V., AND BOWEN, A.J., 1980, A physical model for the transport and sorting of finegrained sediments by turbidity currents: Sedimentology, v. 27, p. 31–46. STOW, D.A.V., AND PIPER, D.J.W., 1984, Deep-water fine-grained sediment: facies models, in Stow, D.A.V., and Piper, D.J.W., eds., Fine-Grained Sediments: Deep-Water Processes and Facies: Geological Society of London, Special Publication 15, p. 611–646. STOW, D.A.V., AND SHANMUGAM, G., 1980, Sequence of structures in fine-grained turbidites: comparison of recent deep-sea and ancient flysch sediments: Sedimentary Geology, v. 25, p. 23–42. TEIXEIRA, C., 1972, Carta Geolo´gica de Portugal (1/500 000): Serviços Geolo´gicos de Portugal, Lisbon. THOMSON, J., HIGGS, N.C., AND COLLEY, S., 1989, A geochemical investigation of reduction haloes developed under turbidites in brown clay: Marine Geology, v. 89, p. 315–330. THOMSON, J., AND WEAVER, P.P.E., 1994, An AMS radiocarbon method to determine the emplacement time of recent deep-sea turbidites: Sedimentary Geology, v. 89, p. 1–7. THOMSON, J., HIGGS, N.C., CROUDACE, I.W., COLLEY, S., AND HYDES, D.J., 1993, Redox zonation of elements at an oxic/post-oxic boundary in deep-sea sediments: Geochimica et Cosmochimica Acta, v. 57, p. 597–595. WEAVER, P.P.E., AND KUIJPERS, A., 1983, Climatic control of turbidite deposition on the Madeira Abyssal Plain: Nature, v. 306, p. 360–363. WEAVER, P.P.E., AND ROTHWELL, R.G., 1987, Sedimentation on the Madeira Abyssal Plain over the last 300 000 years, in Weaver, P.P.E., and Thomson, J., eds., Geology and Geochemistry of the Abyssal Plains: Geological Society of London, Special Publication 31, p. 71–86. WEAVER, P.P.E., ROTHWELL, R.G., EBBING, J., GUNN, D.E., AND HUNTER, P.M., 1992, Correlation, frequency of emplacement and source directions of megaturbidites on the Madeira Abyssal Plain: Marine Geology, v. 109, p. 1–20. WHITMARSH, R.B., AND MILES, P.R., 1995, Models of the development of the West Iberia rifted continental margin at 408309N deduced from surface and deep-tow magnetic anomalies: Journal of Geophysical Research, v. 100, p. 3789–3806. WILSON, T.R.S., THOMSON, J., COLLEY, S., HYDES, D.J., AND HIGGS, N.C., 1985, Early organic diagenesis: the significance of progressive subsurface oxidation fronts in pelagic sediments: Geochimica et Cosmochimica Acta, v. 49, p. 811–822. WILSON, T.R.S., THOMSON, J., HYDES, D.J., COLLEY, S., CULKIN, F., AND SORENSEN, J., 1986, Oxidation fronts in pelagic sediments: diagenetic formation of metal-rich layers: Science, v. 232, p. 972–875. Received 13 October 1995; accepted 7 November 1996.