the saharan debris flow: an insight into the mechanics of long runout submarine debris flows

The Saharan debris ¯ow: an insight into the mechanicsof long runout submarine debris ¯ows

M. J. R. GEE*, D. G. MASSON , A. B. WATTS* and P. A. ALLENà*Department of Earth Sciences, Parks Road, Oxford OX1 3PR, UK (E-mail: [email protected]) Southampton Oceanography Centre, Empress Dock, Southampton SO14 3ZH, UKàDepartment of Geology, Trinity College, Dublin 4, Ireland

ABSTRACT

New 3á5 kHz pro®les and a series of piston cores from the north-west African

margin provide evidence that the Saharan debris ¯ow travelled for more than

400 km on a highly ¯uid, low-friction layer of poorly sorted sediment. Data

suggest that the Saharan debris ¯ow is a two-phase event, consisting of a basal,

volcaniclastic debris ¯ow phase overlain by a pelagic debris ¯ow phase. Both

phases were emplaced on the lower continental rise by a single large debris

¯ow at around 60 ka. The volcaniclastic ¯ow left a thin deposit less than 5 m

thick. This contrasts with the much thicker (over 25 m) deposit left by the

pelagic debris ¯ow phase. We suggest that pelagic sediment, sourced and

mobilized as debris ¯ow from the African continental margin, loaded and

destabilized volcaniclastic material in the vicinity of the western Canaries.

When subjected to this loading, the volcaniclastic material appears to have

formed a highly ¯uid sandy debris ¯ow, capable of transporting with it the

huge volumes of pelagic debris, and contributing to a runout distance

extending over 400 km downslope of the Canary Islands on slopes that

decrease to as little as 0á05°. It is likely that the pelagic debris formed a thick

impermeable slab above the volcanic debris, thus maintaining high pore

pressures generated by loading and giving rise to low apparent friction

conditions. The distribution of the two debris phases indicates that the

volcaniclastic debris ¯ow stopped within a few tens of kilometres after

escaping from beneath the pelagic debris ¯ow, probably because of dissipation

of excess pore pressure when the seal of pelagic material was removed.

Keywords Long runout, low friction, mass wasting, two-phase debris ¯ow.

INTRODUCTION

Catastrophic mass movement events were initial-ly thought to be rare in the ocean basins except onvery steep slopes, deltas or in active seismic areas(Moore, 1961). It is now recognized that giantavalanches and debris ¯ows involving hundredsto thousands of cubic kilometres of material arerelatively common (Hampton et al., 1996). Theoccurrence of giant submarine debris ¯ows onpassive margins is of interest because, in terms ofemplacement mechanics, they are a particularlyenigmatic phenomenon. Numerous examples ofgiant debris ¯ows with extraordinarily long

runouts over slopes that decrease to virtually 0°are now recognized (Embley, 1976, 1980; Jacobi,1976; Bugge et al., 1988; O'Leary, 1993; Masson,1996). This paper focuses on the giant Saharandebris ¯ow on the north-west African margin(Embley, 1976) and proposes a new hypothesisfor its observed long runout. Central to thehypothesis is a two-phase debris ¯ow modelinvolving a raft of relatively coherent materialcarried on a highly ¯uid, low-friction basal layer.An understanding of the distributions of the twolithologically distinctive debris ¯ow phases givesvaluable insights into the mechanics of the ¯ow.Comparison with three well-known subaerial

Sedimentology (1999) 46, 317±335

Ó 1999 International Association of Sedimentologists 317

landslides suggests that they may have hadsimilar emplacement mechanisms. Our new dataalso allow important conclusions to be drawnregarding the anatomy and depositional form ofdebris ¯ows, with implications for the recogni-tion of giant debris ¯ows in the rock record.

PREVIOUS WORK

High-frequency 3á5 kHz echo sounders have beenused for many years to image the super®cialsediments of the deep sea (e.g. Damuth, 1975;Damuth & Hayes, 1977; Jacobi & Hayes, 1992).Initial mapping of the Saharan debris ¯ow wascarried out by Embley (1975, 1976). He used3á5 kHz data and piston cores to identify a long,sur®cial debris ¯ow of calcareous sediment ex-tending 700 km from its source on the north-westAfrican margin (Fig. 1). Volume estimates for this¯ow range from 600 to 1100 km3 (Embley, 1976,1982). More recently, long-range, side-scan sonar(GLORIA) has been used to demonstrate theextent of the mass ¯ow (Simm & Kidd, 1983;Kidd et al., 1985, 1987; Masson et al., 1992).Further detailed regional mapping was carriedout by Jacobi & Hayes (1992) who recognized arange of acoustic facies associated with theSaharan debris ¯ow and reported the existenceof at least two debris ¯ows of calcareous pelagicsediment and numerous turbidite pathways.Strongly prolonged echoes (SPEs) have beenidenti®ed below the Saharan debris ¯ow on3á5 kHz records by Embley (1976) and Jacobi &Hayes (1992). SPEs have been reported frommany areas of the deep ocean and are typically

regarded as a characteristic of turbidity currentpathways (Embley, 1975; Jacobi & Hayes, 1992).

Two sediment sources, the Canary Islands andthe north-west African continental margin, havebeen recognized as the major contributors to mass¯ows on the north-west African margin. TheCanary Islands have been source to large volumesof volcaniclastic material transported into deeperwater by debris avalanche, debris ¯ow andturbidity current processes. Large debris ava-lanches have been reported from Tenerife (Watts& Masson, 1995) and El Hierro (Holcomb & Searle,1991; Masson, 1996), in total involving a fewthousand cubic kilometres of material. In somecases, these avalanches are associated with largeturbidity currents and debris ¯ows, which trans-port sediment across the continental rise and ontothe abyssal plain (Jacobi & Hayes, 1992; Masson,1996). Turbidity currents may deposit sand-richlayers on the lower continental slope and rise,giving horizons that are acoustically character-ized by highly re¯ective layers extending forhundreds of kilometres from island source toabyssal plain.

Slide scars covering 18 000 km2 of sea¯oorhave been documented between latitudes 24°and 26°N on the north-west African continentalmargin, in water depths of about 2000 m (Embley,1982; Jacobi & Hayes, 1992). Up to 1100 km3 ofcalcareous sediment are believed to have beenderived from these scars (Embley, 1982). As aresult of highly productive surface waters, pri-mary sedimentation in this area to the south andeast of the Canary Islands is mainly pelagic, withminimal input from terrigenous sources. Esti-mates for sedimentation rates during theHolocene range from 0á025 to 0á06 mm year)1

(Seibold et al., 1976) to around 0á07 mm year)1

(Schmincke et al., 1995).For this study, we used an extensive 3á5 kHz

database (Fig. 2) to remap the Saharan debris ¯owbased on a new process-oriented interpretativescheme that recognizes three main groups of eightfacies (Table 1). The scheme was designed to mapthe various complex echo types associated withthe Saharan debris ¯ow in detail. Each facies hasa characteristic re¯ection strength, sub-bottompenetration and geometry. The interpretation hasbeen calibrated with piston core data.

3 á5 kHZ OBSERVATIONS

Three main facies groups, each containing two orthree facies types, have been determined

Fig. 1. Location of the Saharan debris ¯ow on the NWAfrican continental margin to the south of the CanaryIslands. Contours are in kilometres below sea level.

318 M. J. R. Gee et al.

Ó 1999 International Association of Sedimentologists, Sedimentology, 46, 317±335

(Table 1). Facies group A is characterized byparallel re¯ectors and penetration of several tensof metres. Facies group B consists of acousticallytransparent layers, with lens-shaped or irregulargeometry. Facies group C consists of ¯at, strongprolonged echoes (SPEs) with little sub-bottompenetration.

Large areas of the continental rise west of theCanary Islands are characterized by parallel

re¯ectors and several tens of metres of penetra-tion on 3á5 kHz pro®les (facies group A). Re¯ec-tors are typically parallel and laterally persistent(facies A1, Fig. 3A and B). Areas characterized bymore intermittent and undulatory re¯ectors canalso be distinguished (facies A2).

3á5 kHz pro®les across the Saharan debris ¯oware characterized by acoustically transparentlayers, often with lens-shaped geometries. The

Fig. 2. Locations of 3á5 kHz pro®les (®ne continuous lines) and piston cores (triangles) used in this study. Numberedbold lines locate illustrated 3á5 kHz sections. Figure located on Fig. 1. Contour interval is 100 m.

Table 1. Classi®cation of 3.5 kHz echo types.

Echo type Description

Group A Strati®ed geometryA1 Continuous, sharp, parallel re¯ectors. Echoes are strong, relatively ¯at and laterally persistent

Penetration is 40 m or moreA2 Intermittent, continuous, parallel re¯ectors. Echoes are medium to strong

Penetration is up to 40 m but often less than A1A3 Subcontinuous, intermittent, subparallel re¯ectors

Sea¯oor is undulatory and has a strong echoPenetration is up to 30 m

Group B Lens geometryB1 Acoustically transparent lenses up to 25 m thick. Unit has a distinct wedging character and

positive relief. Sea¯oor echo is weak and fuzzy, often with a stronger basal echoB2 Acoustically transparent lenses with strong, often irregular sea¯oor and intermittent

sub-bottoms. Generally thinner than B1

Group C Strong echo, poor penetrationC1 Strong, often prolonged sea¯oor echo with no sub-bottom re¯ectors. The echo has a ¯at,

`fuzzy' signatureC2 Strong, very ¯at echo, often prolonged. Penetration is only a few metres with often

a few poorly resolved sub-bottomsC3 Very prolonged echoes in the top 20 m. Poor penetration

Saharan debris ¯ow 319


surface echo of these layers is typically weak andslightly prolonged (facies B1, Fig. 3A). The lack ofsubsurface re¯ections on the right of Fig. 3A canbe contrasted with A1 facies, where strati®ed

re¯ection events can be seen to a depth of over40 m. Mapping of the acoustically transparent B1facies reveals a lobate, 25-km-wide debris ¯ow(Embley, 1976; Jacobi & Hayes, 1992) sourced

Fig. 3. Examples of 3á5 kHz pro®les. See Fig. 2 for locations. (A) Lens-shaped geometry of debris ¯ow (B1 facies)overlying strati®ed slope sediments (A1 facies). (B) Typical strongly prolonged seabed echo (C1 facies) overlyingstrati®ed slope sediments (A1 facies). (C) Rough topography debris ¯ow (B2 facies) with strongly prolonged echoabove and below.



from around 2000 m water depth on the north-west African margin over 200 km south-east of ElHierro and terminating west of El Hierro at 27°50¢N, 21°46¢ W in about 4850 m of water (Figs 1 and3). The thickness of the ¯ow, as observed from3á5 kHz pro®les, varies from 5 to over 25 m (Fig. 4and Masson et al., 1993). The debris ¯ow has awell-de®ned main thalweg, along which debris isover 20 m thick, surrounded by thinner, sheet-like areas of debris between 10 and 15 m thick.The main thalweg of acoustically transparentdebris ¯ow always overlies a re¯ector that has astrongly prolonged echo. Between 20° 36¢ and 21°40¢ W, the base of the acoustically transparentlayer cannot always be resolved clearly by the3á5 kHz system, with the greatest depth observedat about 25 m (Fig. 4). West of 20° 36¢ W, the ¯owdeposit curves slightly south-westwards, where itterminates as a double-lobed snout around asmall bathymetric high (Fig. 5). East of 21°W,3á5 kHz pro®les show a series of B1 lensesonlapping each other, with strongly prolongedechoes beneath each lens (Fig. 6A). B1 facies alsodisplay a smoothing effect, where irregular, un-derlying topography is not re¯ected on the uppersurface of the debris ¯ow (Fig. 6B).

Strongly prolonged echoes (group C facies)surround, and can be traced beneath, large areascharacterized by an acoustically transparent sur-face layer, where penetration of the 3á5 kHzacoustic pulse allows. The commonest group C

facies type, C1, is characterized by a stronglyprolonged and very ¯at sea¯oor echo (Fig. 3B).This echo may consist of either a single ¯at andlaterally continuous re¯ector or a closely spacedpair of re¯ectors. Acoustically transparent B1facies overlying two strongly prolonged re¯ectionevents can be observed over much of the lowerpart of the Saharan debris ¯ow (Fig. 6C). Thecharacter of group C facies changes with slope,having greater relief and prolonged echo onsteeper slopes east of 20° 00¢ W (C3 facies) and amuch ¯atter, less prolonged echo downslope tothe west (C1 and C2 facies). In this latter region,C1 facies occurs in broad, shallow channels(Fig. 7A) spreading out and extending 73 kmfurther downslope than B1 facies. Between someof these channel areas, intermittent strati®ed sub-bottoms are overlain by a very ¯at, laterallycontinuous sea¯oor re¯ector (Fig. 7B). This re-sults from a feather edge of C2 facies overlying anarea of A1 facies.

Several different onlapping relationships havebeen identi®ed between facies groups A, B and C.To the east of 20° 40¢ W along the southern edgeof the B1 facies zone, B1 facies directly onlap A1facies. In the northern and western region, B1facies onlap group C facies which, in turn, onlapgroup A facies. Where B1 facies onlap C or Afacies, the relationship is easily recognizable andmappable (Fig. 3). Between 20° 40¢ W and 21° 30¢W, onlapping relationships are complex, with A,

Fig. 4. Map of Saharan debris ¯ow west of Canaries based on new 3á5 kHz and core data. Depth contouring of thepelagic debris ¯ow phase was possible on account of its acoustic transparency on 3á5 kHz records (B1 facies). Forlocation, see Fig. 1. Bathymetric contour interval is 100 m.



C and B facies being superimposed within a smallarea. East of 21° 00¢ W, on the northern edge of theB1 facies zone, there are positive relief featuresthat have the geometry of B1 facies and the strongprolonged echoes typical of group C facies (faciesB2, Fig. 3C). C facies always onlap A facies, and Bfacies always onlap C or A facies (Fig. 7C).

CORE DATA

Saharan debris ¯ow cores

Cores recovered from the lower part of theSaharan debris ¯ow and core D24 recovered300 km upslope contain correlatable sand-richmass ¯ow units and ®ne-grained pelagic clays,marls and oozes (see Figs 2 and 8). Some sand-rich units are well sorted, graded and haveinternal structures such as planar laminations,wavy laminations and cross-bedding. Othersand-rich units appear largely structureless,consisting of poorly sorted, angular volcaniclas-tic sand set in a ®ne-grained nannofossil clay-

grade matrix. This latter type occurs in a layerabout 1á5 m thick in cores D16 and V32 (volca-niclastic debris ¯ow phase in Fig. 8) which, onthe basis of grain size (Fig. 10) and lithology,appears to correlate with a similar 2á5-m-thicklayer in D24, 300 km upslope. Several cores,however, failed to penetrate through this layerand, therefore, its maximum thickness has notbeen determined.

The volcaniclastic sand and clay units aredistinguished from the well-sorted, graded sandsby their poor sorting, homogeneous texture andlateral extent. In cores V32 and D19, fragile clastsof marl, ranging in thickness from a few centime-tres to over a metre, occur within the poorlysorted sand and clay units. Core V32 contains twoclasts about 50 cm thick, which contain upside-down turbidites interbedded with pelagic marl.These clasts show some signs of disaggregationand, when inverted, match the sequence imme-diately below the base of the poorly sorted muddysand unit.

Cores D18, D21 and D22 show over 8 m offolded white ooze, beige marl and minor

Fig. 5. Detailed map of the debris ¯ow snout showing relationships between the terminations of the pelagic andvolcaniclastic debris ¯ow phases and their interactions with the sea¯oor topography. Locations of cores on transectacross the debris ¯ow snout are indicated. Contour interval is 50 m.



volcaniclastic sand layers. Folded sediments withsimilar geometries and lithologies were alsofound in core D20 at a depth of 1á5±3 m overlyinga poorly sorted, volcaniclastic sand and clay

layer. This sediment is distinctly different interms of colour and geometry from the brown,unfolded hemipelagic marls and clays of thesurrounding lower rise.

Fig. 6. Examples of 3á5 kHz pro®les. See Fig. 2 for locations. (A) Superimposed onlapping debris ¯ow lenses (B1facies) with strongly prolonged basal re¯ectors. (B) Irregular (erosional?) sea¯oor topography smoothed by in®lling bypelagic debris ¯ow (facies B1). (C) Debris ¯ow unit (B1 facies) underlain by paired strongly prolonged echo re¯ectors.



Stratigraphy of debris ¯ow cores

Coccolith assemblages from the debris ¯ow layersand in situ sediments were analysed to establishthe date of debris ¯ow emplacement and the agesof sediments involved in the ¯ow, using themethods of Weaver & Kuijpers (1983) and Weaver

(1994). In the upper sections of several cores, coredisturbance and incomplete recovery is apparent.In the few cores that appear to have sampled acomplete in situ succession above the debris ¯ow(D16, D19, D20), all of the sediments recoveredbelong to oxygen isotope stages 1±3 (i.e. sedi-ments aged from 0 to 50 ka). Although these

Fig. 7. Examples of 3á5 kHz pro®les. See Fig. 2 for locations. (A) Narrow tongue of volcaniclastic debris ¯ow (C1facies) overlying strati®ed sediments (A1 facies). The area of C1 facies is interpreted as a shallow debris ¯ow channel.(B) Thin volcaniclastic unit (C2 facies) overlying strati®ed sediments (A1 facies). (C) Critical relationships betweenfacies A, B and C.



stages cannot be separated on the basis ofcoccolith assemblages alone, the sequence oflithological units (from the top down) of marl,marly clay and marl can be used to infer thatstages 1±3 are all present in these cores (Fig. 9;see Weaver, 1994).

The pelagic debris ¯ow contains folded sedi-ments of isotope stage 5±8 age, although only theupper 8 m of the pelagic debris ¯ow weresampled. From the limited analysis carried outto date, all of the sampled material appears to bein approximate stratigraphic order (Fig. 9).

The volcaniclastic sand and clay units containa near constant coccolith assemblage, indicating amixture of sediments from stages 6/7±13. Withinthe volcaniclastic sand and clay unit, two inver-ted clasts sampled in core V32 and a single clastin core D19 are composed, respectively, of stage 4and 5 and stage 7 and 8 sediments.

Cores V32, D16 and D24 sampled in situsediments beneath the debris ¯ow. The youngestdatable in situ sediments are of oxygen isotope

stage 5 age (Fig. 9). However, in cores V32 andD16, the stage 5 sediments are overlain by a thinin situ pelagic clay layer, which can be inferred tobe of stage 4 age on lithological grounds (seeWeaver, 1994).

Grain size analysis

A quantitative analysis of the volcaniclastic sand-rich units was undertaken to determine grain sizetrends vertically and laterally within each depos-it. Samples were taken from the apparently poorlysorted sand/clay layer (volcaniclastic DFP inFig. 10) and from visibly graded volcaniclasticsand units (turbidite in Fig. 10) and processedwith a CILAS 960 granulometer. Normal frequen-cy curves for the total range (355±0á07 microns)and also for the coarser fraction (355±30 microns)were plotted. Other statistical parameters mea-sured were moment measures of mean andstandard deviation for the 355±30 micronwindow.

Fig. 8. Lithological summary of key piston cores that sampled debris ¯ow sediments. Cores are arranged from distal(D16, left) to more proximal (D24 right). Correlation between cores is based on both lithology and nannofossil stra-tigraphy. Heavy continuous correlation lines show the limits of debris ¯ow sediments. Cores are located on Fig. 2.



Grain size normal frequency plots are clearlybimodal in most samples of the volcaniclasticsand and clay layer, with a broad mode in the ®nesand±coarse silt range and a series of sharp peaksaround 2±3 microns (Fig. 10, cores D16, D19 andD24). The broad mode mostly consists of brokenforaminifera and volcaniclastic glass, while the®ne modes around 2±3 microns comprise mainlycoccoliths with minor quartz and other particles.Vertical trends within the volcaniclastic sand/clay layer (volcaniclastic DFP in Fig. 10) areminimal, with a relatively constant proportionof clay, silt and sand throughout each unit(Fig. 10). Vertical trends for moment measuresreveal a very constant mean grain size, standarddeviation (sorting) and skewness. In cores D17and D20, there is a minor decrease in mean grainsize and an increase in standard deviation at thetop of the volcaniclastic sand/clay unit.

Down¯ow measures of mean grain size andstandard deviation show minor trends. Betweencores D19 and D16, mean grain size decreases fromaround 113±50 microns, and standard deviation

(sorting) values decrease from around � 2á24 to� 0á98 (Fig. 11). The percentage of sediment over355 microns also decreases between D19 and D16from about 38% to 15%. In contrast, it can beobserved that the 10±2 micron fraction increasesfrom around 20% to 40%, whereas the clay fractionremains relatively constant at around 10%.

Comparison with a graded volcaniclastic sandstratigraphically above the unsorted sand/clayunit enables important contrasts to be identi®edin terms of grain size, grading and sorting. Thisgraded sand unit was present in cores D16, D20,D19 and probably D17, although the top 2 m fromcore D17 was not recovered because of coringdif®culties. Grain size analysis shows a well-sorted, grain-supported sediment with well-de-veloped vertical and lateral normal grading (seeFig. 10). The individual normal frequency plotsshow a single, well-developed sand±silt-sizedmode and relatively few ®nes. Vertical normalgrading was well developed in cores D20 andD19, with the mean grain size decreasing from200 microns to less than 30 microns.

Fig. 9. Summary of the stratigraphy of cores shown in Fig. 8. Subdivision of oxygen isotope stages 1±4 is based on acombination of nannofossil and lithological information (see text). Cores D16 and V32 show that emplacement of thedebris ¯ow occurred between stages 3 and 4.



INTERPRETATION

Sediments characterized by parallel re¯ectors andseveral tens of metres of penetration on 3á5 kHzpro®les (A facies) are interpreted as interbedded

pelagic and turbidite sequences (Damuth, 1975;Masson et al., 1992). Cores taken within thisfacies show mainly ®ne-grained sediments, al-though some turbidites have graded sandy bases.

Sediment layers, commonly with lens-like ge-ometries in cross-section and with acousticallytransparent 3á5 kHz signatures (B1 facies), havebeen calibrated using piston core data and resultfrom debris ¯ows composed of pelagic sedimentsourced from the north-west African margin(Embley, 1976; Jacobi & Hayes, 1992). The trans-parent acoustic signature results from the chaoticnon-strati®ed nature of the sediment, and theweak, slightly prolonged surface echo resultsfrom hummocky microtopography and a lack ofsand and silt on the surface of the deposit(Embley, 1975; Jacobi, 1976). Variable depositthickness, steep margins, leveÂes and generalpositive surface relief all indicate a material withstrength, which results in the preservation of ¯owstructures. Buried lenses indicate the passage ofmore than one debris ¯ow pulse, as has beenobserved in subaerial debris ¯ows (Johnson,1970). Facies B2 also has a lens-like acousticsignature and similar geometries to B1 facies, butwith a smoother, strongly prolonged echo on theupper surface. It is interpreted as consisting of thesame material as the B1 facies (see below) withthe major difference that a layer of sand-richmaterial has been deposited on the upper surface.

The white and pale beige, folded pelagicsediment of cores D18, D21 and D22 has clearly¯owed, with the preservation of deformed but

Fig. 10. Summary of grain sizeanalysis of the volcaniclasticdebris ¯ow phase and comparisonwith turbidite data. Relative grainsize percentages, mean grain sizeand the spectra (normal frequencydistributions) are illustrated. Themixed grain size and lack ofgrading in the volcaniclastic DFPare clearly shown. Comparison ofdata between proximal (D24) anddistal (D16) cores shows littleevolution of the unit betweenthese two sites, which are 300 kmapart. See Fig. 8 for core lithologydata and Fig. 2 for core locations.

Fig. 11. Variations in mean grain size and standarddeviation (sorting) for the volcaniclastic debris ¯owphase between core D16 (distal) and core D24 (proxi-mal).



original sediment layering (Masson et al., 1993).Coccolith assemblages show that the sedimentsremain in original stratigraphic order and, there-fore, appear to have been subjected to relativelyminor plastic deformation. We call this carbon-ate-rich sediment the pelagic debris ¯ow phase(pelagic DFP).

Facies characterized by strong prolonged ech-oes (group C facies) that surround the pelagicdebris ¯ow phase are interpreted as the acousticsignal of a volcaniclastic debris ¯ow phase(volcaniclastic DFP). The highly re¯ective andprolonged echo return is thought to indicate arelatively high sand content and a short-wave-length bottom roughness (Jacobi & Hayes, 1992).3á5 kHz and core data indicate that this ¯owmaterial directly underlies the pelagic DFP(Fig. 7C). The subsurface `double' re¯ection eventseen in Fig. 6C may be imaging the upper andlower surfaces of the volcaniclastic DFP. Corescontaining the volcaniclastic DFP are all recov-ered from areas where the sea¯oor echo is strongand prolonged, with the exception of D16, locatedat the margin of the mapped volcaniclastic DFP,which shows a slightly prolonged surface echoand some subsurface penetration.

Grain size analyses of the volcaniclastic DFPindicate many features characteristic of debris¯ow material, such as a bimodal size frequencydistribution, matrix support of clasts and over10% clay content (Coussot, 1992). The presenceof a clay matrix would increase the cohesiveproperties of the sediment and suppress turbu-lence during ¯ow, resulting in a poorly sortedcohesive material with a plastic rheology. Obser-vations of deposit geometries from group C faciescon®rm a lack of strength indicators such asprominent leveÂes and positive topography, indi-cating a highly ¯uid material. However, matrixsupport of the sand fraction indicates that thevolcaniclastic DFP did possess a ®nite but lowshear strength.

The highly ¯uid nature of the volcaniclasticDFP is evident from the complex geometry of thedeposit as a whole, but speci®cally in the distalregion beyond the snout of the pelagic DFP. Here,the variable SPE signatures show how a thin,acoustically re¯ective material has exploitedbroad, shallow channels and left a `wash-over'deposit on bathymetric highs. In channels, theacoustic signal is completely attenuated by per-haps as little as 1 or 2 ms of the volcaniclasticDFP, giving rise to SPEs with no sub-bottompenetration. Where the deposit thins over highs,gently undulating, strati®ed sea¯oor can be seen

below the much ¯atter, strongly prolonged echosignature of the volcaniclastic DFP. Although theevidence strongly suggests a cohesive materialwith a plastic rheology, minor grain settlingappears to have occurred between core D19 andD16, resulting in a slight increase in sorting and adecrease in mean grain size.

The mapped distribution of the volcaniclasticDFP shows that it occurs over a similar butslightly larger area than the pelagic DFP. Coredata and 3á5 kHz mapping indicate that thepelagic DFP almost always overlies the volcanic-lastic DFP. Coccolith assemblages from in situsediments directly beneath and above the volca-niclastic DFP and above the pelagic DFP indicatethat its emplacement occurred close to the oxygenisotope stage 3/4 boundary (around 60 ka; Fig. 8).The emplacement of the pelagic DFP is less wellconstrained, but evidence from cores D20 andD24 indicate that pelagic DFP rests directly uponvolcaniclastic DFP (Figs 7 and 8), although it isrecognized that the direct superposition ofpelagic and volcaniclastic DFP is the result ofdebris ¯ow processes and cannot be taken asabsolute proof that the two phases were emplacedat the same time. Nevertheless, the combinationof constraints that both phases (i) were emplacedat the isotope stage 3/4 boundary; (ii) have a cleardistributional relationship; and (iii) occur indirect contact, is good evidence that the twophases were emplaced during a single event. Inparticular, the occurrence of white nannofossilooze resting directly upon the volcaniclastic DFPin D20 (Fig. 8) provides strong evidence that thetwo debris ¯ow phases are part of a single ¯owevent. The nannofossil ooze is interpreted as aclast derived from the main body of the pelagicDFP which, according to 3á5 kHz data, terminatesin a prominent snout about 5 km upslope. Theclast appears to have been rafted out in front ofthe main body of the pelagic DFP, carried by thevolcaniclastic DFP.

The lithology of clasts obtained from the distalregion of the volcaniclastic DFP indicates deriva-tion from the slope west of the Canaries. In coreV32, coccolith assemblages and grading of tur-bidites in two clasts indicate that both have beeninverted. Both lithological and coccolith assem-blage data con®rm that the inverted clasts and thein situ sediment immediately beneath the volca-niclastic DFP contain the same sedimentarysuccession. This is good evidence for localerosion of the sea¯oor by the volcaniclastic DFP.The marl clast in D19 also appears to be locallyderived but is slightly older than the substrate on



which the debris ¯ow was emplaced, indicatingsubsea¯oor erosion to a depth of around 3 m.Within clasts transported by the debris ¯ow, sandlayers appear to show a greater degree of defor-mation and disaggregation than pelagic layers.This may indicate the instability of sand layerswithin the substrate when subjected to loadingand shearing.

The volcaniclastic graded sand unit (Fig. 10) isinterpreted as a turbidite, as it has features typicalof the deposit of a turbulent mixed-grain-sizesediment suspension, such as normal grading,sorting and cross-bedding. Down¯ow changes indeposit thickness and grain size for this turbiditecontrast with the near constant thickness andgrain size of the volcaniclastic debris ¯ow phase,indicating a more ¯uid ¯ow process dominatedby grain settling. Such a distinction furtheridenti®es the latter as a highly ¯uid, ®ne-graineddebris ¯ow phase.

DISCUSSION

Our investigation of 3á5 kHz and core data indi-cate that the Saharan Debris Flow is a two-phase¯ow event, with a basal volcaniclastic layer andan overlying pelagic layer. The upper pelagic DFPforms a relatively thick deposit with prominentleveÂes and steep margins, features common insubaerial debris ¯ows with signi®cant shearstrength (e.g. Johnson, 1970). Given the longrunout distance of 700 km, the limited degree ofinternal deformation apparent in cores from thepelagic DFP is also remarkable, suggesting arelatively little deformed sediment mass that hasbeen rafted to its current position. Observations ofthe internal deformation from cores can only bemade for the top 8 or 9 m. Therefore, it is likelythat more deformation occurred at lower levels inthe pelagic DFP. An emplacement mechanism inwhich a raft of semi-coherent, plastically deform-ing material was carried along by a basal layerundergoing laminar ¯ow (Bingham ¯ow mecha-nism; Johnson, 1970) is also supported by obser-vations of the surface morphology of the pelagicDFP (Masson et al., 1993).

In contrast to the pelagic DFP, the volcaniclasticDFP occurs as a thin uniform layer and lackssurface topography and steep margins. Its homo-geneous character, suggesting high degrees ofdeformation and mixing, also indicates a materialof low apparent shear strength. Thus, the basalvolcaniclastic DFP appears to be rheologically andlithologically distinct from the upper pelagic DFP.

Origin of volcaniclastic debris ¯ow phase

The origin of the volcaniclastic DFP is clearlyimportant in understanding the overall ¯owprocess. The volcaniclastic component of thedebris ¯ow must have been sourced from theproximity of the Canary Islands. Thus, the con-trast between the volcaniclastic DFP and thepredominant white oozes of the pelagic DFP(indicating derivation from upslope of the CanaryIslands where volcaniclastic material is absent)indicates that the volcaniclastic DFP must havebeen derived from the substrate over which thepelagic DFP travelled. Further evidence for ero-sion of the substrate comes from dating ofcoccolith assemblages within the volcaniclasticDFP. As discussed above, these assemblagesindicate a source older than the contemporarysea¯oor at the time of volcaniclastic DFP em-placement.

Remobilization of turbidites or debris ava-lanches derived from the Canary Islands is themost likely origin for the volcaniclastic DFP. Theuniformity of the volcaniclastic DFP over itsentire area of occurrence suggests derivationfrom a limited source area with rapid achieve-ment of its ®nal, poorly sorted character. Apossible alternative, that it was generatedcontinuously by progressive mobilization of thesubstrate over a wide area, would require aremarkable uniformity of both that substrate andthe mobilization and mixing process thatproduced the volcaniclastic DFP. One possibilityis that the volcaniclastic DFP is made up ofmaterial remobilized from a landslide on the¯ank of the western Canary Islands. Flankcollapse landslides reported from the CanaryIslands typically result in ®elds of blockyvolcaniclastic material and sand-rich turbidites(Watts & Masson, 1995; Masson, 1996). A likelycandidate landslide is the El Julan landslide onEl Hierro, part of which was over-run by theSaharan debris ¯ow (Holcomb & Searle, 1991).

The age range of the coccolith assemblages inthe volcaniclastic DFP indicates that it incorpo-rated sediments as old as 500 ka, indicatingremobilization of material to a considerable depthbelow the contemporary seabed. The 450 kacontinuous record of sedimentation in core D16on the continental rise corresponds to 8 m ofsection, which can be taken as a rough (probablyminimum) indication that the top 8 m of sedi-ment may have been removed from south-west ofthe island ¯anks to form the volcaniclastic DFP.Sedimentation rates closer to the islands are



higher because of upwelling (P. P. E. Weaver,personal communication 1997), indicating thatmore than 8 m of sediment was probably removedto generate the volcaniclastic DFP.

Characteristics of sandy debris ¯owsand the nature of the volcaniclastic DFP

Sandy debris ¯ows typically give rise to massivesand deposits with some mud matrix(Shanmugam, 1996). They may carry large, fragileclasts with a typical planar clast fabric andcommonly show clear evidence for laminar orplug ¯ow. A sandy debris ¯ow has a plasticrheology and represents a continuous spectrum ofprocesses between cohesive and cohesionlessdebris ¯ows (Shanmugam, 1996). Sandy sedimentcan move as debris ¯ow with as little as 2% (byweight) clay content if the sand is ®ne grained,and 19% (by weight) if the sand is coarse grained(Hampton, 1975). Coussot (1992) suggests that aclay content in the range 4±10% will cause apoorly sorted mixture to behave as a viscoplastic¯uid regardless of the maximum particle size.

Sediment samples from the volcaniclastic DFPhave been recovered only beyond the limits of thepelagic DFP; no samples have been obtained frombeneath the pelagic phase. All samples recoveredfrom the volcaniclastic DFP have a clay content ofbetween 10% and 20% and a poorly sorted coarsefraction ranging from silt to coarse sand size. Agenerally homogeneous composition, both later-ally and vertically, and abundant clasts of pelagicsediment are additional evidence for a sandydebris ¯ow process. Beyond the limits of thepelagic DFP, the volcaniclastic DFP would havebeen able to ¯ow in a free and uncon®nedmanner, giving rise to the observed typical sandydebris ¯ow deposit. In this region, the volcanic-lastic DFP appears to have behaved in a highly¯uid but plastic manner, exploiting channels and¯owing around bathymetric highs of only a fewmetres relief, especially in the distal region. Theobservation that the deposit thickens into bathy-metric lows and thins over highs may indicate ahigh-velocity, surge-like ¯ow, very similar to asubaerial pyroclastic surge, or may re¯ect theinterpreted low strength of the material as itdrained into bathymetric lows after the ¯ow hadstopped (Orton, 1996).

Although we have not been able to obtainvolcaniclastic DFP samples from immediatelybeneath the pelagic ¯ow, the similarity in com-position, grain size and sorting of the volcanic-

lastic material recovered from cores D19, takenadjacent to the pelagic ¯ow snout, and D24, takenadjacent to the pelagic ¯ow margin over 200 kmupslope, indicate little downslope change in thenature of the volcaniclastic ¯ow. This suggests auniformity of the volcaniclastic DFP over itsentire area of occurrence, including the areabeneath the pelagic DFP, with no evidence forevolution during ¯ow, even over long distances.Beneath the pelagic DFP, however, the volcanic-lastic DFP would have been subjected to highshearing stresses, and it is reasonable to suggestthat its rheological behaviour may have beensigni®cantly different from that of the uncon®nedvolcaniclastic ¯ow.

The volcaniclastic DFP as a low apparentfriction layer

Our hypothesis is that the Saharan debris ¯owconsisted of a relatively viscous mass of pelagicmaterial carried on a basal volcaniclastic layerwith low apparent friction characteristics. Themost likely explanation for the low apparentfriction behaviour of the volcaniclastic DFP isexcess pore pressure created by the loading andshearing action of the pelagic DFP. Excess porepressures are widely regarded as an importantparameter in submarine landslides (Morgenstern,1967; Hutchinson & Bhandari, 1971; Norem et al.,1990; Hampton et al., 1996). Although moststudies have concentrated on the role of excesspore pressure in the initial slope failure leading tolandsliding, Norem et al. (1990) argued that `thelong runout distance for submarine ¯owslides canbe explained only by high excess pore pressuresduring the whole slide event'. Recent experimen-tal debris ¯ow studies con®rm that elevated porepressures exist at the base of debris ¯ows over thewhole ¯ow event (Iverson, 1997).

Initial failure of the seabed in the source area ofthe volcaniclastic DFP probably occurred as aresult of abrupt and undrained loading by thepelagic DFP. Persistence of excess pore pressuresbeneath the pelagic DFP for the whole ¯ow eventcould have occurred because the over-ridingpelagic DFP provided an effective blanket, pre-venting dissipation of the induced high porepressures. The ®ne-grained component and ho-mogeneous fabric of the volcaniclastic DFP mayalso have reduced the permeability, which wouldnot only have preserved, but prevented, dissipa-tion of excess pore pressures. However, theevidence that the volcaniclastic DFP appearedincapable of travelling far beyond the limits of the



pelagic DFP suggests that the latter mechanism isof lesser importance. Whatever the precise mech-anism, the volcaniclastic DFP would have beenable to form a laterally continuous basal layer,capable of acting as a low apparent friction layerand of supporting and transporting the pelagicDFP. The actual value of excess pore pressuremay ¯uctuate during ¯ow as a result of variationsin shear rate (Sassa, 1988) and also as a result ofparticle size or permeability of the ¯owing mate-rial (Hutchinson, 1986). These ¯uctuations maygive rise to a series of ¯ow pulses advancingdownslope (Iverson, 1997; Major, 1997). Evidencefor this may be preserved as the series ofoverlapping debris ¯ow lenses observed on3á5 kHz records.

A model for the emplacement of the Saharandebris ¯ow

A conceptual model showing the development ofthe Saharan debris ¯ow from an initial singlepelagic debris ¯ow to a two-phase debris ¯ow inwhich the additional basal ¯ow phase wasgenerated by failure and shearing of the substrateis illustrated schematically in Figure 12. Stage 1shows the pelagic DFP before mobilization of thebasal volcaniclastic layer. At this stage, we do notknow whether the debris movement was achievedby distributed deformation throughout the debrismass or whether deformation was alreadyconcentrated near the base of the ¯ow. Aftergeneration, the volcaniclastic DFP (stage 2)supported the pelagic DFP, which it carried as apassive raft. At this point, internal deformation ofthe pelagic DFP would have decreased, possiblyaccounting for the relatively small amount ofdeformation observed within the pelagic DFPdespite a maximum runout of 700 km. Steady-state ¯ow (i.e. when both debris ¯ow phasestravel with the same velocity and ¯ow volumeremains constant) is illustrated in stage 3. In theterminal ¯ow phase, the volcaniclastic DFP isshown ¯owing ahead of the pelagic DFP, re¯ect-ing its lower shear strength and therefore itsability to ¯ow on very low slopes (stage 4). The¯ow of the volcaniclastic DFP from beneath themargins of the pelagic DFP may have continuedafter emplacement of the main debris ¯ow mass,because of residual excess pore pressures. How-ever, it is important to recognize that the volca-niclastic DFP was deposited a relatively shortdistance beyond the pelagic DFP margin,re¯ecting the apparent ef®ciency of the pelagic

DFP in sealing the excess pore pressure withinthe volcaniclastic DFP layer.

Understanding the precise relationship be-tween the volcaniclastic and pelagic DFPs isclearly the key to understanding the modelpresented in Figure 12. In discussing the originof the volcaniclastic DFP, we have already sug-gested that this debris ¯ow phase was derivedfrom a limited source area. This would suggest amechanism in which, after mobilization of thevolcaniclastic DFP, the two debris ¯ow phasesmoved as a single complex unit. A possiblemechanism is one in which the volcaniclasticDFP ¯owed just ahead of the pelagic DFP layingdown a carpet of easily remobilized material. Thisis an attractive hypothesis because the excesspore pressure `reservoir' could have been contin-uously recharged by reloading of a layer ofvolcaniclastic DFP that had escaped from beneaththe pelagic DFP and ¯owed ahead of it, only to beover-ridden again by the pelagic layer. Derivationof the volcaniclastic DFP from a limited areaappears to rule out a ¯ow mechanism in whichprogressive failure of the substrate occurredcontinuously as the ¯ow moved downslope.However, this type of substrate failure may also

Fig. 12. Model of Saharan debris ¯ow mechanics il-lustrating the generation of the volcaniclastic debris¯ow phase and its relationship to the overlying pelagicdebris ¯ow phase.



have occurred to some extent and may have beenthe source of locally derived marl clasts into thevolcaniclastic DFP. This process has beendemonstrated in the nearby Canary debris ¯ow(Masson et al., 1997).

Accretion of material to the base of an activedebris ¯ow has important implications in terms of¯ow runout. It would be possible for debris ¯owsto increase in volume during ¯ow, thereby in-creasing their potential runout. In areas where therate of volume increase as a result of failure of thesubstrate is matched by the rate of depositionbehind the ¯ow head, a debris ¯ow would be ableto maintain a constant volume, even though thedebris deposit left behind would suggest a wan-ing ¯ow. An element of apparent enhancedrunout might be created, in that not all the ¯owmaterial would be transported over the full lengthof the ¯ow path. This effect could have consid-erable signi®cance. In the case of the Saharandebris ¯ow, the volcaniclastic DFP comprisesapproximately one-third of the total ¯ow volumewest of the island of El Hierro and extends theoverall ¯ow length 70 km beyond the limit of thepelagic DFP.

Deposition of debris ¯ow phases

Deposition of the Saharan debris ¯ow might havebeen controlled by a variety of factors, includingdecreasing gradient, thinning of the ¯ow, dissi-pation of pore pressure and obstacles in the ¯owpath. The simplest explanation for deposition isthat the slope gradient diminished to around0á05°, which was suf®cient for the frictionalforces to exceed downslope driving forces.

In terms of ¯ow thickness, a simple Binghammodel of debris ¯ow predicts that deposition willoccur once the overall thickness decreases belowa critical value, as governed by its yield strength(Johnson, 1970). However, if the pelagic debris¯ow was carried as a largely passive raft, then itsthickness and yield strength might not be animportant factor controlling deposition. In thissituation, the thickness of the volcaniclastic DFPmay be the key factor. Although we have littleconstraint on the volcaniclastic DFP thickness,we know that its composition and grain sizeremain relatively constant over the length of the¯ow (from core sites D24 to D16). This suggeststhat no signi®cant extra material was added to itdownslope from core site D24. Thus, if depositionwas occurring, it might have thinned beyond thecritical level at which Bingham ¯ow criteria werepossible.

If high pore pressures in the underlying volca-niclastic DFP supported the pelagic DFP duringtransport, then the dissipation of excess porepressures within the volcaniclastic DFP may havecontributed to deposition. Dissipation intothe subsurface or from around the margins ofthe moving debris mass might both be important.

Another factor that could have slowed thedebris ¯ow is the occurrence of obstacles inthe ¯ow path. 3á5 kHz pro®les from the distalregion of the debris ¯ow show bathymetric highsthat the volcaniclastic DFP ¯owed over or around.These may have had the effect of splitting the¯ow into smaller lobes, which would have had asmaller runout potential.

Comparison with other landslidesand debris ¯ows

We are not aware of any other examples ofsubmarine landslides in which a basal layer withdistinct rheological properties has beendescribed. For the north-west African margin,regional maps produced by Jacobi & Hayes (1982,1984a) show numerous elongated areas charac-terized by strong prolonged echo characterextending away from the base of the continentalslope into over 5000 m water depth. Aroundvolcanic islands such as the Canaries and CapeVerdes, these may indicate volcaniclastic DFPssimilar to those associated with the Saharandebris ¯ow. However, the similarity in echosignature of turbidite pathways and areas ofvolcaniclastic DFP prevents de®nitive recogni-tion of further areas of volcaniclastic DFP.Unsorted sandy volcaniclastic sediments are rec-ognized in several cores from the area of theCanary debris ¯ow but are absent from adjacentareas (Masson et al., 1997; unpublished data). It isnot easy to establish whether these volcaniclas-tics form a discrete basal layer, because theCanary debris ¯ow is entirely composed ofmaterial from the western Canary Island slopes,and thus the clear lithological distinctionbetween the basal layer and the upper raft ofmaterial, as present in the Saharan ¯ow, islacking. However, in one core, a thin layer ofvolcaniclastic sand is observed to coat a 5-m-thickfolded clast, suggesting that the clast may havebeen carried by ¯ow in the volcaniclasticmaterial. Most of the Canary debris ¯ow coresthat sampled volcaniclastic material also occur inthe distal ¯ow area, suggesting a situation similarto that observed in the Saharan ¯ow.



Descriptions of subaerial landslides with two-phase structures similar to that of the Saharandebris ¯ow can be found in the publishedliterature. A basal layer described as `sand andgravel base material which had run ahead of thecoarse debris of the slide' and ®ne-grained`splash' material around the edges of the mainslide were observed in the Frank Slide in Alberta,Canada (Cruden & Hungr, 1986). The `splash'material is interpreted as having been highly ¯uidwith a low strength at the time of emplacement,as it spread out into a thin sheet extending about100 m around the coarse debris. Although noevidence is presented as to the source andmobilization of the ®ne-grained basal layer ofthe Frank Slide, the ¯ow mechanism appears tobe very similar to that of the much larger Saharandebris ¯ow.

The Blackhawk Landslide (Shreve, 1968) has athin layer of muddy sand exposed around itsmargins. The same facies is also seen in twointernal `windows' through to the base of thelandslide and in a number of clastic dykesinjected into the main landslide mass (Johnson,1978). The key observation is that ®ne-grainedbasal material occurs beneath and has ¯owedbeyond the margins of the landslide. Both Johnson(1978) and Melosh (1986) noted that the basalmaterial showed evidence of `a relatively highdegree of ¯uidity during emplacement of thelandslide'. As with the Frank Slide, the overall¯ow mechanism appears to be similar to that ofthe Saharan debris ¯ow.

The Ontake-san debris avalanche, which oc-curred in 1984 in Japan, produced a deposit thatconsists mainly of `large debris blocks¼sur-rounded by a matrix of rock fragments of thesame composition as the blocks' (Voight & Sousa,1994). However, a thin layer of `®ne-grainedsaturated sediments', also described as `a pum-iceous slurry', is found at the base and around themargins of the coarse-grained deposit. Sandvolcanoes associated with the basal layer areevidence for elevated pore pressures within thislayer during avalanche emplacement (Voight &Sousa, 1994). A two-layer model has thus beenproposed for this debris avalanche, with therelatively dry coarse debris carried on a saturatedlayer of ®ner grained material (Sassa, 1988;Voight & Sousa, 1994). The basal layer of theOntake-san debris avalanche appears to be verysimilar to the basal volcaniclastic DFP describedin this paper, in that both deposits consistessentially of poorly sorted muddy sand. Sassa(1988) proposed a model for the Ontake-san

avalanche in which the basal layer was largelygenerated by basal accretion of saturated, ®ne-grained material from the ¯oor of the avalanchepathway. This model is clearly very similar tothat proposed for the Saharan debris ¯ow. Voight& Sousa (1994), while accepting the two-layermodel, propose that the bulk of the basal layerwas derived from a 2-m-thick pumice bed in theavalanche source area and that basal accretionmay have been of lesser importance. However,irrespective of the source of the basal material, thetwo-layer nature of the Ontake-san debris ava-lanche clearly supports our interpretation of theSaharan debris ¯ow deposit.

CONCLUSIONS

The Saharan debris ¯ow west of the island of ElHierro consists of two distinct debris ¯ow phases,with a lower volcaniclastic phase overlain by apelagic phase. We propose that the volcaniclasticphase was generated when the sea¯oor on the¯anks of El Hierro was loaded and mobilized by apelagic debris ¯ow originating from higher on thenorth-west African continental margin in waterdepths of around 2000 m. The sediment remobi-lized by loading may have consisted of volcanic-lastic turbidites or volcanic debris from a ¯ankcollapse on El Hierro, with some intermixed ®ne-clay-grade, pelagic material. Under loading by thepelagic debris ¯ow phase, the volcaniclasticmaterial appears to have formed a highly ¯uidsandy debris ¯ow, capable of transporting with itthe huge volumes of pelagic debris. We suggestthat the pelagic debris formed a thick impermeableslab above the volcanic debris, thus maintaininghigh pore pressures generated by loading andgiving rise to low apparent friction conditions.The distribution of the two debris phases indicatesthat the volcaniclastic debris ¯ow came to a haltwithin a few tens of kilometres after it escapedfrom beneath the pelagic debris ¯ow, probablybecause of dissipation of excess pore pressurewhen the seal of pelagic material was removed.

The long runout of the Saharan debris ¯ow overslopes as low as 0á05° can be explained by a basallayer within which low-friction conditions werecreated and maintained. In the case of the Saharandebris ¯ow, the basal layer consists of poorly sortedvolcaniclastic material ranging from coarse sand toclay grain sizes, and the contrast in lithologybetween the volcaniclastic and pelagic phasesallows the distribution and role of the two phasesto be de®ned clearly. An implication of substrate



failure is that the ¯ow can increase or maintain aconstant volume as it progresses downslope,thereby further increasing total runout potential.

In the Saharan debris ¯ow, the presence ofpelagic and volcaniclastic sediment in the sub-strate, with the probability that these mixed tocreate the volcaniclastic debris ¯ow phase, is thekey factor contributing to long runout. However,it seems certain that it is the poor sorting of thevolcaniclastic phase, rather than its volcaniclasticcomposition, that gives it its apparent low frictioncharacter. We do not know de®nitely whetherpoor sorting was a primary characteristic of thein situ volcaniclastic material or whether it wasachieved by mixing with pelagic sediment de-rived from the substrate during mobilization.However, the presence of marl clasts within thevolcaniclastic debris ¯ow phase and nannofossilevidence indicating a considerable depth oferosion suggest that the poorly sorted materialwas formed by mixing during or shortly afterfailure. In a more general application to theunderstanding of submarine debris ¯ows (andother types of landslide), we would thereforesuggest that any heterogeneous substrate, inwhich a variety of grain sizes was available,could have the potential to fail in a similarmanner. On continental slopes, for example, anysuccession containing coarse-grained turbiditescould meet this criterion. In practice, recognitionof a discrete basal facies in other landslidedeposits may be dif®cult, because a lithologicalcontrast as obvious as that observed in theSaharan debris ¯ow is unlikely to be a commonoccurrence. Much more likely is a situation inwhich both basal and upper facies are derivedfrom one area of slope and contain lithologicallysimilar sediments. In this situation, the occur-rence of a sandy debris ¯ow facies at the base ofthe landslide deposit should, however, still allowrecognition of the two-phase ¯ow described inthis paper. Evidence for such a basal facies in afew other major landslides leads us to believe thatthis may be the key to understanding long runoutin a variety of landslide types.

ACKNOWLEDGEMENTS

We gratefully acknowledge the Master, crew andscienti®c party of RRS Discovery cruise 205 (fordata collection) and Dr G. Kalher (for nannofossildata). Discussions with P. P. E. Weaver have alsobeen useful in formulating some ideas in thispaper. Financial support from the EC MAST II

project MAS2-CT94±0083 (STEAM) is gratefullyacknowledged. M. J. R. Gee acknowledges NERCstudentship GT4/95/252.

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Manuscript received 22 April 1998;revision accepted 4 June 1998.



the saharan debris flow: an insight into the mechanics of long runout submarine debris flows

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