Turbidite Variability and Architecture of Sand-Prone, Deep-Water Slopes Eocene Clinoforms in the Central Basin, Spitsbergen - JSR, 2001.pdf

7/25/2019 Turbidite Variability and Architecture of Sand-Prone, Deep-Water Slopes Eocene Clinoforms in the Central Basin, Sp…

http://slidepdf.com/reader/full/turbidite-variability-and-architecture-of-sand-prone-deep-water-slopes-eocene 1/18

JOURNAL OF SEDIMENTARY RESEARCH, V OL. 71, NO. 6, NOVEMBER, 2001, P . 895–912Copyright 2001, SEPM (Society for Sedimentary Geology) 1527-1404/01/071-895/$03.00

TURBIDITE VARIABILITY AND ARCHITECTURE OF SAND-PRONE, DEEP-WATER SLOPES: EOCENECLINOFORMS IN THE CENTRAL BASIN, SPITSBERGEN

PIRET PLINK-BJORKLUND1,* DONATELLA MELLERE,2 AND RON J. STEEL3

1 Institute of Energy Research, University of Wyoming, Laramie, Wyoming 82071, U.S.A.

e-mail: [email protected] Department of Geology, University of Padova, Padova, Italy

3 Dept. of Geology and Geophysics, University of Wyoming, Laramie, Wyoming 82071, U.S.A.

ABSTRACT: The architecture and turbidite variability within sixwedge-shaped (downslope-thinning), sand-prone slope accumulationsare documented from Eocene shelf margins on Spitsbergen. The Cen-tral Basin formed as a small foreland or piggy-back basin, and thestudied turbidites accumulated mainly on the slope portion of sand-prone clinoforms that developed during depocenter migration and in-filling of the basin. The shelf-margin clinoforms have amplitudes (min-imum water depths) from 100 to 350 meters, and their shelf, slope, andbasin-floor segments are well imaged, and can be walked out alongmany of the mountainsides.

Only a small percentage of the clinoforms are sand prone, and thesedeveloped when sea level occasionally fell to or below the shelf edge.

Of the sand-prone clinoforms, some had their sand budget partitionedmainly out onto the basin floor (basin-floor fans), but most trapped thesand on the slope only. The latter are now visible as downslope-thin-ning wedges, some 2.5–3.5 km in downdip extent. The turbidites withinthis type of clinoform have been examined and classified.

The lower-slope to base-of-slope segment of the studied clinoformcomplexes are dominated by lobes consisting of broad, shallow chan-nels and sheet-like turbidites, becoming heterolithic and muddy out onthe basin floor. Beds on the lower slope vary from thick (up to 4.5meters), ungraded or laminated sandstones, to thinner ungraded sand-stones with coarse cappings. The middle-slope segment of clinoformcomplexes is dominated by narrow channels (chutes) that feed down-slope to progradational chute-mouth lobes. Chutes contain ungradedand laminated sandstone beds up to 3 m thick, whereas the chute-mouth lobes show alternations of thinner, ungraded to laminated or

rippled sandstones. These lobes become more heterolithic and muddydownslope. The shelf-edge to upper-slope segment of clinoform com-plexes is dominated by upward-coarsening and -thickening sheetsandsof steep-fronted shelf-edge deltas. The sandsheets of the delta front canbe traced updip into mouth-bar and distributary-channel sandstones.

It is argued that shelf-margin accretion, represented by the sand-prone slope wedges, was achieved mainly by sand-laden currents thatflooded from the shelf edge as hyperpycnal flows. This hypothesis issupported by: (1) the direct connection between channel and mouthbar systems at the shelf edge, and the turbidites of the slope lobes, (2)the systematic progradational character of the slope lobes, (3) the ab-sence of large-scale slump scars, gullies, or canyons on the slope, and(4) the dominance of a type of turbidite that implies deposition fromsustained flow.

Detailed examination of the architecture of one of the slope wedgesshows that there are unconformities developed within the stratigraphybelow the shelf edge and that these erosional terraces beheaded thedeltas perched on the uppermost slope. The erosion surfaces indicatefall of sea level to this position. Despite the magnitude of this fall (upto 80 meters), the lack of canyons on the slope prevented the construc-tion of basin-floor fans. Such falls of sea level, on non-canyoned slopes,simply promote sand-prone accretion of the shelf margin.

* Present Address: Department of Earth Sciences: Geology, Goteborg University,Box 460, SE-405 30 Goteborg, Sweden.

INTRODUCTION

The Central Eocene Basin of Spitsbergen was a small foreland (Steel etal. 1985) or piggyback (Blythe and Kleinspehn 1998) basin bounded to thewest by an active fold-and-thrust belt (Fig. 1). The basin was asymmetri-cally infilled from rivers draining a rising and eastward-migrating fold-and-thrust belt (Harland 1969; Steel et al. 1985). The basin was at least 100km wide and infilled with latest Paleocene to Eocene coastal plain-shelf-slope-basin-floor deposits that filled an eastward-migrating depocenter withtime. The general stratigraphy and paleogeographic setting for the latestPaleocene–earliest Eocene stages of infilling of the Central Tertiary Basinhave been discussed briefly by Steel et al. (1981) and Steel et al. (1985),and later in more detail by Helland-Hansen (1990, 1992). The eastwardand southeastward migration of the basin depocenter, driven by tectonicloading, created an asymmetric sedimentary succession (Helland-Hansen1990), which was more than 1.5 km thick in the west, thinning to less than600 m in the east. The character of the deltaic and barrier shoreline depositshave been documented in several previous works (e.g., Kellogg 1975; Steel1977; Helland-Hansen 1990), but only few researchers have addressed theslope and basin-floor accumulations (Steel et al. 1981; Nyberg et al. 1995).

Individual phases of linked shelf-slope-basin-floor sand deposition canbe seen on mountainside exposures (10 1 km scale), with the basin-floordeposits developing in water depths of 150–350 m, as judged by the am-plitude of the shelf-slope clinoforms (Figs. 1, 2). Mapping over distancesof 40 km out from the western edge (foredeep) of the basin demonstrates

that the shelf to basin-floor clinothems occur in a shingled stratigraphicpattern, with successively younger clinothems being offset basinwardsthrough time (Fig. 2). Most of the basin-infilling clinothems in the CentralBasin of Spitsbergen are shale-prone throughout (Type 4), or are sand-prone only as far out as their inner-shelf reaches, and shale-prone fromouter-shelf to slope and basin-floor reaches (Type 3). However, there areoccasional clinothems that are sand-prone across the entire shelf, bringingsands onto the slope (Type 2), and even rarer clinothems that give evidenceof sand transport far out beyond the slope, onto the basin floor (Type 1)(Mellere et al. 1997; Plink et al. 1997; Steel et al. 1997). The latter typeof clinoforms belong to the Battfjellet Formation, sharply overlie the Gil-sonryggen Formation shales, and are themselves overlain by marine shalesin younger downlapping shaly clinothems of the Battfjellet Formation.

This communication focuses on the Type 2 clinothems, clinothems thathave sandstones that reach as wedge-shaped bodies from the shelf edge tothe lower slope (Hogsnyta) (Figs. 1, 2) or for short distances onto the basinfloor (Haagfjellet, Fossilfjellet, Finsenfjellet, Bjorsonfjellet, and Semmel-ryggen) (Fig. 1). The sandstone wedges are up to 80 m thick at the shelf edge and pinch out on the lower slope or just beyond, onto the basin floor.The normal dip extension of the sandstone wedges is within in a distanceof ca. 3500–5500 m. These Type 2, slope to basin-floor sandstone wedges(Fig. 2) occur north from Van Mijefjorden and are significantly differentfrom Type 1 clinothems documented in Van Keulenfjorden (Steel et al.2000) (see Fig. 1). The Type 1 clinothems have thick basin-floor fans farbeyond the base of slope, and their coeval slope segments are canyonizedand bypass dominated (Fig. 2).



896 P. PLINK-BJO RKLUND ET AL.

FIG. 1.—The location of the six studied shelf-slope clinoform complexes (1–6) in the Central Basin, Spitsbergen. The broken line at each location represents the shelf-edge position prior to a stratigraphic rise. The arrows show the sediment supply direction and clinoform migration path. The main map shows the Central Basin in relationto the West Spitsbergen fold-and-thrust belt. (1) Haagfjellet, (2) Fossilfjellet, (3) Finsenfjellet, (4) Bjornsonfjellet, (5) Semmelryggen, (6) Hogsnyta, St Storvola, HrHyrnestabben.

The purpose of this paper is to (1) document the internal architecture of the sandstone wedges within the sequence stratigraphic framework; (2) de-

scribe the range of sedimentary facies within the Type 2 clinoform slopewedges; and (3) attempt, in particular, to separate turbidites from debrisflow deposits, as well as to describe the wide diversity of turbidite deposits.The larger-scale development of the basin, as well as the Type 1 clinothemcomplexes, are discussed elsewhere (Steel et al. 2000).

METHODS

The shelf-edge to basin-floor sandstone wedges, investigated by mappingof lateral facies and measuring vertical sections across six mountainsideexposures, are about 2500–3500 m wide and up to 80 m thick (Fig. 1).The excellent quality of outcrops enabled ‘‘walk-out’’ of internal architec-ture (Fig. 2) as well as individual bedding surfaces where the mountainsideswere not too steep, otherwise detailed lateral relationships were determinedfrom helicopter photomosaics. The exposures provide a 2-D picture of

shelf-edge to basin-floor clinothems in an approximate dip direction (fromnorthwest to southeast).

Bed-Scale Study

The bed-scale study of the units within the slope to basin-floor wedgesdocuments a diversity of gravitational processes, and good quality outcropsenable differentiation between slumps, debris flows, and different types of turbidites. In addition to vertical measured sections, some individual unitsand beds were followed in outcrop in a downslope direction for at least afew hundred meters, to document changes along the paleoflow direction.We strongly support the view that it is possible to differentiate turbidites

from debris flows in outcrops on the bases of sedimentary structures, bedgeometry, and bounding surfaces (see discussion in Hiscott et al. 1997;

Shanmugam et al. 1985; Shanmugam et al. 1995; Shanmugam 1996). Inthis paper a deposit is interpreted as a turbidite if we could document (1)normal grading, (2) internal scours, or (3) any traction structures like cur-rent ripples or plane-parallel lamination, together with (4) sheet-like orchannel geometry, and (5) erosional bases. Lack of these features togetherwith (1) shear structures, (2) chaotic fabric, (3) pinch-outs without thinning,and (4) topography-building geometry are used as criteria for debris-flowdeposits. A turbidity current is defined as a fully turbulent noncohesiveflow, and a debris flow as a cohesive flow supported by matrix strengthand dispersive pressure.

This study documents that the range of beds named ‘‘turbidites’’ is muchmore diverse than suggested by the Bouma model (see Bouma 1962; alsoLowe 1982; Pickering et al. 1995; Branney et al. 1990; Kneller et al. 1991;Middleton 1993; Edwards et al. 1994; Kneller and Branney 1995; Kneller1995; and Hiscott et. al 1997). In order to classify this diversity we have

compared existing classifications of Bouma (1962), Lowe (1982) andKneller (1995).

Study of Internal Architecture

The well exposed, seismic-scale outcrops on Spitsbergen are such thatfacies associations could be identified by their paleo-bathymetric positionwithin the clinoforms (Fig. 2A, B) as (1) shelf-edge and upper-slope, (2)middle-slope, or (3) lower-slope to base-of-slope deposits. The three faciesassociations are characterized by a distinct combination of lithofacies, bedgeometry, and occurrence and type of channels.



897 ARCHITECTURE AND TURBIDITE VARIABILITY OF SAND-PRONE, DEEP-WATER SLOPES

FIG. 2.—Clinoform geometries. A) The clinoforms develop from the shelf edge to the basin floor and vary in their height up to 350 meters. Type 1 clinoforms showmainly bypass from the shelf edge to lower slope but have a sandy development from the base of slope out onto the basin floor. Type 2 clinoforms develop a sandy slopewedge but have no significant basin-floor extension. B) Example of Type 2 clinothems on Hogsnyta (loc. 6, Fig. 1). C) Examples of Type 1 and Type 2 clinothems onStorvola (St; Fig. 1).

The studied slope wedges accumulated basinwards of the sandy shelf segment of clinothems and consist of: (1) shelf-edge deltas with steep-fronted delta-front sheets, and distributary channels and mouth bars on theshelf-edge and upper-slope reaches, (2) sandy to heterolithic mound-likelobes, tens to hundreds of meters wide, dissected by low-sinuosity channelsca. 50–100 m wide and up to 5 m deep on the middle-slope reaches, and(3) sandy to heterolithic sheets dissected by channels 50–100 m wide and2–4 m deep on the lower-slope to base-of-slope reaches. In the proximalpart of the basin, next to the thrust front (Fig. 1), there is preservation of only middle-slope (segment 2) and lower-slope to basin-floor (segment 3)

portions. Farther into the basin, on Hogsnyta (Fig. 1), the whole successionfrom shelf edge to lower slope is exposed.

ARCHITECTURE

Following Rich (1951) we use the term clinoform for the surface ex-pression of the shallow-water to deep-water, shelf-slope-basin-floor profile,whereas the term clinothem denotes the rock unit forming along the lengthof the clinoform. Although Rich (1951) restricted the term clinoform to the




‘‘slope’’ part of the bathymetric profile, we use it for the entire sigmoidallyshaped profile.

The studied clinothem complexes consist of a series of sand-prone, par-tially amalgamated clinothems that can be traced from shelf edge to basinfloor. The individual clinothems are 5–10 m thick in their landward reaches.Muddy and heterolithic units (up to 0.5 m thick) that thicken basinwardsseparate the individual clinoforms. In fact, each individual shelf-edge tobasin-floor clinoform consists of series of prograding lobate bodies: (1)

delta lobes at the shelf edge and on upper-slope reaches, (2) minor lobesin the middle-slope reaches, and (3) extensive lobes on the lower-slope tobase-of-slope reaches (Fig. 3). The prograding lobes are separated by thinheterolithic beds that merge at the bottomsets, forming the upward-finingmuddy units that separate individual clinoforms. The prograding sand-prone lobes with their heterolithic to muddy fringes are the ‘‘buildingblocks’’ of the clinothem complexes. Very commonly the topographicallyhighest parts of the prograding lobes are cut by channels, i.e., the uppersurfaces of individual clinoforms are channelized (Figs. 3, 4).

The stacking pattern of clinothems in different clinothem complexesvaries significantly with distance from the driving fold-and-thrust belt andrelated change in water depth (water depth decreases away from the fold-and-thrust belt; see Fig. 1). The four slope wedges closest to the fold-and-thrust belt (Haagfjellet, Fossilfjellet, Finsenfjellet, and Bjornsonfjellet) are

simple basinward-prograding, sand-prone bodies (Fig. 3). The next slopewedge basinwards, Semmelryggen, steps basinwards initially but is land-wards-stepping in its final phases. The Hogsnyta wedge, the most distalwith respect to the fold-and-thrust belt, has the most complex architecturewith several internal unconformities. All the slope wedges prograde ontoolder marine shales and are draped by younger marine shales.

The volumetric partitioning of sediment into different segments of cli-nothems varies in the different wedges (Fig. 3). The wedges closer to thethrust belt (Fig. 1) contain a higher proportion of lower-slope to basin-floordeposits (up to 90%). The youngest turbidite wedge, on Hogsnyta (Fig. 1),has a much higher proportion of middle-slope deposits (ca. 50%) and shelf-edge and upper-slope deposits (ca. 25%).

Hogsnyta

The Hogsnyta wedge (Figs. 3, 4) was deposited into an initial waterdepth of ca. 150 m. This wedge is 70 m thick in its shelf-edge end, and itpinches out before reaching the base of slope, across a distance of ca. 3000m. Hogsnyta wedge can be subdivided into a lower, progradational seriesof clinothems and an upper, aggradational to retrogradational series of cli-nothems. A significant difference between Hogsnyta wedge and the otherwedges is that the shelf-edge reaches are preserved, allowing us to separateerosional unconformities from local channel erosion surfaces. Furthermore,the lower-slope deposits of Hogsnyta thin and fine very markedly withinthe lower slope and pinch out before reaching the base of slope. The lower-slope deposits are generally thinner bedded and finer-grained compared tothe other wedges. Continuous exposures on Hogsnyta, across the wholelength and height of the wedge (Fig. 2B) provide an excellent basis for anarchitecture study.

Progradational Clinothem Series.—The progradational series of cli-nothems consists of two architectural compartments. The oldest compart-ment contains steep-fronted shelf-edge deltas that prograded onto the upperslope, just below the shelf edge (Figs. 3, 4). The deltas, whose tops areseverely eroded, form two upward-coarsening, prograding clinothems, eachca. 10 m thick. The sandy prograding delta fronts here are unusually steep(up to 20 degrees), presumably because they were the first to build downfrom the shelf edge. The sandy delta-front sheets are separated by thinheterolithic to muddy units.

The second compartment contains a series of larger-radius clinothemsthat contain minor lobes and channels on the middle slope, and sheet-liketurbidites on their lower-slope segments (Figs. 3, 4). The upper-slope reach-

es are preserved only within the two youngest clinothems, where they con-sist of steep-fronted deltas with mouth-bar and distributary channels at theirtops. These youngest clinothems are truncated by two unconformities thaton the middle slope pass into an intensely channeled surface. The secondand most prominent erosion surface merges with the older at the landwardend of the outcrop. As the clinoforms prograded basinwards, there was amarked downward shift in their depositional trajectory.

The middle-slope reaches of the clinothems have a high density of chan-

nels, especially along the unconformities. Furthermore, the size and densityof channels is higher in the progradational series of clinothems, than in theaggradational–retrogradational series.

Aggradational Clinothem Series.—The aggradational series of clino-thems occur in a single compartment. The clinothems stack vertically andshow a landward onlap (Fig 4). They are composed of channels and minorlobes on the middle slope, and sheet-like turbidites on the lower slope.Relatively small channels (up to 25 m wide and a few meters deep) occurat the top of individual clinothems. These clinothems onlap the major ero-sion surface at the top of the progradational series (Fig. 4). The uppermostclinothem onlaps the major erosion surface as far back as the shelf-edgeregion (Fig. 4).

Retrogradational Clinothem Series.—The retrogradational series of clinothems are contained in two compartments with wave-influenced deltas

on the shelf edge and upper slope, minor lobes on the middle slope, andsheet-like turbidites on the lower slope. The first compartment has clino-thems that are individually prograding but successively landward-stepping(Fig. 4). The clinothems have stepped significantly landwards relative tothe clinothems of the aggradational series, as shown by their lower-slopereaches now overlying middle-slope reaches (Fig. 3). This compartmenthas the lowest density of channels within the Hogsnyta sandbody complex,i.e., the middle slope is entirely dominated by minor lobes.

The visible part of the second compartment contains prograding clino-thems composed of middle-slope channels and minor lobes and lower-slopesheet-like turbidites. The oldest clinothem of this compartment has steppedlandwards with respect the underlying compartment, as shown by its mid-dle-slope facies directly overlying deltaic mouth bars (Figs. 3, 4). Somesmall channels occur at the top of this clinothem. The youngest clinothemis markedly progradational, and large channels extensively incise its top.

Semmelryggen

The Semmelryggen wedge is 22 m thick at its landward end, and itpinches out on the basin floor, across a distance of ca. 5000 m. The wedgewas deposited in a water depth of ca. 200 m. The Semmelryggen wedgehas a lower progradational series of clinothems and an upper retrograda-tional series of clinothems. As on Hogsnyta, the top of the progradationalseries is intensively channelized (Fig. 3). The lack of preservation of theshelf edge makes it difficult to know if this channel-erosion surface cor-relates into a subaerial unconformity farther landward. The density and sizeof middle-slope channels is generally greater in the progradational seriesof clinoforms. The lower-slope to basin-floor deposits are represented bysheet-like turbidites only; no channels were documented. The retrograda-tion is seen by the landward stepping of the sheet-like turbidites across themiddle-slope deposits of the progradational series.

Haagfjellet, Fossilfjellet, Finsenfjellet and Bjornsonfjellet

These slope wedges are 20–80 m thick in their landward reaches andthey pinch out on the basin floor, across a distance of ca. 3500–400 m.They were deposited in water depths of ca. 250–350 m. These prograda-tional wedges (their exposed extent) consist of middle-slope deposits andlower-slope to basin-floor deposits (Fig. 3). The density and size of chan-nels on the lower-slope to basin-floor reaches of these wedges is generallyhigher than in the more basinward wedges. Especially the Bjornsonfjellet



FIG. 3.—Two-dimensional (dip) architecture of the six studied slope wedges (clinothem complexes) at locations shown in Fig. 1. The original shelf edge is preservand lower-slope architectural elements are shown. Numbers show location of measured sedimentary sections.




FIG. 4.—Architecture of the Hogsnyta slope wedge. The downslope extent of the wedge is about 2.5 km. The clinothems making up the wedge are grouped into an earlyprogradational series, a middle aggradational series, and a late retrogradational series. Note the marked unconformities within and at the top of the progradational series.Location of measured sections is shown.

wedge has a high density of sandstone-filled channels at the top of lower-slope and basin-floor clinothems. Furthermore, the occurrence of thicksandstone beds on the lower slope and basin floor is characteristic of thesefour wedges.

LOWER-SLOPE TO BASIN-FLOOR DEPOSITS

The lower-slope and basin-floor deposits are dominated by sheet-liketurbidite beds, dissected by only a few shallow channels (50–100 m wideand a few meters deep) filled with turbidites similar to those in sheet-likebeds (Table 1, Fig. 5).

Sheet-Like Beds and Channels

Facies 1.1 and 1.2: Thick Sandstone Beds.—Facies 1.1 and 1.2 containfine- to medium-grained sandstone beds, up to 4.5 m thick, with sharp andflat or scoured bases (Fig. 5). Individual beds fine and thin basinwards.Thinner beds of Facies 1.1 (up to 2 m thick) commonly have a sheet-likegeometry whereas the thicker beds (4–4.5 m thick), together with beds of Facies 1.2, fill incisions 50–100 m wide and 1.5–4 m deep. Facies 1.1 and1.2 are most common in the lower-slope to basin-floor segment of clino-thems.

Facies 1.1 consists of ungraded beds (average 0.5–2 m thick), with in-ternal diffuse and discontinuous scour surfaces in places. Dish structures,and ball and pillow structures occur in the 2–4.5 m thick beds (Fig. 5).Facies 1.2 consists of alternating ungraded and plane-parallel or current-ripple laminated sandstone beds (average 2–2.5 m thick) with no verticaltrend of grain size or sedimentary structures. In many places aligned androunded shale clasts occur (Fig. 5). Clast orientation is random, or bedding-parallel and/or imbricated. The shale clasts, at bed bases or within thesandstone beds, are in places aligned above internal scour surfaces (Fig.5).

Interpretation.—Scoured or flat and sharp bases, along with the sheet-like geometry and internal erosion surfaces of ungraded beds of Facies 1.1,

as well as alternation of laminated and ungraded intervals of Facies 1.2,reflect deposition by turbidity flows (Hiscott et al. 1997). The thick bedswith no vertical trend of grain-size or structures, which fine and thin down-slope suggest deposition by gradual aggradation from sustained downcur-rent-decelerating turbidity currents (Kneller 1995; see also Branney andKokelaar 1992; Kneller and Branney 1995).

Constant grain size through the beds of Facies 1.1 indicates a fairly

constant shear stress through time, or rapid fallout of grains that disabledtraction and created an ungraded bed. The capture of fluids within the bedsconfirm rapid deposition by loss of flow capacity. The internal scour sur-faces reflect temporal variations in flow velocity and sediment flux withina single event, as constrained by the discontinuity of the scour surfaces andthe constant grain size above and below the surfaces.

Alternation of traction-deposited (laminated) intervals with ungraded in-tervals in Facies 1.2 indicates deposition triggered by loss of flow com-petence, and through alternation of the fallout rates. The ungraded intervalsreflect rapid fallout rates. The laminated intervals occurred when the ver-tical grain flux intermittently decreased and enabled a distinct interfacebetween the current and the substrate, i.e., a discontinuity of velocity, con-centration, or flow rheology at the base of the current (Arnott and Hand1989; Kneller and Branney 1995), allowing traction and the developmentof plane beds and ripples.

The alignment of clay clasts in Facies 1.2 signifies unrecognizable de-positional boundaries that migrated upwards during sedimentation. Theroundness of aligned clast reflects near-bed flow modification, i.e., bedloadtransport and clast orientation suggests traction deposition (Middleton1993; Johansson and Stow 1995; Kneller and Branney 1995; Hiscott et al.1997; cf. Shanmugam et al. 1995). Alternative interpretations, such as thesettling of clasts through a flow or gliding at the top of high-density inertia-flow layers (Postma et al. 1988), or the laminar flow of debris (Stauffer1967; Lowe 1982, 1988; Shanmugam et al. 1995) are unlikely because of clast fabric, roundness, traction structures and internal erosion surfaces (Ar-nott and Hand 1989; Kneller and Branney 1995; Hiscott et al. 1997).




FIG. 5.—Photographs, representative vertical section, and interpretations of lower-slope facies association.




TABLE 1.— Lower-slope to basin-floor deposits

Facies Textures Structures Bed Cont. Bed Geometry Bed Thick.HydrodynamicInterpretation

1.1 thick ungradedsandstone beds

fine- to medium-grained sandstone

ungraded beds, internal dif-fuse and non-continuousscours; dish, ball and pil-low structures;

sharp and flatbases

beds 2 m thick aresheet-like; beds 2m thick have flatchannel geometry;beds fine and thinbasinwards

up to 4.5 m, aver-age 0.5–2.0 m

steady downslope-de-celerating turbidityflows, long-livedconstant discharge

1.2 thick laminated andungraded sand-stone beds

fine- to medium-grained sandstone;rounded shale clasts

ungraded, plane-parallel orcurrent-ripple laminated;internal scours; shaleclasts occur at bases of sandstone beds and haverandom or plane-paralleland imbricated orientation

scoured bases fill gul lies up to 50 mwide and 5 m deep,average 1.5–4 mdeep; beds fine andthin basinwards

up to 4.5 m, aver-age 2.0–2.5 m

quasi-steady down-slope-deceleratingturbidity flows,long-lived constantdischarge, alternat-ing fallout rate

1.3 ungraded sandstonebeds with coarserlaminated tops


ungraded lower portionswith coarser plane-parallelor current-ripple laminatedtops; ripple laminatedtops; ripples usuallyclimbing

flat or scoured sheet-like geometry;beds fine and thinbasinwards

0.2–1.0 m waxing downs lope-ac-celerating turbidityflows, long-livedconstant discharge

1.4 fine sandstone bedswith siltstonecapping

fine-grained sandstoneto siltstone

ungraded, normally gradedflat or laminated sand-stone intervals with plane-parallel or current-ripplelaminated, very fine-grained sandstone or silt-stone cappings

flat sheet-like geometry;beds fine and thinbasinwards

from 0.05–0.15 mup to 0.25–0.50m

waning downslope-de-celerating turbidityflows, surge-typeepisodic flows

1.5 very fine sandstoneand siltstone beds

very fine grained sand-stone to siltstone/ mudstone

ungraded to laminated, veryflat fine sandstone whichgrades upwards into pla-nar- or current ripple-lami-nated siltstone/mudstone


0 .0 1– 0.1 0 m w an in g d own slo pe -d e-celerating turbidityflows, surge-typeepisodic flows

1.6 deformed sandstonebeds


folds, overturned folds, anddish and ball structures

sharp fill topographic lows up to 4.5 m; aver-age 0.5–1.0 m

water escape andslumping slope in-stability related, epi-sodic

Facies 1.1 is comparable to the massive Ta intervals in turbidites of Bouma (1962), S3 of Lowe (1982), and deposits of depletive (downslope-decelerating) steady turbidity flows accumulated by gradual aggradation of Kneller (1995, also Branney and Kokelaar 1992, Kneller and Branney1995). Facies 1.2 is comparable to S3 interval in turbidites of Lowe (1982)and deposits of depletive quasi-steady turbidity flows of Kneller (1995).

Facies 1.3: Ungraded Sandstone Beds with Coarser LaminatedTops.—Individual beds of Facies 1.3 consist of fine- to medium-grainedsandstone with ungraded lower parts and slightly coarser plane-parallel orcurrent-ripple laminated tops (Fig. 5). In places, thin plane-parallel or cur-rent-ripple laminated intervals dominate the beds. Climbing current ripplesoccur and may be present at the bases of slightly coarser ungraded intervals(Fig. 5). The laminated tops may disappear laterally, scoured by an over-lying sandstone bed. In most cases the erosional contacts are subtle. Theindividual beds, 0.2–1 m thick, have sharp, flat, or scoured bases, and sheet-like geometry. Individual beds fine and thin basinwards.

Interpretation.—The alternation of ungraded and laminated intervals,together with scouring suggest deposition from turbidity flows (Hiscott etal. 1997). The coarsening from the ungraded to the laminated intervals,and the occurrence of climbing ripples at the base of ungraded beds suggesta drop in fallout rates, probably due to a temporal increase in shear veloc-ities rather than a drop in concentration (Kneller and Branney 1995; seealso Hiscott and Middleton 1979). The latter suggests that Facies 1.3 wasdeposited by temporally accelerating flows that decelerated downslope (asseen by the basinward fining and thinning of beds). Intervals with thickerungraded sandstone and climbing ripples denote a quasi-steady state of fluctuating flows. These beds are similar to the deposits of depletive waxingturbidity currents of Kneller (1995).

Facies 1.4 and 1.5: Sandstone Beds with Siltstone Capping.—Facies1.4 and 1.5 consist of a succession of ungraded to normally graded andplane-parallel or current-ripple laminated sandstone intervals, capped bysiltstone or mudstone (Fig. 5). Individual beds are flat based and have a

sheet-like geometry. Beds of Facies 1.4 fine and thin basinwards into bedsof 1.5. Facies 1.4 and 1.5 typically occur on the outer margins of the lower-slope and basin-floor lobes, basinwards from Facies 1.1–1.3. The very finesandstone or siltstone capping in Facies 1.4 is typically only a few centi-meters thick (Fig. 5). Individual beds in Facies 1.4 are 5–15 cm thick inthinly bedded packages and 25–50 cm thick in thickly bedded packages.

The laminated intervals in Facies 1.5 consist of alternating siltstone andmudstone. Individual beds of Facies 1.5 are 1–10 cm thick and occur inpackages up to 1 m thick.

Interpretation.—Normally graded beds are a characteristic product of waning turbidity currents (Bouma 1962, Walker 1967, Lowe 1982, Mid-dleton 1993, Kneller 1995). Waning downslope-decelerating turbidites aredeposited from surge-type short-lived turbidity currents that gradually losetheir capacity to carry sediment (Lowe 1982, Hiscott 1994, Kneller andBranney 1995). The basinward fining and thinning indicates downslopedeceleration of flows. The preservation of siltstone or mudstone intervalsindicates low erosion and entrainment capacity of the flows. Facies 1.4 and1.5 resemble Tabcd and Tcde intervals, respectively, of Bouma turbidites(1962) as well as the products of the waning depletive turbidity currentsof Kneller (1995).

Facies 1.6: Deformed Sandstone Beds.—Facies 1.6 contains beds withfolds, overturned folds, and ball-and-pillow structures. The latter are com-monly partly destroyed by fold structures. The deformed fine- to medium-grained sandstone beds (average 0.5–1.0 m thick) are up to 4.5 m thick(Fig. 5) (Bjornsenfjellet, Finsenfjellet; see Fig. 1). In a few places theprimary structures of the deformed sandstone beds are preserved, showingthat Facies 1.6 occurs in the distal parts of thick Facies 1.1 sandstone bedsor as the lateral expression of laminated and ungraded sandstone beds withshale clasts (Facies 1.2).

Interpretation.—Folds and overturned folds indicate slope-inducedslumping or deformation by the shear stress of overriding currents (e.g.,Nichols 1996). Associated water-escape structures indicate that seepage




liquification by escaping pore fluids may have triggered the slumping. Thehigh sand content of the deformed beds suggests that sandy units weremore collapse-prone than the heterolithic deposits. Association of Facies1.6 with Facies 1.1 and 1.2 suggests an origin from collapsing of channelbanks.

Lateral and Vertical Distribution

The lower-slope to basin-floor turbidites described above fine and thindistally and laterally from thick-bedded, fine-grained sandstone beds (Fa-cies 1.1–1.3) to thin-bedded heterolithic beds (Facies 1.4 and 1.5) withinlow-relief and extensive lobes across a distance of ca. 3000 m. The prox-imal parts of these lobes are dissected by broad channels filled with Facies1.1 and 1.2. In places the channel banks have collapsed (Facies 1.6). Walk-out of outcrops shows that the lower-slope to basin-floor channels are con-nected to the middle-slope channels (Fig. 3).

MIDDLE-SLOPE DEPOSITS

The middle-slope segment of the studied clinothems is dominated bysmall channels (average 50–100 m wide) and associated minor lobes (afew hundred meters across) (Table 2, Fig. 6). The toesets of clinothems inthe middle-slope areas are less muddy than those in the lower-slope areas.

Channels

Facies 2.1: Channeled Sandstones.—Facies 2.1 has a typical channelgeometry. The erosionally based channels are infilled by fine- or medium-grained sandstones with shale and coal clasts (up to 5 cm in diameter) atthe base, overlain by ungraded or normally graded and upward-thinningand -fining, parallel-laminated or ripple-laminated intervals (Fig. 6). Someof the above units fine upwards into a siltstone or mudstone capping, a fewcentimeters to 25 cm thick. Beds within the channels may be ungradedthroughout, ungraded to laminated, or laminated throughout. Thick basalintervals often show multiple scour surfaces marked by lags of shale clasts(Fig. 6). Aligned shale clasts also occur without associated erosion surfaces.The upper parts of laminated intervals often contain small coal fragments.Individual beds are 0.3–3 m thick (average 1.8–2 m thick), and fill channels

50–100 m wide, cut a few meters into other channel beds or into lobateunits of Facies 2.2 or 2.3. Facies 2.1 fines and thins basinwards, and tendsto lose both the upper, laminated and the basal ungraded intervals. Stackedchanneled units form successions up to 20 m thick. The channels appearto be extensive in the dip-parallel succession, suggesting that they are fairlystraight.

Interpretation.—This facies was clearly deposited in channels. The bas-inward thinning and loss of both bases and tops of units within channelsreflect downslope acceleration of depositing turbidity currents (Kneller1995). The alternation of ungraded (high fallout rates) or laminated (lowfallout rates and traction) intervals in thick sandstone beds suggest sus-tained fluctuating flows. The aligned shale clasts mark repeated erosion.These beds are similar to the deposits of the quasi-steady accumulativeturbidity currents of Kneller (1995).

The small size and low sinuosity of the channels, their close relation thedelta front and their position on the slope suggest that these channels aresimilar to the chutes documented from modern slope and delta-front en-vironments (see Prior et al. 1981; Kostaschuk and McCann 1987; Prior andBornhold 1990; Phillips and Smith 1992; Carlson et al. 1992).

Minor-Lobes

Facies 2.2: Lobate Sandstone Units.—Facies 2.2 consists of repeatedalternations of ungraded to laminated, fine-grained sandstone sets or bedsarranged in high-angle (10–15), accretionary wedging units that dip down-slope (Fig. 6). Individual beds, 1–7 cm thick, are ungraded, plane-parallel

or current-ripple laminated. Thicker beds are commonly ungraded or con-tain climbing ripples, whereas thinner beds are typically laminated. Indi-vidual beds thin laterally without significant changes in grain size. Unitsare 0.3–2 m thick and are separated from each other by slight discordances.In places the units diverge towards the channels of Facies 2.1 or morecommonly, are cut by the channels of Facies 2.1. Successive units displaythinning-upward or thickening-upward trends (Fig. 6). Stacked successionsof such sandy units are up to 35 m thick.

Interpretation.—The downslope thinning of ungraded and laminatedbeds without significant fining reflects deposition by downcurrent-acceler-ating waning turbidity flows (see Kneller 1995). The sandy units are lobesat the terminations of the chutes of Facies 2.1, or in some places theyindicate overbanking from chutes that were too small to confine the tur-bidity currents. A levee interpretation is unlikely, because chutes alwayshave an erosional contact with the lobes.

Facies 2.3: Lobate Heterolithic Units.—Facies 2.3 consists of repeatedalternations of ungraded or planar/ripple-laminated, very fine-grained sand-stone, siltstone, and mudstone beds (Fig. 6). Beds, 1–10 cm thick, fineupwards, and thin and fine downslope. Sandstone/shale ratio changes lat-erally into areas dominated by siltstones and mudstones. The heterolithicunits, typically, 0.2–1.5 m thick, display an upwards-thinning or -thick-ening trend and occur as a fine-grained fringe of the sandy lobes of Facies

2.2. Successions of heterolithic units reach 10 m thick.Interpretation.—The basinward fining and thinning of beds within theselobes reflects deposition by waning and downcurrent-decelerating turbidityflows, similar to the Tcde of Bouma (1962) and depletive waning turbidityflows of Kneller (1995).

Facies 2.4: Deformed Sandstones.—Facies 2.4, like Facies 1.6, consistsof folded and deformed beds but passes laterally into the chute-mouth lobesof Facies 2.2 or channeled sandstones of Facies 2.1. It is likely that Facies2.4 represents locally slumped lobe margins or chute banks.


The middle-slope reaches of clinothems expose chutes (Facies 2.1) thatcommonly terminate in minor lobes (Facies 2.2 and 2.3), though in otherplaces the chutes continue onto the lower-slope areas (Facies 1.1 and 1.2).

Upslope towards the shelf edge, the chutes can be seen to cut into delta-front deposits (Facies 3.1–3.5). Successive chute-mouth lobes, whetherstacked progradationally or retrogradationally, tend to be cut by large (50–100 m wide) or small (10 m wide) chutes. Most of the chute-mouth lobesare sand-prone (Facies 2.2), and the heterolithic deposits (Facies 2.3) occurmainly as a fine-grained fringe to the lobes. The character of these lobesalso varies with the character of the chute fill. Upward-thinning, chute-fillsandstones of uniform grain size are generally associated with sandy slopelobes, whereas fining-upward abandoned chute fills are associated withmore heterolithic slope lobes

SHELF-EDGE AND UPPER-SLOPE DEPOSITS

The shelf-edge and upper-slope segments of clinoforms are dominatedby steep-fronted shelf-edge deltas that contain upward-coarsening succes-

sions of delta-front sheets, overlain and incised by mouth bars and distrib-utary channel systems (Table 3, Fig. 7).

Delta-Front Sheets

Facies 3.1 and 3.2: Ungraded and Laminated Sandstone Beds.—Fa-cies 3.1 and 3.2 contain medium- to coarse-grained sandstone beds withsharp and flat or scoured bases. The beds fine and thin basinwards, andcommonly contain coal fragments. The beds have a sheet-like geometry;they occur as steeply inclined foresets (up to 20) 2–10 m thick and pinchout on the upper-slope reaches of clinothems, in a distance of about 600m (Figs. 3, 7).




FIG. 6.—Photographs, representative vertical section, and interpretations of middle-slope facies association.




TABLE 2.— Middle-slope deposits


2.1 channeled sandstonebeds

shale and coalclasts; fine- tomedium-grainedsandstone; silt-stone/mudstonecapping

ungraded or normallygraded intervals andupward thinning andfining plane-parallelor current-ripple lami-nated intervals;

erosional; cutdown a fewmeters;

channel geometry 50–100 mwide; beds fine and thindownslope, and tend tolose both the upper-mostand the lower-most inter-val

0.3–3.0 m; average1.8–2.0 m

quasi-steady downslope-accelerating turbidityflows in chutes, long-lived constant dis-charge

2.2 lobate sandstoneunits

fine-grained sand-stone

repeated alternations of ungraded to plane-parallel or current-rip-ple laminated, fine-grained sandstonebeds

sharp; slight in-ternal discor-dances

intermediate-angle (10–15)divergent accretion unitswith reverse polarity; bedsthin laterally from channelaxes without significantchanges in grain size

beds 0.01–0.07 mbedsets: 0.3–2.0m

waning downslope-ac-celerating turbidityflows in chute-mouthlobes, long-lived con-stant discharge

2.3 lobate heterolithicunits

very fine-grainedsandstone, silt-stone and mud-stone

repeated alternations of ungraded or plane-parallel or current rip-ple-laminated beds

flat low-angle (less than 10) di-vergent accretion sets;beds fine and thin awayfrom channels

beds: 0.01–0.10 mbedsets: 0.2–1.5m

waning downslope-de-celerating turbidityflows in chute-mouthlobes, surge-type, epi-sodic flows

2.4 deformed sandstonewedges


folds, overturned folds,ball and dish struc-tures

sharp wedge like 0.4–3.0 m water escape and slump-ing, slope instabilityrelated lobe collapse,episodic

Individual beds of Facies 3.1, 0.2–1.5 m thick, are ungraded throughout

or contain diffuse and discontinuous internal scour surfaces. Facies 3.2consists of alternating sets of ungraded and plane-parallel or current-ripplelaminated sandstone beds. Laminated intervals are 0.25–2.0 m thick, andungraded intervals are 0.2–2.0 m thick.

Interpretation.—Scoured or flat and sharp bases, along with the sheet-like geometry and internal erosion surfaces of the ungraded beds of Facies3.1, as well as the alternation of laminated and ungraded intervals of Facies3.2, reflect deposition by turbidity flows (Hiscott et al. 1997). The thickhomogeneous ungraded sandstone beds of Facies 3.1 indicate depositionfrom flows sustained at relatively constant discharge, and basinward thin-ning and fining indicates downslope deceleration. Similar thick beds withinternal scour surfaces originated by bed aggradation from fluctuating butsustained flows. Alternation of traction-deposited (laminated) intervals withungraded intervals in Facies 3.2 indicates deposition triggered by loss of competence and through alternation of high and low fallout rates.

Facies 3.3: Very Fine Sandstone and Siltstone Beds.—Facies 3.3 con-sists of beds of ungraded to laminated, very fine sandstone, with cappingsiltstone or mudstone (Fig. 4). Individual beds are flat based, 1–10 cm thickand occur in packages up to 1 m thick. Beds of Facies 1.3 fine and thinbasinwards and occur in lowermost reaches of delta-front foresets and inbottomsets.

Interpretation.—Normally graded beds that fine and thin basinwardsare a product of waning and downslope-decelerating, surge-type turbiditycurrents (Bouma 1962; Lowe 1982; Middleton 1993). These beds representTcde of Bouma (1962) turbidites, and they are the product of waning de-pletive turbidity currents of Kneller (1995).

Facies 3.4: Deformed Beds.—Facies 3.4 contains heterolithic units andsandstone beds (up to 2 m thick) that are folded and otherwise deformed,and pass laterally into steeply dipping beds of Facies 3.1–3.3. They orig-

inated from slumping on the front of the shelf-edge deltas.

Distributary Channels and Mouth Bars

Facies 3.5 and 3.6: Channeled Sandstone Units.—Facies 3.5 and 3.6show channel forms that contain cross-stratified and plane-parallel lami-nated sandstone sets with coal fragments and rounded shale clasts. Unidi-rectional trough cross-stratified sets (5–25 cm thick) at the bases of thechannel fills are overlain by sets of plane-parallel laminated sandstone (Fig.7).

Facies 3.5 consists of medium- to coarse-grained sandstone units 0.3–1.0 m thick in channels 25–50 m wide. Facies 3.6 consists of fine- to

medium-grained sandstone units in channels up to 3–5 m wide and about

1 m deep.Interpretation.—The geometry of Facies 3.5 and 3.6 indicates deposi-

tion within channels. The trough-cross stratification originated from migra-tion of 3-D dunes, and the plane-parallel stratification was formed fromupper-phase plane beds. The association with mouth-bar units suggests de-position in larger (Facies 3.5) and smaller (Facies 3.6) distributary chan-nels.

Facies 3.7: Low-Angle Sets of Coarse-Grained Sandstone.—Facies3.7 contains sets of coarse grained, unidirectional, low-angle trough cross-stratified and plane-parallel stratified sandstone with coal fragments (Fig.7). Units of this facies show a crude upward coarsening within large barforms that pass laterally and downslope into the steeply inclined foresetsof Facies 3.1–3.4. Channels of Facies 3.5 and 3.6 cut into Facies 3.7 inplaces.

Interpretation.—The unidirectional cross stratification and the low-an-

gle plane-parallel stratification reflect deposition by high-velocity tractioncurrents. The upward-coarsening character of the units and the associationwith distributary channels (Facies 3.6) and delta-front deposits (Facies 3.1–3.4) suggest that Facies 3.7 represents deltaic mouth bars.

Facies 3.8: Pebbly Sandstone Beds.—Facies 3.8 contains coarse-grained sandstone beds with distributional inverse or inverse-to-normalgrading. In places the coarsest parts of beds contain clay pebbles. Individualbeds are 0.9–2 m thick and have a sharp base and top. Clay clasts areoriented parallel to flow direction. Facies 3.8 has a restricted extent andoccurs at the most updip reaches of the steeply inclined Facies 3.1–3.4.

Interpretation.—The coarse grain size, inverse grading, and restrictedextent of beds, together with lack of traction structures or imbrication,suggest deposition from cohesionless grain flows (Bagnold 1954; Savage1979, 1983; Nemec 1990). The occurrence of this facies at the top of delta-

front units and the bar-form geometry suggest deposition in mouth bars.

Wave Reworking

Facies 3.9: Upward-Coarsening Units with Wave-Ripples.—Facies3.9 consists of upward-coarsening packages of wave-rippled and plane-parallel laminated very fine- to coarse-grained sandstone units (Fig. 7). Thesandstone units occur in packages up to 10 m thick. Facies 3.9 occurs inthe topographically highest reaches of delta-front sheets in the retrograda-tional series of clinothems (Fig. 3).

Interpretation.—The wave ripples reflect deposition above storm-wavebase. The upwards-coarsening sheet-like units originated from deposition




FIG. 7.—Photographs, representative vertical section, and interpretations of shelf-edge and uppermost-slope facies association.

in a delta-front or mouth-bar environment, in areas subject to frequent wave

reworking.


Shelf-edge deltas that prograded onto the upper slope represent the up-permost reaches of the studied shelf-edge to slope clinothems. The delta-front lobes (up to 10 m thick) consist of upward-thickening and -coarseningunits of thick sandy beds (Facies 3.1 and 3.2) and thin heterolithic beds(Facies 3.3). The heterolithic deposits form a fringe to the sand-prone delta-front lobes (Fig. 3), which show local slump collapse (Facies 3.4). Mouthbars with upwards-coarsening units of traction (Facies 3.7) and grain-flowdeposits (Facies 3.8) occur in the proximal reaches of the delta front and

are cut by distributary channels (Facies 3.5 and 3.6). In some places, the

distributary channels extend down the delta front to connect into middleand lower slope chutes (Fig. 3). In the retrogradational phase of develop-ment the deltas are wave reworked (Facies 3.9).

SYSTEMATIC SLOPE ACCRETION: INDICATION OF HYPERPYCNAL FLOW?

The studied slope accumulations were clearly driven by shelf-edge deltasystems. Most of the flows discharged from the delta mouth accumulatedon the deepwater slope. This is in contrast to Type 1 clinothems, wherecanyons at the shelf edge and on the upper slope funnel sediment beyondthe slope and onto the basin floor (see Steel et al. 2000). The distributary-channel and mouth-bar systems of the studied clinothems indicate that the




TABLE 3.—Shelf-edge and upper-slope deposits


3.1 ungraded sandstonebeds

medium- to coarse-grained sand-stone; coal frag-ments

ungraded beds, internaldiffuse and non-con-tinuous scours

sharp and flat bases sheet-like 0.2–1.5 m steady downslope- decelerat-ing turbidity flows in del-ta-front sheets and inmouth bars, long-livedconstant discharge

3.2 laminated and un-

graded sandstonebeds

fine- to coarse-

grained sandstone

alternation of ungraded,

plane- parallel or cur-rent-ripple laminatedintervals; internalscours

scoured bases sheet-like laminated: 0.25–2.0

m; ungraded;0.2–2.0 m; units2–10 m

quasi-steady downslope- de-

celerating turbidity flows,long-lived constant dis-charge, alternating falloutrates

3.3 very fine sandstoneand siltstone beds

very fine grainedsandstone to silt-stone/mudstone

ungraded to laminated,very fine sandstoneand planar- or currentripple-laminated mud-stone


0.01–0.10 m waning downslope-deceler-ating turbidity flows,surge-type, episodic flows

3.4 deformed sandstonebeds


folds, overturned folds,ball and dish struc-tures

sharp wedge-like 0.4–3.0 m water escape and slumping,lobe collapse, slope insta-bility related, episodic

3.5 channeled coarse-grained sandstonebeds

medium- to coarse-grained sand-stone; coal frag-ments

unidirectional troughcross-stratification,plane-parallel lamina-tion

erosional cut downa few meters

channel geometry; 25–50 m wide; fine andthin upwards anddownslope

u p t o 1 m t ra ct io n d ep os it io n in d is -tributary channels

3.6 small-scale chan-neled sandstonebeds

fine- to medium-grained sand-stone; shaleclasts

plane-parallel lamina-tion or trough-crossstratification

erosional cut downa few tens of centimeters

channel geometry 3–5m wide; beds fine up-wards and downslope

0 .4 –1 .0 m t ra ct io n d ep os it io n in m in ordistributary channels

3.7 low-angle sets of coarse-grainedsandstone

coarse-grainedsandstone; coalfragments

trough-cross and plane-parallel stratification

sharp sheet-like 0.25–1.0 m traction deposition in mouthbars

3.8 pebbly sandstonebeds

coarse-grainedsandstone; claypebbles

distributional inverse;inverse-to-normalgrading

sharp restricted extent 0.9–2.0 m cohesionless grain flows inmouth bars

3.9 upwards-coarseningwave-rippledunits

fine- to coarse-grained sandstone

wave rippled, plane-par-allel laminated

flat sheet-like units 10 m wave reworking in wave in-fluenced delta front

shelf-edge deltas were fluvially dominated during the progradational phasesof the slope-wedge development but were increasingly wave influencedduring the retrogradational phases (Fig. 8). The distributary channels fedthe mouth bars and the sediment-gravity-flow-dominated delta front. Thephysical connection of fluvial channels into the chutes, and lack of evidencefor major delta collapse and slumping suggest that the fluvial feeder system

was dumping sediment directly onto the slope (see also Kostaschuk andMcCann 1987, Bornhold and Prior 1990, Phillips and Smith 1992, Carlsonet al. 1992). This is confirmed by the very systematic and regular accretionof the individual minor lobes as well as the general accretion of the shelf-edge to basin-floor clinothems. The hydrodynamic interpretation of theslope deposits is also consistent with this notion. The chute fills indicatedeposition from quasi-steady downslope-accelerating turbidity currents, i.e.,flows that required rather sustained discharge as well as sediment flux,highly suggestive of hyperpycnal flow input (see Wright et al. 1990, Chikita1990, Zeng et al. 1991, Mulder et al. 1998, Wright et al. 1998, Piper et al.1999). Delta collapse and slumping would have to produce surge-type flowsthat waned and decelerated downcurrent. Even the sandy chute-mouth lobesindicate deposition from downcurrent-accelerating flows. This, however,can be explained by the relatively steep fronts of the lobes. The lower-slope to basin-floor channels that are connected upslope to chutes, as well

as the thick sheet-like sandstone beds, also indicate deposition from sus-tained flows, and again confirm the hyperpycnal input.

Deposits of waning flows are typically mud-prone and form a fine-grained fringe to the delta lobes, as well as chute-mouth lobes and lower-slope to basin-floor lobes. The waning-flow deposits separate individualshelf-edge to basin-floor progradational episodes, i.e., individual clino-forms.

SEA-LEVEL BEHAVIOR AND SYSTEMS TRACTS IN THE SLOPE WEDGES

Sea-level behavior has been reconstructed from the internal character of the Hogsnyta wedge, our main example of a wedge where shelf-edge reach-

es are preserved and now exposed. The progradation of shelf-edge deltasout across the shelf and onto the upper slope could have been caused eitherby a fall of sea level below the shelf edge or by an unusually high sedimentsupply. The geometric relationships between the progradational, aggrada-tional, and retrogradational series of clinoforms, and the occurrence of anunconformity at the top of the progradational series but below the shelf

edge, strongly suggest that deltas were forced across the shelf edge as aresult of falling sea level (Fig. 9.1). Within the progradational series, themarked basinward shift of the facies and the overall downward depositionaltrajectory of the clinothems, the presence of another unconformity, and thehigh density of slope chutes clearly document the continued fall of relativesea level. The two youngest clinothems of the progradational series markthe maximum progradation of the system. The surface of maximum pro-gradation is thus a terrace of subaerial erosion in the uppermost-slope andshelf-edge reaches but is a surface of subaqueous chute erosion on themiddle and lower slope. This surface is therefore the sequence boundary,and the underlying progradational series represent a forced regressive sys-tems tract (Fig. 9.2). The highstand systems tract, landward of the latter,has now been entirely eroded.

In the aggradational series of clinothems, the landward onlap of clino-

thems evidence an initial rise of relative sea level (Fig. 9.3). The aggra-dational series thus represents a poorly developed lowstand wedge systemstract. In the retrogradational series, the backstepping of clinothems and lackof slope channels and erosion surfaces indicate that sea level continued torise. The reappearance of backstepping deltas at the shelf edge suggestsincreased water depth above the shelf edge at this time (Fig. 9.4). Theretrogradational series is part of a transgressive systems tract. The progra-dational character of the youngest clinothem and the extensive channelingat its top indicate either a higher sediment supply or a lower rate of relativesea-level rise prior to (maximum) flooding and deposition of thick marineshales at the top of the Hogsnyta clinothem complex (Fig. 9.5).




FIG. 8.—Block diagram showing schematicrepresentation of the main facies associationsbetween shelf edge and lower slope for thestudied slope wedges. A) Regressive and B)retrogradational phases are shown.

DISCUSSION AND CONCLUSIONS

No Basin-Floor Fans

The Type 2 clinoforms show that sediment was trapped mainly on theslope and that there are no basin-floor fans related to these accumulations.This happens despite the documented (on Hogsnyta) sea-level fall belowthe shelf edge (Fig. 9). Apparently the sea-level fall forced the shelf-edgedeltas to prograde down onto the upper slope, where they became perchedand eventually beheaded. All the sand generated by erosion of the shelf,shelf edge, and uppermost slope was accommodated entirely on the slope,by progradation of middle-slope minor lobes and lower-slope extensivelobes fed by subaqueous chutes. This retention of sediment on the slope,in a high-sediment-supply system, appears to be characteristic of non-can-yoned slopes and is a good illustration of how such shelf margins growbasinwards. The lack of basin-floor fans, even where sea level has fallenbelow the shelf edge, warns against over-simple prediction of deep-watersands.

Hyperpycnal Flows

The hydrodynamic interpretations strongly suggest that the slope wedgeswere fed largely by hyperpycnal flows. The likelihood of slope accretionby hyperpycnal flows is high here because: (1) the waters in this tinyforeland basin, with great deltaic input, are likely to have been of lowsalinity, (2) the rising topography in the fold-and-thrust belt provided highinput of coarse material, (3) river-dominated deltas reached out to the shelf edge, and (4) the slope beyond the shelf edge was relatively steep (ca. 4degrees), allowing sediment-laden river floods to become hyperpycnalflows.

The hyperpycnal nature of the sediment output from the shelf edge mayprovide an explanation for the fact that most of the sediment was dumpedonto the slope instead of continuing to the basin floor. There are two prin-cipal external forces acting on turbidity currents. Gravity propels sediment-laden fluids downslope, and friction between the base of the flow and thesubstrate resists the downslope movement. For a turbidity current to stopwe need to decrease the former or increase the latter.




FIG. 9.—History of the progradational,aggradational, and retrogradational phases of development for the Hogsnyta slope wedge, inrelation to falling and rising relative sea level.




FIG. 10.—The general distribution of sand-prone, heterolithic, and shale-prone lithologies within the six clinothem complexes (slope wedges) studied.

As explained above, the chutes are small, probably too small to confinethe turbidity currents, i.e., the turbidity flows overbank the chutes and dis-perse. Dispersal increases the basal area of the current, thus also increasingfriction and causing the turbidity currents to decelerate. When this happens,chute-related lobes develop from turbidity currents that start to lose theirdensity, wane, and eventually die out. Moreover, in hyperpycnal flows thesediment is mixed with fairly fresh water which has a lower density thanthe marine water. This can cause flow stratification and splitting of the flowinto a dense underflow and a buoyant plume (McLeod et al. 1999). Thefrontal speed of underflows with buoyant interstitial fluid decreases morerapidly than that of currents in which the interstitial fluid has the samedensity as the ambient fluid. These underflows ultimately come to a haltas sedimentation results in a reversal of buoyancy (the bulk density dropsbelow that of the ambient fluid), and the current lifts off. The lifted fluidcarries fine sediment with it (see also Sparks et al. 1993, Carey et al. 1996,Kneller and Buckee 2000).

We also know, however, that some turbidity currents are able to erodesignificantly into the substrate, creating submarine channels and canyons.In fact, sandy (i.e., noncohesive) turbidity currents must have the ability toentrain sediment from the vicinity of the bed, because turbulence alone isnot sufficient to keep these noncohesive materials in suspension (Parker1982). The fact that the turbidity currents of the present study were neverable to erode major channels or canyons implies that they never acceleratedto significantly entrain sediment or reached the ‘‘ignited’’ stable state, inwhich turbidity currents are self-sustained, highly erosive, and competent

to scour out submarine canyons (Parker 1982). Turbidity flows below ig-nitive state are unstable and invariably die out.

The latter idea is further supported by the flow behavior through a rel-ative-sea-level cycle. The chutes were largest and most frequent duringrelative sea-level fall, whereas they were almost nonexistent during theinterval of fastest relative sea-level rise, when the feeding shelf-edge systembackstepped landwards and even less sediment reached the slope.

The dilute nature of turbidity flows may also have been caused by theirsupercritical behavior. Supercritical flows entrain ambient water fromabove and thus need to entrain even more sediment from the substrate inorder to keep or increase their density and ‘‘ignite’’ (Parker 1982).

Water-Depth Estimates from Clinoform Dimensions

The dimensions of shelf-edge to slope clinoforms normally allow anestimate of water depth for any point down the length of the clinoform, if a value for water depth at the shelf edge is added. The amplitude of theclinoform thus provides an estimate of water depth for the base-of-slopearea, uncorrected for compaction. The water depth calculated in this waycan be a serious overestimate if the shelf-edge to slope clinoforms areactually clinoform complexes, as in the present study. In the present casethe general amplitude of the clinoform complex is an overestimate of thereal water depth at the time of sand deposition on the slope, because of the sea-level fall below the edge of the preexisting shelf. Thus, the ampli-tude of the clinoform complex from shelf edge to basin floor gives a mea-




sure of water depth prior to slope deposition, not for the time of slopedeposition.

On Hogsnyta, the amplitude (and presence of subaerial terraces) of theyoungest clinoforms developed during forced regression suggests that thewater depth at the base of the slope was less than 100 m. In contrast, theamplitude of individual clinoforms in the transgressive systems tract (where40 m is added for the water depth at the shelf edge) suggests that themaximum water depth at the base of slope was nearly 200 m.

Channel Types

The channels vary in their character in the different facies associations.(1) Channels on the shelf edge and upper slope are filled with thin suc-cessions of coarse-grained sandstones, show more traction cross-strata, andtypically do not show an abandonment phase (a fine-grained capping). (2)Chutes on the middle and lower slope are larger, fairly steep-sided, areassociated with sandy and heterolithic lobes, and have typically upwards-fining and -thinning character. (3) The channels on the lower slope to basinfloor are relatively broad. The beds infilling the lower-slope channels havethe same character as the sheet-like turbidites outside the channels. Theyare thick ungraded or laminated beds that show only minor basal erosion,suggesting a constructional type of channel, i.e., the channels were filledwith the same sheet-like flows that diverged on the lower slope.

Sand/Shale Ratio

The turbidite wedges are low-angle (up to 4 degrees) sand-prone cli-nothem complexes that contrast with the more common shaly clinothemsthat dominate the basin. The wedges are isolated from each other by shales,and occur only occasionally, because times of sea-level fall below the shelf edge were rare. The individual clinothems that occur within a clinothemcomplex are also sand-prone, but there are also mud-prone deposits fringingand separating the clinothems. These mud-prone segments tend to thinupslope, causing the amalgamation and sand-prone nature of the clinoformcomplexes near the shelf edge (Fig. 10).

ACKNOWLEDGMENTS

Amoco, Conoco, Exxon, Norsk Hydro, Mobil, Phillips, Shell, Statoil, and UPRCfinanced the Wyoming Consortium on Linkage of Facies Tracts (WOLF). JSR re-viewers Carlos Pirmez and Lincoln Pratson improved the manuscript. We thank alsothank Ben Kneller, Lars-Magnus Falt, Richard Moiola, and G. Shanmugam for re-viewing earlier versions of the manuscript.

REFERENCES

ARNOTT, R.W.C., AND HAND, B.M., 1989, Bedforms, primary structures, and grain fabric in thepresence of suspended sediment rain: Journal of Sedimentary Petrology, v. 59, p. 1062–1069.

BAGNOLD, R.A., 1954, Experiments on a gravity-free dispersion of large solid spheres in aNewtonian fluid under shear: London, Royal Society, Philosophical Transactions, v. A 225,p. 49–63.

BLYTHE, A.E., AND KLEINSPEHN, K.L., 1998, Tectonically versus climatically driven Cenozoicexhumation of the Eurasian Plate margin, Svalbard; fission track analyses: Tectonics, v. 17,p. 621–639.

BOUMA, A.H., 1962, Sedimentology of Some Flysch Deposits: A Graphic Approach To Facies

Interpretation: Amsterdam, Elsevier, 168 p.BRANNEY, M.J., KNELLER, B.C., AND KOKELAAR, B.P., 1990, Disordered turbidite facies—a prod-uct of continuous surging density currents: (abstract) International Sedimentological Con-gress, Nottingham, Abstracts, p. 38.

BRANNEY, M.J., AND KOKELAAR, B.P., 1992, A reappraisal of ignimbrite emplacement: progres-sive aggradation and changes from particulate to non-particulate flow during emplacementof high-grade ignimbrite: Bulletin of Volcanology, v. 54, p. 504–520.

CAREY, S., SIGURDSSON, H., MANDEVILLE, C., AND BRONTO, S., 1996, Pyroclastic flows and surgesover water: an example from the 1883 Krakatau eruption: Bulletin of Volcanology, v. 57,p. 493–511.

CARLSON, P.R., POWELL, R.D., AND PHILLIPS, A.C., 1992, Submarine sedimentary features on afjord delta front, Queen inlet, Glacier Bay, Alaska: Canadian Journal of Earth Sciences, v.29, p. 565–573.

CHIKITA, K., 1990, Sedimentation by river-induced turbidity currents: field measurements andinterpretation: Sedimentology, v. 37, p. 891–905.

EDWARDS, D.A., LEEDER, M.R., BEST, J.L., AND PANTIN H.M., 1994, On experimental reflecteddensity currents and interpretation of certain turbidites: Sedimentology, v. 41, p. 347–461.

HARLAND, W.B., 1969, Contribution of Spitsbergen to understanding of the tectonic evolutionof the North Atlantic Region, in Kay, M., ed., North Atlantic Geology and Continental Drift:American Association of Petroleum Geologists, Memoir 12, p. 817–851.

HELLAND-HANSEN, W., 1990, Sedimentation in a Paleocene foreland basin, Spitsbergen: Amer-ican Association of Petroleum Geologists, Bulletin, v. 74, p. 260–272.

HELLAND-HANSEN, W., 1992, Geometry and facies of Tertiary clinothems, Spitsbergen: Sedi-mentology, v. 39, p. 1013–1029.

HISCOTT, R.N., 1994, Loss of capacity, not competence, as the fundamental process governingdeposition from turbidity currents: Journal of Sedimentary Research, v. A64, p.209–214.

HISCOTT, R.N., AND MIDDLETON, G.V., 1979, Depositional mechanics of thick-bedded sandstonesat the base of a submarine slope, Tourelle Formation, Quebec, Canada, in Doyle, L.J. andPilkey, O.H., Jr., eds., Geology of Continental Slopes, SEPM, Special Publication 27, p.307–326.

HISCOTT, R.N., PICKERING, K.T., BOUMA, A.H., HAND, B.M., KNELLER, B.C., POSTMA, G., AND

SOH, W., 1997, Basin-floor fans in the North Sea: sequence stratigraphic models vs. sedi-mentary Facies: Discussion: American Association of Petroleum Geologists, Bulletin, v. 81,p. 662–665.

JOHANSSON, M., AND STOW, D.A., 1995, A classification scheme for shale clasts in deep watersandstones, in Hartley, A.J., and Prosser, D.J., eds., Characterization of Deep Marine ClasticSystems: Geological Society of London, Special Publication 94, p. 221–241.

KELLOGG, H.E., 1975, Tertiary stratigraphy and tectonism in Svalbard and continental drift:American Association of Petroleum Geologists, Bulletin, v. 59, p. 465–485.

KNELLER, B. C., 1995, Beyond the turbidite paradigm: physical models for deposition of tur-bidites and their implications for reservoir prediction, in Hartley, A.J. and Prosser, D.J.,eds., Characterization of Deep Marine Clastic Systems: Geological Society of London, Spe-cial Publication 94, p. 31–49.

KNELLER, B.C., AND B RANNEY, M.J., 1995, Sustained high-density turbidity currents and thedeposition of thick ungraded sands: Sedimentology, v. 42, p. 607–616.

KNELLER, B., AND BUCKEE, C., 2000, The structure and fluid mechanics of turbidity currents: areview of some recent studies and their geological implications: Sedimentology, v. 47, p.62–94.

KNELLER, B.C., EDWARDS, E., MCCAFFREY, W., AND M OORE, R., 1991, Oblique reflection of turbidity currents: Geology, v. 19, p. 250–252.

KOSTASCHUK, R.A., AND MCCANN, S.B., 1987, Subaqueous morphology and slope processes ina fjord delta: Bella Coola, British Columbia: Canadian Journal of Earth Sciences, v. 24, p.52–59.

LOWE, D.R., 1982, Sediment gravity flows: II. Depositional models with special reference tothe deposits of high-density turbidity currents: Journal of Sedimentary Petrology, v. 52, p.279–297.

LOWE, D.R., 1988, Suspended-load fallout rate as an independent variable in the analysis of current structures: Sedimentology, v. 35, p. 765–776.

MCLEOD, P., CAREY, S., AND SPARKS, R.S.J., 1999, Behaviour of particle-laden flows into theocean: experimental simulation and geological implications: Sedimentology, v. 46, p. 523–537.

MELLERE, D., GRANBERG, E., JOHANSEN, S., AND OLSEN, T., 1997, Facies architecture and stratig-raphy of a deltaic-detached turbidite system in the Eocene of Svalbard (Norwegian Arctic):(abstract) American Association of Petroleum Geologists, 1997 Annual Convention, Dallas,Abstracts, p. A81.

MIDDLETON, G.V., 1993, Sediment deposition from turbidity currents: Annual Review of Earthand Planetary Sciences, v. 21, p. 89–114.

MULDER, T., SYVISTSKI, J.P.M., AND SKENE, K.I., 1998, Modeling of erosion and deposition byturbidity currents generated at river mouths: Journal of Sedimentary Research, v. 68, p. 124–137.

NEMEC, W., 1990, Aspects of sediment movement on steep delta slopes, in Colella, A., andPrior, D.B., eds., Coarse-Grained Deltas: International Association of Sedimentologists, Spe-cial Publication 10, p. 29–74.

NICHOLS, R.J., 1996, The liquefaction and remobilisation of sandy sediments, in Hartley, A.J.and Prosser, D.J., eds., Characterization of Deep Marine Clastic Systems: GeologicalSocietyof London, Special Publication 94, p. 63–76.

NYBERG, K., SOLHEIM, B., HELLAND-HANSEN, W., AND NOTTVEDT, A., 1995, Linked coastal andbasin-floor fan systems, Battfjellet Formation, Spitsbergen: punctuated progradations duringoverall sea level rise: (abstract) Norwegian Petroleum Society, Stavanger 1995, Abstracts.

PARKER, G., 1982, Conditions for the ignition of catastrophically erosive turbidity currents:Marine Geology, v. 46, p. 307–327.

PHILLIPS, A.C., AND S MITH, N.D., 1992, Delta slope processes and turbidity currents in pro-

deltaic submarine channels, Queen Inlet, Glacier Bay, Alaska: Canadian Journal of EarthSciences, v. 29, p. 93–101.PICKERING, K.T., CLARK, J.D., RICCI LUCHI, F., SMITH, R.D.A., HISCOTT, R.N. AND KENYON, N.H.,

1995, Architectural element analysis of turbidite systems, and selected topical problems forsand-prone deep-water systems, in Pickering, K.T., Hiscott, R.N., Kenyon, N.H., Ricci Luc-chi, F. and Smith, R.D.A., eds., Atlas of Deep Water Environments: Architectural Style inTurbidite Systems: London, Chapman & Hall, p. 1–10.

PIPER, D.J.W., HISCOTT, R.N., AND NORMARK, W.R., 1999, Outcrop scale acoustic facies analysisand latest Quaternary development of Hueneme and Dume submarine fans, offshore Cali-fornia: Sedimentology, v. 46, p. 47–78.

PLINK, P, MELLERE, D., AND STEEL, R. 1997, Diversity of turbidites in basin floor lobes andchannel/levee systems in deepwater wedges, Central Tertiary Basin, Spitsbergen: Abstract,American Association of Petroleum Geologists, 1997 Annual Convention, Dallas, p. A93.

POSTMA, G., NEMEC, W., AND K LEINSPEHN, K.L., 1988, Large floating clasts in turbidites: amechanism for their emplacement: Sedimentary Geology, v. 58, p. 47–61.




PRIOR, D.B. AND BORNHOLD, B.D., 1990, The underwater development of Holocene deltas, inColella A. and Prior, D.B., eds., Coarse-Grained Deltas, International Association of Sedi-mentologists, 10, p. 75–90.

PRIOR, D.B., WISEMAN W.J. AND BRYANT W.R., 1981, Submarine chutes on the slopes of fjorddeltas: Nature, v. 290, p. 326–328.

RICH, J.L., 1951, Three critical environments of deposition and criteria for recognition of rocksdeposited in each of them: Geological Society of America, Bulletin, v. 62, p. 1–20.

SAVAGE, S.B., 1979, Gravity flow of cohesionless granular materials in chutes and channels:Journal of Fluid Mechanics, v. 92, p. 53–96.

SAVAGE, S.B., 1983, Granular flows down rough inclines—review and extension, in Jenkins,J.T., and Satake, M., eds., Mechanics of Granular Materials: New Models and ConstitutiveRelations, Amsterdam, Elsevier, p. 261–282.

SHANMUGAM, G., 1996, High-density turbidity currents: are they sandy debris flows?: Journalof Sedimentary Research, v. 66, p. 2–10.

SHANMUGAM, G., BLOCH, R.B., MITCHELL, S.M., BEAMISH, G.W.J., HODGKINSON, R.J., DAMUTH,J.E., STRAUME, T., SYVERTSEN, S.E., AND SHIELDS, K.E., 1995, Basin-floor fans in the NorthSea: sequence stratigraphic models vs. sedimentary facies: American Association of Petro-leum Geologists, Bulletin, v. 79, p. 477–512.

SHANMUGAM, G., DAMUTH, J.E., AND M OIOLA, R.J., 1985, Is the turbidite Facies Associationscheme valid for interpreting ancient submarine fan environments?: Geology, v. 13, p. 234–237.

SPARKS, R.S.J., BONNECAZE, R.T., HUPPERT, H.E., LISTER, J.R., HALLWORTH, M.A., MADER, H.,AND PHILLIPS, J., 1993, Sediment-laden gravity currents with reversing buoyancy: Earth andPlanetary Science Letters, v. 114, p. 243–257.

STAUFFER, P.H., 1967, Grain-flow deposits and their implications, Santa Ynez Mountains, Cal-ifornia: Journal of Sedimentary Petrology, v. 37, p. 487–508.

STEEL, R.J., DALLAND, A., KALGRAFF, K., AND LARSEN, V., 1981, The Central Tertiary Basin of

Spitsbergen: sedimentary development of a sheared margin basin, in Kerr, J.W. and Fer-guson, A.J., eds., Geology of the North Atlantic Borderland: Canadian Society of PetroleumGeologists, Memoir 7, p. 647–664.

STEEL, R.J., GJELBERG, J., HELLAND-HANSEN, W., KLEINSPEHN, K., NOTTVEDT, A., AND LARSEN,M.R., 1985, The Tertiary strike-slip basins and orogenic belt of Spitsbergen, in : Biddle,K.T., and Christie-Blick, N., eds., Strike-Slip Deformation, Basin Formation, and Sedimen-tation: SEPM, Special Publication 37, p. 339–359.

STEEL, R.J., MELLERE, D., AND PLINK, P., 2000, Towards prediction of deepwater sands on slopeand basin floor: American Association of Petroleum Geologists, 2000 Annual Convention,New Orleans, Abstracts, p. A142.

STEEL, R., PLINK, P., AND MELLERE, D., 1997, Architecture, stratigraphic distribution and originof Eocene deepwater sandstones, Central Tertiary Basin, Spitsbergen: (abstract) AmericanAssociation of Petroleum Geologists, 1997 Annual Convention, Dallas, Abstracts, p. A111.

WRIGHT, L.D., WISEMAN, W.J., BORNHOLD, B.D., PRIOR, D.B., SUHAYDA, J.N., KELLER, G.H.,YANG, L.S., AND FAN, Y.B., 1998, Marine dispersal and deposition of Yellow River silts bygravity-driven underflows: Nature, v. 332, p. 629–632.

WRIGHT, L.D., WISEMAN, W.J., JR., YANG, Z.S., BORNHOLD, B.D., KNELLER, B.C., PRIOR, D.B.,AND SUHAYDA, J.N., 1990, Processes of marine dispersal and deposition of suspended siltsoff the modern mouth of the Huanghe (Yellow) River: Continental Shelf Research, v. 10,p. 1–40.

WALKER, R.G., 1967, Turbidite sedimentary structures and their relationship to proximal anddistal depositional environments: Journal of Sedimentary Petrology, v. 37, p. 25–43.

ZENG, J., LOWE, D.R., PRIOR, D.B., WISEMAN, W.J., JR., AND BORNHOLD, B.D., 1991, Flow prop-erties of turbidity currents in Bute Inlet, British Columbia: Sedimentology, v. 38, p. 975–996.

Received 6 October 1999; accepted 5 March 2001.

Turbidite Variability and Architecture of Sand-Prone, Deep-Water Slopes Eocene Clinoforms in the Central Basin, Spitsbergen - JSR, 2001.pdf

Documents