anatomy of shelf deltas at the edge a progradion

8/9/2019 Anatomy of Shelf Deltas at the Edge a Progradion

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Anatomy of shelf deltas at the edge of a progradingEocene shelf margin, Spitsbergen

D O N A TELLA M ELLERE* , PI RET PLI N K - BJ ORKLUND and RONALD STEEL

*Department of Geology, University of Padova, Via Giotto 1, 31137 Padova, Italy(E-mail: [email protected])Department of Geology & Geophysics, University of Wyoming, Laramie, 82071 WY, USA

ABSTRACT

Although shelf-edge deltas are well-imaged seismic features of Holocene andPleistocene shelf margins, documented outcrop analogues of these importantsand-prone reservoirs are rare. The facies and stratigraphic architecture of anoutcropping shelf-edge delta system in the Eocene Battfjellet Formation,Spitsbergen, is presented here, as well as the implications of this delta systemfor the generation of sand-prone, shelf-margin clinoforms. The shelf-edge

deltas of the Battfjellet Formation on Litledalsfjellet and Hgsnyta produced a35 15 km, shelf edge-attached, slope apron (70 m of sandstones proximally,tapering to zero on the lower slope). The slope apron consists of distributarychannel and mouth-bar deposits in its shelf-edge reaches, passing downslopeto slope channels/chutes that fed turbiditic lobes and spillover sheets. In thetransgressive phase of the slope apron, estuaries developed at the shelf edge,and these also produced minor lobes on the slope. The short-headedmountainous rivers that drained the adjacent orogenic belt and fed thenarrow shelf, and the shelf-edge position of the discharging deltas, made anappropriate setting for the generation of hyperpycnal turbidity currents on theslope of the shelf margin. The abundance of organic matter and of coalfragments in the slope turbidites is consistent with this notion. Evidence that

many of the slope turbidites were generated by sustained turbidity currentsthat waxed then waned includes the presence of scour surfaces and thickintervals of plane-parallel laminae within turbidite beds in the slope channels,and thick spillover lobes with repetitive alternations of massive and flat-laminated intervals. The examined shelf-edge to slope system, now preservedmainly below the shelf break and dominated by sediment gravity-flowdeposits, has a threefold stratigraphic architecture: a lower, progradationalpart, in which the clinoforms have a slight downward-directed trajectory; athin aggradational zone; and an upper part in which clinoforms backstep uponto the shelf edge. A greatly increased density of erosional channels andchutes marks the regressive-to-transgressive turnaround within the slopeapron, and this zone becomes an angular unconformity up near the shelf edge.

This unconformity, with both subaerial and subaqueous components, isinterpreted as a sequence boundary and developed by vigorous sand deliveryand bypass across the shelf edge during the time interval of falling relative sealevel. The studied shelf-margin clinoforms accreted mostly during falling stage(sea level below the shelf edge), but the outer shelf later became estuarine assea level became re-established above the shelf edge.

Keywords Clinoforms, hyperpycnal flow, shelf-edge deltas, slope lobes,spillover sheets, turbidite channels and chutes.

Sedimentology(2002) 49, 11811206

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INTRODUCTION

Shelf-edge deltas are now a well-imaged seismicfeature of Holocene and Pleistocene shelf margins(Suter & Berryhill, 1985; Berryhill et al., 1986;Matteucci & Hine, 1987; Tessonet al., 1990; Poaget al., 1990; Morton & Suter, 1996; Kolla et al.,

2000). Although there is an extensive literature onthe seismic character and overall geometry ofshelf-margin deltas (McMaster et al., 1970; Win-ker & Edwards, 1983; Tesson et al., 1990, 1993;Field & Trincardi, 1991; Sydow & Roberts, 1994;Kolla et al., 2000), relatively little is known oftheir internal architecture, with only a fewexceptions (e.g. Mayall et al., 1992; DiBona &von der Borch, 1993).

Shelf-edge deltas are typically developed wherethere has been a significant fall in sea level and ahigh supply of sediment. In this setting, deltaic

shorelines can move far out across the shelf, ofteninto a position at or near the break (Colemanet al., 1983; Suter & Berryhill, 1985; Coleman &Roberts, 1988). Although most often associatedwith relative fall in sea level, shelf-edge deltasappear nevertheless to be developed in a range ofrelative sea-level conditions: (1) during falling sealevel (e.g. Trincardi & Field, 1990; Tesson et al.,1993; Sydow & Roberts, 1994; Morton & Suter,1996), in which case the deltas fall into thecategory of forced-regressive or falling-stage pro-grading wedges (Hunt & Tucker, 1992; Plint &Nummedal, 2000); (2) during minimum and early

rise in sea level (e.g. Posamentier et al., 1988;Hart & Long, 1996), also known sequence strati-graphically as lowstand wedges; or (3) duringlater sea-level rise (e.g. Kolla & Perlmutter, 1993),forming part of a transgressive systems tract.

The above type of delta is identified in the geo-logical record in two different ways. In Holoceneand Pleistocene strata, the main criterion is thelocation of the delta with respect to the site of thepresent shoreline and the shelf margin; in thesecases, they are variously termed shelf-perched orshelf-margin deltas (Suter & Berryhill, 1985) or

shelf-edge deltas (Sydow & Roberts, 1994). In oldersuccessions, shelf-margin deltas are difficult toidentify with certainty, because of the difficulty ofpinpointing the position of the shelf-slope break.

An important condition likely to prevent thedevelopment and preservation of shelf-edge del-tas is where sea level falls below the shelf edge fora prolonged period, and canyons on the upperslope become coupled directly with incisedfluvial systems on the shelf (Kolla & Perlmutter,1993). This condition is overcome most easily

when sea level does not fall below the shelf break,or if the fall below the shelf break is short-lived.

In the present study, the facies and architectureof well-exposed, Early Eocene shelf-edge deltasassociated with a shelf break and non-canyonedslope in the Central Basin of Spitsbergen are docu-mented. These deltas were forced slightly beyond

the shelf edge and are now perched on the upper-most slope. Their position indicates a relativesea-level fall of 50100 m relative to the formerhighstand shoreline. A subsequent rise in sea levelcaused the deltas to draw back up onto the shelf.Thefall and rise cycle resulted in thedelivery of nosignificant volumes of sand beyond the toe ofslope, even though sea level fell below the shelfedge. However, the basinward-thinning wedge ofsand (more than 70 m at the shelf edge) has signi-ficant strike extent (more than 15 km). It is arguedbelow that this sand wedge resulted from the lack

of significant incision on the upper slope and fromsediment that was dispersed broadly from an apronof deltas rather than from a point source.

GEOLOGICAL SETTING

The Late PaleoceneEocene Central Basin ofSpitsbergen was a relatively small foreland basinthat formed in front of the developing WestSpitsbergen fold-and-thrust belt (Harland, 1969,1995; Lowell, 1972; Kellog, 1975; Spencer et al.,1984; Steel et al., 1985; Myhre & Eldholm, 1988;

Fig. 1A). Basin infill progressed from west to east,leaving a spectacular record of large-scale (150350 m amplitude) clinoforms that reflect theoverall progradation of the shelf-edge-to-slopesystem (Steel et al., 1985; Helland-Hansen, 1990)(Fig. 2). Sand-prone clinoforms developed byprogradation of deltas across the pre-existing shelfplatform, delivery of sediment beyond the pre-existing shelf break and down onto the slope.Repeated sediment delivery of this type producedsignificant shelf-margin growth. Sediment wassupplied mainly from the growing orogenic belt

along western Spitsbergen, some 2530 km westof the outcrops in this study. The depositionalsystems are well-exposed in three dimensionsalong mountainsides dissected by glacial valleys.The stratigraphic section attains a thickness ofsome 1500 m, with clinothems (sensu Helland-Hansen, 1992) up to 350 m high. Helland-Hansen(1992) described the clinoform geometry, withcoastal plain topsets and turbidite-rich toesets.Although the importance of shelf-edge deltaswas not emphasized, Helland-Hansen (1992) did

1182 D. Mellere et al.

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recognize that the clinoforms were fluvially dri-ven. More recently, Steel et al. (2000) examinedclinoform trends and growth styles at basin scale.Most of the clinothems are shale prone (type 4clinothems of Steelet al., 2000) and can be recog-

nized on the mountainside only by the presence ofheterolithic units of thin-bedded (13 cm thick)sandstones and siltstones that stand out from theuniform slope shales. Others are sand-proneacross much of the storm- and wave-dominatedshelves, but evidently delivered little sand ontothe slope (type 3). Type 2 clinothems were gener-ated by shelf deltas and are sand-prone along bothshelf and slope segments but with insignificantsand volumes on the time-equivalent basin floor.Type 1 clinothems produce basin-floor fans. The

sedimentary wedges of Litledalsfjellet and Hgs-nyta are generated by type 2 clinothem complexes.They form a thick slope wedge that pinches out bydownlap before the base of slope, but attain amaximum thickness of 70 m just below the shelf

break. In contrast to type 1 clinoforms, no slopecanyons or major slope-collapse features are asso-ciated with the type 2 clinoforms (Fig. 3).

METHODS

The sandstone clinothems of Litledalsfjelletand Hgsnyta are part of the same stratigraphicinterval, referred to hereafter as the Reindalenclinothem complex (Fig. 4). The clinothems

Fig. 1. (A) Location and tectonic setting of the studied area. The Central Tertiary Basin is a foreland basin, loaded bythe thrust sheets of the West Spitsbergen Orogenic Belt. The basin is highly asymmetrical and is 90100 km wide.(B) Location of the Litledalsfjellet and Hgsnyta outcrops in Reindalen. The shelf-edge delta wedges are up to 70 mthick, but taper out at the base of slope, some 35 km away from the shelf margin.

Anatomy of a prograding Eocene shelf margin, Spitsbergen 1183



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produce a large basinward-dipping, sandstonewedge that developed on the slope below theshelf edge. The two mountainsides converge at anangle of about 20 and are separated by a5 km wide glacial valley, enabling a partial three-dimensional reconstruction of the Reindalendepositional system (Figs 1 and 4). A total of 24

measured sections, spaced 50500 m apart, havebeen measured on the two mountainsides (Figs 5and 7). Most of the outcrop allows physicalcorrelation. On the steepest cliffs (south-west-ward ends of Litledalsfjellet and Hgsnyta), bedswere correlated by tracing on enlarged photo-mosaics shot from a helicopter (Fig. 6).

A

B

Fig. 3. (A) Clinoform types, as recognized by Steel et al. (2000) within the Battfjellet Formation. (B) Block diagramand profile outlining the shelf-edge terminology used for the studied, sand-prone type 2 clinoforms.

Fig. 2. Block diagram with palaeo-geographic and stratigraphic over-view of the studied successionduring an interval when the low-stand shoreline was at the shelf

edge. The shelf-edge deltas of theBattfjellet Formation prograded ontothe Gilsonryggen shales and werefed by the coastal plain system ofthe Aspelintoppen Formation.




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The choice of datum for the correlation panelsis important. The photomosaics of Litledalsfjelletand Hgsnyta show that the coastal plain strata inthe upper half of the mountainsides are nearlyhorizontal, so that the easterly dip of the under-lying clinothems represents the apparent dip ofthe slope of the depositional system. The corre-lation panels in Figs 5 and 7 are therefore hung

from the lowest sandstone unit containing coastalplain deposits. Minimum water depth at the baseof the Reindalen wedges has been estimated fromthe height between the shelf-slope break down tothe clinothem toes (Fig. 4, graph insert). A cor-rection was then made for compaction and forwater depth at the shelf edge.

The average gradient of the slope segment ofthe clinoforms is 34. The deposits form a shelfedge-to-slope turbidite apron 35 km in downdipextent and 15 km in strike length. Fine- to

medium-grained sandstones are concentrated inthe most proximal reaches (at the shelf-edgepalaeoshoreline, in the uppermost part of theoutcrops) and extend into water depths estimatedat more than 200 m. The lower slope gradienteventually declines to 05and then flattens ontothe shaley basin floor beyond the sandstone wedge.

THE MORPHOLOGICAL/BATHYMETRICPROFILE

Although the shelf-edge and slope depositsdescribed here derive entirely from shelf-edgedeltas, it is misleading and incorrect to view theclinoforms as simply deltaic. The deltas builtacross a pre-existing platform created during anearlier phase of shelf-margin accretion at relativesea-level lowstand. On reaching the shelf edge,

A

B C

Fig. 4. (A and C) General overview of the Litledalsfjellet (A) and Hgsnyta (C) clinoform complexes with location ofthe measured sections (for details, see Figs 5 and 7). Note the flat-lying youngest clinoforms, chosen as datum for the

construction of the correlation panels. The mountains are 700 m high, and the mountain face is 35 km long. Thecoastal plain deposits of the Aspelintoppen Formation form the succession in the upper part of the mountains.(B) Diagram showing the slope angles of the clinoforms in Litledalsfjellet, measured from the most proximal to thedistal reaches. The studied slope wedge (clinoform 1) has a gradient of 34.




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Fig.5.

Litledalsfjelletcross-sectionalpanel(seelocationofthe

measuredsectioninFig.4)withp

ositionofthelogdetailsandphotographsusedinthe

descriptionofthefaciesassoc

iations.




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Fig.6.Photomosaicandinter

pretationofthemostproximalreachesoftheLitledalsfjelletshelf-edge

deltacomplex.




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the new generation of deltas prograded downonto a pre-existing slope. In addition, the deltafront that then draped across the slope did so byaccreting at a steeper angle (downlap) than theslope itself. Because of this and the now greatlyextended length of the slope apron, slope ratherthan delta front terminology is used here to

describe the setting. This view is consistent withthe extensive channelling/erosion and great vol-umes of turbidites that occur on the study slopecompared with normal shelf-delta slopes.

The bathymetric profile so constructed consistsof (1) a shelf with width up to 1520 km andgradient of less than 05; and (2) a slope withlength 35 km and gradient of 34. The gradientat the base of slope decreases gradually, and thebottomset (basin floor) of the clinoform becomessubparallel with the topset. The clinoform setcreated by the gradual migration of the shelf edge

reflects a (decompacted) water depth of some50 m at the shelf break and some 250300 m downto the basin floor during highstand of sea level.

FACIES ASSOCIATIONS

Shelf-edge delta deposits

The shelf-edge delta deposits are characterized bya high sand/shale ratio and consist typically ofcoarsening-upward mouth-bar and distributary

channel units. The shelf-edge units accrete atangles up to 10 onto a slope itself dipping atsome 34. This succession merges downdipdirectly into sediment gravity flow-dominatedlobes and channels of the slope system.

Proximal mouth-bar deposits (A1)

Description. Association A1 consists of: (1) coar-sening-upward sandstone units; (2) sigmoidalbars; and (3) scour-and-fill features. The associ-ation occurs in the most landward and updipreaches of the Litledalsfjellet and Hgsnyta wed-ges (profiles 1 and 13, respectively, in Figs 57).

The coarsening-upward sandstone units consistof clean, well-sorted sandstones, up to 3 m thick,forming stacked intervals up to 10 m thick. Thesandstone units are dominated by a variety ofcross-stratified and ripple-laminated bed sets,

022 m thick, separated by up to 015 cm thickshale interbeds. The shale interbeds commonlycontain a high amount of coal and plant debris,usually in patchy lenses or chaotically distributed.

The typical bedform of the shelf-edge mouthbars is a sigmoidal barform similar to the alluvialflood-generated sigmoidal bars of Mutti et al.(1996). In the Reindalen shelf-edge deltas, suchbars consist of repeated alternations of ungraded/graded to plane-parallel- and ripple-laminatedfine- to coarse-grained sandstone beds (Fig. 8A)

Fig. 7. Hgsnyta cross-sectional panel, modified from Plink-Bjorklund & Steel (2002) (see location of the measuredsection in Fig. 4).




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Fig.8.Overviewanddetailso

fshelf-edgedeltadeposits(location

inFig.5).(A)Distalmouth-barasso

ciation(FaciesAssociationA2).Mostofthedepositsare

thinsetsofripple-toplanar-laminatedsandstones.(BandC)Rive

r-fed,sigmoidalbarsattheshelf-edgedeltamouthonLitledalsfjellet(A

ssociationA1).Note

theslightlandwardback-tilt(s

lightleftwarddipofsomebedsinB)oftheheadofthebars.

(D)OverviewofriverdistributarychannelinH

gsnyta.Thedeposits

arethinlyplanarlaminated,s

uggestingdepositionfrom

flood-dom

inatedhyperconcentratedflows.




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arranged (in section parallel to palaeoflow) inhigh-angle (1520) accretionary sigmoidal sets(Fig. 8B and C). Sets are convex upward, 052 mthick and bounded by erosional surfaces thattruncate the underlying sets. Individual beds andaccretion sets are modified at their tops by wave-ripple laminations. The barforms are lenticular

(up to 2 m thick, some 610 m wide) and aretruncated by flat-lying erosional surfaces or bysmall scour-and-fill features in their upcurrentparts (Fig. 8B). The latter are 041 m deep,065 m wide, infilled by planar-laminated, occa-sionally by trough cross-stratified, medium- tofine-grained sandstones, with abundant rip-upclasts and coal fragments. In a downcurrentdirection, the sigmoidal lenses broaden, the ero-sional bounding surfaces tend to become con-formable, and the beds progressively flatten andare dominated by plane-parallel lamination.

Successive sigmoidal barforms stack downstreamand vertically to build up composite mouth-barbodies, often cut by capping channels.

Interpretation. The association described abovesuggests the presence of an inertia-dominatedriver mouth, controlled by periodic, high-velo-city, flood-dominated river discharge. Sedimentwas dispersed within a turbulent jet onto asteeply dipping slope, producing elongate,steep-fronted mouth bars. Rapid dumping ofbedload and coarse suspended material occurrednear the mouth along with slower accumulation

of suspended material further seaward under-neath the sediment plume (see also Scruton,1960). Both suspended load and bedload sedi-ments become progressively finer with increaseddistance from the distributary mouths. The barcomplex seems to have been built by amalgama-ted flood units characterized by an upward-increasing tabular geometry.

The coarsening-upward sandstone units weredeposited by traction currents, as indicated by thealternating cross-stratification and ripple-lamin-ation. The thin ungraded, climbing-ripple and

flat-laminated beds in the sigmoidal bars mayindicate deposition by frictional freezing of high-density turbidity flows (massive to normal-gradedbeds; Lowe, 1982), followed by combined trac-tional and fall-out processes. The occurrence ofcapping wave ripples indicates that the deposi-tional depth was above the storm wave base. Thesigmoidal bars possibly originated from hydraulicjump at the shelf break, with transformation ofthe hyperconcentrated flows to rapidly depos-ited sediment, because of flow expansion and

deceleration. Energy dissipation through intenseturbulence is thought to have been the maincause of the extensive scouring observed in theupstream terminations of sigmoidal bars. It isinferred here that high-gradient distributary chan-nels carried floods directly to the shelf edge.According to Muttiet al. (1996), the development

of the complete sigmoidal bars, as well as thethick and extensive individual mouth bars, isindicative of poorly efficient flows, i.e. flows thatwere unable to carry most of their sediment loadfurther downslope.

Distal mouth-bar deposits (A2)

Description. Distal mouth bars consist of sand-prone, heterolithic units of sandstones and shalesorganized into coarsening- and thickening-upward motifs, 0515 m thick (Fig. 8). They

develop down the slope for about 1 km from theproximal mouth-bar complexes and overlie themore shale-prone heterolithic slope units(Fig. 8A). Sandstone beds are 00505 m thick.Their basal contacts are soft-sediment deformed,slightly erosive and loaded into underlying shalein places. Sole marks are absent or rare. Mostsandstone beds are current rippled and plane-parallel laminated. Normally graded and inverselygraded sandstone beds also occur, capped byshales. When present, inverse grading usuallycharacterizes the first few centimetres near thebase of the sandstone beds and often presents

shear lamination.Sandstone beds show low to moderate lateral

persistence, pinching out downdip or being cutby slope channels and slump scars. Slump inter-vals, up to 2 m thick but normally less than 1 mthick, are usually seen at the base of the coarsen-ing- and thickening-upward units. Shale inter-beds are millimetres to a few centimetres thickand display current-ripple lamination. Sandstoneabundance varies between 60% and 75%. Alter-nations of slumped, muddy intervals and cleanplane-parallel-laminated sands, such as those

described by Mayall et al. (1992), are also seenbut are not a dominant feature of the shelf edgedescribed here.

In some places, the boundaries between distalmouth-bar units are expressed as large-scalehummocky features, causing a lenticular geom-etry. The hummocks are up to 2 m high, 5070 m in wavelength and convex upward. Theyhave been observed in the proximal part of theLitledalsfjellet wedge, in the lower half of theoutcrop (Fig. 6) and in Hgsnyta (Fig. 7). Around




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the edges of such lenticular bodies, scour-and-fillstructures are evident, and some of these werefilled by upslope accretion. Where there arevertically repeated units of this type, they showan offset or compensatory stacking pattern.

Interpretation. The sand-prone heterolithic units

represent the distal reaches of the mouthbarassociation, deposited by an array of tractional,mass-flow and slump processes, such as would beexpected where the mouth bars draped furtherdown onto the steep upper slope of the shelfmargin. The current-ripple and planar-lamin-ated, fine-grained sandstone beds as well as theungraded to normally graded beds overlain bysiltstones indicate deposition by turbidity cur-rents (Bouma, 1962; Middleton, 1967; Walker,1967; Middleton & Hampton, 1973; Lowe, 1982).Thinner sandstone beds (less than 05 m), with

some normal grading, probably resulted fromshort-lived surge-type flows. The inverse-gradedto ungraded sandstone beds, with basal shearlaminae and sharply capped by shale, may be aproduct of sandy debris flows (Shanmugam &Moiola, 1995; Shanmugam, 1996). Given thescarcity of a mud fraction within the sand beds,the development of viscous flows may have beeninhibited, favouring more frictional and inertia-dominated sandy debris flow.

The hummocky features recorded in placesprobably reflect lobe switching of the mouth bars,particularly those accumulated at an early stage of

the wedge when the depositional slopes weresteep.

Channel-fill deposits (A3)

Description. Channels cross-cut the mouth-barcomplexes. Individual channels are characterizedby erosional relief of a few decimetres to a fewmetres, and they cut the sigmoidal bars of theproximal-bar association (Figs 57). Most of thechannels visible on Litledalsfjellet are dip orien-ted. Thickness/width ratio is about 1:20. Chan-

nels develop lateral wings up to 15 m long. OnHgsnyta, channels are 25100 m wide, with upto 5 m of vertical relief (Fig. 8D). Channel-filldeposits typically show a blocky to fining-upwardvertical trend. They consist of erosional-based,poorly sorted, coarse- to medium-grained, mas-sive or plane-parallel-laminated sandstone beds570 cm thick. The beds infill broad, cross-cut-ting shallow scours (Fig. 8D), paved by abundantshale rip-up clasts, coal and wood fragments.Sandstone beds locally display crude low-angle

stratification or cut-and-fill-type cross-stratifica-tion. The late-stage deposits of channel fillsconsist of thinner sandstone beds (315 cmthick), which are massive to planar and ripplelaminated, and separated by shales. Localizedslumps, clay clasts and plant debris are common.

Interpretation. The close association with thesigmoidal bars, the lack of bioturbation, thetraction structures and the large amount of coaland wood debris suggest that these channels wereriver distributaries. The upward-fining packagesof massive to laminated and cross-stratified chan-nel fills suggest deposition from waning currentsinvolving traction. Considering the dimensionsand, particularly, the thickness of individualchannel fills, which average a few metres ofvertical relief, it seems likely that a network ofshallow streams operated on the delta system,

sporadically enhanced and deepened by floods.The presence of repetitive scour-and-fill unitswas probably produced during rising flood dis-charges in the feeder rivers.

Slope deposits

Deposition on the slope, below the shelf edge,was dominated by the delivery of river-fed massflow deposits that formed a broad slope apron,composed of slope lobes. Slope lobes were fedand dissected by channels or chutes filled withcoarse-grained sandstones and capped by later-

ally extensive sandstone sheets produced by long-lived, river-fed spillovers. Although the hetero-lithic lobes are the dominant association of theslope, smaller lobe systems were also recognizedat the mouth of the chutes and channels or asoverbank deposits from chutes. The four mainfacies associations recognized are slope lobes,channels and chutes, channel and chute-mouthlobes and spillover lobes. Although slumpedunits are moderately common on the slope, inthe present study, they tend to be smaller and lessabundant than those described by Mayall et al.

(1992) from the shelf-edge deltas of the USA GulfCoast. However, it is noted that slump deforma-tion is much more common on the slope in casesof shelf-margin deltas that become deeply incisedby their distributary channels (Steel et al., 2000).

Slope-lobe deposits (B1)

Description. Lobate heterolithic units, some500 m wide and up to 2 km long, dominate theslope segments of the clinothems. The lobes are




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directly linked back upslope to a mouth-barsystem and distributary channels (Figs 57). Theheterolithic units, up to 15 m thick, are coarsen-ing and thickening upward or fining and thinningupward and consist of thin-bedded sandstonesand shales. The units are commonly capped byspillover sandstone sheets and cut by slope

channels and chutes. The heterolithic units showa variety of thin, laterally extensive beds inbroadly convex-up packages. Sandstone/shalecontent varies between 50% and 85%, with thehighest value close to the mouth-bar system.Individual sandstone beds, of 315 cm averagethickness, have diffuse to distinct boundaries andmay pass laterally into thicker amalgamatedsandstone units. Internally, they are plane-paral-lel to current-ripple laminated (Fig. 9A). Struc-tureless ungraded to normal- and inverse-gradedbeds also occur. The latter show sheared laminae

at the bottom (Fig. 9C). Bioturbation is rare exceptfor some small Chondrites traces.

Interpretation. Thin graded and laminated bedsindicate that waning turbidity currents providedthe normal sedimentation for the heterolithicunits, although inverse grading in some bedsmay reflect deposition by sandy debris flows(Shanmugam & Moiola, 1995; Shanmugam, 1996).Turbidity flows emanated from nearby distribu-

tary-channel mouths as hyperpycnal flows. Thesandier turbidites probably reflect channel floodstages, as well as transportation of materiallocally derived from gravity failure of advancingmouth bars. Facies organization, the relativelyhigh sand/shale ratio seen even in the most distalreaches of the system and the general lack of tracefossils indicate that sedimentation rates werevery high.

The coarsening-upward (C-U) and fining-upward (F-U) motifs are the result of lobe progra-dation followed by lobe abandonment or lobe

shifting. In the middle part of the Litledalsfjelletslope, the C-U or F-U motifs are equally common,

Fig. 9. Details of slope-lobe deposits (Facies Association B1, location in Figs 5 and 6). The deposits show a variety ofsedimentary structures from plane-parallel and ripple lamination (A) to normal and inverse grading (B and C),suggesting deposition from waning turbidity flows and possibly from sandy debris flows. Note in (C) the inversegrading and shear laminae at the base of the bed.




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suggesting that the lobe radius remained constantover time. Where lobes prograded in addition togradual shifting, there the C-U motifs dominateover those that fine upwards.

Channel and chute deposits (B2)

Description. Channelled sandstone bodies arepresent on the upper to middle slope reaches ofthe Litledalsfjellet and Hgsnyta clinoform com-plexes (Figs 5 and 7). The bases of such bodies aresharp and erosive into underlying heterolithicstrata (Fig. 10A), and often paved by shale rip-upclasts. The erosional channelled features can be upto 5 m deep, 50200 m wide when seen in a flow-transverse direction and extend downslope formore than 3 km. The channels are wider in theupper slope reaches, where they connect to thedistributary systems, becoming narrower downs-

lope. The channels pass downslope into longitu-dinally, slightly convex-upward sandstone sheets,up to 3 m thick, that may extend for up to 1 km tothe base of the slope (Fig. 6). Locally, bowl-shapedscars were identified at the heads of some chan-nels, cutting the downslope lee-side of mouth bars.

Channel infill consists of poorly sorted, coarse-to medium-grained, massive to planar-laminated,sharp-based sandstone beds infilling scours 033 m deep (Fig. 10). Shale rip-up clasts and largeamounts of organic debris (coal and plant) areparticularly abundant in the lower part of thechannel fill, paving stacked broad and shallow

troughs (Fig. 10A). The channelled units can beoverlain by fining- and thinning-upward hetero-lithic strata comprising ungraded to planar-lam-inated sandstones and siltstones, or by a series ofstacked thinning-upward sandstone beds.

Well-sorted sandstones, massive to pervasivelydeformed by water-escape structures, character-ize the infill of the bowl-shaped channel heads.

Downslope, the channel infills become finergrained and thinner and tend to lose both theupper laminated and the basal ungraded interval.Although some of the channels can be followed

throughout the longitudinal profile of the slope(Fig. 6), others show clear evidence of sinuosityand are found cutting the same spillover sheetunit at different localities.

Interpretation. The dimensions, geometry andposition of the channels within the slope segmentof the clinoforms, and the direct link to the shelf-edge delta system, suggest that most of the chan-nelled units represent the slope continuation of theshelf-edge delta distributaries. The channels were

supplied directly from the river system, as sugges-ted by the coarseness of the deposits, their thick-ness and the large volumes of organic debris andcoal fragments.

The thin, normally graded beds with shalerip-up clasts in upslope reaches indicate depos-ition from waning turbidity currents. The thick

(053 m) ungraded, plane-parallel or ripple-lam-inated beds reflect deposition from sustainedturbidity flows (Kneller, 1995; Kneller & Branney,1995). The thinning- and fining-upward unitswith silty tops possibly denote abandonment ofthe channels, whereas the thinning upwardswithout any significant fining may reflect a moreerosionaldepositional phase (Normark, 1970;Mutti & Normark, 1987; Clark & Pickering, 1996).

Bowl-shaped scars at some of the channelheads are interpreted as slump scars. Associatedchannelled units represent chutes, the deep,

narrow segments of a slump-driven gully system(Prioret al., 1981; Boumaet al., 1991). Chutes arecommon on muddy slopes beyond shelf edgesand are easily recognizable in modern systemsinvestigated with high-resolution seismic reflec-tion methods (Postma et al., 1988; Bornhold &Prior, 1990; Prior & Bornhold, 1990; Soh et al.,1995). On seismic and acoustic profile data,chutes are straight, can be traced from the shelfedge onto the upper and middle slope and maymerge into bulges and lobes on the lower slope.The deposits observed at the downslope end of agully system are pervasively convoluted and very

well sorted, suggesting that the gullies originatedby collapse of unstable, well-sorted delta-frontdeposits that accumulated rapidly at and beyondthe shelf edge, thus provoking abnormal sedi-mentary loading on the upper slope.

Channel/chute-mouth lobe deposits (B3)

Description. This association consists of hetero-lithic units of sandstone and shale forming coar-sening- and convex-upward, wedge-shapedbodies at the mouth of slope channels and chutes

or, in places, diverging from chute channels. Theheterolithic units are commonly up to 3 m thick,but may stack vertically forming bodies up to10 m thick, usually overlain by channels/chutesor spillover lobes. Normally, they are located inthe middle to distal reaches of the slope and canbe followed as progradational packages downs-lope for hundreds of metres (Fig. 5). Successiveunits typically show pinch-out in opposingdirections, probably a compensation feature, withinternal discordances and small-scale basal




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Fig.10.

Detailsofslopedepo

sits(FaciesAssociationB,

location

inFig.5).(A

andB)Slopechannels(AssociationB2)cuttingunderly

ingspilloversheets

(AssociationB4).Notetheim

bricateshalerip-upclastsconcentratedatthebaseofthechannel.Th

espilloversheet-likebeds(C)form

longtabularunits,

wedgingslightlydownslopeandconsistingofthinlylaminatedsan

dstones.Transportedwoodandcoalfragmentsarecommonandareoftenconcentratedatthe

baseofthebeds.




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erosion surfaces (Fig. 11). The thickness of suchwedges is inversely proportional to their lateralextent and decreases progressively downslope.Some coarsening-upward units are followed ver-tically by thin-bedded sandstones and shales.

Although developed on the mud slope, thefacies shows a relatively high sandstone content

(6070%). Sandstone beds, 575 cm thick, arefine grained, massive to low-angle planar lamin-ated and draped by shales (Fig. 11). Current-ripple lamination is subordinate and present onlyin the finest grained heterolithic strata at the baseor top of the coarsening-upward units. Soft-sediment deformation is common, particularlyin the thickest beds. Thicker beds are erosionallybased and composite. Most split into laterally

wedging thinner beds, with a slight but progres-sive increase in shale interbeds. Shale interbeds,up to 20 cm thick, consist of alternations oflaminated mudstone and normally graded tocurrent-ripple-laminated siltstones.

Interpretation. The pinch-out style of the con-

vex-upward bodies, the association with chutesand channels and the relatively high sandstonecontent despite the relatively distal position onthe slope favour the interpretation of the facies aschute and channel-mouth lobes. There are threefeatures consistent with this interpretation: (1)chutes and channels tend to cut the coarsening-upward lobe packages; (2) the downslopedecrease in the height of the convex-upward

Fig. 11. Details of channel/chute-mouth lobes (Facies Association B3) from the middle slope of Litledalsfjellet(location in Fig. 5). Note the gentle wedging of the deposits in apparently opposite directions and the coarsening- andthickening-upward motifs.




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units is associated with the downslope decreasein depth of the channels/chutes and thickness ofthe fills as seen in facies B2; (3) the grain sizeof the deposits is comparable with the grain sizeof the channel/chute fills. The alternation ofmassive and low-angle, planar-laminated sand-stone beds indicates deposition by turbidity

flows. The normally graded siltstone to laminatedmudstone beds reflect deposition from the tailends of waning turbidity currents.

Spillover deposits (B4)

Description. Sandstone sheets that extenddownslope for more than 15 km have beenrecognized in the lower half of the Reindalenwedges (Fig. 6). They consist of plane-parallel-laminated, normally graded to ungraded, fine- tomedium-grained sandstone bed sets up to 2 m

thick. Bed set bases are sharp and slightly erosive,and some show concentrations of layered andimbricate shale rip-up clasts (Fig. 10C). Individ-ual beds are up to 8 cm thick (average 35 cmthick), commonly massive to normal graded,planar laminated, locally current-ripple lamin-ated. The thickest bed sets are amalgamated andshow the highest lateral persistence. Downslope,some split into distinctive sheet units separatedby shale or concentrations of coal and woodfragments. The sheets overlie 15 to 5 m thickcoarsening- and thickening-upward slope-lobeunits (Fig. 10C) and are cut by slope channels

and chutes. Sandstone lobes wedge out bothbasinwards and upslope, where they connect tothe distributary channels and mouth-bar systems(see Fig. 7 between profiles 1 and 2). Theydevelop at abrupt changes in channel directionaccompanied by sediment failure of the channelmargin and/or scour formation.

Interpretation. The thick beds with alternationof ungraded and plane-parallel-laminated inter-vals suggest deposition from sustained turbiditycurrents (Kneller, 1995; Kneller & Branney, 1995).

The connection with the mouth-bar systems andthe presence of abundant coal and wood debristogether with the interpreted sustained nature ofthe flows indicate a direct link to the shelf-edgedelta distributary system. The sheet-like depositsemerge from the main distributaries at abruptchanges in channel direction and are interpretedto represent spillover lobes (Normark & Piper,1991). The basal erosional surface is likely to havebeen cut during rising flood stage, whereas theconsiderable thickness and internal variability of

the unit reflect the sustained nature of theformative flow. The sandiness, thickness andlateral persistence of these deposits suggest thata significant percentage of flows passed throughthe spillover points, causing the remaining flowwithin the slope channels to decelerate rapidly.

Shelf-edge estuary deposits

The upper part of the Reindalen clinoformcomplex is characterized by a succession of slopelobes (association B1) that lead updip to tidal barsand tidal channels developed at the mouth of anestuary. At both localities (Litledalsfjellet andHgsnyta), the clinoform sets in this upper part ofthe wedge are stacked retrogradationally (Fig. 5).The slope lobes of the retrogradational clinoformset are analogous, although thinner bedded andmuddier, to the slope-lobe deposits and the shelf-

edge delta deposits of the progradational clino-form sets described above. However, the numberand thickness of channels/chutes is much less.

Tide-influenced channel deposits (C1)

Description. A belt of channel and bar com-plexes marks the uppermost and most landwardreaches of the wedge on Litledalsfjellet (Figs 5and 6). The channels are bounded by a surface oferosion that can be followed downslope for15 km. Channel margins are step-like and oftendraped by sand-prone, up to 1 m thick slump

deposits (Fig. 12A). Lateral wings, a few tens ofmetres wide, develop in the uppermost channellevels. Channel fills are 38 m thick, and succes-sive units stack vertically to form complexes up to15 m thick. Individual channel fills are thinningand fining upwards and consist of thin-bedded(38 cm thick), structureless to plane-parallel andripple-laminated, medium-grained sandstones(Fig. 12), often draped by siltstone and claystonedrapes. Thicker, trough cross-stratified sand-stones are present mostly at the base, but mayalso occur through the channel unit. The upper-

most parts of the channels typically show cou-plets of plane-parallel- to wave-ripple-laminatedbeds (Fig. 12B), often alternating with shale anddouble shale layers. Palaeocurrent directions aremostly unimodal and oriented landwards.Bioturbation is present with only small forms ofPlanolitesat the base of the sandstone beds.

Interpretation. Channel-fill deposits are inter-preted in terms of normal to hyperconcentratedflood currents that were actively reworked by




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flood tides, as suggested by palaeocurrent direc-tion, the presence of shale and double shaleinterlayers and the association with tidal bardeposits. The step-like margins of the channelssuggest multiple phases of erosion and infill.

Tidal bar deposits (C2)

Description. Planar to trough cross-bedded sand-stones form bars up to 2 m thick that extend up to

30 m upslope with gentle dip angles (1015).The bars overlie and border the flanks of the tidalchannels. Cross-beds are 2030 cm thick, separ-ated by reactivation surfaces (Fig. 13) and trainsof shale rip-up clasts (Fig. 13C). Internal sets aresigmoidal, 38 cm thick and separated by doublemud drapes.

The bars are composite, with reverse tabularcross-beds and climbing ripples, particularlyalong the bottomsets. Change in dune polaritiesand reactivation surfaces are common. Wave

ripples commonly modify the top. Predominantpalaeoflow direction is landward. Trough andplanar cross-beds are, in places, intensely de-formed by water-escape structures. Convolutebedding is abundant, particularly in the mostdownslope reaches of the association, with foldaxes oriented downslope. Bioturbation is rare toabundant, with Planolites, Ophiomorpha andThalassinoidestraces.

Interpretation. The bipolar palaeocurrent direc-tions, double mud drapes and reactivation surfa-ces indicate tidal deposition. Bar migrationadjacent to channels suggests deposition in amainly subtidal depositional setting (Klein, 1970;de Raaf & Boersma, 1971; Visser, 1980; Clifton,1983; Dalrymple et al., 1990). The frequent con-volution and water-escape structures wereformed during deposition, as the axial planes ofthe folds have a preferential downslope direc-tion. Deformation could have been initiated by

Fig. 12. Tidal-dominated, shelf-edge estuarine deposits from the upper part of Litledalsfjellet (see location in Fig. 5).




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Fig.13.Detailsoftheestuarin

ebarsatthetopoftheLitledalsfjelletwedge(locationinFig.5).Theest

uarinebarsoverlietheheadwarden

dofachutechannel

(A).(B)Althoughpalaeocurrentdirectionismostlyorientedlandwards,oppositedirectionsareoftenseeninsuccessivebars.Notetheshalerip-upclasts

separatingsetboundaries(C).




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sediment creep and wave-induced liquefaction(Dalrymple, 1979). However, rapid changes inwater level, together with excavation of scarps bychannel migration and downcutting along theslope, are believed to have been the major localcauses that encouraged liquefaction. The abovefeatures together with evidence for frequent wave

reworking suggest that the channel network anddune field formed at the mouth of an estuary (e.g.Dalrymple et al., 1990; Berne et al., 1993).

SHELF-MARGIN ARCHITECTUREAT LITLEDALSFJELLET

The location of the series of measured profiles onLitledalsfjellet is shown in Fig. 1. Profiles 13(Figs 5 and 14) are only slightly oblique to thestrike orientation of the palaeoslope, whereas

profiles 410 are nearly dip-parallel to the slope.It should additionally be noted that the landwardend of the Litledalsfjellet outcrop (location 1) liessome distance basinwards compared with thelandward end of the Hgsnyta, i.e. a more prox-imal part of the shelf-delta system can be seen onHgsnyta.

The present dip configuration of the studiedstrata is near the original depositional dip, as canbe deduced from the flat-lying attitude of thedelta plain strata in the upper part of the succes-sion on Litledalsfjellet. Figure 5 shows clearlythat the lower two-thirds of the succession is

overall regressive (slope lobes R1 to R9), whereasthe upper third of the succession is overalltransgressive, with successive slope lobes step-ping landwards through time (T1 to T4).

Architecture of the regressive succession

The most conspicuous features of the regressive,lower half of the succession on Litledalsfjellet are(1) the relatively coarse-grained channels andmouth-bar sands at the shelf edge; and (2) the setsof clinoformal units (slope lobes and subaqueous

channels) that stack on the slope and representthe main accretion of the shelf margin (Figs 5and 6).

The prograding slope lobes

The regressive slope succession seen in the cross-section in Fig. 5 consists of nine progradationalclinoform sets of slope lobes (R1 to R9), each witha thickness up to 12 m. Internally, each lobeflattens on the lowermost slope, is steepest on the

middle to upper slope and tends to flatten slightlytowards the shelf edge. The clinoform set isinternally progradational (the lobes build gradu-ally down the slope) and consists of: (1) adownlap surface that has a gradient similar tothe general slope gradient; (2) a muddy butheterolithic lobe fringe that progrades and laps

down onto the underlying lobe; and (3) a sandyamalgamation of shallow channels, chutes andspillover sand sheets that cap the lobe, liketopsets to the prograding heterolithic fringe.

The capping sandy channels and sheets gener-ally lie erosively on the underlying, heterolithicforesets. This relationship, as well as the commonupward-coarsening and -thickening character ofthe latter, strongly suggests that the channelswere feeding the heterolithic lobe fringe, causingit to prograde. Upward-fining and -thinningtrends within the heterolithic deposits indicate

lateral lobe shifting and abandonment of slopesegments. Because of this and the restrictedlateral dimensions of the lobes, lobe-by-lobecorrelation between Litledalsfjellet and Hgsnytais impossible.

In addition to the main elements in the lobes,there are also minor lobes that were lateral to orformed beyond the mouths of chutes or channels.In the latter case, the chutes presumably stoppedshort of, or incised beyond, the lobe fringe. Suchminor lobes have dimensions of only tens ofmetres and are characterized by spectacularwedging of beds in addition to upward-coarsen-

ing trends (Fig. 11).

Hyperpycnal turbidity currents on the slope

It has been argued above, especially in theinterpretation of turbidites from channels andspillover lobes, that the channel systems on theslope were linked directly upslope to the maindistributary channels of the shelf-edge deltas. Theabundance of coal fragments and other organicdebris within the slope turbidites is one of themost explicit signs that many of the flows wereriver fed, rather than derived from slumps on theslope or from the mouth bar. It is suggested thatmany of the flows were quasi-steady, hyperpycnalturbidity currents (Mulder & Alexander, 2001),fed from prolonged river flooding and withsediment concentrations high enough to makethe rivers hyperpycnal (Mulder & Syvitski, 1995,1996), but probably lower than many surge-typeturbidity flows. The following points summarizethe argument for this type of flow in the Litle-dalsfjellet slope turbidites:




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SHE

LF-EDGE

ESTUAR

INE

DEPOSITS

DISTRIBUTARY

CHANNELS

FLOO

DINGS

URFACE

SEQU

ENCEBOUNDARY

REGRESSION

SURFACE

OFFORCED

Fig.14.SystemstractsinLitl

edalsfjellet(A)withoutcropdetails

(B).




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1 Drainage and delivery settings. The moun-tainous drainage area for the rivers feeding outonto the relatively narrow Battfjellet coastal plainlay within the West Spitsbergen Orogenic Belt,only 2030 km to the west of the study area. It isknown that small mountainous rivers of this typeare capable of bringing a relatively large bedload

to the sea (Milliman & Syvitski, 1992). Thelocation of these rivers and deltas, at the shelfedge, would have ensured sand delivery into thedeeper water areas, and the steepness of thefronting slope would have aided the main-tainence of hyperpycnal flow (see Mulder et al.,1998).

2 Character of the deposits. There is evidenceof sustained flow (for periods longer than, e.g.slump-generated flows) within the turbiditedeposits, both in the main slope channels andin the spillover lobes, on the slope. Individual

channelized turbidite bodies contain multipleerosional scours, without giving any impressionof a series of separate flows. Such thick bedswith internal erosion surfaces suggest accumu-lative flow (Kneller, 1995) in the channels,probably reflecting the period of rising floodstage in the feeder rivers. The waxing flowtends to scour and cannibalize its own deposits,as the effect of the rising flood moves steadilydownstream (Mulder et al., 1998; Mulder &Alexander, 2001). Other important features inchannel deposits include thick (>2 m) units ofplane-parallel-laminated sandstone that con-

tain local, laterally restricted scour surfaces.Evidence of sustained flow is also seen inindividual spillover lobes. Such lobes occur assharp-based sand units up to 2 m thick, butinternally consist of thin (


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edge, leaving the distinctive erosional terracesand allowing the slope clinoforms to have aslightly descending geometric trajectory through

time (Fig. 14).

Forced regressive systems tract

The unconformity noted above, the downward-directed growth trajectory of successive slopeclinoforms (R1R7) and the below the shelf edgelocation of this package strongly suggest that thelower part of the succession reflects a falling stageof relative sea level and should therefore bereferred to as a forced regressive or falling stagesystems tract (Hunt & Tucker, 1992; Plint &Nummedal, 2000).

The erosive sandy cappings of clinoform sets,also documented in shelf-edge sand wedges ofthe Late Quaternary Rhone Delta (Tesson et al.,1990) and Late Pleistocene MississippiAlabamaShelf (Sydow & Roberts, 1994), have been arguedto originate by stepped relative sea-level fall,forcing successively younger delta lobes furtherout and further down across the shelf edge.Fluvial downcutting would have continued onthe shelf as a diachronous process during sea-level lowering. The topsets would then havebeen over-ridden and incised by the advancing

and downcutting distributary channels duringthe latter part of sea-level fall (Figs 14 and 15).

Top of the forced regressive systems tract:the sequence boundary

The transition from the regressive (R1R7) to theaggradational (R8 and R9) successions in theLitledalsfjellet wedge is an irregular erosionalsurface formed by the multilateral and multi-storey stacking of distributary channels and

mouth-bar complexes in the updip areas andsubaqueous channels and sheet sandstones indowndip areas (Figs 5 and 14). Erosional escarp-

ments are up to 7 m deep (Figs 5 and 6). Thiserosional surface can be followed downslope forup to 3 km. At the shelf break, the surface is welldefined by angular discordances with the under-lying heterolithic clinoform sets. Downslope, theerosional surface produced by the distributarychannel system merges into an erosional surfaceproduced by gully erosion and then into a moreconformable surface when gully edges merge intodistal lobes. This extensive erosional surface isinterpreted as a sequence boundary. The sea-levelfall caused migration of the youngest delta lobeout onto and beyond the shelf edge.

Facies analyses indicate that the Litledalsfjelletdelta was a fluvial-dominated system. Wavereworking and littoral transport processes werelikely to have been overwhelmed by rapid depo-sitional rates on the delta front (see also Coleman,1978).

One of the interesting aspects of the Reindalenslope system is the lack of canyons. Most of thedelta-front architecture consists of small distrib-utary channels and mouth-bar complexes, butapparently none of the channels had enoughcapacity to produce significant erosion at the

shelf break, deliver sediment into upper slopegullies or canyons and develop sandy lower slopeand basin-floor systems. Twichell & Roberts(1982) noted that canyons occur on the modernNew Jersey slope when the gradient exceeds 3,with spacing between canyons decreasing asthe gradient increases. However, in the olderMiocene clinoforms of the New Jersey margin(Fulthorpe & Austin, 1998), canyons are rarelyobserved, despite clinoform gradients of morethan 3, a situation resembling that of the

Fig. 15. Systems tracts in Hgsnyta (A) (modified from Plink-Bjorklund & Steel, 2002) with outcrop details (B). Notethe sequence boundary in (B). This erosional surface can be followed for 3 km along the outcrop.




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Reindalen wedges. Fulthorpe & Austin (1998)argued that the lack of evidence for canyonsbreaching the clinoform breakpoints can beinterpreted as (1) discharge of major rivers at thebreakpoints occurred outside the studied area; or(2) fluvial discharge at breakpoints occurred toobriefly to produce large, slope-breaching canyons.

The outcrops at Reindalen allow a three-dimen-sional examination of the system, as well as of theshale section beyond the pinch-out of the wedges.Data suggest that most of the river dischargeoccurred at the breakpoints, at places well belowthe shelf edge, judging by the basinward extent ofthe mouth bars in cycles R7R9. The gentlycurved to linear trends of clinoform breakpoints,mostly unmarked by canyon erosion, indicatethat sediment was efficiently distributed alongstrike as stacked slope lobes, although someconfinement also occurred, given the number of

chutes and channels recognized.

Lowstand and transgressive systems tracts

It is possible that the thin set of strata lyingbetweenR8 and R9, which shows an aggradational stackingpattern and represents the most pronounced bas-inward shift of the entire system, canbe designatedas a poorly developed lowstand wedge. The over-lying backstepping lobe system is a transgressivesystems tract, leading back up to the estuarinedeposits on the outer shelf (Figs 5 and 14).

REINDALEN CLINOFORMS: THEIR ROLEIN SHELF-MARGIN ACCRETION

The sand-prone clinoforms on Litledalsfjellet andHgsnyta resulted from both (1) shelf deltasextending out to a shelf-edge position, a locationnecessary for sand delivery into deep water areas;and (2) sand delivery (across the shelf break) ontoa non-canyoned slope, so that most of the sandwas retained as a wedge on the slope and notdispersed beyond the base of slope. This storage

of sand on the slope, up to 70 m thick andextending up to 15 km along strike, was appar-ently achieved by the accumulation of sedimentfrom mostly unconfined, depletive turbidity flows(sensu Kneller, 1995) and minor sandy debrisflows. The slope channel and chute systems seemto pinch out on the slope, attesting to the lowefficient nature of the turbidity currents and thelack of flow initiation on the slope. This type ofsand partitioning in the Reindalen clinothemcomplex (type 2 clinoforms of Steel et al., 2000)

(Fig. 3) contrasts with another situation com-monly seen elsewhere in Battfjellet Formation(type 1 of Steel et al., 2000), in which sanddelivered from the shelf edge was focusedthrough canyons and large channels, i.e. withsignificant slope bypass, to accumulate on deep-water fans beyond the base of slope.

Alternation of type 1 and type 2 clinoforms inthe Eocene Central Basin produced alternatingconditions of basin-floor aggradation and slopeaccretion. Because type 2 clinoforms, such as inReindalen, are much more common than type 1,shelf-margin growth was the key component inbasin infilling. The Eocene shelf margin prograd-ed more than 60 km, albeit with an irregular andpartly aggradational trajectory.

In addition to sand-prone type 1 and 2 clino-forms, there are abundant shale-prone clinoformsthat were major contributors to basin infilling.

Such shaly intervals commonly represent trans-gressive and highstand conditions in each relat-ive sea-level cycle, were critically important asaggradational elements in the stratigraphy andcaused the overall shelf-edge trajectory to have acomponent of stratigraphic rise across the basin.

CONCLUSIONS

Documentation of the architecture of Eoceneshelf-edge deltas has allowed insight into theprocesses of shelf-margin accretion. This has been

particularly useful for that part of the margin thatis least known in the literature, i.e. the part thatlies below the shelf edge. Litledalsfjellet andHgsnyta shelf-edge deltas formed a 70 m thick,15-km along-strike apron built down onto a 34slope. The system was dominated by the out-building of distributary channels and mouth barsat the shelf edge and by density underflows on theslope. It is argued that many of the latter weresustained hyperpycnal turbidity flows because ofthe nature of the river drainage, the shelf-edgelocation of the deltas, the abundance of coaly and

organic materials in the turbidites, as well as thecharacter of many of the channel fills andspillover lobes on the slope.

Channel-fed heterolithic lobes (some 500 mwide, up to 2 km long and 1015 m thick),erosional channels and chutes, channel/chute-mouth lobes and spillover sandstone sheets arethe main architectural elements of the slopeassociation.

Stratal configuration and erosional unconformi-ties suggest that progradation of the shelf-edge




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deltas occurred mostly during forced regressiveand lowstand conditions. There was rapid back-stepping of the system, with re-establishment ofshelf-edge estuaries when sea level rose again toabove the shelf break during transgression.

The relatively low sediment concentration(compared with high-density turbidites) in the

hyperpycnal turbidity currents and the lack ofcanyons below the Reindalen shelf edge produceda largely unfocused deltaic discharge onto theslope. The lack of canyons was critical for pre-venting sediment bypass on the slope and, thus,for the absence of deep-water fans beyond the baseof slope. This situation arose despite the demon-strable fall in sea level to below the shelf edge.

ACKNOWLEDGEMENTS

We thank Amoco, Conoco, Exxon, Mobil, NorskHydro, Phillips, Shell, Statoil and UPRC, whosponsored and enthusiastically supported thisresearch (Wolf Consortium). We are indebted toreviewers Ben Kneller and Mike Mayall, and toChris Fielding, for their constructive comments,which helped to improve the manuscript.

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anatomy of shelf deltas at the edge a progradion

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