Anatomy of shelf deltas at the edge of a prograding Eocene ... · *Department of Geology,...

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

DONATELLA MELLERE*, PIRET PLINK-BJORKLUND� 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 and

Pleistocene shelf margins, documented outcrop analogues of these important

sand-prone reservoirs are rare. The facies and stratigraphic architecture of an

outcropping shelf-edge delta system in the Eocene Battfjellet Formation,

Spitsbergen, is presented here, as well as the implications of this delta system

for the generation of sand-prone, shelf-margin clinoforms. The shelf-edge

deltas of the Battfjellet Formation on Litledalsfjellet and Høgsnyta produced a

3–5 · 15 km, shelf edge-attached, slope apron (70 m of sandstones proximally,

tapering to zero on the lower slope). The slope apron consists of distributary

channel and mouth-bar deposits in its shelf-edge reaches, passing downslope

to slope channels/chutes that fed turbiditic lobes and spillover sheets. In the

transgressive phase of the slope apron, estuaries developed at the shelf edge,

and these also produced minor lobes on the slope. The short-headed

mountainous rivers that drained the adjacent orogenic belt and fed the

narrow shelf, and the shelf-edge position of the discharging deltas, made an

appropriate setting for the generation of hyperpycnal turbidity currents on the

slope of the shelf margin. The abundance of organic matter and of coal

fragments in the slope turbidites is consistent with this notion. Evidence that

many of the slope turbidites were generated by sustained turbidity currents

that waxed then waned includes the presence of scour surfaces and thick

intervals 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 preserved

mainly below the shelf break and dominated by sediment gravity-flow

deposits, has a threefold stratigraphic architecture: a lower, progradational

part, in which the clinoforms have a slight downward-directed trajectory; a

thin aggradational zone; and an upper part in which clinoforms backstep up

onto the shelf edge. A greatly increased density of erosional channels and

chutes marks the regressive-to-transgressive turnaround within the slope

apron, and this zone becomes an angular unconformity up near the shelf edge.

This unconformity, with both subaerial and subaqueous components, is

interpreted as a sequence boundary and developed by vigorous sand delivery

and bypass across the shelf edge during the time interval of falling relative sea

level. The studied shelf-margin clinoforms accreted mostly during falling stage

(sea level below the shelf edge), but the outer shelf later became estuarine as

sea 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, 1181–1206

� 2002 International Association of Sedimentologists 1181

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; Tesson et 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, deltaicshorelines 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 earlyrise 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) orshelf-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 beyondthe shelf edge and are now perched on the upper-most slope. Their position indicates a relativesea-level fall of 50–100 m relative to the formerhighstand shoreline. A subsequent rise in sea levelcaused the deltas to draw back up onto the shelf.The fall and rise cycle resulted in the delivery 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 lackof 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 Paleocene–Eocene 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 (150–350 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 beltalong western Spitsbergen, some 25–30 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.

� 2002 International Association of Sedimentologists, Sedimentology, 49, 1181–1206

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 Steel et al., 2000) and can be recog-nized on the mountainside only by the presence ofheterolithic units of thin-bedded (1–3 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 Høgs-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 shelfbreak. 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 Høgsnyta 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 90–100 km wide.(B) Location of the Litledalsfjellet and Høgsnyta outcrops in Reindalen. The shelf-edge delta wedges are up to 70 mthick, but taper out at the base of slope, some 3–5 km away from the shelf margin.

Anatomy of a prograding Eocene shelf margin, Spitsbergen 1183


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 50–500 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 Høgsnyta), 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 shelfedge. The shelf-edge deltas of theBattfjellet Formation prograded ontothe Gilsonryggen shales and werefed by the coastal plain system ofthe Aspelintoppen Formation.



The choice of datum for the correlation panelsis important. The photomosaics of Litledalsfjelletand Høgsnyta 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 hungfrom 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 3–4�. The deposits form a shelfedge-to-slope turbidite apron 3–5 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 0Æ5� and 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 Høgsnyta (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 theconstruction of the correlation panels. The mountains are 700 m high, and the mountain face is 3–5 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 3–4�.



Fig

.5.

Lit

led

als

fjell

et

cro

ss-s

ecti

on

al

pan

el

(see

locati

on

of

the

measu

red

secti

on

inF

ig.

4)

wit

hp

osi

tion

of

the

log

deta

ils

an

dp

hoto

gra

ph

su

sed

inth

ed

esc

rip

tion

of

the

facie

sass

ocia

tion

s.



Fig

.6.

Ph

oto

mosa

ican

din

terp

reta

tion

of

the

most

pro

xim

al

reach

es

of

the

Lit

led

als

fjell

et

shelf

-ed

ge

delt

acom

ple

x.



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 todescribe 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 15–20 km andgradient of less than 0Æ5�; and (2) a slope withlength 3–5 km and gradient of 3–4�. 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 edgereflects a (decompacted) water depth of some50 m at the shelf break and some 250–300 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 3–4�. 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 Høgsnyta wed-ges (profiles 1 and 1–3, respectively, in Figs 5–7).

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,0Æ2–2 m thick, separated by up to 0Æ15 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 alluvial‘flood-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. Høgsnyta cross-sectional panel, modified from Plink-Bjorklund & Steel (2002) (see location of the measuredsection in Fig. 4).



Fig

.8.

Overv

iew

an

dd

eta

ils

of

shelf

-ed

ge

delt

ad

ep

osi

ts(l

ocati

on

inF

ig.

5).

(A)

Dis

tal

mou

th-b

ar

ass

ocia

tion

(Facie

sA

ssocia

tion

A2).

Most

of

the

dep

osi

tsare

thin

sets

of

rip

ple

-to

pla

nar-

lam

inate

dsa

nd

ston

es.

(Ban

dC

)R

iver-

fed

,si

gm

oid

al

bars

at

the

shelf

-ed

ge

delt

am

ou

thon

Lit

led

als

fjell

et

(Ass

ocia

tion

A1).

Note

the

slig

ht

lan

dw

ard

back-t

ilt

(sli

gh

tle

ftw

ard

dip

of

som

ebed

sin

B)

of

the

head

of

the

bars

.(D

)O

verv

iew

of

river

dis

trib

uta

rych

an

nel

inH

øgsn

yta

.T

he

dep

osi

tsare

thin

lyp

lan

ar

lam

inate

d,

suggest

ing

dep

osi

tion

from

flood

-dom

inate

dh

yp

erc

on

cen

trate

dfl

ow

s.



arranged (in section parallel to palaeoflow) inhigh-angle (15–20�) accretionary sigmoidal sets(Fig. 8B and C). Sets are convex upward, 0Æ5–2 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 6–10 m wide) and aretruncated by flat-lying erosional surfaces or bysmall scour-and-fill features in their upcurrentparts (Fig. 8B). The latter are 0Æ4–1 m deep,0Æ6–5 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 accumulationof 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 andflat-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 Mutti et al. (1996), the developmentof 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, 0Æ5–1Æ5 m thick (Fig. 8). Theydevelop 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 0Æ05–0Æ5 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 presentsshear lamination.

Sandstone beds show low to moderate lateralpersistence, 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 thosedescribed 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, 50–70 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 Høgsnyta (Fig. 7). Around



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 unitsrepresent the distal reaches of the mouth–barassociation, 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 0Æ5 m), withsome 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 ofthe 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 5–7). 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. OnHøgsnyta, channels are 25–100 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 beds5–70 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 (3–15 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



directly linked back upslope to a mouth-barsystem and distributary channels (Figs 5–7). 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 slopechannels 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 3–15 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 laminaeat 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 lobeshifting. 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.



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 Høgsnyta 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, 50–200 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 0Æ3–3 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 shallowtroughs (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 followedthroughout 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(0Æ5–3 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 moreerosional–depositional 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(Prior et al., 1981; Bouma et 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 verywell 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 chutesor, 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



Fig

.10.

Deta

ils

of

slop

ed

ep

osi

ts(F

acie

sA

ssocia

tion

B,

locati

on

inF

ig.

5).

(Aan

dB

)S

lop

ech

an

nels

(Ass

ocia

tion

B2)

cu

ttin

gu

nd

erl

yin

gsp

illo

ver

sheets

(Ass

ocia

tion

B4).

Note

the

imbri

cate

shale

rip

-up

cla

sts

con

cen

trate

dat

the

base

of

the

ch

an

nel.

Th

esp

illo

ver

sheet-

like

bed

s(C

)fo

rmlo

ng

tabu

lar

un

its,

wed

gin

gsl

igh

tly

dow

nsl

op

ean

dcon

sist

ing

of

thin

lyla

min

ate

dsa

nd

ston

es.

Tra

nsp

ort

ed

wood

an

dcoal

fragm

en

tsare

com

mon

an

dare

oft

en

con

cen

trate

dat

the

base

of

the

bed

s.



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(60–70%). Sandstone beds, 5–75 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.



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 turbidityflows. 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 1Æ5 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 mthick. 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 3–5 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 1Æ5 to 5 m thickcoarsening- and thickening-upward slope-lobeunits (Fig. 10C) and are cut by slope channelsand 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 andHøgsnyta), 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 for1Æ5 km. Channel margins are step-like and oftendraped by sand-prone, up to 1 m thick slumpdeposits (Fig. 12A). Lateral wings, a few tens ofmetres wide, develop in the uppermost channellevels. Channel fills are 3–8 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(3–8 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 ofPlanolites at 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



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 to30 m upslope with gentle dip angles (10–15�).The bars overlie and border the flanks of the tidalchannels. Cross-beds are 20–30 cm thick, separ-ated by reactivation surfaces (Fig. 13) and trainsof shale rip-up clasts (Fig. 13C). Internal sets aresigmoidal, 3–8 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 andThalassinoides traces.

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).



Fig

.13.

Deta

ils

of

the

est

uari

ne

bars

at

the

top

of

the

Lit

led

als

fjell

et

wed

ge

(locati

on

inF

ig.

5).

Th

eest

uari

ne

bars

overl

ieth

eh

ead

ward

en

dof

ach

ute

ch

an

nel

(A).

(B)

Alt

hou

gh

pala

eocu

rren

td

irecti

on

ism

ost

lyori

en

ted

lan

dw

ard

s,op

posi

ted

irecti

on

sare

oft

en

seen

insu

ccess

ive

bars

.N

ote

the

shale

rip

-up

cla

sts

sep

ara

tin

gse

tbou

nd

ari

es

(C).



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 wavereworking 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 1–3(Figs 5 and 14) are only slightly oblique to thestrike orientation of the palaeoslope, whereasprofiles 4–10 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 Høgsnyta, i.e. a more prox-imal part of the shelf-delta system can be seen onHøgsnyta.

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 isoverall 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 subaqueouschannels) 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 lapsdown 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 indicatelateral lobe shifting and abandonment of slopesegments. Because of this and the restrictedlateral dimensions of the lobes, lobe-by-lobecorrelation between Litledalsfjellet and Høgsnytais 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:



SH

EL

F-E

DG

EE

ST

UA

RIN

E D

EP

OS

ITS

DIS

TR

IBU

TAR

Y C

HA

NN

EL

SF

LO

OD

ING

SU

RFA

CE

SE

QU

EN

CE

BO

UN

DA

RY

RE

GR

ES

SIO

NS

UR

FAC

E O

F F

OR

CE

D

Fig

.14.

Syst

em

str

acts

inL

itle

dals

fjell

et

(A)

wit

hou

tcro

pd

eta

ils

(B).



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 20–30 km to the west of the study area. It isknown that small mountainous rivers of this typeare capable of bringing a relatively large bedloadto 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. Individualchannelized 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 (<8 cm), alternatingplane-parallel-laminated and massive intervals,with increasing amounts of current-ripplelamination towards the top of the unit. Theseunits generally give an impression of irregularaggradation of the bed in depletive flow condi-tions.

3 Lobe progradation and clinoform growth.The nature, scale and gradual progradation of thesandy turbidite lobes, as well as the systematicgrowth of clinoforms, also suggest that slopesystems were fed directly from distributarychannels at the shelf edge. This is not to say thatslumping of the mouth bars or collapse of seg-ments of the slope was unimportant in producingturbidity currents, but the gradual downslopeprogradation of the slope lobes is more consistentwith sustained flows.

Architecture of the transgressive succession

Prograding slope lobes also dominate the upperpart of the Litledalsfjellet succession. However,these lobes, each capped by sheet sandstonesrather than chute channel complexes, step land-wards through time (T1–T4), illustrating theoverall transgressive character of this part of thesuccession. In the most distal reaches of theslope, lobes have an aggradational stacking archi-tecture (R8, R9 and T1), before they backstep onthe slope (Figs 5 and 14). At the proximal end ofthe transgressive succession, there are well-preserved distributary channels and some associ-ated mouth bars that fed the slope lobes. The topsof mouth-bar units show wave-ripple laminationand broad scours, indicating abandonment andwave reworking of the delta lobes (e.g. Penlandet al., 1988) (Fig. 12).

Landward migration of the baylineand estuarine deposits

As sea level continued to rise, the alluvial valleywas drowned by a landward-migrating bayline,and an estuarine environment formed. The risingsea level produced a landward-thickening wedgeof stacked estuarine channels and bars, with sheetsands and shoals further offshore (Figs 5 and 14).The thickness of the transgressive deposits, up to20 m on the shelf, was probably enhanced by a highgradient of ravinement or by a slow landwardtranslation of the ravinement. The tidal channeland bar system recognized on Litledalsfjellet wasprobably deposited at the mouth of an unbarredestuarine valley (Rahmani, 1988; Dalrymple et al.,1992). No traces of intertidal, central mud basindeposits were recognized in the studied area.

SEQUENCE STRATIGRAPHY

Updip unconformity

The stratal configuration seen in Fig. 5 suggeststhat the erosional surfaces at the top of theyoungest clinoform sets in the progradationallower part of the succession truncate, in a land-ward direction, progressively older lobes, leavingerosional terraces or unconformities near theshelf edge.

This same pattern of updip unconformities isseen even more clearly on Høgsnyta (Figs 7 and15; Plink-Bjorklund & Steel, 2002). This config-uration suggests that sea level fell below the shelf



edge, leaving the distinctive erosional terracesand allowing the slope clinoforms to have aslightly descending geometric trajectory throughtime (Fig. 14).

Forced regressive systems tract

The unconformity noted above, the downward-directed growth trajectory of successive slopeclinoforms (R1–R7) and the ‘below the shelf edge’location 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 Mississippi–AlabamaShelf (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 advancingand 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 (R1–R7) 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 theshelf 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 Høgsnyta (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.



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 R7–R9. 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 ofchutes and channels recognized.

Lowstand and transgressive systems tracts

It is possible that the thin set of strata lying betweenR8 and R9, which shows an aggradational stackingpattern and represents the most pronounced bas-inward shift of the entire system, can be 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 andHøgsnyta 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 storageof 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 beenparticularly useful for that part of the margin thatis least known in the literature, i.e. the part thatlies below the shelf edge. Litledalsfjellet andHøgsnyta shelf-edge deltas formed a 70 m thick,15-km along-strike apron built down onto a 3–4�slope. 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 andorganic 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 10–15 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



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 thehyperpycnal 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.

REFERENCES

Berne, S., Castaing, P., Le Drezen, E. and Lericolais, G. (1993)

Morphology, internal structure, and reversal of asymmetry

of large subtidal dunes in the entrance to Gironde estuary

(France). J. Sed. Petrol., 63, 780–793.

Berryhill, H.L., Suter, J.R. and Hardin, N.S. (1986) Late Qua-

ternary facies and structure, northern Gulf of Mexico. AAPG

Studies in Geology, 23, 289 pp.

Bornhold, B.D. and Prior, D.B. (1990) Morphology and sedi-

mentary processes on the subaqueous Noeick river delta,

British Columbia, Canada. In: Coarse-Grained Deltas (Eds A.

Colella and D.B. Prior), Int. Assoc. Sedimentol. Spec. Publ.,

10, 169–181.

Bouma, A. (1962) Sedimentology of Some Flysch Deposits,

a Graphic Approach to Facies Interpretation. Elsevier,

Amsterdam, 168 pp.

Bouma, A.H., Roberts, H.H., Coleman, J.M. and Prior, D.B.(1991) Delta front gullies as part of mass-movement phe-

nomena: Mississippi River delta front. In: From Shoreline toAbyss: Contribution in Marine Geology in Honor of Francis

Parker Shepard (Ed. R.H. Osborne), SEPM Spec. Publ., 46,98–105.

Clark, J.D. and Pickering, K.T. (1996) Submarine Channels,Processes and Architecture. Vallis Press, Oxford, UK, 231 pp.

Clifton, H.E. (1983) Discrimination between subtidal and

intertidal facies in Pleistocene deposits, Willapa Bay,

Washington. J. Sed. Petrol., 53, 353–369.

Coleman, J.M. (1978) Deltas: Processes of Deposition and

Models for Exploration. Continuing Education Publishing

Co., Louisiana, 102 pp.

Coleman, J.M. and Roberts, H.H. (1988) Sedimentary devel-

opment of the Louisiana continental shelf related to sea-

level cycles. Geo-Mar. Lett., 9, 161–170.

Coleman, J.M., Prior, D.B. and Lindsay, J.F. (1983) Deltaic

influences on shelf edge instability processes. In: The Shelf

Break, Critical Interface on Continental Margins (Eds D.J.

Stanley and G.T. Moore), SEPM Spec. Publ., 33, 121–137.

Dalrymple, R.W. (1979) Wave-induced liquefaction: a modern

example from the Bay of Fundy. Sedimentology, 26, 835–844.

Dalrymple, R.W., Knight, R.J., Zaitlin, B.A. and Middleton,G.V. (1990) Dynamics and facies model of a macrotidal

sandbar complex, Cobequid Bay – Salmon river estuary (Bay

of Fundy). Sedimentology, 37, 577–612.

Dalrymple, R.W., Zaitlin, B.A. and Boyd, R. (1992) Estuarine

facies models: conceptual basis and stratigraphic implica-

tions. J. Sed. Petrol., 62, 1130–1146.

DiBona, P.A. and von der Borch, C.C. (1993) Sedimentary

geology and evolution of an outcropping shelf-margin delta,

Late Proterozoic Wonoka Formation, South Australia.

AAPG Bull., 77, 963–979.

Field, M. and Trincardi, F. (1991) Regressive coastal deposits

on Quaternary continental shelves: preservation and legacy.

In: From Shoreline to Abyss (Ed. R.H. Osborne), SEPM Spec.

Publ., 46, 107–122.

Fulthorpe, C.G. and Austin, J.A. (1998) Anatomy of Rapid

Margin Progradation: three-dimensional geometries of

Miocene clinoforms, New Jersey Margin. AAPG Bull., 82,251–273.

Harland, W.B. (1969) Contribution of Spitsbergen to under-

standing of the tectonic evolution of the North Atlantic

region in North Atlantic, Geology and Continental Drift.

AAPG Mem., 12, 817–851.

Harland, W.B. (1995) The West Spitsbergen Fold Belt: the

result of Late Cretaceous–Paleocene Greenland–Svalbard

convergence? Discussion. Geol. J., 30, 189–195.

Hart, B.S. and Long, B.F. (1996) Forced regressions and low-

stand deltas: Holocene Canadian examples. J. Sed. Res., 66,820–829.

Helland-Hansen, W. (1990) Sedimentation in Paleogene fore-

land basin, Spitsbergen. AAPG Bull., 74, 260–272.

Helland-Hansen, W. (1992) Geometry and facies of Tertiary

clinothems, Spitsbergen. Sedimentology, 39, 1013–1029.

Hunt, D. and Tucker, M.E. (1992) Stranded parasequences and

the forced regressive wedge systems tract: deposition during

base-level fall. Sed. Geol., 81, 1–9.

Kellog, H.E. (1975) Tertiary stratigraphy and tectonism in

Svalbard and continental drift. AAPG Bull., 59, 465–485.

Klein, G. de V. (1970) Depositional and dispersal dynamics of

intertidal sand bars. J. Sed. Petrol., 40, 1095–1127.

Kneller, B. (1995) Beyond the turbidite paradigm: physical

models for deposition and their implications for reservoir

prediction. In: Characterization of Deep Marine ClasticSystems (Eds A.J. Hartley and D.J. Prosser), Geol. Soc.

London Spec. Publ., 94, 31–49.

Kneller, B. and Branney, M.J. (1995) Sustained high-density

turbidity currents and the deposition of thick massive

sands. Sedimentology, 42, 607–616.

Kolla, V. and Perlmutter, M.A. (1993) Timing of turbidite

sedimentation on the Mississippi Fan. AAPG Bull., 77,1129–1141.

Kolla, V., Biondi, P., Long, B. and Fillon, R. (2000) Sequence

stratigraphy and architecture of the Late Pleistocene Lagni-

appe delta complex, northeast Gulf of Mexico. In: Sedimen-tary Responses to Forced Regression (Eds D. Hunt and R.L.

Gawthorpe), Geol. Soc. London Spec. Publ., 172, 291–327.



Lowe, D.R. (1982) Sediment-gravity flows. II. Depositional

models with special reference to the deposits of high-den-

sity turbidity currents. J. Sed. Petrol., 52, 279–297.

Lowell, J.D. (1972) Spitsbergen Tertiary orogenic belt and the

Spitsbergen fracture zone. Geol. Soc. Am. Bull., 83, 3091–

3102.

McMaster, R.L., DeBoer, J. and Ashraf, A. (1970) Magnetic

and seismic reflection studies on continental shelf of Por-

tuguese Guinea, Guinea and Sierra Leone, West Africa.

AAPG Bull., 54, 158–167.

Matteucci, T.D. and Hine, A.C. (1987) Evolution of Cape Fear

terrace: a complex interaction between the Gulf Stream and

a paleo shelf edge delta. Mar. Geol., 77, 185–205.

Mayall, M.J., Yeilding, C.A., Oldroyd, J.D., Pulham, A.J. and

Sakurai, S. (1992) Facies in a shelf-edge delta – an example

from the subsurface of the Gulf of Mexico, Middle Pliocene,

Mississippi Canyon, Block 109. AAPG Bull., 76, 435–448.

Middleton, G.V. (1967) Experiments on density and turbidity

currents. III. Deposition of sediment. Can. J. Earth Sci., 4,475–505.

Middleton, G.V. and Hampton, M.A. (1973) Sediment gravity

flow: mechanics of flow and deposition. In: Turbidity andDeep Water Sedimentation (Eds G.V. Middleton and

A.H. Bouma), Pacific Section, SEPM Short Course Notes, 1,

1–38.

Milliman, J.D. and Syvitski, J.P.M. (1992) Geomorphic/

tectonic control of sediment discharge to the ocean; the

importance of small mountainous rivers. J. Geol., 100,525–544.

Morton, R.A. and Suter, J.R. (1996) Sequence stratigraphy

and composition of Late Quaternary shelf-margin deltas,

Northern Gulf of Mexico. AAPG Bull., 80, 505–530.

Mulder, T. and Alexander, J. (2001) The physical character of

subaqueous sedimentary density flows and their deposits.

Sedimentology, 48, 269–299.

Mulder, T. and Syvitski, J.P.M. (1995) Turbidity currents

generated at river mouths during exceptional discharges to

the world oceans. J. Geol., 103, 285–299.

Mulder, T. and Syvitski, J.P.M. (1996) Climatic and morpho-

logic relationships of rivers: implications of sea-level fluc-

tuations on river loads. J. Geol., 104, 509–523.

Mulder, T., Syvitski, J.P.M. and Skene, K. (1998) Modeling of

erosion and deposition by turbidity currents generated at

river mouths. J. Sed. Res., 68, 124–137.

Mutti, E. and Normark, W.R. (1987) Comparing examples of

modern and ancient turbidite systems: Problems and con-

cepts. In: Marine Clastic Sedimentology: Concepts and Case

Studies (Eds J.K. Leggett and G.G. Zuffa), pp. 1–38. Graham

and Trotman, London.

Mutti, E., Davoli, G., Tinterri, R. and Zavala, C. (1996) The

importance of ancient fluvio-deltaic systems dominated by

catastrophic flooding in tectonically active basins. Soc.Geol. Ital. Mem., 48, 233–291.

Myhre, A.M. and Eldholm, O. (1988) The Western Svalbard

Margin (74–80�N). Mar. Petrol. Geol., 5, 134–156.

Normark, W.R. (1970) Growth patterns of deep-sea fans.

AAPG Bull., 54, 2170–2195.

Normark, W.R. and Piper, S.J.W. (1991) Initiation processes

and flow evolution of turbidity currents: implications for

the depositional record. In: From Shoreline to Abyss: Con-

tribution in Marine Geology in Honor of Francis Parker

Shepard (Ed. R.H. Osborne), SEPM Spec. Publ., 46, 207–

230.

Penland, S., Boyd, R. and Suter, J.R. (1988) The transgressive

depositional systems of the Mississippi delta plain: a model

for barrier shoreline and shelf sand development. J. Sed.

Petrol., 58, 932–949.

Plink-Bjorklund, P. and Steel, R.J. (2002) Sea-level fall below

the shelf-edge without basin floor fans. Geology, 30, 115–

118.

Plint, A.G. and Nummedal, D. (2000) The falling stage systems

tract: recognition and importance in sequence stratigraphic

analysis. In: Sedimentary Responses to Forced Regression(Eds D. Hunt and R.L. Gawthorpe), Geol. Soc. London Spec.

Publ., 172, 1–17.

Poag, C.W., Swift, B.A., Schlee, J.L., Ball, M.M. and Sheet, L.L.(1990) Early Cretaceous shelf-edge deltas of the Baltimore

Canyon Trough: principal sources for sediment gravity de-

posits of the northern Hatteras Basin. Geology, 18, 149–152.

Posamentier, H.W., Jervey, M.T. and Vail, P.R. (1988) Eustatic

control on clastic deposition I – conceptual framework. In:

Sea-Level Changes: an Integrated Approach (Eds C.K. Wil-

gus, B.S. Hastings, C.G.StC. Kendall, H.W. Posamentier,

C.A. Ross and J.C. Van Wagoner), SEPM Spec. Publ., 42,109–124.

Postma, G., Nemec, W. and Kleinspehn, K.L. (1988) Large

floating clasts in turbidites: a mechanism for their

emplacement. Sed. Geol., 58, 47–61.

Prior, D.B. and Bornhold, B.D. (1990) The underwater devel-

opment of Holocene fan deltas. In: Coarse-Grained Deltas

(Eds A. Colella and D.B. Prior), Int. Assoc. Sedimentol.Spec. Publ., 10, 75–90.

Prior, D.B., Bryant, W.R. and Wiseman, W.J. (1981) Submar-

ine chutes on the slopes of fjord deltas. Nature, 290, 326–

328.

de Raaf, J.F.M. and Boersma, J.R. (1971) Tidal deposits and

their sedimentary structures. Geol. Mijnbouw, 50, 479–503.

Rahmani, R.A. (1988) Estuarine tidal channels and nearshore

sedimentation of a late Cretaceous epicontinental sea,

Drumheller, Alberta, Canada. In: Tide-Influenced Sedi-

mentary Environments and Facies (Eds P.L. de Boer, A. Van

Gelder and S.D. Nio), pp. 433–471. D. Reidel Publishing,

Dordrecht.

Scruton, P.C. (1960) Delta building and the deltaic sequence –

Recent Sediments, Northwest Gulf of Mexico. AAPG Symp.,

82–102.

Shanmugam, G. (1996) High-density currents: are they sandy

debris flows? J. Sed. Res., 66, 2–10.

Shanmugam, G. and Moiola, R.J. (1995) Reinterpretation of

depositional processes in a classic flysch sequence (Penn-

sylvanian, Jackfork Group), Ouachita Mountains, Arkansas

and Oklahoma. AAPG Bull., 79, 672–695.

Soh, W., Tanaka, T. and Taira, A. (1995) Geomorphology and

sedimentary processes of a modern slope-type fan delta

(Fujikawa fan delta), Suruga Trough, Japan. Sed. Geol., 98,79–95.

Spencer, A.M., Home, P.C. and Berglund, L.T. (1984) Tertiary

structural development of the western Barents Shelf,

Tromso to Svalbard. In: Petroleum Geology of the North

European Margin (Ed. A.M. Spencer), pp. 199–209.

Norwegian Petroleum Society, Graham and Trotman,

London.

Steel, R.J., Gjelberg, J., Helland-Hansen, W., Kleinspehn, K.L.,Nøttvedt, A. and Rye-Larsen, M. (1985) The Tertiary strike-

slip basins and orogenic belt of Spitsbergen. In: Strike-Slip

Deformation, Basin Formation and Sedimentation (Eds K.T.

Biddle and N. Christie-Blick), SEPM Spec. Publ., 37, 339–

359.

Steel, R., Mellere, D., Plink-Bjorklund, P., Crabaugh, J.,Deibert, J. and Schellpeper, M. (2000) Deltas versus rivers



on a shelf edge: their relative contributions to the growth of

shelf-margins and basin-floor fans (Barremian and Eocene,

Spitsbergen). In: Deep-Water Reservoirs of the World (Eds

P. Weimer, R.M. Slatt, J. Coleman, N.C. Rosen, H. Nelson,

A.H. Bouma, M.J. Styzen and D.T. Lawrence). Transactions

Gulf Coast SEPM Annual Conference, Houston, pp. 981–

1009 (CD-ROM).

Suter, J.R. and Berryhill, H.L. (1985) Late Quaternary shelf-

margin deltas, northwest Gulf of Mexico. AAPG Bull., 69,77–91.

Sydow, J. and Roberts, H.H. (1994) Stratigraphic framework

of a Late Pleistocene shelf-edge delta, northeast Gulf of

Mexico. AAPG Bull., 78, 1276–1312.

Tesson, M., Gensous, B., Allen, G.P. and Ravenne, C. (1990)

Late Quaternary deltaic lowstand wedges on the Rhone

Continental Shelf, France. Mar. Geol., 91, 325–332.

Tesson, M., Allen, G.P. and Ravenne, C. (1993) Late Pleisto-

cene shelf-perched lowstand wedges on the Rhone contin-

ental shelf. In: Stratigraphy and Facies Associations in aSequence Statigraphic Framework (Eds H.W. Posamentier,

C.P. Summerhayes, B.U. Haq and G.P. Allen), Int. Assoc.

Sedimentol. Spec. Publ., 18, 183–196.

Trincardi, F. and Field, M.E. (1990) Geometry, lateral vari-

ation and preservation of downlapping regressive shelf

deposit: eastern Tyrrhenian Sea margin, Italy. J. Sed. Petrol.,

61, 775–790.

Twichell, D.C. and Roberts, D.G. (1982) Morphology, distri-

bution and development of submarine canyons on the

United States Atlantic continental slope between Hudson

and Baltimore Canyons. Geology, 10, 408–412.

Visser, M.J. (1980) Neap-spring cycles reflected in Holocene

subtidal large-scale bedform deposits: a preliminary note.

Geology, 8, 543–546.

Walker, R. (1967) Turbidite sedimentary structures and their

relationship to proximal and distal depositional environ-

ments. J. Sed. Petrol., 37, 25–37.

Winker, C.D. and Edwards, M.B. (1983) Unstable prograda-

tional clastic shelf margins. In: The Shelfbreak: Critical

Interface on Continental Margins (Eds D.J. Stanley and

G.T. Moore), SEPM Spec. Publ., 33, 139–157.

Manuscript received 16 March 2001;revision accepted 15 March 2002.



Anatomy of shelf deltas at the edge of a prograding Eocene ... · *Department of Geology,...

Documents

Transcript of Anatomy of shelf deltas at the edge of a prograding Eocene ... · *Department of Geology,...