Highstand Shelf-margin Delta System

7/31/2019 Highstand Shelf-margin Delta System

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A highstand shelf-margin delta system from the Eocene

of West Spitsbergen, Norway

Carlos A. Uroza , Ronald J. Steel

Department of Geological Sciences, The University of Texas at Austin, 1 University Station, C-1100, Austin, TX 78712, USA

Received 26 March 2007; received in revised form 19 November 2007; accepted 7 December 2007

Abstract

Demonstration of shelf-margin accretion by shelf-edge deltas during rising and highstand of relative sea level has important consequences for

deepwater sand depositional models. Although highstand shelf-edge deltas are conceptually feasible and have been recently argued from

subsurface data, we describe here the first outcrop example, thus providing facies and architectural data on this important category of delta. Deltas

are able to reach the shelf-edge during rising sea level, if one or more of the key conditions of sediment supply, shelf width/gradient, or basinal

processes are such as to allow complete cross-shelf progradation before the onset of delta auto-retreat. Such highstand deltas promote the retention

of high volumes of sand on the aggrading shelf and coastal plain, and thus potentially have a reduced sand budget available for delivery to the

deeper water areas. Clinoform 17, one of a series of eastward-prograding, shelf-margin clinoforms from the Eocene Battfjellet Formation on West

Spitsbergen, contains a sand-rich delta complex sited near the clinoform shelf-slope rollover, and is argued to be a highstand (rising relative sea

level) shelf-margin delta based on: (1) its highly aggradational architecture shown by an unusual (compared to other clinoforms) regressive unit

thickness and its marked stacking of parasequences, (2) coeval accumulation of delta-plain and lagoonal deposits that are well-preserved in the

landward reaches of the same clinoform, and (3) its context within a mappable, longer-term rising shelf-edge trajectory (through 5 clinoforms). It

is likely that the delta reached its shelf-edge location because the shelf was narrow (less than 20 km), and not because of high sediment supply or

relative sea-level fall. The delta system was markedly wave-dominated as might be predicted at a shelf-edge site.The sand-rich, shelf-edge portion of Clinoform 17 consists of (1) a 3035 m thick regressive deltaic unit with offshore mudstones and thin

tempestite layers, wave-dominated delta-front sandstones, and tidalfluvial-distributary channels on the delta topsets, (2) an overlying 1523 m

thick, aggrading-to-transgressive shoreface/barrier unit with associated tidal-inlet/estuarine channel-fill deposits, and (3) an uppermost, b20 m

thick regressive deltaic unit similar to (1). The slope successions of the units described in (1) and (3), beyond and below the shelf-edge, contain

thin upper-slope tempestite sheet sandstones, within an otherwise shale-dominated environment. Neither sandy slope channels nor basin-floor fans

are observed within the otherwise shale-prone deepwater segments of the clinoform.

2008 Elsevier B.V. All rights reserved.

Keywords: Highstand shelf-margin delta; Rising shelf-edge trajectory; Auto-retreat; Rising relative sea level

1. Introduction

Shelf-edge deltas developed during conditions of relative

sea-level fall are well-known from the Pleistocene shelf-margin

in the Gulf of Mexico (Suter and Berryhill, 1985; Sydow and

Roberts, 1994; Morton and Suter, 1996; Roberts et al., 2000),

the Eocene of West Spitsbergen (Mellere et al., 2002; Plink-

Bjrklund and Steel, 2005), the Porcupine Basin offshore

Ireland (Johannessen and Steel, 2005), and the PlioPleistocene

Orinoco delta in Trinidad (Sydow et al., 2003) among others.

Shelf-edge deltas are conventionally associated with low sea

level because falling sea level is known to be an efficient driver

for bringing shorelines entirely across the shelf (Muto and Steel,

2002). However, deltas that crossed the shelf to the shelf-edge

area during rising relative sea level (highstand conditions) have

not been architecturally documented though these have been

conceptually postulated by Burgess and Hovius (1998), imaged

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Corresponding author.

E-mail address: [email protected] (C.A. Uroza).

0037-0738/$ - see front matter 2008 Elsevier B.V. All rights reserved.doi:10.1016/j.sedgeo.2007.12.003
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on seismic data by Bullimore et al. (2005), and described from

subsurface data by Carvajal and Steel (2006).

The two conditions that would most likely promote highstand

deltas at or near the shelf-edge are: (1) high and continuous

sediment flux from supply-dominated deltas (e.g., Einsele, 1996;

Burgess and Hovius, 1998), and (2) narrow shelves (Fleming,

1981; Ito and Masuda, 1988). A narrow shelf causes shelf-transittime to be brief and would allow the delta to reach the shelf-edge

before auto-retreat is enacted (Muto and Steel, 1997, 2002).

Narrow shelf settings would clearly sustain shelf-edge deltas

irrespective of whether sea level was falling or rising. Deltas that

are supply-dominated, and driven to the shelf-edge by high

sedimentflux (rather than by negative accommodation)are ableto

transit even moderately wide shelves under conditions of rising

relative sea level. Such supply-dominated deltas not only accrete

at the shelf-margin, but also have the potential to deliver large

volumes of sand to the deepwater slope and basin-floor areas as

occurs in the Maastrichtian Fox HillsLewis system, SE

Wyoming (Carvajal and Steel, 2006). It should also be noted

that this high-supply condition is likely to have an additional

effect. Even where the shelf-margin prism is wide (and therefore

potential transit distance for deltas is great) the transgressivetransit may take the retreating deltas only a short distance back

across the shelf, i.e., the deltas remain on the outer-shelf platform

site throughout a series of cycles (Burgess and Hovius, 1998;

Burgess andSteel, in press). This situation is in contrast to settings

where sea level plays a larger role, i.e., where accommodation-

drive forces deltas to transgress back across much of the shelf

platform during each half-cycle. We illustrate here a case where

deltas reach a shelf-edge position during rising and highstand sea-

Fig. 1. (A) The Central Tertiary Basin and the West Spitsbergen Orogenic Belt (modified from Blythe and Kleinspehn, 1998). (B) The Van Mijenfjorden Group,

showing the Lower Eocene Battfjellet Formation (modified from Steel et al., 1985). (C) Location of the study area in Van Keulenfjorden, showing the location of themountains Storvola and Hyrnestabben, and the 17 measured profiles from both mountains. See also the approximate shelf-break location for Clinoform 17.

230 C.A. Uroza, R.J. Steel / Sedimentary Geology 203 (2008) 229245


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level conditions, not because of high supply (calculation of shelf-

margin accretion/aggradation rates show that the supply was

relatively low), but because shelf width was less than 20 km.

However, despite the deltas being perched at the shelf-edge there

was little sand delivered down into the deepwater areas beyond, to

judge from the absence of slope channels or basin-floor fans. The

deltas simply aggraded at the shelf-margin.

The Eocene Battfjellet Formation in West Spitsbergen,

Norway (Fig. 1) provides good examples of deltas that prograded

to the shelf-edge during both falling and rising sea-level

conditions. The Battfjellet Formation succession is composed ofabout 20 sand-prone, eastwards prograding, shelf-margin clino-

forms (Steel and Olsen, 2002) each deposited during a time

interval of a few 100 ky (4th-order sequences) (Steel and Olsen,

2002; Petter and Steel, 2006). In this study we selected Clinoform

17, which is located towards the top of the Battfjellet Formation.

This clinoform is markedly aggradational along its topset (Fig. 2),

shows landward-interfingering with delta-plain and lagoonal

deposits and lacks deep channelized erosion, characteristics

typical of shoreline successions that accumulate during relative

rise of base level, i.e., with normal regression (Helland-Hansen

and Martinsen, 1996). This aggradational growth style contrasts

greatly with what is observed in some other clinoforms of the

Battfjellet succession (Fig. 3), where there are much thinner andmore amalgamated progradational units into which there are

multiple fluvial incisions, suggesting much flatter shoreline

Fig. 2. General view of the 3 component units 17A, 17B and 17C within Clinoform 17 at Profile 13 location. The clinoform topset succession here is about 80 m thick.Note the two upward-coarsening parasequences in 17A, and the significant thickness (shoreface and barrier/tidal-inlet succession) of unit 17B.

Fig. 3. Schematic shelf-edge trajectories for 4th-order Clinoforms 1417 (Battfjellet Formation). Figure is not to scale andis vertically exaggerated(slope angle: 34).

Note the initial flat trajectories for Clinoforms 14 and 15 (implying stable to falling relative sea level) prior to rise and transgression. For Clinoforms 16 and 17, the

trajectoryis generallycontinuously rising, implying a lack of sea-level fall. Sketch shows partitioning ontothe shelf, slope and basinfloor. Shelf-edge deltas develop on

theshelfsegment of theclinoforms;slopechannels and basin-floorfansare developed in theearlystage of Clinoforms 14 and possibly 15. Clinoform # 14is an exampleof type 1; # 15 could be either type 1 or 2; and # 16 and 17 are examples of type 3. A type 4 clinoform is not present in this succession.

231C.A. Uroza, R.J. Steel / Sedimentary Geology 203 (2008) 229245


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trajectories or even falling relative sea level (e.g., Mellere et al.,

2003). The aggradational style of clinoform 17 implies that much

of the sediment budget is likely to have been trapped within the

coeval coastal plain and shelf segments of the clinoform, with

correspondingly less potential to deliver sand to the deepwater

slope and basin floor (Fig. 3).

The purpose of this paper is to characterize Clinoform 17 interms of its facies associations, external geometry, internal

architecture and sequence stratigraphy. We will evaluate the role

of waves, tides, and river currents in molding the component

sand-bodies and will argue that the entire clinoform developed

under rising relative sea level.

2. Geological setting

2.1. Structure and stratigraphy

Spitsbergen is the largest island of the Svalbard archipelago,

located to the north of mainland Norway (Fig. 1) and in thenorthwest corner of the Barents shelf (Kellogg, 1975). Early

Eocene transpression, as N Greenland slid northwards past the

Barents shelf, during the opening of the NorwegianGreenland

Sea, formed the West Spitsbergen Orogenic Belt (Harland, 1969;

Steel and Worsley, 1984) with areas of basement uplift, folding

and thrusting along a NNWSSE trending fold- and thrust belt

(Fig. 1A). Regional flexural subsidence produced by loading of

thrust sheets along the orogenic belt formed the Central Tertiary

Basin (Steel et al., 1985; Braathen et al., 1999). During initial

periods of activethrusting, the subsiding area was a foreland basin,

but through continued thrusting it may have taken on the character

of a piggy-back basin in the study area (Blythe and Kleinspehn,

1998). Infilling of this basin occurred with moderate subsidencerates (Schellpeper, 2000), with sediments supplied mainly

transversely from the growing orogenic belt to the west

(Helland-Hansen, 1990). The basin filled by the development

and growth of large-scale clinoforms (200400 m high)

representing a linked coastal plain/shelf/slope/deepwater basin-

floor system that systematically migrated eastwards in the Early

Eocene (Helland-Hansen, 1992; Steel and Olsen, 2002). The

sedimentary succession, latest Paleocene through Early Eocene in

age (Manum and Throndsen, 1986), is about 1.5 km in thickness

and has a lithostratigraphy consisting of the Frysjaodden (deep-

water slope and basin floor), Battfjellet (shoreline and shelf) and

Aspelintoppen (coastal plain and estuarine) Formations (Fig. 1B).

2.2. Battfjellet formation clinoforms

The basin transect transverse to the fold-and-thrust belt along

Van Keulenfjorden (Fig. 1) contains some 20 shelf-margin

clinoforms, which have previously been classified into four

types depending on: (1)the progradational trajectory at the shelf-

break, (2) the degree of fluvial channel incision at the shelf-edge,

and (3) the presence or absence of thick sand at this outermost

part of the shelf (see Steel et al., 2000). Type 1 clinoforms show a

flat-to-downward shelf-break trajectory with marked fluvial

erosion and often collapse features, resulting in sand by-pass and

partitioning into the slope and basin floor (e.g. Crabaugh and

Steel, 2004; Petter and Steel, 2006). Type 2 clinoforms have a

flat or low-angle rising trajectory but lack large channels or deep

erosion at the shelf-edge and deposit shelf-attached turbidite

aprons on the slope, without the development of basin-floor fans

(e.g. Plink-Bjrklund et al., 2001; Mellere et al., 2002).

Clinoform 17, presented here, is an example of Type 3 cli-

noforms where there was significant cross-shelf sand transport,but a generally rising shelf-edge trajectory and consequent

aggradational style of the clinoform topsets resulted in only

modest to negligible volumes of sand being delivered onto the

slope or basin floor (see also Deibert et al., 2003). On the outer

shelf, the deltas were substantially reworked by waves, and

much sand was carried alongshore, because of exposure to open

ocean swell and storm waves. Type 4 clinoforms are entirely

muddy on the outer shelf and deepwater reaches because the

deltas did not reach even the outer-shelf areas of the system.

Clinoform Types 1 and 2 are relatively thin (2050 m) along

their topset reaches, and generally show evidence of stable to

falling relative sea level during development. Type 3 clinoforms,described herein, have thicker (N80 m un-decompacted) and

more aggradational (significant marine mudstone interfingering)

topsets, show a more marked parasequence stacking, lack deep

fluvial incisions, and so are interpreted to have developed with

continuously rising relative sea level (Fig. 3).

Clinoform 17 has an intermittently exposed length of about

15 km from Brogniartfjellet to its distal reaches on Hyrnes-

tabben, and a thickness of about 80 to 100 m (un-decompacted)

in the Storvola and Hyrnestabben areas, respectively. Internally,

Clinoform 17 has three distinctive units (Fig. 2), two of them

(17A and 17C) showing regression with aggradation, whereas

the middle one (17B) is initially progradational but becomes

highly aggradational to slightly backstepping.

3. Methodology

The study transect of Clinoform 17 (Fig. 1) is oriented within

20 of a true depositional-dip section. Fieldwork consisted of:

(1) measuring 18 vertical sedimentary profiles (see Fig. 1C for

location of profiles 117) at progressively downdip locations

along 3 mountainsides (Brogniartfjellet, Storvola and Hyrnes-

tabben), located in the Van Keulenfjorden area of West

Spitsbergen, (2) outcrop photographing from both the ground

and helicopter to delineate the general sand-body geometry, and

to aid in the characterization of the deltas and other relatedfacies, and (3) gamma-ray profiles through part of the

succession in three locations in order to compare the gamma-

ray response with most of the facies associations. The external

geometry and internal architecture of Clinoform 17 were

documented using helicopter photomosaics and a constructed

correlation panel of measured profiles. This panel was hung

from a transgressive mudstone level lying above and parallel to

the shelf platform of Clinoform 17.

4. Early Eocene shelf-margin accretion in Van Keulenfjorden

The Spitsbergen shelf-margin (along the exposed transect)

shows a relatively low accretion rate into the basin. The shelf-



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break prograded at a rate of about 5 km/my and aggraded at a rate

of 192 m/my, together suggesting relatively low sediment-supply

conditions compared to other margins of similar clinoform

amplitude (Carvajal and Steel, 2006; Steel et al., in press). This

low apparent rate of sediment supply strongly suggests that the

Spitsbergen deltaic delivery systems would have required the aid

of negative accommodation (sea-level fall) in order to partitionsignificant sand volumes out beyond the shelf-edge into the

deepwater slope and basin floor. However, for Clinoform 17,

there appear to be no coeval deepwater sands, and rather than sea-

level fall, there is evidence of continuous sea-level rise. It is most

likely that it was mainly the narrowness of the shelf that allowed

the deltas to reach the shelf-edge. For both Clinoform 17 and the

underlying Clinoform 16, the deepwater slope was markedly

mud-prone, with only very thin tempestite beds seen on the distal

mountainside of Hyrnestabben.

5. Facies associations

The three units of Clinoform 17 show some common

features in terms of facies associations. The lower and upper

units (17A and 17C) pass upwards from offshore mudstones at

the base, to a wave-dominated delta front, which is then

truncated by tidally-and-fluvially-influenced distributary chan-

nels towards the top of the succession. However, unit 17B

consists of muddy lagoonal facies in its basal and landward

reaches, followed by a thin transgressive mudstone and

overlying shoreface facies that are truncated by tidal-inlet

deposits with an offshore mudstone capping. These deposits are

fronted by a vertically-aggrading barrier island succession.

Facies associations are organized here in two groups: units 17A

and C, and unit 17B. Each group is in some manner described

from shallower to deeper paleowater depth.

5.1. Facies associations for units 17A and C

5.1.1. Tidally-influenced fluvial-distributary channels

This association overlies (typically erosionally) the upward-coarsening deposits of units 17A and 17C (described below in

Section 5.1.2), and shows a general fining-upwards of grain size

(Figs. 4 and 5). The association in unit 17A consists mainly of

fine- to medium-grained cross-stratified sandstones that have

planar and trough cross-strata with westward orientation,

though south to southeastward-oriented cross-strata are also

found (see Fig. 4). Towards West Storvola, the sandstones are

medium- to fine-grained with mainly eastward-oriented planar

and trough cross-strata (but also some south to southeast-

oriented cross-stratification), lack marine indicators, and show

erosional surfaces with coal debris and mud pebbles within the

cross-stratified beds (Fig. 5). In unit 17C, there is an abundanceof eastward-oriented cross-stratification, structureless beds,

convolute bedding, multiple internal erosional surfaces with

coal debris and mud pebbles, and current-ripple laminae atop

the individual sets of cross-strata.

The erosional bases and general fining-upward tendency of the

individual sandstonebodies high in units17A andC, with abundant

cross-strata, and association with underlying delta-front facies

(see Section 5.1.2 below), suggest distributary channels (Bhatta-

charyaand Walker, 1991; Coleman et al., 1964). At some locations,

especially towards the east of Storvola, an abundance of westward-

oriented paleocurrents in the distributarychannels, and the presence

of subordinate southeastward-oriented cross-strata, suggest that

Fig. 4. Tidally-influenced distributary channel deposits from unit 17A, Profile 16 ( Fig. 1), East Storvola. (A) Cross-stratified sandstone at the base of a tidally-

influenced channel fill (35 cm stick as scale). (B) Tidally-influenced channel eroding the flat to swaley cross-stratified delta-front deposits. (C) Medium-grained

sandstone with landward-oriented cross-strata sets (flood-tide generated dunes) and smaller-subordinated ebb-oriented cross-strata on top (section in C is 90 cm thick.(D) Cross-stratification (bidirectional) located few meters laterally of photo C (section in D is 80 cm thick).



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there was some tidal influence in the channels (Dalrymple and

Choi, 2007). Towards the westward (landward) end of the system,

the distributary channels contain abundant eastward-oriented

trough and planar cross-strata, created by the downdip migration

of3D and 2D dunes (Miall, 1978). Here, we also use the consistent

seawards orientation of paleocurrent data (see Fig. 5) as a signal ofthe fluvial influence (see also Selley, 1968; Long, 1978). The

gamma-ray curves in Figs. 4 and 5 show a blocky character for

both channel types, which is typical of such deposits.

5.1.2. Wave-dominated delta-front deposits

This association, identified in units 17A and 17C, is

characterized by upward-coarsening and thickening successions

(up to 57 m thick) containing a high sand/mud ratio, which are

capped by thick-sand channelized units (see earlier Section

5.1.1). The association overlies transitionally the deposits of the

muddy association described below in Section 5.1.3. The

sandstones are mostly well-sorted and fine-grained, but rangefrom very-fine to lower medium-grained. They show a

predominance of flat lamination, swaley cross-stratification

(2080 cm sets that can be followed 10 s of m laterally), and

wave-ripple lamination (Figs. 6 and 7), though hummocky

cross-strata, current-ripple lamination, scattered Ophiomorpha

burrows and soft-sediment deformation structures (load casts,

pillows and water escapes structures) (Figs. 6 and 7) are also

common. Locally, mud pebbles are found within the hummocky

and swaley cross-stratified facies (Fig. 6). Sets of planar and

trough cross-stratification, associated with sets of wave-ripple

laminae, are also found especially towards the top of the

individual coarsening-upward sand packages. It is also common

to find abundant plant and other organic material within the

finer-grained beds. Gamma-ray response (Fig. 6) shows a

decreasing-up pattern for this association.

The upward-coarsening successions of 17A and C are

interpreted as delta-front deposits, with the corresponding

prodelta deposits immediately below, because of their channe-

lized capping (see Fig. 6 and Section 5.1.1 above), the verticalgrain size trend, the vertical thickness of the whole section (up

to 25 m thick) including the prodelta below (see Section 5.1.3),

and the presence of abundant plant and other organic matter in

the finer-grained beds. The occurrence of hummocky and

swaley cross-stratification indicates the strong influence of

storm waves on the delta front (Dott and Bourgeois, 1982;

Leckie and Walker, 1982; Walker et al., 1983; Swift et al., 1983;

Walker and Plint, 1992). The presence of load-cast, pillows and

water escapes structures associated with the swaley and

hummocky cross-stratified beds may possibly result from the

cyclic effect of storm waves on unconsolidated sediments (see

Molina et al., 1998; Alfaro et al., 2002). Also, the occurrence ofmud pebbles within the hummocky and swaley-stratified facies

may correspond to lag deposits (Kreisa, 1981) created by

storms. Intervals of wave-ripple laminae are a common feature

in the upper part of the individual storm-beds, and may indicate

fair-weather wave reworking. The intervals up to medium-

grained sandstone with sets of planar and trough cross-strata

(also capped by wave-ripple lamination) may indicate more

proximal mouth-bar facies on the delta front.

5.1.3. Prodelta to offshore mudstones with occasional thin-

bedded sandstones

These are abundant at the bottom of unit 17A and also

present at the base of unit 17C (Hyrnestabben location). This

Fig. 5. Fluvial-distributary channel facies without obvious tidal influence, unit 17A, Profile 1, West Storvola. (A) Mudstone pebbles and coal debris at the base of

trough cross-stratified sets (20 cm stick as scale). (B) Erosional contact (highlighted by yellow line) between fluvial-distributary channel and delta-front facies (50 cm

stick as scale).



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association is up to 18 m thick and composed mainly of gray

mudstones, but also contains some thin beds of very-fine to fine

sandstone, in beds and bedsets up to 1 m thick. These thin beds

(usually b40 cm) are ripple-laminated or flat-laminated, or

show alternating rippled (mostly current-ripples, but also wave

ripples in some parts) and flat-laminated intervals. Bioturbation

is mostly represented by Phycosiphon and some Planolites

traces.

Because of its stratigraphic position, fine grain size and

marine-like bioturbation, the association is interpreted as off-

shore/prodelta to shelf deposits, with thin turbidite-like beds (flat-

to-ripple-laminated beds) that are probably tempestites because

Fig. 7. Common sedimentary structures in the delta front: (A) soft-sediment deformation (load casts and water escape structures) at base of the sandy delta-front

succession of unit 17A, Profile 15, Storvola (meter stick as scale). (B), (C) and (D) Symmetrical wave ripples from Profiles 6, 10, and 17 respectively. 20 cm stick asscale in D.

Fig. 6. Upward-coarsening, wave-dominated delta-front facies (including base of distributary channel on top of delta front); Profile 1, West Storvola ( Fig. 1) (vertical

scale in meters). (A) Stacked swaley cross-stratified sandstone sets. (B) An 80 cm thick section of swales alternating with wave-ripple lamination. (C) Soft-sediment

deformation at the bottom of the delta front (50 cm meter stick as scale).



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of their wave-rippled capping (see also Myrow and Southard,

1991, 1996) and their updip association with wave-dominated

shoreline deposits. Reineck and Singh (1972) described similar

laminated sand beds within shelf mudstones from the modern

North Sea and interpreted them as storm deposits.

5.1.4. Upper-slope turbidite-like beds (tempestites)These are found on the mountainside of Hyrnestabben

mostly at the bottom of unit 17A, and they occupy an upper-

slope setting on the clinoform, just basinwards of the shelf-slope

break. Updip, they are also associated with the delta front

described in Section 5.1.2. These beds are composed of sharp-

based, thin sets of flat- to ripple-laminated, fine to very-fine

sandstone, interbedded with mudstones. It is common to find

Phycosiphon trace fossils within the muddy beds (A. Uchmann,

2004, personal communication).

Because of their slope setting and association with an updip

delta front that is storm-wave dominated, these beds are likely to

be storm-wave generated from the shelf-edge area and drivenout onto the upper slope as tempestites during storms (see also

Myrow and Southard, 1991, 1996).

5.2. Facies associations for unit 17B

5.2.1. Fluvial crevasse channels and splays (within coastal

plain deposits)

This association is preserved on west Storvola (see panel in

Fig. 12) and on Brogniartfjellet (west of Storvola). It comprises

relatively thin (11.5 m thick), channelized successions of fine-

grained sandstone with sets of trough cross-strata and current-

ripple lamination towards their top (Fig. 8). Structureless

sandstone is common and soft-sediment deformation is also

present. These channelized units are frequently capped by coal

layers. Within this association, there are also thin sheet-like

packages of fine to very-fine sandstone with current-ripples and

wavy bedding.

The abundant coal layers capping the channels, suggestsdeposition within the coastal plain and rapid abandonment to

areas of vegetation (see Guion, 1984; Fielding, 1985). This, in

turn, strongly suggests that these thin channelized successions

do not reflect distributary channels but are ephemeral crevasse

channels, with their associated crevasse splays (sheet-like sand-

bodies), that occasionally broke out from the distributaries

during floods (Elliott, 1974; Fielding, 1984). The sandstone

bodies of this association are similar in character to those

described by Plink-Bjrklund (2005) for the coastal plain of the

Aspelintoppen Formation on Brogniartfjellet and Storvola.

5.2.2. Lagoonal deposits (brackish-water mudstones)This association is composed of mudstone deposits, which

are rich in carbonaceous matter (thin coal layers, plant and coal

fragments), and is mainly located westwards of the tidal-inlet-

channel deposits of unit 17B (see panel in Fig. 12). Towards the

west of Storvola, they are also interbedded with the fluvial

crevasse channel and crevasse-splay deposits, described above

in Section 5.2.1.

We interpret these deposits as brackish-water lagoonal

mudstones because of their abundant organic content (Reading

and Collinson, 1996) and the absence of marine fauna (D. Van

Nieuwenhuise, 2007, personal communication). They also

Fig. 8. Fluvial channel and crevasse-splay deposits from unit 17B, Profile 2, West Storvola. (A) Trough cross-stratified sandstone at base of channel (85 cm stick asscale). (B) Rippled sandstone probably associated with crevasse-splay deposits. 17 cm stick as scale (lower right of photo).



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occupy a position behind the tidal-inlet-barrier complex of unit

17B (see Fig. 12).

5.2.3. Tidal-inlet deposits

This association is erosionally based, dominates unit 17B in

places and is up to 15 m thick. It is composed of clean, well-to-

moderate sorted, fine to coarse-grained sandstone, with cross-sets that have characteristic and persistent northwestward-

directed paleocurrents (Fig. 9). Individual cross-sets are up to

1.5 m. thick, and some are capped by wave-ripple lamination.

This association is generally in erosional contact with the

underlying shoreface and barrier bar deposits (Figs. 9 and 12).

The gamma-ray response of this association is blocky in

character, with a slight increasing-upward trend towards the top

of the succession (see Fig. 9).

These deposits are interpreted as inlet-channel infill (Hoyt

and Henry, 1967; Kumar and Sanders, 1974; Moslow and Tye,

1985) because of their location between lagoonal and barrier

deposits, their marked and deep channelized erosion andpersistent flood-tidal paleocurrents (northwestward-oriented),

great thickness, and their incision into the wave-generated

shoreface and barrier deposits.

5.2.4. Shoreface

This association, particularly found at the bottom of unit

17B, is up to 9 m thick, and contains some coarsening-upward

packages (up to 35 m thick). Plane parallel-laminated fine-

grained sandstones with wave-ripple capping, and hummocky/

swaley cross-stratified sandstone sets, also capped by wave-

ripple lamination, dominate the association (Fig. 10). Bioturba-

tion is common, with mainly Ophiomorpha burrows (Fig. 10B)

and soft-sediment deformation (mainly load casts) is also

prominent. Some planar and trough cross-strata, wavy lamina-tion, and current-ripple lamination, also occur. Gamma-ray

response, for this association, shows a decreasing upward

pattern (Fig. 10).

We interpret this facies association as shoreface deposits,

based on the following: (1) it is a coarsening-upward facies

succession (above marine mudstones) that was deposited during

coastal progradation (sensu Walker and Plint, 1992), (2)

occasionally there are sharp-based sandstone beds with hum-

mocky cross-strata and wave-ripple lamination, which corre-

spond to storm-beds (see also Dott and Bourgeois, 1982; Walker

et al., 1983), (3) there is an absence of a fluvial feeder landwards

of the association (see Fig. 12), (4) this wave-dominatedsuccession did not prograde as far as the earlier deltaic system.

5.2.5. Sandy barrier complex

This association occurs within unit 17B and it is represented

by an upward-coarsening succession (up to 13 m thick) (Fig. 11)

with similar facies to those described above in Section 5.2.4.

Sandstone is mostly fine-grained, well-sorted, with mainly

Fig. 9. Tidal-inlet-channel facies from unit 17B, Profile 9, Mid Storvola. (A) Channel eroding into the shoreface facies (40 cm stick at lower portion of photo). (B) and(C) Tabular cross-bedding oriented landwards (45 cm stick as scale in B, and 30 cm in C).



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swaley cross-stratified and plane parallel-laminated sets, though

hummocky cross-strata sets are also common. Individual

bedsets are 1040 cm thick and commonly capped by wave-

ripple lamination (Fig. 11). Sets of swaley cross-strata within

this association reach a few m in wavelength. Ophiomorpha

burrows are common, and minor low-angle cross-bedded

sandstones are also found within the succession.

All the features described above are indicative of a wave-

dominated shoreface setting (Walker and Plint, 1992). However,

the position of this association near the shelf-slope break (see

Fig. 11. Barrier-bar facies from the seaward end of unit 17B, Profile 16, East Storvola. (A) Swaley cross-stratification (50 cm stick as scale). (B) Plane-parallel and

wave-ripple lamination in alternation within sandstones. (C) Close-up of middle portion of A showing SCS (30 cm stick as scale). (D) Swaley cross-stratifiedsandstone capped by wave-ripple laminated sandstone. Walking stick scale in B and D: 70 cm.

Fig. 10. Shoreface facies from unit 17B (regressive component), Profile 9, Storvola. (A) Shoreface facies truncated by tidal-inlet channel (see truncation T

highlighted by red line). Outcrop section in A is about 2.2 m. (B) Flat-laminated sandstone with Ophiomorpha burrows (see black pen as scale).



11/17

Fig. 12. Correlation Panel and helicopter photomosaics for Clinoform 17, Battfjellet Formation, West Spitsbergen. The whole clinoform is interpreted to have developed du

located towards the eastern and western ends of Storvola respectively.


12/17

Fig. 12) and its thickened aggradational character (Fig. 11)

suggest that this association is a barrier complex (McCubbin,

1981; Reinson, 1992; Friis et al., 1998) related to the initiation

of transgression in the system. This barrier complex is mostly

developed on eastern Storvola and pinches out into mudstones

between Storvola and Hyrnestabben (see correlation panel in

Fig. 12).

6. Clinoform 17: geometry, architecture, and formative

processes

6.1. Geometry and architecture

The geometric configuration of Clinoform 17 as a whole, as

well as of the component sand-bodies of this deltaic and barrier

lagoon succession, is illustrated in Fig. 12. The clinoform is

mainly seen in a 2-D, near-downdip section. Similarly to the

stratigraphic configuration of the underlying clinoforms (num-

bers 12

16) on Storvola, a shelf-slope break is preserved inClinoform 17. This slope break must occur between the two

mountains, because Storvola contains the shelf-edge deltas,

whereas on Hyrnestabben the same clinoform mostly shows

muddy slope deposits with thin tempestite sands (Fig. 12).

Fig. 12 clearly illustrates the general rising character of the

shoreline trajectory within the progradational trend of the

Clinoform 17 shoreline system (yellow and light red colors),

and the coeval, parallel rising belts of coastal plain/lagoonal

(green colors) and offshore (blue) deposits. In addition tothese topset components of the clinoform, the shale-prone,

upper-slope component of the shelf-margin can also be seen at

the right-hand end of the correlation panel on the mountain

Hyrnestabben. Complicating this overall stratigraphy in

Fig. 12, the three main phases of shoreline development are

evident, namely an early regressive phase (17A), a middle

aggradational to slightly transgressive phase (17B), and a late

regressive phase (17C). Phases 17A and C are thick wave-

dominated delta successions with well-preserved feeder

distributary channels emphasizing the normal character of

shoreline progradation. Phase 17B represents an intervening

interval of initial coastline progradation with aggradation andthen slight transgression, during which an inlet-channel and

barrier/lagoon coast developed. This temporary aggradation/

Fig. 13. Schematic summary of the depositional history of Clinoform 17, including the depositional setting for the three component units. The whole deltaic system isinterpreted to have been deposited under highstand conditions. Figure is vertically exaggerated. Slope angle: 3 4.



13/17

transgression in 17B further emphasizes the overall aggrada-

tional character of Clinoform 17.

Clinoform 17 resembles the geometry of other documented

clinoform successions that developed associated with a rising

shelf-edge trajectory. For example, Bullimore et al. (2005)

documented seismic-scale clinoforms with high-angle positive

shelf-edge trajectory in the Norwegian Sea Molo Formation.Here, they attributed this condition to normal regression in

which coastal plain units were deposited and preserved in the

topset segment of the clinoforms (see Bullimore et al., 2005,

Figs. 6 and 8).

6.2. Variability of depositional processes on the clinoform

Clinoform 17 shows evidence of an interplay of fluvial,

wave, tidal, and storm ebb-surge processes, though the wave

domination on the open coast is clear throughout the record of

the overall regressive-aggradational shelf transit of the deltaic

system. Wave-domination is evidenced both on the delta-frontdeposits of units 17A (Figs. 6 and 7) and 17C, and the shoreface

(Fig. 10) and barrier facies (Fig. 11) of unit 17B. Storm ebb-

surge processes (Mount, 1982; Dott and Bourgeois, 1982;

Cheel, 1991) acted at different times during the regressive

transit of the delta system, and they were responsible for the

deposition of the thin tempestite sand beds within the mud-

prone prodelta to offshore and upper-slope environments, and

probably the hummocky cross-stratified beds. Strong tidal

influence, on the other hand, is seen especially well in the

aggradational to transgressive deposits of unit 17B (Fig. 9), and

in the distributary channels of units 17A (Fig. 4) and 17C.

Fluvial influence is implicit at all times during the coastal

regression and is especially implied by the system reaching the

shelf-break area despite the overall aggradational tendency. The

fluvially-influenced portion of the distributary channels is well-

preserved in 17A (Fig. 5) and 17C.

7. Discussion

7.1. Clinoform 17: a highstand shelf-edge delta system

The characteristics of Clinoform 17 described above show

clearly that:

It is a sand-rich delta system.

The delta system was sited on the shelf-margin, near the

shelf-slope break of the clinoform (Figs. 12, 13 and 14).

The markedly aggradational character of the delta complexduring progradation strongly suggests that relative sea level was

rising during its development (see also Helland-Hansen and

Gjelberg, 1994; Bullimore et al., 2005). There is no evidence of

forced regression (sensu Posamentier et al., 1992) or relative sea-

level fall. The partly aggradational/transgressive middle interval

(unit 17B) in Clinoform 17 further reinforces its overall

aggradational character (Figs. 12 and 13).

Despite its shelf-slope break site, the deltas were apparently

able to deliver only modest volumes of sand out onto the

Fig. 14. The conventional lowstand model (A) vs. the highstand model (B) proposed here for clinoform growth, Battfjellet Formation, West Spitsbergen.



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adjacent deepwater slope, and there was no visible development

of turbidite slope channels, or basin-floor fans (see also Deibert

et al., 2003) despite relatively good exposure on adjacent

mountains.

These features together strongly suggest that Clinoform 17

represents a highstand (rising relative sea level) shelf-edge deltasystem that crossed a relatively narrow shelf (Fig. 13). The

occurrence of highstand shelf-edge deltas is somewhat unusual

in the literature because falling relative sea level is commonly

invoked to account for deltas at the shelf-edge, giving rise to

lowstand shelf-edge deltas (Muto and Steel, 2002; Porbski and

Steel, 2003) (Fig. 14). Clinoform 17 can therefore be used to

test some literature concepts concerning highstand deltas at the

shelf-edge.

7.2. Clinoform 17: an outcrop test of highstand shelf-edge

deltas

7.2.1. Occurrence of shelf-edge deltas during highstand

Although no previous highstand shelf-edge deltas have

been described from outcrops, there have been a number of

suggestions as to the conceptual feasibility of highstand deltas

reaching the shelf-edge. Initial suggestions were made by

Burgess and Hovius (1998), especially for short shelf-transit

distances. Such occurrences were disputed by Muto and Steel

(2001), who emphasized the auto-retreat tendencies of deltas

during rising sea level, especially during long shelf transits.

However, later subsurface descriptions of highstand shelf-

edge deltas (50100 km wide shelves), made by Carvajal and

Steel (2006), concluded that this was possible because of a

supply domination during shelf transit. Also, Hiscott (2003)documented highstand delta lobes (though muddy), sited on

the outer continental shelf to upper slope, from the Late

Quaternary Baram delta of northwestern Borneo. This history,

plus the outcrop evidence from Clinoform 17 herein, suggests

that:

Most deltas can attain a shelf-edge position at lowstand of

sea level, irrespective of shelf width, and generate sands into

deepwater areas. This scenario is the conventional one, is

accommodation driven, and does not require a high fluvial

drive (see also Porbski and Steel, 2006; Yoshida et al., 2007)

(Fig. 14A).Deltas can attain a shelf-edge position at highstand of sea

level if sediment supply is very high, irrespective of shelf width,

and can also generate deepwater sands. This is the supply-

driven scenario (see Carvajal and Steel, 2006).

Even low-supply delta systems, such as has been argued for

Clinoform 17, can reach the shelf-edge as highstand deltas if

shelf width is narrow. Nevertheless slopes are likely to be

muddy, and delivery of deepwater sand is likely to be minimal,

unless the shelf width is reduced severely.

The delivery and accumulation of large volumes of deep-

water sands is most favored by (a) high sediment supply, (b)

deltas drawn to the shelf-edge by falling relative sea level, and

(c) significantly narrow shelves.

7.2.2. Wave influence at the shelf-edge

Other characteristics of highstand deltas have been proposed.

For example, Porbski and Steel (2006) and Yoshida et al.

(2007) have suggested that shelf-edge deltas should normally be

wave-dominated, because waves arriving at the shelf-edge from

the open ocean are large, but tend to become smaller as they

cross the shelf. This was also observed by Sydow et al. (2003) intheir study of reservoirs near the Pliocene Orinoco shelf-edge in

offshore Trinidad. However, Suter and Berryhill (1985); Morton

and Suter (1996); and Roberts et al. (2000) showed their studied

Pleistocene shelf-margin deltas as fluvial-dominated with some

wave-modification. On the other hand, Cummings et al. (2006)

recorded strong tidal signals on the front of Cretaceous deltas

near the Nova Scotia shelf-edge, but showed that this was

probably due to the deltas being sited within a shelf-edge

embayment. Shelf-edges are probably more likely to be

embayed at sea-level lowstand, whereas shorelines on highstand

shelf-edges are more likely to be straight and open, and

therefore wave-dominated. This was also suggested anddocumented by Ainsworth et al. (in press) from subsurface

reservoir data. The outcrop data presented here on Clinoform 17

are consistent with a relationship between rising sea level

(irrespective of the cause of rise) at the shelf-edge and the

occurrence of wave-dominated shorelines.

7.2.3. Enhanced thickness and parasequence development in

highstand shelf-edge delta units

The greater thickness of regressive deltaic units at the shelf-

edge compared to the inner shelf is obvious because of the

greater water depth normally encountered towards the outer

shelf. This increased shelf-edge thickness was suggested to be

further enhanced by rising relative sea level, and the propositionmade that this thickness increase would be accompanied by an

increased number of parasequences in the unit (Porbski and

Steel, 2006). In the study of the Orinoco Pliocene shelf-margin

reservoirs (Sydow et al., 2003), rising relative sea level created

by high subsidence rates caused the wave-dominated sequences

to be unusually thick, and with multiple parasequences (N150 m

in places). This thickness aspect is consistent with the character

of the Clinoform 17 data. Contrary to this, Morton and Suter

(1996) reported a significant thickness reduction in several late

Quaternary deltaic sequences (Gulf of Mexico) that were

formed under a rapid lowering in sea level.

8. Conclusions

Shelf-margin deltas preserved in Clinoform 17 of the Eocene

Battfjellet Formation on Spitsbergen are argued to have

developed during conditions of rising relative sea level

(highstand), because of their thick aggradational and parase-

quence-prone character. The wave-dominated character of the

delta-front and barrier deposits, suggesting an open, straight

coast, is consistent with this. The strong fluvial drive and high

flux of sediment normally associated with highstand shelf-edge

deltas cannot be argued here because the shelf-margin

progradation rate is estimated to have been low. So, the reason

the Clinoform 17 deltas reached the shelf-break was the



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narrowness of the shelf itself. If the shelf had been wider, these

low-supply deltas would have retreated before reaching the

outer shelf, in accordance with auto-retreat principles.

The well-exposed outcrops of Clinoform 17 add to our

knowledge of highstand shelf-edge deltas, and suggest the

following characteristic features:

Deltaic units are unusually thick because they occurred in

deeper water at the shelf-edge and because rising relative sea

level caused aggradation of the body and a rising shelf-edge

trajectory.

The aggradation of the sediment body was achieved by a

vertical stacking of parasequences. This also caused the shelf-

edge delta to contain short-lived shale-prone transgressive

incursions.

The deltas are wave-dominated, as would be expected

because shelf-edge areas are almost always exposed to open

ocean waves. Where coastline morphology is less straight,

signals of tidal influence would be better preserved.Highstand deltas in general deposit and store a large

proportion of their sediment budget on the shelf and coastal

plain, whereas falling-stage deltas deliver more of their

sediment into deepwater areas. Where the shelf-edge deltas

are of high supply type, deepwater sands might still be

expected beyond the shelf-edge. Where supply is low, as in the

studied succession, the deltas reached the shelf-margin because

the shelf was narrow, and the deepwater slope and basin floor

tended to be mud-prone.

Because highstand shelf-edge deltas are aggradational, they

contain no major widespread erosional surface or sequence

boundary, in contrast to falling-stage deltas.

Acknowledgements

We thank the WOLF Consortium (BP, Norsk Hydro, Statoil,

Shell, BHP Billiton, ConocoPhillips, and Pdvsa) for their

financial support to cover the field work and for lively discussion.

Thanks to the Jackson School of Geosciences at the University of

Texas at Austin for their administrative support. Also, the authors

recognize the valuable contribution of Atle Folkestad, Andrew

Petter, Piret Plink-Bjrklund, Alfred Uchmann, and Cristian

Carvajal, through assistance in the field or contribution to this

research. Dr. Marc Edwards is thanked for a constructive review

of an early draft of this paper. We especially acknowledge Drs.Chris Fielding, M. Royhan Gani, and Jesus Soria, for their

constructive criticism in reviewing the manuscript.

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Documents

Highstand Shelf-margin Delta System