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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.
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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.
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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).
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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.
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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.
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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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