GEOLOGIC NOTE
AUTHORS
Hyperpycnal flow variability
and slope organization
on an Eocene shelf margin,
Central Basin, Spitsbergen
Andrew L. Petter and Ronald J. Steel
ABSTRACT
Identification of bypass at the shelf margin is critical to deep-water
exploration. We examine the shelf margin of an early Eocene fourthorder sequence with an attached basin-floor fan in the Spitsbergen
Central Basin. Turbidity currents were fed mainly by hyperpycnal flow emerging from shelf-edge deltas. The life span of any
turbidity current was determined primarily by the sediment concentration of the flow and the duration of the river flood. Highdensity hyperpycnal flows created sand-filled slope-channel complexes 10–15 m (33–49 ft) thick and 100–200 m (328–656 ft)
wide that served as conduits for bypass to the basin floor. Lowdensity hyperpycnal flows were unconfined and deposited heterolithic lobes on the slope. Shelf-margin accretion of about 1.5 km
(0.9 mi) during the falling stage gave way abruptly to bypass in the
early lowstand. Most of the basin-floor fan growth was achieved
after shelf-edge incision and before relative sea level rise. Coastalplain aggradation in the late lowstand sequestered sediment from
the shelf-edge distributaries, effectively diminishing high-density
hyperpycnal flow output. The late lowstand was therefore marked
by a second phase of shelf-margin accretion with only limited bypass
to the basin floor, and a heterolithic, prograding complex downlapped the early lowstand channels. Transgression ultimately led to
the abandonment of the shelf-edge delta complex and the accumulation of mainly mudstone on the margin. The shelf-margin architecture exhibited by this sequence should serve as a type example
of a deep-water feeder system in which hyperpycnal flow is the
primary initiator of turbidity currents for sand accumulation on
the slope and basin floor.
Copyright #2006. The American Association of Petroleum Geologists. All rights reserved.
Manuscript received September 15, 2005; provisional acceptance January 10, 2006; revised manuscript
received April 10, 2006; final acceptance April 24, 2006.
DOI:10.1306/04240605144
AAPG Bulletin, v. 90, no. 10 (October 2006), pp. 1451 –1472
1451
Andrew L. Petter Department of Geological Sciences, University of Texas at Austin,
1 University Station, C-1100, Austin, Texas
78712; petter@mail.utexas.edu
Andrew L. Petter is a Ph.D. aspirant in the
Jackson School of Geosciences, University of
Texas at Austin. He earned an M.S. degree in
geological sciences from the University of Texas
at Austin in 2005, where he focused on the
topic covered in this article. He is currently
studying the Paleogene shelf margin of the Gulf
of Mexico and its relationship to transport of
reservoir-quality sediment to the deep water.
Ronald J. Steel Department of Geological Sciences, University of Texas at Austin,
1 University Station, C-1100, Austin, Texas,
78712
A University of Texas professor, Ron Steel
received his B.Sc. degree in 1967 and his Ph.D.
in 1971 from the University of Glasgow. His
research is aimed primarily at using clastic
sedimentology to address problems in basin
analysis and particularly to decipher the signatures of tectonics, sea level change, and
sediment supply in stratigraphic successions.
ACKNOWLEDGEMENTS
We recognize the financial support of the
WOLF Consortium (BP, Norsk Hydro, ExxonMobil, Shell, BHP Billiton, PDVSA, and ConocoPhillips) for fieldwork. A.L. Petter also thanks
the Jackson School of Geosciences, the University
of Texas at Austin, and ConocoPhillips for additional funding of his studies. We are also appreciative of Stuart Blackwood, Cristian Carvajal,
Piret Plink-Björklund, Alfred Uchmann, and
Carlos Uroza, who contributed to this paper
through discussions with the authors and/or
assistance in the field. Craig Fulthorpe, William
Galloway, Richard Moiola, and Michael Sweeney
provided insightful and constructive reviews
of this paper and are thanked for their efforts.
INTRODUCTION
The Tertiary Central Basin of Spitsbergen (Figure 1) is
well known for spectacular exposures of shelf-margin
and basin-scale clinoforms (Helland-Hansen, 1992; Mellere et al., 2002; Steel and Olsen, 2002; Crabaugh and
Steel, 2004; Plink-Björklund and Steel, 2004; Clark
and Steel, 2006). These clinoforms are composed of
linked facies tracts from the shelf to the basin that recorded the early Paleogene filling of the foreland basin.
Turbidites are well exposed in the basin and can be correlated updip to coeval shallow-marine deposits at the
shelf edge. The present study documents the shelf to
middle-slope deposits of an early Eocene shelf-margin
clinoform from the Tertiary Central Basin of Spitsbergen.
Earlier work on this clinoform (Crabaugh and Steel,
2004; Clark and Steel, 2006) has focused on an attached sand-rich basin-floor fan, making it an ideal candidate on which to examine sand bypass from the shelf
edge and upper slope to the basin floor. This work encompasses the facies and architecture of the shelf, shelf
break, and upper- to middle-slope environments, which
were strongly dominated by fluvial processes while the
sediment delivery system was located on the shelf edge.
In particular, sandy hyperpycnites (fluvially generated
density-flow deposits) are abundant and well exposed
along the clinoform of interest from the upper slope to
the basin floor. Recent workers (Normark and Piper,
1991; Mulder and Syvitski, 1995; Mulder et al., 2003)
have suggested that these quasi-steady density flows may
be much more common than previously believed. The
particular exposures described here, from clinoform 14
(see Steel and Olsen, 2002) on the Storvola mountainside, offer a unique opportunity to examine continuousflow turbidites from their origin at the shelf edge through
deposition on the slope and bypass to the basin floor.
Hyperpycnal Flows
Figure 1. (A) Map of Spitsbergen illustrating the relationship
of the Tertiary Central Basin (CB) and the Tertiary orogenic belt
(WSO). (B) Map of study area showing locations of measured
sections, gross depositional environments, and mountains (e.g.,
Storvola, Hyrnestabben) (created with Generic Mapping Tools
[GMT ] through the Web site, Online Map Creation [Online Map
Creation, 2004]).
1452
Geologic Note
Hyperpycnal flows were originally defined as density
outflows from a river mouth with a density greater than
the ambient fluid in which they flow (Bates, 1953).
Consequently, hyperpycnal flows are bottom-riding
density flows that are thought to occur primarily because of extreme river-flood events. Hyperpycnal flows
can evolve into turbidity currents, and turbidity currents generated in this manner are commonly referred
to as hyperpycnal flows. As turbidity currents, the sediment load of hyperpycnal flows is implied to be suspended particles of medium- or finer grained sand
(Lowe, 1982; Mulder and Syvitski, 1995). Mulder et al.
(2003) stated that turbidity currents generated by hyperpycnal flow should behave as low-density turbidity
currents (sensu Lowe, 1982). However, other authors
(e.g., Mutti et al., 2003) propose a broader spectrum
of flow behaviors for hyperpycnal flow, ranging from
dense flows (including debris flows, hyperconcentrated
flows, and high-density turbidity currents) that would
incorporate flushed, coarse-grained fluvial bed load,
to low-density turbidity currents. Prior and Bornhold
(1989) and Normark and Piper (1991) both suggested
that river bed load could continue beyond the delta
front as an inertial flow. Normark and Piper (1991)
theorized that hyperpycnal flows from high-bed-load
rivers would be more likely to initiate turbidity currents
because of hyperconcentration of discharged sediment
from these rivers. Previous workers have generally documented hyperpycnal flows on the shelf because of present highstand conditions. Momentum changes caused
by the gradient increase at the shelf edge would augment the inertial forces acting on the fluvial bed-load
discharge. The shelf edge is one of two locations along
the depositional profile that show a significant gradient
increase in the downdip direction, the other position
being the coastal prism (Talling, 1998). Therefore, hyperpycnal flows generated from shelf-edge deltas should
be more likely to continue downslope inertially than
previously described shelf examples.
Recognition Criteria
Hyperpycnites range from mud and silt (Mulder et al.,
2001a; Nakajima, 2006) to sand-rich (Plink-Björklund
et al., 2001; Plink-Björklund and Steel, 2004) deposits;
other flood-generated deposits may consist of gravel- to
boulder-size material (Prior and Bornhold, 1989; Mutti
et al., 2003). Theoretically, hyperpycnites should exhibit vertical trends that reflect the waxing-to-waning
energy of the flood hydrograph (Mulder et al., 2001a).
However, Best et al. (2005) have shown temporal fluctuations in flow velocity that are attributed to lateral
shifting of the jet flow across the delta front. This shifting perhaps explains frequently encountered internal discontinuities and absent sections in the waxing-to-waning
trend of hyperpycnites. Hyperpycnites are characterized by upper-stage, plane-parallel lamination, climbing
ripples, and abundant organic matter of terrestrial origin, including leaves and other plant matter that would
not commonly be preserved in mouth-bar deposits (PlinkBjörklund et al., 2001; Plink-Björklund and Steel, 2004).
Internal scours can be onlapped or downlapped by planeparallel lamination in a manner that suggests cut and laterally accreting fill. Additionally, hyperpycnites in slope
channels can be recognized by the connection of the
channel updip with shelf-edge distributary channels.
Conditions Conducive to Formation of Marine Hyperpycnal Flow
The high density of seawater (1.02 g/m3) relative to
fresh water means that quite limited conditions will
be conducive to the formation of marine hyperpycnal
flows (Mulder et al., 2003). Mulder and Syvitski (1995)
calculated a critical sediment concentration threshold
of 36 – 43 kg/m3 (dependent on temperature and salinity of the seawater) to achieve flow bulk density in
excess of this value; however, this value has been challenged by Parsons et al. (2001), who experimentally
created hyperpycnal plumes at concentrations 40 times
lower than calculated by Mulder and Syvitski (1995).
The suspended load of rivers can be related to the
river discharge by a power law (Mulder et al., 2003).
Therefore, increased river discharge can lead to an exponential increase in sediment concentration. Major
floods have the potential to carry more sediment than
the rivers would carry during several years of normal
discharge. Additionally, major floods would pass a large
influx of fresh water into the receiving basin, lowering
the flow bulk density required for the discharge to become hyperpycnal.
Variability of Hyperpycnal Flows
Two types of hyperpycnal flow have been described in
the literature. Wright et al. (1986) observed the collapse of sediment-laden water into a bottom-riding current on the Yellow River delta. They termed these
plunging flows ‘‘low-density’’ hyperpycnal plumes because the density of the flow was too low to entrain
sediment by erosion of the underlying bed. Wright et al.
(1986) proposed a negative feedback mechanism for
low-density hyperpycnal plumes that would lead them
to die abruptly. Deposition of suspended sediment would
cause the sediment concentration to decrease, decelerating the current and forcing further deposition. Lowdensity hyperpycnal flows would most likely behave
as low-density turbidity currents of Lowe (1982).
Wright et al. (1986) also observed channels on the
Yellow River delta front that they hypothesized were
created by high-density hyperpycnal plumes. Highdensity hyperpycnal flows are more energetic than lowdensity hyperpycnal flows and therefore have the potential to erode the underlying substrate and form channels.
These flows require a higher sediment concentration and
are therefore likely to form only during enhanced fluvial
discharge (i.e., major floods). High-density hyperpycnal
flows can potentially have the structure of high-density
turbidity currents (Lowe, 1982), in which turbulence
in the basal zone of the flow is dampened by the hyperconcentration of particles. This dampening of turbulence, however, would decrease the erosive potential of the flow. Therefore, channelized high-density
hyperpycnal flows must initially exist in a state intermediate between low- and high-density flows to be
Petter and Steel
1453
fully turbulent yet still energetic enough to cut channels. Waxing discharge in the initial flood stage of rivers would allow the evolution of hyperpycnal flows
from low-density flows to turbulent, energetic flows
with an intermediate density to true high-density, hyperconcentrated flows. The high energy of high-density
hyperpycnal flows makes them more likely candidates
for long-distance runout and bypass to the basin floor.
Both high- and low-density hyperpycnites are interpreted in the early Eocene shelf-margin succession of
the Central Basin.
Recurrence of Hyperpycnal Flow
Holocene highstand examples of hyperpycnal flow illuminate the frequency of hyperpycnal events, which
can commonly occur at subdecadal to centennial time
scales (Mulder and Syvitski, 1995; Johnson et al., 2001;
Mulder et al., 2001b; Warrick and Milliman, 2003) and
transport a large fraction of the coastal sediment budget
( Warrick and Milliman, 2003). Mulder and Syvitski
(1995) concluded that hyperpycnal discharge was most
likely in so-called ‘‘dirty’’ rivers; i.e., small rivers with
a mountainous drainage basin. Rivers with large, lowrelief drainage basins were found to be unlikely candidates to produce hyperpycnal flows because of the
capacity of the flood plain to capture floodwaters.
In areas with narrow shelves, it was also found that
some hyperpycnal flows could bypass the shelf and be
transported through submarine canyons during highstand as turbidity currents (e.g., Monterrey Canyon,
Johnson et al., 2001; Var Canyon, Mulder et al., 2001b;
Santa Barbara and Santa Monica basins, Warrick and
Milliman, 2003). If deltas are brought directly to the
shelf edge, as what could happen during relative sea level
lowstand, then distributary channels would discharge
directly onto the upper slope, and the fluvial system
would be directly connected to the basin-floor fan via
slope channels. The frequency of hyperpycnal flows
reaching the deep water would be increased greatly.
Order or Lack of Order on the Slope
Deep-water basin-floor architectures have generally
been well documented, and elegant models of turbidite systems illustrating the relationship between basinfloor fans and submarine channels on the basin floor are
well established (Weimer, 1990; Posamentier et al.,
1991; Normark et al., 1993; Mutti et al., 1994; Reading
and Richards, 1994; Bouma, 2004). These models, however, have tended to overlook or oversimplify the shelfedge and slope system that serves as sediment feeder and
1454
Geologic Note
conduit, respectively, for the basin-floor deposits. The
shelf edge and upper slope are a critical factor in the
delivery and dispersal of sand-rich sediments to the deep
basin (Pore˛bski and Steel, 2003). Therefore, these turbidite models are extremely useful as reservoir analogs
and are invaluable in the development and production
phase of hydrocarbon exploitation. Initial exploration
for basin-floor deposits, however, is better facilitated by
models of linked shelf-edge and slope sediment-delivery
systems. The sequence architecture of these systems
should also be reflected in the deep basin.
Lomas (1999) argued against systematic stacking
patterns in slope systems based on his study of a Lower
Cretaceous slope apron in Antarctica. The claim was
made that deposition there was controlled by episodic
depositional events, and that order in slope-apron stacking reflected sediment-transport vectors controlled solely
by structural trends. An objective in the present article
is to document that there can be significant time-space
order in slope architectures during a cycle in which
sediment supply interacts with accommodation change
on the shelf margin. Systematic stacking patterns of
slope architecture are difficult to observe when the slope
or parts of the slope are observed in isolation. However,
slope architectures can become significantly more ordered when viewed within the context of a sequencestratigraphic framework established from the shelf and
shelf edge, despite the episodic nature of depositional
events on the slope. This article will also discuss the
relationship between dynamic depositional processes at
the shelf edge and stacking patterns on the slope.
Vail and Wornardt (1990) showed an early sequencestratigraphic cartoon of systems tracts across the shelf,
slope, and basin floor based on type-log responses. Bypass of sediment down the slope is suggested by the
presence of a sand-prone basin-floor fan and slope fan.
However, slope bypass conduits were not included in
the model, ignoring the connection between the shelf
delivery and basinal sand accumulation. Kolla (1993)
did draw attention to time and space order in the shelfedge–slope-basin system, but again did not discuss the
character of the delivery system other than to include
early lowstand canyons.
LOCATION AND GEOLOGICAL SETTING OF
STUDY SYSTEM
The Tertiary Central Basin of Spitsbergen is located in
the Arctic archipelago of Svalbard (Figure 1), located
on the northwestern edge of the Barents Shelf. The
north-northwest–trending Tertiary Central Basin formed
as a late Paleocene to Eocene foreland basin (Steel
et al., 1985; Helland-Hansen, 1990, 1992) because of
uplift and loading from the western Spitsbergen orogenic belt on the western flank of the basin (Harland,
1969). The proximity of the orogenic belt to the depocenter (generally < 50 km [ < 31 mi]) provided a readily
available source of coarse clastic sediment for the basin fill.
Shallow-marine sandstones of the Tertiary basin
fill make up the Battfjellet Formation, essentially the
shallow-water topsets of a clinoformed, major regressive sequence (Helland-Hansen, 1990, 1992). Individual clinoforms can only be distinguished when sand
prone on the slope. Therefore, clinoforms that are sand
prone on the slope are bundled with underlying and
overlying shales into marine shale-bounded genetic
sequences (sensu Galloway, 1989; see Steel and Olsen,
2002). Resulting sequences are referred to as ‘‘clinoforms’’ despite containing multiple clinoform strata.
The bounding shale layers are inferred to be fourthorder maximum flooding surfaces based on dinocyst
age dates that bound the entire package of clinoforms.
The lowermost clinoform is underlain by a marine shale
dated to be at or near the Paleocene– Eocene boundary
(55 Ma), whereas a marine shale near the top of the
formation was placed within the lower Eocene, giving
a maximum duration for the 25 genetic sequences of
about 6 m.y. Based on this reasoning, the maximum
average span of each genetic sequence is approximately
250 k.y. A single clinoform sequence (clinoform 14 of
Steel and Olsen, 2002) with slope disruption, shelfedge deltas, sand-filled channels on the slope, and an
obvious attached basin-floor fan was chosen for testing
the notion of time-space order on the slope. The outcrop is oriented very slightly oblique to the inferred
depositional-dip profile of the basin.
Clinoform 14 has a threefold stratigraphic architecture, designated (from the base) as 14A, 14B, and
14C (Figure 2). Clinoform 14A is fairly coarse grained,
sand-rich, and ranges from 15 to 25 m (49 to 82 ft)
thick on the shelf and upper slope. The upper slope is
dominated by coeval slope channels and lobes, the
morphologies of which are influenced by a strong fluviodeltaic system at the shelf break. Clinoforms 14B
and 14C are both finer grained, heterolithic, and thinbedded units. Clinoform 14B ranges from 15 to 30 m
(49 to 98 ft) thick, and 14C ranges from 12 to 25 m (39
to 82 ft). These individual sandy components of clinoform 14 are separated by distinct landward-wedging
shale intervals.
METHODS OF DATA COLLECTION
AND CORRELATION
Thirteen stratigraphic sections were measured through
clinoform 14 on the mountain of Storvola, spaced between 75 and 500 m (246 and 1640 ft) apart. Facies
associations were chosen based on the character of the
deposits, location of the sections along the clinoform,
and interpreted depositional environments. Individual
units were distinguished by major shifts in facies associations, generally landward shifts in facies caused by
marine flooding. Physical correlation between some
of the sections was achieved by walking out the beds.
Other sections were correlated using photomosaics
shot from a helicopter or by stacking pattern analysis of
units in the sections. The datum was chosen at the top
of clinoform 14C. This surface is a shale-on-sand contact and is inferred to represent a major marine flooding surface separating genetically distinct clinoforms.
FACIES ASSOCIATIONS
Shelf and Shelf-Edge Deposits
Deposits located on the shelf and shelf-edge segments
of the clinoform are predominantly deltaic in origin
and typically form coarsening-upward, sand-rich units
(summarized in Table 1). The three partitions of clinoform 14 (A, B, and C) distinguish individual deltaic
complexes. The deltas are argued to have been fluvial
dominated because of the sparsity of wave structures
and evidence for rapid deposition, such as climbing
ripples, tractive structures, and slump deposits, as well
as low abundance and diversity of trace fossils. Thus,
individual facies associations are interpreted in terms
of the different elements of a fluvial-dominated delta
complex (Figure 3A). The delta front incorporates the
proximal to distal mouth bars, and the coastal plain
consists of flood-plain and interdistributary-bay deposits and encased fluvial channels. The delta-front association can be seen to change in a basinward direction
from tidal-influenced deposits to deposits lacking tidal
reworking. Evidence for tidal influence is found in
stacked sets of landward-directed sigmoidal strata. This
transition occurred as the deltas evolved from a shelf
position to a shelf-edge position on the preexisting clinoform surface. The shallow water in front of shelf
deltas would have amplified tidal energy, whereas
this was less so for shelf-edge deltas facing deeper water. Consistent with this, tidal influence in delta-front
Petter and Steel
1455
1456
Geologic Note
Figure 2. (A) Photomosaic of Storvola and Hyrnestabben showing large-scale stacking patterns of gross depositional environments in clinoform 14. Note the attached, sand-rich
basin-floor fans (20 – 50 m [66 – 164 ft] thick) on Hyrnestabben. Discontinuous sand bodies on the slope segment of the clinoform are channels that provided conduits, at several
time intervals, for sand bypass to the basin floor. (B) Cartoon illustrating the threefold stratigraphic architecture of clinoform 14, designated 14A, 14B, and 14C, separated by
landward-wedging marine shale layers. Clinoform 14A is the lowermost division and is sand rich and thick bedded. Clinoforms 14B and 14C overlie 14A and are more heterolithic
and thin-bedded units.
deposits is predominant in the western (landward) extent of the clinoform topset, but absent in the eastern
(basinward) extent. In addition to the spatial reorganization of the deltas from shelf to shelf edge, there
were also temporal changes to the deltas that caused
the shelf-edge deltas of clinoforms 14B and 14C to be
significantly more heterolithic and finer grained than
the clinoform 14A shelf-edge deltas (Figure 3A).
Slope Deposits
Slope deposits dip basinward with a gradient of approximately 2–5j and can be seen to pass landward into
more flat-lying shelf deposits. Two facies associations
are recognized on the slope: slope-channel deposits and
slope-lobe deposits (see Table 1; Figure 3B). These facies are downdip equivalents to the distributary-channel
facies at the shelf edge. Slope-channel fills are correlative to at least part of the attached basin-floor fan
(Crabaugh and Steel, 2004; Clark and Steel, 2006).
They are limited to the upper slope on clinoform 14,
but most likely continue downdip out of section. Slope
lobes, although correlative to both basin-floor and slopechannel deposits, are limited to the upper slope. Thick
channel turbidites with upper-stage plane-parallel lamination and climbing ripples indicate sustained traction and aggradation of sand within flows. Examples of
hyperpycnites from clinoform 14 are shown in Figure 4.
Sustained currents are also suggested by lateral accretion of deposits in slope channels. Abundant and wellpreserved terrestrial organic material is present within
slope deposits. Deposits of the attached basin-floor fan
are similar (Crabaugh and Steel, 2004; Clark and Steel,
2006). Evidence for bioturbation on the slope is limited, consisting of occurrences of Arenituba in slope
channels and Arenituba and Phicosiphon in slope lobes
(A. Uchmann, 2004, personal communication).
In combination with the connection of slope channels with shelf-edge distributaries, these observations
suggest a fluvial origin (i.e., hyperpycnal flow) for the
slope deposits (Plink-Björklund et al., 2001; PlinkBjörklund and Steel, 2004). Deposition on the slope
was dominated by intervals of vigorous sand delivery
from shelf-edge deltas during river-flood events. The
disparate styles of deposition between slope channels
and slope lobes were brought about by the variability in
hyperpycnal flows caused mainly by differences in initial density contrast between fluvial discharge and ambient basin water that caused some flows to channelize,
whereas others traveled in an unconfined state.
SHELF-MARGIN AND SLOPE ARCHITECTURE
OF CLINOFORM 14
The shelf margin of clinoform 14, dominated by shelfedge deltas, systematically accreted some 2–3 km (0.6–
1.2 mi) beyond the shelf edge of the previous clinoform
(clinoform 13), before being abandoned during shelf
transgression. Two types of shelf-edge deltas existed on
clinoform 14 at different times during its history. The
early shelf-edge deltas of clinoform 14A had a flat to
falling shoreline trajectory (sensu Helland-Hansen and
Martinsen, 1996). As argued above, the important features of these deltas are distributary channels, observed
on the uppermost slope, that dumped sand directly into
upper-slope channel complexes. This provided a direct
connection between the fluvial system and the basinfloor fan, a characteristic feature of these early deltas.
The shoreline trajectories of the late shelf-edge deltas
(14B and 14C), however, were slightly rising (Figure 5).
Consequently, these deltas were more aggradational,
and bypass was largely replaced by slope accretion and
coastal-plain deposition.
Growth faulting is observed on the uppermost
slope, but appears to be of minor importance. The fault
provides meter-scale expansion of the stratigraphy,
but was probably not a major depocenter. The noted
growth faulting does provide evidence for instability
on the upper slope that is supported by the presence
of a small (2-m [6.5-ft]-high) slump block located
downslope. Therefore, the potential for slope canyons
exists on clinoform 14, but these features are difficult
to infer beneath the presence of scree. Plink-Björklund
and Steel (2004) suggest that slope canyons in the Central Basin could have been initiated through meterscale scouring by hyperpycnal flows.
The upper-slope architecture of clinoform 14A
was highly influenced by vigorous sediment delivery
from shelf-edge deltas. In particular, hyperpycnal discharge from distributary channels directly onto the
upper slope would have provided a large sediment
volume that could have either accreted on the slope
or bypassed to the basin floor. The resulting scenarios of accretion or bypass, as well as the resultant depositional elements, were partially controlled by the
magnitude of the hyperpycnal events and, thus, the
magnitude of the responsible flooding event. Additionally, as floods waxed and waned, the magnitude
of hyperpycnal flow could therefore have varied within individual events, causing an evolution from one
element to another in a single depositional episode
(Figure 6).
Petter and Steel
1457
1458
Geologic Note
Table 1. Compilation of Facies Descriptions (A) and Interpretations (B) from Clinoform 14
A
Facies Association
Grain Size
Shelf and Shelf Edge
Fluvial-dominated delta front
Tidally influenced
88 – 250 mm,
silty toesets
Tidally lacking
62 – 250 mm
Distributary channel
177 – 350 mm
Coastal plain
Prodelta and marine shale
Slope
Slope channel
Slope lobe
< 62 –125 mm, shaly
< 62 mm, shaly
Typical Set and Bed Thickness
Sedimentary Structures
20 – 70-cm (8 –27-in.) sets,
Sigmoidal cross-strata,
75 – 150-cm (29 – 59-in.) beds landward directed
to bidirectional
10 – 40-cm (4 –16-in.) sets,
Flat to low-angle cross-strata;
40 – 75-cm (16 – 29-in.) beds
plane-parallel lamination;
current ripples; slumps
20 – 40-cm (8 –16-in.) sets,
High-angle planar and
100 – 150-cm (39 – 59-in.) beds trough cross-strata;
current ripple caps,
lateral-accretion sets
2 – 10-cm (0.8 – 4-in.) sets
Interbedded; wavy,
5 – 15-cm (2 – 6-in.) beds
indistinct laminations
?
Interbedded to laminated
177 – 350 mm
15 – 75-cm (6 –29-in.) sets,
Upper-stage plane-parallel
15 – 125-cm (6 – 49-in.) beds
lamination; current ripples;
ungraded beds; scours;
lateral accretion sets
< 62 – 250 mm
2 –10-cm (0.8 – 4-in.) sets,
20 – 75-cm (8 – 29-in.) beds
Stacking Pattern
Bioturbation
Notable Features
Coarsening upward, Low
lower part of cycles Ophiomorpha
Limited to landward
segment of clinoform;
erosional surfaces
Coarsening upward, Low to absent
Found basinward
lower part of cycles
of tidal-influenced
delta front
Fining upward,
Originating from Erosional basal surfaces
above delta front
upper surface
Ophiomorpha
Caps uppermost
cycle of 14A
Base of cycles
Blocky to
fining-upward
sets
Thin bedded; current ripples; Coarsening-upward
wavy to flat lamination
units
–
?
Low Arenituba
Mostly absent from
14A (except for top)
Thickness increases
basinward of shelf edge
200–500 mm wide isolated
sand bodies; underlain
by erosional surfaces
with meter-scale scours;
coal chips and plant
fragments common
Low to moderate Sand-shale ratio increases
upward within units;
Phicosiphon
progradational
Arenituba
B
Facies Association
Interpretation
Petter and Steel
Shelf and Shelf Edge
Fluvial-dominated delta front
Tidally influenced
Deposited at the front of falling-stage shelf deltas, based on the location along the clinoform. Cross-strata were created by subaqueous dunes and bars
migrating normal to depositional strike that were generated by tidal reworking of delta-front deposits. Strength of the tidal currents was enhanced by
the relatively shallow-water depths across the shelf, landward of the shelf edge. Erosive surfaces are caused by tidal ravinement or alternatively may
underlie tidally reworked distributary-channel deposits.
Tidally lacking
Deposited at the front of falling-stage shelf-margin deltas. Fluvial domination is evident by the abundance of current ripples and slumping reflective of
rapid deposition. These deposits were less influenced by tidal processes because their shelf-margin position caused them to prograde into deeper
water in which tides were not amplified by the water bottom.
Distributary channel
Distributary-channel deposits of the falling-stage deltas. As the principal sediment feeders of the deltas, distributary-channel deposits would have
been slightly coarser than delta-front deposits. Decreasing current energy caused by gradual channel abandonment led to fining-upward units.
Postabandonment colonization by marine organisms is recorded by Ophiomorpha traces beginning at the upper surface and decreasing in density
downward.
Coastal plain
Terrestrial sediments deposited on the flood plain only during periods of coastal-plain aggradation. The conspicuous absence of coastal-plain deposits
in most of clinoform 14A is suggestive of forced regression.
Prodelta and marine shale Deposited as prodelta silts and shales basinward of the delta front or as shale under open-marine conditions during marine flooding events associated
with transgression or delta-lobe abandonment.
Slope
Slope channel
Turbidites deposited under steady-flow conditions in turbidity current scoured channels. Steady-flow conditions are suggested by lateral accretion (Abreu
et al., 2003), blocky grain-size profiles (Mulder et al., 2003) and thick sets of plane-parallel laminations and ungraded beds. Connection to shelf-margin
distributary channels and abundant material indicates that turbidity currents were generated by hyperpycnal flow during times of river flooding. Individual
units reflect the flood hydrograph (Figure 6; see also Mulder et al., 2003), with the blocky, plane-parallel laminated to ungraded portion deposited
during peak flood and fining-upward ripple caps deposited during waning flow. Waxing flow, when channel-erosion potential was highest, was likely
bypassed.
Slope lobe
Deposited by nonignitive turbidity currents as a point-sourced lobe on the upper slope. The most likely source of the turbidity currents was hyperpycnal
flow based on the presence of coeval distributary channels at the shelf edge. Successive hyperpycnal flows would have avoided the topography created
by prior slope-lobe deposits, resulting in progradation of the lobes.
1459
Figure 3. Comparison of clinoform
14A shelfal (a) and shelf-edge delta facies
(b) with clinoforms 14B and 14C shelfedge delta facies (c). DF = delta front
without tidal influence; DC = distributary
channel; T = tidally influenced delta
front. Shelf deltas exhibit greater tidal
influence than shelf-edge deltas in clinoform 14. Deltas of clinoform 14A are considerably more sand rich and coarser than
the clinoforms 14B and 14C deltas. Cross
section shows the location of measured
sections from this figure.
Indications of Bypass
Attachment of the basin-floor fan to the clinoform is
proof-positive of sand bypass on the slope. In the absence of this piece of information, however, evidence
for bypass can be found within the slope deposits.
1460
Geologic Note
Kneller and Branney (1995) argued that thick, ungraded
sands lacking tractive structures can be deposited slowly
by sustained high-density underflows instead of collapse or freezing of turbidity currents. Deposition
occurs within a highly concentrated basal zone of flow
where transport of sediment is hindered by grain-to-grain
Figure 3. Continued. Comparison of slope-channel (d) and
slope-lobe (e) facies. Sections of high-density hyperpycnites are
indicated on the slope-channel complex section and are punctuated by periods of relative inactivity when low-density hyperpycnites and/or surge-type turbidites caused by shelf-margin
collapse are deposited. Note the coarsening-upward signature
of the slope-lobe deposits caused by progradation of the lobe
over multiple hyperpycnal events.
interactions. Turbidity would be dampened by the high
particle concentration as described by Lowe (1982),
and the basal zone would behave somewhat like a debris flow instead of turbidity current (although the behavioral analogy with debris flow is not exact). The sediment concentration increases toward of the base of the
flow, until grain interactions reach a critical threshold
at which the yield strength of the flow exceeds the applied shear stress from the overlying flow, and transport
of sediment ceases.
The basal hyperconcentrated zone is continually
resupplied with sediment by a downward flux of momentum from the overlying, turbid flow. This continuous downward flux of sediment forces the hyperconcentration and deposition of sediment (Lowe, 1982).
The most interesting and perhaps most important part
of this process, however, is that the continuous overlying turbidity current will bypass sediment downslope
simultaneous with the local deposition on the slope.
This raises important issues regarding the timing of
basin-floor deposition, as well as the connectivity of
slope-channel and basin-floor sands. These questions
should be confronted when continuous-flow turbidity
currents are suspected as a mechanism for transport
into a deep-water setting.
Ungraded sand beds, some on the order of 1 m
(3.3 ft) thick, are found in the inferred axis of a slopechannel complex on clinoform 14 (see description of
slope-channel facies association, Table 1). These beds
are associated with shale- and silt-filled scours originating from bed boundaries that cut down into the
ungraded sands. The shale and silt beds in the scours
are deformed in a manner indicating that they had been
draped over the scour and subsequently slumped down
the scour surface. The fine-grained nature of these deposits suggests that they are not related to the scourcreating current. Instead, the sediment transported by
the scouring current was bypassed down the slope
and potentially to the basin-floor fan. Note that the
hyperconcentrated flow discussed above is not erosional
because the basal zone of flow is laminar. Therefore,
the scouring current must have been a turbidity current
Petter and Steel
1461
without a well-developed basal zone of hyperconcentration. Because the hyperconcentrated basal zone is
nonerosional and therefore cannot entrain sediment,
continuous downward flux of grains from an overlying,
quasisteady flow is required to sustain the current for
long runout.
Slope Channels
The average width of slope-channel complexes on
clinoform 14 is estimated to be approximately 100 –
200 m (328– 656 ft). Slope-channel complexes range
in thickness from 10 to 15 m (33 to 49 ft), but individual channel cuts are approximately 1 –5 m (3– 16 ft)
deep (Figure 7). Internal scours in the channels are less
than 50 cm (20 in.) deep (Figure 4A). Stacked channel
complexes appear to be shingled basinward, although
this may possibly be an artifact of the two-dimensional
aspect of the exposure. Slope-channel complexes are
generally capped by 1 – 2 m (3 – 6 ft) of slope-lobe deposits reflective of diminishing discharge and shelfedge delta abandonment.
Slope channels on clinoform 14 are believed to
have been generated by point-sourced, relatively confined jet flow coming from shelf-edge distributary channels. These flows had erosive potential that allowed
them to erode the underlying substrate and cut channels into the upper slope; i.e., high-density hyperpycnal
flows. In addition to the initial accelerative feedback
from entrainment of eroded material, confinement
in the self-generated channels allowed the high-density
hyperpycnal flow to continue inertially for a longer distance than an unconfined jet flow, which would tend
to be dispersive (Bates, 1953). Bypass to the basin floor
likely occurred as high-density turbidity currents flowed
through the slope channels.
Slope Lobes
Thin-bedded, heterolithic, and nonchannelized deposits on the slope are interpreted as slope lobes. Slope
lobes on clinoform 14 are also point sourced from
Figure 4. (A) Plane-parallel lamination filling decimeter-scale
scour. Note the hammer for scale (circled). (B) Stacked channels
in a slope-channel complex. (C) Alternation of plane-parallel
stratification and ripples in hyperpycnite bed. Note the pen for
scale.
1462
Geologic Note
Petter and Steel
Figure 5. Correlation of measured sections along clinoform 14. The panel is oriented slightly oblique to depositional dip with the basin to the right (see Figure 1 for section map).
The attached basin-floor fan lies approximately 4 km (2.5 mi) to the southeast of the most basinward measured section.
1463
Figure 6. (A) Flood hydrograph and associated sequence
for hyperpycnal flow that remains fully turbulent. (B) Flood
hydrograph and associated sequence for hyperpycnal flow in
which turbulence is suppressed
at peak discharge. ‘‘A’’ divisions
are deposited from low-density
hyperpycnal flow; the ‘‘B’’ divisions are deposited from fully
turbulent high-density hyperpycnal flow, whereas the ‘‘C’’
divisions are deposited from
high-density hyperpycnal flow
with suppressed turbulence
caused by grain interactions.
The subscripts D and E refer to
depositional and erosional flows,
respectively. See also Mulder
et al. (2001a).
shelf-edge distributary channels during times of hyperpycnal flow. However, slope lobes are the result of
low-density hyperpycnal flows instead of the highdensity hyperpycnal flows that formed the slope channels. Low-density hyperpycnal flows, as described above,
are created by the collapse of a sediment-laden water
column into a bottom-riding current. The collapse occurs at a plunge front some distance from the mouth of
the distributary channels (Wright et al., 1986). Prior to
plunging, therefore, the discharge behaves as an un1464
Geologic Note
confined jet flow (Bates, 1953) and spreads out downstream. Upon plunging, the resulting hyperpycnal flow
runs both within and outside any channels that may
exist. The effects of dispersion and initial low sediment
concentration cause the current to die and the flow to
be deposited as a lobe-shaped body on the slope. Systematic shingling of individual low-density hyperpycnites causes the formation and progradation of the slope
lobe element. The slope lobes generally have a steeper
gradient (nondecompacted average of approximately
Figure 7. Interpreted photograph of an early lowstand
slope-channel complex. Bracketed bars indicate periods of
strong hyperpycnal activity and
channel cutting in the complex.
The slope-channel complex is
overlain by a second cycle of
slope-lobe deposits and minor
channels that occur lateral to
a separate slope-channel complex.
4–5j) than the slope-channel surfaces (nondecompacted average of approximately 2–3j) that reflects the
progradational nature of the depositional element.
Bornhold (1989) are very similar, suggesting that unconfined low-density turbidity currents were the chief
transport process to the slope lobes.
Relationship between Slope Channels and Lobes
SEQUENCE-STRATIGRAPHIC FRAMEWORK
Note that slope channels and slope lobes were coeval
elements on clinoform 14. Alternation between slopechannel and slope-lobe deposition was caused by variability in the magnitude of hyperpycnal activity and
could have switched within single flood events as discharge waxed and waned. Prior and Bornhold (1989)
described a Holocene fan delta in British Columbia in
which hyperpycnal flow was attributed to be a major
factor in the development of the fan delta. They noted
sand and gravel transported a significant distance from
the river mouth, as well as chutes, indicative of highly
erosive currents. Gravel and coarse sand deposits were
generated by inertial bed-load flow during vigorous
flood-stage discharge at the river mouth. Accompanying
suspended load formed both high-density turbidity currents capable of scouring the chutes as well as unconfined low-density turbidity currents that deposited
thinly interbedded sands, silts, and clays in the interchute setting. The chutes and associated coarse deposits of Prior and Bornhold (1989) are comparable
to the slope channels of clinoform 14; therefore, it
is likely that the slope channels were also formed by
high-density flows. Likewise, the slope-lobe facies of
clinoform 14 and the interchute deposits of Prior and
Clinoform 14 consists of a sand-prone sediment wedge
encased within marine shales and represents a period
of significant shoreline progradation across the shelf
followed by marine flooding. Regression across the
entire shelf was most likely forced by base-level fall;
therefore, the clinoform can be considered as resulting from a single base-level cycle, probably of fourth
order (a few hundreds of thousands of years; Steel and
Olsen, 2002). Superimposed upon the fourth-order cyclicity are smaller scale, shallowing-upward cycles, interpreted as fifth-order cycles. The character of fifthorder deltaic cycles differs between clinoform 14A and
clinoforms 14B and 14C (Figure 3A). These differences are attributed to the transition from falling-stage
and early lowstand systems tracts (clinoform 14A) to
late lowstand systems tract (clinoforms 14B and 14C).
Fifth-order deltaic cycles in the falling stage and early
lowstand are sand rich and medium to thick bedded
and range in thickness from 1 to 5 m (3 to 16 ft). Fifthorder deltaic cycles in the late lowstand are heterolithic and thin bedded and are thicker than the fallingstage and early lowstand cycles, ranging from 5 to 15 m
(16 to 49 ft) thick. In addition to a greater proportion of shale and silt than the falling-stage and early
Petter and Steel
1465
Figure 8. Sequence evolution of clinoform 14:
(A) falling-stage systems
tract; (B) early lowstand
systems tract; (C) late lowstand systems tract; (D) transgressive systems tract.
lowstand cycles, the sand fraction of the late lowstand
cycles is also much finer (62 to 177 mm as opposed to
125 to 250 mm).
Fourth-Order Sequence
Falling-Stage Systems Tract: Delta Progradation and
Shelf-Margin Accretion
Sediments deposited during forced regression of the
fluvial-dominated delta across the shelf to the shelf
edge form the fourth-order falling-stage systems tract
1466
Geologic Note
(Figure 8A). These deposits make up most of clinoform 14A on the shelf. Individual cycles in the fallingstage systems tract have a flat to downward shoreline
trajectory (sensu Helland-Hansen and Martinsen, 1996),
foreshortened stratigraphy (Posamentier and Morris,
2000) reflective of regressive erosion, and a distinct
lack of coastal-plain deposits (Hunt and Tucker, 1992),
all of which are indications of forced regression. However, stacking of fifth-order cycles in the falling-stage
systems tract caused a component of aggradation in
addition to progradation, which has not previously
Figure 8. Continued.
been documented for this systems tract. The observed
aggradation was a consequence of the basal transgressive parts of the cycles, which caused minor phases of
retreat of the falling-stage deltas. The resulting architectural framework of the fourth-order falling-stage
systems tract consists of stacked delta cycles.
Forced regression brought deltas to the shelf break
repeatedly during the falling-stage systems tract and
likely dumped hyperpycnal flows onto the upper slope.
Bypass of sand to the basin floor might have occurred
at these times but was probably limited. Instead, sedi-
ment accreted on the slope, leading to shelf-margin
accretion of at least 1.5 km (0.9 mi). The volume of
sediment stored by shelf-margin accretion makes bypass during this time unlikely. Clinoform progradation
gave way abruptly to slope bypass toward the end of
the falling-stage systems tract. Several reasons for the
transition from progradation to bypass are possible.
Gradual steepening of the shelf edge caused by rapid
progradation could have increased the gradient at the
shelf edge to a critical condition for turbidity current
ignition. Alternatively, shelf-edge steepening could have
Petter and Steel
1467
led to slope failure and the creation of a slope canyon
that would have served as a conduit for bypass to the
basin floor. Little evidence for a large canyon exists on
clinoform 14; however, the basal surface of the canyon
would most likely lie within fine-grained sediments that
are generally covered by scree in Spitsbergen.
A third cause for the initiation of bypass could
have been high-order climatic change impacting precipitation in the hinterland and transformation of hyperpycnal potential (Mutti et al., 1994). The early Eocene
is generally regarded as the warmest epoch of the Cenozoic (Miller et al., 1987; Wing et al., 1991; Zachos et al.,
2001) and was most likely ice free. Several rapid global
warming events have been shown to have occurred in
the latest Paleocene and early Eocene (Kennett and Stott,
1991; Thomas and Zachos, 2000; Wing, 2000; Zachos
et al., 2001; Hollis et al., 2005), possibly as a result of
immense gas hydrate dissociation from near-sea-floor
sediments (Dickens et al., 1995). Additionally, astronomically induced climatic forcing of sediment supply
may have been more significant during greenhouse
epochs because of the reduced magnitude and rate of
eustatic sea level fluctuations ( Van der Zwan, 2002).
It might be argued that increased precipitation in
the hinterland could have created sufficient fluvial
discharge to generate high-density hyperpycnal flows,
whereas decreased precipitation in arid times would
have led to diminished potential for hyperpycnal flow.
However, increased precipitation in the Arctic is generally associated with periods of warming that led to
widespread vegetation. The increased vegetation cover
would have caused a decrease in sediment runoff in the
hinterland. Instead, high-density sediment pulses should
be associated with semiarid, sparsely vegetated periods
with infrequent flash floods or monsoonal conditions
instead of humid, well-vegetated periods. Molnar (2001)
showed that the frequency-magnitude distribution of
large-magnitude floods can be greater in semiarid instead of humid climates. Consequently, incision rates
and bed-load–transport rates would also be greater in
semiarid climates. Early Eocene aridification of northern Europe and the North Sea has been noted by Jolley
and Widdowson (2005) based on palynological assemblages in the Balder Formation. They argue for explosive
volcanism (evidenced by abundant ash layers in the same
interval) associated with rifting of the North Atlantic
as a potential cause for the rapid environmental degradation. Egger et al. (2005) argue for monsoonal conditions along the European Tethys margin based on clay
mineral and planktonic assemblages of a Paleocene–
Eocene Tethyan bathyal slope section in Austria. They
1468
Geologic Note
also discovered a marked increase in detrital quartz
and feldspar in the section, suggesting the enhancement
of continental runoff associated with the monsoonal
conditions.
More humid conditions may be associated with
high sediment supply caused by increased weathering,
but transport of this sediment will be evenly distributed temporally. Therefore, sediment concentrations
at any given time will be relatively low. Alternatively,
the sediment supply of semiarid climates will be transported at high concentrations during very short intervals. As a result, the potential for high-density hyperpycnal events (and thus, ignition of turbidity currents)
should be higher during times of aridity. This reasoning
suggests that basin-fill models incorporating sediment
supply as an input should be reexamined to reflect the
importance of sediment concentration.
Early Lowstand Systems Tract: Sediment Bypass and
Hyperpycnal Flows
The initiation of high-volume sediment bypass to the
basin floor marks the transition from the falling-stage
systems tract to the lowstand systems tract (Figure 8B).
The fourth-order lowstand systems tract can be divided
into early and late stages. The early stage of the lowstand systems tract is restricted to clinoform 14A and
represents a time of significant sediment bypass across
the shelf and to the basin-floor fan. On the shelf, the
initiation of early lowstand is marked by the base of a
prominent shelf-edge distributary channel that extends
down onto the uppermost slope. A second cycle, resulting from minor backstepping or avulsion of the shelfedge delta, overlies the basal distributary channel in the
early lowstand.
At least two slope-channel and slope-lobe complexes are observed in the early lowstand systems tract,
each connected updip to one of the fifth-order fallingstage to early lowstand cycles on the shelf. Shelf-margin
accretion would have been minor during the early lowstand because a large fraction of the sediment budget was
bypassed to the basin-floor fan, and in fact, the shelf edges
of both early lowstand cycles are nearly coincident. As
discussed previously, however, deposition and bypass in
the slope channels occurred contemporaneously. Therefore, slope-channel fill and basin-floor deposits of the
early lowstand are correlative and potentially connected.
Late Lowstand Systems Tract: Renewed Shelf-Margin Accretion
The uppermost fifth-order cycle of clinoform 14A differs from the lower cycles in 14A because of its heterolithic, thin-bedded character in the delta front and by
the coastal-plain accumulation of the shoreline deposits landward (Figure 8C). Therefore, this cycle is interpreted as a fifth-order late lowstand cycle and marks
the base of the fourth-order late lowstand systems tract.
The fourth-order late lowstand systems tract includes
clinoform 14B across the shelf and downdip to the lower
slope, as well as the lower part of clinoform 14C from
the outer shelf to the lower slope. On the lower slope,
the late lowstand strata consist of stacked backstepping levee-channel complexes. The late lowstand strata
also form a prograding wedge of thinly bedded slopelobe deposits that drape down the upper to middle
shelf. The shelf-edge delta complex of the late lowstand
accreted basinward by approximately 0.5 km (0.3 mi).
Internal bypass surfaces in the late lowstand deltaic and
lobe deposits are marked by beds of medium-grained
sand and are correlative with levee-channel complexes
on the lower slope (Figure 2). Bypass of the shelf edge
to the lower slope was potentially fed by major hyperpycnal events during the late lowstand. Short-lived
forced regressions could have also punctuated the relative sea level rise of the late lowstand, similar to the
minor transgressions of the falling stage. Most importantly, the bypass surfaces in the late lowstand deltaic
and lobe deposits illustrate that the late lowstandprograding wedge is not a younger lowstand element
than the levee-channel complex. Instead, the two elements are roughly coeval.
The late lowstand is marked by a significant coastalplain accumulation. By contrast, coastal-plain deposition was missing from the falling-stage systems tract.
A major implication of increased coastal-plain storage
would be to decrease sediment discharge at the river
mouth, particularly during flood stage. Suppression
of hyperpycnal events should result from the diminished sediment concentration. Most of the late lowstand slope deposits are classified as low-density hyperpycnites, suggesting that storage of sediment on
the coastal plain could have decreased the frequency
of high-density hyperpycnal flows with the potential
for ignition. The sediment budget for the late lowstand, most of which was stored on the shelf and upper slope as opposed to basin-floor storage during the
early lowstand, reflects this speculation.
Saller et al. (2004) described a clinoform set from
Indonesia in which early lowstand deltas were linked
to sand-rich slope and basin-floor deposits, whereas
late lowstand deltas were associated downdip with
mud-rich accumulations. The differences between early
and late lowstand deep-water systems were attributed
to eustatic-controlled shelf-sediment budgets. The am-
plitude and rate of relative sea level change in the Indonesian setting (Pleistocene icehouse system; Fischer,
1981) presented by Saller et al. (2004) was most likely
significantly greater than in Spitsbergen during the Eocene (greenhouse system; Fischer, 1981). This leads
to noteworthy differences in clinoform-stacking patterns, particularly during the late lowstand, when rapid
icehouse flooding allowed little progradation by late
lowstand shelf-margin deltas. Relative sea level rise
during the early Eocene proceeded slowly enough to
allow significant shelf-margin progradation during the
late lowstand.
Transgressive Systems Tract: Retreat across Shelf Platform
The upper part of clinoform 14C is cut by a prominent and widespread erosional surface on the outer
shelf that is interpreted as a transgressive ravinement
surface and marks the base of the fourth-order transgressive systems tract (Figure 8D). The transgressive ravinement surface truncates the underlying late lowstandprograding wedge. The basinward termination of the
transgressive ravinement surface constrained the position of the latest lowstand shelf break, thereby determining the amount of shelf-margin progradation
(0.5 km; 0.3 mi) during the fourth-order base-level
cycle. The transgressive systems tract is composed of
sand-rich estuarine and barrier bar deposits marked by
landward-prograding sets.
IMPORTANCE OF HYPERPYCNAL FLOW FOR
DEEP-WATER RESERVOIR CHARACTERIZATION
AND EXPLORATION
Hyperpycnal flow should be considered as a major alternative scenario to sediment and slope failure and
storm processes for the generation of turbidity currents.
The process of turbidity current initiation significantly
impacts reservoir distribution, quality, and connectivity. Net-to-gross sand will be greater for high-density
than low-density hyperpycnites because of the higher
energies associated with the denser flows. As discussed
above, high-density hyperpycnites are more prevalent in the early lowstand systems tract, whereas lowdensity hyperpycnites dominate the late lowstand of
clinoform 14. Separation of buoyant hypopycnal plumes
from coeval hyperpycnal flows has been hypothesized for submarine glacial-melt outflow (Lønne, 1997)
and presumably should also occur for hyperpycnal
Petter and Steel
1469
flow generated from fluvial systems. The implication of this would be to separate much of the finer
fraction into the hypopycnal plume, leaving the hyperpycnal flow cleaner than classical surge-type turbidity
currents formed from delta-front instability. The cleanness of hyperpycnites versus surge-type turbidites,
however, remains to be tested by comparison of recent
deposits.
As mentioned previously, long-lived, continuousflow turbidity currents, such as hyperpycnal flows, have
the potential to deposit sediment in slope channels
while simultaneously bypassing sand to the basin floor.
Therefore, simple stratigraphic trapping of early lowstand basin-floor sands may not be a reasonable evaluation when slope channels are also filled during the
early lowstand. This scenario of early lowstand channel fill differs from conventional sequence-stratigraphic
models in which slope channels are filled during the
late lowstand.
Conventional sequence-stratigraphic models have
also placed the timing of the late lowstand-prograding
wedge after the deposition of a late lowstand toe-ofslope channel-levee complex. The channel-levee complex of clinoform 14, however, is time-equivalent to
the late lowstand wedge (clinoform 14B) and was therefore likely fed by a pulse of hyperpycnal activity that
generated turbidity currents strong enough to bypass
the delta front and slope lobes that served as the depocenter for most of the late lowstand sediments.
Clear bypass surfaces in the late lowstand shelf-margin
deposits attest to the contemporaneous timing of the
channel-levee complex in clinoform 14.
Obviously, basin-floor deposits fed by hyperpycnal flow will be point-sourced submarine fans (sensu
Reading and Richards, 1994) ranging from mud rich to
gravel rich, depending on the sediment caliber available at the river mouth. Slope hyperpycnites will be
clustered in and around slope channels. The key to exploring for hyperpycnite reservoirs, therefore, would
seem to be to locate the point source for hyperpycnal
flow, i.e., shelf-margin distributary channels. The payoff for locating these deposits should, in theory, be
thicker bedded deep-water reservoirs than those formed
by classical surge-type turbidites because of the longer
duration of individual hyperpycnal flows. However, it
is unclear from this study whether the small rivers that
are prone to generating hyperpycnal flow can deliver
sufficient volumes of sand to create economically viable
reservoirs. For an in-depth discussion of clinoform 14
basin-floor fans, see Crabaugh and Steel (2004) and
Clark and Steel (2006).
1470
Geologic Note
CONCLUSION: INSIGHTS TO SHELF-MARGIN
AND SLOPE ARCHITECTURE AND
CONTRIBUTION OF HYPERPYCNAL
FLOW TO BASIN DEPOSITS
Documentation of the architecture of a single clinoform in the Tertiary Central Basin of Spitsbergen has
provided insights on the relationship between the shelf
edge, slope, and basin floor.
1. Understanding of the shelf edge is critical for deciphering order in slope deposits. The stacking pattern of shelf-edge delta cycles is reflected in the
architecture of the coeval slope and should be recognizable on the basin floor.
2. Hyperpycnal flow can dominate transport of
reservoir-quality sediment to deep water in certain
settings. In addition, the variability of hyperpycnal
flow will create variability in deep-water depositional elements. Slope channels and slope lobes, for
instance, resulted from high- and low-density hyperpycnal flow, respectively. Thus, the initial density
contrast between the delta inflow and the ambient
basin water will control the depositional character
on the slope in deep-water systems dominated by
hyperpycnal flow.
3. The connection of distributary channels at the shelf
edge to the basin-floor fan, via slope channels, is
critical for bringing hyperpycnal flow to the deep
basin. Flow emerging from shelf-edge distributary
channels follows a steeper gradient than flow down
shelf delta fronts because of the coincidence of the
shelf edge and coastal prism. The basinward gradient increase in a shelf-margin setting makes ignition
of long-lived turbidity currents from fluvial discharge more likely.
4. Forced regression enhances the odds of generating
hyperpycnal flow, not only by bringing distributary
channels to the shelf edge, but also by limiting floodplain sediment storage. The sediment concentration
at the river mouth during floods is thus increased.
As sea level rises, however, aggradation will occur
behind the delta, reducing the amount of sediment
available at the river mouth to produce hyperpycnal
flow.
5. Last, the climatic setting of a river’s drainage basin may influence the chances of hyperpycnal flow.
Semiarid basins with periodic, large-magnitude flash
floods may provide the ideal scenario for hyperpycnal flow to bypass to the basin floor, particularly
during forced regression.
REFERENCES CITED
Abreu, V., M. Sullivan, C. Pirmez, and D. Mohrig, 2003, Lateral
accretion packages (LAPs): An important reservoir element in
deepwater sinuous channels: Marine and Petroleum Geology,
v. 20, p. 631 – 648.
Bates, C. C., 1953, Rational theory of delta formation: AAPG
Bulletin, v. 37, p. 2119 – 2162.
Best, J. L., R. A. Kostaschuk, J. Peakall, P. V. Vallard, and M. Franklin,
2005, Whole flow field dynamics and velocity pulsing within
natural sediment-laden underflows: Geology, v. 33, p. 765 – 768.
Bouma, A. H., 2004, Key controls on the characteristics of turbidite
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