Sedimentology (2005) 52, 391–428
doi: 10.1111/j.1365-3091.2005.00703.x
Stacked fluvial and tide-dominated estuarine deposits in highfrequency (fourth-order) sequences of the Eocene Central Basin,
Spitsbergen
PIRET PLINK-BJÖRKLUND
Göteborg Dynamic Stratigraphy Group, Department of Earth Sciences: Geology, Göteborg University, Box
460, SE-405 30 Gothenburg, Sweden (E-mail: piret@geo.gu.se)
ABSTRACT
Eighteen coastal-plain depositional sequences that can be correlated to
shallow- to deep-water clinoforms in the Eocene Central Basin of
Spitsbergen were studied in 1 · 15 km scale mountainside exposures. The
overall mud-prone (>300 m thick) coastal-plain succession is divided by
prominent fluvial erosion surfaces into vertically stacked depositional
sequences, 7–44 m thick. The erosion surfaces are overlain by fluvial
conglomerates and coarse-grained sandstones. The fluvial deposits show
tidal influence at their seaward ends. The fluvial deposits pass upwards into
macrotidal tide-dominated estuarine deposits, with coarse-grained riverdominated facies followed further seawards by high- and low-sinuosity tidal
channels, upper-flow-regime tidal flats, and tidal sand bar facies associations.
Laterally, marginal sandy to muddy tidal flat and marsh deposits occur. The
fluvial/estuarine sequences are interpreted as having accumulated as a series
of incised valley fills because: (i) the basal fluvial erosion surfaces, with at least
16 m of local erosional relief, are regional incisions; (ii) the basal fluvial
deposits exhibit a significant basinward facies shift; (iii) the regional erosion
surfaces can be correlated with rooted horizons in the interfluve areas; and (iv)
the estuarine deposits onlap the valley walls in a landward direction. The
coastal-plain deposits represent the topset to clinoforms that formed during
progradational infilling of the Eocene Central Basin. Despite large-scale
progradation, the sequences are volumetrically dominated by lowstand
fluvial deposits and especially by transgressive estuarine deposits. The
transgressive deposits are overlain by highstand units in only about 30% of
the sequences. The depositional system remained an estuary even during
highstand conditions, as evidenced by the continued bedload convergence in
the inner-estuarine tidal channels.
Keywords Coastal plain, estuarine, fluvial, incised valleys, sediment partitioning, tide-dominated.
INTRODUCTION
The Aspelintoppen Formation in the Eocene
Central Basin of Spitsbergen consists of alternating fluvial and estuarine deposits that can be
‘walked out’ downdip into coeval shallow- to
deep-marine sequences (Fig. 1) of the Battfjellet
Formation. The coastal-plain succession is continuously exposed in 500 · 5000 m scale mountainside exposures (Fig. 2) that are oriented
2005 International Association of Sedimentologists
approximately parallel to the depositional dip.
The exposures thus provide excellent constraints
on downdip facies transitions as well as the
vertical stacking of the fluvial/estuarine depositional sequences. There are relatively few places
where the relationships between marine and
alluvial/coastal plain depositional systems can
be examined in continuous exposures like
this (Shanley & McCabe, 1991, 1993, 1994, 1995;
Shanley et al., 1992; Hettinger et al., 1993;
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Legaretta et al., 1993; Wright & Marriott, 1993;
Gibling & Bird, 1994; Gibling & Wightman, 1994;
Aitken & Flint, 1995 and Olsen et al., 1995; Plint
et al., 2001).
Unlike most of the existing sequence stratigraphic models for coastal-plain successions (e.g.
Shanley & McCabe, 1991, 1993; Olsen et al.,
1995), the Aspelintoppen Formation of the Eocene Central Basin preserves mainly stacked
lowstand fluvial and transgressive estuarine
deposits. In contrast to current estuarine or
valley-fill models of Zaitlin et al. (1994), or most
of the coastal-plain models (Shanley & McCabe,
1993; Olsen et al., 1995), little or no sediment
accumulated on the coastal plain during highstands of sea-level (see also Plint et al., 2001)
despite the overall progradational character of the
succession. Estuaries have been generally considered to typify transgressive coastlines (e.g. Boyd
et al., 1992), and estuaries that occupy drowned
valleys are extremely common along modern
transgressive coasts (Dalrymple et al., 1992).
However, the Spitsbergen database discussed
below shows that estuaries can persist even
during the (early) highstand (see also Allen &
Posamentier, 1993).
Estuaries are defined in this paper according to
Dalrymple et al. (1992) as ‘a seaward portion of a
drowned valley system which receives sediment
from both fluvial and marine sources and which
contains facies influenced by tide, wave and
fluvial processes’. The estuaries documented in
the Eocene Central Basin are interpreted as tidedominated macrotidal estuaries. Despite the
Fig. 1. Spitsbergen archipelago is situated in Norwegian Arctic (A). Location of Eocene Central Basin of Spitsbergen,
east of West Spitsbergen Orogenic Belt (B). The basin was infilled with south-eastward-migrating clinothems (C and
D). The broken lines represent the shelf-edge position of individual clinothems and the arrows show the general
sediment supply direction and clinoform migration path (C). Coastal-plain sequences on Brogniartfjellet are coeval
with shallow- to deep-marine clinothems on Storvola, and coastal-plain sequences on Storvola with shallow- to
deep-marine clinothems on Hyrnestabben (D and E).
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Fig. 2. Seismic-scale coastal plain, shelf and slope clinoforms on (A) Storvola and (B) Brogniartfjellet. The
mountainside exposure on Storvola is ca 5 km long, and on Brogniartfjellt ca 6 km long. Both exposures are northwest–south-east oriented.
abundance of modern macrotidal estuaries and
the high preservation potential of estuarine
deposits (Demarest & Kraft, 1987; Dalrymple
et al., 1992), ancient examples are still rare (e.g.
Clifton, 1982; Galloway & Hobday, 1983). Studies
of modern macrotidal estuaries (Amos & Long,
1980; Lambiase, 1980a,b; Bartsch-Winkler &
Ovenshine, 1984; Amos & Zaitlin, 1985; Harris
& Collins, 1985; Allen & Rae, 1988; BartschWinkler, 1988; Woodroffe et al., 1989, 1993;
Dalrymple et al., 1990) show them to be highly
efficient sediment traps, with landward transport
of large volumes of sediment. The presence of net
landward movement of sediment from outside of
the estuary mouth (averaged over a period of
several years) is one of the primary features that
distinguish estuaries from delta distributaries
where the net transport is seaward.
The main aims of this paper are: (i) to describe
the facies and stratal architecture of the coastal-
plain fluvial and tidal estuarine deposits in the
Eocene Central Basin of Spitsbergen; (ii) to demonstrate that the fluvial/estuarine sequences were
deposited in incised valleys, and the coastalplain succession was formed by a series of
vertically stacked incised valley fills; and (iii) to
show that lowstand, and especially transgressive
deposits volumetrically dominate the fourth
order (ca 100 000–300 000 years) coastal-plain
sequences in this basin.
GEOLOGICAL SETTING
During the Tertiary, Spitsbergen was situated
along a transform fault margin (the Hornsund
Fault Zone) separating the North American and
Eurasian plates as the Lomonosov Ridge and the
Norwegian-Greenland Sea rift began to open
(Steel et al., 1981; Steel & Worsley, 1984; Steel
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et al., 1985; Teyssier et al., 1995; Braathen et al.,
1999). Transpression in the Palaeocene to Eocene
formed the West Spitsbergen Orogenic Belt with
areas of basement uplift, and folding and thrusting along a NNW–SSE trending zone (Fig. 1).
Regional flexural subsidence produced by loading
from the orogenic belt formed the Central Tertiary
Basin (Braathen et al., 1999). This basin was a
foreland basin during the initial period of active
thrusting, but became a piggy-back basin (Blythe
& Kleinspehn, 1998) on account of foreland
propagation of thrusting. The Lomfjorden and
Billefjorden fault zones east of the Central Basin
probably represent late stage reactivation of deepseated reverse faults (Braathen et al., 1999).
The Eocene Central Basin was asymmetrically
infilled by rivers draining a rising and eastwardmigrating Western Spitsbergen fold-and-thrust
belt (Harland, 1969; Eldholm et al., 1984). The
eastward and south-eastward migration of the
basin depocentre, driven by tectonic loading,
created an asymmetric sedimentary succession
(Helland-Hansen, 1990) that is more than 1Æ5 km
thick in the west, thinning to <600 m in the east.
The basin is one of the few basins in the world
that preserves shallow to deepwater clinoforms at
a seismic scale in large mountainside exposures
(Kellogg, 1975; Helland-Hansen, 1992). The clinoforms, and their equivalent ‘rock units’ referred to
as clinothems (Rich, 1951) reflect the outgrowth
of the basin margin driven by sediment supply
from the fold-and-thrust belt to the west (Fig. 1).
Each clinoform surface is a time line and represents the morphologic profile extending from the
coastal plain to the marine shelf and down into
the deeper water slope and basin-floor environments. The term ‘clinoform’ is used here in a
broader sense than first defined by Rich (1951),
i.e. for the entire length of the time line. The
topset of the clinoforms (coastal-plain facies belt)
belongs to the Aspelintoppen Formation, and the
foreset and bottomset (shelf, slope and basin-floor
facies belts) to the Battfjellet Formation (Fig. 3).
The Battfjellet clinoforms sharply overlie the
Gilsonryggen Member shales (Fig. 3). The characteristics of some of these clinoforms have been
discussed by Steel et al. (2000), a classification of
clinoform types was outlined by Mellere et al.
(2002), and characteristics of some shelf-margin
delta types documented by Plink-Björklund &
Steel (2005). The slope reaches of the clinoforms
have been discussed in Plink-Björklund et al.
(2001) and Plink-Björklund & Steel (2002), and
initiation of turbidites by river effluent in PlinkBjörklund & Steel (2004).
Fig. 3. Palaeocene and Eocene stratigraphy in Tertiary
Central Basin of Spitsbergen (Steel et al., 1985).
The Aspelintoppen Formation thus constitutes
the continental counterpart of a thick overall
regressive shelf, slope and basin-floor succession
(Kellogg, 1975; Steel et al., 1981; Steel & Worsley,
1984; Steel et al., 1985). It is >500 m thick in
places, has an abrupt and/or interfingering relationship with the shelf, slope and basin-floor
parts of the clinoforms belonging to the Battfjellet
Formation. The Aspelintoppen Formation is generally aggradational and mud-prone, with a sand/
mud ratio of ca 0Æ25. It comprises fresh/brackishwater shales, alternating with non-marine and
marine brackish-water sandstones, coals and siltstones (Steel et al., 1981).
FACIES ASSOCIATIONS
The sedimentology of the Aspelintoppen Formation coastal-plain deposits was documented
on two mountainsides, Brogniartfjellet and
Storvola, along northern shores of Van Keulenfjorden (Fig. 1). The mountainsides have a
northwest to southeast orientation, and provide
a roughly dip-orientated section. Each of the
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Fig. 4. Thin sand-prone erosionally based levels that can be followed across the coastal-plain facies belt, divide the coastal-plain succession into depositional
sequences.
Stacked fluvial and tide-dominated estuarine deposits in high-frequency sequences
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Fig. 5. A representative measured section through fluvial deposits of FA 1, and palaeocurrent measurements derived
from cross-strata. Numbers by the measured section refer to facies (see Table 1). (A) Channel fills commonly have
multiple erosion surfaces (dashed lines). (B and E) Coarse-grained, cross-stratified sandstones erosionally overlie
very-fine-grained muddy tidal sandstones. Fluvial channels are typically filled with (C) cross-stratified coarse- to
medium-grained sandstones and conglomerates rich in lithic fragments (D). Coal fragments are common (F). Lowangle-cross-stratification occurs in places (G) and soft-sediment deformation is common (H). Coal horizons (I) and
plant roots (J) are common at the top of channel-fill units.
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Fig. 6. Measured sections and palaeocurrent measurements from a fluvial channel fill, ‘walked out’ in a dip-parallel
section on Brogniartfjellet. Dashed lines identify erosion surfaces.
Fig. 7. Fluvial channel fills can be ‘walked out’ into coeval tidally influenced deposits further seawards (southeast).
Example from Storvola.
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Fig. 8. A representative measured section through tidally influenced fluvial deposits of FA 2. Numbers by the
measured section refer to facies (see Table 1). Tidal influence is marked by subordinate landward-oriented palaeocurrent directions. The medium- to fine-grained sandstones contain bimodal cross-strata (A), a variety of high- and
low-angle compound cross-strata (B and E). Tidally influenced fluvial deposits occur as channel-fill units and have
multiple erosion surfaces (dashed lines in C). Sigmoidal cross-strata or mud drapes are rare (D). Coal horizons (F) are
rather common at the top of channel-fill unit.
mountainsides exposes coastal-plain deposits in
upper parts of the mountains (Fig. 2). The
coastal-plain deposits on Brogniartfjellet are
time-equivalent to shelf and slope deposits on
Storvola, and coastal-plain deposits on Storvola
have their contemporary shelf succession on the
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Stacked fluvial and tide-dominated estuarine deposits in high-frequency sequences
next mountain to the southeast, Hyrnestabben
(Fig. 1D and E).
Viewed from a distance, the coastal-plain facies
belt on Brogniartfjellet and Storvola is aggradational and mud-prone, although the whole basin
infill progrades to the south-east/east (Fig. 1). The
only obvious stratigraphic markers are thin sandprone, erosionally based levels that can be
‘walked out’ in a downdip direction across the
coastal-plain facies belt (Fig. 4).
Sedimentary facies were studied by using
measured vertical sections, ‘walking out’ stratigraphic levels downdip, and interpretation of
photomosaics taken from a helicopter. The documented sedimentary facies are grouped into seven
facies associations based on textures, sedimentary
structures, geometry, palaeocurrent indicators,
lateral facies transitions and position of the facies
on the dip-orientated transect. The individual
sedimentary facies are summarized in Table 1,
and Figs 5–19.
Facies Association 1: fluvial deposits
Facies Association (FA) 1 consists of lenticular
sandbodies that are 2–16 m thick and that can be
traced for up to 6–7 km across the NW–SEoriented mountainside outcrops (Fig. 4). The
sandbodies are dominated by granule conglomerates and coarse-grained sandstones (Fig. 5C, D).
These lenticular sandbodies typically have multiple basal and internal erosion surfaces, and the
amount of downcutting is in most cases equal to
the thickness of the sandbody itself, i.e. 2–16 m
(Figs 5A,B and 6). The sandbodies are volumetrically dominated by trough-cross-stratified,
coarse-grained, lithic sandstones (Facies 3), and
in most places based by crudely bedded (Facies 2)
and trough-cross-stratified (Facies 1) lithic conglomerates, coal fragments (up to 5 cm in diameter), and occasional clay chips (Figs 5 and 6,
Table 1). The conglomerates and coal fragments
typically occur above the internal erosion surfaces. The cross-set thickness varies from 0Æ1 to
0Æ9 m, but is most commonly 0Æ20 to 0Æ35 m.
The trough-cross-stratified sandstones (Facies
3) are interbedded with planar cross-stratified
sandstones (Facies 4), low-angle cross-stratified
sandstones (Facies 5), and plane parallel-laminated sandstones with parting lineations (Facies
6). Most of the channel fills do not fine upwards
significantly, but are capped by 0Æ01–0Æ5 m thick
very fine- to fine-grained rippled sandstones
(Facies 13), and plane parallel-laminated sandstones and mudstones (Facies 17). In some places,
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fine-grained rippled sandstones (Facies 13), and
plane parallel-laminated sandstones and mudstones (Facies 16) occur adjacent to the lenticular
sandbodies.
Palaeocurrent directions derived from crossstrata and ripples vary within individual sandstone bodies by up to 90, but generally only 40–
50. The mean palaeocurrent direction varies
between 100SE and 170SE, and most of the
currents are in the range 90E to 180S (Figs 5 and
6).
Water-escape structures, soft-sediment deformations and overturned cross-beds are common
in thicker units (Fig. 5H). Leaves, wood and other
plant fragments, rooted horizons or coal layers
occur occasionally at the top of, and adjacent to
the channel-shaped units (Fig. 5I and J).
Interpretation
The lithic conglomerates and coarse-grained sandstones, the extensive unidirectional (towards the
SE) cross-stratification, together with the channelshape, significant basal erosion, wood and plant
fragments, coal and root horizons, and lack of
marine trace fossils suggest deposition in fluvial
channels. The trough-cross-stratified conglomerates and sandstones (Facies 1 and 3) were deposited in 3D dunes and in bars. The associated
crudely bedded conglomerates represent lag
deposits related to the migration of the channel
thalweg and dunes (Kleinhans et al., 2002). The
planar cross-stratified sandstones (Facies 4), lowangle cross-stratified sandstones (Facies 5), and
plane parallel-laminated sandstones with parting
lineations (Facies 6) were formed as 2D dunes,
scour fills, and upper flow-regime plane beds
respectively. The ripple-laminated sandstones
(Facies 13), and plane-parallel-laminated sandstones and mudstones (Facies 17) were deposited
on adjacent floodplains or mark abandonment of
channels.
The multiple erosion surfaces indicate repeated
episodes of channel incision and infill. The lack
of lateral accretion beds, the coarse grain size of
the channel fills, the low abundance of overbank
deposits and relatively low palaeocurrent variability suggest that the channels had relatively low
sinuosity (Bridge et al., 2000). The abundant
water escape structures and associated soft-sediment deformation indicate high rates of deposition, consequent water escape and bed collapse,
or frequent slumping of channel banks (Owen,
1996). The overturned cross-strata mark high bedshear stresses, suggesting high river-current shear
strengths (Owen, 1996).
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Fig. 9. Depositional sequences 12–15 on the western side of Brogniartfjellet. See Fig. 3 for location. Arrows on the
right side indicate landward stepping facies shifts. The photo is a helicopter-taken photomosaic.
Facies Association 2: tidally influenced fluvial
deposits
Facies Association 1 can be walked-out into FA
2, by following the channel-shaped units towards the south-east along the mountainsides
(Fig. 7). FA 2 consists of lenticular coarse- to
fine-grained sandbodies, 1Æ5–15Æ0 m thick, with
multiple internal erosion surfaces. The amount
of erosion is similar to the thickness of the
sandbodies. The sandbodies contain upwardsfining successions of bipolar cross-stratified
sandstones (Facies 7), compound cross-stratified
sandstones (Facies 9), low-angle bipolar crossstratified sandstones (Facies 8), plane-parallellaminated sandstones with parting lineations
(Facies 6), and rare sigmoidal cross-stratification
(Facies 10; Fig. 8 and Table 1). Cross-set thicknesses in simple cross-strata vary between 0Æ2
and 0Æ4 m, whereas in compound cross-strata
between 0Æ3 and 1Æ0 m. Bipolar ripple-laminated
very-fine-grained sandstones (Facies 14) cap the
lenticular sandbodies.
There is a wide range of compound (or
inclined) cross-stratification in sandstones of FA
2 (Fig. 8B and E). In places the cross-strata are
steeply dipping (25–30) with reactivation surfaces, occasional mud drapes, or ripples climbing
up the inclined surfaces. In other places, compound cross-strata are characterized by low-angle
(5–15) dipping surfaces with decimetre-scale
cross-strata climbing up or down these surfaces.
Occasionally, the decimetre-scale cross-sets may
be overlain by asymmetric ripples or mud drapes.
The high-angle compound cross-stratification is
more common than the low-angle compound
cross-stratification in FA 2. Mud drapes or
organic debris drapes are rare. The sigmoidal
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Fig. 10. Depositional sequences 13–19 on eastern side of Brogniartfjellet. See Fig. 3 for location. Arrows on the right
side mark landward- and seaward-stepping facies shifts. See Fig. 9 for key.
cross-stratification is rare, and recognized by
downcurrent transitions in foreset angle from
gently dipping to more steeply dipping and back
to gently dipping, accompanied by increasing to
decreasing cross-strata thickness within the sets,
bounded by reactivation surfaces (Fig. 8D).
Palaeocurrent directions derived from crossstrata and ripples group into two sectors, northwest and southeast directions. The south-easterly
palaeocurrents vary within individual sandstone
bodies by up to 90, but generally only 40–50
(Figs 7 and 8). The mean south-easterly palaeocurrent direction varies between 124 and
152SE, and most of the currents are in the range
110 to 170SE. The north-westerly palaeocurrents vary within individual sandstone bodies by
up to 100, but generally only 30–40 (Figs 7 and
8). Mean north-westerly palaeocurrent direction
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Fig. 11. Depositional sequences 20–29 on Storvola. See Fig. 3 for location. Arrows on the right side mark landwardand seaward-stepping facies shifts.
varies between 260 and 330WNW, and most of
the currents are in the range 260 to 350WNW.
The palaeocurrent directions derived from plane
parallel-laminated intervals vary between 120
and 170SE.
Water-escape structures, soft-sediment deformations and overturned cross-beds are common
in thicker units. In places leaves and wood
fragments, rooted horizons or coal seams occur
at the top of the lenticular units (Fig. 8F).
Interpretation
The channel shape, extensive trough-crossstratification, coarse grain size, bipolar, but
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Fig. 12. Lateral relationship between tidally influenced fluvial deposits and high-sinuosity tidal channel deposits on
mountainside outcrop. Palaeocurrent measurements derived from cross-strata are shown in black, whereas the dipdirection of master surfaces (lateral accretion surfaces) is shown in grey.
dominantly south-eastwards palaeocurrents,
compound cross-strata, occasional mud drapes,
together with the walked-out transition from
fluvial channels, presence of plant roots and lack
of marine fossils suggest deposition in tidally
influenced fluvial channels (e.g. Allen, 1991),
where the fluvial channel changes into a tidal–
fluvial channel below the upstream tidal limit
(Dalrymple et al., 1992).
The bipolar cross-strata (Facies 7 and Facies 8)
indicate reversing currents of approximately
equal strength, whereas the compound crossstrata (Facies 9) reflect a dominant south-eastward current and a subordinate north-westward
current. The reactivation surfaces on steeply
dipping cross-strata indicate that a bedform leeside is changed into the stoss-side by reversals in
flow direction (e.g. Boersma, 1969; Boersma &
Terwindt, 1981a,b; Allen & Homewood, 1984;
Shanley et al., 1992), or migration of superimposed bedforms (McCabe & Jones, 1977;
Dalrymple, 1984; Shanley et al., 1992).
The compound cross-stratification with asymmetric ripples climbing up the lee faces of
high-angle cross-strata indicate a much stronger
dominant current (e.g. Allen, 1980). The
decimetre-scale cross-strata separated by lowangle inclined set boundaries show that the
dominant and subordinate currents did not differ
that greatly in strength (see Allen, 1980). The low
abundance of the low-angle compound crossstrata indicates that the subordinate (i.e. northwesterly) current rarely achieved high velocity,
and the depositional environment was dominated
by the south-easterly river current. The occasional mud drapes were deposited during slackwater periods. Individual cross-sets in these
facies are interpreted as tidal bundles, deposited
and modified in response to neap–spring–neap
tide fluctuations (e.g. Allen, 1980; Boersma &
Terwindt, 1981a,b; Dalrymple, 1984).
The sigmoidal beds with increasing to decreasing foreset angle and cross-strata thickness are
interpreted as representing as acceleration changing to full vortex flow conditions, followed by
deceleration within a single ebb tide (sigmoidal
beds in Shanley et al., 1992; see also Boersma &
Terwindt, 1981b; Allen & Homewood, 1984;
Kreisa & Moiola, 1986; Uhlir et al., 1988). The
reactivation surfaces bounding the sigmoidal
beds have also here been interpreted as resulting
from reversals of flow directions (Boersma, 1969;
Boersma & Terwindt, 1981a,b; Allen & Homewood, 1984; Shanley et al., 1992). The reactivation surfaces within the individual sigmoidal
beds have been attributed to the migration of
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Fig. 13. A representative measured section through high-sinuosity tidal channel deposits of FA 3. Numbers by the
measured section refer to facies (see Table 1). Palaeocurrent measurements derived from cross-strata are shown in
black, whereas the dip-direction of master surfaces (lateral accretion surfaces) is shown in grey. Low-angle compound cross-strata (A and F) and lateral accretion sets (B and C) dominate the facies association. Bimodal cross-strata
(D) occur in places. Organic debris drapes and mud drapes (E, H and I) are ubiquitous. Individual beds in accretion
sets are rippled, plane-parallel laminated or structureless (G). Small-scale soft sediment deformation is common (J).
Both, plant roots (H and L) and marine/brackish trace fossils (escape burrows on the photo K) occur in FA 3.
superimposed bedforms (McCabe & Jones, 1977;
Dalrymple, 1984; Shanley et al., 1992).
The plane parallel-laminated sandstones (Facies 6) were deposited during upper-flow-regime
conditions when south-easterly directed river
currents dominated. The bipolar ripple-laminated sandstones (Facies 10) represent abandonment of channels or were deposited in
interchannel areas.
Fluvial deposits: lateral and vertical
transitions
Fluvial deposits of FA 1 comprise laterally continuous units of relatively thin fluvial conglomerates and coarse sandstones above regional
erosion surfaces (Figs 9–11). Most of the fluvial
channel fills are 4–5 m thick, whereas some reach
16 m locally (sequence 14 in Fig. 9; sequences 14
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Fig. 14. Lateral relationship between high-sinuosity tidal channel deposits and coeval low-sinuosity tidal channel
deposits further to the southeast (seawards). The latter can be ‘walked out’ into upper-flow-regime tidal flats and tidal
sand bars even further to the south-east. Palaeocurrent measurements derived from cross-strata are shown in black,
and the dip direction of inclined master surfaces and lateral accretion surfaces in grey.
Fig. 15. A representative measured section through low-sinuosity tidal channel deposits of FA 4. Numbers by the
measured section refer to facies (see Table 1). Low-angle compound cross-stratification (A, E and F) dominates, but
high-angle compound cross-stratification (B), low-angle cross-stratification (C), and bipolar cross-strata (D) also occur
in these erosionally based (H), very-fine grained sandstones. Both plant fragments and marine/brackish trace fossils
occur (G and I). Pl – Planolites, Sk – Skolithos, Te – Teichichnus.
and 19 in Fig. 10; sequence 29 in Fig. 11). Individual fluvial channels can be ‘walked out’ across
the whole coastal-plain length, i.e. for 6–7 km
along the individual mountainsides (Fig. 3), and
they occur with vertical spacing of 7–44 m
(Figs 9–11).
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2005 International Association of Sedimentologists, Sedimentology, 52, 391–428
Stacked fluvial and tide-dominated estuarine deposits in high-frequency sequences
407
Fig. 16. A representative measured section through upper-flow-regime tidal flat deposits of FA 5. Numbers by the
measured section refer to facies (see Table 1). Plane-parallel lamination with current lineations (D) dominates, but
occasional trough-cross-stratification (A, B and C), and occasional sigmoidal stratification (A) also occur. Erosion
surfaces with several tens of centimetres of erosion are rather common (E).
Fig. 17. Tidal sand bars are large elongate bedforms that can be ‘walked out’ for several kilometres along dipdirection mountainside exposures. Palaeocurrent measurements derived from cross-strata are shown in black,
whereas dip-direction of master surfaces (where measured) is shown in grey.
Tidally influenced fluvial deposits of FA 2
occur at the south-eastward (seaward) ends of the
fluvial channels (sequences 19, 20 in Fig. 11), or
where the tops of the fluvial channel fills grade
into tidally influenced fluvial deposits (sequences
14 and 15 in Fig. 9; sequence 14 in Fig. 10;
sequences 23, and 29 in Fig. 11). The vertical
change into tidally influenced deposits marks the
landward migration of the bayline, and occurred
when fluvial channels were submerged below the
mean tidal limit.
Besides the described fluvial intervals, the FA 1
and 2 also occur, at the landward end of the tidaldominated intervals that are discussed in more
detail below (sequence 16 in Fig. 10; sequences
20, 21, 23, 27 in Fig. 11).
Facies Association 3: high-sinuosity tidal
channels
Facies Association 2 can be walked out into FA 3,
by following the channel-shaped units towards
the south-east along the mountainsides (Fig. 12).
FA 3 consists of lenticular sand-prone bodies,
0Æ6–4 m thick and 8–25 m wide. The association
typically has a basal erosion surface, but the
amount of erosion is only 0Æ6–1Æ0 m (in rare cases
up to 1Æ5 m). The lenticular bodies consist of
compound cross-stratified fine-grained sandstones with bipolar dip directions in adjacent
sets (Facies 9), inclined heterolithic strata (Facies
12), and bipolar cross-stratified fine-grained sandstones (Facies 7; Table 1; Fig. 13). Grain size in
FA 3 does not exceed fine sand. Mud drapes are
ubiquitous through the FA 3.
A wide range of compound (or inclined) crossstratification (Facies 9) occurs in FA 3, similar to
FA 2. However, in FA 3 the compound-crossstrata are dominated by broad, low-angle (5–15)
dipping surfaces with smaller sets of cross-stratification or ripple cross-lamination dipping up or
down the bedding surfaces (Fig. 13A and F). The
smaller sets of cross-strata are in many places
overlain by asymmetric ripples and/or mud
drapes. Cross-stratification with reactivation surfaces is also common, and the reactivation surfaces are typically covered with single or double
mud drapes or coal drapes (Fig. 13E, H, and I).
The inclined heterolithic strata (Facies 12) in
0Æ4–1Æ0 m thick sets, consist of low-angle (5–10)
inclined or sinusoidal beds, 1–5 cm thick. The
individual sinusoidal beds consist of sandstonemudstone couplets. The sandstone intervals are
structureless, plane parallel- or ripple-laminated.
2005 International Association of Sedimentologists, Sedimentology, 52, 391–428
408
P. Plink-Björklund
Fig. 18. A representative measured section through tidal sand bar deposits of FA 6. Numbers by the measured
section refer to facies (see Table 1). Palaeocurrent measurements derived from cross-strata are shown in black, and
dip direction of inclined master surfaces in grey. Trough-cross-stratified sandstone sets with bimodal palaeocurrent
directions in adjacent depositional units (A, E and F) dominate. Occasional sigmoids (A and B), and high-angle
compound cross-strata (C) occur. In seaward reaches of the outcrop belts the tidal sand bar deposits erosionally cover
wave-reworked intensively bioturbated sandstones (A and D).
2005 International Association of Sedimentologists, Sedimentology, 52, 391–428
Stacked fluvial and tide-dominated estuarine deposits in high-frequency sequences
409
Fig. 18. Continued.
The sandstone intervals are in places erosionally
based (2–10 cm of erosion). The mudstone intervals drape the sandstone intervals (Fig. 13B and
G). The sandstone–mudstone couplets can be
traced from the upper parts of the inclined strata
to the basal portions. The dip of the inclined
heterolithic strata is oriented in a direction
normal to the palaeocurrent indicators derived
from cross-stratification (see below). Bipolar
cross-stratified sandstones (Facies 7) are finegrained and their foresets and bottomsets are
commonly draped with single or double mudstone or coal drapes (Fig. 13D). The cross-sets are
only 0Æ1–0Æ2 m thick, and in places climbing at
the top of the inclined heterolithic sets.
Palaeocurrent directions derived from crossstrata and ripples group into two modes, northwest
and southeast. The south-easterly palaeocurrents
vary within individual sandstone bodies by
up to 110, but generally up to 80 (Figs 12–14).
Mean south-easterly palaeocurrent directions
vary between 100 and 170SE, and most of the
currents are in the range 90E to 180S. The northwesterly palaeocurrents vary within individual
sandstone bodies by up to 130, but generally up
to 80 (Fig. 13). Mean north-westerly palaeocurrent directions vary between 280 and 345NW,
and most of the currents are in the range 270 to
350NW. Most of the inclined heterolithic sets
dip towards 10–90NE and 180–270SW.
Water-escape structures, and decimetre-scale
soft-sediment deformation are common (Fig. 13J).
In places plant and wood fragments, rooted
horizons or coal layers occur at the top of the
lenticular units (Fig. 13H and L). In other places
Skolithos, Planolites, rare Teichichnus burrows,
or escape burrows are found (Fig. 13K).
Interpretation
The channel-shape, extensive bipolar compoundcross-stratification, ubiquitous single and double
mud drapes, and marine/brackish trace fossils
suggest deposition in tidal channels. The inclined
heterolithic strata are interpreted as lateral accretion deposits on tidal point bar surfaces, and
indicate that FA 3 was deposited in high-sinuosity tidal channels. Rooted horizons and coal
layers mark abandonment of the channels.
The bipolar cross-strata (Facies 7) indicate
reversing currents of approximately equal
2005 International Association of Sedimentologists, Sedimentology, 52, 391–428
410
P. Plink-Björklund
Fig. 19. A representative measured section through marginal tidal mixed- to mud-flat and marsh deposits of FA 7.
Numbers by the measured section refer to facies (see Table 1). Flaser, wavy and lenticular bedding (A, B, C, E and G)
dominate the locally bioturbated (F) marginal tidal flat deposits that represent an inter-tidal succession of decreasing
current energy. Coaly mudstones and coals (B) rich in plant fragments, roots, stems in growth position, in places
reworked by pedogenic processes (C, H, I and J) are interpreted as the uppermost, supratidal parts of tidal mud flats
and marshes. Pl – Planolites, Sk – Skolithos, Te – Teichichnus.
strength, whereas the compound cross-strata (Facies 9) reflect the existence of a dominant and a
subordinate current. Compared to the high-angle
compound-cross-strata in FA 2, the low-angle
compound-cross-strata indicate that the dominant and subordinate currents did not differ that
greatly in strength (see Allen, 1980). The lowangle ‘master’ bedding surfaces are formed by
more intense reworking of the dominant-current
bedforms by the subordinate current. The reworking is directly proportional to the strength of the
subordinate current (Dalrymple et al., 1990). The
transition from low-angle compound cross-strata
into smaller sets of cross-strata, into current
ripples and mud drapes reflects decelerating
current velocity (e.g. Shanley et al., 1992), and
is interpreted as reflecting a tidal cycle.
The persistent mud drapes that can be traced
from the upper surface of the lateral accretion sets
into the deepest parts of the channel, combined
with inclined sandstone beds that are ripplelaminated or plane-parallel-laminated through-
2005 International Association of Sedimentologists, Sedimentology, 52, 391–428
Unimodal:
80–180
Unimodal:
80–180
Crudely normally graded,
structureless, or imbricated
Trough-cross-stratified, in places
overturned cross-strata and soft
sediment deformation. Multiple
erosion surfaces lined with
coal clasts or clay chips
Planar-cross-stratified, in places
soft-sediment deformed
Low-angle (<10) cross-stratification,
in places soft-sediment deformed
Plane-parallel lamination,
parting lineations
Trough- or planar-cross-stratified,
in places overturned cross-strata
and soft sediment deformation.
Multiple erosion surfaces lined
with coal clasts or clay chips
Granules and occasional
pebbles, coal clasts
(<5 cm), clay chips
Fine- to coarse-grained
sandstones, coal clasts
(< 4 cm), occasional
clay chips
Fine- to coarse-grained
sandstones, coal clasts
(<4 cm), occasional
clay chips
Fine- to medium-grained
sandstones, coal clasts
and clay chips
Fine-grained sandstones,
coal clasts and
clay chips
Fine- to coarse-grained
sandstones, occasional
mud drapes, coal clasts
and occasional clay chips
2: Crudely
bedded
conglomerate
3: Trough
cross-stratified
sandstone
4: Planar
cross-stratified
sandstone
5: Low-angle
cross-stratified
sandstone
6: Plane-parallel
laminated
sandstone
7: Bipolar
cross-stratified
sandstone
Bimodal:
80–180
and
250–350
Unimodal:
80–180
Trough-cross-stratified, in places
soft-sediment deformed
Granules and occasional
pebble, coal clasts
(<5 cm), clay chips
1: Trough
cross-stratified
conglomerate
2005 International Association of Sedimentologists, Sedimentology, 52, 391–428
In places wood and plant
fragments, occasional roots
and coal layers at the top. In
other places Skolithos,
Planolites, rare Teichichnus,
or escape burrows
In places wood and plant
fragments. Occasional
roots and coal layers
at the top
Wood, leaves, rooted
horizons (up to 1 m
long roots). Occasional
coal layers at the top
Wood, leaves, rooted
horizons (up to 1 m
long roots). Occasional
coal layers at the top
Wood
Wood
Palaeocurrent
directions
Trace fossils and biota
Structures
Textures
Facies
Table 1. Sedimentary facies.
0Æ2–0Æ4 m thick crosssets close to the base
of erosionally based
lenticular units, and
above internal
erosion surfaces
0Æ05–0Æ2 m thick,
at the base of
erosionally based
lenticular units,
line internal erosion
surfaces
0Æ1–0Æ9 m thick,
most commonly
0Æ20–0Æ35 m thick
cross-sets in
erosionally based
lenticular units
0Æ6–15Æ5 m thick
0Æ2–0Æ9 m thick
cross-sets in
erosionally based
lenticular units
0Æ6–15Æ5 m thick
0Æ2–0Æ3 m thick
cross-sets
interbedded with
high-angle crossstrata, in erosionally
based, lenticular
units in FA 1
0Æ2–0Æ3 m thick beds
interbedded with
high-angle crossstrata, in erosionally
based, lenticular
units or in 0Æ5–1Æ5 m
thick tabular units
0Æ1–0Æ4 m thick crosssets in erosionally
based 0Æ3–5Æ0 m thick
lenticular units
Occurrence
Migration of 3D
dunes in channels,
bidirectional currents
of equal strength
Upper-flow-regime
plane beds
Downstream migration
of 2D dunes,
oblique migration
of bars in channels,
unidirectional current
Scour fills, or washedout dunes,
unidirectional current
Downstream migration
of 3D dunes, oblique
migration of
longitudinal bars
in channels,
unidirectional current
Lag deposits,
longitudinal bars,
unidirectional current
Migration of 3D dunes
in channels,
unidirectional current
Interpretation
Stacked fluvial and tide-dominated estuarine deposits in high-frequency sequences
411
15: Bimodal
rippled
sandstone beds
with mud
drapes
Very fine- to fine-grained
sandstones separated
by mud drapes
Bimodal
80–180
and
250–350
Bimodal
Bipolar dip directions in adjacent
ripple sets, separated by mud drapes;
in places convoluted and soft
sediment deformed
Unimodal:
80–180
Bimodal,
cross-strata
280–360
150–180,
master surfaces
70–140
220–280
Bimodal
110–170
and
240–350
Asymmetrical ripples
Fine- and very-fine-grained Inclined (at angles between 10–15),
internally structureless, plane-parallel
sandstones bounded by
laminated or ripple-laminated
mudstone laminae
sandstone strata, separated
by mud drapes
Very fine- to fine-grained Asymmetrical ripples
sandstone, and siltstones
14: Bimodal ripple- Very fine- to fine-grained
sandstone, and
laminated
siltstones, occasional
sandstone
mud drapes
13: Ripplelaminated
sandstone
12: Inclined
heterolithic
strata
Medium- to coarse11: Trough-crossstrata on inclined grained sandstones
master surfaces
Fine- to mediumgrained sandstones
10: Sigmoidal
cross-stratified
sandstone
Bimodal
80–180
and
250–350
Very-fine- to coarse-grained
sandstones, in places
mud drapes or organic
debris drapes
9: Compound
cross-stratified
sandstone
(1) high-angle (25–30) cross-strata
with reactivation surfaces and mud
drapes, (2) high-angle (25–30)
cross-strata with smaller sets of
cross-strata or ripples climbing
up or down the surfaces (3) low-angle
(5–15) cross-strata with smaller sets
of cross-strata or ripples climbing
up or down the surfaces
Downcurrent transition in foreset
angle from gently dipping to more
steeply dipping and back to gently
dipping, accompanied by increasing
to decreasing cross-strata thickness
within the sets, bounded by
reactivation surfaces.
Inclined master surfaces with
superimposed cross-strata
Bimodal
Low-angle (< 10) cross-stratification,
in places soft-sediment deformed,
in places ubiquitous reactivation surfaces
Fine- to medium-grained
sandstones, occasional
mud drapes, coal clasts
and clay chips
8: Low-angle
bipolar crossstratified
sandstone
Occurrence
Interpretation
Acceleration changing
to full vortex flow
conditions, followed
by deceleration
within a single tide
Unimodal current
ripples deposited
during abandonment
of channels, or in
overbank environment
Bimodal current
ripples deposited
during abandonment
of channels, or in
overbank environment
Migration of ripples
from bidirectional
currents with muddraping during
ensuing slack-water
periods
Tidal bundles in
lateral accretion
sets (point bars)
Cross-sets: 0Æ05–0Æ5 m, Migration of 3D
dunes on laterally
depositional units:
accreting bar surfaces
5–15 m
0Æ2–0Æ4 m thick sets
in erosionally based
1Æ5–5Æ0 m thick
lenticular units
Sandstones:
0Æ01–0Æ1 m;
mudstones: to
0Æ005 m in 0Æ3–3Æ0
thick units
0Æ001–0Æ05 m thick
Wood, leaves, rooted
ripple sets, in flathorizons (up to 1 m long
roots). Occasional coal layers based tabular units
that cap lenticular
units
0Æ001–0Æ05 m
In places wood and plant
thick ripple sets,
fragments. Occasional roots
in flat-based tabular
and coal layers at the top
units that cap
lenticular units
Flat-based sheet-like
In places Skolithos,
units 0Æ3–3Æ0 m thick.
Planolites, in other
Ripple sets: a few cm
places plant roots,
a few tens of cm long
Skolithos, Planolites,
rare Teichichnus,
or escape burrows
In places wood and
plant fragments, or
occasional roots and
coal layers at the top
In places wood and plant
fragments, or occasional
roots and coal layers
at the top
0Æ1–0Æ3 m thick cross- Scour fills, or
washed-out dunes,
sets, in erosionally
based lenticular units reworking by
subordinate current,
bidirectional current
0Æ3–1Æ0 m thick cross- Downcurrent
In places wood and plant
migration of 3D
sets in erosionally
fragments, or occasional
dunes and oblique
based 0Æ6–5Æ0 m
roots and coal layers
thick lenticular units migration of bars in
at the top. In other
channels,
places Skolithos,
bidirectional current,
Planolites, rare Teichichnus,
mud drapes from
or escape burrows
slack-water periods
Palaeocurrent
directions
Trace fossils and biota
Structures
Textures
Facies
Table 1. Sedimentary facies.
412
P. Plink-Björklund
2005 International Association of Sedimentologists, Sedimentology, 52, 391–428
Flat-based sheet-like
units 0Æ2–1Æ0 m thick.
Laminae: a few mm
0Æ005–0Æ1 m thick beds,
in flat-based, sheet-like
about 0Æ5 m thick units,
that cap lenticular units
Flat-based units 0Æ1–9Æ0 m
thick, Coal layers:
0Æ02–0Æ2 m
In places Skolithos,
Panolites, In other places
plant roots, a few tens
of cm long
Wood, leaves, rooted
horizons (up to 1 m long roots),
occasional coal layers.
Leaves, stems in growth
position, plant roots a few
tens of cm to 1 m long
Very low-amplitude sinusoidal
lamination, in places convoluted, soft
sediment deformed
Parallel stratification, occasional ripples
Laminated or structureless
mudstones with coal layers
Very fine- and fine-grained
sandstones and mudstones
Very fine- to fine-grained
sandstones alternating
with mudstone
Mudstones, organic
debris, coal
16: Sinusoidally
laminated beds
17: Plane-parallel
laminated
heterolithics
18: Organic-rich
mudstones
and coals
Migration of ripples
at angles exceeding
the stoss slope angles,
mud-draping during
ensuing slack-water
periods
Alternating deposition
from traction currents
and suspension,
channel abandonment,
or overbank
Mudstones deposited
from suspension; coals
formed by rapid
accumulation of plant
fragments and peats
Stacked fluvial and tide-dominated estuarine deposits in high-frequency sequences
413
out, suggest a relatively even shear velocity
distribution within the channel (Shanley et al.,
1992). Such a velocity distribution is markedly
different from that normally encountered in fluvial systems (see also Thomas et al., 1987; Rahmani, 1988; Smith, 1988; Nio & Yang, 1991). The
fine interlamination of sandstones and mudstones in tidal point bars is interpreted as
reflecting deposition during single tidal cycles
(Bridges & Leeder, 1976).
The ubiquitous occurrence of mud drapes
reflects deposition within the turbidity maximum
zone (McCave, 1979; Jouanneau & Latouche,
1981; Dalrymple et al., 1990; Allen, 1991; Dalrymple, 1992). Deposition of fine-grained sediment occurs due to abrupt changes in shear
velocity and is promoted by flocculation of clay
particles due to salinity mixing of fluvial fresh
water and marine waters (Nichols & Biggs, 1985).
Facies Association 4: low-sinuosity tidal
channels
Facies Association 3 can be walked out into FA 4
by following the lenticular bodies towards the
south-east along the mountainsides (Fig. 14). FA 4
consists of lenticular sand-prone bodies, 0Æ6–4 m
thick and 10–15 m wide. The association is
typically erosionally based, but the amount of
erosion is only 0Æ6–1Æ0 m (in rare cases up to 1Æ8 m,
Fig. 15H). The lenticular bodies consist of compound cross-stratified fine-grained sandstones
with bipolar dip directions in adjacent sets (Facies
9), bi-directional cross-stratified sandstones
(Facies 7), and low-angle cross-stratified sandstones (Facies 8; Table 1; Fig. 15). Mud drapes are
ubiquitous, but somewhat less abundant than in
the FA 3. Grain size does not exceed fine sand.
The compound cross-stratified sandstones (Facies 9) are in most places fine grained. In places
the cross-strata are steeply dipping (25–30) with
reactivation surfaces and mud drapes, or smaller
cross-strata or ripples climbing up the inclined
surfaces. In most places, the compound crossstrata are characterized by low-angle (5–15)
dipping surfaces with smaller sets of cross-stratification or ripple cross-lamination climbing up or
down the bedding surfaces (Fig. 15A–C, E and F).
The bipolar cross-stratified sandstones (Facies 7)
are fine grained (Fig. 15D). Cross-sets are typically 0Æ1–0Æ2 m thick. The low-angle (<10) crossstrata (Facies 8) display reactivation surfaces and
a slightly sigmoidal shape, but lack mud drapes.
Palaeocurrent directions derived from crossstrata and ripples group into two modes, north-
2005 International Association of Sedimentologists, Sedimentology, 52, 391–428
414
P. Plink-Björklund
west and south-east directions. The variation
of palaeocurrent directions within individual
lenticular bodies is less compared with FA 3.
The south-easterly palaeocurrents vary within
individual sandstone bodies by up to 70, but
generally only 40–50. Mean south-easterly palaeocurrent directions vary between 105 and
150SE, and most of the currents are in the range
100 to 160SE. The north-westerly palaeocurrents vary within individual sandstone bodies by
up to 80, but generally only 40–50. Mean northwesterly palaeocurrent directions vary between
270 and 310NW, and most of the currents are in
the range 280 to 310NW (Figs 14 and 15).
Water-escape structures, and decimetre-scale
soft-sediment deformation are common. In places
plant and wood fragments, rooted horizons or
coal layers occur at the top of the lenticular units
(Figs 14 and 15). In other places, Skolithos,
Planolites, rare Teichichnus burrows, or escape
burrows are found (Fig. 15).
Facies Association 4 has many similarities with
FA 3, as both associations are characterized by a
dominance of very fine- to fine-grained sandstones with bipolar palaeocurrent indicators,
ubiquitous mud drapes, occurrence of roots as
well as marine trace fossils. The difference is that
FA 4 lacks inclined heterolithic strata, has a
narrower distribution of palaeocurrents, and mud
drapes are somewhat less common.
Interpretation
The channel-shape, extensive bipolar compoundcross-stratification, ubiquitous single and double
mud drapes, and marine/brackish trace fossils
suggest deposition in tidal channels. Rooted
horizons and coal layers mark abandonment of
the channels. The lack of lateral accretion beds,
together with lower variability of palaeocurrent
directions suggests lower sinuosity of tidal channels in FA 4 compared to otherwise very similar
FA 3. The bipolar cross-strata (Facies 7) and
compound cross-strata (Facies 9) are interpreted
similarly to FA 3. The low-angle cross-strata (8)
are interpreted as reflecting intense reworking by
the subordinate current similar to low-angle
compound-cross-strata. The intense reworking
may have removed the mud drapes.
Facies Association 5: upper-flow-regime tidal
flats
Facies Association 4 can be walked out into FA 5
by following the lenticular bodies to the southeast along the mountainside (Fig. 14). FA 5
consists of plane parallel-laminated sandstones
(Facies 6), and occasional trough cross-stratified
sandstones (Facies 7) and sigmoidal cross-stratified sandstones (Facies 10; Fig. 16, Table 1).
The plane parallel-laminated sandstones (Facies 6) show parting lineations, and consist of
fine-grained sandstones. Individual beds, 0Æ2–
0Æ4 m thick, are sandy throughout, and in places
erosionally based (Fig. 16A, B, C and D). In other
places, the beds are composed of a lower set or
two of trough cross-stratified sandstones, or are
capped by a set of sigmoidal beds (Fig. 16E). The
beds are grouped into 0Æ5–1Æ2 m thick depositional units that can be followed for more than a
kilometre along the mountainside exposures
(Fig. 17).
The 0Æ1–0Æ2 m thick sets of trough cross-stratified
sandstones (Facies 7) are fine to medium grained,
and have occasional coal fragments in foresets. The
sigmoidal cross-strata (Facies 10) display reactivation surfaces, but lack mud drapes. Individual sets
are up to 0Æ3 m thick, and consist of fine- to
medium-grained sandstones.
Palaeocurrent indicators derived from parting
lineations, sigmoidal beds and cross-strata are
towards the north-west. The palaeocurrent directions vary within individual sandstone bodies by
only up to 20, and the mean palaeocurrent
direction varies between 260SW and 350NW,
and most of the directions are in the range
240SW to 350NW. The south-easterly palaeocurrents from cross-strata are in the range of
110SE to 190SW. Palaeocurrent directions show
reversals in adjacent units (1–3 m thick), rather
than in adjacent sets.
Interpretation
Plane parallel-laminated fine-grained sandstones
with parting lineations and strongly dominant
north-westerly palaeocurrent directions, together
with sand-prone cross-strata, indicate upper flow
regime traction deposition by flood currents. The
plane parallel-laminated intervals formed in
maximum tidal flow velocities, whereas the
cross-strata at the base or top developed during
accelerating or waning flow conditions (e.g. Kreisa & Moiola, 1986).
Upper flow regime (UFR) tidal sand flats have
been considered diagnostic for macrotidal estuarine environments, as they have only been
reported from the axial portions of modern tidedominated macrotidal estuaries (Hamilton, 1979;
Lambiase, 1980a,b; Dalrymple et al., 1990; Dalrymple, 1992). Current speeds that exceed
2 m s)1, and water depths <2–3 m have been
2005 International Association of Sedimentologists, Sedimentology, 52, 391–428
Stacked fluvial and tide-dominated estuarine deposits in high-frequency sequences
measured in modern environments, where upper
flow regime plane beds form in fine sand (Dalrymple et al., 1990).
Most of the cross-stratified sandstones were
deposited by slower ebb currents, whereas the
sigmoidal beds show landwards palaeocurrent
directions. The reversal of palaeocurrent directions in adjacent depositional units (several
metres thick) rather than in adjacent cross-sets
suggests that flood and ebb currents used slightly
different paths for several tidal cycles.
Facies Association 6: tidal sand bars
Facies Association 5 can be walked out into FA 6
by following the plane-parallel laminated sandbodies to the southeast along the mountainsides
(Figs 14 and 17). FA 6 consists of quartz-rich,
well-sorted, medium- to coarse-grained sandstones. FA 6 is dominated by about 5–15 m thick
inclined master surfaces with superimposed decimetre-scale cross-strata (Facies 11). In places,
sigmoidal beds (Facies 10) and compound-crossstrata (Facies 9) occur (Figs 17 and 18, Table 1).
The dip of the inclined master surfaces is
oriented in a direction at 40 to 60 to the
palaeocurrent indicators derived from superimposed cross-stratification. These large inclined
beds are erosionally based. The amount of erosion
varies from a few decimetres to several metres,
and in rare cases up to 10 m. The erosionally
based inclined beds are organized into depositional units that thin updip and downdip, and
they can be walked out on mountainside exposures for 1–2 km (Fig. 17).
The superimposed cross-sets are 0Æ05–0Æ5 m
thick and display bipolar palaeocurrent directions in adjacent depositional units, a few decimetres to a few metres thick (Fig. 18). This is in
contrast to FA 2, 3 and 4 where the palaeocurrent
directions are bipolar in adjacent sets. Grain-size
tends to vary together with the palaeocurrent
reversals in the depositional units. Sigmoidal
beds (Facies 10) are more common in FA 6 than in
other facies associations. Compound-cross-strata
(Facies 9) are less common, and represented by
high-angle cross-strata with reactivation surfaces,
or high-angle cross-strata with smaller sets of
cross-strata or asymmetric ripples dipping up the
inclined surfaces.
Palaeocurrent directions derived from crossstrata group into two sectors, north-west and
south-east. The north-westerly currents dominate over the south-easterly currents. The northwesterly palaeocurrents vary within individual
415
sandstone bodies by up to 50, but generally
only 30 (Figs 17 and 18). Mean north-westerly
palaeocurrent directions vary between 280 and
350NW, and most of the currents are in the
range 280NW to 360N. The south-easterly
palaeocurrents vary within individual sandstone
bodies by up to 30 (Figs 16 and 18). Mean
south-easterly palaeocurrent directions vary between 150 and 170SE, and most of the
currents are in the range 150SE to 180S.
Palaeocurrent directions from the inclined master surfaces dip towards 70NE–140SE and
220SW–280NW (Figs 17 and 18).
When walked out further southeast, FA 6 cuts
into wave-reworked and heavily bioturbated
deposits (Fig. 18). The wave-reworked deposits
are characterized by wave ripples, low-angle (<5)
cross-strata, swaley cross-strata, and very intensive bioturbation (including Ophiomorpha, Terebellina). When walked out towards the northwest,
upper-flow-regime tidal flats (FA 5) overlie FA 6
(Fig. 17). In other places, FA 6 is covered by coal
layers, coaly mudstones or by FA 7.
Facies Association 6 has been documented only
in some of the sequences (sequences 13, 15, 16
and 18 on Brogniartfjellet, and sequences 20, 21
and 24 on Storvola, Figs 10 and 11), as they occur
in the most south-eastward (seaward) end of the
coastal-plain exposures.
Interpretation
The erosional base, extensive trough-cross-stratification, lateral accretion, coarse grain-size, bipolar palaeocurrent directions in adjacent
depositional units, together with the significant
height and length of the sandbodies suggests
deposition as tidal sand bars in a subtidal environment. Such large tidal bars are characteristic
of the seaward portions of most macrotidal environments (Hayes, 1975; Harris, 1988; Dalrymple &
Zaitlin, 1989; Dalrymple et al., 1990). Tidal sand
bars have been described from estuaries, tidedominated delta fronts and tide-dominated shallow-marine settings (sand ridges; Swift, 1975;
Dalrymple et al., 1990). A lack of bioturbation
and mud drapes in the tidal bars is most probably
due to the rapid migration of these large bedforms
(Amos et al., 1980; Yeo & Risk, 1981).
The reversals of palaeocurrents in adjacent sets
indicate that flood and ebb currents used slightly
different paths (see also Dalrymple et al., 1990;
Dalrymple, 1992). The more quartz-rich and
better-sorted character of the sandstones indicates
an active marine sediment supply by flood
currents.
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P. Plink-Björklund
Facies Association 7: mixed to muddy tidal
flats and marshes
Facies Association 7 occurs lateral to the described FA 2–6. FA 7 is volumetrically dominated
by heterolithic beds of ripple-cross-laminated
fine- to very fine-grained sandstones, separated
by mud drapes (Facies 15). In some places the
asymmetrical ripples have opposite dip directions in adjacent sets (Table 1, Fig. 19A, B, C, E
and G). In other places the ripples are unidirectional. The mud drapes vary in thickness from a
millimetre to 0Æ5 cm. In places the rippled heterolithic beds are dominated by sandstone, or the
sandstone-mudstone content is equal, and in
other places the mudstone dominates the heterolithic beds. Sandstone-prone beds with aggradational asymmetric ripples (Facies 16) are also
common. These beds vary from climbing ripples
to vertically aggrading ripples, separated by thin
mud drapes (Fig. 19H).
Small-scale deformation is common in both
types of rippled beds (Fig. 19D). In places the
rippled beds are bioturbated (Skolithos, Planolites, Fig. 19F), in other places they are intensively rooted or reworked by pedogenic
processes. Locally, whole tree leaves (Fig. 19J)
occur and fragments of plant stems are found in
the growth position (Fig. 19H and I).
Facies 15 and Facies 16 grade laterally and
vertically into 0Æ3–9 m thick organic-rich mudstones (Facies 18). The mudstones commonly
contain root traces, and are rich in tree leaves,
plant stems in growth position and up to 20 cm
thick coal layers (Fig. 19B). Characteristically, the
rippled beds of Facies 15 and Facies 16 become
finer-grained upwards and grade into organic-rich
mudstones, or they become coarser-grained upwards and are cut by tidal channels or bars. In
places lenticular sandstone-prone bodies (Facies
9), 0Æ5–1 m thick occur. These lenticular bodies
are typically based by a lag of mud clasts and
consist of compound-cross-stratified fine-grained
sandstones.
Palaeocurrent directions derived from ripplecross-lamination are very widely spread.
Although there are two dominant groups (towards
280NW and towards 110SE), the directions vary
by almost 360 (Fig. 19).
Interpretation
The succession of bipolar rippled heterolithic
beds with mud drapes (Facies 15) is interpreted
as flaser, wavy and lenticular bedding, according
to the sandstone/mudstone ratio (see Reineck &
Wunderlich, 1968). Flaser, wavy and lenticular
bedding with marine/brackish trace fossils typically indicates deposition from reversing tidal
currents (Reineck & Wunderlich, 1968; Reineck &
Singh, 1980) and represents an inter-tidal succession of decreasing current energy (Dalrymple,
1992). Rippled sands are deposited during maximum tidal flow, whereas mud drapes form
during ensuing slack-water periods.
Low-angle sinusoidal lamination (Facies 16)
has been documented as formed by ripples that
climb at angles exceeding the stoss slope angles of
ripples (Allen, 1968; Yokokawa et al., 1995). This
type of lamination has also been called ‘sinusoidal ripple-lamination’ (Jopling & Walker, 1968) or
‘draped lamination’ (Ashley et al., 1982). Increasing angles of climb in this type of ripple lamination indicate that the ratio of the vertical bed
aggradation rate to the downstream ripple migration rate increases (Ashley et al., 1982), i.e. the
ratio of deposition between suspended bed material and traction bed load increases (Jopling &
Walker, 1968), compared to flaser and wavy
lamination beds.
Organic-rich mudstone and coals (Facies 18),
rich in plant fragments and roots and in places
reworked by pedogenic processes are interpreted
as the uppermost, supratidal parts of tidal mud
flats and marshes. The small-scale channelled
units of compound cross-stratified sandstones
(Facies 9) indicate deposition in tidal gullies that
crossed the tidal flats (Dalrymple et al., 1991).
The great variability of palaeocurrents suggests
that the tidal currents tend to become shorelineperpendicular in FA 7. FA 7 is interpreted as
intertidal to supratidal mixed to mud flats and
marshes that rim the margins of the higher energy
tidal environments described in FA 2–6.
Tidal deposits: lateral and vertical transitions
Tidal deposits (FA 4–7) stratigraphically overlie
the basal fluvial channels, and form the bulk
volume of the coastal-plain depositional sequences in the Eocene Central Basin. The fluvial
and tidally influenced fluvial deposits (FA 1 and 2)
are restricted to the most north-westward (landward) portions of the tidally dominated intervals
(Fig. 20). They pass to the south-east into high- (FA
3) and low-sinuosity (FA 4) tidal channel deposits,
and then into upper flow-regime tidal flat deposits
(FA 5) and at the seaward end of the profile into
tidal sand bars (FA 6). The low-energy tidal mixed
to mud flats (FA 7) occur lateral (marginal) to the
described axial succession (Fig. 20).
2005 International Association of Sedimentologists, Sedimentology, 52, 391–428
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417
Fig. 20. Seaward facies transitions, grain-size trends and palaeocurrent directions indicate deposition in a tidedominated estuary. The estuary model is redrawn after Dalrymple et al. (1992). Palaeocurrent measurements derived
from cross-strata are shown in black, and dip direction of inclined master surfaces in grey.
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418
P. Plink-Björklund
Grain-size decreases along the described axial
profile towards the south-east, from conglomerates and coarse-grained fluvial sandstones (FA 1)
to medium- and fine-grained tidally influenced
fluvial sandstones (FA 2), and into fine- to very
fine-grained sandstones with ubiquitous mud
drapes in high-sinuosity tidal channels (FA 3;
Fig. 20). Grain size also decreases towards the
northwest from the seaward end of the profile,
from coarse- to medium-grained tidal sand bars
(FA 6), to medium- and fine-grained UFR tidal
flats (FA 5), and into fine- to very fine-grained and
mud-drape-rich low- and high-sinuosity tidal
channels (FA 4 and 3). These grain-size trends
indicate bedload convergence in the high-sinuosity tidal channel setting (Fig. 20). Along with this
central bedload-transport convergence, large
quantities of suspended load fines also accumulated in the form of ubiquitous mud drapes.
Alternative depocentres for the suspended load
are the inter- to supra-tidal marginal tidal flats
and marshes (FA 7; Fig. 20), where significant
volumes of mud also accumulated (Figs 9–11).
Palaeocurrent indicators show a dominant
south-eastward (seaward) directed current in the
tidally influenced fluvial channels (FA 2), currents of equal strength in the high- and lowsinuosity tidal channel (FA 3 and 4), and a
dominantly north-westward (landward) directed
current on the UFR tidal flats (FA 5) and in tidal
sand bars (FA 6; Fig. 20).
The described facies transitions, together with
grain-size and palaeocurrent trends indicate
deposition in tide-dominated estuaries (Dalrymple et al., 1992). Modern examples of tidedominated estuaries include Cobequid Bay and
the Salmon River (Dalrymple & Zaitlin, 1989;
Dalrymple et al., 1990, 1991), the Severn River,
England (Hamilton, 1979; Harris & Collins, 1985),
and the South Alligator River, northern Australia
(Woodroffe et al., 1989, 1993). Most modern tidedominated estuaries are macrotidal, although
tidal-dominance can occur at much smaller tidal
ranges if wave action is limited or the tidal prism
is large (Hayes, 1979; Davis & Hayes, 1984). The
presence of UFR tidal flats (FA 5), however
strongly suggests macrotidal (i.e. tidal range is
higher than 4 m, Davies 1964) conditions (Dalrymple et al., 1992) in the Spitsbergen estuaries.
Tide-dominated estuaries are characteristically
comprised of: (1) river-dominated headward
channels (FA 1 and 2 in this paper), (2) a mixed
energy middle portion within a tidal channel (FA
3 and 4 in this paper), (3) a tide-dominated outer
portion with UFR tidal flats and tidal sand bars
(FA 5 and 6 in this paper), and (4) marginal tidal
flats and marshes (FA 7 in this paper).
The tidally influenced fluvial deposits (FA 2)
indicate that rivers became tidally influenced
below the mean high tide, but the net sediment
transport was still seawards due to the long-term
dominance of river flow over tidal currents. This
portion is in a setting analogous to the ‘inner
straight’ portions of tidal channels in Dalrymple
et al. (1992); Fig. 20). The sediment is further
transported into the meandering portion of the
tidal channel (FA 3), where the highest frequency
of mud drapes together with high sinuosity
indicate that in this location the total energy
minimum occurs, and tidal and fluvial energy is
approximately equal (Ashley & Renwick, 1983;
Dalrymple & Zaitlin, 1989; Woodroffe et al., 1989,
1993; Dalrymple et al., 1992). This setting is
similar to the ‘meandering’ portion of a tidal
channel of Dalrymple et al. (1992), where deposition of fine-grained populations of both fluvialand marine-derived sediment occurs. Basinwards
from the meandering tidal channel, in the lowsinuosity tidal channel (FA 4) the tidal energy is
slightly higher, as witnessed by the lower frequency of mud drapes. This setting is similar to
the ‘outer straight’ portion of the tidal channel in
Dalrymple et al. (1992). Further seawards, the
tidal channel broadens into a funnel, and marinederived sediments are deposited in UFR tidal flats
(FA 5) and in tidal sand bars (FA 6). The UFR
tidal flats (FA 5) indicate that the flood-tidal
energy maximum lies landward from the tidal
sand bars (FA 6) (Harris, 1988; Dalrymple &
Zaitlin, 1989; Dalrymple et al., 1990, 1992). The
tidal maximum occurs in this headward portion
of the estuarine funnel, because the incoming tide
is progressively funnelled into a smaller crosssectional area (Myrick & Leopold, 1963; Wright
et al., 1975; ‘hypersynchronous’ behaviour of
Nichols & Biggs, 1985). The mixed- to muddytidal flats and marshes (FA 7) reflect gradually
decreasing tidal energy and gradually shallowing
depositional environment away from the axis of
the estuary (Fig. 20).
COASTAL PLAIN RESPONSE TO
SEA-LEVEL CHANGES
The Eocene Central Basin infill is overall progradational and consists of clinoforms that migrated
to the south-east (Fig. 1). The coastal-plain facies
belt, however, is very aggradational and mudprone (Fig. 2), except for the fluvial channels
2005 International Association of Sedimentologists, Sedimentology, 52, 391–428
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419
Fig. 21. Basal erosion surfaces (dashed lines), overlain by fluvial deposits, cut more than 16 m (A) into tidal deposits
of previous sequences. Lateral to fluvial incisions, palaeosols occur (B). Fluvial deposits that overlie outer estuarine
tidal sand bar deposits (C), or tidal flat deposits (D) mark a significant seaward shift of depositional facies.
associated with regional erosion surfaces. These
erosion surfaces that cut across the whole coastal
plain (Fig. 4) are used to divide the coastal-plain
succession into a series of depositional sequences
that are 7–44 m thick. Eighteen depositional
sequences were studied along the Van Keulenfjorden transect on two different mountainsides
over a distance of some 15–20 km. The coastalplain succession on Brogniartfjellet (sequences
12–19; Figs 9 and 10) is broadly time-equivalent
with shelf to slope clinothems on Storvola and
basin-floor clinothems on Hyrnestabben. The
coastal-plain succession on Storvola (sequences
20–29; Fig. 11) is broadly time equivalent with
shelf to slope clinothems on Hyrnestabben
(Fig. 1).
Each of the stratigraphic sequences consists of
(1) lowstand deposits (FA 1 and 2) just above the
major erosion surface, (2) transgressive deposits
with landward-stepping, estuarine deposits (FA
1–7), and (3) highstand deposits with aggradational to seaward-stepping, estuarine deposits
(FA 1–4, 7). The third segment has been documented in only about 30% of the sequences
(Figs 9–11).
Sequence boundaries
The basal erosion surface of each sequence erodes
into older estuarine deposits with local erosional
relief of up to 16 m (Fig. 21A). In places, where
the erosion surfaces are exposed in a more
oblique view, abundant rooted horizons and
horizons intensively reworked by pedogenic processes (e.g. ferruginous features, very intensive
bioturbation by roots and loss of all sedimentary
structures, Fig. 19B) occur lateral to the deepest
incisions (sequence 22 in Fig. 11; Fig. 21B). The
full lateral extent of the erosion surfaces is
unknown.
Interpretation
The prominent and widespread erosion surfaces
that are overlain by coarse-grained fluvial channel
deposits, are interpreted as sequence boundaries,
because (1) these erosion surfaces are traceable in
a dip-direction across the whole coastal-plain
facies belt, (2) the depth of incision is significant
(local erosional relief up to 16 m), and (3) these
erosion surfaces are overlain by fluvial deposits
that mark a significant seaward facies shift
(Fig. 21C and D). Sequence boundaries formed
during the fall of relative sea-level to its lowest
position. Extensive rooting and pedogenesis
occurred in time-equivalent interfluve segments
adjacent to the deeper incisions.
Lowstand deposits
The most prominent, more-or-less continuous
sandstone levels that stand out on the photomosaics (Fig. 21A) are the levels where fluvial (FA 1)
and tidally influenced fluvial (FA 2) channels
occur across the entire coastal plain (Figs 9–11).
The fluvial channels (FA 1) grade seawards and
upwards into tidally influenced fluvial deposits
(FA 2, Figs 7 and 9–11). The fluvial units are in
some sequences up to 16 m thick, but in most
sequences 3–5 m thick.
Fine-grained overbank deposits adjacent to the
fluvial channels are documented only in a few
places (sequence 12 in Fig. 9, sequence 17 in
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420
P. Plink-Björklund
Fig. 10, sequences 20, 27 and 28 in Fig. 11), and
are characterized by thin-bedded, ripple-laminated very-fine- to fine-grained sandstones, and
plane-parallel-laminated sandstones and mudstones.
Rooted horizons and coals commonly occur at
the tops of fluvial channel fills. In strike-oriented
outcrops, multiple soils and root horizons occur
lateral to the channel margins (Fig. 21B). Most of
the documented rooted horizons in the coastalplain succession are limited to the lowstand
deposits, especially where the fluvial channels
grade upwards into tidally influenced fluvial
channels.
At the base of some sequences, there is a single
fluvial or tidally influenced fluvial channel
(sequences 12, 13, 14 and 15 in Fig. 9; sequences
13, 14, 15 and 18 in Fig. 10; and sequences 21, 25,
26, 27 and 28 in Fig. 11). In other sequences,
multiple vertically stacked fluvial or tidally
influenced fluvial channels occur, separated by
an aggradational muddy interval or flooding
surface (sequence 16 in Fig. 10; and sequences
20, 22, 23, 24 and 28 in Fig. 11).
Fig. 22. During lowstand fluvial deposits accumulated
on the coastal plain (A), in incised valleys eroded
during sea-level falls. River channels were tidally
influenced in their seaward ends. Rising relative sealevel drowned the valleys and landward-stepping
estuarine deposits accumulated (B and C). The transgressive estuarine deposits young landwards as they
onlap. During HST the inner parts of the valleys started
filling ‘in situ’ (D), and marshes developed over larger
areas, as rate of sea-level rise decreased. Gradually the
inner estuarine deposits shifted seawards, and the
youngest HST deposits covered the oldest transgressive
deposits.
Interpretation
The fluvial and tidally influenced fluvial channels reflect deposition in rivers that became
tidally influenced at their the river mouths, below
the mean high tide (Fig. 22A). The successions of
coarse-grained fluvial and tidally influenced fluvial deposits that shift abruptly seawards across
older inner- and outer-estuarine deposits of previous sequences are interpreted as lowstand
deposits. The upwards transition from fluvial to
tidally influenced fluvial deposits indicates
deposition during the initial relative sea-level
rise. Associated rooted horizons and pedogenically reworked levels suggest that areas between
the active channel belts, underwent prolonged
subaerial exposure. (Figs 9–11).
In those sequences that show vertically stacked
fluvial or tidally influenced fluvial channels,
separated by mud-prone flooding intervals, the
basal channel fill is represented as early lowstand
fill, whereas the upper channel deposits belong to
late lowstand. This kind of intra-lowstand flooding surface has been described also from contemporaneous shallow- to deep-marine portions of
the Eocene clinoforms. Such intra-lowstand
flooding is hitherto not widely recognized, but it
has been suggested that it marks the sea-level rise
back above the shelf-edge (Plink-Björklund &
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421
Fig. 23. Estuarine deposits of the landward-stepping facies onlap (white arrow) the basal fluvial deposits in a
landwards direction (A). In some sequences the landward-stepping deposits have a basal by tidal ravinement surface
that cuts through the underlying estuarine deposits (B and C). In places the tidal ravinement surface merges with the
sequence boundary. In other sequences, and dominantly in more landward reaches of the sequences, the tidal
ravinement surface occurs higher up in the sequence, and inner-estuarine or marginal facies occur above the fluvial
deposits and below the outer-estuarine deposits (D).
Steel, 2005). The rising sea-level would have
drowned incised river mouths or distributary
channels (at the shelf edge), and caused temporary sand storage in the fluvial system, until the
fluvial systems aggraded, and deltaic systems
prograded back to the shelf edge again to form
the late lowstand deltas (Plink-Björklund & Steel,
2005). Similar drowning of river mouths, that
causes temporary sand storage within the fluvial
system, has been described from modern rivers.
Transgressive surface
At the top of the fluvial channel deposits there is
a marked transgressive surface that signifies
valley drowning and the landward migration of
the bayline, separating the fluvial lowstand
deposits from the estuarine transgressive deposits. The transgressive surface separates fluvial
aggradation from landward-stepping estuarine
deposition in seaward portions of the coastal
plain. On the more landward reaches of the
coastal plain, fluvial aggradation may continue
across the transgressive surface. In these cases the
transgressive surface separates amalgamated,
lowstand fluvial deposits from more aggradational fluvial deposits, interbedded with floodplain, marsh or tidal flat deposits (sequences 13,
14 in Fig. 9; sequences 25, 27, 28 in Fig. 11).
Transgressive deposits
The fluvial channels are generally overlain by
tide-dominated estuarine deposits (Figs 9–11).
These tidally dominated estuarine deposits have
a landward-stepping character as they can be seen
to onlap the basal fluvial deposits in a landward
direction (Fig. 23A; sequence 16 in Fig. 10). At
the landward ends of the exposed estuaries (i.e.
inner estuarine segments), the landward-stepping
is implied by a landward shift of the tidally
influenced fluvial channels, and high- and
low-sinuosity tidal channels, i.e. the tidally
influenced fluvial channels are replaced by the
high-sinuosity tidal channels, and the high-sinuosity tidal channels are replaced by the lowsinuosity tidal channels at gradually higher
stratigraphic levels (Figs 9–11). At the seaward
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422
P. Plink-Björklund
end of the exposed estuaries (outer estuarine
segment), the landward stepping is implied by
the landward shift of UFR tidal flats and tidal
sand bars.
The tidal sand-bar deposits rest on highly
erosional basal tidal ravinement surfaces that in
some cases cut through the whole underlying
estuarine and fluvial deposits, and merge with the
sequence-bounding erosion surface (sequence 13
in Fig. 10, and sequence 21 in Figs 11, 23B and
C). In most sequences the tidal ravinement surfaces occur within the landward-stepping estuarine
deposits, some distance above the basal fluvial
segment (Fig. 23D). In some sequences the tidal
ravinement occurs on several stratigraphic levels
(sequence 20 in Fig. 11). In most cases, the tops of
the tidal sand bars mark the landward- to seaward-stepping turnaround, i.e. the most landward position of outer-estuarine deposits.
The transgressive deposits volumetrically dominate the depositional sequences. In contrast to
the lower, fluvial segments, the middle transgressive segments are mud-prone, and the only thick
and prominent sands are those that occur as tidal
sand bars (Fig. 17). The high- and low-sinuosity
tidal channels are also sandy, but the sands are
rather thin (0Æ6–1Æ5 m thick) and discontinuous.
The majority of the mudstones were deposited
on marginal tidal flats and marshes. The fluvial
overbank deposits grade seawards and upwards
into tidal sand to mud flats. This gradual vertical
transition from unipolar rippled and plane parallel-laminated beds into more rhythmically bedded ripple-laminated beds with mud drapes, and
finally into bipolar lenticular, wavy or flaser
bedded units is especially well seen in sequence
28 on Storvola (Fig. 11). Depositional units of the
marginal sand to mud flats and marshes have a
coarsening-up character in the transgressive segment, as gradually higher energy parts of the
marginal tidal flats cover the lower-energy areas.
Tidal-flat deposits in seaward portions of the
estuaries (outer estuarine segment) tend to be
more bioturbated that those in more landward
portions (inner estuarine segment). The latter are
more rich in coal layers, rooted horizons, and
plant debris. Coal layers are more abundant in the
lowest parts of the landward-stepping segments.
Interpretation
The landward-stepping estuarine succession that
accumulated above the lowstand deposits is
interpreted as transgressive estuarine deposits.
A relative sea-level rise is indicated by a landward shift of all depositional environments, a
landward onlap, and the development of tidal
ravinement surfaces (Fig. 22B and C).
Maximum flooding surface
The maximum flooding surface is recognized by
the turnaround from landward-stepping to seaward-stepping successions in the seaward reaches of the coastal plain. In sequences where the
outer estuarine tidal sand bars are present, the
maximum flooding surface coincides with the top
of the tidal bar deposits (sequences 13, 14, 15, 16
and 18 in Fig. 10; sequences 20, 21 and 24 in
Fig. 11). In the landward reaches of the coastal
plain, the highstand deposits are missing or the
maximum flooding surfaces have not been recognized.
Highstand deposits
The highstand deposits, where present (sequences 13, 14, 15, 16 and 18 in Fig. 10,
sequences 20, 21 and 24 in Fig. 11), comprise
sedimentary facies similar to those in the underlying transgressive deposits. The highstand segment differs from the transgressive segment by (1)
a seaward-shift of inner-estuarine facies, (2) more
extensive root horizons and coal layers, and (3) a
higher proportion of marsh deposits.
The seaward shift is mainly indicated by
replacement of outer-estuarine facies with innerestuarine facies (sequences 15 and 18 in Fig. 10,
sequences 20 and 21 in Fig. 11), or low-sinuosity
tidal channels by high-sinuosity tidal channels
(sequences 16 and 18 in Fig. 10, sequence 20 in
Fig. 11) in successive vertically stacked deposits.
The coaly mudstones and coal layers of marshes
are volumetrically more significant in the seaward-stepping segments compared to the landward-stepping segments. Coal layers become
more abundant towards the top of the seawardstepping segments.
The presence of high-sinuosity tidal channels,
where fluvial-derived and marine-derived deposits converge, indicates continued bedload convergence into the inner estuary, and strongly
indicates that deposition still occurred in an
estuarine environment, rather than on a tidedominated delta plain. Dalrymple et al. (1992)
argued that the presence of net landward movement of sediment derived from outside the estuary mouth (averaged over a period of several
years) is one of the primary features that distinguish estuaries from delta distributaries, where
the net sediment transport is seawards. Similar
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Stacked fluvial and tide-dominated estuarine deposits in high-frequency sequences
seaward-stepping of facies in tide-dominated
estuaries, with the relative distribution of facies
remaining essentially constant has been documented from the South Alligator River estuary
(Woodroffe et al., 1989, 1993; see also Harris,
1988) and the Gironde Estuary (Allen & Posamentier, 1993) since the end of the Holocene
transgression. In a tidally influenced delta, sands
fine seawards unidirectionally, in contrast to the
estuaries where sands are coarsest at the head and
mouth (see Harris, 1988; Dalrymple et al., 1991).
Interpretation
The seaward-stepping estuarine successions indicate decreasing rates of sea-level rise, and are
assigned to the highstand systems tract
(Fig. 22D). The higher abundance of coals and
coaly mudstones together with the seaward-shift
of inner-estuarine facies strongly suggest that
accommodation in the inner portions of the
estuaries was infilled, and marshes developed
over larger areas. Similar regressive phases of
estuary development in the Gironde Estuary
(Allen & Posamentier, 1993), and the South
Alligator River estuary (Woodroffe et al., 1989,
1993; see also Harris, 1988) occurred during
highstand of sea-level. Highstand estuaries were
also predicted from conceptual models (Dalrymple et al., 1992).
INCISED VALLEYS?
The estimation of the total depth of erosion, i.e.
eventual incised valley depth, is not completely
clear, because the outcrops are dominantly dipparallel. The database, however, strongly suggests
423
that the estuarine deposits backfilled incised
valleys, eroded during the preceding sea-level
falls. The criteria for recognizing incised valleys
outlined by Zaitlin et al. (1994) can be recognized
here: (i) the basal erosion surfaces are regional
incisions; (ii) the basal fluvial deposits exhibit a
significant basinward facies shift; (iii) the base of
the incised valleys can be correlated with rooted
horizons in the interfluve areas; and (iv) the
estuarine infills onlap landwards the valley walls.
Moreover, tide-dominated estuaries require confinement in a narrow, funnel-shaped geometry
(see Dalrymple et al., 1992; Zaitlin et al., 1994),
suggesting that the whole thickness of the individual depositional sequences may have been
confined within the valleys. This gives an
approximate estimate for the minimal valley
depth for individual sequences, assuming that
the valley fills were simple, i.e. each valley was
filled completely during one lowstand-transgressive-highstand sequence (see Rahmani, 1988;
Wood & Hopkins, 1989; Zaitlin et al., 1994).
The thickness of the depositional sequences
varies from 7 to 44 m. The database also shows
that the thickest sequences tend to be associated
with sequence boundaries with the most local
erosional relief (Fig. 24). Correlation with shallow- and deep-marine portions of the same
clinoforms shows that such thick coastal-plain
sequences with high local erosional relief produced coeval slope or basin-floor sandstone bodies (e.g. sequence 14). Thin sequences with little
basal erosional relief, on the other hand, generally
did not produce sands beyond the shelf edge.
This implies that the incised valleys of the
Eocene Central Basin coastal-plain were at least 7
to 44 m deep, as it is not known if the valleys
Fig. 24. The thickest sequences
tend to be associated with the sequence boundaries that have most
local erosional relief. Such coastalplain sequences (e.g. sequence 14)
were more likely to produce during
falling stage and lowstand coeval
slope or basin-floor sandstone bodies.
2005 International Association of Sedimentologists, Sedimentology, 52, 391–428
424
P. Plink-Björklund
were filled completely by each of the depositional
sequences. The sequences that have highstand
systems tracts, were obviously filled at least in
their landward reaches, whereas the sequences
that lack highstand deposits may not have been
filled during a single sea-level fall-to-rise cycle,
unless their highstand portions were eroded
during following sea-level fall, or the highstand
deposits are located further seawards (see below).
SEDIMENT PARTITIONING ON THE
COASTAL PLAIN
Falling-stage deposits have not been clearly
identified in the coastal-plain deposits of the
Aspelintoppen Formation; although they could
be locally present but difficult to differentiate
from lowstand deposits. The fluvial deposits are
assigned to the lowstand, rather than to the falling
stage, because they become characteristically
tidally influenced upwards, indicating a baselevel rise and bayline migration up the valley.
Assigning the amalgamated fluvial channel-fill
deposits to the lowstand systems tract is consistent with existing stratigraphic practice (e.g.
Shanley & McCabe, 1993; Zaitlin et al., 1994;
Olsen et al., 1995), although the recognition of
the intra-lowstand flooding in some of the
sequences is unusual. The lowstand was the
main phase of fluvial deposition in the Eocene
Central Basin. As the bayline moved significantly
up the valley during the early transgressive time,
the clearly identifiable transgressive surface
formed that separates fluvial deposits from overlying estuarine deposits (Fig. 22A). However, in
landward reaches of the coastal plain, the transgressive surface occurs within the fluvial succession, separating amalgamated channels below
from more aggradational channels above (see also
Shanley & McCabe, 1993; Olsen et al., 1995; Plint
et al., 2001), as seen in sequence 14 (Figs 9 and
10), where early transgressive tidally influenced
fluvial deposits in Fig. 9 correlate into estuarine
deposits in Fig. 10. By the time of the late
transgressive systems tract, the estuaries had
migrated further up-valley, leaving behind a
landwards onlapping and landwards younging
transgressive succession (Fig. 22C). This implies
that the oldest transgressive estuarine deposits
are found close to the incised valley mouth,
whereas in the landward reaches of an incised
valley the estuarine deposits belong to the late
transgressive systems tract. The transgressive
systems tract was the main phase of coastal-plain
aggradation in the Eocene Central Basin of Spitsbergen. During early stages of highstand time,
when sea-level rise rate decreased, the estuaries
stopped migrating landward, and started filling
in situ (see Dalrymple et al., 1992). The earliest
highstand deposits are found in the inner parts of
the incised valleys, and the highstand deposits
gradually prograded seawards (Fig. 22D). Highstand deposits, where present, are documented
only in the outer-estuarine reaches of the
sequences, as inner estuarine reaches of sequences 13 and 14 in Fig. 9 consist of lowstand
and transgressive tracts, whereas outer-estuarine
reaches of the same sequences in Fig. 10 also
contain highstand deposits. This probably reflects
a lack of accommodation in the inner-estuarine
reaches (see Plint et al., 2001). The maximum
flooding surface is easy to identify in outer
estuarine reaches, where it separates early transgressive deposits from highstand deposits and
thus the seaward facies shift is clear, as opposed
to the inner estuarine reaches, where the maximum flooding surface separates the latest transgressive deposits from the earliest highstand
deposits.
CORRELATION TO MARINE SEQUENCES
The detailed correlation of coastal-plain sequences
to their shallow- and deep-marine counterparts is
not the objective of this paper. However, it should
be mentioned that the depocentre shifted along the
coastal-plain to deepwater clinoforms during
the relative sea-level cycles. As explained above,
the falling-stage deposits are insignificant on the
coastal plain, whereas on the contemporary shelf
and slope falling stage deltas accumulated, as the
sediment was fed through the incising valleys.
Lowstand deposits were also mainly partitioned
beyond the shelf edge in lowstand deltas, although
aggradation of fluvial deposits within incised
valleys also occurred. The intra-lowstand flooding
surface is traceable across the shelf edge and onto
the slope, where it separates the falling-stage and
early lowstand deposits from the prograding late
lowstand wedge (Plink-Björklund & Steel, 2005).
During the transgression, however, the depocentre
was on the coastal plain and inner shelf, as
sediments were fed into the estuaries both from
the fluvial and marine sources. Earliest highstand
deposits were partitioned on the coastal plain, but
the depocentre moved out onto the shelf, as
highstand deltas occur within the shelf segments
of some of the clinoforms. These shelf deltas are
2005 International Association of Sedimentologists, Sedimentology, 52, 391–428
Stacked fluvial and tide-dominated estuarine deposits in high-frequency sequences
slightly younger than the highstand deposits in the
described estuaries.
CONCLUSIONS
Eighteen, fourth-order coastal-plain depositional
sequences, 7–44 m thick, were documented in the
Eocene Central Basin of Spitsbergen. The depositional sequences consist of basal fluvial deposits,
covered by tide-dominated estuarine deposits.
Coastal-plain aggradation occurred due to estuarine infilling of incised valleys, cut during relative sea-level falls.
The incised valleys were filled with: (1) lowstand fluvial deposits, (2) transgressive estuarine
deposits, and (3) highstand estuarine deposits. The
transgressive deposits volumetrically dominate
the coastal-plain succession, as transgressive times
were the main phase of coastal plain aggradation.
The transgressive estuarine deposits are oldest in
seaward reaches of the incised valleys, and young
and onlap landwards. The highstand deposits,
preserved on the coastal plain, are interpreted as
estuarine rather than deltaic, because of continued
bedload convergence in the inner-estuarine environment. Highstand deltas developed further seawards, and were slightly younger compared to the
coastal-plain estuaries.
ACKNOWLEDGEMENTS
Conoco, Mobil, Norsk Hydro, PDVSA, Phillips,
Shell, Statoil and UPRC have financed this work
as a part of the Wyoming Consortium on Linkage
of Facies Tracts (WOLF). The Swedish Foundation for International Cooperation in Research
and Higher Education (STINT) provided a postdoctoral research stipend for Piret Plink-Björklund. Brian Zaitlin and Robert Brenner reviewed
an earlier version of the manuscript and are
gratefully acknowledged.
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