Faint traces of high Arctic glaciations: an early Holocene ice-front
fluctuation in Bolterdalen, Svalbard
IDA LØNNE
Lønne, I. 2005 (August): Faint traces of high Arctic glaciations: an early Holocene ice-front fluctuation in
Bolterdalen, Svalbard. Boreas, Vol. 34, pp. 308–323. Oslo. ISSN 0300-9483.
Raised marine beach gravel at 62 m a.s.l. in Bolterdalen indicates that the inner part of Adventfjorden, central
Spitsbergen, was ice-free shortly before 10 025+160 yr BP. A glacier advanced across the regressive, frozen
beach terraces and into shallow water, 58 m above the present sea level, where a small wave-influenced icecontact delta was formed, 9775+125 yr BP. Maximum ice-front position was reached 9625+95 yr BP, 7 km
outside the present ice margin. The advance was climatically forced and of several decades’ duration, as seen
from abundant molluscs growing in the prograding foreset beds. Today, the beaches appear as a continuous
regressive sequence with no geomorphic evidence of the former ice margin. Sedimentological studies show,
however, that a thin (1 m) deformation till was emplaced, the substrate was subglacially sheared to a depth of
1 m, and elongated clasts in the beach gravel were reoriented in an ice-flow parallel direction. The glacial
deposits and structures, formed within 200 m from the ice front, highlight some important aspects of subglacial
to ice-marginal processes in permafrost terrain. As the dead ice melted, the released debris was redistributed into
thin sediment sheets down to 40 m a.s.l., which means that the postglacial meltwater-controlled reworking lasted
c. 500 years. Similar isolated depocentra may be a key for future identifications of former ice margins at high
latitudes.
Ida Lønne (e-mail: [email protected]), Villaveien 21, NO-1440 Drøbak, Norway; received 22nd October 2004,
accepted 10th March 2005.
Glaciers in polar regions have a lower impact on
geomorphic activity compared to ice masses at midlatitudes. Erosional and depositional rates are lower,
and depositional landforms and sediment sheets are
thinner and have a lower preservation potential during
the deglaciation. Studies along modern retreating ice
margins on Svalbard, 76–81 N (Bennett et al. 1996;
Etzelmüller et al. 1996; Glasser et al. 1999; Lyså &
Lønne 2001), show a deglaciation dynamics strongly
imposed by the polar night conditions, which includes
factors such as continuous permafrost, low and strongly
seasonal solar radiation, low sediment transfer and
slow but efficient de-icing of ice-cored moraines.
Glacial debris is redistributed in a step-wise fashion,
forming a landscape predominated by fluvial and
colluvial morphologies, where traces of former ice
margins are few and disputable (Lønne & Lyså in
press).
The submarine deposition is affected by these
extreme conditions but to a smaller extent, and the
onshore versus offshore reconstructions of glacial
history have therefore been contradictory (Landvik
et al. 1998, 2005). Although ice-contact deposits and
till sheets below sea level are thinner (e.g. Sexton et al.
1992; Whittington et al. 1997) than in areas with
comparable relief further south, marine sediments are
less susceptible to postglacial reworking compared to
onshore debris. Ancient glacier activities on Svalbard
are therefore chiefly based on the marine record. The
areas between the marine limit and present sea level
play a key role; it is here that most of the onshore
stratigraphy and chronology have been obtained, and
where marine processes can be studied in detail.
Another factor that complicates the reconstruction of
former ice-front movements in this region is the high
frequency of glacial surges. A surge-type glacier is an
ice mass with a cyclic flow regime that is unrelated to
climatic forcing (Meier & Post 1969). Although estimated incidence of surge-type glaciers on Svalbard
varies (90% in Lefauconnier & Hagen 1991, 36% in
Hamilton & Dowdeswell 1996 and 13% in Jiskoot et al.
1998), such cyclic behaviour most certainly affected
glaciers also in the past. The two major challenges
with respect to the mapping of former glacier activity on
Svalbard are related to the recognition of ice margins in
the field, and the identification of the ice-front positions
that were related to surge advances, which should be
avoided in palaeoclimatic reconstructions.
This article examines field observations from
Bolterdalen, central Spitsbergen (Fig. 1), where an
apparently continuous sequence of marine beach
terraces descends from 62 m to 20 m a.s.l. However, a
small outcrop section, 270 m long and 0–5 m thick,
shows that the beach progradation was interrupted by a
glacier that advanced into the fjord and deposited a
small, wave-influenced, ice-contact delta and retreated
while the glacioisostatic uplift proceeded. Sediments left
by the terrestrial part of the glacier were redistributed
DOI: 10.1080/03009480510012971 # 2005 Taylor & Francis
BOREAS 34 (2005)
Glacier fluctuation in Bolterdalen, Svalbard
309
Fig. 1. Location of the study area in Bolterdalen, central Spitsbergen. Encircled numbers are the local marine limits (altitude in metres above
sea level).
during the deglaciation, and today this ice margin can
hardly be identified. Reconstruction of the syn- and postsedimentary processes provides new insight into the
factors that control formation and decay of ice-marginal
deposits and glacial landforms at high latitudes.
Setting
The marine limit (ML) at the mouth of Adventfjorden
is an abrasion platform at c. 70 m a.s.l. (Fig. 1) partly
covered by scree deposits and solifluction. The platform descends eastwards into Adventdalen, where it
becomes gradually more diffuse due to lower wave
energy and larger input of debris along the valley
slopes, and is at c. 65 m a.s.l. at Isdammen, 9 km east of
the fjord mouth. Soltvedt (2000) found a ML at 63 m
a.s.l. at the mouth of Endalen.
Bolterdalen is a 13-km-long, north–south trending
valley, where the inner part is occupied by Scott
Turnerbreen (Fig. 1). Ice-cored moraines from the
Little Ice Age (LIA) maximum, at the turn of the 19th
century, form a concentric zone along the ice front
(Lønne & Lauritsen 1996; Sletten et al. 2001). The
valley floor is occupied by a braided river and colluvial
debris covers the valley walls, leaving no evidence of
former ice margins outside the LIA maximum.
A small hill at the mouth of Bolterdalen is c. 300 m
broad, 600 m long and reaches 110 m a.s.l. The bedrock
surface is inclined towards the northwest and terminates in a 10-m-high cliff (Fig. 2). A series of narrow,
elongate, strike-parallel (NW–SE) bedrock ridges of
Cretaceous shale (Major et al. 2000) are draped by
matrix-supported diamicton. Scattered blocks, predominantly subrounded sandstones with a spherical
shape, occur on the surface down to 62 m, where
marine terraces descend to the present fluvial plain
(Fig. 3).
Description of the outcrop section
0
The sediments, exposed in a 270-m-long outcrop (A–A
in Fig. 2), are disturbed from 0–70 m (referring to the
horizontal scale in Fig. 4), but well preserved from
70–270 m. Five sedimentary units are distinguished.
310
Ida Lønne
BOREAS 34 (2005)
0
0
0
0
Fig. 2. The study area in Bolterdalen with location of the outcrop section A–A and the topographic profiles B–B , C–C and D–D . Shown
are clast fabric analyses from the upper till (unit 4) and beach gravel (unit 2), and the average orientation of normal faults in the foreset beds
of unit 3.
BOREAS 34 (2005)
Glacier fluctuation in Bolterdalen, Svalbard
311
Fig. 3. Surface morphology and sediment structures are shown by the aerial photograph (excerpt from S90-1241, Norwegian Polar Institute).
0
Arrows indicate the direction of sediment reworking, and A–A shows the location of the outcrop section that developed after 1991. Some of
the largest blocks on the surface are encircled. White numbers are altitudes in metres above sea level.
312
Ida Lønne
BOREAS 34 (2005)
0
Fig. 4. Generalized stratigraphy along profile A–A (west side of the ditch; see Fig. 2) with the five sedimentary units and location of four
radiocarbon-dated samples (ages are given in yr BP). Shown are clast fabric analyses from the upper till (unit 4) and beach gravel (unit 2),
and the average orientation of normal faults in the foreset beds of unit 3.
Unit 1 – clast-rich diamicton
Unit 2 – upwards coarsening foreset succession
Unit 1 is a firm, clast-rich diamicton (up to 50 cm thick)
above the bedrock (Fig. 4), with an upper contact to
clayey mud that varies from sharp to gradational
(Fig. 5). The unsorted, polymodal texture has a clayey
matrix and the coarse fraction is fine pebbles to blocks,
but little sand. The clasts, commonly striated, are
subangular to well rounded, and the shape varies from
spherical to disc- and iron-shaped. The thin, poorly
exposed and still frozen diamicton is difficult to
examine, but the lag along the ditch provides some
information on its clast composition. Dark shales and
lighter fine-grained sandstones from the Cretaceous and
Tertiary substrate in Bolterdalen dominate, although
the presence of exotic clasts is striking: well-sorted
reddish sandstone, quartzite, quartzitic conglomerate,
reddish gneiss and reddish granite. None of these
lithologies are exposed in Bolterdalen (Major et al.
2000), and must therefore have been brought here by an
active glacier.
The diamicton is overlain by clast-rich mud that coarsens upwards to sandy foreset beds with well-sorted
gravel on top, reaching a thickness of 3.2 m and a
lateral extent of 80 m (Fig. 4).
The mud is finely laminated, but thin interbedded
debrisflow deposits, slump structures, intraclasts and
lenses of seaweed (Fig. 5) suggest highly variable
depositional conditions. Scattered, isolated clasts
throughout the unit vary from well rounded to subangular. Shales and light-coloured sandstones from the
local bedrock dominate, but exotic clasts also occur. A
few individuals of the mollusc Mya truncata are found,
but no foraminifera.
The sandy foreset beds, dipping at 14 , are well
sorted, and stratified sand and sandy gravel alternate
with wave-worked debris with scattered pebbles and
granules, separated by nearly planar erosional surfaces
characteristic for shore-face facies (Komar 1976;
Massari & Parea 1988). The gravelly foreset shows
BOREAS 34 (2005)
Glacier fluctuation in Bolterdalen, Svalbard
313
Unit 3 – delta
Fig. 5. Sedimentary log 1 (for location, see Fig. 4).
upwards transition from spherical to disc-shaped clasts
with rolling fabric (a(t)b(i); Fig. 6A, B), interpreted as
pebbly foreshore facies, with a transition to welldrained beach gravel (Lønne & Nemec 2004). The
upper, well-sorted pebbly beach is up to 1 m thick and
descends from 62 m to 58 m. A radiocarbon date
obtained from a Mya truncata in the mud yielded 10
025+160 yr BP (T-13882; Fig. 4, Table 1).
Unit 2 is a regressive beach succession deposited on
a north-dipping bedrock surface with a thin discontinuous cover of glacial diamicton at a water depth
15 m. The beach was prograded towards the northeast
during a fall in sea level from 62 to 58 m.
Unit 3 is a 3 to 5-m-thick and 120 m long package of
northwest-dipping foreset beds, coarsening upwards
from mud to gravel (Fig. 4). This succession is
distinguished from the foreset of unit 2 by a conspicuous change in facies at 140 m: the sandy part thickens
to 3 m and contains a large number of Mya truncata in
growth position. The gravelly top becomes gradually
more unsorted, with an increased grain size variation
and more spherical clasts (cf. Figs 6C, 7A).
The basal muddy part alternates between laminated
facies, mudflows and gravelly debrisflow deposits, and
has a gradual upper contact to the sandy foreset (Fig. 4).
Poorly sorted debrisflow beds, some with diamictic
composition, are intercalated from 180 m, where also
outsized clasts and sporadic diamictic intraclasts increase in number. Some rod-shaped particles in vertical
position have clearly been dropped from floating ice.
The upper gravel pinches out at 180 m and shows
scour-fills and a few high density turbidites (Fig. 6D).
The base of the topset is fairly subhorizontal in the
southern part, typical for traditional Gilbert-type deltas
(Nemec et al. 1999; Postma 1990). The more
pronounced interfingering towards the north indicates
that the stream-transported debris was influenced by
waves, which is also supported by the sigmoidal foreset
geometry along the east side of the ditch (Fig. 7D). The
grain size variation increases from c. 180 m, and wave
sorted sandy facies with local scour-fills are here
intercalated with thin gravelly turbidites and debrisfall
gravel (Fig. 6D). A short wedge-shaped accumulation
of coarse-grained debris, onlapping the upper delta
front, is interpreted as a mouth bar deposited in a
period with high fluvial sediment input. A considerable
sedimentation rate is shown by a 60-cm-long escape
trace formed by a Mya truncata (Fig. 7A, B). The
depositional conditions varied seasonally, as seen from
strongly bioturbated sandy foreset beds that reflect low
sediment input and open sea (Fig. 7D), whereas the
regular occurrence of up to 10-cm-thick horizons of
clay-rich mud (Figs 6C, 7C) represent periods with icebounded fjord. Shallow scours and chute-fill backsets
(Nemec 1990; Lønne et al. 2001) are exposed on the
east side of the ditch (Fig. 7D), but are difficult to
discriminate from listric thrust faults. A profound
change in sediment architecture at c. 200 m is related to
a shift in subglacial conditions and is described below
as part of unit 4.
The large numbers of thick-walled Mya truncata in
unit 3 have an average length of 5–8 cm, with periostracum and siphons well preserved (Fig. 7E). Further
downslope, Hiatella arctica, Macoma calcarea and
Chlamys islandica are found at 150 and 240 m (Fig. 4),
respectively. Towards the north and upper part of the
succession, the number and size of Mya truncata
decline slightly to an average length of 2–3 cm. Two
radiocarbon dates yielded ages of 9775+125 yr
314
Ida Lønne
BOREAS 34 (2005)
Fig. 6. A. Lateral facies variations in unit 4 and the underlying beach gravel of unit 2 (c. 70 m in Fig. 4). B. Overlay drawing of A showing
that the deformation structures are a combination of thrust and shear deformation, and penetrate some decimetres into the substrate. Note the
alignment of elongated particles, parallel to the sense of tectonic deformation, in both units 2 and 4. C. Detail from the upper diamicton at
c. 180 m. D. Segment (just to the left of the outcrop in C) showing unsorted, coarse-grained debris dumped in front of the advancing glacier
and forming a small mouth bar. E. Well-sorted pebbly beach faces with disc-shaped clasts (see location in C) that were picked up by the
advancing glacier, transported along the glacier sole and deposited as part of the till.
Glacier fluctuation in Bolterdalen, Svalbard
BOREAS 34 (2005)
315
Table 1. Radiocarbon dates from the Bolterdalen section (for location of the samples, see Fig. 4). All dates have been corrected for a marine
reservoir age of 440 years (Mangerud & Gulliksen 1975).
Lab. ref. no.
Material
Age (14C yr BP)
Unit, facies and location
(metres along horizontal
scale in Fig. 4)
T-13882
Mya truncata
whole shell
Mya truncata
in growth position
Mya truncata
in growth position
Whalebone
10 025+160
Unit 2, mud 125 m
T-13883
T-13884
T-16768
and 9625+95 yr BP (T-13883 and T-13884; Fig. 4,
Table 1).
A >2-m-long whalebone, protruding from the frozen
delta foreset beds, was found 1.5 m below the base of
unit 4 (210 m in Figs 4, 7F). The bone’s lower surface
has a crust of sorted fine pebbles, mollusc shells,
seaweed and siphon fragments, which suggests that it
was frozen to a well-drained beach and later eroded and
transported to the delta front. The upper surface is
draped by a 5 to 6-cm-thick horizon of laminated sandy
mud with scattered pebbles (Fig. 7F) that indicate low
input of sediment and ice-covered fjord just after
deposition. The overlying succession consists of
slightly coarser-grained, sandy foreset beds, without
molluscs, deposited by rapid infilling of the slump scar
and restoration of the delta’s slope profile when ablation started in spring. A radiocarbon date of the bone
yielded 9855+155 yr BP (T-16768; Fig. 4, Table 1),
and suggests a time-lag between the age of the whale
and unit 3.
Unit 3 is a wave-influenced, ice-contact delta built
towards the northeast, and graded to a sea level of 58 m
above the present. The deltaic phase is attributed to an
advancing glacier, as the transition from underlying
beach facies is gradual and the deltaic foreset beds
become more unsorted in a downslope direction.
Clay-rich mud horizons on the upper delta front and
ice-rafted debris are related to winter sea ice conditions.
Unit 4 – diamicton
Unit 4 is 0 to 1-m-thick, exposed from 70 to 270 m
(Fig. 4), with a sharp but complex lower boundary,
eroded top and highly variable lithological and textural
composition. The diamicton contains red sandstones
and some microscopic shell fragments.
The southernmost exposure of unit 4 has a lower,
matrix-supported diamictic part, overlain by wellsorted, clast-supported pebble gravel, with disc-shaped
cobbles, originally deposited as beach facies, and
overlain by homogeneous medium sand (Fig. 6A, B).
The diamictic material must have been in subglacial
transport, while beach gravel that is only accessible up
to 62 m altitude was presumably picked up from the
9775+125
Unit 3, delta foreset, 160 m
9625+95
Unit 3, delta foreset, 250 m
9855+155
Unit 3, delta foreset, 210 m
substrate by ice-marginal thrusting and thus only
displaced a few tens of metres. The sandy portions are
interpreted as shoreface facies eroded from underlying
beach succession. Alternatively, beach facies may have
been formed also on the south side of the bedrock hill
(Fig. 1), which would increase the transport distance to
600–800 m.
The thickest part of unit 4 (Fig. 6C) shows a similar
mixture of matrix-supported diamictic debris and
sorted material, erosively above the sandy delta foreset,
and with a distinct but complex lower boundary. Small
pockets of glacially transported clasts in the otherwise
clast-poor delta-sand indicate transport in front of the
advancing glacier. The deformation shows a combination of subhorizontal shear and ice-marginal thrusting, and a strong correspondence between clast
orientation and deformation structures. In this area, the
foreset beds show no evidence of dewatering or
tectonic disturbances.
A detail from 190 m (Fig. 7A) shows diamictic
debrisflows, deposited proglacially on the delta top,
and later deformed by the overriding glacier to become
part of unit 4. A 5 to 10-cm-thick strongly sheared
horizon at the base of unit 4 is interpreted as a
glaciotectonite formed by subglacial shearing of preexisting sediments (see Benn & Evans 1996). Other
segments of the diamicton comprise matrix-supported,
spherical to iron-shaped clasts, with upflow imbrication
and well-developed orientation towards the northwest
(Fig. 2). This ice-flow direction corresponds to the
re-oriented beach gravel, normal faults and glaciotectonic thrusts.
Details from c. 200 m include dewatering on the
leeside of clasts ploughed into the fine-laminated sandy
substrate, which shows that the sandy surface was
unfrozen when overridden. The glacier must, however,
have passed over a terrestrial, frozen sediment surface,
as overturned folds along the base of the diamicton
contain inclusions of angular, laminated, sandy intraclasts that evidently were eroded in a frozen state. A
series of normal faults in the foreset beds, with a
displacement of a few centimetres, are overlain by
intensely sheared sediments with dewatering structures
(Fig. 8A). The whole succession is truncated by listric
316
Ida Lønne
BOREAS 34 (2005)
Fig. 7. A. Delta foreset and the deformation till emplaced on the unfrozen delta (at c. 190 m). The delta topset is eroded here, and the
deformations have a sharp and well-defined base. B. Pebbles aligned along a thrust fault in the delta sand. The deformation continues into the
till. C. Turbiditic sandy foreset beds interfingering with clay-rich mud horizons from periods with sea-ice conditions. D. Sigmoidal forset
beds in unit 3 that are truncated by thrust faults with almost negligible displacement. Inset shows strongly bioturbated delta front sand. E. An
8-cm-long Mya truncata from unit 3, in growth position, with periostracum and siphon well preserved. F. A whalebone in unit 3 (trowel is
20 cm long). Location is shown in Fig. 4.
BOREAS 34 (2005)
Glacier fluctuation in Bolterdalen, Svalbard
317
Fig. 8. A. Bedding architecture and facies from 200–270 m, and the relative age relationship between foreset beds and tectonic deformation
(1–4). B. Thrust and shear deformation from the upper fan foreset and overlying 5-cm-thick till. C. Normal faults, triggered by the advancing
glacier, are truncated by ice-marginal thrust faults that are associated with intense dewatering of the well-sorted sandy foreset facies.
thrust faults (Fig. 8B, C). These sediments were
unfrozen when pushed northwards by the advancing ice
front, and subsequently overridden and sheared. The
rather complex depositional and structural architecture
suggest that the glacier was lifted from the sediment
surface, presumably by tidewater (present tidal range is
c. 1 m). The ice front must have been thin, and the
cross-cutting thrust faults indicate that the terminus
fluctuated seasonally.
Unit 4 is interpreted as a deformation till (see
Boulton 1987; Hart & Boulton 1991; Hart 1995; Hart
& Rose 2001) with large textural and structural variations. The glacier, eroding, depositing and deforming
the substrate, was at least partially temperate and
advanced towards the northwest.
Unit 5 – marine facies
Unit 5 is recognized as a thin horizon of sandy wave
ripples above unit 4, c. 50 m a.s.l. (Fig. 4). These
postglacial marine deposits correlate to a marine
erosional surface, whose upper limit has been difficult
to establish. The maximum altitude, however, is limited
by the 58-m sea level at the ice-front advance. A series
of beaches formed after the glacier retreat is partly cut
into the ice-contact deposits and partly depositional in
character.
Surface morphology and sediment distribution
An aerial photograph from the study area (Fig. 3)
shows that the sediment surface has a distinct character,
with a pronounced fan-shaped depocenter that is
described as unit 6. Grooves and sediment flows indicate that the volume of debris along the crest of the hill,
and particularly towards the northern tip, was larger
than elsewhere, and have smoothed out the bedrock and
beach relief down to an altitude of c. 57 m a.s.l.
The bedrock morphology from 110 to 62 m has an
uneven debris cover that varies texturally from washed
gravel to matrix-supported diamicton, with clasts up to
2.5 m. The boulders are subrounded with spherical to
blocky shape and are embedded in diamictic debris
with a washed and reworked surface, as seen from the
network of decimetre-deep grooves.
The marine limit is at 62 m a.s.l., as inferred from the
exposed beach facies in unit 2. Morphometric profiles
from the east-facing part of the ridge (Fig. 2) show a
small break at this level. It is difficult to exclude that
sea level was higher because the palaeobeach on the
east side of the hill was more or less parallel to the
strike of the bedrock, and anthropogenic activity has
hampered the direct link between surface morphology
and sediment units. However, 62 m a.s.l. corresponds
well with the eastwards drop in ML along Adventdalen
palaeofjord (Fig. 1). The area between 62 and 58 m
consists of a regressive beach sequence along the ditch,
318
Ida Lønne
whereas the eastern part of the hill shows wave-washed
bedrock.
Discussion
Palaeoclimatic inferences based on former ice margins
depend strongly on correct identification of climatically
controlled events. This recognition is most sensitive for
areas with limited data, and particularly in high latitudes where glacier accumulation rates are low
(Paterson 1994; Hodkins 1997) and the glacier’s
response to climatic cooling is slower and shorter
compared to mid-latitudes. It is well known that tidewater glaciers may fluctuate independently of climate
(Meier & Post 1987; Alley 1991), but the ice front in
Bolterdalen advanced at least 1 km on land before
moving into the fjord.
Surge advances are common along the modern ice
fronts on Svalbard, and probably also affected former
ice masses. The mechanisms that trigger glacial surges
are not fully understood (Hamilton & Dowdeswell
1996; Harrison & Post 2003), but it is becoming
increasingly apparent that the Svalbard surges have a
character that is different from surges elsewhere. The
rates of ice flow and terminus advance are lower, and
the active phase duration and run-out distance are
longer (Dowdeswell et al. 1991), which render it even
more complicated to distinguish between dynamically
and climatically forced fluctuations. The most extended
active surge phase recorded is from Fridtjovbreen
(Bellsund area, Fig. 1) that surged for over 12 years;
the duration of the ice-front advance period was 5 years
and the area affected by syn-surge sedimentation was
4–5 km2 (Lønne in prep.).
The ice-front advance deposits in Bolterdalen cover
200 m in ice-flow parallel direction. The presence of
long, thick-shelled molluscs in growth position, and
only scattered juvenile individuals, indicates that the
event lasted several decades. This is supported by
the regular presence of clay-rich sediments close to the
former sea level that reflect periods with frozen fjord,
and a pronounced age difference between the oldest
and youngest part of the delta (Fig. 4). It is therefore
concluded that the succession reflects a climatically
controlled fluctuation of a glacier that occupied
Bolterdalen.
Early Holocene glacial event and palaeoenvironmental
implications
Unit 1 is a subglacial unit with exotic lithologies.
Reddish sandstones occur in Devonian and Early
Carboniferous part of the stratigraphic record and
quartzite and quartzitic conglomerates are also
common in Early Carboniferous strata (Dallmann
1999). The Early Palaeozoic record is not mapped east
of Bolterdalen (Dallmann et al. 2002), but may possibly
be brought to the surface along some major tectonic
BOREAS 34 (2005)
lineaments in the glaciated areas. The palaeoenvironmental implication is that the till was emplaced by eastflowing ice. The gradual contact to the beach succession
indicates deposition during the Late Weichselian retreat,
followed by glacioisostatically forced beach progradation towards the northeast (Fig. 9A). The radiocarbon
date from unit 2 is correlated to the beach gravel at 62 m
a.s.l. and suggests that the age of the marine limit is
10 025 yr BP or slightly older.
The subsequent ice-front advance is the first record
of Early Holocene glacier expansion in the Adventdalen area. Four distinct phases are distinguished. The
environmental conditions during each of these provide
new insight into the dynamics related to glacier growth
and decay, and how glacial relief and deposits were
nearly removed during the postglacial time.
Phase 1: Ice-front re-advance across a terrestrial
surface. – The Bolterdalen glacier advanced northwards over a bedrock that probably still hosted dead ice
from the deglaciation a few hundred years earlier. At
this time, the ice front may still have occupied the head
of the Adventdalen palaeofjord, as seen from reddish
blocks of ice-rafted sandstone in unit 2. Reddish
sandstones in the upper diamicton were more likely
entrained from glacial debris left onshore. The advancing ice front proceeded across the regressive beach
terraces (Fig. 9B) that were frozen, except for an active
layer formed in the summer. The sharp base of the till
above unit 2 (Fig. 6A) was formed as a combination of
subglacial erosion and thrust faults. The heterogeneous
textural mixture of sorted sand, beach gravel and
massive diamicton cannot have formed by subglacial
abrasion, particle by particle, but is rather a result of
frozen rafts incorporated into the till (see Cuffey &
Alley 1996; Alley et al. 1997). Unit 4 has a sharp
lithological contrast to the underlying, apparently
undisturbed foreset (Fig. 6A); however, the mapped
clast orientation and bedding structures (Fig. 6B)
demonstrate that thrust and shear deformation extended
into the substrate.
Phase 2: Progradation of a shallow marine ice-contact
delta, overridden by the glacier. – The thickness of the
ice-contact delta (unit 3, Fig. 9B) suggests that a
subglacial drainage system with a consistent depositional axis existed throughout phase 2. This basal
meltwater supply was presumably controlled by the
glaciological conditions upglacier where the ice was
thicker and at the pressure melting point.
The delta has a typical shoal-water profile (Postma
1990), with a short delta slope (3–5 m vertical thickness) composed of sediment gravity flow deposits
intercalated with wave-worked strata with a smooth
transition to prodelta mud. The foreset beds, onlapping
the distal facies of unit 2, are capped by a waveinfluenced delta topset. The fluvial sediment input must
have been relatively low and the topset formed at some
distance from the ice front. Mouth bars indicate that the
BOREAS 34 (2005)
Glacier fluctuation in Bolterdalen, Svalbard
319
Fig. 9. Reconstructed evolution of the sedimentary succession in Bolterdalen. The glacial event is divided into four phases (1–4, see text for
discussion).
320
Ida Lønne
meltwater supply shifted seasonally and increased as
the glacier approached the fjord. Today, rivers are
running 2–3 months of the year (Førland et al. 1997)
and the fjord heads are frozen 6–8 months. Horizons of
intense bioturbation that occur regularly along the
outcrop (e.g. Fig. 7D, inset) are related to periods with
negligible fluvial input and open sea, whereas the clayrich horizons formed when the fjord was ice-covered.
As the sediment input to the fjord increased, wave
reworking declined, and the matrix-rich sediment
wedges in the lower topset from 180–220 m can be
characterized as a prograding wave-influenced ‘dumpmoraine’, onlapping the delta front (Fig. 6D). Proglacial push and thrust in front of the advancing glacier
triggered resedimentation and slumping. The absence
of turbiditic delta front lobes, which are common in
ice-contact deltas (Lønne 2001; Lønne et al. 2001), are
attributed to the low sediment input, wave-reworking
and the release of very small-volume sediment flows
along the short delta slope. A large portion of the sand
at the delta front was deposited from suspension.
Deformation of the substrate was moderate and
displacement so small that the overridden foreset beds
appear undisturbed in the field. Detailed mapping of the
deformation structures shows, however, that pervasive
tectonic deformation reached a maximum depth of 1 m
below the till base. Elongated particles have been
reoriented, a few drag-folds are found, and, locally,
subglacial shear has disturbed all primary sedimentary
structures (Fig. 7A).
Phase 3: Submarine fan facies formed by a tidal
influenced terminus. – The input of meltwater-delivered
debris persisted throughout phase 3, but as the thin
ice front moved into deeper water, the glacier sole was
lifted from the substrate, presumably at high tide. The
fluctuating buoyancy line is shown by cross-cutting
geometries of diamictic debris, sedimentary packages
with dewatering structures and glaciotectonic thrust
faults from c. 200 m and folded foreset beds (Fig. 8A).
The fluvial, delta topset facies are absent (Fig. 9C),
presumably due to increased water depth and higher
sediment input, or higher advance rate. All the meltwater sediment was redistributed to the delta slope,
which genetically is thus a submarine ice-contact fan or
a grounding-line fan (cf. Lønne 1995, 2001). Such facies
often possess larger variations in grain-size and sorting
compared to deltas (Benn & Evans 1998; Lønne 1997;
Nemec et al. 1999), as also demonstrated by comparing
the facies from phases 2 and 3.
Fine-grained deposits are rare or absent on the upper
part of delta front because of the high energy conditions here (Nemec 1990; Postma 1990). The regular
presence of clay-rich horizons in unit 3 is therefore
attributed to periods with ice-covered fjord. The high
input of subglacial stream sediment during the deltaic
stage, particularly when fjord ice breaks up in the
spring, would cover the winter mud and increase its
BOREAS 34 (2005)
preservation potential. The thicker and more frequent
mud layers in the fan foreset are related to increased
progradation rate and relative decline in wave working,
whereas the lack of mud in unit 2 reflects wave erosion
rather than non-deposition.
Ice-rafted debris indicates that blocks of ice were
breaking off the thin tidewater ice front. Seasonal
terminus fluctuations triggered pulses of submarine
push of sediments along the north-dipping sea floor and
normal faulting, associated by repeated events of
proglacial thrusting and subglacial shearing (Fig. 8).
The maximum ice-front position at 270 m was reached
at 9625+95 yr BP (Fig. 9C).
The large spatial variability in unit 4 was formed by
a combination of several factors. Glacial debris,
transported from wet-based upglacier portions of the
glacier, was transferred into the colder snout that
moved over a substrate with a drop of 20 m along the
discussed transect. The geometrical organization has a
twofold deformational structure, but indicates only one
single episode of glaciotectonic deformation. Subglacial shear was spread as thrusts towards the ice margin,
both in the ice and the substrate, and proglacial thrustblock moraines (Hambrey & Huddart 1995; Huddart &
Hambrey 1996) propagated downslope in front of the
advancing glacier. These were truncated, partly incorporated by the overriding glacier and transformed to a
deformation till. There is no dramatic change in
deformation style as the glacier moves from frozen
beach gravel, via intertidal topset, to a submarine
foreset substrate.
The base of unit 4 is texturally sharp, but structurally
composite. The pronounced lithological contrasts
suggest that the till base was a direct continuation of
the ice/bed interface (see Hart 1995; Piotrowski &
Tulaczyk 1999), but with additional small-scale
deformation that continued into the substrate (Murray
1997; van der Meer et al. 2003). Smaller ramps and
irregularities along the till base suggest that the tectonic
strain varied, as a result of the thin ice and surface
dipping away from the glacier. The boundary to the
frozen surface of unit 2 is quite sharp, but here, too,
thrust planes continue into the substrate. The top of the
unfrozen delta (unit 3) shows larger variations, with
local thickening of the till due to downslope-directed
thrusts (e.g. Fig. 6C). Clasts are locally transported into
the sandy intertidal part of the delta along such tectonic
surfaces (a to d, in Fig. 7A). The base of this clast
transport correlates to the base of a sharp décollement
on the outer delta front, which may represent the
boundary between winter-frozen intertidal delta top and
the water-saturated foreset below.
Phase 4: Final glacier retreat, sediment redistribution
and glacioisostatic uplift. – The glacier calved rapidly
back from the fjord, whereas the thin onshore portion
inferably stagnated and broke up into isolated blocks
that were subject to passive down-wasting (Fig. 9D).
BOREAS 34 (2005)
There are no remnants of ice-cored moraines or
hummocky sediment ridges today, which suggests that
these never formed or were later eroded. Debris
released from the ice was redistributed by glacial
meltwater and the surface of the upper till was modified
and partly removed, and bedrock and beach terraces
were covered by a thin sheet of reworked debris (unit 6,
Fig. 9D). Such subaerial glacigenic sediment-flow
deposits commonly occur along down-wasting terrestrial ice margins (Lawson 1988; Bennett et al. 2000).
Palaeoclimatic implications
The 100-km-long Isfjorden basin was evidently
deglaciated between 12.3 and 11 kyr BP as seen from
radiocarbon dates from the lake Linnévannet at the
mouth of the fjord (Mangerud et al. 1992) and Kapp
Ekholm at the fjord head (Mangerud & Svendsen
1992). By inference, the outer parts of the tributaries
were also partly ice-free; however, the chronology for
this period is poorly documented (Svendsen et al.
1996). A tentative age for the marine limit at the mouth
of Adventfjorden is thus c. 11 kyr BP or slightly older.
The fjord floor has an average sediment thickness of
12 m (Elverhøi et al. 1995) and rises evenly towards the
east with no evidence of distinct ice-front stillstand
positions (unpublished seismic data, UNIS). The
deglaciation succession along the valley floor is
covered by prograding delta facies that reach c. 80 m
thickness at the fjord head (Johansen et al. 2003). The
withdrawal of ice in the Adventdalen palaeofjord, east
to Bolterdalen, appears to have been relatively uniform
and spanned about a thousand years in time, with an
average glacioisostatic uplift rate of c. 0.8 m/100 years,
rising to 1.6 m/100 years from 10 to 9.8 kyr BP. Similar
rates were obtained from the inner part of Isfjorden
(Salvigsen 1984).
The glacial event in Bolterdalen is attributed to a
brief climatic event, bracketed by the two radiocarbon
dates 9.8 and 9.6 kyr BP. Two other glaciers along the
coasts of Isfjorden were advancing in the same time
period; Aldegondabreen advanced after 9980+120 yr
BP (T-6290; O. Salvigsen, pers. comm. 2004) and
Esmarkbreen advanced shortly after 9500+120 yr BP
(T-6286; Salvigsen et al. 1990). These glacier events
are thought to be of short duration. Thermophilous
mollusc taxa from central Spitsbergen indicate warmer
water from 9.5 kyr BP (Salvigsen et al. 1992).
The retreat of marine outlet glaciers is controlled by
calving, strongly influenced by water depth and thinning of the ice. The deglaciation dynamics changes
dramatically, however, when the glaciers become landbased, and for glaciers whose mass balance is largely
controlled by summer ablation the mean air temperature during the ablation season and elevation are critical
factors (Paterson 1994; Benn & Evans 1998).
The shortage of water in polar deserts strongly
influences the rate and mode of sediment transfer.
Glacier fluctuation in Bolterdalen, Svalbard
321
Today, the average annual precipitation in Adventdalen
is 190 mm and a considerable proportion falls as snow
(Førland et al. 1997). More importantly, the drainage of
water into the site in Bolterdalen is negligible and
glacial meltwater was therefore critical for the redistribution of sediment during the Early Holocene.
Reworking commenced as the ice-front retreated 9.6 yr
BP and proceeded until the glacioisostatic sea level fall
had reached 40 m, c. 9 kyr BP (Lønne & Nemec 2004).
Dead ice may thus have survived for about 500 years at
sea level. It is interesting to compare with the melting
rates from a south-facing cirque glacier at Hiorthfjellet
(Fig. 1). At this site, 500–900 m a.s.l., glacier ice may
have survived for 5000 years, as sediment reworking
proceeded until c. 6 kyr BP (Lønne & Nemec 2004).
The deglaciation in Bolterdalen was characterized by
a thin tidewater glacier where sediments were chiefly
transported by meltwater, as seen from the large
proportion of the substrate-derived debris in the
deformation till. No material from the adjacent mountain walls reached the site. The terminus eventually
broke up into isolated blocks, aided by the meltwater,
without the formation of ice-cored moraines. The
release and reworking of debris occurred more or less
contemporaneously.
Identification of terrestrial ice margins in
polar regions
Ice-cored moraines, which commonly disintegrate to
form hummocky moraines (Bennett & Boulton 1993;
Hambrey et al. 1997), are rarely seen in the Svalbard
landscape, and apparently have a low preservation
potential. The low debris content in the glaciers is one
important factor. If the proglacial surface is frozen,
water from the melting ice will follow the ground
surface or base of the active layer, and a larger portion
of the glacial sediments will be reworked. Examples
from Arctic Canada (Dyke 1993) demonstrate that if
the meltwater volume is high, all sediments may be
removed and only fluvial channels may delineate the
former ice margin. However, in areas with restricted
meltwater, as in Bolterdalen, thin overlapping sheets of
redistributed glacial debris is the only morphology left
behind. Such isolated depocentra that reflect ancient
periods of glacial meltwater activity (e.g. Lægreid
1999; Soltvedt 2000; Lønne & Nemec 2004) may play
an important part in the reconstruction of former
terrestrial ice margins in semi-arid to dry polar regions.
The highest beach terraces in Bolterdalen appear as
morphologically well preserved, but they were
obviously eroded and the sediments were incorporated
in the glacier sole. The conclusive evidence of
subglacial processes has been from detailed mapping
on a bed-scale, and the conform ice-flow parallel orientation of thrust and shear deformation and clast-fabric
322
Ida Lønne
in till and beach gravel. There is no dramatic change in
deformation style between the previously frozen and
unfrozen sediments at this site, but the till has a sharper
base along the presumably frozen ground surface
compared to the unfrozen delta where water escape
structures are cut by thrust faults.
Conclusion
The present study of a continuous succession of icefront advance and retreat deposits in Bolterdalen, well
constrained in space and time, provides new information on the nature of ice-marginal processes in extreme
polar settings.
Four distinct phases are distinguished: (1) Ice-front
advance over a frozen sediment surface; (2) advance
into shallow marine water with formation of a waveinfluenced ice-contact delta; (3) continued advance and
deposition of a submarine fan foreset in front of a
partly floating tidewater front, and (4) final ice-front
retreat where subglacial till deposits and sediment rich
ice left onshore were subject to reworking and redistribution into thin sediment sheets.
The advancing glacier was evidently thin, but thick
enough to deform frozen beach facies and unfrozen
deltaic sediments down to a depth of 1 m. Subglacially
transported debris and sediments from the substrate
were remoulded to a 1-m-thick deformation till, which
is in strong contrast to the few modern observations
available that are only a few centimetres thick
(Piotrowski et al. 2001).
The formation of an ice-front advance delta fed by
subglacial meltwater appears incompatible with a coldbased terminus. There are observations from glaciers
on Svalbard where meltwater produced in thicker,
temperate ice upglacier is transported through the coldbased margin (Hodkins 1997). However, the glacier in
Bolterdalen was not frozen to the substrate: the icefront moved northwards, it fluctuated seasonally and
the subglacial till includes beach deposits that have
been transported some tens of metres in an ice-flow
parallel direction.
Acknowledgements. – The fieldwork was funded by the University
Centre in Svalbard. Radiocarbon-dating was done at the Radiological Laboratory in Trondheim. O. Salvigsen allowed an unpublished radiocarbon date (T-6290) to be included, and the Norwegian
Polar Institute kindly permitted publication of the aerial photograph
S90-1241. The manuscript was reviewed by D. Huddart and J. van
der Meer.
References
Alley, R. B. 1991: Sedimentary processes may cause fluctuations of
tidewater glaciers. Annals of Glaciology 15, 119–124.
Alley, R. B., Cuffey, K. M., Evenson, E. B., Strasser, J. C., Lawson,
D. E. & Larson, G. J. 1997: How glaciers entrain and transport
basal sediment: physical constraints. Quaternary Science Reviews
16, 1017–1038.
BOREAS 34 (2005)
Benn, D. I. & Evans, D. J. A. 1996: The interpretation and classification of subglacially-deformed materials. Quaternary Science
Reviews 15, 23–52.
Benn, D. I. & Evans, D. J. A. 1998: Glaciers and Glaciation. 734 pp.
Arnold, London.
Bennett, M. R. & Boulton, G. S. 1993: A reinterpretation of Scottish
‘hummocky moraine’ and its significance for the deglaciation of
the Scottich Highlands during the Younger Dryas or Loch
Lomond Stadial. Geological Magazine 139, 301–318.
Bennett, B. R., Huddart, D., Glasser, N. F. & Hambrey, M. J. 2000:
Resedimentation of debris on an ice-cored lateral moraine in the
high-Arctic (Kongsvegen, Svalbard). Geomorphology 35, 21–40.
Bennett, M. R., Huddart, D., Hambrey, M. J. & Ghienne, F. 1996:
Moraine development at the high-arctic valley glacier Pedersenbreen, Svalbard. Geografiska Annaler 78A, 209–222.
Boulton, G. S. 1987: Sediment deformation beneath glaciers: rheology and geological consequences. Journal of Geophysical
Research 92, 9059–9082.
Boulton, G. S., van der Meer, J. J. M., Hart, J., Beets, D., Ruegg, G.
J. J., van der Wateren, F. M. & Jarvis, J. 1996: Till and moraine
emplacement in a deforming bed surge – an example from a
marine environment. Quaternary Science Reviews 15, 961–987.
Cuffey, K. & Alley, R. B. 1996: Is erosion by deforming subglacial
sediments significant? (Toward till continuity). Annals of
Glaciology 22, 17–24.
Dallmann, W. K. (ed.) 1999: Lithostratigraphic Lexicon of Svalbard. Review and Recommendations for Nomenclature Use.
Upper Palaeozoic to Quaternary Bedrock. Norsk Polarinstitutt,
Tromsø, 318 pp.
Dallmann, W. K., Ohta, Y., Elvevold, S. & Blomeier, D. (eds.)
2002: Bedrock map of Svalbard and Jan Mayen. Norsk Polarinstitutt Temakart No. 33.
Dowdeswell, J., Hamilton, G. & Hagen, J. O. 1991: The duration
of the active phase of surge-type glaciers: contrasts between
Svalbard and other regions. Journal of Glaciology 37, 86–98.
Dyke, A. S. 1993: Landscapes of cold-centred late Wisconsinian
ice caps, Arctic Canada. Progress in Physical Geography 17,
223–247.
Etzelmüller, B., Hagen, J. O., Vatne, G., Ødegård, R. S. & Sollid,
J. L. 1996: Glacier debris accumulation and sediment deformation
influenced by permafrost, examples from Svalbard. Annals of
Glaciology 22, 53–62.
Elverhøi, A., Svendsen, J. I., Solheim, A., Andersen, E. S., Milliman, J. D., Mangerud, J. & Hook LeB, R. 1995: Late Quaternary
sediment yield from the high Arctic Svalbard area. Journal of
Geology 103, 1–17.
Førland, E. J., Hanssen-Bauer, I. & Nordli, P. Ø. 1997: Climate
statistics and longterm series of temperature and precipitation at
Svalbard and Jan Mayen. Det Norske Meteorologiske Institutt
Report 21/97, 72 pp.
Glasser, N. F., Bennett, M. R. & Huddart, D. 1999: Distribution of
glaciofluvial sediment within and on the surface of a high arctic
valley glacier: Marthabreen, Svalbard. Earth Surface Processes
and Landforms 24, 303–318.
Hambrey, M. J. & Huddart, D. 1995: Englacial and proglacial
glaciotectonic processes at the snout of a thermally complex
glacier in Svalbard. Journal of Quaternary Science 10, 313–326.
Hambrey, M. J., Huddart, D., Bennett, M. R. & Glasser, N. F. 1997:
Genesis of ‘hummocky moraines’ by thrusting in glacier ice:
evidence from Svalbard and Britain. Journal of Geological
Society, London 153, 623–632.
Hamilton, G. S. & Dowdeswell, J. A. 1996: Controls on glacier
surging in Svalbard. Journal of Glaciology 42, 157–168.
Hart, J. K. 1995: Subglacial erosion, deposition and deformation
associated with deformable beds. Progress in Physical Geography 19, 173–191.
Hart, J. K. & Boulton, G. S. 1991: The interrelationship between
glaciotectonic deformation and glaciodeposition. Quaternary
Science Reviews 10, 335–350.
BOREAS 34 (2005)
Hart, J. & Rose, J. 2001: Approaches to the study of glacier bed
deformation. Quaternary International 86, 45–58.
Harrison, W. D. & Post, A. S. 2003: How much do we really know
about glacier surging? Annals of Glaciology 26, 1–6.
Hodkins, R. 1997: Glacier hydrology in Svalbard, Norwegian High
Arctic. Quaternary Science Reviews 16, 957–973.
Huddart, D. & Hambrey, M. J. 1996: Sedimentary and tectonic
development of a high-arctic thrust moraine complex: Comfortlessbreen, Svalbard. Boreas 25, 227–243.
Jiskoot, H., Boyle, P. J. & Murray, T. 1998: The incidence of glacier
surging in Svalbard: evidence from multivariate statistics.
Computers in Geoscience 24, 387–399.
Johansen, T. A., Digranes, P., Schaack, M. van & Lønne, I. 2003:
Seismic mapping and modeling of near-surface sediments in polar
areas. Geophysics 68, 566–573.
Komar, P. D. 1976: Beach Processes and Sedimentation. 417 pp.
Prentice Hall, Englewood Cliffs, N.J.
Landvik, J. Y., Bondevik, S., Elverhøi, A., Fjeldskaar, W.,
Mangerud, J., Siegert, M. J., Salvigsen, O., Svendsen, J. I. &
Vorren, T. O. 1998: The last glacial maximum of Svalbard and
the Barents Sea area: ice sheet extent and configuration.
Quaternary Science Reviews 17, 43–75.
Landvik, J. Y., Ingólfsson, Ó., Mienert, J., Lehman, S. J., Solheim,
A., Elverhøi, A. & Ottesen, D. 2005: Rethinking Late Weichselian ice sheet dynamics in coastal NW Svalbard. Boreas 34, 1–19.
Lawson, D. E. 1988: Glacigenic resedimentation: classification
concepts and application to mass-movement processes and
deposits. In Goldtwait, R. P. & Matsch, C. L. (eds.): Genetic
Classification of Glacigenic Deposits, 147–169. A. A. Balkema,
Rotterdam, 294 pp.
Lefauconnier, B. & Hagen, J. O. 1991: Surging and calving glaciers
in eastern Svalbard. Norsk Polarinstitutt, Meddelelser 116, 130 pp.
Lyså, A. & Lønne, I. 2001: Moraine development at a small
High-Arctic valley glacier: Rieperbreen, Svalbard. Journal of
Quaternary Science 16, 519–529.
Lægreid, A. K. 1999: Postglacial Sedimentation in a High-Arctic
Valley: Slope Processes and Geomorphic Development in
Endalen, Spitsbergen. Cand. Scient. thesis, University of Bergen
and University Courses on Svalbard, 135 pp.
Lønne, I. 1995: Sedimentary facies and depositional architecture
of ice-contact glaciomarine systems. Sedimentary Geology 98,
13–43.
Lønne, I. 1997: Sedimentology and depositional history of an
early Holocene ice-contact submarine fan: the Egge-Lyngås
‘end-moraine’, southern Norway. Norsk Geologisk Tidsskrift 77,
137–157.
Lønne, I. 2001: Dynamics of marine glacier termini read from
moraine architecture. Geology 29, 199–202.
Lønne, I. & Lauritsen, T. 1996: The architecture of a modern pushmoraine at Svalbard as inferred from ground-penetrating radar
measurements. Arctic and Alpine Research 28, 488–495.
Lønne, I. & Lyså, A. In press: Deglaciation dynamics following the
Little Ice Age on Svalbard: implications for shaping of landscapes
at high latitudes. Geomorphology.
Lønne, I. & Nemec, W. 2004: High-arctic fan delta recording
deglaciation and environment disequilibrium. Sedimentology 51,
553–589.
Lønne, I., Nemec, W., BLikra, L. H. & Lauritsen, T. 2001: Sedimentary architecture and dynamic stratigraphy of a marine icecontact system. Journal of Sedimentary Research 71, 922–943.
Major, H., Haremo, P., Dallmann, W. K. & Andresen, A. 2000:
Geological map of Svalbard, 1 : 100.000, sheet C9G Advetndalen,
Norsk Polarinstitutt, Tromsø.
Mangerud, J. & Gulliksen, S. 1975: Apparent radiocarbon ages of
recent marine shells from Norway, Spitsbergen, and Arctic
Canada. Quaternary Research 5, 263–273.
Mangerud, J. & Svendsen, J. I. 1992: The last interglacial–glacial
period on Spitsbergen, Svalbard. Quaternary Science Reviews 11,
633–664.
Glacier fluctuation in Bolterdalen, Svalbard
323
Mangerud, J., Bolstad, M., Elgersma, A., Helliksen, D., Landvik,
J. Y., Lønne, I., Lycke, A. K., Salvigsen, O. & Svendsen, J. I.
1992: The last glacial maximum on western Svalbard. Quaternary
Research 38, 1–31.
Massari, F. & Parea, G. C. 1988: Progradational gravel beach
sequences in a moderate to high-energy microtidal marine
environment. Sedimentology 35, 881–913.
Meier, M. F. & Post, A. 1969: What are glacier surges? Canadian
Journal of Earth Sciences 6, 807–817.
Meier, M. F. & Post, A. 1987: Fast tidewater glaciers. Journal of
Geophysical Research 92 (B9), 9051–9058.
Murray, T. 1997: Assessing the paradigm shift: deformable glacier
beds. Quaternary Science Reviews 16, 995–1016.
Nemec, W. 1990: Aspects of sediment movement on steep delta
slopes. In Colella, A. & Prior, D. B. (eds.): Coarse Grained
Deltas. International Association of Sedimentologists International Publication 10, 29–73.
Nemec, W., Lønne, I. & Blikra, L. H. 1999: The Kregnes moraine in
Gauldalen, west-central Norway: anatomy of a Younger Dryas
proglacial delta in a palaeofjord basin. Boreas 28, 454–476.
Paterson, W. S. B. 1994: The Physics of Glaciers. 480 pp. Pergamon, Oxford.
Piotrowski, J. A., Mickelson, D. M., Tulaczyk, S., Krzyszkowski, D.
& Junge, F. W. 2001: Were deforming subglacial beds beneath
past ice sheets really widespread? Quaternary International 86,
139–150.
Piotrowski, J. A. & Tulaczyk, S. 1999: Subglacial conditions under
the last ice sheet in northwest Germany: ice-bed separation
and enhanced basal sliding? Quaternary Science Reviews 18,
737–751.
Postma, G. 1990: Depositional architecture and facies of river and
fan deltas: a synthesis. In Colella, A. & Prior, D. B. (eds.): Coarse
Grained Deltas. International Association of Sedimentologists
International Publication 10, 13–27.
Salvigsen, O. 1984: Occurrence of pumice on raised beaches and
Holocene shoreline displacement in the inner Isfjorden area,
Svalbard. Polar Research 2, 107–113.
Salvigsen, O., Elgersma, A., Hjort, C., Lagerlund, E., Liestøl, O. &
Svensson, N.-O. 1990: Glacial history and shoreline displacement
on Erdmannflya and Bohemanflya, Spitsbergen, Svalbard. Polar
Research 8, 261–273.
Salvigsen, O., Forman, S. & Miller, G. 1992: Thermophilous
molluscs on Svalbard during the Holocene and their paleoclimatic
implications. Polar Research 11, 1–10.
Sexton, D. J., Dowdeswell, J. A., Solheim, A. & Elverhøi, A. 1992:
Seismic architecture and sedimentation in northwestern Spitsbergen fjords. Marine Geology 103, 53–68.
Sletten, K., Lyså, A. & Lønne, I. 2001: Formation and disintegration
of a modern high-arctic ice-contact system, Scott Turnerbreen,
Svalbard. Boreas 30, 272–284.
Soltvedt, S. 2000: The Sedimentation Pattern of the Last Deglaciation and Postglacial Phase in a High-Arctic Valley: Endalen,
Spitsbergen. Cand. Scient. thesis, University of Bergen and
University Courses on Svalbard, 144 pp.
Svendsen, J. I., Elverhøi, A. & Mangerud, J. 1996: The retreat of the
Barents Sea Ice Sheet on western Svalbard margin. Boreas 25,
244–256.
van der Meer, J. J. M., Menzies, J. & Rose, J. 2003: Subglacial
till: the deforming glacier bed. Quaternary Science Reviews 22,
1659–1685.
Whittington, R. J., Forsberg, C. F. & Dowdeswell, J. A. 1997:
Seismic and side-scan sonar investigations of recent sedimentation
in an ice-proximal glacimarine setting, Kongsfjorden, north-west
Spitsbergen. In Davies, T. A., Bell, T., Cooper, A., Josenhans, H.,
Solheim, A., Stoker, M. S. & Stravers, J. A. (eds.): Glaciated
Continental Margins. An Atlas of Acoustic Images, 175–178.
Chapman & Hall, London.