Glacial geology and deglaciation chronology of the
area between inner Nordfjord and
Jostedalsbreen
Strynefjellet, western Norway
NORALF RYE, ATLE NESJE, RUNE LIEN, LARS HARALD BLIKRA, OLIANNE EIKENÆS,
PER AUDUN HOLE & INGRID TORSNES
Rye, N., Nesje, A., Lien, R., Blikra, L. H., Eikenæs, 0., Hole, P. A. & Torsnes, 1.: Glacial geology and deglaciation chronology of
the area between inner Nordfjord and Jostedalsbreen - Strynefjellet, western Norway. Norsk Geologisk Tidsskrift, Vol. 77, pp.
51-63. Oslo 1997. ISSN 0029-196X.
A lower limit of blockfields is inferred to indicate the maximum heights and thus thickness of the Late Weichselian iee sheet in the
inner Nordfjord region. lee movements in this area have been topographically controlled during the entire Weichselian glaciation.
Prominent lateral moraines de1imit the Younger Dryas valley glaciers in inner Nordfjord. Subsequent to the Younger Dryas
Chronozone, the glaciers retreated rapidly due to calving in the fjord and climatic amelioration. In a later phase of deglaciation, in
all probability around the early and middle part of the Preboreal Chronozone, an iee eentre east of Strynefjellet dominated, while
the Jostedalsbreen area is thought to have played a minor role as a eentre of iee dispersal. The final deg1aciation was dominated
by vertically down-wasting iee remnants in the lake basins and tributary valleys. Terminal moraines in front of several outlet
glaciers of Jostedalsbreen beyond the 'Little lee Age' moraines indicate a climatic deterioration at the end of the Preboreal
Chronozone.
N. Rye, Department of Geology, University of Bergen, A/legt. 41, N-5007 Bergen; A. Nesje, Department of Geography, University of
Bergen, Breiviken 2, N-5035 Bergen-Sandviken; R. Lien, Statens Vegvesen, P.O. Box 608, 9800 Vadsø; L. H. Blikra, Geological
Survey of Norway, P.O. Box 3006, N-7002 Trondheim; O.
N-0301 Oslo; P. A. Hole, Statoil, N-5020 Bergen;
l.
Eikenæs,
Norges vassdrags- og energiverk, P.O. Box 5091 Majorstua,
Torsnes, Mø//esvingen 2, 0854 Oslo, Norway.
Introduction
Recent investigations in the area between ioner Nord­
fjord and Jostedalsbreen (breen = glacier) - Strynefjellet
(fjellet= mountain) (Rye et al. 1984, 1987; Nesje 1984;
Lien 1985; Hole 1985; Blikra 1986; Nesje et al. 1987,
1991, Nesje & Kvamme 1991; Nesje 1992; Dahl & Nesje
1992; Nesje & Dahl 1992; McCarroll & Nesje 1993;
Torsnes et al. 1993) make it possible to present a deglaci­
ation history from the Late Weichselian maximum up to
the present in this part of western Norway.
Marine terraces have previously been mapped by Kal­
dhol (1912). Rye (1963, 1978) and Fareth (1970, 1987)
described the deglaciation in the middle and inner parts
of Nordfjord, while the areas between Jostedalsbreen and
Strynefjellet have been less well known. Stokke (1982)
mapped the Quatemary deposits in the valley bottom in
Stryn and Hjelledalen.
Bedrock
The bedrock in Nordfjord consists of Precambrian
gneisses and granites which are l 000-1800 million years
old. The rocks are divided into two main units; the
Fjordane Complex (at surface) and the Jostedal Complex
(at depth). Ioner Nordfjord belongs to the Jostedal Com­
plex. The dorninating rocks are banded gneiss and
granitic gneiss. The tectonic history of the bedrock in
ioner Nordfjord is complicated and the rocks have been
subject to several deformation phases during the Precam­
brian and Caledonian orogenies.
Landforms
The landscape of ioner Nordfjord (Figs. l, 2) has evolved
from a plateau landscape which was developed close to
sea level during the Mesozoic and during a subsequent
uplift in the Tertiary. At present, remnants of this
plateau landscape can be seen as rather smooth, undulat­
ing sumrnit areas along Nordfjord, gently sloping from
1800-2000 m at Jostedalsbreen to 400-600 m at the
coast (Fig. 3). During the Quatemary glaciations, the
plateau landscape was exposed to glacial erosion. During
the ice-free interglacial periods, fluvial and avalanche
activity modified the landscape. The valleys in ioner
Nordfjord have topographic features characteristic of
glaciated areas: steep valley sides with U-shaped trans­
verse profiles, bedrock basins occupied by lakes or filled
in with sediments, rock thresholds, cirques, and hanging
valleys. The fjord and the main valleys in ioner Nord­
fjord are good examples of topographic features, whose
direction and glaciated form were controlled by fracture
zones more so than in the homogeneous and resistant
areas. In the main valleys there are both wide, deep
basins and short, narrow gorges where rivers form water­
falls.
52
N. Rye et al.
NORSK GEOLOGISK TIDSSKRIFT
77 (1997)
SUNNMØRE
Skjlk
o
Fig.
10
20
30km
l. Location map of Nordfjord.
One branch of the Nordfjord, Oldedalen (dalen =
valley), stretches about 20 km southward to Brigsdalen
where two outlet glaciers from Jostedalsbreen, Brigsdals­
breen and Melkevollsbreen are situated. Another branch,
Lodalen, reaches 15 km southeast where the three valleys
Bødalen, Nesdalen and Kjenndalen coalesce. In these
valleys the glacier outlets from Jostedalsbreen; Bø­
dalsbreen, Ruteflotbreen and Kjenndalsbreen are lo­
cated. Strynedalen branches north from Nordfjord,
swings east, continuing about 20 km to the east, where
the valleys Hjelledalen/Videdalen and Erdalen coalesce.
Glomsdalen is a north-south oriented, hanging valley to
Strynedalen at the eastern end of Strynevatnet ( vat­
net = lake). The valleys Skjerdingsdalen, Grasdalen and
Sunndalen are tributary valleys to Hjelledalen.
Late Weichselian glacier extent, glacial
maximum and ice movements
Mapping of blockfields in the mountain areas has
demonstrated that the weathering limit is found to be
about 1750 m a.s.l. at Strynefjellet, descending to about
1500 m a.s.l. between inner Nordfjord and Sunnmøre
(Nesje et al. 1987; Rye et al. 1987, Fig. 4). The lower
limit of the blockfields is interpreted by Nesje et al.
(1987) to represent the upper limit of the Late Weichse­
lian maximum ice sheet. At that time, the glacier front
reached the edge of the continental shelf off the Møre
coast (Andersen 1979, 1981; Rokoengen 1979; Bugge
1980).
Resistance to the concept of Late Weichselian nuna­
taks has none the less persisted, most recently by Fol­
lestad (1990), who stresses the consistent pattern of
striations and till fabrics in the Nordmøre region, sug­
gesting ice movement largely independent of the local
terrain and an ice surface above the level of mountain
summits. He notes the presence of erratics and tills in
same blockfields reaching a level of at least 600-700 m.
However, Pollestad seems to have included al­
lochthonous blockfields and boulder-rich till in his
blockfield definition, and consequently his blockfield
limit is 200-300 m lower than the autochthonous
blockfield boundary described by Nesje et al. ( 1987) in
the same region. In a recent paper, Larsen et al. (1995)
conclude that the lower limit of the blockfield cannot be
taken as the upper glacial surface during the Weichselian
max1mum.
Glacial striations show that ice movements in this area
have been topographically controlled throughout the
Late Weichselian glaciation, especially during the late
phases (Fig. 5). The oldest striations in the area are
found on Langvasseggi (1600 m a.s.l.). Striations north­
west of Videdalen and Djupvassegga (1500 m a.s.l.) show
ice movements toward the northwest crossing Videdalen
and Grasdalen towards Geiranger during the most exten­
sive phase of the last glaciation (Blikra 1986). The tribu­
tary valley Glomsdalen has two northern pass-points at
1360 and 1400 m a.s.l. toward Holedalen and Hellesylt,
while Skjerdingsdalen has a pass-point of 1200 m a.s.l.
toward Flydalen and Geiranger. Hole (1985) and Blikra
(1986) mapped glacial striations in these pass-points
(Fig. 5), showing ice movements from south to north.
Lateral moraines deposited during the Younger Dryas
Chronozone, indicating an ice surface around 800 m
a.s.l. in Holedalen and around 900 m a.s.l. in Flydalen
(Kalstad 1993). This indicates that the striations mapped
in the pass-points must be older than the Younger Dryas
Chronozone. These northerly striations might possibly
represent the Late Weichselian maximum ice movement.
In two pass-points between Glomsdalen and Skjerdings­
dalen (1340 and 1260 m a.s.l.) striations show glacier
NORSK GEOLOGISK TIDSSKRIFT
Fig. 2.
77 (1997)
Location map of inner Nordfjord and the Geirangerfjord area.
Glacial geology. Nordfjord-Jostedalsbreen
53
54
N. Rye et al.
NORSK GEOLOGISK TIDSSKRIFT
2500
77 (1997)
Jostedals·
breen
2000
I
Ålfotbreen
1500
"'
'O
�
;;'
1000
500
o
I
·500
:;
""'
Longitudinal profile of Nordfjord showing altitude of
mountains on the northem and southem sides of the fjord
together with the fjord bottom. The submarine part is
adapted from Giskeødegaard (1983).
Fig. 3.
Cl
1 00
50
movements towards the west and southwest (Fig. 5). The
northern pass-point is probably so high that no active ice
movement from Skjerdingsdalen to Glomsdalen took
place during the Younger Dryas Chronozone.
The other pass-points between these two valleys are
lower and situated further south, which favour an active
ice movement from east to west. In addition, Glomsdalen
has a pass-point toward a valley between Flo and
Holedalen (1160 m a.s.l.). Glacier movements towards
the southwest out of Glomsdalen, as demonstrated by
glacial scouring, probably took place both during the
Late Weichselian glacial maximum and the Younger
Dryas Chronozone (Fig. 5) (Hole 1985).
Younger Dryas glaciation
The investigated area is situated at the borderline of the
inland ice during the Younger Dryas Chronozone. Be-
2400
Distance
(km)
yond the inland ice, local glaciation, especially lateral­
frontal moraine, was common (e.g., Reite 1967; Fareth
1970, 1987; Mangerud et al. 1979; Larsen et al. 1984).
Evidence of Younger Dryas local glaciation, especially
lateral-frontal moraine, is found in the cirque valleys
south of Holedalen (east of Hellesylt).
Fareth (1970, 1987) mapped the extent of the Nord­
fjord glaciers during the Younger Dryas readvance on
the basis of lateral moraines. In Stryn, lateral moraines
from this advance are deposited in Vikadalen (760-800
m a.s.l.) and Staurnibba (1084 m a.s.l.) (Fig. 5). On the
basis of these moraines, Fareth (1970, 1987), in his
reconstruction, indicates an ice surface at 1000-1100 m
a.s.l. in Stryn. In Olden and Loen he placed the ice
surface at 1100-1200 m a.s.l. In Olden two sets of lateral
moraines 1170 and 1120-1130 m a.s.l. on the plateau
north of Sisiliekruna may also be correlated to the
Younger Dryas lateral moraines at Skarsteinfjellet fur­
ther to the north. On the eastern side of Oldedalen,
lateral moraines at altitudes from 1080 to 1200 m a.s.l.
E
w
2200
2000
E
�
1800
Q)
"C
1600
-.;:::
1400
2
<(
1200
1ooo
8oo
mountain
peaks
_
_Skåla
_Snønipa
_
8/ockfie/d boundary
G.Jer dea ksa
1
Lodalskåpa
_Brenibba
Melheimnibba
t Sisiliekruna
Jostedals-breen
Younger Dryas glacier surface
L----St ryn
Olden
Loen
Hjelle
600 4-------��-----r------�-------.------�-------,.-------�----__,
40
1 o
20
30
o
Distance (km)
Fig. 4.
The blockfield boundary in ioner Nordfjord.
NORSK GEOLOGISK TIDSSKRIFf
77 (1997)
Glacial geology. Nordfjord-Jostedalsbreen
20km.
10
o
�------��Contour inter val 200 m
Leg end:
=
'"...,
Marginal moraine depos�ed by fjord
or valley glaciers
n
Mounds and ridges
- Glacial striation, movement towards the observation point
Glaciofluvial depos� with terrace slope
-- Crossing glacial striations, increasing number of
-... ticks indicate increasing relative age.
B
Glaciofluvial deposit
1120
Bs
Glaciolacustrine deposit
••
Esker
�
Fig. 5.
Submarin marginal moraine
Glacial geology of inner Nordfjord and adjacent district.
eoo
Pass·point
Contour interval
55
56
NORSK GEOLOGISK TIDSSKRIFT
N. Rye et al.
77 (1997)
6. Reconstruction of the valley glaciers in inner Nord­
fjord during formation of the early Preboreal moraines.
Fig.
are mapped and correlated with the Younger Dryas
marginal moraines (Fig. 5). In Fosdalen, east of Loen,
lateral moraines at 1020-1100 m a.s.l. (Lien 1985; Nesje
& Dahl 1992), indicate a somewhat lower ice surface
than that suggested by Fareth.
Extrapolation of the Younger Dryas lateral moraines
from the head of the fjord toward Jostedalsbreen this
gives an ice surface of 1300-1400 m a.s.l. in the inner
parts of the valleys, suggesting that fairly steep glacier
falls may have existed along the western margin of the
Jostedalsbreen plateau. North of Flo at Strynevatnet, the
pass-point toward Hellesylt is situated at 540 m a.s.l. In
his reconstruction, Fareth (1970, 1987) placed a Younger
Dryas valley glacier through the pass-point to a location
4 km north of Hellesylt, where Giskeødegaard (1983)
mapped a 100 m high submarine frontal deposit. This
reconstruction is not, however, based on any lateral
moraines. In Holedalen, Kalstad (1993), however,
mapped prominent lateral moraines around 800 m a.s.l.
(Fig. 5). In addition, marginal moraines were deposited
by local glaciers in the small cirques south of Holedalen.
At the mouth of the Geirangerfjord, Giskeødegaard
(1983) mapped a submarine end moraine, and in the
valley sides further to the east (Fig. 5) are several promi­
nent lateral moraines which can be correlated with this
end moraine. In Flydalen the corresponding lateral
moraines are situated 900 m a.s.l., showing that there
was possibly no active glacier moving across the pass­
point from Skjerdingsdalen to Geiranger during the
Younger Dryas Chronozone. Lateral moraines north of
Dalsnibba, descending from 1300 to 1000 m a.s.l. at a
distance of l km, show that the glacier had a steep
surface profile toward Geiranger.
Marginal moraines in the valley bottom and the valley
sides in the Tafjord area (Fig. 5) were deposited during
the late Younger Dryas Chronozone and early Preboreal
Chronozone (Eikenfæs 1991).
The ice surface along the Stryn valley was probably ca.
1300 m a.s.l. south of Glomsdalen during the Younger
Dryas. With pass-points at 1360-1400 m a.s.l. to the
north, no active ice movement toward Holedalen could
have taken place. On the other hand, the pass-points
NORSK GEOLOGISK TIDSSKRIFT
Table l.
Glacial geology. Nordfjord-Jostedalsbreen
77 (1997)
57
Radiological datings. For other radiological datings in this area, see Fig. 12 in Rye et al. (1987).
No.
Lab. no.
2
3
4
5
6
7
T-5812
T-5811
T-5606B
T-5810A
T-616
T-4839
T-4234
Matr.
14C-age
Refr.
Peat
9340± 130
8080± 60
9260± 140
9030± 100
9390± 200
8810± 130
2450± 40
(1987)
(1986)
Kvarnme (1984)
Lien (1985)
Fareth (1970)
Nesje (1984)
Peat
Peat
Peat
Shells
Limnic sediments (gyttja)
Coal (forest)
north of Glomsdalen and Skjerdingsdalen were domi­
nated by local glaciers merging with the valley glaciers.
In Glomsdalen this might have resulted in a fairly hori­
zontal ice surface. During this period glaciers drained
from Skjerdingsdalen to Glomsdalen (pass-point 1260 m
a.s.l.) and from Glomsdalen to the valley between Flo
and Hellesylt (pass-point 1160 m a.s.l.) (Fig. 5).
In the mountain areas between Oldedalen, Lodalen
and Strynedalen, extensive areas are presently covered by
cirque- and plateau glaciers. During the Younger Dryas
these glaciers probably had outlet glaciers merging with
the valley glaciers from the inland ice, but no deposits
have been found which can confirm or refute this, be­
cause these areas lay above the equilibrium-line altitude
(ELA) at that time.
During the Younger Dryas the ice surface reached
about 1600 m a.s.l. at Strynefjellet, based on an interpre­
tation of glacial scouring. At this time, the glacier front
was located at Anda-Lote, in the middle part of Nord­
fjorden. When the Preboreal end moraines at Olden,
Loen and Stryn were deposited, the glacier surface
reached about 1200-1300 m a.s.l. in the inner parts (Fig.
6). Steep glacier falls may have existed between Jostedals­
breen and the valley glaciers. This illustrates that the
elevation of the ice surface in inner parts of the valleys
shows little variation while the response at the glacier
fronts was extensive. This 'hinging'-effect is clearly
demonstrated in a smaller scale at the recent glaciers,
e.g., the Bødalsbreen glacier (Lien 1985). In Fosdalen, an
easterly tributary to Loen, Nesje & Dahl ( 1992) calcu­
lated the equilibrium-line depression at 425 m during the
Younger Dryas. In Innvik (Fig. 1), Dahl & Nesje (1992)
calculated the Younger Dryas ELA depression at about
500 m. They also calculated that the winter precipitation
was reduced to about 60% of the present values.
Deglaciation after the Younger Dryas
Chronozone
Subsequent to the Younger Dryas glacier advance, a
period of rapid retreat of the Nordfjord glacier occurred,
probably as a result of extensive calving in the deep
fjords and in the deep lake, Hornindalsvatn. A radiocar­
bon date on peat west of Stryn gives a minimum age of
the deglaciation of the pass-point Markane, between
Rye et al.
Blikra
Kvamme & Randers
(1982)
Innvikfjorden and Lake Hornindalsvatn (Fig. l) of
9340 ± 130 14C yr BP (Table l, T-5812). Marginal de­
posits are mapped at the mouths of Strynedalen, Lodalen
and Oldedalen (Fig. 5). These marginal deposits are
relatively prominent and were deposited during the early
part of the Preboreal Chronozone (Fig. 6). In Stryn, two
ice-marginal deposits are mapped, the Vinsrygg moraine
and the Årheim terrace (Fareth 1970, 1987; Nesje 1984;
Rye et al. 1987). In Loen, a submarine and a supra­
marine terminal moraine were deposited (Lien 1985). In
Olden two marginal deposits at Melheim-Løken and
Eide are mapped.
In previous work, Preboreal marginal deposits have
been thought to indicate a climatic deterioration (e.g.,
Vorren 1973; Bergstrøm 1975), white recent investiga­
tions (Kjenstad & Sollid 1982; Sollid & Reite 1983; Anda
1984; Rye et al. 1987) include glaciodynamic principles as
a possible explanation for their formation. Thus, as a
result of rapid calving in Nordfjord, the glacier had a
steep and dynamically unstable profile at the front. As
the three valley glaciers in Stryn, Loen and Olden be­
came grounded on bedrock thresholds or in a narrow
portion of the valley, the frontal retreat ceased or became
strongly reduced. In order to achieve dynamic stability,
the glaciers advanced in order to adjust their surface
profile to the new dynamic conditions. The glaciody­
namic theory as a model for these marginal deposits
makes it impossible to correlate them chronostratigraph­
ically. According to the glaciodynamic model, the mar­
ginal deposits at Stryn (Vinsrygg) were probably
deposited somewhat earlier than the deposits in Olden
and Loen.
Stryn
East of Stryn, Fareth (1970, 1987) mapped prominent
marginal moraines deposited by a valley glacier in
Strynedalen (Fig. 5). In the valley sides to the east,
lateral moraines 200-300 m below the Younger Dryas
lateral moraines have been mapped and correlated with
the marginal moraines at Stryn, which he termed the
Vinsrygg moraines. Fareth (1970, 1987), in his recon­
struction, placed the icefront during the Vinsrygg event
across Nordfjorden northeast of Innvik. Recently, the
corresponding ice-marginal deposit has been found on
the eastern side of the Stryn bay. Ir. our opinion, there-
58
NORSK GEOLOGISK TIDSSKRIFT
N. Rye et al.
fore, the glacier front was a calving front located in the
bay at Stryn when the Vinsrygg moraine was formed.
Fareth ( 1970, 1987) interpreted a terrace lying 69 m
a.s.l. at Øvreeide, in the western end of Strynevatnet as
an ice-frontal deposit (Fig. 5), which he correlated to the
Eide moraine in Olden. Stokke ( 1980) reinterpreted the
deposit to be a kame terrace built up to 65 m a.s.l. (Fig.
5). No lateral moraines are recognized in connection with
this deposit.
Lateral moraines thought to have been deposited during
the Preboreal Chronozone have been mapped in Sun­
ndalen, Hjelledalen, and Glomsdalen between 900 and
1200 m a.s.l. (Fig. 5) (Rye et al. 1984; Hole 1985). The
deposits are, however, too scattered to justify finn corre­
lations with any frontal deposits at Stryn. In G1omsdalen,
Skjerdingsdalen, Grasdalen, Hjelledalen and Erdalen there
are deposits (mounds and ridges) showing that the last part
of deglaciation took place as vertical down-wasting of
stagnant ice remnants in Strynevatnet and Hjelledalen
(Rye et al. 1984, 1987; Nesje 1984; Hole 1985).
Hjelledalen
At Hjelle, glacioftuvial terraces with a relatively horizon­
tal surface were built up to 75 m a.s.l. (Stokke 1982; Rye
et al. 1984) (Fig. 5). Kaldhol (1912) described these
terraces noting large boulders along the edge of the
terrace toward Strynevatnet which he interpreted as a
result of an ice remnant in the Strynevatnet basin during
terrace deposition. Thus, the terrace slope toward Stryn­
evatnet was believed to be an ice-contact slope. At
Vollsnes, northeast of Hjelle, a section in another terrace
built up to at least 70 m a.s.l. shows glacioftuvial gravel
and stones with westward dipping foresets. Covering the
foreset beds is l - 2 m of horizontal, laminated fine sand
and silt at the top. In the eastem part of the terrace,
segments and blocks of bedded glacioftuvial material
have been tilted and thrust by glaciotectonic activity.
About 3-4 km further east in Hjelledalen several
terraces were built up to about 75 m a.s.l., mainly along
the northem valley side (Fig. 5). Rye et al. ( 1984), Nesje
( 1984), and Hole ( 1985) concluded that the glacier in
Strynevatnet melted down vertically. The terraces at
Hjelle were probably deposited into a freshwater lake
with a communicating water level to the sea further west
while there still was an ice remnant in the Strynevatnet
basin. The deposits at Vollsnes show that the valley
glacier in Hjelledalen advanced across the glacioftuvial
terrace eroding the top and leaving a layer of till. The
fine-grained sediments on top were deposited as the
glacier retreated. The tilted units of glacioftuvial material
are related to another glacier advance shortly afterwards.
Stokke (1982) interpreted the terraces in the eastem part
of Hjelledalen to be remnants of a glacioftuvial deposit
filling the entire valley. The evidence indicating glacier
advances and subsequent sedimentation of fines shows,
however, that this cannot have been the case. If
Hjelledalen were filled up with glacioftuvial deposits, one
77 (1997)
would expect the terraces in the eastem part to be higher
than in the west, but this is not the case. The terraces are
interpreted to be lateral deposits built up along down­
wasting ice remnants in the valley bottom. Such a rapid
down-wasting could have been the result of deglaciation
of the pass-point at Videdalen-Strynefjellet.
At Skora a small esker and a relatively large glacioftuvial
lateral terrace are present (Fig. 5). Deep meltwater chan­
nels in the terrace are evidence for lateral meltwater
drainage. The terrace is interpreted to be a lateral deposit
to an ice remnant in Hjelledalen deposited after deglacia­
tion of the pass-point at Strynefjellet-Videdalen. At the
pass-point in Videdalen, eskers and hurnmocky moraine
were deposited.
Breidalen
In Breidalen the pass-point at Djupvatnet (1160 m a.s.l.)
toward Geiranger was deglaciated while there still was ice
to the east. As a result, a glacier-dammed lake developed
in the eastem part of Djupvatnet, where well-sorted fine
sand and silt was deposited. At Langvatnet further east in
Breidalen, eskers and hummocky moraine were formed,
indicating a vertically down-wasting, dynamically inactive
glacier.
Glomsdalen
In Glomsdalen Hole (1985) mapped lateraljsublateral
glacioftuvial deposits at different altitudes ranging altitudi­
nally between 1200 m in the north and 300 m in the south
(Fig. 5). Lateral glacioftuvial terraces were deposited in two
smaller tributary valleys east of Glomsdalen at altitudes
between 1200 and 900 m. On the valley bottom in the
northem part of Glomsdalen there are hummocky deposits
with a complex composition (Hole 1985). Further south
glacioftuvial deposits as eskers and kames are common
(Fig. 5). In the mouth of Glomsdalen fine material has been
deposited in a glacier-dammed lake about 400 m a.s.l. The
pass-points in the north end of Glomsdalen at 1360 and
1400 m a.s.l. must have been deglaciated at the end of the
Younger Dryas Chronozone, while a glacier was still
moving across the pass-point (1260 m a.s.l.) from Skjerd­
ingsdalen in the east. As this pass-point was itself deglaci­
ated, the last part of the deglaciation in G lomsdalen took
place as vertical down-wasting. Finally, the ice surface
sloped up-valley, and a glacial lake was formed in the
southem part. Small glacioftuvial lateral terraces 300 m
a.s.l. at the mouth of Glomsdalen were deposited along the
margin of the ice remnant in the Strynevatn basin (Hole
1985). At Strynevatnet a glacioftuvial terrace was formed
at a water level of about 75 m a.s.l.
Grasdalen f Skjerdingsdalen
Blikra (1986) mapped glacial striations indicating that
glaciers moved from Breidalen over Oppljosvatn and
NORSK GEOLOGISK TIDSSKRIFT
Glacial geology. Nordfjord-Jostedalsbreen
77 (1997)
down Grasdalen (Fig. 5), possibly during the Younger
Dryas Chronozone. At the same time, ice flowed across
Grasdalsvatnet through a pass-point 1440 m a.s.l. toward
Videdalen/Hjelledalen in the south. During this period
the ice surface must have been about 1600 m a.s.l. in the
area around Oppljosvatnet. Below 1500 m a.s.l. indica­
tions of glacial meltwater drainage are recognized in the
form of rounded boulders and stones in pass-points and
along mountain slopes (Blikra 1986). The pass-point
south at Oppljosvatn ( 1440 m a.s.l.) was one of the first
to be deglaciated. The closing of this pass-point may
have led to the deposition of a lateral moraine 1420 m
a.s.l. south of Oppljosvatn.
Southerly glacial striations in the pass-point between
Djupvassegga and Oppljosegga (1340 m a.s.l., Fig. 5)
may have been formed at the end of the Younger Dryas
or early Preboreal Chronozones before the pass-point
was closed. When the pass-point northwest of Grasdals­
vatnet (1240 m a.s.l.) was closed, all ice passing Oppljos­
vatnet flowed down Grasdalen. In this phase of
deglaciation the glacier in Skjerdingsdalen melted down
vertically. Lateral deposits between 600 and 1000 m a.s.l.
indicate that the glacier surface was somewhat lower in
the south than in the north. This demonstrates that the
valley glacier flowing through Grasdalen was not very
dynamically active. The presence of hummocky ablation
till in the northern part of Skjerdingsdalen demonstrates
that the glacier was almost dynamically inactive at the
final stage of deglaciation. When the glacier in Skjerd­
ingsdalen had melted down, a valley glacier still flowed
down Grasdalen. A prominent lateral-frontal moraine
deposited where Grasdalen and Skjerdingsdalen coalesce
illustrates this. The Grasdalen valley glacier was not
connected to the Videdalen/Hjelledalen glacier at this
time. This is proved by glaciolacustrine sediments de­
posited in the southern part of Skjerdingsdalen. The lake
was dammed by the down-wasting glacier in Videdalen/
Hjelledalen.
Glaciofluvial lateral terraces east of Grasdalsvatnet
show that the lake was dammed in a period during the
melting of the valley glacier. When the pass-point north­
west of Grasdalsvatnet was deglaciated, the glacier
moved through Grasdalen. This is demonstrated by
glacial striations east of Grasdalsvatnet (Fig. 5). When
the pass-points at Oppljosvatnet were deglaciated, the
deglaciation in Grasdalen was characterized by vertical
down-wasting (Blikra 1986). A radiocarbon date from
peat in Skjerdingsdalen gives a minimum deglaciation
age of 8080 ± 60 14C yr BP (Table l, T-5811).
Sunndalen
No evidence of lateral meltwater drainage is recognized in
Sunndalen, but possible deposits may have been removed
by subsequent Holocene avalanche activity. No terraces
were built up to the marine limit at the mouth of
Sunndalen. This indicates that the pass-point to Sunndalen
59
(1240 m a.s.l.) was deglaciated earlier than the sea could
penetrate into Hjelledalen. The lower part of Sunndalen
was probably covered by an ice remnant in connection with
the glacier occupying Hjelledalen, thus preventing deposi­
tion of a terrace at the mouth of Sunndalen. In Sunndalen
a radiocarbon date on peat yielded 9260 ± 140 14C yr BP
(Table l, T-5606 B) giving a minimum date of the
deglaciation in this area (Kvamme 1984).
Erdalen
In Erdalen Nesje (1984) mapped glaciofluvial deposits at
altitudes ranging from 690 to 130 m a.s.l. (Fig. 5). He
interpreted these as lateral/sublateral terraces deposited
along the edge of a down-wasting glacier with a relatively
horizontal surface. In Vetledalen, a tributary valley to
the northeast of Erdalen, a glaciofluvial terrace was
deposited 690 m a.s.l., dammed by a valley glacier in
Erdalen. In the distal part of the terrace, the glacier
occupying Erdalen pushed up a lateral moraine. At that
time, there was no connection between the glacier in
Erdalen and the glacier in Vetledalen. Nesje (1984) ex­
plained this by ice supply from an ice culmination east of
the Jostedalsbreen plateau. According to radioecho
soundings on Jostedalsbreen (Sætrang & Holmquist
1987), the subglacial bedrock pass toward Vetledalen is
dose to 1500 m a.s.l., while the pass-point to Erdalen is
between 1300 and 1400 m a.s.l. As a result, Vetledalen
did not receive ice from the east, while the lower pass­
point toward Erdalen allowed ice to drain across the
pass-point into Erdalen. If the glacier movement down
Erdalen was climatically determined, one might assume
that the Vetledalsbreen glacier had merged with the
glacier in Erdalen. This indicates that Jostedalsbreen did
not play a dominant role in ice supply to Erdalen in this
late phase of the deglaciation, except during the Erdalen
event (Holmquist 1987, p. 27).
Loen
At the fjord bottom 1.5 km west of the Loen village,
there is a 15 m high submarine ridge covered by boulders
(Fig. 5). No lateral moraines have been found attributed
to this ridge, but it is interpreted as a terminal moraine
deposited by a valley glacier in Lodalen (Lien 1985). On
the northern side of the village of Loen a 4-5 m high
terminal moraine was deposited. South of Loen a lateral
moraine was deposited starting at the valley bottom and
terminating at a level of 230 m a.s.l. on the valley side.
This lateral moraine is correlated to the terminal moraine
north of the Loen village. At Sæten, in the northwestern
part of Lovatnet, a glaciofluvial terrace was formed
86-89 m a.s.l. The terrace is regarded as a frontal
deposit, built up during a halt in the deglaciation. A
radiocarbon date on peat east of Loen gives a minimum
age for the terminal moraines at Loen of 9030 ± 100 14C
yr BP (Table l, T-5810 A).
60
NORSK GEOLOGISK TIDSSKRIFT
N. Rye et al.
77 (1997)
120 �------.
Legend:
100
E
Q)
"O
:::s
-
1!1
80
•
Younger Dryas
•
Hornindai/Utfjorden
+
cc
•
·... .
•
Loen
60
•
o
• •
c
�
1111
c
c
"
1111
Stryn
;;
C(
The Nor moraines
Tapes
40
Olden
20
o
o
20
60
40
BO
100
120
Distance (km)
Fig. 7.
Shoreline diagram for Nordfjord showing the Younger Dryas and Tapes marine levels in Nordfjord. Adapted from Fareth (1987) and Rye et al. (1987).
In the tributary valleys of Lodalen; Fosdalen, Breng
and Austerdalen (Fig. 5), lateral moraines were deposited
at levels between 300 and 1120 m a.s.l. The lateral
moraines at the mouth of Fosdalen 460-500 m a.s.l. can
probably be correlated with the terminal moraine in Loen.
The other lateral moraines have a scattered distribution,
and it is therefore difficult to correlate them with a
particular terminal moraine in Loen. Bødalen, east of
Lodalen, has a marked 'break' in the longitudinal profile
580 m a.s.l. (Fig. 5). In the lower part of the valley, below
500 m a.s.l., eskers and lateral/sublateral glaciofluvial
deposits have been mapped (Lien 1985). A lateral deposit
in Nesdalen indicates that at this time, an active glacier
was still moving down Bødalen merging with the glacier
in Lodalen. Eskers and lateral deposits in the lower part
of Bødalen show that in a later phase the Bødalen glacier
was isolated from the glacier in Lodalen, probably at the
'break' 580 m a.s.l. In the lower part of Bødalen the ice
melted down vertically as deglaciation in Lodalen pro­
ceeded. In Nesdalen a significant, narrow terraced mar­
ginal moraine is situated in the eastern valley side (Fig. 5).
It is located 600m a.s.l. descending to 550 m a.s.l. at the
valley bottom further to the south. Beneath this terrace
the thick till cover is heavily gullied. The terrace is
interpreted to represent the upper surface of a valley
glacier in Lodalen in a late phase of the deglaciation when
the connection to the glacier in Nesdalen had been cut off.
At that time the ice surface in Lodalen at the mouths of
Nesdalen and Bødalen was at least 600 m a.s.l. The last
ice remnants in Nesdalen melted down as stagnant ice.
As a result of the steep valley sides between 100 and
1300 m a.s.l. in Kjenndalen, no deposits from the deglaci­
ation phase have been preserved. The deglaciation in
Nesdalen and Bødalen indicates that ice supply from
Jostedalsbreen through Kjenndalen may have been main­
tained longer than in Nesdalen and Bødalen. At the
mouth of Bødalen a glaciofluvial terrace is built up to l 00
m a.s.l. This terrace has probably been deposited in a
glacier-dammed lake (Lien 1985).
Olden
About 2 km south of the village of Olden, a large
marginal deposit at Melheim-Løken is presently
damming Lake Floen. About l km south another mar­
ginal deposit was formed at Eide in front of Oldevat­
net (Fareth 1970, 1987). Shells found at Håheim,
proximal to the Eide moraine are radiocarbon dated to
9390 ± 200 14C yr BP (Table l , T-616), thus giving a
minimum age of the Eide and Melheim/Løken
moraines.
In Sundsdalen east of Oldevatnet lateral moraines,
which probably can be correlated to the Melheim/Løken
moraine, were deposited 1040 m a.s.l. (Fig. 5). As a result
of the steep valley sides, no other lateral moraines are
recognized in Oldedalen, and those which exist are too
scattered to make any reliable correlations. Terraces at
Åbrekk and Melkevoll at the inner part of Oldedalen
(Fig. 5) reach as high as 86 and 100m a.s.l., respectively,
thus indicating a rapid retreat of the glacier in Oldedalen.
Degree of rock surface weathering
The degree of rock surface weathering was measured at
sites in Oldedalen and Brigsdalen and on an altitudinal
transect from Loen to Skåla (McCarroll & Nesje 1993).
The Schmidt hammer was useful only for distinguishing
sites covered during 'the Little lee Age' from those
deglaciated during the Lateglacial and early Holocene.
Roughness of granitic augen gneiss surfaces was
quantified from profiles measured in the field using a
micro-roughness meter and profile gauge. In the western
valley slope below Skåla there is a significant increase in
surface roughness above a distinct trimline at ca. 1350 m.
However, there was no significant increase above the
higher blockfield boundary at ca. 1560 m. The vertical ice
limits at Skåla await 10Be and 26Al exposure dating
(Brook et al. in prep.).
NORSK GEOLOGISK TIDSSKRIFT
77 (1997)
Glacial geology. Nordfjord-Jostedalsbreen
w
E
8
o.:
ai
9
o
o
o
E 10
f!
Local
glaciation
as
Q)
>.
c:
o
11 .
-e
as
o
o
:c 12
as
a:
l
13
o
Fig. 8.
20
40
60
80
Distance (km)
100
120
140
Time/distance diagram for the deglaciation of Nordfjord.
Sea level
Fareth (1970, 1987) made a shoreline diagram on the basis
of marine terraces in Nordfjord. This was later somewhat
modified by Rye et al. (1987, Fig. 7). By extrapolating the
Younger Dryas shoreline, this reaches 11O m a.s.l. at the
inner part of Oldedalen and Lodalen. The terraces at
Melkevoll and Bødal were built up to 100 m a.s.l. This
shows that the deglaciation of the inner fjord- and valley
areas subsequent to the Younger Dryas was very rapid, or
that there has been a period of stable sea level, as recorded
at Sunnmøre (Lie and Lømo 1981; Lie et al. 1983; Svendsen
1985; Svendsen & Mangerud 1987). Lien (1985) discussed
the possibility that the terrace 100 m a.s.l. in Bødalen may
have been deposited in a glacier-dammed lake with an ice
remnant in Lovatnet, but it has not been possible to
determine which alternative is correct. So far, no sign of
vertical down-wasting of the glacier in Oldedalen has been
found. The lack of terraces at the marine limits at the
mouth of Erdalen (Nesje 1984) and Sunndalen is inter­
preted to be a result of down-wasting ice remnants in the
basin of Strynevatnet, while higher-lying pass-points were
already deglaciated.
The Erdalen event
At Vetledalssetra a prominent terminal moraine is de­
posited across the valley bottom in Erdalen, built up by
a glacier draining down Erdalen (Fig. 5). A radiocarbon
date on gyttja proximal to the terminal moraine gave a
minimum age of 8810 ± 130 14C yr BP (Table l , T-4839,
Nesje 1984). The end moraine is situated about l km distal
to the maximum extent of Erdalsbreen during the 'Little
lee Age'.
This end moraine is deposited during readvances of
glaciers at the Jostedal plateau, caused by a short climatic
deterioration, which probably happened very late in the
61
Preboreal Chronozone. (Rye et al. 1987). During this stage
the mean ELA depression has been calculated to 325 m.
On the proximal side of the terrace at Melkevoll is an
end moraine deposited by Melkevollsbreen. In addition,
two closely spaeed end moraines are deposited at the
mouth of Brigsdalen, a tributary valley to Oldedalen.
These two end moraines are deposited by Brigsdalsbreen
(Pedersen 1976). At both localities, the end moraines are
located well beyond the moraines of the 'Little lee Age',
formed at about AD 1750. Just above the 'break' in
profile in Bødalen, prominent marginal moraines were
deposited 600 m a.s.l., at a position 1-1. 5 km beyond the
maximum extent of Bødalsbreen during the Little lee
Age. A radiocarbon date on coal from a forest fire
proximal to the moraine ridges gives a minimum age of
2450 ± 40 14C yr BP (Table l , T-4234) (Kvamme &
Randers 1982).
Thus, in front of Erdalsbreen, Bødalsbreen, Brigsdals­
breen and Melkevollsbreen terminal moraines have been
mapped distally of the maximum extent of these glaciers
during the 'Little lee Age'. Terminal moraines of possi­
bly the same age have been mapped east of Jostedals­
breen also (Elgersma & Nesje 1978; Aa & Sønstegaard
1987), and they are probably the result of a short-lived
climatic deterioration and reactivation of glaciers on the
Jostedalsbreen plateau at the end of the Preboreal
Chronozone. This glacier readvanee has been termed the
Erdalen event (Nesje 1992).
The Holocene
Lithostratigraphic and palaeobotanical studies show that
during the Hypsithermal (ca. 8000-6000 yr BP) the ELA
was about 400-500 m higher than at present. As a result,
Jostedalsbreen probably disappeared entirely during that
period. The glacier formed again about 5300 yr BP. The
first significant glacier advance occurred between 3700 and
3100 yr BP. The ELA intersected the mean modem
elevation five times from ca. 2600 yr BP to the present.
The 'Little lee Age'
Torsnes et al. (1993) calculated the modem and 'Little lee
Age' ELA of 20 valley glaciers from Jostedalsbreen. Using
an accumulation area ratio (AAR) of 0.6 ± 0.05 yielded
a mean 'Little lee Age' ELA depression of 70 m. By using
lichenometric evidence, Bickerton & Matthews (1993)
demonstrated that the maximum 'Little lee Age' advances
at seven outlet glaciers occurred beween AD 1741 and
1863.
Conclusions
In Nordfjord, the fronts of the Younger Dryas fjord/valley
glaciers were located in the fjord at Sandane, at
62
NORSK GEOLOGISK TIDSSKRIFT
N. Rye et al.
Anda/Lote, and at the western end of Homindalsvatn.
Prominent lateral moraines along the fjord delimit the
glacier during this stage, the altitude of the glacier at the
fjord head in inner Nordfjord being 1000-1100 m above
sea level. Beyond and above the Younger Dryas valley
and fjord glaciers in Nordfjord, local glaciers were
formed, of which the glacier covering the Ålfoten area
was the most extensive. The Younger Dryas ELA depres­
sion in middle and inner Nordfjord has been calculated
at 450 ± 50 m, while the winter precipitation was reduced
to about 60% compared to the present.
During the early Preboreal, the fjord glaciers retreated
rapidly to the head of the fjord (Fig. 8), forming two ice
marginal deposits in each of the main valleys in inner
Nordfjord. The final deglaciation in lower-lying valleys
in Loen and Stryn was characterized by vertical down­
wasting in the bedrock basins presently occupied by the
lakes Lovatnet and Strynevatnet, respectively. Terminal
moraines located up to l km beyond the 'Little lee Age'
moraines surrounding Jostedalsbreen, dated at 9100 ±
200 14C yr BP, have been termed the Erdalen event after
the type site in Erdalen. The mean ELA depression
during this stage has been calculated to 325 m, while the
mean winter precipitation was reduced to about 70% of
modem values. During the 'Little lee Age', marginal
moraines were formed in front of the outlet glaciers from
Jostedalsbreen and by local cirque glaciers. The most
representative 'Little lee Age' ELA depression in the
Jostedalsbre region is calculated to be 150 m (Nesje et al.
1991), while Torsnes et al. (1993) calculated the average
'Little lee Age' ELA depression of 20 outlet glaciers from
the Jostedalsbre ice cap as 70 m by means of the AAR
approach.
- This work is part of 'Fjordaneprosjektet' led by N. Rye
with financial support from the University of Bergen. The work was also made
possible by financial support from the Norwegian State Power System (Statkraft)
in connection with the planning of hydroelectric power development in the inner
Nordfjord area (Breheimen-Stryn). Parts of the investigations were carried out in
connection with this planning. A. Reite kindly read the manuscript critically and
suggested many improvements. Comments by reviewers improved the clarity of
the paper. E. King corrected the English text of the final manuscript. E. Lier and
J. Ellingsen drew the figures and E. Loodtz typed the manuscript. To these
persons and institutions we proffer our sincere thanks.
Acknowledgements.
Manuscript received May 1995
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