Deglaciation of Fennoscandia
Arjen P. Stroeven a, b, *
, Clas H€attestrand a, b
, Johan Kleman a, b
, Jakob Heyman a, b
,
Derek Fabel c
, Ola Fredin d, e
, Bradley W. Goodfellow b, f, g
, Jonathan M. Harbor a, b, h
,
John D. Jansen a, b, i
, Lars Olsen d
, Marc W. Caffee h, j
, David Fink k
, Jan Lundqvist a, b
,
Gunhild C. Rosqvist a, b, l
, Bo Str€omberg a, b
, Krister N. Jansson a, b
a
Geomorphology and Glaciology, Department of Physical Geography, Stockholm University, Sweden
b
Bolin Centre for Climate Research, Stockholm University, Sweden
c
SUERC-AMS, Scottish Universities Environmental Research Centre, East Kilbride Scotland, UK
d
Geological Survey of Norway, Trondheim, Norway
e
Department of Geography, Norwegian University of Science and Technology, Trondheim, Norway
f
Department of Geological Sciences, Stockholm University, Sweden
g
Department of Geology, Lund University, Sweden
h
Department of Earth, Atmospheric, and Planetary Sciences, Purdue University, West Lafayette, USA
i
Institute of Earth and Environmental Science, University of Potsdam, Germany
j
Department of Physics and Astronomy/Purdue Rare Isotope Measurement Laboratory, Purdue University, West Lafayette, USA
k
Australian Nuclear Science and Technology Organization, PMB1, Menai, Australia
l
Department of Earth Science, University of Bergen, Norway
a r t i c l e i n f o
Article history:
Received 1 May 2015
Received in revised form
28 August 2015
Accepted 14 September 2015
Available online xxx
Keywords:
Fennoscandian Ice Sheet
Deglaciation
Glacial geomorphology
Geochronology
Ice sheet dynamics
a b s t r a c t
To provide a new reconstruction of the deglaciation of the Fennoscandian Ice Sheet, in the form of
calendar-year time-slices, which are particularly useful for ice sheet modelling, we have compiled and
synthesized published geomorphological data for eskers, ice-marginal formations, lineations, marginal
meltwater channels, striae, ice-dammed lakes, and geochronological data from radiocarbon, varve,
optically-stimulated luminescence, and cosmogenic nuclide dating. This is summarized as a deglaciation
map of the Fennoscandian Ice Sheet with isochrons marking every 1000 years between 22 and 13 cal kyr
BP and every hundred years between 11.6 and ﬁnal ice decay after 9.7 cal kyr BP.
Deglaciation patterns vary across the Fennoscandian Ice Sheet domain, reﬂecting differences in climatic
and geomorphic settings as well as ice sheet basal thermal conditions and terrestrial versus marine
margins. For example, the ice sheet margin in the high-precipitation coastal setting of the western sector
responded sensitively to climatic variations leaving a detailed record of prominent moraines and other
ice-marginal deposits in many fjords and coastal valleys. Retreat rates across the southern sector differed
between slow retreat of the terrestrial margin in western and southern Sweden and rapid retreat of the
calving ice margin in the Baltic Basin. Our reconstruction is consistent with much of the published
research. However, the synthesis of a large amount of existing and new data support reﬁned reconstructions
in some areas. For example, the LGM extent of the ice sheet in northwestern Russia was
located far east and it occurred at a later time than the rest of the ice sheet, at around 17e15 cal kyr BP.
We also propose a slightly different chronology of moraine formation over southern Sweden based on
improved correlations of moraine segments using new LiDAR data and tying the timing of moraine
formation to Greenland ice core cold stages.
Retreat rates vary by as much as an order of magnitude in different sectors of the ice sheet, with the
lowest rates on the high-elevation and maritime Norwegian margin. Retreat rates compared to the climatic
information provided by the Greenland ice core record show a general correspondence between
retreat rate and climatic forcing, although a close match between retreat rate and climate is unlikely
because of other controls, such as topography and marine versus terrestrial margins. Overall, the time
slice reconstructions of Fennoscandian Ice Sheet deglaciation from 22 to 9.7 cal kyr BP provide an
* Corresponding author.
E-mail address: arjen.stroeven@natgeo.su.se (A.P. Stroeven).
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journal homepage: www.elsevier.com/locate/quascirev
http://dx.doi.org/10.1016/j.quascirev.2015.09.016
0277-3791/© 2015 The Authors. Published by Elsevier Ltd. This is an open access article under the CC BY-NC-ND license (http://creativecommons.org/licenses/by-nc-nd/4.0/).
Quaternary Science Reviews xxx (2015) 1e31
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10.1016/j.quascirev.2015.09.016
important dataset for understanding the contexts that underpin spatial and temporal patterns in retreat
of the Fennoscandian Ice Sheet, and are an important resource for testing and reﬁning ice sheet models.
© 2015 The Authors. Published by Elsevier Ltd. This is an open access article under the CC BY-NC-ND
license (http://creativecommons.org/licenses/by-nc-nd/4.0/).
1. Introduction
Melting of the Greenland and Antarctic ice sheets, and the threat
of accelerated melt in response to future climate warming, has
ﬁrmly positioned ice sheet deglaciation processes and rates on the
global research agenda (Warrick and Oerlemans, 1990; Briner et al.,
2009; Church et al., 2013; Stokes et al., 2014). This is because an
important implication of accelerated ice sheet melt, in addition to
ice sheet mass loss through calving, is an expected rise in global
mean sea level, with spatial variations around that mean (Milne
et al., 2009; Kopp et al., 2010; Slangen et al., 2014) and resulting
challenges for coastal land use. The current condition of the
Greenland and Antarctic ice sheets, grown out of highlands but also
covering extensive lowlands and subglacial basins below sea level,
is similar to the situation of the former Laurentide and Fennoscandian
ice sheets at their last maximum positions, and implies ice
sheet retreat with margins extending offshore. Future retreat patterns,
if recent trends persist, will likely differ starkly for margins
that are predominantly terrestrial and those that are terminating in
a marine environment. The latter are prone to destabilization and
run-away effects through sea level rise and margin thinning
(Hughes, 1975; Favier et al., 2014). This insight has been gained
from the dynamics and deglaciation histories of the former
Northern Hemisphere ice sheets (Kleman and Applegate, 2014) and
from measurements and modelling pertaining to the Greenland
and Antarctic ice sheets (Joughin et al., 2014; Rignot et al., 2014).
At the height of glaciation, during the global Last Glacial
Maximum (LGM, 26.5-20 thousand years ago [cal kyr BP]; Clark
et al., 2009b), a considerable portion of the Northern Hemisphere
landmass above 60 N was ice-covered (Denton and Hughes, 1981).
Reconstructions of the maximum extent and the timing of initial
retreat of these Northern Hemisphere ice sheets has been a
research focus for the last 175 years (Agassiz, 1840; Torell, 1872,
1873; Jackson and Clague, 1991). The ﬁrst deglaciation reconstructions
were entirely based on geomorphological and sedimentological/stratigraphical
evidence for glaciation. In the absence
of a reliable dating technique, the pace of deglaciation was initially
inferred from the correlation between sequences of silty light- and
clayey dark-coloured sediment couplets. These ‘varves’ formed
during summer and winter seasons, respectively, through ice sheet
melt, runoff, and proglacial sedimentation. Extensive varve deposits
are typically exposed between highest shore lines and the
present coasts, and can be used for dating of the ice recession. This
is because the age of the ﬁrst varve overlying the formerly subglacial
terrain (typically bedrock or till), denotes the age of deglaciation
and therefore the former position of the ice sheet margin
(De Geer, 1884, 1912, 1940; Sauramo, 1918, 1923). In a series of
seminal studies on the deglaciation of the Fennoscandian Ice Sheet,
De Geer (1884, 1896, 1912, 1940) developed the Swedish Time Scale
(STS) varve chronology (Liden, 1938; Wohlfarth et al., 1995). In the
past three decades many studies have reﬁned the STS (Str€omberg,
1985a, b, 1989, 1990; 1994; Kristiansson, 1986; Cato, 1987;
Andren, 1990; Brunnberg, 1995; Wohlfarth et al., 1995, 1998;
Hang, 1997; Lindeberg, 2002), eventually resulting in a correlation
of the STS with the Greenland GRIP and NGRIP ice core record
layer-counting chronology (Andren et al., 1999, 2002; Stroeven
et al., 2015). These attempts to correlate varve- and ice core
chronologies have, however, revealed that hundreds of varves are
missing in the STS, thus exposing a key shortcoming of this indirect
dating technique (Andren et al., 2002).
With the advent of radiometric dating techniques (Bard and
Broecker, 1992), in particular radiocarbon (Anderson et al., 1947;
Arnold and Libby, 1949), the timing of maximum glacier extent,
has typically been constrained by the ﬁrst occurrence of living
matter in proglacial lakes dammed by the ice margin (yielding ages
older than the maximum ice extent) and in lakes dammed by the
end moraine once the ice margin had retreated from its maximum
extent (yielding ages younger than the maximum ice extent).
Dating the initiation of ice-free conditions using radiocarbon has
been the dominant dating-driven ice sheet reconstruction method,
and an abundance of minimum age constraints has permitted
detailed ice-sheet wide retreat reconstructions (e.g., Dyke et al.,
2003; Gyllencreutz et al., 2007).
There are a number of limitations associated with radiocarbon
dating in formerly glaciated regions (Hajdas, 2008). Critically, there
is dearth of datable organic material in many locations because
deglaciation occurred in polar deserts. Given the inevitable delay in
organic growth following deglaciation, 14
C dates provide minimum
limiting ages on deglaciation. In addition, the precision of radiocarbon
dating is compromised by the potential incorporation of
young carbon contaminants, incorporation of old carbon in the
depositional environment (marine reservoir or hard water effects;
Snyder et al., 1994), and variations in the atmospheric radiocarbon
concentration over time. These combined effects produce similar
radiocarbon ages for samples that were deposited hundreds of
years apart (radiocarbon dating plateaux). Because of these potential
pitfalls, considerable effort has been devoted to the
improvement of sample preparation methods and calibration of the
radiocarbon chronology (Bard et al., 1990, 1997; Wohlfarth et al.,
1995; Reimer et al., 2009, 2013). The best radiocarbon age determinations
come from environments where terrestrial macrofossils
have been used to constrain the age model (Barnekow et al.,
1998).
During recent decades two new dating techniques have
emerged, based on the burial of sand through optically-stimulated
luminescence (OSL) and the exposure of quartz-bearing clasts and
bedrock through measuring concentrations of cosmogenic nuclides.
In each case, datable material is abundant in pro-glacial and
glacial environments.
The OSL method is based on the build-up of a luminescence
signal in quartz grains that are shielded from sunlight through
burial (Rhodes, 2011). Exposure to sunlight deletes any previous
luminescence dose (bleaches the quartz grain). Hence, OSL can be
applied to date the burial of quartz grains (feldspar is also routinely
measured) given that two crucial conditions are met: 1) during
transport the grains are exposed to sunlight for a duration sufﬁcient
to become bleached and; 2) the sample has not been re-exposed
(Huntley et al., 1985; Aitken, 1998). Whereas the latter condition
can usually be veriﬁed in stratiﬁed sediments, partial-bleaching is a
major obstacle when dating glacial sediments, commonly resulting
in an over-estimation of the depositional age of the landform
(Fuchs and Owen, 2008; Alexanderson and Murray, 2012b). OSL is
therefore typically applied in settings where these conditions are
more easily met, such as where aeolian, ﬂuvial, or lacustrine
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sedimentation has occurred. Single-grain approaches account for
partial bleaching through evaluating the suitability of individual
grains in a sample, which strongly improves the reliability of OSL
(Murray and Wintle, 2000).
Cosmogenic nuclide surface exposure dating is applied to
samples taken from bedrock or boulders chosen for the information
they provide on deglaciation (Gosse and Phillips, 2001). Again, the
preferred mineral is quartz, in which four nuclides are produced
through exposure to cosmic rays; one stable (21
Ne) and three
radioactive (10
Be, 26
Al, and 14
C) nuclides. Beryllium-10 has been by
far the most reliable nuclide (Portenga and Bierman, 2011) and has
been extensively used in recent decades to construct glacial chronologies
around the globe (Stone et al., 2003; Balco and Schaefer,
2006; Ivy-Ochs et al., 2006; Rinterknecht et al., 2006; Hein et al.,
2010; Heyman, 2014; Rother et al., 2014; Stroeven et al., 2014).
Typically, samples are extracted from boulders on end moraine
crests, although samples from other landforms (Stroeven et al.,
2011) and of bedrock (Fabel et al., 2004; Li et al., 2005) have been
shown to also yield useful deglaciation ages. The reliability of
cosmogenic nuclide exposure dating for yielding accurate ages of
deposition and deglaciation is based on the assumption that the
sampled boulder/bedrock surface has been: 1) shielded from cosmic
rays prior to the last deglaciation and therefore contains no
inherited nuclides; and 2) continuously exposed to the full ﬂux of
cosmic rays since deglaciation with no shielding from sediment,
snow or vegetation. A breach of assumption 1 would cause erroneously
old ages, whereas a breach of assumption 2 would cause
erroneously young ages. An evaluation of these assumptions has
shown that for sample groups with large age scatter, prior exposure/inheritance
is typically less common than incomplete
exposure/post-glacial shielding (Heyman et al., 2011). Suites of
samples from individual landforms can be statistically analysed to
potentially identify anomalous dates and thereby determine accurate
landform ages (Applegate et al., 2010, 2012; Heyman, 2014).
An ever increasing availability and handling efﬁciency of
remotely sensed data (aerial photographs, satellite imagery, and
LiDAR; Smith et al., 2006) has heralded a resurgence of ice sheet
maximum and retreat reconstructions from landforms (Andersen,
1979, 1980, 1981; Lundqvist, 1986, 1994; Boulton and Clark,
1990a, c; Lundqvist and Saarnisto, 1995; Kleman et al., 1997,
2010; Andersen and Pedersen, 1998; Lindstr€om et al., 2000;
Ehlers and Gibbard, 2004; Margold et al., 2013). This has led to
regional compilations of ﬂow traces (striations, eskers, till lineations,
bedrock lineations, basal till fabrics, meltwater channels; e.g.,
H€attestrand, 1998; H€attestrand and Clark, 2006a) and their inclusion
in ice sheet-wide analyses (Kleman et al., 1997; Boulton et al.,
2001). Kleman et al. (1997, 2006) grouped coherent patterns of ice
ﬂow traces of the same age into ﬂow trace fans (map representations
of glacial landform swarms) and combined the undated
stacked record of subglacial ice ﬂow traces with dated ice marginal
successions to produce a reconstruction of Fennoscandian Ice Sheet
evolution over a glacial cycle. LiDAR scanning has produced recent
orders-of-magnitude increases in the resolution of elevation data
over landscape scales (Dowling et al., 2013). In this compilation we
take advantage of LiDAR data to advance our understanding of ice
sheet marginal retreat, particularly over southern Sweden (see
5.2.).
Geophysical- and ice sheet-modelling are increasingly used to
derive ice sheet reconstructions. These are independent methodologies,
the results of which can be evaluated against ﬁeld evidence
(Davis et al.,1999; Lambeck,1999; Napieralski et al., 2007). As an ice
sheet grows and decays, it transfers a shifting load onto Earth's
crust which responds through elastic and visco-plastic deformation.
As the crust rebounds following maximum glaciation (termed
glacial isostatic adjustment: GIA), its effect is recorded, by shifting
relative sea levels. Hence, from large sets of shoreline displacement
curves it is possible to separate the effects of eustatic sea level rise
(through ice sheet melting) and isostatic rebound, which can then
be used to inversely deduce the history of the ice load (Lambeck
et al., 1998). Because far-ﬁeld effects of deglaciation, for example
in Antarctica, will have direct, predictable, but uneven inﬂuence on
regional sea level, such as around Scandinavia, Earth-response
models need to be global (Slangen et al., 2014).
Ice sheet models are the most holistic way of addressing the
temporal evolution of glaciation for a particular ice sheet (Denton
and Hughes, 1981, 2002; Budd and Smith, 1982; Payne et al.,
1989; Huybrechts, 1993; Marsiat, 1994; Boulton et al., 1995;
Holmlund and Fastook, 1995; Hubbard, 1999; Siegert et al., 2001;
Kleman et al., 2002; Marshall et al., 2002; Boulton and Hagdorn,
2006; Clason et al., 2014; Seguinot et al., 2015). This is because
the extents and thicknesses of ice sheets are calculated over small
time increments and over large spatial scales in response to
changes in climate (mass balance) forcing. The choice of climate
forcing and the conversion of climate to mass balance remains the
largest limitation in ice sheet modelling, which necessitates calibration
of ice sheet models against ﬁeld evidence (Li et al., 2007;
Napieralski et al., 2007; Seguinot et al., 2014). It is with the aim
of providing targets for the evaluation of ice sheet modelling output
that we present a new reconstruction of the deglaciation of the
Fennoscandian Ice Sheet. Using the new reconstruction we revisit
important questions concerning the inﬂuence of ice sheet dynamics
and paleoclimate forcing on the ice sheet margin history,
the pace of retreat for different ice sheet sectors, and the inﬂuence
of topography on deglaciation patterns and rates. Ice sheet models
will ultimately yield the most comprehensive answers to these
questions, when properly tuned against the presented deglaciation
reconstruction, and provide a framework with which to query the
future behaviour of contemporary ice sheets.
At the time of the global LGM a contiguous ice mass covered
northern Europe from off-shore western Ireland to onshore
northwestern Taimyr Peninsula, on the eastern fringes of the Kara
Sea (Svendsen et al., 2004, Fig. 1). From a dynamic perspective, this
ice mass consisted of three ice sheets, each of which responded
individually to external forcing (geothermal heat, climate, sea level,
GIA), and which were amalgamated for a relatively brief period of
the total ice sheet duration. Following the global LGM the BritishIrish
Ice Sheet separated from the Fennoscandian Ice Sheet
offshore of southern Norway (Clark et al., 2012) and the Barents Sea
Ice Sheet unzipped from the Fennoscandian Ice Sheet offshore
northern Norway (Bjarnadottir et al., 2014). The focus of our study
is the retreat of the Fennoscandian Ice Sheet following its isolation
from these other ice masses (Fig. 1). Several attempts to establish
the deglaciation chronology of the Fennoscandian Ice Sheet precede
our efforts (Fig. 2). We present a new reconstruction, which
incorporates an abundance of publications during the past 15 years
that contain new geomorphological and geochronological data.
This is speciﬁcally aimed at delivering calendar-year time-slice
representations of ice sheet extents for use by ice sheet modellers.
2. Data
2.1. Geomorphology
We begin with a description of the key indicative ice-marginal
and subglacial landforms on which the deglaciation reconstruction
is primarily based. These include, in a progression from proglacial/ice
marginal to subglacial, ice-dammed lakes, marginal
meltwater channels, ice-marginal formations (moraines and glaciﬂuvial
deposits), eskers, lineations, and striae. We then review the
geochronological tools available for deglaciation reconstructions
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(e.g., Hughes et al., 2011), shortly review their strengths and pitfalls,
and present the data included in our deglaciation reconstruction.
2.1.1. Ice-dammed lakes
Where ice margins block the natural drainage of ice-free
catchments, water ponding may lead to the formation of icedammed
lakes. Such lakes are inherently unstable and drain catastrophically
when the ice dam fails. This occurs when the lake
hydrostatic pressure exceeds the ice overburden pressure at the
lake outlet, when a retreating ice margin exposes lower terrain, or
through overtopping. Glacial lakes have formed with dimensions
over many orders of magnitude, from the Baltic Ice Lake
(349,000 km2
; Jakobsson et al., 2007) to the numerous intermediate-
and small-scale ice-dammed lakes impounded between the
Scandinavian Mountains and the retreating western margin of the
decaying ice sheet (Lundqvist, 1972; Kleman, 1992, Fig. 3). Evidence
of former ice-dammed lakes such as shorelines (erosional/depositional),
perched deltas, and spillway (overﬂow) channels are useful
tools for reconstructing the ice marginal retreat pattern in areas
formerly covered by cold-based ice (Fr€odin, 1913; Lundqvist, 1973;
Jansson, 2003). Fig. 3 shows the post-Younger Dryas extent of icedammed
lakes in Fennoscandia, compiled from available sources
(Lundqvist, 1972, 1973; Melander, 1977; Ulfstedt, 1981; Borgstr€om,
1989; Longva and Thoresen, 1991).
2.1.2. Marginal meltwater channels
Water produced during ice sheet surface melting predominantly
runs off the surface and along the ice sheet margin where it abuts
higher ground. While ﬂowing along the ice margin, streams erode
the ground surface at the junction with the ice and form marginal
meltwater channels (Borgstr€om, 1989; Mannerfelt, 1945, 1949;
Syverson and Mickelson, 2009). Marginal meltwater channels are
typically tens of meters deep, meters wide, and hundreds of meters
long, and usually form in subparallel down-slope sequences.
Importantly, the slope and orientation of marginal channels occur
at oblique angles to the hillslope topography into which they are
eroded, and they provide a record of retreating ice margins that is
independent of other deglacial landforms (Mannerfelt, 1945;
Lundqvist, 1973; Borgstr€om, 1989; Kleman, 1994; Greenwood
et al., 2007; Margold et al., 2011). This is because, in contrast to
eskers and lineations, marginal meltwater channels form also
during deglaciation under cold-based conditions (Kleman, 1992;
Dyke, 1993; H€attestrand and Stroeven, 2002; Jansson et al., 2002).
Meltwater landforms have therefore been used in our reconstruction
primarily where the ﬁnal deglaciation occurred under coldbased
conditions (Kleman, 1992; Kleman et al., 1997, 2006;
Kleman and H€attestrand, 1999) (Fig. 4).
2.1.3. Ice-marginal formations
Along the margins of ice sheets, there are several processes by
which sediment exits the ice and becomes part of the glacier
foreland (Boulton et al., 1985). The sediment frequently becomes
concentrated in ice-marginal formations, including end moraines
and glacioﬂuvial deposits that generally mirror the shape and position
of former ice margins (Fig. 5). These formations occur
commonly along the entire Fennoscandian Ice Sheet margin and
indicate either interruptions in ice sheet retreat or re-advances
following the LGM. Particularly extensive ice-marginal formations
Fig. 1. The Eurasian ice sheet at the Last Glacial Maximum (LGM) as portrayed by Svendsen et al. (2004). This ice sheet complex consisted of the amalgamation of three separate ice
sheet centers, the British-Irish Ice Sheet in the west, the Barents Sea Ice Sheet in the east, and the Fennoscandian Ice Sheet, the object of our study, in the center (red box; Figs. 4e7,
9, 10, 12a). (For interpretation of the references to colour in this ﬁgure legend, the reader is referred to the web version of this article.)
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characterize the eastern and southern ice sheet limits, through
Russia, the Baltic countries, Poland, Germany, Denmark, and into
Norway (Fig. 6). Hence, series of these formations can be traced
inwards from local glacial maximum positions to Younger Dryas
positions (Fig. 5), which mark the last ice sheet-wide interruption
in margin retreat before complete deglaciation. Where icemarginal
formations are punctuated by gaps of non-deposition or
meltwater stream erosion, they can often be extrapolated to each
Fig. 2. Four reconstructions of the deglaciation pattern of the Fennoscandian Ice Sheet by a) Lundqvist (1986), b) Lundqvist and Saarnisto (1995), c) Kleman et al. (1997), and d)
Boulton et al. (2001).
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other assuming lateral ice sheet continuity. To guide pre-Younger
Dryas deglaciation patterns of the Fennoscandian Ice Sheet, we
have compiled moraine positions from maps (Fig. S1,
Supplementary dataset). Post-Younger Dryas ice-marginal formations
are much rarer and so only guide deglaciation patterns
regionally, and they can indicate both interruptions of the ice
margin retreat and re-advances up to late Preboreal, 10,500 cal
years BP (Sveian et al., 1979).
2.1.4. Eskers
Eskers are ridges of coarse-grained sorted sediment deposited in
meltwater tunnels at the base of an ice sheet. They can be single
ridges or form networks of several parallel ridges. Eskers can be
short (hundreds of meters) and straight but more typically are long
and winding and can extend for hundreds of kilometers and be tens
of meters high (De Geer, 1897; Lundqvist, 1979; Storrar et al., 2014).
Because eskers are such recognizable and sizeable landforms, they
have been accurately mapped from aerial photographs, and reliable
esker maps exist for individual countries (Lundqvist, 1959) as well
as for larger regions such as northern Fennoscandia (Nordkalott
Project, 1986). For our deglaciation reconstruction of the Fennoscandian
Ice Sheet, we present an ice sheet-wide esker map for
shield areas, where they are abundant (Fig. 5). The map is compiled
from existing publications (Nordkalott Project, 1986; Niemel€a et al.,
1993; H€attestrand, 1998; Bargel et al., 1999; H€attestrand and Clark,
2006a; NGU, 2014) and managed in ArcGIS. The esker pattern on
Fig. 3. Glacial lakes in Fennoscandia used to constrain the ice margin retreat pattern. The Baltic Ice Lake existed during the Younger Dryas, until it ﬁnally drained at its northwestern
extremity, Mount Billingen, at 11,620 cal years BP (Stroeven et al., 2015). Additional smaller glacial lakes existed before the Younger Dryas, but we have made no attempt to make a
systematic inventory of them, because we rely predominantly on ice-marginal formations to guide our deglaciation reconstruction. The position of the ﬁnal narrow ice ridge joining
residual ice in southern Norway with the main dome in the north (cf. 10.2 cal kyr BP, Fig. 9), is constrained by glacial lakes having been dammed in opposite directions. The period
with ice-dammed lakes in existence along the mountain backbone is short, around 600 years. The drainage of some of the lakes has been traced and dated in the clay varve
chronology (De Geer, 1940; Borell and Offerberg, 1955; F€oz€o, 1980; Str€omberg, 1989).
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Fig. 5 provides a generalized pattern because of the relatively small
scales of the source maps, and short eskers (including subglacially
engorged eskers) are therefore not included or used for the
deglaciation reconstruction. Eskers that formed over the sedimentary
bedrock areas west, south, and east of the Baltic Sea are
generally smaller and shorter than those over shield areas, and
esker compilations generally cover only minor areas (e.g., Rattas,
2007). As eskers have been shown to form within limited distances
of the contemporaneous ice margin (Hebrand and Åmark,
1989; Kleman et al., 1997), and because water ﬂow directions
follow the overburden pressure regime, we use the direction of
eskers to guide the overall shape of former ice sheet margins. In our
reconstruction, ice sheet margins are always drawn perpendicular
to the esker long-axes. It should be noted, however, that there are
areas where eskers could not be used for our reconstruction
because they did not form during the last deglaciation but, rather,
formed during earlier deglaciations. This is the case, for example, in
extensive areas of northern Sweden (Lagerb€ack and Robertsson,
1988), and Finland (Johansson and Kujansuu, 1995), where eskers
of a pre-LGM deglaciation (Helmens et al., 2000) are cross-cut by
younger esker systems, are covered by tills, and have kettle-holes
with interstadial sediments, all which indicate that these eskers
escaped erosion during the last deglaciation through sustained
cold-based conditions. In these areas, other meltwater landforms,
such as ice-dammed lake traces and marginal meltwater channels,
were used to reconstruct retreat of the Fennoscandian Ice Sheet
during the last deglaciation.
2.1.5. Lineations
The most widely utilised subglacial landform for ice sheet reconstructions
is the glacial lineation (Fairchild, 1907; Linton, 1963;
Punkari, 1982; Boulton and Clark, 1990a, b; Kleman, 1992; Clark,
1993; Kleman et al., 1997). Lineations are elongated landforms
that form parallel to ice ﬂow and are usually referred to as drumlins.
Because lineations can be formed through depositional and
erosional processes, they may be comprised of diamicts, sorted
sediments, and/or bedrock (cf. review by Stokes et al., 2011).
Although larger landforms may have formed during multiple glaciations
(H€attestrand et al., 2004), and although later generations of
lineations do not necessarily erase older lineations (Kleman, 1992),
the association of lineations with other deglacial landforms implies
that most of these inform the ice ﬂow direction, and therefore the
ice surface slope, just prior to deglaciation. Lineations are a
particularly useful complement to the directional information
contained in eskers because they often form in swarms hundreds of
kilometers in extent and may contain tens of thousands of elements
(H€attestrand et al.,1999, 2004; Dellgar Hagstr€om, 2006; Clark et al.,
2009a). We have employed the lineation database of Kleman et al.
(1997, Fig. 3) in our reconstruction of the last deglaciation of the
Fennoscandian Ice Sheet.
2.1.6. Striae
Striae represent the ﬁnest-scale imprint of ice ﬂow on bedrock.
On outcrops where more than one set of striae are preserved, their
cross-cutting relationships may reveal the evolution of ice ﬂow
directions (Erdmann, 1868; Lundqvist, 1969) and indicate the ice
ﬂow direction closest in time to deglaciation. Lineations and striae
both record the ice-ﬂow direction at the time of formation, and
could therefore be expected to yield the same information
regarding ice ﬂow-evolution. In reality, there are important differences
in the information provided by the two data types (Kleman,
1990). Lineations are typically formed from a glacial deposit, and
their spatial arrangement means that the continuity and extent of a
ﬂow pattern can be visually judged. Striae, on the other hand, are
purely erosional bedrock forms and constitute detailed point data
even though large collections of striae observations, with less
precision than for lineations, still can give a visual imprint of ﬂow
patterns. Importantly though, the maximum “time depth” is larger
for striae than for lineations. This is because a rock outcrop typically
provides facets or steps that are sheltered during later ice ﬂow and
may therefore locally preserve older striae. No corresponding local
protection mechanisms exists for lineations, except for protection
from erosion under cold-based conditions (Kleman et al., 2002),
and so preservation of older directional information decreases
more directly as a function of subsequent ice ﬂow velocity and
duration.
An important property of the composite striae record is that the
locally youngest striae may, in some places, indicate deglacial ice
ﬂow directions in areas lacking lineation swarms. We have used
compilations of striae (ﬁrst pioneered by Sefstr€om, 1836) to extract
the youngest ice ﬂow direction in areas where such information is
otherwise absent (Ljungner, 1943). These areas are predominantly
along the Gulf of Bothnia (Fig. 6) where the youngest sets of striae
indicate the re-advance of an ice lobe (Lundqvist, 2007) and in
northwestern Sweden where meltwater landforms and striae can
be used to construct ﬁnal deglaciation ice ﬂow directions in areas
characterized by cold-based ice (Kleman, 1990).
2.2. Chronology
No dating technique is applicable in every ﬁeld setting, whether
it be due to limits on the materials available to date or the timescale
spanned by the method itself. Consequently, a range of geochronological
tools are employed and the most important methods in
the Fennoscandian context are as follows.
Fig. 4. The shrinkage of the cold-based core area of the Fennoscandian Ice Sheet
during deglaciation from its local LGM maximum position (Kleman, 1992; Kleman
et al., 1997, 2006; Kleman and H€attestrand, 1999; H€attestrand and Clark, 2006b). The
outer blue envelope represents the inferred minimum cold-based extent at LGM. The
innermost envelope represents areas inferred to have had cold-based conditions until
local deglaciation. The intermediate envelope shows how ice streams in northern
Norway, Finland, and the collapse event following a surge in the Gulf of Bothnia
(Str€omberg, 1989; Lundqvist, 2007; Kleman and Applegate, 2014; Greenwood et al., in
press), extended wet-based conditions into the ice sheet in a corridor-like pattern
during the decay phase. At any given point in time, the border zone between warmand
cold-based conditions was probably mosaic-like in sheet ﬂow areas (Kleman et al.,
1999; Kleman and Glasser, 2007).
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2.2.1. Radiocarbon dating
Radiocarbon dating of organic material has traditionally been
the key chronological tool for deﬁning the timing of deglaciation
(e.g., Dyke, 2004). With glacier retreat, new terrain becomes ice free
and available for the production, storage, and preservation of
organic material in pro-glacial sedimentary archives. With a halflife
of 5730 ± 40 years for 14
C, which limits its application to
about the last ~50e40 kyr, radiocarbon dating provides chronological
constraint on the Fennoscandian Ice Sheet deglaciation (c.
24e10 kyr).
We have compiled a database of 335 published 14
C ages of
relevance for the deglaciation of the Fennoscandian Ice Sheet
(Table 1, Fig. 7, Supplementary dataset). The 14
C ages are primarily
derived from basal sediment in lakes and peat cores and include
measurements on both bulk sediment and terrestrial macrofossils.
One-third of the 14
C dates in our compilation, mainly from the
western and southeastern sectors of the Fennoscandian Ice Sheet,
are derived from sub-till sediment samples. Organic material in
these samples pre-dates the glacial advance to the LGM conﬁguration
and subsequent retreat to the sample site, and they therefore
represent maximum ages of deglaciation.
All radiocarbon ages have been calibrated using OxCal 4.2
Fig. 5. Prominent eskers and ice-marginal positions in the area evacuated by the Fennoscandian Ice Sheet since the LGM. Note that the esker data only covers shield areas, where
eskers form continuous esker chains that are useful for reconstructing the deglaciation. Eskers in non-shield areas have been omitted since they are short, infrequent, and add little
to the deglaciation information provided by the generally rich record of ice-marginal positions in these areas. The overwhelming majority of mapped eskers date from the last
deglaciation (Nordkalott Project, 1986; Niemel€a et al., 1993; H€attestrand, 1998; Bargel et al., 1999; H€attestrand and Clark, 2006a; NGU, 2014). For ice-marginal formations in the
southern and eastern sectors of the Fennoscandian Ice Sheet, we have relied primarily on literature of the past 15 years (Fig. S1, Supplementary dataset). Where published data have
been in conﬂict, we have employed minimum-complexity assumptions for spatial (and thereby chronological) correlation of data, assessing spatial, morphological, chronological,
and glaciological relationships and probabilities. For southern Sweden (black box, Fig. 11) we have made amendments to existing moraine maps through interpretation of LiDARgenerated
DEM data. Landforms and ages (in cal kyr BP) from Sollid et al. (1973), Houmark-Nielsen and Kjær (2003), Demidov et al. (2006), H€attestrand and Clark (2006b), Rise et al.
(2006), Raukas et al. (2010), Mangerud et al. (2011), Saarse et al. (2012), Marks (2012), Lasberg and Kalm (2013), Anjar et al. (2014), Bjarnadottir et al. (2014), Briner et al. (2014),
Rinterknecht et al. (2014), Stokes et al. (2014), and Svendsen et al. (2015).
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(Bronk Ramsey, 2009) and the Intcal13 curve (Reimer et al., 2013).
All marine samples were corrected for a marine reservoir effect, by
applying a correction of 300e800 years, according to the original
publications (Supplementary dataset), for the presence of old carbon
derived from the marine environment (Mangerud and
Gulliksen, 1975).
2.2.2. Varves, Swedish Time Scale (STS)
The strength of the clay varve chronology, if the varve
measuring sites are closely spaced, is that retreat of the ice margin
can be resolved more accurately than with any other correlation
method, regardless of whether the varve chronology is ﬂoating in
time or is of calendar-year quality. Fig. 8 exempliﬁes the level of
detail that can be achieved using clay varve correlations. Isochrons
are drawn on the basis of clay varve correlations using the age of
the oldest varve. In this way the STS should be internally robust for
2400 years of Late Glacial time (Table 2, -2250 to þ140 STS varve
years or 12,340 to 9950 cal yrs BP; Bergstr€om, 1968; F€oz€o, 1980;
Kristiansson, 1986; Str€omberg, 1989, 1990, 1994, 2005;
Brunnberg, 1995; Wohlfarth et al., 1998). To obtain a complete
varve chronology, connecting the Late Glacial varve sequence with
postglacial varves that extend to the present (Cato,1987), over 1300
Fig. 6. Map of the Fennoscandian Ice Sheet deglaciation domain with ice-marginal formations and place names mentioned in the main text.
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sites have been measured. Despite these efforts, based on AMS 14
C
dates on terrestrial macrofossils embedded in the varves
(Wohlfarth,1996), 14
C-dated marker horizons in Swedish lacustrine
deposits, central-European tree-ring chronologies, and Greenland
ice core records (Bj€orck et al., 1996), it has been shown that hundreds
of varves are missing, most probably in the postglacial section
of the STS. Estimates of the number of missing varves have
varied over time but are generally about 700e900 (Str€omberg,
1994; Andren et al., 2002).
We use the catastrophic drainage of the Baltic Ice Lake, 35 years
before the start of the Holocene (Andren et al., 2002), as an event
that can be used to tie the STS to ice core records (Andren et al.,
1999, 2002; Bj€orck et al., 2001), and we speciﬁcally explore its
link to the NGRIP ice core record (Stroeven et al., 2015). The NGRIP
ice core record has been layer-counted across the Younger Dryas/
Preboreal (Holocene) transition, yielding an age of 11,700 ± 99 cal
years b2k (Walker et al., 2009; Rasmussen et al., 2014), or 11,650 cal
years BP. Detailed statistical comparisons between the 14
C record in
tree rings and the 10
Be record in ice cores across this boundary
(Muscheler et al., 2008, 2014) imply that the ice core record may be
65 years too old. The best estimate of the start of the Holocene is
11,585 cal years BP and the drainage, being 35 years older (Andren
et al., 2002), is therefore pinned to 11,620 cal years BP (Stroeven
et al., 2015). The timing of the Baltic Ice Lake drainage occurred
at STS -1530 (Bj€orck et al., 2001; Andren et al., 2002) or 10,770
varve years BP, which instils an age difference of 850 years between
the two annual records (Table 2). Fig. 8 shows the constraints that
the STS varve record offers to ice marginal positions between
13,390 cal years BP (North of Vimmerby; Kristiansson, 1986;
Wohlfarth et al., 1998), and 9950 cal years BP (Pautr€ask;
Bergstr€om, 1968) even though our connection between the robust
part of the record and the Kristiansson-Wohlfarth section of the
record (black series in Fig. 8) remains challenging.
Table 1
Publications with radiocarbon dates included in the Supplementary dataset, including the number of samples (N).
Publications N Publications N
Aas and Faarlund (1988) 1 Larsen et al. (2006) 5
Abrahamsen and Readman (1980) 1 Larsen et al. (2014) 1
Alm (1993) 2 Lasberg and Kalm (2013) 6
Alstadsæter (1982) 2 Liiva et al. (1966) 1
Andersen (1975) 1 Linden et al. (2006) 4
Antonsson et al. (2006) 1 M€oller et al. (2013) 8
Arppe and Karhu (2010) 3 Nese and Lauritzen (1996) 2
Bang-Andersen (2003) 2 Noe-Nygaard and Heiberg (2001) 1
Bennike and Jensen (1995) 2 Nydal et al. (1972) 1
Berglund (1995) 5 Olsen (1997) 2
Berglund (2005) 2 Olsen (2000) 1
Berglund et al. (1976) 3 Olsen (2002) 2
Bergman et al. (2004) 1 Olsen (2004, unpublished) 1
Bergman et al. (2005) 2 Olsen et al. (1996) 5
Bergstrøm (1975) 2 Olsen et al. (2001) 40
Bitinas et al. (2002) 3 Olsen et al. (2013b) 18
Bj€orck and Digerfeldt (1982a) 2 Putkinen and Lunkka (2008) 4
Bj€orck and Digerfeldt (1982b) 1 Repo and Tynni (1967) 1
Bj€orck and Digerfeldt (1986) 1 Repo and Tynni (1969) 2
Bj€orck and Digerfeldt (1991) 1 Repo and Tynni (1971) 4
Bj€orck and M€oller (1987) 1 Richardt (1996) 1
Blake and Olsen (1999) 5 Rinterknecht et al. (2006) 47
Corner et al. (2001) 3 Rosen (2005) 4
Digerfeldt (1979) 1 Rosen et al. (2001) 2
Donner et al. (1978) 3 Rotnicki and Borowka (1995) 2
Dreimanis and Zelcs (1995) 15 Rubensdotter (2006) 3
Eilertsen et al. (2005) 16 Saarse et al. (2009) 2
Ek (2004) 1 Saarse et al. (2012) 1
Eronen (1976) 2 Sandgren et al. (1999) 1
G€ottlich et al. (1983) 1 Segerstr€om and von Stedingk (2003) 4
Håkansson (1970) 1 Seidenkrantz and Knudsen (1993) 1
Håkansson (1975) 2 Seirien_e et al. (2006) 2
Håkansson (1978) 1 Sepp€a and Birks (2002) 1
Håkansson (1982) 1 Sepp€a and Weckstr€om (1999) 1
Håkansson (1987) 3 Sepp€a et al. (2004) 2
Hammarlund et al. (2004) 1 Sepp€a et al. (2012) 3
Heikkil€a and Sepp€a (2003) 1 Shemesh et al. (2001) 1
Heinsalu and Veski (2007) 1 Snyder et al. (1997) 2
Helmens et al. (2000) 1 Snyder et al. (2000) 1
Hillden (1979) 3 Stancikait_e et al. (2008) 3
Houmark-Nielsen and Kjær (2003) 3 Stankowska and Stankowski (1988) 2
Jensen et al. (2002) 2 Svedhage (1985) 1
Johnsen et al. (2010) 1 Svensson (1989) 4
Johnson and Ståhl (2010) 5 Tolonen and Ruuhij€arvi (1976) 1
Klovning and Hafsten (1965) 1 Valen et al. (1996) 2
Korsager et al. (2003) 1 Vorren (1978) 2
Kramarska (1998) 2 Vorren and Alm (1999) 1
Krog and Tauber (1974) 2 Vorren et al. (1988) 2
Lagerlund and Houmark-Nielsen (1993) 1 Vorren et al. (2013) 5
Larsen et al. (1999) 1 Wohlfarth et al. (1999) 1
Zernitskaya et al. (2007) 1
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2.2.3. Optically-stimulated luminescence (OSL) dating
OSL dating enables direct dating of sediment deposition and
burial, and it can potentially yield accurate minimum ages of
deglaciation. However, in practice it appears that some OSL ages are
older than the expected deglaciation age for sites in Fennoscandia
(Alexanderson and Murray, 2012b; Johnsen et al., 2012). Because
this is probably due to incomplete bleaching of, especially, subglacial
till and proximal glacioﬂuvial sediment samples
(Alexanderson, 2007; Alexanderson and Murray, 2012a), most of
the 138 OSL samples that are part of the deglaciation reconstruction
(Table 3), concern minimum age constraints for distal glacioﬂuvial,
lacustrine, and eolian sediment samples (Fig. 7; Larsen et al., 1999,
2006, 2014; Strickertsson and Murray, 1999; Lyså et al., 2001, 2011,
2014; Houmark-Nielsen, 2003; Houmark-Nielsen and Kjær, 2003;
Kjær et al., 2003a; Kortekaas and Murray, 2007; Kortekaas et al.,
2007; Johnsen et al., 2010; Lüthgens et al., 2011; Alexanderson and
Murray, 2012b).
2.2.4. Cosmogenic nuclide exposure dating
Cosmogenic surface exposure dating of glacial landforms and
deposits has become a key tool for deﬁning glacier and ice sheet
chronologies. We present a compilation of published and new 10
Be
(n ¼ 786) and 26
Al (n ¼ 74; eight of these have no corresponding
10
Be measurements) exposure ages for the area covered by the LGM
Fennoscandian Ice Sheet (Table 4, Fig. 7; Supplementary dataset).
The exposure ages are derived from sampled bedrock surfaces
(n ¼ 284), glacial boulders (n ¼ 474), and cobble/pebble/sediment
(n ¼ 36). All samples with potential importance for the deglaciation
chronology have been included, including samples from bedrock
surfaces that have been preserved under cold-based ice (Fabel et al.,
2002; Stroeven et al., 2002b; Linge et al., 2006a; Darmody et al.,
2008). Most of the new, previously unpublished samples
(n ¼ 132) are from northern Sweden and Norway (n ¼ 84), plus
some from the Kola Peninsula, Russia (Fig. 6; n ¼ 15), Finland
(n ¼ 17), and east-central (n ¼ 4), west-central (n ¼ 5), and
Fig. 7. Locations of optically-stimulated luminescence, radiocarbon, and cosmogenic isotope samples included in the chronological database that is part of the deglaciation
reconstruction (Tables 1, 3 and 4; Supplementary dataset). Cosmogenic isotope samples have resulted in both 10
Be and 26
Al dates.
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southwestern Sweden (n ¼ 9). These new samples have been prepared
following standard procedures (Fabel et al., 2002, 2006;
Stroeven et al., 2002a, 2002b, 2011, in review) and were
measured at PRIME Lab, ANSTO, and SUERC over the years
2000e2008 (see Supplementary dataset).
Data for exposure age calculations has been compiled from the
original publications and recalculated with a consistent production
rate and scaling scheme. We use the reported sample thickness,
sample density (adopting 2.65 g cmÀ3
where not otherwise stated),
and topographic shielding and we assume zero surface erosion. An
important reﬁnement in our compilation is 10
Be standardization.
With the 10
Be AMS standard calibration of Nishiizumi et al. (2007),
it became clear that previously assumed isotope ratios of AMS
standards, and reported 10
Be concentrations based on those
standards, deviate by up to 14% from the new Nishiizumi et al.
(2007) standard. We have made an effort to track the applied
standardization for all 10
Be samples and we regard our dataset as
the best available for the Fennoscandian Ice Sheet region. All
exposure ages have been calculated using a modiﬁed version of the
CRONUS calculator (Balco et al., 2008), with the nuclide speciﬁc
LSDn production rate scaling scheme (Lifton et al., 2014), the
regional Scandinavian reference 10
Be production rate of 3.95 ± 0.10
atoms gÀ1
yrÀ1
(Stroeven et al., 2015), and the corresponding
reference 26
Al production rate of 26.71 ± 1.60 atoms gÀ1
yrÀ1
.
Exposure ages for the CRONUS scaling schemes (Balco et al., 2008)
and the two LSD scaling schemes (Lifton et al., 2014), using the
regional Scandinavian 10
Be reference production rate, are listed in
the Supplementary dataset. The exposure ages of the various
Fig. 8. a) Compilation of the Late Glacial clay varve chronologies in Sweden and Finland that are used in this study (Table 2). The Swedish Time Scale (STS) is expressed as varve
years before (negative values) and after (positive values) the ”zero-year” in the revised and corrected STS (Bergstr€om, 1968; Str€omberg, 1985b, 1989, 1990, 1994, 2005; Kristiansson,
1986; Cato, 1987; Brunnberg, 1995; Wohlfarth et al., 1998). The framework may be considered as a two-part chronology, one with “glacial” varves, deposited before a ”zero-year” (De
Geer, 1940), and one with “postglacial” varves, deposited after this “zero-year” (Liden, 1938). Hence, the varve ±0 has been an “anchor point” for all past revisions of the STS. The age
of the zero-year varve, STS ±0, is c. 10,090 cal years BP based on the adopted correlation of the STS with the Greenland ice core record (Stroeven et al., 2015, Table 2). The ice
recession lines of Str€omberg (1994) in west-central Sweden have been corrected, and the “ﬂoating” varve chronology by Kristiansson (1986) has been correlated with the STS, as
proposed by Wohlfarth et al. (1993) and Brunnberg (1995). The ice-marginal lines in Finland are based on Str€omberg (1990, 2005), which include new varve measurements north of
the Second Salpausselk€a Moraine (Ss II). The post-Younger Dryas recessional lines of the varve chronology are considered robust. However, the chronology is less robust between
STS c. -2300 and À2600, i.e. during the early part of the Younger Dryas, because ice front oscillations may have eroded already deposited varves. b) Locally, short varve sequences
require a dense pattern of varve measuring sites to get reliable correlations between varve diagrams. Because, for the area shown here, the distance between measuring sites had to
be restricted to only a few kilometres, the resulting pattern of ice margin retreat becomes highly detailed. Modiﬁed from Str€omberg (1989).
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scaling schemes generally agree well and all CRONUS scaling
scheme exposure ages are within 7.1% of the LSD scaling exposure
ages.
3. Methodology
The outlines of ice sheet retreat isochrons are primarily based on
the pattern of eskers and ice-marginal positions (Fig. 5). Retreat
isochrons that are drawn perpendicular to esker long-axes (Kleman
et al., 2006) tend to produce smooth contours. Isochrons based on
ice-marginal positions, however, account for the irregular and often
lobate nature of ice margins as expressed in remnant deglacial
landforms and commonly yield highly irregular retreat contours.
Some of the margins are smoother interpreted positions between
more detailed dated ice-marginal formations to allow for a
consistent 1000 year contouring. As far as we know, the retreat
pattern never violates these principles of reconstruction and
‘youngest sets’ of lineations and striae typically conform to the
presented retreat pattern. While eskers form an almost continuous
pattern in the shield areas they are scarce in the mountains (Fig. 5),
probably because of persistent cold-based conditions. In areas
characterised by cold-based deglaciation, ice sheet retreat is
reconstructed with the use of lateral channels (Kleman, 1992;
H€attestrand, 1998) and ice-dammed lakes. Where thawing
occurred shortly before the retreating ice margin passed through
these cold-based regions, faint lineations in the till sheet and glacial
striae on bedrock are also used to reconstruct ice sheet retreat
(Clarh€all and Kleman, 1999; Harbor et al., 2006).
The deglaciation timing is almost entirely based on published
constraints and correlations. However, the quality of, and detail
provided by, the constraints differ across the glaciated domain. For
ice-marginal positions older than 13 cal kyr BP, the chronology is
largely constrained by radiocarbon, OSL, and cosmogenic nuclidederived
ages (Fig. 7; Supplementary dataset). For the
17e13 cal kyr BP deglaciation of southern Sweden, we also couple
the radiocarbon and cosmogenic nuclide data for ice-marginal
positions to climatic events recorded in Greenland ice cores (for a
detailed explanation, see section 5.2). For ice-marginal positions
younger than 13 cal kyr BP, we use the timing of the retreat of the
Fennoscandian Ice Sheet from its Younger Dryas position in
southern Sweden and Finland as a starting point to build the
chronology. When the ice sheet retreated from the northern tip of
Mount Billingen, the Baltic Ice Lake (Fig. 3) catastrophically drained
to the Kattegatt (Fig. 6; Lundqvist, 1921; Johansson, 1926). The
event has been dated to 11,620 ± 100 cal years BP using radiocarbon
and by correlating this event in the STS with the Pleistocene/Holocene
boundary in the NGRIP ice core (Stroeven et al., 2015). We
use the internally-consistent varve record between c. 12,340
(Korsberga; Str€omberg, 1994) and 9950 (Pautr€ask; Bergstr€om,
1968) cal years BP (Table 2, Fig. 8) to guide the pace of retreat for
this part of the record. Str€omberg (1990, 2005) connected the varve
record in southern Finland, across Åland (Fig. 6), to the STS, thereby
allowing further geochronological control on ice-marginal positions
(Fig. 8). Final deglaciation in the Sarek Mountains of northwestern
Sweden (Fig. 6) occurred after 9.7 cal kyr BP, in general
agreement with the clay varve record and ages derived using
radiocarbon and cosmogenic nuclides.
Hence, the method of reconstructing isochrons used here differs
between pre- and post-Younger Dryas periods. For regions that
were deglaciated before the Younger Dryas, ages of ice-marginal
positions were derived from published data. In cases where published
ages for mapped ice-marginal formations were inconsistent,
groups of samples with consistent ages, or consistent ages using
different dating techniques, were considered more reliable than
individual ages. In the ﬁnal step of reconstructing time slices that
were 1 kyr apart, some mapped ice-marginal formations were used
directly while others were visually interpolated from adjacent icemarginal
formations, generally assuming a steady retreat rate, but
giving due consideration to topography. For regions that were
deglaciated after the Younger Dryas, clay varve chronology was
used to construct retreat isochrons. This was achieved by adopting
the principle that isochrons extrapolated away from established
ice-marginal positions (Fig. 8) should always be perpendicular to
youngest deglaciation ice ﬂow traces, as indicated by eskers, glacial
lineations, and striations, and using constraints provided by icedammed
lakes and meltwater channels. The pace of retreat as
derived from the youngest part of the varve record was used to
construct the last two isochrons, and the ﬁnal age of deglaciation
was cross-checked against published deglaciation ages.
4. Results
4.1. Deglaciation overview
We present a deglaciation map of the Fennoscandian Ice Sheet
Table 2
Conversion of the Swedish Time Scale (STS) to cal yrs BP.
STS (Fig. 8)a
Varve yrs BPb
cal yrs BPc
þ140 9100 9950
±0 9240 10,090
À1400 10,640 11,490
À1500 10,740 11,590
À1530 10,770 11,620
À2200 11,440 12,290
À2250 11,490 12,340
À3300 12,540 13,390
a
The Swedish Time Scale, STS, was proposed by Gerard De Geer (1935, 1940). He
deﬁned clay varves deposited before the zero-year, ±0, as glacial varves (negative
values), and varves deposited after the zero-year as postglacial varves (positive
values).
b
Conversion to varve yrs BP is based on a connection of the ﬂoating STS to the
present by Cato (1987), using 9240 varve years for the zero-year varve, relative to
1950 (BP).
c
Conversion to cal yrs BP is based on the connection of the STS to the NGRIP
Greenland ice core chronology using the timing of the Baltic Ice Lake Drainage (STS
c. -1530; Andren et al., 2002) which correlates to c. 11,620 cal years BP in the ice core
record (Stroeven et al., 2015), thus requiring a further revision of c. 850 years.
Table 3
Publications with OSL dates included in the Supplementary dataset, including the number of samples (N).
Publications N Publications N
Alexanderson and Henriksen (in press) 12 Larsen et al. (1999) 14
Alexanderson and Murray (2012b) 30 Larsen et al. (2006) 6
Houmark-Nielsen (2003) 7 Larsen et al. (2014) 8
Houmark-Nielsen and Kjær (2003) 6 Livingstone et al. (2015) 2
Johnsen et al. (2012) 7 Lüthgens et al. (2011) 7
Kjær et al. (2003a) 7 Lyså et al. (2011) 4
Kortekaas and Murray (2007) 2 Lyså et al. (2014) 24
Kortekaas et al. (2007) 1 Strickertsson and Murray (1999) 1
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10.1016/j.quascirev.2015.09.016
with isochrons marking every 1000 years between 22 and
13 cal kyr BP and every hundred years between 11.6 and 9.7 cal kyr
BP (Fig. 9; Video, Supplementary dataset). Abundant literature attests
to the difﬁculties in resolving the dynamic behaviour of the
Fennoscandian Ice Sheet during the Younger Dryas chronozone,
although two to three extensive end moraine belts indicate
standstills and re-advances of the ice sheet along most of its margin
during this period (Rainio et al., 1995; Lundqvist, 2004; Mangerud
et al., 2011; Putkinen et al., 2011). The onset of the Younger Dryas in
Scandinavia is delayed by 100 years relative to the Greenland ice
core record (GS-1; 12.8e11.7 cal kyr BP; Lohne et al., 2013). Hence,
we denote the net retreat distance during the Younger Dryas as an
ice-marginal zone spanning 12.7e11.6 cal kyr BP, rather than a
series of individual and speciﬁc ice marginal positions. The ice
marginal zone straddles the last extensive zone of end moraines
(including the Ra Moraine system in Norway, the Middle-Swedish
end moraine zone, the Salpausselk€a I and II moraines in Finland,
and the Koitere and probably Rugozero moraines in Russia; Figs. 5
and 6) and it divides a period of retreat interspersed with standstills
and re-advances since the LGM from a period with fewer interruptions
in retreat up to ﬁnal deglaciation in the northwestern
Swedish Mountains.
The radiocarbon, OSL, and cosmogenic 10
Be and 26
Al exposure
ages included in our deglaciation reconstruction are presented in
the Supplementary dataset. Of the 335 radiocarbon samples, 223
yield minimum age constraints for deglaciation because they are
part of the post-glacial environment and 112 yield maximum age
constraints for deglaciation because they are from sub-till samples.
Calibrated radiocarbon ages range from 7.2 ± 0.1 cal kyr BP to
27.7 ± 0.1 cal kyr BP for minimum age-constraint samples and from
16.3 ± 0.7 cal kyr BP to 33.5 ± 0.5 cal kyr BP for sub-till samples.
While 125 OSL samples are from sediment layers that were
deposited contemporaneously with or postdate the timing of local
ice extent, and yield minimum age constraints for deglaciation, 13
are from sub-till samples, yielding maximum age constraints for
deglaciation. OSL ages range from 5.9 ± 0.7 kyr to 131.0 ± 8.0 kyr for
minimum age-constraint samples, and from 15.1 ± 1.3 kyr to
25.3 ± 1.6 kyr for maximum age-constraint samples. The recalculated
10
Be (26
Al) exposure ages range from 1.1 ± 0.3 kyr to
456 ± 20 kyr (7.8 ± 0.5 kyr to 107 ± 4.6 kyr) with 74% (59%) of the
ages falling in the time window between 9 kyr and 22 kyr. The new
10
Be and 26
Al exposure ages from Sweden, Norway, and Finland are
generally similar to previously published exposure ages, with some
sites yielding well-clustered exposure ages in the deglaciation age
range and others, primarily located in the north and at high altitude,
yielding ages signiﬁcantly older than the last deglaciation due
to cosmogenic inheritance.
4.2. Deglaciation history and dynamics
We have subdivided the Fennoscandian Ice Sheet domain into
four sectors (Fig. 10) to guide our presentation of the deglaciation
history and dynamics of the entire ice sheet. This subdivision is
guided by topography and ice sheet deglaciation dynamics.
4.2.1. The western sector
The Norwegian shelf was deglaciated between the local LGM
and 14e15 cal kyr BP (Andersen, 1979, 1981; Sollid and Torp, 1984).
Because of the high-precipitation coastal setting, the ice sheet
margin in this sector responded rapidly to climatic variations. The
most distinct climate variation, the Younger Dryas cold interval,
produced the most laterally-continuous moraines (Lundqvist,1990;
Andersen et al., 1995a, 1995b). Stratigraphical evidence for Younger
Dryas re-advances of c. 40e50 km have been reported from
southwestern to northern Norway (Mangerud, 1977; Andersen
et al., 1995b; Bergstrøm et al., 2005; Mangerud et al., 2011). There
is limited evidence for ice marginal positions during the Allerød
that could shed light on the pre-Younger Dryas ice sheet geometry
and illustrate to what extent the ice sheet had retreated. Estimates
of ice sheet retreat rely on observations of Allerød sediments,
predominantly marine sediments, that were overrun by the
Younger Dryas ice sheet (Lohne et al., 2007; Mangerud et al., 2011).
Mangerud (1977), for example, reports an estimated minimum
retreat of 40 km in the Bergen area (Fig. 6), which implies extensive
ice free coastal areas in western Norway during the Allerød. Similar
inferences come from the Stavanger area (Fig. 6) and are summarized
by Lohne et al. (2007). Observations from the south coast of
Norway and the Oslo area reveal less extensive Younger Dryas readvances
of <18 km (Sørensen, 1992; Bergstrøm, 1995). It is
thought that the more maritime setting of the western margin, and
its proximity to an Atlantic moisture supply, may explain why readvances
on the western margin were more extensive than
Table 4
Publications with cosmogenic dates included in the Supplementary dataset, including the number of10
Be and26
Al samples.
Publications 10
Be 26
Al Publications 10
Be 26
Al
Alexanderson and Fabel (2015) 8 e Linge et al. (2006a) 41 e
Anjar et al. (2014) 23 e Linge et al. (2006b) 6 e
Briner et al. (2014) 34 e Linge et al. (2007) 25 e
Brook et al. (1996) 4 4 Mangerud et al. (2013) 34 e
Darmody et al. (2008) 2 2 Matthews et al. (2008) 3 3
Fabel et al. (2002) 13 2 Nesje et al. (2007) 16 e
Fabel et al. (2006) 19 17 Paasche et al. (2006) 2 e
Fjellanger et al. (2006) 11 e Rinterknecht et al. (2004) 9 e
Goehring et al. (2008) 69 e Rinterknecht et al. (2005) 41 e
Goehring et al. (2012) 8 e Rinterknecht et al. (2006) 95 e
Goodfellow et al. (2014) 2 2 Rinterknecht et al. (2012) 5 e
Harbor et al. (2006) 3 2 Rinterknecht et al. (2014) 21 e
H€attestrand et al. (2004) 1 e Shakesby et al. (2008) 5 5
Heine et al. (2009) 6 e Stroeven et al. (2002a) 3 e
Houmark-Nielsen et al. (2012) 35 e Stroeven et al. (2002b) 4 3
Jansen et al. (2014) 19 e Stroeven et al. (2006) 3 e
Johnsen et al. (2009) 6 e Stroeven et al. (2011) 12 2
Johnsen et al. (2010) 6 e Stroeven et al. (2015) 10 e
Larsen et al. (2012) 17 e Svendsen et al. (2015) 18 e
Li et al. (2005) 14 e Tschudi et al. (2000) 4 e
Li et al. (2008) e 1 This study 127 31
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10.1016/j.quascirev.2015.09.016
elsewhere. In contrast to both younger and older moraines that
occur in many valleys, the Younger Dryas moraines can frequently
be traced over intervening uplands, with a notable exception for
mid-Norway where the Younger Dryas ice margin is poorly mapped
in forested and alpine areas. Prominent moraines and associated
ice-marginal formations deposited after the Younger Dryas,
observed in fjords and coastal valleys (Sveian et al., 1979; Corner,
1980; Andersen et al., 1981; Nesje and Rye, 1990; Rye et al., 1997),
reﬂect standstills or minor re-advances which record a response of
the dynamic western margin of the Fennoscandian Ice Sheet to the
Preboreal oscillation (PBO). Accurate cartographic representation of
the dynamic ﬂuctuations at the western ice sheet margin is
exceedingly difﬁcult at the ice sheet-scale because of the relatively
high degree of topographic complexity over short spatial scales.
Hence, we have simpliﬁed our depiction as one of slow but steady
retreat over the course of the deglaciation.
4.2.2. The southern sector
The uniform retreat from the local LGM conﬁguration in the
southern sector, including the Main Stationary Line in Denmark, the
Brandenburg in northern Germany, and the Lezno in Poland
(Houmark-Nielsen and Kjær, 2003; Marks, 2012; Rinterknecht
et al., 2014) was interrupted by two Young Baltic advances
creating the East Jylland and Bælthav ice-marginal formations in
Denmark (Fig. 6; Stephan, 2001; Kjær et al., 2003a, b). We follow
Houmark-Nielsen and Kjær (2003) in correlating the East Jylland
with the Frankfurt and Poznan ice-marginal formations in Germany
and Poland, and the Bælthav ice-marginal formation with the
Fig. 9. Deglaciation pattern and chronology for the Fennoscandian Ice Sheet. Ice margins are given at 1-kyr intervals before the Younger Dryas (12.7e11.6 cal kyr BP), that is from
local LGM to 13 cal kyr BP. Post-Younger Dryas margins are shown at 100-year intervals (11.6e9.7 cal kyr BP). Ice margins south and east of the Baltic Sea follow established icemarginal
formations (Figs. 5 and 6) and show lobate outlines related to bedrock type, local topography, and the formation of ice streams. The less lobate ice margin on the
Scandinavian Peninsula (set in shield rocks) is partly a result of the reconstruction technique, and partly reﬂects a difference in ice dynamics. Ice-marginal formations did not form
during the rapid ﬁnal retreat. Consequently, the ice margin is reconstructed transverse to the youngest documented ﬂow traces (predominantly eskers; Fig. 5), and is also constrained
by marginal meltwater channels and ice-dammed lakes.
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Pomeranian ice-marginal formation in Germany (Fig. 6). The Young
Baltic advances show a highly lobate ice margin including a
northerly ﬂow direction in some inter-island straits. A comparison
between the 16 and 14 cal kyr BP ice margin conﬁgurations in Fig. 9
clearly shows the remarkable contrast in retreat rates of the slow
terrestrial margin retreat in western and southern Sweden and the
rapid retreat of the calving ice margin in the Baltic Basin following
the Young Baltic advances (Duphorn et al., 1979). This may be
explained by the 16 cal kyr BP position representing an overextended
ice margin, related to the previous surge in the southernmost
Baltic Basin. A thin surge lobe would be highly susceptible
to rapid calving and break-up.
4.2.3. The eastern sector
The local LGM was attained as early as 19 cal kyr BP in the
southern part of the eastern sector (and has been correlated with
the East Jylland-, Frankfurt-, and Poznan ice-marginal formations)
and as late as 17 cal kyr BP in the northern part of the eastern sector
(Kalm, 2012; Larsen et al., 2014; Lyså et al., 2014), which broadly
correlates with the Bælthav- and Pomeranian ice-marginal formations
(e.g., Houmark-Nielsen and Kjær, 2003). From its local LGM
extent, the ice margin retreated into a depression spanning the
entire length from northern Germany to Lake Onega in western
Russia (Fig. 6; Rainio et al., 1995). This led to the formation of a
number of ice-dammed lakes in Poland, Lithuania, Latvia, Estonia
and Russia, and later the formation of the Baltic Ice Lake (Fig. 3). The
latter had its outlets in the €Oresund and Storebælt depressions
(Figs. 3 and 6) across a landbridge between Denmark and Sweden.
Although the Baltic Ice Lake persisted until the end of the Younger
Dryas, at approximately 11,620 ± 100 cal years BP (Stroeven et al.,
2015), it may have drained once previously at c. 13.0 cal kyr BP
(Bj€orck, 2008; Bj€orck et al., 1996; Boden et al., 1997).
4.2.4. The central sector
4.2.4.1. Southern Sweden. The deglaciation of southern Sweden
before the onset of the Younger Dryas was characterised by a predominantly
terrestrial and slow (<150 m yrÀ1
) ice marginal retreat.
This retreat was repeatedly interrupted by still-stands and minor
re-advances (Lundqvist and Wohlfarth, 2001), as indicated by a
series of ice-marginal formations that are particularly well developed
on the western side of the peninsula (Fig. 11). In contrast,
there is only one signiﬁcant ice-marginal formation on the eastern
side of the peninsula; the Vimmerby Moraine (Agrell et al., 1976;
Malmberg Persson et al., 2007; Johnsen et al., 2009), and there
has been no conclusive correlation between it and the ice-marginal
formations on the western side. LiDAR elevation data, which has
recently become available over southern Sweden (Dowling et al.,
2013), has enabled the identiﬁcation of these ice-marginal formations
over longer lateral extents than could previously be identiﬁed.
For example, the Vimmerby Moraine has a clear morphological
continuation to the west and can be connected to the Berghem
Moraine (Fig. 11). This physical evidence supports what has previously
been suspected by others (e.g., Anjar et al., 2014), and it
provides, for the ﬁrst time, a strong spatial link between the
deglaciation chronologies on the western and eastern sides of
southern Sweden (see 5.2.). A well-developed pattern of esker
systems and ubiquitous glacial lineations indicate that basal conditions
were mostly warm-based in southern Sweden during
deglaciation. However, parts of southern Sweden probably
remained cold-based up until ﬁnal deglaciation, resulting in limited
subglacial erosion, as indicated by numerous pre-LGM glacial deposits
in the area (Lemdahl et al., 2013), and extensive ribbed
moraine ﬁelds (Kleman and H€attestrand, 1999; M€oller, 2010).
4.2.4.2. Central Sweden and southeastern Norway. Here the icemargin
retreat was mostly terrestrial. The topography ranges
from high-relief, over the plateaus and mountains of south-central
Norway, to low-relief in central Sweden. The ice sheet was generally
wet-based and left a ubiquitous record of striae, till lineations
and eskers, which enables an accurate reconstruction of the ice
margin retreat pattern compared with other sectors, although
dating constraints are scarce from the inner areas (Fig. 7). The
deglacial landform pattern reveals that the southwestern sector of
the Fennoscandian Ice Sheet retreated towards the mountains of
south-central Norway while the southeastern sector continued its
northward retreat, yielding a narrow ice ridge (saddle) at
10.2 cal kyr BP that connected residual ice in south-central Norway
to the main ice sheet remnant in the north. This saddle was an
obstruction to meltwater drainage and large ice-dammed lakes
consequently formed in Østerdalen (Nedre Glåmsjø ice lake) and
adjacent valleys in Norway (Fig. 3), the drainage of which has been
described by Garnes and Bergersen (1980) and Longva and
Thoresen (1991). During this ice sheet conﬁguration the Ljungan-,
Ljusnan-, and the central J€amtland complex of ice-dammed lakes
began to develop in west-central Sweden, on the east side of the
bordering topographic ice sheet saddle (Lundqvist, 1969, 1972,
1973; Borgstr€om, 1989).
4.2.4.3. Southern and central Finland. The most conspicuous glacial
landforms are three lobate ridges, from outer to inner e the Salpausselk€a
(Ss) I, II, and III moraines, respectively (Fig. 6). Ss I and II
are considered to be of Younger Dryas age (Donner, 1978, 2010;
Rainio et al., 1995; Rinterknecht et al., 2004) and are therefore
synchronous with the Younger Dryas moraines in Norway, Sweden
and Russia (Rainio et al., 1995). From the pattern of lineations
immediately distal to the ridges, retreat of the Fennoscandian Ice
Sheet just prior to the deposition of the Ss I and II displayed a nonlobate
structure. The well-developed drumlin zones proximal to
the Salpausselk€a moraines indicate that laterally well-constrained
ice streams developed early in the Younger Dryas, feeding the
fan-like ﬂow structure of advancing ice sheet lobes. There is
stratigraphical indication that a substantial re-advance of the ice
Fig. 10. The Fennoscandian Ice Sheet area divided into four geographical and icedynamical
sectors which are discussed individually in the text. The Central sector is
further subdivided into ﬁve regions because of internal differences in ice dynamics
during deglaciation.
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margin, up to 50 km (Rainio, 1993), occurred as part of the building
of the Ss I Moraine. The two feeder ice streams (Punkari, 1995)
transported large volumes of ice and were probably instrumental in
shifting the dispersal centre of the ice sheet to the west during the
Younger Dryas (Kleman et al., 1997). Areas between these ice
streams, such as coastal Ostrobothnia, were subjected to slow and
probably cold-based sheet ﬂow, which preserved older landforms
and organic deposits (Niemel€a and Tynni, 1979; Salonen et al.,
2008; Pitk€aranta et al., 2014).
4.2.4.4. The Baltic Basin. The Baltic Basin is around 1500 km long
and bathymetrically separated into many sub-basins (Fig. 1).
Although water depths remain generally shallower than 100 m,
three of the basins reach depths of 250e460 m. Retreat of the ice
sheet mostly occurred with the ice margin orientated perpendicular
to the long axis of the Baltic Basin. Ice sheet retreat deviates
from this pattern only in the Gulf of Bothnia where it retreated
obliquely to the basin long axis towards the location of the ﬁnal ice
remnants in the northern Swedish mountains (Andren, 1990). In
the absence of published regional ﬂow traces on the Baltic Sea ﬂoor,
the ice dynamics in the basin have been mostly inferred from clay
varves, and striae and other landforms yielding ice-marginal positions
and ﬂow traces on adjacent coasts (H€ornsten, 1964;
Bergstr€om, 1968; F€oz€o, 1980; Lundqvist, 1987; Str€omberg, 1989).
In contrast to North America, where rapid calving in the interior
basin of Hudson Bay divided the ice sheet in discrete smaller
remnants (Dyke, 2004), no such separation seems to have occurred
in Fennoscandia.
Retreat along the Baltic Basin started with the decay or collapse
of the extended terminal lobe of the Bælthav Young Baltic ice
stream, at around 16.5 cal kyr BP. This lobe extended northwards
into the Storebælt and €Oresund straits, and was sufﬁciently thin for
its ﬂow pattern to be governed by the modest topographic relief of
the southern Baltic Basin, where present-day water depths remain
<100 m. Retreat through the southern and central Baltic Basin, up
to the Younger Dryas position, took approximately 4000 years. The
Fig. 11. Major end moraine belts in southern Sweden mapped from LiDAR data (Swedish national elevation model “GSD-H€ojddata, grid 2þ”, with a 2 m horizontal resolution). The
moraine pattern correlates well with earlier maps (e.g., Lundqvist and Wohlfarth, 2001), although we demonstrate, for the ﬁrst time, that there is a morphological continuation of
the Berghem Moraine across southern Sweden to the Vimmerby Moraine. The age assignments (in cal kyr BP) are based on radiocarbon dates (summarized in Lundqvist and
Wohlfarth, 2001), cosmogenic isotope dates (Larsen et al., 2012; Anjar et al., 2014), and, primarily, correlations with the Greenland ice core event stratigraphy (Rasmussen
et al., 2014; see also Table 5, Fig. 13).
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lobate moraines east and south of the Baltic Basin are indicative of
ice streaming in the deeper parts of the basin, particularly east of
Gotland (Fig. 6). During the Younger Dryas, the ice margin appears
to have stabilized around 150 km south of Åland, although its
precise location remains elusive (Noormets and Floden, 2002).
North of Åland, the rate of retreat sharply increased, and it appears
that at least one major surge occurred in the Gulf of Bothnia, at
around 10.8 cal kyr BP (Sandegren, 1929; Str€omberg, 1989;
Lundqvist, 2007; Kleman and Applegate, 2014). This was possibly
caused by a sudden and widespread change in subglacial conditions
from cold- to warm-based (Str€omberg, 1989; Kleman and
Applegate, 2014). A relatively deep basin in the seaﬂoor south of
Umeå (Fig. 6) may have been the initiation point for this surge, as
evidenced by a convergence of striae on the coast towards this
depression, and new geomorphological data from the central Gulf
of Bothnia ﬂoor showing typical signs of ice streaming, such as
highly elongated glacial lineations (Greenwood et al., in press).
Perhaps coeval to surging in the Gulf of Bothnia, the ice sheet
margin in Finland surged as well. Initially, retreat proceeded towards
the northwest, but it was interrupted by a standstill and readvance
indicated by the Central Finland Ice-Marginal Formation,
CFIMF (Rainio et al., 1986). The CFIMF (Figs. 5, 6 and 9) is limited to
Central Finland. It consists of large glacioﬂuvial deposits (plateaux
and ridges) and end moraines. Stratigraphical evidence, differing
patterns of striae and eskers on proximal and distal sides, and its
lobate outline, are all consistent with an ice margin advance
through surging. Inwards from this location, ice margin retreat
proceeded towards the northwest, and appears to have been rapid
and uninterrupted by dynamic events, such as renewed ice
streaming. The ice sheet retreated onshore at around 10 cal kyr BP.
4.2.4.5. The ﬁnal deglaciation of northern Fennoscandia.
Although deglacial landforms and deposits are common in northern
Fennoscandia, inherited deglacial landforms and deposits are
regionally important (Lagerb€ack and Robertsson, 1988; Kleman,
1990; H€attestrand, 1998). On many western and northern uplands
only a Late Quaternary weathering mantle is present (Goodfellow
et al., 2014), capping a non-glacial bedrock morphology (Kleman
and Stroeven, 1997; Goodfellow et al., 2008). Although repeated
glaciations on the Kola Peninsula, Russia, have left almost no lithostratigraphical
record (Niemel€a et al., 1993), there is a discernible
deglaciation record from meltwater landforms (H€attestrand and
Clark, 2006a, 2006b; H€attestrand et al., 2007). In much of northern
Fennoscandia, the Late Weichselian till is a thin (1e3 m) veneer
atop older glacial and non-glacial deposits (Nordkalott Project,
1986; Kleman et al., 2008). The pre-Late Weichselian glacial landscape
in northern Sweden and Finland is the most important
inherited element in the Fennoscandian glacial landscape. It consists
of widespread Veiki moraine (dead ice topography), as well as
drumlins and eskers indicating a regional ice ﬂow towards the
southeast (Hoppe, 1959; Lagerb€ack, 1988; Kleman, 1992;
H€attestrand, 1998; Kleman et al., 2006). The deglacial landform
imprint of mainly lineations and eskers is locally well-developed,
such as in northern Norway and throughout northeastern Sweden,
where this landform assemblage can be traced inland to the
position of the last ice remnant in the northern Swedish mountains
(Nordkalott Project, 1986; H€attestrand et al., 1999; Dellgar
Hagstr€om, 2006; Stroeven et al., 2011). Widespread deglaciation
landforms also occur inland from the Bothnian coast (Fig. 5). In this
area, abundant De Geer moraines indicate deglaciation at a calving
ice margin (Hoppe, 1948). Towards the interior areas of northern
Sweden, deglacial landforms are less developed, especially inland
of a series of subglacial ﬂuvial gorges cut at the highest coastline
(Jansen et al., 2014), and they are often superimposed on landforms
of Early Weichselian age. Glacial lakes constrain 10.1e9.7 cal kyr BP
marginal positions on the eastern side of the Scandinavian Mountains
and, during the earliest Holocene, most of the river valleys in
northern Sweden were locations for repeated drainage events in
response to ice margin failures (Elfstr€om,1987). The last remnant of
the Fennoscandian Ice Sheet vanished after 9.7 cal kyr BP in the
eastern Sarek Mountains of northern Sweden (Lundqvist, 1986;
Kleman, 1990; Kleman et al., 1997; Boulton et al., 2001).
5. Discussion
The deglaciation chronology of the Fennoscandian Ice Sheet
presented here is consistent with the last two published deglaciation
chronologies by Kleman et al. (1997) and Boulton et al. (2001),
and adds considerable new detail. Our reconstruction is based on
additional geomorphological constraints, including many publications
addressing the distribution and chronology of marginal positions
in the eastern sector (Fig. S1, Supplementary dataset), which
were poorly covered, and a ﬁrst compilation of eskers across the
glaciated domain (Fig. 5). Compared to the previous reconstructions,
there are now more abundant radiometric constraints,
including cosmogenic nuclide constraints which were
unavailable at the time (Supplementary dataset). Hence, our
reconstruction is much improved for the southern and eastern ice
sheet sectors. Below, we discuss some of these aspects and analyse
their implication for ice sheet dynamics.
5.1. Uncertainty in the timing of ice sheet deglaciation
The position of the LGM ice margin around the perimeter of the
Fennoscandian Ice Sheet is relatively well established (Ehlers and
Gibbard, 2004), even if opinions can diverge on when the local
maximum position was attained. However, one area where
considerable uncertainties persist regarding the LGM extent of the
ice sheet is in northwestern Russia. Here, different reconstructions,
even those published in the last 10 years, place the LGM ice margin
hundreds of kilometres apart. We have largely adopted the LGMreconstruction
suggested by Larsen et al. (2014) and Lyså et al.
(2014). They base their suggested ice marginal outline and chronology
on a recent mapping of marginal moraines by Fredin et al.
(2012) and a new dataset of OSL and radiocarbon ages from
below and above LGM sediments (Lyså et al., 2014; see
Supplementary dataset). Combined, these data indicate that the
LGM ice margin was located further east than previously suggested
(Fredin et al., 2012; Larsen et al., 2014; Lyså et al., 2014), and peaked
at a later time than for the rest of the ice sheet perimeter, at around
17e15 cal kyr BP (Larsen et al., 1999; Lyså et al., 2014).
Fig. 12 shows the deviation of individual sample ages from the
reconstructed deglaciation age for that location. The scatter in ages
is signiﬁcant (Fig. 12bee) and, consequently, a large number of
sample ages do not match the reconstructed deglaciation chronology.
Of the radiocarbon ages, only 12% (0%) of the minimum
(maximum) ages overlap within uncertainties with the reconstructed
age whereas 71% (2%) reside within a time window of
±2 kyr around the reconstructed deglaciation age. Of the OSL ages,
29% (31%) of the minimum (maximum) ages overlap within uncertainties
with the reconstructed age and 54% (31%) reside within
a time window of ±2 kyr around the reconstructed deglaciation
age. Of the 10
Be (26
Al) exposure ages, 27% (28%) overlap within
uncertainties with the reconstructed age and 51% (41%) reside
within a time window of ±2 kyr around the reconstructed deglaciation
age. These numbers help illustrate the relatively poor ﬁt of
individual sample ages to the reconstructed deglaciation age and
highlight the challenges associated with dating ice margin retreat
using these techniques. The radiocarbon ages yield better results
than the cosmogenic and OSL ages. Individual sample ages show
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more disparity with the reconstructed age in the early part of the
deglaciation than in the latter part of the deglaciation. However,
this is not generally true for cosmogenic ages because inheritance is
a problem for many samples collected from surfaces preserved
under cold-based non-erosive ice (e.g., Fabel et al., 2002; Stroeven
et al., 2002b; Goehring et al., 2008).
The spatial and chronological uncertainties for the deglaciation
reconstruction vary across the paleo-ice sheet domain. Generally,
the uncertainties are lowest for the post-Younger Dryas deglaciation,
which is well constrained by geomorphology and the clayvarve
chronology. Distal to the Younger Dryas margin the uncertainties
increase progressively with the largest uncertainties
attained for the eastern and southeastern sectors of the LGM
conﬁguration. It is difﬁcult to estimate the uncertainties because
the reconstruction is based on a range of geomorphological and
chronological data with highly different characteristics and problems
potentially causing errors in the reconstruction. While the
uncertainty for the post-Younger Dryas deglaciation is in the range
of 100e500 years, the uncertainty in the earlier part of the deglaciation
chronology is much higher, perhaps 500e2000 years.
5.2. A climatic imprint on Fennoscandian ice margin behaviour: a
southern Sweden case study
We consider the GRIP/GISP2/NGRIP harmonized Greenland ice
core record of Rasmussen et al. (2014), dated by annual layer
counting, as the best calendar-year record of the Younger Dryas
climatic event, which left a prominent glacial geological record in
Fennoscandia (Lundqvist, 1990; Andersen et al., 1995a). According
to Rasmussen et al. (2014), this climatic event spans the period
12.8e11.7 cal kyr BP in the Greenland ice core record (Fig. 13). The
same record also shows ﬁve cold phases in the period leading up to
the Younger Dryas (18e12.8 cal kyr BP), which is largely coincident
with Greenland interstadial 1 and is the time during which
southern Sweden became deglaciated (Hillefors, 1969; Berglund,
1976; Lundqvist and Wohlfarth, 2001). The deposition of the
Fig. 12. Age differences between the reconstructed deglaciation age (Fig. 9) and individual cosmogenic (10
Be), radiocarbon (14
C), and OSL ages (Supplementary dataset). a)
Interpolated deglaciation reconstruction including 10
Be, 14
C, and OSL sample locations within the local LGM domain. b) Deviation from the reconstructed deglaciation age for 10
Be
exposure ages (sample age minus reconstructed age). c) Deviation from the reconstructed deglaciation age for 14
C ages (sample age minus reconstructed age). d) Deviation from the
reconstructed deglaciation age for OSL ages (sample age minus reconstructed age). e) Deviation of >10 kyr from the reconstructed deglaciation age for OSL, 14
C, and 10
Be ages
(sample age minus reconstructed age).
A.P. Stroeven et al. / Quaternary Science Reviews xxx (2015) 1e31 19
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patchy Halland Coastal moraine belt, which is laterally conﬁned to
about 75 km, was followed by the deposition of four morecontinuous
and laterally-extensive end moraine belts, the
G€oteborg, Berghem-Vimmerby, Trollh€attan, and Levene (Table 5,
Figs. 5, 6 and 11). Each of these ﬁve moraine belts appears to have
generally formed in response to a stand-still or re-advance of the
Fennoscandian Ice Sheet. We have not identiﬁed evidence, such as
deviating striae patterns or lobate margin outlines, which would
indicate that these moraines are a result of internal ice sheet dynamics
(surge moraines). An exception is the central part of the
Vimmerby Moraine, which has a lobate outline with a splaying
pattern of glacial lineations on its proximal side. However,
considering the clear lateral continuation of the Vimmerby Moraine
to the Berghem Moraine, ice-dynamic oscillation appears to have
occurred on only a minor part of the continuous BerghemVimmerby
ice-marginal formation. Hence, the South Swedish
moraines may offer the possibility to directly link the proxy climatic
record of Greenland to southern Swedish glacial geology.
The Halland Coastal Moraine belt (Caldenius, 1942; Fernlund,
1993) is the oldest in Sweden. It differs from the other moraine
belts in southern Sweden because, rather than being a semicontinuous
single or double ridge, it is a zone up to 15 km wide
with numerous parallel moraines. Individual segments are rarely
more than a few kilometers long and some resemble De Geer, or
washboard, moraines. The Halland Coastal Moraine belt is estimated
to have an age of 18e16 cal kyr BP based on radiocarbon ages
(Lundqvist and Wohlfarth, 2001), and 17.0 ± 0.9 to 16.8 ± 1.0 cal kyr
BP according to recent cosmogenic isotope dating (Anjar et al.,
2014; Larsen et al., 2012). Within error margins, these ages correspond
with the onset of Greenland interstadial GS-2.1a at 17.4 ka
(Rasmussen et al., 2014).
The G€oteborg Moraine (Hillefors, 1975) is a well-deﬁned icemarginal
formation and, where it runs parallel to the west coast of
Sweden, it also approximates the marine limit. Consequently, there
may be some dynamic control on the location of this moraine
segment. The 700 years preceding the GI-1d event comprises the
warm B€olling phase (GI-1e; Fig. 13), during which climaticallycontrolled
moraine formation appears unlikely. We therefore
assign the G€oteborg Moraine a youngest possible age of 14.7 cal kyr
BP (Table 5), which marks the ﬁnal cold stage of stadial GS-2.1. It has
been suggested that the G€oteborg Moraine continues eastward into
a hummocky moraine zone on the southern Swedish highlands
(e.g., M€oller, 2010; Anjar et al., 2014) and similarly that other westSwedish
ice marginal formations continue eastwards into hummocky
moraine equivalents on the southern Swedish highlands
(e.g., Lundqvist and Wohlfarth, 2001). However, the new LiDAR
coverage of southern Sweden leads us to question this linkage
eastwards of the G€oteborg Moraine. Using this data we ﬁnd that,
Fig. 13. Subset of the Greenland ice core event stratigraphy (from Rasmussen et al., 2014, Fig. 1), with, indicated, the proposed correlation of the south Swedish end moraine belts
with ice core cold stages (Fig. 11). Not indicated, but mentioned in the text, is evidence of the response of the Fennoscandian Ice Sheet to the “11.4 ka event”, or the Preboreal
oscillation, in Norway and in Finland (Ss III). (For interpretation of the references to colour in this ﬁgure legend, the reader is referred to the web version of this article.)
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where the G€oteborg Moraine laps onto the southern Swedish
highlands hummocky moraine zone, it appears as one or two
distinct moraine ridges superimposed on a hummocky moraine
zone that lacks clear proximal and distal borders. Hence, there are
no morphologically-based arguments to support a genetic relationship
between them.
The most logical correlation for the Berghem-Vimmerby
Moraine is the GI-1d cold phase at 14.0 cal kyr BP (Fig. 13). There
is, however, a discrepancy between the varve chronology in Fig. 8
and the age assignment of the Vimmerby Moraine in Fig. 11. According
to the varve chronology, deglaciation from the Vimmerby
Moraine occurred at c.13.5 cal kyr BP, whereas absolute dating and
our Greenland ice sheet correlation indicate that the combined
Berghem-Vimmerby line is no younger than 14.0 cal kyr BP. Presently,
these two data sets appear irreconcilable and, because of the
difﬁculties in extending the varve chronology across the Younger
Dryas zone, we have chosen to adhere to the radiocarbon chronology
in southeastern Sweden, rather than the varve chronology.
New clay varve locations south of the Younger Dryas zone, not yet
connected to the STS and where tephra (Vedde ash) has been found,
may help to resolve this issue (MacLeod et al., 2014).
The Trollh€attan Moraine is smaller and less continuous than the
Berghem-Vimmerby and G€oteborg moraines. We consider a correlation
between the relatively small Trollh€attan Moraine and a
short-lived and low-magnitude cold phase (event GI-1c2; Fig. 13),
to be most plausible. Cold event GI-1c2 postdates the GI-1d cold
phase by about 400 years (Table 5).
The Levene Moraine and its extension in the Oslofjord area, the
Onsøy Moraine (Fig. 6; Sørensen, 1992), is the youngest of the four
ice-marginal formations. These moraines are especially important
because of their lateral extent (>250 km). They form a narrow belt
located 5e20 km outside Younger Dryas-age ice-marginal formations
(Berglund, 1979; Hillefors, 1979; Lundqvist and Wohlfarth,
2001). The Levene Moraine likely reﬂects the ice sheet margin
reacting to a short-lived climatic deterioration of regional signiﬁcance,
shortly before the Younger Dryas event. The regional character
of this climate event is further illustrated by ice margin
deposits of 13.1e13.2 cal kyr BP in Central Norway about 20 km
outside local Younger Dryas moraines (Olsen et al., 2013a). A
comparison with the harmonized Greenland ice core record indicates
that it correlates with a short, sharp, cold phase in
Greenland interstadial 1 (event GI-1b; Table 5, Fig. 13) at around
13.2 cal kyr BP.
Our chronology is based on the premise that there is a correlation
between the southern Swedish ice-marginal formations and
the harmonized Greenland ice core cold stages. If this is correct, we
argue that the correlation scheme as presented in Table 5 is the
most likely and also ﬁts within 200 years of the error margins of
available radiocarbon ages for the moraines, as summarized in
Lundqvist and Wohlfarth (2001). This chronology for the southern
Swedish ice-marginal formations is younger than recently suggested
by Larsen et al. (2012) and Anjar et al. (2014) on the basis of
cosmogenic exposure ages. However, sampling of the G€oteborg
Moraine for the Larsen et al. (2012) study was done on the portion
of this moraine that approximates the marine limit and which is
closely juxtaposed to, and possibly merges with, the Halland
Coastal Moraine belt. This portion of the moraine appears to have
been dynamically controlled and the ice margin may have persisted
at this location throughout GS-2.1a (Fig. 13). The Anjar et al. (2014)
age assignments for some locations in southern-most Sweden are
incompatible with age assignments for ice marginal positions in
northern Germany and Denmark. OSL, 14
C, and other cosmogenic
isotope data indicate that the ice margin at this time was positioned
close to the northern coasts of Germany and Poland (Marks, 2012),
and in eastern Denmark (Houmark-Nielsen and Kjær, 2003),
respectively. Our reconstruction conforms to this latter, more
extensive, data set.
5.3. Climatic and dynamic controls on moraine formation: the
Younger Dryas case study
From the record of ice-marginal formations, we infer that the
ice sheet margin periodically halted in its retreat or re-advanced.
An important question when utilizing such ice-marginal formations
for the tracing of retreat patterns, is what controlled the
interruption in retreat? The two end member answers are full
climatic control, with essentially unchanged ice dynamics, and
dynamic control, irrespective of climatic trends. Interruptions in
retreat through dynamic controls typically result in ice margin
advances as part of surges. Moraines that form from surges are
well-known from the periphery of Vatnaj€okull, Iceland, and other
major surging glaciers (Evans and Rea, 1999; Benediktsson et al.,
2010; Schomacker et al., 2014). The Younger Dryas cold period
resulted in an almost continuous belt of ice-marginal formations,
which we analyse in terms of the controls behind the margin
response.
We regard the Younger Dryas moraines in Norway and Sweden
as classical examples of the effect of climatically controlled
ice margin responses, i.e. they mark standstills, sometimes with
minor to modest re-advances immediately prior to moraine formation.
In Sweden, there is no evidence in the landform record
for ice streaming proximal to the Younger Dryas moraines or for
major lobation (which is an indicator of strong lateral velocity
gradients caused by ice streaming; Stokes and Clark, 1999). The
extreme topographic relief of Norway renders simultaneous
surging in a large number of outlets as an extremely unlikely
scenario.
The situation with regards to the Salpausselk€a moraines in
Finland is, however, more complex. The Ss I and Ss II moraines
(Fig. 5) are also of approximately Younger Dryas age, with Ss I likely
being formed at the start of the Younger Dryas cold period (Rainio
et al., 1995). Thus, they are broadly linked to a major cold event and
have traditionally been seen as correlatives to the Younger Dryas
moraines in Norway and Sweden. Ss II continues northeastward in
Table 5
Southern Swedish moraine chronology from radiocarbon and varve dating (Lundqvist and Wohlfarth, 2001) and cosmogenic nuclide dating (Larsen et al., 2012; Anjar et al.,
2014). We infer a slightly different chronology (this study) based on a one-to-one matching with cold phases in the harmonized Greenland ice core record of Rasmussen et al.
(2014). Ages expressed in cal kyr BP.
Moraine Lundqvist and Wohlfarth (2001) Anjar et al. (2014) Larsen et al. (2012) Rasmussen et al. (2014) This study
Levene 13.4 13.8 ± 0.8 13.6 ± 0.9 13.2 (GI-1b) 13.2
Trollh€attan 14.2e13.4 13.6 (GI-1c2) 13.6
Berghema
14.4e14.2 14.0 (GI-1d) 14.0
Vimmerbya
14.6e14.5 14.0
G€oteborg 15.4e14.5 16.1 ± 0.9 16.2 ± 0.9 15.0e14.7 (GS-2.1a) 14.7
Halland Coastal 18.0e16.0 17.0 ± 0.9 16.8 ± 1.0 17.4e16.0 17.4e16.0
a
In this study the Berghem Moraine is correlated with the Vimmerby Moraine.
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the Koitere Moraine and probably also further in the Rugozero
Moraine (Fig. 6; Rainio et al., 1995). This continuity, over a 1000 km
of ice margin distance, and with, at least, the Koitere Moraine
fronting a sector of the ice sheet with little evidence for fast ice ﬂow
dynamics, supports the inference of climatic control for the formation
of Ss II. The occurrence of paired deltas with an elevation
difference of 26e28 m, immediately north of the Ss II, and in all
likelihood reﬂecting the late Younger Dryas drainage of the Baltic
Ice Lake, ﬁrmly anchors the Ss II to a Younger Dryas age (Donner,
1978, 2010; Saarnisto and Saarinen, 2001). Hence, we also regard
the Ss II as a robust part of the circum-Fennoscandian Younger
Dryas ice-marginal formation.
We note that Ss I, in contrast to Ss II, only fronts the two major
lobes in southern Finland and does not continue in a northeasterly
direction (Fig. 5). Although it appears that the Ss I is of early
Younger Dryas age (Rainio et al., 1995), we suspect that its formation
is not directly related to the Younger Dryas climatic cooling.
Based on the pattern of landforms proximal to the Ss I, we instead
infer that it represents a surge moraine, located at the outermost
position of ice sheet lobes that indicate the development of two ice
streams. In contrast, the striae record immediately distal of the Ss I
indicates that sheet ﬂow immediately precede the Ss I event. This
indicates that fast ice ﬂow that formed the feeder ice streams
commenced shortly before the Younger Dryas, during the late
Allerød.
We propose that pronounced warming and surface melting
during the Allerød warm period triggered the onset of ice streaming,
which resulted in the formation of Ss I as a surge moraine. We
further consider it plausible that this phase of ice streaming
continued throughout the Younger Dryas cold period, preventing
fast collapse of the lobes, and that Ss II, and possibly also Ss III, were
formed during climatically-controlled standstills during the retreat
from the non-climatically controlled Ss I. In summary, our view,
represented in the deglaciation map (Fig. 9), is that the Ss I and Ss II
moraines fall within the Younger Dryas chronozone, and that only
Ss II and Ss III are climatically controlled moraines and, in that
sense, Ss II is directly correlative to the Norwegian and Swedish
Younger Dryas moraines.
Ss III is a c. 10 km-wide ice-marginal formation composed of end
moraines, glacioﬂuvial ridges, and deltas. After formation of the Ss I
and Ss II during the Younger Dryas, a rapid deglaciation
commenced and persisted for c. 130 years until, in south-western
Finland, it was interrupted by slow retreat, standstill, and readvance
of the ice margin during c. 160 years (Str€omberg, 2005).
A re-advance of the ice margin is indicated by the end moraine
stratigraphy (Salonen, 1990). No retarded retreat of the ice margin
has been recorded in Sweden (Brunnberg, 1995), even though large
glacioﬂuvial deposits 20e25 km south of Stockholm were suggested
to be correlatives (Nilsson, 1968). A more likely counterpart
to Ss III is the Jaamankangas-Pielisj€arvi formation in northern
Karelia (Figs. 5 and 6; Rainio et al., 1995).
We regard this temporary halt in retreat, and limited readvance,
to be the result of colder conditions during the shortlived
Preboreal oscillation (PBO) or “11.4 ka event” (Rasmussen
et al., 2014, Fig. 13). Bj€orck et al. (1997) ﬁrst proposed a formation
of Ss III during the PBO, a short-lived event that commenced 300
years after the end of the Younger Dryas, even though there was a
mismatch between the timing of the event in the ice core (GRIP),
dendrochronology, and varve chronologies (Bj€orck et al., 1996).
However, with the new timing of the PBO in the Greenland ice core
event stratigraphy of Rasmussen et al. (2014), the PBO is dated at
11.47e11.35 cal kyr BP, only about 130e250 years after the end of
the Younger Dryas, and is therefore well aligned with varve evidence
for the age of Ss III.
5.4. Bedrock control on the behaviour of ice sheet margins: a
northwestern Russia case study
Ice sheet margins in the southern and northeastern sectors of
the Fennoscandian Ice Sheet differ in style from ice margins in
shield bedrock areas. In the latter, margins tend to be slightly
curved, such as is the case with most of the Younger Dryas ice
margin. This contrasts with the pre-Younger Dryas ice-marginal
formations deposited in non-shield areas, which reﬂect ice margins
that were much more topographically controlled and therefore
show a more lobate pattern (Kalm and Gorlach, 2014). In the
southern sector, such lobes typically extended tens of kilometres
beyond adjacent ice margins, with a maximum of up to 100 km for
ice lobes southeast of the Baltic Basin. In the eastern sector, ice
lobes extended several hundreds of kilometres beyond the adjacent
ice margin. There were three major lobes of this type located in the
Dvina, Vologda, and Rybinsk basins (Fig. 6), and these are uniquely
long for the Fennoscandian Ice Sheet. This situation is analogous to
the southern margin of the Laurentide Ice Sheet, where ice marginal
retreat started at 23 kyr (Ullman et al., 2015), with the
exception of two ice lobes hundreds of kilometres long (Des Moines
lobe and James lobe; Clayton and Moran, 1982), which attained
their LGM maximum position long after adjacent ice margins,
because of ice-dynamical factors rather than climatic drivers. For
the ice lobes in NW Russia, Larsen et al. (2014) argued that the
advances were due to ice-bed decoupling caused by a combination
of successive damming of pro-glacial lakes and ice overriding
waterlain sediments and tills with low shear strength. This enabled
fast-ﬂowing low-gradient ice lobes to expand into the exceptionally
wide valleys and basins of the northwestern Russian Plain. We note
also that these three major ice lobes at the eastern Fennoscandian
Ice Sheet margin extended beyond major basins (the White Sea,
Lake Onega and Lake Ladoga; Fig. 6). These basins provided the
low-shear sediments and probably acted as conduits for fast ice
ﬂow, feeding the lobes with ice from interior parts of the ice sheet,
possibly as a result of thawing at the ice sheet bed during early
deglaciation. We infer that these dynamic ice sheet responses led to
an overstretching of the ice sheet surface proﬁle as it created the ice
lobes. Given the extremely low ice surface gradients, and the warm
climate following these expansions, the ice-marginal retreat
following the expansion occurred uninterrupted by stillstands and
possibly included local areal down-wasting.
5.5. Retreat rate across different sectors
Retreat rates for different sectors of the ice sheet for the post-
16 cal kyr BP evolution differ drastically. Total retreat distances
towards the two ﬁnal retreat centers (Figs. 9 and 14) for the
southern and eastern sectors of the ice sheet during this period
were on the order of 1500 km, whereas retreat distances on the
high-elevation and maritime Norwegian margin during the same
period amounted to roughly 150 km. Moreover, the western margin
was also much closer to a state of equilibrium during overall ice
sheet retreat as witnessed by several standstills and concomitant
pulses of moraine formation in response to minor climatic deteriorations,
such as occurred during the Preboreal (Sveian et al.,
1979; Corner, 1980; Sollid and Torp, 1984). No correlative moraines
exist on the eastern ﬂank; here, only the Younger Dryas was a
climatic cooling of sufﬁcient magnitude to temporarily halt the
recession.
In Fig. 14 we compare ice sheet retreat rates along three
southern and eastern transects since local LGM. A clear, and
climatically governed, pattern is apparent in all three proﬁles.
Retreat rates were variable and high, mostly in the 50e200 m yrÀ1
range, between 17 and 13 cal kyr BP, very slow or absent retreat
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during the Younger Dryas, and rapid retreat of 200e1600 m yrÀ1
until ﬁnal deglaciation. Retreat rates were highest where the ice
sheet margin terminated offshore and retreated across a bed with a
reverse slope (deepening water up-ice) and over bathymetric
basins.
5.6. Correlation across the Baltic Basin
Inside the Younger Dryas position, correlation of ice margins
across the Baltic Basin is fairly straightforward, using highprecision
clay varve archives (Str€omberg, 1990, 2005). Distances
across the water are also shorter than in the southern Baltic Basin.
Outside the Younger Dryas zone, correlations across the Baltic Basin
are more difﬁcult to make (De Geer, 1935). This is because of
chronological uncertainties and the topographical complexity of
the basin ﬂoor. The ice-marginal formations in coastal Estonia,
Latvia, and Lithuania indicate a lobe extending southwards in the
Baltic Basin during the 16e14 cal kyr BP period. On the opposing
coast, in eastern Sweden, eskers and striae indicate a SWeNE oriented
ice margin during the same period, i.e. a calving bay
conﬁguration that is difﬁcult to reconcile with the evidence from
the eastern side. We suggest that a southward-directed lobe was
conﬁned to the deeper eastern part of the southern Baltic Basin and
that the ice margin had a pronounced change of direction between
Gotland and the Swedish mainland. The overall conﬁguration of
closely-spaced ice-marginal formations in southern Sweden between
c 17 and 13 cal kyr BP and widely-spaced ice marginal formations
in the same period on the other side of the Baltic Basin,
indicates that southern Sweden was a hinge point which ﬁxed the
swinging gate of the Baltic Basin deglaciation (cf. with
reconstruction of the deglaciation pattern in the Ross Sea Basin by
Conway et al., 1999).
6. Conclusions
Understanding ice sheet deglaciation processes and rates is of
considerable importance given current and projected melting of
the Greenland and Antarctic ice sheets. We have compiled and
synthesized published geomorphological data for eskers, icemarginal
formations, lineations, marginal meltwater channels,
striae, ice-dammed lakes, and geochronological data from radiocarbon,
varve, optically-stimulated luminescence, and cosmogenic
nuclide dating, to provide reconstructions of ice sheet extents for
deglaciation of the Fennoscandian Ice Sheet, in the form of
calendar-year time-slices. This is summarized as a deglaciation map
of the Fennoscandian Ice Sheet (Fig. 9), with isochrons marking
every 1000 years between 22 and 13 cal kyr BP and every hundred
years between 11.6 and ﬁnal ice decay after 9.7 cal kyr BP. We estimate
that the uncertainty for the post-Younger Dryas deglaciation
is 100e500 years, while the uncertainty in the earlier part of the
deglaciation chronology is perhaps 500e2000 years.
The ice margin morphology in plan view varies between sectors;
shield areas tend to produce straight or slightly curved ice margins,
even where there is a considerable topographic relief. Ice margins
in non-shield areas, on the other hand, tend to be highly lobate
because of the strong topographic steering that shallow topographic
depressions exert on low-gradient marginal areas on
deformable beds and sediments with high pore-water pressures.
An extreme example of the latter is in NW Russia, where ice lobes
hundreds of kilometers long extended to their farthest LGM
Fig. 14. Ice sheet retreat rates (in m yrÀ1
) along three southern and eastern transects since local LGM. Although retreat rates along the three proﬁles differ in detail, a slow or absent
retreat during the Younger Dryas (12.8e11.7 cal kyr BP) is apparent in all three proﬁles. The peak retreat rate at 10.3e10.4 cal kyr BP in proﬁle A was probably caused by a retreat of
the ice margin across a bed with a reverse slope (deepening water up-ice) at the eastern side of the Gulf of Bothnia. Similarly, the very rapid retreat seen in proﬁle B, across the Gulf
of Bothnia at 10.5e10.6 cal kyr BP, was related to retreat in deep water, and possibly also reﬂects a rapid disintegration of a post-surge ice lobe.
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position well into the deglaciation period for the ice sheet as a
whole.
Deglaciation patterns vary across the Fennoscandian Ice Sheet,
reﬂecting both differences in climatic and geomorphic settings as
well as differences in ice sheet basal thermal conditions and
terrestrial versus marine termination. The ice sheet margin in the
high-precipitation coastal setting of the western sector responded
sensitively to climatic variations leaving a detailed record of
prominent moraines and ice-marginal deposits in many fjords and
coastal valleys. The southern sector includes a major contrast in
retreat rates of the slow terrestrial margin retreat in western and
southern Sweden and the rapid retreat of the calving ice margin in
the Baltic Basin. Our reconstructions are consistent with much of
the published research. However, the synthesis of a large amount of
existing and new data support reﬁned reconstructions in some
areas. For example, (i) the LGM extent of the ice sheet in northwestern
Russia was located far east and it peaked at a later time
than the rest of the ice sheet perimeter, at around 17e15 cal kyr BP;
and (ii) using new LiDAR data over southern Sweden to correlate
moraines, and coupling the timing of moraine formation to cold
periods in the Greenland ice core, we propose a slightly different
chronology.
Retreat rates vary by as much as an order of magnitude in
different sectors of the ice sheet, with the lowest rates on the highelevation
and maritime Norwegian margin. Retreat rates compared
to the climatic information provided by the NGRIP ice-core record
show a general correspondence between retreat rate and climatic
forcing, although a close match between retreat rate and climate is
unlikely because of other controls, such as topography and marine
versus terrestrial termination. Thus, although many of the ice
marginal features can be interpreted as straightforward responses
to climate variations as recorded in ice cores, in some cases the
relationships are more complex. For example, we conclude that
rapid warming and surface melting during the Allerød warm period
triggered the onset of ice streaming upstream of the Salpausselk€a I
moraine in Finland, and that it represents a surge moraine.
Overall, the time slice reconstructions of Fennoscandian Ice
Sheet deglaciation from 22 to 9.7 cal kyr BP (Fig. 9; provided as
shape ﬁles in the Supplementary dataset) provide an important
dataset for understanding the contexts that underlie spatial and
temporal patterns in retreat of the Fennoscandian Ice Sheet, and are
an important resource for testing and reﬁning ice sheet models.
Future work is needed to improve the chronological control in areas
where dating is sparse, and to test the new interpretations provided
in parts of this paper.
Acknowledgements
We thank Chris Clark and an anonymous reviewer for useful
constructive criticism that improved the clarity of the paper.
Swedish Research Council Grants G-AA/GU 12034-300 and G-AA/
GU 12034-301 to Stroeven and National Science Foundation Grants
OPP-9818162 and OPP-0138486 to Harbor provided partial support
for this work. The Swedish Research Council (2007-15) and Granholmsstiftelse
at Stockholm University generously supported
Stroeven to host the ﬁrst of a couple of deglaciation workshops at
the Tarfala Research Station in late MarcheApril 2007. Jansen was
supported by a UK NERC fellowship (NE/EO14143/1). We thank
Helena Alexanderson, John Clague, Timothy Johnsen, Yingkui Li,
and Kurt Lambeck for their contribution to the 2007 workshop. The
ﬁeld data has been collected with the help of Ingmar Borgstr€om
(2000), Magnus Johansson (2001), Torbj€orn Dahlgren (2002), Karin
Ebert (2002), Angelica Feurdean (2002), Barbara Wohlfarth (2002),
Svante Bj€orck (2003), Eva Sahlin (2003), Vasily Kolka (2004), and
Alexandru T. Codilean (2009). We received considerable support at
three accelerator facilities (PrimeLab, ANSTO, and SUERC), and we
are most grateful to staff at these facilities. Fig. 2c was reprinted
from the Journal of Glaciology with permission of the International
Glaciological Society.
Appendix A. Supplementary data
Supplementary data related to this article can be found at http://
dx.doi.org/10.1016/j.quascirev.2015.09.016.
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