Dating the onset of LGM ice surface lowering in the High Alps
Christian Wirsig a
, Jerzy Zasadni b, *
, Marcus Christl a
, Naki Akçar c
, Susan Ivy-Ochs a
a
Laboratory of Ion Beam Physics, ETH Zürich, Otto-Stern-Weg 5, 8093 Zürich, Switzerland
b
Faculty of Geology, Geophysics and Environmental Protection, AGH University of Science and Technology, al. A. Mickiewicza 30, 30-059 Krakow, Poland
c
Institute of Geological Sciences, University of Bern, Baltzerstrasse 1þ3, 3012 Bern, Switzerland
a r t i c l e i n f o
Article history:
Received 17 December 2015
Received in revised form
30 April 2016
Accepted 4 May 2016
Available online 18 May 2016
Keywords:
10
Be
Deglaciation
Zillertal
Mont Blanc
a b s t r a c t
The general chronological outline of the Last Glacial Maximum (LGM) in the Alps is well known after
more than two centuries of ice age research in Europe. Current studies focus on resolving the details of
the deglaciation process after the LGM. Particularly few data in this context are available from the High
Alps. Here we report chronological constraints on the onset of deglaciation in two different study areas
from the Western and Eastern Alps: the Mont Blanc (Italy) and Zillertal Alps (Austria). We sampled 32
bedrock and boulder surfaces from high elevation ridges beneath the glacial trimline. The oldest 10
Be
exposure ages of ~18.5 ka correspond to the initiation of lowering of the LGM ice surface. Identical ages
from both study areas suggest synchronous decay of the LGM glaciers in the accumulation zones of the
Western and Eastern Alps. Thus, the High Alpine ice surface lowered roughly synchronously to the
downwasting of the glacier tongues in the forelands that was completed 19e18 ka. In both study areas,
ages of numerous samples that are consistently 2e3 ka younger than 18.5 ka suggest the persistance of
remnant ice patches at high elevations in the Lateglacial accumulation zones approximately until the
Bølling-Allerød interstadial.
© 2016 Elsevier Ltd. All rights reserved.
1. Introduction
The Alps are the largest mountain range in Central Europe. They
stretch in a wide arc from southern France in the west to Slovenia in
the east over a distance of approximately 1200 km. The Alpine
divide (Fig. 1A) at its center is often more than 3000 m high. It
controls precipitation patterns and thereby sketches the outline of
climate zones in neighboring regions: maritime climate to the west,
Mediterranean climate to the south and more continental climate
to the northeast (Veit, 2002, and references therein). Today only the
highest areas above ~2500e3000 m are occupied by glaciers and
permafrost (Hoelzle,1996; Evans, 2006). In contrast, during the Last
Glacial Maximum (LGM e 26.5e19 ka; Clark et al., 2009) the Alps
were almost completely covered by ice (Fig. 1A). The accumulation
zones within the High Alps sustained an extensive network of
transections glaciers that filled the main valleys (Penck and
Brückner, 1901/1909; J€ackli, 1962; van Husen, 1987). Whereas a
few dispersive centers of ice accumulation dominated the Central
Swiss Alps (Florineth and Schlüchter, 1998; Kelly et al., 2004; Bini
et al., 2009), glaciers in the Eastern Alps and southern side of the
Western Alps were more strongly controlled by topography (Mayr
and Heuberger, 1968; van Husen, 1987). Large ice masses in the
main valleys fed outlet piedmont lobes which extended into the
northern and southern Alpine forelands (Penck and Brückner,1901/
1909).
The maximum dimension of ice cover in the High Alps is well
recorded in the glacial trimline (Ballantyne, 1997) that is visible in
the topography as the boundary between an ice-moulded downslope
area and a frost-shattered upslope zone (Florineth and
Schlüchter, 1998; Kelly et al., 2004). In large-scale morphology,
the trimline is commonly discernible as a truncation of valley-side
spurs showing a more or less sharp transition (30e100 m) between
the frost-weathered ar^ete zone above and the gentle ice-moulded
ridges below (Florineth, 1998; Hippe et al., 2014; Wirsig et al.,
2016). It is generally assumed in the Alps that the trimline represents
the maximum elevation of the active LGM ice surface (subaerial/subglacial
boundary) (Florineth and Schlüchter, 1998; Kelly
et al., 2004; Cossart et al., 2012). This is in contrast to the
increasing body of evidence in the British Isles (Fabel et al., 2002;
Ballantyne and Stone, 2015; and references therein) and Scandinavia
(e.g. H€attestrand and Stroeven, 2002) that relates the late
Pleistocene periglacial trimline to an englacial thermal boundary* Corresponding author.
E-mail address: jerzyzasadni@geol.agh.edu.pl (J. Zasadni).
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Quaternary Science Reviews 143 (2016) 37e50
between cold-based ice above and erosive warm-based ice below.
In particular, glacial erratic boulders perched on periglacial blockfields
far above the trimline show young LGM or post-LGM exposure
ages; thus the periglacial zone had to have been overridden by
cold-based ice during the last glaciation (Fabel et al., 2002;
Ballantyne and Stone, 2015). No such evidence has been presented
for the Alps.
The timing of the transition of the Alps from the LGM to its
present state has been intensely studied (Schlüchter and
R€othlisberger, 1995; Ivy-Ochs et al., 2004; Ivy-Ochs, 2015, and
Fig. 1. A) The Alps during the Last Glacial Maximum with locations mentioned in the text. The red outlines mark the locations of the two study areas: MB ¼ Mont Blanc,
ZA ¼ Zillertal Alps. The blue arrows show the course of the Dora Baltea and Inn Glacier to their piedmont lobes. Blue dashed lines show main Alpine water divide and extent of the
Aosta (Dora Baltea) and Inn Valley systems. The ice extent is from Ehlers and Gibbard (2004, http://www.qpg.geog.cam.ac.uk/lgmextent.html, accessed 20.11.2015). B, C) Simplified
geomorphological maps of the MB and ZA study areas, respectively. Samples are arranged by elevation within each sampling site. (For interpretation of the references to colour in
this figure legend, the reader is referred to the web version of this article.)
C. Wirsig et al. / Quaternary Science Reviews 143 (2016) 37e5038
references therein). Most available data are from the forelands,
where fluctuations of the ice margin are better recorded than
elevation changes of the ice surface in the High Alps. The outermost
positions in the Italian morainic amphitheaters were occupied until
24e23 ka (Monegato et al., 2007; Gianotti et al., 2008, 2015;
Ravazzi et al., 2012; Fontana et al., 2014; Scapozza et al., 2015).
Withdrawal from the maximum positions of major lobes in the
northern Swiss foreland (Rh^one, Aare, Reuss) appears to have
occurred simultaneously (Ivy-Ochs et al., 2004; Akcar et al., 2011;
Reber et al., 2014; Graf et al., 2015; Ivy-Ochs, 2015). Subsequently,
a second LGM advance phase is recognized on both sides of the Alps
ending around 21e19 ka, followed by retreat of the glacier lobes
from the forelands (Lister, 1988; Monegato et al., 2007; Starnberger
et al., 2011; Ravazzi et al., 2012; Fontana et al., 2014; Reber et al.,
2014). A re-calculation and re-interpretation of exposure ages
revealed a similar chronology in the Italian Maritime Alps (Federici
et al., 2012, 2016; Ivy-Ochs, 2015). In Austria, it has been suggested
that inner-Alpine valleys were ice free as early as 19 ka (Klasen
et al., 2007; Reitner, 2007) due to a catastrophical collapse of the
foreland tongues (van Husen, 2000; Reitner, 2013).
Comparably little chronological information is available from
the High Alps. Wirsig et al. (2016) report that the highest ice surface
in the Oberhasli (CH) appears to have been maintained until 23 ± 1
ka. Severe ice surface lowering initiated at Simplon pass (CH) and in
the Oberhasli at ~ 18 ka (Dielforder and Hetzel, 2014; Wirsig et al.,
2016). Other passes in Switzerland - Grimsel, Gotthard and Albula show
no signs of deglaciation before 16e14 ka (Kelly et al., 2006;
B€ohlert et al., 2011; Hippe et al., 2014). The authors suggest that
the time lag to the retreat of ice from the forelands was caused by
local ice that persisted on the passes, which remained in the
accumulation area during most of the Lateglacial. The first major
glacier re-advance during the Gschnitz stadial was dated to 17e16
ka (Ivy-Ochs et al., 2006a). Mass movement deposits underlying the
Gschnitz till are evidence of previously ice-free valleys within the
Alps (van Husen, 1997). Therefore, the LGM glaciers had disappeared
before the Gschnitz stadial (Kerschner et al., 1999;
Schmidt et al., 2012).
This study aims to contribute to the development of a more
detailed understanding of the deglaciation chronology in the High
Alps. In particular, we investigate whether ice surface lowering at
the centers of ice accumulation was synchronous to the retreat of
piedmont lobes from the forelands. We address this question by
interpretation of exposure ages from boulder and bedrock surfaces
from locations at slightly lower elevations than the LGM ice surface
in two study areas: one at the southern side of the Mont Blanc in
the Western Alps, the other in the Zillertal Alps in the Eastern Alps
(Fig. 1A). These areas are subject to rather different climate regimes
today: while the Mont Blanc area is dominated by prevailing
Westerlies, Zillertal is under a stronger continental influence. Did
deglaciation in the two areas, which are separated by 600 km W - E
and oriented towards different sides of the main Alpine divide, start
at the same time?
2. 10
Be surface exposure dating
Due to the predictable build-up with time, the concentration of
cosmogenic nuclides such as 10
Be in a rock surface carries information
of how long it has been exposed to cosmic rays. This is
routinely used to date the deposition age of landforms such as
moraines, fluvial terraces or landslides (Gosse and Phillips, 2001;
Bierman et al., 2002; Ivy-Ochs and Kober, 2008, and references
therein). The 10
Be concentration in previously glacier-covered
bedrock has been studied to determine the time since the ice disappeared
(e.g. Ivy-Ochs, 1996; Kelly et al., 2004; Ivy-Ochs and
Briner, 2014; Hughes et al., 2016a).
2.1. Sampling strategy
The sample campaigns in both study areas were prepared by a
thorough pre-analysis of possible sites. Based on field evidence,
interpretation of maps and DEMs and the survey of existing reconstructions
of the LGM ice surface we identified sample site
candidates. We aimed for sites 1) that were as close as possible
below the reconstructed LGM ice surface, 2) that were outside the
known extent of any Lateglacial glacier readvance, 3) where
topography reduces the risk of long-lasting snow or local ice fields,
4) that contain bedrock outcrops or boulders of a suitable (i.e.
quartz-bearing) lithology and 5) that are accessible both from the
perspectives of logistics and safety.
Applying these criteria led to the targeting of two types of
landforms: 1) narrow ridges just below the trimline (Fig. 2A) and
2) glacially smoothed high elevation flat areas (Fig. 2B). Within a
trough, narrow ridges below truncated spurs are ideal sampling
spots to address the questions posed in this study. The relative
position to the LGM ice surface is well-constrained due to the
proximity to the trimline. Furthermore, an ice surface at comparable
elevation can only be attained by a glacier system of LGMlike
size, which excludes interference caused by Lateglacial stadials.
It has to be examined carefully, however, whether glaciers
from nearby tributary valleys or cirques could have extended over
the sampling spots (Ivy-Ochs, 1996; Kelly et al., 2004; Ivy-Ochs
et al., 2007; Hippe et al., 2014). In contrast, high-elevation areas
that were clearly below the LGM ice surface (judged by
morphology and position in the valley) qualify as sampling site
candidates only if there is no accumulation zone above that could
support a glacier reaching out onto the site. Keeping this in mind,
we focused on convex areas, in particular ridges and bedrock
highs within the glacial trough. As the complete ridge or bedrock
high was buried by LGM ice in these cases, there is no local
trimline evidence. Therefore one has to rely on a wider-scale ice
surface reconstruction to estimate how thick the ice cover at the
sample site was. Furthermore, it has been suggested that local
patches of ice can persist in similar settings, high elevation
transfluence passes and trough-shoulders, throughout the Lateglacial
until the Bølling-Allerød interstadial (B€ohlert et al., 2011;
Wirsig et al., 2016). This type of ice cover is also conceivable on
ridges below the trimline, if they are not as narrow as in Fig. 2A
(Aiguille du Ch^atelet site).
Within each site candidate, suitable sampling spots were
chosen to minimize cover by annual snow. If boulders of sufficient
size (>1 m) were present they were preferred as they are thought
to stick out of the snow. Otherwise we targeted the top of the
highest bedrock outcrops, preferably close to steep drops in
elevation (Fig. 2C and D). To constrain the uncertainty introduced
by unknown erosion rates, severely weathered rock surfaces
were avoided. Some quartz veins displayed glacial polish that
indicates negligible weathering. In the absence of clearly polished
surfaces we selected bedrock that formed a coherent surface
consistent with being ice-moulded. Quartz veins were sampled
where available. We collected the uppermost 1.0e4.0 cm thick
layer of the rock surface with hammer, chisel and a batterypowered
saw.
2.2. Sample preparation, measurement and age calculation
The method of Beryllium sample preparation is comprehensively
described in the literature (e.g. Kohl and Nishiizumi, 1992;
Ivy-Ochs, 1996). We use grain sizes <1 mm of quartz separates.
The addition of a commercial 9
Be carrier is followed by quartz
dissolution in HF and Be purification using ion-exchange columns
and selective precipitation. The 10
Be/9
Be ratio is measured by
C. Wirsig et al. / Quaternary Science Reviews 143 (2016) 37e50 39
accelerator mass spectrometry (AMS) on the 600 KV TANDY system
at the Laboratory of Ion Beam Physics (LIP) at ETH Zürich
against the in house standard S2007N (Christl et al., 2013) that is
calibrated against the 07KNSTD standard (Nishiizumi et al., 2007).
We correct for a long-term average full chemistry procedural
blank of 10
Be/9
Be ¼ (3.6 ± 2.6) * 10À15
. Ages were calculated using
the production rate of Balco et al. (2009): 3.87 ± 0.19 at/g/a at sea
level and high latitude (SLHL) and 'Lm' scaling, with the CRONUSEARTH
online calculator (Balco et al., 2008). Erosion rates of
0.1 mm/ka were applied for bedrock, 1 mm/ka for boulder samples.
This takes into account the observed difference in surface
weathering. The reported errors propagate uncertainties from
AMS standard reproducibility, counting statistics, standard mean
error of samples, blank correction and the local production rate
value. These external errors are used to compare absolute ages to
independent chronologies. Internal errors ('int.') exclude the uncertainty
of the local production rate. All errors are reported on
the 1s level.
Following the argument given in Wirsig et al. (2016), we apply a
snow correction of 50 cm for 6 months per year to all bedrock
samples. To reflect our sample preference for exposed locations on
bedrock high, these values are chosen low compared to average
snow cover in the study areas in the past decade which are >100 cm
for 6 months per year based on reports of local ski areas. We
exclude samples from the Aiguille du Ch^atelet site from this
correction, because the steep wall of the ice-moulded truncated
spur does not offer any space for snow accumulation (see Fig. 2A).
Furthermore, no snow correction is applied to boulder samples. The
correction factor is calculated following eq. (3.76) in Gosse and
Phillips (2001) with a snow density of 0.3 g/cm3
and an attenuation
length for fast neutrons in snow of 109 g/cm2
(Zweck et al.,
2013; Delunel et al., 2014; Dunai et al., 2014). This reduces spallogenic
production rates by ~6.4% (cf. Schildgen et al., 2005).
3. Study areas
This work has been performed in two study areas (Fig. 1). One is
an area of 20 km * 25 km on the Italian side of the Mont Blanc
massif. The other is the Zillertal Alps in Austria, a 30 km * 35 km
area of northward trending valleys. The two areas are approximately
600 km apart on the Alpine divide, mainly in an W-E direction.
Their orientation with respect to the main mountain chain
puts them on opposite sides of the weather divide: the Mont Blanc
study area to the south, Zillertal Alps to the north. At an average of
6.3 C and 1139 mm annual precipitation, Courmayeur (1261 m)
near the Mont Blanc is approximately 1 C warmer and 30% moister
than Ginzling (1259 m, 5.2 C, 896 mm) near Mayrhofen in Zillertal
(http://en.climate-data.org/, accessed 16.12.2015). Whereas the
precipitation in Courmayeur is distributed evenly over the year,
Ginzling in Zillertal is characterized by wetter summers and drier
winters with 40e50 mm precipitation per month on average.
3.1. Mont Blanc
The Mont Blanc massif is located in the Western Alps at the
borders of France to the NW, Italy to the SE and Switzerland to the E.
The study area is at its south-eastern flank that consists mainly of
granites, gneisses and mica schists forming very steep walls
(Pfiffner, 2014, and references therein). The Veni and Ferret Valleys
(Fig.1B) mark the contact to the metasedimentary rocks connecting
to the massif towards the south that form lower and gentler terrain.
Crystalline rocks crop out at only three places on the adjacent range
of hills to the immediate south. The crystalline summit of Testa
Bernada (2534 m) is part-way up the Ferret Valley. Monte Chetif
(2343 m) and the slope at La Saxe at the end of Veni and Ferret
Valleys, respectively, are separated by the town of Courmayeur at
the head of the Aosta Valley (Fig. 1B). Here, at the junction of the
Fig. 2. Pictures from sites in the Mont Blanc study area. A) Aiguille du Ch^atelet: an ice-moulded truncated spur. View is downhill from the tributary into the Veni trough valley. Note
the adverse slope at the spur end. B) Glacially smoothed high elevation area between the sampling sites La Saxe and Testa Bernada. C) Sample Cou5 at site La Saxe. D) Samples Cou10
and Cou11 at site Testa Bernada. Note humans on each picture for scale.
C. Wirsig et al. / Quaternary Science Reviews 143 (2016) 37e5040
Veni and Ferret Valleys is where the Dora Baltea River drains towards
the south.
Approximately 30% of the massif's 550 km2
area are glacierized
today due to its high elevation that includes Europe's highest
summit, the Mont Blanc at 4810 m (Deline et al., 2012). The area has
received high interest since the beginning of the study of Quaternary
glaciations (de Saussure, 1779e1796). Penck and Brückner
(1901/1909) devised a general framework of Alpine glaciations
that included the attribution of the Ivrea Amphitheater to the
maximum extent of the Dora Baltea Glacier. Gianotti et al. (2008,
2015) conducted intensive research at the Ivrea Amphitheater
which was the terminus of the LGM as well as several previous
Quaternary glaciations. The LGM ice surface around the Mont Blanc
massif itself has been described by multiple authors through
trimline mapping with diverging results. Porter and Orombelli
(1982) and Coutterand and Buoncristiani (2006) reconstruct the
ice surface at Courmayeur at 2600 m, 200 m higher than Kelly et al.
(2004) and Bini et al. (2009). The offset remains constant up to the
heads of the steep tributary valleys at the Mont Blanc SE face.
Lateglacial deposits are rarely encountered on the Aosta Valley
floor, as this is thickly covered by alluvium and rockfall deposits
(Porter and Orombelli,1982). Gianotti et al. (2008, 2015) discuss the
evidence for seven locations of Lateglacial stadial advances along
the Aosta Valley, upvalley of the Ivrea Amphitheater. A constraint
on the ice surface within our study area is only available for two of
the potential terminal positions, however. Based on the upper limit
of erratic boulders along the Ferret and Veni Valleys, Porter and
Orombelli (1982) reconstruct the ice surface of an advance with a
terminal position close to Courmayeur. The limit of the smaller
Planpincieux advance is determined by mapping of moraines and
outwash deposits (Porter and Orombelli, 1982). Surface exposure
ages from a roche moutonnee 80 km downvalley from Courmayeur
are 17e16 ka (Gianotti et al., 2008, 2015), synchronous to ages
determined for the Gschnitz stadial (Ivy-Ochs et al., 2006a). The ice
surface in our study area for this readvance is unknown.
We selected three sampling sites in the Mont Blanc study area
(Fig. 1B). The Aiguille du Ch^atelet (2525 m) is located at the
southern wall of the Mont Blanc, 750 m above the current terminus
of Miage Glacier in the Veni Valley (Fig. 2A). It is a narrow ridge on a
truncated spur in between the Fr^eney and Brouillard Glacier. Its
smooth western slope has clearly been moulded by the LGM Veni
Glacier moving from W to E. Trimlines on adjacent spurs at
~2650 m indicate that Aiguille du Ch^atelet was covered by
100e150 m of ice during the LGM. It has an adverse slope at its end
and a local relief to the adjacent glacier beds of >400 m. Glaciers of
this thickness in the steep tributary valleys can be excluded during
any of the Lateglacial readvances based on estimated terminal
positions (Gianotti et al., 2008, 2015). A line of boulders uphill on
the ridge potentially marks the extent of a Lateglacial glacier
advance well outside the reach of the end of the spur where we
sampled. Aiguille du Ch^atelet is therefore expected to have become
and continue being exposed once the LGM ice surface in Veni Valley
lowered by more than ~150 m. Here we took four bedrock samples
(Cou1-4) from the crest of the sharp ridge a few meters from the
top towards the direction of the tributary valleys.
The site at La Saxe is on top of a steep bedrock cliff towering to
the east of Courmayeur at the beginning of the Ferret Valley
(Fig. 1B). The top is a comparably flat, regular surface with
numerous roches moutonnees. It was covered by more than 350 m
of ice during the LGM (Porter and Orombelli, 1982; Kelly et al.,
2004; Coutterand and Buoncristiani, 2006). Here we collected
bedrock samples from two different elevations, separated by a
farming area with marked human impact on the landscape. Samples
Cou6,7 from the lower elevation range at ~1900 m are 700 m
above the present Aosta valley floor and 600 m above the
reconstructed ice surface of the Courmayeur advance (Porter and
Orombelli, 1982). The locations of Cou8,9 are at ~2000 m. We targeted
the top of the highest well-preserved roches moutonnees
that stood 1e2 m above the surrounding topography (Fig. 2C).
Following the same ice-moulded high elevation area to the east,
the third sample site is a granite outcrop at the foot of Testa Bernada.
Depending on which LGM ice surface reconstruction is consulted,
this spot was covered by 50 m (Kelly et al., 2004; Bini et al.,
2009) to 250 m of ice (Porter and Orombelli, 1982; Coutterand and
Buoncristiani, 2006). The general morphology is smooth, but
irregular with indentations where snow fields persist until late
summer (Fig. 2B). Despite a considerable degree of weathering, the
bedrock preserved general stoss and lee features. We collected two
samples (ZA10,11) where the bedrock stood 1e2 m above the surrounding
topography (Fig. 2D). We did not sample any quartz veins
or polished rock surfaces at Mont Blanc.
3.2. Zillertal Alps
The Zillertal Alps massif is located in the Eastern Alps between
Brenner Pass to the W and the Hohe Tauern mountain range to the E
(Fig. 1A). The main crest of the massif is formed by a range of peaks
with summits above 3000 m that is part of the main Alpine water
divide and constitutes the Austrian - Italian border. The highest
peaks are Hochfeiler (3509 m) and Olperer (3476 m) that form a
relief of ~2000 m to the adjacent valley floors. The northern slope of
the Zillertal Alps is dissected by a system of four principal tributary
valleys: Tuxertal, Zemmtal, Stillupgrund and Zillergrund that
merge in one point near Mayrhofen (Fig. 1C). From there the Ziller
River drains towards the north into the Inn River.
The highest part of the Zillertal Alps is built of weatheringresistant
rocks of the crystalline core of the Tauern Window, Zentralgneis
(Frisch, 1980; Vesela et al., 2008). In Tuxertal, a variety of
metasedimentary and metavolcanic rocks stacked in the Venediger
nappes outcrop in the south, whereas the north is formed by
Bündnerschiefer of the Glockner nappes (Froitzheim et al., 2008,
and references therein). The marked difference in relief and preservation
of glacial erosional landforms (troughs, cirques, ar^etes) is
due to the resistant gneisses found in the southern part of the study
area compared to weaker lithologies encountered towards the
north (Penck and Brückner, 1901/1909; Bobek, 1933).
Geomorphological mapping of the area at a scale 1: 10,000 was
performed over the course of ten years by J. Zasadni as an assignment
from the Geological Survey of Austria (Geologische Bundesanstalt
Wien, GBA). Particular emphasis was put on mapping of
glacial landforms and trimlines that enabled a reconstruction of the
LGM ice surface and identification of Lateglacial stadial extents
(Zasadni, 2010, 2011). A simplified version of the geomorphological
map is shown in Fig. 1C. The focus of this study is on the southern
part of the Zillertal Alps, quoted many times as a text-book example
of High Alpine glacial relief (e.g. Penck and Brückner, 1901/1909;
Heuberger, 1968; Louis and Fischer, 1979). Whereas Tuxertal and
Stillupgrund are simple glacial troughs, Zemmtal and Zillergrund
constitute a dendritic network of well-developed troughs with
tributary hanging valleys (Penck, 1905).
Together with the Hohe Tauern mountain range, the Zillertal
Alps formed a dominant center of accumulation of the Eastern
Alpine transection glaciers during the Last Glacial Maximum (van
Husen, 1987). The LGM Ziller Glacier joined the Inn Glacier system
that extended into the northern Alpine foreland. The Inn Valley
system is a type region of the ‘Würm’ glaciation in the Alps as
defined by the Subcommission of European Quaternary Stratigraphy
(Chaline and Jerz, 1984). Original localities of the Lateglacial
stadials Gschnitz and Egesen are located in this area (Kerschner,
1986; Ivy-Ochs et al., 2006b, and reference therein). In the
C. Wirsig et al. / Quaternary Science Reviews 143 (2016) 37e50 41
Hopfgarten area, part of the Inn Valley system 50 km downvalley
from the Zillertal Alps, morphological and sedimentary evidence
indicates a phase of early Lateglacial ice decay of the Inn Glacier
that followed the full LGM condition and preceded the Gschnitz
stadial (Reitner, 2007).
The LGM ice surface in the study area has been reconstructed by
several authors (Penck and Brückner, 1901/1909; Mayr and
Heuberger, 1968; van Husen, 1987). Moraines near Mayrhofen
and in the middle section of Tuxertal were tentatively correlated to
the Lateglacial Gschnitz readvance (Penck and Brückner, 1901/
1909; Bobek, 1932). Recent mapping results (e.g. Zasadni, 2014)
indicate that large dendritic Gschnitz glaciers almost entirely
occupied four principal tributary valleys around Mayrhofen
(Fig. 3A). Moraines that were attributed to the Egesen stadial are
mostly on the mountain slopes and cirques; the glaciers with the
largest accumulation areas reached the valley floors, e.g. from
Pfitscherjoch to Zamsergrund (Bobek, 1932; Heuberger, 1977).
Climate reconstructions from the Spannagel cave in Tuxertal based
on speleothem records yield detailed information on local paleotemperatures,
in particular during the Holocene (Vollweiler et al.,
2006; Fohlmeister et al., 2013), but also for the more distant past
(Sp€otl and Mangini, 2006, 2007).
We selected six sampling sites in the Zillertal study area. Differences
in elevation to the LGM ice surface are based on an
interpolation of mapped trimline data (Fig. 1C) which indicate an
ice surface elevation above Mayrhofen and in the lower section of
tributary valleys of ca. 2200 m and 2300 m, respectively. Three
sample sites are on different ridges close to the confluence area at
Mayrhofen (Fig. 3A): Gams (ZA26-28), Dristner (ZA16-20) and
Ahorn (ZA1,2). The LGM ice surface is tightly constrained by truncated
spurs on each ridge that mark the local trimline. We aimed to
collect samples at the highest elevation with suitable bedrock
outcrops, since no boulders were available. The Gams and Ahorn
sites are relatively broad, covered with low vegetation that is
interrupted by several meter high outcrops of weathered Zentralgneis.
Glacial polish is preserved on quartz veins that we preferably
sampled. A lateral moraine in the forest ~1000 m below the
Ahorn site is correlated to the Gschnitz readvance (Zasadni, 2014).
The Dristner site is on the sharpest of the three ridges. The two
topmost samples (ZA17,16) are from a sparsely vegetated area less
than 100 m below the trimline where glacially moulded Zentralgneis
forms a steep and narrow ridge. Three additional samples
were taken in steps of ~100 m altitude (ZA18-20) to look at the rate
of ice surface lowering. The ridge broadens with decreasing altitude
and vegetation becomes denser.
Two sites, Am Flach and Tettensjoch, are on opposite slopes on
the same ridge in Tuxertal. The sites were covered by ca. 100 m of
ice during the LGM as inferred from nearby trimline elevation
points (Fig. 1C). The upstream SW slope of the Am Flach site is
gentle and flat. It hosts 14 erratic Zentralgneis boulders with an aaxis
size between 0.7 m and 14 m. We sampled four of the largest
well-preserved boulders (ZA3A,4,5,8; Fig. 3B) and two bedrock
outcrops (polished quartz veins, ZA6,7). At Tettensjoch on the NE
side of the ridge pointing downstream towards Mayrhofen, steps of
fractured bedrock form a steeper and narrower slope than at Am
Flach, sometimes preserving a roche moutonnee shape. Here we
targeted the sharp edge of the ridge and bedrock knobs next to
steep drops for collecting four bedrock samples (ZA29-32).
Sample site Ameiskopf is located 18 km upvalley in Zemmtal.
Similar to the Aiguille du Ch^atelet site in the Mont Blanc study area,
it is an ice-moulded truncated spur at the end of a ridge that separates
two tributary cirques. The elevation difference to the adjacent
cirque floors is 200e300 m. Here the orientation of the
sampling site is different than at Aiguille du Ch^atelet, however. The
Ameiskopf site is located directly beneath the spur towards the
main trough, not behind the spur towards the cirques. The troughside
of the spur where we sampled is narrow, but flat and smooth.
The bedrock there shows abundant evidence of onion-type exfoliation
weathering. We therefore targeted spots with quartz veins
that preserved glacial polish. The two samples (ZA24,25) were
taken only few centimeters from the steep and deep drop that
limits the side of the flat area.
4. Results and interpretation
Sample information and measured 10
Be concentrations of the 32
Fig. 3. Sample positions in the Zillertal Alps. A) Ahorn, Dristner and Gams sites with respect to the reconstructed LGM and Gschnitz stadial ice surfaces. Red dots are sample positions,
red lines moraines, white lines reconstructed ice surfaces. The photograph is taken from Penken mountain (2095 m), view to the SSW. Ice flow was out of the valleys towards the
confluence at Mayrhofen. The Am Flach and Tettensjoch sites are just out of view in Tuxertal. Ameiskopf is located upvalley in Zemmtal. B) Sampled erratic boulders at the Am Flach
site. Boulders were transported 3 km from H€ollensteinkar where their source lithology occurs (red layer); Ahorn gneiss (AG; boulders: ZA3A, ZA4 and ZA8) and H€ollenstein gneiss (HG:
boulder ZA5). Note person on ZA4 for scale. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)
C. Wirsig et al. / Quaternary Science Reviews 143 (2016) 37e5042
samples are summarized in Table 1. Tables 2 and 3 list calculated
exposure ages with and without snow correction and associated
uncertainties for the Mont Blanc and Zillertal Alps study area,
respectively. The snow corrected exposure ages will be used for the
interpretation and discussion. First we present the age patterns
within each study area, then point out implications in the discussion
section.
4.1. Mont Blanc
The obtained ages from ten samples in the Mont Blanc study
area range from 14.2 ± 1.0 ka to 18.5 ± 1.1 ka (Table 2, Fig. 4A). The
two oldest ages are from Cou3 and Cou1 at the Aiguille du Ch^atelet
site on the northern side of Veni Valley. Due to the sharp ridge,
adverse slope and high elevation difference to the adjacent glacier
beds, we consider there is no possibility of temporary ice cover.
Furthermore, there is no place for annual snow to accumulate on
the narrow ridge crest (Fig. 2A). The age of Cou3 (18.5 ± 1.1 ka) is
therefore interpreted as the onset of ice surface lowering in Veni
Valley. The younger age of Cou4 must relate to spalling of the
original surface that was undetected due to the generally weathered
character of the ridge crest.
Our interpretation of trimline evidence at the Aiguille du
Ch^atelet is within the range of previous reconstructions. Fig. 4A
shows the relative position of the sampling site to the LGM ice
surface constrained by the trimlines on the adjacent truncated
spurs. On the spur to the east, we place the trimline at 2620 m, 50 m
higher than some authors (Kelly et al., 2004; Bini et al., 2009), but
150 m lower than others (Porter and Orombelli, 1982; Coutterand
and Buoncristiani, 2006). On the southern slope of Ferret Valley,
our trimline point at 2500 m at Testa Bernada displays the same
pattern of differences to previous reconstructions. Since we do not
generate a comprehensive reconstruction of the LGM ice surface in
the Mont Blanc study area, we calculate the elevation difference of
individual sample spots to the ice surface in the map of Bini et al.
(2009), because it appears more similar to our interpretation. Due
to the constant offset between different reconstructions, the relative
distribution of elevation differences to the LGM ice surface is
conserved if another reconstruction is selected as reference.
The two samples at Testa Bernada (Cou10,11) yield congruent
ages of ~16.4 ka. Likewise, the four ages from La Saxe (Cou6-9) are
very consistent at 14.7e15.8 ka. These ages from the southern side
of the Ferret Valley show an evident mismatch to the onset of LGM
ice surface lowering at the Aiguille du Ch^atelet site of 2e3 ka. A
delayed lowering of the LGM ice surface by 2e3 ka in the Ferret
Valley, compared to the Veni Valley, is not possible, however; due
to their immediate confluence at Courmayeur, the ice surfaces of
the LGM trunk glaciers in the two valleys depended on the position
of the same glacier tongue in the Aosta Valley and must behave
synchronously. Samples from the Testa Bernada site should therefore
reflect exposure simultaneous to or even before ~18.5 ka
(Cou3), because they were closer to the ice surface than the Aiguille
du Ch^atelet. The age mismatch indicates temporary cover by ice or
sediment after the LGM ice surface dropped. Indeed, the irregular
topography on the La Saxe ridge with rounded indentations (Fig.
2B) offers plenty of opportunities for local patches of ice to form
and persist, given sufficiently cold conditions as for example during
the Lateglacial. Despite our careful selection of sample sites we
Table 1
Sample list and measured 10
Be concentrations.
Sample ID Latitude Longitude Elevation Thickness Topographic
shielding factor
10
Be conc.
WGS 84 m a.s.l. cm 105
at/gqtz
a
Mt. Blanc
Cou1 45.7958 6.8843 2500 4.0 0.967 4.75 ± 0.14
Cou2 45.7958 6.8843 2497 1.0 0.959 4.34 ± 0.17
Cou3 45.7961 6.8841 2488 3.0 0.958 4.89 ± 0.15
Cou4 45.7961 6.8844 2477 2.0 0.888 3.46 ± 0.16
Cou6 45.8079 6.9739 1913 2.0 0.992 2.74 ± 0.17
Cou7 45.8080 6.9740 1912 3.0 0.987 2.56 ± 0.10
Cou8 45.8099 6.9800 2016 2.0 0.986 2.71 ± 0.11
Cou9 45.8100 6.9792 2027 4.0 0.990 2.84 ± 0.12
Cou10 45.8243 7.0091 2402 3.5 0.992 3.94 ± 0.16
Cou11 45.8241 7.0091 2401 2.5 0.996 3.96 ± 0.31
Zillertal Alps
ZA1 47.1334 11.8781 2076 2.5 0.968 3.54 ± 0.20
ZA2 47.1346 11.8674 1917 1.5 0.992 2.84 ± 0.15
ZA3A 47.1374 11.8698 2176 3.5 0.995 3.39 ± 0.33
ZA4 47.1371 11.7361 2168 1.5 0.992 3.31 ± 0.18
ZA5 47.1361 11.7356 2143 2.0 0.992 3.37 ± 0.16
ZA6 47.1377 11.7363 2172 1.5 0.971 3.63 ± 0.12
ZA7 47.1376 11.7360 2165 2.0 0.975 3.15 ± 0.13
ZA8 47.1366 11.7373 2192 4.0 0.994 3.40 ± 0.16
ZA16 47.1276 11.8342 2122 2.5 0.945 3.29 ± 0.16
ZA17 47.1276 11.8349 2130 1.5 0.914 3.46 ± 0.18
ZA18 47.1298 11.8350 2039 2.0 0.995 3.09 ± 0.18
ZA19 47.1306 11.8338 1988 2.5 0.978 2.98 ± 0.13
ZA20 47.1325 11.8325 1834 2.0 0.968 2.47 ± 0.12
ZA24 47.0231 11.6671 2465 2.0 0.972 3.78 ± 0.12
ZA25 47.0231 11.6671 2463 1.5 0.978 3.92 ± 0.17
ZA26 47.1281 11.7901 2223 1.5 0.988 3.73 ± 0.13
ZA27 47.1282 11.7899 2221 1.0 0.985 3.25 ± 0.24
ZA28 47.1285 11.7909 2172 1.5 0.954 3.05 ± 0.13
ZA29 47.1474 11.7536 2126 1.0 0.987 3.35 ± 0.19
ZA30 47.1474 11.7538 2115 1.0 0.944 3.46 ± 0.13
ZA31 47.1478 11.7539 2108 1.5 0.995 3.24 ± 0.17
ZA32 47.1479 11.7542 2089 1.5 0.988 3.68 ± 0.58
a
Measured against standard 07KNSTD (Nishiizumi et al., 2007), corrected for full process blank of (3.6 ± 2.6) * 10À15 10
Be/9
Be.
C. Wirsig et al. / Quaternary Science Reviews 143 (2016) 37e50 43
cannot exclude that these ice patches grew thick and large enough
to cover even the highest bedrock outcrops.
The age difference between the Testa Bernada and La Saxe sites
possibly reflects the rate of ice surface lowering at the head of the
Aosta Valley. This would indicate lowering by 300e450 m in ~1 ka.
On the other hand, the lower elevation makes the La Saxe site more
likely to be covered during Lateglacial readvances. While the
Courmayeur readvance was well below the sampling sites, the ice
surfaces of other readvances with terminal positions further down
the Aosta Valley are not known. At Testa Bernada, considering the
elevation difference of 800 m from the sample spots to the valley
floor, it appears highly unlikely that any Lateglacial readvance in
Ferret Valley affected the samples. Cover and erosion during Lateglacial
advances at La Saxe, however, would decrease the measured
nuclide concentrations and result in apparently younger ages.
Chronological data to reference our results are available mostly
from the intensive studies of the Ivrea Amphitheater that defined
the terminus of the Dora Baltea Glacier during the LGM. 10
Be surface
exposure ages of 23.8 ± 1.7 ka and 20.1 ± 3.0 ka are interpreted
to date the moraine stabilization at the initiation of glacier retreat
(Gianotti et al., 2008; recalculated in Gianotti et al., 2015). The
amphitheater was completely abandoned by the Dora Baltea
Glacier before 17.7e16.8 ka (14.2 ± 1.5 14
C ka (Schneider, 1978)).
Before 17e16 ka, the glacier tongue had retreated more than 30 km
before depositing a frontal moraine in Torredaniele (Gianotti et al.,
2015). Our ages correspond well to this framework: ice surface
lowering in the accumulation area starting at 18.5 ka followed after
the initiation of, but presumably synchronous to, the retreat of
Table 2
Sample type and ages from the Mt. Blanc study area, arranged by difference in elevation to the reconstructed LGM ice surface and grouped to sampling sites.
Sample ID D elevation to LGMa
Elev. Type Site
(elevation of ice
surface)
10
Be exposure ageb
no correction incl. snow corr.c
m m a.s.l. ka ± (int.) ka
Cou10 À48 2402 Bedrock Testa Bernada 15.4 ± 1.0 0.6 16.4 ± 1.0
Cou11 À49 2401 Bedrock (2450 m) 15.3 ± 1.4 1.2 16.3 ± 1.5
Cou1 À80 2500 Bedrock 17.8 ± 1.0 0.5
Cou2 À83 2497 Bedrock Aiguille du Ch^atelet 16.1 ± 1.0 0.7
Cou3 À92 2488 Bedrock (2580 m) 18.5 ± 1.1 0.6
Cou4 À103 2477 Bedrock 14.2 ± 1.0 0.7
Cou9 À353 2027 Bedrock 14.5 ± 0.9 0.6 15.4 ± 1.0
Cou8 À364 2016 Bedrock La Saxe 13.8 ± 0.9 0.6 14.7 ± 0.9
Cou7 À468 1912 Bedrock (2380 m) 14.1 ± 0.9 0.5 15.0 ± 0.9
Cou6 À467 1913 Bedrock 14.8 ± 1.2 0.9 15.8 ± 1.3
a
Elevation difference to the LGM ice surface as reconstructed by Bini et al. (2009).
b
Using NENA production rates of 3.87 ± 0.19 at/g/a (Balco et al., 2009), 'Lm' scaling, erosion rates of 0.1 mm/ka for bedrock and 1 mm/ka for boulder samples. Rock density is
2.65 g/cm3
.
c
Assuming 50 cm snow for 6 months per year.
Table 3
Sample type and ages from the Zillertal Alps study area, arranged by difference in elevation to the reconstructed LGM ice surface and grouped to sampling sites.
Sample ID D elevation to LGMa
Elev. Typeb
Site
(elevation of ice
surface)
10
Be exposure agec
no correction incl. snow corr.d
m m a.s.l. ka ± (int.) ka
ZA26 À27 2223 Bedrock, qtz, p Gams
(2250 m)
15.9 ± 1.0 0.6 17.0 ± 1.0
ZA27 À29 2221 Bedrock 13.9 ± 1.2 1.0 14.8 ± 1.3
ZA28 À78 2172 Bedrock 14.0 ± 0.9 0.6 14.9 ± 1.0
ZA17 À120 2130 Bedrock, p Dristner
(2250e2200 m)
17.0 ± 1.2 0.9 18.1 ± 1.3
ZA16 À128 2122 bedrock, qtz 15.8 ± 1.1 0.8 16.9 ± 1.2
ZA18 À201 2039 Bedrock, qtz 15.0 ± 1.1 0.9 16.0 ± 1.2
ZA19 À232 1988 Bedrock, qtz 15.3 ± 1.0 0.7 16.3 ± 1.1
ZA20 À366 1834 Bedrock 14.3 ± 1.0 0.7 15.2 ± 1.0
ZA24 À96 2465 Bedrock, qtz, p Ameiskopf
(2560 m)
14.0 ± 0.8 0.4 14.9 ± 0.9
ZA25 À98 2463 Bedrock, qtz, p 14.4 ± 0.9 0.6 15.3 ± 1.0
ZA8 À88 2192 Boulder, 0.9 m Am Flach
(2280 m)
15.2 ± 1.1 0.7
ZA3A À104 2176 Boulder, 2.4 m 15.3 ± 1.7 1.5
ZA6 À108 2172 Bedrock, qtz, p 16.5 ± 1.0 0.6 17.6 ± 1.0
ZA4 À112 2168 Boulder, 1.8 m 14.8 ± 1.1 0.8
ZA7 À115 2165 Bedrock, qtz, p 14.4 ± 0.9 0.6 15.4 ± 1.0
ZA5 À137 2143 Boulder, 4.5 m 15.4 ± 1.1 0.8
ZA1 À114 2076 Bedrock Ahorn
(2190 m)
17.4 ± 1.3 1.0 18.6 ± 1.4
ZA2 À273 1917 Bedrock, qtz 15.2 ± 1.1 0.8 16.2 ± 1.2
ZA29 À134 2126 Bedrock Tettensjoch
(2260 m)
15.3 ± 1.2 0.9 16.3 ± 1.2
ZA30 À145 2115 Bedrock, qtz, p 16.6 ± 1.0 0.6 17.7 ± 1.1
ZA31 À152 2108 Bedrock 14.9 ± 1.1 0.8 15.9 ± 1.1
ZA32 À171 2089 Bedrock, qtz 17.2 ± 2.9 2.8 18.4 ± 3.1
a
Elevation difference to the LGM ice surface as reconstructed based on adjacent trimline elevation points.
b
‘qtz’ indicates that a quartz vein was sampled, ‘p’ indicates a polished surface, ‘xx m’ gives the boulder height.
c
Using NENA production rates of 3.87 ± 0.19 at/g/a (Balco et al., 2009), 'Lm' scaling, erosion rates of 0.1 mm/ka for bedrock and 1 mm/ka for boulder samples. Rock density is
2.65 g/cm.3
.
d
Assuming 50 cm snow for 6 months per year.
C. Wirsig et al. / Quaternary Science Reviews 143 (2016) 37e5044
significant distances of the Dora Baltea Glacier tongue.
4.2. Zillertal Alps
The obtained ages from 22 samples in the Zillertal Alps range
from 14.8 ± 1.1 ka to 18.6 ± 1.4 ka (Table 3). We refrain from the
interpretation of ZA32 due to the large uncertainty of 3.1 ka. The
oldest ages are from three of the five sites near Mayrhofen (see
Fig. 4B): 18.6 ± 1.4 ka at Ahorn (ZA1),18.1 ± 1.3 ka at Dristner (ZA17)
and 17.7 ± 1.1 ka at Tettensjoch (ZA30). We interpret these as
minimum ages for the onset of ice surface lowering at the Zillertal
confluence basin, which is supported by the good agreement of
maximum ages between the individual study sites of ~18 ka. It
appears that this is a somewhat delayed reaction to the phase of
early Lateglacial ice decay in the Inn Valley of 21.2e19.5 ka (Reitner,
2007) and approximately synchronous first signs of recession of
other LGM glaciers in the Eastern Alps (van Husen, 1997; Reuther
et al., 2011; Starnberger et al., 2011). The interpretation of our
ages is difficult to align with the hypothesis that the lower Inn
Valley was ice-free prior to 18 ka (Reitner, 2007, and references
therein) as this implies an almost complete melting of the Zillertal
tongue compared to the thickness of >1500 m attained during the
LGM.
Ages appear to decrease with elevation and therefore distance to
the LGM ice surface at some sites, e.g. Dristner and Ahorn. It is thus
tempting to interpret the younger ages at lower elevations as
tracing the rate of ice surface lowering. Arguing on the basis of
individual sites is not correct, however. With the exception of
Ameiskopf (ZA24, 25), all sites are strongly connected as they
depend on the elevation of the same ice surface in the confluence
area near Mayrhofen (Figs. 3A and 4B); they were affected by
simultaneous ice surface lowering. Therefore, not individual study
sites, but all samples (without ZA24, 25) must be compared to judge
if ages decrease with greater elevation distance to the LGM ice
surface, as would be expected if they were determined by the rate
of ice surface lowering. This is clearly not the case: while the Gams
Fig. 4. Surface exposure ages in ka in context with the reconstructed LGM ice surface in the A) Mont Blanc and B) Zillertal Alps study regions. The view is an orthogonal projection of
topography. Vertical scale is 1.5x exaggerated. Extent of projections is indicated in the small maps in the bottom right corners that show the outline of the study areas as in Fig. 1. To
include all sample sites, two planes of projection are portrayed in panel A. In panel B, the Ameiskopf sample site in the upper Zemmtal Valley is not visible directly. Blue tint and
arrows indicate reconstructed LGM ice surface and flow lines. Violet shows ice-moulded ridges, spurs and high elevated flat areas (compare Fig. 1). Exposure ages are arranged by
elevation within each sampling site. Red font marks the oldest age at each site. Asterisks show erratic boulder samples. (For interpretation of the references to colour in this figure
legend, the reader is referred to the web version of this article.)
C. Wirsig et al. / Quaternary Science Reviews 143 (2016) 37e50 45
site was covered by merely 30e70 m of ice and yielded a mean age
of 15.6 ± 1.0 ka, the Ahorn site was underneath 110e270 m of ice
but yielded a mean age of 17.0 ± 0.9 ka. We instead observe that the
geomorphological position of the sample site, in particular with
respect to the possibility to locally sustain remnant ice patches, is
the stronger control on exposure ages. The correlation of age and
elevation at individual sites presumably reflects that the ridges
become narrower at higher elevations and can therefore sustain
only thinner or no ice. This is particularly true at Dristner and
Ahorn. In addition, during the Gschnitz stadial at 17e16 ka (IvyOchs
et al., 2006a) glaciers readvanced to a presumed terminal
position at Mayrhofen (Penck and Brückner, 1901/1909). Lateral
moraines on the hillslopes testify that ice margins were ~1000 m
below our sample sites (Fig. 3A). Consequently, most of the 1500 m
thick LGM ice must have disappeared before the Gschnitz stadial.
The influence of the geomorphological position of the study site
is exemplified at Am Flach. As indicated by the name of the hillslope
that translates to ‘where it is flat’, local topography is exceptionally
smooth considering the altitude of 2140e2190 m. The ages of five
boulder and bedrock samples are consistently around 15.2 ka, only
ZA6 gives an age of 17.4 ± 1.0 ka. The Tettensjoch site is in the
immediate vicinity, a ridge on the opposite side of the same hill at
similar elevation (2090e2130 m), but steeper and narrower. Here
the ages of four samples range from 18 to 16 ka. If by ~18 ka the LGM
ice surface dropped sufficiently to expose the bedrock at Tettensjoch,
it must have deposited the boulders at Am Flach simultaneously.
In contrast to Tettensjoch, however, the flat high-elevation
area at Am Flach has been conducive to the formation of local ice
patches. The younger ages are therefore interpreted to reflect the
melting of these remnant patches of ice that persisted on the flat
landform and the associated complex exposure of the sampling
sites.
Both ages at Ameiskopf are around 15 ± 1 ka. As the site is
located further upvalley than the other locations, the position well
inside the Gschnitz accumulation area potentially entails a greater
susceptibility to ice cover during that period. It is conceivable that
the slow rate of LGM ice decay kept the ice surface sufficiently high
to cover the site until the Gschnitz stadial. In order to accomplish
this, an ice thickness of 600 m in Zamsergrund would have to be
sustained. Alternatively, a relatively thin ice cover on the Zamsergrund
trough shoulder covering the samples would yield identical
results (Wirsig et al., 2016).
5. Discussion
We first outline the implications from the combined dataset of
the Mont Blanc and Zillertal Alps study sites concerning a) the
potential influence of (glacial) erosion on the calculated exposure
ages, b) synchronicity of ice surface lowering across the High Alps
and c) temporary cover of the sample sites after the retreat of the
LGM ice. Finally these results will be put into the context of an
Alpine-wide and global deglaciation chronology.
5.1. Implications from both study areas
Rock surface erosion can potentially influence the measured
nuclide concentrations in two ways: 1) insufficient erosion during
the LGM would cause inheritance of pre-LGM nuclides and 2) postLGM
erosion would remove the rock surface with the highest
nuclide concentration yielding too young ages (Ivy-Ochs et al.,
2007). The lack of ages older than 18.5 ka in our data sets suggests
that LGM glacial occupation was able to completely erode the
pre-LGM 10
Be inheritance (Kelly et al., 2006; Wirsig et al., 2016).
This shows that erosive, warm-based ice operated very close to the
trimline (~30 m at the Gams site, Zillertal). To test our data sets for
post-LGM erosion, we compare different sample types. Polished
sample surfaces indicate negligible erosion since the disappearance
of ice. Likewise, quartz veins are more resistant to weathering than
the surrounding rock. If post-LGM erosion had a strong influence on
our results, samples of these types would be expected to yield
significantly older ages than the remaining samples. Indeed, three
of the four oldest ages in the Zillertal Alps are from polished rock
surfaces (ZA6,17,30). On the other hand, though, the surface of ZA1
with the absolutely oldest age was not polished. Furthermore,
young ages of ~15 ka from other polished surfaces (ZA7,24,25)
indicate that post-LGM erosion does not dominantly determine the
exposure ages. We therefore conclude that differences in the ages
of individual samples are caused by their different exposure histories,
not by differences in pre-/or post-LGM erosion or loss of the
sample surface.
First signs of ice surface lowering are registered at the Mont
Blanc at 18.5 ± 1.1 ka, in Zillertal at 18.6 ± 1.4 ka. Considering the
difference in location of the two study areas this level of agreement
is quite remarkable. Today the Mont Blanc area in the Western Alps
and the Zillertal in the Eastern Alps are subject to rather different
climate regimes: while the Mont Blanc area is dominated by prevailing
Westerlies, Zillertal is under a stronger continental influence.
Today even on the regional level, the impact of higher
temperature and lower precipitation to the south of Mont Blanc
causes glaciers to be significantly smaller than to the north (Deline
et al., 2012). From this perspective, synchronous lowering of the
LGM ice surface in the two study regions could not have been
assumed. During the LGM, however, a southward shift of the Polar
Front caused dominating southerly atmospheric circulation for
central Europe (Florineth and Schlüchter, 2000; Luetscher et al.,
2015). Since the differences in precipitation and climate would
have been smaller across the Alps during the LGM than they are
today, synchronous behavior of the glacier systems in east and west
was more likely. Nevertheless, the uncertainties of the dating
approach leave room for undetected asynchrony.
The oldest ages in this study of ~18.5 ka were only determined at
few sample locations, however. In contrast, most of the remaining
boulders and bedrock surfaces, 23 of 32 samples, were exposed for
16.4e14.8 ka. Since the unexpectedly young ages were not caused
by erosion, we propose that local patches of ice persisted for part of
the Lateglacial at high altitudes in both study areas. In the proximity
of the Zillertal Alps the Gschnitz ELA has been estimated to be
at 1860e1950 m (Gross et al., 1977; Kerschner et al., 1999; Zasadni,
2014). According to this, most (20 of 22) Zillertal Alps samples are
from the Gschnitz accumulation zone where the build-up of local
ice patches is plausible, if allowed by topography. The same is valid
for the samples in the Mont Blanc study area due to their location at
high elevations. Temperatures in the Alps were considerably colder
than today until the Bølling-Allerød interstadial, e.g. summers were
8e10 colder during Gschnitz (Kerschner, 2009; Schmidt et al.,
2009), which enabled the accumulation of local patches of ice
that persisted at high altitudes for at least part of the Lateglacial if
topography allowed. This has been proposed in particular at
concave settings such as High Alpine passes or trough shoulders
(Ivy-Ochs et al., 2006b; B€ohlert et al., 2011; Hippe et al., 2014;
Wirsig et al., 2016). Considering the culmination of evidence, it
can be stated tentatively that the phenomenon of ice patches persisting
after the LGM at high elevations might have occurred across
the Alps. Consistent minimum exposure ages of ~15 ka at all sites
indicate that the ice patches disappeared before or at the onset of
the Bølling-Allerød interstadial.
5.2. The deglaciation chronology of the High Alps in context
The best chronological constraint on the LGM fluctuations in the
C. Wirsig et al. / Quaternary Science Reviews 143 (2016) 37e5046
foreland lobes in the Alps is from the Tagliamento amphitheater
(Fig. 5; Monegato et al., 2007; Fontana et al., 2014). The first
Tagliamento glacial pulse e the Santa Margherita phase - occurred
between 26.5 and 23 ka. The glacier readvanced within the
amphitheater between 23 and 21 ka during the Canodusso phase
and it remained there through the Remanzacco recessional phase
that ended at 19 ka (Monegato et al., 2007). A culmination of evidence
that emerged during the last decade (Fig. 5) demonstrates
that many foreland lobes reached their maximum position before
23e22 ka, thus during Greenland Stadial 3 (Rasmussen et al., 2014)
or the global LGM as proposed by Hughes and Gibbard (2015). An
early advance, which would correlate to the Santa Margherita
phase of the Tagliamento Glacier (26.5e23 ka), has also been dated
in other Italian morainic amphitheaters (e.g. Ravazzi et al., 2012;
Scapozza et al., 2015). This advance could have its equivalent in
the Maritime Alps in the Durance Valley (Jorda et al., 2000) and the
Gesso Glacier (Federici et al., 2012, 2016). Similarly, in the northern
Swiss foreland, exposure ages pinpoint the initiation of withdrawal
from the maximum positions at 24 ± 2 ka at the Rh^one glacier (IvyOchs
et al., 2004) and at 22 ± 2 ka at the Reuss glacier (Reber et al.,
2014; Ivy-Ochs, 2015). The presented exposure ages from the
trimlines at Mont Blanc and in the Zillertal Alps show that the
lowering of the ice surface was certainly delayed with respect to the
abandonment of the maximum positions of piedmont glaciers at
many places around the Alps at 26e22 ka (Fig. 5).
Numerous chronological constraints on the fluctuation of the
foreland lobes suggest abandonment of the maximum positions or
recessional stages which overlap in time with the Remanzacco
phase (>19 ka, Monegato et al., 2007): 18.6e17.9 ka for the Oglio
Glacier in the Iseo amphitheater (Ravazzi et al., 2012); 19.8e17.6 ka
for the Traun Glacier (van Husen, 1997) and 19.7 ± 1.4 ka for the
Isar-Loisach Glacier (Reuther et al., 2011); in the northwest
18.6 ± 0.9 ka for the Reuss Glacier (Reber et al., 2014) and 18 ± 1 ka
for the Linth-Rhein Glacier (Lister, 1988). Initial lowering of the ice
surface in the High Alps therefore appears to have occurred broadly
synchronous (within the resolution of the dating method) to
marked retreat of the piedmont lobes and ice-free forelands soon
after 19 ka (van Husen, 1987; Lister, 1988; Schlüchter and
R€othlisberger, 1995; Monegato et al., 2007; Reitner, 2007;
Lauterbach et al., 2012; Schmidt et al., 2012; Baroni et al., 2014;
Ravazzi et al., 2014).
Delayed ice surface lowering in the High Alps with respect to the
initial retreat from the maximum position of glacier lobes in the
forelands (26e22 ka) can be explained by its location in the upper
reaches of the glacier systems, more than 100 km from the piedmont
lobe termini. In those settings a smaller and mitigated change
of ice surface elevation in reaction to the fluctuation in the glacier
terminal zones is reasonable. Because our samples are not located
directly at the trimline there could have been undetected initial
lowering of ice surface elevation before 18.5 ka. The oldest ages
obtained in the Mont Blanc area (Cou3) and in the Zillertal Alps
(ZA1,6,17,30) are located 90 and on average 120 m below the
trimline, respectively. Assuming that the glacial trimline in the Alps
represents the subaerial to subglacial boundary, this means that
these locations were exposed for the first time when glaciers had
lost about 10% of their maximal thickness in the valleys which was
Fig. 5. Ages correlated to the LGM in the Alps. The map shows the results of this study in the context of a compilation of published ages by Ivy-Ochs (2015) that were correlated to
the end of the LGM in the Alps. The type color indicates in which context an age has been interpreted by the authors. The LGM ice extent is from Ehlers and Gibbard (2004, http://
www.qpg.geog.cam.ac.uk/lgmextent.html, accessed 20.11.2015).
C. Wirsig et al. / Quaternary Science Reviews 143 (2016) 37e50 47
~1000 m (Fig. 3). As shown by Wirsig et al. (2016), exposure ages
from a high elevation trimline site in the central Swiss Alps can
record stepwise lowering after full LGM conditions with the first
sign of ice surface lowering about 50 m at 23 ± 0.8 ka, tentatively
linked to the initial foreland glacier retreat after reaching of the
maximum LGM positions. However, more evidence in other locations
in the Alps is needed to confirm this hypothesis.
It cannot however be excluded that some portion of the terrain
above the trimline was covered by cold-based, non-erosive ice and
that the trimline in the Alps represents an englacial thermal
boundary. Nevertheless, this hypothesis would be hard to prove or
disprove in the Alps using exposure ages. For example, in the British
Isles the blockfield-mantled summit plateaus are clearly distinguishable
morphologically from the glacially eroded ice-stream
valley slopes (Ballantyne, 1997; Ballantyne and Stone, 2015). The
periglacial zone in the Alps is completely different; the zone above
the glacial trimline comprises sharp ar^etes and steep rock walls.
Deposition and/or preservation of datable erratic boulders is
therefore unlikely. Dating the bedrock surfaces above the trimline
has been attempted previously in the Alps: B€ohlert et al. (2008)
presented ages ranging from 40 to 2 ka from steep rock walls and
ar^etes located far above the LGM ice surface in the Mont Blanc area.
This showed that in both cases (subaerial to subglacial boundary
and englacial thermal boundary) one could equally obtain preLGM,
LGM and post-LGM ages.
Whether or not there was an englacial thermal boundary,
retreat of the glacier lobes after 19 ka in the foreland and the onset
of loss of the major part of ice in the High Alps at 18.5 ka occurred
nearly synchronously. This implies prominent reorganization of
Alpine ice cover from vast 100e300 km long transection glaciers
supplying the foreland glacier lobes (van Husen, 1987; Bini et al.,
2009) to one order of magnitude shorter local mountain glaciers
during the Gschnitz stadial advance (van Husen, 1997; Ivy-Ochs
et al., 2006a; Ivy-Ochs, 2015). Therefore, most of the LGM ice volume
in the High Alps must have disappeared before 17e16 ka,
because at this time Gschnitz glaciers advanced into ice-free terrain
(Ivy-Ochs et al., 2008). This left very little time, ca. 1e2 ka, for
complete deglaciation of 1000e1500 m of LGM ice in the High
Alpine valleys and implies rapid ice surface lowering rates
(Schlüchter, 1988).
The onset of ice surface lowering in the High Alps is potentially
synchronous to the L€angsee warm oscillation at 18.5e18.1 ka that
preceded the return of severe climate cooling during Heinrich 1
event and the Gschnitz stadial glacier advances (Schmidt et al.,
2012). This warming triggered a glacier mass balance change
above a critical threshold that no longer allowed support of the full
LGM glacial geometry condition in the Alps (Reitner, 2007). This
crisis could also be amplified with precipitation starving associated
with cessation of the southward shifted trajectory of North Atlantic
storm track in the Alps at 19 ka recorded in Sieben Hengste speleothems
(Luetscher et al., 2015). The onset of ice surface lowering
in the Alps is associated with accelerated retreat of glaciers in mid
latitudes at 19e18 ka (Schaefer et al., 2006; Clark et al., 2009),
where a globally increased rate of deglaciation after full LGM
conditions contributed to the 19 ka meltwater pulse (Lambeck
et al., 2000, 2014). Across the Eurasian Ice Sheet the timing of
advance and retreat were remarkably spatially variable. However,
the peak in its total area and volume occurred at 21e20 ka and
initial retreat occurred soon after 19 ka (Hughes et al., 2016b and
references therein). Similar conclusions can be drawn from the
LGM evolution of the British-Irish component of this ice sheet,
which at 19 ka was in crisis with considerable ice loss and dynamic
thinning (Clark et al., 2012). Exposure ages located in the center
zone of the LGM Welsh Ice Cap show rapid thinning after the LGM
at 19 ± 1 ka and exposure of summits in North Wales as nunataks
(Hughes et al., 2016a). The presented results strengthen the evidence
of remarkable climatic change and glacier reaction in the
mid-latitudes at 19e18 ka that was not seen in isotopic records
from high latitude Greenland ice cores (Schaefer et al., 2006).
6. Conclusions
The oldest exposure ages from bedrock and erratic boulder
surfaces from high elevation ridges just beneath the glacial trimline
in the High Alps are 18.5 ± 1.1 ka in the Mont Blanc area (Western
Alps) and 18.6 ± 1.4 ka in the Zillertal Alps (Eastern Alps). They
indicate that the ice surface in the High Alps likely remained close
to its maximum elevation throughout full LGM conditions when
glacier tongues fluctuated on the Alpine forelands (26.5e19 ka).
Trimline positions were finally abandoned synchronous with significant
retreat of the piedmont lobes soon after 19 ka, presumably
during the L€angsee warm oscillation (18.5e18.1 ka). Identical ages
from both study areas, as well a similar results obtained in the
Swiss Alps (Dielforder and Hetzel, 2014; Wirsig et al., 2016) suggest
synchronous initiation of LGM ice surface lowering in the accumulation
zones of transection glacier systems across the Alps.
Decline of Alpine ice cover during the early Lateglacial was
extremely rapid; most of the ice in the LGM accumulation centers
must have disappeared in a short time (~1e2 ka) before local
Gschnitz valley glaciers started to advance into ice-free terrain in
response to Heinrich 1 stadial conditions (17e15.4 ka).
The simultaneous ice-surface lowering and complete abandonment
of the forelands by the piedmont lobes soon after 19 ka
was likely amplified by moisture starvation due to northward shift
of the North Atlantic storm track at 19 ka (Luetscher et al., 2015).
Acknowledgements
This work was funded by the Swiss National Fonds (SNF), project
2-77099-11 to S. Ivy-Ochs and by AGH-UST statutory grant No.
15.11.140.626 to J. Zasadni. All members of the Laboratory of Ion
Beam Physics group at ETH are recognized for their commitment
and support of this project. Special thanks go to A. Madella and C.
Sp€otl for help during fieldwork.
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