Introduction
Glaciers are sensitive indicators of climate change as their length, area and
volume respond primarily to variations in air temperature and precipitation
(e.g. Oerlemans, 2001). In general small and steep alpine glaciers show
faster response (∼ 10 years) to climate fluctuations than large, less
inclined ice bodies (∼ 100 years) (Holzhauser, 1997). In the European
Alps, glaciers have undergone major variations at glacial-interglacial timescales as they greatly expanded during the last glacial period and contracted
dramatically during the last deglaciation (Ivy-Ochs et al., 2008). While
traces of the large expansions of the last glacial period are well preserved
even at very low elevations (Ravazzi et al., 2014), evidence for the
subsequent smaller Holocene glacier variations is most often overridden by
the Little Ice Age expansion (LIA; 14th–19th century). Nevertheless organic
fragments (e.g. wood, peat bogs) found recently in forefields of retreating
glaciers provide information on the lower altitude limits of past glacial
extents, demonstrating that glaciers in the Alps were smaller during the
mid-Holocene than they are today (Hormes et al., 2001; Joerin et al., 2006, 2008;
Nicolussi and Patzelt, 2000; Porter and Orombelli,
1985).
The end of the Younger Dryas (11.7 kyrs BP) is generally considered to mark
the onset of the Holocene. At that time conditions favourable to glaciers
persisted in the European Alps until 10.5 kyrs BP (Ivy-Ochs et al., 2009)
when a period of climatic warming started, culminating between 6 and
9 kyrs BP during the so-called Northern Hemisphere Climatic Optimum
(Vollweiler et al., 2006). At this time the Northern Hemisphere summer
insolation and solar irradiance reached maximum levels during the Holocene
(Berger and Loutre, 1991; Stuiver et al., 1998). Specifically, this climatic
optimum was characterized by three particularly warm phases at 9.2,
7.45–6.65 and 6.20–5.65 kyrs BP (Joerin et al., 2008). A climate that was
generally unfavourable for glacier advances persisted at least until
6.8 kyrs BP (Luetscher et al., 2011). Between 5.3 and 3.3 kyrs BP the
changed climatic conditions marked the beginning of the Neoglaciation in the
European Alps (Magny and Haas, 2004; Ivy-Ochs et al., 2009). During this new
phase glaciers showed larger variations which culminated in three large LIA
expansions (14th, 17th and 19th centuries) (Holzhauser et al., 2005),
followed by an ongoing phase of intense glacier waning (Zemp et al., 2006).
At the end of summer 1991, the 5300-year-old Tyrolean Iceman mummy emerged
from the ablating ice field of the Tisenjoch, a saddle at 3210 m near the
Italian–Austrian border in the eastern Alps (Seidler et al., 1992). The
excellent state of preservation of the Tyrolean Iceman provides strong
evidence for the minimum coverage of this ice field at this time, which has
only recently been surpassed. This discovery suggests that past atmospheric
temperatures characterizing warm phases such as the Roman (250 BC–400 AD)
and the Medieval (950–1250 AD) periods may have never exceeded that of the
current time in this sector of the Alps (Baroni and Orombelli, 1996).
However, a recent chironomid-based summer air temperature reconstruction from
Lake Silvaplana, in the nearby Upper Engadine (Switzerland), documented that
the 20th century and the Medieval period (from 1030 AD, start of the
dataset, until 1260 AD) were both 1 ∘C warmer than the modern
climate reference period (1961–1990) in the eastern Alps (Larocque-Tobler
et al., 2010). Clearly, information about the glaciation of the eastern Alps
before the LIA remains sparse (Nicolussi and Patzelt, 2000).
In this context ice cores can provide useful information. In the western Alps, an ice core extracted from Colle Gnifetti (4450 m, Monte Rosa,
Italian–Swiss border) provided evidence for more than ∼ 10 kyr-old ice
in its lower section (Jenk et al., 2009), suggesting a continuous glaciation
of at least the highest locations of the western Alps throughout the
Holocene. In 1991, at the time of the discovery of the Tyrolean Iceman, only
pollen analyses (Bortenschlager et al., 1992) were performed on the ice in
which this mummy was embedded for ∼ 5.3 kyrs at the Tisenjoch, which
is now completely deglaciated. This is unfortunate because this now melted
ice had the potential to be as old as the Tyrolean Iceman himself, and may
have preserved a unique snapshot of additional information of the past
environmental conditions experienced by the Alpine populations during the
mid-Holocene.
In 2010, we formulated the hypothesis that an ice core record encompassing
the time of the Tyrolean Iceman was embedded in the upper reaches of the Alto
dell'Ortles (3859 m), the main glacier of Mt. Ortles (3905 m, Italy), which
is the highest mountain of South Tyrol, located 37 km from the Tisenjoch
(Gabrielli et al., 2010). This idea was based on the following observations:
(i) Alto dell'Ortles is wind-exposed and located in a rain shadow (Schwarb,
2000), and thus is likely characterized by a low accumulation rate compared
to the average of the Alps; (ii) the upper exposed margins of Alto
dell'Ortles show laminated ice layers down to bedrock; and (iii) the
concomitant significant thickness and moderate tilt (8–9∘) of the
upper part of Alto dell'Ortles may be indicative of minimal basal flow and a
frozen ice/bedrock interface, as it could have also been expected at this
elevation in this area (Suter et al., 2001).
Until recently, alpine ice core records have been obtained only from the
western Alps (Barbante et al., 2001, 2004; Jenk et al., 2009; Legrand et al.,
2003; Preunkert et al., 2000, 2001; Schwikowski et al., 1999a, b; Van de Velde et al., 2000a; Wagenbach et al., 1988) because of
their high elevation and the consequent common occurrence of the cold firn
zone (Golubev, 1975; Haeberli and Alean, 1985), which is more likely to retain
climatic and environmental signals (Eichler et al., 2001). In contrast,
because of their lower elevation, the highest glaciers in the eastern Alps
were assumed to be entirely within the temperate firn zone, and thus
unsuitable for preserving intact ice core records (Oerter et al., 1985).
However, Suter and co-workers modelled the firn temperature in the Alps and,
based on altitude and exposure, suggested that cold firn should exist above
3400 m in northerly aspects and above 4150 m on southern slopes (Suter et
al., 2001).
Modern climatic conditions in the eastern Alps are very unusual because since
∼ 1980 summer air temperatures have shown a step increase of about
2 ∘C at high elevations (Gabrielli et al., 2010; Auer et al., 2006).
Consequently the extensive summer meltwater percolation through the shallow
(< 10 m) temperate firn layers, recently observed on Mt. Ortles,
could be a relatively recent phenomenon that intensified after ∼ 1980
(Gabrielli et al., 2010). Interestingly, ice that preserved climatic and
environmental signals was already found in 2003 below temperate firn on the
Quelccaya ice cap in the Peruvian Andes, where recent and strong meltwater
percolation did not affect the climatic signal embedded within the deepest
impermeable ice layers (Thompson et al., 2006). We likewise speculated that a
climatic signal might still be preserved within the deep ice layers of the
glacier Alto dell'Ortles (Gabrielli et al., 2010).
During the autumn of 2011, working within the framework of an international
program aimed at studying past and present climate conditions in the Alps
(“Ortles Project”, www.ortles.org), we extracted the first ice
cores drilled to bedrock in the eastern Alps from Alto dell'Ortles
(Gabrielli et al., 2012). Here we present glaciological observations (a
digital elevation model, englacial characteristics, glacial dynamic and
bedrock topography of the drilling site) and various dating techniques (based
on tritium, beta emissions, 137Cs, 210Pb and 14C) that
allowed a timescale for the Mt. Ortles ice cores to be obtained. We show that the
Alto dell'Ortles cores contain records at millennial timescales, and that the
bottom ice dates back ∼ 7 kyrs BP. Combining ice core and
glaciological observations, we discuss these findings in light of the state
of the knowledge of the glaciation of the eastern Alps during the Holocene.
The drilling site
General characteristics
The Alto dell'Ortles glacier covers the northwestern side of Mt. Ortles,
which gently slopes (8–9∘) from near the summit for ∼ 300 m,
then flows on steeper bedrock into two major tongues down to 3018 m
(Fig. 1). According to some recent lidar measurements (see Sect. 2.2.2),
the total surface area was 1.12 km2 in 2011 and 1.07 km2 in
2013, of which ∼ 10 % constitutes the upper gentle plateau. The ice
core drilling campaign was conducted during September and October 2011 on a
small col (3859 m; 10∘32′′ 34, 46∘30′′ 25) between the
summit of Mt. Ortles (3905 m) and the Vorgipfel (3845 m) (Figs. 1–2). At
the drilling site, the bedrock is at ∼ 75 m of depth (Gabrielli et
al., 2010, 2012) and the current accumulation rate (2011–2013) is
∼ 800 mm water equivalent (w.e.) per year.
(a) Geographic location of Mt. Ortles. (b) Map of
the Alto dell'Ortles glacier (South Tyrol, Italy), including the area (box)
where the drilling operation was conducted during September–October 2011.
(c) Detailed map of the drilling site, including (i) the specific
locations where the four cores were extracted and (ii) the traces of the
detailed ground penetration radar (GPR) survey performed in July 2013.
Comparison of terrestrial photographs of Mt. Ortles taken from the
summit of Gran Zebrù (3851 m) (a) during the years 1900–1930
(http://www.montagnedifoto.com/, last access 16 March 2016) and
(b) on the 4 July 2010 (photo: Roberto Seppi). The triangle in panel
(b) shows the position of the 2011 drilling site. The four symbols
(+) indicate the reference points used for co-registering the two photos
(ESRI ArcMap™) before estimating the thickness
variation at the drilling site during this period.
Over the last 3 decades (1980–2009) the reconstructed average summer
(JJA) air temperature was -1.6 ∘C, ∼ 2 ∘C higher
than during the previous 115 years, with a peak of +2 ∘C during
the summer of 2003 (Gabrielli et al., 2010). In 2011, englacial temperatures
provided firm evidence for the concomitant presence of a temperate firn
portion, deep cold ice layers and a frozen bedrock. In fact, thermistors
located within the firn indicated temperatures at or near the pressure
melting point, while those positioned in the ice (below the firn ice
transition at ∼ 30 m depth) clearly demonstrate negative temperatures
at 35 m (-0.4 ∘C), 55 m (-1.8 ∘C) and at 75 m
(-2.8 ∘C) close to bedrock, confirming the presence of cold ice
(Gabrielli et al., 2012). We concluded that this glacier probably represents
a unique remnant of the colder climate prior to ∼ 1980, which has
since been shifting from a cold to a temperate state.
Current dynamic
Elevation changes
Comparison of terrestrial photographs of Alto dell'Ortles taken in 2010 and
during the period from 1900 to 1930 (Fig. 2), suggests a thinning of 8–10 m
at the drilling site. Comparison of digital terrain models (DTMs) obtained
from topographic maps created in 1962 and 1984 (obtained by aerial
photogrammetry and provided by the Istituto Geografico Militare and Province
of Bolzano, respectively) and from lidar surveys in 2005 and 2013 (provided
by the Province of Bolzano and by the Institute of Atmospheric and
Cryospheric Sciences, University of Innsbruck; Galos et al., 2015) indicate
a major thinning at the drilling site from 1962 to 1984
(-25.0 ± 4.7 m). From 1984 to 2005 thickening prevailed
(+10.5 ± 7.9 m), followed by minor elevation changes from 2005 to
2013 (-0.7 ± 1.0 m). More extended analyses covering the upper 50 m
of the glacier Alto dell'Ortles, which includes the drilling site, indicate
that thinning was widespread from 1962 to 1984 (-9.5 ± 4.7 m) in the
upper part of this glacier, followed by rather stationary conditions
(-1.5 ± 0.3 m) between 1984 and 2013.
Reconstructed flow lines and borehole displacements over time. The
surface topography of the drill site was obtained from a lidar survey
conducted during the 2011 campaign (note that the contour lines of the
drilling dome are visible). Inset: the displacement between 5 October 2011
and 7 September 2012 is shown in red, while the shift between 7 September 2012
and 1 July 2013 is in green (borehole no. 3 only). The values indicate the
displacement (in metres) measured by GPS during these two periods.
These observations indicate that (i) the upper part of Alto dell'Ortles was
subject to significant elevation changes during the last decades; (ii) the
drilling site itself experienced even larger elevation changes; and
(iii) these elevation changes are not directly linked to atmospheric changes
(e.g. summer atmospheric warming and surface glacier ablation). It is indeed
remarkable that while the site thinned during the relatively cold period
between the 1960s and 1980s, most glaciers at lower altitude expanded in this
geographic area (Carturan et al., 2013 and references therein). Local
elevation changes at the drilling site likely result from the interplay of
glacier dynamics and spatial variability of ablation and, notably at this
high-elevation site, snow accumulation and redistribution by the wind.
Surface and internal dynamics
A borehole displacement of 3.2 m yr-1 over 1.7 years
(5 October 2011–1 July 2013) was determined at the glacier surface by means of differential GPS
measurements (Fig. 3). However, the glacial flow was not constant during this
period as it varied between 3.7 m yr-1 (5 October 2011–7 September 2012)
and 2.6 m yr-1 (7 September 2012–1 July 2013), suggesting a
seasonal variability characterized by a higher flow in summer than in winter.
The direction of the measured displacement of the boreholes and the
variability of the glacial flow during the observed periods are consistent
with (i) glacier flow lines originating from the southern flank of
Mt. Ortles, whose summit is located 270 m uphill from the drill site
(Fig. 3); and (ii) basal sliding of the glacier bed lubricated by summer
meltwater, perhaps percolated from the outcrops of bedrock located uphill
from the drilling site. The surface flow lines have been inferred from a DTM
obtained with a lidar survey performed during the 2011 ice core drilling
campaign, using the flow accumulation tool of ESRI
ArcMap™. Importantly, this analysis also shows
that the drilling site was located at or in very close proximity to the ice
divide (as derived from the surface topography, Fig. 3).
In order to infer the internal dynamics of the glacier, inclinometric
measurements were performed in borehole no. 2 43 days after the end of the
drilling operation (5 October–17 November 2011). Further measurements were
not possible because of the rupture of the pipe (or the formation of an
internal ice lens) at 25 m depth. Uncertainty of this measure is
±6 mm/25 m. A cumulative displacement of 277 mm (2.4 m yr-1)
relative to the bottom part of the inclinometer was observed on the glacier
surface (Fig. 4). We also note that the glacier flow decreased linearly with
depth, which is inconsistent with the velocity fields typically recorded
within glaciers frozen at the bed (Paterson, 1999). While the relative
inclinometric measurement does not necessarily imply a net basal sliding of
the drilling site, it does indicate that currently ice layers located next to
bedrock are dynamically active (38 mm (0.3 m yr-1) at 65 m; 18 mm
(0.2 m yr-1) at 70 m). This information is important in order to
evaluate the age of the basal ice (see discussion in Sect. 6.2).
Relative cumulative displacement of borehole no. 2 along the two
axes (A: 340∘ N; B: 70∘ N) during the 43 days after the end
of the 2011 drilling operation. The cumulative displacement is relative as it
uses the bottom portion of the inclinometer as a reference (not the bedrock).
Bedrock topography
Ground penetrating radar (GPR) was used to determine the bedrock topography
(e.g. Binder et al., 2009; Moran et al., 2000) and to infer information about
possible englacial features (e.g. Blindow and Thyssen, 1986; Konrad et al.,
2013). GPR profiles of 50 MHz were collected with a GSSI SIR 3000 system during
July 2013. This spatial survey focused on a region of 50 × 50 m,
including the four 2011 boreholes with an inter-profile distance of 4 m
(Fig. 1c). Based on the available continuous snow/firn/ice density data from
a snow pit and from borehole no. 2 (BH2), a 1-D velocity function was derived
from the correlation between snow/ice density and dielectric permittivity by
Kovacs et al. (1995). The two-way travel times (TWTs) chosen were converted to
depth with the 1-D velocity function, and interpolated to continuous surfaces.
The top images illustrate the GPR X-profiles (constant
y coordinate) and Y-profiles (constant x coordinate) through the
locations of boreholes no. 1–4 (see also Fig. 6). Two-dimensional slices of
the interpolated bedrock surfaces (thick grey line) and a possible englacial
layer (thin grey line) are shown. The thick blue line indicates the glacier
surface. Black dots show the logged depths for boreholes no. 1–4.
The reflection horizon at TWTs of 800–900 ns (∼ 75–85 m) was
interpreted as bedrock. For BH1, BH2 and BH3 there is a good correspondence
(within 1 m) between the GPR-derived ice thicknesses and the ice core
lengths (Fig. 5). The GPR-derived ice depth for BH4 indicated that the ice
core drilling was stopped ∼ 15 m from bedrock. Since the applied snow/ice
density–dielectric permittivity correlation is valid for dry polar
conditions, this confirms a dry and cold glacial body at the drilling site.
Figure 5 shows orthogonal profile slices through the four boreholes according
to the coordinate system displayed in Fig. 6.
Bedrock contours and surface topography of the 2011 ice core drill
site. Data were obtained during the 2013 GPR and lidar surveys, respectively.
Bedrock contours are shown in green, while the surface topography is displayed
in black. The positions where the four cores were extracted in 2011 are shown
as filled red circles. The GPR X-profiles (constant
y coordinate) and Y-profiles (constant x coordinate) through the
locations of boreholes no. 1–4 are also shown (see also Fig. 5).
A continuous internal layer was identified 20–40 m above the bedrock
(Fig. 5). The internal layer could be tracked throughout the investigated
area, which suggested an isochronical origin. Two close melt ice layers at a
depth of about 45 m corresponded well with this spatially continuous
englacial reflection. The inclination of this melt layer is consistent with
the higher (+200 mm w.e.) modern annual snow accumulation (1000 mm w.e.)
observed on the left side (Fig. 5, X-profiles) of the drilling site at
3830 m (Festi et al., 2015). It is likely that surface snow was blown away
from the ice ridge and redeposited further down the slope. Thus the steep
isochrone suggests that the oldest ice of the glacier can be found on the
right side (Fig. 5, X-profiles), which is under the ice divide.
Alternatively, it indicates that the oldest ice can be found at higher depth
resolution below the ice divide.
Detailed bedrock topography at the drill site, as obtained by spatial
interpolation of GPR point measurements, is illustrated in Fig. 6. We note
that the boreholes reached bedrock near a ∼ 10 m step located between
the drilling site and the Vorgipfel (Figs. 1, 5, and 6). However, this feature is
only on one side of the drilling site, and thus we conclude that while this
morphological feature may have facilitated the in situ retention of
old bottom ice, it is unlikely to have caused a complete dynamical entrapment
and the consequent formation of fossil ice decoupled from the upper
stratigraphic sequence, which is also implied by the ice flow observed near
bedrock (Fig. 4).
Ice core characteristics
The four cores were drilled on Alto dell'Ortles within ∼ 10 m
(Fig. 1), reaching final logged depths of 73.53 m (no. 1), 74.88 m (no. 2),
74.83 m (no. 3) and 61 m (no. 4). The drilling of core no. 4 was stopped
for technical issues at a depth ∼ 15 m above bedrock as determined by
GPR (see previous section). It has been stored intact for possible future
complementary analysis. Likewise, the drilling of core no. 1 stopped at
73.53 m just short of bedrock because of technical issues. On the other
hand, it was obvious that cores no. 2 and no. 3 did reach bedrock since
further penetration was not possible and damage to the drill cutters was
observed. The comparable final logged depths of core no. 1, no. 2 and no. 3
(73.53, 74.88 and 74.83 m) are in close agreement with the glacier thickness
(75 m) as determined by GPR before (2009) (Gabrielli et al., 2010) and after (2013) the
drilling operation (see previous section).
Consistent with the geology of the Mt. Ortles summit, a few limestone rock
particles and pebbles were observed in the deepest sections of cores no. 1,
no. 2 and no. 3 within ∼ 1 m of bedrock, providing visible evidence
for the bed material in the bottom ice. In core no. 2 a single large pebble
(∼ 1 cm) was observed 2.77 m above bedrock. This was either entrained
from the bed or from the glacier surface because of the short distance
(∼ 250 m) of the drilling site from the rock outcrops of the
Mt. Ortles summit.
Physical characteristics of the Mt. Ortles cores. The temperate
firn portion is enclosed in red shading, while the cold ice is in blue.
(a) Borehole no. 3 temperatures recorded 43 days after the end of
the drilling operations (from Gabrielli et al., 2012). (b) Virtual image of
core no. 1 reconstructed from 360∘ Televiewer visual scanning of
borehole no. 1. (c) Red component of the RGB digital signal obtained
by means of visual scanning. High values indicate higher light reflection.
(d) Densities of the Mt. Ortles ice cores no. 1, no. 2 and no. 3.
Density measurements of cores no. 1, no. 2 and no. 3 indicate a firn/ice
transition at ∼ 30 m depth (Fig. 7), with a measured average ice
density of 882 kg m-3. The air bubbles entrapped within the ice layers
throughout the core range from a few millimetres to less than ∼ 1 mm in
diameter. Elongated air bubbles (up to 10–15 mm) are widespread through the
cores, confirming that flow is a significant component of the ice dynamics of
this drilling site (Gabrielli et al., 2012).
Direct observations of the three cores were performed and borehole images
were obtained from a 360∘ continuous-imaging scan of borehole no. 1
(Optical Televiewer, Advanced Logic Technology, Luxembourg). As light
reflectance is determined by the concentration and size of air bubbles, this
technique highlights the core layers and the presence of ice lenses (melt
layers). Horizontal to tilted (10–20∘) bubble-free ice lenses are
present throughout the entire lengths of the cores, where the angles may
reflect, at least in part, the tilt of the basal slope. The cumulative
thickness of the ice lenses constitutes ∼ 20 % of the entire length
of the firn portion (when expressed in ice equivalent), ∼ 15 % of the
glacier ice between 30 and 55 m and ∼ 5 % between 55 and 65 m
depth (Gabrielli et al., 2012).
Fission products determined in the Mt. Ortles cores no. 1, no. 2
and no. 3. The temperate firn portion is enclosed in red shading, while the
cold ice is in blue. Tritium is depicted by blue triangles, beta emissions by
red dots (BPCRC) or red crosses (LGGE) and 137Cs by green diamonds. Beta
emissions in cores no. 2 and no. 3 are determined at different resolutions
depending on the sample mass available.
Measured 210Pb activity (left y axis, logarithmic scale) and
calculated age (right y axis) vs. depth (in metres of water equivalent)
relationship in the Mt. Ortles core no. 2. The fit is also reported for
210Pb activity (logarithmic, black) and age (linear, blue); the latter
is shown with its 1σ confidence interval (in blue).
Borehole images of core no. 1 also suggest strong ice layer thinning from
60 to 65 m depth to the basal ice (Fig. 7). In this case we interpret the low
reflectance of this basal ice as a consequence of the bubbles shrinking due
to the overburden pressure over thousands of years (see Sect. 4.3). The
digital red index associated with the borehole images (this expresses
numerically the red component of each pixel that displays colours as a
combination of red, green and blue; RGB colour code) suggests that this
transition does not occur abruptly (as expected in the case of a physical
hiatus) but instead over ∼ 5 m, perhaps indicating a continuous change
of the physical properties of the ice over time. This observation is also
important in order to evaluate the possible presence of a decoupled fossil
ice portion near the bedrock.
Ice core samples analysis
Fission products
Beta activity and tritium
Depending on the ice mass available, analyses of beta activity (Byrd Polar
and Climate Research Center, BPCRC) and tritium (University of Bern and
University of Venice) were performed with various degrees of continuity and
resolution in various sections of the cores using established methods (Maggi
et al., 1998; Schotterer et al., 1998; van der Veen et al., 2001) (Fig. 8).
Three sections from core no. 2 (2.75, 22.95 and 41.27 m) were reanalysed for
beta activity at the Laboratory of Glaciology and Geophysics of the
Environment (LGGE) in Grenoble, France, by means of a Berthold LB770-2
gas-flow proportional counter (Ar/CH4 gas) (Magand, 2009; Pourchet et
al., 2003; Vimeux et al., 2008), showing variations that are consistent with
those determined at BPCRC (Fig. 8).
A well-defined peak of beta and tritium activity can be observed at 41 m
depth, indicating the 1963 radioactive horizon resulting from nuclear weapon
testing, which is consistent with 210Pb dating (see below). The match
between the tritium (entrained in the ice matrix) and beta emission (emitted
by ions such as 90Sr) peaks in core no. 2 suggests that
post-depositional effects due to meltwater percolation were negligible from
the time of the deposition of this radioactive layer (1963) until it was
entrained below the firn–ice transition. This is also a preliminary
indication that the chemical stratigraphy in the firn and in the ice was
preserved before the onset of the exceptional current warming (1980 in this
area) and the linked surface melting and meltwater percolation (Gabrielli et
al., 2010, 2012). Scrutiny of the well-resolved beta record in core no. 2
indicates a secondary beta emission peak at 44 m depth, likely resulting
from the widely reported 1955–1958 thermonuclear tests (Gabrieli et al.,
2011).
Evidence of the radioactive fallout from the 9 March 2011 Fukushima nuclear
plant accident is observed in the shallow spring–summer 2011 layer (Fig. 8).
In this case, while tritium levels remain low, beta activity reaches values
that are comparable to the residual radioactivity released by the ice layers
contaminated by the atmospheric nuclear tests of the 1950s and the beginning
of the 1960s. Glaciological evidence for the Fukushima radioactive fallout has
already been reported from Tibetan Plateau (Wang et al., 2015) and Arctic
(Ezerinskis et al., 2014) snow samples. Our data from the eastern Alps are
consistent with the clearly detectable Fukushima radioactive fallout widely
observed in Europe (Masson et al., 2011), including the northern Italian city
of Milan (Clemenza et al., 2012).
An additional peak of beta activity can be observed in the Mt. Ortles cores
at ∼ 28 m of depth (Fig. 8), near a thick ice lens immediately above
the firn ice transition (∼ 30 m). This signal does not seem to be
directly related to the 1986 Chernobyl radioactive fallout as its timing is
inconsistent with the 210Pb age determined at this depth
(1979.5 ± 3; see Sect. 4.2). It is more likely that percolating summer
meltwater transported Chernobyl radionuclides through the porous temperate
firn from their actual deposition layers (see next section) down to
∼ 28 m where at least part of the water refroze and the radionuclides
were accumulated. This process also suggests that the chemical stratigraphy
is not well preserved in the firn portion of the Mt. Ortles cores.
137Cs
The three sections analysed for beta emissions at LGGE (see previous section)
were also analysed by a very low background germanium planar detector at the
Laboratoire Souterrain de Modane (LSM – 4800 water equivalent) in France
(Loaiza et al., 2011). These measurements were below the detection limit with
the exception of the section at 41.27 m (1963), which showed a non-decay-corrected activity of 0.035 ± 0.003 (Bq kg-1) for 137Cs.
This result is consistent with 137Cs values typically linked to the 1963
radioactive deposition
(United Nations Scientific Committee on the Effects of Atomic Radiation, 2000).
The undetected 137Cs in the 22.95 m depth layers (∼ 1986–1987
according to 210Pb dating, see next section) is notable, taking the large 137Cs quantities released in the European atmosphere
during the 1986 Chernobyl accident into account
(United Nations Scientific Committee on the Effects of Atomic Radiation,
2000). This provides additional support for the idea of the occurrence of
post-positional effects due to abundant multi-year meltwater percolation
through the firn during the recent warm summers (Gabrielli et al., 2010).
The very low 137Cs values measured in Mt. Ortles snow within the top
2.75 m (spring–summer 2011) after the Fukushima accident are consistent with
the low levels measured in aerosols over Europe, which is 3 to 4 orders of
magnitude lower than activity levels encountered after the Chernobyl event
(Mietelski et al., 2014; Povinec et al., 2013). Therefore, it is more likely
that the beta activity detected in the Mt. Ortles shallow 2011 snow layer was
the product of long-range transport of other Fukushima-derived radionuclides,
especially131I (Clemenza et al., 2012; Mietelski et al., 2014; Povinec
et al., 2013; Lin et al., 2015). Given the short half-life of 131I
(8 days), it might be expected that this layer could be only a short-term
glaciological reference.
210Pb dating. Determinations of 210Pb activity
and calculated ages for the upper part of the Mt. Ortles ice core no. 2
(0–58.67 m). Note the notation used for ages in yrs b2012.
Top depth
Bottom depth
Top depth
Bottom depth
Activity in 2011
Uncertainty
Age
Lower age -1σ
Upper age +1σ
Age
(m)
(m)
(m w.e.)
(m w.e.)
(mBq kg)
(mBq kg-1)
(yrs b2012)
(yrs b2012)
(yrs b2012)
(AD)
0.00
2.75
0.00
1.46
149.4
3.9
1.5
1.4
1.6
2010.5
2.75
5.56
1.46
3.02
120.4
3.4
4.3
3.9
4.6
2007.7
5.56
8.31
3.02
4.66
69.5
2.6
7.2
6.6
7.8
2004.8
8.31
11.09
4.66
6.33
82.7
2.7
10.2
9.3
11.1
2001.8
11.09
13.85
6.33
8.08
111.3
3.1
13.3
12.1
14.5
1998.7
13.85
16.64
8.08
9.89
60.4
2.2
16.6
15.1
18.0
1995.4
16.64
19.44
9.89
11.84
37.0
1.5
20.0
18.2
21.7
1992.0
19.44
22.25
11.84
13.92
21.6
1.2
23.6
21.6
25.7
1988.4
22.25
25.05
13.92
16.20
49.3
2.1
27.6
25.2
30.1
1984.4
25.05
27.88
16.20
18.52
37.8
1.7
31.8
29.0
34.6
1980.2
27.88
30.68
18.52
20.88
33.9
1.6
36.1
32.9
39.3
1975.9
30.68
33.52
20.88
23.24
26.0
1.3
40.4
36.8
43.9
1971.6
33.52
36.34
23.24
25.83
32.2
3.5
44.9
40.9
48.8
1967.1
36.34
39.15
25.83
28.40
36.1
2.0
49.6
45.2
54.0
1962.4
39.15
41.92
28.40
30.94
22.8
1.7
54.2
49.4
59.0
1957.8
41.92
44.74
30.94
33.52
18.7
1.9
58.9
53.6
64.1
1953.1
44.74
47.48
33.52
36.03
14.6
1.1
63.5
57.9
69.1
1948.5
47.48
50.27
36.03
38.59
12.6
2.5
68.1
62.1
74.2
1943.9
50.27
53.07
38.59
41.16
14.9
1.3
72.8
66.3
79.3
1939.2
53.07
55.88
41.16
43.74
7.9
0.8
77.5
70.6
84.4
1934.5
55.88
58.67
43.74
46.29
7.6
0.7
82.1
74.8
89.5
1929.9
14C analyses of the particle organic fraction (WIOC)
obtained from the four sections (tubes) of the Mt. Ortles ice cores no. 1 and
no. 3. Except for Sect. 103b, the samples were analysed in three
subsamples (top, middle, bottom). 14C determination in
Sect. 105b (core no. 1) refers to a larch leaf that was found in the ice.
Samples reported in bold are those also included in Table 3. Note the
notation used for calibrated ages in yrs b2012.
Core
Tube
Measure
Top
Bottom
WIOC
F14C
14C age
Cal age
µcal age
µcal age
σ
no.
no.
depth (m)
depth (m)
(µg)
(yrs BP)
(yrs cal BP)
(yrs cal BP)
(yrs b2012)
(years)
1
98b
WIOC
68.26
68.49
17.11
0.971 ± 0.024
236 ± 199
(-4–461)
279
341
167
1
98b
WIOC
68.49
68.73
17.86
0.911 ± 0.021
749 ± 185
(550–902)
732
794
163
1
98b
WIOC
68.73
68.96
15.06
0.900 ± 0.024
846 ± 214
(562–974)
824
886
192
98b WIOCb
68.26
68.96
0.927 ± 0.025
609 ± 217
(331–790)
595
657
205
3
102
WIOC
70.87
71.14
7.98
0.784 ± 0.043
1955 ± 451
(1395–2431)
2011
2073
517
3
102
WIOC
71.14
71.35
7.15
0.867 ± 0.065
1146 ± 602
(537–1720)
1253
1315
620
3
102
WIOC
71.35
71.57
13.28
0.818 ± 0.032
1614 ± 314
(1187–1921)
1590
1652
347
102 WIOCb
70.87
71.57
0.823 ± 0.027
1565 ± 264
(1262–1818)
1521
1583
286
1
103b
WIOC
71.8
72.48
10.37
0.932 ± 0.037
569 ± 320
(156–903)
517
579
291
1
105b
Larch leaf
73.25
73.25
68a
0.728 ± 0.006
2550 ± 65
(2500–2752)
2612
2674
101
3
106
WIOC
73.73
74.02
10.91
0.628 ± 0.031
3737 ± 397
(3593–4787)
4173
4235
523
3
106
WIOC
74.02
74.24
11.50
0.568 ± 0.030
4544 ± 424
(4623–5715)
5178
5240
530
3
106
WIOC
74.24
74.47
18.47
0.481 ± 0.020
5879 ± 334
(6354–7156)
6742
6804
365
a Pure C extracted after combustion.
b Combined values from the three subsamples of tubes 98b and
102.
Alignment of the stable isotopic profiles from Mt. Ortles cores
no. 1, no. 2 and no. 3 by depth. Vertical bars indicate the points used to
tie cores no. 1 and no. 3 to core no. 2. The latter is used as the reference.
210Pb
210Pb activity was determined continuously in core no. 2 between 0 and
58.67 m depth at a sample resolution of ∼ 2.80 m length (Table 1)
using an established method (Gäggeler et al., 1983; Eichler et al.,
2000). The age–depth relationship was derived from the slope of the linear
regression of the logarithmic 210Pb activity as a function of depth
in metres of water equivalent (Fig. 9). The y axis intercept (109 ± 14 mBq kg-1)
corresponds to the 210Pb activity at the surface of the Alto dell'Ortles
glacier, and is comparable to values typically observed
(∼ 85 ± 10 mBq kg-1) in high-altitude glaciers in the Alps
(Eichler et al., 2000). At a depth of 41.92 m, the calculated age of
54 ± 5 years (1953–1963, 1σ range) is consistent with the 1963
beta and tritium activity peak found at 41 m (see previous sections). The
age of the lowest sample (bottom depth 58.67 m) dated by 210Pb is
82 ± 7 years (1923–1937, 1σ range).
14C analysis
Four large (∼ 1 kg) samples from Mt. Ortles cores no. 1 (68.96 and
72.48 m) and no. 3 (71.57 and 74.47 m) (Table 2) were selected for 14C
dating using a method based on 14C determination in the water-insoluble
organic carbon fraction (WIOC) of the aerosols in the ice (Jenk et al., 2006, 2007, 2009; Sigl et al., 2009). Core no. 2 was not
sampled because sufficient ice volume was not available. Each ice section was
divided into three subsamples (top, middle and bottom), which were filtered
separately and analysed. For the section at 72.48 m from core no. 1, the
amount of filtered WIOC from the first subsample was estimated to be
insufficient, and therefore the three subsamples were filtered together,
resulting in a total ice volume exceeding our standard dimensions, possibly
introducing a larger blank because of the modified treatment (i.e. increased
potential for contamination due to a higher number of steps during sample
processing).
14C analyses were conducted using the compact radiocarbon Accelerator Mass Spectrometer (AMS) system
“MICADAS” at the University of Bern (LARA laboratory). For details about
sample preparation, WIOC separation, blank correction and calculation of
F14C, see Uglietti et al. (2016). The conventional 14C ages were
calibrated using OxCal v4.2.4 software (Bronk Ramsey and Lee, 2013) with the
IntCal13 calibration curve (Reimer et al., 2013). Dates are provided as
radiocarbon ages (yrs BP), calibrated radiocarbon ages (yrs cal BP) or as
years before the date of sampling, with 2012 being the closest approximation
of a full year (yrs b2012 = 2012–1950 + yrs cal BP). Average
values (μ) are presented with a 1 sigma (σ) uncertainty.
The F14C results obtained for the three subsamples (top, middle, bottom) of
section 98b at 68.96 m (core no. 1) and 102 at 71.57 m (core no. 3),
respectively, were combined to obtain the most reliable ages representative
of the mid-depth of these sections (mean value of the three subsamples with
1σ standard error) with the data-supported assumption of a negligible
age spread within these two sections. The resulting mean calibrated ages were
595 ± 205 yrs cal BP and 1521 ± 286 yrs cal BP,
respectively (Table 2). On the other hand, the three subsamples from the
deepest section of core no. 3 (74.47 m) dated with an age of
4173 ± 523, 5178 ± 530, and 6742 ± 365 yrs cal BP,
respectively, indicate very strong glacier thinning close to bedrock, and were
thus treated as individual dating horizons. The analysis of the single sample
from the section at 72.48 m (core no. 1) provided an age of
517 ± 291 yrs cal BP that would represent a chronological inversion
with respect to all the other 14C ages. Considering the above-mentioned
increased risk of contamination during sample preparation this age value was
disregarded.
14C analyses of a larch leaf found at 73.25 m depth in core no. 1 were
performed by AMS at the National Ocean Sciences Accelerator Mass Spectrometry
facility at the Woods Hole Oceanographic Institution. This provided an
additional, conventionally derived 14C age of
2612 ± 101 yrs cal BP, which is stratigraphically and
chronologically consistent with the WIOC 14C ages of 1521 ± 286
and 4173 ± 523 yrs cal BP obtained from core no. 3 at 71.57 and
74.02 m of depth, respectively (Table 2). Stratigraphic consistency also
remains valid when a common depth scale for the three cores is adopted (see
next section).
Discussion and implications
Age of the bottom ice
Dating of the ice filtered from the Mt. Ortles cores with the 14C method
employing WIOC provides evidence for a bottom ice age of
6.7 ± 0.4 kyrs cal BP. The determination of 14C in a larch leaf
in core no. 1 by traditional methods provides an absolute and accurate
timeline (2.6 ± 0.1 kyrs cal BP) that is chronologically consistent
with the 14C ages of the ice samples (Fig. 11). While the exceptional
thinning of the ice layer thicknesses below 58 m is difficult to explain,
the monotonicity of the chronological empirical curve (Fig. 11) and of its
derivative, together with the gradual change of the red index between
∼ 60 and ∼ 70 m (Fig. 7), suggests that the thinning process
operated consistently, at least at a millennial timescale. Thus, the overall
ice core stratigraphy is likely continuous over millennia.
Nevertheless, a still unexplained physical process of thinning, or
alternatively an unrecognized stratigraphic centennial-scale hiatus between
58 and 68 m of depth, must have taken place. Similar situations might also
have been observed in other non-polar-latitude and high-altitude glaciers where
the ice core records obtained were considered to be continuous (Thompson et
al., 1995). We note that continuity of the ice core records is relative to
the timescale considered as, by their own nature, ice cores are constituted
by wet deposition occurring intermittently at different timescales
(e.g. meteorological (snow events),
seasonal (wet–dry seasons), decadal (droughts caused by recurring patterns of
ocean-atmosphere climate variability etc.)), and thus stratigraphic hiatuses
are the norm rather than the exception. However, despite the inherent
hiatuses, when considering the appropriate timescale (millennial in our
case), continuity of the ice core records can be assumed.
The preservation of early Holocene ice in the Alto dell'Ortles glacier is
probably due to at least two factors. First, the location of the drill site
in a rain shadow and its exposure to wind scour result in relatively low snow
accumulation. The medium-term snow accumulation at the drilling site was
estimated to be 850 mm w.e. yr-1 between 1963 and 2011. This latter value
was obtained by applying the Nye model to correct for firn compaction and ice
thinning down to the well-dated 1963 radioactive peak at 41 m of depth; this
is consistent with the current 2011–2013 observations of
∼ 800 mm w.e. yr-1. Second, frozen bedrock (current value,
-2.8 ∘C) did not allow basal melting and a significant basal flow
(see next section), thus preserving the oldest bottom ice in situ.
Last glacial age ice has been observed at several high-altitude–low-latitude
drill sites such as Huascarán (Peruvian Andes) (Thompson et al., 1995),
Sajama and Illimani (Bolivian Andes) (Thompson et al., 1998; Ramirez et al.,
2003) and Guliya (Western Tibetan Plateau) (Thompson et al., 1997). In
contrast, ice core records from the western Alps typically extend over just a
few centuries (Barbante et al., 2004; Preunkert et al., 2001; Schwikowski et
al., 1999a; Van de Velde et al., 2000b). However, the ice embedded in the
deepest layers of Alto dell'Ortles dates to the demise of the Northern
Hemisphere Climate Optimum (NHCO), and is among the oldest ice discovered in
the European Alps, exceeded in age only by the ice more than 10 kyrs old
retrieved at Colle Gnifetti (Jenk et al., 2009). We surmise that Last Glacial
Maximum ice does not exist in the Mt. Ortles record because of the lack of
the stable isotopic depletion (4 to 5 ‰) characteristic of such ice in
the bottom of Alto dell'Ortles (Fig. 10). Bottom ice from several other low-latitude–high-altitude drill sites such as Dasuopu (Himalaya) (Thompson et
al., 2000) and Tsambagarav (Western Mongolia) (Herren et al., 2013) behaves
in the same way, and extends back several millennia to the mid-Holocene.
Dynamic of the bottom ice
When considering the location of the drill site in the uppermost part of the
glacier Alto dell'Ortles, the current glacier flow (3.2 m yr-1 at the
surface, ≥ 0 m yr-1 at the base) and the age of the Mt. Ortles
ice cores (6.7 kyrs cal BP at the bottom), one may wonder why such old ice
deposited on the summit of Alto dell'Ortles was not quickly removed during
the Holocene. Here we demonstrate that the only possible answer is that the
observed significant ice flow must be a very recent phenomenon, and that
consequently a much slower basal flow has typically been common since the NHCO.
To study the origin of the Mt. Ortles bottom ice quantitatively, and to verify
its consistency with the local geography (e.g. the core layers cannot
originate from a location beyond the margins of the glacier), we have
employed a simple bi-dimensional dynamic model that estimates the
lower limit of the distance covered by a single glacier layer over
time, under the null hypothesis of an unchanged past dynamic of the glacier
from the current conditions. We have therefore assumed a linear variation of the glacier velocity
with depth (Vx) (as determined by means of current
inclinometric measurements) between a negligible value (0 m yr-1 at
the bedrock, current lower limit value) and 2.6 m yr-1 (current lower
value recorded at the surface by GPS). For modelling the vertical glacier
flow (Vz) we have employed two approaches: (i) conservative use of a Nye
model (not shown), and (ii) more realistically, a linear combination of
exponential functions that interpolate the empirical chronological timelines
obtained along the depth profile (Fig. 12).
The results are generally consistent with our ice core sections originating
uphill from the drill site along the flow line (Fig. 3). For instance, the
1963 radioactive ice layer (at 41 m) would originate from a location that is
at least 90 m uphill of the drill site (using both the conservative and
realistic Vz). However, several thousand-year-old ice (below 70 m of depth)
would originate from an unrealistic minimum distance that is 300–500 m
uphill (using the conservative and realistic Vz, respectively), thus much
larger than the distance between the drill site and the origin of the flow
lines (Fig. 3). As this result contradicts the null hypothesis of unchanged
dynamic conditions over time, we conclude that the formation and preservation
of very old ice at the Alto dell'Ortles drill site was possible, only when the
flow velocity of the ice layers near the base was much lower throughout the
time since the formation, when compared to the values measured today. In
other words, the significantly positive glacier flow Vx near the base, which
was recorded in the present day, must be the result of a very recent change in
the ice dynamics at the drill site.
We interpret this result as being possibly indicative of a large-scale dynamic
change, probably involving the entire Alto dell'Ortles. However, this dynamic
variation is likely not caused by changes in the slope of this glacier that,
according to our large-scale comparisons of the DTMs employed (not shown),
seem negligible, even under the action of the strong ablation occurring since
1980, especially in terms of calving at the lowermost margins of Alto
dell'Ortles. Instead, we speculate that this dynamic variation may be a
consequence of two possible alternatives or concomitant factors: (i) recent
summer meltwater influx from the bed outcrops uphill from the basal portion
of the drilling site may be lubricating the glacier/bedrock interface. As
shown for cold-based ice caps and ice sheets in Svalbard and Greenland, there
is evidence that seasonal meltwater can reach bedrock and change the ice
velocity (Bartholomew et al., 2010; Dunse et al., 2015). This would be
consistent with the observed seasonal changes in the surface velocity, and
with the quasi-linear profile of Vx obtained with the inclinometer
(Fig. 4). (ii) Changes in the plastic behaviour of the cold portion of the glacier as a
consequence of the ongoing thermic transition from polythermal to temperate
conditions may play a role. This latter hypothesis would be consistent with (a) a negligible
basal flow, (b) long-term changes in the vertical thermal profile (perhaps
particularly significant since the end of the LIA) and (c) the observed
significant elevation changes of the drilling site during the last century.
Horizontal displacement of a hypothetical glacier layer sinking
from the surface to bedrock according to the vertical dynamic that follows
the empirical age–depth relationship developed in this work.
The Alto dell'Ortles glaciation during the Holocene
A marked stable isotopic enrichment observed in the speleothem record from
Spannagel Cave (2500 m, Austria) during the NHCO (Vollweiler et al., 2006)
suggests that this period was likely the warmest during the entire Holocene
in the sector of the Alps where Alto dell'Ortles is located. This is
consistent with the minimum thickness of the Upper Grindelwald Glacier in
Switzerland between 9.2 and 6.8 kyrs BP (Luetscher et al., 2011) and with
the minimum extent of the Tschierva Glacier in the adjacent Mt. Bernina group
(Joerin et al., 2008), which indicates that the equilibrium line altitude
(ELA) was 220 ± 20 m higher than the 1985 reference level (2820 m).
Today the estimated ELA on Alto dell'Ortles lies at about 3300 m as inferred
from in situ observations and from the ELA obtained from the nearby glacier
of Vedretta della Mare (Carturan, 2016). Under the current warm climatic
conditions the drilling site (3859 m) is polythermal, characterized by
temperate firn and cold basal ice. Similarly to the adjacent Mt. Bernina
group, the ELA was almost certainly higher than today on Alto dell'Ortles
during the NHCO, and thus it is certainly possible that this glacier was
entirely under a temperate regime throughout its thickness. This would imply
the occurrence of basal melting/sliding and thinning of the glacier and the
quick removal of bottom ice during the NHCO. This mechanism might explain the
absence of ice older than 6.7 kyrs BP at the drill site. Remarkably, a
chironomid-based air temperature reconstruction from another nearby site from
Schwarzsee ob Sölden (2796 m, Austria) suggests that it was
∼ 2 ∘C warmer than today during the 7.9–4.5 kyr BP period
(Ilyashuk et al., 2011). This finding, when combined with recent field
evidence, supports the hypothesis of temperate conditions during the NHCO,
and indicates the possibility of a complete deglaciation of the drilling site
during the warmest phase of the NHCO.
At the end of the NHCO temperatures started to decrease (Vollweiler et al.,
2006), probably causing a reversal of the thermal regime of Alto dell'Ortles
from temperate to polythermal, and thus allowing the accumulation of cold ice
on frozen bedrock. As inferred from European palaeolake levels, a short
increase in precipitation at ∼ 7 kyrs BP (Magny, 2004) could also
have contributed higher accumulation and ice thickening recorded/inferred at
∼ 6.8 kyrs BP for small and climatically sensitive glaciers such as
Upper Grindelwald (Luetscher et al., 2011) and Alto dell'Ortles (this work).
Although high precipitation did not persist during the mid-Holocene,
progressively more favourable glacial conditions characterized the eastern Alps at the end of the NHCO. While there was at least a warm spell
(4.4–4.2 kyrs BP) in this area (Baroni and Orombelli, 1996), glaciers
extended in general to lower elevations, including the Tisenjoch (3210 m)
where the Tyrolean Iceman was buried in snow and ice since 5.3 kyrs BP.
Today, due to strong atmospheric summer warming, Alto dell'Ortles is
transitioning back from a polythermal to a temperate state. Basal sliding
conditions that could have occurred only during the NHCO are likely to be
soon, or are already, fully restored, with important and immediate
consequences for the dynamic of the entire glacier.