Haa2001b

Haa2001b
Journal of Glaciology, Vol. 47, No. 159, 2001
Surface properties and processes of perennial Antarctic
sea ice in summer
Christian Haas,1 David N. Thomas,2 Jo«rg Bareiss3
1
Alfred Wegener Institute for Polar and Marine Research, Columbusstrasse, P.O. Box 120161, D-27568 Bremerhaven, Germany
2
School of Ocean Sciences, University of Wales Bangor, Menai Bridge, Anglesey LL59 5EY,Wales
3
Department of Climatology, Faculty of Geography/Geosciences, University of Trier, D-54286 Trier, Germany
ABSTRACT. Ice-core and snow data from the Amundsen, Bellingshausen and Weddell
Seas, Antarctica, show that the formation of superimposed ice and the development of seawater-filled gap layers with high algal standing stocks is typical of the perennial sea ice in summer. The coarse-grained and dense snow had salinities mostly below 0.1ù. A layer of fresh
superimposed ice had a mean thickness of 0.04^0.12 m. Gap layers 0.04^0.08 m thick extended
downwards from 0.02 to 0.14 m below the water level. These gaps were populated by diatom
standing stocks up to 439 mg L^1 chlorophyll a. We propose a comprehensive heuristic model of
summer processes, where warming and the reversal of temperature gradients cause major
transformations in snow and ice properties. The warming also causes the reopening of incompletely frozen slush layers caused by flood^freeze cycles during winter. Alternatively, superimposed ice forms at the cold interface between snow and slush in the case of flooding with
negative freeboard. Combined, these explain the initial formation of gap layers by abiotic
means alone. The upward growth of superimposed ice above the water level competes with a
steady submergence of floes due to bottom and internal melting and accumulation of snow.
1. INTRODUCTION
Sea ice in the Southern Ocean extends from a summer minimum of about 46106 km2 in March to a winter maximum
of about 206106 km2 in October (Gloersen and others,
1992). This strong seasonal variability has impacts on processes and interactions in the ocean, atmosphere and the
marine biosphere. The rapid sea-ice formation in winter is
well understood, and occasionally referred to as the ``pancake cycle'' (Lange and others,1989): in the turbulent supercooled waters of the marginal sea-ice zone, frazil ice forms
and subsequently consolidates into pancake ice. The pancakes grow further by rafting and ad-freezing, and aggregate into larger ice floes. The frazil origin of most sea ice in
Antarctica is reflected in its fine orbicular granular crystal
texture (Eicken and Lange, 1989), typical of the majority of
ice cores drilled in the Weddell Sea (Gow and others, 1987;
Drinkwater and Haas, 1994; Eicken, 1998), the Bellingshausen, Amundsen and Ross Seas (Jeffries and Weeks,1993;
Jeffries and others,1997) and in East Antarctica (Worby and
others, 1998). Once the ice cover is closed and waves are
absent, freezing takes place in open leads or underneath
the floes primarily by congelation. These ice types are
generally marked by a columnar crystal texture.
A further notable and widespread process of Antarctic
sea-ice formation is the freezing of snow slush at the snow/ice
interface after surface flooding. A thick snow cover depresses
the ice surface below the water level, such that sea water floods
the ice surface and subsequently refreezes together with the
water-saturated snow. This process is referred to as ``flood^
freeze cycling'' (Fritsen and others,1998). The resulting ``snow
ice'' has a fine granular texture, too, and can only be distinguished from granular ice of frazil origin by its negative 18O
due to the large meteoric snow fraction (Lange and others,
1990). The contribution of snow ice to the total sea-ice volume
may be as large as 25% (Eicken and others,1994; Jeffries and
others,1997).
The processes described above are most important in
the cold season from early fall to late spring, when air temperatures are below the freezing temperature of sea water.
In summer, temperature gradients in the ice and snow are
frequently zero or even positive. Consequently, most ice deteriorates by bottom and lateral melting due to heating from
above and below (Martinson and Iannuzzi, 1998). Before
melting completely, the rotten ice becomes extremely porous and develops a honeycomb-like structure. Nevertheless,
some sea ice survives the Antarctic summer season, but little
is known about its surface processes and properties.
Casually, ``superimposed ice'' has been observed at the
snow/ice interface of drifting pack ice (Ackley and others,
1979; Gow and others,1987; Jeffries and others,1994; Eicken,
1998), and frequently on fast ice (Kawamura and others,
1993, 1997). This has a large-grained, polygonal granular
(PG) texture, and forms as a result of snow metamorphosis
during ``melt^freeze cycling'' (Colbeck,1997). Single crystals
can have diameters of several centimetres, as a result of
rapid grain coarsening in wet snow (Colbeck,1987).
Occasionally, an almost continuous ``gap''or high-porosity
layer is found at, or just below, the water level, overlain by a
thin solid layer of ice. These gap layers host active biological
assemblages, providing optimum conditions of light and nutrient replenishment (Kottmeier and Sullivan, 1990; Garrison
and Buck, 1991; Ackley and Sullivan, 1994; Fritsen and others,
1994; Lytle and Ackley, 1996). Haas (1998) has inferred that
this gap layer is widespread in the Bellingshausen and
Amundsen Seas, because the apparent electrical ice conduc613
Journal of Glaciology
Fig. 1. Photograph of a tilted floe broken by the ship, showing
the surface layer, gap and rotten ice structure typically found in
summer.Total thickness of the floe is about 1.00 m.
tivity was significantly higher than in other seasons as a result
of sea water filling the gaps.
Ackley and others (1979) and Ackley and Sullivan (1994)
explain the gap layers, also referred to as ``freeboard layers''
(Ackley and Sullivan, 1994; Fritsen and others, 1994), as a
result of a biophysical feedback process at the onset of summer warming. Due to increased temperatures and the seasonal rise in solar radiation, brine drainage from the ice
surface triggers a salinity and nutrient flux to the layer
below the surface. There, an algal bloom occurs, increasing
radiation absorption due to the dark colour of the algae,
leading to further melt and the development of a rotten ice
layer. It should be noted here that this mechanism requires
thin snow in order to allow for significant radiation transmission, and needs rather high algal standing stocks (Zeebe
and others,1996), as well as a positive ice freeboard.
Here, we report results from comprehensive physical
and biological investigations of ice and snow properties of
perennial ice in summer 1994 in the Bellingshausen and
Amundsen Seas, and in summer 1997 in the Weddell Sea.
During both studies, large amounts of superimposed ice as
well as numerous surface gap layers were observed. Figure 1
shows a photograph of a floe fragment tilted upright by an
ice-breaker. The rotten sea ice with a honeycomb-like upper
layer, a surface gap and a solid layer of mostly superimposed
ice overlain by snow can be distinguished. We will particularly focus on the ice layer above the gap, which we sampled
along horizontal surface profiles with a high vertical resolution. Based on our data, we propose a heuristic model of
the various processes at the surface of Antarctic sea ice from
the onset of summer melt until the beginning of fall freezeup. These could provide a link between summer properties
and features described in previous studies, as well as alternative explanations for the gap-layer formation. The processes also have profound consequences for remote-sensing
studies of Antarctic sea ice in summer.
2. MATERIAL AND METHODS
Snow-property measurements and ice coring were performed
during two expeditions of RV Polarstern to the Bellingshausen
and Amundsen Seas in February 1994 (Haas and Viehoff,
1994; Haas and others, 1996) and to the Weddell Sea in January and February 1997 (Haas and others, 1998). En route,
2 hourly visual ice observations were performed and docu614
Fig. 2. Map of the research areas in the Amundsen and Bellingshausen Seas and Weddell Sea in 1994 and 1997, respectively,
showing the locations of the sampled floes (circles). Open circles
indicate floes with PG ice at the snow/ice interface.The mean
February ice concentration derived from satellite passive-microwave Special Sensor Microwave/Imager data (provided on
CDs by the U.S. National Snow and Ice Data Center, Boulder,
Colorado) is also shown for the respective year.
mented as proposed by Worby and Allison (1999). The
locations of the sampled floes, as well as the mean ice coverage
in February 1994 and 1997, are shown in Figure 2. In 1997 the
ship crossed the Weddell Sea twice along a southeast^northwest transect, covering a wide range of ice regimes. Most floes
were visited by helicopter while the ship was underway. In
total, 33 floes were sampled in 1997, and 15 in 1994.
On almost every floe, electromagnetic (EM) ice-thickness measurements were carried out using a Geonics EM31
with 5 m spacing along profiles of 50 to >1000 m length
(Haas and others, 1997; Haas, 1998). In 1994, most EM
soundings were accompanied by drillhole measurements
for validation and to determine the ice freeboard. In 1997,
only a few drillhole measurements were carried out and
the freeboard was surveyed with a laser levelling device
along the EM profiles.
2.1. Snow measurements
Snow thickness was measured with a spacing of 1m on
representative 50^100 m subsections along the ice-thickness
profiles with the aid of a ruler stick.Vertical profiles of snow
temperature, wetness, density, salinity and grain-size were
measured in up to three snow pits. Both thin snow on level
ice and thick snow around pressure ridges were sampled.
Snow temperature was measured with a Pt-100 thermometer at vertical spacings of 0.05^0.10 m, depending on snow
thickness. In parallel, snow wetness was measured with a
TOIKKA Snow Fork (Helsinki, Finland). Snow wetness in
percentage liquid-water content is calculated from fork
resonator measurements of the complex snow permittivity
(Shivola and Tiuri,1986).
After visual examination of the snow stratigraphy,
samples were taken from the main layers with 0.05 m diameter, 0.5 L steel tubes. In the cold laboratory on board the
ship, snow density was measured by weighing these samples.
After visual grain-size estimation using a millimetre grid,
the samples were melted and subsequently their salinity was
measured with a Wissenschaflich-Technische-Werksta«tten
GmbH LF181 conductivity meter which was calibrated to
Haas and others: Summer processes of perennial Antarctic sea ice
display salinity. If the salinity was <0.1ù, the instrument
gave a reading of 0.0ù.
2.2. Ice coring and analysis
From each floe, two 0.09 m diameter ice cores were drilled
from level ice with a gas-motor-driven Kovacs Enterprise ice
auger after the snow had carefully been removed. Such cores
are referred to as ``main cores''. In 1997, the top sections of the
cores were drilled mostly by hand until the gap layer was
encountered (see below). This ``surface core'' was retained
from the core tube without turning it upside down, to avoid
contamination of the ice with sea water and brine. After
recovery, the surface cores were stored upright in pre-cooled
cold boxes to minimize melting, brine drainage and contamination during the helicopter transport to the ship's cold
laboratory (<^25³C). Then cores of the remaining ice were
taken and stored horizontally in plastic tubes for transfer to
the ship.
The second core was immediately cut into 0.1m sections
and allowed to melt slowly for later chlorophyll-a measurements by standard fluorometric procedures (Thomas and
others,1998).
When surface gaps were encountered at the main coring
site in 1997, up to four additional surface cores were drilled a
few metres apart to determine the extent and abundance of
these features. The thicknesses of the ice layers, the gap
layers, as well as the freeboard height were measured. Water
from the gaps was sampled for salinity. Ice crystals and slush
were sieved out using a mesh size of 1mm2.
In the cold laboratory, detailed analyses of ice texture
were performed on thick and thin sections of the cores (e.g.
Lange, 1988). The cores were cut according to stratigraphic
units. The surface cores were cut with a high vertical resolution of 0.01^0.03 m.While two-thirds of the segments were
subsequently melted for salinity measurements, the rest was
transferred to the home laboratory. On a limited number of
these samples, 18O concentrations were determined as
described by Eicken and others (1994) and Eicken (1998).
3. RESULTS
3.1. General ice conditions
In February 1997, the ice extent in the Weddell Sea was rather
high (Fig. 2), with highly deformed and thick ice in the southeastWeddell Sea (Haas and others,1999; Table 2, shown later).
The sampled floes were located across the entire Weddell Sea
(Fig. 2), and were visited during a period of 42 days under
varying climatic conditions (Fig. 3). Air temperatures measured on board Polarstern ranged between 0³ and ^5³C until
day 55, when values dropped below ^10³C when Polarstern
transited a second time to the southeast, where minimum
temperatures were as low as ^22³C. However, at that location
a slight cooling had already commenced by day 32, according
to data logged by a drifting buoy in that region (Fig. 3b). The
rather cold conditions started on day 49, about 7 days before
the arrival of Polarstern. To distinguish between different
climatic and ice regimes, the 1997 dataset is subdivided into
three periods defined as summer (days 19^27) and early fall
(days 56^61) in the southeast, and as summer in the northwest
Weddell Sea (days 46^55; Figs 2 and 3). The 40³ W meridian
is taken as the boundary between the southeast and northwest
Weddell Sea.
Fig. 3. Air temperatures recorded on board Polarstern along the
cruise track in 1997 (a) and by a drifting buoy (Argos ID 8059)
at about 40³ W, 73.2³ S in the southeast Weddell Sea (b).
In February 1994 in the Bellingshausen and Amundsen
Seas, air temperatures ranged mostly between 0³ and ^4³C.
Only close to the coast were lower temperatures (^5³ to
^8³C) encountered. With respect to ice thickness (Haas,
1998; Table 2, shown later), we distinguish between the
Bellingshausen Sea and Amundsen Sea.
In late January 1997 (days 19^27) the ship operated in
closed pack ice in the southeast Weddell Sea, with ice concentrations mostly higher than 0.9 and floe diameters up to
500 m. During the second phase of the expedition in the inner
marginal ice zone floe sizes exceeded 100 m only on days 53
and 54 as well as 57 and 58, while ice concentration varied
between 0.5 and 1.0. The ice in the northwest Weddell Sea
was characterized by a mixture of thick deformed and thin
level floes, indicating the presence of old and young ice,
respectively (Lange and Eicken,1991). The latter carried only
a thin snow cover and their surfaces appeared greyish. In
1994, most floes were located well within the perennial ice
zone (Haas and Viehoff, 1994; Morris and others, 1998),
although their size was mostly <100 m.
Numerous surface ponds were observed throughout the
southeast and northwestWeddell Sea, mostly originating from
ridge loading (Ackley and Sullivan, 1994). However, aerial
photographs (C. Haas, unpublished information) showed
that they covered only 0.3% of the ice surface. The mean salinity of 19 pond-water samples was 19.4 7.1 (range 8.3^30.9),
indicating a mixture of sea water and melted snow and ice.
3.2. Snow properties
Table 1 summarizes the average snow thicknesses in each
region calculated from the thickness-profile data of each
floe. The snow layer was generally much thicker in the
Bellingshausen and Amundsen Seas than in the Weddell
Sea, where the thinnest snow was observed in the northwest.
There, in contrast, winter observations of Eicken and others
(1994) and Massom and others (1997) report the highest
snow thicknesses of the Weddell Sea. Clearly, our findings
indicate strong snowmelt in the northwest during summer,
which will be discussed later in more detail.
Figure 4 shows typical vertical profiles of snow tempera615
Journal of Glaciology
Fig. 4. Typical vertical snow profiles of (a) temperature
obtained on level ice and (b) wetness (percentage liquid
water by volume) obtained on level ice and ridge flanks in
three different regions in 1997. Note different scale in upper
right graph.
ture and wetness. The differences in air temperatures (Fig.
3) were clearly reflected in the snow (Fig. 4a). While temperatures in the southeast Weddell Sea were higher at the
surface than at the snow/ice interface (positive gradients),
positive as well as negative gradients were found in the
northwest Weddell Sea. In both summer situations, temperatures at the snow/ice interface were close to or above
the freezing point of sea water (^1.8³C). In contrast, strong
negative gradients ranging from ^8³ to ^66³C m^1 were
observed in the southeast Weddell Sea in early fall, 75% of
these being less than ^25³C m^1. Depth hoar was found in
every snow pit. The ice surface had cooled below ^5³C at
several sites, and dendritic ice growth underneath the surface layer into the gap layer was observed (see below).
Figure 4b shows that snow wetness seldom exceeded 4%
This means that during most of the sampling period the snow
was in the pendular regime where liquid water cannot freely
drain but is held at the triple junctions of ice grains and air.
This regime promotes snow coarsening by the growth of
grains into clusters. Snow wetness increased with depth on
most floes and was highest at the bottom (Fig. 4b). This was
most pronounced in the southeast Weddell Sea in fall.
Of all 169 salinity samples, only 31 had a salinity equal to
or higher than 0.1ù (Table 1). Only three samples which
were obtained from the snow/ice interface had salinities
higher than 1ù in 1997. These summer values are significantly lower than measurements in other seasons, where
mean snow salinities range between 1ù and 16ù, with
more than 10% of all measurements ranging from 10ù up
to 40ù (Eicken and others, 1994; Massom and others, 1997;
Sturm and others,1998; Worby and others,1998).
Snow density ranged between 130 and 518 kg m^3, with
means of 300^400 kg m^3. Grain-sizes ranged between 1 and
10 mm, with mean values of 2.5^3.1mm (Table 1). Apart from
the depth hoar observed in the southeast Weddell Sea in fall,
and fine new snow that mostly melted within 1 or 2 days, the
snowpack consisted of coarse and very coarse grain clusters
(seasonal-snow-on-the-ground classification 6cl/6mf; Colbeck
and others, 1990). The snowpack was very icy at the bottom.
Just above the snow/ice interface, grain-size increased rapidly
and the grains were partially frozen onto the underlying ice.
Therefore, the snow/ice interface was very rough on the centimetre scale. As the ice underneath was mostly superimposed
ice (see below), the rough surface was a transition zone of
snow being metamorphosed into ice.
Observed grain-sizes and densities are similar to values
reported by others for snow which was subject to melting
phases. However, they are much higher than values of cold
Antarctic winter snow (Eicken and others, 1994; Massom and
others,1997; Sturm and others,1998; Worby and others,1998).
3.3. Ice properties
Table 2 summarizes all EM and drillhole thickness measurements. Note that these include both level and deformed ice,
the fraction of the latter being rather high (Haas and others,
1999). The thickness of undeformed, level ice is 0.4^1.3 m (cf.
Figs 5 and 6). However, mean thicknesses in the southeast
Weddell Sea were much greater than those recorded in previously published winter data (Eicken and others, 1994;
Massom and others, 1997), and are much closer to values
reported by Strass and Fahrbach (1998). In the northwest
Weddell Sea, thicknesses were comparable to those measured
for undeformed second-year ice in spring in the same region
(Lange and Eicken, 1991). The ice in the Bellingshausen and
Amundsen Seas is some of the thickest Antarctic sea ice
reported (Haas,1998).
Table 1. Mean snow properties
Southeast Weddell (summer)
Northwest Weddell (summer)
Southeast Weddell (early fall)
Bellingshausen Sea
Amundsen Sea
0.32 0.16
0.16^0.56
9
0.51 0.11
0.39 0.75
10
Snow thickness (m)
Range
N (floes)
0.24 0.02
0.21^0.26
5
0.11 0.14
0.04^0.58
14
0.20 0.12
0.07^0.42
10
Salinity (ù)
Range
N (samples)
0.04 0.09
0^0.2
5
0.17 1.00
0^8.6 (1 sample 41)
81
0.14 0.45
0^2.4 (2 samples 41)
46
0.02 0.08
0^0.36 (3 samples 4 0.1)
37
Density ( kg m3)
Range
N (samples)
ND
368 64
218^480
81
313 72
130^496
46
391 71
212-518
37
Grain-size (mm)
Range
N (samples)
ND
3.1 1.2
1^5.5
81
2.6 1.2
1^10.0
46
2.5 1.0
1^4.0
11
ND, not determined.
616
Haas and others: Summer processes of perennial Antarctic sea ice
Table 2. Mean ice properties
Total thickness (m)
Range
Freeboard
Range
Percentage of gap features
Surface ice layer thickness (m)
Range
Surface layer draft (m)
Range
Gap thickness (m)
Range
PG ice thickness (m)
Range
Texture classes (%)
Orbicular granular
Columnar
PG
Mixed/others
PG ice salinity (ù)
Southeast Weddell
(summer) 11 floes
Northwest Weddell
(summer) 14 floes
Southeast Weddell
(early fall) 8 floes
Bellingshausen Sea
6 floes
Amundsen Sea
9 floes
1.56 0.31
(n ˆ 5, EM)
1.24^1.91
0.07 0.05
0^0.14
73
0.07 0.04
0.02^0.15
0.02 0.02
0^0.07
0.08 0.09
0.01^0.27
0.04 0.03
0^0.09
1.52 0.75
(n ˆ14, EM)
0.26^2.86
0.08 0.05
0^0.17
86
0.20 0.06
0.10^0.29
0.14 0.05
0.08^0.23
0.04 0.03
0.01^0.09
0.08 0.06
0.01^0.25
1.92 1.39
(n ˆ 8, EM)
0.41^4.30
0.07 0.05
0^0.13
38
0.15 0.05
0.10^0.16
0.14 0.01
0.13^0.15
0.05 0.06
0.01^0.09
0.08 0.01
0.07^0.08
2.37 1.56
(n ˆ 6, drilling)
1.22^4.98
5.11 2.32
(n ˆ 5, drilling)
3.34^9.14
ND
ND
ND
ND
ND
ND
ND
ND
0.10 0.05
0.03^0.16
0.12 0.07
0.03^0.25
66
28
2
3
0.68 0.40
53
33
10
4
0.61 1.06
83
11
6
0
3.78 3.25
68
18
3
11
0.86 1.26
Notes: All listed variables except ice thickness were determined from surface or main ice cores. Texture classes derived from main ice cores only. Mean total
thickness was calculated from EM and drillhole profiles.
ND, not determined.
Figure 5 shows high-resolution vertical cross-sections
typical for floes encountered in the three ice regimes in the
Weddell Sea, obtained from thickness drilling. While the
floes in Figure 5a and c were rather level and much larger
than 100 m, the floe in Figure 5b was only 40 m in diameter
and comprised an old pressure ridge. All three profiles show
the typical sequence of a snow layer above a surface ice layer
overlying a slush- or sea-water-filled gap. Underneath, the
main ice was encountered, which was often extremely rotten
and porous (Figs 1 and 6b).
The frequency of occurrence, the thicknesses and drafts
of the surface ice layers, as well as the gap thicknesses, are
summarized in Table 2. In total, gap layers were observed
on 64% of all 33 sampled floes in 1997. The surface ice layer
in the southeast Weddell Sea was much thinner than in the
other regions, and its underside was at or just below sea level
(Fig. 5a; Table 2). This is similar to earlier summer observations of a crusty snow layer above the slush (Lytle and
Ackley, 1996). Two samples of gap water had salinities of
14ù and 17ù. In the other two regions, by contrast, the
thick surface ice layer extended on average 0.14 m into the
water, with most of its thickness not rising above the water
level. The mean salinity of 37 gap-water samples was
29.3 2.6ù (range 20.6^31.3ù). As with the pond water,
the low gap-water salinities indicate that melting was taking
place within or adjacent to the gaps.
The lower portion of Figure 5a^c shows vertical profiles
of texture and salinity of the main and some surface cores, as
well as four d18O profiles. All cores consisted of superimposed
PG ice at the top. The typical PG ice texture with isometric
grains with planar boundaries can be seen in the thin- and
thick-section photographs in Figure 6. Grain-size often
decreased with depth. In the thick-section photographs (Fig.
6b and c), PG ice is less clearly recognizable once the grainsize becomes smaller than the sample thickness. Therefore,
fractions of PG ice given below are probably underestimates
of the real amount. The downward-decreasing grain-sizes
also make it difficult to distinguish it from underlying finegrained ice with higher salinities occurring in some instances
(Figs 5 and 6). That ice is probably of snow-ice origin.
PG ice contributed significantly to the overall surface
core lengths (see, e.g., Fig. 5b). It was found at the top of
76% of the 33 main cores taken in 1997, and of 73% of the 15
cores taken in 1994 (Fig. 2). Table 2 summarizes the mean PG
ice-layer thicknesses. The smallest values occurred in the
southeast Weddell Sea in summer; they doubled there by
early fall. The greatest thicknesses were observed in the
Bellingshausen and Amundsen Seas. The amount of PG ice
growth during summer is roughly comparable with fractions
of snow ice formed during flood^freeze cycles in the cold
season (Eicken and others,1994; Jeffries and others,1997).
Figures 5 and 6 also show typical vertical high-resolution salinity and d18O profiles of the respective cores. Salinities of all surface and main cores are presented in Figure
7.They ranged between 3 and 4 for the main ice underneath
the gap. In contrast, salinities were very low in the surface
cores, and mostly zero for PG ice (Table 2). Only immediately above the gap were high salinities sometimes measured, coinciding with a fine-grained granular texture.
Additionally, in some locations in the southeast Weddell
Sea in fall, PG ice salinities rose at the top (see also Table
2), when the freeboard was less than about 0.01m, or where
the gap had started to refreeze.Then, the surface cores often
revealed a C-shaped salinity profile (e.g. at 4 m in Fig. 5b),
comparable to the salinity distribution of young nilas ice in
winter. Apparently, ice formation within the gap leads to
some upward brine expulsion.
The d18O values of the surface ice layers were extremely
small (Figs 5 and 7). At the very top, they were more like values
for snow, ranging between ^3ù and ^28ù in the Weddell Sea
in winter (Eicken and others, 1994). With increasing depth,
they slowly became more typical of sea ice, at 0^2.5ù, indicating the presence of snow ice (Eicken and others, 1994; Jeffries
and others,1997).
617
Journal of Glaciology
Fig. 5.Top:Vertical cross-sections typical for floes encountered
in the southeast and northwest Weddell Sea in summer (a, b)
and in the southeast Weddell Sea in fall (c), obtained from
thickness drilling (see (a) for legend). Bottom: Salinity,
d18O and texture for typical surface and main cores drilled
along the profiles above at locations indicated by thick vertical
lines. Data are plotted at centre position of individual core
segments with respect to the water level (z ˆ0.0 m).
Two other common observations are worth mentioning
(Fig. 6): First, PG ice was often underlain by a thin layer of
columnar ice, either immediately below, as in Figure 6a, or
further below, as in the case of dendritic new ice growing into
the underlying gap (Fig.6c).This columnar ice layer indicates
the former presence of liquid water, in contrast to the finegrained granular snow ice forming from slush during flood^
freeze cycles. Secondly, in the southeast Weddell Sea in fall,
surface ice with small freeboard often had a C-shaped salinity profile (Fig. 6b). Below the surface of this ice, salt-water
droplets were observed within the fresh superimposed ice,
indicating downward melting of sea water which had flooded
the ice surface.The salt water caused the salinity increase at the
618
top, and destroyed the PG texture of the ice, making the PG
origin of the ice difficult to recognize. This observation was
common on very rotten, thin ice, where the main sea ice underneath had melted almost completely and the floes consisted
only of the surface ice layer (Fig. 6b). These features indicate
the final stage of an ice floe in late summer, when it would either
soon deteriorate completely or become restrengthened once
subject to fall freezing. The rotten, highly porous structure of
the ice underneath the gap (Fig. 6b) permits large convective
transport of nutrients, salt and heat at the onset of fall freeze-up
(Fritsen and others,1994; Lytle and Ackley,1996).
Finally, in Figures 5^7, note that PG and very low-salinity ice occurred below the water level. As this ice origi-
Haas and others: Summer processes of perennial Antarctic sea ice
Fig. 6. Photographs of a thin section (a:1994) and two thick sections (b, c:1997) of three typical surface cores, taken with the
samples between crossed polarizers. In addition, (b) shows the rotten ice between the gap and the bottom, and arrows mark saltwater droplets penetrating into the superimposed ice.The salinity profile of each core, plotted at the centre position of individual core
segments with respect to the water level (z ˆ0.0 m, y axis), as well as the texture interpretation are also shown.
Fig. 7. Salinity and d18O profiles of all surface and main cores obtained in 1994 and 1997 plotted at centre position of individual
core segments with respect to the water level (z ˆ0.0 m).
nated from the refreezing and metamorphosis of snow
above sea level, it must have moved downward relative to
the water level during summer.
Figure 8 presents chlorophyll-a concentrations of the
upper metre of all cores obtained in the Weddell Sea. Unfortunately, the averaging over 0.1m long segments does not
allow detailed study of surface concentrations. However, it
can be seen that chlorophyll-a concentrations were generally
highest at 0.1^0.3 m depth below the surface, i.e. in the ice
surrounding the gap. This demonstrates the importance of
the surface-layer/gap structure for the primary productivity
in perennially ice-covered regions. The highest concentrations of up to 439 g L^1 were found in the dendritic new ice
forming from refreezing gaps in the southeast Weddell Sea in
fall. Similar observations have been reported by Fritsen and
others (1994).
4. DISCUSSION
4.1. Snowmelt and refreezing
We have presented numerous snow and ice data obtained
from perennial Antarctic sea ice. Although the dataset is heterogeneous, several common features emerge which can be
considered typical for the summer situation. Most importantly, there is considerable surface melting in the Antarctic
619
Journal of Glaciology
are close to the freezing point of sea water. This results in
the refreezing of fresh meltwater further down at the snow/
ice interface, and finally in the formation of superimposed
ice with its PG texture. However, the measured low snow
wetnesses could also indicate that superimposed ice may
form just by growth and merging of grain clusters. The processes described have also been observed on Arctic fast ice
(Holt and Digby, 1985; Onstott, 1992; Barber and others,
1998), although melting is much stronger there and more
continuous. Therefore, the snow and successively the superimposed and underlying ice layers melt completely.
Generally, superimposed-ice formation should affect
every spot of an ice floe, including level ice and pressure
ridges. However, depending on the amplitude of melting
and on the amount of meltwater, the water will drain to the
lowest topographic parts of a floe where the thickest superimposed ice would then be formed. When melting is not
strong enough to produce freely draining water, as indicated
by our wetness measurements, lateral differences in the
amount of superimposed-ice formation might be caused by
heterogeneous snow properties. Clearly, the local variability
in snow thickness, stratigraphy, grain types and iciness
(Massom and others, 1997; Sturm and others, 1998) will
result in different light penetration and scattering, causing
variable intensities of internal melt.
Fig. 8. Chlorophyll-a concentrations of ice cores obtained in 1997,
plotted at centre depth of individual core segments with respect to
the ice surface. For clarity, data from below 1m were omitted.
pack ice. The amount of melt is much less than in the Arctic,
where the snow melts completely and the ice surface is
extensively covered with melt ponds. Nonetheless, internal
snow melting on Antarctic sea ice is strong enough to induce
extensive snow metamorphism and the formation of superimposed ice. In the northwest Weddell Sea, where the thickest snow occurs in winter, melting was strongest and resulted
in both the thinnest residual snow cover and the thickest surface ice layer. Average superimposed ice thicknesses of 0.08^
0.25 m correspond to late-winter snow thicknesses of 0.24^
0.75 m if a snow density of 330 kg m^3 is assumed. The low
snow salinities show that there was considerable meltwater
percolation through the snow, transporting salt downwards.
When measured, however, the snow was in the pendular
regime. Nevertheless, there must have been episodic events
of enhanced surface melt and percolation, for example
during periods of warm-air advection (Massom and others,
1997; Morris and others, 1998). Snowmelt and wetness can
also be much higher earlier in the season, when the amount
of incoming solar radiation is highest. Very likely, wetness
varies diurnally with changing solar elevation, resulting in
melt^freeze cycling (Fig. 3b).
Although surface melting was regarded as atypical in
earlier studies (Andreas and Ackley,1982; Jeffries and others,
1994), the regional and temporal distribution of our observations and those of Fritsen and others (1994) and Lytle and
Ackley (1996) suggests that it is actually quite common, not
only in the marginal ice zone, but also within the inner pack.
Melting occurs even at air temperatures below 0³C
through absorption of solar radiation within the snow
(Colbeck, 1989; Launiainen and Cheng, 1998). Temperature
gradients in a melting snowpack on sea ice are positive, i.e.
temperatures decrease towards the ice surface, where they
620
4.2. Gaps
The other common summer feature is the widespread
occurrence of slush- or water-filled porous layers or gaps
underneath a surface ice layer. As outlined in the introduction, their origin has been explained as the result of summer
warming and the subsequent increase in brine volume, which
fosters a biophysical feedback mechanism (Ackley and
others, 1979; Ackley and Sullivan, 1994). In fact, the high salinities near the top of first-year ice (Maykut and Untersteiner,
1971; Eicken,1992) will result in large brine volumes once the
ice warms up. However, the results of Zeebe and others (1996)
show that a positive feedback mechanism by absorption of
solar radiation would only be effective with rather high
chlorophyll-a concentrations in the gap of >150 mg chl a m^2
and a snow cover thinner than 0.05 m. Although our snow
and chlorophyll data suggest that absorption by algae might
become significant in the course of the summer, these conditions are unlikely to prevail at the onset of summer warming.
For example, Gu«nther and Dieckmann (1999) observed no
accelerated rise in nutrients or chlorophyll in the upper ice
layers until the end of December in a time-series study on
snow-covered fast ice.
However, the formation of gap layers can also be
explained by purely abiotic processes. First, the widespread
formation of snow-ice layers during flood^freeze cycles in
the cold season provides a mechanism to form layered structures with strong salinity gradients (Fritsen and others, 1998;
Maksym andJeffries, 2000). Although some of the saline brine
might be expelled downwards during flood^freeze events
because the underlying ice is warm and permeable (Maksym
and Jeffries, 2000), it is likely that some high-salinity layer
remains at the former slush/ice interface on completion of the
freeze cycle. These high saline layers would reopen as soon as
ambient temperatures rise. As outlined by Eicken and others
(1995) and Maksym and Jeffries (2000), due to energetic constraints, the slush layer will only incompletely congeal if
Haas and others: Summer processes of perennial Antarctic sea ice
flooding and freezing commence after early September.
Thus, there is a high probability that slush-filled gaps exist
already at the end of winter. Where these are present, superimposed ice will form on top of snow ice overlying a gap, as
shown in Figures 5 and 6.
Secondly, the surface-ice/gap structure might be formed
due to melting of the snow above a slush layer of flooded
snow. At negative freeboard, surface flooding is most probable and widespread in summer, as the ice is most permeable
and the warm snow allows for long-range lateral flow of the
sea water at the snow/ice interface. Initially, however, the
temperature of the saline slush layer will be close to the
freezing point of sea water. The less dense and fresh meltwater percolating through the overlying snow onto the cold
slush surface will therefore refreeze at the interface, and will
thicken by the addition and refreezing of more meltwater
from above. The resulting structure was often observed in
the southeastWeddell Sea in summer and is shown in Figure
5a. In this case, a snow-ice layer between the superimposed
ice and the gap was missing (Figs 5a and b and 6a).The situation is comparable to the formation of ice below under-ice
melt ponds in the Arctic (Eicken, 1994), where an ice layer
forms at the boundary between a fresh- and a salt-water
layer both at their freezing point.
Once the porous or gap layers have formed, the gaps are
widened both by absorption of solar radiation and by advection of sea water from the floe edges. Absorption might be
enhanced by high chlorophyll-a concentrations which could
develop in the gap habitat once it had formed. Slight
increases of the gap water temperature above its freezing
point lead to melting of the slush within it or of the adjacent
ice, as indicated by our gap-water salinity measurements.
While in this case the gap water is diluted and might still
be in phase equilibrium with its environment, advection of,
and exchange with, more saline sea water from the floe
edges leads to further melting, finally resulting in the rotten,
honeycomb-like ice structure. This will be particularly
important closer to the ice edge, where water motion within
the floes can be induced by swell penetrating into the ice.
Clearly, as also shown by nutrient data (e.g. Fritsen and
others,1994; Thomas and others,1998), at the end of summer
the permeability of the ice is high enough to allow for considerable exchange with the surrounding sea water. Another
source of energy available for melting results from the
downward flux of latent heat released by meltwater refreezing during superimposed-ice formation. As with the underice melt ponds, a small fraction of this heat is consumed by
melting at the very bottom of the surface ice layer.
4.3. Summer surface processes
Our measurements were made while moving through both
space and time. While the data from the southeast Weddell
Sea represent conditions at the beginning of summer, the
northwest Weddell Sea data are typical of the peak of the
ablation season. As indicated by low air temperatures and
the refreezing of gaps, measurements in the southeast Weddell
Sea in fall show the transition back to winter conditions. By
summarizing the observations in these regions, we suggest a
heuristic model of processes on perennial Antarctic ice from
late spring to early fall, shown in Figure 9.
The late-spring situation describes the preconditioning
of the ice surface before the onset of surface melt. Freeboard
is either positive (Fig. 9a) or negative. In the case of negative
freeboard, flood^freeze cycling with the formation of snow
ice may be present (Fig. 9c), or the snow may remain dry
because the underlying ice is impermeable and flooding
from the sides does not reach the interior of floes through
the cold snow (Fig. 9b). Either way, the salinity at the top
of the ice is high.
In summer, the ice and snow temperature gradients
reverse. As a consequence, the porosity at the top of the ice
increases drastically (Fig. 9a), incompletely frozen slush
layers widen (Fig. 9c), or the ice surface becomes widely
flooded (Fig. 9b). Simultaneously, snow metamorphoses and
meltwater percolates downward, and superimposed ice forms
above the water level, on top of either ice or slush. Generally,
the summer season is also characterized by desalination of
the original sea or snow ice below the gap. Among these processes, the ``freeboard layer'' mechanism (Ackley and others,
1979; Ackley and Sullivan,1994) is just one possibility related
to the porosity increase (Fig. 9a).
Finally, as air temperatures drop in fall, and snow and
ice temperature gradients become negative again, superimposed-ice formation and widening of gaps cease. Depending
on the amount of slush left in the gaps, they will congeal
from top to bottom, developing either a columnar or granular texture. This final freezing, with all the consequences of
high primary productivity and large vertical heat fluxes
(Fritsen and others,1994; Lytle and Ackley,1996), concludes
the summer season. Following continued accumulation of
snow, the ice will begin to undergo the well-known processes
of flood^freeze cycling.
4.4. Submergence
One more process complicates the transformations described
above (Fig. 9): over summer, the ice slowly submerges further
into the water due to changes in the isostatic equilibrium.
First, during summer the ice melts from below. Depending
on the amount of ocean heat flux and on the initial ice thickness, a thinning of 0.2^0.6 m can occur during summer
(Martinson and Iannuzzi, 1998). Secondly, mass is added as
either snow or rain to the ice surface. According to data presented by Eicken and others (1994), snow accumulation from
December to February amounts to 0.12^0.16 m at the continental coast of the southeast Weddell Sea, and is much higher
in the northwest. Finally, submergence is also due to internal
melt, leading to increased brine volumes and therefore higher
bulk densities. While density increases of 10^20 kg m^3 are
possible (based on calculations using the equations of Cox
and Weeks (1983)) for warming of the ice to close-to-melting
temperatures, even higher bulk density rises are likely for the
ice^water mixture of the very rotten, honeycomb-like ice.
The amount of submergence in relation to changes in
these variables can easily be deduced by differentiating the
equation for ice freeboard zfb
zfb ˆ zi
‰i …zi ‡ s †zs Š=water
…1†
for snow and ice thickness zs and zi, as well as for ice density
i . Characteristic values are given inTable 3. As can be seen,
the three variables in total could yield submergence distances of >0.1m during the summer months. Thus, drafts
of the surface ice layer of 0.14 m (Table 2) can be explained
by submergence alone, if the superimposed ice has formed
immediately above a slush layer. On the other hand, as
621
Fig. 9. Schematic drawing of summer surface processes described in the text.
Journal of Glaciology
622
Haas and others: Summer processes of perennial Antarctic sea ice
Table 3. Sensitivity of ice freeboard dzfb to changes in ice and
snow thickness dzi and dzs and in ice density di, as calculated
by differentiating Equation (1) for each respective variable
Ice thickness zi
Snow thickness zs
(i ˆ 920 kg m^3) (s ˆ 350 kg m^3)
dzi
dzfb
dzs
dzfb
m
m
m
m
^0.2
^0.4
^0.6
^0.02
^0.04
^0.06
0.1
0.15
0.2
^0.03
^0.05
^0.07
Ice density i
(zi ˆ 0.5 m)
(zi ˆ1 m)
di
dzfb
di
dzfb
kg m^3
20
30
40
m
^0.01
^0.01
^0.02
kg m^3
20
30
40
m
^0.02
^0.03
^0.04
Note: For all calculations, a water density of 1024 kg m3 has been assumed.
explained earlier, the gaps may form already at some depth
below the water level.
The data presented in Figures 5^7 show that the low surface ice salinities and d18O ratios are not changed much
when the ice submerges. However, once freeboard becomes
very small or even negative, as in Figures 5c and 6b, flooding might again take place, either causing a new cycle or
ending superimposed-ice formation completely at the end
of summer. Then, salinities at the very top are raised and
eventually salt-water droplets may melt themselves into the
superimposed ice, destroying the original PG texture (Fig.
6b). Thus, during summer there is a competition between
upward superimposed-ice formation and downward ice
submergence. Ultimately, the latter takes over towards the
end of the ablation season unless it is balanced by ice growth
at the ice underside.
Internal snowmelt and superimposed-ice formation will
take place both in the perennial and in the seasonal ice zone.
In the latter, however, submergence and surface flooding are
so strong that no significant amounts of superimposed ice
can form above the water level even before new flooding
takes place and the ice finally deteriorates.
5. CONCLUSIONS
We have presented ice-core and salinity data obtained from
perennial Antarctic sea ice during summer. The data show
the widespread occurrence of snow metamorphosis and
superimposed-ice formation on drifting pack ice. We have
also discussed the formation of sea-water-filled gaps typical
for summer sea ice. Both features are linked to each other
only indirectly, although both are generally related to a thick
snow cover. As a prerequisite, temperature gradients have to
reverse in the snow and ice compared to the cold season. Our
results reveal that submergence, well known to occur during
the cold season and manifested in the extensive formation of
snow ice, continues during summer. Thus, throughout the
year, the behaviour of sea ice in the Antarctic is opposite to
that of its counterpart in the Arctic Basin, where the ice generally rises upwards by bottom freezing in winter and surface
melting in summer. While summer ice formation mainly
takes place at the underside of Arctic ice by freezing of
under-ice ponds, ice grows only at the surface in Antarctica.
Superimposed-ice formation is important in many
respects. First, no mass is removed from the ice surface by
run-off even if snow melts.This, and the thermodynamic consequences of melting and refreezing for the energy balance,
have to be considered in numerical modelling studies. In contrast, snowmelt then contributes to summer ice formation.
Second, the resulting layer is mechanically very strong, and
extends the lifetime of a particular floe by preventing fracturing and deterioration of the otherwise rotten ice. Therefore,
for example, solar radiation input into the upper water layer
is reduced. The low-salinity, superimposed-ice layer floats in
water with sub-freezing temperatures, thus melting only
slowly.Third, the transformation of light-scattering snow into
clear, well-transmitting ice considerably changes the optical
properties of the ice surface. This and the coarsening of snow
reduces the albedo, as was obvious from the greyish appearance of many floes.
Although we lack quantitative data of surface roughness
and porosity of superimposed ice, the fresh, bubbly ice along
with the coarse snow in the pendular regime will strongly
increase radar backscatter.This is well known from the Arctic
(Holt and Digby, 1985; Onstott, 1992). Around Antarctica
also, Drinkwater (1998) and Morris and others (1998), for
example, show higher radar backscatter of perennial ice in
summer than in winter. Haas (2001) links the seasonal cycle
of European Remote-sensing Satellite scatterometer signatures with the summer processes discussed here. Rapid
increases of backscatter between November and December
mark the onset of superimposed-ice formation. This increase
is sporadically interrupted by sudden signal drops indicating
stronger melt events and snow wetting. The slow signal
decline from March onwards corresponds to the transition
to flood^freeze cycle processes. Thus, radar backscatter can
be taken to distinguish between warm- and cold-season
regimes. The seasonal backscatter cycle is observed in all perennially ice-covered regions around Antarctica, showing the
widespread occurrence of the processes described here.
By increasing light transmission and supporting the persistence of sea-water-filled gaps, the surface ice layer is of great
importance for establishing high biological activity, and in
particular high algal standing stocks within summer sea ice
(Fig. 8). In fact, these particular ice habitats support the greatest biological activities in pack ice recorded in the Southern
Ocean, comparable to those found in land-fast bottom ice or
platelet assemblages (Thomas and others,1998). Our heuristic
model links the physical processes leading to several habitats
formerly described by others. In fact, it appears that the infiltration layers (Syvertsen and Kristiansen, 1993), as well as the
highly porous or ``freeboard layers'' (Ackley and others, 1979;
Garrison and Buck, 1991; Ackley and Sullivan, 1994; Fritsen
and others,1994,1998) observed between late spring and early
fall, correspond to specific stages of the processes outlined
here. Thus, the gap layers (summer B and C in Fig. 9) might
well have developed from superimposed-ice formation on top
of what was an infiltration layer before (late-spring B in Fig. 9).
Our observations are biased by problems with sampling
rotten summer ice, and by the thick-section analyses. The
former leads to the loss of brine from the ice, on the one hand,
and the contamination of fresh ice with salt water, on the
other. The latter problem results in underestimates of the
amount of superimposed PG ice, which cannot easily be distinguished from granular snow ice if the grain-size is small.
However, the low salinity and d18O of much of the finegrained ice shows that it also is solely of meteoric origin.
The question whether superimposed-ice formation is a continuous, diurnal or episodic event could not be answered, as
the measured snow wetness was relatively low. Although we
introduced abiotic processes for the initial formation of gap
layers, the role of biophysical feedback mechanisms in further
developing these features remains unclear and deserves quan623
Journal of Glaciology
tification. Therefore, more detailed and quantitative future
studies of summer processes are a challenge to sampling
methods and strategies. This would also require the permanent monitoring of transformations in ice properties and of
the meteorological and oceanographic forcing at fixed
locations (e.g. at a drift station during melt onset).
ACKNOWLEDGEMENTS
We would like to express our sincere gratitude to captains
and crews as well as to the scientific cruise leaders H. Miller
andW. Jokat of RV Polarstern during ANT11/3 and ANT14/3
for their support of our work. Sampling also benefited
greatly from the care and companionship of the helicopter
teams. J. Askne kindly lent us the snow fork. M. Steffens
helped in the field. Constructive reviews by S. Gerland and
S. Ackley, as well as editorial remarks by M. Sturm,
improved the paper considerably. We also acknowledge discussions with G. Dieckmann and H. Eicken. The work was
partly supported by the Nuffield Foundation, U.K. Natural
Environment Research Council (GT9/2894, GR9/3309) and
the Deutscher Akademischer Austauschdienst/British Council
(ARC Programme).
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MS received 20 March 2000 and accepted in revised form 16 August 2001
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