Oceanogr.. 41(5), 1996,966-976
0 1996, by the American Society of Limnology and Oceanography, Inc.
Limnol.
Effects of glacial meltwater inflows and moat freezing on
mixing in an ice-covered antarctic lake as interpreted from
stable isotope and tritium distributions
Laurence G. Miller
U.S. Geological Survey, Water Resources Division,
Menlo Park, California
94025
GeorgeR. Aiken
U.S. Geological Survey, Water Resources Division,
Boulder, Colorado
80303
Abstract
Perennially ice-covered lakes in the McMurdo Dry Valleys have risen several meters over the past two
decadesdue to climatic warming and increased glacial meltwater inflow. To elucidate the hydrologic responses
to changing climate and the effects on lake mixing processes we measured the stable isotope (al80 and 6D)
and tritium concentrations of water and ice samples collected in the Lake Fryxell watershed from 1987
through 1990. Stable isotope enrichment resulted from evaporation in stream and moat samples and from
sublimation in surface lake-ice samples. Tritium enrichment resulted from exchange with the postnuclear
atmosphere in stream and moat samples. Rapid injection of tritiated water into the upper water column of
the lake and incorporation of this water into the ice cover resulted in uniformly elevated tritium contents
(> 3.0 TU) in these reservoirs. Tritium was also present in deep water, suggesting that a component of bottom
water was recently at the surface. During summer, melted lake ice and stream water forms the moat. Water
excluded from ice formation during fall moat freezing (enriched in solutes and tritium, and depleted in IsO
and 2H relative to water below 15-m depth) may sink as density currents-to the bottom of the lake. Seasonal
lake circulation, in response to climate-driven surface inflow, is therefore responsible for the distribution of
both water isotopes and dissolved solutes in Lake Fryxell.
interactions between antarctic lakes and their surrounding watersheds and, specifically, how climate-driven
changes in the inflow of water and solutes effect vertical
mixing processes and the estimate of the age of basin
refilling in Lake Fryxell.
Lake levels in the McMurdo Dry Valleys, the largest
of the polar desert oases on the antarctic coast, have risen
several meters since measurements began in the 1970s
(Chinn 1993). Lake Bonney, in the Taylor Valley, rose
13.5 m since it was fortuitously surveyed in 1903 during
the Scott expedition. Other Dry Valley lakes, including
Fryxell, continue to rise an average of 0.15 m yr- l, suggesting that the present climate is changing toward increased energy available for meltwater production (Wharton et al. 1992; Chinn 1993). Decreased ice-cover thickness in Lake Hoare, adjacent to Lake Fryxell, also suggests
a response to increased summer temperatures and melting
(McKay et al. 1985; Wharton et al. 1992).
Water in closed-basin antarctic lakes is predominantly
meteoric in origin (Ragotzkie and Friedman 1965; Matsubaya et al. 1979). Freezing, evaporation, sublimation,
and mixing of water from different reservoirs alter the
isotopic and chemical composition of meteoric water,
resulting in a distribution of water isotopes and dissolved
solutes that provides a distinctive signature of lake processes. The stable isotope ratios of water (P*O and 6D)
are conservative, such that only mixing and physical processessuch as evaporation, sublimation, and freezing will
alter their values. Tritium (3H as HTO) is a useful tracer
of water that has been in recent contact with the atmosphere, both because of its 12.3-yr half-life and because
Seasonal or perennial ice covers protect many polar
lakes from freezing completely in winter and represent a
balance between cold winter temperatures and warm
summer temperatures (Doran et al. 1996). In closed-basin
antarctic lakes, perennial ice covers restict vertical mixing
in the water column and influence the biogeochemical
processesthat occur there and in the sediments (e.g. Lawrence and Hendy 1985, 1989; Green et al. 1988). Vertical
profiles of dissolved solutes in ice-covered lakes have
been previously applied in estimating the lifetime of restricted mixing with one-dimensional diffusion models
and interpreted as basin refilling ages(Wilson 1964; Hendy et al. 1977; Lawrence 1982; Lawrence and Hendy
1985; W. B. Lyons pers. comm.). The purpose of this
study is to examine the effects of changing climate on the
Acknowledgments
We thank R. Smith, D. McKnight, M. Brooks, E. Furlong,
R. Quadry, R. Van Etten, B. Howes, and the staff at Berg Field
Center, McMurdo, for invaluable assistance in the field. We
especially thank the crews of VXE-6 for safe conduct through
a harsh environment. M. Huebner and D. White provided stable
isotope analyses, and R. Michel analyzed tritium via scintillation counting. We thank S. Drenkard, R. Weppernig, and P.
Schlosser for providing tritium analyses by 3He ingrowth. The
manuscript was greatly improved by suggestions from T. Torgersen, W. Carrothers, R. Michel, B. Lyons, D. McKnight, and
P. Mulholland. Special thanks to J. Garcia.
The W. M. Keck Foundation provided generous support for
establishment of the Lamont-Doherty Earth Observatory (LDEO) helium isotope facility.
966
Hydrology
of Lake Fryxell
of the peak in atmospheric tritium concentrations resulting from aboveground nuclear testing in 1950s and
early 1960s. Here, we use measurements of 6180, 6D, and
3H in water samples collected from the possible sources
and reservoirs in the Lake Fryxell watershed to examine
the mixing processes that control lake circulation in response to climate-driven variation in surface inflow.
Common features of many antarctic closed-basin lakes
(including Fryxell, Bonney, and Vanda) include the formation of perimeter moats in summer and increasing
dissolved solute concentrations with depth. Attempts to
explain the origin of saline bottom water in closed-basin,
antarctic lakes have focused on relict seawater (Hendy et
al. 1977; Masuda et al. 1988) evaporative concentration
of a pre-existing water body (Lawrence and Hendy 1985;
Torii et al. 1989), and input of high-salinity groundwaters
or modified surface waters (Green and Canfield 1984;
Green et al. 1989). One or more of these processes may
account for the distribution of dissolved solutes in icecovered coastal ponds (Matsubaya et al. 1979; Bird et al.
199 1) and in several dry valley lakes in the McMurdo
Sound region of southern Victoria Land. Distinguishing
between these processes is critical to understanding regional paleoclimate because the evaporative concentration hypothesis has been used to estimate the timing of
desiccation in several McMurdo Dry Valley lakes (Wilson
1964; Hendy et al. 1977; Lawrence 1982; Lawrence and
Hendy 1985; W. B. Lyons pers. comm.) These studies all
assume that following desiccation and basin refilling, a
solute concentration gradient existed between underlying
brine and overlying dilute water, and that molecular diffusion was responsible for the total solute mass now present in the upper water. However, if solute mass were
added to the surface, then the ages calculated would be
overestimates of the true age of the basin refilling. Conversely, if solute mass were added to the deep water of
the lake, then calculated ageswould be underestimates of
the true age of basin refilling. This paper presents evidence
for the addition of dissolved solutes to the deep water of
the lake via advection of modified surface water.
Site description
Lakes of the McMurdo Dry Valleys region of western
Antarctica are in a relatively ice-free region sustained by
the precipitation shadow of the Transantarctic mountains
and by the flow of persistently low-humidity air masses
originating on the polar plateau (Green et al. 1989). Lake
Fryxell lies near the eastern end of Taylor Valley at 77”37’S,
163’8’E and is the lowest (elevation, 18 m) in a series of
ice-covered lakes extending to Lake Bonney (elevation,
57 m) at the foot of Taylor Glacier. Lake Fryxell occupies
a closed basin between Canada Glacier to the west and
New Harbor on McMurdo Sound to the east (Fig. 1); it
drains a watershed area of 230 km2 (Lawrence and Hendy
1985). As many as 14 ephemeral meltwater streams drain
into the lake from the surrounding alpine glaciers during
summer, and the terminus of Canada Glacier is in direct
contact with the water column of the lake at its western
end.
967
%.--.-T.&p:
1.000
CONKIIJR
,NTuwAL
M’LES
Z,MM METERS
50 METERS WITH SUPFl EMFNTAW
COPiTOURS
DATUM IS hIEM4 5EA LEVEL
AT 25 METER INTtRVALS
Fig. 1. Map of the Lake Fryxell basin showingdrainagefrom
Commonwealth and other alpinc glaciersand the proximity of
Canada Glacier. Water-column sampling station occupied in
1987-l 989- x . Moat samples were collected between Huey
Creek and CanadaStream.
During summer, an ice-free moat commonly forms as
a result of surface input and local melting of the ice cover.
The combined sources of water maintain a positive water
balance in the lake, as seen by a 2-m rise in lake level
since 197 1 (Chinn 198 1, 1993). Surface-ice ablation is
the main water loss (30-40 cm yr-l; Henderson et al.
1965). The ice-cover thickness of Lake Fryxell has varied
little over time, ranging from 4.0 kO.5 m in 1963 and
1987-l 99 1 (Hoare et al. 1965; this study) to a minimum
of 3.5a0.5 m in 1986 (Clow et al. 1988), suggesting that
at present, freezing and ablation are in balance, but the
water budget is not in hydrologic steady state because
inputs exceed ablation plus moat evaporation.
The water column of Lake Fryxell was vertically stable
over the sampling period 1987-1990. Conductivity increased regularly with depth and gradients were consistent
from year to year (Smith et al. 1993). Temperature was
similarly consistent, ranging from 0.4”C below the ice
surface to 2.6”C at 18 m (max depth, 19 m) and reaching
a middepth maximum of 3.4”C between 9- and 10.5-m
depth with little seasonality. Aiken et al. (199 1) calculated
in situ density of the lake water from temperature and
conductivity measured in 1987, which, corrected for the
partial molar volume of dissolved solutes and gases(O,,
Miller and Aiken
968
Table 1. Morphometry and hypsography of Lake Fryxell
(estimated from bathymetry given by Lawrence (1982)).
Zone
Lake
Moat
Surface water
Deep water
Contour
interval*
(ml
o-3
3-6
6-9
9-12
12-15
15-18
>18
o-1.5
5-8
>15
vol.+
Area?
(x 1O-6 m2) (X 1O-6 m3)
17.1
7.14
4.27
11.3
8.39
3.26
5.36
2.33
2.99
1.24
1.47
0.75
0.35
0.23
2.15
1.44
9.80
3.60
1.82
0.75
* Relative to piezometric water surface.
t Surface area of top of contour interval, except Moat surface
calculated as area between shoreline and 1.5-m depth contour.
$ Volume = (lower area)(interval) + [(upper area - lower
area)(interval)/2].
Ar, N,), ranged from 1.00026 g cm-3 at 5 m to 1.006 12
g cm-3 at 18 m. They compared the average vertical
density gradient with density gradients in the thermocline
region of temperate stratified lakes. Using a relationship
between density gradient and vertical eddy diffusion coefficient (Quay 1977), Aiken et al. estimated an average
vertical eddy diffusivity of K, = 5.9 + 1.2 x 1Om5cm -2 s- l
for the water column (5-l 8 m) of Lake Fryxell. This value
was 6 + 1 times the tracer diffusion coefficient of water at
the in situ temperature (Li and Gregory 1974) and was
the basis for assuming that transport of dissolved solutes
in the water column was diffusion controlled.
The chemical and stable isotope mass balances we calculated are based on the bathymetric map provided for
1979 by Lawrence (Lawrence 1982; Lawrence and Hendy
1985). Integrated surface area and volume were determined for 3-m contour intervals, assuming that the 1979
and 1989 surface elevations were similar (Table 1). This
ignores the approximately 0.6-m rise in lake level since
1979 (Chinn 1993). The estimated volume error is > 20%
and is a result of both the uncorrected effect of surface
elevation on water depths and the sparse areal coverage
of the original bathymetric survey.
Methods
Sampling- Water-column samples were collected during summer 1987- 1989 by peristaltic pumping at < 1 liter
min-l through Tygon tubing. Water-column depths reported were relative to the piezometric water level (0.2
m below ice surface), which introduced an uncorrected
spatial error due to seasonal and yearly changes in lake
surface elevation. Ice-cover sample depths reported were
relative to the ice surface. Samples for stable isotopes
(al80 and 6D) were collected unfiltered and stored in 20-
ml glass vials with polyseal screwcap closures. Stream
and moat samples were collected by directly filling glass
vials after several rinses with the sample. Glacier, snow,
and lake-ice samples were melted in air before filling the
vials and stored unfrozen. Samples for tritium (3H) analyses were collected unfiltered in 2-liter glass bottles by
overflowing several bottle volumes with a tube placed in
the bottom to reduce contact with the atmosphere and
stored unfrozen. Additional water-column tritium samples were collected by slowly pumping (qO.2 liter min-l)
upward through vertically held sections of 0.95-cm-o.d.
copper tubing and sealing the tubing with pinch clamps.
Lake-ice tritium samples were collected in butyrate
tubes, thawed in an argon atmosphere, and syphoned
through vertically held sections of copper tubing as above.
A moat sample was collected in a 2-liter glass bottle and,
by keeping the outlet tubing away from the atmosphere,
siphoned with minimal air contact through the copper
tubing and stored unfrozen.
Sediments were cored in 1990 by driving a lined corer
(2.5-cm i.d.) into the bottom at locations in the lake corresponding to water depths of 8, 10, and 18 m. Upon
retrieval, the liner was separated from the corer and sectioned on shore by extruding intervals of the core into
gas-powered squeezers(Reeburgh 1967). Filtered (0.4 pm)
pore waters were collected in plastic syringes and transferred to scintillation vials for transport, unfrozen, to
Menlo Park for analysis.
Analytical techniques -Stable isotopes of water were
determined by CO2 equilibration (Epstein and Mayeda
1953) for 180, and zinc reduction (Kendall and Coplen
1985) for 2H and are reported vs. VSMOW. Stable isotope
compositions are reported in standard delta (6) notation
using units of per mil (7~) where
&180,
D)
VSMOW
=
(R,,rnp~RVSMOW
-
ljx
1,000.
(1)
R is the ratio of 180 *. 160 or 2H : ‘H. The 2-o analytical
precision was &O. lY& for al80 and + 1.57~ for 6D. Tritium
was analyzed both by liquid scintillation counting after
electrolytic enrichment (Thatcher et al. 1977) and by 3He
ingrowth (Clarke et al. 1976). Tritium concentrations are
reported in tritium units (TU), equal to 1 tritium atom
per 1018hydrogen atoms, and are corrected for decay to
the date of sampling. Precision (2 a) of tritium analyses
was + 0.3 TU at the lowest levels of detection (0.3 TU)
with counting and +O. 15 TU with 3He ingrowth. Chloride was analyzed on pore-water and water-column samples by ion chromatography (Oremland et al. 1987) after
1OO-fold dilution with deionized water.
Results
Stable isotope compositions of all water samples from
the Lake Fryxell watershed are presented in Table 2 and
Fig. 2, along with the global meteoric water line (6D = 8
k al80 + 10; Craig and Gordon 1965). Values of al80
ranged from - 24 to - 33o/oo
and values of 6D ranged from
- 200 to - 25 8a/oo.
The lightest samples were found in deep
969
Hydrology of Lake Fryxell
Table 2. Isotopic composition of water (vs. VSMOW) in the Lake Fryxell watershed.
Sample
Water column
Water
Moat and stream
Moat
Moat at creek
Huey Creek
Fryxell stream
Pore water
Sed. core, 8 m
Date
collected
Depth*
W-0
6’80
@d
18Dec87
18 Dee 87
18 Dee 87
18 Dee 87
4 Dee 88
4 Dee 88
4 Dee 88
4 Dee 88
4 Dee 88
26 Nov 89
26 Nov 89
26 Nov 89
26 Nov 89
26 Nov 89
26 Nov 89
26 Nov 89
26 Nov 89
26 Nov 89
26 Nov 89
5
7.5
11.5
18
5
8.5
10
15
18
5
6.5
8
9.5
11
12.5
14
15.5
17
18
-30.2
-31.0
-31.2
-32.0
-30.5
-31.3
-31.7
-32.1
-32.3
-31.1
-31.1
-31.3
-31.6
-31.6
-32.0
-32.1
-32.3
-32.3
-32.2
-232
-241
-240
-249
-233
-242
-243
-246
-244
-242
-243
-244
-246
-245
-249
-251
-253
-251
-250
2 Jan 88
4 Dee 88
15 Dec88
11 Dee 89
14 Dee 89
22 Dee 88
25 Dee 88
24 Nov 89
14 Dee 89
0.5
0.5
0.5
0.5
0.5
0.1
0.1
0.1
0.1
-25.8
-27.1
-26.8
-26.8
-28.4
-29.9
- 30.0
-30.5
-29.5
-211
-218
-216
-217
-228
-238
-240
-240
-234
- 30.4
-30.5
- 30.4
-30.3
-29.8
-29.4
-31.0
-31.1
- 30.9
-31.2
-30.7
-28.7
-30.0
-30.5
-29.5
-31.4
-30.4
-30.6
-31.1
-28.9
-31.4
-240
-241
-241
-239
-239
-235
-245
-245
-243
-244
-244
-236
-239
-241
-240
-247
-244
-244
-246
-238
-249
19 Dee 90
water§
9 Dee 90
0.05 11
9 Dee 90
0.15
19 Dee 90
0.36
19 Dee 90
0.48
Sed. core, 10 m
19 Dee 90
water
19 Dee 90
0.05
19 Dee 90
0.15
19 Dee 90
0.25
19 Dee 90
0.35
19 Dee 90
0.46
19 Dee 90
0.57
Sed. core, 18 m
18 Dee 90
water
18 Dee 90
oozell
18 Dee 90
0.35
18 Dee 90
0.12
18 Dee 90
0.22
18 Dee 90
0.32
18 Dee 90
0.42
18 Dee 90
0.52
18 Dee 90
0.62
* Depth below piezometric water level.
t Depth below ice surface.
# Water underlying ice cover.
0 Overlying water collected with core.
I( Depth below sediment-water interface.
ll Organic surface collected with core.
Sample
Ice
Glacier ice
Snow
Surface ice
Moat ice
Ice core 6
Ice core 9
Ice core 10
Ice core 11
Ice core 12
Ice core 15
Ice core 16
Date
collected
Depth?
Cm)
6’80
(~4
24 Dee 87
24 Nov 89
24 Nov 89
18 Nov 89
16 Dee 88
16 Dee 88
16 Dee 88
16 Dee 88
16 Dee 88
11 Dee 89
30 Nov 89
30 Nov 89
2 Dee 89
2 Dee 89
2 Dee 89
2 Dee 89
2 Dee 89
2 Dee 89
2 Dee 89
2 Dee 89
2 Dee 89
2 Dee 89
2 Dee 89
2 Dee 89
2 Dee 89
2 Dee 89
2 Dee 89
2 Dee 89
6 Dee 89
6 Dee 89
6 Dee 89
6 Dee 89
8 Dee 89
8 Dee 89
8 Dee 89
8 Dee 89
8 Dee 89
9 Dee 89
9 Dee 89
9 Dee 89
9 Dee 89
9 Dee 89
9 Dee 89
9 Dee 89
9 Dee 89
9 Dee 89
9 Dee 89
9 Dee 89
0
0
0
0
0
0
0
1.0
0
0
4.2
water*
0.25
0.75
1.25
1.75
2.25
2.75
3.25
3.6
water
3.6
0.25
0.75
1.25
1.75
2.25
2.75
2.5
3.25
3.6
water
0.75
1.75
2.25
3.25
3.75
0.25
0.75
1.25
1.75
2.25
2.75
3.25
3.75
4.25
4.75
water
-32.3
-31.0
-32.8
-32.1
-26.8
-26.1
-24.2
- 26.4
-25.7
-26.3
-27.8
- 30.6
-24.2
-25.7
-25.5
- 26.0
-25.8
-26.2
-26.5
-27.6
- 30.4
-27.5
-25.5
-25.6
-25.8
-25.5
-27.1
-26.3
-26.8
-26.5
-27.6
- 30.4
-26.6
-26.6
-26.2
-27.0
-26.7
-26.4
-26.4
-26.8
-26.6
-26.7
-27.5
-26.7
-27.3
-27.4
-26.9
-29.5
-249
-247
-258
-255
-219
-214
-202
-214
-210
-217
-226
-242
-208
-213
-215
-217
-215
-215
-219
-222
-242
-226
-212
-213
-214
-213
-222
-215
-218
-219
-222
-242
-219
-218
-218
-218
-217
-217
-216
-218
-217
-218
-223
-219
-222
-222
-221
-237
970
Miller and Aiken
Water column -Stable isotope composition of water
decreased with depth (Fig. 3A,B). Data from 1988 were
fit to a logarithmic equation, and 1989 data were best fit
to a linear equation. The 1989 6D data were systematically 4o/oo
lighter than the 1988 and 1987 data, except that
18-m water (1987-l 989) agreed within the analytical error. No systematic difference existed in the multiyear al80
profiles, except that 1989 samples collected just below
the ice cover were lighter. Water-column tritium data
from 1989 and 199 1 were combined to produce one profile of 3HH0 vs. depth (Fig. 3C). Both L-DE0 and USGS
data were used in this plot and the observed agreement
between methods is excellent. Tritium concentrations
ranged from 4.0 TU at 5 m to below the detection limits
(co.3 TU) at 12 and 18 m with elevated concentrations
in four samples at 14-, 15.5-, and 17-m depths.
-210
-220
g-230
s
-240
-250
-260
-34
-32
-30
-28
-26
-24
6’“O (%o)
Fig. 2. P*O vs. 6D of all waters sampled in the study: Cllake water column; A- moat; 0 -streams; e -pore water; + Canada Glacier; O-lake ice; v-surface lake ice; + -snow.
Shown also is the global meteoric water line (GMWL) and the
regional evaporation trend line (slope = 5.6). Error bars show
+l SD.
levels of the water column, fresh snow at lake level, and
ice from Canada Glacier. Heavier samples were associated with sublimation of lake ice and evaporation of the
moat.
-32
Ice cover-Stable isotope composition of water decreased with depth in the ice cover (Fig. 5) with the lightest
values at the bottom of the ice cover corresponding to
3HH0 (TU)
6D (%o)
8’80 (%o)
-33
Pore water-Mean pore-water values of stable isotopes
and the isotope compositions of overlying water collected
with the cores were heavier than those of the water-column samples collected in the center of the lake at depths
corresponding to the sediment-water interface (Table 3,
Fig. 4A,B). Pore-water chloride concentrations, however,
were similar to corresponding lake-water values and
showed no trend with depth (Fig. 4C), suggesting that the
upper sediments were open to exchange with lake water.
-31
-30 -255
-250
-245
-240
-235
-230
0
1.0
2.0
3.0
4.0
5.0
12 lb..16”
18 -
A
Fig. 3. Vertical profiles of 6’*0, 6D, and 3HH0 vs. depth in the water column: O- 1987; A- 1988; Cl- 1989. Solid lines are
log (1988) and linear (1989) fits to the stable isotope data and polynomial (1987-1989) fit to 3HH0 data. Tritium samples from
1989 measured by 3He ingrowth (Cl-L-DEO), all others by scintillation counting (O-USGS). Error bars show Ifl 1 SD.
Hydrology of Lake Fryxell
971
Table 3. Stable isotope composition (%) and chloride concentration (mM) of bottom water
and sediment pore water in Lake Fryxell. (Not determined--d.)
89
10
18
Pore water?
Bottom water*
Depth
Cm)
6’80
6D
-31.2f0.2
-243&l
(n = 3)
-31.4kO.3
(n = 3)
-32.3kO.3
(n = 3)
(n = 3)
-244f2
(n = 3)
-248+3
(n = 3)
Cl-$
6D
6180
-240+ 1
(n = 5)
-242+4
(n = 5)
-243+4
(n = 9)
-30.320.3
(n = 5)
-30.7kO.8
(n = 5)
-30.4LO.9
(n = 9)
28.4
43.0
115
Clnd
66.4f4.0
(n = 5)
1lOk4.8
(n = 6)
* Mean f 1 SD of samples collected 1987-l 989.
t Mean + 1 SD of all samples collected 1990.
I#I1987 Water column Cl- from Aiken et al. 1996.
0 Depth of core.
Discussion
equilibrium freezing of surface waters (Craig and Gordon
1965). Values of both al80 (Fig. 5A) and 6D (Fig. 5B)
were heaviest at the ice surface, presumably due to sublimation of ice. Generally constant values were observed
between the surface and bottom of the ice cover. Tritium
values in the ice were greater than in corresponding samples collected in the water column just below the ice cover
(Table 4). Tritium concentrations in the deepest ice samples increased from 4.0 TU at the center of the lake to
6.5 TU at a site near Canada Glacier and the outlet of
Canada Stream.
6’80 (%o)
-34
-32
6D (%o)
-30
$8
-260
B
20
-
+s
Oki’0
A
v
I
1 ,
overlying water
,y’ 7’
l-w IJiP
/
%I+
50
-230 0
1
I
C
I
‘,
150
100
,A,
‘,V’
I
overlying water
l-o-l
\
‘0
El I
$30
E.
-
I
0 loo
o"40
-
/
O\
q \
50 -
,
I,
Chloride (mM)
-240
-250
I
10 -
Water budget -Because Fryxell is a terminal lake, its
surface level readily responds to changes in the water
budget, particularly meltwater input. Since 197 1 the surface level of the lake has increased an average 10.5 cm
yr-l, although some yearly decreases in elevation were
also measured (Chinn 1993). Surface-ice ablation determined in Lake Fryxell (30-40 cm yr-I, Henderson et al.
1965) is similar in nearby Lake Vanda and Lake Hoare
(30-35 cm yr- l, Chinn 1993; Clow et al 1988) and is
s,
60 -
Fig. 4. Vertical profiles of 6l8O, 6D, and Cl vs. depth in sediment pore waters: O-core at 8 m; Cl-core at 10 m; 0 -core at
18 m. Open symbols above sediment-water interface represent water collected during coring process; solid symbols are watercolumn samples collected separately. Dashed lines connect 18-m data. Error bars show + 1 SD.
972
-32
Miller and Aiken
-30
6’80
-28
(9&I)
-26
-24
-22
'50
-240
6D (YAo)
-220
-230
-210
-200
Table 5. Stable isotope composition and tritium concentration of the well-mixed reservoirs and water masses in Lake
Fryxell (mean + 1 SD of samples collected in 1987-l 989). (Not
determined - nd.)
Glacier
Stream
Moat
Surface
(5-8 m)
Deep
(15-18 m)
Fig. 5. Vertical profiles of 6’*0 and 6D vs. depth in selected
cores from the ice cover: 0 -ice core 6; O-ice core 9; O-ice
core 15; A-ice core 16; + -moat ice. Solid symbols are underlying water. Error bars show + 1 SD.
likely constant with time. Using 30 cm yr-l ablation and
Lawrence’s ( 1982) estimates of surface area (7.1 X 1O6m2)
and volume (47 x lo6 m3), we calculate a steady-state (ignoring lake level increases) residence time of water as
r=
volume
ablation x area’
(2)
yielding a whole-lake residence time of 22 yr. This is a
maximum residence time because evaporation of water
from the moat is not included and because we chose the
lower estimate for ablation. It is not a true whole-lake
residence time, however, because there is much less than
complete mixing in the water column. Rather, this is
considered an estimate of how quickly water entering the
lake at the surface is mixed in the upper water column,
frozen in the ice cover, and advected through the ice to
the surface. For an ice-cover thickness of 4.5 m and ab-
Table 4. Tritium concentration of the ice cover and surface
water below the ice (samples collected in 1989).
Location
IC-9
IC-9, WC?
IC-10
IC-15
IC-15, WC/
Depth*
b-d
3.5-3.7
3.8
3.5-3.7
0.5-l .o
1.5-2.0
2.0-2.5
3.0-3.5
3.5-4.0
5.0, 5.5
6.12kO.15
4.00+0.15
6.53kO.15
4.61 f0.18
4.79f0.17
3.63kO.15
2.72f0.15
3.95kO.15
3.75kO.15
(12= 2)
~ * Depth below ice surface, except WC depth below piezometric water level. WC depth is -0.2 m lower.
t Mean k 1 SD of concurrent below-ice water samples.
a’*0
W)
-32.0+0.9
(n = 3)
-3O.OkO.5
(n = 3)
-26.6f0.6
(n = 4)
-31.2kO.l
(n = 3)
-32.3kO.l
(n = 6)
;i
-251+6
(n = 3)
-238+3
(n = 3)
-216+3
(n = 4)
-243+2
(n = 3)
-249f3
(n = 6)
(;“u)
nd
3.lkO.3
(n4=22)
’
3.7&O. 1
(n = 5)
0.4kO.3
(n = 3)
lation of 30 cm yr - l, a transit time of water through the
ice cover is calculated as
thickness
r=
(3)
ablation ’
yielding an ice-cover residence time of 15 yr. These calculations, coupled with the observations of elevated tritium concentrations in surface lake water and in the ice
cover relative to incoming streams (Tables 4, 5), suggest
that much of the new water entering the lake is rapidly
cycled through the upper water column and the overlying
ice.
Sources of water- Liquid water has continuously occupied the Lake Fryxell basin for at least 1,000 yr (Lawrence and Hendy 1985; Green et al. 1988, 1989; W. B.
Lyons pers. comm.). The possible sources of water entering the lake are direct input of surface meltwater from
Canada Glacier to the ice cover, the proglacial streams
and their hyporheic zones (McKnight and Andrews 1993),
the moat, permafrost melting (Stuiver et al. 198 l), and
groundwater (S. Tyler pers. comm.). The measured water
reservoirs are distinguished by their isotope compositions
(Table 5), although information is lacking for permafrost
6D and groundwater in general.
Canada Glacier ice is a source of dilute water with the
lightest stable isotope composition in the watershed (Fig.
2, Table 5). Canada Glacier is 2-3?&~lighter in 180 than
the other glaciers draining into Lake Fryxell (Matsubaya
et al. 1979; Stuiver et al. 198 1). Tritium was not measured; however, glacier ice away from accumulation zones
should contain tritium only in surface samples.
Stream water is likewise chemically dilute compared
to lake water (Green et al. 1989) but is enriched in both
180 and 2H due to evaporation (Table 5). The trend of
the evaporation line (Fig. 2) is described by a slope of
5.6, consistent with low humidity control of the process
(Gat and Tzur 1967). Stream-water tritium concentrations are similar to expected ambient concentrations resulting from exchange with the atmosphere.
Hydrology of Lake Fryxell
Moat water is chemically similar to stream water. Concentrations of S042- and Cl- in moat water collected in
1987 (35.0 and 269 PM; L. Miller unpubl. data) were
14& 1% greater than the volume weighted S042- and Clconcentrations of all streams measured in 1982 (3 1.O and
235 IAM, Green et al. 1988). However, moat water is
greatly enriched in 180 and 2H compared to stream water
(Table 5) and lies further along the evaporation line (Fig.
2). Moat water is also enriched in tritium over stream
water (Table 5), but it is unclear whether enrichment is
a result of greater contact with the atmosphere, evaporation, or selective storage of water with higher tritium
content (i.e. memory effect of higher tritium concentrations in the past decade).
The three remaining reservoirs - permafrost, hyporheic
zone water, and groundwater- were not sampled in this
study. However, surface permafrost al80 measured during the Dry Valley drilling project ( 1975) was similar to
glacier al80 in the vicinities of Commonwealth and Canada Glaciers (cores 11 and 12, Stuiver et al. 198 1). For
example, the 6180 of permafrost from 1O-m depth in both
cores was -32%~. Permafrost melting is considered an
unkown but possible source of subsurface water to nearby
Lake Hoare (S. Tyler pers. comm.). The hyporheic and
saturated zones are environments where solutes (including aerosols and weathering products) are likely to be
dissolved into water, and both zones offer restricted contact with the atmosphere. Unfortunately, the extent and
chemical and isotopic compositon of these subterranean
waters are unknown.
Lake chemistry and isotopes-Surface lake water is
chemically similar to moat and stream water (Green et
al. 1989; Aiken et al. 199 1; McKnight et al. 1993). The
stable isoltope composition of surface lake water is likely
a result of equilibrium freezing of moat or stream water,
resulting in heavier ice and lighter residual water (Craig
and Gordon 1965). The equilibrium fractionation factors
for freezing in Lake Fryxell are 1.004 for al80 and 1.020
for 6D (CU
factors derived from Table 2). Yearly variations
in the stable isotope composition of surface water (Fig.
3) can then be explained by fluctuations in the amount
of “new” input (heavier 6180and 6D) if the annual amount
of freezing is constant. The resulting fluctuation in stable
isotope composition of ice (lym for 6180 and loo/o0for 6D)
is less than the variability observed in the ice cover values
(Fig. 5).
Enhanced vertical mixing in the upper water column
is evident from nearly uniform concentrations of dissolved solutes between the bottom of the ice cover and
6-7-m depth (Aiken et al. 199 1, 1996; Smith et al. 1993;
McKnight et al. 1993). This effect is also seen in the
tritium profile (Fig. 3C). Stream and moat water entering
the lake has an estimated density of 1.00004 g cm-3 at
1.5”C (derived from stream chemistry of Green et al.
1989), which would allow it to sink only to 5 m. During
the melting period, surface water is transported from the
edges of the lake toward the center by horizontal eddy
diffusion (KX = 10-3-10-1 cm2 s-l, Quay 1977; Coleman
and Armstrong 1983) and by horizontal advection of wa-
973
ter down gradient (1 .Ocm s-l, Ragotzkie and Likens 1967;
McKnight and Andrews 1993). Enhanced vertical mixing
therefore may result from both local convection and shear
below the ice cover.
During the freezing period, surface water may move
laterally under the ice in the opposite direction, toward
the shoreline (Gade et al. 1974; Welch and Bergmann
1985). Freezing of surface water to the bottom of the ice
cover in winter causes salt exclusion and local convection
of the upper water column (Canfield et al. 1983) and
increased concentration of salts toward shore (Ferris et
al. 199 1) in seasonally ice-covered lakes. This cryogenic
concentration of solutes in the nearshore surface zone
may contribute to the formation of bottom water in Lake
Fryxell during fall.
The vertical structure between 8 and 15 m is described
by regularly increasing density with depth resulting from
nearly uniform temperature (3.2+0.2”C, n = 20) and increasing dissolved solute concentrations (Aiken et al. 199 1;
Smith et al. 1993). Biogeochemical processesare complex
within this region of the water column, including calcium
carbonate precipitation between 8 and 9 m (Lawrence
and Hendy 1989). Overall, this region is characterized by
nearly linear profiles of dissolved conservative constituents, including stable isotopes of water (Fig. 3A,B), leading to a uniformly high density gradient and low K,. This
middle region is the most stable vertical zone in the water
column.
Below 15-m depth, the water column is less stable.
Deep water has elevated concentrations of major and
minor elements (Lawrence and Hendy 1985; Green et al.
1989; Aiken et al. 1991; McKnight et al. 1993; Smith et
al. 1993); however, the density gradient between 15 and
18 m is less steep, reflecting a smaller vertical gradient
in dissolved solute concentrations (Aiken et al. 199 1).
The deep lake water is isotopically lighter than water
higher in the column (Fig. 3A,B). In addition, tritium is
present in water collected at 14-, 15.5-, and 17-m depths
(Fig. 3C), suggesting that some component of bottom
water was in contact with the postnuclear atmosphere
and that this signal has not decayed to below the detection
limit for 3H. This tritiated deep water cannot be explained
by diffusion from either above or below the zone.
Deep water is isotopically lighter than pore water (Table
3; Fig. 4A,B). To further distinguish sources of bottom
water, the 1989 water-column stable isotope values are
plotted against water-column chloride data (Aiken et al.
1996) in Fig. 6, along with the 1990 pore-water samples
collected at 10 and 18 m. The linear trends in watercolumn values suggest end-member mixing between surface and deep water. Pore waters plot off the trend, suggesting two possibilities: if pore water is the source of
solutes and water isotopes to the bottom water, then some
fractionation in stable isotope ratios should occur during
transport across the sediment-water interface, or the pore
water is not a mixing end-member for water isotopes.
Advection of bottom water-To explain the bottomwater distribution of stable isotopes and tritium, we propose that modified surface water, enriched in solutes,
974
Miller and Aiken
tagged with tritium, and depleted in 180 and 2H, is injected at 15-l 8-m depth. Bottom water (> 15 m) presently
has a density > 1.00523 g cm-3, hence surface water such
as stream or moat water would have to be significantly
concentrated to sink as density currents to 15-m depth.
Two mechanisms for solute enrichment are considered
feasible: dissolution of soluble salts during subsurface
transport and salt exclusion during fall freezing of the
moat. Both mechanisms rely on density currents to transport modified surface water along the basin boundary to
the deepest part of the lake.
Stream water entering the moat has a well-defined
chemical and isotopic signature, whereas water transported in the hypohreic zone (4 times the cross-sectional
stream area; McKnight and Andrews 1993) may be more
enriched in solutes. Seepagemeters in Lake Hoare sampled groundwater transported through the shallow sediments of the lake; this water was enriched in chloride and
0.5%0 lighter in I80 than the ambient lake water at the
same site (S. Tyler pers. comm.).
If we assume that moat water is essentially stream water
and begin with volume-weighted average major element
concentrations (Green et al. 19SS), we calculate a moat
density of 1.000042 g cm-3 at 1.5”C. If freezing were 90%
efficient at excluding dissolved solutes (Canfield et al.
1983), then an enrichment
factor of [0.00523/
(0.000042 x 0.9)] or 138 is required to increase moatwater density to 15-m water density. This could occur
only when the moat is nearly completely frozen and < 1%
of its volume remains as a saline brine, presumably just
below the moat ice. The resulting moat ice would be
isotopically heavier than the previous moat and the residual brine would be depleted in the stable isotopes of
t
-I
-230
10 m pore water
- B
P
18 m pore water-
-240
2
s
1989 water column
-250
r = 0.9574
-260
0
20
40
60
80
100
120
Chloride (mM)
water. Beginning with moat isotope composition (Table
5), water remaining after freezing would follow a Rayleigh
curve, depending on the isotopic fractionation factor cy
and the fraction of water remaining (f):
6 = 1,000(~+1) - I),
(4)
Fig. 6. Dissolved chloride vs. 6180 and 6D in water-column
and pore-water samples. Solid lines are linear regressions of
1989 water-column data (Cl) where Cl- (Aiken et al. 1996) increasesregularly with depth. Pore waters are from cores collected
in 1990 at 10-m (0) and 18-m (A) water depths.
resulting in brine compositions of 6180 = - 43.4?& and
6D = - 299o/oofor a = 1.004 and 1.020, repectively (from
Table 2), and forf = 0.0 1. Tritium may fractionate twice
as much as 2H. This implies a tritium fractionation factor
cy= 1.040 and results in a calculated decrease in tritium
of 16.8%. Advection of 1% of the moat volume as brine
with this isotopic composition would add 2 x lo4 m3 of
water, enriched in tritium and depleted in stable isotopes,
to the bottom water (15-18-m depth) annually. This is
calculated as an isotope mass balance:
= c3v3.
CJ, -I- c,v,
(5)
could supply the observed tritium in the bottom water
over the four decades of elevated surface tritium deposition. It is likely that this process is not at steady state
(i.e. there may be years when little or no brine is formed
or the brine formed is not dense enough to sink to the
bottom). In years when a larger moat develops due to
increased glacial meltwater inflows, greater production of
high density brines may occur during moat freezing. Over
time, this input could result in significant and measurable
depletion of bottom-water stable isotopes and contribute
to the present distribution of water isotopes and dissolved
solutes in the water column.
The seasonal input of glacial meltwater to the lake surface and of modified surface water to the deep-water column promotes stability in the water column by increasing
Cl, C2, and C3, are the isotope ratios (180 : 160, 2H : ‘H)
and tritium concentration of the initial water at 15 m,
the moat, and the final composition
at 15 m, and V1, V2,
and V3, are the volumes of 15-18-m water, the volume
of added brine, and the final volume. This annual addition (N 1% of the deep-water volume) would result in
decreases in 6l 8O and 6D of 0.12 and 0.6%0 and an increase
in tritium of 0.04 TU. Although the input function of
tritium in Antarctica is not well known, this mechanism
the solute concentration gradient. This, in turn, affects
any regional climatic inferences based on estimates of the
basin refilling age. If the concentration gradient reflects
both diffusive and advective processes, then conceptually
advection
would tend to steepen the gradient over time,
and diffusion would tend to flatten the gradient. Calcu-
Hydrology of Lake Fryxell
lated diffusion-cell ages, therefore, may be underestimates
of the age of basin refilling if the concentration gradient
has a strong advective component.
It is unclear what fraction of dissolved solutes in the
bottom water originates by the above mechanism vs.
transport from the sediments. To evaluate the relative
contribution
of surface brines to the deep-water solute
flux we must construct mixing models using time-dependent geochemical tracers with known input functions. For
example, measurement of the anthropogenic gases CFC11 and CFC-12 in the water column may help to identify
. the extent and timing of bottom-water advection. Determining lateral gradients in below-ice concentrations of
solutes, 3H, and stable isotopes of water (al80 and 6D)
during fall freezing of the moat and throughout winter
would demonstrate the significance of our proposed hypothesis and link surface phenomena (e.g. changes in
meltwater input as a function of climate change) to bottom-water formation processes.
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