ocenography

Notes
application to lakes and seas. Acta R. Sot. Sci.
Litt. Gothob. Geophys. 1: l-26.
WILLIAMS, G. P. 1969. Water temperature during the
melting of lake ice. Water Resour. Res. 5: 11341138.
Limnol. Oceanogr., 36(2), 1991, 335-342
0 1991, by the AmericanSociety of Limnology
and Oceanography,
335
Submitted: 21 September 1989
Accepted: 5 September 1990
Revised: 20 October 1990
Inc
Similarity of whole-sediment molecular diffusion coefficients in
freshwater sediments of low and high porosity
Abstract- Whole-sediment
molecular diffusion coefficients (0,) for tritiated water in pore
waters of various lakes were determined experimentally by adding ‘H,O to the overlying water
of asphyxiated (without bioirrigation)
and unasphyxiated cores and measuring the resulting
pore-water profiles after a period of time. Our
objectives were to determine the relationship between D, and Do (the diffusion coefficient in pure
water) in sediments with a wide range of porosities and organic contents and to examine the
influence of bioitigation
on solute transport and
on the predictability of 0,.
We found that Do/D, did not change as much
as expected with increasing porosity, i.e. in lowporosity sediments the average DJD, was 1.8 +O. 1
and in high-porosity sediments it was 1.520.2.
We also found that the effect of fauna1 activity
on the predictability of D, was only significant in
sediments with high (14,000 ind. m-z) invertebrate populations. This result means that in most
freshwater sediments, the sediment diffusion coefficient can be predicted reliably from the molecular diffusion coefficient at in situ temperature.
Measurement of vertical fluxes of substances across the sediment-water interface
is important in both freshwater and marine
research. These flux measurements are useful in studying diagenesis below the sediAcknowledgments
We thank N. Hafkamp, his crew, and P. Kieskamp
for technical assistance, V. St.Louis, H. Roon, A. Furutani, B. Miskimin, and D. Hamilton were helpful in
the field and in the laboratory. 0. van Tongeren was
helpful with the mathematical modeling of the tritium
diffusion profiles. Valuable criticism of the manuscript
was provided by G. Brunskill. The comments of M.
Rutgers van der Loeff and two anonymous reviewers
are greatly appreciated.
This research was partially supported by NSERC
grants OGPGPOlO and STRGP036 and the Department of Fisheries and Oceans, Canada.
men&water interface as well as its effects on
the chemistry of the overlying water. Vertical fluxes of solutes between sediments and
the overlying water can be calculated from
Fick’s first law as described by Bemer (1980):
J, = -#D,(dc/dx)
(1)
where J, is the flux in mol cm-’ SC’, rb the
porosity of surface sediment, D, the “wholesediment molecular diffusion coefficient” for
the diffusing substance in cm2 s-l, dcldx the
initial concentration gradient in the sediment in mol cmS3 cm-‘, and x the sedimentary depth measured positively downward.
When making flux estimates, the porosity
and concentration gradients of solutes in the
sediments are usually measured, but D, is
usually estimated. D, must be estimated because while diffusivities of solutes in particle-free water (D,) are well known (Li and
Gregory 1974), the effect of tortuosity on
D, is not. Tortuosity (6) is defined as the
ra-atioof the distance (dl) that a diffusing species must travel through the interstitial space
in a porous medium to the linear distance
(dx) that it would travel in particle-free medium (0 = dl/ti).
Therefore, D, is always
CL),, and the relationship between them has
been described mathematically
by Bemer
(1980) as
D, = D,/B2.
(2)
In addition to molecular diffusion, fauna1
activity can affect the vertical solute transport in the sediment (Aller 1982; Rutgers van
der Loeff et al. 1984; Krantzberg 1985).
Thus, the effective or apparent diffusion coefficient (0,) of animal-containing
sediment
336
Notes
is the sum of 0, and the increased diffusion
due to fauna1 activity (Di):
In the Dutch lakes profundal sediment
cores (50 cm long, 7-cm diam) were collected with a modified Jenkin surface mud
De = D, + Di.
(3) sampler. Littoral sediment cores (40 cm
Our main objectives here were to inveslong, 5- or 7-cm diam) were taken with a
tigate Do/D, of sediments with a wide range cylindrical,
stainless-steel, bottom corer
of porosities and the influence of bioiniequipped with a steel cutter head and a sepgation on solute transport across the sedi- arate Perspex core liner. The sampler is
ment-water interface. To do so, we used forced into the bottom by hand and then
tritiated water to examine the variation of closed with rubber stoppers. In the North
D, in minimally disturbed cores of sedi- American and Norwegian lakes, cores (Perments from a variety of freshwater systems.
spex core tubes, 5-cm diam x 15 cm long)
D, values were also measured via asphyxiwere taken by a diver (see Rudd et al. 1986).
ated sediment cores (Rutgers van der Loeff
The effective diffusion coefficient (D,) was
et al. 1984). These measurements allowed
estimated as described by Rudd et al. (1986)
the calculation of Do/D, and Do/De in sev- with slight modifications for the Dutch lakes.
eral different kinds of freshwater sediments.
After retrieval from the lake, cores were
In this study, D, is defined as the diffusion
taken to the lab and placed at in situ temcoefficient in pure water. In contrast, D, is perature for several hours to be sure that
measured in pore water and includes the the entire core was at the same temperature.
effects of tortuosity as well as any other fac- Tritiated water (3H,0) was then added to
tor decreasing diffusion. There is no way to the surface water (1 O-40-cm column) of dumeasure 0 directly, but it is thought to be plicate or triplicate cores. The top stoppers
related to porosity and to be the main de- were sealed tightly and covered with Vasterminant of D, (Eq. 2). We have measured
eline to prevent evaporative loss of 3H,0.
D, and examined how it differs among sed- The cores were then incubated for 12-24 h
iment types and whether current theory ex- at in situ temperature in the dark. During
plains what we found.
incubation, the surface water was stirred
D, also includes Di, and the calculation
gently by rotating the cores in a l-r-pm roof Di from D, - D, (Eq. 3) allowed an esti- tary shaker (ELA), by gentle periodic stirmate of the fractional contribution of sed- ring (Norway, Adirondacks), or with a moiment macrofauna to the apparent diffusion
tor-driven
impeller
(The Netherlands,
Denmark) (Sweerts et al. 1989). The stirring
coefficient (DJD,).
Cores were obtained from Lake Vechten,
procedures of the overlying water introLoosdrecht, Maarseveen I, Tongbersven,
duced a temperature-dependent,
diffusive
mean boundary layer roughly estimated to
and Gerritsfles in the Netherlands. Lake
be z (mm) = 1.0 - 0.02 x T (“C) (Sweerts
Vechten is mesotrophic and monomictic,
et al. 1989). After incubation the cores were
Loosdrecht is composed of several shallow
eutrophic lakes, and Tongbersven and Gersliced at 0.5-l -cm intervals. The slices were
ritsfles are oligotrophic
moorland pools.
centrifuged (10 min at 20,000-30,000
x g)
Cores were also obtained from oligotrophic
to separate pore water from sediment.
Lake Kalgaard in Denmark and from Lakes
Concentrations of 3Hz0 in the pore-water
227, 302S, and 223 in the Experimental
samples at depth x and time t (C,,,), and in
Lakes Area (ELA, Northwestern Ontario).
the initial (C,) and final (C,) surface-water
L223 is small and oligotrophic, L302S is samples, were determined by scintillation
mesotrophic, and L227 is small and eu- counting spectrometry. The initial concentrophic. In addition, cores were obtained
tration of 3H,0 in the surface water of the
from Howatn
and Lille Hovvatn Lakes
cores did not decrease by > 10% during in(southern Norway),
Woods, Sagamore,
cubation. D, and D, for all but the Dutch
Darts, and Big Moose Lakes in the Adironlakes were calculated with the finite differdack Mountains (upper New York State), ence model described by Rudd et al. (1986).
and Crystal Lake (northern Wisconsin)-all
This model uses the measured porosity proare small and oligotrophic.
file and a finite volume of overlying water
Notes
as in the experiment so that the change in
C, to C, is included.
In the Dutch lakes we used a constantsource computer model based on the error
function described by Duursma and Hoede
(1967):
C,,, = Co x erfc[x/2(Dt)ya]
(4)
where erfc is the complementary error function, C,, the initial 3H,0 concentration in
the surface water of the core, and D is D, in
unasphyxiated cores and 0, in asphyxiated
sediments. This simplification was possible
as the porosity of the cores was nearly constant with depth (Table 1). The two models
gave similar results. A correction was made
in the program for the decrease of 3H20 in
the overlying water during incubation (C,
was adjusted similar to equation 3.32 of
Crank 1975). For each core the experimental data were compared with the model for
different trial values of diffusion coefficients. The best estimate of 0, or 0, was
determined by least-squares. The depth (x)
of each slice corresponding with its average
3H,0 concentration [C&,1 was corrected iteratively for the nonlinear diffusion profile
(effective depth x~). This correction was
needed because in a nonlinear profile the
average concentration over a layer is not
equal to the concentration at the average
depth of the layer. The effect of such a correction is the same as application of the leastsquares method to the mean concentrations
as derived by integration of the expected
profile. Using this method we measured D,
of 3H,0 in agar (99% water at 18°C and
found it to be 1.96kO.7 x 10e5 cm2 ssl) (n
= 3) which is within 5% of the theoretical
value (2.07 x 10m5 cm2 s-l; Wang et al.
1953).
After incubations were completed, cores
were sliced into 0.5-cm intervals, and porosity was determined from weight loss on
drying (60°C). The specific gravity of water
was taken as 1.O. The specific gravity of the
inorganic part of the sediment was taken as
2.6 (based on specific gravity measurements
from seven European lakes), and the specific
gravity of the organic part of the sediment
was taken as 1.2 (Hakanson and Jansson
1983). For this calculation organic matter
337
was measured by weight loss on ignition (3
h at 550°C) of the dried sediments.
Carbon content of the dried and homogenized material of the top 4 cm of a sediment core was determined with a CHN analyzer (Carlo Erba Strumentazione elemental
analyzer model 1106).
To examine the effect of bioturbation on
transport across the sediment-water interface, we stopped bioirrigation
by the asphyxiation technique (Rutgers van der Loeff
et al. 1984); the surface water of duplicate
cores was allowed to become anoxic for 1
week before determination of 0,. The procedure stopped all bioirrigation
in cores
taken in winter and summer. Macrofaunal
densities were determined
for nonasphyxiated sediments by counting the average number of individuals from the top 4
cm of triplicate sediment cores.
Two general types of sediments were
identified on the basis of water and organic
content, i.e. littoral sediments of low porosity and low organic content (sand) and
littoral and profundal sediments of high porosity and a high organic content (silt, floe,
and peat, Table 1). Of the inorganic material
of the low-porosity,
organic-poor
sediments, 90% consisted of particles > 105 pm.
Of the inorganic material of the high-porosity, organic-rich sediments, 90% consisted of particles < 16 pm (unpubl.).
The average variation of the 3H20 diffusion measurements in duplicate or triplicate sediment cores was 14%. Core-to-core
variations found in other radioisotope diffusion studies of sediments have been 15%
(Duursma and Bosch 1970), 25% (Goldhaber et al. 1977), and 14% (Dicke 1986).
Temperature was an important determinant of D, or D,, as it was for D, in particlefree water (Fig. 1). The effect of temperature-induced
viscosity changes on 3H20
diffusion in dilute solutions is described by
D, = (0.0525T + 1.099) x 10e5 cm2 s-’
(5)
where T is “C. This formula is derived from
the Stokes-Einstein relationship for the diffusion coefficient of water as a function of
temperature (Li and Gregory 1974), and the
diffusion coefficients (Do) for 3H20 (Wang
et al. 1953). All of the 0, and 0, values,
10
10
3
3
3
3
3
3
3
3
3
3
3
3
3
2
2
1
2
24
3
3
1.50
1.50
4
3.50
1.50
2.50
2.50
5
2
5
2
2
2
3
3
6
Vechten
Vechten
Vechten
Vechten
Loosdrecht
Loosdrecht
Loosdrecht
Loosdrecht
Loosdrecht
Loosdrecht
Loosdrecht
Loosdrecht
Loosdrecht
Loosdrecht
Loosdrecht
Gerritsfles
Tongbersven
Kalgaard
Maarseveen I
Maarseveen I
L227
L227
302s
302s
223
Big Moose
Woods
Sagamore
Darts
Hovvatn
Hovvatn
L. Hovvatn
L. Hovvatn
Crystal
Crystal
Crystal
Crystal
Cl-ptZll
Depth (ml
Lake
Table I_ Sediment diffision
Silt
Silt
Sand/clay
Sand/clay
Peat
Peat
Peat
Peat
Silt
Silt
Silt
Sand
Sand
Sand
Sand
Floe
Floe
Sand
Sand
Silt
Silt
Silt
Sand
Sand
Sand
Sand
Floe
Floe
Plot
Floe
Floe
Floe
Floe
Sand
Sand
Sand
Sand
Sand
Sediment
parameters.
5
7
14
19
5
14
18
19
5
10
28
5
10
19
28
12
12
13
14
6
22
22
13
23
22
15
21
14
17
14
14
14
14
16
16
16
16
16
Temp.
P3
7.7
7.7
0.2
0.2
28.5
28.5
28.5
28.5
37.3
37.3
37.3
0.4
0.4
0.4
0.4
0.1
1.2
1.2
1.3
15.7
1.2
2.2
26.8
19.4
18.1
18.1
-
org. c
m4
0.93
0.93
0.42
0.44
0.94
0.95
0.93
0.94
0.95
0.95
0.95
0.52
0.46
0.48
0.50
0.81
0.94
0.51
0.56
0.90
0.93
0.93
0.56
0.56
0.53
0.63
0.86
0.60
0.59
0.97
0.97
0.96
0.97
0.77
0.74
0.41
0.43
0.66
0.96490
0.96-0.90
0.46-0.40
0.46-0.42
0.96-0.9 1
0.96-0.91
0.96-0.91
0.96-0.9 1
0.97-0.94
0.97-0.94
0.97-0.94
0.56-0.48
0.50-0.44
0.63-0.43
0.534.46
0.85-0.78
0.95-0.92
0.53-0.49
0.57-0.55
0.91AJ.88
0.97-0.90
0.97-0.90
1.36
1.44
1.83
2.10
1.36
1.83
2.04
2.10
1.36
1.62
2.57
1.36
1.62
2.10
2.51
1.75
1.75
1.78
1.83
1.41
2.25
2.25
1.78
2.31
2.25
1.86
2.18
1.83
1.99
1.83
1.83
1.83
1.83
1.94
1.94
1.94
1.94
1.94
1.2kO.l
1.2kO.l
l.OkO.4
l.OkO.2
0.950.2
l.lkO.1
1.3kO.2
1.4kO.l
0.8kO.l
1.0t0.1
1.6
0.7kO.2
0.9kO.2
1.2t-0.1
1.4
1.1
1.1
1.1
l.lkO.2
l.OkO.1
1.7+0.2
1
3.1 kO.4
1.3
1.3
1.3
1.3
1.7
1.7
1.3
1.2
1.2
1.3
1.4
1.4
1.1
1.4
1.7
1.8
1.1+0.1
1.2kO.l
0.9
1.1
0.9
1.2
1.4
1.5
0.8kO.l
l.OkO.2
1.6
0.7
1.0
1.2
1.6
1.1
1.2
1.8
2.1
1.5
1.7
1.6
1.5
1.7
1.6
1.6
2.0
1.8
1.8
1.8
1.6
1.6
1.6
1.7
1.4
1.3
1.4
-
1.8
1.7
1.4
1.3
1.1
1.6
1.6
1.5
1.4
1.3
1.4
1.8
1.4
1.1
1.1
-
1.2
1.2
2.0
1.9
1.5
1.5
1.5
1.4
1.7
1.6
1.6
2.0
1.6
I.8
1.6
-
B
3
Notes
3.2
vJ 2.4 “E
*
+
. * Do
0 Ds (asphyxiated)
+ D,bunsphyxiated)
1.6-
5
;
0.8 0.0 !
0
I
I
20
10
temperature (“C)
I
30
Fig. 1. Diffusion coefficients for ‘H,O measured in
the pore waters of various freshwater sediments at
varying temperatures and in asyphyxiated and unasphyxiated cores (Table 1). Line is the predicted value
for the diffusion coefficient of “H,O in particle-free
water (D,, Eq. 5).
except one, fell below the predicted line for
D, (Fig. 1) and generally increased with increasing temperature.
The importance of temperature is apparent in sediments from a 3-m site in Lake
Lcosdrecht (Fig. 2), where D, and D, increased by a factor of > 2 in the warm summer months (28”C), compared to winter values (3”(Z), consistent with Eq. 5. Bioirrigation
339
had no effect, and 0, and D, values were
essentially the same (Fig. 2). At a deep-water (10 m) site in Lake Vechten, where the
temperature ranges from only 3” to 7°C 0,
and D, changed very little.
In Lake Vechten sediments (1 O-m depth),
Chaoborus larvae were the dominant macrofauna, with densities varying between 400
and 1,200 ind. m-* during the year. In sediments from Lake Loosdrecht (3-m depth),
Chironomidae larvae and Oligochaeta were
the dominant macrofauna, varying from 500
to 1,600 ind. mm2throughout the year. Lake
227 was the only lake in which macrofaunal
activity appeared to increase D, over D,. In
this eutrophic Canadian Shield lake, bioirrigation due to high densities of Chironomidae larvae (14,000+4,000
ind. m-*) resulted in a very high value of D, (3.1 kO.4
x 1O-5 cm2 s-l, Table 1).
By comparing D, and D, (as measured in
sediments) with D, (as measured in pure
water; Wang et al. 1953) for a given temperature, we could examine the effects of
factors other than temperature (tortuosity,
viscosity, bioirrigation)
(Table 1). In highorganic sediments, 0,/D, values averaged
1.5kO.2 (*SD) over a porosity range 0.6-
Fig. 2. Diffusion coefficients in unasphyxiated (De-closed symbols) and asphyxiated (D,-open
sediment cores from Lake Loosdrecht and Lake Vechten in 1986.
symbols)
Notes
340
2.2
+ +
1.0 !
0.4
0.8 ]I
•I WA
+ Do’%
. cl
11
0.6
I
I
-I
0.6
0.8
1.0
porosity
0.0
0.1
0.2
0.3
-log porosity
0.4
Fig. 3. Do/D, and Do/D, values as a function of
porosity for various lake sediments (Table 1).
Fig. 4. Relation between log F and -log porosity
for various lake sediments (Table 1, r2 = 0.87).
0.95. The DO/D, values of the same sediments were not significantly
different
(1.4kO.2). In the low-organic sediments,
average DO/D, values (1.8 +O. 1) were also
not significantly different from DO/D, values
(1.6kO.3) over a porosity range of 0.410.77. For all DO/D, and DO/D, values together, the mean in high-organic sediments
was 1.4kO.2 and in low-organic sediments
it was 1.7kO.3.
Although there were differences in DO/D,
and DO/D, as noted above, the overall range
of mean values was very small (1.4-1.8).
For the asphyxiated data only, the relationship with porosity was o2 = -0.734 + 2.17,
r = -0.71, n = 21. For all data it was I!?~=
-0.474 + 1.91, r = -0.43, n = 53 (Fig. 3).
This change in DO/D, with porosity was statistically significant but was smaller than
hypothesized for marine sediments based
on indirect measurements of d2 and assuming that DO/D, = 13~(Eq. 2) (Manheim and
Waterman 1974; Lerman 1978; Ullman and
Aller 1982) (Fig. 3).
In our experiments the movement of 3H20
into the sediment was primarily by molecular diffision with little contribution from
bioirrigation.
It appears that up to 1,600
ind. rnp2 (Chironomidae,
Chaoborus, Oligochaeta) have only a slight effect on solute
movement (Lake Loosdrecht, Lake Vechten), but 14,000 ind. mm2(Lake 227) do have
a large impact. Macrofaunal increases of
water or solute movement across the sediment-water interface have generally been
demonstrated in more densely populated
lake (Fisher 1982; Krantzberg 1985; Mati-
soff et al. 1985; Tatrai 1986) and marine
(Rutgers van der Loeff et al. 1984; Dicke
1986) sediments. For example, Dicke (1986)
measured an integrated annual mean increase in 3H,0 movement by bioturbation
of 3.4 X in the marine sediments of Boknis
Eck with high densities of Mollusca, Polychaeta, and Crustacea. In the sediments of
a tidal flat in the Dutch Waddenzee fauna1
activity increased 3H,0 transport by 1.85.4 x (Sweerts unpubl.).
Another outcome of our experiments was
the weak relationship between DO/D, and 6
(Fig. 3). Our results suggest that in highporosity sediments, D, might be decreased
by increased viscosity of the pore water due
to dissolved or gellike organic compounds.
This effect would have an opposing effect
on the expected decrease in apparent tortuosity at higher porosities. Navari et al.
(1970) found that the diffusion rate of oxygen decreases in water containing various
organic compounds. Revsbech (1989) has
found that 4 D, was 5 1% of DOin a diatom
biofilm that had an extracellular matrix of
polysaccharides and a porosity of 0.9 5. The
lack of much decrease in DO/D3in the highporosity sediments might also be related to
how very small grains are dispersed in highporosity sediments.
In marine sediments 19~has been estimated from measurements of differences in
electrical resistivity in separated pore water
(R,) and in whole sediment(R) (Archie 1942;
Andrews and Bennett 198 1). The ratio of
these two resitivities is called the formation
factor Q and
Notes
F = R/R,.
(6)
Tortuosity squared is then estimated from
the geometrical relationship (Ulmann and
Aller 1982):
t12= C#J
x F.
(7)
F is empirically
related to porosity (Manheim and Waterman 1974; Bemer 1980).
There is a significant relationship between
log F and -log 4 from our measurements
of DO/D, and DO/OS(Fig. 4). It differs from
the marine results, however, in that the slope
of the line is 1.2 compared to 1.8 for the
marine case (Bemer 1980). This result suggests that F, if it could be measured, would
be a good predictor of DO/D, in freshwater
sediments at low porosities and would tend
to underestimate
it at higher porosities
(> 0.9). The marine data sets include mostly
low-porosity sediments (we do not know of
any study investigating sediments with porosities > 0.9). If it is true that high-porosity
freshwater sediments have increased DO/D,
due to the elevated viscosity in the pore
water, it could explain part of the lower slope
of our data.
Jean-Pierre R.A. Sweertsl
Limnological Institute
Rijksstraatweg 6
363 1 AC Nieuwersluis
The Netherlands
Carol A. Kelly
Department of Microbiology
University of Manitoba
Winnipeg R3T 2N2
John W. M. Rudd
Ray Hesslein
Freshwater Institute
50 1 University Crescent
Winnipeg, Manitoba R3T 2N6
Thomas E. Cappenberg
Limnological Institute
Rijksstraatweg 6
I Present address: Delft Hydraulics, Water Resourcesand Environment Division, P.O. Box 177,2.600
MH Delft. The Netherlands.
341
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Submitted: 17 June 1988
Accepted: 15 March 1990
Revised: 19 November 1990
36(2), 1991, 342-354
0 1991, by the American of Limnology and Oceanography, Inc.
The post- 1979 thermohaline structure of the Dead Sea
and the role of double-diffusive mixing
Abstract- After centuries of meromixis (yearround stratification with a permanent halocline),
the Dead Sea has passed through two distinct
stages in the last decade: first a 4-yr meromictic
stage and then a holomictic stage. In the first
stage, classic one-dimensional processes dominated. In the second stage,three different regimes
operated in a seasonal cycle: salt precipitation in
spring and early summer, double-dil-htsive mixing in late summer and autumn, and vertical mixing in winter. During the second (holomictic) stage
the Dead Sea as a whole also underwent secular
changes: a gradual change in the salt composition
of its brines, an increase of salt concentration,
and a gradual heating.
Not many large natural hypersaline bodies of water exist today from which evaporites precipitate. Yet there are many examples from the geological past in which
sea-level changes resulted in a concentration of saline water in isolated basins to the
point at which evaporites were deposited,
sometimes in massive layers. Very little is
Acknowledgments
Thanks to skipper Moti Gonen of the RV Tiulit and
his crew, who handled the fieldwork. Valuable insight
was gained during discussions with J. R. Gat. This
paper benefited from remarks by C. H. Mortimer.
Comments by A. Wuest improved the manuscript.
This work was supported in part by a grant from the
Israeli Ministry of Energy and Infrastructure.
known about the hydrographic processes
that may cause deposition of evaporites; the
present work is an attempt to observe how
a hypersaline column of Dead Sea water
behaves under circumstances that could be
similar in some respects, for example, to
those in the Mediterranean during the Miocene.
Until 1979, the dead sea had been stably
stratified (meromictic) for several centuries.
Since the historical 1979 overturn of its water column (Steinhorn 1985), there have
been two distinct stages: a 4-yr meromictic
stage from the beginning of 1979 to the end
of 1982 and a holomictic stage from the
beginning of 1983 on. The period covered
here (1979-1988), except for winter 1980
(see Anati et al. 1987, figure 4) was characterized by a continuing water deficit, that
is, by conditions favoring holomixis marked
by regular winter overturn with consequent
changes in bottom-water properties.
The post-1979 water column is composed, to a first approximation, of two main
layers (Steinhorn 1985; Anati et al. 1987):
an upper layer between the surface and 1530-m depth and a lower layer extending to
the bottom (at -320-m depth). The thin
transition layer is disregarded here, as its
complex fine structure and its very inter-