LI, HONG-CHUN, TEH-LUNG KU, LOWELL D. STOTT, AND

Limnol. Oceanogr., 42(2), 1997, 230-238
0 1997, by the American Society of Limnology
and Oceanography,
Inc.
Stable isotope studies on Mono Lake (California). 1. al80 in lake sediments as
proxy for climatic change during the last 150 years
Hong-Chun Li, Teh-Lung Ku, and Lowell D. Stott
Department of Earth Sciences, University
of Southern California,
Los Angeles, California 90089-0740
Robert F. Anderson
Lamont-Doherty
Earth Observatory, Columbia University,
Palisades, New York 10964
Abstract
Mono Lake is a hypersaline, alkaline lake in the Mono Basin located #atthe easternbase of the Californian Sierra
Nevada. Its lake-level history since 1912 has been recorded instrumentally, showing the decline of lake-surface
elevations initiated by the 1941 artificial diversion of stream inflow. We have made high-resolution oxygen isotopic
measurementson the total carbonate fraction of lake sediments and shown that the S’*O record parallels the lakelevel fluctuations rather well. The measurementswere carried out on sedimentsthat had been leached with deionized
water to isolate the isotopic signals of the calcium carbonate from those of pore water and water-soluble carbonate
salts in the sediment. Extending the S’“O record back in time, we found that lake level changed markedly during
the past 150 yr, reflecting climatic variations and resultant runoff fluctuiitions. Lake levels were high around 1845,
1880, and 1915 and low around 1860, 1900, and 1933. This study demonstratesthat closed-basin lake sediment
S180provides an effective means of probing past precipitation variations in arid to semiarid regions such as the
Great Basin in the western United States.
Numerous attempts have been made to reconstruct paleolake levels on the basis of geological and stable isotopic
evidence from which paleohydrological
conditions have
been inferred. Studies on closed-basin lakes in semiarid and
tropical zones have contributed greatly to paleoclimatic research, particularly for the Late Pleistocene and the Holocene (e.g. Street-Perrott and Harrison 1985; Chivas et al.
1986; Gasse et al. 1991). Water entering closed-basin lakes
via rain and runoff leaves by evaporation. Hence, fluctuations in lake level and volume should serve to indicate
changes in the moisture budget associated with climatic
change. Mono Lake, located in eastern California about 480
km north of Los Angeles (Fig. l), is the only terminal lake
in the Mono Basin. Its present ratio of surface area (150
km2) to catchment area (2,072 km2) is -7 : 100, in contrast
to the ratio of -33 : 100 for Lake Russell during the last
glacial period (Russell 1889; Lajoie 1968; Mono Lake Committee 1981). Since 1912, the levels of Mono Lake have
been measured by the U.S. Geological Survey (USGS) and
the Los Angeles Department of Water and Power (LADWP)
(Blevins et al. 1987). Since its historic high in 1915, the
water level within the lake has fallen as a result of climatic
changes. The lake level fall has been exacerbated by the
diversion of inflow to the lake by LADWP ‘since 1941 and
has caused a further increase in salinity and alkalinity. DurAcknowledgments
We thank D. E. Hammond and S. P. Lund for insightful discussions. D. R. Cayan kindly made available to us his unpublished data
on Nevada City precipitation and streamflow. Comments made by
L. V. Benson and an anonymous reviewer greatly helped our presentation.
This work was supported by NSF grants ATM 93-03587 (T.-LX.)
and OCE 93-14192 (L.D.S.), and by NOAA grant NA16-RC-0084
(R.EA.).
230
ing the diversion period the lake level fell - 13 m (from an
elevation of 1,956 m to its historic low stand of 1,942 m)
between 194 1 and 1982. During the high runoff years between 1982 and 1986 (a wet period caused by strong 19821983 El Nifio), the lake level rose by -2.4 m to 1,944.6 m
before falling again.
The fluctuations in lake level and volume of a closedbasin lake such as Mono Lake may be recorded in the 6’*0
of authigenic carbonates deposited within the lake sediments. This is feasible because the lake S180 monitors the
hydrologic balance between the isotopically distinct inflow
and outflow. For example, Benson (1994) monitored the 6’*0
of lake waiter from 1985 to 1992 in Pyramid Lake, Nevada,
another closed-basin lake in the Great Basin. His results
showed that although the surface-water &*O exhibits seasonal variations, reflecting varying amounts of fresh water
discharged to and lost from the lake, the yearly S180 trend
over the study period correlates well with that of the lakevolume change (lake-water oxygen isotopic values become
lighter as volume increases). This observation suggests that
when the lake undergoes a volume change, the hydrological
balance is the most important factor influencing the lake
6’*0, although other factors such as stratification, relative
humidity, ,air temperature, and the atmospheric @*O may
also play a role. If carbonate minerals form in a lake in
isotopic equilibrium with lake water, their 6’*0 values can
be used toI reconstruct lake volume fluctuations. This approach to lake history investigations forms the basis of several previous studies (Lister et al. 1991; Phillips et al. 1992;
Johnson et al. 1991). However, it would be reassuring if a
comparison could be made between the reconstructed lakelevel changes with instrumental records. Such a comparison
can be conducted for Mono Lake because, as mentioned
above, its lake-level record is available for the past 80 yr
through the measurements of LADWP and USGS (Blevins
Stable isotopes in Mono Luke
231
lake volume (km3) (R2 = 0.99). Therefore, volume change
can be directly read from lake-level change. We describe
‘lake-level fluctuations reconstructed from S180 records of
carbonates deposited within the lake. From these records we
infer lake-volume fluctuations.
Methods
Fig. 1. Map showing Mono Basin and sample locations (adapted from the Mono Lake Committee 198 1).
et al. 1987). In this paper, we describe our efforts in the use
of the sediment S’*O record to reconstruct lake-level fluctuation and climatic change in the Mono Basin during the
last 150 yr.
Mono Lake hydrology
The moisture input to the hydrologically closed Mono Basin is principally controlled by the amount of precipitation
it receives, which, in a larger scheme, is tied to the meteorology of the Great Basin (Houghton 1969; National Research Council 1987). Precipitation in the Great Basin is
dominated by Pacific winter storms. Winter snows and rainfalls feed the stream runoff from the Sierra Nevada, which
constitutes -70% of the water input to Mono Lake. The
remaining 30% is roughly divided between direct precipitation onto the lake surface and groundwater seepage (Blevins et al. 1987). As a result, surface runoff and rainfall in
the Mono Basin assume a linear relationship (Fig. 2a). Because output of the lake water is by evaporation only and
the annual rate of evaporation is relatively constant (Blevins
et al. 1987), lake volume also varies linearly with runoff to
the lake (Fig. 2b). Figure 2 shows that closed-basin lakevolume fluctuations could effectively reflect changes in basinwide precipitation, because evaporation, being highly temperature dependent, remains rather constant on annual-todecadal time scales. The relationships shown in Fig. 2 provide a framework
to probe paleoprecipitation
and
paleohydrology in a closed basin. Namely, if a means can
be found to reconstruct the paleolake volume (or level), then
past changes in the surface runoff and precipitation may be
inferred from the reconstruction. According to the measurements of volume and elevation of Mono Lake (Blevins et
al. 1987), there is a strong linear correlation between lake
volume and lake level: lake level (m) = 1,929.9 + 4.8515
Sample collection and core chronologies-Core ML9 1
FC3 was collected in 1991 from the deepest part of Mono
Lake at 39 m (Fig. 1) with a rectangular freeze corer similar
to the type described by Crusius and Anderson (1991). Sediments were kept frozen before returning to the laboratory
for processing. The core is 76 cm long and exhibits fine
laminations throughout (Fig. 3). Its chronology can be determined from the distributions of 210Pband 239+240Pu
in an
adjacent freeze core (ML91 FC6) taken on the same day and
at the same location. The laminar features of the two cores
are correlative and indicate similar sedimentation histories.
The average sedimentation rate of the cores as determined
from 210Pband 239+240Pu
is 0.65 cm yr-’ (Fig. 4a,b). There is
a change in the sedimentation rate beginning at -35 cm
depth in the core (Fig. 4a), from -0.75 cm yr-1 above to
0.4 cm yr-1 below. The 35 cm horizon corresponds to the
time when substantial changes in runoff discharge to the lake
occurred due to the stream diversion by LADWF? The 210Pb
profile does not allow annual resolution nor does it give
overall chronology to better than + 10%. X-ray diffraction
(XRD) and scanning electron microscope (SEM) analyses
show that major minerals of the sediment consist of calcite,
halite, and quartz. Downcore XRD analyses (at l-cm intervals) showed that the carbonate phase in the sediments of
the core is mostly calcite, except near 39 cm where aragonite
is the dominant phase. As Li (1995) has shown, calcite and
aragonite are the primary mineral precipitates in Mono Lake
sediments, with aragonite mostly occurring in surface sediments from water depths shallower than 16 m in the lake.
Because FC3 core was taken from the deepest part of Mono
Lake at 39 m, aragonite content in the sediments of this core
should be very low.
Stable isotope measurements-Core ML91 FC3 was sectioned at 0.4-cm intervals under frozen conditions and
- 100-400 pg of the sediment from each interval was sampled for 6180 analyses. The material, after oven-drying overnight at 7O”C, was converted to CO, in 100% phosphoric
acid at 9O”C, using an “autocarb” device attached to a VG
Prism II triple-collecting mass spectrometer. Replicate analyses of a calcite standards run together with the samples
gave reproducibilities of t0.26%0 for 6’*0. The method of
S1*O analysis for water samples is the CO,-H,O equilibrium
method (Epstein et al. 1953). The isotope ratios are reported
as the per mil deviation from the PDB standard for carbonate
material and from the SMOW standard for water samples.
The measurements were first made on CO, released from
acidification of 182 bulk sediment samples. Although the
6180 evolution trend of these unwashed, bulk samples correlates well with the historic lake volume changes, the al80
values cannot be explained by calcite precipitation in isotopic equilibrium with the lake water. Therefore, we then
Li et al.
0
20
40
60
80
100
120
Rainfall in Mono Basin (cm yi’)
washed salmples were oven-dried overnight at 70°C and
weighed again to determine the salt content. Table 1 shows
the S180 vailues of the washed samples.
Relative to the washed samples (8180 from +0.3 to
-6.0%0), the unwashed bulk samples show systematically
higher 618Clvalues (+ 13.0 to -4.5%0), the differences being
well beyond the analytical errors. Heavy S180 values of the
unwashed samples may be caused by Na,CO, that formed
as the samples were dried. The aI80 of the rapidly forming
Na,CO, is expected to be close to that of dissolved bicarbonate ions whose 6180 values are -8%0 heavier relative to
the equilibrated water (Craig 1965; O’Neil et al. 1969). Li
(1995) has noted that the 6180 offset between the washed
and unwashed samples becomes larger when the Na,CO,
concentration increases and when the CaCO, concentration
decreases in the bulk sediments. The NqCO, concentration
reflects roughly the carbonate alkalinity of porewater and the
CaCO, concentration is related to the input of Ca2+; both are
functions of lake volume. This explains the anticorrelation
between cV0 and lake-volume change, even though the S180
of unwashed samples does not reflect that of lake water.
Results and discussion
Y. = -0.1345
+ 1.0811x
R2 = 0.83
-,.,2J-J-LJ,,
0
0.1
0.2
0.3
Runoff to Mono lake (km3 yr-‘)
Fig. 2. [a.] Correlation between rainfall and stream runoff in
the Mono Basin from 1935 to 1985, at the Gem Lake Station located 20 km south of Mono Lake in the Sierra Nevada. [b.] Correlation between stream runoff to Mono Lake and Mono Lake volume change from 1935 to 1985, based on measurements made by
LADWP; the volume change refers to the volume difference between two consecutive years. Data source: Blevins et al. 1987.
made 61x0 analyses on 97 sediment samples (selected from
the 182 samples) after they had been leached with deionized
water buffered at a pH of 7 in order to dissolve the interstitial
salts. The leaching was done as follows. About 25 ml of
water was mixed with a known quantity (-0.5 g) of the
sample. The mixture was shaken, centrifuged, and the supernatant discarded. The leaching procedure was repeated
three times until pH of the leachate was close to 7. The
Carbonate precipitation in sur$ace water-Authigenic
carbonates in lake sediments can be inorganic or biogenic.
Although direct experimental data are lacking, available empirical evidence suggests that oxygen isotopic equilibrium
between lake water and lacustrine carbonates exists (e.g.
Gasse et al. 1987; Turner et al. 1983; Johnson et al. 1991).
The carbonate aI80 is a function of both temperature and
lake-water S180. The 6180 of the washed sample (Scalcltc)in
the topmost layer (representing the deposit of summer 1991)
of core ML91 FC3 is - 1.05%0. XRD analyses showed that
the carbonate phase in this sample is calcite.
We have measured the S180 of Mono Lake surface water
(a,,,,,) collected in summer 1991 to be -O.l%, (SMOW)
(Table 2). Assuming water-carbonate isotopic equilibrium,
we calculate a temperature of 22°C for the lake carbonate
formation. ‘The paleotemperature equation used, ~=calcite
- &,,,,,,
= 3.95 - 0.232T (“C), is an average of the rather tight results of previous work on the calcite-water system (Epstein
et al. 1953; O’Neil et al. 1969; Grossman and Ku 1986). The
calculated temperature indicates that the calcite precipitated
from Mono Lake surface waters during warm seasons because only during June-October do surface waters reach
temperatures of 22°C (R. Jellison unpubl. data 1991). This
observation is consistent with studies carried out on Walker
Lake (Newton and Grossman 1988), Lake Ontario (Schelske
and Hodell 1991), and several other lakes (Kelts and Hsii
1978; McKenzie 1985) all showing a summer precipitation
record. In #spring, streams supply Ca2+ to the lake. Precipitation of CaCO, may occur in the summer and autumn as a
result of increase of Ca2+ supply, loss of CO, from the epilimnion by photosynthesis and/or gas exchange with the atmosphere, increase of water temperature, and increase of nucleation surface due to enhanced biomass.
PO in authigenic carbonates as proxy for lake-level jluctuation-For relatively small lakes such as Mono Lake, the
,
Fig 3. Photograph of the freeze core ML91 FC3 showing the laminated nature of rhc bedmxma.
Top of the core is at upper left. Scales shown on the sediments are in centime!ers.
lake-water 61a0 depends largely on the difference in volume
and rS”O between the input and output waters, the summer
temperature remaining relatively constant. The lake 61x0 of
a closed-basin lake progressively gets heavier as evaporation
proceeds until an isotopic steady state is achieved. Thus 6180
of input waters (surface runoff + direct precipitation) is always lighter than that of the lake water. For example, the
mean 6’80 of input waters to Pyramid Lake is about - 10%0
@MOW) as opposed to -OS%0 for the lake water (Benson
1994). In Walker Lake, Nevada, the input water has an average 6’“O of ~ 14.1%0, whereas the lake SIX0 is 2X%:, (Newton and Grossman 1988). The larger the amount of freshwater inflow relative to evaporation, the lighter the lake SIXO
will become. We have measured the 6’“O of mput warerb
(streams, snow melt, and springs) to Mono Lake and obtained an average of 14.6 -+ 1.0%~ (Table 2). The following
estimation shows the response of Mono Lake water S’80 to
changes in the lake level. The 1982-1983 El Niiio event
caused Mono Lake to rise 2 m, increasing the lake storage
(volume) by -12% (Blevins et al. 1987). It is assumed that
prior to the event Mono Lake water had a ,YxO value of
-O.l%o. Mixing this water with the input water that had a
PO value of 14.6%0 would lower the lake 6180 by about
1.7%-a signal that is readily detectable.
Our estimation ignores the vapor exchange across the
lake-atmosphere interface, which can play a very important
234
Li et al.
Excess 2’aPb (dpm g-l)
1
10
20
-t
(a)
-t
30
40
50
60
E
-+-
f
700
239+24(&(dpm g“)
30F-,i--
+++I-
Fig. 4. [a.] Excess 210Pbdistribution in core ML91 FC6, collected on the same day from the site where core ML91 FC3 was
retrieved. Average sedimentation rate estimated by fitting data from
the top 35 cm is 0.75 cm yr-I, and that by fitting data from 35 to
profile on ML91 FC6. The peak
56 cm is 0.4 cm yr-I. [b.] 279+240Pu
at 19.5 cm is assumed to correspond to year 1963, yielding an
average sedimentation rate of 0.7 cm yr-I for the top 20 cm of the
role in affecting lake al80 and forcing it to reach a steadystate value, as 618Oof the atmospheric vapor has a relatively
constant value. For a stabilized lake level or when lake volume changes slowly, lake S180 is mainly controlled by vapor
exchange and there will be no correlation between #*O and
lake volume. Only when the lake level experiences relatively
rapid fluctuations does the vapor exchange effect diminish,
so that S180 and lake volume covary (Lister et al. 1991; Cat
1995; Li 1995). Mono Lake experienced large, rapid changes
in volume during the period under study allowing us to use
the S180 record to reconstruct lake volume changes.
In Fig. 5 we compare the sediment al80 record from Mono
Lake with the historic lake-level record. The historic Mono
Lake levels from 19 12 to 199 1 were recorded by LADWF
and USGS (Blevins et al. 1987; USGS Water Resour. DataCalif. V. l-5, 1986-1991). In Fig. 5, the relative scales for
the two ordinates are set arbitrarily to facilitate the comparison that shows a general anticorrelation between lake sediment 6180 and lake level. However, we emphasize that one
cannot read the lake volume directly from the aI80 record,
which has been pointed out by Benson (1994) in his Pyramid
Lake study. Hydrological balance in a closed basin may persist for several years leading to a near-stable lake level. As
Lister et al. (1991) and Benson et al. (1996) have stated, the
S180 of a closed-basin lake depends on the rate of change in
lake volume. When the rate of change in lake volume is
small and the lake level is stabilized, the S180 of lake water
will approach a steady-state value, the magnitude of which
is controlled by vapor exchange across the lake-atmosphere
interface and the S180 of input water but is independent of
lake volume. Furthermore, the rate of change in lake volume
is also a function of initial lake volume. For a given input
flux, the smaller the initial lake volume, the larger the rate
of lake-volume change will be. Therefore, a return to a former lake level due to changes in climatic conditions does
not mean the al80 should return to the former value associated with. the former lake level (Benson 1994).
Quantitative reconstruction of lake volume (level) must
be obtained from a dynamic isotope mass-balance model.
Nevertheless, we use al80 change to indicate lake-level fluctuation, that is, a sharp (say a change of >1%0 over 10 yr)
decrease in. aI80 reflects a rapid rise in lake level, and a sharp
increase in S18Orepresents a strong decline of lake level. A
period (> 30 yr) of relatively constant S180 may result from
achievement of a hydrologic and isotopic steady state within
the closed-basin lake. In addition, aI80 values of individual
data points may be affected by seasonal lake stratification
and inhomogeneity of water temperature and S180. As Benson (1994) pointed out in his Pyramid Lake study, freshwater
input during high-streamflow years (e.g. 1986) lowered the
S180 in surface water (epilimnion) by 1.0%0 relative to that
in deep water (hypolimnion). In dry years (e.g. 1991), however, al80 of the epilimnion was enriched by 0.5%0 due to
surface evaporation. Hence, it is better to use a 5-yr running
average aI80 curve to infer lake-level fluctuations.
Large offsets between the two curves exist around 1933
and the interval 1972-1980-two
time periods each near the
end of a prolonged drop in lake level. In Fig. 5, a marked
decrease in S180 from 1900 to 1910 implies a strong transgression in lake level that was previously demonstrated by
Harding (unpubl. rep. 1962). The al80 values reached a minimum value at the historic high stand near 19 15. A strong
decline of lake level between 1920 and 1933 is reflected in
a sharp increase in the S180 values. In addition to evaporation, those heavy S’*O values may also have been caused by
intensified stratification of aI80 in the lake water. The lakelevel drops reflected in the aI80 values prior to 1933 (Fig.
5) signify a prolonged period of dry climate in the Mono
Basin, which could have led to the higher epilimnion 6180
values that are recorded by the carbonates. Lake level increased from 1933 to 1935 and then remained relatively stable, corresponding to a wet climatic regime from around
1935 to 1943 (see Fig. 7). The 1.8%0 decrease in Si80 be-
235
Stable isotopes in Mono Lake
Table 1. The measured SIR0 data of washed samples in core ML91 FC3.
Sample
Depth
(mm)
Year
6’80
(%o PDB)
FC3-4
FC3-5
FC3-6
FC3-7
FC3-8
FC3-9
FC3-10
FC3-11
FC3-12
FC3-13
FC3-14
FC3-15
FC3-16
FC3-17
FC3-23
FC3-24
FC3-25
FC3-27
FC3-28
FC3-29
FC3-34
FC3-35
FC3-36
FC3-37
FC3-38
FC3-39
FC3-42
FC3-44
FC3-46
FC3-47
FC3-48
FC3-49
FC3-53
FC3-57
FC3-58
FC3-59
FC3-61
FC3-63
FC3-66
FC3-67
FC3-68
FC3-69
FC3-73
FC3-74
FC3-75
FC3-76
FC3-78
FC3-80
10.0
12.0
14.5
18.0
23.5
29.5
34.5
39.5
44.5
49.5
54.5
59.5
64.5
69.5
99.0
106.5
114.0
125.0
129.0
133.0
151.0
154.0
159.0
163.0
167.0
171.0
183.0
191.0
199.0
203.0
207.0
211 .o
227.0
243.0
247.0
251.0
259.0
267.0
279.0
283.0
287.0
291.0
307.0
311.0
315.0
319.0
327.0
335.0
1989.4
1989.1
1988.6
1987.9
1987.1
1986.4
1985.7
1985.1
1984.4
1983.7
1983.1
1982.4
1981.7
1981.1
1976.8
1975.8
1974.3
1973.8
1973.3
1970.9
1970.5
1969.8
1969.3
1968.7
1968.2
1967.7
1966.1
1965.0
1963.9
1963.4
1962.9
1962.3
1960.2
1958.1
1957.5
1957.0
1955.9
1954.9
1953.3
1952.7
1952.2
1951.7
1949.5
1949.0
1948.5
1947.9
1946.9
1945.3
-2.24
-1.71
- 1.45
-1.59
-2.29
-1.96
-1.63
-1.70
-2.76
-1.54
-2.07
-1.81
-1.05
-1.72
-2.03
-1.58
-0.37
-0.85
0.30
-0.50
-3.00
-2.50
-1.78
-1.57
-1.26
- 1.46
-0.50
- 1.20
-0.58
-0.14
-0.85
-0.58
-1.50
-1.39
-2.13
-0.85
-2.00
-2.80
-3.15
-3.37
- 1.70
-2.65
-2.94
-2.25
-2.01
-2.10
-3.30
-3.50
tween 1933 and 1945 may be caused by a combination of
three processes: increased freshwater discharge between
1933 and 1935; vertical mixing of isotopically light hypolimnion water with heavy epilimnion water inherited from a
previous period of lake stratification; and approaching a
lighter steady-state S80 value after 1935 when the lake level
was stabilized. Beginning in 1941 when the stream diversion
started in the Mono Basin, the Mono Lake had experienced
an almost continuous decrease in water level until 1982. The
Sample
FC3-82
FC3-86
FC3-87
FC3-88
FC3-89
FC3-91
FC3-96
FC3-97
FC3-100
FC3- 105
FC3-106
FC3-108
FC3-109
FC3-111
FC3-113
FC3-115
FC3-116
FC3-117
FC3-118
FC3-119
FC3-120
FC3-124
FC3-126
FC3-128
FC3- 129
FC3-134
FC3-137
FC3-138
FC3-141
FC3-145
FC3-146
FC3-147
FC3-151
FC3-152
FC3-155
FC3-157
FC3- 159
FC3-161
FC3-163
FC3-165
FC3-166
FC3-167
FC3-168
FC3-171
FC3-173
FC3-176
FC3-177
FC3-181
FC3-182
Depth
(mm>
343.0
359.0
363.0
367.0
374.0
388.0
411.0
418.0
433.0
460.0
470.0
481.0
485.0
493.0
501.0
509.0
513.0
517.0
521.0
525.0
529.0
542.5
548.5
554.5
557.5
573.0
582.5
585.5
595.0
610.0
614.0
618.0
634.0
638.0
650.0
658.0
666.0
674.0
682.0
690.0
694.0
698.0
702.0
714.0
722.0
734.0
738.0
754.0
758.0
Year
6’80
(%o PDB)
1943.3
1939.3
1938.3
1936.6
1935.0
1931.3
1926.0
1924.0
1920.7
1914.0
1912.0
1908.8
1907.8
1905.8
1903.8
1901.8
1900.8
1899.8
1898.8
1897.8
1896.8
1893.7
1892.2
1890.7
1890.0
1886.0
1883.7
1883.0
1880.5
1876.6
1875.6
1874.6
1870.6
1869.6
1866.6
1864.6
1862.6
1860.6
1858.6
1856.6
1855.6
1854.6
1853.6
1850.6
1848.6
1845.6
1844.6
1840.6
1839.6
-4.50
-2.01
-2.58
-2.59
-1.37
-1.59
-4.28
-3.46
-4.50
-5.99
-5.89
-4.51
-4.59
-4.30
-3.80
-3.50
-2.92
-2.82
-2.44
-2.75
-2.87
-3.70
-2.87
-3.65
-3.96
-3.28
-3.50
-4.34
-5.50
-5.16
-3.59
-3.21
-3.95
-3.43
-2.96
-2.67
-2.50
-3.00
-3.50
-4.54
-3.54
-3.17
-3.82
-4.27
-4.50
-5.70
-3.72
-4.37
-4.36
6180 value continuously increased from 1945 to 1963 due to
evaporation prevailing over input. Decreases in PO at the
intervals of 1965-1969 and 1981-1986 are attributed to increases in lake level during these two periods. An apparent
offset between the two curves exists around the interval
1972-1980, i.e. lake level decreases while the PO also decreases. The cause for the offset in the interval 1972-1980
remains unclear.
The Fig. 5 comparison indicates that the WO of authi-
Li et al.
236
Table 2. S’*O measurements in waters from the Mono Lake Basin.
Sample
No.
Location
M-l
M-2
M-3
M-4
M-5
M-6
M-7
M-8
M-9
M-10
M-11
M-12
M-13
M-14
M-15
M-16
M-17
M-18
M-19B
Ellery Lake
Saddle Bag Lake
Saddle Bag Lake
Lee Vining Creek Hwy 395
Lundy Lake
Mill Creek
Wilson Creek Hwy 167
Wilson Creek lower level
Grant Lake
Rush Creek Hwy 395
Walker Creek Hwy 395
Rush reek lower level
Mono County Park
Mono County Park
Mono County Park
Old Marina
Old Marina
Navy Beach
Navy Beach
Sample type
Lake water
Lake water
Snow
Creek water
Lake water
Creek water
Creek water
Creek water
Lake water
Creek water
Creek water
Creek water
Spring water
Spring water
Spring water
Spring water
Spring water
Warm spring water
Mono lake water
Distance*
M-0
16
22
22
3.8
13
10
5.2
3.5
12.5
8.5
7.3
2.2
0.02
0.02
0.2
0.03
0.03
0.05
0
Elevation
W-0
2,907
3,048
3,048
2,067
2,714
2,256
2,048
2,013
2,173
2,09 1
2,073
1,964
1,946
1,946
1,946
1,946
1,946
1,946
1,945
S’“0
(%o, SMOW)
- 14.4
- 14.3
-12.1
- 14.3
- 14.7
-15.8
-14.6
-15.8
-13.7
-13.8
-16.0
-15.3
-16.2
-14.4
-17.5
-14.5
-14.0
-14.1
-0.1
* From the Mono Lake shore in 1991, at 1,944.6 m elevation.
genie carbonates from a closed-basin lake can be used as a
proxy for lake-level fluctuations in the past. The S80 record
suggests that for the past 150 yr or so, Mono Lake has had
three major cycles of lake-level fluctuations before the tributary diversion begun in 1941, with low stands occurring
around 1860, 1900, and 1933 and high stands around 1845,
1880, and 19 15. After 1940, pulses of high input of freshwater occurred around 1943, 1953, 1969, and 1983, as registered by the light S80 peaks superimposed on the continuous decline of lake level due to the stream diversion.
----OS-
measured 6”O
-
5-yr running ave.
-
lake level
0
,
I /I I
1840
1860
,_I
1880
1900
1920
1940
1960
1980
2000
Year (A.D.)
Fig. 5. Comparison of measured SIX0 and their 5-yr running
averages in washed samples of core ML91 FC3 with the historically
measured Mono Lake level. Scales for the two ordinates are set
arbitrarily to facilitate the comparison. S180 decrease corresponds
to lake level increase. Solid arrows indicate periods of high runoff
and high lake stands. Open arrows indicate periods of low runoff
and low lake stands. Note that the lake level continuously declined
between 1941 and 1981 largely due to the stream diversion, with
secondary peaks denoting high runoff periods during this interval.
Paleo-lake level, paleoprecipitation, and streamjlow history---We think the good correlation between S180 changes
and historic lake-level variations strongly support the notion
that sediment proxies can be used to extend the lake-level
histories back in time. However, we would like to know
whether lake-level variations on time scales of 10°.5-102 yr
reflect primarily climate signals of precipitation (P) or of
evaporation (E). Variations in the size of a closed-basin lake
are mainly controlled by changes in the effective wetness (P
- E). Although climatic variables such as precipitation, temperature, humidity, cloudiness, and wind speed all affect the
moisture budget, it can be argued that precipitation is the
dominant factor. This argument gains support from the
monthly records of precipitation, runoff, and evaporation
shown in Fig. 6 for a closed-basin lake in Oregon (Hostetler
1995). The records show that on an annual basis, river runoff
correlates well with precipitation (with slight delays in the
“peaking”’ of runoff due partly to melting of winter snows)
but not with evaporation, which remains rather constant. As
evaporation is highly temperature dependent, one may surmise that relatively constant average annual temperatures
impart little to runoff variations. Another reason for relatively constant evaporation in the Great Basin is probably
due to the cloudless skies that prevail in warm seasons when
evaporation occurs (Benson 1981). Hence, we submit that
for the Mono Basin changes in the effective wetness mainly
reflect changes in precipitation, at least on interannual-todecadal time scales. In the Mono Basin, relationships between lake volume, precipitation, and surface runoff, as illustrated in Fig. 2, also delineate this situation. These relationships are generally characteristic of arid and semiarid
regions where groundwater discharge is small compared to
streamflow. In summary, when precipitation rate increases
in a closed basin, runoff and lake storage in the basin will
increase and the lake S1*O will be depleted.
237
Stable isotopes in Moio Lake
6
Silvies
600
400
4c
b
1 c
s
200
Ed
0
;z"zi -200
.o z
gy
-400
:$ g
-600
JP
-800
6
i
-8
-10
-2
-4
-6
06F
04
i3'
8
&
Evaporation
4 c
1840
1860
1880
1900
1920
1940
1960
1980
2000
Year (A.D.)
1980 1981 1982 1983 1984
1985 1986 1987
YEAR
Fig. 6. Monthly records of precipitation, runoff, and evaporation in Harney-Malheur Lake, Oregon, 1980-1986. Vertical dashed
lines are for demarcations of water years. The surface elevation
graph shows both the simulated (solid line) and measured (triangles)
lake level for the same period (after Hostetler 1995).
In Fig. 7 we plotted the records of Mono Lake sediment
S’*O, the precipitation anomalies of Los Angeles and the
Mono Basin, and the precipitation and streamflow anomalies
of Nevada City (located at the western slope of the Sierra
Nevada, -200 km northwest of Mono Lake; see Fig. 1). As
we mentioned earlier, precipitation in the Great Basin is
dominated by Pacific winter storms. Li (1995) has shown
that the rainfall pattern in the Mono Basin is similar to those
of the longer records kept for Nevada City and Los Angeles.
The mean precipitations at these locations differ: 37.6 cm
yr- 1 in Los Angeles (1878-1991), 55 cm yr-1 in the Mono
Basin (1926-1985), and 135.7 cm yr-I in Nevada City
(1864-1994). However, the precipitation anomalies, i.e. deviations from the mean values, show a similar pattern that
can be matched by the 6180 variability in Mono Lake sediments (Fig. 7). The sediment aI80 also records the stream
diversion effect, hence the post-1941 values are distinctly
heavier. There appears to exist a phase shift of a few years,
with the precipitation curves leading the S’*O curve, due
probably to the water residence time of about 25 yr in the
Fig. 7. [a.] Comparison of SIX0 variations in Mono Lake sediments (washed) with precipitation and streamflow anomalies at Nevada City, California. [b.] Comparison of SIX0 variations in Mono
Lake sediments (washed) with precipitation anomalies in Los Angeles and Mono Basin. Data sources: Nevada City precipitation and
streamflow anomalies from D. R. Cayan pers. comm.; Mono Basin
precipitation anomalies from Blevins et al. 1987; Los Angeles precipitation anomalies from data on file at the National Climatic Data
Center, Asheville, North Carolina and compiled by the National
Oceanic and Atmospheric Administration.
lake (Blevins et al. 1987). The matching of the cycles tends
to break down prior to 1900. Whether this breakdown, as in
the case of the phase lag, is a consequence of age uncertainties associated with the sediment record remains to be
ascertained.
Conclusion
This study shows that the oxygen isotopic composition of
carbonate sediments in closed-basin saline lakes provides a
semiquantitative means of probing past precipitation and surface runoff variations in arid to semiarid regions such as the
Great Basin. Most of the authigenic carbonates deposited in
Mono Lake are found to form in surface water during warm
seasons of the year. In view of the small changes in interannual summer temperatures in the region, the effect of temperature on the carbonate 6180 is relatively constant. The
hydrological history of a hypersaline closed-basin lake (e.g.
Mono Lake) can be well deciphered by the #*O record obtained from the lake’s inorganically precipitated calcium carbonates. On annual-to-decadal time scales, the effective wetness (P - E) of a closed-basin lake is mainly reflected by
238
Li et al.
changes in precipitation. Therefore, a detailed study on sedimentary S180 records of closed-basin lakes offers a new
approach to acquiring high-resolution time series records of
precipitation variability. The feasibility of the approach has
been demonstrated in the S180 record of Mono Lake sediment, showing three major cycles of climatically induced
lake-level fluctuations during the past 150 yr.
References
BENSON, L. V. 1981. Paleoclimatic
significance of lake-level fluctuations in the Lahontan Basin. Quat. Res. 16: 390-403.
-.
1994. Behavior of the stable isotopes of oxygen and hydrogen in the Tiuckee River-Pyramid Lake surface-water system. 1. Data analysis and extraction of paleoclimatic information. Limnol. Oceanogr. 39: 344-355.
L. D. WHITE, AND R. RYE. 1996. Carbonate deposition,
Pyiamid Lake Subbasin, Nevada: 4. Comparison of the stable
isotope values of carbonate deposits (tufas) and the Lahontan
lake-level record. Palaeogeogr. Palaeoclimatol. Palaeoecol.
122: 45-76.
BLEVINS, M. L., E. L. COUFAL, L. M. WATANABE, M. R. GARCIA,
AND C. J. FOLEY. 1987. Mono Basin geology and hydrology.
LADWP, Aqueduct Div.-Hydrol. Sect. 108 p.
CHIVAS, A. R., I? DE DECKKER, M. NIND, D. THIRIET, AND G. WATSON. 1986. The Pleistocene paleoenvironmental record of lake
Buchanan: An atypical Australian playa. Palaeogeogr. Palaeoclimatol. Palaeoecol. 54: 21-43.
CRAIG, H. 1965. The measure of oxygen isotope paleotemperatures, p. 9-130. In E. Tongiorgi [ed], Stable isotopes in oceanographic studies and paleotemperatures. CNR, Italy.
CRUSIUS, J., AND R. E ANDERSON. 1991. Core compression and
surficial sediment loss of lake sediments of high porosity
caused by gravity coring. Limnol. Oceanogr. 36: 1021-1031.
EPSTEIN, S., R. BUCHSBAUM, H. A. LOWENSTAM, AND H. C. UREY.
1953. Revised carbonate-water isotopic temperature scale.
Bull. Geol. Sot. Am. 64: 1315-1326.
GASSE, E, AND OTHERS. 1987. Biological remains, geochemistry
and stable isotopes for the reconstruction of environmental and
hydrological changes in the Holocene lakes from North Sahara.
Palaeogeogr. Palaeoclimatol. Palaeoecol. 60: l-46.
AND OTHERS. 199 1. A 13,000-year climate record from
wektern Tibet. Nature 353: 742-745.
GAT, J. R. 1995. Stable isotopes of fresh and saline lakes, p. 139165. In A. Lerman et al. [eds.], Physics and chemistry of lakes.
Springer.
GROSSMAN, E. L., AND T-L. Ku. 1986. Oxygen and carbon isotope
fractionation in biogenic aragonite: Temperature effects. Chem.
Geol. 59: 59-74.
HOSTETLER, S. 1995. Hydrological and thermal response of lakes
to climate: Description and modeling, p. 63-82. In A. Lerman
et al. [eds.], Physics and chemistry of lakes. Springer.
HOUGHTON, J. G. 1969. Characteristics of rainfall in the Great Ba-
sin. Ph.D. thesis, Univ. Oregon. 292 p.
AND W. J. SHOWERS. 199 1. Paleoclimate of the past 4000 years as Lake Turkana, Kenya,
based on the isotopic composition of authigenic calcite. Palaeogealgr. Palaeoclimatol. Palaeoecol. 85: 189-l 98.
KELTS, K., ,%ND K. J. Hsii. 1978. Freshwater carbonate sedimentation, p. 295-323. In A. Lerman [ed.], Lakes: Chemistry, geology, physics. Springer.
LAJOIE, K. R. 1968. Lake Quaternary stratigraphy and geologic
history of Mono Basin, eastern California. Ph.D. thesis, Univ.
California, Berkeley. 379 p.
LI, H.-C. 1995. Isotope geochemistry of Mono Basin, California:
Applications to paleoclimate and paleohydrology. Ph.D. thesis,
Univ. Southern California, Los Angeles. 244 p.
LISTER, G. S., K. KELTS, K.-Z. CHEN, J.-Q. Yu, AND E NIESSEN.
1991. Lake Qinghai, China: Closed-basin lake levels and the
oxygen isotope record for ostracoda since the latest Pleistocene. P.aleogeogr. Palaeoclimatol. Palaeoecol. 84: 141-162.
MCKENZIE, J. A. 1985. Carbon isotopes and productivity in the
lacustrine and marine environment, p. 99-118. In W. Stumm
[ed.], Chemical processes in lakes. Wiley.
MONO LAKE: COMMITTEE. 1981. Mono Lake guidebook. Kutsavi
Press. 113 p.
NATIONAI, R.ESEARCH COUNCIL. 1987. The Mono Basin ecosystem:
Effects of changing lake level. NRC.
NEWTON, M’. S., AND E. L. GROSSMAN. 1988. Late Quaternary
chronology of tufa deposits, Walker Lake, Nevada. J. Geol. 96:
417-434.
O’NEK, J. R.., R. N. CLAYTON, AND T. K. MAYEDA. 1969. Oxygen
isotope fractionation in divalent metal carbonates. J. Chem.
Phys. 51: 5547-5558.
PHIIXPS, E M., AND OTHERS. 1992. A reconstruction of the rcsponse of the water balance in Western United States lake basins to climatic change. New Mexico Water Resour. Res. Inst.
WRRI IRep. 26. V. 1.
RUSSELL, I. C. 1889. Quaternary history of Mono Valley: California. U.S. Geol. Surv. 8th Annu. Rep.
SCHELSKE, C. L., AND D. A. HODELL. 1991. Recent changes in
productivity and climate of Lake Ontario detected by isotopic
analysis of sediments. Limnol. Oceanogr. 36: 961-975.
STREET-PERI~OTT, E A., AND S. F! HARRISON. 1985. Lake levels
and climate reconstruction, p. 291-340. In A. D. Hecht [ed.],
Paleocli mate analysis and modeling. Wiley.
TURNER, J. V., I? FRITZ, P E KARROW, AND B. G. WARNER. 1983.
Isotopic: and geochemical composition of marl lake waters and
implications for radiocarbon dating of marl lake sediments.
Can. J. Earth Sci. 20: 599-615.
JOHNSON, T. C., J. D. HALFMAN,
Received: 13 July 1995
Accepted: 25 June 1996
Amended: 23 October 1996