Climate-driven hydrologic transients in lake sediment records: calibration of groundwater conditions using 20th Century drought

Quaternary Science Reviews 21 (2002) 605–624

Climate-driven hydrologic transients in lake sediment records: calibration of groundwater conditions using 20th Century drought Joseph J. Donovana,*, Alison J. Smithb, Valerie A. Panekc, Daniel R. Engstromd, Emi Itoe a

Department of Geology and Geography, West Virginia University, Morgantown, WV 26506-6300, USA b Department of Geology, Kent State University, Kent, OH 44242, USA c CH2M Hill, Portland, OR 97232, USA d St. Croix Watershed Research Station, Science Museum of Minnesota, Marine on St. Croix, MN 55047, USA e Limnological Research Center, Department of Geology and Geophysics, University of Minnesota, Minneapolis, MN 55455, USA

Abstract The effect of 20th-century drought on groundwater and lakes in a climatically sensitive area of the Northern Great Plains (Grant County, western Minnesota) was investigated by analysis of lake sediments and historical air photos, in conjunction with groundwater flow simulations. Drought caused an observed 4.0 to 5.1 m lake-head decline between 1923 and 1938; all but three either dried completely or declined to o1.7 m depth. From the deepest (Elk Lake), a 210Pb-dated sediment core was analyzed for ostracode species distribution and geochemistry of C. Rawsoni shells (d18O, d13C, Mg/Ca, Sr/Ca). Mg/Ca and d13C increased during drought at the same time that salinity-tolerant ostracodes thrived, suggesting increasing salinity. However, d18O decreased during drought, anti-correlative with Mg/Ca. A numerical model for transient response of groundwater flow to drought, modeled after that of 1923–38, was constrained by drought-end recessional strand lines observed on air photos of one lake, with strand line elevations inferred from bathymetry. Simulated lake-water-level declines in the first 15 drought years were consistent with strand line elevations. The rate of decline was exponential at drought onset, slowing as water in lakes and wetlands receded to below land surface and lake evaporation declined. A near-steady state was attained between 40 and 50 simulation years after onset of drought, with the water table from 1–5 m below the level of dry lakebeds and only the two deepest lakes still holding water. They were, however, greatly constricted in area and depth. Simulated evaporation fluxes were reduced by over 50% within 15 years, despite increased ET rates, due to lake area constriction and water table declines. Both relate to land surface morphology beneath and around lakes. Model results suggest that lake-bed elevations exert control on the rate and ultimate level to which groundwater is depressed by drought. The deepest lakes in any region will tend to become the focus of groundwater flow during sustained drought once lakes around it dry out. The unexpected finding of decreasing shell d18O in response to drought is interpreted as some combination of source and/or evaporation–reduction effects, such as seasonality of recharge/precipitation or reduced ET flux from wetlands that seep back into lakes. r 2002 Published by Elsevier Science Ltd.

1. Introduction Interpreting drought impacts on landscape and human affairs is dependent on accurate reconstruction of past drought records. Such reconstruction frequently uses proxies of climate change from the lake-sediment record, such as diatoms, pollen, and ostracodes (Bartlein and Whitlock, 1993; Fritz et al., 1993; Smith et al., 1997), stable isotopes (Xia et al., 1997), trace metals (Engstrom and Nelson, 1991), and mineralogy (Last, 1989). Well-dated lake-sediment sequences have been *Corresponding author. Tel.: +1-304-293-5603/4308; fax +1-304293-6522. E-mail address: [email protected] (J.J. Donovan). 0277-3791/02/$ - see front matter r 2002 Published by Elsevier Science Ltd. PII: S 0 2 7 7 - 3 7 9 1 ( 0 1 ) 0 0 0 4 2 - 7

demonstrated to record Holocene climatic variations and inferred droughts with high precision (Dean and Megard, 1993). While individual proxies can be diagnostic of climate, entire sets of proxies frequently do not respond to climate variations uniformly; different proxies seem to be sensitive to hydrologic factors as well as climate (Fritz, 1996). While the exact causes of these factors are not well understood, one frequently cited cause is the influence of groundwater on lake chemistry and hydrology (Rosen, 1995). Lake–groundwater interaction fundamentally influences lake water chemistry and resulting mineralogy of endogenic sediments (Sanford and Wood, 1995). Because lakes exchange both solutes and water with groundwater, it is generally unjustified to

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treat lakes as simple ‘‘closed basins’’, even for extremely saline cases (Wood and Sanford, 1990). On the other hand, the specific groundwater mechanisms causing perturbations in lake proxy response have not been fully elucidated. Relatively few investigations have compared climate history as recorded by lake-sediment proxies to the history of groundwater conditions around lakes. Opportunities to do so are limited by data scarcity; few groundwater datasets for both lakes and contiguous aquifers are collected systematically, and records of paleohydrology are commonly difficult to decipher. Frequently, paleohydrogeologic conditions relating to lakes must be inferred through groundwater flow modeling (Anderson and Munter, 1981; Winter, 1983; Sacks et al., 1992). 1.1. Purpose and study area The purpose of this paper is to examine the magnitude of hydrologic effects on lakes and groundwater for a major 20th century drought spanning 1923–38, recorded by lake sediment proxies in the Northern Great Plains (NGP). We examine the historical response of a NGP lake-aquifer system to this severe ‘‘Dust Bowl’’ drought to assess how and to what degree hydrologic response to drought was reflected in sediment proxies of a local deep lake. An additional goal is to examine the response of surface and groundwater to this drought, employing modeling based on the 1930s water level record from one of these lakes. Extension of these model results to a mid-Holocene drought of greater magnitude is presented in Smith et al. (this volume). The study area is the glaciated landscape around the Elk Lake chain, Grant County, western Minnesota (Fig. 1). The area is chosen because: (1) its two major streams constitute strong hydrogeologic boundaries; (2) the sub-humid climate (60–65 cm/yr mean average precipitation, or MAP) is thought to undergo periodic droughts of prolonged duration; and (3) at least one lake in this area has a continuous Holocene lake-sediment record (Smith, 1991). Our approach is to: (1) observe and employ modern (1995) hydrologic conditions to quantify parameters describing lake–groundwater interaction; (2) numerically simulate both modern groundwater conditions and those inferred for the Dust Bowl drought; and (3) compare results of the latter model to a well-dated lake-sediment proxy record for this drought. Whereas models of groundwater flow can be a productive tool for examining lake-aquifer systems, their results require corroboration with field evidence (calibration). Here, we will use as a basis for calibration observations of lake strandlines recorded in 1938 aerial photographs, near the end of the Dust Bowl arid interval. The study area (Fig. 1) lies along the modern prairie– forest border, approximately 120 km north of the

Fig. 1. Location of Elk Lake showing lakes and streams, as well as downstream hydrograph stations. Cross section A–A0 refers to Fig. 3.

Minnesota River. Several of its lakes are sufficiently deep to be perennial. Land surface is hummocky and underlain by glacial sediments. Its western and eastern boundaries are the Pomme de Terre and Chippewa rivers (Fig. 2), tributaries of the Minnesota River that are gaining streams in this, their headwater areas. They possess few tributaries and discharge primarily baseflow except in periods immediately following storms or snowmelt (when they carry runoff) and during prolonged drought (when they may tend to dry up completely). Like so much of the glaciated NGP, this landscape has poorly continuous surface drainage; precipitation and runoff tend to focus on upland

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Fig. 2. Distribution of lakes, wetlands, wells, streamflow stations, and potentiometric surface of the confined sand aquifer (latter after Delin, 1986).

depressions and wetlands, then slowly drain to the major streams as groundwater discharge (Winter, 1986). Lakes interact predominantly with shallow groundwater within either highly permeable outwash or less permeable fine-grained (silty) till. The deepest lake in the area is Elk Lake (1995 maximum depth 9.1 m), one of a linear northwest-trending series of five, from Elk Lake South at the southeast to Turtle Lake at the northwest (Figs. 1 and 2). This series, referred to as the Elk Lake chain, lies in an erosional linear depression of glacial origin. During very wet periods, lakes rise to full stage and spill to lower lakes in the chain via surface overflows. In ‘‘normal’’ and dry periods, the lakes are connected to each other only by groundwater. No permanent streams supply or drain these lakes. At all times, groundwater seeps between adjacent lakes, driven by water level differences. Turtle Lake (lowest in the chain) spills during wet periods to the Pomme de Terre River, but during dry periods, its only water outlet is out-seepage into alluvium. Principal mechanisms for groundwater discharge from till of the NGP include evapotranspiration, lateral flow to wetlands and streams, and loss to deeper aquifers. The relative importance of these fluxes is

influenced by topography, geology, aridity, and hydraulic properties. Irregular or hummocky topography is conducive to ponding of lakes and wetlands that tends to focus lateral groundwater flow in dry climates (Winter, 1976). Anderson and Munter (1981) demonstrated that groundwater inflow to lakes is seasonal and transient, with recharge creating peripheral groundwater mounds that dissipate as water discharges into lakes. 1.2. Regional geology and hydrogeology The study area has a rolling upland topography with about 30 m of maximum relief, just west of the Alexandria moraine, a north-trending regional icestagnation moraine of the Des Moines lobe. Nearsurface geology consists of poorly exposed Pleistocene glacial sediments 100 to 150 m thick overlying deeply buried crystalline rock (Delin, 1986). Fig. 3 shows an east–west cross-section (A–A0 ) from Panek (1996), depicting the upper portion of Quaternary sediments from well logs in the vicinity of Elk Lake. Glacial outwash occupies low-lying channels of the two major rivers on either side of the area, while the shale-rich New Ulm till (Wisconsin age) is extensive on uplands between

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Fig. 3. Cross section of geology beneath the Elk Lake chain (see Fig. 1 for location).

these rivers. The New Ulm till lies at the surface in the study area and interacts closely with lakes. In its type area 200 km southeast of Elk Lake, the New Ulm till is 30 m thick, with an upper yellow-brown oxidized zone 6 m thick over gray unoxidized till (Matsch, 1972). In the study area, its upper weathered portion (commonly described as ‘‘yellow clay’’ in drillers’ logs) is from 7 to 18 m deep and overlies at least 5 m of unoxidized gray clay. The color change is diagnostic of the weathered to unweathered transition. Beneath the New Ulm lies one of two confined sand aquifers, the Erdahl and Barrett aquifers of Delin (1986), 21–30 and 40–50 m deep, respectively, in this area. The till is one of three shallow (o50 m depth) waterbearing units in the area (Fig. 3), along with outwash and the Barrett and Erdahl sand units (‘‘confined sand aquifer’’ of Delin (1986)). There are few wells developed in weathered till due to its moderate hydraulic conductivity, and thus the best indicators of its water table elevation are stages of lakes and wetlands. The confined sand aquifer is poorly continuous (Soukup et al., 1984); its water levels beneath lakes are lower than modern lake stage by 1.5–15 m, thus any modern leakage between lakes and shallow groundwater to the deeper confined aquifers is downward. Delin (1986) estimated downward leakage south of Hoffman at 1–9 (average 2.5) cm/yr, based on hydraulic tests at five wells. Outwash underlies the entire length of the Pomme de Terre and Chippewa valleys (Delin, 1986). This outwash is coarse sand and gravel from 30 to 50 m thick, underlain by ‘‘blue clay’’ (till or lacustrine deposits). It is unclear whether the Barrett (lower) or Erdahl (upper) aquifers are stratigraphically continuous with this outwash, although Delin (1986) speculated the Barrett

aquifer intersects Pomme de Terre outwash west of Hoffman (Fig. 3). The Elk Lake chain lies in the only prominent break in the till upland, a linear topographic depression mapped as a ‘‘tunnel valley’’(Mooers, 1989). Tunnel valleys are erosional channels formed by paleoflow beneath thinning ice margins during glacier stagnation or retreat (Patterson, 1997) and typically contain wetlands, underfit streams, and heterogeneous esker deposits. Tunnel valley deposits can be very permeable. They generally trend oblique to the regional topographic slope and are unevenly incised, with scour pools and transverse sand ridges, as form lakes and inter-lake barriers in the Elk chain. Hydrologic budgets for the Pomme de Terre and Chippewa basins were compiled by Cotter and Bidwell (1966) and Cotter et al. (1968). Delin (1986) modeled groundwater flow in the Barrett and Erdahl aquifers. Fig. 2 shows the distribution of lakes/wetlands and the potentiometric surface of the Barrett sands, which trend to the south-southwest at a gradient from 0.0005 to 0.002. Delin (1986) inferred there may be regional groundwater flow through these sands between the Chippewa (higher) and Pomme de Terre rivers. Annual water level fluctuations of head are 1–2 m in the deeper confined aquifers and o1 m in outwash (Delin, 1986). 1.3. Hydraulic conductivity and recharge in shallow glacial till In the NGP, weathered till of moderate permeability is widely observed to a depth of from 3 to 15 m. The hummocky flat topography overlying many till deposits limits runoff, but till permeability, often relatively low, tends to focus infiltration and recharge in wetlands and

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dry depressions (‘‘potholes’’). Although fine grained, weathered till is generally higher in hydraulic conductivity (K) than apparent from sediment grain size, due to closely spaced fractures and leached carbonates in the matrix (Grisak and Cherry, 1975). At greater depth, weathering diminishes and K is much lower. The transition from weathered (brown) to underlying unweathered (gray) till may be either abrupt or gradual. The depth of active groundwater flow in clayey till was observed to range from 5 to 10 m beneath an Ontario site by Ruland et al. (1991), a typical range compared to other investigations (Keller et al., 1988; Fortin et al., 1991). K estimates for weathered till from borehole tests averages 104–106 cm/s; unweathered till averages at least 1–2 orders of magnitude lower (Hendry, 1982; Cravens and Ruedisili, 1987). These results support the widely stated observation that the top of unweathered till may be considered an effective base for shallow groundwater flow (Hendry, 1988). While vertical velocities to great depth in till are generally quite low, lateral flow is substantially higher in rate due to the role of coarser-grained heterogeneities and to fractures (Simpkins and Bradbury, 1992; In contrast to K; few results are reported for recharge rates to till aquifers. Daniels et al. (1991) estimated recharge at 3.5–4.7 cm/yr in northern Indiana (average precipitation 96 cm/yr) using tritium in vadose water. Negligible recharge to the water table was observed in local-scale vadose measurements in semi-arid North Dakota (Schuh et al., 1993), but substantial lateral inflow indicated recharge was occurring in surrounding locations.

2. Methods For this study, thickness and depth of aquifers were determined from subsurface data, mainly well logs (Larson, 1976; Soukup et al., 1984; Minnesota Geological Survey, file data). Streamflow measurements were taken 5 times during 1994–1995 at five stations along the two rivers (Fig. 2) with an analog Marsh-McBirney magnetic-flux current meter and wading rod. The stations are on the upgradient and downgradient margins of the study area. Flow was estimated by velocity–area integration, with error of approximately 720%. Historic and modern daily discharges were compiled from USGS stations at Appleton (Pomme de Terre R.) and Milan (Chippewa R.) downstream of Elk Lake. Using the watershed boundaries of Cotter and Bidwell (1966), drainage catchment areas were estimated for both rivers upstream from these streamflow stations. Air photos were examined for indications of 20th century hydrologic history. Photos at scale 1 : 20,000 and 1 : 40,000 were collected both from humid years when lakes were full (1965, 1972, 1982, 1984, 1991) and

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from droughts when their water levels were depressed (1938, 1958). The 1 : 20,000 1938 photos (USDA series BIW-3, 8/31/38 and BIW-2, 9/4/38), taken near the end of the 1923–38 drought, were especially striking, showing a number of lakes and streams to be nearly or completely dry. Some dry lake beds exhibited parallel recessional strandlines from the years prior to 1938 as lake levels were dropping. In July 1995, a rotating-laser-beacon level survey was performed with respect to a fixed datum on the upper Chippewa River at station CH1 (Fig. 2) to accurately determine potentiometric levels in lakes, wetlands, and stream channels for one reference time. Maximum closure error was 0.13 m. Lake bathymetry was measured in August 1994 for Elk South, Elk, Spring, Round, Turtle, Thompson and Church lakes, profiling along cross-lake transects spaced ca. 50 m apart with a continuous-recording fathometer, accurate to 70.3 m. Resulting bathymetric contours were overlain on the 1938 photos to estimate 1938 lake level and pre-1938 strandline elevations. Table 1 shows the lake depths and stages measured for 1994 and calculated for 1938, assuming negligible change in lake bottom elevation between the two dates. Four lakes (Spring, Church, Elk, and Elk South) were surveyed for locations of lake-bed seepage inflow using a ‘‘mini-piezometer’’, a 2-m long, 2-cm diameter stainlesssteel probe capable of withdrawing water from lake-bed sediments by pumping or suction (Winter et al., 1988). Sample locations were selected at spacings of 40–100 m along lake perimeters, in water 0.4–1.4 m deep. Identification of groundwater inflow could be done by temperature, as groundwater is substantially colder (9– 121C) than lake water in summer months. The sampling method identifies locations of concentrated groundwater inflow (springs and seeps) in littoral areas. A 2 m long, 12-cm diameter core of recent sediments was collected from Elk Lake using a modified Livingston piston corer in March 1994. The core was taken in 9.1 m of water, just north of the deepest part of the lake. The core was sampled in 1-cm continuous increments; these were homogenized and split for analysis of ostracode species distribution, carbonate and organic carbon, non-combustible residue, and stable and traceelement geochemistry. Ostracode shells were isolated by sample washing through sieves, and hand-picking under a binocular microscope, and identified following the methods described in Vance et al. (1997). Ostracode counts by species based on prevailing taxonomy (Delorme, 1970, 1971) were converted into valves per gram of sediment. Hand-selected carbonate shells of late-instar Candona rawsoni (well preserved throughout the short core) were picked from each interval for chemical and isotopic analysis. Ten to twenty shells of cleaned late-instar juveniles were reacted with ultra-pure H3PO4 for stable

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Table 1 Comparison of lake depths and stages based on bathymetry and air photo analysis, 1994 and 1938 Lake

1994–1938 Lake surface elevation (m)

Elk Lake South Elk Lake Spring Lake Round Lake Turtle Lake Thompson Lake Church Lake Torgerson Lake

08/01/94 (surveyed)

09/01/38 (air photo)

368.74 367.22 365.08 364.95 364.96 376.59 371.00 371.56

363.8 363.3 360.0 360.4 361.0 372.6 367.0 o367.2

Lake head decline (m)

Max. 1994 depth (m)

Max. 1938 depth (m)

Sediment bottom elevation (m)

Modeled 1938 head decline (m)

4.85 3.96 5.05 4.58 3.98 3.97 4.00 >4.3

5.2 9.1 5.6 5.6 5.6 7.6 4.4 4.3

0.3 5.2 0.6 1.1 1.7 3.6 0.4 dry

363.6 358.1 359.4 359.3 359.3 369.0 366.1 367.9

4.62 4.03 3.78 3.76 3.71 3.11 3.03 4.37

isotope analysis using a Finnigan MAT delta E triple collector mass spectrometer (method modified after McCrea, 1950), and the residue used for inductively coupled plasma mass spectrometry (ICP-MS) analysis of selected cation concentrations (Xia et al., 1997). The analytical precision for d18O was 0.2m and for d13C was 0.13m. The precision of the ICP-MS determinations was about 75% for Mg, Ca, and Sr. The upper 100 cm of core was dated by 210Pb using 210 Po-distillation and alpha spectroscopy methods. Sediment samples from 24 depth intervals were spiked with 209Po, distilled at 5501C, and the polonium isotopes plated onto silver planchets (Eakins and Morrison, 1978). Activities were measured for 1–6 days with Sidepleted surface barrier detectors and an Ortec alpha spectroscopy system. Supported 210Pb was estimated by averaging the asymptotic activity below 100 cm, and unsupported 210Pb was calculated by difference and converted into interval mean ages, assuming a constant rate of 210Pb supply. Ages between analyzed intervals were estimated by linear interpolation, and dates before 1850 were extrapolated assuming a mean pre-European sedimentation rate of 0.10 g/cm2/yr1. Precision in dates is o73 years to about 25 cm depth (calculated date 1923), below which it deteriorates to 712 years (date 1877). All dates since 1905 are relatively precise.

3. Results 3.1. Climate history Instrumental climate data for the 20th century (Fig. 4) include a histogram of monthly precipitation (a); a time series of mean annual precipitation (MAP) (b); mean annual precipitation minus lake evaporation, or lake P  E (c); and both instrumental (d) Palmer Hydrologic Drought Index (PHDI) and (e) Palmer Drought Severity

Index (PDSI) values reconstructed from tree ring observations; (Cook et al., 1999). Most (>75%) precipitation occurs in April to September, centered about a June maximum. MAP (mean 63 cm/yr) ranged from a low of 24.9 cm (1976) to 86.6 cm (1984). The PDSI is an index of prolonged abnormal moisture deficiency or excess and is calculated from weekly precipitation total and average temperature, regional constants for available water capacity and content of the soil, and antecedent values of the index (Alley, 1985). Values of PDSI of 2 indicate moderate drought; 3 is severe drought; and 4 is extreme drought. The instrumental PHDI and tree-ring PDSI records indicate droughts in the 1930s, 1950s and 1970s, but by far that of greatest magnitude was from 1922 to 1938, with below-zero PHDI over the entire period and below 3 from 1931 to 37, the most severe portion of this drought. Dendroclimatology shows partial recovery starting in 1938. Pan evaporation (period 1964–90) averaged 102 cm/yr (lake evaporation 76 cm/yr), with a marginally significant inverse correlation (R ¼ 0:43) between P and E; that is, years of below-average precipitation were also dry (negative P  E) periods. 3.2. Elk Lake short core record Geochemical proxies analyzed on C. rawsoni shells included Sr/Ca, Mg/Ca, d18O, and d13C (Fig. 5). Also shown in Fig. 5 are abundances of ostracode species dominating the modern Elk Lake assemblage: the benthic C. rawsoni, C. ohioensis, and Limnocythere itasca, and the nektic Physocypria globula. Trace elements and isotopes all responded sensitively to drought (1923–1938). Sr/Ca decrease and Mg/Ca increase, characteristic of response ascribed to increasing salinity and aragonite precipitation in waters of Mg/ Ca>3 (Engstrom and Nelson, 1991). The stable isotopes also clearly respond to drought but hetero-

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behavior suggests normal response to the 1923–38 drought for all but d18O. The drought years were marked by a fluctuation in abundance of all ostracodes. As salinity rose in Elk Lake (tracked by the Mg/Ca values), the euryhaline C. rawsoni and also L. itasca increased in abundance. Following an apparent decrease in TDS after 1940, species such as C. ohioensis increased in abundance. A decrease in abundance of all species after 1950 may be due to changes to the littoral zone induced by new land use (house construction, recreation) and farming practices around the lake. 3.3. Recharge rate estimates

Fig. 4. Instrumental climatic data: monthly precipitation means (a); mean annual precipitation with 3 yr backward moving average (b); calculated mean lake P  E (c); instrumental Palmer Hydrologic Drought Index (PHDI); and reconstructed Palmer Drought Severity Index (PDSI) from tree ring data (Cook et al., 1999). The precipitation data are for Morris MN; the P  E results employ Wauseca MN Class A evaporation pan rates corrected for lake evaporation using a 0.73 ‘‘pan factor’’; and the PDSI data are for west-central MN (45–471N; 96–981W).

geneously: d18O decreases while d13C increases. The d18O response is counterintuitive for a drought condition: drought might be expected to induce higher 18O in lake water due to evaporative concentration, although there are many other factors related to source and seasonality that influence the 18O budget (Smith et al., this volume). Thus, the observed geochemical proxy

Recharge is approximately equal to stream baseflow, provided both aquifer underflow to deep aquifers and phreatic evapotranspiration are minorFcircumstances not universally attained. But under suitable conditions, its value may be estimated by separation of stream hydrographs (Rutledge and Daniel, 1994). For the period June 1994–July 1995, continuous flows for the Chippewa and Pomme de Terre stations at Milan and Appleton, respectively, were used to interpolate synthetic hydrographs for stations in the study area where only sporadic flow measurements had been taken (CH2, CH6, PT1, and PT4; Fig. 6; see Fig. 3 for location). According to convention (Meyboom, 1961), on an exponentially scaled flow hydrograph, linear segments indicate pure recession (e.g., baseflow) without runoff, as between August 1994 and February 1995. Such baseflow was separated from runoff following rainy periods by backward extrapolation of log-linear segments. The resulting baseflow hydrographs (e.g., the bold line on Fig. 6) were integrated numerically to yield cumulative annual baseflow (Table 2), an estimate of cumulative aquifer recharge. In addition to the stations where flows were measured sporadically, baseflows were calculated and integrated for the continuous Appleton and Milan flow records downstream, yielding a largerscale baseflow estimate. Integrated baseflow was divided by calculated upstream drainage sub-areas, expressing these as spatially uniform recharge rates (Table 2). The average of all areal recharge estimates at all scales is 2.7 cm/yr (7.4  103 m/day). The average of downstream large-scale estimates (5.6 cm/yr) is substantially higher than for upstream local-scale estimates (1.7 cm/ yr) based on individual measurements. Variance between estimates is lower for the Pomme de Terre stations (higher flow) than the Chippewa. For the period 1994– 95 (MAP 70.7 cm/yr, 10.5 cm/yr wetter than the 100 year norm), and thus recharge rate estimates may be somewhat higher than the long-term average. These are approximate estimates of ‘‘modern’’ (1990s) recharge during a humid period. They are subject to considerable uncertainty, whose principle sources are

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Fig. 5. C. rawsoni shell geochemistry (Sr/Ca, Mg/Ca, d18O, and d13C) and distribution of C. rawsoni, C. ohioensis, L. itasca, and P. globula in the Elk Lake core.

Fig. 6. Flow hydrographs of the Pomme de Terre River, showing baseflow separation for the Appleton station and flows for Appleton and PT1.

lack of continuous stream records in the study area; uncertainty in separation of baseflow from runoff; summer ET in the stream valleys and other wetlands; and lack of knowledge regarding underflow to deeper aquifers. Reducing these uncertainties would take substantial data collection for a long period, and even with a strong base of hydrologic data, many uncertaintiesFin underflow, for exampleFcould remain large. These estimates made using interpolated baseflows data are based on few data and are of uncertain accuracy, but are consistent with those estimated in regional field studies of Cotter et al. (1968) and Delin (1986). 3.4. Lake bathymetry and strand lines Late summer 1938 air photos for Elk South, Elk, Spring, Church, Olson, and Torstenson are shown in Figs. 7 and 8, with 1994 bathymetric contours, in meters

below 1994 lake stages, superposed on several lakes. Along the Elk Lake bottom is a submerged sand bar or spit projecting from a peninsula on the northeast side of the lake; its proximity to the coring location required the latter to be moved due to encountering coarse sand at 3– 4 m depth. Table 1 gives estimated 1994 and 1938 lake surface elevations, maximum depths, and hydraulic head change between the two dates for these and other lakes. 1938 lake levels were from 4.0 to 5.1 m lower than 1994 for all lakes; a number (e.g., Torgerson) were completely dry, while all but Elk (5.2 m) and Thompson (3.6 m) were within 1.7 m of being dry. Seven identifiable strand lines are visible at the northeast end of Elk Lake South in the 1938 photo (Fig. 9), although they could not be resolved in the bathymetric survey or were not preserved during rising of the lake. In late summer 1938, Elk Lake South had declined to a shallow (o0.3 m) pool less than 30% of its

J.J. Donovan et al. / Quaternary Science Reviews 21 (2002) 605–624 Table 2 Recharge rates calculated by baseflow integration from 6/94 to 5/95 at different scales Flow station

Chippewa River Appleton PT-4 PT-1 Model area (PT4-PT1) Average (Chippewa stations) Pomme de Terre River Milan CH-6 CH-2+CH3B Model Area (CH6-CH2) Average (Pomme de Terre stations) Average (both rivers)

Annual baseflow (m3)

Catchment Recharge area rate (m2) (cm/yr)

7.4  107 2.2  107 1.7  107 4.9  106

2.3  109 9.4  108 8.4  108 9.8  107

3.16 2.29 1.98 4.96 3.10

3.9  108 7.7  106 1.0  107 2.4  107

4.8  109 7.6  108 5.8  108 1.7  108

8.05 1.02 1.73 1.40 2.35 2.72

1994 area. These strand lines are interpreted as annual lake stands between about 1931 and 1938 in the latter portion of the drought. Lake levels were interpolated for these strand lines using their bathymetrically determined elevations, assuming minimal change in bottom sediment elevation between 1938 and 1994. These water levels (from 2.9 to 4.9 m below 1994 lake level; Table 3) are interpreted as an estimate of the water table decline rate during the latter half of the drought. 3.5. Historical flows in the Pomme de Terre and Chippewa rivers Fig. 10 shows cumulative annual flows for the Pomme de Terre River (period 1931–1999, Appleton station) and for the Chippewa River (period 1937–1999, Milan station). The Pomme de Terre station is 82 km southsouthwest of the study area, near its confluence with the Minnesota River, and the Chippewa station about 10 km SE of Appleton, also near its confluence with the Minnesota. While these two stations are both quite distant from Elk Lake and of no direct use in estimating recharge rates, these flows give an indication of the high variability in streamflow and of the degree to which baseflow is depressed during droughts. The hydrographs indicate wet periods of extremely high flows from 1984 to 88 and from 1992 to 98, including the period of measurements for this study. No comparably wet interval occurred earlier. The Chippewa drainage historically exhibited somewhat higher baseflows and much higher peak flows than the Pomme de Terre, reflecting its larger catchment area upstream from this location (540,200 vs. 232,900 ha). The period 1931– 36 was a severe drought. In the Pomme de Terre, flow at

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most times between 1931 and 1937 was o1 m3/s, and always nearly dry in summer, o0.1 m3/s. Data collection did not start at the Chippewa station until 10/37 but after the drought ‘‘broke’’ in 1937–38 its discharge was similarly low to the Pomme de Terre, or even dry; relatively low flows (generally o10 m3/s at snowmelt, 51 m3/s in summer) persisted through 1942 in both streams. At the time of the air photos (Figs. 7–9, 8/31/ 1938 to 9/4/1938), streamflows ranged from 0.2 to 0.5 m3/s in the Chippewa and from 0.2 to 0.4 m3/s in the Pomme de Terre. However, these stations are far downstream of Elk Lake; in the study area, the Chippewa was observed to be primarily dry, showing only standing intermittent pools in wetlands. The Pomme de Terre near Elk Lake carried an apparently minor flow. Despite the incomplete period of record, it is clear that flows during the drought were greatly reduced or even dry. Baseflow was either zero or far less than potential evaporation for lakes and wetlands in the stream channels. Therefore, it is believed that these baseflows would not have been representative of groundwater recharge, but were reduced below recharge rates by the effects of evaporation. Baseflows during humid years, however, are more prominent and thought to yield an estimate of groundwater recharge. 3.6. Groundwater observations Fig. 11 shows the distribution of seepage flowing into lake beds of Elk, Spring, Elk South, and Church lakes, as sampled with a mini-piezometer probe. In Elk and Spring lakes, groundwater inflow from lakes higher in the chain was observed in near-shoreline sediments. Chemistry of inflow was similar to that of up-gradient lakes. Systematic head differences between adjacent lakes (1.5, 2.1, and 0.13 m moving down the chain from Elk South to Round) as well as drillers’ logs indicate that lake out-seepage occurs through relatively permeable weathered till with sand and gravelly layers (Panek, 1996). In addition to in-seepage from upgradient lakes, isolated springs occur in the littoral zone of lakes parallel to the orientation of the tunnel valley, without apparent spatial pattern. Elk Lake South appears to occupy a groundwater divide between the Pomme de Terre and Chippewa rivers, but was very close to the elevation of the Chippewa River in 1994 and groundwater inflow to the lake from the southeast was confirmed by sampling. 3.7. Steady-state analysis of 1994–95 conditions A numerical model of both steady state and transient groundwater flow was developed for the shallow groundwater in this area using MODFLOW (McDonald and Harbaugh, 1988). The objectives were (a) to

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Fig. 7. Air photos of Elk Lake (left) and Elk Lake South (right) on 9/4/1938, with superposed 1994 bathymetry and maximum depths in 1994 and 1938 (x’s).

obtain a representative steady-state solution describing ‘‘modern’’ (1994) shallow groundwater conditions; and (b) to expose this calibrated aquifer to simulated transient stress (reduced recharge, increased ET) similar to that of the 1923–37 drought. The steady-state

simulation employed spatially distributed values for K; recharge rate (R), evaporation rate (ET), land surface elevation, and aquifer geometry. The transient simulation required all these parameters plus specific yield and stress history. Flow solutions are controlled by bound-

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Fig. 8. Air photos of Church, Olson, and Torgerson lakes on 9/4/38, with superposed 1994 bathymetry for Church Lake and maximum depths in 1994 and 1938 (x’s).

ary conditions, as imposed, for example, by streams or lakes, groundwater divides, or physical aquifer boundaries. To constrain the steady-state simulation, 1994–95 baseflow estimates, from field measurements, were used to approximate recharge rates, and potentiometric levels from lakes were used for calibration. Key assumptions include: (1) a single layer 18-m thick surface-aquifer geometry, with no loss to underlying confined sand aquifers; (2) a time-invariant, but spatially heterogeneous, conductance, independent of hydraulic head;

(3) a simplified 3-zone (till, outwash, and tunnel-valley deposits) heterogeneity model; (4) spatially uniform recharge; and (5) dependency of phreatic (but not vadose) evapotranspiration on water table depth below land surface. A 113 row  119 column north–south grid was constructed with uniform spacing of 100 m (Fig. 12). Conductivity zones were employed to represent shallow weathered till (Zone 1), outwash aquifers (Zones 2 and 3 for Chippewa and Pomme de Terre, respectively),

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Fig. 9. Air photo of strandlines at the north end of Elk Lake South, 9/4/1938.

Table 3 Strandline elevations for Elk Lake South on air photos of 9/4/38 Year after 1923

0 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15

Date

1923 1924 1925 1926 1927 1928 1929 1930 1931 1932 1933 1934 1935 1936 1937 1938

Elk Lake S. Strandline decline (m)

Modeled decline (m)

0 (assumed)

0.00 0.79 1.32 1.75 2.12 2.43 2.73 3.00 3.24 3.46 3.66 4.03 4.33 4.59 4.81 5.01

2.96 3.17 3.66 3.81 3.99 4.18 4.85

tunnel-valley deposits (Zone 4) and lakes (Zone 5). The single unconfined MODFLOW layer (LAYCON type 2) employed non-varying conductanceFthat is, declines in water level did not influence aquifer conductance. Phreatic evapotranspiration (ET) was employed using a linear attenuation function based on water table depth, from maximum rate at land surface to zero below an extinction depth of 3 m. For lakes, the maximum ET value was used. Aquifer top (e.g., land surface) elevation was interpolated from one arc-second DEM values and aquifer bottom was set to a uniform 18 m below this. Slight modifications to bottom elevations were made, to set streambed and lake-bottom elevations to fieldmeasured values.

Fig. 10. Times series of total annual stream discharge rates for the Appleton (Pomme de Terre; period 1931–1999) and Milan (Chippewa, period 1931–1999) gauging stations, 82 and 90 km south of the study area, respectively.

Recharge was applied on Zone 1 (till) at 8  105 m/ day, the average value for all scales from 1994–95 integrated baseflow estimates (Table 2). For precipitation falling on lakes, a much higher value of 1.67  103 m/day (i.e., average MAP) was employed. For Zones 2, 3, and 4 (outwash and tunnel valley deposits), 1.6  104 was used, twice the value of that for till, in consideration of generally sandy soils and low-lying topography subject to focused recharge. The east–west trending north and south boundaries of the model are approximate flowlines, and so no-flux boundaries were employed here (Fig. 12). Fixed-head boundaries were set along the two rivers at surveyed 1994 stage elevations. In addition to providing local hydrologic base levels, these boundaries induced flow through the outwash aquifers into and out of the model area at the prevailing gradient of these two streams. K zones were as summarized in Table 4. Zone 5 (lakes) was set at very high K to project near-zero head gradients across the lakes. The K of Zone 1(till) was varied to achieve a hydraulic head depth no greater than 18 below land surface; a value of 0.2 m/day resulted, close to the value of 0.3 m/day suggested for horizontal till K in a nearby area by Delin (1986). Local values of Zone 4 (tunnel valley sediments between lakes) were iteratively adjusted until the simulated hydraulic heads of all lakes was within 71 m of their surveyed 1994 stage. These calibrated K’s for Zone 4 between each lake are in general 40 to 80 m/day except for a very high value (300 m/day) required to match lake heads between Turtle, Round, and Spring lakes, a result of field conditions (very full lakes during the very rainy 1994). This observed head difference may have been an unnatural condition, as water may have been spilling between lakes via concealed subsurface drains. Therefore, the higher K’s in Zone 4 are probably artifacts.

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Fig. 11. Lakes, wetlands, seeps, and locations of lake-bed seepage in the Elk Lake chain.

Table 4 shows calibrated parameters for K; ET, and R: Critical assumptions for the steady-state model were (a) recharge rates estimated from averaged baseflows of 1994–95, (b) till K that yielded heads near ground surface, (c) adjustment of tunnel valley K between lakes to yield lake levels in agreement with 1994 field measurements; and (d) no leakage into or out of deeper aquifers than till, outwash, or tunnel valley sediments. Table 5 shows the fluid mass balance for steady-state simulation, which agrees to 70.08% between discharge and recharge. Fig. 13 shows calibration between observed and modeled steady-state lake-head elevations (R2 ¼ 0:74), for which highest residuals are for shallow upland wetlands and lowest errors (o1 m) for deep lakes. Fig. 14 (top) shows the steady-state water table;

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the considerable surface irregularities reflect ET in lakes and wetlands. Elk Lake lies downgradient of Elk South and captures its seepage outflow as well as recharge from farmed land and wetlands around it. The Elk Lake chain effectively focuses discharge from till into the Pomme de Terre drainage, driven by the slight evaporative character (negative P  E) of lakes. These results must be qualified. First, steady-state conditions are only an approximation of groundwater– lake interaction, which is transient at time scales from daily to greater than decadal. Second, the recharge rates (and thus K of Zone 1) based on 1994–95 streamflows are certainly only approximate and likely overestimates of average 20th century recharge. Third, the model provides a highly simplified conceptualization of a much more complex hydrogeologic system. However, the steady-state results are consistent with the known locations of groundwater discharge observed in Fig. 11. While there is uncertainty in the precise location of groundwater divides, the catchment areas surrounding lakes and the general configuration of the water table surface are thought to realistically depict timeaveraged shallow groundwater flow under humid (1994–95) conditions. 3.8. Transient simulation of sustained drought based on 1923–37 Transient simulation of a drought modeled after that of 1923–37 was performed starting from the 1994 steady-state heads, decreasing recharge and increasing ET to a time-uniform drought stress. The simulation ran for 50 years and was treated simplistically as a binary

Fig. 12. Distribution of K and boundary conditions for groundwater flow simulation.

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Table 4 Model parameters for steady state and transient groundwater flow simulations Zone

Steady state K (m/day)

1. 2. 3. 4.

Till Chippewa outwash Pomme de Terre outwash Tunnel valley ElkS-Chip Round-Spring Elk-ElkS Spring-Elk Turtle-Pomme de Terre Turtle-Round 5. Lakes

Sy

0.2 300 300 80 300 40 40 40 300 9000

R (m/day)

Transient Max ET (m/day)

5

R (m/day)

3

5

Max ET (m/day)

0.05 0.10 0.10

8  10 1.6  104 1.6  104

1  10 1  103 1  103

4  10 3  105 3  105

1  103 1  103 1  103

0.05 0.05 0.05 0.05 0.05 0.05 1.00

1.6  104 1.6  104 1.6  104 1.6  104 1.6  104 1.6  104 1.7  103

1  103 1  103 1  103 1  103 1  103 1  103 1.9  103

3  105 3  105 3  105 3  105 3  105 3  105 1.2  103

1  103 1  103 1  103 1  103 1  103 1  103 2.0  103

Table 5 Fluid mass balance for steady state and transient (years 1, 15, and 50) simulations Steady state Inflow (m3/day) Change in storage ET Chippewa specified outflow Fixed head cells Recharge Total flux Mass balance error

Outflow (m3/day)

0

Year 1 transient

Year 15 transient

Year 50 transient

Inflow (m3/day)

Inflow (m3/day)

Inflow (m3/day)

Outflow (m3/day)

46,367 21,863

18,654 23,961

20,716

13,835 15,144

42,615

42,579 0.08%

75,346

change from humid to drought recharge and evaporation parameters. Annual stress periods were employed, each converging to 104 m maximum head difference; results (potentiometric levels for lakes, potentiometric surface, and groundwater velocity vectors; see calibration points in Fig. 12) were tabulated after each stress period. Results were robust under varying values of stress period length, convergence parameter, and number of time steps within each stress period. Recharge rates were estimated based on average reduction in precipitation from 1923 to 1937, reduced by half from the steady-state value to 4  105 m/day in Zone 1; 8  105 m/day in Zones 2, 3, and 4; and to 0.0012 in Zone 5; the lake-zone value is the average MAP for the drought period. Maximum ET for lakes was increased to 0.0020 m/day but neither its value nor its extinction depth (103 m/day, 3 m) were changed for land areas. A specific yield of 0.10 was employed for outwash, 0.05 for other glacial sediments, and unity for lakes. For lake beds that had become dry due to water level decline, iterative adjustment was required to reduce recharge rates from lake values (i.e., MAP) to the values for Zones 2, 3, and 4, when simulated water table reached

Outflow (m3/day)

3640 24,001 2000 49,350

75,351 0.01%

17,397 7182 28,219

Outflow (m3/day)

727 10,387 2000 15,964

28,351 0.47%

17,503 6050 24,280

6813 2000 15,606

24,419 0.57%

one meter below local land surface. ET required no adjustment as its flux was attenuated according to depth below water table, reaching zero at 3 m. Thus, as drought proceeded and water levels in and around lakes declined to below the lake bed, first ET gradually declined; then, at depth 1 m, recharge was reduced to land-area values, reflecting revegetation. Therefore, water table declines to below lakebeds tended to diminish the effect of lakes on the hydrologic budget. Boundary conditions in streams were altered from steady-state values. The fixed-head boundary for the Pomme de Terre River was retained but lowered 1 m below steady-state heads; this reflects the assumption that this river continued to convey surface drainage from upgradient and focus groundwater despite the drought. The Chippewa fixed-head condition, on the other hand, was removed, based on the observation that it was dry in 1938 and that it was higher in elevation than the Pomme de Terre. Both interpretations are supported by 1938 photo evidence. To simulate slow recession of Chippewa outwash, a discharge flux of 2000 m3/day was applied at its southern boundary, corresponding to groundwater drainage out of the

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Fig. 13. Calibration of lake heads for steady-state model.

Fig. 14. Hydraulic heads for steady-state modern (1994) groundwater simulation (top) and for transient simulation of drought in 1938 after 15 years (bottom). Vertical exaggeration 50 : 1.

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model area to a discharge area farther south at the prevailing gradient (0.0025) and transmissivity of the Chippewa outwash aquifer. This outwash is discontinuous to the north, so no inflow from that direction was employed. In the simulation, a number of lakes and wetlands dried out during the first 12 years, upon which recharge zones were adjusted from lake (Zone 5) to land values in years 2, 4, 5, 6, 9, and 12. After year 12, only small wetland areas in the centers of Elk, Elk South, Spring, Turtle, Round, Church, Olson, and Thompson remained wet; other wetlands were completely dry. These simulation results are in concert with 1938 air photo observations. Only three surviving lakes (Elk, Thompson, and Turtle) had water deeper than 1.7 m. Resulting transient heads after 15 years of simulated drought (corresponding to 1938) are shown in Fig. 14 (bottom). This surface is a muted facsimile of its steadystate counterpart (top figure), indicating less ET and shallower flow gradients. The water tables decline from pre-drought values was from 3.03 to 4.62 m for lakes (Table 1) and from 0.6 to 8.5 m for the till aquifer on uplands (Fig. 15). Highest aquifer head declines oc-

curred along the divide between the two streams west of Thompson Lake. The lowest drawdowns are in areas distant from lakes and/or close to the Pomme de Terre River. The 1938 Chippewa outwash water table occurs from 3.0 to 6.2 m below its streambed, and thus little ET is lost from this channel. Elk, Turtle, and Thompson Lakes still lose water to ET, but at a much smaller rate due to reduced lake area. The transient head decline for the model agrees well with the trend for the measured Elk South strandlines, with differences of up to 0.63 m (Table 3); this, plus the agreement between Elk South strand line elevations and the modeled lake head (Fig. 16), provide support that model results are robust. Transient head decline for lakes for the 50 year simulation period (Fig. 16) was initially rapid but slowed exponentially. By ca. 40 years, the head decline approaches a new nearly steady-state condition. At 50 years, simulated reduction in lake heads ranged from 6.59 (Turtle) to 7.86 m (Elk South) in the lake chain; only Elk remained above ground, and only by 1.2 m. Thompson (5.2 m) was also still saturated. All other lakes were not only dry but also had declined to several meter below land surface. The fluid mass balance

Fig. 15. Spatial variations of drought-induced change in hydraulic head (meters) after 15 years simulation.

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Fig. 16. Modeled transient changes in lake hydraulic head with time during drought simulation. Large triangles are the Elk Lake South strandline water levels. The aquifer results are shown for a location 1.2 km west of Thompson Lake (top curve).

Fig. 17. Reduction in simulated total evaporative flux for lakes, wetlands, and outwash within the study area during drought.

(Table 5) indicates that recharge flux had steadily declined from 23,961 m3/day at steady state to 6050 at the end of the 50 years simulation. Similar decline was observed for ET (21,863–6813 m3/day). These suggest that ET and recharge were reduced by 69% and 75%, respectively, over pre-drought values. Most of this decline took place in the first 20 years as lakes declined and shrank.

4. Comparison of sediment proxies to model results Both geochemical and ostracode records indicate progressive changes in lake chemistry from about 1923 to 1940 210Pb age. Populations of the three benthic species as well as of Mg/Ca and d13C values all increased

synchronously and gradually as drought continued, suggesting that the lake underwent progressive increase in TDS and eutrophication as the lake shallowed. However, the species assemblage never became dominated by euryhaline or halobiont taxa. Thus while water salinity may be inferred to have increased as suggested by the Mg/Ca proxy, it probably never exceeded about 1400 mg/l TDS. Clearly, the lake was neither at chemical nor at hydraulic steady state at the end of the drought (1938), and was continuing to salinize and decline at a steep rate. Such changes would have continued had the drought persisted longer in time. There is an apparent 2–3 year lag in the proxy response with respect to the drought. This lag time may be real, reflecting time for groundwater recharge to reach the lake, but is also within the limits of

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instrumental uncertainty (e.g., counting error) in the 210 Pb sediment age chronology (73 yr). P. globula decreases in abundance during the 1930s years, consistent with increasing salinity and a reduction of littoral habitat (Fig. 5); however, abundance stays low following the drought. The decline of P. globula during the drought years is understood, but its low productivity following the drought is not. One possibility is that subsequent changes in land use around the lake reduced plant growth and habitat. The most unexpected aspect of the proxy record is the pronounced (2m) negative trend in d18O between about 1923 and 1940 (Fig. 5). This is in spite of the observation that MAP was 5.4 cm lower from 1923–40 than for the 20th century as a whole, and that the peak of the drought (1931–36) was 12.0 cm/yr drier than average (Fig. 4). Lake evaporation was likely higher than normal in this period; thus why did ostracode d18O progressively decline in this period, rather than increase in concentration as would have been the case if evaporative concentration alone controlled its value? There are at least two viable mechanisms for the decline in d18O: (a) variations in either proportions or d18O of water source(s) supplying the lake and (b) a net negative shift in the magnitude of isotopic fractionation caused by evaporation. These two mechanisms are termed ‘‘source effect’’ and ‘‘evaporation effect’’. A decline resulting from ‘‘source effect’’ might reflect changes in either the principal components of the fluid mass balance for the lake (e.g., the ratio of precipitation plus runoff to groundwater inflow) or the d18O of any of these sources. Potential causes of changes in source d18O include seasonality of precipitation and/ or extent of evapotranspiration prior to water reaching the lake. Such a variation in source water d18O due to seasonality was inferred by Gosselin et al. (1997) to explain seasonal changes in both directions for lakes in western Nebraska. As another example, summer ET during humid years has been observed to cause preconcentration of solutes and d18O in wetlands that seep back into Church and Olson lakes (Smith et al., this volume). A decline resulting in ‘‘evaporation effect’’ is in fact one of the results of drought simulation. Phreatic evaporation flux (but not rate) from both lakes and groundwater gradually declined during drought from a high of 21,800 m3/day to a nearly steady-state flux of 6800 m3/day (Fig. 17). The causes of this reduction were (1) decrease of free-water lake surface area resulting from lake level decline, and (2) water table lowering to sufficient depth below land surface to limit ET peripheral to lakes and wetlands. While modeled lake evaporation rates were higher during this simulated drought than during humid years, the model-simulated net evaporation flux at the end of

drought is substantially lower than during wet, highwater periods. Approximately 50% of this change occurs in the first 10 years of drought, and the initial rate of reduction is rapid. This timing of response is in agreement with the observed C. rawsonii proxy response (Fig. 5). 5. Summary and conclusions 1. The proxy indicators in the 20th century core responded to 1923–38 drought in a fashion indicating higher salinity (increasing Mg/Ca, increasing C. rawsoni); however, d18O became lighter during drought, rather than increasing due to increased evaporative concentration. From this, it is inferred that one or more mechanisms, either source- or evaporation-related, acted to decrease net d18O in spite of increased lake-water evaporation. Likely candidates for this light-isotope concentration include seasonality of precipitation and decrease in net evapotranspiration surrounding lakes. 2. A portion of the 15-year 1930s drought record is preserved in 1938 air photos of lake stage declines. For 1938 (end of drought), all lakes had declined from 4.0 to 5.1 m from full-stage levels. All but three were less than 1.7 m in maximum depth or completely dry, as was the Chippewa River and its modern wetlands. The rate of water table decline is indicated by a series of recessional strand lines ascribed to 1931–38. 3. Numerical modeling of groundwater flow to simulate transient response to drought employed a reduced recharge rate (half the 1994–95 rates) and increased lake ET (10% higher than 1994). Model results were calibrated to 15-year declines of nine lake levels (1923–38) as well as to the air-photo strand line levels (1931–38). Agreement between flow model and the air photo lake level record is acceptable. Simulated aquifer response to drought indicates that during the first 10 years, there is rapid exponential water table decline, leveling off as lakes and wetlands dry up and reduce areas and fluxes of ET. Long-sustained drought (e.g. beyond the duration of the 1930s drought) is projected to be capable of inducing up to 6–9 m decline in deep lakes, reaching equilibrium in 40 to 50 years as the water table beneath most lakes receded beyond reach of ET. Modeled ET and recharge fluxes were calculated to have declined by 52% and 71%, respectively, by the first 15 years of drought and by 69% and 75%, respectively, by year 50. 4. Reduction of ET flux during early drought is identified as a potential cause of isotopic decline in lake water d18O, inferred from C. rawsoni shell measurements. This mechanism is based on reduced pre-concentration of d18O in wet-period

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recharge to fringing wetlands that seeps back into Elk Lake. 5. In examinations of paleohydrology, landscape and lake-sediment records may contribute invaluable corroboratory evidence. As demonstrated in the shoreline/air photo records and quantified by model simulations, a switch from humid to sustained uniform drought conditions has a progressive transient effect on lake levels, caused by continued removal of groundwater from storage, mainly by evaporation but also by aquifer drainage near rivers. The geochemical proxy and ostracode species record confirm that salinity also increased in this period. This water level decline and salinity increase was still continuing in 1938 and would have proceeded further had drought persisted; however, simulations suggest an approximate hydraulic steady state would have been reached within about 50 years. The leveling off of drought effects is ascribed primarily to water table declines beneath lakes and wetlands to beyond the reach of evapotranspiration. This drought ultimately induces a reduction in ET flux, even though lake evaporation rates well exceed precipitation during arid years. This attainment of steady state is, therefore, a landscape control, specifically related to the elevations of lake beds that are subaqueous during humid intervals. One intriguing aspect of these Holocene lakes (age ca. o11,000 calendar years B.P.) is that, on a geological timeframe, these sediment basins are rapidly filling and were much lower in elevation in earlier Holocene time. That is, the lakes were much deeper in earlier Holocene time, and this would have altered the characteristics of their control over hydrologic response to drought. This study thus suggests that deeper sediment basins within these lakes could have resulted in a greatly differing response to drought. This paleohydrologic hypothesis is examined in the companion paper (Smith et al., this volume).

Acknowledgements This research was funded by the National Science Foundation (NSF) Hydrologic Sciences grant EAR-9304741. The stable isotope analyses (Ito) were funded in part by grants from the NSF (BIR-9014277 and EAR-9406183) and the University of Minnesota. John Carney and Joan Puller (KSU), Brian Haskell and Eric Gong (UMN), Coburn Wightman (WVU), and Red Westrum (Elk Lake, MN) helped in the field and lab. This is a Paleoecological Reconstructions of Aridity: Interdisciplinary Research InitiativE (PRAIRIE) contribution, and LRC contribution #580.

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