The influence of seasonal precipitation and temperature regimes on lake levels in the northeastern United States during the Holocene

Quaternary Research 65 (2006) 44 – 56 www.elsevier.com/locate/yqres

The influence of seasonal precipitation and temperature regimes on lake levels in the northeastern United States during the Holocene Bryan Shuman a,*, Jeffrey P. Donnelly b a

University of Minnesota Department of Geography and Limnological Research Center, 414 Social Science Building, Minneapolis, MN 55455, USA b Woods Hole Oceanographic Institution, Department of Geology and Geophysics, Woods Hole, MA 02543, USA Received 27 July 2005

Abstract AMS-dated sediment cores combined with ground-penetrating radar profiles from two lakes in southeastern Massachusetts demonstrate that regional water levels rose and fell multiple times during the Holocene when the known climatic controls (i.e., ice extent and insolation) underwent unidirectional changes. The lakes were lowest between 10,000 and 9000 and between 5500 and 3000 cal yr B.P. Using a heuristic moisture-budget model, we explore the hypothesis that changes in seasonal precipitation regimes, driven by monotonic trends in ice extent and insolation, plausibly explain the multiple lake-level changes. Simulated lake levels resulting from low summer precipitation rates match observed low lake levels of 10,000 – 9000 cal yr B.P., whereas a model experiment that simply shifts the seasonality of the modern Massachusetts precipitation regime (i.e., moving the peak monthly precipitation from winter to summer) produces levels that are ¨2 m lower than today as observed for 5500 – 3000 cal yr B.P. The influence of the Laurentide ice sheet could explain dry summers before ca. 8000 cal yr B.P. A later shift from a summer-wet to a winterwet moisture-balance regime could have resulted from insolation-driven changes in the influence of the Bermuda subtropical high. Temperature changes probably further modified lake levels by affecting snowmelt and transpiration. D 2005 Published by University of Washington. Keywords: Lake levels; Holocene; Climate; Moisture balance; Northeastern United States; Ground-penetrating radar; Moisture-budget modeling

Introduction Temporal and spatial patterns of moisture balance can provide clues to the evolution of late-Quaternary climates (e.g., Street and Grove, 1979; Harrison, 1989; Webb et al., 1993a,b; Shuman et al., 2002; Harrison et al., 2003). In the northeastern United States (‘‘New England’’) and adjacent Canada, pollen and lake-level data demonstrate that moisture balance increased and declined multiple times during the Holocene with conditions varying on multi-millennial and shorter time-scales (Fig. 1) (Webb et al., 1993a; Lavoie and Richard, 2000; Newby et al., 2000; Almquist et al., 2001; Shuman et al., 2001, 2004; Muller et al., 2003). By contrast, two progressive (unidirectional) trends were important for shaping conditions during the Holocene: the retreat of the Laurentide ice sheet (Dyke and

* Corresponding author. Fax: +1 612 624 1044. E-mail addresses: [email protected] (B. Shuman), [email protected] (J.P. Donnelly). 0033-5894/$ - see front matter D 2005 Published by University of Washington. doi:10.1016/j.yqres.2005.09.001

Prest, 1987; Barber et al., 1999; Shuman et al., 2002) and the decline in seasonal insolation contrast (Berger and Loutre, 1991; Webb et al., 1993b). Here, we explore the apparent inconsistency between monotonic forcing trends (Figs. 1A, B) and short climatic fluctuations (Figs. 1D –F). We test the hypothesis that progressive shifts in synoptic climatic patterns gave rise to non-linear moisture-balance responses. In particular, we analyze the potential effects of changes in the seasonality of temperature and precipitation on net moisture availability. Today, annual precipitation minus evaporation (P
E) is positive across the eastern United States, but the seasonal distribution of precipitation varies. In the southeastern US, P
E peaks in the summer months, whereas in the northeast, P
E peaks in the winter (Shinker, 1999; Mock, 1996). During the last 2000 years, a southward expansion of spruce populations into northern New England (Webb et al., 1983; Spear et al., 1994), a negative shift in lake water isotopic values (Huang et al., 2002), and a change in Quebec fire regimes (Carcaillet and Richard, 2000) indicate an increase in winter

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Figure 1. New England climates and controls: (A) the seasonal distribution of insolation, represented as a percent change in the difference between summer and winter (Berger and Loutre, 1991), (B) the area of the Laurentide ice sheet (LIS) as a fraction of its area during the last glacial maximum (LGM) (Dyke and Prest, 1987; Barber et al., 1999), (C) the GISP2 Greenland ice core oxygen isotope record of North Atlantic regional temperatures (Stuiver et al., 1995), (D) annual mean temperatures in southern New England inferred by matching modern analogs to fossil pollen data from Rogers Lake, Connecticut (grey) (Davis, 1969), and Winneconnett Pond, Massachusetts (black) (Suter, 1985), (E) hydrogen isotope ratio of sedimentary palmitic acid from Crooked Pond, Massachusetts, which is a record of the hydrogen isotope ratios in the lake water, and (F) the lake-level record from Crooked Pond with its uncertainty in grey (Shuman et al., 2001). At the bottom, letters represent the New England pollen zones that Deevey (1939) and Davis (1969) used to infer the regional climate history. Zones B and C2 represent dry conditions (see also Shuman et al., 2004).

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and spring precipitation. Summer precipitation, however, does not contribute to recharge because P
E is negative during the summer (Fig. 2a). A precipitation increase of ¨20 mm/month is required for summer P
E to become zero. Changes in the timing of precipitation, therefore, can result in significant changes in lake levels, even if the annual precipitation rate does not change (see also Vassiljev, 1998; Vassiljev et al., 1998). Furthermore, isotopes and macrofossil assemblages (Huang et al., 2002; Davis et al., 1980), as well as pollen data (Webb et al., 1993b), indicate that annual mean temperatures in New England were warmer during the mid-Holocene than today by 3– 5-C (Fig. 1). Such temperature changes likely affected total transpiration, and the amount and timing of groundwater recharge related to snowmelt (Stewart et al., 2004; Knowles and Cayan, 2002; Vassiljev et al., 1998). We hypothesize that atmospheric circulation changes, triggered first by the decline of the Laurentide ice sheet (Shuman et al., 2002) and later by the decline in seasonal insolation contrast (Bartlein et al., 1998), not only changed the annual rates of precipitation and evaporation, but also the seasonal distribution of precipitation. Our test of this hypothesis involves two parts, combining new lake-level data with moisture-budget modeling. First, to provide observations for comparison with the modeling, we refine the lake-level history of a highly-permeable, sandy aquifer in southeastern Massachusetts where rapid infiltration enhances the sensitivity to different seasonal precipitation regimes. Then, we use a simplified water-budget model to compare the magnitudes of paleohydrologic change resulting from different climatic regimes. In doing so, we build upon similar modeling studies, such as Vassiljev et al. (1998) in Sweden and Filby et al. (2002) in Minnesota, and our results show differences from those in Europe and the central United States where modern and past precipitation and temperature regimes differ from those in New England. Study area

precipitation and the establishment of the current winter-wet regime in New England. At ca. 8000 –3000 cal yr B.P., the amplified influence of the Bermuda subtropical high due to higher-than-modern summer insolation (COHMAP, 1988; Bartlein et al., 1998) may have shifted the summer-wet regime of the southeastern US further north. Around 11,000 –8000 cal yr B.P., pollen samples from Massachusetts closely match those from the upper Midwest today indicating dry and potentially summer-wet conditions in New England during the Early Holocene (Webb et al., 1993a). Climate model output, however, indicates that the Early Holocene could have been a time of dry summers (Webb et al., 1998). Consequently, the seasonal distribution of precipitation in New England was probably not the same throughout the Holocene, and needs to be better understood. A progressive shift from a summer-wet to winter-wet regime could have had hydrologic consequences because of the relative importance of winter over summer precipitation for groundwater recharge. Positive P
E early in the year today in southern New England results in infiltration of a large percentage of winter

Our analysis focuses on lakes that intersect the water table of the Plymouth – Carver aquifer in southeastern Massachusetts. Our results have broad regional implications, but lakes in different hydrologic settings (i.e., in impermeable till and bedrock) probably did not respond to hydroclimatic changes in the same manner as the lakes of highly permeable Plymouth – Carver aquifer. Late-Pleistocene outwash sands and gravels of the Wareham and Carver pitted plains and kames near Plymouth, Massachusetts, compose most of the 3.63  108 km2 aquifer, which is located along the Atlantic coast to the northwest of Cape Cod (Fig. 3). The sands and gravels sit upon impermeable tills and bedrock, and have a mean horizontal hydraulic conductivity of 6.63  10
2 cm per second (Hansen and Lapham, 1992). Precipitation provides nearly all of the observed recharge of about 1680 m3 per day per km2 (6.09  105 m3 per day in total). There is little surface flow and precipitation infiltrates rapidly. Saturated thickness ranges from 6 to 60 m, with a mean of 32.6 m and a mean specific yield of 0.16. Consequently, the aquifer

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Figure 2. Flow chart with the key variables in a simple hydrologic-budget model for the Plymouth – Carver aquifer of southeastern Massachusetts. Inputs, net radiation, precipitation, and temperature, are in italics and heavy boxes. Plots of monthly variables depict (a) the control scenario (i.e., Massachusetts climatic values) and two scenarios with similar annual precipitation rates as each other: (b) a high summer (modern Minnesota) precipitation rate (dark grey) and (c) a low summerprecipitation rate (light grey and dashed line). Temperature and radiation inputs were the same for all three. Snow cover is represented as its equivalent in mm of water and represents the sum of snowfall during the current month (a temperature-dependent function of precipitation) and existing snow cover minus snowmelt (a temperature-dependent function of snow cover). Discharge rate, not shown, is a linear function of water table elevation (see Eq. (3) in Appendix A). Values represent the 35th year (equilibrium) of a simulation using the same inputs for all years.

contains approximately 1.89  109 m3 of water, with 6.3  105 m3 per day of water discharging from the aquifer to streams, rivers, and the ocean. Newby et al. (2000) and Shuman et al. (2001) showed that, in the past 15,000 years, water levels in the Plymouth –Carver aquifer were lowest between 11,000 and 8000 cal yr B.P. Water levels then rose, but were low again from ca. 5400 to 3000 cal yr B.P., with the wettest conditions of the Holocene during the last 2000 years (Fig. 1). The data match well with patterns observed in New Hampshire (Shuman et al., 2005) and Quebec (Lavoie and Richard, 2000; Muller et al., 2003). Across the northeastern United States and adjacent Canada

general agreement exists among lake level records, although some differences exist due to climatic and hydrologic differences (see Almquist et al., 2001; Shuman et al., 2002 for reviews). Methods In the first part of our study, we present new data on the magnitude and timing of lake level changes in the Plymouth – Carver aquifer, which we then use, in the second part of the study, for comparison with model simulations. The depths of unconformities located near shore in lacustrine sediments

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Figure 3. Site map and aerial photos of Crooked and New Long Ponds showing the location of GPR profiles. Contours of the regional water table (modified from Hansen and Lapham, 1992) plotted at 1.5-m intervals (except in northwest corner where an intermediate 1 m interval is used). Lakes are depicted in grey with only Crooked Pond (CP) and New Long Pond (NLP) labeled. The lower map shows the location of the Plymouth – Carver aquifer (dark grey) and Makepeace Cedar Swamp (MCS; Newby et al., 2000) in southeastern Massachusetts with the white box indicating the location of the upper water table map.

provide the approximate magnitude of past lake level declines (see Digerfeldt, 1986; Dearing, 1997; Shuman et al., 2001). Lake-level changes are the primary explanation for the unconformities (Shuman et al., 2001); local topography and small fetch limits the exposure of these lakes to wind, which could also affect sediment accumulation. A lack of evidence for sub-aerial exposure (e.g., soil development; sediment oxidation) indicates that water levels did not fall below the level of the unconformities; water levels were not significantly higher than the unconformities because wave energy in shallow water is needed to account for associated sand deposits. The depths of peat layers (influenced by different processes than lake sediment) at near-by Makepeace Cedar Swamp (Newby et al., 2000) (Table 1) support our interpretations. To detect unconformities, we collected ground-penetrating radar (GPR) (Moorman, 2002; Shuman et al., 2005) profiles from Crooked Pond where Shuman et al. (2001) studied a transect of cores and from an additional lake, New Long Pond (41.85-N, 70.68-W, 31 m), where we also collected a 10-cm diameter vibra-core. To do so, we used a GSSI SIR-2000 GPR with a 200 mHz antenna, which was towed in an inflated raft behind a small row boat. The range was set between 300 and

600 ns depending on the transect, with an average penetration of >8 m through the very low-conductivity water and fine sediments. New Long Pond is a small (¨10 ha) closed kettle lake located 4.8 km south of Crooked Pond (Fig. 3). AMS radiocarbon dates on organic mater and macrofossils from the New Long Pond core confirm the timing of the inferred lakelevel changes. We used a heuristic box model (Fig. 2) to compare the observed magnitudes of lake-level change with the simulated magnitudes of water-level change that would result in the Plymouth –Carver aquifer from changes in growing-season length, in the seasonal distribution of precipitation, and in the annual precipitation rate. We validated the model by simulating conditions from AD 1960 –1969 before conducting a series of sensitivity analyses to the factors of interest. The model is not hydrologically detailed, and thus, cannot provide specific estimates of past moisture balance change that account for local groundwater controls on the lake levels (e.g., Almendinger, 1990; Donovan et al., 2002; Filby et al., 2002). However, the high permeability of the aquifer enables us to obtain first order patterns by treating the aquifer (including the lakes) as a simple container filled with sand. Most simply, storage of water

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Table 1 New Long Pond AMS radiocarbon samples 14

Lab code

Depth below water surface (cm)

Material

Median

1-sigma range

OS-44431

229 – 230

Wood fragment

4890 T 200

5623

OS-44427

309 – 310

Organics (>500 Am)

8140 T 200

9066

OS-42079

321 – 322

Organics (180 – 500 Am)

8390 T 45

9424

OS-42273

332 – 337

Organics (>500 Am)

9180 T 170

10,369

OS-44313

351 – 352

Tsuga cone scale

9450 T 45

10,683

OS-42080

473 – 476

Organics (>500 Am)

11,700 T 70

13,626

OS-42080

530 – 543

Organics (180 – 500 Am)

11,150 T 70

13,118

5375 – 5331 5770 – 5453 5890 – 5789 8831 – 8776 8884 – 8852 8919 – 8896 9326 – 8930 9352 – 9340 9400 – 9355 9340 – 9326 9355 – 9352 9487 – 9400 10,580 – 10,184 10,636 – 10,615 10,615 – 10,580 10,744 – 10,636 11,035 – 11,031 13,682 – 13,482 13,826 – 13,764 13,193 – 12,991 13,312 – 13,292

in the model aquifer equals any existing storage plus monthly recharge minus monthly outflow (see Appendix A for details). Lake-level results Crooked Pond Like the analyses of cores by Shuman et al. (2001), the GPR profiles from Crooked Pond (CP) show two unconformities (1 and 2) located near shore. Sediment units below the unconformities have been truncated at the bright reflectors 1 and 2. Unconformity 1 extends to a depth of >325 ns or approximately 5.8 m; unconformity 2 extends to >250 ns (5.0 m). The bright reflectors that mark the unconformities correlate with sand layers in near-shore cores. The reflectors that extend across the pond from unconformity 2 correspond with an interval of low loss-on-ignition in the cores from the center of the basin (Shuman, 2003; Shuman et al., 2001). Correlations among cores demonstrate that the unconformities formed (1) between 11,000 and 8000 cal yr B.P. and (2) between 5500 and 3000 cal yr B.P., and that the widespread uppermost sediments are younger than 3000 –2000 cal yr B.P. (Fig. 4). New Long Pond GPR profiles of New Long Pond (NLP) show similar patterns of sediment accumulation to those from CP (Fig. 5). Two prominent reflectors (1 and 2) at depths of 200 and 100 ns (3.8 and 2.5 m) appear to truncate older sediment units. A 10-cm diameter vibracore collected 75 m from shore in 1.1 m of water contains the complete sequence of sediment units shown in the GPR profile at that location (Fig. 5). Two AMS radiocarbon analyses of samples near the base of the core date to 11,700 T 70

C yr B.P.

Cal yr B.P.

Prob.

0.079 0.732 0.189 0.088 0.05 0.036 0.737 0.02 0.069 0.14 0.018 0.841 0.969 0.031 0.233 0.75 0.017 0.833 0.167 0.958 0.042

and 11,150 T 70 14C yr B.P., and have the greatest probability (>0.7) of calibrating to one-sigma ranges of 13,190 –12,990 and 13,680– 13,480 cal yr B.P. based on CALIB 4.4 (Stuiver et al., 1998) (Table 1). The stratigraphic reversal of these two dates demonstrates the possibility of some reworking of coarse organic matter (e.g., charcoal) and the need for caution (and multi-site replication) with respect to the specific timing of events. A 10-cm thick sand layer, which correlates well to unconformity 1, interrupts the fine lacustrine silts of the core at 380 cm below the modern lake surface. An AMS radiocarbon sample from the base of the sand dates to 9180 T 170 14C yr B.P. (10,580 –10,180 cal yr B.P.) An AMS sample from peaty silt above the sand layer dates to 8390 T 45 14C yr B.P. (9490 – 9400 cal yr B.P.) Unconformity 1, therefore, dates to 10,580 – 9400 cal yr B.P., which is consistent with the age of the lower unconformity (1) at CP (Table 2). The dates from NLP refine the chronology for the driest conditions placing them at the same time as peak pine (Pinus) pollen abundance in cores throughout the region (ca. 11,000 –9500 cal yr B.P.) (see Newby et al., 2000; Shuman et al., 2004). An AMS radiocarbon age of 4890 T 200 14C yr B.P. (5770 – 5450 cal yr B.P.) dates a decline in the organic content of the core that correlates to the reflectors forming the base of unconformity 2 (Fig. 5). The age is consistent with the age of the upper unconformity (¨5500 cal yr B.P.) at CP (Table 2). Regional synthesis Past water-level changes were similar across the aquifer based on data from CP and NLP, as well as Makepeace Cedar Swamp (Newby et al., 2000). Peat deposits at Makepeace Swamp interrupt lacustrine silts and muds when unconformities formed near-shore at CP and NLP around 11,600 –8200 cal yr

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Figure 4. GPR profile and cores from Crooked Pond. Diagram of core stratigraphies and locations modified from Shuman et al. (2001). The core and GPR transects are nearly parallel but some minor differences exist. Dashed lines mark correlations between cores based on pollen horizons: second spruce (Picea) peak (S2) at 12,600 cal yr B.P., heath (Ericaceae) peak at 11,000 cal yr B.P., high oak (Quercus) (O) at 8000 cal yr B.P., hemlock (Tsuga) decline (HD) at 5400 cal yr B.P., hemlock rise (HR) at 3000 cal yr B.P., and the ragweed (Ambrosia) rise (RR) at 200 cal yr B.P. Numbers in the GPR profile mark unconformities.

B.P. and 5400– 3000 cal yr B.P. (Table 2). The lower-thanmodern elevations of peat deposits at Makepeace Cedar Swamp (Newby et al., 2000) confirm that the stratigraphic changes at CP and NLP were not shaped by other processes (e.g., redistribution of sediments by changes in wind strength). The magnitude of the inferred water-level decline is similar at all three sites, which are located at high elevations in the aquifer. Based on these data, the water table in the Plymouth – Carver aquifer fell around 11,600 cal yr B.P. to >4 m below the modern level. The water table then rose >1 m after 8000 cal yr

B.P. before dropping again to 2.5 –4 m below modern levels from ca. 5500– 3000 cal yr B.P. Modern levels may be the highest of the Holocene and were probably reached by about 3000– 2000 cal yr B.P. Modeling results The retreat of the Laurentide ice sheet and seasonal insolation changes likely affected multiple aspects of the regional climate conditions that maybe important for explaining the lake-level

50 B. Shuman, J.P. Donnelly / Quaternary Research 65 (2006) 44 – 56 Figure 5. GPR profile and core stratigraphy from New Long Pond. White box on the GPR profile shows the approximate core location. Numbers in the GPR profile mark unconformities. Arrows note the position and most probable 1-sigma age ranges of calibrated AMS radiocarbon dates (Table 1). Black arrows are labeled; for clarity, the in-order ages marked by grey arrows are not shown (see Table 1). Note that depth in the core is not linearly related to time (depth) in the GPR profile because of density differences among silts, sands, and water.

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Table 2 Age bounds for unconformities and peat deposits indicative of low water levels Crooked Ponda

Makepeace Cedar Swampb

New Long Pond

Depth of Unc. Age above

Age below

Depth of Unc. Age above

Unc. 2 >4.0 m age source

3340 – 3000 ca. 5400 14C (D) TSD (D&C)

>2.5 m

Unc. 1 >4.8 m age source

9000 – 8770 11,140 – 10,200 >4.0 m AMS (I) POL (I)

Age below

ca. 3000 6000 – 5200 Interpolated AMS

Depth of peat Age above Peat 2 2.5 – 1.40c m

9500 – 9400 10,780 – 9890 Peat 1 4.4 – 3.5 m AMS AMS

ca. 3000c POL (B)

Age below 5730 – 5480 14C (A)

8020 – 7880 11,690 – 11,350 14C (A) AMS (A)

Letters in parentheses (A, B, D, C, I) indicate which core was conventionally radiocarbon dated (14C), AMS radiocarbon dated (AMS), or dated by pollen stratigraphic correlation to either regional features (POL) or to a radiocarbon age in another core at the same site. ‘‘TSD’’ is a special case of ‘‘POL’’ indicating the well-dated Tsuga decline. ‘‘Unc.’’ means unconformity. a Data from Shuman et al. (2001). b Data from Newby et al. (2000) but see Shuman et al. (2002, 2004) for revised interpretation. c Denotes depth and age of the beginning of peat expansion, indicative of raised water levels.

record. We conducted various sensitivity experiments to reflect possible paleoclimate scenarios inferred from other data (Fig. 1) or from climate model output (e.g., Webb et al., 1998). Validation To test the adequacy of our model (Fig. 2), we simulated water storage from AD 1960 to 1970 using observed temperatures and precipitation for Massachusetts climate division 1 (NCDC, 1994) as input (Fig. 6). The experiment started with total storage equal to the equilibrium value based on the average temperature and precipitation values for Massachusetts climate division 1 for AD 1950– 1995. The simulation produced annual water table fluctuations of approximately 1 m with peak levels in

Figure 6. Model output for a simulation of conditions during the 1960s. Observed temperature and precipitation (top) were used as inputs to produce modeled recharge and water-table elevations (bottom). Simulated water-table elevations (solid line) appear consistent with observations (dashed line and grey bar). Gray bar shows the observed magnitude of drought-related water-table decline in AD 1965 – 66 near Crooked and New Long Ponds based on a map of drawdown in Hansen and Lapham (1992); dashed line shows water table elevation data from an observation well near Plymouth, Massachusetts approximately 7.5 km northwest of the study ponds (Hansen and Lapham, 1992).

the spring and minimum levels in fall and winter, which is consistent with historical observations (Hansen and Lapham, 1992). A drought beginning in 1964 caused a simulated 2.5 m decline in mean water table elevation, which is consistent with observations and other simulations of the drought that estimate a 3– 5 m decline across much of the aquifer for 2 years (Hansen and Lapham, 1992). Mean water levels were lower than those simulated by using the average inputs because precipitation during the 1960s averaged 72 mm/yr less than the 1950– 1995 average. Sensitivity to growing season length Three simulations were run to equilibrium using modern average precipitation values for Massachusetts climate division 1, but with different monthly temperatures to assess the effect of growing season length. First, AD 1950– 1995 average conditions, in which mean temperatures were below freezing for 2 months, were used for the control run (Fig. 7). Second, as may have happened at ca. 4000 cal yr B.P. (Fig. 1), temperatures were increased by 3-C per month, which decreased the proportion of the year that is above freezing by 2 months (Fig. 7A, dashed line). Third, AD 1950– 1995 average temperatures from eastcentral Minnesota (climate division 6) represented a more seasonal climate (as inferred from modern analogs for Massachusetts pollen samples from ca. 11,000 to 8000 cal yr B.P.). In this third case, mean temperatures drop below freezing for a long fraction of the year (4.5 months) (Fig. 7A, solid line). The results depend on changes in the timing of snowmelt and the total annual amount of evapotranspiration. A long above-freezing period (Fig. 7A, dashed) caused snow to melt early and increased recharge in the coldest winter months because not all precipitation fell as snow. Snowmelt was, therefore, reduced. Peak recharge had less volume and occurred earlier in February than in the control. Consequently, the simulated water levels were ¨50 cm below the control levels in spring and summer because although evaporation and outflow rates were similar to the control, the starting point following peak recharge was lower. The reverse occurred under the highly seasonal Minnesota temperature regime (Fig. 7A, solid), which resulted in little winter recharge because precipitation fell as snow. Recharge peaked in March at about

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Figure 7. Modeled sensitivity of the Plymouth – Carver aquifer to changes in (A) growing season length, (B) monthly precipitation distribution, and (C, D) annual precipitation. The control results (based on AD 1950 – 1995 average monthly temperatures and precipitation for Massachusetts climate division 1) are shown in light grey for comparison to all other experiments. Dashed and solid black lines indicate the recharge and water table values resulting from similarly marked temperature and precipitation inputs shown at the top. Monthly precipitation was not varied from the control in the growing season experiments (A), but monthly temperatures were adjusted to Minnesota values (solid black lines) and by adding 3-C per month (dashed). Monthly temperatures and annual precipitation rates were not changed from the control in the precipitation distribution experiments (B), but the precipitation was spread evenly through the months (dashed) and shifted by 6 months from the control (solid black lines). Temperatures were not changed in the annual precipitation experiments, but annual precipitation rate was decreased to 1050 mm/yr by lowering all values by 8 mm/month (C, dashed), to 792 mm/yr by lowering summer rates (C, solid black lines), and to 792 mm/yr by using Minnesota/Wisconsin monthly rates (D, solid black lines).

twice the control volume, thus raising the baseline for spring and summer drawdown. Additionally, because evapotranspiration was reduced, 667 mm/yr of precipitation was stored as compared to 639 and 635 mm/yr in the control and longer growing season simulations. The mean water level rose about 1 m higher than in the control under the Minnesota temperature regime.

evaporated rather than contributing to recharge. High monthly precipitation values in the summer eliminated negative P
E values, and allowed for minimal summer recharge, but this did not make up for the lower-than-control level of recharge earlier in the year. In both cases, simulated water levels were about 2 m lower than the control. Sensitivity to annual precipitation rate

Sensitivity to seasonal precipitation regime To assess the role of the seasonal distribution of precipitation on lake levels, we compared the control simulation with two others that kept mean annual precipitation (1149 mm/yr) constant but shifted the distribution (Fig. 7B). In the first, precipitation was spread evenly through the year at 95.8 mm/ month (Fig. 7B, dashed). In the second, a peak in precipitation equivalent to that observed today in November and December (¨110 mm/month) was shifted to June and July with January and February experiencing the observed minimum (¨77 mm/ month) (Fig. 7B, solid). Both cases reduced total recharge. By shifting winter precipitation to summer, peak precipitation

Webb et al. (1993a) inferred from pollen data that precipitation was low in New England during the Early Holocene. High pine pollen percentages in Early Holocene samples match those in modern samples from the upper Midwest, and indicate that the precipitation rate was approximately 800 mm/yr, >200 mm/yr lower than today. To test the inference made from pollen data, we ran several simulations that used different annual precipitation rates. First, we reduced the control values by a small (<10%) amount, 8 mm/month, to ¨1050 mm/yr to evaluate the impact of a small, but evenly distributed change in annual precipitation rate (Fig. 7C, dashed line). Second, we used a low annual rate of 792 mm/yr but with

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winter monthly values like those found in New England today and with extremely low summer values of 5 mm/month (Fig. 7C, solid line). This second case is consistent with reduced summer precipitation rates in climate model simulations of 9000 cal yr B.P. (i.e., Webb et al., 1998). Third, we used an average of AD 1950– 1995 precipitation values for adjacent Minnesota and Wisconsin climate divisions (MN6 and WI1) (NCDC, 1994), which has an annual rate of 792 mm/yr with peak precipitation during June, July, and August (¨105 mm/ month) and low monthly rates during December, January, and February (29 – 18 mm/month) (Fig. 7D). The Minnesota/ Wisconsin case represents the match between modern upper Midwest pollen samples and early-Holocene fossil pollen samples in New England (Webb et al., 1993a). All three cases caused a significant decline in simulated water levels. Simulations using Minnesota precipitation values (consistent with pollen-based inferences for early-Holocene conditions in New England; Webb et al., 1993a,b) decreased the total amount of recharge severely by both reducing total inputs and by reducing critically important winter precipitation. Water levels dropped ¨20 m below the control (Fig. 7D). By reaching a similarly low annual precipitation rate, but keeping winter precipitation near to modern values (Fig. 7C, solid), the second scenario reduced water levels by only 6 m (and by 5m when coupled with Minnesota temperatures; not shown). Reducing annual precipitation by only ¨10% to 1050 mm/yr (Fig. 7C, dashed), lowered the water level by about 4 m, which

Figure 8. Two possible Holocene climate scenarios. Model output from two series of simulations (solid and dashed lines) that show non-linear water-level responses to linear trends in summer (light grey) and winter (dark grey) precipitation rates. Dashed lines represent a scenario with the additional effect of a change in temperature regime, consistent with a shift from modern Minnesota temperatures at ¨10,000 cal yr B.P. to temperatures that were 1.4-C higher at ¨4000 cal yr B.P. than in Massachusetts today. Water table elevations represent annual mean values. ‘‘LIS’’ means Laurentide Ice Sheet, which was probably a major control on atmospheric circulation over North America before 8000 cal yr B.P. Annual precipitation rate is held constant at 1150 mm/yr from 8000 cal yr B.P. to today in both scenarios.

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is consistent with the observed early-Holocene lake levels. Together with the analysis of sensitivity to the monthly distribution of precipitation, these results confirm that the sensitivity to a small change in winter precipitation equals the sensitivity to a large change in summer precipitation because recharge depends heavily on winter precipitation (see Figs. 7A –D). Discussion The GPR profiles show that lake levels in the Plymouth – Carver aquifer of southeastern Massachusetts changed repeatedly during the Holocene with the driest conditions around 10,000 – 9000 and 5500– 3000 cal yr B.P. Moisture balance shifted in a non-linear fashion despite the long-term, monotonic changes in climatic forcing (Fig. 1). The observed fluctuations may have resulted from a progression of atmospheric circulation patterns that caused changes in annual precipitation (Webb et al., 1993a,b, 1998; Almquist et al., 2001; Shuman et al., 2002), but our simulations demonstrate how non-linear responses to monotonic trends could have arisen since 8000 cal yr B.P. without significant change in annual precipitation rate (Fig. 8). Like Vassiljev (1998) and Filby et al. (2002), we find winter precipitation is important for lake recharge and that interactions between temperature and precipitation, via processes such as the timing of snowmelt, can shape lake level changes. However, unlike in northern Europe and Minnesota, lake level changes in Massachusetts over the past 8000 yr do not require changes in annual precipitation rate and some significant lake level changes were likely the product of reductions in summer moisture. Early Holocene (11,600 – 8000 cal yr B.P.) water levels were 4– 5 m below modern levels. Based on our model simulations, these levels imply at least a small decrease in total annual precipitation because a seasonal redistribution of modern precipitation rates cannot account for the observed decline (compare Figs. 7B and C). Furthermore, if annual precipitation rate had decreased to ¨800 mm/yr as inferred from pollen data by Webb et al. (1993a), then most of the decline in precipitation must have occurred in the summer months because a decline in winter precipitation would produce too large a drop in water levels (compare Figs. 7C and D). Low annual precipitation regimes with peak rainfall in the summer, such as in Minnesota today (Fig. 7D), or with an even distribution of precipitation throughout the year are inconsistent with our lake level estimates. Our simulations show that such precipitation regimes would draw water levels down far more than observed (by >>5 m) because they would substantially reduce groundwater recharge. Lower-than-modern annual precipitation rates with a high proportion falling during the winter (Fig. 7C) do match our lake levels results and are consistent with earlyHolocene lake-water isotope values at Crooked Pond that were more negative than those during the mid-Holocene (Huang et al., 2002) (Fig. 1). Early-Holocene conditions were likely drier than during the late-Pliestocene because temperatures had increased, causing 1) snowmelt to occur early in the year and 2) a prolonged season of transpiration and low recharge.

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By 8000 cal yr B.P., total recharge could have increased and lake levels risen as observed if annual precipitation increased because summer precipitation increased from low early-Holocene levels (Fig. 7C). The result holds even if accompanied by a limited decline in winter precipitation rate (Fig. 8) because slightly positive summer P
E enabled some summer recharge. A positive shift in lake-water isotopic values (Fig. 1) matches with this inference and indicates that precipitation under the new regime peaked in the summer or was at least more evenly distributed through the year than today. If so, lake levels may not have risen to modern levels (i.e., like our analyses of precipitation distribution in Fig. 7B). Such a scenario matches the observation that sediments from 8000 to 6000 cal yr B.P. extend further shoreward than Early Holocene sediments but not as far as recent sediments. (In order for sediments to expand shoreward, the ‘‘sediment limit’’, e.g. Digerfeldt, 1986, must rise—and this limit is largely determined by lake level, see Dearing, 1997). A transition from a summer-wet regime to a winter-wet regime beginning around 8000 cal yr B.P. (Fig. 8) would result in stable or declining aquifer water-levels followed by a rise in levels as precipitation first became evenly distributed throughout the year (i.e., intermediate winter precipitation rates but negative P
E in summer) and then increased in the winter (adding significantly to recharge). An initial decline in water levels would be strengthened if the growing season lengthened (affecting snowmelt and transpiration) (Fig. 8) as would be expected by 5500– 3000 cal yr B.P. given high pollen- and isotope-inferred temperatures (Fig. 1). (A small-magnitude decline in water levels between 8000 and 4000 cal yr B.P. is consistent with Shuman et al.’s, 2001, interpretation that unconformity 2 at CP resulted from a combination of nearly static lake-levels and sediment in-filling). Over the past 3000 years, a decline in isotopic values (Huang et al., 2002), the southward expansion of spruce populations (Webb et al., 1983), and a change in fire regimes (Carcaillet and Richard, 2000) support a possible increase in winter precipitation and a decline in temperatures. Therefore, the high late-Holocene lake levels are consistent with the establishment of today’s positive winter P
E regime. Possible paleoclimate scenarios Early in the Holocene, the presence of the Laurentide ice sheet likely shaped storm tracks, controlled the position of high pressure (and atmospheric subsidence), and prohibited the northward advection of subtropical moisture (COHMAP, 1988; Bartlein et al., 1998; Shuman et al., 2002). Consequently, New England probably received less precipitation than it does today, but probably also a different seasonal distribution of precipitation. A storm track south of the ice sheet could have provided winter precipitation, but high pressure, anomalous northnortheasterly flow, and atmospheric subsidence downstream of the ice sheet could have reduced summer precipitation, as shown in CCM1 simulations (Kutzbach et al., 1998; Webb et al., 1998). Consequently, temperatures (and mean annual

precipitation) in New England could have been similar to those observed in Minnesota today, as estimated from fossil pollen and isotope data (Webb et al., 1993a; Huang et al., 2002), but the seasonal distribution of precipitation was probably heavily weighted towards winter (like Oregon or Labrador today). After the retreat of the Laurentide ice sheet, insolation became the dominant control on North American climate patterns, but the seasonal contrast in insolation decreased progressively (COHMAP, 1988). As a result, the influence of the Bermuda subtropical high was probably amplified at first and then waned. Such a circulation change would have increased and then gradually reduced the northward advection of moist subtropical air during the summer. Winter precipitation rates may have initially declined as ice-sheet-related storm tracks changed, but as winter insolation (and temperatures) increased, greater winter precipitation was probably facilitated. Consequently, New England climates may have shifted from a summer-wet regime shortly after 8000 cal yr B.P. to a winterwet regime by 3000 cal yr B.P. The change in the seasonality of precipitation and temperature would have raised water-levels by 8000 cal yr B.P. (if total precipitation increased from the early-Holocene), decreased water levels at ca. 5400 cal yr B.P. (because precipitation was evenly distributed throughout the year and the growing season was long), and then raised them substantially after 3000 cal yr B.P. (as winter recharge increased again) (Fig. 8). Conclusions Reproducible lake-level histories in southeastern Massachusetts show low levels from 11,000 to 8000 cal yr B.P., intermediate levels from 8000 to 5500 cal yr B.P., low levels again from 5500 to 3000 cal yr B.P., and high levels after 3000 cal yr B.P. Our simple water-budget model demonstrates that such Holocene water-level fluctuations could result from complex seasonal responses to progressive changes in regional climate controls and, thus, in synoptic climate patterns. As the climate shifted between winter-wet and summer-wet regimes, recharge varied non-linearly. More specifically, we infer that, in southern New England, early-Holocene lake levels were low because the presence of the Laurentide ice sheet created dry conditions with peak precipitation in the winter. After the influence of the ice sheet diminished by 8000 cal yr B.P., lake levels rose because annual precipitation increased as summer precipitation became more important. By 5500 cal yr B.P., levels fell again, however, as the growing season lengthened and precipitation had become more evenly distributed throughout the year. Water levels have risen substantially in the past 3000 years because winter P
E has became positive and adds to spring recharge. Acknowledgments Research was supported by a grant (ATM-0402308) from the NSF Earth System History program to B. Shuman. The Ocean and Climate Change Institute at the Woods Hole

B. Shuman, J.P. Donnelly / Quaternary Research 65 (2006) 44 – 56

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Oceanographic Institution provided support for J. Donnelly, and the Ida and Cecil Green Foundation provided a grant to J. Donnelly for the purchase of ground penetrating radar equipment. P. J. H. Richard and another anonymous reviewer provided helpful discussion of the manuscript. We also benefited from the assistance of E. Bryant, W. DeAndrea, J. Hou and P. Newby in the field, and from useful comments on the text from P. Bartlein, T. Webb III, and J. Williams.

freezing fraction of the month when temperature was below zero (no transpiration or open water). Following Darcy’s Law, water flows out of the aquifer at a rate ( QV) linearly proportional to the modern rate ( Q), based on the ratio between the modern observed and calculated groundwater head (H and H V respectively).

Appendix A. Box model description

We calculate the groundwater head by dividing the stored volume of water by the area of the aquifer to calculate a watercolumn height. Changes in head relate to changes in water table height based on the mean specific yield of 0.16 (Hansen and Lapham, 1992). We plot changes in water table height as a deviation from the mean observed saturated thickness. Total storage, calculated monthly, reaches equilibrium after about 35 years. Simplifications in the model include (1) using a constant Bowen ratio (between latent and sensible heating) when apportioning net radiation to calculate monthly evapotranspiration rates, (2) applying a simplified effect of temperature on transpiration (i.e., reduced to zero when temperatures are below freezing), (3) allowing the infiltration of all precipitation without runoff, and (4) not considering variation in evaporative loss due to changes in the net radiation caused by orbital change (Berger and Loutre, 1991) or in wetland and lake area. Such simplifications create uncertainty, but do not prohibit evaluation of the potential for different seasonal regimes to affect aquifer volume.

Our model represents a sand-filled bucket with a drain in which water accumulates to an equilibrium level determined by the balance between monthly inputs and the rate of outflow (Fig. 2). The model subtracts monthly evaporation (E), based on net radiation, and snow accumulation (Sa) from monthly precipitation (P) and snowmelt (Sm) to calculate recharge (Rc) to the system. Rc ¼ P
Sa þ Sm
E

ð1Þ

Snow accumulates without adding to recharge when monthly temperatures drop below freezing. However, not all precipitation during a given (cold) month falls as snow; the proportion of the monthly precipitation that accumulates as snow depends upon the number of days per month below freezing based on daily temperatures estimated by linear interpolation between monthly means. Snowmelt occurs when monthly temperatures rise above freezing but is not instantaneous. The fraction of the accumulated snow that melts during a given month is proportional to the number of days in the month above freezing. If E is greater than P, no water is added to recharge, except by snowmelt. However, negative recharge (in summer months when E > P) is only possible for a fraction (15%) of the aquifer, primarily wetlands and lakes, where evapotranspiration can draw water from the water table. We based the fraction on the extent of wetlands and lakes across the land area of the aquifer today. Monthly precipitation, temperature, and net radiation values were inputs (Fig. 2). Evaporation calculations follow the equation used by Kutzbach (1980)    1 R E¼ ð2Þ 1þB L where B is the Bowen ratio, R is net monthly radiation in Wm
2, and L is the latent heat of evaporation, 0.96 Wm
2 per month per millimeter of water, based on 2.5  106 J/kg. Monthly values of R in Massachusetts were determined from NCEP Reanalysis Project data (Kistler et al., 2001). We used a Bowen ratio of 1.05, based on an empirical assessment of the model, which is consistent with the ratio of sensible and latent heating over land from Sellers (1965): 0.875 for 40– 50-N and 1.580 for 30 – 40-N. Based on the annual balance of precipitation and outflow (P
E equals outflow) and an annual net radiation estimate of 80 Wm
2, B would equal 1.005. E was artificially set to zero when negative (assuming that the precipitation values input into the model account for any condensation). E was also reduced proportionally to the sub-

QV ¼

HV Q H

ð3Þ

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