Environmental influences on speleothem growth in southwestern Oregon during the last 380 000 years

Earth and Planetary Science Letters 279 (2009) 316–325

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Earth and Planetary Science Letters j o u r n a l h o m e p a g e : w w w. e l s e v i e r. c o m / l o c a t e / e p s l

Environmental influences on speleothem growth in southwestern Oregon during the last 380 000 years Vasile Ersek a,⁎, Steven W. Hostetler b, Hai Cheng c, Peter U. Clark a, Faron S. Anslow a,1, Alan C. Mix d, R. Lawrence Edwards c a

Department of Geosciences, Oregon State University, Corvallis, OR 97331, USA U.S. Geological Survey, Department of Geosciences, Oregon State University, Corvallis, OR 97331, USA Department of Geology and Geophysics, University of Minnesota, Minneapolis, MN 55455, USA d College of Oceanic and Atmospheric Sciences, Oregon State University, Corvallis, OR 97331, USA b c

a r t i c l e

i n f o

Article history: Received 18 October 2008 Received in revised form 6 January 2009 Accepted 6 January 2009 Available online 12 February 2009 Editor: P. DeMenocal Keywords: speleothem Oregon U-th dating climate modeling paleoclimate

a b s t r a c t The growth of carbonate formations in caves (speleothems) is sensitive to changes in environmental conditions at the surface (temperature, precipitation and vegetation) and can provide useful paleoclimatic and paleoenvironmental information. We use 73 230Th dates from speleothems collected from a cave in southwestern Oregon (USA) to constrain speleothem growth for the past 380 000 years. Most speleothem growth occurred during interglacial periods, whereas little growth occurred during glacial intervals. To evaluate potential environmental controls on speleothem growth we use two new modeling approaches: i) a one-dimensional thermal advection–diffusion model to estimate cave temperatures during the last glacial cycle, and ii) a regional climate model simulation for the Last Glacial Maximum (21000 years before present) that assesses a range of potential controls on speleothem growth under peak glacial conditions. The two models are mutually consistent in indicating that permafrost formation did not influence speleothem growth during glacial periods. Instead, the regional climate model simulation combined with proxy data suggest that the influence of the Laurentide and Cordilleran ice sheets on atmospheric circulation induced substantial changes in water balance in the Pacific Northwest and affected speleothem growth at our location. The overall drier conditions during glacial intervals and associated periods of frozen topsoil at times of maximum surface runoff likely induced drastic changes in cave recharge and limited speleothem growth. This mechanism could have affected speleothem growth in other mid-latitude caves without requiring the presence of permafrost. © 2009 Elsevier B.V. All rights reserved.

1. Introduction Speleothems are important archives of terrestrial paleoclimate because they can preserve high temporal resolution environmental records derived from chemical and isotopic proxies, and can be precisely dated using uranium-series disequilibrium (Richards and Dorale, 2003; McDermott, 2004). Changes in speleothem growth can also provide an important constraint on paleoenvironmental conditions because their formation requires water availability and elevated CO2 levels in the soil zone above the cave, which in turn are dependent on local water balance, biomass, and temperature (Dreybrodt, 1980; Dorr and Munnich, 1989; Dreybrodt, 1999; Fairchild et al., 2006). Accordingly, the growth of many speleothems at nonglaciated mid-

⁎ Corresponding author. Tel.: +1 541 737 2627; fax: +1 541 737 1200. E-mail address: [email protected] (V. Ersek). 1 Present address: Department of Earth and Ocean Sciences, University of British Columbia, Vancouver, Canada BC V6T 1Z4. 0012-821X/$ – see front matter © 2009 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2009.01.008

and high-latitudes slowed or ceased altogether at various times during the Quaternary because of some combination of increase in the areal extent of permafrost or glaciers, a decrease in partial pressure of soil (pCO2), or changes in water availability (Thompson et al., 1974; Harmon et al., 1977; Gordon et al., 1989; Lauritzen et al., 1990; Baker et al., 1995; Spötl et al., 2002; Genty et al., 2003; Brook et al., 2006). Distinguishing among these potential controls from information on growth changes alone, however, is generally unconstrained, requiring additional proxy information or modeling. Turgeon and Lundberg (2001, 2004) studied two flowstones and one 5-cm long stalagmite from Oregon Caves National Monument (OCNM). A chronology established by 14 thermal ionization mass spectrometry (TIMS) 230Th ages indicated that growth of these three samples occurred primarily during early to mid-interglacial periods, with one interval of deposition during marine isotope stage (MIS) 10. They proposed that growth ceased during glacial intervals because the glaciation limit was lower, favoring permafrost conditions at the cave site. Vacco et al. (2005) studied the δ18O record from a 26.5 cm long stalagmite from OCNM and identified a cold interval

V. Ersek et al. / Earth and Planetary Science Letters 279 (2009) 316–325

that is synchronous with the Younger Dryas (12.5–11.7 ka) period recognized elsewhere in the Northern Hemisphere. The climate of the Pacific Northwest (PNW) is strongly influenced by atmosphere-ocean interactions that originate in the Pacific Ocean (Taylor and Hannan, 1999), implying that understanding of paleoclimate mechanisms at our study site has broad significance to understanding past variations of Pacific storm tracks where they enter North America. We present 46 new 230Th ages from five OCNM stalagmites to further constrain the growth history of speleothems at this location. These ages support previous work (Turgeon and Lundberg, 2001, 2004) in indicating that times of speleothem growth largely correspond to warm (interglacial) conditions. We also find, however, that several growth episodes occurred during cold (glacial) intervals and we investigate possible factors affecting speleothem growth during glacial conditions. Many studies of speleothem growth have relied on regional or global climate records to infer possible environmental controls on speleothem formation (Thompson et al., 1974; Harmon et al., 1977; Gordon et al., 1989; Baker et al., 1995; Spötl et al., 2002; Genty et al., 2003), but this approach can be complicated by uncertainties in the chronologies of the paleoproxies used (e.g. Baker et al., 1993), particularly for periods beyond the limit of radiocarbon dating. In this study, we develop two new approaches to evaluate controls on speleothem growth by combining paleoproxy data with a regional climate model to simulate the climate of the PNW at the peak of the last global glaciation, 21000 years ago (21 ka), and with a thermal advection–diffusion model to calculate the sub-surface temperature evolution at the OCNM site through the last glacial cycle.

2. Regional setting OCNM is situated in the Klamath Mountains of SW Oregon (42°05′ N, 123°25′W), approximately 75 km east of the Pacific coast (Fig. 1). The Klamath Mountains were formed by tectonic accretion along an active continental margin and are divided into four arcuate lithic belts bounded by east-dipping thrust faults (Snoke and Barnes, 2006). The OCNM system is developed in a small Triassic marble lens and has several subhorizontal levels developed along fractures and faults at an elevation of 1150 to 1270 meters. The mean annual air surface temperature (MAT) at OCNM for the period 1981-2006 was 8.8 °C ± 0.7 °C (OCNM internal document). The active cave level is ∼ 60 m below the surface and the cave temperature is approximately constant over the whole year at ∼7 °C. The modern vegetation in the area is

317

dominated by Pseudotsuga menziesii (Douglas-fir) and Abies concolor (white fir). Historical patterns and seasonal variations in temperature in SW Oregon track variations in ocean temperature, resulting in coherent changes of PNW climate west of the Cascade Range. Regional climate responds to contrasts in ocean–continent heating that lead to strong seasonality, with relatively warm and dry summers associated with high pressure over the North Pacific, and cool and wet winters associated with an intensification of the Aleutian low pressure cell (Fig. 1). Along the coast, seasonal temperature extremes are moderated by strong summer ocean upwelling, driven by northerly winds. At interannual-to-decadal timescales, the climate of the PNW demonstrates significant covariance with large-scale dynamic oscillations such as the tropical El Niño-Southern Oscillation, and the highlatitude Pacific Decadal Oscillation (Mantua et al., 1997). Over longer timescales, the geologic record clearly demonstrates significant millennial and orbital variability in Pacific sea surface temperatures (SST) (Ortiz et al., 1997; Mix et al., 1999; Barron et al., 2003). Variability at these timescales perhaps corresponds to sustained spatial patterns in the ocean-atmosphere system similar to those that occur at interannual and decadal timescales. However, other large-scale controls also influenced the climate of the Pacific Northwest, including changes in atmospheric greenhouse gases, the seasonal and latitudinal distribution of solar insolation, the size of North American ice sheets (Bartlein et al., 1998; Hostetler and Bartlein, 1999), and long-distance teleconnections to North Atlantic temperatures through their control on waves in the westerly wind system (Hostetler et al., 1999). The cave hydrologic system responds quickly to atmospheric precipitation. The cave passages have clear seasonal cycles in drip rates, with dry passages during summer and high drip rates during winter and spring. Data from an automatic driplogger installed in the cave show that for at least one cave site, there is a ∼one-month lag between the time of first significant winter-season surface precipitation and reactivation of drips. As a consequence of the strong seasonality of precipitation, historical calcite deposition at OCNM is significantly reduced or terminated during summer when cave drip rates are minimal or nonexistent.

3. Controls on speleothem growth Speleothems are secondary cave carbonates formed by precipitation of calcium carbonate from supersaturated seepage waters (Dreybrodt, 1980; White, 1988; Dreybrodt, 1999). Speleothem growth is largely

Fig. 1. Sea-level pressure climatology (1948–2007) for western North America and eastern Pacific Ocean. A. Winter season (December to February). B. Summer season (June to August). The location of Oregon Caves National Monument is marked by an X in both graphs. Data retrieved and plotted from the NCEP Reanalysis website available at http://www.cdc.noaa.gov/.

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determined by the climate régime above the cave because the chemical kinetics of the carbonate system are a function of water availability, soil pCO2 and temperature (Kaufmann, 2003). On orbital timescales at mid- and high-latitudes temperature is commonly the principal controlling factor of speleothem growth because it determines the physical state of water (liquid vs. ice) and controls biologic productivity (thus influencing soil pCO2 and carbonate dissolution rate). Plant respiration and decay of organic matter in organic-rich soils supply CO2 to percolating waters, leading to elevated pCO2 in the infiltrating solutions. When this slightly acidic water comes in contact with the carbonate bedrock, dissolution occurs until the water approaches saturation with respect to CO2, and its dissociation products HCO3 − and CO3 − (White, 1988). The driving force of speleothem precipitation is the difference between pCO2 in the cave atmosphere and pCO2 in the soil microenvironment and in percolating waters. Because cave pCO2 is usually lower than that of the infiltrating water, dripwaters are generally out of equilibrium with respect to cave pCO2, and the dissolved CO2 is released to the cave atmosphere. Through this process, the dripwater becomes supersaturated with respect to calcium carbonate and carbonate minerals (typically calcite) precipitate to form stalactites, stalagmites, and flowstones, commonly known as speleothems (Moore, 1952). Thus, speleothem growth depends on a steady supply of water that is enriched in CO2, HCO3 − and CO3 −. In certain cases, calcite supra-saturation can also occur due to evaporation or due to common ion effects (Atkinson, 1983; Harmon et al., 1983). 4. Materials and methods 4.1. Chronology Changes in the growth rate of any individual stalagmite may reflect some combination of variability in localized factors along the water path as well as climate (Baker et al., 1996). Accordingly, we collected five stalagmites from different locations in OCNM, as well as include

the three samples from the Turgeon and Lundberg (2004) study. This level of replication in many samples helps to account for possible localized factors, such as random changes in groundwater flow paths, in modulating the climate signal. We sampled our five stalagmites (Fig. 2) by sectioning them parallel to the growth axis. Subsamples of 30 mg to 200 mg for 230Th dating were drilled perpendicular to the growth direction. In addition to sampling within intervals of similar calcite petrography to determine growth variability, we also targeted areas of visible changes in the color and texture of the stalagmites to determine whether these petrographic changes were associated with changes in growth. Stalagmite OCNM02-1 is 26.5 cm long, up to 8 cm in diameter and XRD analyses of powdered samples indicate a composition of 99.9% calcite (Vacco et al., 2005). Between 15 and 16.5 cm from the top, the stalagmite has a distinctly darker interval, with no dissolutional features identified either below or above this area (Fig. 2). Stalagmite OCNM02-2 is 29 cm long, up to 14.5 cm in diameter and is formed of several intervals of large prismatic calcite crystals bounded by intervals of darker calcite that include visible discontinuities (Fig. 2). Stalagmite OCNM05 is 32 cm long, up to 12.5 cm wide, and comprises two coalesced stalagmites (OCNM05-1A and OCNM05-1B) (Fig. 2) that became connected after the dripwater location moved slightly sometime after 119 ka and before 57 ka. Sample OCNM07-1 is 5 cm long and 4 cm in diameter, yellow-brown in color and is formed of columnar calcite crystals (Fig. 2). Thirty-eight 230Th ages were measured at the Minnesota Isotope Laboratory, University of Minnesota on a Thermo-Finnigan Element equipped with a double-focusing sector-field magnet in reversed NierJohnson geometry and a single MasCom multiplier in peak-jumping mode. Eight additional small-size (100 mg or less) samples (OCMN021–15.26 to 16.50 in Table 1) were analyzed on a multi-collector inductively coupled plasma mass spectrometer (MC-ICP-MS, ThermoFinnigan-Neptune) using a MasCom multiplier behind the retarding potential quadrupole (RPQ) in peak-jumping mode; the Neptune has higher sensitivities for both uranium and thorium (close to one order

Fig. 2. Images of speleothems discussed in text. From left to right, these are OCNM05-1A and 1B, OCNM02-2, OCNM02-1 and OCNM07-1. Black lines indicate the locations of samples for 230Th dating, with the exception of the bracket on stalagmite OCNM02-1 which shows the short growth interval during the last glacial period where we had samples at closely spaced intervals (Table 1).

V. Ersek et al. / Earth and Planetary Science Letters 279 (2009) 316–325

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Table 1 230 Th dating results for OCNM speleothems Sample

Depth

238

U

232

Th

230

Th / 232Th −6

δ234U⁎

230

Th / 238U

230

Th age (yr BP)

230

Th age (yr BP)

δ234UInitial⁎⁎

number

(mm)

(ppb)

(ppt)

(atomic × 10 )

(measured)

(activity)

(uncorrected)

(corrected)

(corrected)

OCNM02-1-4 OCNM02-1-0.85 OCNM02-1-1.8 OCNM02-1-2.95 OCNM02-1-3.5 OCNM02-1-4.7 OCNM02-1-6.3 A4⁎⁎⁎ OCNM02-1-8.1 OCNM02-1-10 A7⁎⁎⁎ A9⁎⁎⁎ A12⁎⁎⁎ A13⁎⁎⁎ A8⁎⁎⁎ A10 ⁎⁎⁎ A11⁎⁎⁎ A6 ⁎⁎⁎ A3⁎⁎⁎ OCNM02-1-15.26 OCNM02-1-15.46 OCNM02-1-15.55 OCNM02-1-16.15 OCNM02-1-16.30 A5⁎⁎⁎ OCNM02-1-16.49 OCNM02-1-16.49b OCNM02-1-16.50 A2⁎⁎⁎ A1⁎⁎⁎ OCNM05-1A-4.4 OCNM05-1A-5-7.4 OCNM05-1A-12.2 OCNM05-1A-20.6 OCNM05-1A-5-23.7 OCNM05-1A-29.6 OCNM05-1A-5-38 OCNM05-1B-2.9 OCNM05-1B-8 OCNM05-1B-02 OCNM05-1B-6 OCNM05-1B-5 OCNM05-1B-4 OCNM05-1B-3 OCNM051B-2 OCNM051B-1 OCNM05-1B-21 OCNM07-1-0.5 OCNM07-1-2.45 OCNM07-1-5 OCNM07-1-6.2 OCNM02-2-2.4 OCNM02-2-4.1 OCNM02-2-5.5 OCNM02-2-15.8 OCNM02-2-18 OCNM02-2-19.3 OCNM02-2-19.7 OCNM02-2-22.1

3.0 7.5 14.5 23.0 38.0 44.0 61.5 71.5 78.5 102.0 109.5 120.9 120.9 120.9 133.5 133.5 133.5 138.5 148.6 152.6 154.6 155.5 161.5 163.0 163.6 164.9 164.9 165.0 171.6 244.6 44.0 74.0 122.0 206.0 237.0 296.0 308.0 29.0 80.0 142.0 143.2 152.7 164.2 168.2 187.0 184.8 189.3 5.0 24.5 50.0 62.0 24.0 41.0 55.0 158.0 180.0 193.0 207.0 221.0

90.9 ±0.1 108.0 ± 0.4 101.9 ± 0.2 118.7 ± 0.2 128.5 ± 0.2 67.3 ± 0.1 95.9 ± 0.2 115.0 ± 0.2 86.2 ± 0.2 85.5 ± 0.3 134.2 ± 0.2 126.9 ± 0.3 127.6 ± 0.1 126.1 ± 0.1 159.2 ± 0.3 164.8 ± 0.2 151.1 ± 0.2 103.1 ± 0.2 170.9 ± 0.3 141 ± 1 110.9 ± 0.2 98.2 ± 0.3 136.2 ± 0.4 104.4 ± 0.3 95.3 ± 0.1 92.2 ± 0.2 118 ± 1 105.2 ± 0.3 208.8 ± 0.3 81.0 ± 0.2 227.4 ± 0.3 238.9 ± 0.5 170.1 ± 0.2 179.7 ± 0.3 249.4 ± 0.5 312.4 ± 0.3 241.2 ± 0.5 206.1 ± 0.2 167.8 ± 0.2 245.2 ± 0.3 192.8 ± 0.2 176.5 ± 0.2 152.4 ± 0.2 95.3 ± 0.1 88.7 ± 0.1 83.5 ± 0.1 114.5 ± 0.1 289.6 ± 0.5 322.2 ± 0.4 372.9 ± 0.4 307.4 ± 0.5 259.0 ± 0.4 253.2 ± 0.3 95.1 ± 0.2 133.1 ± 0.2 114.5 ± 0.2 129.3 ± 0.2 198.0 ± 0.4 134.7 ± 0.2

190 ± 20 290 ± 20 190 ± 20 1200 ± 20 870 ± 20 800 ± 20 280 ± 20 150 ± 20 280 ± 20 840 ± 20 630 ± 10 480 ± 10 880 ± 10 1830 ± 10 2310 ± 10 2590 ± 10 2450 ± 20 1320 ± 10 430 ± 20 8450 ± 90 9100 ± 30 9810 ± 110 8970 ± 100 2560 ± 30 2900 ± 20 17780 ± 70 3790 ± 40 5420 ± 60 190 ± 20 260 ± 20 884 ± 5 46 ± 20 1211 ± 16 397 ± 11 48 ± 16 639 ± 6 35 ± 16 437 ± 14 801 ± 15 735 ± 6 742 ± 23 532 ± 12 722 ± 18 644 ± 22 580 ± 19 1061 ± 18 351 ± 14 2919 ± 35 426 ± 14 3533 ± 38 604 ± 15 15140 ± 54 3064 ± 11 6066 ± 19 2013 ± 12 30220 ± 184 9676 ± 31 13197 ± 43 8852 ± 26

53 ± 7 66 ± 6 140 ± 20 52 ± 1 106 ± 3 72 ± 3 330 ± 20 720 ± 100 338 ± 20 131 ± 4 339 ± 8 470 ± 10 263 ± 4 130 ± 1 139 ± 1 131 ± 1 126 ± 2 163 ± 2 830 ± 40 50 ± 1 61 ± 1 59 ± 1 93 ± 3 317 ± 5 259 ± 2 42 ± 1 247 ± 3 162 ± 2 13375 ± 1400 4220 ± 330 3195 ± 20 65000 ± 28000 1760 ± 20 5730 ± 160 65000 ± 22000 6150 ± 60 90000 ± 42000 169 ± 6 170 ± 3 472 ± 6 480 ± 20 670 ± 20 490 ± 10 830 ± 30 970 ± 30 560 ± 10 2600 ± 100 44 ± 1 460 ± 20 111 ± 1 530 ± 10 35.2 ± 0.4 829 ± 4 204 ± 1 875 ± 6 59 ± 1 208 ± 1 247 ± 1 250 ± 1

105.5 ± 2.1 105.3 ± 3.5 100.9 ± 2.2 91.1 ± 2.6 92.3 ± 2.4 94.7 ± 2.5 87.1 ± 2.9 89.1 ± 3.3 102.1 ± 2.8 94.3 ± 4.1 102.5 ± 2.1 111.9 ± 3.7 110.4 ± 1.8 105.4 ± 1.8 129.4 ± 2.1 122.0 ± 1.5 99.0 ± 2.0 134.0 ± 3.0 107.6 ± 2.5 106.4 ± 4.3 107.3 ± 2.3 132.0 ± 2.9 123.3 ± 3.2 123.3 ± 3.2 109.4 ± 3.0 104.8 ± 2.4 118.3 ± 5.4 112.7 ± 2.9 89.4 ± 2.5 130.7 ± 6.9 114.0 ± 1.4 109.4 ± 2.3 112.5 ± 1.4 115.3 ± 1.9 112.3 ± 2.9 111.2 ± 1.2 156.5 ± 2.9 177.7 ± 1.2 198.8 ± 1.3 229.1 ± 1.4 196.3 ± 1.4 205.3 ± 1.4 183.5 ± 1.4 158.2 ± 1.4 153.4 ± 1.2 171.0 ± 1.4 181.3 ± 1.3 138.0 ± 1.8 138.8 ± 1.3 124.2 ± 1.4 119.7 ± 1.3 174.8 ± 1.8 152.7 ± 1.5 98.2 ± 2.2 109.1 ± 1.9 56.6 ± 2.1 72.4 ± 2.1 95.1 ± 2.2 92.7 ± 2.0

0.0067 ± 0.0005 0.0107 ± 0.0005 0.0161 ± 0.0006 0.0314 ± 0.0006 0.0435 ± 0.0007 0.0513 ± 0.0013 0.0583 ± 0.0010 0.0584 ± 0.0015 0.0658 ± 0.0009 0.0782 ± 0.0013 0.0965 ± 0.0011 0.1062 ± 0.0015 0.1096 ± 0.0010 0.1146 ± 0.0010 0.1224 ± 0.0011 0.1251 ± 0.0010 0.1237 ± 0.0013 0.1262 ± 0.0014 0.1271 ± 0.0016 0.1802 ± 0.0013 0.3012 ± 0.0022 0.3592 ± 0.0025 0.371 ± 0.011 0.4712 ± 0.0033 0.4778 ± 0.0030 0.4923 ± 0.0051 0.4793 ± 0.0032 0.5065 ± 0.0033 0.7381 ± 0.0040 0.8060 ± 0.0060 0.7523 ± 0.0025 0.7507 ± 0.0029 0.7613 ± 0.0016 0.7677 ± 0.0019 0.7626 ± 0.0027 0.7623 ± 0.0023 0.8026 ± 0.0031 0.0217 ± 0.0002 0.0493 ± 0.0003 0.0856 ± 0.0009 0.1126 ± 0.0005 0.1217 ± 0.0004 0.1407 ± 0.0005 0.3413 ± 0.0009 0.3828 ± 0.0010 0.4329 ± 0.0010 0.4891 ± 0.0012 0.0268 ± 0.0002 0.0365 ± 0.0002 0.0640 ± 0.0002 0.0626 ± 0.0003 0.1245 ± 0.0015 0.6073 ± 0.0022 0.7898 ± 0.0038 0.8013 ± 0.0034 0.944 ± 0.012 0.9405 ± 0.0051 0.9980 ± 0.0049 0.9971 ± 0.0043

610 ± 50 1000 ± 50 1550 ± 60 3120 ± 70 4360 ± 80 5170 ± 140 5950 ± 110 5950 ± 200 6650 ± 100 8010 ± 100 9920 ± 130 10890 ± 170 11270 ± 110 11870 ± 120 12440 ± 120 12820 ± 110 12960 ± 150 12790 ± 160 13220 ± 170 19300 ± 180 34430 ± 300 4120 ± 400 43400 ± 1600 58600 ± 600 60800 ± 600 63600 ± 900 60360 ± 700 65460 ± 600 120700 ± 1300 131200 ± 2600 119300 ± 760 119600 ± 960 121900 ± 570 123200 ± 700 122400 ± 1000 122800 ± 720 123900 ± 1100 1970 ± 20 4520 ± 30 7810 ± 80 10690 ± 50 11520 ± 40 13700 ± 60 37680 ± 130 43490 ± 160 49640 ± 170 57320 ± 200 2540 ± 20 3500 ± 20 6330 ± 20 6200 ± 30 12100 ± 150 80000 ± 450 134500 ± 1400 135300 ± 1200 232000 ± 10000 217000 ± 4000 241000 ± 5000 243000 ± 4000

460 ± 70 810 ± 70 1420 ± 70 2400 ± 200 3870 ± 150 4320 ± 260 5740 ± 120 5860 ± 160 6420 ± 120 7300 ± 230 9600 ± 140 10620 ± 170 10780 ± 150 10820 ± 250 11420 ± 240 11710 ± 260 11780 ± 290 11900 ± 240 13040 ± 180 17720 ± 810 32300 ± 1100 38600 ± 1800 41700 ± 1800 58030 ± 760 58640 ± 720 58500 ± 2800 59540 ± 780 64130 ± 920 120600 ± 1300 131000 ± 2600 119210 ± 760 119580 ± 960 121760 ± 570 123120 ± 700 122300 ± 1000 122700 ± 720 123900 ± 1100 1920 ± 40 4400 ± 90 7740 ± 90 10600 ± 80 11440 ± 70 13580 ± 100 37520 ± 180 43330 ± 190 49330 ± 270 57240 ± 210 2280 ± 180 3460 ± 30 6090 ± 180 6150 ± 50 10700 ± 800 79700 ± 500 132900 ± 1600 135000 ± 1200 225000 ± 10000 215000 ± 4000 240000 ± 5000 241000 ± 4000

106 ± 2 106 ± 3 101 ± 2 92 ± 3 93 ± 2 96 ± 2 88 ± 3 91 ± 3 104 ± 3 96 ± 4 105 ± 2 115 ± 4 114 ± 2 109 ± 2 134 ± 2 126 ± 2 102 ± 2 139 ± 3 112 ± 3 112 ± 5 118 ± 3 147 ± 3 139 ± 4 145 ± 4 129 ± 4 124 ± 3 140 ± 6 135 ± 3 126 ± 4 190 ± 10 160 ± 2 153 ± 3 159 ± 2 163 ± 3 159 ± 4 157 ± 2 222 ± 4 179 ± 1 201 ± 1 234 ± 1 202 ± 1 212 ± 1 191 ± 1 176 ± 2 173 ± 1 197 ± 2 213 ± 2 139 ± 2 140 ± 1 126 ± 1 122 ± 1 180 ± 2 191 ± 2 143 ± 3 160 ± 3 107 ± 5 133 ± 4 187 ± 5 183 ± 4

λ230 = 9.1577 × 10− 6 y− 1, λ234 = 2.8263 × 10− 6 y− 1, λ238 = 1.55125 × 10− 10 y− 1. ⁎δ234U = ([234U/238U]activity − 1) × 1000. ⁎⁎δ234Uinitial was calculated based on 230Th age (T), i.e., δ234Uinitial = δ234Umeasured × eλ234xT. Corrected 230Th ages assume the initial 230Th/232Th atomic ratio of 4.4 ± 2.2 × 10− 6 except for stalagmite OCNM02-1 where based on the isochron technique (Vacco et al., 2005) the ages were calculated with a ratio of 12 ± 3 × 10− 6. The errors are arbitrarily assumed to be 50%. The error is 2σ error. ⁎⁎⁎Ages reported in Vacco et al. (2005).

of magnitude higher than the Element), with 1 to 2% combined ionization and transmission efficiency for uranium and thorium. The chemical procedures used to separate uranium and thorium for 230Th dating are similar to those described in Edwards et al. (1987). The procedures of characterizing multiplier and instrumental approaches were analogous to those described in Cheng et al. (2000) and Shen et al. (2002). The Neptune measurements are similar to those described by Shen et al. (2002) for the Element, with the following

two differences: (a) the higher ionization/transmission efficiency, and (b) we do not peak jump through background positions because the retarding potential quadrupole removes any significant tailing. 4.2. Climate model simulations We used a regional climate model to simulate the climate of the region during the Last Glacial Maximum (LGM, 21 ka) and the present.

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The regional climate model comprises the RegCM2 atmosphere (Giorgi et al., 1993) and LSX, a full physics land-surface model (Thompson and Pollard, 1995) with six soil layers. The model domain covers most of North America at a 50-km horizontal grid space and includes the North American ice sheets. Boundary conditions for the regional model simulations were obtained from global simulations of the 21 and 0 ka climates that were run with the GENESIS v. 2.3 general circulation model with appropriate large-scale boundary conditions which include continental ice sheets, atmospheric composition (180 ppmV CO2 for LGM and 365 ppmV CO2 for control), and specified monthly SSTs (LGM from Hostetler et al. (2006) and present from Levitus (1982)). The boundary conditions for the RegCM (vertical profiles of temperature, humidity, winds and SSTs) are updated every 6 h of simulation. Further details about the models and the modeling methodology are given elsewhere (Hostetler and Bartlein, 1999; Hostetler et al., 2000, 2006). 4.3. Cave temperature modeling As an additional method to evaluate the role of temperature over a longer time scale we applied a one-dimensional thermal advection– diffusion model to calculate the sub-surface temperature evolution at OCNM through the last glacial cycle. Inputs to the model are the geothermal heat flux from below, and air temperature and perturbations in solar radiation at the surface. Movement of heat in the subsurface is governed by the supply of energy from the earth's interior and the supply of energy from the surface, which depends on the surface energy budget as the advection of heat due to water percolating into the cave. The advection–diffusion equation representing heat transport can be written as: AT AT A2 T ρm c m + wθρw cw =k 2 +Φ At Az Az

ð1Þ

where T is temperature, t is time, ρm is the density of marble (2700 kg m− 3), cm is the specific heat capacity of marble (2639 J K− 1 kg− 1), z is the vertical spatial coordinate, w is advection velocity of percolating water assuming a 30-day surface-to-cave transit time (1.897 × 10− 5 m s− 1), θ is the effective porosity of the marble (0.001), ρw is density of water (1000 kg m− 3), cw is the specific heat capacity of water (4181 J K− 1 kg− 1), k is thermal conductivity (3.0 W m− 1 K− 1), and Φ represents heat sources (orbital insolation anomalies and geothermal heat fluxes). At the topmost grid cell, we allow for a mixture of 80% soil (ρ = 1300 kg m− 3, c = 2254 J K− 1 kg− 1, k = 0.6 W m− 1 K− 1) and 20% water. We use finite differences to implement Eq. (1) over a 1000 m domain with a 1-m vertical grid spacing and a 10-year time step. A small geothermal heat source at the bottom is specified as 90 mW m− 2. The record of variations in SSTs of the California Current from ODP 1020 in the northeast Pacific (Lat: 41°0.051'N, Long: 126°26.024'W) (Herbert et al., 2001) is a surrogate for regional temperature variability at millennial scales and longer. This site lies 200 km to the west of OCNM, and present-day SSTs at the site are within ∼ 2 °C of the MAT at OCNM, reflecting the strong influence of Pacific SSTs on the climate at OCNM. We adjust this SST record uniformly by this 2 °C such that the mean temperature for the last 1000 years equals the current air temperature in the cave. The topmost cell is set to this time-varying temperature as a boundary condition. We also test the possibility that changes in direct solar inputs due to changes in orbital parameters could have led to warmer or colder periods that are reflected in higher or lower cave temperatures. To do this, the time-varying anomaly in June solar energy flux at 40°N (Berger and Loutre, 1991) is applied as an additional heat source to the topmost grid cell. The radiation anomaly in reality is modulated by cloud-cover, surface albedo, and possible shading by vegetation. We cannot quantify these factors, so we allow for a varying strength of the solar radiation signal by sensitivity tests in which we vary the range of

Fig. 3. Age versus depth plot for the stalagmites discussed in text (left panel). Most error bars in the present study are smaller than the symbol size (Table 1). Dashed lines are shown to facilitate comparison between dated stalagmites, and are not intended to indicate continuous growth between individual ages. Samples with high 232Th concentrations that resulted in age inversions (Table 1) are not plotted. Also shown (from left to right) are the 230Th ages obtained by Turegon and Lunderg (2004) (TL (2004)), the global benthic δ18O stack (marine isotopes stages labelled) (Lisiecki and Raymo, 2005), the atmospheric CO2 record from the Vostok ice core (Petit et al., 1999), and summer insolation at 45°N (Laskar et al., 2004). Shaded areas highlight the interglacial periods when most speleothem growth occurred.

the base radiation anomaly from 10% to 70% of the total anomaly. We then use the radiation range in combination with the present-day cave temperature to select those modeled time histories of cave temperature that simulated modern temperatures of 7 to 8.8 °C.

5. Results 5.1. Speleothem growth history We constrain the growth history for each of our five stalagmites from the age-depth model provided by our 46 new 230Th dates as well as the 13 dates previously published in Vacco et al. (2005) (Fig. 3, Table 1). These data constrain two general features of stalagmite growth: intervals of no-to-slow growth and intervals when most growth occurred. The combination with the 17 230Th ages on three speleothems from Turgeon and Lundberg (2004) clearly demonstrates that most growth occurred during interglaciations (Fig. 3). The 17 new 230Th dates from stalagmite OCNM02-1 are all in stratigraphic order, ranging from 0.46 ka to 64 ka (Fig. 3, Table 1). When combined with the 13 previously published 230Th dates for this stalagmite (Vacco et al., 2005) (Table 1), these ages suggest that the majority of the speleothem (94%) grew during MIS 5e and the Holocene (Fig. 3). The remaining 6% of the formation is a darkercolored interval, approximately 1.5 cm thick, which occurs between these two interglacial intervals (Figs. 2 and 3). The youngest age from the MIS 5e interval (120.6 ka at 171 mm) occurs immediately adjacent to an age of 64.1 ka at 165 mm marking the start of the dark-colored interval, suggesting an interval of no-to-slow growth between MIS 5e and MIS 4 (Fig. 3). The contact is paraconformable, and the absence of dissolutional features at this contact suggests that water availability was likely the limiting factor in reducing growth. Ten ages from the dark-colored interval range from 17.7 ka to 64.1 ka, with half of them concentrated around 60 ka (Table 1). These ages indicate that at least some deposition occurring during the Last Glacial Interval (LGI — MIS 2, 3 and 4), but growth was either episodic in response to changing

V. Ersek et al. / Earth and Planetary Science Letters 279 (2009) 316–325

environmental conditions above the cave, or was continuous, but very slow, and the 230Th dates average large intervals of time. The eight 230Th dates from stalagmite OCNM02-2 yielded ages between 10 ka and 241 ka, with 87% of the speleothem vertical growth occurring during interglaciations (MIS 1, 5a, 5e, and 7) and 13% growing during glaciations (MIS 2, 3, 4, and 6) (Fig. 3). The speleothem calcite that grew during the last interglaciation is characterized by a large diameter and coarsely crystalline structure, suggesting abundant water supply and a thick water film on the speleothem surface (Dreybrodt, 1999). Of particular interest is the fact that growth began at 135 ± 1.2 ka, a time when atmospheric CO2 (Petit et al., 1999) and summer insolation at the cave site (Laskar et al., 2004) were still near glacial levels (Fig. 3). The early warming suggested from this interval of growth, however, is similar to records elsewhere that indicate an early start to the last interglaciation (Stirling et al., 1998; Henderson and Slowey, 2000; Spötl et al., 2002; Gallup et al., 2002). There is a visible discontinuity in the speleothem layer structure between late MIS 6 and MIS 7, indicating dissolution of the older layers, which suggests that water undersaturated with respect to calcite was entering the cave at this location during some of the glacial interval of MIS 6. Growth occurred during MIS 7, but an apparent age inversion (Table 1), likely due to contamination by detrital clays as indicated by the high 232Th concentrations of sample OCNM02-2-18 (Table 1), prevents an assessment of growth rates in this interval. The seven 230Th dates for OCNM05-1A range from 119 ka to 123 ka, with all ages in stratigraphic order except for sample OCNM05-1-20.6. Accordingly, this speleothem grew exclusively during MIS 5e. The ten 230 Th dates from stalagmite OCNM05-1B indicate that 81% of stalagmite vertical accumulation occurred during the present interglaciation and the remaining 19% occurring during the last glaciation (Table 1, Fig. 3). The four 230Th dates for OCNM07-1 range from 2.3 ka to 6.1 ka, with all ages in stratigraphic order, indicating that all of this formation grew during the present interglaciation (Table 1). In summary, the compilation of all of our new data indicates that 92% of speleothem vertical accumulation occurred during interglaciations and 8% during glaciations. Although growth of any given individual speleothem was not necessarily continuous through the entirety of any given interglaciation, suggesting some internal hydrological variability in controlling growth, a primary climatic control is clearly indicated when including the results of all five samples (Fig. 3). The combination of our new ages with previously published 230Th dates from OCNM (Turgeon and Lundberg, 2004; Vacco et al., 2005) further supports this strong relation between speleothem growth and climate for the past 380 ka, indicating that interglacial environmental conditions (warm climate, high soil pCO2) were favorable to speleothem formation. However, speleothem growth also occurred intermittently during cold (glacial) periods (MIS 3,4 and 6), demonstrating that at these times temperatures were above 0 °C, processes in the soil were sufficiently active to maintain high pCO2 and promote carbonate dissolution, and water was dripping in the cave.

5.2. Temperature modeling Fig. 4 shows the results of temperature modeling for radiation values ranging between 10% and 70% and with the presence and absence of percolating water. Only the time series with values of 10% and 30% perturbations with dripwater present, and the 30% perturbation with no dripwater approximate the modern cave temperature of ∼7 °C. More importantly, the cave temperature does not drop below the freezing point of water during the entire 100-kyr modeled interval, except for short periods in the extreme 70% perturbation of radiation intensity. Accordingly, these results suggest that cave temperature remained above the freezing point of water throughout the LGI.

321

Fig. 4. Results of cave temperature modeling for the last 100 ka. The stippled line represents the adjusted SST record (see text for explanation) and the continuous line shows the average temperature of the model runs that produced modern cave temperatures of 7 to 8.8 °C. The shaded gray area represents one standard deviation of the average modeled temperatures.

5.3. Climate modeling With the exception of a few months during the spring, our RegCM simulations produce monthly LGM surface air temperatures at OCNM that are about 4 °C colder than present (Fig. 5). The colder temperatures are the result of colder LGM SSTs, lower trace-gas concentrations, and regional atmospheric circulations that are forced by the presence of the North American ice sheets. The influence of the ice sheets on atmospheric flow also induces substantial changes in the water balance at OCNM. Topographic and thermal effects from the ice sheet, accompanied by anticyclonic flow along the southern margin of the Cordilleran ice sheet, produces drier conditions over the Pacific Northwest in all but a few months when the LGM is as wet or slightly wetter than present (Fig. 5). For present day conditions the simulated 2-meter MAT at OCNM is ∼ 10 °C, which is about one degree warmer than the observed value of ∼ 9 °C. At 4.5 m soil depth, the mean-annual ground temperature (MAGT) in the control simulation is 8 °C; the 2 °C difference between ground and air temperature is similar to the observed difference. The simulated MAT for the LGM is 7 °C. Soil temperatures are colder throughout the soil zone than those of the present in response to surface cooling and changes in the surface energy balance (Fig. 6A and C). Deeper soil temperatures, however, remain well above freezing. At the lowest level of the soil zone (4.5 m), the LGM MAGT is ∼ 6 °C, which is about 2 °C colder than that of the present and 1 °C colder than the LGM MAT. Liquid water is reduced substantially at the LGM relative to the control throughout the soil zone (Fig. 6D), whereas from November through April there is a substantial increase in frozen soil water during the LGM (Fig. 6E). The reduction in liquid water in part reflects the overall drier surface conditions during the full glacial, but it also suggests an additional mechanism for limiting deep-soil moisture levels. The simulated annual average precipitation at OCNM is 2.2 mmd− 1 and 3.6 mmd− 1 for the LGM and present, respectively (Fig. 6F and G). In contrast to the 39% reduction in precipitation, LGM runoff is 74% greater than that of the present. This counterintuitive result reflects the interplay of snow melt dynamics and the presence of frozen soils at the LGM. During the LGM, the uppermost soil-zone layers begin to freeze in November and continue to cool into the early winter. Subsequent winter snow, which is roughly double that of the present, accumulates and insulates the frozen ground well into the early spring. The snow also plays a role in suppressing the rise of

322 V. Ersek et al. / Earth and Planetary Science Letters 279 (2009) 316–325 Fig. 5. Monthly anomalies of 2-meter air temperature and precipitation for the LGM and present as simulated by the RegCM regional climate model. The anomalies are for 5-yr averages of each simulation. The location of the cave site is shown by an X on December maps. The location of the LGM North American ice sheets is outlined on each map.

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Fig. 6. Simulated seasonal cycle of soil temperature and moisture at OCNM obtained from the 6-layer soil representation in the LSX surface physics model. The values plotted were obtained by inverse-distance weighting of the four model grid cells nearest to the OCNM location. The total depth of the soil zone is 4.5 m; the depth scale on the ordinate axes is scaled relative to the mid-layer depths (0.025, .10, .25, .55, 1.25, and 3.0 m). Seasonal cycle of soil temperature (oK) for the LGM (A), present (B) and the 21 ka-0 ka anomaly (C). Anomalies of fractional volume of liquid soil water (D) and frozen soil water (E, note different soil depth axis). Monthly time series of 2-meter temperature (TS2), precipitation (Precip), snow water equivalent (SWE), surface runoff (Runoff), and a runoff coefficient (R / (P + SWE) for the LGM (F) and present (G). Temperature is plotted on the left ordinates, SWE, runoff, and the runoff coefficient are plotted on the inside right ordinate and precipitation is plotted on the outside right ordinate. The error bars are one standard deviation (1σ) of the time series over the length of the model runs (8 years).

spring air temperatures in that the high-albedo snow surface reflects solar radiation that otherwise would be available for soil heating (TS2 in Fig. 6F and G). As the snowpack melts, the relatively impermeable soil layer limits infiltration and most of the melt water runs off (Fig. 6F). This dynamic is illustrated by the time series of the ratio of runoff to precipitation and snow water equivalent which peaks at ∼ 1.0. In contrast, the control simulation exhibits no seasonal freezing of the soil zone, less winter snowpack, much smaller runoff ratios (maximum of ∼ 0.2), and thus more infiltration of water into the soil. Together with general LGM drying, the seasonal cycle of the surface water balance, in particular the rapid runoff over soils still frozen in the springtime, provides a feedback mechanism to limit

recharge to the deeper soil and reduce or eliminate drip water in the cave. Warm-season precipitation associated with the Mediterranean climate of the region (wet winter and dry summer) is not sufficient to overcome the soil moisture deficit during the LGM winter and spring. 6. Discussion 6.1. Temperature influence on speleothem growth Permafrost formation above OCNM during glacial intervals can prevent water percolation in the cave and thus stop speleothem formation (Turgeon and Lundberg 2001, 2004). For permafrost to form,

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the MAT should be at most 0 °C, but usually a MAT limit of −3 to −5 °C is required for discontinuous permafrost formation and even lower temperatures are necessary for continuous permafrost (French, 1976; Washburn,1980). Given that the average MAT at OCNM is 8.8 °C ± 0.7 °C, a minimum temperature depression of 9 °C and more likely of at least 10 °C would be necessary to sustain permafrost. The computed LGM temperature of the cave based on the thermal advection-diffusion model is 2.5 °C, a 4.5 °C cooling relative to present for a 4 °C depression in air temperature. Under our assumptions, this represents the plausible extreme in LGM cooling. The absence of freezing temperatures at the LGM is robust to a range of specified changes in radiation intensity. The temperature computations and climate model simulations thus support the assertion of ice-free conditions in the cave, and the analysis of the simulated thermal and hydrologic dynamics in the climate model output provide a mechanism and support for a limited supply of drip water and thus of speleothem growth at the LGM. Similar but less extreme dynamics likely operated during times of intermediate ice volume, atmospheric CO2 concentrations, and SSTs, thus allowing for the intermittent occurrence of some growth at other times during the LGI (Fig. 3). Proxy data from OCNM and the surrounding region support the modeled temperatures. Pollen data suggest that the MAT during MIS 2 was lower by 3–5 °C in southern Oregon and northern California (Thompson et al., 1993; Heusser, 1998; Briles et al., 2005) while a survey of alpine periglacial features indicate that permafrost during the LGI only occurred in limited high-altitude areas of the Cascade Range in Oregon and northern California (Péwé, 1983). 6.2. Precipitation and vegetation influence on speleothem growth Climate model simulations suggest that precipitation rates decreased significantly during the LGM relative to modern (Fig. 6), indicating that water balance was likely an important control on speleothem growth. In particular, maximum reduction in LGM precipitation occurs during the winter months, which is the wet season today when percolation waters enter the cave. Because the LGM changes are largely induced by topographic and thermal effects from the ice sheet, modulation of precipitation by ice-sheet size may be an important factor in controlling the timing of speleothem growth. In addition to the direct effect of water availability on speleothem growth, moisture supply, combined with temperature, also induces changes in biologic productivity, thus influencing soil pCO2 and carbonate dissolution rate. For example, a pollen sequence from Carp Lake, WA (45°55'N, 120°53'W) indicates that for the past 125 ka, vegetation changes at this site were closely related to changes in ice volume, solar insolation, and atmospheric CO2 (Whitlock and Bartlein, 1997), while a pollen record from Bolan Lake, ∼10 km SW of OCNM, indicates that the surrounding area was covered by parkland vegetation during late MIS 2, similar to the modern subalpine environment (Briles et al., 2005). Such a change in vegetation would reduce biomass, and thus soil respiration rates over the cave site, inducing a decrease in soil pCO2 and limiting dissolution of the carbonate bedrock. We thus expect that changes in ice volume, solar radiation, and atmospheric CO2 influence speleothem growth at OCNM through their influence on water balance, seasonal snow cover, and vegetation type. This relation is supported by our data, which demonstrate that greatest growth occurred during interglaciations (minima in ice volume, maxima in solar radiation and atmospheric CO2), and little or no growth occurred during glaciations (Fig. 3). Both OCNM02-1 and OCNM05-1B have growth phases during MIS 3 corresponding to intermediate ice volumes and relatively high summer insolation (Fig. 3). This coincides with growth phases identified during this period in a British speleothem (Baker et al., 1995) and with speleothem growth in the Austrian Alps (Spötl and Mangini, 2002), perhaps suggesting that these MIS 3 boundary conditions favored hemispheric-wide speleothem growth at mid-latitudes.

7. Conclusions Speleothem growth at OCNM provides a first-order indication of past climate conditions because it requires the presence of liquid water in the karst system, and vegetation cover above the cave. The combination of 27 previously published 230Th ages (Turgeon and Lundberg, 2004; Vacco et al., 2005) and 46 new ages from five stalagmites indicates that maximum speleothem growth at OCNM occurred during interglaciations, and little growth occurred during glacial intervals. To evaluate whether speleothem growth was interrupted by a temperature decrease sufficient to establish permafrost, by changes in water availability, or by some combination of these factors, we used a regional climate model to evaluate the magnitude and direction of changes in environmental variables (temperature, precipitation, soil moisture), and an energy balance model to assess the probability of freezing temperatures in the cave. The energy balance model results suggest that for the past 100 ka cave temperature remained above 0 °C. Similarly, the regional climate model simulation of LGM climate indicates that while surface soils froze seasonally, deeper soils remained unfrozen throughout the year. Accordingly, these results do not support the presence of permafrost as a mechanism to reduce speleothem growth during glaciations. On the other hand, the regional climate model suggests that from November to April the topsoil was frozen, including during the time of maximum surface runoff in the spring, which combined with the overall drier LGM climate provided a feedback mechanism that limited recharge to the deeper soil horizons and reduced or eliminated drip water in the cave. Warm-season precipitation associated with the Mediterranean climate of the region was not sufficient to overcome the soil moisture deficit during the winter and spring. This mechanism may have also been an important factor in limiting speleothem growth at other cave sites from mid- and high-latitudes that were affected by relatively large temperature changes on glacial-interglacial timescales, but not severe enough to allow permafrost formation. Our modeling results, which are supported by proxy data, thus indicate that OCNM climate and speleothem growth are sensitive to changes in precipitation and runoff associated with changes in ice-sheet size and radiative forcing. Our finding that speleothem growth was suppressed during glacial intervals without requiring the presence of permafrost resolves an apparent conflict that would have required very large temperature changes on land in spite of a modest (∼3 °C glacial-interglacial change) temperature depression inferred for the nearby North Pacific Ocean (Ortiz et al., 1997).

Acknowledgments We thank John Roth, Elizabeth Hale and Deana DeWire for facilitating our research at OCNM, Xianfeng Wang for help with 230 Th dating, and Sarah Shafer, Joe Stoner, Nick Pisias, Bogdan Onac and four anonymous reviewers for comments and suggestions that helped improve this manuscript. This research was supported by the National Science Foundation Paleoclimate Program and the U.S. Geological Survey.

References Atkinson, T.C., 1983. Growth mechanisms of speleothems in Castleguard Cave, Colombia lcefield, Alberta, Canada. Arct. Alp. Res. 15, 523–536. Baker, A., Smart, P.L., Ford, D.C., 1993. Northwest European palaeoclimate as indicated by growth frequency variations of secondary calcite deposits. Palaeogeogr. Palaeoclimatol. Palaeoecol. 100, 291–301. Baker, A., Smart, P.L., Lawrence Edwards, R., 1995. Paleoclimate implications of mass spectrometric dating of a British flowstone. Geology 23, 309–312. Baker, A., Smart, P.L., Edwards, R.L., 1996. Mass spectrometric dating of flowstones from Stump Cross Caverns and Lancaster Hole, Yorkshire: palaeoclimate implications. J. Quat. Sci. 11, 107–114.

V. Ersek et al. / Earth and Planetary Science Letters 279 (2009) 316–325 Barron, J.A., Heusser, L., Herbert, T., Lyle, M., 2003. High-resolution climatic evolution of coastal northern California during the past 16,000 years. Paleoceanography 18, 1–20. Bartlein, P.J., Anderson, K.H., Anderson, P.M., Edwards, M.E., Mock, C.J., Thompson, R.S., Webb, R.S., Webb, T.I., Whitlock, C., 1998. Paleoclimate simulations for North America over the past 21,000 years: features of the simulated climate and comparisons with paleoenvironmental data. Quat. Sci. Rev. 17, 549–585. Berger, A., Loutre, M.-F., 1991. Insolation values for the climate of the last 10 million years. Quat. Sci. Rev. 10, 297–317. Briles, C.E., Whitlock, C., Bartlein, P.J., 2005. Postglacial vegetation, fire, and climate history of the Siskiyou Mountains, Oregon, USA. Quat. Res. 64, 44–56. Brook, G.A., Ellwood, B.B., Railsback, L.B., Cowart, J.B., 2006. A 164 ka record of environmental change in the American Southwest from a Carlsbad Cavern speleothem. Palaeogeogr. Palaeoclimatol. Palaeoecol. 237, 483–507. Cheng, H., Adkins, J.F., Edwards, R.L., Boyle, E.A., 2000. 230Th dating of deep-sea corals. Geochim. Cosmochim. Acta 64, 2401–2416. Dorr, H., Munnich, K.O., 1989. Downward movement of soil organic matter and its influence on trace-element transport (210Pb, 37Cs). Radiocarbon 31, 655–663. Dreybrodt, W., 1980. Deposition of calcite from thin film of calcareous solutions and the growth of speleothems. Chem. Geol. 29, 89–105. Dreybrodt, W., 1999. Chemical kinetics, speleothem growth and climate. Boreas 28, 347–356. Edwards, L.R., Chen, J.H., Wasserburg, G.J., 1987. 238U-234U-230Th-232Th systematics and the precise measurement of time over the past 500,000 years. Earth Planet. Sci. Lett. 81, 175–192. Fairchild, I.J., Smith, C.L., Baker, A., Fuller, L., Spötl, C., Mattey, D., McDermott, F., E.I.M.F., 2006. Modification and preservation of environmental signals in speleothems. Earth Sci. Rev. 75, 105–153. French, H.M., 1976. The Periglacial Environment. Longman, London. 309 pp. Gallup, C.D., Cheng, H., Taylor, F.W., Edwards, R.L., 2002. Direct determination of the timing of sea level change during termination II. Science 295, 310–313. Genty, D., Blamart, D., Ouahdi, R., Gilmour, M., Baker, A., Jouzel, J., Van-Exter, S., 2003. Precise dating of Dansgaard–Oeschger climate oscillations in western Europe from stalagmite data. Nature 421, 833–837. Giorgi, F., Marinucci, M.R., Bates, G.T., 1993. Development of a second-generation regional climate model (RegCM2). Part I: boundary-layer and radiative transfer processes. Mon. Weather Rev. 121, 2794–2813. Gordon, D., Smart, P.L., Ford, D.C., Andrews, J.N., Atkinson, T.C., Rowe, P.J., Christopher, N.S.J., 1989. Dating of late Pleistocene interglacial and interstadial periods in the United Kingdom from speleothem growth frequency. Quat. Res. 31, 14–26. Harmon, R.S., Ford, D.C., Schwarcz, H.P., 1977. Interglacial chronology of the Rocky and Mackenzie Mountains based upon 230Th-234U dating of calcite speleothems. Can. J. Earth Sci. 14, 2543–2552. Harmon, R.S., Atkinson, T.C., Atkinson, J.L., 1983. The mineralogy of Castleguard Cave, Columbia lcefields, Alberta, Canada. Arct. Alp. Res. 15, 503–516. Henderson, G.M., Slowey, N.C., 2000. Evidence against northern-hemisphere forcing of the penultimate deglaciation from U-Th dating. Nature 402, 61–66. Herbert, T.D., Schuffert, J.D., Andreasen, D., Heusser, L., Lyle, M., Mix, A., Ravelo, A.C., Stott, L.D., Herguera, J.C., 2001. Collapse of the California current during glacial maxima linked to climate change of land. Science 293, 71–76. Heusser, L., 1998. Direct correlation of millennial-scale changes in western North American vegetation and climate with changes in California Current system over the past ~60 kyr. Paleoceanography 13, 252–262. Hostetler, S.W., Bartlein, P.J., 1999. Simulation of the potential responses of regional climate and surface processes in western North America to a canonical Heinrich event. In: Clark, P.U., Webb, R.S., Keigwin, L.D. (Eds.), Mechanisms of Global Climate Change at Millennial Time Scales. American Geophysical Union, Washington, D.C. Hostetler, S.W., Clark, P.U., Bartlein, P.J., Mix, A.C., Pisias, N.G., 1999. Atmospheric transmission of North Atlantic Heinrich events. J Geophys. Res. 104, 3947–3952. Hostetler, S.W., Bartlein, P.J., Clark, P.U., Small, E.E., Solomon, A.E., 2000. Simulated interactions between Lake Agassiz and the Laurentide Ice Sheet 11,000 years ago. Nature 405, 334–337. Hostetler, S.W., Bartlein, P.J., Holman, J.O., 2006. Atlas of Climatic Controls of Wildfire in the Western United States. US Geological Survey Scientific Investigations Report 2006-5139. (http://pubs.usgs.gov/sir/2006/5139). Kaufmann, G., 2003. Stalagmite growth and palaeo-climate: the numerical perspective. Earth Planet. Sci. Lett. 214, 251–266. Laskar, J., Robutel, P., Joutel, F., Gastineau, M., Correia, A.C.M., Levrard, B., 2004. A long term numerical solution for the insolation quantities of the Earth. Astron. Astrophys. 428, 261–285. Lauritzen, S.-E., Løvlie, R., Moe, D., Østbye, E., 1990. Paleoclimate deduced from a multidisciplinary study of a half-million-year-old stalagmite from Rana, northern Norway. Quat. Res. 34, 306–316.

325

Levitus, S., 1982. Climatological Atlas of the World Ocean. NOAA Professional Paper No. 13. 173. Lisiecki, L.E., Raymo, M.E., 2005. A Pliocene–Pleistocene stack of 57 globally distributed benthic δ18O records. Paleoceanography 20. doi:10.1029/2004PA001071. Mantua, N.J., Hare, S.R., Zhang, Y., Wallace, J.M., Francis, R.C., 1997. A Pacific interdecadal climate oscillation with impacts on salmon production. Bull. Am. Meteorol. Soc. 78, 1069–1079. McDermott, F., 2004. Paleo-climate reconstruction from stable isotope variations in speleothems: a review. Quat. Sci. Rev. 23, 901–918. Mix, A.C., Lund, D.C., Pisias, N.G., Bodén, P., Bornmalm, L., Lyle, M., Pike, J., 1999. Rapid climate oscillations in the northeast Pacific during the last deglaciation reflect Northern and Southern Hemisphere sources. In: Clark, P.U., Webb, R.S., Keigwin, L.D. (Eds.), Mechanisms of Global Climate Change at Millennial Time Scales. American Geophysical Union, Washington, D.C. Moore, G.W., 1952. Speleothem— a new cave term. Nat. Speleo. Soc. News 10, 32–33. Ortiz, J., Mix, A., Hostetler, S., Kashgarian, M., 1997. The California current of the last glacial maximum: reconstruction at 42°N based on multiple proxies. Paleoceanography 12, 191–205. Petit, J.R., Jouzel, J., Raynaud, D., Barkov, N.I., Barnola, J.-M., Basile, I., Bender, M., Chappellaz, J., Davis, M., Delaygue, G., Delmotte, M., Kolyakov, V.M., Legrand, M., Lipenkov, V.Y., Lorius, C., Pépin, L., Ritz, C., Saltzman, E., Stievenard, M., 1999. Climate and atmospheric history of the past 420,000 years from the Vostok ice core, Antarctica. Nature 399, 429–436. Péwé, T.L., 1983. The periglacial environment in North America during Wisconsin time. In: Porter, S.C. (Ed.), Late-Quaternary environments of the United States. University of Minnesota Press, Minneapolis. Richards, D.A., Dorale, J.A., 2003. Uranium-series chronology and environmental applications of speleothems. Rev. Mineral. Geochem. 52, 407–460. Shen, C.-C., Edwards, R.L., Cheng, H., Dorale, J.A., Thomas, R.B., Moran, S.B., Weinstein, S.E., Edmonds, H.N., 2002. Uranium and thorium isotopic and concentration measurements by magnetic sector inductively coupled plasma mass spectrometry. Chem. Geol. 185, 165–178. Snoke, A.W., Barnes, C.G., 2006. The development of tectonic concepts for the Klamath Mountains province, California and Oregon. In: Snoke, A.W., Barnes, C.G. (Eds.), Geological studies in the Klamath Mountains Province, California and Oregon. . The Geological Society of America Special Paper 410. Geological Society of America, Boulder. Spötl, C., Mangini, A., 2002. Stalagmite from the Austrian Alps reveals Dansgaard– Oeschger events during isotope stage 3: implications for the absolute chronology of Greenland ice cores. Earth Planet. Sci. Lett. 203, 507–518. Spötl, C., Mangini, A., Frank, N., Eichstadter, R., Burns, S., 2002. Start of the last interglacial period at 135 ka: evidence from a high Alpine speleothem. Geology 30, 815–818. Stirling, C.H., Esat, T.M., Lambeck, K., McCulloch, M.T., 1998. Timing and duration of the Last Interglacial: evidence for a restricted interval of widespread coral reef growth. Earth Planet. Sci. Lett. 160, 745–762. Taylor, G.H., Hannan, C., 1999. The Climate of Oregon: From Rain Forest to Desert. Oregon State University Press, Corvallis. Thompson, S.L., Pollard, D., 1995. A global climate model (GENESIS) with a land-surfacetransfer scheme (LSX). Part 1: Present-day climate. J. Climate 8, 732–761. Thompson, P., Schwarcz, H.P., Ford, D.C., 1974. Continental Pleistocene climatic variations from speleothem age and isotopic data. Science 184, 893–895. Thompson, R.S., Whitlock, C., Bartlein, P.J., Harrison, S.P., Spaulding, W.G., 1993. Climatic changes in the western United States since 18,000 yr B.P. In: Wright, H.E.J., Kutzbach, J.E., Webb, T.I., Ruddiman, W.F., Street-Perrott, F.A., Bartlein, P.J. (Eds.), Global Climates Since the Last Glacial Maximum. University of Minnesota Press, Minneapolis. Turgeon, S., Lundberg, J., 2001. Chronology of discontinuities and petrology of speleothems as paleoclimatic indicators of the Klamath Mountains, Southwest Oregon, USA. Carbonates and Evaporites 16, 153–167. Turgeon, S.C., Lundberg, J., 2004. Establishing a speleothem chronology for southwestern Oregon— climatic controls and growth modeling. In: Sasowsky, I.D., Mylroie, J. (Eds.), Studies of Cave Sediments. Physical and Chemical Records of Paleoclimate. Kluwer, New York. Vacco, D.A., Clark, P.U., Mix, A.C., Cheng, H., Edwards, R.L., 2005. A speleothem record of Younger Dryas cooling, Klamath Mountains, Oregon, USA. Quat. Res. 64, 249–256. Washburn, A.L., 1980. Geocryology. A Survey of Periglacial Processes and Environments. John Wiley & Sons, New York. 406 pp. White, W.B., 1988. Geomorphology and Hydrology of Karst Terrains. Oxford University Press, New York. Whitlock, C., Bartlein, P.J., 1997. Vegetation and climate change in northwest America during the past 125 kyr. Nature 388, 57–61.