Rapid climate change and no-analog vegetation in lowland Central America during the last 86,000 years

Quaternary Science Reviews 38 (2012) 63e75

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Rapid climate change and no-analog vegetation in lowland Central America during the last 86,000 years Alexander Correa-Metrio a, b, *, Mark B. Bush a, Kenneth R. Cabrera c, Shannon Sully a, Mark Brenner d, David A. Hodell e, Jaime Escobar f, g, Tom Guilderson h a

Department of Biological Sciences, Florida Institute of Technology, 150 W. University Blvd, Melbourne, FL 32901, USA Instituto de Geología, Universidad Nacional Autónoma de México, Ciudad Universitaria, México D.F. 04520, Mexico Escuela de Geociencias, Universidad Nacional de Colombia, Sede Medellin, Colombia d Department of Geological Sciences and Land Use and Environmental Change Institute (LUECI), University of Florida, Gainesville, FL 32611, USA e Godwin Laboratory for Palaeoclimate Research, Department of Earth Sciences, University of Cambridge, Cambridge CB2 3EQ, UK f Departamento de Ciencias Biológicas y Ambientales, Universidad de Bogotá Jorge Tadeo Lozano, Bogotá D.C., Colombia g Center for Tropical Paleoecology and Archaeology, Smithsonian Tropical Research Institute (STRI), Panama h Center for Accelerator Mass Spectrometry, Lawrence Livermore National Laboratory, Livermore, CA 94551, USA b c

a r t i c l e i n f o

a b s t r a c t

Article history: Received 19 October 2011 Received in revised form 29 January 2012 Accepted 30 January 2012 Available online 22 February 2012

Glacialeinterglacial climate cycles are known to have triggered migrations and reassortments of tropical biota. Although long-term precessionally-driven changes in temperature and precipitation have been demonstrated using tropical sediment records, responses to abrupt climate changes, e.g. the cooling of Heinrich stadials or warmings of the deglaciation, are poorly documented. The best predictions of future forest responses to ongoing warming will rely on evaluating the influences of both abrupt and long-term climate changes on past ecosystems. A sedimentary sequence recovered from Lake Petén-Itzá, Guatemalan lowlands, provided a natural archive of environmental history. Pollen and charcoal analyses were used to reconstruct the vegetation and climate history of the area during the last 86,000 years. We found that vegetation composition and air temperature were strongly influenced by millennial-scale changes in the North Atlantic Ocean. Whereas Greenland warm interstadials were associated with warm and relatively wet conditions in the Central American lowlands, cold Greenland stadials, especially those associated with Heinrich events, caused extremely dry and cold conditions. Even though the vegetation seemed to have been highly resilient, plant associations without modern analogs emerged mostly following sharp climate pulses of either warmth or cold, and were paralleled by exceptionally high rates of ecological change. Although pulses of temperature change are evident in this 86,000-year record none matched the rates projected for the 21st Century. According to our findings, the ongoing rapid warming will cause no-modern-analog communities, which given the improbability of returning to lower-thanmodern CO2 levels, anthropogenic barriers to migration, and increased anthropogenic fires, will pose immense threats to the biodiversity of the region. Ó 2012 Elsevier Ltd. All rights reserved.

Keywords: Paleoclimatology Climate change Paleoecology Ecological change Central America Last Glacial Maximum Heinrich stadials

1. Introduction Orbital precession, with a periodicity between w19,000 and 23,000 years, is arguably the pacemaker for temperature and precipitation changes in tropical America over long time scales (Hooghiemstra et al., 1993; Leyden et al., 1994; Baker et al., 2001; Bush et al., 2002; Clement et al., 2004). Precession is known to have * Corresponding author. Instituto de Geología, Universidad Nacional Autónoma de México, Ciudad Universitaria, Mexico D.F. 04520, Mexico. Tel.: þ52 155 4518 7651; fax: þ52 55 5622 4281. E-mail address: [email protected] (A. Correa-Metrio). 0277-3791/$ e see front matter Ó 2012 Elsevier Ltd. All rights reserved. doi:10.1016/j.quascirev.2012.01.025

caused climate changes that triggered past migrations and reassortments of tropical biota (Baker et al., 2001; Bush et al., 2002; Groot et al., 2011). The gross scale of Neotropical climatic and ecological change during the last ice age is understood within this context (e.g. Bush and Colinvaux, 1990; Bush et al., 1990, 2004; van der Hammen, 1991; Islebe and Hooghiemstra, 1997; Mayle et al., 2000; Bush and Silman, 2004; Hooghiemstra and van der Hammen, 2004; LozanoGarcia et al., 2005). However, ecosystem responses to millennialscale climate changes and coherence with observed changes in the North Atlantic Ocean remain poorly understood. During the Quaternary, climate changes in the North Atlantic probably affected intermediate and low latitudes through

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rearrangements in the atmospheric and oceanic systems (Peterson et al., 2000; Sánchez Goñi et al., 2002; Lea et al., 2003; EPICA Community Members, 2006; Lynch-Stieglitz et al., 2007; Hodell et al., 2008; Ziegler et al., 2008). Cold stadials and massive ice release in the North Atlantic slowed Atlantic Meridional Overturning Circulation (AMOC), reducing heat transport to the Northern Hemisphere and promoting heat accumulation in the Southern Hemisphere. Conversely, warm interstadials enhanced AMOC heat transport and heat accumulation at high northerly latitudes. This alternating redistribution of heat in the two hemispheres is known as the bipolar see-saw (EPICA Community Members, 2006). Through time, strength of AMOC and its associated interhemispheric heat flows have been associated with migration of the Intertropical Convergence Zone (ITCZ), which largely controls the seasonal distribution of tropical precipitation. This atmospheric feature is the ascending limb of a larger circulation formed by the Hadley cells, and although it never reaches northern Central America, the moisture that ascends from it provides wet season rains in the region (Waliser and Gautier, 1993; Waliser et al., 1999). Changes in the inter-hemispheric temperature gradient during the late Quaternary affected the mean position and the intra-annual migration of the ITCZ. Cold conditions in the North Atlantic kept convective activity further south in the tropics, whereas warm conditions facilitated its migration to more northerly positions (Peterson et al., 2000; Wang et al., 2004; Hodell et al., 2008). The Atlantic Warm Pool (AWP) exerts a further influence on Central American climate, acting as a convective source and adding to the flow of moisture and heat onto land during summer (Wang et al., 2006; Wang and Lee, 2007). Weakening (strengthening) of AMOC would be expected to reduce (expand) the geographic extent of the AWP, decreasing (increasing) precipitation and temperature regimes in the Central American lowlands and Caribbean (Hodell et al., 2008). An understanding of the past linkages among these atmospheric and oceanic features and Central American climate can provide insights that will enable creation of regional climatic scenarios associated with global change. Ecological responses of tropical ecosystems to environmental change are generally reported as rates of assemblage change caused by climatic variability, which take place over various time scales. Climate changes caused by the slow, steady pace of precession amount to a “press” that exerts constant pressure over extended time periods (e.g. Bush and Colinvaux, 1990; Leyden et al., 1994). In contrast, fast-acting “pulses” of rapid climate change, such as Heinrich events during the last glacial (Correa-Metrio et al., in press), cause ecosystems to respond quickly and adapt to new conditions. Given the connections between tropical climate and AMOC anomalies induced by DansgaardeOeschger cycles and Heinrich Stadials (HS) (Peterson et al., 2000; Hodell et al., 2008), the last glacial provides a context in which to investigate the importance of climatic “pulses” and “presses” on Neotropical vegetation. The most valuable paleoecological records reflect “real-time” migration of climate-sensitive species as they maintain equilibrium with their bioclimatic envelope in response to climate change (Williams and Jackson, 2007). The best settings in which to document such responses are characterized by high biodiversity and short migration distances, e.g. tropical lowland environments that abut steep mountain slopes (Colwell et al., 2008). Records of climate-driven vegetation migration are preserved as pollen spectra in lake sediments. It is rare, however, to find low-elevation tropical lakes that experienced continuous, high sedimentation rates throughout the last glacial cycle. The paleoenvironmental archive contained in the sediments of Lake Petén-Itzá, in lowland northern Guatemala, thus provides a unique opportunity to address

both climatic and ecological questions at high temporal resolution (Hodell et al., 2006). Here we used pollen from Lake Petén-Itzá to investigate the effects of DansgaardeOeschger cycles and HSs on the lowland vegetation of Central America. In conjunction with studies of modern pollen (Correa-Metrio et al., 2011a), the fossil pollen data provided a means to refine the quantification of temperature changes in the Central American lowlands during the last glacial. 2. Setting and background 2.1. Regional setting The Yucatan Peninsula is influenced by the descending limb of a Hadley cell that is centered at w20 N (Waliser et al., 1999). Consequently, the air is warm and dry for much of the year. Whereas mean annual air temperature is relatively constant from year to year, w25  C, precipitation varies from 900 to 2500 mm/yr, with a regional mean of w1600 mm/yr (Deevey et al., 1980). During the boreal summer, northward migration of the ITCZ provides the moisture and atmospheric instability that generates strong convective rains. The Caribbean Low Level Jet transports moisture from the Caribbean AWP and produces most of the precipitation that falls on the Yucatan Peninsula between June and October (Mestas-Nuñez et al., 2007). During winter, even though moisture from the Caribbean is diverted southward, polar air masses bring light sporadic rains into the area (Bradbury, 1997). Lake Petén-Itzá is located in the lowlands of northern Guatemala (16 550 N, 89 500 W), at an elevation of w110 m above sea level (Fig. 1). The lake is formed by a series of karstic solution basins and has a maximum depth of w165 m (Hodell et al., 2006; Mueller et al., 2010). Direct rainfall, runoff, and subsurface groundwater provide inputs to Lake Petén-Itzá, a waterbody that lacks a surface outflow (Hodell et al., 2008). The nearest highlands lie w60 km to the east, in Belize, and w100 km to the south, in Guatemala, and display an altitudinal temperature gradient that spans 5e6  C. The modern climate of the Yucatan lowlands exhibits a strong north-tosouth precipitation gradient, from w500 to 3200 mm/yr over a distance of w400 km (Correa-Metrio et al., 2011a). These steep climatic gradients provide an ideal setting in which to investigate vegetation response to both abrupt and persistent, long-term climate changes, i.e. “pulses” and “presses.” Additionally, given the great depth of Lake Petén-Itzá, it did not desiccate even during the driest climate periods of the late Quaternary, and thus contains a long record of continuous sediment deposition (Hodell et al., 2008; Mueller et al., 2010). Furthermore, the lake remains thermally stratified through much of the year, at least at present, and hypoxic/anoxic conditions in deep-water create ideal conditions for preservation of pollen and other organic microfossils. These factors made Lake Petén-Itzá an excellent target for palynological investigation of climate change. 2.2. Peten-Itza Scientific Drilling Project Following detailed seismic surveys in 1999 and 2002 sediment cores were retrieved in 2006 from seven deep-water locations in Lake Petén-Itzá, using the GLAD-800 drilling platform (Hodell et al., 2006, 2008; Mueller et al., 2010). Site PI-6, in 71 m of water, was selected for analysis because of its w95% sediment recovery over the last w86,000 years and its relatively high mean rate of sediment accumulation, w88 cm per millennium (Hodell et al., 2008). Lithologic variability, expressed by density and magnetic susceptibility (MS), showed gypsum precipitation during lake lowstands, i.e. dry episodes, and deposition of clays during periods of high lake level, i.e. periods of more abundant rainfall (Hodell et al., 2008). The

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Fig. 1. Geographic location of Lake Petén-Itzá (solid star) relative to elevation of the Yucatan Peninsula and adjacent mountains. Inserts: a) Lake Petén-Itzá (solid star) and Colombia Basin (hollow star) in the context of Central America; b) Bathymetric map of Lake Petén-Itzá showing the location of coring sites. PI-6 is shown by a star (from Hodell et al., 2008).

During the Last Glacial Maximum (21  2 ka) (Mix et al., 2001), the vegetation was dominated mainly by Pinus, Quercus and Myrica (Bush et al., 2009; Correa-Metrio et al., in press), and the changes that occurred within this chronozone were relatively unimportant

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sediment lithology indicated reduced lake levels during Greenland stadials and high lake levels during Greenland interstadials. These findings were consistent with southerly migration of the ITCZ during times of enhanced meltwater discharge from the Laurentide ice mass and sea ice cover in the North Atlantic (Peterson et al., 2000; Chiang and Bitz, 2005; Hodell et al., 2008). An age-depth model for core PI-6 was originally developed using AMS 14C dates from 21 depths in core PI-6 and nearby core PI3 (Hodell et al., 2008). Ages from the latter section were projected onto depths in PI-6 by inter-core correlation, using the highresolution magnetic susceptibility records. The PI-6 chronology was refined in this study (Fig. 2), using 18 new AMS 14C dates from core PI-6, and depth-correlated cores PI-2 and PI-3 (Correa-Metrio et al., in press). All AMS 14C dates were run on samples of terrestrial organic matter to avoid hard-water-lake error (Deevey and Stuiver, 1964), which can confound dates on bulk organic matter from water bodies in this karst area (Hodell et al., 1995; Curtis and Hodell, 1996). Dates were calibrated using Oxcal-Intcal09 (Reimer et al., 2009). The chronology for the last w43 ka (All ages are calibrated and expressed as ka ¼ thousands of years before present) was derived by linear interpolation between selected data points, which are fairly evenly distributed along the uppermost w45 m of the core (Fig. 2). Beyond the reach of radiocarbon dating, three ash layers were identified and their ages used to derive the chronological model (Hodell et al., 2008): Congo tephra (53  3 ka at 52.48 cm), Guasal1 (c. 55 ka at 53.5 m), and Los Chocoyos tephra (84  0.5 ka at 79.99 m). Pollen analysis showed that regional vegetation during the last glacial consisted of a mix of tropical and temperate species that coexisted throughout most of the time period (Correa-Metrio et al., in press). Pine-oak forests that dominated during glacial time were replaced by herb- and shrub-dominated vegetation during HSs.

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Age (calibarted years BP) Fig. 2. Chronology for core PI-6. Ages are calibrated AMS 14C dates on terrestrial organic matter from core PI-6 (hollow triangles), core PI-3 (hollow diamonds), and core PI-2 (solid circles). Dates from cores PI-3 and PI-2 were projected onto core PI-6 by high-resolution, inter-core correlation of the magnetic susceptibility records (after Correa-Metrio et al., in press; and Hodell et al., 2008).

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compared with HSs (Correa-Metrio et al., in press). Overall, the most dramatic vegetation change of the last 86,000 years was that which occurred at the PleistoceneeHolocene transition. Previous palynological study of surface sediments from 81 lakes on the Yucatan Peninsula and in the mountains of Guatemala and Mexico allowed the elucidation of pollen-vegetation-climate relationships in the area (Correa-Metrio et al., 2011a). Temperature strongly influences modern pollen assemblages, allowing construction of a pollen-temperature transfer function that used a novel technique called Synthetic Assemblages (SyAs). Because pollen types used in constructing the transfer function were the same as those found in sediment samples at depth in the core, it was possible to reconstruct temperature changes through time. Additionally, modern biogeographic patterns were distinguishable in the modern pollen spectra using multivariate techniques (Correa-Metrio et al., 2011a). Therefore, the modern data set proved to be suitable for quantifying similarity between fossil and modern pollen assemblages through the use of the modern-analog technique (MAT, sensu Overpeck et al., 1985). 3. Methods Four hundred and forty-five samples from PI-6 core were analyzed for pollen and charcoal, yielding a mean sampling resolution of 190 years, i.e. approximately one tree generation (Hartshorn, 1978). Samples for pollen analysis (0.5 cm3) were prepared according to standard protocols (Faegri and Iversen, 1989) and gravimetrically separated to concentrate pollen and spores (Krukowski, 1988). A Lycopodium clavatum pellet with approximately 18,500 spores was added to each sample to allow calculation of pollen concentration (grains per cm3) (Stockmarr, 1972). Samples were analyzed at magnifications of 400 and 1000, using a transmitted light microscope. To avoid pollen counts dominated by a few, very abundant, primarily anemophilous taxa, Cyperaceae, Moraceae, Pinus, and Quercus were quantified, but excluded from the pollen sum (after Birks and Birks, 1980). Counts were made until a pollen sum of 200 grains or 2000 Lycopodium spores (w10% of the sample) were enumerated. This sampling effort and exclusion of anemophilous taxa was consistent with the modern pollen survey that successfully replicated the biogeographic and climatic patterns of the Yucatan Peninsula and adjacent mountains of Guatemala and Mexico (Correa-Metrio et al., 2011a). Pollen data were expressed as percentage of the pollen sum. A 0.5-cm3 sediment sample from each depth analyzed for pollen was processed to recover charcoal particles. A digital picture of the sample was taken with a stereomicroscope (Clark, 1988) and the number of pixels covered by charcoal particles was counted using ImageJ (Rasband, 2005). Charcoal area (cm2) was standardized by analyzed volume and expressed in terms of concentration (cm2/cm3). Detrended correspondence analysis (DCA) (Hill and Gauch, 1980) was performed to evaluate vegetation turnover through time, and its association with environmental factors. In the region of the Yucatan Peninsula and adjacent mountains, the rescaled space generated by the DCA scores reflected the ecological envelope defined by the communities represented in the pollen assemblages (Correa-Metrio et al., 2011a). Consequently, Euclidean distance among DCA scores of fossil samples represents differences in composition and structure of pollen assemblages in standard deviations (SD), with 50% of vegetation turnover occurring within one SD (Gauch, 1982). Thus, ecological change was calculated as the Euclidean distance between contiguous samples, calculated using the first four DCA axes (Orlóci et al., 2006; Urrego et al., 2009). Rates of ecological change were derived by dividing the distance between two adjacent samples by the time elapsed between them.

Modern analogs were evaluated by calculating the squaredchord distance between fossil pollen spectra and those derived from modern mudewater interface samples (Overpeck et al., 1985, 1992). To the extent possible, the lakes selected as modern analogs lay within relatively undisturbed vegetation (Correa-Metrio et al., 2011a), minimizing the impacts of human disturbance on our analyses. The modern pollen spectra showing the minimum distance to a given fossil sample was considered as the most likely analog. For North America, minimum squared-chord distances >0.15 (e.g. Overpeck et al., 1985) or >0.35 (e.g. Gill et al., 2009) have been considered indicators of no-analog vegetation. In this study, we compared the three youngest samples to the modern pollen spectra. Mean distance between these samples and the modern assemblages was 0.31, which we adopted as a conservative threshold for no-analog assemblages. Charcoal concentration was regressed against magnetic susceptibility in core PI-6 (Hodell et al., 2008), DCA Axis 2 scores of the PI-6 pollen data, summer-winter insolation difference at 17 N, and d18Oice isotopic data from the NGRIP record (NGRIP Members, 2004; Andersen et al., 2007). Insolation difference was calculated by subtracting winter mean (from December 21st to March 21st) from summer mean (from June 21st to September 21st), as calculated with Analyseries 2.0 (Paillard et al., 1996). The regression was fitted through a Poisson generalized linear model (Gelman and Hill, 2007) and did not consider interactions among independent variables. This technique is only appropriate for count data, and charcoal concentration usually results in quantities below one. Therefore, charcoal data were taken to mm2/l by multiplying by 1000 and rounding to the nearest whole value. The other independent variables were used in their original scale. The temperature-climate transfer function developed by Correa-Metrio et al. (2011a) was applied to the fossil pollen sequence to produce temperature estimates from 85.5 ka to present. The approach uses non-parametric regressions that predict specific pollen percentages as a function of a given environmental parameter (Correa-Metrio et al., 2011a). Having predicted percentages for 30 taxa that proved to be responsive to temperature, synthetic pollen assemblages were constructed for the temperature gradient from 13.5 to 26.3  C, at increments of 0.25  C, yielding a total of 52 SyAs. Subsequently, fossil samples were compared to the SyAs using Canberra distance, and selecting for the particular time slice in which the temperature that minimized the dissimilarity index (see Correa-Metrio et al., 2011a for further details). Cross-validated error, using modern samples for temperature estimations, was w1.45  C. All statistical processing was performed using R (R Development Core Team, 2009), especially packages paleoMAS version 2.0-1 (Correa-Metrio et al., 2011b) and vegan version 1.17-3 (Oksanen et al., 2009). 4. Results Pollen counts varied between 45 and 3050 grains (mean 745 grains, 99% of the samples had counts between 258 and 2154 grains). Three samples (0.7% of the total) had pollen counts below 200 grains (44, 97, and 198 grains), all of them located in sections of the core composed almost exclusively of carbonates, which usually contain very low pollen concentrations (Heusser and Stock, 1984). However, given that the sampling effort was consistent throughout all the samples (at least 10% of the volumetric sample when low pollen concentrations were present), all samples were considered representative of the conditions under which they were deposited. A total of 177 pollen types and 21 spores were identified. Tropical lowland and montane taxa, as well as temperate elements, coexisted throughout the entire PI-6 record, but displayed alternating dominance (Fig. 3). The pollen-temperature transfer function was

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based on 30 taxa that had significant responses to temperature in the modern pollen study (Correa-Metrio et al., 2011a). This subgroup of pollen taxa represented between 23.7 and 95.8% of pollen counts (mean 70.1%, between 35.8 and 91.0% of the total counts for 95% of the samples). Charcoal concentrations varied between 0 and 89.9 cm2/cm3 (mean 1.30, with 95% of the observations between 0 and 5.43 cm2/cm3). Axes 1 and 2 of the DCA of samples were 2.29 and 1.57 standard deviations (SD) of species turnover in length, and had eigenvalues of 0.22 and 0.11, respectively. Tropical elements such as Brosimum, Bursera, Cecropia, Ficus, Moraceae, and Trema had the highest scores on DCA Axis 1, while temperate and montane elements, e.g. Juglans, Myrica, Quercus, and Pinus, were associated with the lowest scores. DCA Axis 2 scores seem to have been negatively associated with moisture availability because known mesic taxa such as Juglans, Melastomataceae, Myrica, Quercus, Sapium, and spores had negative scores, while drought-associated taxa Acacia, Amaranthaceae, Ambrosia, Asteraceae, Byrsonima, and Hymenaea had higher scores. Although Axes 1 and 2 sample scores were correlated significantly with MS, the correlation with Axis 1 was very low (0.12, p < 0.05, d.f. ¼ 441), while that of Axis 2 was much stronger (0.41, p < 0.001, d.f. ¼ 441). DCA-derived rates of ecological change varied between 0.008 and 35 SD/100 years (mean 0.42, with 95% of the observations between 0.03 and 2.06 SD/100 years). Peaks of ecological change occurred mostly during HSs and between 21 and 9 ka (Fig. 4). Our data were consistent with recent analyses that show that Greenland stadials associated with Heinrich events (also known as HSs) show multiple climatic stages (Sánchez-Goñi and Harrison, 2010). In the Peten-Itza record, these stages were clearly marked by shifts in the relative abundance of Quercus, Myrica, Pinus, Celtis, Acacia, Asteraceae, Ambrosia, and Poaceae (for details see Correa-Metrio et al., in press). Minimum squared-chord distance provided a measure of similarity between modern and past pollen assemblages. Very high values identified periods with assemblages that appeared to be without modern analog between w85.5 and 81, w17 and 10, and 4.2 and 0.4 ka (Fig. 4). Other isolated peaks of minimum squaredchord distance occurred at 60, 49, 39, 31, and 24 ka, coinciding with HSs 6 to 2. The lowest minimum squared-chord distances occurred between w62 and 50, and 9 and 6 ka, indicating a high likelihood that the vegetation of those times had modern analogs in the region. The generalized linear model to explain charcoal concentration as a function of other proxies was significant. The residual deviance, 1,525,626 on 437 d.f., was less than the null deviance of 1,657,087 on 441 d.f., implying that the independent variables explained a large proportion of variance in the dependent variable. Furthermore, all independent variables included in the multivariate analysis were significant with p-values lower than 0.001 (Table 1). Residuals of the fitted regression were distributed homogeneously along the predicted values, and showed a distribution close to normal (SOM Fig. 1). Whereas the coefficients for MS, DCA Axis 2 scores, and summer-winter insolation differences were positive, the relationship between charcoal and d18Oice from GISP2 was negative. As expected, independent sediment variables that served as climate and environmental proxies, i.e. MS (rainfall), insolation difference (seasonality), DCA Axis 2 (vegetation), and oxygen isotopes in the Greenland ice core (changes in AMOC), all played a role in determining fire frequency, inferred from charcoal concentration. Paleotemperature inferences revealed that the mean annual temperature w85.5 ka was at least 2.5e3.5  C cooler than today. Thereafter, there was a long-term decline of at least 1.5  C over the next 65 ka (Fig. 5). The pollen-derived temperature estimate suggests a cooling of at least 4e5  C relative to modern during the

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LGM. Superimposed upon the long-term trajectory of falling temperatures, were a number of shorter-term, abrupt warmings of w0.5e1  C. These warming events coincided with Greenland interstadials, and were followed by gradual cooling that paralleled the progression of Greenland stadials. Although these warmings were apparent and systematic through the record, they have to be interpreted cautiously as they were less than the estimated error for the transfer function, and chronological uncertainties prevent definite conclusions in this sense. Sharp cooling that exceeded the estimated error of the model were evident during HSs, when further temperature declines of least w1.5e2.5  C occurred. Though our data indicated that cold conditions prevailed in the Yucatan until w15.5 ka, rapidly changing temperatures characterized the record from this point on. Indeed, the highest rates of temperature change occurred between 15.5 and 10 ka. The PleistoceneeHolocene temperature transition occurred in two steps. A 3- C warming occurred between w15.5 and 13.5 ka, followed by a reversal of w2  C. A second began ca 12 ka, causing Table 1 Estimated coefficients for a Poisson generalized linear model relating charcoal concentration (cm2/l) to other proxies. MS: magnetic susceptibility (SI  106) (Hodell et al., 2008); DCA Axis 2: Axis 2 scores of Detrended Correspondence Analysis for fossil pollen samples from core PI-6; Insolation: Summer-winter insolation difference (W/m2) (Paillard et al., 1996); Greenland: d18Oice record from core NGRIP (&) (NGRIP Members, 2004). All independent variables were significant at a level <0.001. Variable

Estimate

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Intercept MS DCA Axis 2 Insolation Greenland

5.157 0.006 0.232 0.017 0.055

0.0708 0.0001 0.0146 0.0002 0.0015

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temperatures to rise from mean values of 20  Ce24  C at 10 ka. As modern values for the region were c. 25  C, essentially modern temperatures were established by 10 ka. 5. Discussion 5.1. Moisture availability and fire frequency The DCA ordination of the fossil pollen data from core PI-6 produced a first axis that corresponded to the balance between temperate and montane vs. tropical taxa. A clear distinction was evident between Pleistocene and the Holocene vegetation in terms of Axis 1 scores, with Holocene pollen samples lying at the “most tropical” end of Axis 1 (Fig. 4). The lowest scores of this axis characterized the LGM chronozone, when there was maximum temperature depression on a global scale (CLIMAP, 1981; Mix et al., 2001; Ballantyne et al., 2005). During the LGM peak, a mesic temperate forest occupied the area. During HSs 5 to 1, DCA Axis 1 scores showed small surges, probably a consequence of increases in Acacia and Celtis percentages that resulted from low moisture availability (Fig. 4) (Correa-Metrio et al., in press). Similarly, from 18.5 to 10.5 ka, rising scores on DCA Axis 1 reflected the growing influence of tropical dry forest taxa, mainly Acacia, Bursera, Celtis, Mimosa, and Sapium (Figs. 3 and 4). High and significant negative correlations between DCA Axis 2 and MS, and the distribution of taxa along the axis, suggested that Axis 2 reflected a gradient of decreasing moisture availability. Similar to MS, DCA Axis 2 score peaks (positive in this case) show a systematic trend and coincide with Greenland stadials (Fig. 6). These findings reinforce prior inferences of wet conditions in Central America and the circum-Caribbean during Greenland interstadials (Peterson et al., 2000; Hodell et al., 2008). Despite the

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Fig. 5. Pollen-based temperature reconstruction from core PI-6, using Synthetic Assemblages. Left panel: Distribution of each taxon along the observed modern temperature gradient (Correa-Metrio et al., 2011a); abundances were scaled to a uniform size for illustration purposes. Gray shapes represent relative abundance of pollen across the temperature gradient at 0.5  C temperature increments, although 0.25- C increments were used for the calculations. Right panel: Temperature reconstruction. Dark (light) red represents smaller (larger) distances of fossil samples from pollen synthetic (ideal) assemblages shown in the left panel. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

overall resemblance between DCA Axis 2 scores and MS, pollen data displayed a stronger response to HSs than did the MS data. Although both DCA Axis 2 and MS were apparently related to precipitation, they were proxies for different gradients. Whereas MS was strongly associated with runoff and lake level, DCA Axis 2 reflected soil moisture deficit. The occurrence of the highest peaks of DCA Axis 2 scores during HSs suggested the prevalence of extremely dry conditions, which were also inferred from changes in the composition of the pollen assemblages (Correa-Metrio et al., in press). Profound drying during HSs probably resulted from major disruptions of the AMOC (Peterson et al., 2000; Hodell et al., 2008; González and Dupont, 2009; Correa-Metrio et al., in press). The charcoal record from PI-6 revealed that fire was an inconsistent (non-stationary) force on the Yucatan landscape through the last 86 ka (Fig. 4). Modern meteorological studies indicate that fires are most likely to occur during droughts, especially when lightning follows a protracted period of low relative humidity, i.e. during the first convective storms of the wet season (Hodanish et al., 1997). Statistical analysis revealed that fire frequency, as reflected by charcoal concentration, was influenced by the synergy of a variety of factors. High scores on DCA Axis 2, which reflected dry conditions, were positively associated with high fire frequency. Similarly, increased insolation seasonality was also positively associated with charcoal concentration. With respect to the NGRIP record (NGRIP Members, 2004), high (low) temperatures in Greenland seem to have promoted low (high) fire frequency. However, contrary to our initial expectations, the relationship between fire frequency and MS, which was positively associated with rainfall, was positive. High MS (wetter periods) values may have been caused by strong convective activity causing ignition of

accumulated biomass at the end of the dry season and inducing erosive pulses into the lake. Other factors that may have accounted for this result were: 1) different plant assemblages may have had variable levels of flammability; 2) seasonality and convective regimes may have varied independently to produce non-stationary fire frequencies through time; 3) periods of extreme drought (low MS) may be times when fuel accumulation is suppressed, thereby limiting fires; and 4) difficulties to identify precisely the magnetic susceptibility values associated with each sample analyzed for charcoal. Three periods of high fire activity (w85.5e81 ka, 60e50 ka, and 10.5e5.5 ka) coincided with peaks in difference between summer and winter insolation (Fig. 6), i.e. strong seasonality. These time periods also coincided with high DCA Axis 2 scores, aligning with vegetation assemblages adapted to drier conditions. Nevertheless, the period of high insolation seasonality between w31 and 24 ka, though marked by low DCA scores, showed relatively low charcoal activity. This time interval was suggested to have been dominated by low annual precipitation, and by winter rains associated with cold fronts penetrating from the north due to intensified westerly activity (Bradbury, 1997; Hodell et al., 2008; Bush et al., 2009). Although overall rainfall and lake level may have fallen at this time, the winter moisture would have reduced dry season drought and flammability. Expansion of the Laurentide Ice Sheet was likely to have caused frequent outbursts of polar air, bringing winter rains into Central America and reducing the seasonal difference in moisture availability (Bradbury, 1997; Hodell et al., 2008; Bush et al., 2009). Between w50 and 10.5 ka, most of the severe droughts revealed by DCA Axis 2 and MS were associated with HSs (Fig. 6). Disruptions

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of the AMOC probably shortened the rainy season by reducing the size of the AWP, and causing further southward displacements of the ITCZ. Conversely, during interstadials, the number of dry months per year was probably reduced, suppressing flammability and/or reducing lightning ignition. With the exception of the period between w31 and 24 ka, orbitally-driven phases of high seasonality were probably an important pacemaker of fire frequency in the Yucatan. With the onset of the Holocene, ca 10 ka, increasing fire activity linked to rising temperature was evident once more. High temperatures and seasonality probably allowed greater buildups of fuel and intensified convective storms providing ignition at the beginning of the wet season. Causes of natural fires became largely irrelevant after about 5 ka, when human activities overwhelmed climate as a factor influencing fire (Leyden, 2002). Alternatively, between 18.5 and 14.5 ka, DCA Axis 2 scores were high and charcoal concentrations were relatively low. This time interval was marked by severe droughts, which probably suppressed flammability. The almost complete shutdown of the AMOC during HS1 (McManus et al., 2004), in conjunction with meltwater discharge into the Gulf of Mexico, probably caused this drought. 5.2. Temperature change At w85.5 ka, temperatures were 2.5e3.5  C colder than present. There was then a long-term progressive cooling of at least 1.5  C that lasted until 22 ka. The temperature change paralleled a change of the same magnitude in Caribbean sea surface temperature inferred from the Colombia Basin (Schmidt et al., 2004) (Figs. 1

and 7), reflecting the progressive cooling throughout the circumCaribbean. Superimposed on this long-term trend were fluctuations of 1e3  C, which coincided roughly with isotopic variations in the NGRIP ice core record from Greenland (NGRIP Members, 2004; Andersen et al., 2007) (Fig. 7). Ponding of tropical heat in the southern tropical oceans during Greenland stadials caused increased (decreased) temperatures in the southern (northern) Neotropics (Peterson et al., 2000; Chiang and Bitz, 2005; Barker et al., 2009). This global “see-saw” mechanism (EPICA Community Members, 2006; Barker et al., 2009) apparently accounted for temperature changes in the Central American lowlands. Greenland stadials in Yucatan were cold, with abruptly cooler episodes evident at Petén-Itzá, in antiphase with the Antarctic temperature record (Jouzel et al., 2007) (Fig. 7). During HSs, further temperature declines on the Yucatan were probably the result of synergistic changes in oceanic heat transport and declining cloud cover over the Peninsula caused by southward displacement of the ITCZ (Peterson et al., 2000; Hodell et al., 2008), which may have increased nocturnal black-body radiation. The magnitude of temperature decline in the Central American lowlands during the LGM has remained controversial (Bush and Colinvaux, 1990; Leyden et al., 1993; Bush et al., 2004). Temperature during the LGM chronozone, estimated from core PI-6, was w20  C. Nevertheless, the coldest period throughout the record began at w22 ka and lasted until w14.5 ka, with a mean temperature of w19  C (Figs. 5 and 7). The new data from Petén-Itzá thus indicated that maximum cooling of the air during the LGM in Yucatan was between 4 and 5  C, similar to the cooling estimated

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for the Panama lowlands (Bush and Colinvaux, 1990). In contrast, studies in the Central American highlands of downslope vegetation migration and glacial descent suggested a cooling of between w5 and 8.5  C (e.g. Martin, 1964; Islebe and Hooghiemstra, 1997; Lozano-Garcia and Ortega-Guerrero, 1998; Lachniet and VazquezSelem, 2005; Lozano-Garcia et al., 2005). The difference between estimates of cooling in the highlands and lowlands may be a consequence of lower CO2 partial pressure and lower cloud cover at high altitudes, leading to greater loss of heat through radiation (Bush and Silman, 2004). Sea surface temperatures in the western tropical Atlantic (Guilderson et al., 1994; Rühlemann et al., 1999) and Gulf of Mexico (Williams et al., 2010) started to increase before HS1 at ca 18.5 ka. The fossil pollen data from Petén-Itzá, however, suggested warming began at w15.5 ka, coinciding with the end of HS1 (SánchezGoñi and Harrison, 2010). This difference in timing of the onset of warming may have resulted from heat ponding in southern latitudes due to the shutdown of the AMOC (McManus et al., 2004), with the Yucatan Peninsula remaining cold as a consequence of heat loss due to low cloud cover. In fact, DCA Axis 2 and MS from core PI-6 showed that this period was dominated by extremely dry conditions, and therefore low cloud cover, which persisted from w18.5 to 14.7 ka. The Bølling-Allerød onset was clearly reflected in the pollen and MS records, and was marked by a change toward wetter conditions rather than abrupt warming. Complete resumption of AMOC probably caused expansion of the AWP, increasing summer precipitation and reducing seasonality over the peninsula. As in other records from the Circum-Caribbean, (e.g. Guilderson et al., 1994; Lea et al., 2003), temperature at Petén-Itzá showed a decrease that coincided with the Younger Dryas. The 2  C cooling was accompanied by extremely dry conditions, reflected by MS, DCA Axis 2, and pollen assemblages, which suggested the presence of xeric vegetation (Figs. 3 and 4). Temperatures rose progressively to 24  C around 10 ka, when tropical forest elements first became

dominant in the pollen spectra. Establishment of tropical forest in the area was the result of both warmer and wetter conditions, as opposed to rising temperatures alone. 5.3. Modern analogs, and rates of ecological and temperature change Despite the significant cooling that occurred during HSs, rates of temperature change during the glacial were low, between 0.5 and 0.5  C/100 years. However, temperature oscillations during the Lateglacial and the deglaciation represented the most rapid rates of change yet reported from Neotropical Central America, and the temperature shifts induced unprecedented ecological responses. Temperature changes of about 2 and 2  C/100 years characterized the inception and termination of the LGM, respectively. The temperature decline occurred between w24 and 23.5 ka, and was probably associated with HS2, whereas the temperature increase corresponded to the beginning of deglaciation. Excluding two possible outliers, rates of temperature change during the deglacial ranged between 2.5 and 2  C/100 years. According to these findings, only the most conservative IPCC scenario for AD-2100 Central America, which suggests a warming of a 2  C relative to present (Christensen et al., 2007), was experienced in the Yucatan Peninsula during the last 85,500 years. During glacial times, pollen assemblages lacking modern analogs (sensu Williams and Jackson, 2007) were episodic. During HSs 6 to 3, there were characteristic declines in representation of Quercus and Pinus, while Acacia and Poaceae increased, creating nomodern-analog savanna landscapes (Figs. 3 and 4) (Correa-Metrio et al., in press). Indeed, no-modern-analog plant associations formed primarily during HSs, calling for caution when interpreting temperature estimations. The contrast between high rates of ecological change, but low rates of climatic change during HSs 6 to 3, suggested that no-analog assemblages were the result of noanalog climates. Slowly changing temperatures imply that

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community reassortments were caused by the novel climates, as opposed to vegetation assemblages resulting from climatevegetation disequilibrium. Increased fire frequency during HSs probably contributed to rapid ecological turnover during these periods of extremely cold and dry conditions. In contrast, high rates of ecological turnover during the deglaciation were probably driven by high rates of temperature change, and therefore, vegetationclimate disequilibrium. At that time, no-modern-analog vegetation may have been caused by the inability of populations to migrate quickly enough to keep pace with climate change, as has been described for the Younger Dryas in temperate North America (Huntley and Webb, 1989; Shuman et al., 2002). 5.4. Revisiting the vegetation and climates of the LGM and the deglacial periods in the Yucatan Peninsula Earlier interpretation of pollen in core 80-1 from Lake Quexil, northern Guatemala, suggested that climate on the Yucatan Peninsula during the LGM was extremely dry and cold (Leyden et al., 1993, 1994; Huang et al., 2001). Progressively warmer and moister conditions, leading to forest expansion, were inferred for the deglacial period (Leyden et al., 1993, 1994; Leyden, 1995). Nevertheless, pollen and MS data from core PI-6 have brought these findings into question (Hodell et al., 2008; Bush et al., 2009). Given the proximity of Lake Quexil to Lake Petén-Itzá (a distance of <10 km), it is parsimonious to assume near contemporaneity in major vegetation changes between these records. Indeed, the pollen records from Lakes Quexil and Petén-Itzá are broadly similar, but if the two core chronologies are accepted, there are large discrepancies in timing of corresponding vegetation changes between the two sites (Fig. 8). In Lake Quexil core 80-1, the transition from cold and dry conditions of the deglacial to warm and wet climate of the Holocene occurred at w12 ka (Leyden et al., 1993). In the Lake Petén-Itzá

PI-6 record, the transition took place at w10.5 ka, a temporal discrepancy of w1500 years. The LGM (21  2 ka B.P., after Mix et al., 2001) pollen assemblages in Lake Petén-Itzá were similar to those identified by Leyden et al. (1993) for the end of the pollen zone they called the Interstadial, Quexil pollen zone IS in Fig. 8. These time discrepancies probably resulted from two factors: 1) sedimentation rate in core 80-1 might not have been continuous, as assumed by Leyden et al. (1993, 1994), and more importantly, 2) the oldest date in the Quexil chronology was derived from an aquatic mollusc shell, and was too old because of hard-water-lake error (Bush et al., 2009). The robust Petén-Itzá PI-6 chronology indicates that the Quexil record requires adjustment. Four major events evident in the two pollen records were used to correlate the profiles (Fig. 7) and thereby adjust the Quexil chronology. The new Quexil chronology requires re-evaluation of the previously accepted timing for regional climatic changes during the Late Pleistocene (e.g. Leyden, 1984, 1995; Leyden et al., 1993, 1994; Huang et al., 2001). 6. Conclusions Data from core PI-6, collected in Lake Petén-Itzá indicated that both climate “pulses” and “presses” drove ecological changes on the Yucatan Peninsula over the last 86,000 years. Long-term “presses,” probably associated with precession, drove the steady decline in temperature from ca 86 to 20 ka, and were probably responsible for the fire frequency pattern seen in this record. “Pulses,” in contrast, came from HSs and the abrupt temperature changes associated with the deglaciation. These short, sharp pulses caused rapid ecological changes and formation of no-modern-analog plant associations. About 50 ka, the climate of the Yucatan lowlands apparently flipped from being largely dominated by insolation dynamics to being associated with the dynamics of the North Atlantic Ocean.

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Responses in Central America to DO cycles and HS (Peterson et al., 2000; González et al., 2008; Hodell et al., 2008; González and Dupont, 2009) were identified in the ecological dynamics of the Yucatan Peninsula, and especially in pollen-based climate estimates from core PI-6. Positive temperature anomalies (0.5e1  C) occurred during Greenland interstadials, whereas stadials were marked by cold, dry conditions and high fire frequency. HSs were coincident with vegetation turnover (Correa-Metrio et al., in press), and cooling of between w1.5 and 2.5  C. These major coolings were antiphased with the most important warmings inferred from the Antarctic Dome C record (Jouzel et al., 2007). Core PI-6 temperature estimates reinforce the tight relationship between the global see-saw mechanism (EPICA Community Members, 2006) and climate in Central America. From w86 ka to present, fire frequency on the Yucatan Peninsula was highly variable, though general relationships with other factors were identified. Changes in summer-winter insolation differences seemed to have been the main “press” for fire activity. Correlation between charcoal concentration and higher insolation seasonality was probably blurred by the relatively stronger influence of North Atlantic Ocean dynamics between 50 and 10 ka. Greenland stadials and local factors associated with vegetation structure and precipitation provided “pulses” of fire. These peaks of fire activity occurring during Greenland stadials, probably reflected seasonality changes and promoted the establishment of fireassociated vegetation. Fire frequency, inferred from charcoal concentration, was probably the result of interactions among local and hemispheric factors, not all of them identified in this study. Temperature estimates from core PI-6 showed a w5  C cooling during the LGM, an inference consistent with previous qualitative estimates (Bush et al., 2009) and estimates from the Panamanian lowlands (Bush and Colinvaux, 1990). These findings prompted a reinterpretation of results from Lake Quexil (Leyden, 1984, 1995; Leyden et al., 1993, 1994; Huang et al., 2001), which found to be were consistent with the PI-6 data, once chronological issues were reconciled. Climate projections for AD-2100 in Central America suggest a 2e5  C warming relative to present (Christensen et al., 2007). Of the inferred temperature changes on the Yucatan Peninsula for the last 86,000 years, only those of the deglaciation, a warming of w3.5  C in 1500 years between 15.5 and 14 ka, come close to those projected for the coming century. Even so, this rate of increase is still almost an order of magnitude less than the projected temperature change (Fig. 4). The Petén data indicate that past nomodern-analog vegetation are not necessarily tied to the magnitude of climate change, but rather to the rate at which climate change occurs. This finding is consistent with what is known of nomodern-analog vegetation in other regions (Huntley and Webb, 1989; Shuman et al., 2002). Thus, it is improbable that species migrations will keep pace with the rapid climate changes predicted for the region in the near future. Formation of no-modern-analog communities is almost certain, and local extinctions are probable. On the Yucatan Peninsula, where migration distances between lowlands and adjacent highlands are relatively short, uplands may prove to be critical bioclimatic spaces for populations of many lowland taxa in the future. In the past, no-modern-analog vegetation assemblages caused by HSs were replaced by plant associations similar to those that existed before the HS events occurred (Fig. 2). An exception occurred after HS1, because the event was part of a longer-term climate trend, i.e. the transition toward the Holocene. With this retrospective view in mind, we predict that the rapid warming pulse of the 21st century will produce no-modernanalog communities that will last for the duration of the event. Given the improbability of returning to lower-than-modern CO2 levels, existence of anthropogenic barriers to migration, and the

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increased probability of fire associated with human activity, it is very unlikely that plant communities will return to their preindustrial state. Acknowledgments We are grateful to H. Hooghiemstra, an anonymous reviewer, and J.S. Carrión, handling editor, whose comments strengthened the manuscript. We thank all individuals who participated in the field and laboratory work during the Lake Petén-Itzá Scientific Drilling Project. We are also grateful to the numerous agencies and individuals in Guatemala who provided assistance to the project. We also thank our many collaborators from University of Florida, University of Minnesota (Minneapolis/Duluth), Geoforschungszentrum (Potsdam), Swiss Federal Institute of Technology (Zurich), Université de Genève, as well as the personnel of DOSECC. The cores are archived at LacCore (National Lacustrine Core Repository), Department of Geology and Geophysics, University of Minnesota-Twin Cities and we thank Kristina Brady, Amy Myrbo and Anders Noren for their assistance in core description and curation. Anders Noren kindly sampled the cores for this study. This project was funded by grants from the US National Science Foundation (ATM-0502030), the International Continental Scientific Drilling Program, the Swiss Federal Institute of Technology, and the Swiss National Science Foundation. This is publication 66 of the Florida Institute of Technology Institute for Research on Global Climate Change. Appendix. Supplementary data Supplementary data related to this article can be found online at doi:10.1016/j.quascirev.2012.01.025. References Andersen, K.K., Bigler, M., Clausen, H.B., Dahl-Jensen, D., Johnsen, S.J., Rasmussen, S.O., Seierstad, I., Steffensen, J.P., Svensson, A., Vinther, B.M., Davies, S.M., Muscheler, R., Parrenin, F., Röthlisberger, R., 2007. A 60000 year Greenland stratigraphic ice core chronology. Climate of the Past Discussions 3, 1235e1260. Baker, P.A., Rigsby, C.A., Seltzer, G.O., Fritz, S.C., Lowenstein, T.K., Bacher, N.P., Veliz, C., 2001. Tropical climate changes at millennial and orbital timescales on the Bolivian Altiplano. Nature 409, 698e701. Ballantyne, A.P., Lavine, M., Crowley, T.J., Liu, J., Baker, P.B., 2005. Meta-analysis of tropical surface temperatures during the Last Glacial Maximum. Geophysical Research Letters 32, L05712. Barker, S., Diz, P., Vautravers, M.J., Pike, J., Knorr, G., Hall, I.R., Broecker, W.S., 2009. Interhemispheric Atlantic seesaw response during the last deglaciation. Nature 457, 1097e1102. Birks, H.J.B., Birks, H.H., 1980. Quaternary Palaeoecology. University Park Press, Baltimore. Bradbury, J.P., 1997. Sources of glacial moisture in Mesoamerica. Quaternary International 43/44, 97e110. Bush, M.B., Colinvaux, P.A., 1990. A pollen record of a complete glacial cycle from lowland Panama. Journal of Vegetation Science 1, 105e119. Bush, M.B., Colinvaux, P.A., Wiemann, M.C., Piperno, D.R., Liu, K.-B., 1990. Late Pleistocene temperature depression and vegetation change in Ecuadorian Amazonia. Quaternary Research 34, 330e345. Bush, M.B., Correa-Metrio, A., Hodell, D.A., Brenner, M., Anselmetti, F.S., Aristegui, D., Muller, A.D., Curtis, J.H., Burton, C., Gilli, A., 2009. Re-evaluation of climate change in Lowland Central America during the Last Glacial Maximum using new sediment cores from Lake Petén Itzá, Guatemala. In: Vimeux, F., Sylvestre, F., Khodri, M. (Eds.), Past Climate Variability from the Last Glacial Maximum to the Holocene in South America and Surrounding Regions. Springer, pp. 113e128. Bush, M.B., Miller, M.C., de Oliveira, P.E., Colinvaux, P.A., 2002. Orbital forcing signal in sediments of two Amazonian lakes. Journal of Paleolimnology 27, 341e352. Bush, M.B., Silman, M.R., 2004. Observations on Late Pleistocene cooling and precipitation in the lowlands Neotropics. Journal of Quaternary Science 19, 677e684. Bush, M.B., Silman, M.R., Urrego, D.H., 2004. 48,000 years of climate and forest change from a biodiversity hotspot. Science 303, 827e829. Chiang, J.C.H., Bitz, C.M., 2005. Influence of high latitude ice cover on the marine Intertropical Convergence Zone. Climate Dynamics 25, 477e496.

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