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Quaternary Science Reviews xxx (2012) 1e13
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Quaternary Science Reviews
journal homepage: www.elsevier.com/locate/quascirev
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 xxx
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).
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
0277-3791/$ e see front matter Ó 2012 Elsevier Ltd. All rights reserved.
doi:10.1016/j.quascirev.2012.01.025
Please cite this article in press as: Correa-Metrio, A., et al., Rapid climate change and no-analog vegetation in lowland Central America during the
last 86,000 years, Quaternary Science Reviews (2012), doi:10.1016/j.quascirev.2012.01.025
2
A. Correa-Metrio et al. / Quaternary Science Reviews xxx (2012) 1e13
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
Please cite this article in press as: Correa-Metrio, A., et al., Rapid climate change and no-analog vegetation in lowland Central America during the
last 86,000 years, Quaternary Science Reviews (2012), doi:10.1016/j.quascirev.2012.01.025
A. Correa-Metrio et al. / Quaternary Science Reviews xxx (2012) 1e13
3
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
0
10
Depth (m)
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.
20
30
40
0
10,000
20,000
30,000
40,000
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).
Please cite this article in press as: Correa-Metrio, A., et al., Rapid climate change and no-analog vegetation in lowland Central America during the
last 86,000 years, Quaternary Science Reviews (2012), doi:10.1016/j.quascirev.2012.01.025
4
A. Correa-Metrio et al. / Quaternary Science Reviews xxx (2012) 1e13
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
Please cite this article in press as: Correa-Metrio, A., et al., Rapid climate change and no-analog vegetation in lowland Central America during the
last 86,000 years, Quaternary Science Reviews (2012), doi:10.1016/j.quascirev.2012.01.025
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Fig. 3. Pollen percentage diagram of selected taxa from core PI-6, Lake Petén-Itzá, lowland Guatemala.
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A. Correa-Metrio et al. / Quaternary Science Reviews xxx (2012) 1e13
5
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
Age (ka)
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0.0 0.5 1.0 1.5 -0.5 0.0 0.5 1.0 0
20 40 60 80
5
10
15
0.5 1.0 1.5 2.0
sq M
u in
di are im
sta d- um
nc cho
e rd
E
(S c co
D ha log
/1 n i
00 ge cla
ye
ar
s)
co Ch
nc ar
en co
tra al
tio
n
D
CA
A
xi
s2
A
ge
(k
a)
D
CA
A
xi
su Ma
sc g
ep ne
tib tic
ili
ty
A. Correa-Metrio et al. / Quaternary Science Reviews xxx (2012) 1e13
s1
6
0.2
0.6
0.4
0
10
HS1
20
HS2
30
HS3
40
HS4
50
HS5
60
HS6
70
80
a
b
0.0 0.5 1.0 1.5 -0.5 0.0 0.5 1.0 0
c
d
20 40 60 80
5
e
10
15
0.5 1.0 1.5 2.0
f
0.2
0.4
0.6
Fig. 4. Pollen, magnetic susceptibility, and charcoal data from core PI-6. Gray bands in panels a, b and c show Heinrich Stadials (HS, Sánchez-Goñi and Harrison, 2010). a) and b) DCA
Axes 1 and 2 scores from pollen analysis, respectively. c) Magnetic susceptibility (from Hodell et al., 2008). d) Charcoal concentration. e) Rates of ecological change based on the first
four DCA axes from pollen analysis. f) Minimum squared-chord distance between fossil pollen samples from core PI-6 and modern pollen samples from Correa-Metrio et al. (2011b);
dashed line shows threshold for no-modern analogs (0.31). Previously published data are reproduced here using the updated core chronology.
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
Std. error
z value
Pr(>jzj)
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
72.85
39.95
15.82
77.19
36.80
<2e-16
<2e-16
<2e-16
<2e-16
<2e-16
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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A. Correa-Metrio et al. / Quaternary Science Reviews xxx (2012) 1e13
7
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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A. Correa-Metrio et al. / Quaternary Science Reviews xxx (2012) 1e13
10
20
1.0
HS1
30
HS2
HS3
40
50
HS4
60
HS5
70
80
HS6
0.5
Wetter
conditions
a
-34
GI1
GI2
GI12
GI3 GI5 GI7 GI8 GI10
GI14
GI16
GI15 GI17
GI4
GI9 GI11
GI13
GI6
GI18
GI21
GI19 GI20
-36
-38
b
-40
-44
GS15
2.0
c
1.5
-42
GS3
Higher fire
frequency
GS2
GS14 GS16 GS17 GS18
GS4 GS6 GS8 GS10 GS12
GS5 GS7 GS9 GS11
GS13
GS19 GS20 GS21
1.0
0.5
2
Charcoal (ln(mm 2/mm3))
Warmer
18
-0.5
δ OIce(‰)
0.0
Δ Insolation (W/m )
DCA Axis 2 scores (SD)
Age (ka)
0
0.0
140
120
100
80
60
d
More seasonality
0
10
20
30
40
Age (ka)
50
60
70
80
Fig. 6. Fire frequency and DCA Axis 2 from core PI-6. Numbers preceded by GI and GS show Greenland interstadials and stadials, respectively. a) DCA Axis 2 scores from pollen
analysis of Petén-Itzá. b) Oxygen isotopes (d18O) from Greenland core NGRIP (from 60 ka to present from Andersen et al., 2007; before 60 k from NGRIP Members, 2004). c) Charcoal
concentration from core PI-6 (trend softened using a 3-point moving average). d) Insolation difference between summer and winter; Insolation data from Analyseries 2.0 (Paillard
et al., 1996). Gray bands in panels a, b and c show Heinrich Stadials (HS, Sánchez-Goñi and Harrison, 2010); gray bands in panel d show periods of high insolation seasonality.
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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A. Correa-Metrio et al. / Quaternary Science Reviews xxx (2012) 1e13
9
Age (ka)
10
20
HS1
0
30
HS2 HS3
40
50
HS4
HS5
60
70
80
HS6
A
-10
B
23
C
28
SST (oC)
-5
26
21
-34
19
D
0
10
20
30
40
Age (ka)
50
60
70
-42
18
-38
δ OIce(‰)
Temperature (oC)
Temperature (oC)
0
80
Fig. 7. Temperature estimates from core PI-6. Gray bands show Heinrich Stadials (HS, Sánchez-Goñi and Harrison, 2010). A) Temperature anomaly from Dome C, Antarctica (Jouzel
et al., 2007). B) Sea surface temperature, core ODP 999A, Colombia Basin (Schmidt et al., 2004). C) Mean annual air temperature reconstruction from core PI-6. D) Oxygen isotopes
(d18O) from Greenland core NGRIP (from 60 ka to present from Andersen et al., 2007; before 60 k from NGRIP Members, 2004). Insolation data from Analyseries 2.0 (Paillard
et al., 1996).
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
Please cite this article in press as: Correa-Metrio, A., et al., Rapid climate change and no-analog vegetation in lowland Central America during the
last 86,000 years, Quaternary Science Reviews (2012), doi:10.1016/j.quascirev.2012.01.025
150
50 100
15
45
20
60
5 15 20 40
300
200 400
15 45
10 20
e
el
as
to
m
at
M
ac
or
ea
ac
e
ea
e
M
an
th
ac
ar
A
m
bs
ea
100
H
er
Q
ue
rc
us
Pi
nu
s
ea
e
el
as
to
m
at
M
ac
or
ea
ac
e
ea
e
po Q
lle ue
n xi
zo l
ne
s
M
an
th
ac
ar
A
m
H
er
Pi
nu
s
50
bs
A. Correa-Metrio et al. / Quaternary Science Reviews xxx (2012) 1e13
Q
ue
rc
us
10
48
100 300
10
8
I
Lg1
Lg2
16
Quexil age (ka)
G2
24
30
Petén-Itzá age (ka)
G1
20
32
IS
40
40
50
150
50 100
15
45
20
60
5 15 20 40
100
300
200 400
30 60
10 20
48
300 600
Fig. 8. Selected taxa of pollen percentage diagrams from cores Lake Quexil 80-1 (left Leyden et al., 1993; Leyden et al., 1994) and Lake Petén-Itzá PI-6 (right; this study).
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
11
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.
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