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For example, a detailed study of the Greenland ice-core records

(Taylor et al., 1997) suggests that the transition from the end of

the Younger Dryas to the Holocene warm period may have

been completed in a series of warming steps lasting no more

than a few decades, most probably in less than 50 years (Alley

and Clark, 1999). About half of the warming was concentrated

into a single short step that lasted no more than 15 years. In

contrast to the abrupt warming at the start of the “Bølling,”

the cooling from the warm “Bølling” to the Younger Dryas

appears to have been a more gradual process, involving a series

of steps and minor climatic oscillations extending over a period

in excess of 1,500 years.

The extent to which abrupt climate shifts during the LGIT

were globally synchronous, or not, is pivotal to our understanding

of global climate mechanisms. Detailed comparisons between

the Antarctic and Greenland ice cores show that Antarctic climate

changes were significantly out of step with those in the North

Atlantic (Figure Q10) and provide another example of the bipolar

climate seesaw described above. The data suggest that cooling

(the Antarctic Cold Reversal) occurred in Antarctica during the

period of warming in the North Atlantic region referred to as

the “Bølling-Allerød Interstadial.” They also suggest that the

severely cold Younger Dryas Stadial was a period of rapid

warming in Antarctica. In addition, Ridgwell et al. (2003) have

shown that the majority of the changes in atmospheric carbon

dioxide occurred in three major pulses that are synchronous with

rapid warming in Antarctica (Figure Q10).

Possible causes of abrupt climate oscillations during

the LGIT

It is believed that deglaciations were dominated by orbitally-

induced ice sheet and associated climatic changes. The colder

episodes, especially the very pronounced Younger Dryas cool-

ing (Dansgaard et al., 1993), are periods imprinted on this trend

of warmer interglacial conditions. For example, marine records

from all parts of the North Atlantic during the Younger Dryas

show significant increases in freshwater and ice-rafted debris

input. This has led to the suggestion that the Younger Dryas

was caused by a sudden influx of cold fresh meltwater from

the North American and European Ice Sheets, which was suffi-

cient to cap deepwater formation in the North Atlantic. Another

suggestion is that meltwater diverted through the Gulf of

St. Lawrence was enough to stop deepwater formation (Berger

and Jansen, 1995; Clark et al., 2001; Carlson et al., 2007;

Carlson, 2008). It seems almost axiomatic that the driving

Figure Q9 Possible “deep-water oscillatory system” explaining glacial and interglacial Dansgaard-Oeschger cycles (adapted from Maslin et al.,

2001). Meltwater in either the North Atlantic or Southern ocean can start the system. The meltwater caps the sites of deep water formation in that

particular hemisphere and reduces the deepwater formation. This increases heat piracy to the other hemisphere, warming that hemisphere, which

instigates ice sheet melting after hundreds or thousands of years, thus reversing the whole process. The system is quasi-periodic because of its

oscillatory nature. The stochastic element observed in the records (Alley et al., 2001) comes from the different internal dynamics of each ice sheet.

The link between the Dangaard-Oeschger 1,500 year cycles and the Heinrich events is through their successive effect on sea level. At a certain

point, sea level has risen enough to undercut the Laurentide Ice Sheet and precipitate a full Heinrich event.

852 QUATERNARY CLIMATE TRANSITIONS AND CYCLES

mechanism for the Younger Dryas episode, for example, must be

some form of internal mechanism, such as oceanic circulation

change, since this episode of very pronounced cooling occurred

at a time of summer radiation maximum in Milankovitch calcu-

lations. There is also a new theory that suggest an asteroid

impact may have caused the Younger Dryas event (Firestone

et al., 2007) and this may explain why a Younger Dryas event

is not seen at other earlier Terminations (Carlson, 2008).

On top of this are climate records of Greenland and Antarc-

tic ice cores showing an anti-phase between the North and

Southern Hemispheres through the LGIT. Prominent among the

emerging ideas is Broecker’s(1998) suggestion of a “bipolar

seesaw behavior in thermohaline circulation” as a mechanism

that explains this anti-phase relationship, as discussed above.

Recent modeling work has shown that the anti-phase hemi-

spheric climate during the LGIT can be produced from sim-

ple alternation of the deep ocean circulation driven by the

Younger Dryas meltwater event in the North Atlantic (Knorr

and Lohmann, 2003; Stocker, 2003; Weaver et al., 2003;

Carlson, 2008).

Conclusions

Based on our present understanding of Quaternary climate change,

large climatic transitions have occurred on the timescale of indivi-

dual human lifetimes, for example, the end of the Younger Dryas

and various D-O cold events. Many other substantial shifts in

Figure Q10 (a) GRIP d

O “temperature proxy” showing timing of the Bølling / Allerød (B/A), and Younger Dryas (YD) intervals, the Antarctic Dome

C dD “temperature proxy” record highlighting the relative timing of the Antarctic Cold Reversal (ACR), and observed atmospheric CO

changes as

gray-filled circles, and the “residual” CO

record as empty circles, formed by subtracting the effects of re-growth of the terrestrial biosphere, sea

level and reef buildup (see Ridgwell et al., 2003). Both data sets are now plotted as change relative to LGM values. (b) Rate of change of observed

(thick gray line) and “residual” (dashed red line)CO

since the LGM. The timing of the sharp spikes corresponds to discernable “jumps” in the

observed CO

record and correlates with significant warming in Antarctica (adapted from Ridgwell et al., 2003).

QUATERNARY CLIMATE TRANSITIONS AND CYCLES 853

climate took at most a few decades or centuries e.g., Heinrich

events and events during the LGIT. Despite a huge amount of data,

many problems remain in understanding the mechanisms control-

lingthese rapid climate changes. Both external and internal forcing

play a part. Of fundamental importance, however, is the discovery

of the central role of the deepwater system in controlling past rapid

global climatic changes (Stocker, 2003) and this should be of

central concern in possible future climate change (Maslin, 2004a).

Mark Maslin

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Cross-references

Albedo Feedbacks

Antarctic Bottom Water and Climate Change

Astronomical Theory of Climate Change

Binge-Purge Cycles of Ice Sheet Dynamics

Carbon Dioxide and Methane, Quaternary Variations

Climate Change, Causes

Climate Forcing

Dansgaard-Oeschger Cycles

Eccentricity

Glaciations, Quaternary

Heinrich Events

Holocene Climates

Ice cores, Antarctica and Greenland

Last Glacial Maximum

Last Glacial Termination

Millennial Climate Variability

Mountain Uplift and Climate Change

North Atlantic Deep Water and Climate Change

Obliquity

Plate Tectonics and Climate Change

Precession, Climatic

SPECMAP

Sun-Climate Connections

Thermohaline Circulation

Weathering and Climate

Younger Dryas

QUATERNARY CLIMATE TRANSITIONS AND CYCLES 855

QUATERNARY VEGETATION DISTRIBUTION

Introduction

All aspects of Quaternary vegetation distributions, from the

range and abundance of individual species, to the global posi-

tion, extent, and composition of biomes, have changed exten-

sively and repeatedly in response to the orbitally-paced

oscillations in climate, atmospheric CO

, and ice extent that

are characteristic of the Quaternary. At centennial and longer

timescales, the distribution and abundance of plant species clo-

sely track shifts in climate, despite the rooted habit of indivi-

dual plants, via the dispersal of propagules to newly

favorable regions, local changes in population abundances,

and the dying out of populations in unfavorable locations. Dis-

tributions of vegetation formations and biomes emerge from

these species-level responses. The history of Quaternary vege-

tation distributions is thus indirectly a history of global and

regional climate change, and paleo-vegetational records are an

important source of information about Quaternary environ-

ments. As an active component of the Earth system, shifts in

the global vegetation affect the biogeochemical and biogeophy-

sical fluxes between the land surface and atmosphere as well as

the functioning of the global carbon cycle. Moreover, the pre-

sent genetic structure and distribution of populations, species,

and communities is in part a legacy of late-Quaternary species

migrations and colonization history.

Primary sources of information for Quaternary vegetation his-

tory are the fossil pollen records collected from lake, wetland,

and near-shore marine sediments, supplemented extensively by

plant macrofossil, phytolith, and stable carbon isotope records.

The present genetic structure of plant populations provides

additional information about colonization routes and refugia.

Detailed information about Quaternary vegetation

distributions is mainly limited to the past 25,000 years, due

to the limited length of most terrestrial records and because

radiocarbon dating, the primary source of age estimates

for late-Quaternary organic sediments, becomes increasingly

uncertain after 40,000 years as radioactivity declines to back-

ground levels. The accuracy of radiocarbon ages varies over

time, due to secular variations in atmospheric radiocarbon

concentrations; all radiocarbon ages here have been conver-

ted to calendar years unless otherwise indicated. The past

40,000 years include the Last Glacial Maximum, last glacial-

interglacial transition, and current interglacial and so may serve

as a model for vegetation distributions during previous glacial-

interglacial sequences. Individual records can span one million

years or more, allowing longer-term reconstructions of vegeta-

tion history for certain locales.

Vegetation processes and dynami cs

At Quaternary time scales (10

–10

years), vegetation dyn-

amics are mainly forced by climate change, and plant taxa

have primarily accommodated late-Quaternary climate change

via local changes in abundance and regional to continental

shifts in distribution (Huntley and Webb, 1989). These biogeo-

graphic shifts can be dramatic: for example, Picea (spruce),

which today is abundant in boreal and alpine regions, at the

Last Glacial Maximum was apparently most abundant in the

eastern interior of the USA (Figure Q11), a geographic shift

of approximately 2,000 km. The maps of spruce also illustrate

that although migration is an important component of Quatern-

ary vegetation dynamics, within-range changes in abundance

are equally significant. The range of spruce (approximated

by the 5% pollen contour in Figure Q11) has been relatively

stable during the middle to late Holocene, but the area of high-

est abundance has shifted from western to eastern Canada

(Figure Q11). These internal shifts in abundances are the result

of changing rates of reproduction, growth, and mortality within

and among populations, mediated by climate and disturbance

regimes. The changes in extent and abundance differ among

plant taxa, and the associations among taxa have changed with

time. Species such as Fagus grandifolia (American beech) and

Tsuga canadensis (eastern hemlock) that today are closely

associated, had differing late-glacial distributions and migration

histories (Figure Q11). Conversely, plant associations existed in

the past that no longer exist today (indicatedby high dissimilarities

in Figure Q11), and may have grown under climatic conditions

unlike any found today (Williams et al., 2001). Species move-

ments along altitudinal gradients in response to climate change

also display individualistic behavior, rather than wholesale shifts

in vegetation zones.

The available evidence indicates that vegetation is in

dynamic equilibrium with climate at millennial time scales

(Webb, 1986). Climatic reconstructions based on pollen-

response surfaces are generally consistent with other climatic

proxies and general circulation model simulations and can

successfully predict pollen abundances for other plant taxa

(Prentice et al., 1991). Multi-proxy sediment records with evi-

dence of abrupt climate change typically show that terres-

trial vegetation composition responded to such changes with

response times on the order of decades to a century (Tinner

and Lotter, 2001). Holocene migration rates based upon pollen

data range from 10 to 1,000 m yr

–1

. However, these migration

rates are at the edge of or beyond those predicted by current

dispersal theory (Clark et al., 2001), and may be over-estimates

in cases where outlier populations invisible to the pollen record

are nucleating range expansions. Disturbance agents (fire,

flood, windthrow, extreme climate events) contribute to and

accelerate vegetation change by increasing plant mortality and

opening up gaps for colonization.

At subcontinental to global scales, the individualistic

responses of plant taxa to climate change cause the position,

extent, and species composition of plant functional types (PFTs)

and biomes to change over time. PFTs and biomes are useful

units of vegetation at global scales, because they are based upon

structural features of the vegetation and so are not constrained

by intercontinental differences in evolutionary history and floris-

tic lists (Prentice et al., 2000). Moreover, the number of plant

morphological and functional types is limited relative to the

number of species, so biomes can persist despite individualistic

species behavior.

Vegetation history

Plio-Pleistocene transition to the last glacial maximum

(ca. 2.5 million to 21,000 y

BP)

The major floristic elements of the Quaternary vegetation were

in place by the end of the Tertiary, with few new species or

morphological changes apparent in the fossil record. Vegeta-

tion distributions during the Plio-Pleistocene transition changed

in response to the onset of widespread glaciation in the Northern

Hemisphere at ca. 2.45 million y

BP, decreased global mean tem-

peratures, and increased climate variability. Vegetation change

856 QUATERNARY VEGETATION DISTRIBUTION

Figure Q11 Rows 1–3: Mapped pollen percentages for spruce (Picea), hemlock (Tsuga), and beech (Fagus) in northern and eastern North America

for selected 1,000-year time periods since the Last Glacial Maximum. Each row header indicates the age in y

BP. Maps are based upon fossil pollen

records from the North American Pollen Database, and consist of percent pollen abundances calculated relative to a summed pollen count and

interpolated to a 50-km grid. Such maps are widely interpreted to represent past shifts of plant distributions, and analyses of modern

pollen-vegetation relationships have shown that such maps capture contemporary patterns of plant abundances. Row 4: Mapped dissimilarities

between fossil pollen assemblages and a dataset of modern and 500-y

BP pollen samples. High dissimilarities indicate that the fossil pollen

assemblage has no close counterpart among modern pollen samples, and is interpreted to indicate plant communities unlike any today.

QUATERNARY VEGETATION DISTRIBUTION 857

was most pronounced at high latitudes, where during the rela-

tively warm Middle Pliocene, coniferous forests had grown to

the Arctic coast in regions today occupied by tundra; while in

the Southern Hemisphere, beech grew in coastal areas of Antarc-

tica (Thompson and Fleming, 1996). Despite the drastic climate

changes during the Plio-Pleistocene transition, there is little

evidence of significant global plant extinctions, although local

extinctions were widespread (Watts, 1988).

For the past 700,000 years, vegetation history has been

paced by the pronounced 100,000-year cycle in climate, atmo-

spheric CO

, and global ice volume. Vegetation distributions at

previous interglacials are broadly similar to the current intergla-

cial, with expansions in the abundance and range of thermophi-

lous taxa, but just as climates differed among interglacials, so

too did the distribution and composition of the vegetation.

The Eemian Interglacial in Europe (130–107 thousand y

BP),

which is the last interglacial prior to the current one and is par-

ticularly well documented, is characterized by high pollen

abundances of oak (Quercus), hazel (Corylus), alder (Alnus),

beech (Fagus), hornbeam (Carpinus), and other temperate

woody taxa, with differences among sites likely reflecting inter-

regional differences in Eemian climates. After the last intergla-

cial ended, the proportion of treeless vegetation increased, with

high pollen abundances of grass (Poaceae), sagebrush (Artemi-

sia), chenopods (Chenopodiaceae), pine (Pinus), and juniper

(Juniperus) reported from records spanning the last glacial per-

iod (Allen et al., 1999). On both sides of the North Atlantic,

the proportion of woody vegetation increased during Heinrich

events and interstadials (Grimm et al., 1993). In western North

America, high abundances of oak, willow (Salix), and

alder pollen during interglacials alternated with high abundances

of fir (Abies), pine, and Taxaceae/Cupressaceae/Taxodiaceae

(key taxa include redwood, cedar, and cypress) pollen during

glacial periods (Adam, 1988;Heusser,2000).

Last Glacial Maximum (LGM; 21,000 yBP)

By the last glacial maximum (LGM), enough well-dated paleo-

vegetation records are available to systematically map global

vegetation distributions, although tropical regions remain

underrepresented. LGM vegetation distributions (Figure Q12a)

formed in response to climates that were colder and drier

than present. Atmospheric CO

concentrations that were

80 ppm lower than pre-industrial levels may have amplified

the ecological impact of decreased moisture availability. Tree-

less conditions predominated in high-latitude regions not bur-

ied by ice, with tundra intergrading into steppe in central

Asia and China (Figure Q11a) and steppe vegetation in North

America along the margin of the Cordilleran and Laurentide

Ice Sheets, in southern Europe and Northern Africa, in southern

Africa, and in Australasia. Deserts extended to the east coast

of China in now-forested regions. In contrast, lake level records

from the western interior of North America indicate that

full-glacial climates were wetter than present, and conifer

woodlands grew widely in areas that today are desert or

shrubland-steppe (Thompson and Anderson, 2000). Northern

Hemisphere temperate and boreal forests were mainly confined

to the southeastern USA, southeastern China, and Japan, and

included a large component of cold-tolerant evergreen taxa.

The difference between LGM and modern climate and

vegetation in the tropics is less pronounced than in higher lati-

tudes, but neither tropical climates nor vegetation were stable

during the late Quaternary (Bush et al., 2001). Tropical lowland

rain forests were somewhat diminished in extent, although

the extent of the decrease and degree of fragmentation has

been controversial. In Africa, tropical rain forest was parti-

ally replaced by tropical seasonal forest, whereas in southern

Amazonia and Australasia, grasslands and savanna encroached

on areas today occupied by tropical rain forest, subtropical

forest, and savanna (Kershaw, 1988; Behling, 2002). The upper

treeline was as much as 1,000 m lower than present, as the

elevational ranges of upland types moved downslope, and

upland evergreen forests were replaced by steppe and sclero-

phyll shrubland.

Recent explanations for LGM tropical treeline depressions

have focused upon the effects of increased aridity and low

concentrations in addition to lowered temperatures (Jolly

and Haxeltine, 1997). Tropical plant communities above tree-

line are dominated by plants utilizing the C

photosynthetic

pathway, which enables maintenance of high intracellular

concentrations despite low ambient CO

concentrations.

plants, conversely, are adversely affected by increased

photorespiration and decreased water-use efficiency under low

ambient CO

concentrations and the resulting competitive

disadvantage may have depressed tropical treelines. In some

areas, however, full-glacial C

communities were replaced by

communities during deglaciation, indicating that the physio-

logical effects of varying CO

concentrations can be overridden

by local climatic controls (Huang et al., 2001). The importance

of varying CO

concentrations for Quaternary vegetation

dynamics remains under debate.

Deglaciation (14,000–8,000 yBP)

Global vegetation distributions were fundamentally trans-

formed during deglaciation as glacial climates warmed to

interglacial conditions during a period of higher-than-present

seasonal contrasts in insolation. Forest cover increased, thermo-

philous taxa expanded upslope and northwards, and vegetation

formations characteristic of the Holocene (e.g., taiga) emerged

from the reshuffling of species distributions. In northwestern

Canada and northern Eurasia, the northern treeline reached

its northernmost position during the early Holocene, advancing

near to the Arctic coastline. It then retreated to its present posi-

tion by 4,000–3,000 y

BP (MacDonald et al., 2000 ). Montane

forest developed in central Africa by 10,000 y

BP (Jolly et al.,

1997) and podocarp forests in New Zealand by 13,000 y

(McGlone, 1988). In Europe and eastern North America,

thermophilous plant taxa began migrating north between

14,000 and 13,000 y

BP. In Africa, C

plants diminished in

abundance (Street-Perrott et al., 1997).

This period is further marked by abrupt changes in atmo-

spheric and oceanic circulation that were superimposed upon

the general transition from glacial to interglacial climates, and

vegetation distributions closely tracked these rapid climate

events. During the Younger Dryas Chronozone (12,900–

11,600 y

BP), plant communities adjacent to the North Atlantic

reverted to cold-tolerant assemblages, switching to tundra in

the European Alps and to alder shrublands in Maritime Canada

(Mayle et al., 1993; Ammann et al., 2000). Vegetation changes

consistent with a reversion to lower temperatures have been

reported in Alaska, the Pacific Northwest, and at individual

sites in South America and the Indo-Pacific region (Peteet,

1995). In other regions, vegetation shifts coincident with the

Younger Dryas suggest variations in moisture balance rather

than temperature. In eastern North America, pine rapidly

858 QUATERNARY VEGETATION DISTRIBUTION

Figure Q12 Global vegetation maps produced by the Biome 6000 working group (Prentice et al., 2000; Bigelow et al., 2003), version 4.1 for the Last

Glacial Maximum (18,000 radiocarbon y

BP; ca. 21,000 calendar yBP), mid-Holocene (6,000 yBP), and present (0 yBP), based upon systematic application

of a standardized classification algorithm to pollen and macrofossil data. The present-day vegetation reconstructions closely agree with

independent maps of the vegetation. Ice sheet extent and changes in sea level are not shown. Paleovegetational reconstructions for India, the

Indo-Pacific, South America, and Australia will be included in future versions of the Biome 6,000 synthesis. Biome codes: BORF, Boreal Forest; DESE,

Desert; DRYT, Dry Tundra; STEP: Steppe / Grassland/ Dry Shrubland; SAVA, Savanna and Open Woodland; TEMP, Temperate Forest; TROP, Tropical

Forest; TUND, Tundra; WRMT, Warm Temperate Forest.

QUATERNARY VEGETATION DISTRIBUTION 859

migrated westwards into the Great Lakes region, and spruce

abundances declined (Shuman et al., 2002). Tropical African

pollen records indicate that the Younger Dryas coincided with

increased aridity. Vegetation distribution shifts during the

Younger Dryas thus appear to have varied in direction among

regions and continents as changes in climate varied among

these regions.

Middle Holocene (7,000–4,000 yBP)

Global vegetation distributions during the mid-Holocene were

similar to present, although differences persisted regionally

(Figure Q11b). The position of the northern treeline was asym-

metric relative to its present position, approximately 25 km

north in northwestern Canada, 200–300 km north in central

Siberia, and up to 200 km south in eastern Canada (Prentice

et al., 2000; Bigelow et al., 2003). Temperate forests in eastern

North America, Europe, and China were poleward of their

present position, with shifts on the order of 100 km and as

much as 800 km in northeastern China. In New Zealand, south-

ern beech (Nothofagus) forests expanded (McGlone, 1988).

These patterns suggest warmer-than-present temperatures in

middle and upper latitudes.

Other distributional differences between the mid-Holocene

and present vegetation were likely due to interregional variations

in moisture balance. These differences are greatest in northern

Africa, where steppe, xerophytic woodlands, and even warm

mixed forest grew across a broad area now occupied

by Saharan desert (Figure Q12b). Lake level highstands

and modeling studies show that the African monsoon was

enhanced relative to present (COHMAP Members, 1988). In

the Columbian Andes, xeric vegetation types increased in extent

during the middle Holocene although vegetation distributions in

Chile suggest wetter-than-present conditions (Marchant et al.,

2002). Steppe expanded at the expense of forest in the northern

Midwest but was reduced in central Asia relative to present.

Consequences of Quaternary vegetation change

Earth system effects and feedbacks

The terrestrial biosphere is an integral component of the Earth

system, and changes in vegetation distributions likely affected

atmospheric circulation and the functioning of global energy

and biogeochemical cycles. At high latitudes, changes in surface

albedo associated with shifts in treeline may have amplified the

orbitally-induced variations in temperature (Foley et al., 1994).

In Northern Africa, the mid-Holocene expansion of steppe at

the expense of desert may have enhanced monsoonal strength

by increasing rates of evapotranspiration (Kutzbach et al.,

1996). At the LGM, decreased vegetation cover in mid- and

high-latitudes may have amplified cooling, due to an albedo

feedback, whereas decreased evapotranspiration and a weakened

hydrological cycle may have enhanced tropical aridity in the tro-

pics (Levis et al., 1999). Furthermore, the LGM decrease in vege-

tation cover may have increased the source area for dust

emissions and increased atmospheric LGM dust loadings (Harri-

son et al., 2001), thereby indirectly affecting the Earth’s radiative

budget and increasing rates of nutrient transport. The effects of

vegetation change upon glacial-interglacial carbon budgets are

large but uncertain, with recent estimates ranging from 500 to

1,100 Pg carbon lost from the terrestrial biosphere at the LGM,

compared to present-day stocks of 2,000 Pg C. (Maslin, 2003).

Increases in atmospheric methane concentrations from 350 ppbv

at the LGM to 700 ppbv at the beginning of the Holocene may be

due to expansion in tropical and boreal wetlands, or increased

rates of wetland CH

emissions due to increased CO

levels

(Kaplan, 2002).

Evolutionary and genetic impacts

The close similarities between late Tertiary and Pleistocene fau-

nas and the corresponding lack of widespread extinction and

macroevolutionary change during the Quaternary is somewhat

surprising given that the signal feature of the Quaternary is its

extensive and repeated large climatic fluctuations. Apparently,

plant taxa have accommodated Quaternary climate change

primarily by ecological rather than evolutionary processes.

Bennett (1997) has argued that individualistic species behavior

effectively decouples microevolutionary and macroevolutionary

processes, as orbitally-paced reorganizations of species associa-

tions remove coevolutionary adaptations.

Nevertheless, the repeated range shifts characteristic of the

Quaternary are highly likely to have affected species’ genetic

structure, as individual populations have propagated or gone

extinct (Davis and Shaw, 2001). Whether Quaternary vegetation

dynamics tended to foster increased or decreased genetic diver-

sity is not yet known. The genetic structure observed in present

populations and species is in part a legacy of the biogeographic

responses to the last deglaciation. For example, haplotype distri-

butions within European white oaks (Quercus spp., subgenus

Lepidobalanus) and North American whitebark pines have been

convincingly related to likely colonization routes and glacial

population distributions (Petit et al., 2002; Richardson et al.,

2002). Differences in genetic structures among modern popula-

tions thus may help refine inferences about vegetation history.

Quaternary environmental variability may have helped

shape the modern latitudinal gradient in species richness and

other biogeographic patterns of species diversity. Historical

explanations for high levels of tropical species diversity include

focused upon two contrasting hypotheses: (a) long-term stabi-

lity of tropical climates encouraging persistence of tropical

rainforests and low rates of extinction versus (b) high levels

of endemism caused by repeated fragmentation of rainforests

during glacial periods (Hooghiemstra and van der Hammen,

1998). The extent of Amazonian rain forest fragmentation

remains controversial, but lowland rainforests appear to have

persisted as intact, albeit diminished, bodies of vegetation

(Colinvaux et al., 2000), weakening the endemism hypotheses.

John Jack W. Williams

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Cross-references

Carbon Dioxide and Methane, Quaternary Variations

Carbon Isotopes, Stable

Holocene Climates

Holocene Treeline Fluctuations

Last Glacial Maximum

Last Glacial Termination

Late Quaternary-Holocene Vegetation Modeling

Monsoons, Quaternary

Paleobotany

Palynology

Pollen Analysis

Quaternary Climate Transitions and Cycles

Younger Dryas

QUATERNARY VEGETATION DISTRIBUTION 861