Quaternary Biogeography and Climate Change

Quaternary Biogeography and Climate Change

Quaternary Biogeography and Climate Change JL Blois and JE Williams, University of California, Merced, CA, USA r 2016 Elsevier Inc. All rights reserve...

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Quaternary Biogeography and Climate Change JL Blois and JE Williams, University of California, Merced, CA, USA r 2016 Elsevier Inc. All rights reserved.

Glossary Albedo The reflecting power of a surface. Ancient DNA (aDNA) DNA that has survived in the remains of ancient organisms. Meridional overturning circulation A global oceanic circulation pattern wherein surface waters in the high latitudes are cooled, thereby becoming denser; this dense water sinks and flows toward the equatorial regions. In

Introduction The Quaternary, which started 2.588 million years ago (Ma) and encompasses both the Pleistocene (2.588 Ma to 11.7 thousand years ago (ka)) and the Holocene (11.7 ka to the present), is characterized by substantial climatic and biogeographic change. Climatically, this time period is known as the ‘Ice Ages’ because the earth has cycled between a series of cold glacial and warmer interglacial periods (Hays et al., 1976). A significant amount of biogeographic change also occurred, including the evolution of Homo sapiens in Africa and subsequent dispersal to the rest of the world, and the evolution of modern biotic communities (Blois and Hadly, 2009; Henn et al., 2012). Finally, this time period also encompasses the Anthropocene, a proposed new epoch within the Quaternary characterized by significant human impacts on natural processes (Zalasiewicz et al., 2011). Here, we discuss Quaternary climate change and biogeography, primarily using examples from North America over the past 21 000 years. This well-studied time period encompasses the most recent transition from a glacial to an interglacial period (Denton et al., 2010). It also saw the spread of humans from Eurasia into North and South America (Achilli et al., 2013; Hamilton and Buchanan, 2010; Henn et al., 2012), the nearglobal extinction of large megafauna (e.g., Barnosky et al., 2004; Koch and Barnosky, 2006), and significant population-level changes in vegetation and the surviving smaller animals (e.g., Blois et al., 2010; Gill et al., 2009; Williams et al., 2004). Many of these biogeographic changes happened, at least in part, due to climate change (e.g., Barnosky et al., 2004; Lorenzen et al., 2011; Williams and Jackson, 2007). Thus, the Quaternary and the last 21 000 years in particular provide baselines for expected rates and magnitudes of climate change in normal conditions, as well as for how populations, species, communities, and ecosystems have responded to those changes.

Quaternary Climate Change Orbital Variation and Quaternary Glacial–Interglacial Cycles The Quaternary occurred at the end of a long-term cooling trend that followed the warm early Eocene climatic optimum

Encyclopedia of Evolutionary Biology, Volume 3

tropical and subtropical regions around the world these waters eventually mix with other waters, becoming less dense, and they return to the surface to ultimately flow toward the higher latitudes and complete the circuit. No-analog assemblages Combinations of species observed at particular places and times in the past that are not seen anywhere on the landscape today.

roughly 50 Ma. Thus, while the glacial–interglacial cycles that characterize the Quaternary encompass both cold and warm times, even the warm periods are colder than most times in the Cenozoic (the past 65 million years) (Figure 1; Zachos et al., 2001). This long-term cooling trend facilitated the development of extensive ice sheets in both the southern and northern hemispheres. While permanent ice sheets have been in place in Antarctica for tens of millions of years (Kennett, 1977; Zachos et al., 2001), the start of the Quaternary, 2.588 Ma, roughly coincided with the development of permanent ice sheets in the northern Hemisphere (Ruddiman and Raymo, 1988). One of the most striking physical features of the Quaternary is the periodic growth and decline of the northern hemisphere ice sheets, which provide visible evidence of the accompanying climatic changes that occurred throughout the Quaternary. These glacial–interglacial cycles are driven primarily by changes in the position of the earth relative to the sun, which fluctuates due to orbital variations (Raymo et al., 2006). Together, orbital variations influence the amount and distribution of solar radiation across the earth and the seasonal contrasts of solar radiation, pacing the build up and dissipation of large ice sheets at the poles (Lisiecki and Raymo, 2005; Raymo et al., 2006; Ruddiman, 2008; Zachos et al., 2001). Orbital variations have a number of consequences for both marine and terrestrial environments and physical landscapes, including interrelated changes to ice sheets, sea levels, CO2 and other atmospheric gases, temperature, and aridity. Ice sheet buildup is spurred initially by orbitally forced decreases in summer insolation which causes summer cooling, amplified (at least in North America) by changes in vegetation from forest to tundra (Kageyama et al., 2004). Subsequent feedbacks between land and ice through changes to albedo, changes in the altitude of ice sheets, and restriction of freshwater flows into the adjacent oceans helps amplify the buildup of the ice sheets (Kageyama et al., 2004). Increases in the extent and volume of ice on land correspond to decreases in sea levels, which were dramatically lower during glacials, for example, reaching around 120 m below present-day sea level (Denton et al., 2010). Similar to the role of insolation in glacial inception, deglaciation is initiated when summer insolation increases. However, increases in summer insolation do not always cause glacial terminations; other factors are also important (Denton




Quaternary Biogeography and Climate Change



Cretaceous Paleogene Boundary



Benthic δ18O (‰)

1 Plio-Pleistocene Glaciations

4 Pleistocene Pliocene

Colder Paleocene 60

Eocene 50


Oligocene 40 30 Age (millions of years ago)


Miocene 20


0 Holocene

−380 Warmer





Colder Pleistocene 700

600 500 400 300 200 Age (thousands of years before 1950 on EDC2 timescale)



0 Holocene




Colder Pleistocene 21 (c)

δ18O (‰)




15 12 9 6 Age (thousands of years before 2000 on GICC05 timescale)



Figure 1 Three views of climate change across the Cenozoic. (a) Global cooling from the early Cenozoic to the present inferred from changes in oxygen isotopes (δ18O), modified from Blois, J.L., Hadly, E.A., 2009. Mammalian response to Cenozoic climatic change. Annual Review of Earth and Planetary Sciences 37, 181–208 and adapted from Zachos, J., Pagani, M., Sloan, L., Thomas, E., Billups, K., 2001. Trends, rhythms, and aberrations in global climate 65 Ma to present. Science 292, 686–693, with permission from American Association for the Advancement of Science. (b) Glacial–interglacial variation over the past 750 000 years, inferred from changes in deuterium (δD) throughout the EPICA Dome C ice core in Antarctica (Augustin et al., 2004). (c) Transition from the Last Glacial Maximum to the present, inferred from δ180 from the NGRIP ice core in Greenland (Andersen et al., 2004; Rasmussen et al., 2006).

et al., 2010). One factor appears to be continental ice volume: when continental ice volume nears its maximum, enough fresh meltwater flows from the continental ice sheets into the North Atlantic to decrease North Atlantic meridional overturning circulation. This initiates a cascade of global effects to both oceans and atmospheres that causes atmospheric CO2 levels to rise high enough to sustain an interglacial (Denton et al., 2010). Sea levels also rise during transitions from glacial to interglacial states as the large amount of water formerly

locked up as ice on land drains into the ocean (Lambeck et al., 2002). Changes to the ice sheets and sea levels during deglaciation affect the physical space available to terrestrial and marine species; in some places (e.g., northern North America), new physical space becomes available to terrestrial species as ice sheets melt and retreat northward, whereas in other locations physical environments are lost as sea levels rise (e.g., coastlines). The converse is true of the physical changes that occur during glacial inceptions.

Quaternary Biogeography and Climate Change Each glacial and interglacial period was unique – in magnitude of warming and cooling, in duration, and in the variation in short-term climate events overlain on the glacial– interglacial cycles (Augustin et al., 2004; Petit et al., 1999). However, several general patterns emerge. Glacial–interglacial cycles in the early Pleistocene were characterized by highfrequency cycles of relatively low magnitude, and cycles were symmetric; that is, duration of the glacial–interglacial transitions (entry into interglacial) was relatively similar to that of the interglacial–glacial transitions (termination of glacial) (Ruddiman and Raymo, 1988; Tziperman and Gildor, 2003). However, the periodicity and magnitude of glacial–interglacial cycles changed between 1 Ma and 600 ka (Figure 1): cycles became longer (less frequent, on the order of 100 000 years duration), and the amplitude of each cycle increased (Lisiecki and Raymo, 2005; Ruddiman and Raymo, 1988). Since then, the duration of the glacial portion of the cycle has been between 80 000 and 120 000 years in length (Denton et al., 2010). In addition, the rates of change at the initiation and termination of glacials became asymmetric: initiation of glacial periods was slow, characterized by an oscillating buildup, but termination of the glacial period was rapid (Denton et al., 2010; Tziperman and Gildor, 2003).

Rapid Climate Change and the Last Glacial–Interglacial Transition A majority of the biogeographic data available for the Quaternary are from the most recent glacial–interglacial transition, so we focus on this transition to illustrate the decadal to millennial-scale climatic variability that accompanies the more general glacial to interglacial transition (Figure 1). We also illuminate specific patterns and mechanisms of shorter-term variation that occurred at earlier glacial–interglacial transitions (e.g., Raymo et al., 1998). The global climate transition from the Last Glacial Maximum (LGM) ca. 21 ka to the start of the Holocene interglacial period 11.7 ka was marked by collapse and retreat of the ice sheets, substantial rises in sea levels, and relatively rapid global warming (Denton and Hughes, 2002). However, these changes did not occur smoothly and monotonically but were instead marked by starts and reversals, and periods of rapid change followed by periods of little change. For example, ice sheet retreat and corresponding changes in sea levels were discontinuous (Carlson and Winsor, 2012). In North America, the Cordilleran and Laurentide ice sheets covered northern North America, and an additional ice sheet still occurs on Greenland. Portions of these ice sheets reached their maximum extent as early as 30 ka and persisted at maximal extent until roughly 19 ka (Carlson and Winsor, 2012; Clark et al., 2009), at which point they extended south of the 49th parallel into the United States. The most rapid rates of ice sheet retreat occurred around 19 ka and during the Bølling–Allerod warm period (14.7–12.9 ka), though final collapse of the Laurentide did not occur until 8.2 ka (Carlson and Winsor, 2012). Sea level rose in spurts corresponding to changes in the ice sheets, and finally leveled off around 6 ka (Carlson and Winsor, 2012). Altogether, sea levels rose roughly 120 m across the last glacial–interglacial transition (Denton and Hughes, 2002).


Global temperatures rose 3–5 1C on average from the LGM to the present (Jansen et al., 2007), but the transition was marked by substantial spatial and temporal variability. Among the most significant climatic changes during deglaciation were two periods of rapid warming with an intervening cold interval, all of which occurred within roughly 3000 years (Figure 1; Steffensen et al., 2008). Following initial deglaciation, rapid warming occurred around 14.7 ka during the entry into the Bølling–Allerod period. For example, a Greenland ice core showed that around 9 1C degrees of warming occurred within 70 years (Severinghaus and Brook, 1999). The causes of this event are still being investigated, but warming was likely related to reduced amounts of freshwater runoff into the North Atlantic (Liu et al., 2009; Denton et al., 2010). Following this warm period, the climate system entered into the Younger Dryas (12.9–11.7 ka) when large pulses of freshwater into the North Atlantic reduced meridional overturning circulation and thus global heat transport (McManus et al., 2004). The end of the Younger Dryas was marked by rapid warming into the Holocene interglacial, on the order of 5–10 1C in Greenland in as little as a few decades (Dansgaard et al., 1989; Severinghaus and Brook, 1999; Taylor et al., 1997). Both of these rapid climate changes have been detected in diverse proxy records around the world (Clark et al., 2009) and were truly global events, though the details differ between the northern and southern hemispheres and the magnitude of climate change was variable across space and time. Changes were also seen in many other aspects of the climate system, such as temperature, aridity, and the seasonality of both (e.g., Bartlein et al., 1998).

Holocene Variability Climates also varied during the Holocene interglacial period, but the magnitude of change was generally less than during deglaciation (Figure 1; Mayewski et al., 2004; Shuman, 2012; Viau et al., 2006). Mayewski et al. (2004) identified at least six periods of rapid climate change occurring throughout the Holocene of varying durations. Each event experienced significant enough change that it was detected in global arrays of paleoclimate proxy records (though not every site showed every change). The first period (9–8 ka) was a cool period in the northern Hemisphere. It represented a lagged response to orbital forcings at the end of the Pleistocene and corresponded to the final collapse of the Laurentide ice sheet into the North Atlantic Ocean at 8.2 ka (Alley et al., 1997). Low latitude regions generally saw increased aridity during this time. Subsequent Holocene periods of rapid climate change were also characterized generally by polar cooling and tropical aridity (Mayewski et al., 2004). The primary exception is the most recent period of climate change, 600–150 years BP, when the polar regions were cool but low latitudes wet. The most significant Holocene climate changes involved changes in aridity rather than temperature.

Quaternary Biogeography Overall, the environmental variation that characterized the Quaternary – both in climates and in physical environments – significantly affected the biological diversity of the globe (e.g., Blois and Hadly, 2009). Understanding these impacts and how


Quaternary Biogeography and Climate Change

Phenological change

Physiological change

Behavioral change

Primarily recorded in living systems

Recorded in fossil or living systems

Abundance change

Genetic change Range shift

Morphologic change

Extinction Primarily recorded in fossil systems

Speciation Intercontinental Immigration

Functional turnover

Biotic turnover

Figure 2 Different responses that species may exhibit to climate change, and the systems (modern and/or fossil) in which they are typically studied. For simplicity, only the main pathways between responses are indicated. Responses discussed in the text are highlighted in teal color. Modified with permission from Blois, J.L., Hadly, E.A., 2009. Mammalian response to Cenozoic climatic change. Annual Review of Earth and Planetary Sciences 37, 181–208.

organisms have responded to past climate changes is important for designing effective conservation strategies that optimize the ability of extant species to persist in the face of significant future climate changes. Below, we review some of the biogeographic changes that occurred in response to Quaternary climate change (Figure 2; Blois and Hadly, 2009), among them abundance changes, genetic structure and diversity changes, range shifts, speciation, and extinction. These changes are not independent of one another, nor are they the only responses that can be exhibited by species (e.g., physiological or morphological responses are also seen in response to climate change; Millien et al., 2006; Pörtner and Farrell, 2008); however, they are the responses most

often detected in fossil systems. We then examine how changes within single populations and species scale up to biogeographic responses detected at the assemblage and ecosystem levels.

Population and Species Responses to Climate Change Abundance change One of the first responses that a species shows to climate change is changes in abundance (Lundberg et al., 2000). Because abundance changes determine the size and/or location of populations within a species distribution, they can drive other observed responses to climate change such as genetic

Quaternary Biogeography and Climate Change

change, morphological change, and range shifts (Figure 2; Blois and Hadly, 2009; Hewitt, 2000). As climate changes and the abundance of a species decreases or increases, the species’ ability to withstand perturbations will similarly decrease or increase (Johnson, 1998; Kiessling and Aberhan, 2007; Nogués-Bravo et al., 2008; Purvis et al., 2000). Abundance is a function of the amount and availability of suitable habitat present (Andrewartha and Birch, 1954; del Monte-Luna et al., 2004). How abundance changes as climate changes is a complex function of the interaction between climate and/or habitat change, life history strategies, habitat preferences of a species, and interactions with other species (e.g., Forchhammer et al., 2002). Although a single species has a general range of conditions within which it can survive (e.g., its environmental niche sensu; Grinnell, 1917; Soberón, 2007), the specific shape of environmental fluctuations (e.g., magnitude, duration, and rate of climate change), combined with intraspecific variation in tolerances within populations (such as between adults and juveniles), can result in many possible outcomes: sustained populations, periodic extirpationrecolonization events, or permanent extirpation (Jackson et al., 2009). In the fossil record, abundance changes appear to be largely individualistic responses that vary according to the environmental niche of each species (Blois et al., 2010; Terry et al., 2011; Williams et al., 2004). Paleo-records have also shown that rapid climate change can affect population abundances almost immediately, at times leading to rapid ecosystem restructuring. For example, Yu (2007) determined that forested vegetation in New Jersey responded to the Bølling–Allerod warming event with a lag of at most 200 years (Figure 1), with the cold-tolerant Picea declining and the warm-tolerant Pinus increasing. Responses to subsequent rapid cooling during the Younger Dryas (Figure 1) again occurred almost immediately (e.g., decreased abundances of warm-tolerant oak). Similarly, rapid vegetation response to climate changes led to the restructuring of communities elsewhere in northeastern North America, where pine, hemlock, and spruce abundances declined only to be replaced by another dominant species (e.g., Shuman et al., 2009). Changes in abundance have also been seen in mammals in response to late Quaternary climate changes (e.g., Blois et al., 2010; Grayson, 2006), though limits of dating precision and the taphonomy of many mammalian fossil deposits do not typically allow accurate assessment of mammalian responses to rapid climate change.

Genetic change The advent of ancient DNA (aDNA) has identified significant genetic changes that occurred within species across the past several millennia. Under optimal preservation conditions such as in permafrost settings, aDNA as old as 50 000–65 000 years old can be recovered (Willerslev et al., 2003), though in some exceptional cases recovery of much older aDNA is possible (e.g., Orlando et al., 2013). Ancient DNA provides the opportunity to study the impact of climate change on the structure and genetic diversity within past populations and species (e.g., Hadly et al., 2004; Hofreiter et al., 2004b; Hofreiter and Stewart, 2009). Genetic change is a result of population-level processes such as recombination, mutation, selection, random genetic drift, and gene flow (Charlesworth


et al., 2003), all of which take place in species over tens to thousands of years. Most of these processes may be influenced by climate, either directly such as when climate selects for or against key traits or indirectly through abundance or habitat change, which influences gene flow and genetic drift. For example, aDNA has been used to determine changes in the relative abundance of ancient populations (Chan et al., 2006; Hadly et al., 2004; Lorenzen et al., 2011), whether populations became connected or remained isolated during periods of climate change (Hadly et al., 2004, 1998; Hofreiter et al., 2004a; Teacher et al., 2011), and whether and which clades survived across periods of climate change (Brace et al., 2012; Foote et al., 2013; Fulton et al., 2013). Species reservoirs of genetic diversity and their potential for genetic change may become increasingly important for species to successfully adapt to anthropogenic climate change (e.g., Jump and Penuelas, 2005).

Range shifts One of the most prominent and well-supported responses that species have to climate change is range shifts, both today and in the past (Lyons, 2003; Ordonez and Williams, 2013; Sexton et al., 2009; Tingley et al., 2009; Waltari et al., 2007). Species ranges are limited at broad scales by environmental factors such as temperature, precipitation, and relative humidity (Brown et al., 1996; Eronen and Rook, 2004). As climate changes and environmental gradients shift, species may track their preferred environments at varying directions and rates according to each species physiological limitations (Graham et al., 1996; Tingley et al., 2009; Walther et al., 2002). At one end of the scale, changes in climate may result in only minimal adjustments to species ranges, particularly for species that are not tightly controlled by climate or are located in regions with a large amount of topographic and thus microclimatic complexity (Ackerly et al., 2010; Loarie et al., 2009). On the other hand, changes in climate may result in large range shifts, particularly in areas where climate velocity is large and for species that have good dispersal capabilities (Loarie et al., 2009; Ordonez and Williams, 2013; Schloss et al., 2012). Shifts in species ranges are a direct result of changes in abundance (Figure 2), and occur when existing populations are extirpated and/or new populations are formed just beyond the edges of ranges (Figure 3; Walther et al., 2002). Range expansion is due to episodes of colonization into favorable habit (Walther et al., 2002), and the rate at which a species colonizes new habitat is dependent upon many factors such as climate change, dispersal capability, and microclimate (Lesser and Jackson, 2013). At the very broadest scale, species can move into entirely different continents, such as when humans crossed the Bering land bridge into North America (Achilli et al., 2013; Goebel et al., 2008). These events are more rare than within-continent range shifts and are often accompanied by speciation or biotic turnover. Studies on extant species have illuminated the processes that may underlie present and past range shifts. Not all species alter their ranges in response to climate change, but the species that do are typically ones that live at higher latitudes or elevations and are often found near the edge of their physiological limits (Hickling et al., 2006; McCain and King, 2014; Walther et al., 2002). The magnitude and direction of range shift depends on the species life history characteristics


Quaternary Biogeography and Climate Change

North b






X c


South T1

T2 Past

T3 Present

Figure 3 Potential range shifts of a species as a response to climate change across three time periods. In each panel from time 1 (T1) to time 3 (T3), the geographic extent remains the same but the climatic gradient across space shifts from colder (blue) to warmer (red). At T1, the species is able to persist in a range of environments (gray filled circles), though some of them are unsampled (X) so the range limits (both geographic and climatic, indicated by the dashed line) appear to be narrower than they are. They are not able to persist in other environments (unfilled circles). With increasing global temperatures (e.g., from T1 to T3) individuals from the northern leading edge of the distribution will colonize new localities in search of habitats with suitable climates. As temperatures continue to increase from T2 to T3, previously unsuitable habitats become suitable as local climates warm, while previously suitable habitats become unsuitable. Finally, when temperatures reach a maximum (T3), populations at the southern trailing edge of the distribution become extirpated, due to the loss of suitable habitat. Also, the entire species range has shifted in geographic space to the north. Shifts have occurred in climate space as well (i.e., temperatures at the southern localities are warmer at T3 than anything observed in T1). Shifts in climate space arise from, potentially, adaptation to warmer climate regimes in the south, but adaptation was not as strong as empirically observed because some localities from warm places were unsampled in T1, creating the perception of a narrower climatic niche.

such as dispersal ability, their sensitivity to the environment, and the amount of climate change that occurs at a local scale (McCain and King, 2014). Species that have narrow tolerance ranges are affected more by climate change than species that can tolerate varied climates and habitats (Liow et al., 2009; McCain and King, 2014). The location and distribution of species are also a function of biotic interactions. Direct evidence shows that biotic interactions such as competition can regulate the extent of a species range (Sexton et al., 2009). Additionally, species ranges are smaller at lower latitudes, suggesting that increased biotic interactions and competition in areas with higher species diversity, i.e., the tropics, can dictate the size of a species distribution (Brown et al., 1996; Rapoport, 1982). However, the extent to which biotic interactions are detected in the fossil record is limited (Blois et al., 2014). Information about past species localities has been fairly well preserved in the fossil record and these data have been aggregated in public databases (i.e., FAUNMAP) and museum repositories (Brewer et al., 2012; Uhen et al., 2013), facilitating reconstruction of species ranges at various times in the past (Graham et al., 1996; Lyons, 2003). For example, between the Pre-Glacial to Glacial and Glacial to Holocene time periods, North American mammals shifted their ranges more northward than they did southward (Graham et al., 1996; Lyons, 2003). Plant populations also generally expanded northward across the late Quaternary. Davis and Shaw (2001) provided evidence of this northward shift by

examining spruce pollen percentages in lake sediments from eastern North America and determined that spruce had a more southerly range approximately 21.5 thousand years ago compared to 500 years ago. Similar changes occurred in many other plant genera (Davis, 1983; Huntley and Webb, 1989). The advent of paleo-species distribution modeling (SDM) and paleoclimate simulations has facilitated correlation of past changes in species distributions with climate change (Nógues-Bravo, 2009; Svenning et al., 2011; Varela et al., 2011). Paleo-SDMs have provided insight into, for example, the potential location of glacial refugia (Gavin et al., 2014; Svenning et al., 2008), how past distributions changed as climate changed (e.g., Nogués-Bravo et al., 2008; Ordonez and Williams, 2013; Rodriguez-Sanchez et al., 2010), and the role of climate in past extinction events (Lorenzen et al., 2011; Nogués-Bravo et al., 2008).

Speciation The many glacial–interglacial cycles of the Quaternary may have driven speciation within many plants and animals by shifting climate zones, forming and/or removing environmental barriers, and fragmenting populations (Barnosky, 2005; Hewitt, 1996). Habitat fragmentation can result in populations that are isolated from one another; if isolation persists for an extended period of time, substantial genetic divergence could occur between populations and result in speciation (Barnosky, 2005; Hewitt, 2000; Kadereit et al.,

Quaternary Biogeography and Climate Change 2004; Stewart et al., 2010). However, whether and for how long a population is isolated depends on the interaction between the nature of the barrier and the species environmental tolerances (Hewitt, 2000, 1996). For example, if a species is adapted to cooler climates it will occupy refugia during warmer climate cycles (which are typically short in duration during the Quaternary), whereas species that are adapted to warmer climates will occupy refugia during the longer glacial periods; these differences in isolation time can also affect the timing of divergence between species (Gavin et al., 2014; Kadereit et al., 2004; Stewart et al., 2010). Because so many climate fluctuations occurred, the rate of speciation during the Quaternary should be elevated compared to similar time periods without glacial–interglacial cycles (Avise and Walker, 1998). However, evidence has been provided both for (Jakob et al., 2007) and against (Barber and Jensen, 2012) climate-induced speciation. The glacial and interglacial cycles led to the isolation of populations for relatively short periods of time (at most, approximately 100 thousand years), which resulted in intraspecific genetic differences. In some cases, these genetic differences led to speciation events (Brochmann and Brysting, 2008; Galbreath and Cook, 2004; Jakob et al., 2007). However, the rates of speciation during the Quaternary period are not higher than previous time periods, at least for mammals (Barnosky, 2005; Bobe and Behrensmeyer, 2004), perhaps because the duration of isolation was not long enough to facilitate speciation in these groups. The reconnection of fragmented populations during favorable conditions may have allowed for increased gene flow, which resulted in lower than expected speciation rates for the Quaternary period (Barnosky, 2005; Klicka and Zink, 1997; Zink et al., 2004).

Extinction A major extinction event occurred during the Quaternary between approximately 50 ka and 10 ka (Barnosky et al., 2004; Koch and Barnosky, 2006). During this time, 97 of the 150 largest mammalian genera went extinct across all continents (Barnosky et al., 2004; Lyons et al., 2004; Roberts et al., 2001), though extinction severity was spatially heterogeneous. Africa was not affected as heavily as the other continents and only lost 18% of its genera, whereas continents such as North America and South America lost 76% and 86% of their genera, respectively (Nógues-Bravo et al., 2010). Megafaunal extinction was also heterogeneous through time (Barnosky et al., 2004; Hofreiter and Stewart, 2009; Stuart and Lister, 2012). Many hypotheses about the cause of the megafaunal extinction have been postulated, but the two main hypotheses are climate change and human interactions (Barnosky et al., 2004; Koch and Barnosky, 2006; Lyons et al., 2004). Humans clearly impacted megafaunal populations and caused some species to go extinct, but the exact mechanisms of human impact are still unclear. Potential human drivers of extinction include habitat degradation and fragmentation by practices such as deforestation and fire (Burney et al., 2004; Robinson et al., 2005), overhunting (Brook and Johnson, 2006; Lyons et al., 2004; Martin, 1967), and pandemic disease introduced by humans (Alroy, 2001; Rothschild and Laub, 2006). Of these drivers, the leading hypothesis is overhunting (Barnosky et al., 2004; Haynes, 2013; Koch and Barnosky, 2006; Lyons et al., 2004; Martin, 1967). Several studies have examined the timing of the megafaunal extinctions on each


continent and found a strong correlation between human arrival and megafaunal extinction (e.g., North America and Australia; Barnosky et al., 2004; Johnson et al., 2013). Human arrival has also been correlated with the extinction of other groups such as birds (Steadman, 1995). The ability of humans to severely impact megafauna through overhunting is thought to result from the naivety or lack of coevolution of animals with humans on most continents and the slow reproductive rates of large mammals (Haynes, 2013; Lyons et al., 2004; Zuo et al., 2013). However, many of the extinct megafauna show no signs of being hunted by humans (Grayson, 1984; Koch and Barnosky, 2006) and some species may have coexisted with humans for thousands of years before going extinct (Lima-Ribeiro and Diniz-Filho, 2013; Wroe et al., 2004). Another alternative is that extinction could have occurred due to indirect effects of hunting (Lyons et al., 2004), such as habitat change or ecological cascades resulting from extinction of just a few key megafaunal species (Gill et al., 2009). All of these factors contribute uncertainty about the total impact of overhunting on megafaunal extinction (Sandom et al., 2014). Climate change at the end of the last glacial period has also been hypothesized as a cause of megafaunal extinction, perhaps by reducing the types and quality of habitat available to the animals and thus reducing the resources necessary to support species (Koch and Barnosky, 2006). For example, the megafaunal extinction rate was positively correlated with the magnitude of climate change on all continents except for South America, which only had a moderate amount of climate change during the late Quaternary period (Nógues-Bravo et al., 2010). Some species went extinct before the local arrival of humans and the timing of extinction was correlated with significant climate changes in some systems, ruling out humans as the cause of extinction for those taxa (Barnosky, 1986; Stuart and Lister, 2012; Wroe et al., 2013). Despite the apparently clear link between extinction and climate, the climate change hypothesis has had a difficult time explaining why only large mammals were affected and not other groups such as marine organisms or plants (Lyons et al., 2004; Martin, 1984), though size-dependent differences in reproductive rates may have been a factor (Johnson, 2002; Zuo et al., 2013). Also, large mammals that were nocturnal, arboreal, and lived in inaccessible forests (e.g., mammals that were not as exposed to human hunting) were not as drastically affected by the megafaunal extinction (Koch and Barnosky, 2006). Finally, climate change hypotheses have a difficult time explaining why there was an increased extinction rate for large mammals during the LGM and not in previous glacial periods (Lyons et al., 2004). Although many studies have proposed either humans or climate change as the culprit for the late Quaternary megafaunal extinctions, recent studies show that the mechanism of extinction for many species is likely a combination of both human impacts and climate change (Haynes, 2014; Lorenzen et al., 2011; Nogués-Bravo et al., 2008; Prescott et al., 2012). These studies postulate that climate change altered species abundances and distributions at the end of the last glacial period, such that the smaller populations were more susceptible to overhunting from humans (e.g., Nogués-Bravo et al., 2008). Since some continents had high extinction rates with low climate change (Nógues-Bravo et al., 2010) and other species went extinct with no evidence of human hunting or interference, extinction mechanisms are likely location- and/or species-specific.


Quaternary Biogeography and Climate Change

Community Responses to Climate Change The combined effect of changes within individual species across the late Quaternary led to significant changes in community attributes such as diversity, evenness, and composition. For example, the small mammal assemblage in northern California showed significant losses in diversity across the last glacial–interglacial transition, even though no small mammal species went globally extinct (Blois et al., 2010). These changes were similar to those seen in other small mammal assemblages across the west (e.g., Grayson, 2006; Lyman, 2014) and were postulated to be due to climate change favoring more generalist species (Blois et al., 2010). Another significant biogeographic change was the appearance and disappearance of ‘no-analog’ assemblages. Small mammals occurred in no-analog assemblages, where species that today live in very different and spatially separated localities coexisted in the past (e.g., Graham et al., 1996). Noanalog assemblages have also been seen in plant communities and were linked to climate changes. For example, peak dissimilarity in plant assemblages occurred at the same time as peak dissimilarity in climates, in both Alaska and eastern North America (Williams and Jackson, 2007) and dissimilarity in northeast pollen assemblages spiked at times of significant climatic transitions (Shuman et al., 2009). Similarly, spatial dissimilarity among plant assemblages in eastern North America peaked from 14–11 ka. These relationships were primarily due to climatic factors (Blois et al., 2013) and not to biotic interactions (Blois et al., 2014). Despite the clear evidence for the impact of climate change on assemblage structure and diversity, biotic interactions (at least cross-trophic interactions) clearly also were a factor. For example, no-analog conditions were closely correlated with the functional loss of megafauna from the system, indicating that herbivory by megafauna was a key driver maintaining plant assemblages (Gill et al., 2009, 2012).

Conclusion Overall, the Quaternary encompassed a time of significant climatic and biogeographic change. While species and communities responded to a multitude of drivers, the physical and climatic changes that affected the earth system over the past several million years were primary. However, much of the data for both climatic and biogeographic processes in the Quaternary is drawn from the past few millennia – from the LGM to the present. Although this perspective has provided a detailed understanding of biogeographic responses to deglaciation and global warming, comparatively less is known about biogeographic responses to environmental change during the thousands and millions of years prior to the LGM, and across other periods of rapid climate changes. Additionally, much of the focus has been on biogeographic responses of a few groups, such as pollen or mammals, due to the different nature of the fossil record between groups. Today, however, biogeographic patterns and processes are being altered by extensive human impacts, so much so that a new geological epoch has been proposed, the Anthropocene (Zalasiewicz et al., 2011). The extent to which ecological processes today are

altered compared to the past is unknown, and a question that can only be answered by combining paleo- and present-day data. Integrating paleobiogeographic data from multiple components of our biotic systems, understanding responses of all components across periods of rapid warming events, and establishing relevant biogeographic baselines for past processes all represent key frontiers in paleobiogeography and avenues that are particularly relevant for conservation for present-day species.

See also: Biogeography, Evolutionary Theories in. Responses to Climate Change, Evolution and

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