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2 Lessons from Past Warm Worlds Alfred Wegener’s concept of continental drift, reformulated in the modern theory of plate tectonics, arose in part as a way to explain the geographic distribution of paleoclimate indicators in ancient rocks. Permo-Carboniferous (~300-million-year-old) glacial deposits in distinctly nonpolar regions of present-day Africa, South America, and Australia rectify to polar latitudes when the ancient supercontinent of Gondwana is reconstructed. Continental drift transported them through broad climate belts—humid tropics, arid subtropics, moist and cool temperate zones, and cold and arid polar regions (Box 2.1). Paleoclimate reconstructions, however, reveal that although paleogeography and the plate tectonics that control continental configurations are important, they are not the major determinant of climate change. Global warm climates have prevailed when large continents covered the poles, and deep “snowball Earth” gla - ciations occurred when there apparently were no polar continents. Instead, it appears that the greenhouse gas content of the atmosphere was the key factor in determining whether a particular interval of Earth’s past was an icehouse or a greenhouse. Although most deep-time greenhouse climates occurred when there were distinctly different continental configurations, and thus are not direct analogues for the future, past warm climates and abrupt transi- tions into even hotter states (known as hyperthermal events; Thomas et al., 2000) provide important insights into how physical, biogeochemical, and biological processes operate under warm conditions more analogous to what is anticipated for the future than the moderate and stable climates of the Holocene (past 10,000 years) or the relatively warm interglacials of 26
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27 LESSONS FROM PAST WARM WORLDS BOX 2.1 Continental Drift and Climate Plate tectonics has been rearranging Earth’s configuration of continents ever since the plates on Earth became rigid approximately 2.5 billion years ago (Figure 2.1). On long (millions of years) timescales, the movement of tectonic plates—and the continents that ride upon them—has strongly influenced Earth’s distribution of solar insolation, ocean and atmospheric circulation, and carbon cycling between the Earth’s deep and shallow reservoirs, thereby profoundly impacting global climate, sea level, and the overall planetary ecology. The arrangement of the continents through time is most reliable for the past 800 million years, the period for which the chronostratigraphic tools necessary for reconstructions are available. The global views presented in Figure 2.1 show how the continents on Earth’s surface may have appeared during four intervals of time that are noted throughout this report: the unipo- lar glaciated Pennsylvanian (300 million years ago [Ma]), mid-Cretaceous (105 Ma), Eocene (50 Ma), and mid-Pliocene (3 Ma). The major transitions between climatic icehouse and greenhouse con- ditions are ultimately most probably driven by the deep Earth processes of plate tectonics, as a function of the long-term balance between CO2 degassing at spreading centers and the conversion of atmospheric CO2 to mineral carbon through long-term silicate weathering and oceanic carbon- ate formation (Berner, 2004). For example, the eruptions of large igneous provinces in the mid-Cretaceous and the subduction of the carbonate-rich tropical Tethys Sea in the early Cenozoic are the most likely cause of the high-CO2 equilibrium climates of the Cretaceous and Eocene greenhouses. Conversely, uplift of the Himalayas and Tibetan Plateau associated with “docking” of the Indian subcontinent with Asia (~40 Ma), and the evolu- tion of vascular land plants in the early Paleozoic (~450 Ma), led to the sequestration of atmospheric CO2 through enhanced weathering of silicate minerals (Ruddiman, 2007; Archer, 2009). continued
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28 BOX 2.1 Continued FIGURE 2.1 Continental configurations for the Pennsylvanian (upper left), mid-Cretaceous (upper right), Eocene (lower left), and mid-Pliocene (lower right). Topography was defined on the basis of digital elevation maps of modern Earth from the U.S. Geological Survey; colors portray climate and vegetation distribution based on a synthesis of all geological literature relevant to each time slice. SOURCE: Courtesy R.C. Blakey, Colorado Plateau Geosystems. UNDERSTANDING EARTH’S DEEP PAST
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29 LESSONS FROM PAST WARM WORLDS the Pleistocene (past 2 million years). The following sections describe the insights provided by understanding past warm periods, including the role of greenhouse gases in controlling—or “forcing”—global warming; the impact of warming on ice sheet stability, sea level, and oceanic and hydrological processes; and the consequences of global warming for eco - systems and the global biosphere. CLIMATE SENSITIVITY TO INCREASING CO2 IN A WARMER WORLD Fundamentally, Earth’s climate results from the balance between absorbed energy from the sun and radiant energy emitted from Earth’s surface, with changes to either component resulting in a forcing of the climate system. The net forcing of the climate system over geological time caused episodes of warming and cooling that are coincident with greenhouse and icehouse climates, respectively. Most projections indicate that, by the end of this century, climate forcing resulting from increased CO2 will be at least of the same magnitude as that experienced in the early Cenozoic (during the late Eocene, ~34 Ma) (Figure 2.1), and possibly analogous to estimates for the Cretaceous Period (~80-120 Ma)—probably one of the times of greatest radiative forcing since the evolution of animals (Hay, 2010). Climate sensitivity—the equilibrium warming resulting from a dou- bling of atmospheric carbon dioxide relative to preindustrial levels of CO2—provides a measure of how the climate system responds to external forcing factors and is also used to compare global climate model outputs to understand why different models respond to the same external forcings with different outputs. Climate sensitivity to CO2 strongly influences the magnitude of warming that Earth will experience at any particular time in the future (Box 2.2). The magnitude of climate sensitivity and Earth’s surface temperature are determined by a myriad of short-term (human timescales) and long-term (thousands to tens of thousands) interactions and feedbacks (e.g., water vapor, cloud properties, sea ice albedo, snow albedo, ice sheet and terrestrial biome distribution, ocean-atmosphere CO2 interaction, and silicate weathering). As noted above, synthesis of the various estimates of Earth’s climate sensitivity for the past 20,000 years has lead to the general conclusion that sensitivity most probably lies in the range of 1.5 to 4.5°C (IPCC, 2007), with some recent projections suggesting that the value may be even as high as 6-8°C (Hansen et al., 2008; Knutti and Hegerl, 2008). However, estimates of equilibrium climate sensitivity averaged over tens to hundreds of millen - nia (i.e., long term) and extending back for 400 million years are minimally between 3 and 6°C (Royer et al., 2007). For the most recent period of global
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30 UNDERSTANDING EARTH’S DEEP PAST FIGURE 2.2 Estimated atmospheric pCO2 for the past 45 million years (late Eocene through Miocene) calculated using all available stable carbon isotopic values of diunsaturated alkenones in deep-sea sediments. Values of CO2aq were translated to atmospheric pCO2 using Henry’s Law and a range of dissolved phosphate values and sea surface temperatures for each site, and a salinity of 35 parts per thousand. The dark gray shaded region shows the range of maximum to intermediate esti- mates, and the dashed line represents minimum estimates. The uncertainty in pCO2 estimates ranges from ~20 percent for the Miocene to 30 to 40 percent for the Paleogene. The broad pale red band (pCO2 values of 600-1,100 parts per million by volume) encompasses most of the CO2 concentration range for nonmitigation emission scenarios projected for the end of this century (figure 10.26 in IPCC, 2007); the dark red band (values of 800-1,000 ppmv) corresponds to the Intergovernmen - tal Panel on Climate Change (IPCC) A2 “business-as-usual” scenario. SOURCE: Modified after Pagani et al. (2005). warming, the middle Pliocene (~3.0-3.3 Ma), climate sensitivity may have been as high as 7-9.6°C ± 1.4°C per CO2 doubling (Pagani et al., 2010). Such values, which are well above short-term climate sensitivity estimates based on more recent paleoclimate and instrumental records, indicate that long-term feedbacks operating at accelerated timescales (decadal to cen - tennial) promoted by global warming can substantially magnify an initial temperature increase. As Earth moves toward a warmer climate state, it is important to under- stand the extent to which climate sensitivity will change due to processes
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31 LESSONS FROM PAST WARM WORLDS BOX 2.2 Why Does Climate Sensitivity Matter? For any particular increase in atmospheric CO2 (and other greenhouse gases), a system with high climate sensitivity to CO2 will warm more in the future than a world with low climate sensitivity. Thus, if the climate sensitiv- ity is high, restricting future global warming will require a larger reduction in future CO2 emissions than if climate sensitivity is lower. Comparison of emission scenarios for the period until 2100, calculated for a range of CO2 stabilization targets (Figure 2.3A) and the corresponding equilibrium global average temperature increases (Figure 2.3B; IPCC, 2007), based on the Intergovernmental Panel on Climate Change (IPCC) range of climate sensitivities (2 to 4.5°C), illustrates the impact of fossil carbon emissions on future surface temperatures and the extent of reductions required to limit the warming to ≤2°C relative to preindustrial conditions. Even if anthropogenic carbon emissions to the atmosphere are reduced, CO2 levels will continue to increase for a century or more because the removal of CO2 from the atmosphere by natural processes of carbon seques- tration (e.g., CO2 absorption by the surface ocean, CO2 fertilization of ter- restrial vegetation) is slow (Archer et al., 2009). Consequently, temperature increases may continue for several centuries until equilibrium temperatures are reached, especially for higher CO2 stabilization targets. However, even if climate sensitivity is at the lower end of the possible range, global tem- perature increases of ≥2°C will be reached with CO2 stabilization levels of 450-550 parts per million by volume. Given that equilibrium temperature increases may be protracted, emissions could continue to increase into the middle of this century (Figure 2.3; Caldeira et al., 2003). However, if the climate sensitivity is 4.5°C or greater, then a significant and immediate reduction in CO2 emissions—to levels ultimately below those of the present day—is required to stay below a target warming of 2°C. continued
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32 UNDERSTANDING EARTH’S DEEP PAST BOX 2.2 Continued World CO2 emissions (GtCO2/yr) A Year FIGURE 2.3 Global CO2 emissions and equilibrium global average temperature increases for a range of target CO2 stabilization levels. (A) Measured (1940 to 2000) and projected (colored shading; 10th to 90th percentile) global CO2 emissions for the range of IPCC emission scenarios and associated stabilization CO2 levels indicated by roman numerals (ppm CO2-eq). (B) Corresponding relationship between the different that have not operated in recent icehouse climate regimes. One simple example is to consider a warm world with no sea ice at either pole—as CO2 increases, the sea ice albedo feedback is removed and therefore this nega- tive feedback’s contribution to climate sensitivity is absent. In addition, it is important to determine the potential for nonlinear responses that are specific to a greenhouse or transitional world, and whether such responses would enhance climate sensitivity. For example, the destabilization of con- tinental ice sheets resulting from warming of polar regions can potentially lead to a decrease in deep-water formation, thereby affecting global ocean circulation, stratification, and carbon cycling, leading to higher climate sen- sitivity than indicated by present estimates. In sufficiently warm climates, even water vapor has a nonlinear dependence on temperature, and this can introduce new and potentially rapid feedbacks, operating at a subdecadal scale, into the climate system. Destabilization of methane and its release
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33 LESSONS FROM PAST WARM WORLDS E quilibrium global average temperature increas e above preindustrial (°C) B G HG concentration s tabilis ation level (ppm C O2 -eq) CO2 stabilization targets shown in (A) and equilibrium global average temperature increase above preindustrial levels. Colored regions for each stabilization target were calculated for a range of climate sensitivity (2–4.5°C) and “best estimate” climate sensitivity of 3°C (blue solid line in middle of shaded area). SOURCE: IPCC (2007, Figure 5.1, page 66). into the atmosphere in response to warming— through either the melting of terrestrial permafrost reservoirs or the dissolution of subseafloor clath- rate deposits—would dramatically increase greenhouse gas contents in the atmosphere. Hence, an initial warming from greenhouse gases released by burning fossil fuels could end up releasing even more greenhouse gases from natural sources, exacerbating the original warming of the atmosphere. TROPICAL AND POLAR CLIMATE STABILITY AND LATITUDINAL TEMPERATURE GRADIENTS IN A WARMER WORLD With more than half of Earth’s surface lying within 30° latitude of the equator, the response of tropical climates to increased greenhouse gas forcing is critically important. Modern observational data (Ramanathan and
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34 UNDERSTANDING EARTH’S DEEP PAST Collins, 1991) suggest that western tropical Pacific sea surface temperatures rarely exceed ~30-32°C, and this has led to speculation that the Earth’s tropics have a “thermostat” that limits maximum sea surface temperatures. Explanations about how such a thermostat might work have included the buildup of clouds that reflect heat back into space (Ramanathan and Collins, 1991), evaporative cooling (Hartmann and Michelsen, 1993; Pierrehumbert, 1995), winds, or an increase in transport of heat out of the tropics by ocean currents (Clement et al., 1996; Sun and Liu, 1996). Most of these studies have used present-day data to explain surface temperature regulation, although there are artifacts in these datasets that call into question the robustness of the observed trends (Clement et al., 2010). Paleoclimate reconstructions of tropical temperatures during past greenhouse times, however, document sea surface temperatures that were much warmer than modern tropical maxima—possibly as high as 42°C—and thus were probably not thermo- statically “regulated” (Bice et al., 2006; Came et al., 2007; Pearson et al., 2007; Trotter et al., 2008; Kozdon et al., 2009). The discovery of a giant Paleocene snake fossil in South America (Head et al., 2009; Huber, 2009; although see discussions by Makarieva et al., 2009; Sniderman, 2009), as well as other terrestrial paleotemperature indicators such as paleoflora leaf-margin analysis and stable isotope com - positions of biogenic apatites and soil minerals (Fricke and Wing, 2004; Tabor and Montañez, 2005; Passey et al., 2010), further suggests anoma- lously high continental temperatures (~30-34°C) for the terrestrial tropics of past warmer worlds. Additionally, coupled climate model simulations with large radiative forcings and/or paleoclimate simulations for elevated greenhouse gases do not produce a thermostatic regulation of tropical temperatures (e.g., Boer et al., 2005; Poulsen et al., 2007b; Cherchi et al., 2008), suggesting that the tropical warming in response to greenhouse gas forcing is neither moderated nor local in its impacts (Xie et al., 2010). Such deep-time paleoclimate studies have documented that tropical surface temperatures during past greenhouse periods were not thermostatically regulated by the negative feedback processes that operate in the current icehouse climate system, further illustrating how knowledge of deep-time warm periods is fundamental to understanding Earth’s climate system. There is also abundant evidence for anomalous polar warmth during past greenhouse periods (e.g., middle Cretaceous to Eocene, Pliocene; see Figure 2.4) associated with reduced equator-to-pole temperature gradients (e.g., Huber et al., 1995; Crowley and Zachos, 2000; Hay, 2010; Miller et al., 2010). To date, climate models have not been able to simulate this warmth without invoking greenhouse gas concentrations that are notably higher than proxy estimates (Figure 2.4; Bice et al., 2006). This has prompted mod- eling efforts to explain high-latitude warmth through vegetation (DeConto et al., 1999), clouds (Sloan and Pollard, 1998; Abbot and Tziperman, 2008;
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Eocene Cretaceous 55–48 Myr ago 45–39 Myr ago j Modern k g 40 d b j g h k 30 f n e p SST (°C) d m c 20 b a p 10 a S –80 –60 –40 –20 0 20 40 60 80 N Latitude FIGURE 2.4 Zonal mean surface temperature (°C) as a function of latitude for the Eocene and Cretaceous. For the Eocene, the solid lines represent second-order polynomials excluding triangle data, with error bars representing the range of variation. Dashed lines represent deep-sea temperatures. For the Cretaceous, the dotted line is the modern observed zonal surface temperature, while the symbols indicate a compilation of empirical data for specific periods. SOURCES: Eocene—modified from Bijl et al. (2009), reprinted by permission of Macmillan Publishers Ltd.; Cretaceous—courtesy K. Bice, personal communication, 2010. 35
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36 UNDERSTANDING EARTH’S DEEP PAST Kump and Pollard, 2008), intensified heat transport by the oceans (Barron et al., 1995; Korty et al., 2008), and increased tropical cyclone activity (Sriver and Huber, 2007; Fedorov et al., 2010). The ability to successfully model a reduced latitudinal temperature gradient state, including anomalous polar warmth, presents a first-order check on the efficacy of climate models as the basis for predicting future greenhouse conditions. Since significant changes in tropical and polar surface temperatures and pole-to-equator temperature gradients occurred in the past, and could occur in a future warmer world, it is imperative to understand the mecha- nisms and feedbacks that lead to such changes and their consequences for atmospheric and oceanic circulation (Hay, 2008). The fundamental mis - match between model outputs, modern observations, and paleoclimate proxy records discussed above, however, may indicate some very impor- tant deficiencies in scientific knowledge of climate and the construction of climate models (e.g., Huber, 2008). Resolution of this disparity, as well as an improved understanding of the anthropogenic signal in observational data, can likely be obtained by analysis of paleoclimate records from past warm worlds. HYDROLOGICAL PROCESSES AND THE GLOBAL WATER CYCLE IN A WARMER WORLD Earth’s hydrological processes—including precipitation, evaporation, and surface runoff—are susceptible to, and play a critical role in, both past and future climate change (Pierrehumbert, 2002). Large-scale atmo- spheric processes determine the general position of climate zones and the intensity of precipitation and storms; the intertropical convergence zone is a region of significant rainfall, while large regions of atmospheric subsidence lead to dry desert regions. Regional hydroclimates, such as the Southwest Indian and the East Asian summer monsoons, which affect nearly half of Earth’s human population, are highly sensitive to distal climate changes and to mean warming (Sinha et al., 2005; Wang et al., 2005) via teleconnections (e.g., changes in high-latitude surface tempera- tures or Arctic sea ice extent impact lower-latitude climate through atmo- spheric processes). Overall, the vapor-holding capacity of the atmosphere increases substantially with increased global mean temperatures if there is no change in the relative humidity. Consequently, climate models for global warming predict an intensified hydrological cycle and, on a global scale, enhanced precipitation (IPCC, 2007). Observations over the past few decades indicate that precipitation has increased faster (~7 percent per degree of surface warming; Wentz et al., 2007) than that predicted by models (1-3 percent per degree of surface warming; Zhang et al., 2007). Although the reasons for this substantial
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52 Multiple hypoxia events—Mediterranean Thickness (cm) 50 25 0 75 125 100 Ocean Anoxic Event 1B, Western North Atlantic Thickness (cm) 50 25 0 75 125 100 FIGURE 2.11 Examples of ancient hypoxic episodes in a Plio-Pleistocene drill core from the Mediterranean. Black bands in the upper photo (from Ocean Drilling Program Site 964) are “sapropels”—layers rich in organic carbon—formed when the surface waters of the Mediterranean abruptly warmed, became fresher, and ceased to circulate as they do today. Warming and high nutri - ent supply led to blooms of algae and bacteria, preserved as a layer of organic carbon on the seafloor. Lower photo shows a core from the western North Atlantic (from ODP Site 1049) in which a similar layer rich in organic carbon was deposited during an event 112 million years ago that was broadly analogous to the Mediterranean hypoxic events. SOURCE: Images courtesy Integrated Ocean Drilling Program Science Services.
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53 LESSONS FROM PAST WARM WORLDS continue to increase well above the levels of the Pleistocene interglacials and as the geographic distributions of climate zones change. For example, higher CO2 levels are expected to saturate the CO2 fertilization effect, resulting in a shift of the terrestrial biosphere from a net sink to a net source of carbon sometime within this century (Cao and Woodward, 1998; Cox et al., 2000). Furthermore, as the surface oceans warm and become less alkaline with increasing atmospheric CO2, carbonate-bearing animals will be strongly impacted (e.g., see Box 2.6), further perturbing biota-climate feedbacks compared with those reconstructed from the recent past. The deep-time geological record, in particular the record of warm periods of higher atmospheric pCO2 and including the transitions into and out of these periods, has the potential to yield unique insights into the nature and rate of biotic response to climate perturbation as well as into the biota-climate feedbacks accompanying global warming. For example, the mid-Paleozoic “greening” of continents, marked by the evolution and spread of vascular land plants (Gensel and Andrews, 1987; Beerbower et al., 1992), records a large-scale natural experiment in the climatic effects of vegetation—reflecting the contrast between a largely unvegetated pre- Devonian world compared with one that was heavily vegetated—that has been linked to major changes in atmospheric CO2 and a vastly dif- ferent hydrological regime (Algeo et al., 1995, 2001). Another example of the potential of the deep-time record is provided by the repeated major restructuring and turnover within terrestrial floral communities that occurred in step with recurrent shifts in surface temperature, precipita- tion levels, seasonality, and soil moisture during the demise of the Late Paleozoic Ice Age at ~295-260 Ma, the vegetated Earth’s only analogue of a CO2-forced icehouse-to-greenhouse transition (see Box 2.7). More recently, the gradual but extreme warming (perhaps up to >30 to 42°C in the tropics) of the early Eocene greenhouse (Box 2.8) may have triggered a major tropical vegetation die-off, with substantial changes in evapotranspiration fluxes, precipitation, albedo, surface temperature, and carbon feedbacks (Huber, 2008). During the transient global warming and short-term aridity of the Paleocene-Eocene Thermal Maximum (PETM) major restructuring among terrestrial biomes resulted in expansion in the latitudinal range of subtropical and tropical rainforests (Wing et al., 2005). Oxidation of the terrestrial biosphere at the Paleocene-Eocene boundary may have released several gigatons of carbon into the atmosphere, sub - stantially amplifying the existing greenhouse warming and its climate effects (Kurtz et al., 2003). The potential vulnerability of modern biotic communities to cata- strophic disruption (Jackson et al., 2001; Chase and Leibold, 2003) is an issue designated as one of the “grand challenges” in the environmental sciences (NRC, 2001). Globally, current extinction rates are estimated to
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54 UNDERSTANDING EARTH’S DEEP PAST BOX 2.6 Impact of Past and Future Climate Change on Coral Reefs Healthy coral reef ecosystems develop under a relatively narrow range of ocean temperatures and chemistry (Kleypas et al., 1999) and are there- fore sensitive indicators of environmental conditions. Global change models predict that reef systems, with their abundant biodiversity, will be exposed to higher ocean temperatures and increasingly more acidic waters in the next century (Hoegh-Guldberg et al., 2007; see Figure 2.12). Indeed, research suggests that global climate change has already caused steep declines in coral growth on reef systems around the world (Hoegh- Guldberg, 1999). Culturing experiments with corals in acidified waters show that skeleton growth drops as acidity increases and, in extreme cases, coral colonies can lose their skeletons completely and grow as soft-bodied anemone-like animals (Fine and Tchenov, 2007). In fact, ocean acidification may vie with global warming as the most severe threat to marine ecosystems (Hoegh-Guldberg et al., 2007; De’ath et al., 2009). Reef systems, however, are intrinsically complex structurally and ecologically, making it difficult to evaluate the likely impact of future global change on modern reefs based solely on studies of present-day systems. The geologic record of fossil reef evolution provides opportunities to study the response of reef ecosystems to past episodes of increased global temperatures and ocean acidification. The coral reef crisis occurring in modern oceans may be the sixth such major reef crisis recorded in the past 500 million years of marine metazoan evolution. Four of the previous five metazoan reef crises appear to have been driven by greenhouse gas-forced global warming that was probably associated with ocean acidification (Veron, 2008; Kiessling and Simpson, 2010). At least three of these reef crises were associated with massive re- lease of greenhouse gases into the oceans and atmosphere, leading to pCO2 increases analogous to—or perhaps even greater than—those anticipated for Earth’s future. For example, major reef crises during the Early Jurassic and during the Cretaceous were associated with massive releases of vol- canic CO2 to the atmosphere that led to global warming, oceanic anoxia, and quite likely ocean acidification (Knoll et al., 1996; Svensen et al., 2007; Hermoso et al., 2009). One of the major reef crises occurred at the same time as the best-documented case of greenhouse gas-induced ocean acidi- fication in the geological record, the Paleocene-Eocene Thermal Maximum (PETM) of 56 Ma (described in more detail in the next chapter). Although coral-algal reefs began to decline throughout the Tethyan region in the early Eocene due to the development of very warm (~30-35°C) tropical sea surface temperatures (Scheibner and Speijer, 2008) (Figure 2.13), PETM extinction rates indicate that ocean acidification must have been a major continued
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LESSONS FROM PAST WARM WORLDS FIGURE 2.12 Extant examples of reefs from the Great Barrier Reef that are used as analogues for the ecological structures anticipated for atmospheric CO2 values of (A) 380 ppmv, (B) 450-500 ppmv, and (C) >500 ppmv (Hoegh- Guldberg et al., 2007). Scenario C corresponds to a +2°C increase in sea temperature. The atmospheric CO 2 and temperature increases shown are those for the scenarios and do not refer to the particular locations photographed. SOURCE: Photographs by and with permission of Ove Hoegh-Guldberg, Global Change Institute, University of Queensland. 55
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56 UNDERSTANDING EARTH’S DEEP PAST BOX 2.6 Continued causative factor along with global warming (Kiessling and Simpson, 2010). Notably, the deep-time record of this major reef crisis uniquely captures the consequences on larger-scale marine ecosystems that might be anticipated with the future loss of reefs. Thanetian- E. Eocene L. Thanetian Paleo- Area Selandian Clim. Optimum (55.2-56.3 Ma) latitude (56.3-60 Ma) (50-55.2 Ma) 43°N N. Calcareous Alps, to W. Carpathians 32°N Italy, Greece N. Adriatic platform, 38°N Pyrenees 20°N Egypt, Oman to 12°N 5°N NW India, Somalia to 0° 5°N to NE India, Tibet 5°S Individual corals Small patch reefs Coral-Algal reefs Larger foraminifera banks FIGURE 2.13 Paleogene reef history from the Mediterranean area and southern Asia. Coral-algal reefs that are widespread in the early Paleo- gene largely disappeared during the peak of greenhouse warming in the Early Eocene Climatic Optimum, and were replaced by carbonate mounds formed by larger benthic foraminifera (nummulite banks). SOURCE: Modified from Scheibner and Speijer (2008).
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57 LESSONS FROM PAST WARM WORLDS BOX 2.7 Climate-Driven Restructuring of Late Paleozoic Tropical Forests Integration of climate proxy records with tropical paleobotanical archives from the Late Paleozoic shows repeated climate-driven ecosystem restructuring of paleotropical flora in step with climate and pCO2 shifts, illustrating the biotic impact associated with past CO2-forced turnover to a permanent ice-free world (Montañez et al., 2007; DiMichele et al., 2009). Wetland flora—consisting of ferns and pteridosperms, sphenopsids, and lycopsids—was rapidly replaced in the earliest Permian by dryland flora that diversified in the now seasonally dry habitats created by an abrupt shift from ever-wet to semiarid conditions. Tree and fern-rich floras reappeared during wetter, cooler conditions of the subsequent glaciation at ~285 Ma characterized by lowered pCO2 (Figure 2.14). Such dramatic floristic changes occurred with each climate transition during the final stage of the Late Paleozoic Ice Age. The fact that these temporally successive floras tracked climatic condi- tions and contained progressively more evolutionarily advanced lineages suggests that evolutionary innovation occurred in extrabasinal areas and was revealed by climate-driven floral migration into lowland basins. One such scenario occurred during the return to cold conditions at the close of the Early Permian when unique seed-plant assemblages, not observed again until the Late Permian (conifers) and Mesozoic (cycads), migrated into low- land basins. Climate transitions drove macroevolution in the oceans as well, including significant changes in marine invertebrate biodiversity coincident with the appearance of a diverse array of early terrestrial vertebrate lineages and major restructuring of floral biomes (Clapham and James, 2007). continued
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58 UNDERSTANDING EARTH’S DEEP PAST BOX 2.7 Continued
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59 LESSONS FROM PAST WARM WORLDS FIGURE 2.14 Floral abundance patterns (A and B) in the paleotropics during the latest Pennsylvanian through Middle Permian, plotted against (C) estimated pCO2 (blue line) and paleo-sea surface temperatures (red band). Periods of glaciation or widespread cooling in the high southern latitudes are shown by blue bars. Top panel (A) shows the temporal distribution of typical latest Carboniferous wetland floras (ferns and pteridosperms, sphenopsids, and lycopsids). The middle panel (B) illus - trates the temporal distribution of dryland floras (conifers, callipterids and other seed plants) that diversified in seasonally dry habitats. The short-term intercalation of the two floras—at the likely millennial scale— occurred with the return of wetland floras in the mid Early Permian transient glaciation under cooler and wetter conditions brought on by significantly lowered pCO2 and renewed glaciation. SOURCE: Modified after Montañez et al. (2007).
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60 UNDERSTANDING EARTH’S DEEP PAST BOX 2.8 Biome Distributions in the Cretaceous-Early Eocene Hothouse The discovery more than a century ago of coal seams, fossil forests, and fossil leaves of warm temperate trees above the Arctic Circle on the west coast of Greenland, more than 1,000 km north of the modern tree line, was an early indicator of anomalous warmth at high latitudes in the past. That arctic regions more than 50 Ma had been forested as far north as there was land, despite a continental configuration similar to that of the present day, became increasingly more apparent with the discovery of hundreds of similar sites on arctic islands and across arctic Asia and North America (Spicer et al., 2008). Northern hemisphere Cretaceous-early Eocene polar forests were fundamentally different from present-day boreal forests, which do not grow north of the Arctic Circle, as they are dominated by deciduous conifers related to the bald cypress and dawn redwood and by a variety of deciduous broadleaf trees. Leaf margin analysis of Paleocene-Eocene floras from arctic Canada on Axel Heiberg Island (at 78ºN) yield mean annual temperatures of 10º ± 2ºC (Basinger et al., 1994), in striking contrast to the modern mean annual temperatures of minus 30ºC. The fossil floral record indicates that in this warmer world, both subtropical and tropical rainforests had greatly expanded latitudinal ranges (Figure 2.15). Subtropical conditions in the polar Arctic are further indicated by the occurrence of early Eocene crocodiles, turtles, and snakes on Ellesmere Island at 80ºN at this time (Dawson et al., 1976; Markwick, 2007). Sub- sequent discovery of fossil mammals and plants, related to contemporary biotas in France and Wyoming, confirmed the hypothesis that Arctic Canada at this time was part of a warm temperate land connection between Europe and North America (Hickey et al., 1983). The recent and surprising dis- covery of the aquatic fern Azolla in 47 Ma Eocene sediments in an ACEX Integrated Ocean Drilling Program (IODP) core in the middle of the Arctic Ocean (Brinkhuis et al., 2006) adds an almost surreal element to this vignette of crocodile-infested subtropical swamp forests on the shores of a warm, fresh arctic ocean covered with floating aquatic plants. Studies of such ice-free, high-latitude, deep-time analogues are important scientific windows into how the Arctic ecosystem might operate in the absence of permanent sea ice or in fully deglaciated conditions. In a world with forested poles and tropical midlatitudes, the nature of the equatorial realm is a serious question. Recent estimates of sea surface tem- peratures for the Late Cretaceous to Eocene tropics, based on well-preserved marine microfossils, suggest that temperatures may have exceeded 35-40ºC (Huber, 2002; Norris et al., 2002; Pearson et al., 2007)—the absence of equatorial coral reefs may have been because seawater was too hot. Ample evidence from equator to pole shows that the last greenhouse was a very dif- ferent place from today, and that the composition and distribution of biomes were wholly different from the present—it was not just a warmer world, but rather a completely different world from the present day.
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LESSONS FROM PAST WARM WORLDS FIGURE 2.15 Reconstruction showing the large latitudinal range of tropical rainforests during the Paleocene greenhouse. SOURCE: Morley (2000), courtesy of John Wiley & Sons, Inc. 61
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62 UNDERSTANDING EARTH’S DEEP PAST be at least two orders of magnitude higher than the long-term average (Hassan et al., 2005), a rate potentially commensurate with the largest mass extinctions of the geological past (Sepkoski, 1996; Bambach, 2006). Modeling future biodiversity losses and their effects on the Earth’s eco- systems and climate, however, is inherently difficult (Botkin et al., 2007), making it imperative to assess the outcome of equivalent “natural experi - ments” in the geological record (NRC, 1995; Myers and Knoll, 2001). The five major, and dozens of minor, mass extinctions of the past half-billion years (Sepkoski, 1996; Bambach, 2006) offer unique insights regarding eco- system susceptibility and response to environmental stress, the potential for ecological collapse, and the mechanisms of ecosystem recovery (Benton and Twitchett, 2003; Bottjer et al., 2008). Furthermore, the integration of paleontologic, stratigraphic, and geochemical records for many intervals of the past half-billion years have revealed the variable character of past biotic turnovers and mass extinction events (e.g., Boxes 2.4, 2.6, 2.7, 2.8), which differ in regard not only to severity but also to duration, selectivity, and the nature of environmental stresses (e.g., the transition out of super- greenhouse conditions into Ordovician glaciation [Trotter et al., 2008]; the Early to Middle Triassic radiations [Payne et al., 2004]; the nannoplankton crisis and foraminiferal turnovers of the Cretaceous ocean anoxic events [Leckie et al., 2002]; Eocene-Oligocene faunal extinction and immigration [Kobashi et al., 2001; Ivany et al., 2004]). Most importantly, the geologi - cal record uniquely captures past climate-ecological interactions that are fully played out and thereby archive the impact, response, interaction, and recovery from past global warming and major climate transitions.