Paleoclimate

Here we briefly review something of what is known about the evolution of climate over Earth history. Fig. 12.12 lists standard terminology for key periods of geologic time. Study of paleoclimate is an extremely exciting area of research, a fascinating detective story in which scientists study evidence of past climates recorded in ocean and lake sediments, glaciers and icesheets, and continental deposits. Proxies of past climates are myriad, and to the uninitiated at least, can be bizarre (packrat middens, midges...), including such measurements as the isotopic ratios of shells buried in ocean sediments, thickness and

FIGURE 12.12. The names and dates of the key periods of geologic time. The unit of time is millions of years before present (My), except during the Holocene, the last 10,000 years (10k y).

FIGURE 12.12. The names and dates of the key periods of geologic time. The unit of time is millions of years before present (My), except during the Holocene, the last 10,000 years (10k y).

density of tree rings, chemical composition of ice, and the radioactivity of corals. Moreover new proxies continue to be developed. What is undoubtedly clear is that climate has been in continual change over Earth history and has often been in states that are quite different from that of today. However, it is important to remember that inferences about paleoclimate are often based on sparse evidence,7 and detailed descriptions of past climates will never be available. Musing about paleoclimate is nevertheless intellectually stimulating (and great fun), because we can let our imaginations wander, speculating about ancient worlds and what they might tell us about how Earth might evolve in the future. Moreover, the historical record challenges and tests our understanding of the underlying mechanisms of climate and climate change. With such a short instrumental record, paleoclimate observations are essential for evaluating climate variations on timescales of decades and longer. One can be sure that the laws of physics and chemistry (if not biology!) have not changed over time, and so they place strong constraints on what may or may not have happened.

7One should qualify this statement by recognizing the key role of observation in paleoclimate. There are some "hard" paleo observations. For example the glacial terminus (moraines) of North America are incontrovertible evidence of the extent of past glaciations. As our colleague Prof. Ed Boyle reminds us, "Ice Ages would be polite tea-party chit-chat were it not for geologists climbing mountains in muddy boots.''

Theory and modeling of paleoclimate and climate change is still rudimentary. This is in part because we must deal not only with the physical aspects of the climate system (difficult enough in themselves) but also with bio-geochemical transformations, and on very long timescales, geology and geophysics (cf. Fig. 12.1). Understanding biogeochem-istry is particularly important because it is often required to appreciate and quantify the proxy climate record itself. Moreover, because greenhouse gases, such as H2O, CO2, and CH4, are involved in life, we are presented with a much more challenging problem than the mere application of Newton's laws of mechanics and the laws of thermodynamics to the Earth. There are many ideas on mechanisms driving climate change on paleoclimate timescales, only a few on which there is consensus, and even when a consensus forms, there is often little supporting evidence.

Here we have chosen to focus on those aspects of the paleoclimate record for which, it seems to us, there is a broad consensus and are less likely to be challenged as new evidence comes to the fore. In Section 12.3.1 we review what is known about the evolution of climate on the billion year timescale, and then in Section 12.3.2, focus in on the last 70M y or so. Warm and cold climates are discussed in Sections 12.3.3 and 12.3.4, respectively. We finish by briefly reviewing the evidence for glacial-interglacial cycles and abrupt climate change (Section 12.3.5) and, very briefly, global warming (Section 12.3.6).

12.3.1. Climate over Earth history

Earth has supported life of one form or another for billions of years, suggesting that its climate, although constantly changing, has remained within somewhat narrow limits over that time. For example, ancient rocks show markings that are clear evidence of erosion due to running water, and primitive life forms may go back at least 3.5B y. One might suppose that there is a natural ''thermostat'' that ensures that the Earth never gets too warm or too cold. One might also infer that life finds a way to eke out an existence.

It is clear that some kind of thermostat must be in operation because astrophysicists have concluded that 4B y ago the Sun was burning perhaps 25-30% less strongly than today. Simple one-dimensional climate models of the kind discussed in Chapter 2 suggest that if greenhouse gas concentrations in the distant past were at the same level as today, the Earth would have frozen over for the first two thirds or so of its existence.8 This is known as the faint early Sun paradox (see Problem 4 at the end of the chapter). A solution to the conundrum demands the operation of a thermostat, warming the Earth in the distant past and compensating for the increasing strength of the Sun over time. If the thermostat involved carbon, an assumption that perhaps needs to be critically challenged but is commonly supposed, then we must explain how CO2 levels in the atmosphere might have diminished over time.

On very long timescales one must consider the exchange of atmospheric CO2 with the underlying solid Earth in chemical weathering. Carbon is transferred from the Earth's interior to the atmosphere as CO2 gas produced during volcanic eruptions. This is balanced by removal of atmospheric CO2 in the chemical weathering of continental rocks, which ultimately deposits the carbon in sediments on the sea floor; see the schematic Fig. 12.13. It is remarkable that the rate of input of CO2 by volcanic activity and the rate of removal by chemical weathering has remained so closely in balance, even though the input and output themselves are each subject to considerable change.

Volcanic activity is unlikely to be part of a thermostat, because it is driven by heat

8Indeed there are hints in the paleoclimate record that Earth may have come close to freezing over during several periods of its history (most likely between 500M and 800M y ago), to form what has been called the ''snowball Earth.''

FIGURE 12.13. Carbon from the Earth's interior is injected in to the atmosphere as CO2 gas in volcanic eruptions. Removal of atmospheric CO2 on geological timescales is thought to occur in the chemical weathering of continental rocks, being ultimately washed into the ocean and buried in the sediments. The two processes must have been in close, but not exact balance, on geological time scales.

FIGURE 12.13. Carbon from the Earth's interior is injected in to the atmosphere as CO2 gas in volcanic eruptions. Removal of atmospheric CO2 on geological timescales is thought to occur in the chemical weathering of continental rocks, being ultimately washed into the ocean and buried in the sediments. The two processes must have been in close, but not exact balance, on geological time scales.

sources deep within the Earth that cannot react to climate change. Chemical weathering of rocks, on the other hand, may be sensitive to climate and atmospheric CO2 concentrations, because it is mediated by temperature, precipitation, vegetation, and orographic elevation and slope, which are closely tied together (remember the discussion in Section 1.3.2). So, the argument goes, if volcanic activity increased for a period of time, elevating CO2 levels in the atmosphere, the resulting warmer, moister climate might be expected to enhance chemical weathering, increasing the rate of CO2 removal and reducing greenhouse warming enough to keep climate roughly constant. Conversely, in a cold climate, arid conditions would reduce weathering rates, leading to a build-up of atmospheric CO2 and a warming tendency. Scientists vigorously debate whether such a mechanism can regulate atmospheric CO2; it is currently very difficult to test the idea with observation or models.

Whatever the regulatory mechanism, when the fragmentary paleoclimate record of atmospheric CO2 levels and temperatures is pieced together over geologic time, a connection emerges. Figure 12.14 shows a synthesis of evidence for continental glaciation plotted along with estimates of atmospheric CO2 (inferred from the geological record and geochemical models) over the past 600My. Such reconstructions are highly problematical and subject to great uncertainty. We see that CO2 levels in the atmosphere were thought to have been generally much greater in the distant past than at present, perhaps as much as 10 to 20 times present levels 400-500M y ago. Moreover, glaciation appears to occur during periods of low CO2 and warm periods in Earth history seem to be associated with elevated levels of CO2. That the temperature and CO2 concentrations appear to co-vary, however, should not be taken as implying cause or effect.

Factors other than variations in the solar constant and greenhouse gas forcing must surely also have been at work in driving the changes seen in Fig. 12.14. These include changes in the land-sea distribution and orography (driven by plate tectonics), the albedo of the underlying surface, and global biogeochemical cycles. One fascinating idea—known as the Gaia hypothesis—is that life itself plays a role in regulating the climate of the planet, optimizing the environment for continued evolution. Another idea is that ocean basins have evolved on geological timescales through continental drift, placing changing constraints on ocean circulation and its ability to transport heat meridionally. For example, Fig. 12.15 shows paleogeo-graphic reconstructions from the Jurassic (170My ago), the Cretaceous (100My ago), and the Eocene (50M y ago). We have seen in Chapters 10 and 11 how the circulation of the ocean is profoundly affected by the geometry of the land-sea distribution and so we can be sure that the pattern of ocean circulation in the past, and perhaps its role in climate, must have been very

FIGURE 12.14. (a) Comparison of CO2 concentrations from a geochemical model (continuous line) with a compilation (Berner, 1997) of proxy CO2 observations (horizontal bars). RCO2 is the ratio of past atmospheric CO2 concentrations to present day levels. Thus RCO2 = 10 means that concentrations were thought to be 10 times present levels. (b) CO2 radiative forcing effects expressed in Wm-2. (c) Combined CO2 and solar radiance forcing effects in Wm-2. (d) Glaciological evidence for continental-scale glaciation deduced from a compilation of many sources. Modified from Crowley (2000).

FIGURE 12.14. (a) Comparison of CO2 concentrations from a geochemical model (continuous line) with a compilation (Berner, 1997) of proxy CO2 observations (horizontal bars). RCO2 is the ratio of past atmospheric CO2 concentrations to present day levels. Thus RCO2 = 10 means that concentrations were thought to be 10 times present levels. (b) CO2 radiative forcing effects expressed in Wm-2. (c) Combined CO2 and solar radiance forcing effects in Wm-2. (d) Glaciological evidence for continental-scale glaciation deduced from a compilation of many sources. Modified from Crowley (2000).

different from that of today. It has been hypothesized that the opening and closing of critical oceanic gateways—narrow passages linking major ocean basins—have been drivers of climate variability by regulating the amount of water, heat, and salt exchanged between ocean basins. This, for example, can alter meridional transport of heat by the ocean and hence play a role in glaciation and deglaciation. There are a number of important gateways.

Drake Passage, separating South America from Antarctica, opened up 25-20My ago, leaving Antarctica isolated by what we now call the Antarctic Circumpolar Current (Fig. 9.13). This may have made it more difficult for the ocean to deliver heat to the south pole, helping Antarctica to freeze over. However this hypothesis has timing problems. Ice first appeared on Antarctica 35M y ago, before the opening of Drake Passage, and the most intense glaciation over Antarctica occurred 13My ago, significantly after it opened. Uplift of Central America over the past 10M y closed a deep ocean passage between North and South America to form the Isthmus of Panama about 4 million years ago. Before then the Isthmus was open, allowing the trade winds to blow warm and possibly salty water between the Atlantic and the Pacific. Its closing could have supported a Gulf Stream carrying tropical waters polewards, as in today's climate, possibly enhancing the meridional overturning circulation of the Atlantic basin and helping to warm northern latitudes in the Atlantic sector, as discussed in Chapter 11. Finally, it has been suggested that closing of the Indonesian seaway, 3-4My ago, was a precursor to East African aridification.

12.3.2. Paleotemperatures over the past 70 million years: the á18O record

Let us zoom into the last 70My period of Fig. 12.14. The paleorecord suggests that

FIGURE 12.15. Paleogeographic reconstructions for (top) the Jurassic (170My ago), (middle) the Cretaceous (100M y ago), and (bottom) the Eocene (50My ago). Panthalassa was the huge ocean that in the paleo world dominated one hemisphere. Pangea was the supercontinent in the other hemisphere. The Tethys Sea was the body of water enclosed on three sides (and at times, almost four sides) by the generally ''C-shaped'' Pangea.

FIGURE 12.15. Paleogeographic reconstructions for (top) the Jurassic (170My ago), (middle) the Cretaceous (100M y ago), and (bottom) the Eocene (50My ago). Panthalassa was the huge ocean that in the paleo world dominated one hemisphere. Pangea was the supercontinent in the other hemisphere. The Tethys Sea was the body of water enclosed on three sides (and at times, almost four sides) by the generally ''C-shaped'' Pangea.

over the last 55M y there has been a broad progression from generally warmer to generally colder conditions, with significant shorter-term oscillations superimposed. How can one figure this out? Some key supporting evidence is shown in Fig. 12.16 based on isotopic measurements of oxygen. Sediments at the bottom of the ocean provide a proxy record of climate conditions in the water column. One key proxy is 518O—a measure of the ratio of two isotopes of oxygen, 18O and 16O—which is recorded in sea-bed sediments by the fossilized calcite shells of foraminifera (organisms that live near the surface or the bottom of the ocean). It turns out that the 518O in the shells is a function of the 518O of the ocean and the temperature of the ocean (see Appendix A.3 for a more detailed discussion of 518O). The record of 518O over the last 55M y indicates a cooling of the deep ocean by a massive 14°C. In other words, deep ocean temperatures were perhaps close to 16°C (!!) compared to 2°C as observed today (cf. Fig. 9.5). If over this period of time the abyssal ocean were ventilated by convection from the poles as in today's climate (note how temperature surfaces in the deep ocean thread back to the pole in Fig. 9.5), then one can conclude that surface conditions at the poles must also have been very much warmer. Indeed, this is consistent with other sources of evidence, such as the presence of fossilized remains of palm trees and the ancestors of modern crocodiles north of the Arctic Circle 60My ago.

To explain such a large cooling trend, sustained over many millions of years, one needs to invoke a mechanism that persists over this enormous span of time. Following on from the discussion in Section 12.3.1, at least two important ideas have been put forward as a possible cause. Firstly, it has been suggested that the balance in Fig. 12.13 might have changed to reduce CO2 forcing of atmospheric temperatures over this period due to (a) decreased input of CO2 from the Earth's interior to the

FIGURE 12.16. A compilation of <5180 measurements from the fossilized shells of benthic foraminifera analyzed from many sediment cores in the North Atlantic over 70M y. Modified from Miller et al (1987).

climate system, as the rate of sea-floor spreading decreases over time, reducing volcanic activity, and (b) increased removal of CO2 from the atmosphere due to enhanced physical and chemical weathering of unusually high-elevation terrain driven by tectonic uplift. Secondly, it has been suggested that poleward ocean heat transport progressively decreased because of changes in the distribution of land and sea, and the opening up and closing of gateways, as briefly discussed in Section 12.3.1.

Whatever the mechanisms at work, as shown in Fig. 12.14, the paleorecord suggests that, over the past 100My or so, the Earth has experienced great warmth and periods of great cold. We now briefly review what ''warm'' climates and ''cold'' climates might have been like.

12.3.3. Greenhouse climates

In the Cretaceous period Earth was a ''greenhouse world.'' There were no ice caps and sea level was up to 100-200 m higher than present, largely due to the melting of all ice caps and the thermal expansion of the oceans that were much warmer than today. The great super continent of Pangea had begun to break apart, and by 100My ago one can already recognize present-day continents (Fig. 12.15). High sea level meant that much of the continental areas were flooded and there were many inland lakes and seas. Indeed, the meaning of the word Cretaceous is ''abundance of chalk,'' reflecting the widespread occurrence of limestone from creatures living in the many inland seas and lakes of the period. Broad-leaved plants, dinosaurs, turtles, and crocodiles all existed north of the Arctic Circle.

It is thought that the Cretaceous was a period of elevated CO2 levels—perhaps as much as five times preindustrial concentrations (see Fig. 12.14)—accounting in part for its great warmth. CO2 forcing alone is unlikely to account for such warm poles where temperatures were perhaps 25°C warmer than today. One proposed explanation is that the oceans carried much more heat poleward than today, rendering the poles warmer and the tropics colder. There is speculation that the deep ocean may have been much warmer and saltier than at present, possibly due to convection in the tropics and/or subtropics triggered by high values of salinity, much as observed today in the Eastern Mediterranean. Indeed, the configuration of the continents may have been conducive to such a process; the presence of the Tethys Sea (see Fig. 12.15 middle) and a large tropical seaway extending up in to subtropical latitudes, underneath the sinking branch of the Hadley cell bringing dry air down to the surface, could have increased evaporation and hence salinity to the extent that ocean convection was triggered, mixing warm, salty water to depth. But this is just speculation. It is very difficult to plausibly quantify and model this process and efforts to do so often meet with failure.

Another major challenge in understanding the paleorecord in the Cretaceous is the evidence that palm trees and reptiles were present in the interior of the continents. Crocodiles and (young) palm trees are not frost resistant, indicating that temperatures did not go below freezing even during the peak of winter at some latitudes north of 60° N and in the middle of continents, away from the moderating effects of the ocean. Models, however, simulate freezing conditions in the continental interiors in winter even when CO2 is increased very dramatically. Note the large seasonal change in temperature in the interior of continents observed in the present climate (Fig. 12.2). Perhaps lakes and small inland seas helped to keep the interior warm.

12.3.4. Cold climates

Most of the time in the last 1My, Earth has been much colder than at present, and ice has encroached much further equator-ward (Fig. 12.14). The glacial climate that we know the most about is that at the height of the most recent glacial cycle—the last glacial maximum (LGM) between 18 and 23k y ago, during which ice sheets reached their greatest extent 21k y ago. Reconstruction of climate at the LGM was carried out in the CLIMAP project9 employing, in the main, proxy data from ocean sediments. Thick ice covered Canada, the northern United States (as far south as the Great Lakes), northern Europe (including all of Scandinavia, the northern half of the British Isles, and Wales) and parts of Eurasia. The effect on surface elevation is shown in Fig. 12.17, which should be compared to modern conditions shown in Fig. 9.1. Where Chicago, Glasgow, and Stockholm now stand, ice was over 1 km thick. It is thought that the Laurentide Ice sheet covering N. America had roughly the volume of ice locked up in present-day Antarctica. Sea level was about 120-130 m lower than today. Note that the coastline of the LGM shown in Fig. 12.17 reveals that, for example, the British Isles were connected to Europe, and many islands that exist today were joined to Asia and Australia. Most of the population lived in these fertile lowlands, many of which are now under water. Ice sheets on Antarctica and Greenland extended across land exposed by the fall of sea level. Moreover, sea ice was also considerably more extensive, covering much of the Greenland and Norwegian Seas, and persisted through the summer. In the southern hemisphere, Argentina, Chile, and New Zealand were under ice, as were parts of Australia and South America.

Figure 12.17b shows the difference between average August SST centered on the LGM and August SST for the modern era. Many details of this reconstruction have been challenged, but the broad features are probably correct. The average SST was 4°C colder than present and North Atlantic SSTs were perhaps colder by more than 8°C. It appears that low latitude temperatures were perhaps 2°C lower than today. Winds at the LGM were drier, stronger, and dustier than in the present climate. Ice sheets, by grinding away the underlying bedrock, are very efficient producers of debris of all sizes, which gets pushed out to the ice margin. At the LGM, windy, cold, arid conditions existed equatorward of the ice. Winds scooped up the finer-grained debris, resulting in great dust storms blowing across the Earth's surface with more exposed shelf areas. Indeed, glacial layers in ice cores drilled in both Greenland and Antarctica carry more dust than interglacial layers. Forests shrank and deserts expanded. Today the N. African and Arabian deserts are key sources of dust; at the LGM deserts expanded into Asia. One very significant feature of glacial climates evident in the paleorecord is that they exhibited considerably more variability than warm climates. For example, in an event known as the Younger Dryas, which occurred about 12k y ago, the climate warmed only to suddenly return to close to LGM conditions for several hundred years; see Fig. 12.23 and the discussion in Section 12.3.5.

Key factors that may explain the dramatically different climate of the LGM are the presence of the ice sheets themselves, with their high albedo reflecting solar radiation back out to space, and (see below) lower levels of greenhouse gases. It is thought that the pronounced climate variability of glacial periods suggested by the paleorecord may have been associated with melting ice producing large inland lakes that were perhaps cut off from the oceans for hundreds of years, but which then intermittently and perhaps suddenly discharged into the oceans. It has been argued that such sudden discharges of buoyant fluid over the surface of the northern N. Atlantic could have had a significant impact on

9CLIMAP (Climate: Long-range Investigation, Mapping and Prediction), was a major research project of the 1970s and 1980s, which resulted in a map of climate conditions during the last glacial maximum based on proxy data from ocean sediments.

Glacial Maximum Pennsylvania

FIGURE 12.17. (a) CLIMAP reconstruction of elevation at the Last Glacial Maximum (LGM). The white (black) areas represent terrain with a height in excess of (less than) 1.5 km and are indicative of ice-covered areas. The depth of the ocean is represented with a grey scale (dark is deep). The white contour marks the 4 km deep isobath. This figure should be compared with Fig. 9.1. Note the modification of the coast line relative to the modern, due to the 120 m or so drop in sea level. (b) August SST at LGM (from CLIMAP) minus August SST for the modern climate (°C). The brown areas represent negative values, the green areas positive values.

FIGURE 12.17. (a) CLIMAP reconstruction of elevation at the Last Glacial Maximum (LGM). The white (black) areas represent terrain with a height in excess of (less than) 1.5 km and are indicative of ice-covered areas. The depth of the ocean is represented with a grey scale (dark is deep). The white contour marks the 4 km deep isobath. This figure should be compared with Fig. 9.1. Note the modification of the coast line relative to the modern, due to the 120 m or so drop in sea level. (b) August SST at LGM (from CLIMAP) minus August SST for the modern climate (°C). The brown areas represent negative values, the green areas positive values.

the strength of the ocean's meridional overturning circulation and its ability to transport heat polewards.

12.3.5. Glacial-interglacial cycles

The left frame of Fig. 12.18 shows the 518O record over the past 2.5 million years, recorded in the calcite of foraminifera in sediments of the subpolar North Atlantic. Before about 800k y ago, one observes remarkable oscillations spanning 2M y or so, with a period of about 40k y. After 800k y ago the nature of the record changes and fluctuations with longer periods are superimposed. These are the signals of great glacial-interglacial shifts on a roughly 100k y timescale. There have been about 7 such

FIGURE 12.18. Left: <518 O over the last 2.5 million years recorded in the calcite shells of bottom dwelling foraminifera in the subpolar North Atlantic. Shown is the average of tens of <518 O records sampled from various marine sediment cores (Huybers, 2006). Values are reported as the anomaly from the average <518 O over the past million years. More negative values (rightward) indicate warmer temperatures and less ice volume. Right: <518 O of ice over the last 50 k y measured in the GISP2 ice-core (Grootes and Stuiver, 1997). In contrast to the <518 O of marine shells, less negative values in the <518 O of ice indicate warmer atmospheric temperatures, in this case in the vicinity of Greenland.

FIGURE 12.18. Left: <518 O over the last 2.5 million years recorded in the calcite shells of bottom dwelling foraminifera in the subpolar North Atlantic. Shown is the average of tens of <518 O records sampled from various marine sediment cores (Huybers, 2006). Values are reported as the anomaly from the average <518 O over the past million years. More negative values (rightward) indicate warmer temperatures and less ice volume. Right: <518 O of ice over the last 50 k y measured in the GISP2 ice-core (Grootes and Stuiver, 1997). In contrast to the <518 O of marine shells, less negative values in the <518 O of ice indicate warmer atmospheric temperatures, in this case in the vicinity of Greenland.

cycles, during which temperate forests in Europe and North America have repeatedly given way to tundra and ice. Ice has periodically accumulated in the North American and Scandinavian areas until it covered hills and mountains to heights of 2-3 km, as was last observed at the LGM (see Fig. 12.17) and today only in Greenland and Antarctica.

Such glacial-interglacial signals are not limited to the North Atlantic sector. Qualitatively similar signals are evident in different kinds of paleorecords taken from around the world, including deep sea sediments, continental deposits of plants, and ice cores. These reveal a marked range of climate on Earth, cycling between glacial and interglacial conditions.

In particular, ice cores taken from glaciers yield local air temperature,10 precipitation rate, dust, and direct records of past trace gas concentrations of CO2 and CH4. The deepest core yet drilled (> 3 km), from Antarctica, records a remarkable 700ky history of climate variability shown in Fig. 12.19. The core reveals oscillations of Antarctic air temperature, greenhouse gas concentrations which more-or-less covary with a period of about 100k y. Note, however, that the oscillations do not have exactly the same period. Of the six or seven cycles seen in the Antarctic record, the two most recent have a somewhat longer period than the previous cycles.

The 100ky signals evident in Figs. 12.18 and 12.19 are thought to be representative of climate variability over broad geographical regions. Scientists vigorously debate whether, for example, changes over Antarctica led or lagged those over Greenland, or whether CO2 changes led or lagged temperature changes. This is very difficult to tie down because of uncertainty in the precise setting of the "clock" within and between records. Here we simply state that at zero order the low frequency signals seem to covary over broad areas of the globe, strongly suggestive of global-scale change.

The oscillations seen in Fig. 12.19 have a characteristic "saw-tooth" pattern, typical of many records spanning glacial-interglacial cycles, with a long period of cooling into the glacial state followed by rapid warming to the following interglacial. Abrupt increases in CO2 occur during the period of rapid ice melting. Superimposed on the sawtooth are irregular higher frequency oscillations (to be discussed below). Typically, the coolest part of each glacial period and the lowest CO2 concentrations occur just before the glacial termination. Temperature fluctuations (representative of surface conditions) have a magnitude of about 12°C and CO2 levels fluctuate between 180 and 300 ppm. The Antarctic dust record also confirms continental aridity. Dust transport was more prevalent during glacial than interglacial times, as mentioned in Section 12.3.4. Finally, it is worthy of note that present levels of CO2 (around 370 ppm in the year 2000; cf. Fig. 1.3) are unprecedented during the past 700kys. By the end of this century levels will almost certainly have reached 600 ppm.

Milankovitch cycles

It seems that climate on timescales of 10ky-100kys is strongly influenced by variations in Earth's position and orientation relative to the Sun. Indeed, as we shall see, some of the expected periods are visible in the paleorecord, but direct association (phasing and amplitude) is much more problematical. Variation in the Earth's orbit over time—known as Milankovitch cycles11—cause changes in the amount and distribution of solar radiation

10Note that 18O/16O ratios in ice cores have the opposite relationship to temperature than that of 18O/16O ratios in CaCO3 shells (see Appendix A.3). Snow produced in colder air tends to have a lower <518O value than snow produced in warmer air. Consequently, the <518O value of glacial ice can be used as a proxy for air temperature, with low values indicating colder temperatures than higher values (see Fig. 12.18).

Milutin Milankovitch (1879—1958), the Serbian mathematician, dedicated his career to formulating a mathematical theory of climate based on the seasonal and latitudinal variations of solar radiation received by the Earth. In the 1920s he developed improved methods of calculating variations in Earth's eccentricity, precession, and tilt through time.

FIGURE 12.19. Ice-core records of atmospheric carbon dioxide (left) and methane (middle) concentrations obtained from bubbles trapped in Antarctic ice. Values to 400 k y ago are from Vostok (Petit et al, 1999), whereas earlier values are from EPICA Dome C (Siegenthaler et al, 2005; Spahni et al, 2005). (right) S D concentrations from EPICA Dome C (EPICA community members, 2004) measured in the ice, as opposed to the bubbles, are indicative of local air temperature variations, similar to S18 O of ice measurements. A rightward shift corresponds to warming.

200 220 240 260 280 400 500 600 700 -440 -420 -400 -380

C02 (parts per million) CH4 (parts per billion) § d (parts per thousand)

FIGURE 12.19. Ice-core records of atmospheric carbon dioxide (left) and methane (middle) concentrations obtained from bubbles trapped in Antarctic ice. Values to 400 k y ago are from Vostok (Petit et al, 1999), whereas earlier values are from EPICA Dome C (Siegenthaler et al, 2005; Spahni et al, 2005). (right) S D concentrations from EPICA Dome C (EPICA community members, 2004) measured in the ice, as opposed to the bubbles, are indicative of local air temperature variations, similar to S18 O of ice measurements. A rightward shift corresponds to warming.

reaching the Earth on orbital timescales. Before discussing variations of the Earth's orbit over time, let us return to ideas introduced in Chapter 5 and review some simple facts about Earth's orbit around the Sun and the cause of the seasons.

Imagine for a moment that the Earth travelled around the Sun in a circular orbit, as in Fig. 12.20a (left). If the Earth's spin axis were perpendicular to the orbital plane (i.e., did not tilt), we would experience no seasons and the length of daytime and nighttime would never change throughout the year and be equal to one another. But

(b) Obliquity

(b) Obliquity

(c) Precession

FIGURE 12.20. (a) The eccentricity of the Earth's orbit varies on 100k y & 400k y timescales from (almost) zero, a circle, to 0.07, a very slight ellipse. The ellipse shown on the right has an eccentricity of 0.5, vastly greater than that of Earth's path around the Sun. (b) The change in the tilt of the Earth's spin axis—the obliquity—varies between 22.1°and 24.5° on a timescale of 41k y. The tilt of the Earth is currently 23.5°. (c) The direction of the Earth's spin vector precesses with a period of 23k y.

now suppose that the spin axis is tilted as a constant angle, as sketched in Fig. 5.3, and, moreover, that the direction of tilt in space is constant relative to the fixed stars. Now, as discussed in Section 5.1.1, we would experience seasons and the length of daytime would vary throughout the year. When the northern hemisphere (NH) is tilted toward the Sun, the Sun rises high in the sky, daytime is long, and the NH receives intense radiation and experiences summer conditions. When the NH tilts away from the Sun, the Sun stays low in the sky, the daytime is short and the NH receives diminished levels of radiation and experiences winter. These seasonal differences culminate at the summer and winter solstices. In modern times, the longest day of the year occurs on June 21st (the summer solstice) and shortest day of the year on December 21st (the winter solstice) (see Fig. 5.4). The length of the day and night become equal at the equinoxes. Thus we see that seasonality and length of day variations are fundamentally controlled by the tilt of the Earth's axis away from the orbital plane. This tilt of the Earth's axis away from the orbital plane is known as the obliquity (see Fig. 12.20b). It varies between 21.1° and 24.5° on about 41ky timescales; at the present time it is 23.5°. Obliquity affects the annual insolation in both hemispheres simultaneously. When the tilt is large, seasonality at high latitudes becomes more extreme but with little effect at the equator.

The Earth's orbit is not exactly circular, however. As shown in Figs. 5.4 and 12.20, Earth moves around the Sun following an elliptical path; the distance from the Sun varies between 153 million km at perihelion (closest distance of the Earth to the Sun) and 158 million km at aphelion (farthest distance between the Earth and Sun). As can be seen in Fig. 5.4, in modern times the Earth is slightly closer to the Sun at the NH winter solstice. Winter radiation is slightly higher than it would be if the Earth followed a perfectly circular orbit. Conversely, at the NH summer solstice the Earth is slightly farther

FIGURE 12.20. (a) The eccentricity of the Earth's orbit varies on 100k y & 400k y timescales from (almost) zero, a circle, to 0.07, a very slight ellipse. The ellipse shown on the right has an eccentricity of 0.5, vastly greater than that of Earth's path around the Sun. (b) The change in the tilt of the Earth's spin axis—the obliquity—varies between 22.1°and 24.5° on a timescale of 41k y. The tilt of the Earth is currently 23.5°. (c) The direction of the Earth's spin vector precesses with a period of 23k y.

away from the Sun, and so NH summer radiation is slightly lower than it would be if the Earth followed a perfectly circular orbit. This is a rather small effect, however, because the Earth-Sun distance only varies by 3% of the mean. Nevertheless the eccentricity of the Earth's orbit around the Sun (see Fig. 12.20a), a measure of its degree of circularity, enhances or reduces the seasonal variation of the intensity of radiation received by the Earth. The eccentricity varies with periods of about 100ky and 400ky. It modulates seasonal differences and precession, the third important orbital parameter.

Precession measures the direction of the Earth's axis of rotation, which affects the magnitude of the seasonal cycle and is of opposite phase in the two hemispheres. Earth's spin axis precesses at a period of 27ky with respect to the fixed stars. However, this is not the climatically relevant period because the direction of the major axis of Earth's eccentric orbit also moves. Thus climatologists define the climatic precession as the direction of Earth's spin axis with respect to Earth's eccentric orbit. This has a period of about 23k y. Today the rotation axis points toward the North Star, so setting the dates during the year at which the Earth reaches aphelion and perihelion on its orbit around the Sun (see Fig. 5.4). At the present time, perihelion falls on January 3rd, only a few weeks after the winter solstice, and so the northern hemisphere winter and southern hemisphere summer are slightly warmer than the corresponding seasons in the opposite hemispheres.

We discussed in Chapters 5 and 8 those factors that control the annual-mean temperature as a function of latitude and in particular the importance of the latitudinal dependence of incoming solar radiation. This latitudinal dependence is critically modulated by orbital parameters. Because of their different periodicity (see Fig. 12.21), the composite variations in solar radiation are very complex. They are functions of both latitude and season, as well as time.

Eccentricity Precession

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