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Palaeoclimate
Coordinating Lead Authors:
Eystein Jansen (Norway), Jonathan Overpeck (USA)
Lead Authors:
Keith R. Briffa (UK), Jean-Claude Duplessy (France), Fortunat Joos (Switzerland), Valérie Masson-Delmotte (France), Daniel Olago (Kenya), Bette Otto-Bliesner (USA), W. Richard Peltier (Canada), Stefan Rahmstorf (Germany), Rengaswamy Ramesh (India), Dominique Raynaud (France), David Rind (USA), Olga Solomina (Russian Federation), Ricardo Villalba (Argentina), De’er Zhang (China)
Contributing Authors:
J.-M. Barnola (France), E. Bauer (Germany), E. Brady (USA), M. Chandler (USA), J. Cole (USA), E. Cook (USA), E. Cortijo (France), T. Dokken (Norway), D. Fleitmann (Switzerland, Germany), M. Kageyama (France), M. Khodri (France), L. Labeyrie (France), A. Laine (France), A. Levermann (Germany), Ø. Lie (Norway), M.-F. Loutre (Belgium), K. Matsumoto (USA), E. Monnin (Switzerland), E. Mosley-Thompson (USA), D. Muhs (USA), R. Muscheler (USA), T. Osborn (UK), Ø. Paasche (Norway), F. Parrenin (France), G.-K. Plattner (Switzerland), H. Pollack (USA), R. Spahni (Switzerland), L.D. Stott (USA), L. Thompson (USA), C. Waelbroeck (France), G. Wiles (USA), J. Zachos (USA), G. Zhengteng (China)
Review Editors:
Jean Jouzel (France), John Mitchell (UK)
This chapter should be cited as:
Jansen, E., J. Overpeck, K.R. Briffa, J.-C. Duplessy, F. Joos, V. Masson-Delmotte, D. Olago, B. Otto-Bliesner, W.R. Peltier, S. Rahmstorf, R. Ramesh, D. Raynaud, D. Rind, O. Solomina, R. Villalba and D. Zhang, 2007: Palaeoclimate. In: Climate Change 2007: The Physical Science Basis. Contribution of Working Group I to the Fourth Assessment Report of the Intergovernmental Panel on Climate Change [Solomon, S., D. Qin, M. Manning, Z. Chen, M. Marquis, K.B. Averyt, M. Tignor and H.L. Miller (eds.)]. Cambridge University Press, Cambridge, United Kingdom and New York, NY, USA.
Palaeoclimate Chapter 6
Table of Contents
Executive Summary .................................................... 435
6.1 Introduction ......................................................... 438
6.2 Palaeoclimatic Methods ................................... 438
6.2.1 Methods – Observations of Forcing and Response...................................................... 438 6.2.2 Methods – Palaeoclimate Modelling ................... 439
6.3 The Pre-Quaternary Climates ...................... 440
6.3.1 What is the Relationship Between Carbon Dioxide and Temperature in this Time Period?.... 440 6.3.2 What Does the Record of the Mid-Pliocene Show? ........................................... 440 6.3.3 What Does the Record of the Palaeocene-Eocene Thermal Maximum Show? ................................... 442
6.4 Glacial-Interglacial Variability
and Dynamics ...................................................... 444
6.4.1 Climate Forcings and Responses Over Glacial-Interglacial Cycles ................................... 444 Box 6.1: Orbital Forcing ................................................... 445 Box 6.2: What Caused the Low Atmospheric Carbon Dioxide Concentrations During Glacial Times? ................................................... 446 6.4.2 Abrupt Climatic Changes in the Glacial-Interglacial Record .................................. 454 6.4.3 Sea Level Variations Over the Last Glacial-Interglacial Cycle ..................................... 457
6.5 The Current Interglacial ................................. 459
6.5.1 Climate Forcing and Response During the Current Interglacial ........................................ 459 Box 6.3: Holocene Glacier Variability ................................. 461 6.5.2 Abrupt Climate Change During the Current Interglacial .............................................. 463 6.5.3 How and Why Has the El Niño-Southern Oscillation Changed Over the Present Interglacial? ............................................ 464
6.6 The Last 2,000 Years .......................................... 466
6.6.1 Northern Hemisphere Temperature Variability ..... 466 Box 6.4: Hemispheric Temperatures in the ‘Medieval Warm Period’ ..................................... 468 6.6.2 Southern Hemisphere Temperature Variability ............................................................. 474 6.6.3 Comparisons of Millennial Simulations with Palaeodata ................................................... 476 6.6.4 Consistency Between Temperature, Greenhouse Gas and Forcing Records; and Compatibility of Coupled Carbon Cycle-Climate Models with the Proxy Records ............................................... 481 6.6.5 Regional Variability in Quantities Other than Temperature................................................. 481
6.7 Concluding Remarks on
Key Uncertainties ............................................... 483
Frequently Asked Questions
FAQ 6.1: What Caused the Ice Ages and Other Important Climate Changes Before the Industrial Era? ............ 449
FAQ 6.2: Is the Current Climate Change Unusual Compared to Earlier Changes in Earth’s History? ................... 465
References ........................................................................ 484
Supplementary Material
The following supplementary material is available on CD-ROM and in on-line versions of this report. Appendix 6.A: Glossary of Terms Specifi c to Chapter 6
Palaeoclimate Chapter 6
that any of these regional signals were coherent at the global scale, or are capable of explaining the majority of global warming of the last 100 years.
- Glaciers in several mountain regions of the Northern Hemisphere retreated in response to orbitally forced regional warmth between 11 and 5 ka, and were smaller (or even absent) at times prior to 5 ka than at the end of the 20th century. The present day near-global retreat of mountain glaciers cannot be attributed to the same natural causes, because the decrease of summer insolation during the past few millennia in the Northern Hemisphere should be favourable to the growth of the glaciers.
- For the mid-Holocene (about 6 ka), GCMs are able to simulate many of the robust qualitative large-scale features of observed climate change, including mid- latitude warming with little change in global mean temperature (<0.4°C), as well as altered monsoons, consistent with the understanding of orbital forcing. For the few well-documented areas, models tend to underestimate hydrological change. Coupled climate models perform generally better than atmosphere-only models, and reveal the amplifying roles of ocean and land surface feedbacks in climate change.
- Climate and vegetation models simulate past northward shifts of the boreal treeline under warming conditions. Palaeoclimatic results also indicated that these treeline shifts likely result in significant positive climate feedback. Such models are also capable of simulating changes in the vegetation structure and terrestrial carbon storage in association with large changes in climate boundary conditions and forcings (i.e., ice sheets, orbital variations).
- Palaeoclimatic observations indicate that abrupt decadal- to centennial-scale changes in the regional frequency of tropical cyclones, floods, decadal droughts and the intensity of the African-Asian summer monsoon very likely occurred during the past 10 kyr. However, the mechanisms behind these abrupt shifts are not well understood, nor have they been thoroughly investigated using current climate models.
How does the 20th-century climate change compare with the climate of the past 2,000 years?
- It is very likely that the average rates of increase in CO 2 , as well as in the combined radiative forcing from CO 2 , CH (^4) and N 2 O concentration increases, have been at least five times faster over the period from 1960 to 1999 than over any other 40-year period during the past two millennia prior to the industrial era. - Ice core data from Greenland and Northern Hemisphere mid-latitudes show a very likely rapid post-industrial era increase in sulphate concentrations above the pre- industrial background. - Some of the studies conducted since the Third Assessment Report (TAR) indicate greater multi-centennial Northern Hemisphere temperature variability over the last 1 kyr than was shown in the TAR, demonstrating a sensitivity to the particular proxies used, and the specific statistical methods of processing and/or scaling them to represent past temperatures. The additional variability shown in some new studies implies mainly cooler temperatures (predominantly in the 12th to 14th, 17th and 19th centuries), and only one new reconstruction suggests slightly warmer conditions (in the 11th century, but well within the uncertainty range indicated in the TAR). - The TAR pointed to the ‘exceptional warmth of the late 20th century, relative to the past 1,000 years’. Subsequent evidence has strengthened this conclusion. It is very likely that average Northern Hemisphere temperatures during the second half of the 20th century were higher than for any other 50-year period in the last 500 years. It is also likely that this 50-year period was the warmest Northern Hemisphere period in the last 1.3 kyr, and that this warmth was more widespread than during any other 50- year period in the last 1.3 kyr. These conclusions are most robust for summer in extratropical land areas, and for more recent periods because of poor early data coverage. - The small variations in pre-industrial CO 2 and CH concentrations over the past millennium are consistent with millennial-length proxy Northern Hemisphere temperature reconstructions; climate variations larger than indicated by the reconstructions would likely yield larger concentration changes. The small pre-industrial greenhouse gas variations also provide indirect evidence for a limited range of decadal- to centennial-scale variations in global temperature. - Palaeoclimate model simulations are broadly consistent with the reconstructed NH temperatures over the past 1 kyr. The rise in surface temperatures since 1950 very likely cannot be reproduced without including anthropogenic greenhouse gases in the model forcings, and it is very unlikely that this warming was merely a recovery from a pre-20th century cold period. - Knowledge of climate variability over the last 1 kyr in the Southern Hemisphere and tropics is very limited by the low density of palaeoclimatic records.
Chapter 6 Palaeoclimate
- Climate reconstructions over the past millennium indicate with high confi dence more varied spatial climate teleconnections related to the El Niño-Southern Oscillation than are represented in the instrumental record of the 20th century.
- The palaeoclimate records of northern and eastern Africa, as well as the Americas, indicate with high confi dence that droughts lasting decades or longer were a recurrent feature of climate in these regions over the last 2 kyr.
What does the palaeoclimatic record reveal about feedback, biogeochemical and biogeophysical processes?
- The widely accepted orbital theory suggests that glacial- interglacial cycles occurred in response to orbital forcing. The large response of the climate system implies a strong positive amplification of this forcing. This amplification has very likely been influenced mainly by changes in greenhouse gas concentrations and ice sheet growth and decay, but also by ocean circulation and sea ice changes, biophysical feedbacks and aerosol (dust) loading.
- It is virtually certain that millennial-scale changes in atmospheric CO 2 associated with individual antarctic warm events were less than 25 ppm during the last glacial period. This suggests that the associated changes in North Atlantic Deep Water formation and in the large-scale deposition of wind-borne iron in the Southern Ocean had limited impact on CO 2.
- It is very likely that marine carbon cycle processes were primarily responsible for the glacial-interglacial CO 2 variations. The quantification of individual marine processes remains a difficult problem.
- Palaeoenvironmental data indicate that regional vegetation composition and structure are very likely sensitive to climate change, and in some cases can respond to climate change within decades.
Chapter 6 Palaeoclimate
reduce age uncertainty, and palaeoclimatic interpretations must take into account uncertainties in time control. There continue to be significant advances in radiometric dating. Each radiometric system has ranges over which the system is useful, and palaeoclimatic studies almost always publish analytical uncertainties. Because there can be additional uncertainties, methods have been developed for checking assumptions and cross verifying with independent methods. For example, secular variations in the radiocarbon clock over the last 12 kyr are well known, and fairly well understood over the last 35 kyr. These variations, and the quality of the radiocarbon clock, have both been well demonstrated via comparisons with age models derived from precise tree ring and varved sediment records, as well as with age determinations derived from independent radiometric systems such as uranium series. However, for each proxy record, the quality of the radiocarbon chronology also depends on the density of dates, the material available for dating and knowledge about the radiocarbon age of the carbon that was incorporated into the dated material.
6.2.1.4 How Can Palaeoclimatic Proxy Methods Be Used to Reconstruct Past Climate Dynamics?
Most of the methods behind the palaeoclimatic reconstructions assessed in this chapter are described in some detail in the aforementioned books, as well as in the citations of each chapter section. In some sections, important methodological background and controversies are discussed where such discussions help assess palaeoclimatic uncertainties. Palaeoclimatic reconstruction methods have matured greatly in the past decades, and range from direct measurements of past change (e.g., ground temperature variations, gas content in ice core air bubbles, ocean sediment pore-water change and glacier extent changes) to proxy measurements involving the change in chemical, physical and biological parameters that reflect
- often in a quantitative and well-understood manner – past change in the environment where the proxy carrier grew or existed. In addition to these methods, palaeoclimatologists also use documentary data (e.g., in the form of specific observations, logs and crop harvest data) for reconstructions of past climates. While a number of uncertainties remain, it is now well accepted and verified that many organisms (e.g., trees, corals, plankton, insects and other organisms) alter their growth and/or population dynamics in response to changing climate, and that these climate-induced changes are well recorded in the past growth of living and dead (fossil) specimens or assemblages of organisms. Tree rings, ocean and lake plankton and pollen are some of the best-known and best-developed proxy sources of past climate going back centuries and millennia. Networks of tree ring width and density chronologies are used to infer past temperature and moisture changes based on comprehensive calibration with temporally overlapping instrumental data. Past distributions of pollen and plankton from sediment cores can be used to derive quantitative estimates of past climate (e.g., temperatures, salinity and precipitation) via statistical methods calibrated against their modern distribution and associated climate
parameters. The chemistry of several biological and physical entities reflects well-understood thermodynamic processes that can be transformed into estimates of climate parameters such as temperature. Key examples include: oxygen (O) isotope ratios in coral and foraminiferal carbonate to infer past temperature and salinity; magnesium/calcium (Mg/Ca) and strontium/ calcium (Sr/Ca) ratios in carbonate for temperature estimates; alkenone saturation indices from marine organic molecules to infer past sea surface temperature (SST); and O and hydrogen isotopes and combined nitrogen and argon isotope studies in ice cores to infer temperature and atmospheric transport. Lastly, many physical systems (e.g., sediments and aeolian deposits) change in predictable ways that can be used to infer past climate change. There is ongoing work on further development and refinement of methods, and there are remaining research issues concerning the degree to which the methods have spatial and seasonal biases. Therefore, in many recent palaeoclimatic studies, a combination of methods is applied since multi-proxy series provide more rigorous estimates than a single proxy approach, and the multi-proxy approach may identify possible seasonal biases in the estimates. No palaeoclimatic method is foolproof, and knowledge of the underlying methods and processes is required when using palaeoclimatic data. The field of palaeoclimatology depends heavily on replication and cross-verification between palaeoclimate records from independent sources in order to build confidence in inferences about past climate variability and change. In this chapter, the most weight is placed on those inferences that have been made with particularly robust or replicated methodologies.
6.2.2 Methods – Palaeoclimate Modelling
Climate models are used to simulate episodes of past climate (e.g., the Last Glacial Maximum, the last interglacial period or abrupt climate events) to help understand the mechanisms of past climate changes. Models are key to testing physical hypotheses, such as the Milankovitch theory (Section 6.4, Box 6.1), quantitatively. Models allow the linkage of cause and effect in past climate change to be investigated. Models also help to fill the gap between the local and global scale in palaeoclimate, as palaeoclimatic information is often sparse, patchy and seasonal. For example, long ice core records show a strong correlation between local temperature in Antarctica and the globally mixed gases CO 2 and methane, but the causal connections between these variables are best explored with the help of models. Developing a quantitative understanding of mechanisms is the most effective way to learn from past climate for the future, since there are probably no direct analogues of the future in the past. At the same time, palaeoclimate reconstructions offer the possibility of testing climate models, particularly if the climate forcing can be appropriately specified, and the response is sufficiently well constrained. For earlier climates (i.e., before the current ‘Holocene’ interglacial), forcing and responses cover a much larger range, but data are more sparse and uncertain, whereas for recent millennia more records are
Palaeoclimate Chapter 6
available, but forcing and response are much smaller. Testing models with palaeoclimatic data is important, as not all aspects of climate models can be tested against instrumental climate data. For example, good performance for present climate is not a conclusive test for a realistic sensitivity to CO 2 – to test this, simulation of a climate with a very different CO 2 level can be used. In addition, many parametrizations describing sub-grid scale processes (e.g., cloud parameters, turbulent mixing) have been developed using present-day observations; hence climate states not used in model development provide an independent benchmark for testing models. Palaeoclimate data are key to evaluating the ability of climate models to simulate realistic climate change. In principle the same climate models that are used to simulate present-day climate, or scenarios for the future, are also used to simulate episodes of past climate, using differences in prescribed forcing and (for the deep past) in configuration of oceans and continents. The full spectrum of models (see Chapter 8) is used (Claussen et al., 2002), ranging from simple conceptual models, through Earth System Models of Intermediate Complexity (EMICs) and coupled General Circulation Models (GCMs). Since long simulations (thousands of years) can be required for some palaeoclimatic applications, and computer power is still a limiting factor, relatively ‘fast’ coupled models are often used. Additional components that are not standard in models used for simulating present climate are also increasingly added for palaeoclimate applications, for example, continental ice sheet models or components that track the stable isotopes in the climate system (LeGrande et al., 2006). Vegetation modules as well as terrestrial and marine ecosystem modules are increasingly included, both to capture biophysical and biogeochemical feedbacks to climate, and to allow for validation of models against proxy palaeoecological (e.g., pollen) data. The representation of biogeochemical tracers and processes is a particularly important new advance for palaeoclimatic model simulations, as a rich body of information on past climate has emerged from palaeoenvironmental records that are intrinsically linked to the cycling of carbon and other nutrients.
6.3 The Pre-Quaternary Climates
6.3.1 What is the Relationship Between Carbon Dioxide and Temperature in this Time Period?
Pre-Quaternary climates prior to 2.6 Ma (e.g., Figure 6.1) were mostly warmer than today and associated with higher CO 2 levels. In that sense, they have certain similarities with the anticipated future climate change (although the global biology and geography were increasingly different further back in time). In general, they verify that warmer climates are to be expected with increased greenhouse gas concentrations. Looking back in time beyond the reach of ice cores, that is, prior to about
1 Ma, data on greenhouse gas concentrations in the atmosphere become much more uncertain. However, there are ongoing efforts to obtain quantitative reconstructions of the warm climates over the past 65 Myr and the following subsections discuss two particularly relevant climate events of this period. How accurately is the relationship between CO 2 and temperature known? There are four primary proxies used for pre-Quaternary CO 2 levels (Jasper and Hayes, 1990; Royer et al., 2001; Royer, 2003). Two proxies apply the fact that biological entities in soils and seawater have carbon isotope ratios that are distinct from the atmosphere (Cerling, 1991; Freeman and Hayes, 1992; Yapp and Poths, 1992; Pagani et al., 2005). The third proxy uses the ratio of boron isotopes (Pearson and Palmer, 2000), while the fourth uses the empirical relationship between stomatal pores on tree leaves and atmospheric CO 2 content (McElwain and Chaloner, 1995; Royer, 2003). As shown in Figure 6.1 (bottom panel), while there is a wide range of reconstructed CO 2 values, magnitudes are generally higher than the interglacial, pre-industrial values seen in ice core data. Changes in CO 2 on these long time scales are thought to be driven by changes in tectonic processes (e.g., volcanic activity source and silicate weathering drawdown; e.g., Ruddiman, 1997). Temperature reconstructions, such as that shown in Figure 6.1 (middle panel), are derived from O isotopes (corrected for variations in the global ice volume), as well as Mg/Ca in forams and alkenones. Indicators for the presence of continental ice on Earth show that the planet was mostly ice-free during geologic history, another indication of the general warmth. Major expansion of antarctic glaciations starting around 35 to 40 Ma was likely a response, in part, to declining atmospheric CO 2 levels from their peak in the Cretaceous (~100 Ma) (DeConto and Pollard, 2003). The relationship between CO 2 and temperature can be traced further back in time as indicated in Figure 6.1 (top panel), which shows that the warmth of the Mesozoic Era (230–65 Ma) was likely associated with high levels of CO 2 and that the major glaciations around 300 Ma likely coincided with low CO 2 concentrations relative to surrounding periods.
6.3.2 What Does the Record of the Mid-Pliocene Show?
The Mid-Pliocene (about 3.3 to 3.0 Ma) is the most recent time in Earth’s history when mean global temperatures were substantially warmer for a sustained period (estimated by GCMs to be about 2°C to 3°C above pre-industrial temperatures; Chandler et al., 1994; Sloan et al., 1996; Haywood et al., 2000; Jiang et al., 2005), providing an accessible example of a world that is similar in many respects to what models estimate could be the Earth of the late 21st century. The Pliocene is also recent enough that the continents and ocean basins had nearly reached their present geographic configuration. Taken together, the average of the warmest times during the middle Pliocene presents a view of the equilibrium state of a globally warmer world, in which atmospheric CO 2 concentrations (estimated
Palaeoclimate Chapter 6
to be between 360 to 400 ppm) were likely higher than pre- industrial values (Raymo and Rau, 1992; Raymo et al., 1996), and in which geologic evidence and isotopes agree that sea level was at least 15 to 25 m above modern levels (Dowsett and Cronin, 1990; Shackleton et al., 1995), with correspondingly reduced ice sheets and lower continental aridity (Guo et al., 2004). Both terrestrial and marine palaeoclimate proxies (Thompson, 1991; Dowsett et al., 1996; Thompson and Fleming, 1996) show that high latitudes were significantly warmer, but that tropical SSTs and surface air temperatures were little different from the present. The result was a substantial decrease in the lower-tropospheric latitudinal temperature gradient. For example, atmospheric GCM simulations driven by reconstructed SSTs from the Pliocene Research Interpretations and Synoptic Mapping Group (Dowsett et al., 1996; Dowsett et al., 2005) produced winter surface air temperature warming of 10°C to 20°C at high northern latitudes with 5°C to 10°C increases over the northern North Atlantic (~60°N), whereas there was essentially no tropical surface air temperature change (or even slight cooling) (Chandler et al., 1994; Sloan et al., 1996; Haywood et al., 2000, Jiang et al., 2005). In contrast, a coupled atmosphere-ocean experiment with an atmospheric CO 2 concentration of 400 ppm produced warming relative to pre-industrial times of 3°C to 5°C in the northern North Atlantic, and 1°C to 3°C in the tropics (Haywood et al., 2005), generally similar to the response to higher CO 2 discussed in Chapter 10. The estimated lack of tropical warming is a result of basing tropical SST reconstructions on marine microfaunal evidence. As in the case of the Last Glacial Maximum (see Section 6.4), it is uncertain whether tropical sensitivity is really as small as such reconstructions suggest. Haywood et al. (2005) found that alkenone estimates of tropical and subtropical temperatures do indicate warming in these regions, in better agreement with GCM simulations from increased CO 2 forcing (see Chapter 10). As in the study noted above, climate models cannot produce a response to increased CO 2 with large high-latitude warming, and yet minimal tropical temperature change, without strong increases in ocean heat transport (Rind and Chandler, 1991). The substantial high-latitude response is shown by both marine and terrestrial palaeodata, and it may indicate that high latitudes are more sensitive to increased CO 2 than model simulations suggest for the 21st century. Alternatively, it may be the result of increased ocean heat transports due to either an enhanced thermohaline circulation (Raymo et al., 1989; Rind and Chandler, 1991) or increased flow of surface ocean currents due to greater wind stresses (Ravelo et al., 1997; Haywood et al., 2000), or associated with the reduced extent of land and sea ice (Jansen et al., 2000; Knies et al., 2002; Haywood et al., 2005). Currently available proxy data are equivocal concerning a possible increase in the intensity of the meridional overturning cell for either transient or equilibrium climate states during the Pliocene, although an increase would contrast with the North Atlantic transient deep-water production decreases that are found in most coupled model simulations for the 21st century
(see Chapter 10). The transient response is likely to be different from an equilibrium response as climate warms. Data are just beginning to emerge that describe the deep ocean state during the Pliocene (Cronin et al., 2005). Understanding the climate distribution and forcing for the Pliocene period may help improve predictions of the likely response to increased CO 2 in the future, including the ultimate role of the ocean circulation in a globally warmer world.
6.3.3 What Does the Record of the Palaeocene- Eocene Thermal Maximum Show?
Approximately 55 Ma, an abrupt warming (in this case of the order of 1 to 10 kyr) by several degrees celsius is indicated by changes in 18 O isotope and Mg/Ca records (Kennett and Stott, 1991; Zachos et al., 2003; Tripati and Elderfield, 2004). The warming and associated environmental impact was felt at all latitudes, and in both the surface and deep ocean. The warmth lasted approximately 100 kyr. Evidence for shifts in global precipitation patterns is present in a variety of fossil records including vegetation (Wing et al., 2005). The climate anomaly, along with an accompanying carbon isotope excursion, occurred at the boundary between the Palaeocene and Eocene epochs, and is therefore often referred to as the Palaeocene-Eocene Thermal Maximum (PETM). The thermal maximum clearly stands out in high-resolution records of that time (Figure 6.2). At the same time, 13 C isotopes in marine and continental records show that a large mass of carbon with low 13 C concentration must have been released into the atmosphere and ocean. The mass of carbon was sufficiently large to lower the pH of the ocean and drive widespread dissolution of seafloor carbonates (Zachos et al., 2005). Possible sources for this carbon could have been methane (CH 4 ) from decomposition of clathrates on the sea floor, CO 2 from volcanic activity, or oxidation of sediments rich in organic matter (Dickens et al., 1997; Kurtz et al., 2003; Svensen et al., 2004). The PETM, which altered ecosystems worldwide (Koch et al., 1992; Bowen et al., 2002; Bralower, 2002; Crouch et al., 2003; Thomas, 2003; Bowen et al., 2004; Harrington et al., 2004), is being intensively studied as it has some similarity with the ongoing rapid release of carbon into the atmosphere by humans. The estimated magnitude of carbon release for this time period is of the order of 1 to 2 × 10^18 g of carbon (Dickens et al., 1997), a similar magnitude to that associated with greenhouse gas releases during the coming century. Moreover, the period of recovery through natural carbon sequestration processes, about 100 kyr, is similar to that forecast for the future. As in the case of the Pliocene, the high-latitude warming during this event was substantial (~20°C; Moran et al., 2006) and considerably higher than produced by GCM simulations for the event (Sluijs et al.,
- or in general for increased greenhouse gas experiments (Chapter 10). Although there is still too much uncertainty in the data to derive a quantitative estimate of climate sensitivity from the PETM, the event is a striking example of massive carbon release and related extreme climatic warming.
Chapter 6 Palaeoclimate
Figure 6.2. The Palaeocene-Eocene Thermal Maximum as recorded in benthic (bottom dwelling) foraminifer (Nuttallides truempyi) isotopic records from sites in the Antarctic, south Atlantic and Pacifi c (see Zachos et al., 2003 for details). The rapid decrease in carbon isotope ratios in the top panel is indicative of a large increase in atmospheric greenhouse gases CO 2 and CH 4 that was coincident with an approximately 5°C global warming (centre panel). Using the carbon isotope records, numerical models show that CH 4 released by the rapid decomposition of marine hydrates might have been a major component (~2,000 GtC) of the carbon flux (Dickens and Owen, 1996). Testing of this and other models requires an independent constraint on the carbon fluxes. In theory, much of the additional greenhouse carbon would have been absorbed by the ocean, thereby lowering seawater pH and causing widespread dissolution of seafloor carbonates. Such a response is evident in the lower panel, which shows a transient reduction in the carbonate (CaCO 3 ) content of sediments in two cores from the south Atlantic (Zachos et al., 2004, 2005). The observed patterns indicate that the ocean’s carbonate saturation horizon rapidly shoaled more than 2 km, and then gradually recovered as buffering processes slowly restored the chemical balance of the ocean. Initially, most of the carbonate dissolution is of sediment deposited prior to the event, a process that offsets the apparent timing of the dissolution horizon relative to the base of the benthic foraminifer carbon isotope excursion. Model simulations show that the recovery of the carbonate saturation horizon should precede the recovery in the carbon isotopes by as much as 100 kyr (Dickens and Owen, 1996), another feature that is evident in the sediment records.
Chapter 6 Palaeoclimate
Box 6.1: Orbital Forcing
It is well known from astronomical calculations (Berger, 1978) that periodic changes in parameters of the orbit of the Earth around the Sun modify the seasonal and latitudinal distribution of incoming solar radiation at the top of the atmosphere (hereafter called ‘insolation’). Past and future changes in insolation can be calculated over several millions of years with a high degree of confidence (Berger and Loutre, 1991; Laskar et al., 2004). This box focuses on the time period from the past 800 kyr to the next 200 kyr. Over this time interval, the obliquity (tilt) of the Earth axis varies between 22.05° and 24.50° with a strong quasi-periodicity around 41 kyr. Changes in obliquity have an impact on seasonal contrasts. This parameter also modulates annual mean insolation changes with opposite effects in low vs. high latitudes (and therefore no effect on global average insolation). Local annual mean insolation changes remain below 6 W m–2. The eccentricity of the Earth’s orbit around the Sun has longer quasi-periodicities at 400 and around 100 kyr, and varies between values of about 0.002 and 0.050 during the time period from 800 ka to 200 kyr in the future. Changes in eccentricity alone modulate the Sun-Earth distance and have limited impacts on global and annual mean insolation. However, changes in eccentricity affect the intra-annual changes in the Sun-Earth distance and thereby modulate significantly the seasonal and latitudinal effects induced by obliquity and climatic precession. Associated with the general precession of the equinoxes and the longitude of perihelion, periodic shifts in the position of solstices and equinoxes on the orbit relative to the perihelion occur, and these modulate the seasonal cycle of insolation with periodicities of about 19 and about 23 kyr. As a result, changes in the position of the seasons on the orbit strongly modulate the latitudinal and sea- sonal distribution of insolation. When averaged over a season, insolation changes can reach 60 W m–2^ (Box 6.1, Figure 1). During peri- ods of low eccentricity, such as about 400 ka and during the next 100 kyr, seasonal insolation changes induced by precession are less strong than during periods of larger eccentricity (Box 6.1, Figure 1). High-frequency variations of orbital variations ap- pear to be associated with very small insolation changes (Bertrand et al., 2002a). The Milankovitch theory propos- es that ice ages are triggered by min- ima in summer insolation near 65°N, enabling winter snowfall to persist all year and therefore accumulate to build NH glacial ice sheets. For ex- ample, the onset of the last ice age, about 116 ± 1 ka (Stirling et al., 1998), corresponds to a 65°N mid-June in- solation about 40 W m–2^ lower than today (Box 6.1, Figure 1). Studies of the link between or- bital parameters and past climate changes include spectral analysis of palaeoclimatic records and the iden- tification of orbital periodicities; pre- cise dating of specific climatic transi- tions; and modelling of the climate response to orbital forcing, which highlights the role of climatic and biogeochemical feedbacks. Sections 6.4 and 6.5 describe some aspects of the state-of-the-art understanding of the relationships between orbital forcing, climate feedbacks and past climate changes.
Box 6.1, Figure 1. (Left) December to February (top), annual mean (middle) and June to August (bottom) latitu- dinal distribution of present-day (year 1950) incoming mean solar radiation (W m–2). (Right) Deviations with respect to the present of December to February (top), annual mean (middle) and June to August (bottom) latitudinal distribu- tion of incoming mean solar radiation (W m–2) from the past 500 kyr to the future 100 kyr (Berger and Loutre, 1991; Loutre et al., 2004).
Palaeoclimate Chapter 6
Box 6.2: What Caused the Low Atmospheric Carbon Dioxide Concentrations During Glacial Times?
Ice core records show that atmospheric CO 2 varied in the range of 180 to 300 ppm over the glacial-interglacial cycles of the last 650 kyr (Figure 6.3; Petit et al., 1999; Siegenthaler et al., 2005a). The quantitative and mechanistic explanation of these CO 2 variations remains one of the major unsolved questions in climate research. Processes in the atmosphere, in the ocean, in marine sediments and on land, and the dynamics of sea ice and ice sheets must be considered. A number of hypotheses for the low glacial CO 2 concentra- tions have emerged over the past 20 years, and a rich body of literature is available (Webb et al., 1997; Broecker and Henderson, 1998; Archer et al., 2000; Sigman and Boyle, 2000; Kohfeld et al., 2005). Many processes have been identified that could potentially regulate atmospheric CO 2 on glacial-interglacial time scales. However, the existing proxy data with which to test hypotheses are relatively scarce, uncertain, and their interpretation is partly conflicting. Most explanations propose changes in oceanic processes as the cause for low glacial CO 2 concentrations. The ocean is by far the largest of the relatively fast-exchanging (<1 kyr) carbon reservoirs, and terrestrial changes cannot explain the low glacial values be- cause terrestrial storage was also low at the Last Glacial Maximum (see Section 6.4.1). On glacial-interglacial time scales, atmospheric CO 2 is mainly governed by the interplay between ocean circulation, marine biological activity, ocean-sediment interactions, seawater carbonate chemistry and air-sea exchange. Upon dissolution in seawater, CO 2 maintains an acid/base equilibrium with bicarbon- ate and carbonate ions that depends on the acid-titrating capacity of seawater (i.e., alkalinity). Atmospheric CO 2 would be higher if the ocean lacked biological activity. CO 2 is more soluble in colder than in warmer waters; therefore, changes in surface and deep ocean temperature have the potential to alter atmospheric CO 2. Most hypotheses focus on the Southern Ocean, where large volume- fractions of the cold deep-water masses of the world ocean are currently formed, and large amounts of biological nutrients (phos- phate and nitrate) upwelling to the surface remain unused. A strong argument for the importance of SH processes is the co-evolution of antarctic temperature and atmospheric CO 2. One family of hypotheses regarding low glacial atmospheric CO 2 values invokes an increase or redistribution in the ocean alkalinity as a primary cause. Potential mechanisms are (i) the increase of calcium carbonate (CaCO 3 ) weathering on land, (ii) a decrease of coral reef growth in the shallow ocean, or (iii) a change in the export ratio of CaCO 3 and organic material to the deep ocean. These mecha- nisms require large changes in the deposition pattern of CaCO 3 to explain the full amplitude of the glacial-interglacial CO 2 difference through a mechanism called carbonate compensation (Archer et al., 2000). The available sediment data do not support a dominant role for carbonate compensation in explaining low glacial CO 2 levels. Furthermore, carbonate compensation may only explain slow CO 2 variation, as its time scale is multi-millennial. Another family of hypotheses invokes changes in the sinking of marine plankton. Possible mechanisms include (iv) fertilization of phytoplankton growth in the Southern Ocean by increased deposition of iron-containing dust from the atmosphere after being carried by winds from colder, drier continental areas, and a subsequent redistribution of limiting nutrients; (v) an increase in the whole ocean nutrient content (e.g., through input of material exposed on shelves or nitrogen fi xation); and (vi) an increase in the ratio between carbon and other nutrients assimilated in organic material, resulting in a higher carbon export per unit of limiting nutrient exported. As with the first family of hypotheses, this family of mechanisms also suffers from the inability to account for the full amplitude of the reconstructed CO 2 variations when constrained by the available information. For example, periods of enhanced biological production and increased dustiness (iron supply) are coincident with CO 2 concentration changes of 20 to 50 ppm (see Section 6.4.2, Figure 6.7). Model simulations consistently suggest a limited role for iron in regulating past atmospheric CO 2 concentration (Bopp et al., 2002). Physical processes also likely contributed to the observed CO 2 variations. Possible mechanisms include (vii) changes in ocean tem- perature (and salinity), (viii) suppression of air-sea gas exchange by sea ice, and (ix) increased stratification in the Southern Ocean. The combined changes in temperature and salinity increased the solubility of CO 2 , causing a depletion in atmospheric CO 2 of perhaps 30 ppm. Simulations with general circulation ocean models do not fully support the gas exchange-sea ice hypothesis. One explanation (ix) conceived in the 1980s invokes more stratification, less upwelling of carbon and nutrient-rich waters to the surface of the Southern Ocean and increased carbon storage at depth during glacial times. The stratification may have caused a depletion of nutrients and carbon at the surface, but proxy evidence for surface nutrient utilisation is controversial. Qualitatively, the slow ventilation is consistent with very saline and very cold deep waters reconstructed for the last glacial maximum (Adkins et al., 2002), as well as low glacial stable carbon isotope ratios (^13 C/ 12 C) in the deep South Atlantic. In conclusion, the explanation of glacial-interglacial CO 2 variations remains a difficult attribution problem. It appears likely that a range of mechanisms have acted in concert (e.g., Köhler et al., 2005). The future challenge is not only to explain the amplitude of glacial-interglacial CO 2 variations, but the complex temporal evolution of atmospheric CO 2 and climate consistently.
Palaeoclimate Chapter 6
Crucifix and Hewitt, 2005). Changes in biogeochemical cycles thus played an important role and contributed, through changes in greenhouse gas concentration, dust loading and vegetation cover, more than half of the known radiative perturbation during the LGM. Overall, the radiative perturbation for the changed greenhouse gas and aerosol concentrations and land surface was approximately –8 W m–2^ for the LGM, although with
significant uncertainty in the estimates for the contributions of aerosol and land surface changes (Figure 6.5). Understanding of the magnitude of tropical cooling over land at the LGM has improved since the TAR with more records, as well as better dating and interpretation of the climate signal associated with snow line elevation and vegetation change. Reconstructions of terrestrial climate show strong spatial
Figure 6.4. The concentrations and radiative forcing by (a) CO 2 , (b) CH 4 and (c) nitrous oxide (N 2 O), and (d) the rate of change in their combined radiative forcing over the last 20 kyr reconstructed from antarctic and Greenland ice and firn data (symbols) and direct atmospheric measurements (red and magenta lines). The grey bars show the reconstructed ranges of natural variability for the past 650 kyr (Siegenthaler et al., 2005a; Spahni et al., 2005). Radiative forcing was computed with the simplified expres- sions of Chapter 2 (Myhre et al., 1998). The rate of change in radiative forcing (black line) was computed from spline fits (Enting, 1987) of the concentration data (black lines in panels a to c). The width of the age distribution of the bubbles in ice varies from about 20 years for sites with a high accumulation of snow such as Law Dome, Antarctica, to about 200 years for low-accumulation sites such as Dome C, Antarctica. The Law Dome ice and fi rn data, covering the past two millennia, and recent instrumental data have been splined with a cut-off period of 40 years, with the resulting rate of change in radiative forcing shown by the inset in (d). The arrow shows the peak in the rate of change in radiative forcing after the anthropogenic signals of CO 2 , CH 4 and N 2 O have been smoothed with a model describing the enclosure process of air in ice (Spahni et al., 2003) ap- plied for conditions at the low accumulation Dome C site for the last glacial transition. The CO 2 data are from Etheridge et al. (1996); Monnin et al. (2001); Monnin et al. (2004); Siegenthaler et al. (2005b; South Pole); Siegenthaler et al. (2005a; Kohnen Station); and MacFarling Meure et al. (2006). The CH 4 data are from Stauffer et al. (1985); Steele et al. (1992); Blunier et al. (1993); Dlugokencky et al. (1994); Blunier et al. (1995); Chappellaz et al. (1997); Monnin et al. (2001); Flückiger et al. (2002); and Ferretti et al. (2005). The N 2 O data are from Machida et al. (1995); Battle et al. (1996); Flückiger et al. (1999, 2002); and MacFarling Meure et al. (2006). Atmospheric data are from the National Oceanic and Atmospheric Administration’s global air sampling network, representing global average concentrations (dry air mole fraction; Steele et al., 1992; Dlugokencky et al., 1994; Tans and Conway, 2005), and from Mauna Loa, Hawaii (Keeling and Whorf, 2005). The globally averaged data are available from http://www.cmdl.noaa.gov/.
Chapter 6 Palaeoclimate
Frequently Asked Question 6. What Caused the Ice Ages and Other Important Climate Changes Before the Industrial Era?
Climate on Earth has changed on all time scales, including long before human activity could have played a role. Great prog- ress has been made in understanding the causes and mechanisms of these climate changes. Changes in Earth’s radiation balance were the principal driver of past climate changes, but the causes of such changes are varied. For each case – be it the Ice Ages, the warmth at the time of the dinosaurs or the fluctuations of the past millennium – the specific causes must be established individually. In many cases, this can now be done with good confidence, and many past climate changes can be reproduced with quantitative models. Global climate is determined by the radiation balance of the planet (see FAQ 1.1). There are three fundamental ways the Earth’s radiation balance can change, thereby causing a climate change: (1) changing the incoming solar radiation (e.g., by changes in the Earth’s orbit or in the Sun itself), (2) changing the fraction of solar radiation that is reflected (this fraction is called the albedo – it can be changed, for example, by changes in cloud cover, small particles called aerosols or land cover), and (3) altering the long- wave energy radiated back to space (e.g., by changes in green- house gas concentrations). In addition, local climate also depends on how heat is distributed by winds and ocean currents. All of these factors have played a role in past climate changes. Starting with the ice ages that have come and gone in regu- lar cycles for the past nearly three million years, there is strong evidence that these are linked to regular variations in the Earth’s orbit around the Sun, the so-called Milankovitch cycles (Figure 1). These cycles change the amount of solar radiation received at each latitude in each season (but hardly affect the global annual mean), and they can be calculated with astronomical precision. There is still some discussion about how exactly this starts and ends ice ages, but many studies suggest that the amount of sum- mer sunshine on northern continents is crucial: if it drops below a critical value, snow from the past winter does not melt away in summer and an ice sheet starts to grow as more and more snow accumulates. Climate model simulations confirm that an Ice Age can indeed be started in this way, while simple conceptual models have been used to successfully ‘hindcast’ the onset of past glacia- tions based on the orbital changes. The next large reduction in northern summer insolation, similar to those that started past Ice Ages, is due to begin in 30,000 years. Although it is not their primary cause, atmospheric carbon di- oxide (CO 2 ) also plays an important role in the ice ages. Antarctic ice core data show that CO 2 concentration is low in the cold gla- cial times (~190 ppm), and high in the warm interglacials (~ ppm); atmospheric CO 2 follows temperature changes in Antarctica with a lag of some hundreds of years. Because the climate changes at the beginning and end of ice ages take several thousand years,
most of these changes are affected by a positive CO 2 feedback; that is, a small initial cooling due to the Milankovitch cycles is subsequently amplified as the CO 2 concentration falls. Model sim- ulations of ice age climate (see discussion in Section 6.4.1) yield realistic results only if the role of CO 2 is accounted for. During the last ice age, over 20 abrupt and dramatic climate shifts occurred that are particularly prominent in records around the northern Atlantic (see Section 6.4). These differ from the gla- cial-interglacial cycles in that they probably do not involve large changes in global mean temperature: changes are not synchro- nous in Greenland and Antarctica, and they are in the opposite direction in the South and North Atlantic. This means that a major change in global radiation balance would not have been needed to cause these shifts; a redistribution of heat within the climate system would have sufficed. There is indeed strong evidence that changes in ocean circulation and heat transport can explain many features of these abrupt events; sediment data and model simula- tions show that some of these changes could have been triggered by instabilities in the ice sheets surrounding the Atlantic at the time, and the associated freshwater release into the ocean. Much warmer times have also occurred in climate history – during most of the past 500 million years, Earth was probably completely free of ice sheets (geologists can tell from the marks ice leaves on rock), unlike today, when Greenland and Antarc- tica are ice-covered. Data on greenhouse gas abundances going back beyond a million years, that is, beyond the reach of antarc- tic ice cores, are still rather uncertain, but analysis of geological
FAQ 6.1, Figure 1. Schematic of the Earth’s orbital changes (Milankovitch cycles) that drive the ice age cycles. ‘T’ denotes changes in the tilt (or obliquity) of the Earth’s axis, ‘E’ denotes changes in the eccentricity of the orbit (due to variations in the minor axis of the ellipse), and ‘P’ denotes precession, that is, changes in the direction of the axis tilt at a given point of the orbit. Source: Rahmstorf and Schellnhuber (2006).
(continued)
Chapter 6 Palaeoclimate
Figure 6.5. The Last Glacial Maximum climate (approximately 21 ka) relative to the pre-industrial (1750) climate. (Top left) Global annual mean radiative influences (W m–2) of LGM climate change agents, generally feedbacks in glacial-interglacial cycles, but also specifi ed in most Atmosphere-Ocean General Circulation Model (AOGCM) simulations for the LGM. The heights of the rectangular bars denote best estimate values guided by published values of the climate change agents and conversion to radiative perturbations using simplifi ed expressions for the greenhouse gas concentrations and model calculations for the ice sheets, vegetation and mineral dust. References are included in the text. A judgment of each estimate’s reliability is given as a level of scientific understanding based on uncertainties in the climate change agents and physical understanding of their radiative effects. Paleoclimate Modelling Intercomparison Project 2 (PMIP-2) simulations shown in bottom left and right panels do not include the radiative infl uences of LGM changes in mineral dust or vegetation. (Bottom left) Multi-model average SST change for LGM PMIP-2 simulations by five AOGCMs (Community Climate System Model (CCSM), Flexible Global Ocean-Atmosphere-Land System (FGOALS), Hadley Centre Coupled Model (HadCM), Institut Pierre Simon Laplace Climate System Model (IPSL-CM), Model for Interdisciplinary Research on Climate (MIROC)). Ice extent over continents is shown in white. (Right) LGM regional cooling compared to LGM global cooling as simulated in PMIP-2, with AOGCM results shown as red circles and EMIC (ECBilt-CLIO) results shown as blue circles. Regional averages are defined as: Antarctica, annual for inland ice cores; tropical Indian Ocean, annual for 15°S to 15°N, 50°E to 100°E; and North Atlantic Ocean, July to September for 42°N to 57°N, 35°W to 20°E. Grey shading indicates the range of observed proxy estimates of regional cooling: Antarctica (Stenni et al., 2001; Masson-Delmotte et al., 2006), tropical Indian Ocean (Rosell-Mele et al., 2004; Barrows and Jug- gins, 2005), and North Atlantic Ocean (Rosell-Mele et al., 2004; Kucera et al., 2005; de Vernal et al., 2006; Kageyama et al., 2006).
cold and salty deep water (Adkins et al., 2002). Most of the deglaciation occurred over the period about 17 to 10 ka, the same period of maximum deglacial atmospheric CO 2 increase (Figure 6.4). It is thus very likely that the global warming of 4°C to 7°C since the LGM occurred at an average rate about 10 times slower than the warming of the 20th century. In summary, significant progress has been made in the understanding of regional changes at the LGM with the development of new proxies, many new records, improved understanding of the relationship of the various proxies to climate variables and syntheses of proxy records into reconstructions with stricter dating and common calibrations.
6.4.1.3 How Realistic Are Results from Climate Model Simulations of the Last Glacial Maximum?
Model intercomparisons from the first phase of the Paleoclimate Modelling Intercomparison Project (PMIP-1), using atmospheric models (either with prescribed SST or with simple slab ocean models), were featured in the TAR. There are now six simulations of the LGM from the second phase (PMIP-2) using Atmosphere-Ocean General Circulation Models (AOGCMs) and EMICs, although only a few regional comparisons were completed in time for this assessment. The radiative perturbation for the PMIP-2 LGM simulations available for this assessment, which do not yet include the effects of vegetation or aerosol changes, is –4 to –7 W m–2.
Palaeoclimate Chapter 6
These simulations allow an assessment of the response of a subset of the models presented in Chapters 8 and 10 to very different conditions at the LGM. The PMIP-2 multi-model LGM SST change shows a modest cooling in the tropics, and greatest cooling at mid- to high latitudes in association with increases in sea ice and changes in ocean circulation (Figure 6.5). The PMIP-2 modelled strengthening of the SST meridional gradient in the LGM North Atlantic, as well as cooling and expanded sea ice, agrees with proxy indicators (Kageyama et al., 2006). Polar amplification of global cooling, as recorded in ice cores, is reproduced for Antarctica (Figure 6.5), but the strong LGM cooling over Greenland is underestimated, although with caveats about the heights of these ice caps in the PMIP-2 simulations (Masson- Delmotte et al., 2006). The PMIP-2 AOGCMs give a range of tropical ocean cooling between 15°S to 15°N of 1.7°C to 2.4°C. Sensitivity simulations with models indicate that this tropical cooling can be explained by the reduced glacial greenhouse gas concentrations, which had direct effects on the tropical radiative forcing (Shin et al., 2003; Otto-Bliesner et al., 2006b) and indirect effects through LGM cooling by positive sea ice-albedo feedback in the Southern Ocean contributing to enhanced ocean ventilation of the tropical thermocline and the intermediate waters (Liu et al., 2002). Regional variations in simulated tropical cooling are much smaller than indicated by MARGO data, partly related to models at current resolutions being unable to simulate the intensity of coastal upwelling and eastern boundary currents. Simulated cooling in the Indian Ocean (Figure 6.5), a region with important present-day teleconnections to Africa and the North Atlantic, compares favourably to proxy estimates from alkenones (Rosell-Mele et al., 2004) and foraminifera assemblages (Barrows and Juggins, 2005). Considering changes in vegetation appears to improve the realism of simulations of the LGM, and points to important climate-vegetation feedbacks (Wyputta and McAvaney, 2001; Crucifix and Hewitt, 2005). For example, extension of the tundra in Asia during the LGM contributes to the local surface cooling, while the tropics warm where savannah replaces tropical forest (Wyputta and McAvaney, 2001). Feedbacks between climate and vegetation occur locally, with a decrease in the tree fraction in central Africa reducing precipitation, and remotely with cooling in Siberia (tundra replacing trees) altering (diminishing) the Asian summer monsoon. The physiological effect of CO 2 concentration on vegetation needs to be included to properly represent changes in global forest (Harrison and Prentice, 2003), as well as to widen the climatic range where grasses and shrubs dominate. The biome distribution simulated with dynamic global vegetation models reproduces the broad features observed in palaeodata (e.g., Harrison and Prentice, 2003). In summary, the PMIP-2 LGM simulations confirm that current AOGCMs are able to simulate the broad-scale spatial patterns of regional climate change recorded by palaeodata in response to the radiative forcing and continental ice sheets of the LGM, and thus indicate that they adequately represent
the primary feedbacks that determine the climate sensitivity of this past climate state to these changes. The PMIP- AOGCM simulations using glacial-interglacial changes in greenhouse gas forcing and ice sheet conditions give a radiative perturbation in reference to pre-industrial conditions of –4.6 to –7.2 W m–2^ and mean global temperature change of –3.3°C to –5.1°C, similar to the range reported in the TAR for PMIP- (IPCC, 2001). The climate sensitivity inferred from the PMIP- LGM simulations is 2.3°C to 3.7°C for a doubling of atmospheric CO 2 (see Section 9.6.3.2). When the radiative perturbations of dust content and vegetation changes are estimated, climate models yield an additional cooling of 1°C to 2°C (Crucifix and Hewitt, 2005; Schneider et al., 2006), although scientific understanding of these effects is very low.
6.4.1.4 How Realistic Are Simulations of Terrestrial Carbon Storage at the Last Glacial Maximum?
There is evidence that terrestrial carbon storage was reduced during the LGM compared to today. Mass balance calculations based on 13 C measurements on shells of benthic foraminifera yield a reduction in the terrestrial biosphere carbon inventory (soil and living vegetation) of about 300 to 700 GtC (Shackleton, 1977; Bird et al., 1994) compared to the pre-industrial inventory of about 3,000 GtC. Estimates of terrestrial carbon storage based on ecosystem reconstructions suggest an even larger difference (e.g., Crowley, 1995). Simulations with carbon cycle models yield a reduction in global terrestrial carbon stocks of 600 to 1,000 GtC at the LGM compared to pre-industrial time (Francois et al., 1998; Beerling, 1999; Francois et al., 1999; Kaplan et al., 2002; Liu et al., 2002; Kaplan et al., 2003; Joos et al., 2004). The majority of this simulated difference is due to reduced simulated growth resulting from lower atmospheric CO 2. A major regulating role for CO 2 is consistent with the model-data analysis of Bond et al. (2003), who suggested that low atmospheric CO 2 could have been a significant factor in the reduction of trees during glacial times, because of their slower regrowth after disturbances such as fire. In summary, results of terrestrial models, also used to project future CO 2 concentrations, are broadly compatible with the range of reconstructed differences in glacial-interglacial carbon storage on land.
6.4.1.5 How Long Did the Previous Interglacials Last?
The four interglacials of the last 450 kyr preceding the Holocene (Marine Isotope Stages 5, 7, 9 and 11) were all different in multiple aspects, including duration (Figure 6.3). The shortest (Stage 7) lasted a few thousand years, and the longest (Stage 11; ~420 to 395 ka) lasted almost 30 kyr. Evidence for an unusually long Stage 11 has been recently reinforced by new ice core and marine sediment data. The European Programme for Ice Coring in Antarctica (EPICA) Dome C antarctic ice core record suggests that antarctic temperature remained approximately as warm as the Holocene
