6.4 Glacial-Interglacial Variability and Dynamics
6.4.1 Climate Forcings and Responses Over Glacial-Interglacial Cycles
Palaeoclimatic records document a sequence of glacial-interglacial cycles covering the last 740 kyr in ice cores (EPICA community members, 2004), and several million years in deep oceanic sediments (Lisiecki and Raymo, 2005) and loess (Ding et al., 2002). The last 430 kyr, which are the best documented, are characterised by 100-kyr glacial-interglacial cycles of very large amplitude, as well as large climate changes corresponding to other orbital periods (Hays et al., 1976; Box 6.1), and at millennial time scales (McManus et al., 2002; NorthGRIP, 2004). A minor proportion (20% on average) of each glacial-interglacial cycle was spent in the warm interglacial mode, which normally lasted for 10 to 30 kyr (Figure 6.3). There is evidence for longer interglacial periods between 430 and 740 ka, but these were apparently colder than the typical interglacials of the latest Quaternary (EPICA community members, 2004). The Holocene, the latest of these interglacials, extends to the present.
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 seasonal distribution of insolation. When averaged over a season, insolation changes can reach 60 W m–2 (Box 6.1, Figure 1). During periods 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 appear to be associated with very small insolation changes (Bertrand et al., 2002a).
The Milankovitch theory proposes that ice ages are triggered by minima in summer insolation near 65°N, enabling winter snowfall to persist all year and therefore accumulate to build NH glacial ice sheets. For example, the onset of the last ice age, about 116 ± 1 ka (Stirling et al., 1998), corresponds to a 65°N mid-June insolation about 40 W m–2 lower than today (Box 6.1, Figure 1).
Studies of the link between orbital parameters and past climate changes include spectral analysis of palaeoclimatic records and the identification of orbital periodicities; precise dating of specific climatic transitions; 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.
The ice core record indicates that greenhouse gases co-varied with antarctic temperature over glacial-interglacial cycles, suggesting a close link between natural atmospheric greenhouse gas variations and temperature (Box 6.2). Variations in CO2 over the last 420 kyr broadly followed antarctic temperature, typically by several centuries to a millennium (Mudelsee, 2001). The sequence of climatic forcings and responses during deglaciations (transitions from full glacial conditions to warm interglacials) are well documented. High-resolution ice core records of temperature proxies and CO2 during deglaciation indicates that antarctic temperature starts to rise several hundred years before CO2 (Monnin et al., 2001; Caillon et al., 2003). During the last deglaciation, and likely also the three previous ones, the onset of warming at both high southern and northern latitudes preceded by several thousand years the first signals of significant sea level increase resulting from the melting of the northern ice sheets linked with the rapid warming at high northern latitudes (Petit et al., 1999; Shackleton, 2000; Pépin et al., 2001). Current data are not accurate enough to identify whether warming started earlier in the Southern Hemisphere (SH) or Northern Hemisphere (NH), but a major deglacial feature is the difference between North and South in terms of the magnitude and timing of strong reversals in the warming trend, which are not in phase between the hemispheres and are more pronounced in the NH (Blunier and Brook, 2001).
Greenhouse gas (especially CO2) feedbacks contributed greatly to the global radiative perturbation corresponding to the transitions from glacial to interglacial modes (see Section 6.4.1.2). The relationship between antarctic temperature and CO2 did not change significantly during the past 650 kyr, indicating a rather stable coupling between climate and the carbon cycle during the late Pleistocene (Siegenthaler et al., 2005a). The rate of change in atmospheric CO2 varied considerably over time. For example, the CO2 increase from about 180 ppm at the Last Glacial Maximum to about 265 ppm in the early Holocene occurred with distinct rates over different periods (Monnin et al., 2001; Figure 6.4).
Box 6.2: What Caused the Low Atmospheric Carbon Dioxide Concentrations During Glacial Times?
Ice core records show that atmospheric CO2 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 CO2 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 CO2 concentrations 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 CO2 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 CO2 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 because terrestrial storage was also low at the Last Glacial Maximum (see Section 6.4.1). On glacial-interglacial time scales, atmospheric CO2 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, CO2 maintains an acid/base equilibrium with bicarbonate and carbonate ions that depends on the acid-titrating capacity of seawater (i.e., alkalinity). Atmospheric CO2 would be higher if the ocean lacked biological activity. CO2 is more soluble in colder than in warmer waters; therefore, changes in surface and deep ocean temperature have the potential to alter atmospheric CO2. 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 (phosphate 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 CO2.
One family of hypotheses regarding low glacial atmospheric CO2 values invokes an increase or redistribution in the ocean alkalinity as a primary cause. Potential mechanisms are (i) the increase of calcium carbonate (CaCO3) weathering on land, (ii) a decrease of coral reef growth in the shallow ocean, or (iii) a change in the export ratio of CaCO3 and organic material to the deep ocean. These mechanisms require large changes in the deposition pattern of CaCO3 to explain the full amplitude of the glacial-interglacial CO2 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 CO2 levels. Furthermore, carbonate compensation may only explain slow CO2 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 fixation); 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 CO2 variations when constrained by the available information. For example, periods of enhanced biological production and increased dustiness (iron supply) are coincident with CO2 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 CO2 concentration (Bopp et al., 2002).
Physical processes also likely contributed to the observed CO2 variations. Possible mechanisms include (vii) changes in ocean temperature (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 CO2, causing a depletion in atmospheric CO2 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 (13C/12C) in the deep South Atlantic.
In conclusion, the explanation of glacial-interglacial CO2 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 CO2 variations, but the complex temporal evolution of atmospheric CO2 and climate consistently.