thropogenic transients ( 4 , 7 ). Firn air sampling allows one to collect 1,000 liters or more of isotopically unfractionated air as old as 90 yr. These samples can be used to measure concentrations of trace organic compounds, O2/N2 ratios, and the isotopic composition of greenhouse gases, all measurements that are far more difficult to make on ice core samples.

Glacial–Interglacial Changes in Atmospheric Chemistry Recorded in Ice Cores

Selected climate records are summarized in Fig. 4 and Fig. 5 , covering the periods from 0 to 350 kyr and from 0 to 100 kyr before the present (B.P.), respectively. The d18O of calcitic foraminifera from deep sea sediments is a proxy indicator for ice volume. The dD or d18O of ice from ice cores is a proxy indicator of temperature in the area of the ice core. Recent inversions of borehole temperature data conclude that the glacial–interglacial temperature change in Greenland was about 20°C ( 39 , 40 ). The dust content of ice cores is a proxy for dryness of the source areas. High dust contents are generally attributed to dry source areas from which dust is readily suspended into the atmosphere. CO2 and CH4 concentrations are measured in bubbles of polar ice as described above. June solar insolation at 65° N latitude is plotted at the bottom of the figure.

On a broad scale, the atmospheric CO2 concentration is clearly linked to climate change. Lorius et al. ( 41 ) argued that the glacial–interglacial CO2 change (with a small contribution from CH4) was responsible for a 2°C temperature change when all feedbacks are considered. Obviously this number is highly uncertain but it is very likely that increasing CO2 concentrations in air contributed to global warming during glacial terminations 1 and 2. The link between CO2 and higher-frequency climate changes is weak. For example, there is no minimum in CO2 corresponding to the cold period (Glacial Stage 5d) at about 110 kyr B.P. Similarly, CO2 does not rise during the interstadial events of the past 35 kyr (ref. 42; Fig. 5 ).

Many factors have been invoked to explain the glacialinterglacial change in the CO2 concentration of air. The cause of this change was vigorously debated for a decade beginning in 1982. Recently attention to this subject has diminished, more because the protagonists are exhausted than because the issue is resolved. It is widely accepted that the pCO2 of the atmosphere is regulated by the pCO2 of surface seawater, because about 99% of the CO2 in the ocean/atmosphere system resides in the oceans. In his classic paper initiating the debate, Broecker ( 43 ) noted that lower glacial temperatures would cause the CO2 concentration of air to fall, whereas higher salinities would cause CO2 to rise. A glacial ocean temperature decrease of 1.5°C would cause CO2 to fall by 20 ppmv. This change would be roughly offset by a salinity increase of 1‰. Recent controversial results showing that equatorial temperatures may have increased by up to 5°C suggest that the glacial–interglacial change in sea surface temperatures may have been much larger than Broecker initially estimated. If this is correct, the temperature change may have contributed 20–30 ppm of the glacial–interglacial change in atmospheric pCO2 after correcting for the compensating effect of salinity (the temperature dependence of pCO2 on sea surface temperature is about 13 ppmv/°C; ref. 43).

Clearly other factors must contribute to the glacialinterglacial difference in atmospheric pCO2. These factors must act either by decreasing the total CO2 concentration of surface seawater or by increasing the surface alkalinity. Changes in ocean circulation, the nutrient/carbon ratio of organic matter, and varying rates of nutrient utilization in Antarctic waters have been invoked to account for the changing surface water total CO2 concentrations. Removal of CaCO3 during sea level rise, changing ratios of CaCO3 to organic carbon in biogenic debris falling out of the surface ocean, and changes in the calcium carbonate compensation depth in the deep ocean have been invoked to change the alkalinity of surface waters.

There have been large changes in the CH4 concentration of air during the last ˜200 kyr ( Fig. 4 and Fig. 5 ). CH4 changes are very closely linked with changes in climate ( 29 , 44 ). CH4 differs from CO2 in two important ways. First, CO2 changes are about five times more important than those of CH4 in driving glacial–interglacial temperature changes ( 41 ). Second, CH4 changes, which are believed to be caused mainly by changes in emission rates, are somewhat better understood and serve much better as a climate proxy. Chappellaz et al. ( 27 ) generally interpreted changes in atmospheric CH4 as reflecting increases in the extent of low-latitude wetlands due to increased precipitation and perhaps temperatures.

Recent records of atmospheric CH4 variations over the last 200 kyr ( Fig. 4 and Fig. 5 ) reveal several very interesting features. As noted above, there is a large glacial–interglacial change, with Antarctic CH4 concentrations varying between glacial levels of 350 ppbv and interglacial levels of 700 ppbv. There are changes of somewhat smaller magnitude associated with climate changes linked to the 20-kyr climate cycles associated with precession. Examples are CH4 maxima associated with warm periods at 80 kyr and 100 kyr B.P. Third, there are large CH4 variations associated with interstadial events of the last 40 kyr. These events, recorded most dramatically in Greenland ice cores (e.g., ref. 30; Fig. 5 ), reflect rapid warming over Greenland, slow cooling, and rapid cooling back to baseline glacial values. During the last 40 kyr at least, interstadial warmings are accompanied by increases of about 150 ppbv in the CH4 concentration of air. Chappellaz et al. ( 30 ) invoked wetter and warmer conditions in the tropics to explain the CH4 increases. Support for this idea comes from the fact that the longer interstadial events, at least, are recorded in Antarctica, illustrating their global nature ( 45 ), and are sometimes coincident with deposition of cave deposits in Botswanaland, which Holmgren et al. ( 46 ) linked to increased tropical precipitation. Finally, the CH4 concentration of the atmosphere was surprisingly variable during the Holocene. Methane was at a maximum early in the Holocene, fell to a broad minimum about 5 kyr B.P., and has risen during the last 3 kyr ( 22 ). Only the abrupt minimum at about 8.2 kyr B.P. has been directly linked to a climate event, namely an abrupt cooling that occurred at that time. The causes of the mid-Holocene minimum and subsequent slow rise are topics of current investigation.

This research was supported by grants from the New England Regional Center of the National Institute of Global Environmental Change, Department of Energy (Subagreement 901214-HAR), and the Office of Polar Programs, National Science Foundation.

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