Geol. 656 Isotope Geochemistry Lecture 36 Spring 2003 STABLE ISOTOPES IN PALEOCLIMATOLOGY II CARBON ISOTOPES, OCEAN CIRCULATION, AND CLIMATE In the previous lecture, we noted the need for feedback mechanisms to amply the Milankovitch sig- nal, which is the primary driving force of Quaternary climate oscillations. We observed that the Milankovitch variations change only the distribution of solar energy received by the Earth, not the total amount. If this were the only factor in climate change, we would expect that the glaciation in the southern and northern hemispheres would be exactly out of phase. This, however, is not the case. Thus there must be feedback mechanisms at work capable of producing globally synchronous climate variation. Broecker (Broecker, 1984 and subsequent papers) argued that one of these was the deep cir- culation of the ocean. The role of surface ocean in climate is well understood: for example, the south-flowing California Current keeps the West Coast of the U.S. relatively dry and the coastal regions cooler than they would otherwise be. The role of the deep, or thermohaline, circulation of the oceans is less obvious, but perhaps no less important. Whereas the surface ocean circulation is wind-driven, the deep circu- lation is driven by density, which is in turn controlled by temperature and salinity. In the present ocean, most deep ocean water masses “form” in high latitudes. Once these deep- water masses form, they do not return to the surface for nearly a thousand years. The principal site of deep water formation is the Southern Ocean where the Antarctic Intermediate Water (AAIW) is formed in the Antarctic Convergence and Antarctic Bottom Water (AABW), the densest of ocean wa- ter masses, is formed in the Weddell Sea. A lesser amount of deep water is also formed in the North Atlantic during winter south of Iceland; this water mass is called North Atlantic Deep Water (NADW). The only “deep water” formed at intermediate latitudes is Mediterranean Intermediate Water (MIW), which sinks as a result of evaporation in the Mediterranean increasing salinity and hence density. MIW, however, is a smaller water masses than the others, and furthermore has only intermediate density. Formation of deep water thus usually involves loss of thermal energy by the ocean to the atmosphere. Therefore, the present thermohaline circulation of the oceans keeps high latitude climates milder than they would otherwise be. In particular, energy extracted from the At- lantic Ocean water in the formation of NADW keeps the European climate relatively mild. We saw in Lecture 28 that d13C is lower in deep water than in surface water (Figure 28.5). This re- sults from biological cycling: photosynthesis in the surface waters discriminates against 13C, leaving the dissolved inorganic carbon of surface waters with high d13C, while oxidation of falling organic particles rich in 12C lowers d13C in deep water: in effect, 12C is “pumped” from surface to deep water more efficiently than 13C. d13C values in the deep water are not uniform, varying with the “age” of the deep water: the longer the time since the water was at the surface, the more enriched it becomes 12 13 in C and the lower the d C. Since this is also true of total inorganic carbon and nutrients such as PO4 13 13 and NO3, d C correlates negatively with nutrient and SCO2 concentrations. NADW has high d C be- cause it contains water that was recently at the surface (and hence depleted in 12C by pho- tosynthesis). Deep water is not formed in either the Pacific or the Indian Oceans; all deep waters in those oceans flow in from the Southern Ocean. Hence deep water in the Pacific, being rather “old” has low d13C. AABW is a mixture of young NADW, which therefore has comparably high d13C, and recirculated Pacific deep water and hence has lower d13C than NADW. Thus these water masses can be distinguished on the basis of d13C. Examining d13C in benthic foraminifera in cores from a variety of locations, Oppo and Fairbanks (1987) concluded that production of NADW was lower during the last glacial maximum and increased to present levels in the interval between 15000 and 5000 years ago. Figure 36.1 shows an example of data from core RC13-229, located in the South Atlantic. d13C values increase as d18O increases. As we saw in the previous lecture, d18O in marine carbonates is a measure of glacial ice volume and climate. As the climate warmed at the end of the last glacial interval, d13C values in bottom water in the 266 4/28/03 Geol. 656 Isotope Geochemistry . Lecture 36 Spring 2003 4.5 South Atlantic increased, reflecting a in- crease in the proportion of NADW relative 4 more to AABW in this region. From d13C varia- tions in Mediterranean and Central Atlantic 3.5 cores, Oppo and Fairbanks (1987) also con- 3 cluded that the production of MIW was d18O greater during the last glacial maximum. Ice Volume 2.5 Thus the mode of ocean circulation appar- ently changes between glacial and intergla- 2 less cial times; this change may well amplify 0.6 the Milankovitch signal. Further evidence for the role of North At- 0.4 lantic Deep Water was provided by a 1992 more study by Charles and Fairbanks. Working 0.2 with a high resolution core (i.e., high sedi- 13 mentation rate) from the Southern Ocean, d C 0 they found d13C increased dramatically just -0.2 NADW -0.4 less -0.6 0 5 10 15 20 25 Age, ka Figure 36.1. Variation in d18O and d13C in benthic foraminfera from core RC13-229 from the eastern South Atlantic. d13C data suggest the proportion of NADW in this region increased as the climate warmed. Data from Oppo and Fairbanks (1987). over 12,000 years ago (Figure 36.2), indicating a greatly increased flux of NADW to the region. The in- crease in NADW occurred just as ice volume was begin- ning to decrease rapidly (as judged from increasing sealevel in Barbados). The rather sudden melting of the continental ice caps is hard to explain by the slowly increasing Northern Hemisphere insolation at that time. If production of NADW suddenly started, it would have warmed the North Atlantic climate, accelerating melting of glaciers. Thus NADW produc- tion may represent a positive feedback amplifying the primary “Milankovitch” signal. However, why Figure 36.2. Calculated northern hemi- NADW production should suddenly increase dramati- sphere insolation changes (due to orbital cally remains a mystery. changes), glacial melt water discharge 18 (calculated from sealevel rise rates deter- THE TERTIARY d O RECORD mined from 14C dating of fossil coral reefs Imbrie’s (1985) analysis suggests that the climate off Barbados), and d13C in bethic foramin- system’s response to Milankovitch forcing has changed fera in core RC11-83 (41°36’ S, 94° 48’ E) significantly even over the last 800,000 years. The from the Southern Ocean. The d13C data present glacial-interglacial cycles began only 2 mil- are a measure of the proportion of NADW lion years ago, yet the Milankovitch forcing must in Circumpolar Deep Water. European have been present before that. This suggests the cli- pollen assemblages are also shown. From Charles and Fairbanks (1992). 267 4/28/03 Geol. 656 Isotope Geochemistry Lecture 36 Spring 2003 mate system’s response has changed even more drasti- cally over the past few mil- lion years. In addition to ocean circulation and atmos- pheric CO2, the positions of land masses relative to the poles and elevation of land masses may play an im- portant role in global cli- mate. All these factors have varied during the Tertiary, so it is perhaps not surprising that there have been signifi- cant climatic changes through the Tertiary. These changes have been recorded by d 18O in deep-sea sedi- ments. Figure 36.3. Variation in d18O in planktonic (a) and benthic (b) fora- Figure 36.3 shows the Ter- minifera over the past 60 million years in 3 DSDP cores from the 18 tiary d O variations in ben- South Pacific. Note that d 18O is relative to PDB, rather than thic and planktonic fora- SMOW. This is conventional for carbonates. After Shackleton and minifera recorded in three Kennett (1975). DSDP cores from the South Pacific. The data show an increase in d18O through this time. Because extensive northern hemi- sphere glaciation only began in the Pleistocene or late Pliocene, variations in d18O through most of the Tertiary are thought to primarily reflect temperature changes rather than changes in ice volume, though the latter, mainly in Antarctica, were also important. The d18O record therefore testifies to gradual cooling through the Tertiary. Superimposed on the general increase in d18O are some important “events” in which d18O changes more rapidly. Going backward through time, these include the shifts that mark the onset of Pleisto- cene glaciation, the rapid increase in d18O from mid-Miocene through Pliocene, and the nearly 1‰ in- crease at or near the Eocene-Oligocene boundary, and a more steady decrease, amounting to nearly 2‰ during the Eocene. Comparison of cores from different parts of the ocean shows that the changes are globally synchronous. Studies of spatially distributed cores suggest that global temperatures were some 2° C warmer dur- ing the Eocene that at present. Perhaps more significantly, the latitudinal gradient in temperature may have been only half the present one. This suggests oceanic and atmospheric circulation was dif- ferent from the present, and on the whole much more efficient at transporting heat from equator to poles. Why this was so remains unclear.