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6.16 Tracers of Past Ocean Circulation J 6.16 Tracers of Past Ocean Circulation J. Lynch-Stieglitz Columbia University, Palisades, NY, USA 6.16.1 INTRODUCTION 433 6.16.2 NUTRIENT WATER MASS TRACERS 434 6.16.2.1 Carbon Isotopes 434 6.16.2.1.1 Controls on d13C of oceanic carbon 434 6.16.2.1.2 Carbon isotope ratios in benthic foraminifera 435 6.16.2.1.3 Cadmium in benthic foraminifera 437 6.16.2.1.4 Cadmium in seawater 437 6.16.2.1.5 Cd/Ca in foraminifera 438 6.16.2.2 Ba/Ca 439 6.16.2.3 Zn/Ca 439 6.16.3 CONSERVATIVE WATER MASS TRACERS 439 6.16.3.1 Mg/Ca in Benthic Foraminifera 439 6.16.3.2 Pore-water Chemistry 439 6.16.3.3 Oxygen Isotopes in Benthic Foraminifera 440 6.16.3.4 Carbon Isotope Air–Sea Exchange Signature 441 6.16.4 NEODYMIUM ISOTOPE RATIOS 441 6.16.5 CIRCULATION RATE TRACERS 441 6.16.5.1 Radiocarbon 441 6.16.5.2 231Pa/ 230Th Ratio 442 6.16.5.3 Geostrophic Shear Estimates from d18O in Benthic Foraminifera 442 6.16.6 NON-GEOCHEMICAL TRACERS OF PAST OCEAN CIRCULATION 443 6.16.7 OCEAN CIRCULATION DURING THE LAST GLACIAL MAXIMUM 443 6.16.8 CONCLUSIONS 447 ACKNOWLEDGMENTS 448 REFERENCES 448 6.16.1 INTRODUCTION flow patterns and mixing of subsurface water masses. Information about how the ocean circulated Several decades ago it was realized that during the past is useful in understanding changes chemistry of the shells of benthic foraminifera in ocean and atmospheric chemistry, changes in (carbon isotope and Cd/Ca ratios) carried an the fluxes of heat and freshwater between the imprint of the nutrient content of deep-water ocean and atmosphere, and changes in global wind masses (Shackleton, 1977; Broecker, 1982; Boyle, patterns. The circulation of surface waters in the 1981). This led rapidly to the recognition that the ocean leaves an imprint on sea surface tempera- water masses in the Atlantic Ocean were arrayed ture, and is also inextricably linked to the patterns differently during the last glacial maximum than of oceanic productivity. Much valuable infor- they are today, and the hypothesis that the glacial mation about past ocean circulation has been arrangement reflected a diminished contribution inferred from reconstructions of surface ocean of low-nutrient North Atlantic deep water temperature and productivity, which are covered (NADW) (Curry and Lohmann, 1982; Boyle and in separate chapters. Here the focus is on the Keigwin, 1982). More detailed spatial reconstruc- geochemical tracers that are used to infer the tions indicated a shallow nutrient-depleted water 433 434 Tracers of Past Ocean Circulation mass overlying a more nutrient-rich water mass in nutrient tracer reflects the increased contribution the glacial Atlantic. These findings spurred from the regeneration of organic material as the advances not only in geochemistry but in water mass ages, it also carries a signature of the oceanography and climatology, as workers in initial nutrient content of the source regions. In these fields attempted to simulate the inferred today’s Atlantic, the contrast between high- glacial circulation patterns and assess the nutrient Southern Ocean source water and low- vulnerability of the modern ocean circulation to nutrient NADW is strong, and the residence time of changes such as observed for the last ice age. deep water in the Atlantic Ocean is short, allowing While the nutrient distributions in the glacial little additional accumulation of nutrients from Atlantic Ocean were consistent with a diminished sinking organic material. In this case the nutrient flow of NADW, they also could have reflected an tracers can be used to document the relative increase in inflow from the South Atlantic and/or a mixtures of Northern and Southern water masses. shallower yet undiminished deep-water mass. Clearly, tracers capable of giving information on 13 deep-water flow rate, rather than nutrient content 6.16.2.1.1 Controls on d C of oceanic alone, were needed. Differences between surface carbon water (measured on planktonic foraminifera) and During photosynthesis, organisms preferen- deep-water (measured on coexisting benthic tially take up the lighter isotope of carbon (12C) foraminifera) radiocarbon concentrations pro- increasing surface ocean d13C. When the 13C- vided the first rate constraint (Broecker et al., depleted organic matter is decomposed, this CO2 1988; Shackleton et al., 1988). Reduced amounts release decreases the d13C of seawater DIC at of protactinium relative to the more particle- depth (Deuser and Hunt, 1969; Craig, 1970; reactive thorium in the glacial Atlantic suggested Kroopnick, 1985). If this process were the only that deep water was exported from the Atlantic one responsible for the distribution of d13Cin during glacial times (Yu et al., 1996). More the ocean, d13C of oceanic carbon would recently, density gradients in upper waters have decrease by ,1.1‰ for every 1 mmol kg21 been used to infer changes in the upper ocean increase in oceanic PO4 (Broecker and Maier- return flow that compensates the deep-water Reimer, 1992). Reflecting this biological cycling, export (Lynch-Stieglitz et al., 1999b). as a deep water mass with a homogeneous source Many of these tracers of paleo-ocean flow have ages (such as in the deep Indian and Pacific been applied to all of the ocean basins, and have 13 oceans) the d CandPO4 increase in this been extended in time throughout the Neogene. biologically expected ratio (Figure 1). However, Despite this progress, a consistent picture of the the discrimination against 13C during photosyn- circulation of the ocean during even the last ice thesis increases from about 19‰ in the warm age has yet to emerge. While circulation tracers surface ocean to about 30‰ in the Antarctic, suggest a rearrangement of water masses in the 13 which means that the PO4 and d C changes Atlantic, there is still considerable disagreement associated with the uptake and regeneration of about the water masses and circulation in the rest organic material can vary regionally. of the World Ocean. Some of this arises from still Carbon isotope fractionation during air–sea insufficient data coverage, but some is the result of exchange is also an important factor in determin- conflicting information from the various deep ing the isotopic composition of seawater. If the circulation tracers. In this chapter, we examine in CO2 in the atmosphere were in isotopic equili- more detail the methods used to reconstruct past brium with the dissolved inorganic carbon (DIC) ocean circulation, which will illuminate the source in the ocean, the DIC would be enriched in 13C of some of this conflicting information. relative to the atmospheric CO2 by ,8‰ at 20 8C (Inoue and Sugimura, 1985). This equilibrium fractionation depends on the temperature of 6.16.2 NUTRIENT WATER MASS TRACERS equilibration, with the oceanic carbon more 6.16.2.1 Carbon Isotopes enriched relative to the atmospheric value by about 0.1‰ per degree of cooling (Mook et al., Measuring the carbon isotope ratio in the calcite 1974). If the surface ocean were everywhere in tests of bottom dwelling (benthic) foraminifera is isotopic equilibrium with atmospheric CO2, perhaps the most widespread method for recon- the 30 8C range in ocean temperatures would structing the distribution and properties of deep- cause a 3‰ range in oceanic d13C with higher water masses. The carbon isotope ratios in the values where surface temperatures are cold, and foraminifera reflect that ratio in seawater. In turn, lower values in warm surface waters. This range is the carbon isotope ratio in the deep sea is primarily similar to the magnitude of d13C change induced controlled by the regeneration of 13C-poor organic by biological processes. However, isotopic equili- material and has a distribution much like a nutrient brium is rarely reached, because the timescale for such as phosphate in the modern ocean. While a carbon isotope equilibration between the ocean Nutrient Water Mass Tracers 435 1.5 d 13 C d 13 as =0 C as NADW =0.5 1.0 WSBW CDCDWW 0.5 C (‰ PDB) 13 d Biology 0.0 d 13 C as =–0.5 Air–sea exchange –0.5 1.0 1.5 2.0 2.5 3.0 µ –1 PO4 ( mol kg ) 13 Figure 1 Deep water (.2,000 m) d C and PO4 data from the modern ocean after Lynch-Stieglitz et al. (1995). The 13 13 lines indicate a “Redfield slope” and d Cas indicates the deviation from the mean ocean Redfield slope for d C and PO4 (Broecker and Maier-Reimer, 1992). While the deep Indian and Pacific values follow the slope expected for biologic processes, deep waters which form in the North Atlantic and Southern Ocean have, respectively, lower and higher d13C due to air–sea exchange. 13 and atmosphere (,10 years for a 50 m deep mixed By definition, water with d Cas of 0‰ has the layer) is about 10 times longer than chemical same air–sea exchange signature as mean ocean 13 equilibration between atmospheric CO2 and DIC deep water. A positive value of d Cas means that 13 (Broecker and Peng, 1974; Tans, 1980). This the water has a higher d Cas than mean ocean deep isotopic equilibration time is significantly longer water (more of an influence of air–sea exchange at than surface waters stay in one place in the ocean. cold temperatures/less at warm temperatures), 13 The exchange of atmospheric CO2 and surface where as a negative d Cas value implies less d13 ocean aqueous CO2 has the potential to leave C enrichment due to air–sea exchange than for 13 mean ocean deep water.
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