As first noted by Heinrich, 1988, glacial age

Abstract. Down water column traverses of core top weights for three planktonic species confirm Lohmann's [1995] relationship between foraminifera shell weight loss and bottom water carbonate ion content. However, they also suggest that the initial shell thickness varies with growth habitat and that the offset between bottom water and pore water carbonate ion concentration varies even on small space scales. IntroductionLohmann [1995] makes a case that the weight of whole foraminifera shells picked from a narrow size range (355-415 Ixm) provides a measure of the extent of dissolution and in so doing has the potential to serve as a paleocarbonate ion proxy. He further points out that this method is not only likely more reliable than the size index proxy [Broecker and Clark, 1999] but also that it is applicable over a wider range of water depth. We agree with Lohmann and therefore have undertaken down core studies using both approaches. In conjunction with these studies we have attempted to evaluate two important assumptions behind Lohmann's method. The first is that the initial weight of planktonic shells of a given species and diameter is independent of growth conditions, in other words, that the thickness of the shell walls is always the same. The second is that the slope of the shell weight for a given species versus pressure-corrected carbonate ion concentration in bottom water is the same for all locales. Since much of the dissolution documented by Lohmann [1995] occurs in the sediment pore waters, one might suspect the relationship to vary from one locale to another. The reason is that the ACO3 = between bottom water and pore water might differ from place to place depending on the rain rate of organic carbon. In order to shed light on these questions we have made weight measurements on three planktonic species (Globigerina sacculifer, Pulleniatina obliquiloculata, and Neogloboquadrina dutertrei) from core top samples coveting a range of water depths at three locales in the tropical ocean (the Ceara Rise in the Atlantic, the Ninetyeast Ridge in the Indian, and the Ontong Java Plateau in the Pacific, see Figure 1). Methods and ResultsWet sediment is passed through a 63-1xm sieve. The coarse fraction is then soaked for 2 hours in a sodium hexametaphosphate solution and then subjected to an 8-s duration ultrasonic treatment. The material is then dried and passed through a pair of sieves in order to isolate the 355-415 vim size fraction. Then, 50 whole shells of each of the three species are picked and weighed using a microbalance. The mean whole-shell weights established in this way are listed in Table 1 There is, however, a suggestion for P. obliquiloculata that a dogleg in dissolution rate occurs below 90 Ixmol kg -1, but as two of the three data points supporting this steepening come from eastern equatorial Pacific cores rather than from the Ontong Java locale, until we gain an understanding of Ninetyeast Ridge-Ontong Java Plateau weight difference for P obliquiloculata, we cannot be confident that the do...

[1] On the basis of measurements of the relative amounts of CaCO 3 in the less than 20-mm and the greater than 20-mm size fractions in open ocean core tops, we find that the coccoliths contribute about half the calcite present in late Holocene deep sea sediments which have experienced little or no dissolution. Although this ratio is of importance to the understanding of the ocean's CaCO 3 cycle, we can find only a few quantitative estimates of their relative contribution to currently forming marine sediments. As dissolution of foraminifera calcite takes place more rapidly than that of coccolith calcite, coccoliths dominate the CaCO 3 in sediments which have experienced sizable dissolution. Although coccoliths contribute 40-60 of the CaCO 3 in tropical sediments, higher-latitude sediments and those adjacent to continental margins often have larger proportions of coccolith CaCO 3 .Citation: Broecker, W., and E. Clark (2009), Ratio of coccolith CaCO 3 to foraminifera CaCO 3 in late Holocene deep sea sediments, Paleoceanography, 24, PA3205,