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The Pacific stations, with the exception of MANOP M, lie below the lysocline. It is immediately apparent that fractional dissolution at these sites, with the exception of the California borderland, is greater than at the Atlantic sites. Estimates range from 70 to 100 percent (except for one low estimate for MANOP S). The lower value at the California borderland site (Berelson et al., 1987) is surprising, and contrasts to the conclusion of Berger (1970), who found that 85 percent of foraminifera shells reaching the sediments there dissolved, even at depths as shallow as a few hundred meters on the slope of the Santa Barbara basin. The Pacific results argue for the importance of dissolution below the saturation horizon. The results tend to indicate that dissolution is occuring deep enough within the sediments to be measured by limited-resolution pore water profiles. At all the MANOP sites, estimates are based both on the particle rain/burial and particle rain/diagenetic flux pairs, and they agree well (again, the low estimate based on NO3- profiles at MANOP S is an exception). In addition, the pore water based estimates are calculated from organic matter oxidation stoichiometry, and so do not include dissolution as a result of undersaturated bottom waters. If the sediment trap based estimates of rain rates do not underestimate the true rain rate of CaCO3, then we must conclude that, even at these sites below the lysocline, dissolution due to metabolic acids is an important process.
The driving force for biogenic silica dissolution in marine sediments is similar to that for CaCO3: pore waters are undersaturated with respect to the solid phase present. There are two important differences, however: seawater is universally undersaturated with respect to amorphous silica (Hurd, 1972), and pore waters rarely reach saturation; and the rate of dissolution is significantly slower for biogenic silica, occurring on a time scale of years to tens of years in sediments, compared to minutes to days for CaCO3. The result of these differences is that sediment mixing plays a more important role in determining silica preservation.
In laboratory experiments, it has been shown that the dissolution of opaline tests in seawater depends on the concentration of opal surface exposed to the seawater, on several details of surface characteristics, and on the degree of saturation of the water with respect to amorphous silica (Hurd and Theyer, 1975; Hurd, 1983). Marine pore waters rarely reach saturation with respect to amorphous silica (about 1000 µM at 3° C: Hurd and Theyer, 1975). However, pore waters do approach asymptotic dissolved silica concentrations in the upper 50 to 150 cm of the sediments. The values vary from little over 100 µM in sediments underlying unproductive surface waters (Bender et al.,
TABLE 10.8 Biogenic Silica Mass Balances: Percent Preservation
1985/86; Sayles, 1979), to values of up to 580 µM below the equatorial upwelling region in the eastern Pacific (Jahnke et al., 1982b), and to values in excess of 750 µM in the siliceous oozes of the southern ocean (Sayles, 1981). Schink et al. (1975) have explained this variability in asymptotic pore water values in terms of a kinetic model based on the rate law of Hurd (1972). According to this model, the asymptotic [SiO2] results from a balance between supply to the sediments and loss to overlying water. The balance depends on the rain rate of opal to the sea floor, the rate of opal dissolution, and the bioturbation mixing coefficient. Opal burial efficiency must also depend on these factors.
The data for calculating biogenic silica burial efficiencies are limited, but enough exist to illustrate its variability. In Table 10.8, we list calculations for five sites from the eastern equatorial Pacific, the California borderland, and the Antarctic continental shelf. Fractional dissolution varies substantially, from about 90 percent in eastern equatorial Pacific sediments to values as low as 25 percent on the Antarctic continental shelf. Clearly, the highest preservation rates occur on the Antarctic shelf (>50 percent of the rain is preserved). This is consistent with the fact that about 75 percent of the opal accumulation in marine sediments occurs in the deep ocean surrounding Antarctica and on the Antarctic continental shelf (Ledford-Hoffman et al., 1986).
The study of organic carbon diagenesis is complicated by the wide variety of compounds grouped under the heading. Its bulk reactivity varies with the source of the organic matter, as different compounds have differing susceptibilities to degradation in sediments (e.g., Emerson and Hedges, 1988); in addition, the composition of or-