Seawater as the source of minor elements in black shales, phosphorites and other sedimentary rocks

CHEMICAL GEOLOGY ISOTOPE GEOSCIENCE

ELSEVIER

Chemical Geology 114 (1994) 95-114

Seawater as the source of minor elements in black shales, phosphorites and other sedimentary rocks David Z. Piper U.S. Geological Survey, 345 Middlefield Road, Menlo Park, CA, 94025, USA

(Received June 2, 1993; revision accepted October 15, 1993)

Abstract

Many of the minor elements in seawater today have a concentration-depth profile similar to that of the biologically essential nutrients, NO~- and PO 3-. They show a relative depletion in the photic zone and enrichment in the deep ocean. The difference between their surface- and deep-ocean values, normalized to the change in PO 3-, approaches the average of measured minor-element :P ratios in marine plankton, although individual analyses of the latter show extreme scatter for a variety of reasons. Despite this scatter in the minor-element analyses of plankton, agreement between the two sets of data shows unequivocally that an important marine flux of many minor elements through the ocean is in the form of biogenic matter, with a composition approaching that of plankton. This interpretation is further supported by sediment studies, particularly of sediments which accumulate in shelfslope environments where biological productivity in the photic zone is exceptionally high and organic carbon contents of the underlying sediment elevated. The interelement relations observed tbr some of these sediments approach the average values of plankton. These same interelement relations are observed in many marine sedimentary rocks such as metalliferous black shales and phosphorites, rocks which have a high content of marine fractions (e.g., organic matter, apatite, biogenic silica and carbonates). Many previous studies of the geochemistry of these rocks have concluded that local hydrothermal activity, and/or seawater with an elemental content different from that of the modern ocean, was required to account for their minor-element contents. However, the similarity in several of the minor-element ratios in many of these formations to minor-element ratios in modern plankton demonstrates that these sedimentary rocks accumulated in environments whose marine chemistry was virtually identical to that seen on continental shelf-slopes, or in marginal seas, of the ocean today. The accumulation of the marine fraction of minor elements on these ancient sea floors was determined largely by the accumulation of organic matter, settling from the photic zone and with a composition of average plankton. A second marine fraction of minor elements in these rocks accumulated through precipitation and adsorption from seawater. The suite of elements in this fraction reflects redox conditions in the bottom water, as determined by bacterial respiration. For example, high Mn, high C r + V and high Mo concentrations, above those which can be attributed to the accumulation of planktonic matter, characterize accumulation under bottom-water oxidizing, denitrifying and sulfate-reducing conditions, respectively. [EW]

0009-2541/94/$07.00 © 1994 Elsevier Science B.V. All rights reserved SSDI 0009-2541 ( 93 ) E0240-T

96

D.Z. Piper / Chemical Geology 114 (1994) 95-114

1. Introduction Marine sediments are complex mixtures of detrital debris and seawater-derived material (Goldberg, 1963a); each contributes minor elements to the total sediment inventory. The detrital fraction consists of terrigenous material, introduced to the oceans via riverine and eolian transport, and marine volcanic debris. The seawater-derived fractions accumulate as: ( 1 ) metal oxides, hydroxides or sulfides and adsorbed phases (the hydrogenous fraction); and (2) CaCO3, opal and organic matter (the biogenic fraction). The composition of the detrital fraction reflects provenance, whereas the composition of the marine fractions reflects the conditions of deposition. For example, planktonic debris accumulating on the sea floor delivers minor elements directly to the sediments (the biogenic sediment fraction), possibly in an amount that is a function of primary productivity in the photic zone. Partial oxidation of settling planktonic debris within the water column also releases minor elements in the water column but, perhaps more importantly, controls the precipitation-dissolution reactions of minor elements in seawater (the hydrogenous fraction). This study examines the role of biological productivity in the photic zone to the accumulation of minor elements in sedimentary rocks, through an examination of the m o d e m environment. Many sedimentary rocks have lost this signal through diagenetic alteration immediately following deposition, by alteration during burial at depth, or via weathering during subaerial exposure. The rocks most likely to retain the depositional signal are those with relatively high contents of marine fractions, such as metalliferous shales and phosphorites, rocks which accumulated in continental seas of the ocean's margins, under conditions of high primary productivity in the photic zone (McKelvey et al., 1959). I show that the minor-element content in the marine fraction of these and most sedimentary rocks can be explained by accumulation from an ocean whose seawater chemistry was virtually the same

as that of the modern ocean. No exotic seawater chemistry or minor-element source was required. That is, the content of minor elements, above levels which can be assigned to a detrital fraction, is attributable to: ( 1 ) the accumulation ofbiogenic debris; and (2) the precipitation and/ or adsorption of minor elements from seawater under oxic to denitrifying to sulfate-reducing conditions. The apparently unusual composition of sedimentary deposits, enriched in such minor elements as the heavy metals, is often attributed to a local hydrothermal minor-element input (Hulbert et al., 1992). This theory has long been advocated by marine geologists (Arrhenius and Bonatti, 1965; Dickens and Owen, 1993) as one of several elemental sources for marine sediments. The discovery of ihot brine pools in the Red Sea (Degens and Ross, 1969) and of hydrothermal springs along the East Pacific Rise (Corliss et al., 1979), often in association with massive metal-sulfide deposits (Edmond et al., 1979 ), provides support. The importance of hydrothermal activity cannot be overemphasized; it is a major source of Mn and other minor elements to the global marine budget of minor elements (Von D a m m et al., 1985a), in addition to a terrestrial source. Thus, the oceans are enriched in minor elements via both marine hydrothermal activity and continental denudation (J.-M. Martin and Maybeck, 1979). However, it is unlikely that a hydrothermal source can be identified as an immediate source of minor elements for the vast majority of sedimentary deposits, even though a local hydrothermal source can be clearly identified in a few sedimenthosted, heavy-metal deposits (Goodfellow, 1987; MacIntyre, 1992 ).

2. Biological (planktonic) source The minor-element content of plankton has been measured by several researchers. Measurements vary considerably (J.H. Martin and Knauer, 1973; Eisler, 1981; Collier and Ed-

D.Z. Piper / Chemical Geology 114 (1994) 95-114

mond, 1984) due to inconsistencies introduced during sample collection and analysis, but also to natural factors such as variations among species, geographic variability, growth rates and seawater pH (Eisler, 1981 ). Of the data considered in this study (Fig. 1 ), the care with which the samples were collected and analyzed (J.H. Martin and Knauer, 1973; Collier and Edmond, 1984; Sherrell, 1989) requires that the variations are real, rather than artifacts of sampling and analysis. The results show that we can never expect to obtain a single value for the concentration of any minor element in different samples of plankton. We will have to settle for an average, or "best" estimate. To dismiss these data as unreliable, simply because they show a variability, is analogous to discarding all measurements of primary productivity in Antarctica because summer values do not agree with winter values. Given the amount of variation, I have chosen to use the estimates of Brumsack (1986), rather than introduce a new set of values (Table 1 ). I might add that significant variations also are reported for the contents of NO~-, PO43- and Si(OH)4 in plankton (Sverdrup et al., 1942; Collier and Edmond, 1984), the major nutrients (Redfield et al., 1963 ) whose functions to plankton growth are well established. The magnitude of minor-element variations (Fig. 1 ) can easily discourage further consideration of the data and their application to geochemical investigations of sedimentary rocks which accumulated in ocean-margin basins. However, two important aspects of the marine environment require that the contribution of organic debris to the accumulation of minor elements be examined in these rocks. Firstly, rates of primary productivity throughout the ocean margin are higher than in the oceanic environment (Berger et al., 1988 ), resulting in a greater flux of organic matter to the sea floor (Calvert et al., 1991 ). Secondly, a greater percentage of the organic matter produced in the photic zone accumulates on the sea floor of ocean margins, due to shallower water depths (Reimers and Suess, 1983). As noted by Collier and Edmond (1984), sediments of the ocean margins should represent a major, even dominant sink for many minor

97

elements in the oceans today and, thus, in ancient oceans as well. Seawater profiles of minor elements (i.e. concentration vs. depth) offer a critical evaluation of the selected values of minor-element contents in plankton and settling organic matter. The profiles are similar to the profiles for NO3-, PO aand Si ( O H ) 4 ( Fig. 2 ). Nitrate and PO 3- are the limiting nutrients to phytoplankton productivity in the ocean today (Sverdrup et al., 1942; Broecker and Peng, 1982; McElroy, 1983; Codispoti, 1989), and Si(OH)4 may be the limiting nutrient in some ocean margin areas (Dugdale and Goering, 1970; Nelson et al., 1981 ). They are extracted from seawater by phytoplankton in the photic zone, i.e. during photosynthesis, and returned to the ocean at depth via bacterial respiration. Only a small percentage of the initial particulate form survives to the sea floor (Broecker and Peng, 1982), and only a fraction of that (in the range of 1-5% of surface productivity) is incorporated in pelagic sediments (Calvert et al., 1991). The absolute seawater concentrations of these nutrients and the shapes of their depth profiles are, consequently, largely a function of photosynthesis in the photic zone, bacterial respiration throughout the water column, and water-column advection. As a result of the cycling of N O £ and PO 3between the photic zone and the deep ocean, their concentrations in the surface water are relatively low and even approach zero; they increase sharply below the photic zone to the thermocline, and then remain relatively constant at depths below ~ 500 m (Fig. 2). Their distributions in seawater can be expressed by the following concentration ratio: 6-seawater [ N O ; ] = NO3(deep ..... ) -- NO3-(photi ..... ) 3-3P O 4 (deep ocean ) - - P O 4 ,~photic zone )

(

1)

The ratio is 15:1, equal to the average N : P atomic ratio of plankton (Sverdrup et al., 1942; Redfield et al., 1963; Broecker and Peng, 1982). This similarity demonstrates the dominant role of the biological cycle to the marine budgets of these nutrients.

D.Z. Piper/Chemical Geology 114 (1994) 95-114

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D.Z. Piper / Chemical Geology 114 (1994) 95-114

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Table 1 Content of minor elements (in mass units) in marine plankton and seawater Element

Organic matter" (ppm)

Element: P ratio of organic matter

~-Seawater at O-2000-m depth

Seawaterconcentration at2000-m depth (ppb)

Cd Cr Cu Mo Ni Se V Zn

12 2 11 2 7.5 3 3 110

1.6" 10 -3 0.26"10 -3 1.40"10 3 0.26" 10 -3 1.0" 10 -3 0.40" 10 -3 0.40" 10 -3 14.00- 10 -3

1.3" 10 -3 0.95"10 3 1.1.10-3 0.0 5.1" 10 -3 1.5" 10 -3 1.1" 10 -3 5.8" 10 -3

0.10 b 0.21 c 0.23 d 10.20 e 0.47 f 0.13 g 1.80 h 0.39 f

a-hSources of data: a-Brumsack ( 1986 ); b-Boyle et al. ( 1976 ); c-Murray et al. ( 1983 ); d-Boyle et al. ( 1977 ); e-Collier ( 1985 ); f-Bruland ( 1983); g-Measures et al. ( 1983); h-Collier (1984). Fig. 1. Relations between Cu and: (a) Cd; (b) Ni; and (c) Zn, in plankton samples and suspended-matter samples collected from the Pacific Ocean. Phytoplankton and zooplankton values are from J.H. Martin and Knauer ( 1973 ) and include samples collected in Monterey Bay, California, and the open ocean. Mixed plankton values are samples collected in the open ocean by Collier and Edmond (1984). The data for suspended matter (average of 2 values from the photic zone) are from the western boundary of the California Current (Sherrell, 1989). The trends for A-seawater, extrapolated into the field for plankton, are the ratios of the ~-values listed in Table 1 and discussed in the text. The plankton lines represent concentration ratios suggested by Brumsack (1986).

Minor elements that exhibit a similar depth profile in the oceans are also interpreted as being biologically reactive (Broecker and Peng, 1982 ). Cd, a highly toxic element at relatively low concentrations, nonetheless exhibits a depth profile (Fig. 2 ) that parallels the profile of PO 3- (Boyle et al., 1976), support for an interpretation that it is incorporated in plankton, along with PO43and NO~-, rather than merely being surface adsorbed. The profiles of Cu, Ni and Zn also are similar to the PO43- profile (Bruland, 1983 ). Cu does not approach zero in the photic zone, having a higher concentration in seawater than the demand for plankton growth. The V (Collier, 1984) and Cr (Murray et al., 1983 ) profiles also allow for only slight depletions of these two elements in the photic zone. The closer similarity of the Zn and Ni profiles to the Si (OH)4 than to the PO43- profile (Sclater et al., 1976; Bruland, 1983 ) suggests that they are incorporated in the siliceous fraction of plankton. The ZnmSi(OH)4 relation in seawater (Broecker and Peng, 1982) corresponds to a Zn content in opal of ~ 6 0 ppm, which approaches its content in plankton of 110 ppm

(Table 1 ). However, the average measured content of Zn in the siliceous fraction of plankton, 5-10 ppm (J.H. Martin and Knauer, 1973), indicates that Zn and probably Ni are associated dominantly with the soft parts of organisms, as is Cd. Mo, which has the highest concentration of the heavy metals in seawater (Goldberg, 1963b), exhibits a uniform concentration with ocean depth (Fig. 2), although it too is taken up by plankton (Table 1 ). Its uptake by plankton in the photic zone and remineralization at depth by bacteria are simply insufficient to measurably effect its vertical distribution in the water column. Its concentration in plankton is one-sixth that of Cd, but its concentration in seawater is approximately 100-fold greater. Thus, oxidation of organic matter at depth increases its concentration above its surface-water value by ~ 2%, somewhat less than the precision of the measurement. This is similar to the inability to detect the change in Ca 2+ (Broecker and Peng, 1982), forced by phytoplankton precipitation of C a C 0 3 in the photic zone, for example by coccolithophorids, and its dissolution at depth.

D.Z. Piper / Chemical Geology 114 (1994) 95-114

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Fig. 2. Depth profiles of" ( a ) nutrients; and ( b ) minor elements in seawater, from Bruland ( 1 9 8 0 ) and Collier ( 1 9 8 4 , 1 9 8 5 ).

The ~-seawater values of the minor elements should equal element:P ratios in plankton, similar to ~-seawater[NO~- ], if minor-element distributions in the deep ocean are influenced solely by metabolic processes and advection. This seems to be the case for Cd (Table 1 ). Also, its relation with P in sediment-trap samples (Kremling and Streu, 1993) and filtered suspended matter (Sherrell, 1989) suggests transport by settling organic particles. The ~-seawater values for Cr, Ni, Se and V are slightly higher and for Zn lower than the mean element: P ratios of plankton (Table 1 ). The ~-seawater value for Cu, based only on the uppermost 2000 m of seawater, also resembles the Cu: P ratio of plankton, although it is 2.5-fold higher than the plankton ratio when the Cu concentration in bottom water is used in the calculation, rather than the concentration at 2000-m depth (Fig. 2).

The differences between the ~-seawater values and the element:P ratios likely reflect several processes. Boyle et al. ( 1977 ) explained the distribution of Cu in the deep ocean through a combination of biologic and nonbiologic processes, where the latter involve scavenging of Cu throughout the water column by settling particles (Balistrieri et al., 1981; Li, 1981; Clegg and Sarmiento, 1989) and recycling of Cu into the bottom waters from oxic pelagic sediment (Fischer et al., 1986). These and other nonmetabolic processes in the water column might also explain the apparently high 6-seawater values for Cr, Ni, Se and V, but their marine chemistries have not been examined as thoroughly as that of Cu. Therefore, while the absolute values selected as the metal contents of plankton and settling organic debris (Table 1 ) can be debated, the importance of minor-element transport in the ocean

D.Z. Piper / Chemical Geology 114 (1994) 95-114

via settling organic matter is unquestionable. The profiles for the uppermost 2000 m of seawater would seem to support the selected values for plankton. Whether the metal content of planktonic debris changes drastically during its settling through a few hundred meters of the water column is difficult to establish. Fisher and Wente ( 1993 ) concluded that Am, Ag and Sn are retained sufficiently long, ~'even by decomposing cells, to suggest that phytoplankton sinking ... at rates o f a 100 m / d a y would effectively transport these metals hundreds o f meters ..."

to the sea floor. Alternatively, surface adsorption might increase the minor-element load of sinking organic matter. The content of minor elements in plankton (Brumsack, 1986) and the seawater profiles seem to provide conflicting evidence. Cd should be weakly adsorbed (Clegg and Sarmiento, 1989); consistency between its 6-seawater value and the Cd:P ratio of plankton supports this interpretation. Ni and Zn should be more strongly adsorbed than Cd. In the case of Ni, its 5-seawater value is greater than its element: P ratio in plankton; in the case of Zn, its ~seawater value is less than its plankton value. My interpretation is that the minor-element content of settling organic matter is determined more by the actual uptake of minor elements by phytoplankton in the photic zone than by fractional dissolution (oxidation) and surface adsorption onto settling organic matter. This interpretation is stronger for shallow margin environments than for the deep-ocean oligotrophic environments such as the Sargasso Sea. Ocean margins typically exhibit high productivity in the photic zone and a high flux of planktonic debris to the sea floor. Despite the inherent difficulties of determining an optimal content for plankton, its high minor-element content requires that the biogenic cycle of minor elements in seawater be considered in any explanation of the minor-element content of marine sediments and sedimentary rocks. One way of reducing the uncertainty of absolute contents in plankton is to examine minorelement:minor-element ratios. In the case of Zn

101

and Cu (Fig. 1c), single element concentrations vary by approximately 100-fold, whereas the Zn: Cu ratios vary by I 0-fold, at the 90% confidence level. The advantages of using this approach, rather than ~-seawater values, are." ( 1 ) PO 3- often is not reported; and (2) PO~- geochemistry is quite different from that of the minor elements considered in this study. Studies of the minor elements in the modern ocean and in ancient sediments show that the minor elements are likely retained as sulfides and incorporated in the residual organic matter, particularly in rocks which have retained high contents of organic matter. The biogenic PO 3- fraction is retained as carbonate fluorapatite, possibly resulting in fractionation of PO43- and minor elements during early diagenesis. Average interelement relations in surface sediments from the Namibian Shelf (Calvert and Price, 1983 ) potentially offer an opportunity to determine the flux of minor elements to the sea floor by settling organic matter. This shelf experiences very high productivity and low accumulation of terrigenous material (Calvert and Price, 1983 ). The Cu-Ni relation closely resembles the relation in plankton (Fig. 3 ); Zn is low, but only by approximately a factor of 5; Cd was not measured in these sediments; and Mo scatters to both the high and low side of the plankton value (Fig. 4a). Minor-element contents in the soluble fraction of surface sediment from basins of the California Borderlands (Bruland et al., 1974) also plot in or near the field of plankton (Fig. 3). Their accumulation was attributed largely to anthropogenic factors, except for the Soledad Basin (identified by arrow in Fig. 3). The consistency between Cu: Ni: Zn values in the sediments from these basins and in plankton suggests, however, that a significant fraction of the minor elements might have accumulated as biogenic debris. Additionally, the minor-element contents of the detrital fraction (my interpretation of the fraction reported as the resistant fraction) in these sediments closely approach the minor-element contents of the World Shale Average (Wedepohl, 1969-1978 ), suggesting that the soluble fraction

102

D.Z. Piper / Chemical Geology 114 (1994) 95-114 I

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D.Z. Piper / Chemical Geology 114 (1994) 95-114

measured by Bruland et al. indeed represents the marine fraction. Cretaceous black shales collected at Deep Sea Drilling Project ( D S D P ) sites in the Atlantic Ocean (Dean et al., 1984; Brumsack, 1986 ) also have minor-element contents in excess of a detrital contribution (Fig. 3). Brumsack (1986) attributes the high values to sulfide precipitation under sulfate-reducing conditions, whereas Dean et al. (1984) suggest a multistage process. However, it is possible that primary productivity in the photic zone alone provided the bulk of the minor elements to these sediments. The minorelement: Cu ratios for the marine fractions of Ni, Zn (Fig. 3) and Mo (Fig. 4) resemble plankton values; a single Cd: Cu value is lower than that of its A-seawater value and plankton value (Fig. 3a), but it is within the range of measured plankton values (Fig. la). Dean et al. (1984) give an accumulation rate for Zn in sections of DSDP Core 530 that contain this black shale. They report a maximum accumulation rate of 2.4/tg cm -2 yr -1 Zn and an average of 0.53 #g cm -2 yr - t Zn for those sections containing > 5% black shale. Assuming a Zn concentration in plankton of 110 ppm, the average Zn content of the sediments requires an average accumulation rate for organic matter of 4.8 mg cm -2 yr -t organic matter, the accumulation rate necessary to deliver the total marine fraction of Zn to the sea floor. I use the average rather than the maximum because of possible concentration anomalies developed during diagenesis (Dean et al., 1984). The average represents as little as 2.5% of surface productivity in highly productive areas of the ocean today (Chavez and Barber, 1987 ) and is at the low end

103

of the approximate amount of organic matter accumulating on the sea floor in such areas (Reimers and Suess, 1983). Thus, a moderate rate of primary-productivity in the photic zone, based on current marine values, is adequate to give the high concentrations ofZn, Cd, Cu and Mo in the sediment. The high concentrations reflect a low bulk sediment accumulation rate rather than an unusually high accumulation rate for the individual elements. One difficulty of making such calculations for sediments from many margin areas of the ocean today is the uncertainty in the calculations of the marine fraction of minor elements. The accumulation rate of detritus is often large, as in the Cariaco Trench, or primary productivity is very small, as in the Black Sea. In such cases, the marine fraction can represent only a few percent of the total minor-element inventory.

3. Hydrogenous source

This discussion of the redox conditions of deposition and the problems of recognizing those conditions from the minor-element composition of sediments emphasizes the redox conditions of bottom waters, in contrast to the redox conditions of sediment pore water. Redox in both environments is determined by bacterial respiration via oxygen respiration (02 reduction), followed by denitrification (NO~- reduction), and eventually SO j - reduction, Oxygen respiration: 1 0 C H 2 0 + 1002 + 10CaCO3 ~ 20HCO~- + 10Ca 2 +

(2)

Fig. 3. Relations between Cu and: (a) Cd; (b) Ni; and (c) Zn, in the marine fraction of sediments and sedimentary rocks. Concentrations represent: ( 1 ) the soluble fraction of sediment in the California Borderland samples (Bruland et al., 1974); (2) the fraction in excess of a detrital contribution (using A1203 as an estimate of the detrital fraction ) in the samples from the San Gregorio Formation (Piper, 1991 ), Monterey Formation (Piper and Isaacs, 1994a), Atlantic Ocean black shale (Dean et al., 1984; Brumsack, 1986), Namibian Shelf sediments (Calvert and Price, 1983), Phosphoria Formation (Medrano and Piper, 1992 ), and U.S. Midcontinent Pennsylvanian black shales (Coveney and Glascock, 1989; Hatch and Leventhal, 1992; Schultz and Coveney, 1992); and (3) the bulk sediment fraction of Japanese limestone samples (Aizawa and Akaiwa, 1992). The Pennsylvanian black shales are divided into the Heebner-type (open circles) and Mecca-type (open circles with cross). The arrow in frames (b) and (c) identifies the Soledad Basin sample of the California Borderland. The two lines in each frame are the same as those in Fig. I.

D.Z. Piper / Chemical Geology 114 (1994) 95-114

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Mo ( p p m ) Fig. 4. Relations between Mo and: (a) Cu; and (b) V : C r ratio, in the marine fraction o f sediments and sedimentary rocks. Sources o f data are given in Fig. 3. Lines o f best fit for the four sets o f data in (b) all have a correlation less than 0.5, but all are positive.

Denitrification:

Organic matter, expressed here in its simplest form as CH20, is better represented by the formula:

10CH 20 + 8NO~- + 2CACO3 --, 4N 2 + 12HCO~- + 4 H 2 0 + 2 C a 2+

(3)

Sulfate reduction: 10CH20+ 5SO42- + 5CACO3 ---+ 5HS- + 15HCO~- + 5Ca 2+

(4)

(CH20)106 (NH3) ts(H3PO4)(minor elements) o.ox It is the primary reducing agent in the oceans, and 02, NO3- and SO~- are the most important

D.Z. Piper / Chemical Geology 114 (1994) 95-114

electron acceptors. The hierarchy of these reactions is determined by the reaction yielding the greatest free energy (Froelich et al., 1979). Thus, reaction (2) proceeds until 02 decreases to a concentration at which denitrification yields equal free energy, and so on. As O2 is renewed in bottom water by advection, the type of respiration which occurs within the bottom water of a basin represents a balance between the rate at which organic matter settles out of the photic zone and bottom-water residence time. Several of the minor elements are also reduced under these changing conditions of bacterial respiration (Table 2 ). For example, Cr is present in the oxic environment as CrO42-, but it is reduced to the less soluble 3 + valence state under conditions of denitrification (Murray et al., 1983), possibly by some form of the following reaction: 3 C H 2 0 + 2 H 2 0 + 5Ca 2+ + 2HCO~+ 4CRO42- ~ 4Cr ( O H ) 3 + 5 CaCO3

(5)

However, the contribution of Cr and other minor elements to the oxidation of organic matter is negligible, owing to their low concentrations in seawater (Table 2 ). Cr is present at a concentration of ~ 4.10-6 mmol kg- t; vs. 1.5 mmol kg-1 for 02, 0.04 mmol kg-l for NO3- and 28 mmol kg-l for SO42-. Oceanographic research in recent years has greatly enhanced our understanding of the precipitation of minor elements under the various redox conditions in seawater. The work of Goldberg and Arrhenius (1958) and Mero (1965), which showed the potential economic value of ferromanganese deposits accumulating under oxic conditions of the pelagic environment, was followed by a series of studies defining the distribution of minor elements in open ocean seawater (Boyle et al., 1976, 1977; Sclater et al., 1976; Bruland, 1980, 1983; DeBaar et al., 1988), in pore waters of pelagic sediments (Klinkhammer et al., 1982 ), and in the solid phases of these sediments (Cronan and Tooms, 1969; Calvert and Price, 1977; Elderfield et al., 1981; Piper, 1988 ).

105

This work clearly indicates that the hydrogenous fraction of minor elements constitutes a major component of sedimentary deposits only under conditions of extremely slow bulk-sediment accumulation rates of a few tenths of a mg cm-2 yr - - l A further series of studies of elemental distributions in the Black Sea, the Cariaco Trench (off Venezuela), Framvaren Fjord (Norway), Saanich Inlet (British Columbia) and the Baltic Sea (Degens and Ross, 19'74; Kremling, 1983; Jacobs et al., 1985, 1987; DeBaar et al., 1988; Izdar and Murray, 1989; Emerson and Huested, 1991 ) has shown that sediments accumulating in basins in which sulfate reduction occurs in the bottom waters are a major sink for such elements as Mo (Bertine and Turekian, 1973; Emerson and Huested, 1991), Se (Cutter, 1982; Cutter and Bruland, 1984) and U (Veeh, 1967; Kolodny and Kaplan, 1973; Klinkhammer and Palmer, 1991 ), even though this environment is quite limited in the oceans today. Since Mo is found in seawater at a significantly higher concentration than any other heavy metal (Collier, 1985), its enrichment in these sediments, is the defining diagnostic for the identification of sulfate-reducing conditions in the bottom water at the time of deposition (Jacobs et al., 1987; Emerson ar.d Huested, 1991 ). This is clearly demonstrated by the distribution of Mo in sediment from the Cariaco Trench (Jacobs et al., 1987). Its accumulation rate is relatively high, 0.5/~g c m - 2 y r - ~, and equals its availability, ~ 15% of the bottom water flux (Emerson and Huested, 1991 ). For a bottomwater residence time of 100-200 yr (Jacobs et al., 1987), a sultate-recluclng portion oI the water column of ~ 1000 m, and a seawater Mo content of l0 ppb (Collier, 1985), the flux of seawater Mo into the basin is also ~ 5 - 1 0 / ~ g c m - 2 yr ~. The proximity of the Cariaco Trench to major rivers of northern South America causes a high bulk sediment accumulation rate of ~ 100 mg cm -2 yr -~ (W.E. Dean, pets. commun., 1993 ). As a result of the high influx of terrigenous material, the absolute concentration of Mo in the Cariaco Trench sediment is only ~ 50-100 ppm. A similar calculation for the Mo data from

D.Z. Piper / Chemical Geology 114 (1994) 95-114

106

Table 2 Half-cell reactions assuming standard-state conditions Eh

Half-cell reactions

a~

Seawater concentration (mol kg-') or partial pressure

6

Co2+=2.04.10 -11

[1], [2]

0 2 = 0 . 2 atm

131

(V) 1.099

3Cot,q)+4H20(l)--,CosO4(,)+2e

0.805

2HzOo)- O20o+4e-+4H~',~o

0.704

2+ +2H2Ooj_~MnO2ts)+2e-+4H~-aq ) Mn(aq)

6

Mn2+=2.5.10-1°

[3], [4]

0.702

I(aq)+3HzO(o-*IO~(aq)+6e-+6H~aq)

3 4

l =5.0"10 -9 IO - = 4.5" 10 -"

[1], [3]

0.698

N2~)+6H2Oo)~ 2NO~(~o + 10e-+ 12I-I(*,~

3

NO3-=3.9.10 -s N2~) =0.8 atm

[3], [5]

0.545

Cr(OH)3(~)+H2°l~)~CrO~q)+3e-+5H~q)

4

CrO2- =4.04.10 9

[2], [6], [71

0.448

2-2--- +2H(aq) SeO3(.q)+H2Oo)-*SeO4(aq)+2e

4

~Se=2.22-10 -9

[3], [8], [9]

0.320

3+ ~ CeOz(,)+le Ce(.q)+2H2Oo)

0.296

Fe2~q) + 3H2Oo)~ Fe (OH)3(,) + l e - + 3H~-~q)

0.153

Cu+-Cu2++e -

0.013

UOz(~) + 2HCO~-(aq) ~ UO2 (CO3)2~-~q) + 2e- + 2H ~'~q)

0.025

2+

--

+8H(aq)

Reference ~

+

+

4

SeO~- =0.92-10 9

9

Ce3+=l.0.10 -~

[3], [10]

6

Fe2+ = 1.20.10 -9

[3], [4]

S C u = 3 . 5 9 . 1 0 -9, Cu + = 1.80-10 -9

[1], [3]

4.0 4.5

HCO~- = 2.47" !O -3 UO2(CO3)2- = 1.26" 10 -8

[3], [5], [11]

HSe~-,1) + 3H2Oo) ~ SeO2~q) + 6e- + 7H~-~q)

3.5 4

~ S e = 2 . 2 2 . 1 0 -9 SeO~- = 1.11-10 -9

[8]

-0.018

2+ +3H2Oo)~Fe203(s ) + 2 e - +6H(aq) + 2Fe(~q)

6

see above

-0.040

V204(~) + 4H20(1) ~ 2H2VO£(~q) + 2e- + 4H ~'~q)

5.6

H2VO£ =4.0"10 -8

[2], [13]

-0.055

H S ( ,-, o + 4 H 2 O o ) - ~ S O , ~ K oz-+ S e

3.5

HS-=7.75"10

[31, 15]

4

SO~- =2.8-10 -2

-0.055

+ C u 2 S ( ~ ) + 4 H z O ( l ) ~ S,O 4z-( a q ) + 2 C u (+a q ) + 8 e - +8H(aq)

4 2.5

SO]- =2.8"10 2 Cu+ =3.59"10 -9

[1], [3] [12]

-0.137

C d S ( s ) + 4 H 2 O ( l ) ~ S O a (2 -a q ) + C d t 2+ ~q)+8e

4 6

SO 2 = 2 . 8 - l 0 --2

Cd2+ =6.94" 10 -~°

[1], [3] [5], [14]

--

+ +4H(aq)

- + 9 H ( ,+ o

-

+ +8H(aq)

-zS,

-0.162

22+ - +8H(a.q) + ZnSts)+4H2Oo)~SOa(~q)+Zn(aq)+8e

4

SO~ =2.8.10 ..:2

6

Zn2+ = 5.97" 10-9

-0.170

2MoO2~)+2HzOo)~MoO~(~.q)+2e

4.5

MoO2-=l.I.10

- 0 . I75

2- +2SO4(aq)+24H(~q) 2+ MoSz(~)+I2H2Oo)~MoO~(,q)+I8e

4.5 5

SO42- =2.8"10 MOO2- =1.1"10 -~

[3], [15], [16]

-0.188

' 2- +Nl(aq) .2+ + 8 e - +8H~'~q) NIS(s)+4H20(I)~SO4(aq)

4 6

SO 2 =2.8' 10 -~ Ni2+ =8.01" 10-9

[ 1 ], [3], [5], [12]

-0.209

2- + Fe(aq) 2+ +8e-+8H~-~q) FeS(s) + 4H20(t) ~ SO4(,q)

4

SO]- =2.8-10 -2

[1], [3], [4]

6

Fe 2+ =1.20-10-9

-0.250

ReOz(~)+ 2H2Oo)~ReOxo,q) + 3e- + 4H~-aq)

4.5

ReO£ =4.5.10 - u

-

÷ +4H(a,a)

[1], [3], [5] 7

2

[3]. [15]

[3], [ 17], [18]

Standard free energies of formation (AG~-) are from Latimer ( 1953 ) and Wagman et al. ( 1982 ). Total metal-ion, gas and nonmetal-ion concentrations (mol k g - J ) are those of seawater at ~ 2000-m depth. Activity coefficients were calculated by using the Debye-Hiickel equation; a-values are from Nordstrom and Munoz (1985) or estimated from their table 7-4. The pH was taken as 7, although it is close to 8 under conditions of oxygen respiration and 7.5 under conditions of denitrificaton. References give sources of information on the metal contents in seawater and thermodynamic constants of reactions. Eqs. 2-4 are written as half-cell reactions, in bold type. aValues of a, the effective ionic diameter used in the Debye-Hiickel equation, are listed in the order of their appearance in the chemical equation. bReferences: [1 ]-Bruland (1983); [2J-Latimer (1953); [ 3 ] - W a g m a n et al. (1982); [4]-Landing and Bruland (1987); [5]-Broecker and Peng ( 1982); [6J-Murray et al. ( 1983); [7]-Elderfield ( 1970); [8J-Measures et al. (1983); [ 9 ] - C u t t e r and Bruland ( 1984); [ 10]-DeBaar et al. ( 1985 ); [ 11 ]-Langmuir ( 1978 ); [ 12 ]-Jacobs et al. (1985); [ 13 ]-Collier ( 1984); [ 14 ]-Boyle et al. ( 1976 ); [ 15 ]-Collier ( 1985 ); [ 16 ]-Emerson and Huested ( 1991 ); [ 17 ]-Anbar et al. ( 1992); [ 18]-Koide et al. (1986).

D.Z. Piper / Chemical Geology 114 (1994) 95- I 14

the Black Sea (W.E. Dean, pets. commun., 1993 ) also shows concurrence between availability and accumulation rate, although the Mo accumulation rate is 5-fold less in the Black Sea. Involvement of the entire sulfate-reducing portion of the water column to the accumulation of the hydrogenous fraction of Mo, in both the Cariaco Trench and the Black Sea, suggests that precipitation a n d / o r adsorption onto settling particulate matter, rather than diffusion across the benthic boundary, is the dominant mechanism of sedimentation. The profile of solid Mo in the near-surface sediments further supports this interpretation. It shows no gradient in the upper few tens of centimeters of sediment (Calvert, 1990; W.E. Dean, pets. commun., 1993). By contrast, Mo exhibits a well-defined profile in the Gulf of California sediment, where its accumulation is rather clearly attributable to diffusion across the benthic boundary (Brumsack, 1986 ). The amount of Mo in the Gulf of California sediment is, however, lower than in the Cariaco Trench and Black Sea sediment, in the case of the Cariaco Trench by approximately an order of magnitude. This suggests a much lower accumulation rate by diffusion than by precipitation from the water column. The marine fraction of other minor elements in the Cariaco Trench that might support this interpretation are, unfortunately, only weakly enriched above the detrital fraction, owing to their lower concentrations in seawater (Table 1 ) and their dilution in the sediment by the high accumulation rate of the detrital fraction. Bacterial oxidation of organic matter proceeds by denitrification [reaction (3) ] between the extremes of 02 respiration and sulfate reduction (Froelich et al., 1979 ). It occurs at intermediate depth in the water column of all basins which experience sulfate reduction in the bottom water (Emerson et al., 1979). Denitrification also occurs in the bottom water of Santa Barbara Basin of the southern California Borderland (Sholkoritz and Gieskes, 1971 ), in Darwin Bay in the Gal~ipagos Islands (Richards and Broenkow, 1971 ), and in limited areas of the oxygen minimum zone (OMZ) of the open ocean (Calvert and Price, 1971 ; Codispoti, 1980).

107

Although the distribution of minor elements in this environment has not been so thoroughly examined, it is clear that the rare-earth elements (REE) are removed by adsorption (DeBaar et al., 1985, 1988; German et al., 1993) and returned to solution in the sulfate-reducing environment; as noted above, CrO42- is reduced to the less soluble C r 3+ (Murray et al., 1983); and V should be reduced to a less soluble valence state (Table 2) than that of H2VO£, its valence state under fully oxic conditions (Sadiq, 1988 ). However, the distribution of V in the water column of several basins, in which sulfate reduction occurs in the bottom water, gives somewhat equivocal results (Emerson and Huested, 1991 ). Ideally, the reduction of V occurs at the low-Eh boundary of denitrification, whereas the reduction of Cr occurs at the upper boundary. This difference could result in a highly variable Cr:V ratio within the marine fraction of sediments that accumulate under conditions of bottom-water denitrification. It could be much greater than the seawater value of 0.12 in sediments that accumulate under mildly denitrifying conditions, but approach the seawater value in sediments that accumulate under strongly denitrifying to sulfate-reducing conditions, for example in sediments from the Cariaco Trench. Unfortunately, a normative calculation cannot resolve the marine fractions of V and Cr in these sediments with sufficient accuracy to test this hypothesis, a sediment leaching study may be appropriate. Minor-element distributions under the different redox conditions in the modern environments also demonstrate other problems of any simple equilibrium thermodynamic approach. These problems have been discussed at some length in the studies mentioned and in an earlier paper (Piper and Isaacs, 1994a). They include the lack of information on the actual stable dissolved and particulate phases of several of the elements in seawater and sediments; the adsorption of dissolved species onto organic- and inorganic-particulate phases, rather than dissolution/precipitation as a simple compound in response to Eh properties of seawater; and incorporation of seawater-dissolved species into sediments by diffusion from bottom water, perhaps

108

D.Z. Piper / Chemical Geology 114 (1994) 95-114

independent of bottom-water Eh conditions. Despite these problems, the distributions of several of the minor elements in the different modem marine environments suggest that equilibrium between dissolved and particulate phases in the marine environment is at least approached, if not attained. The distribution of Mo in the water column and sediments of the Black Sea and Cariaco Trench, of Cr in the OMZ of the eastern Pacific Ocean, and of Mn in the pelagic environment support this interpretation. As noted earlier, the marine fraction of other minor elements is often only weakly detected in sediments, above the detrital signal. They might otherwise provide confirmation for this interpretation. Unlike today, however, ancient oceans often inundated the continents, creating shallow marine seas hundreds of kilometers in width. This, alone, limited the availability and accumulation of otherwise diluting terrigenous debris. We might expect sediments from such environments to be more enriched in many of the minor elements.

4. M i n o r - e l e m e n t

contents -- Ancient sediments

The identification of a biological or hydrogenous source for minor elements in marine sedimentary deposits now exposed on land depends on the extent of dilution by the detrital fraction, the retention of the minor elements during early diagenesis, and the extent of alteration during the extremes of burial and subaerial weathering. Clearly, these factors will vary in different formations and for many formations be severe enough to mask the signal of deposition. The effect of dilution by terrigenous debris might be overcome through a leaching procedure (Chester and Hughes, 1967; Tessier et al., 1984; Belzile et al., 1989), or by a normative (Leinen and Pisias, 1984; Piper and Isaacs, 1994a, b) or statistical approach (Dymond, 1981; Calvert, 1990). The effects ofdiagenesis (Sholkovitz, 1973; Shaw et al., 1990) are more difficult to identify and remove, as are the effects of weathering (Carswell and Gulbrandsen, 1955 ). The task is further complicated by the dynam-

ics of modern O2-depleted waters and the coarseness of most sampling schemes of the ancient record. The chemical properties of bottom water change quite dramatically over a period of seasons to years (Sholkovitz and Gieskes, 1971; Anderson and Devol, 1973; Codispoti, 1980) and over longer time periods as well, in response to changes in seawater advection. Such short-term changes surely occurred in the past, and it is doubtful whether any sampling scheme of sedimentary rocks can isolate them. Sampling of ancient sediments inevitably averages ocean conditions over hundreds to thousands of years of deposition. For a sedimentary rock that accumulated at the modest rate of lO mg cm -2 yr-~, a sample of only 1 c m 3 represents several hundred years of deposition. Thus, a single sample might display the geochemical signal of several different conditions, the sum of hundreds of years, rather than the detail of season-to-season and year-to-year events. As a result, an individual sample might contain a record of conflicting bottom-water redox conditions. It might have both high Mo and high REE contents, even though such a co-occurrence is excluded by the geochemistry of these minor elements in modem marine environments. Keeping these difficulties in mind, I consider several deposits. They range in age from Tertiary to Paleozoic. Within the phosphatic-enriched Monterey Formation of California, U.S.A. (Piper and Isaacs, 1994a) and San Gregorio Formation of Baja California, Mexico (Piper, 1991 ), both of Tertiary age, interelement relations for the marine fractions of Cu, Ni and Zn approach the values for plankton and/l-seawater (Fig. 3 ). Cd, although lower in the Monterey Formation than the estimated value for plankton (Brumsack, 1986), is similar to actual plankton measurements (Fig. 1). Also, the accumulation rate of these minor elements (Piper, 1991; Piper and Isaacs, 1994a) required a rate of primary productivity comparable to that on the Peru Shelf, higher than for the Atlantic Ocean Cretaceous shales, but in the range of the modern ocean. Interelement relations and calculations of accumulation rate for Cd, Cu, Zn, and possibly Ni in the Phosphoria Formation, Idaho (Piper and

D.Z. Piper / Chemical Geology 114 (1994) 95-114

Medrano, 1994), of Permian age, are also consistent with the interpretation of a biogenic origin (Fig. 3 ). The marine fractions of Cd, Cu, Ni and Zn in Permian and Triassic limestones of Japan (Aizawa and Akaiwa, 1992) cannot be estimated with certainty. Assuming a mean SiO2 content in detritus of 58.4% (Wedepohl, 1969-1978 ), the SiO2 contents of these limestones limit the detritus content to between 0.04% and 2.5% (Aizawa and Akaiwa, 1992, table 3 ). These are upper limits for detritus, as SiO2 probably had a biogenic source in these limestones, as well as a detrital source. The A1203 contents, which were not reported, provide a more reliable estimate of the detrital fraction. However, assuming Cd, Cu, Ni and Zn concentrations in detritus of 0.2, 35, 44 and 100 ppm, respectively (Wedepohl, 19691978; Piper, 1991; Piper and Isaacs, 1994a), these elements clearly have higher concentrations in the limestones than supplied by detritus (Fig. 3). Aizawa and Akaiwa (1992) determined the acetic acid-soluble and -insoluble fractions (their table 6 ). Unfortunately, this procedure does not separate the marine from the detrital fraction. They report both apatite and dolomite in the insoluble residues of virtually all samples. Biogenic silica is also probably present in the insoluble residues, either as opal-CT or quartz (Isaacs, 1980). All of these marine phases might host a portion of the nondetrital, minor-element inventory of the rocks. They attribute this soluble Cd content to a local hydrothermal source. It seems even more likely that the nondetrital fractions of Cu, Ni and Zn would have had a hydrothermal origin. Sulfides of these three elements represent major fractions of metal-sulfide deposits of the East Pacific Rise (Edmond et al., 1979). They also represent a major fraction of the metal inventory of hydrothermal fluids (Von Damm et al., 1985a, b), unlike Cd which is a minor component. Rather than a local hydrothermal source, the interelement relations for all these elements in the limestones (Fig. 3 ) suggest that marine biogenic debris, which had a minor-element stoichiometry approaching the composition of modern

109

plankton, was the nondetrital source of the elements. The averages of the element: Cu ratios for the 7 limestone formations plot along the extrapolated trends for plankton and A-seawater. These interelement relations demonstrate that the absolute concentration of Cd, or of any other single element, reveals very little about the depositional environment of these and other rocks; its ratio with other minor elements might identify minor-element sources, and its rate of accumulation might offer independent confirmation of those sources. Minor elements in Pennsylvanian metalliferous shales have also been attributed to hydrothermal activity (Coveney and Glascock, 1989) and/or an exotic seawater chemistry (Hatch and Leventhal, 1992 ), within a marine environment of high primary productivity. The interpretation of normal seawater as the source (Holland, 1979 ) seems to have been dismissed by most geologists. However, the composition of these deposits supports the interpretation that seawater, with a minor-element composition closely approaching the composition of modern seawater, was the sole nondetrital source. First, consider the rocks in which Mo is < 100 ppm (Heebner-type shales in Schultz and Coveney, I992). These rocks must have accumulated in the absence, or near absence of sulfate-reducing conditions in the bottom-water. The Cd: Cu ratios (Fig. 3) are close to the values of plankton. The average Zn: Ni: Mo relations of approximately 55: 7: 2 also approximate the plankton values of 110: 7.5 : 2 (Table 2 ), although their ratios with Cu are larger than the values for plankton. On balance, however, these interelement relations suggest accumulation as biogenic debris with a composition of modern plankton. The Mo content in these shales is indeed similar to that of the Cariaco Trench sediments, and the Mo:Cu ratio (Fig. 4a) suggests a slight enrichment of Mo above the plankton value. As bulk sediment accumulation rates probably were several orders of magnitude less than in the Cariaco Trench, i.e. dilution by detritus was small, the diffusive flux from bottom water might have been significant. Alternatively, the bottom water was sulfate reducing. It would have been so,

110

D.Z. Piper / Chemical Geology 114 (I994) 95-114

however, for very brief periods, as shown below. First, consider the evidence for bottom-water denitrification during the deposition of the Heebner-type shales. The Heebner-type shales and the phosphatic formations discussed above all have much higher contents o f V ( 160-1100 ppm) and Cr (360-910 ppm), relative to Cu (Table 1 ), than can be attributed to the accumulation of plankton. Cr and V must have precipitated from seawater under bottom-water conditions of denitrification (Table 2). The Mo contents of these rocks and the shift in their Mo: Cu ratios away from the plankton value toward the seawater value (Fig. 4a) suggest that the environment might have been sulfate reducing, as I noted above, for brief periods of time. Most of the sediment M o : C u values are shifted only slightly toward the value for seawater, given that the mean M o : C u ratio in seawater is approximately 250-fold higher than the ratio in plankton (Table 1 ). Although I have no firm evidence, these slight shifts probably represent diffusion of Mo into the sea floor, rather than sulfide precipitation. The relation between Cr and V would seem to support this interpretation. Cr is much more enriched in the Heebner-type shales and phosphatic formations than is V, when compared to their seawater concentrations (Table 1). This enrichment, expressed as the ratio: [Cr: V]sedimenl [Cr:V] . . . . . t e r is greater in the rocks with < 100 ppm Mo, but virtually one in rocks with higher Mo values (Mecca-type shales). The concurrence between the Cr:V ratios in the Mo-enriched rocks with the ratio in modern seawater (Fig. 4b) is strong evidence of their source. My interpretation is that Cr and V were almost totally removed from seawater, but only under bottom-water conditions sufficient to precipitate major amounts of Mo, i.e. under bottom-water conditions that were sulfate reducing [reaction (4)] for a considerable period of time. It is not surprising that the Mo content in the Mecca-type shales reaches values of 2260 ppm; it could be much higher. Had the M o : V ratio of seawater (10:1.8) been con-

served, as was the Cr:V ratio, the Mo content would have been > 10,000 ppm. Clearly, the bottom water was not sulfate reducing during the total time represented by the accumulation of even the Mecca-type shales. The "considerable-period-of-time" probably represented <20% of their depositional history. Information on sediment accumulation rates and detailed sampling might allow us to examine the relative duration of bottom-water sulfate-reducing conditions, residence time of the bottom water, primary productivity in the photic zone of the water column, thickness of that portion of the water column that was sulfate reducing, and accumulation rate of the detrital fraction, the approach taken to explain the minor-element content of sediment from the Sea of Japan (Piper and Isaacs, 1994b). But sediment accumulation rates cannot always be determined, and the ability to resolve short-term oceanographic events in the ancient record is limited because of post-depositional alteration and relative coarseness of sampling.

5. Conclusions

I have emphasized the importance of settling organic matter to the accumulation of several minor elements in marine sediments and sedimentary rocks. Planktonic organic matter carries minor elements to the sea floor directly, but it also drives redox reactions which determine the suite of elements that precipitates from bottom water. Thus, the rate of primary productivity in the photic zone determines the rate of settling organic matter through the water column and it, in turn and in conjunction with the rate of advection throughout the water column, determines the type of bacterial respiration in the bottom water. Thus, the biologic cycle is the single dominating influence to the minor-element composition of the marine fraction in these and virtually all sedimentary rocks with high contents of marine phases such as organic matter, apatite, silica, calcite and dolomite. In the modern ocean O2-depleted environments have a very limited extent. If this was true

D.Z. Piper / Chemical Geology 114 (1994) 95-114

in ancient oceans, the accumulation of biogenic debris on ancient sea floors becomes a more compelling explanation for the high contents of most minor elements in sedimentary rocks than adsorption and/or precipitation reactions. This is particularly true for sediments that accumulated in the relatively shallow environments of continental seas and ocean margins, where today high primary productivity is the rule and where a relatively large fraction of the organic matter produced in the photic zone accumulates on the sea floor.

Acknowledgments This paper was considerably improved by suggestions made on an early draft by C. Isaacs and G. Briet and two anonymous reviewers. Gail Piper proofed the penultimate draft, but I alone am responsible for the interpretations and conclusions, however they are received.

References Aizawa, S. and Akaiwa, H., 1992. Cadmium contents of Triassic and Permian limestones in central Japan. Chem. Geol., 98:103-110. Anbar, A.D., Creaser, R.A., Papanastassiou, D.A. and Wasserburg, G.J., 1992. Rhenium in seawater: confirmation of a generally conservative behavior. Geochim. Cosmochim. Acta, 56: 4099-4103. Anderson, J.J. and Devol, A.H., 1973. Deep water renewal in Saanich Inlet, an intermittently anoxic basin. Estuarine Coastal Mar. Sci., 1: 1-10. Arrhenius, G. and Bonatti, E., 1965. Neptunism and vulcanism in the ocean. In: M. Sears (Editor), Progresses in Oceanography, Vol. 3. Pergamon, London, pp. 7-22. Balistrieri, L.S., Brewer, P.G. and Murray, J.W., 1981. Scavenging residence times of trace metals and surface chemistry of sinking particles in the deep ocean. Deep-Sea Res., 28: 101-121. Belzile, N., Lecomte, P. and Tessier, A., 1989. Testing readsorption of trace elements during partial chemical extractions of bottom sediments. Environ. Sci. Technol., 23: 1015-1020. Berger, W.H., Fischer, K., Lai, C. and Wu, G., 1988. Ocean carbon flux: global maps of primary production and export production. In: C.R. Agegian (Editor), Biogeochemical Cycling and Fluxes between the Deep Euphoric Zone and Other Oceanic Realms. NOAA (Natl. Ocean. Atmos.

111

Adm.) Under Sea Res. Prog., Rep. 88-1, pp. 131-176. Bertine, K.K. and Turekian, K.K., 1973. Mo in marine deposits. Geochim. Cosmochim. Acta, 37:1415-1434. Boyle, E.A., Sclater, F. and Edmond, J.M., 1976. On the marine geochemistry of cadmium. Nature (London), 263: 42-44. Boyle, E.A., Sclater, F.R. and Edmond, J.M., 1977. The distribution of dissolved copper in the Pacific. Earth Planet. Sci. Lett., 37: 38-54. Broecker, W.S. and Peng, T-El., 1982. Tracers in the Sea. EIdigio, Palisades, N.Y., 690 pp. Bruland, K.W., 1980. Oceanographic distribution of cadmium, zinc, nickel, and copper, in the North Pacific. Earth Planet. Sci. Lett., 47: 176-.198. Bruland, K.W., 1983. Trace elements in seawater. In: J.P. Riley and R. Chester (Editors), Chemical Oceanography, Vol. 8. Academic Press, London, pp. 157-220. Bruland, K.W., Bertine, K., Koide, M. and Goldberg, E.D., 1974. History of metal pollution in southern California coastal zone. Environ. Sci. Technol., 8: 425-432. Brumsaek, H.J., 1986. The inorganic geochemistry of Cretaceous black shales (DSDP Leg 41) in comparison to modern upwelling sediments from the Gulf of California. In: C.P. Summerhayes and N.J. Shackleton (Editors), North Atlantic Paleoceanography. Geol. Soc. London, Spec. Publ., 21: 447-462. Calvert, S.E., 1990. Geochemistry and origin of the Holocene sapropel in the Black Sea. In: V. Ittekkot et al. (Editors), Facets of Modern Biogeochemistry, Festschrift fiir E.T. Degens. Springer, Berlin, pp. 326-352. Calvert, S.E. and Price, N.B., 1971. Upwelling and nutrient regeneration in the Benguela Current, October, 1968. Deep-Sea Res., 18: 505-5123. Calvert, S.E. and Price, N.B.. 1977. Geochemical variations in ferromanganese nodules and associated sediment from the Pacific Ocean. Mar. Chem., 5: 43-74. Calvert, S.E. and Price, N.B.. 1983. Geochemistry of Namibian shelf sediments. In: E. Suess and J. Thiede (Editors), Coastal Upwelling, Its Sediment Record. NATO (N. Atlantic Treaty Org.) Conf. Set. Mar. Sci., Vol. IV-10a. Plenum, New York, N.Y., pp. 337-375. Calvert, S.E., Karlin, R.E., Toolin, L.J., Donahue, D.J., Southons, J.R. and Vogel, J.S, 1991. Low organic carbon accumulation rates in Black Sea sediments. Nature (London), 350: 692-695. Carswell, L.D. and Gulbrandsen, R.A., 1955. Uranium in phosphate - - Northwest phosphate variations in composition of phosphate rock with depth in the Conda Mine (Idaho). In: Geologic Investigations of Radioactive Deposits, Semiannual Progress Report. U.S. Geol. Surv., Trace Element Invest. Rep., TEI 540:192. Chavez, F.P. and Barber, R.T., 1987. An estimate of new production in the equatorial Pacific. Deep-Sea Res., 34: 12291244. Chester, R. and Hughes, MJ., 1967. A chemical technique for the separation of ferromanganese nodules, carbonate

112

D.Z. Piper / Chemical Geology 114 (I 994) 95-114

minerals, and absorbed trace elements from pelagic sediment. Chem. Geol., 2: 249-262. Clegg, S.L. and Sarmiento, J., 1989. The hydrolytic scavenging of metal ions by marine particulate matter. Prog. Oceanogr., 23:1-21. Codispoti, L.A., 1980. Temporal nutrient variability in three different upwelling regimes. In: F.A. Richards (Editor), Coastal Upwelling. Am. Geophys. Union, Washington, D.C., pp. 209-220. Codispoti, L.A., 1989. Phosphorus vs. nitrogen limitation of new and export production. In: W.H. Berger, et al. (Editors), Productivity of the Oceans: Present and Past. Wiley, New York, N.Y., pp. 377-394. Collier, R.W., 1984. Particulate and dissolved vanadium in the North Pacific Ocean. Nature (London), 309:441-444. Collier, R.W., 1985. Molybdenum in the northeast Pacific Ocean. Limnol. Oceanogr., 30:1351-1354. Collier, R. and Edmond, J., 1984. The trace element geochemistry of marine biogenic particulate matter. Prog. Oceanogr., 13:113-199. Corliss, J.B., Dymond, J., Gordon, E.I., Edmond, J.M., Von Herzen, R.P., Ballard, R.D., Green, K., Williams, D., Bainbridge, A., Crane, K. and Van Andel, Tj.H., 1979. Submarine thermal springs on the Galapagos Rift. Science, 203: 1073-1083. Coveney, Jr., R.M. and Glascock, M.D., 1989. A review of the origins of metal-rich Pennsylvanian black shales, central U.S.A., with an inferred role for basinal brines. Appl. Geochem., 4: 347-367. Cronan, D.S. and Tooms, J.S., 1969. The geochemistry of manganese nodules and associated pelagic deposits from the Pacific and Indian Oceans. Deep-Sea Res., 16: 335339. Cutter, G.A., 1982. Selenium in reducing waters. Science, 217: 829-831. Cutter, G.A. and Bruland, K.W., 1984. The marine biogeochemistry of selenium: A re-evaluation. Limnol. Oceanogr., 29:1179-1192. Dean, W.E., Arthur, M.A. and Stow, D.A.V., 1984. Origin and geochemistry of Cretaceous deep-sea black shales and multicolored claystones, with emphasis on Deep Sea Drilling Project Site 530, southern Angola Basin. In: W.W. Hay, J.C. Sibuet, et al. (Editors), Initial Reports of the Deep Sea Drilling Project, Vol. 75. U.S. Gov. Print. Off., Washington, D.C., pp. 819-844. DeBaar, H.J.W., Bacon, M.P., Brewer, P.G. and Bruland, K.W., 1985. Rare earth elements in the Pacific and Atlantic Oceans. Geochim. Cosmochim. Acta, 49:1943-1959. DeBaar, H.J.W., German, C.R., Elderfield, H. and van Gaans, P., 1988. Rare earth element distributions in anoxic waters of the Cariaco Trench. Geochim. Cosmochim. Acta, 52: 1203-1219. Degens, E.T. and Ross, D.A., 1969. Hot Brines and Recent Heavy Metal Deposits in the Red Sea. Springer, New York, N.Y.. 600 pp. Degens, E.T. and Ross, D.A., 1974. The Black Sea - - Geol-

ogy, Chemistry, and Biology. Am. Assoc. Pet. Geol. Mem., 20: 633pp. Dickens, G.R. and Owen, R.M., 1993. Global change in manganese deposition at the Cenomanian-Turonian boundary. Mar. Geores. Geotech., 11: 27-43. Dugdale, R.C. and Goering, J.J., 1970. Nutrient limitation and the path of nitrogen in Peru Current production. Anton Bruun Rep., 5: 3-8. Dymond, J., 1981. Geochemistry of Nazca Plate surface sediment; an evaluation of hydrothermal, biogenic, detrital, and hydrogenous sources. In: L.D. Kulm, et al. (Editors), Nazca Plate: Crustal Formation and Andean Convergence. Geol. Soc. Am. Mem., 154: 133-173. Edmond, J.M., Measures, C.I., Mangum, B., Grant, B., Sclater, F.R., Collier, R., Hudson, A., Gordon, C.I. and Corliss, J.B., 1979. On the formation of metal-rich deposits at ridge crests. Earth Planet. Sci. Lett., 46:19-30. Eisler, R., 1981. Trace Metal Concentrations in Marine Organisms. Pergamon, London, 687 pp. Elderfield, H., 1970. Chromium speciation in seawater. Earth Planet. Sci. Lett., 9: 10-16. Elderfield, H., Hawkesworth, C.J., Greaves, M.J. and Calvert, S.E., 1981. Rare earth element geochemistry of oceanic ferromanganese nodules and associated sediments. Geochim. Cosmochim. Acta, 45:513-528. Emerson, S.R. and Huested, S.S., 1991. Ocean anoxia and the concentrations of molybdenum and vanadium in seawater. Mar. Chem., 34: 177-196. Emerson, S., Cranston, R.E. and Liss, P.S., 1979. Redox species in a reducing t~ord: equilibrium and kinetic observations. Deep-Sea Res., 26A: 859-878. Fischer, K., Dymond, J., Lyle, M., Soutar, A. and Rau, S., 1986. The benthic cycle of copper: evidence from sediment trap experiments in the eastern tropical North Pacific Ocean. Geochim. Cosmochim. Acta, 50:1535-1543. Fisher, N.S. and Wente, M., 1993. The release of trace elements by dying marine phytoplankton. Deep-Sea Res., 40: 671-694. Froelich, P.N., Klinkhammer, G.P., Bender, M.L., Luedtke, N.A., Heath, G.R., Cullen, D. and Dauphin, P., 1979. Early oxidation of organic matter in pelagic sediments of the eastern equatorial Atlantic, suboxic diagenesis. Geochim. Cosmochim. Acta, 43:1075-1090. Froelich, P.N., Arthur, M.A., Burnett, W.C., Deakin, N., Hensley, V., Jahnke, R., Kaul, L., Kim, K.-H., Roe, K., Soutar, A. and Vathakanon, C., 1988. Early diagenesis of organic matter Peru continental margin sediments: phosphorite precipitation. Mar. Geol., 80: 309-343. German, C.R., Holliday, B.P. and Elderfield, H., 1993. Reply to the comment by E.R. Sholkovitz on "Recycling of rare earth elements in the suboxic zone of the Black Sea". Geochim. Cosmochim. Acta, 56:4309-4313. Goldberg, E.D., 1954. Chemical scavengers of the sea. J. Geol., 62: 249-265. Goldberg, E.D., 1963a. Mineralogy and chemistry of marine sedimentation. In: F.P. Shepard (Editor), Submarine Geology. Harper & Row, New York, N.Y., pp. 436-466.

D.Z. Piper / Chemical Geology 1I4 (1994) 95-114 Goldberg, E.D., 1963b. The oceans as a chemical system. In: M.N. Hill (Editor), The Sea, Vol. II. Wiley, New York, N.Y., pp. 3-25. Goldberg, E.D. and Arrhenius, G.O.S., 1958. Chemistry of Pacific pelagic sediments. Geochim. Cosmochim. Acta, 13: 153-212. Goodfellow, W.D., 1987. Anoxic stratified oceans as a source of sulfur in sediment-hosted stratiform Zn-Pb deposits (Selwyn Basin, Yukon, Canada). In: N. Clauer and S. Chaudhuri (Editors), Isotopes in the Sedimentary Cycle. Chem. Geol. (Isot. Geosci. Sect.), 65:359-382 (special issue). Hatch, J.R. and Leventhal, J.S., 1992. Relation between inferred redox potential of the depositional environment and geochemistry of the Upper Pennsylvanian (Missourian) Stark Shale Member of the Dennis Limestone, Wabaunsee County, Kansas, U.S.A. In: P.A. Meyers, L.M. Pratt and B. Nagy (Guest-Editors), Geochemistry of Metalliferous Black Shales. Chem. Geol., 99:65-82 (special issue). Holland, H.D., 1979. Metals in black shales - - a reassessment. Econ. Geol., 74: 1676-1680. Hulbert, L.J., Came, R.C., Gr6goire, D.C. and Paktunc, D., 1992. Sedimentary nickel, zinc, and platinum-group element mineralization in Devonian black shales at the Nick property, Yukon, Canada: A new deposit type. Explor. Min. Geol., 1: 39-62. Isaacs, C.M., 1980. Diagenesis in the Monterey Formation examined laterally along the coast near Santa Barbara, California. Ph.D. Thesis, Standford University, Stanford, Calif., 329 pp. Izdar, E. and Murray, J.W., 1989. Black Sea Oceanography. NATO (N. Atlantic Treaty Org.) ASI (Adv. Stud. Inst.) Ser. 351,487 pp. Jacobs, L., Emerson, S. and Skei, J., 1985. Partitioning and transport of metals across the 02/H25 interface in a permanently anoxic basin: Framvaren Fjord, Norway. Geochim. Cosmochim. Acta, 49: 1433-1444. Jacobs, L., Emerson, S. and Huested, S.S., 1987. Trace metal chemistry in the Cariaco Trench. Deep-Sea Res., 34: 965981. Klinkhammer, G. and Palmer, M.R., 1991. Uranium in the oceans: Where it goes and why. Geochim. Cosmochim. Acta, 55: 1799-1806. Klinkhammer, G.P., Heggie, D.T. and Graham, D.W., 1982. Metal diagenesis in oxic marine sediments. Earth Planet. Sci Lett., 61: 211-219. Koide, M., Hodge, V.F., Yang, J.S., Stallard, M., Goldberg, E.D., Calhoun, J. and Bertine, K.K., 1986. Some comparative marine chemistries of rhenium, gold, silver and molybdenum. Appl. Geochem., l: 705-714. Kolodny, Y. and Kaplan, I.R., 1973. Deposition of uranium in the sediment and interstitial water of an anoxic l]ord. In: E. Ingerson (Editor), Proceedings of the Symposium on Hydrogeochemistry and Biogeochemistry, Vol. 1. Clark, Washington, D.C., pp. 418-442. Kremling, K., 1983. The behavior ofZn, Cd, Cu, Ni, Co, Fe,

113

and Mn in anoxic Baltic waters. Mar. Chem., 13:87-108. Kremling, K. and Streu, P., 1993. Saharan dust influenced trace element fluxes in deep North Atlantic subtropical waters. Deep-Sea Res., 40:1155-1168. Landing, W.M. and Bruland, K.M., 1987. The contrasting biogeochemistry of iron and manganese in the Pacific Ocean. Geochim. Cosmochim. Acta, 51: 29-44. Langmuir, D., 1978. Uranium solution-mineral equilibria at low temperatures with applications to sedimentary ore deposits. Geochim. Cosmochim. Acta, 42: 547-569. Latimer, W.M., 1953. Oxidation Potentials, Prentice Hall, New York, N.Y., 392 pp. Leinen, M. and Pisias, N., 1984. An objective technique for determining endmember composition and for partitioning sediments according to their sources. Geochim. Cosmochim. Acta, 48: 47-62. Li, Y.H., 1981. Ultimate removal mechanisms of elements from the oceans. Geochim. Cosmochim. Acta, 45: 16591664. MacIntyre, D.G., 1992. Geological setting and genesis of sediment exhalative barite and barite-sulfide deposits, Gataga district, northeastern British Columbia. Explor. Min. Geol., 1: 1-20. Martin, J.H. and Knauer, G.A., 1973. The elemental composition of plankton. Geochim. Cosmochim. Acta, 37: 1639-1653. Marlin, J.-M. and Maybeck, M., 1979. Elemental mass-balance of material carried by major world rivers. Mar. Chem., 7: 173-206. McElroy, M.B., 1983. Marine biological controls on atmospheric CO2 and climate. Nature (London), 302: 328329. McKelvey, V.E., Williams, JS., Sheldon, R.P., Cressman, E.R., Cheney, T.M. and Swanson, R.W., 1959. The Phosphoria, Park City and Shedhorn Formations in the western phosphate field. U.S. Geol. Surv., Prof. Pap., 313: A 1A47. Measures, C.I., Grant, B.C., Mangum, B.J. and Edmond, J.M., 1983. The relation of the distribution of selenium IV and VI in three oceans to physical and biological processes. In: C.S. Wong, E. Boyle, K. Bruland, J.D. Burton and E.D. Goldberg (Editors), Trace Metals in Seawater. NATO (N. Atlantic Treaty Org.) Conf. Ser., Mar. Sci., Vol. VI-9. Plenum, New York, N.Y., pp. 73-83. Medrano, M.A. and Piper, D.Z., 1992. A normative-calculation procedure used to determine mineral abundances in rocks from the Montpelier Canyon Section of the Phosphoria Formation, Idaho: A tool in deciphering the minor-element geochemistry of sedimentary rocks. U.S. Geol. Surv., Bull. 2023, 23 pp. Mero, J., 1965. Marine Mineral Resources of the Sea. Elsevier, Amsterdam, 312 pp. Murray, J.W., Spell, B. and Paul B., 1983. The contrasting geochemistry of manganese and chromium in the eastern tropical Pacific Ocean, In: C.S. Wong, E. Boyle, K. Bruland, J.D. Burton and E.D. Goldberg (Editors), Trace Metals in Sea Water. NATO (N. Atlantic Treaty Org.)

114

D.Z. Piper / Chemical Geology 114 (1994) 95-114

Conf. Ser.. Mar. Sci., IV-9. Plenum, New York, N.Y., pp. 643-669. Nelson, D.M., Goering, J.J. and Boisseau, D.W., 1981. Consumption and regeneration of silicic acid in three coastal upwelling systems. In: F.A. Richards (Editor), Coastal Upwelling. Am. Geophys. Union, Washington, D.C., pp. 242-256. Nordstrom, D.K. and Munoz, J.L., 1985. Geochemical Thermodynamics. Benjamin-Cummings, Menlo Park, N.J., 477 pp. Piper, D.Z., 1988. The metal oxide fraction of pelagic sediment in the equatorial North Pacific Ocean: A source of metals in ferromanganese nodules. Geochim. Cosmochim. Acta, 52: 2127-2145. Piper, D.Z., 1991. Geochemistry of a Tertiary phosphate deposit: Baja California Sur, Mexico. Chem. Geol., 92: 283316. Piper, D.Z. and Isaacs, C.M., 1994a. Geochemistry of minor elements in the Monterey Formation, CA: Seawater chemistry of deposition. In: C.M. Isaacs and J. Rullk~tter (Editors), Analogy of a Source Rock: The Cooperative Monterey Organic Geochemistry Study. Am. Assoc. Pet. Geol. Stud. Geol. (in press). Piper, D.Z. and Isaacs, C.M., 1994b. Minor elements in Quaternary sediment from the Sea of Japan: A record of surface-water productivity and intermediate-water redox conditions. Geol. Soc. Am. Bull. (submitted). Piper, D.Z. and Medrano, M.D., 1994. The geochemistry of the Phosphoria Formation at Montpelier Canyon, Idaho: The environment of deposition. U.S. Geol. Surv., Bull. 2023B (in press). Redfield, A.C., Ketchum, B.H. and Richards, F.A., 1963. The influence of organisms on the composition of seawater. In: M.N. Hill (Editor), The Sea, Vol. II. Wiley, New York, N.Y., pp. 26-77. Reimers, C.E. and Suess, E., 1983. Spatial and temporal patterns of organic matter accumulation on the Peru continental margin. In: E. Suess and J. Thiede (Editors), Coastal Upwelling, Its Sediment Record. NATO (N. Atlantic Treat Org.) Conf. Ser., Mar. Sci., Vol. IV-10b. Plenum, New York, N.Y., pp. 311-345. Richards, F.A. and Broenkow, W.W., 1971. Chemical changes including nitrate reduction in Darwin Bay, Gal~ipagos Archipelago, over a two month period, 1969. Limnol. Oceanogr., 126: 758-765. Sadiq, M., 1988. Thermodynamic solubility relationships of inorganic vanadium in the marine environment. Mar. Chem., 23: 87-96.

Schultz, R.B. and Coveney, Jr., R.M., 1992. Time-dependent changes for midcontinent Pennsylvanian black shales, U.S.A. In: P.A. Meyers, LM. Pratt and B. Nagy (GuestEditors), Geochemistry of Metalliferous Black Shales. Chem. Geol., 99:83-100 (special issue ). Sclater, F.R., Boyle, E. and Edmond, J.M., 1976. On the marine geochemistry of nickel. Earth Planet. Sci. Lett., 31: 119-125. Shaw, T.J., Gieskes, J.M. and Jahnke, R.A., 1990. Early diagenesis in differing depositional environments: The response of transition metals in pore water. Geochim. Cosmochim. Acta, 54: 1233-1246. Sherrell, R.M., 1989. The tracer metal geochemistry of suspended oceanic particulate matter. Ph.D. Dissertation, Mass. Inst. Technol., WHOI (Woods Hole Oceanogr. Inst. ), Woods Hole, Mass., 211 pp. Sholkovitz, E.R., 1973. Interstitial water chemistry of Santa Barbara Basin sediments. Geochim. Cosmochim. Acta, 37: 2043-2073. Sholkovitz, E.R. and Gieskes, J.M., 1971. A physical-chemical study of the flushing of the Santa Barbara Basin. Limnol. Oceanogr., 16: 479-489. Sverdrup, H.U., Johnson, M.W. and Fleming, R.H., 1942. The Oceans. Prentice-Hall, Englewood Cliffs, N.J., 1087 PP. Tessier, A., Campbell, P.G.C., Auclair, J.C. and Bisson, M., 1984. Relations between the partitioning of trace metals in sediments and their accumulation in the tissues of the freshwater mollusk Elliptio complanata in a mining area. Can. J. Fish. Aquat. Sci., 41: 1463-1472. Veeh, H.H., 1967. Deposition of uranium from the ocean. Earth Planet. Sci. Lett., 3:145-150. Von Damm, K.L., Edmond, J.M., Grant, B., Measures, C.I., Walden, B. and Weiss, R.F., 1985a. Chemistry of submarine hydrothermal solutions at 21 °N, East Pacific Rise. Geochim. Cosmochim. Acta, 49: 2197-2220. Von Damm, K.L., Edmond, J.M. and Grant, B., 1985b. Chemistry of submarine hydrothermal solutions at Guaymas Basin, Gulf of California. Geochim. Cosmochim. Acta, 49: 2221-2237. Wagman, D.D., Evans, W.H., Parker, V.B., Schumm, R.H., Halow, I., Bailey, S.M., Churney, K.L. and Nutall, R.L., 1982. The NBS tables of chemical thermodynamic properties. J. Phys. Chem., Ref. Data: Suppl. 2. Wedepohl, K.H. (Editor), 1969-1978. Handbook of G~ochemistry, Vols. I-IV. Springer, Berlin.