Late Minoan Jar ca. 1450-1400 bce




Lawrence Berkeley Laboratory University of California Berkeley, CA 94720


Office of Naval Research, Asian Office 7-23-17 Roppongi, Minato-ku Tokyo 106, Japan

Accepted March 30, 1993 Modified for the WWW April 1997 After: Chemical Geology, vol. 113, p. 297-317 (1994)
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The black shale facies contains several discrete chemical groupings of manganese, iron and vanadium in visually similar rocks. These groups are related to four thermodynamic zones based on changing pE and pH, are identified from a collection of about 200 Paleozoic and Mesozoic finely-laminated, low-calcic {[Ca] < 0.4%} black shales. The low-calcic nature of this data set permits the simplifying assumption that CaCO3 is not present as a significant complicating factor. The relative variations in the concentrations of Mn and Fe can be used as indicators of pE-pH conditions. In group 1 (oxic conditions: high pE), Mn and Fe occur as Fe(III) and Mn(III, IV) oxides and appear in relatively high concentrations (average concentrations 1300 and 56000 ppm respectively). In group 2 (anoxic: nitrate-reducing to sulfate-reducing at intermediate pH), Mn is reduced to Mn++ as is evidenced by low Mn concentration (310 ppm average); in group 2, Fe remains bound as oxides or sulfides with concentrations comparable to those in group 1 (52000 ppm). In group 3, both Mn and Fe are reduced and relatively soluble reflecting anoxic, but non-sulfate reducing pE, at intermediate to low pH. This solubility and Mn and Fe species instability is shown in lower average concentrations 170 and 23000 ppm respectively in group 3 sediments. Group 4 black shales have V concentrations > 300 ppm, averaging 1500 ppm and indicating deposition under conditions of high organic preservation. Mn and Fe concentrations are low - the Mn concentration is less than 260 ppm and averages 80 ppm; Fe is less than 36000 and averages 19000 ppm. The high V concentration with the relatively low Mn and Fe content suggests low pH and possible methanogenesis at low pE. This study demonstrates the wide variety of black shale depositional environments from oxic to methanic. Accordingly, black shales can not all be assigned to anoxic, sulfidic environments without additional chemical analysis and interpretation.

1. Introduction

Black shales and the black shale facies, visually identified by color, are generally considered to record anoxic sedimentary environments (Pettijohn, 1949; Vine and Tourtelot, 1970). Black sediments are deposited under a wide diversity of conditions: fresh water to estuarine to marine, incorporating varying organic productivity and sedimentary pH. Even under modern conditions, anoxic sediments may be formed (1) under oxic bottom water where the organic content is sufficient to deplete overlying oxygen to produce anoxic conditions in the sediments or (2) where anoxia already exists in the water column. These two conditions may be distinguished by the presence (oxic bottom water, >0.2 mL/L O2) or absence of bioturbation (anoxic bottom water, < 0.2 mL/L O2) (Rhoades and Morse, 1971). The chemistry of oxic and anoxic waters differs vastly (Richards, 1965), particularly with respect to trace metal content (Bacon et al., 1980; Jacobs and Emerson, 1982; Jacobs et al., 1985, 1987; Landing and Bruland, 1987). Therefore, the geochemistry of unbioturbated black shales should differ significantly from those formed under more oxic conditions. Under modern, well-oxygenated estuarine and marine waters, unbioturbated anoxic sediments are rare (Rhoades and Morse, 1971; Byers, 1977; Savrda and Bottjer, 1987), however, unbioturbated sediments (i.e. graptolite facies) are common in the Paleozoic (Berry and Wilde, 1978) and at certain times in the Mesozoic (Hallam, 1987). This demonstrates that wide-spread oceanic anoxia was common during previous warmer climates. The importance of widespread oceanic anoxia during these times is supported by Ce anomaly studies of Wright et al. (1987), and Wang et al. (1986). Accordingly, the geochemical characteristics of black shales may record variations in environmental conditions of deposition.

The black to green color of "black" shales can be caused by the presence of common microcrystalline sulfide compounds such as FeS (black) and FeS2 (green) (Berner, 1981a), which indicate reducing conditions in the sediment. However, black minerals, such as Mn oxides indicative of Mn(IV) conditions, also can impart a dark color to shales. Accordingly, black color is not always an indicator of reducing conditions. Additionally, black shales have been found to contain widely varying concentrations of organic matter. However, organic matter concentration or C/S ratios in black shales are not necessarily indicative of either oceanic or early deposition chemical environment. Berner and Raiswell (1983, Raiswell and Berner, 1986) use carbon/sulfur as environmental indicators. However, as they point out, techniques for determining C and S require "fresh," unweathered samples, such as core material, to yield reliable results due to the mobility of C and S. As such, samples are difficult to find in the field and are limited to recent samples; most paleontologically-dated material from well-documented collections are not acceptable for such C/S analyses. Raiswell and Berner (1987) state that the carbon content of rocks from the Paleozoic and Mesozoic may be indicative of age or maturity, rather than depositional setting. Recent studies have suggested that even in the Cretaceous, preservation of organic matter is not necessarily indicative of anoxic conditions (Pederson and Calvert, 1990). Thus, lithologically-described black shales can be deposited under a variety of environmental conditions, which must be identified through means rather than simple color or even the preservation of organic matter. The ambient redox conditions are important in determining thermodynamic stability fields (Garrels and Christ, 1965, Stumm and Morgan, 1970, Lee and Sillén, 1959) during deposition and early diagenesis (Berner, 1981a,b). However, pH is also important, particularly in sediments where the pH buffer systems operative in sea water may be modified or superseded by diagenetic processes. In particular, open ocean black shales with low rates of deposition offer an opportunity to determine the paleo-chemical environment during deposition and early diagenesis. During slow deposition, such sediments are more likely to be in equilibrium with the chemistry of the overlying waters. Upon continued deposition, compaction and early diagenesis, fine grained materials develop low porosity and permeability so that the remaining chemical signature is basically preserved. Subsequent geologic processes producing either thermal or chemical metamorphism or metasomatism can modify the chemical signature to some degree (Haack et al., 1984). However, such changes should be connected with individual or groups of newly mobile elements and may be seen as modifications of the initial chemical pattern set at the end of diagenesis. Considering these factors, the chemistry of the mobile elements in the shale should be a valid indicator of ambient pE-pH conditions.

The bulk chemical composition of shales reflects (1) their initial (pre-erosional) source rocks for minerals stable in the depositional environment (Wood, 1980; Wood et al., 1979; Thompson et al., 1980; Quinby-Hunt et al., 1989b, 1991); (2) the chemistry of the depositional environment as a function of early diagenetic processes and (3) post-lithofaction processes such as metamorphism, hydrothermal alteration, etc. Only a large data set from a wide variety of well-characterized black shales, analyzed using modern, readily comparable analyses could assure that the data were comparable, and that the many different depositional environments resulting in marine black shales were represented. We have examined the first factor, immobile elements, in Quinby-Hunt and Wilde (1991). The following is an attempt to discern the chemistry of the depositional environment from mobile elements using a thermodynamic model as the indicator of the ambient oceanographic conditions during deposition. This approach is an extension of that used by Krumbein and Garrels (1952) where pE-pH fences bound specific natural environments. We have selected a series of finely-laminated, non-bioturbated, low-calcic marine black shales to determine if differing environmental conditions, as a function of redox conditions, plate-tectonic depositional setting (Hay, 1981; Dickinson et al., 1983), organic productivity, etc. can be discerned in the black shales from the concentration of major mobile elements. For open-ocean conditions, use of low-calcic black shales limits the interference of CaCO3 and its properties. The thermodynamic properties of both Mn and Fe carbonates were considered. Interpretation of the calcareous black shale facies will be the subject of another paper. (Quinby-Hunt and Wilde, 1996)

2. Black Shale Samples

Samples used in this study (Table 1), described in detail in Quinby-Hunt et al. (1989a), are predominantly Early Paleozoic fossiliferous shales. Included as well, however, are dark shales from British Carboniferous sequences and Jurassic (mostly Liassic) black shales representing later major black shale depositional intervals. Our total collection contained some 400 field-identified "black" shales. For this report, we limited the discussion to samples containing less than 1% CaCO3 and greater than 50000 ppm Al. This study specifically excluded shaley limestones or limy shales (Pettijohn, 1949, p.291), in order to focus on the composition of the argillaceous black shale facies as noted above.

3. Procedures 3.1 Chemical Analysis

The concentration of a suite of about 40 elements was determined at Los Alamos National Laboratory, using automated neutron activation analysis (Minor et al. 1982). The advantages of using a single technique and single analytical laboratory and our methods of data handling, detection limits are described in Quinby-Hunt et al., (1989a,b).

3.2. Statistical Methods

The frequency distribution of the concentrations of each of the elements was determined (Quinby-Hunt et al., 1989a). The frequency distributions, as percent, were plotted versus the midpoints of distribution bins used in determining the distribution.

A number of elements: Na, Cl, Ca, Mn, Fe, V, Co, Zn, As, Se, Br, Mo, In, Sb, Ba, W, and U (Quinby-Hunt et al., 1989a) showed multimodal distributions suggesting mobility under the various depositional environments experienced by black shales. Thermodynamic processes that would effect mobility during deposition and pre- metamorphic diagenesis are oxidation-reduction potential (Eh or pE) and pH (Stumm and Morgan, 1970). A biological factor, productivity/organic preservation in sediments, affects both pH and pE. Of the mobile elements, the concentrations of Mn and Fe are sensitive to redox conditions that prevailed during deposition or pre- metamorphic diagenesis. V is an indicator of organic preservation (Brumsack, 1986).

3.2.1 Concentrations of Environmentally Indicative Mobile Elements.

The concentrations of three of the elements with multimodal distributions, Mn, Fe and V, are used to distinguish among different depositional and early diagenetic environments.

Manganese concentrations vary considerably - from 15-3610 ppm. The distribution is trimodal. The primary mode is 150-200 ppm; secondary modes occur at 50 and 1500 ppm (Quinby-Hunt et al., 1989a). The average concentration, 350 ppm, is lower than the 465-850 ppm reported for the more oxic shale composites (Turekian and Wedepohl, 1961; Gromet et al., 1984; Taylor and McLennan, 1985).

The iron distribution is also wide (4600-100000 ppm), and bimodal. The primary mode is 50000 ppm; the secondary mode is at 7000 ppm. The average Fe concentration, 36200 ppm, is lower than those observed in the more oxic shale composites (40000-50000 ppm; Turekian and Wedepohl, 1961; Gromet et al., 1984; Taylor and McLennan, 1985).

The average vanadium content in these black shales is 470 ppm, considerably greater than the average 130-150 ppm reported for more oxic shale composites (Turekian and Wedepohl, 1961; Taylor and McLennan, 1985). The frequency distribution of V is bimodal: the vast majority (the primary mode) of the samples cluster at 150 ppm, however, the samples making up the secondary mode average 1500 ppm (Quinby-Hunt et al., 1989a). Shales enriched in V have been reported in the Dictyonema shale in Balto/Scandinavia (Vine and Tourtelot, 1980; Bjørlykke, 1974; Gee, 1980; and Wilde et al. 1989); the Kupferschiefer in the Permian (Wedepohl, 1964) and even some deposits in the Cretaceous (Brumsack, 1986). The low-calcic shales can be separated into two groups based on V concentrations: those that fluctuate about 150 ppm and those that are enriched with V concentrations > 800 ppm.

3.4. pE-pH Diagrams

The stability fields described in pE-pH diagrams for common marine oxides, sulfides and carbonates depend on a number of variables, including pH, pE, cation and anion concentrations, and the particular equilibrium expression chosen. It should be emphasized that pE-pH diagrams may vary depending on conditions, therefore those diagrams constructed for examining ore solutions for economic geologic situations are not applicable to normal marine environments. By carefully defining variables and ranges of concentrations encountered in the marine regime, it is possible to construct pE-pH diagrams with realistic stability fields. Such diagrams may be useful for understanding sediment compositions for those elements in quasi-chemical equilibrium during early diagenesis. pE-pH diagrams for Mn (Figure 1) and Fe (Figure 2) dominant species were calculated using the methods of Garrels and Christ (1965) and Lee and Sillén (1959); Gibb's free energies were those cited in Garrels and Christ and Brookins (1987). The pH range used was 5-9 bracketing natural marine to estuarine conditions. This includes range for aerobic sea water, 7.8-8.5; that of anaerobic sea water, but usually pH 6.8 or greater (Skirrow, 1975; Grasshoff, 1975; Jacobs et al., 1985), and that of interstitial waters, usually greater than pH 5, (Baas Becking et al., 1960). Certain alkaline hypersaline estuaries may have pH values in excess of 9 (Krumbein and Garrels, 1952), but such examples would be rare in the geologic record or would produce carbonates and evaporites not covered in our sample set. This range of pH includes that cited by Lewan (1984; Baas Becking et al., 1960) for marginal and open marine sediments. The pE range is the stability field of water, where neither oxygen or hydrogen gas is generated. Diagrams were calculated at 25oC and 1 atm. No significant variations at the scale of our diagrams are expected in the thermodynamic constants between 0o and 50oC and up to tens of atmospheres pressure (Evans and Garrels, 1958; Garrels and Christ, 1965; Lewan, 1984).

The equation defining the reduction of NO3- to N2 (Stumm and Morgan, 1970):

2NO3- + 6H+ + 4 e- => N2 + 3H2O, (1)

is an effective boundary between oxic and anoxic conditions and is independent of dissolved NO3- concentration. The equations describing the reduction of SO4= to H2S or HS- (Garrels and Christ, 1965):

SO4= + 10H+ + 8 e- => H2S + 4H2O (2)

and SO4= + 10H+ + 9 e- => HS- + 4H2O (3)

(also independent of dissolved sulfur concentration) approximate the "sulfate-sulfide fence" of Krumbein and Garrels (1952). Both of these redox boundaries vary with pH inclining to low pE with increasing pH. The sulfate- sulfide boundary roughly corresponds to the occurrence of sulfide minerals, whose occurrence depends on both the cationic and dissolved sulfur concentrations. The maximum value for total dissolved sulfur is defined by the oceanic SO4= concentration of the modern sea water concentration at 35 ppt or 28mM/kg (Morris and Riley, 1966) with an activity coefficient of 0.12 (Garrels and Christ, 1965, p.103).

The mean carbonate concentration in sea water depends on atmospheric carbon dioxide pressure. The Phanerozoic content of atmospheric carbon dioxide varied about 18 times from a maximum in the Cambrian to a minimum in the modern world (Berner, 1990). As the total carbonate concentration (Sigma CO2) varied so significantly, pE-pH diagrams were calculated at two concentrations of Sigma CO2. The first concentration was the average observed Sigma CO2 in the modern ocean: 2.2 mM (Skirrow, 1975) with an activity coefficient of 0.47 (Garrels and Christ, 1965, p.93) with a resulting log(Sigma CO2) = -3. The second concentration calculated was estimated using the late Cambrian atmospheric concentration of CO2, roughly 18 times present, the maximum projected during the Phanerozoic (Berner, 1990): log(Sigma CO2) = -1.7. As noted below, this variation was important for the Mn diagram, but not for the Fe diagram. The carbonate mineral stability fence is independent of pE and for marine conditions is located at higher pH.

The Mn concentration for oxic sea water used in the calculations was that of the Pacific Ocean just above the oxygen minimum zone (Martin and Knauer, 1984; Figure 3) - ~10-8; in the interstitial waters of anoxic sediments and anoxic basins, [Mn++] can be - ~10-5 (Jacobs and Emerson, 1982; Jacobs et al., 1987; Figure 3). The concentration boundaries were based on the solubilities of Mn(OH)2:

Mn(OH)2 + 2H+ => Mn++ + 2H2O, (4)


Mn2O3 + 6H+ + 2 e- => 2Mn++ + 3H2O, (5)


Mn3O4 + 8H+ + 2 e- => 3Mn++ + 4H2O, (6)


MnO2 + 4H+ + 2 e- => Mn++ + 2H2O, (7)

and MnCO3:

MnCO3 + H+ => Mn++ + HCO3 (8)

(Bricker, 1964; Garrels and Christ, 1965; Brookins, 1987). Other oxides were found to be even more soluble. Calculations assumed that the activity coefficient for Mn++ is 1; in actuality it probably is considerably lower, and dissolution may occur at higher pE for a specific pH. Rhodochrosite (MnCO3) solubility occurs within the pE-pH field of interest, therefore two "fences" are presented in Figure 1. One at the higher pH (average observed Sigma CO2, log(Sigma CO2) = -3) for the modern ocean and at lower pH (estimated atmospheric concentration : log(Sigma CO2) = -1.7) for the late Cambrian. Under the conditions presented here, alabandite (MnS) is not stable. Alabandite has been observed in the Landsport Deep, a permanently anoxic basin in the Baltic Sea (Suess, 1979). Suess attributes its occurrence to the high productivity conditions, resulting in supersaturation in the interstitial waters as the result of excessive terrestrially-dominated input of Mn to this basin.

In calculating the pE-Ph diagram for Fe (Figure 2), the concentration of Fe++ used was that of oxic sea water for the Pacific Ocean just above the oxygen minimum zone - 10-9 (Gordon et al., 1982; Figure 4). The concentration observed in anoxic basins can be as high as 320 nM as observed by Jacobs and others (1987) in the Cariaco Trench (Figure 4). In determining the solubility of Fe++, the dissolution of Fe2O3:

Fe2O3 + 2H+ + 2 e- => 2Fe++ + 3H2O, (9)

and Fe3O4:

Fe3O4 + 8H+ + 2 e- => 3Fe++ + 4H2O (10)

(Garrels and Christ, 1965; Brookins, 1987) are shown as these compounds remained insoluble to lowest pE. In Figure 2, a [Fe++] of 10-6 is shown to show an upper limit of dissolved iron concentrations in anoxic sea water. Figure 2 shows two windows of solubility for Fe++, one just above sulfate-reduction, another in the low pH region of the sulfide zone. Siderite (FeCO3) is either soluble or unstable relative to the Fe oxides and sulfides under the conditions of our calculations, therefore only one diagram is presented.

4. Discussion

4.1 Mn and Fe Redox-pH Zonation in Black Shales

. Iron and manganese occur in different oxidation states, in numerous minerals with varying solubilities over the range of pE and pH in open ocean and interstitial sediment waters (Garrels, 1960; Garrels and Christ, 1965; Boström, 1967; Burns and Burns, 1979; Murray, 1979; Berner, 1981a,b; Brookins, 1987; Stumm and Morgan, 1970; Figures 1 and 2). The relationships among the statistical distributions of Mn, Fe, and V (Figure 5) suggest a four fold chemical clustering: Group 1 - high Mn, high Fe, low V; Group 2 - low Mn, high Fe, low V; Group 3 - low Mn, low Mn, low V; and Group 4 - low Mn, low Fe, high V. The three boundaries are respectively the nodal points between high and low Mn, high and low Fe, and high and low V.

The origin of the four groups is explained by examining them with respect to a pE-pH diagram of Mn and Fe species constructed for conditions observed in the open ocean and in the underlying sediments (Figure 6). The Mn solid/Mn++ and Fe solid/Fe++ boundaries in Figure 6 reflect modern concentrations. The basic assumption is that a relatively high elemental concentration reflects a position in stable fields for that element and that significantly lower concentrations indicate deposition under conditions where the species is unstable and is dissolving. The combined pE-pH diagram is subdivided into redox zones based on the dominant oxidant. It can then be used to describe the aqueous conditions postulated during deposition of the shales before the system became closed as the result of further deposition. In some cases, the redox zone boundaries or fences (Krumbein and Garrels, 1952) essentially coincide with those of the stability fields for the marine conditions, for example: the O2 to NO3- oxidant fence coincides with the lower limit of the Mn oxide stability field and the NO3- to SO4= oxidant fence coincides roughly with the onset of the FeS2 stability field. On the other hand, the Fe oxide stability field is inclined in the middle of the nitrate oxidant zone. Thus, the clustering of the elements in the groups representing Fe, Mn and V species stabilities may not relate each to a single oxidant-defined zone.

4.1.1. Oxic Conditions: Zone I.

Under modern atmospheric concentrations, oxygen is the primary oxidizing agent and plankton-derived organic matter as the primary reducing agent in the oceans. The following equation (Richards, 1965) indicates the relationship among phytoplankton, nutrients and oxygen in sea water:

[(CH2O)106(NH3)16(H3PO4)] + 138 O2 =>

106 CO2 + 122 H2O + 16 HNO3 + H3O4. (11)

This relationship has been used by Froelich et al. (1979); Wilde (1987); and Shaffer (1989) as the basis of redox reactions in both the water column and in sediments. Under oxic conditions, hematite, goethite, MnO2-type and some Mn2O3 minerals are stable (Berner, 1981a; Figures 1 and 2). Fe and Mn concentrations in the sediments might be expected to be high so long as reducing conditions do not develop during early diagenesis. The solubility of Fe and Mn in oxic sea water has been observed to be exceedingly low, of the order of 10-8-10-9 M/kg in the north Pacific (Gordon et al., 1982; Martin and Knauer, 1984; Figures 3 and 4).

If the water column and bottom conditions during deposition and early diagenesis were primarily oxic, the sediments that collect to form black shales will contain relatively high concentrations of Mn and Fe, due to the high insolubility of their oxides (Group 1: Figure 5). This group of samples corresponds to the upper mode of the Mn frequency distribution. It contains an average of 1300 ppm Mn, 56000 ppm Fe and 130 ppm V (Table 2). The average Mn and Fe concentrations for these samples are even higher than the average concentrations of Mn (850 ppm) and Fe (47200, 50000 ppm) in the general composites of Turekian and Wedepohl (1961) and Taylor and McLennan (1985). The conditions these sediments experienced during deposition and early diagenesis permitted the oxidation of organic matter, therefore the V concentration, 130 ppm, may be considered a baseline concentration for V in marine black shales. This concentration is comparable with the V concentration of 130-150 ppm reported for shales in Turekian and Wedepohl (1961) and Taylor and McLennan (1985). Shales observed with Mn concentrations high enough to be placed in this group never contained V in concentrations greater than 300 ppm.

Group 1 samples fall into the oxic zone (Zone I, Figure 6). They include samples from the Tremadoc in Wales, in particular from the Upper Lingula Flags (Portmadoc, Wales) known to be formed in shallow marine water from their fossil content. From Dob's Linn, the oxic samples were deposited either during the height of glaciations during latest Ordovician time when oceanic environments were well-ventilated (Brenchley and Newall, 1984) or during shoaling in the depositional environment that occurred during convolutus, sedgwickii, and maximus graptolite zones of the early Silurian. Several samples from the Carboniferous also fell in this group.

4.1.2. Nitrate-Reducing Conditions: Zone II.

If most of the oxygen in the water column is consumed, ~<0.2 mL/L) and organic matter is present and nitrate reduction begins (Wilde, 1987; Froelich et al. 1979; Shaffer, 1989; Murray and Kuivila, 1990):

[(CH2O)106(NH3)16(H3PO4)] + x NO3- =>

106 CO2 + y H2O + z [NO2-, N2, NH3] + H3O4 (12)

( x, y, z depend upon conditions). This equation defines the boundary between oxic and anoxic conditions. At pH 8 nitrate-reduction occurs at ~pE of 9.3 (~Eh 0.55, Figure 1). The figure shows that at the pH of sea water, Mn oxides will begin to dissolve concomitant with nitrate reduction. Oceanic conditions under which nitrate is the primary oxidizing agent have been termed the nitratic or nitrate-reducing zone (Wilde, 1987). The conditions for Wilde's nitratic zone and Berner's (1981a) corresponding zone in sediments, the sub-oxic, non-sulfidic zone, extend from the onset of nitrate-reduction to the beginning of sulfate reduction (see Figures 1 and 2). Such conditions are observed in the modern ocean, although not extensively, due to the low concentrations of nitrate. One example observed is in the eastern tropical Pacific off Peru (Anderson, 1982). In sediments from the eastern equatorial Atlantic, Froelich et al. (1979) observed that, as oxygen becomes undetectable, NO3- concentrations decline to zero. After the decline in nitrate concentrations begins, the dissolved Mn+2 concentration increased in the pore waters. Fe+2 concentrations remained undetectable until after nitrate became undetectable. The pH fluctuated between 7.6 and 7.8. Sulfide remained undetectable.

Mn-reduction: Zone IIa. At an pE roughly comparable to that of nitrate reduction in open ocean water, del-MnO2 (IV) is reduced to Mn2O3 (III). At only slightly lower pE, both species are reduced to Mn++ (Figure 1). Under these conditions, concentrations of dissolved Mn++ rise (Murray and Kuivila, 1990), consequently, concentrations of Mn in the sediments are expected to decline. Sediments precipitated under, and continuing to spend early diagenesis under, conditions wherein Mn++ is soluble and Fe2O3 is stable (Zone IIa, figure 6) would result in shales with concentrations such as those in black shale group 2 (Figure 5, Table 2). These contain lower concentrations of Mn (310 ppm average) than in Zone I, but have an average concentrations of Fe (52000 ppm) and V (140 ppm), close to those observed in the oxic zone. V is present at background levels. Assignment of the group 2 shales to this zone is ambiguous as under sulfate-reducing conditions, Fe sulfides are stable and would also be present in high concentrations in the sediments. These samples would also be included in Group 2 (figure 5), and but would have formed under the conditions represented by Zone III (Figure 6). Zone III samples are discussed below in section 4.1.3. Sulfate-Reduction and the problem of assigning group 2 samples to a particular zone is addressed in section 4.2.1. Group 2 Black Shales: Nitratic or Sulfidic?.

Thus sediments deposited Zone IIa conditions would be depleted in Mn, but contain as much or more Fe as is observed under oxic conditions. Berner (1981a) notes that under nitratic (Mn-reducing) conditions, glauconite, and other iron silicates, siderite, vivianite and rhodochrosite, but no sulfides, are stable. Within the range of pH of ocean water, rhodochrosite (MnCO3) may also precipitate in the nitratic zone. However, as can be seen in Figure 1, rhodochrosite would not precipitate to as high concentrations as that of d-MnO2 and Mn2O3 in the Mn oxide zone. In the oceanic water column, even under the high Sigma CO2 conditions expected for the early Paleozoic, little rhodochrosite would be expected to form. Certainly under modern oceanic conditions even the presence of high levels of sulfides is insufficient to cause more than a slight decline in the oceanic concentration of Mn++. However, in regions of high productivity with either high levels of interstitially generated CO2 or unusually high levels of Mn++ and low salinity, such as is observed in the Landsdorf Deep of the Baltic Sea (Suess, 1979), rhodochrosite may precipitate. Under the conditions observed by the low-calcic black shales the low concentrations of Mn demonstrate that Mn is more soluble.

Fe-reducing Conditions: Zone IIb. As the oxygen, Mn(IV), Mn(III), and oxidized-nitrogen species are consumed, at pH 7.5-8, a pE of about -0.85 to -2.5 (Eh of about -0.05 to -0.15) is required for reduction of Fe2O3 (III) to Fe++ (Figure 2). In modern anoxic marine basins and fjords, such as the Cariaco Trench and the Saanich Inlet, both Fe and Mn are dramatically more soluble (Jacobs et al., 1985, 1987), than in the open oxygenated ocean, even under conditions of the very low oxygen at the oxygen minimum (Martin and Knauer, 1984 and Gordon et al., 1982) (Figures 3 and 4). In the presence of Fe-reducing bacteria in the sediments, Fe may be reduced and solubilized by microorganisms (Canfield, 1989).

The third group of black shale samples (Group 3, Figure 5) contains lower concentrations of both Mn and Fe (Table 2); the average Mn concentration is 170 ppm. The average Fe concentration is 23000 ppm: less than half the Fe concentration in samples in the oxic or Mn-reducing zones. These sediments were deposited after Fe- reduction, prior to SO4=-reduction and fall in redox Zone IIb (figure 6). Shales formed under the Zone IIb Fe- reducing conditions also have low V concentrations, indicating that (1) organic matter was oxidized, either in the water column or during early diagenesis, (2) the sediment system was not sufficiently open to the bottom waters to permit accumulation of V in tetrapyrroles or (3) according to Lewan's theory (1984), if organic matter is present, Ni++ competes successfully for sites in the tetrapyrrole (see below Section 4.1.4 Methane Stability, Zone IV).

Samples in Group 3 include all those from the Tremadoc in Levis (Quebec, Canada) and Belgium, the sample from Schaghticoke (New York, USA) and that from northern Norway, some from Bolivia, and all but one of those from the upper middle Ordovician from the Jutland Klippe (New Jersey, USA). Some Dob's Linn samples from the early Silurian cyphus and gregarius graptolite zones were deposited in or spent early diagenesis under Fe- reducing conditions. The vast majority of the samples from the latest Ordovician anceps and persculptus zones were deposited or spent early diagenesis under Fe-reducing conditions, at a time when glaciation had ceased and deep oceans no longer were ventilated (Brenchley and Newall, 1984; Berry and Wilde, 1978). As no anceps samples have high V content (see below), conditions must not have been sufficient for tetrapyrrole formation or preservation. Such conditions are expected from the Scottish Borderland's paleogeographic position in the Iapetus Ocean away from planetary divergences and high surface organic productivity (Wilde, 1991). Some Carboniferous samples also are included in this group.

4.1.3. Sulfate-Reduction, Zone III: Pyritization.

When SO4= becomes the primary oxidizing agent (Wilde's sulfatic zone, Berner's sulfidic) (Froelich et al., 1979; Wilde, 1987; Shaffer, 1989), HS- and H2S are available for coordination or precipitation:

[(CH2O)106(NH3)16(H3PO4)] + 53 SO4= =>

106 CO2 + 16NH3 + 53S= + H3O4 + 106 H2O. (13)

The anoxic marine pH is frequently lower, averaging 6.8-7.4 (Skirrow, 1975; Grasshoff, 1975; Emerson et al., 1983, Jacobs et al., 1985), vs 7.8-8.5 observed for the open oxic ocean (Skirrow, 1975). Initially, in the sulfatic zone, dissolved Fe concentrations in the water column remain high and in the sediments remain low, as insufficient sulfides will be available for precipitation (apparently group 3, Zone IIb black shales). Dissolution of Fe has been observed in the sediments of Boca de Quadra fjord, Alaska (Sugai, 1987), Cariaco Trench (Jacobs et al., 1987), Saanich Inlet (Jacobs and Emerson, 1982), Framvaren fjord (Jacobs et al., 1985), the Black Sea (Spencer and Brewer, 1971) and Chesapeake Bay (Eaton, 1979). As sulfides are generated, Fe sulfides precipitate or are formed by direct reaction with iron oxides in the sediments (Canfield and Berner, 1987). In the Cariaco Trench (Figure 4) and Framvaren Fjord, a decline is observed in the dissolved Fe concentration (Jacobs et al., 1985, 1987).

In Berner's (1981a) sulfidic zone, pyrite, marcassite, rhodochrosite, and alabandite are stable. Reduction of organic matter by sulfates or sulfate-reducing bacteria has been observed (Jørgensen, 1982; Crill and Martens, 1987; Oremland and Taylor, 1978). In sediments, sulfide generated during sulfate-reduction reacts with reactive Fe oxides resulting in Fe sulfides, prior to the increase of H2S in the pore waters (Canfield, 1989). Berner (1984) notes that refractory organic matter is stable in this zone, although sulfate-reduction is still energetically favored (Froelich et al., 1979). In fact, Raiswell and Berner (1986) report a linear correlation between preserved organic carbon and pyrite. They note that the C/S ratio is >1.4 in samples from the Devonian (development of vascular land plants) and more recent. However, the C/S from the Ordovician and Silurian is significantly lower, <0.75. They (Raiswell and Berner, 1986) suggest that prior to the Devonian, more reactive organic matter was buried: algae vs vascular land plants (Meybeck, 1982; Lyons and Gaudette, 1979). Thus, in the suite of low-calcic black shales in this study, the amount of organic matter preserved in this zone is probably low. Only six of the post-Devonian Carboniferous samples investigated fell into group 3. None of the Jurassic samples did.

4.1.4 Methane Stability: Zone IV

After SO4= is consumed, thermodynamically the aqueous system falls below the sulfide-stability line on the pE- pH diagrams (Figures 1, 2, 6). Just below the sulfide-stability line (Eo[HS->SO4=] = 0.252 eV), methane is stable (Eo[CH4->HCO3-] = 0.227 eV). Once sulfate is sufficiently depleted, methanogenesis can occur (Martens and Berner, 1974; Oremland and Taylor, 1978; Froelich et al., 1979; Berner, 1981a,b; Kuivila et al., 1989):

[(CH2O)106(NH3)16(H3PO4)] =>

53 CO2 + 53 CH4 + 16 NH3 + H3O4. (14)

Energetically, although methane is stable at pE only slightly below sulfate-reduction, in the oceanic water column, such large concentrations of sulfate are available that methanogenesis is highly improbable. In the sediments, when the system is only quasi-open, the sulfate concentration is fixed at sea water values in the pore waters. Methanogenesis can occur when sulfate is depleted by bacterial reduction (Martens and Berner, 1974; Oremland and Taylor, 1978; Froelich et al., 1979; Berner, 1981a,b; Kuivila et al., 1989). In the methane-stablility zone, sulfides are available (Howarth and Teal, 1980; Howarth and Jørgensen, 1984; Crill and Martens, 1987). However, as can be seen from Figure 6, in Zone IV both Fe++ and Mn++ are soluble. It is not unreasonable to assume that under those conditions that SO4= is consumed, methanogenesis might occur. This process could be similar to processes of dolomite formation that occur in deep-sea sediments (Kastner, 1983; Burns et al., 1988, Morrow and Ricketts, 1988). Kastner suggested that deep-sea dolomitization only occurs under low- to no-sulfate conditions.

Vanadium as an Inorganic Indicator of Organic Carbon Preservation. If organic matter is not oxidized, highly- reducing conditions prevail. Berner and Raiswell (1983, Raiswell and Berner, 1986) use carbon/sulfur ratios as environmental indicators. However, as they point out, techniques for determining C and S require "fresh" unweathered samples, such as core material, to yield reliable results due to the mobility of C and S. As such samples are difficult to find in the field and are limited to recent samples; most paleontologically-dated material from well-documented collections are not acceptable for such C/S analyses. The major black shale sedimentary intervals occur in the Paleozoic and Mesozoic. Thus the carbon content of these old rocks may be more indicative of age or maturity rather than that of depositional environmental conditions (Raiswell and Berner, 1987) and another indicator must be found. Vanadium can be used as an indicator of the preservation of planktonic organic C under certain conditions. Brumsack (1986) reported that V correlates with organic carbon (r=0.95) for Cretaceous black shales.

Vanadium occurs in organic matter as V-tetrapyrrole complexes (Lewan and Maynard, 1982), which are best preserved under low pH, highly-reducing conditions. V occurs naturally as V(V), V(IV) and V(III). Under most natural pH conditions, V appears as the vanadyl cation, VO++ (Breit, 1988). Planktonic organic matter contains compounds that are precursors to tetrapyrroles which form strong covalent bonds with the vanadyl cation (and Ni++); VOH++ may also form such complexes based on ligand field calculations (Lewan and Maynard, 1982; Lewan, 1984; Sawlowicz, 1985). These bonds are thermally stable, resistant to strong acids and inert to cation exchange reactions (Lewan and Maynard, 1982). Organic matter derived from marine plankton are known to concentrate V. The highest contents of V in fossil fuels occurs in deposits that were originally plankton or algae (Breit, 1988 Lewan, 1980); bitumens derived from plankton are elevated in V content; those derived from higher plants are not enriched. Plankton accumulate V, primarily in chlorophyll, a precursor to tetrapyrroles (Yen, 1975; Lewan and Maynard, 1982; Breit, 1988). Vanadyl tetrapyrrole complexes (Lewan and Maynard, 1982) and organic matter are best preserved under anoxic, aphotic conditions; organic matter can be oxidized even under anoxic conditions by nitrate and sulfate (Froelich et al., 1979), although refractory organic matter can remain intact even in the presence of sulfate-reducing bacteria. Vanadyl tetrapyrrole complexes are even more easily oxidized than most bulk organic matter of the organic matter that settles with them (Lewan and Maynard, 1982), therefore the presence of high concentrations of V suggests conditions appropriate for the preservation of less refractory organic matter.

A low pH may also be necessary for the enhancement of V concentration (Lewan and Maynard, 1982; Lewan 1984). The VO++ and V(OH)++ cations are stable only at pH < 6.5-7 (Lewan, 1984). At pH 7 VO++ is stable to pE ~-3.4 (Eh ~ -0.2 eV) at pH 7, V(OH)++ is stable to pE of < -6.8 (Eh < -0.4 eV) at pH ~6.7. At pH > 7, the stable entities are VO(OH)2, V(OH), V4O124-, V2O74-, and VO43-. These compounds are not as available for substitution into the porphyrin ring.

Vanadium content, therefore, can indicate the quantity of organic C that was originally compacted. Thus, in shales formed under oxic conditions, organic matter and the tetrapyrroles would have been oxidized and the V content represents a baseline inorganic contribution (average ~140 ppm). Vanadyl tetrapyrrole content should be high in shales formed with organic matter that descended through anoxic waters or was incorporated into anoxic sediments rapidly. In order for the V concentration to be enriched over that found in planktonic matter (generally ~25 ppm; Knauss and Ku, 1983), it is necessary that dissolved metals in interstitial waters in open sediment systems are available to metallate the tetrapyrroles. Lewan and Maynard (1982) have demonstrated that only in an open sediment system is sufficient V available to provide the metals necessary for enrichment. It is not clear whether the concentrations of V in sea water (modern oxic ocean, ~ 1 mg/kg; Morris, 1975; Zhou et al., 1982; Huizenga and Kester, 1982) are sufficient to supply the high concentrations of V observed, for example, in black shales from Balto-Scandia (Wilde et al., 1989). The vanadium found in the Balto-Scandic samples could have been supplemented by volcanic exhalations also (Breit, 1988), which is consistent with the vanadium enrichment observed near active ocean ridges (Hodkinson et al., 1986; Marchig et al., 1986).

Vanadium enhancement, Methanogenesis in Black Shales?. If conditions are anoxic (insufficient oxidants such as O2, nitrate or sulfate are available to consume or support consumption of organic matter) in the water column and during early diagenesis AND if marine productivity is sufficiently high that organic matter settles to the bottom, then organic matter may be preserved in the sediments. If the sediment system is quasi-open, then conditions may favor metallation of tetrapyrroles resulting in enhanced V concentrations (Lewan and Maynard, 1982). Lewan (1984; Lewan and Maynard, 1982) suggests that low pH, low pE conditions are necessary for preservation of vanadyl tetrapyrroles. These conditions would be met below the sulfate-reduction line at pH less than 7. In Figure 6, these conditions include Zone IV and part of zone III. However, in the low-calcic black shales, if V concentrations exceed 300 ppm, the concentrations of both Fe and Mn are very low. V concentrations are higher than 260 ppm only if Fe and Mn are less than 300 ppm and 37000 ppm respectively. Therefore, the low-calcic black shales containing high concentrations of V, can be assigned to zone IV. The zone IV boundaries would expand to include much of the HS- region of the pE-pH diagram as sulfides are precipitated or as sulfate becomes less available for reduction to sulfide.

The classic example of shales formed under conditions favoring tetrapyrrole formation and preservation are the Dictyonema Shales of Balto-Scandia and New Brunswick (Wilde et al., 1989). Many of these samples contained thousands of ppm of V; all contained greater than 700 ppm. A second group of high V samples contained from 320-800 ppm V. One sample from the Bolivian Tremadoc was in this group, no other samples from the Tremadoc were. Another was from the middle Ordovician in eastern Pennsylvania. The 300-800 ppm V group also includes a few samples from the latest Ordovician persculptus zone at Dob's Linn, and single samples from the cyphus, convolutus, acuminatus, wilsoni, and gracilis graptolite zones. Another such sample was from the Jurassic, from the Lias at Whitby, a likely oil-source rock (Gad et al., 1969). Although not part of this study, several samples from the oil-bearing rocks of the Monterey shale contained high concentrations (400 - 1300 ppm). The conditions for formation of V-rich, Fe- and Mn-poor black shales (low pE) are not observed in the oxygenated, modern ocean. In modern sediments, however, sufficient organic matter can accumulate to consume the nitrate and sulfate available in interstitial waters (Kuivila et al., 1989, 1990). Clearly, if productivity and rate of deposition are high enough, conditions will favor preservation of organic matter. Under the slow rates of deposition in the open ocean, most organic matter will have been oxidized while falling through the water column or during early diagenesis. However, it is conceivable that conditions for formation of V-rich, Fe- and Mn-poor black shales could develop in the paleo-ocean, given high productivity, slow circulation, and anoxic bottom water in a restricted anoxic basin or in sediments overlain by sulfate-reducing waters.

4.2 Ambiguous Assignments

4.2.1. Group 2 Black Shales, Nitratic or Sulfidic?

Redox-related solubility conditions for Fe in group 2 samples are not unambiguous. At intermediate pH, at pE lower than nitrate reduction, both ferric and ferrous iron species are stable throughout the natural pE range so that the Fe concentrations should remain relatively constant. Thus low Mn, high Fe, and low V samples could be from either Zone IIa (Nitrate-reducing) or Zone III (Sulfate- reducing). Finer discrimination of this group may require other elements such as Zn, Mo, Pb, etc., which preferentially form sulfides and thus would have low concentrations in Zone II and significantly higher concentrations in Zone III. Our data set analyzed by INAA, unfortunately does not contain enough of these elements to confidently separate group 2 samples. Due to the ubiquitous presence of sulfate-reducing bacteria and the low concentrations of oxidized-nitrogen compounds in sea water and in sediments compared to the concentration of oceanic sulfate, it is likely that the Group 2 shales were deposited in Zone III. A possible discrimination factor could be color as the Nitratic zone iron oxides, limonite and hematite, are red to brown, whereas Sulfatic zone iron sulfides would be darker in color.

Samples that occur in Group 2 include all but one of the Liassic samples. The samples from Dob's Linn are from the early Silurian convolutus, cyphus, atavus, and acuminatus graptolite zones. Samples from the cyphus, atavus, and acuminatus zones, were deposited under post-glaciation warming conditions, in relatively deep water (Brenchley and Newall, 1984). They accumulated under conditions that fluctuated between Mn-reducing and Fe- reducing. Most of the samples from the middle Ordovician from eastern Pennsylvania and all those from the Martinsburg formation were placed in this group. Those Welsh Tremadoc samples not in group 1, occur in group 2. Apparently they were deposited under conditions that fluctuated between the oxic and mildly anoxic, which is to be expected as they were deposited under relatively shallow waters.

4.2.2. Group 3 Black Shales, Zones IIb or IV?

The group 3 black shales have been placed in zone IIb. Shales with this chemical signature (low Mn, low Fe and low V) could, in fact, occur at lower pE or at higher pH than is shown for zone IIb under some conditions. First, if productivity is low, or the organic matter produced is readily oxidized, sediments may have a low V signature, even at very low pE, because insufficient organic matter is present to concentrate V. Secondly, under conditions of very high productivity but also high rate of sedimentation, sufficient organic matter could be preserved, but may not provide the quasi-open system, so that V as tetrapyrroles (Lewan and Maynard, 1982) is diluted yielding a "low" bulk vanadium concentration. This situation is not likely under open-ocean conditions with low rates of deposition, however in estuarine or near-shore situations, such conditions probably occur. In our finely-laminated samples, which had low rates of deposition, such a situation is not likely. Group 3 black shales might also occur if there are insufficient sulfides present to precipitate pyrite in zone III. Under marine conditions, such a situation would be highly unlikely. Thus, for the finely-laminated, marine black shales examined in this data set only the possibility that there was insufficient productivity to provide a concentrating agent for V is a reasonable concern.

5. Summary

The combination of thermodynamic analyses of dominant phases using pE-pH diagrams with concentrations of dominant mobile elements for the black shale facies provides a consistent explanation of environmental conditions. Modal plots of Fe, Mn, and V concentrations in sediments identify four chemical groups which may be related to stability fields in pE-pH diagrams. For each group, relatively high concentrations reflect stable fields and relatively low concentrations indicate instability and dissolution for defined marine conditions. Four distinct redox zones based on a sequence of available oxidants, from oxic to methanogenic, can be postulated that are bounded by thermodynamic defining equations. Reduction of nitrate to nitrogen and solubility of Mn oxides define the boundary between oxic and anoxic conditions. The anoxic zone is further subdivided by the stability field of iron oxides, indicating more reducing conditions within the nitratic zone and the stability field of iron sulfides which indicates sulfate-reduction. The stability field of vanadium compounds suggests methanogenesis. Three pH zones are apparent. From pH 7 to 9, the pH is buffered in the open ocean and provides the initial pH for marine interstitial waters. Above about pH 9, Fe and Mn chemically precipitated carbonates are stable as a function of varying carbonate content. Below pH 7, sedimentary processes dominate over oceanic buffers with oxidation of organic matter increasing interstitial carbon dioxide above that in equilibrium with the ocean and lowering the pH.

The wide range of pE-pH conditions considered reasonable for black shale deposition of the four chemically discrete groups suggest that lithologically-defined black shales are not always absolute indicators of narrowly- defined anoxic conditions. Use of additional elements may permit subdivision of these four groups with additional environmental sensitivity.

Although this discussion was based on chemical analyses of field described "black" shales; the fundamental concepts are applicable to any shale regardless of color, organic carbon or carbonate content. The possibility of black "oxic" shales as well as brown or red "anoxic" shales counter to geologic field wisdom suggests that color or organic matter content alone be abandoned as an unambiguous indicator of oxidation-reduction conditions or of organic preservation. "There are black shales and there are black shales".


. We thank George Breit for his helpful discussions on vanadium. We thank Robert A. Berner for his comments and discussions. W. B. N. Berry offered continuous encouragement to publish a redox discussion of black shales. The junior author acknowledges the support of B.-D. Erdtmann and the von Humboldt Stiftung while he was in residence in Berlin in 1989-90. This is contribution MSG-91-001 of the Marine Science Group of the University of California



Anderson, J.J., 1982. The nitrite-oxygen interface at the top of the oxygen minimum zone in the eastern tropical Pacific. Deep-Sea Res. 29: 1193-1201.

Baas Becking, L.G.M., Kaplan, I.R. and Moore, D., 1960. Limits of the natural environment in terms of pH and oxidation potential. J Geol., 68: 243-284.

Bacon, M.P., Brewer, P.G., Spencer, D.W., Murray, J.W., and Goddard, J., 1980. Lead-210, polonium-210, manganese and iron in the Cariaco Trench. Deep-Sea Res., 27: 119-135.

Berner, R.A., 1981a. A new geochemical classification of sedimentary environments. J. Sed. Petr., 51: 359-365.

Berner, R.A., 1981b. Authigenic mineral formation resulting from organic matter decomposition in modern sediments. Forschr. Miner., 59: 117-135.

Berner, R.A., 1984. Sedimentary pyrite formation: an update. Geochem. Cosmochim. Acta, 48: 605-615.

Berner, R.A., 1990. Atmospheric carbon dioxide levels over Phanerozoic time. Science, 249: 1382-1386.

Berner, R.A. and Raiswell, R., 1983. Burial of organic carbon and pyrite sulfur in sediments over Phanerozoic time: a new theory. Geochim. Cosmochim. Acta 47: 855-862.

Berry, W.B.N. and Wilde, P., 1978. Progressive ventilation of the oceans - an explanation for the distribution of the Lower Paleozoic black shales. Amer. J. Sci., 278: 257-275.

Bjørlykke, K., 1974. Depositional history and geochemical composition of Lower Paleozoic Epicontinental sediments from the Oslo Region. Norske Geologisk Undersøkelse, 305: 1-81.

Boström, K., 1967. Some pH-controlling redox reactions in natural waters. In: (W. Stumm, Session Chairman) Equilibrium Concepts in Natural Water Systems, Advances in Chemistry Series, No. 67, American Chemical Society, New York. pp. 286-311.

Breit, G., 1988. Vanadium -- Resources in fossil fuels. U.S. Geological Survey, Bulletin 1877. Denver, CO.

Brenchley, P.J. and Newall, G., 1984. Late Ordovician environmental changes and their effect on faunas. In: Bruton, D.L. (ed.) Aspects of the Ordovician System, Palaeontological Contribution from the University of Oslo, No 295, 65-79.

Bricker, O.P., 1964. Stabiity relations in the system Mn-O2-H2O at 25oC and One Atmosphere Total Pressure. Ph.D. Thesis. Harvard University, Cambridge, Massachusetts.

Brookins, D.G., 1987. Eh-pH Diagrams for Geochemistry. Springer-Verlag, Berlin.

Brumsack, 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: North Atlantic Palaeoceanography (C.P. Summerhayes and N.J. Shakleton, eds.), Geological Special Publication No. 21, Blackwell Scientific, Oxford. pp.447-462

Burns, R.G. and Burns, V.M., 1979. Manganese oxides. In: Marine Minerals (R.G. Burns, ed.), Mineralogical Society of America, Short Course Notes, Vol. 6, Mineralogical Society of America, Washington DC. pp. 1-46.

Burns, S.J., Baker, P.A., and Showers, W.J., 1988. The factos controlling the formation and chemistry of dolomite in organic-rich sediments: Miocene Drakes Bay Formation, California. In: Sedimentrology and geochemsitry of Dolostones (V. Shulka and P.A. Baker, eds.) SEPM Spec. Pub. No. 43, pp. 41-52.

Byers, C.W., 1977. Biofacies patterns in euxinic basins, a general model. In: Deep-water Carbonate Environments (H.E. Cook and P. Enos, eds.) Soc. Econ. Paleontologists, Spec. Pub. No. 25, pp. 5-17.

Canfield, D.E., 1989. Reactive iron in marine sediments. Geochim. Cosmochim. Acta, 53: 619- 632.

Canfield, D.E., and Berner, R.A., 1987. Dissolution and pyritization of magnetite in anoxic marine sediments. Geochim. Cosmochim. Acta, 51: 645-659.

Crill, P.M. and Martens, C.S., 1987. Biogeochemical cycling in an organic-rich coastal marine basin. 6. Temporal and spatial variations in sulfate reduction rates. Geochim. Cosmochim. Acta, 51: 1175-1186.

Eaton, A., 1979. The impact of anoxia on Mn fluxes in the Chesapeake Bay. Geochim. Cosmochim. Acta, 43: 429-432.

Emerson, S, Jacobs, L, and Tebo, B., 1983. The Behavior of trace metals in marine anoxic waters: Solubilities at the oxygen-hydrogen sulfide interface. In: Trace Metals in Sea Water (C.S. Wong, E.Boyle, K.W. Bruland, J.D. Burton, and E.D. Goldberg, eds.). Plenum Press, New York. pp.579-608.

Evans, H.T. Jr., and Garrels, R.M., 1958. Thermodynamic equilibria of vanadium in aqueous systems as applied to the interpretation of the Colorado Plateau ore deposits. Geochim. Cosmochim. Acta, 15: 131-149.

Froelich, P.N., Klinkhammer, G.P., Bender, M.L., Luedtke, N.A., Heath, G.R., Cullen, D., Dauphin, P., Hammond, D., Hartman, B. and Maynard, V., 1979. Early oxidation of organic matter in pelagic sediments of the eastern equatorial Atlantic: Suboxic diagenesis. Geochim. Cosmochim. Acta, 43: 1075-1090.

Gad, M.A., Catt, J.A., and Le Riche, H.H., 1969. Geochemistry of the Whitbian (Upper Lias) sediments of the Yorkshire Coast. Proc. of the Yorkshire Geol. Soc., 37: 105-139.

Garrels, R.M., 1960. Mineral Equilibria at Low Temperature and Pressure. Harper & Bros, New York.

Garrels, R.M. And Christ, C.L., 1965. Solutions, Minerals, and Equilibria. Freeman, Cooper and Co, San Francisco.

Gee, D.G., 1980. The Dictyonema-bearing phyllites at Nordaunevoll, eastern Trondelag, Norway. Norske Geologisk Tiddskrift, 61: 93-95.

Gordon, R.M., Martin, J.H., and Knauer, G.A., 1982. Iron in north-east Pacific waters. Nature 299: 611-612.

Grasshoff, K., 1975. The hydrochemistry of landlocked basins and fjords. In: Chemical Oceanography, 2nd ed. (J.P. Riley and G. Skirrow, eds.), 2: 455-597. Academic, London.

Gromet, L.P., Dymek, R.F., Haskin, L.A., and Korotev, R.L., 1984. The "North American shale composite": Its compilation, major and trace element characteristics. Geochim. Cosmochim. Acta, 48: 2469-2482.

Haack, U., Heinrichs, H., Boness, M., and Schaneider, A., 1984. Loss of metals from pelites during regional metamorphism. Contrib. Mineral. Petrol. 85: 103-115.

Hallam, A., 1987. Mesozoic marine organic-rich shales. In: Brooks, J. and Fleet, A.J., eds., Marine Petroleum Source Books, Geological Society Spec. Pub. 26, 251-261.

Hay, W.W., 1981. Sedimentological and geochemical trends resulting from the breakup of Pangea. Oceanogr. Acta, 4 (suppl.): 135-147.

Hodkinson, R., Cronan, D.S., Glasby, G.P., and Moorby, S.A., 1986. Geochemistry of marine sediments from the Lau Basin, Havre Trough, and Tonga-Kermedec Ridge. New Zealand J. Geology and Geophysics, 29: 335-344.

Howarth, R.W. and Jørgensen, B.B., 1984. Formation of 35S-labelled elemental sulfur and pyrite in coastal marine sediments (Limfjorden and Kysing fjord, Denmark) during short term 35SO4 reduction measurements. Geochim. Cosmochim. Acta, 48: 1807-1818.

Howarth, R.W. and Teal, J.M., 1980. Energy flow in a salt marsh exosystem: the role of reduced inorganic sulfur compounds. Amer. Nat., 116: 862-872.

Huizenga, D.L. and Kester, D.R., 1982. The distribution of vanadium in the northwestern Atlantic ocean. EOS, 63: 990.

Jacobs, L. and Emerson, S., 1982. Trace metal solubility in an anoxic fjord. Earth Planet. Sci. Lett. 60: 237-252.

Jacobs, L., Emerson, S. and Huested, S.S., 1987. Trace metal geochemistry in the Cariaco Trench. Deep-Sea Res. 34: 956-981.

Jacobs, L., Emerson, S., and Skei, J., 1985. Partitioning and transport of metals across the O2/H2S interface ina permanently anoxic basin: Framvaren Fjord, Norway. Geochim. Cosmochim. Acta 49: 1433-1444.

Jørgensen, B.B., 1982. Mineralization of organic matter in the sea bed - the role of sulfate reduction. Nature, 296: 643-645.

Kastner, M., 1983. Origin of dolomite and its spatial and chronological distribution - A new insight. AAPG Bull., 67: 2156.

Knauss, K. and T.-L. Ku, 1983. The elemental composition and decay-series radionuclide content of plankton from the east Pacific. Chem. Geol. 39:125-145.

Krumbein, W.C. and Garrels, R.M., 1952. Origin and classification of chemical sediments in terms of pH and oxidation-reduction potentials. J. Geol., 60: 1-33.

Kuivila, K.M., Murray, J.W., Devol, A.H., and Novelli, P.C., 1989. Methane production, sulfate reduction, and competititon for substrates in the sediments of Lake Washington. Geochim. Cosmochim. Acta, 53: 409-416.

Kuivila, K.M., Murray, J.W., and Devol, A.H., 1990. Methane production in the sulfate-depleted sediments of two marine basins. Geochim. Cosmochim. Acta, 54: 403-411.

Landing, W.M. and Bruland, K.W., 1987. The contrasting biogeochemistry of iron and manganese in the Pacific Ocean. Geochim. Cosmochim. Acta, 51: 29-43.

Lee, T.S. and L.G. Sillén, 1959. Chemical Equilibrium in Analytical Chemistry. Interscience, New York, 317pp.

Lewan, M.D., 1984. Factors controlling the proportionality of vanadium to nickel in crude oils. Geochim. Cosmochim. Acta 48: 2231-2238.

Lewan, M.D. and Maynard, J.B., 1982. Factors controlling enrichment of vanadium and nickel in the bitumen of organic sedimentary rocks. Geochim. Cosmochim. Acta 46: 2547-2560.

Lyons, W.B. and Gaudette, M.E., 1979. Sulfate reduction and the nature of organic matter in estuarine sediments. Org. Geochem., 1: 151-155.

Marchig, V. Erzinger, J. and Heinze, P.M., 1986. Sediment in the black smoker area of the East Pacific Rise (18.5oS). Earth Planet. Sci. Lett., 79: 93-106.

Martens C.S. and Berner, R.A., 1974. Methane production in the interstitial waters of sulfate- depleted sediments. Science, 185: 1167-1169.

Martin, J.H. and Knauer, G.A., 1984. VERTEX: Manganese transport though oxygen minima. Earth Planet. Sci. Lett. 67: 35-47.

Meybeck, M., 1982. Carbon, nitrogen, and phosphorus transport by world rivers. Amer. J. Sci., 282: 401-450.

Minor, M.M., Hensley, W.K., Denton, M.M., and Garcia, S.R., 1982. An automated activation analysis system. J. Radioanalytical Chem. 70: 459-471.

Morris, A.W., 1975. Dissolved molybdenum and vanadium in the northeast Atlantic Ocean. Deep-Sea Res., 22: 49-54.

Morris, A.W. and Riley, J.P., 1966. The bromide/chlorinity and sulphate/chlorinity ratio in sea water. Deep-Sea Res., 13: 699-705.

Morris, J.C. and Stumm, W., 1967. Redox equilibria and measurements of potentials in the aquatic environment. In: (W. Stumm, Session Chairman) Equilibrium Concepts in Natural Water Systems, Advances in Chemistry Series, No. 67, American Chemical Society, New York. pp. 270- 285.

Morrow, D.W. and Ricketts, B.D., 1988. Experimental investigation of sulfate inhibition of dolomite and mineral analogues. In: Sedimentrology and geochemsitry of Dolostones (V. Shulka and P.A. Baker, eds.) SEPM Spec. Pub. No. 43, pp. 25-38.

Murray, J.W., 1979. Iron oxides. In: Marine Minerals (R.G. Burns, ed.), Mineralogical Society of America, Short Course Notes, Vol. 6, Mineralogical Society of America, Washington DC. pp. 47-98.

Oremland, R.S. and Taylor, B.F., 1978. Sulfate reduction and methanogenesis in marine sediments. Geochim. Cosmochim. Acta, 42: 209-214.

Pederson, T.F. and Calvert, S.E., 1990. Anoxia vs. productivity: What controls the formation of organic-carbon-rich sediments and sedimentary rocks? Am. Assoc. Pet. Geol. Bull., 74: 454-466.

Pettijohn, F.J., 1949. Sedimentary Rocks. Harper and Brothers, New York, 526pp.

Quinby-Hunt, M.S., Wilde, P., and Berry, W.B.N. 1989a. Elemental geochemistry of low-calcic black shales - statistical comparison with other shales. In: Metalliferous Black Shales and Related Ore Deposits, U.S. Geological Survey Circular 1037, R.I. Grauch and J.S. Leventhal, eds., pp. 8- 15.

Quinby-Hunt, M.S., Wilde, P., and Berry, W.B.N. 1989b. Use of trace metal discrimination diagrams to determine provenance in metalliferous black shales. 28th International Geological Congress, Washington, DC, July. MSG-89-006.

Quinby-Hunt, M.S. and Wilde, P. 1991. The provenance of low-calcic black shales. Mineralium Deposita, v. 26, p. 113-121.

Raiswell, R. and Berner, R.A., 1986. Pyrite and organic matter in Phanerozoic normal marine shales. Geochim. Cosmochim. Acta, 50: 1967-1976.

Raiswell, R. and Berner, R.A., 1987. Organic carbon losses during burial and thermal maturation of normal marine shales. Geology, 15: 853-856.

Rhoades, D.C. and Morse, J.W., 1971. Evolutionary and ecologic significance of oxygen- deficient marine basins. Lethaia, 4: 413-428.

Richards, F.A., 1965. Anoxic basins and fjords. In: Chemical Oceanography, 1st ed. (J.P. Riley and G. Skirrow, eds.), 1: 611-645. Academic, London.

Savrda, C. and Bottjer, D.J., 1987. The exaerobic zone, a new oxygen-deficient marine biofacies. Nature, 327: 54-56.

Sawkowicz, Z., 1985. Significance of metalloporphyrins for the matal accumulation in the copper- bearing shales from the Zechstein copper deposits (Poland). Mineral. Polonica, 16: 35-42.

Shaffer, G., 1989. A model of biogeochemical cycling of phosphorus, nitrogen, oxygen, and sulfur in the ocean: one step toward a global climate model. J. Geophys. Res., 94(C2): 1979-2004.

Skirrow, G., 1975. The dissolved gases - Carbon dioxide. In: Chemical Oceanography, 2nd ed. (J.P. Riley and G. Skirrow, eds.), 2: 1-192. Academic, London.

Spencer, D.W. and Brewer, P.G., 1971. Vertical advection diffusion and redox potential as controls on the distribution of manganese and other trace metals dissolved in waters of the Black Sea. J. Geophys. Res., 76: 5877-5892.

Stumm, W. and Morgan, J.J., 1970. Aquatic Chemistry. Wiley-Interscience, New York. 780 pp.

Suess, E., 1979. Mineral phases formed in anoxic sediments by microbial decomposition of organic matter. Geochim. Cosmochim. Acta, 43: 339-352.

Sugai, S.F., 1987. Temporal changes in the sediment geochemistry of two south east Alaskan fjords. Deep-Sea Res., 34: 913-925.

Taylor, S.R. and McLennan, S.M., 1985. The Continental Crust: Its Composition and Evolution. Blackwell Scientific Publication, Oxford. 312pp.

Thompson, R.N., Morrison, M.A., Mattey, D.P., Dickin, A.P., and Moorbath, S., 1980. An assessment of the Th-Hf-Ta digram as a discriminant for tectonomagmatic classifications and in the detection of crustal contamination of magmas. Earth Planet. Sci. Letts., 50: 1-10.

Turekian, K.K. and Wedepohl, K.H., 1961. Distribution of the elements in some major units of the Earth's crust. Geol. Soc. Amer. Bull., 72: 175-191.

Vine, J.D. and Tourtelot, E.B., 1970. Geochemistry of black shale deposits - a summary report. Econ. Geol., 65: 253-272.

Wang, Y.L., Liu, Y.-G. and Schmitt, R.A., 1986. Rare earth element geochemistry of South Atlantic deep sea sediments: Ce anomaly change at ~54 MY. Geochim. Cosmochim. Acta, 50: 1337-1355.

Wedepohl, K.H., 1964. Untersuchen am Kupferschiefer in Nordwestdeutschland; Ein Beitrag zur Deutung der Genese bituminoser Sedimente. Geochim. Cosmochim. Acta, 28: 305-364.

Wilde, P., 1987. Model of progressive ventilation of the Late Precambrian-Early Paleozoic ocean. Amer. J. Sci., 287: 442-459.

Wilde, P., 1991. Oceanography in the Ordovician. In Advances in Ordovician Geology, C. R. Barnes and S. H. Williams eds., Geological Survey of Canada Paper, 90-9, p.283-298.

Wilde, P., Quinby-Hunt, M.S., Berry, W.B.N., and Orth, C.J., 1989. Palaeo-oceanography and biogeography in the Tremadoc (Ordovician) Iapetus Ocean and the origin of the chemostratigraphy of Dictyonema flabelliforme black shales. Geol. Mag., 126: 19-27.

Wright, J., Schrader, H., and Holser, W.T., 1987. Paleoredox variations in ancient oceans recorded by rare earth elements in folssil apatite. Geochim. Cosmochim. Acta, 51: 631-644.

Wood, D.A., 1980. The application of a Th-Hf-Ta Diagram to problems of tectonmagmatic classification and to establishing the nature of crustal contamination of basltic lavas of the British Tertiary volcanic province. Earth Planet. Sci. Letts, 50: 11-30.

Wood, D.A., Joron, J.-L., and Treuil, M., 1979. A Re-appraisal of the use of trace elements to classifiy and Discriminate between magma seiries erupted in different tectonic settings. Earth Planet. Sci. Letts., 45: 326-336.

Yen, T.F., 1975. Chemical aspects of metals in native petroleum. In: The Role of Trace Metals in Petroleum, T.F. Yen, ed. Ann Arbor Science Publishers, Inc., Ann Arbor, MI, pp. 1-30.

Zhou, J.Y., McDuff, R., and Murray, J.W., 1982. The distribution of vanadium, chromium and manganese in the northwest Pacific. EOS, 63: 989-990.

Table 1 Summary of samples for study of elemental chemistry of black shales
LiassicEngland, UK10WBNB
CarboniferousWesphalianWales, UK18RAR
SilurianLandoveryScotland, UK7WBNB
LandoveryScotland, UK20SMCU
LandoveryWales, UK2WBNB
LandoveryNew Brunswick, Canada1WBNB
LandoveryMaine, USA1WBNB
OrdovicianAshgillScotland, UK38WBNB
AshgillScotland, UK17SMCU
upper-middleNew York, USA1WBNB
upper-middleNew Jersey, USA7UCMP
middleWales, UK1WBNB
middlePenna, USA8LP
middlePenna, USA3UCMP
TremadocLevis, Canada6WBNB
TremadocWales, UK14WBNB
TremadocNew York, USA1WBNB
TremadocNew Brunswick, Canada5SMCU
* Sources:
WBNB: William B. N. Berry, University of California, Berkeley
SMCU: Sedgwick Museum, Cambridge University (mainly from theBulman collection)
UCMP: Museum of Paleontology, University of California, Berkeley
LP: Lucien Platt, Bryn Mawr University
RAR: Robert A. Raisewell, Leeds University
LUC: Lund University Collection, Sweden
BCPC: Bolivian California Petroleum Company, La Paz, Bolivia

Table 2. Average Concentrations of REDOX Indicators in Low-Calcic Shales
(Concentrations in ppm)

Group 1 OxicGroup 2 Mn-SolubleGroup 3 Mn,Fe-SolGroup 4 V-High

Mn > 800Mn < 750Mn < 750Mn < 750
Fe > 37500Fe > 37500Fe < 37500 Fe < 37500
V < 320V < 320V < 320V > 320