Trace metal records of regional paleoenvironmental variability in Pennsylvanian (Upper Carboniferous) black shales

Chemical Geology 206 (2004) 319 – 345 www.elsevier.com/locate/chemgeo

Trace metal records of regional paleoenvironmental variability in Pennsylvanian (Upper Carboniferous) black shales Anna M. Cruse a,*, Timothy W. Lyons b b

a U.S. Geological Survey, P.O. Box 25046, MS977, DFC, Denver, CO 80225, USA Department of Geological Sciences, University of Missouri, Columbia, MO 65211, USA

Abstract Regional geochemical differences within a laterally continuous, cyclic Pennsylvanian (Upper Carboniferous) shale in midcontinent North America are interpreted in light of models of glacioeustatic forcing and new views on water-column paleoredox stability and trace-metal behavior in black shale environments. Specifically, we characterize differences in transition metal (Fe, Mn, Mo, V, Ni, Zn, Pb and U) concentrations in black shales of the Hushpuckney Shale Member of the Swope Limestone in Iowa and equivalent black shale beds of the Coffeyville Formation in Oklahoma. Although C – S – Fe systematics and uniform 34S-depleted isotope ratios of pyrite indicate pervasive euxinic deposition (anoxic and sulfidic bottom waters) for these shales, regional variations can be inferred for the efficiency of Mo scavenging and for the rates of siliciclastic sedimentation as expressed in spatially varying Fe/Al ratios. Black shales in Iowa show Mo enrichment roughly five times greater than that observed in coeval euxinic shales in Oklahoma. By contrast, Fe/Al ratios in Oklahoma shales are as much as five times greater than the continental ratio of 0.5 observed in the over- and underlying oxic facies and in the coeval black shales in Iowa. Recent work in modern marine settings has shown that enrichments in Fe commonly result from scavenging in a euxinic water column during syngenetic pyrite formation. In contrast to Fe, the concentrations of other transition metals (Mo, V, Ni, Pb, Zn, U) are typically more enriched in the black shales in Iowa relative to Oklahoma. The transition metal trends in these Paleozoic shales are reasonably interpreted in terms of early fixation in organic-rich sediments due to euxinic water-column conditions. However, regional variations in (1) rates of siliciclastic input, (2) organic reservoirs, including relative inputs of terrestrial versus marine organic matter, and (3) additional inputs of metals to bottom waters from contemporaneous hydrothermal vents are additional key controls that lead to geographic variation in the extent of metal enrichments preserved in ancient organic-rich sediments. Published by Elsevier B.V. Keywords: Trace metals; Iron; Molybdenum; Sulfur; Hydrothermal

1. Introduction

* Corresponding author. Tel.: +1-303-236-9379; fax: +1-303236-3203. E-mail addresses: [email protected] (A.M. Cruse), [email protected] (T.W. Lyons). 0009-2541/$ - see front matter. Published by Elsevier B.V. doi:10.1016/j.chemgeo.2003.12.010

Anoxia in the water column arises from the complex interplay of (1) restricted lateral circulation in the deep-water column, (2) thermohaline stratification, and (3) levels of primary productivity and export of organic carbon from surface waters, such

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that respiratory oxygen demand in lower levels of the water column exceeds renewal. The interplay of these processes results in a tripartite layering of water-column chemical conditions, with well-oxygenated surface waters separated from anoxic or anoxic – sulfidic (euxinic) bottom waters by a narrow transition zone known as the chemocline. The chemocline is a critical interface between two fundamentally different chemical environments. This interface is a site of enhanced chemical and biological processes, including redox cycling of transition metals (German et al., 1991; Jørgensen et al., 1991; Lewis and Landing, 1991; Shaw et al., 1994; Spencer et al., 1972; Van Cappellen et al., 1998), enhanced precipitation of authigenic apatite (Burnett et al., 1982), and increased biological activity associated with enhanced nutrient cycling (Mullins et al., 1985). Impingement of the chemocline on a basin margin with attendant redox cycling of metals has been proposed as a key factor in the development of certain ore deposits, such as Mn (Force and Cannon, 1988). Recent research in the Black Sea, the world’s largest anoxic basin, has highlighted the dynamic nature of the chemocline, with natural and possible anthropogenic vertical fluctuations of tens of meters on time scales of hours (Kempe et al., 1990) to decades (Murray et al., 1989). Other studies have revealed a robust, integrated geochemical record of natural fluctuations in the Black Sea on 102- to 103-year time scales (Lyons et al., 1993; Repeta, 1993) and striking paleoredox variation spanning glacial – interglacial transitions in a variety of Quaternary silled marginal marine settings such as the Black Sea and the Cariaco Basin, Venezuela (Arthur and Dean, 1998; Lyons et al., 2003; Wilkin and Arthur, 2001; Yarincik et al., 2000). Recently, a number of workers have challenged the Black Sea paradigm of a deep, highly stratified, restricted environment for black shale deposition, noting instead the comparatively shallow water depths and the time-varying stability in water-column stratification for many oxygen-deficient settings (Murphy et al., 2000; Sageman and Lyons, 2003; Sageman et al., 2003). An ancient record of dynamic redox conditions is preserved in vertical transitions from gray to black shale in Pennsylvanian cyclothems of midcontinent

North America. Although these shales have been the subject of extensive work (e.g., Coveney and Glascock, 1989, and references therein), fundamental questions remain. For example, the specific depositional environment is a matter of debate, with interpretations ranging from nearshore, shallow-water environments (Coveney et al., 1991; Zangerl and Richardson, 1963) to offshore, deep-water settings (Heckel, 1977, 1991; Heckel and Hatch, 1992). Similarly, whereas these midcontinent shales have long been recognized as containing elevated concentrations of a broad suite of transition metals (Coveney and Glascock, 1989; Coveney et al., 1987; Vine and Tourtelot, 1970), the mechanisms responsible for these enrichments remain debated. Early workers advocated models wherein metal mineralization occurred during sedimentation, through interactions between organic matter and seawater or through direct precipitation of sulfide minerals from anoxic waters (e.g., Coveney and Glascock, 1989). These models have been modified to include the possibility that metal concentrations were augmented by epigenetic interactions with meteoric waters and migrating basinal brines (Coveney, 1992; Coveney and Glascock, 1989; Coveney et al., 1987). Beyond regional interest, unraveling the mechanisms responsible for metal enrichment in these units has general relevance in recognizing the processes responsible for the accumulation or release of transition metals from sediments under conditions of dynamic water-column redox. Here, we apply new paradigms for the controls on C – S –Fe accumulation in marine sediments developed from our work in modern anoxic basins such as the Black Sea, Effingham Inlet (an anoxic fjord in British Columbia) and the Cariaco Basin (e.g., Hurtgen et al., 1999; e.g., Lyons, 1997; Lyons et al., 2003), as well as from recent studies on the sequestration of redox-sensitive metals in sediments (Crusius et al., 1996; Morford and Emerson, 1999), to understand the processes that control metal enrichments in Pennsylvanian cyclothems. Within this context, we interpret the regional and temporal geochemical variations recorded in an individual gray – black (cyclothemic) shale sequence. Specifically, high-resolution sampling of the Upper Pennsylvanian (Missourian) Hushpuckney Shale Member of the Swope Limestone in Iowa and stratigraphically equivalent black shales of the

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Coffeyville Formation in Oklahoma is used to illustrate geographic variations in the processes that control metal enrichments preserved in ancient organicrich sediments that must also be considered when using such enrichments as proxies for bottom-water redox conditions.

2. Geologic setting The Pennsylvanian System of midcontinent North America is characterized by well-developed cycles of shale and limestone commonly known as cyclothems with thicknesses in the range of meters to tens of meters (Crowell, 1978; Heckel, 1977, 1986; Wanless and Shepard, 1936). Although autocyclic processes may contribute to the observed patterns (Ferm, 1975), especially near the basin margins, a globally synchronous distribution of upper Paleozoic cyclic strata has fostered a widely held, but still controversial, opinion emphasizing eustatic control driven by the waxing and waning of Gondwanan glaciers (Crowell, 1978; Heckel, 1980; Wanless and Shepard, 1936)—possibly on Milankovitch time scales (Heckel, 1986). Some studies have stressed regional variations in North American cyclothems and thus have argued for a major tectonic signal linked to concurrent Apppalachian – Ouachita orogenesis. Such controls on local sea level, whose relative strengths would vary as a function of paleoceanographic position, may strongly overprint the record of glacioeustatic (climatically driven) forcing (de Klein, 1992; de Klein and Kupperman, 1992). The classic ‘‘Kansas-type’’ cyclothem, as defined by Heckel (1977), is depicted in Fig. 1. Thin ( < 1 m) laminated, black, organic- and metal-rich, phosphatic shales of the offshore shale facies can be correlated throughout the Midcontinent, with surface and subsurface distributions ranging between 20,000 and 100,000 km2 (Coveney and Glascock, 1989). Paleoenvironmental reconstructions of the black shale facies and associated gray shales have been at the center of the highly polarized debate, with interpretations ranging from the deepest offshore facies (Heckel, 1977) to shallow nearshore (Zangerl and Richardson, 1963). Furthermore, these sequences were commonly exposed to subarial weathering during sea-level lowstands, as recorded by paleosols

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capping the uppermost outside shale units and meteoric cements in the regressive limestones (Goebel et al., 1989; Heckel, 1983). These secondary processes must be considered in any rigorous model for metal distributions. The black shales of the Swope cyclothem are the focus of this study. The Swope cyclothem is represented by the Swope Limestone of the Kansas City Group, within which the Hushpuckney Shale Member is the offshore shale facies (Fig. 1). Samples of the Hushpuckney Shale Member and a black shale bed of the Coffeyville Formation of the Skiatook Group were sampled in Iowa and Oklahoma from cores IRC and C-TW-1, respectively (see Fig. 2 for geographic locations). The Coffeyville Formation in Oklahoma was thought to be equivalent to the stratigraphic interval in southern Kansas lying between the Checkerboard Limestone and the Dennis Limestone, which includes the Swope Limestone (Oakes, 1952; Zeller, 1968). Although the Coffeyville Formation was defined from a type location near Coffeyville, Kansas (Schrader and Haworth, 1905), the name has been used almost exclusively in Oklahoma (Watney and Heckel, 1994). Heckel (1992) and Watney and Heckel (1994), however, redefined the Coffeyville Formation as the Coffeyville Group in southern Kansas and northern Oklahoma and correlated this group with the top of the Pleasanton Group and base of the Kansas City Group of northern Kansas, Missouri and Iowa. Within the redefined Coffeyville Group, Watney and Heckel (1994) interpreted the dark (organic-rich) Lower Tacket and Upper Tacket Shales of the Tacket Formation in northern Oklahoma and southern Kansas as being equivalent to the Mound City Shale Member of the Hertha Limestone and the Hushpuckney Shale Member of the Swope Limestone, respectively. Although the core descriptions of C-TW-1 published in 1988 do not specifically delineate individual members of the Coffeyville Formation (Hemish, 1988), Heckel (1992; his Fig. 5) established that the second black shale bed above the Checkerboard Limestone— which is the black shale sampled in this study—is equivalent to the Hushpuckney Shale Member of the Swope Limestone. Lithologically, the sampled intervals are closely similar and contain the gray shale – black shale –gray shale sequence contained in the classic cyclothem

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Fig. 1. Generalized stratigraphic section of a ‘‘Kansas-type’’ Pennsylvanian cyclothem from the North American midcontinent, including the stratigraphic units under consideration in this study. Cyclothem figure is modified from Heckel (1977); stratigraphic names are from Heckel and Watney (2002).

model (Heckel, 1977). The black shale section in core C-TW-1 is approximately 40% thicker than that in core IRC. The Hushpuckney Shale Member is currently deeper in the subsurface in Iowa—31 m versus 8 m for the top of the uppermost gray – black shale boundary—but this difference in depth is not enough to yield a large difference in thermal-maturity levels of the organic matter. Detailed organic geochemical analyses were not part of this study, but work by Hatch and Newell (1999) on samples from core IRC indicated that that the organic matter in this core is generally immature with respect to oil generation. No coals beds are observed in either core.

3. Sampling and analytical methods The drill cores were sampled at centimeter-scale intervals emphasizing transitions between gray and black lithologies. Sedimentological and paleontological models for water-column paleoredox indicate that oxic and anoxic depositional environments can be distinguished by the presence of laminated versus bioturbated sediments (Rhoades and Morse, 1971; Savrda and Bottjer, 1986, 1987, 1991; Savrda et al., 1984). The gray – black shale boundaries, which we assume to correspond to transitions between oxic (gray shale) and anoxic (black shale)

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Fig. 2. Location map showing state boundaries, subarial distribution of the Missourian outcrop belt (stippled area), and the locations of the C-TW-1 and IRC drill-core sites. Base map modified from Heckel (1980). Core IRC was drilled in the SE SE SW sec. 20, T. 67N., R. 41 W., Fremont County, IA. Core C-TW-1 was drilled in the SE1/4 SW1/4 NE1/4 NE1/4 SE1/4 sec. 6, T. 22 N., R. 13 E., Tulsa County, OK (Hemish, 1988).

water-column conditions and corresponding differences in organic carbon accumulation, were determined based on the initial and final appearances of laminations in X-radiographs of the drill core. Therefore, some samples that may have elevated organic carbon concentrations (>2.0 wt.% Corg) are considered to be in the gray shale facies because they lack laminations in the X-radiographs. Assuming that (1) the transitions between gray and black shales are linked to changing sea level (local or global) and (2) that the black, organic-rich shales represent the deepest water facies, the under- and overlying gray shales are considered to be transgressive and regressive facies, respectively. Samples were selected from the cores to emphasize the boundaries between the gray and black shales. The samples were crushed to fine powders using a SPEX 8000 mixer/mill and an alumina ceramic vial set. The

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outer edges of the samples were cleaned with double deionized (DD) H2O to avoid contamination that might have been associated with the drilling process. Sample powders were stored in polyethylene vials prior to analysis. Total and inorganic carbon (TIC) concentrations were determined by liberating the carbon through combustion at 950 jC and by acidification with 2 N HCl, respectively. The amount of CO2 generated by each method was then quantified by coulometric titration. Organic carbon (Corg) concentrations were calculated by difference between the total and inorganic carbon concentrations. Analyses of pure CaCO3 standards were reproducible to within 0.1 wt.% by this method. Total reduced inorganic sulfur (TRIS) concentrations were measured using the chromium reduction method described in Zhabina and Volkov (1978) and Canfield et al. (1986). TRIS, in general, comprises pyrite, iron monosulfide phases and elemental sulfur (Canfield et al., 1986). However, because iron monosulfides and elemental sulfur are typically transient species over geologic time scales (Berner, 1970; Hurtgen et al., 1999), TRIS in ancient shale samples can be considered to represent pyrite sulfur (Spy) in the absence of other base metal sulfides. Given the order of magnitude difference between the concentrations of Fe and metals such as Zn which also form sulfides, this assumption is valid for our samples. Replicate analyses of freshly ground, pure pyrite standards yielded sulfur recoveries that were consistently greater than 96 wt.%. The extent to which the original total reactive iron reservoir has been transformed to pyrite has traditionally been assessed by calculating degree-of-pyritization values (DOP; Berner, 1970; Raiswell et al., 1988). DOP is expressed as: Fepy ; Fepy þ FeHCl

ð1Þ

where Fepy is the pyrite-iron concentration calculated from measured values of Spy. FeHCl is the remaining, unsulfidized portion of the ‘‘reactive’’ iron reservoir that is extracted by boiling for 1 min in 12 N HCl (Berner, 1970; Raiswell et al., 1988). Extracted iron was analyzed spectrophotometrically using the ferrozine method of Stookey (1970). The standard devia-

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tion for replicate iron extractions from a given sample was less than 9.5% of the mean, which results in an error in calculated DOP values of only 1%. Recent work has refined our understanding of Fe reactivity and the primary depositional controls on DOP values in modern and ancient oxygen-deficient settings (Canfield et al., 1996; Lyons et al., 2003; Poulton and Raiswell, 2002; Raiswell and Canfield, 1998). In this model, high DOP values (>0.7) in sediments that have not been mineralized epigenetically are unambiguously interpreted as indicators of euxinic deposition, while low values point to oxic, organic-deficient

environments. Intermediate values are less straight forward to interpret and can reflect euxinic deposition under conditions of rapid siliciclastic accumulation or an oxic to dysoxic, organic-rich setting (Lyons et al., 2003). The stable sulfur isotopic composition of FeS2 was measured by combusting aliquots of Ag2S that were precipitated from H2S liberated using the chromium reduction method (Canfield et al., 1986; Newton et al., 1995; Zhabina and Volkov, 1978). Combustion was conducted on-line in an elemental analyzer in the presence of V2O5 as a catalyst for a quantitative

Fig. 3. Section of core from drillhole C-TW-1 in northeastern Oklahoma. Plotted are concentrations of (A) organic carbon (Corg; diamonds) and inorganic carbon (TIC; triangles), (B) pyrite-sulfur (Spy), (C) degree-of-pyritization values, and (D) d34Spy versus depth. The stippled area corresponds to the black shale facies containing laminations indicative of deposition under an anoxic/euxinic water column.

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Table 1 Geochemical data for core C-TW-1, Oklahoma Sample

Deptha (cm)

Corgb (wt.%)

289.3-14GR 289.3-10GR 289.3-9GR 289.3-8GR 289.3-7GR 289.3-6GR 289.3-5GR 289.3-4GR 298.3-3GR 289.3-2GR 289.3-1GR

8789.65 8807.00 8808.20 8809.60 8810.50 8812.20 8813.20 8814.10 8815.05 8816.17 8818.40

0.81 0.91 0.90 0.88 0.98 0.94 0.94 1.03 1.16 1.16 1.76

TICb (wt.%)

Spyb (wt.%)

0.31 0.18 0.16 0.16 0.12 0.15 0.13 0.02 0.07 0.62 0.75

0.10 0.04 0.04 0.05 0.04 0.05 0.05 1.03 1.55 2.02 1.62

Regressive gray – black shale boundary: 8820.12 cm depthc 289.3-1B 8821.15 6.65 0.64 3.21 289.3-2B 8822.85 10.28 0.30 5.72 289.3-3B 8824.20 10.39 0.19 6.37 289.3-4B 8826.00 9.35 0.68 8.46 289.3-5B 8827.50 11.95 0.59 6.11 289.3-6B 8828.70 11.70 0.50 6.73 289.3-7B 8829.95 12.25 0.36 4.57 289.3-8B 8831.50 10.04 0.63 6.06 289.3-9B 8833.40 10.68 0.64 6.86 289.3-12B 8841.68 13.71 0.51 2.36 289.3-13B 8846.93 7.42 0.02 12.91 289.3-15B 8857.03 8.95 0.74 9.77 294.0-19B 8919.95 9.43 0.32 3.69 294.0-17B 8931.25 10.45 0.78 3.77 294.0-16B 8938.00 5.92 0.36 4.37 294.0-15B 8944.15 5.18 0.34 1.77 294.0-14B 8945.70 294.0-13B 8947.40 4.64 0.32 1.56 294.0-12B 8949.65 6.17 0.41 1.22 294.0-9B 8951.55 5.43 1.24 1.29 Transgressive gray – black shale 294.0-1GT 8955.30 294.0-2GT 8956.45 294.0-3GT 8957.55 294.0-4GT 8958.60 294.0-5GT 8959.70 294.0-6GT 8960.70 294.0-7GT 8961.55 294.0-8GT 8962.50 294.0-9GT 8963.65 294.0-10GT 8965.00 294.0-11GT 8966.50 294.0-12GT 8968.25 294.0-13GT 8969.90 294.0-14GT 8971.30 294.0-15GT 8972.70

boundary: 8954.51 cm depthc 6.08 1.39 2.90 5.51 1.24 2.47 4.77 1.25 1.58 3.00 1.52 1.86 1.97 1.61 1.45 1.96 1.69 1.24 2.20 1.78 1.24 2.30 1.90 1.11 2.37 1.82 1.33 2.65 1.70 1.39 2.97 1.52 1.22 2.02 1.59 1.21

d34Sb (x , CDT)

 0.34 0.64 2.4  0.18

 22.5  24.0  26.0  24.1

 33.7  32.9  26.5  33.9  36.3

 29.6

 31.0  28.2  24.5  23.4  22.7  23.4  21.3  27.1  26.8

Fet (wt.%)

FeHClb (wt.%)

Fepyb (wt.%)

DOPb

3.78 3.18 3.04 2.98 2.99 3.03 3.09 4.70 4.29 4.09 3.75

2.48 1.85 1.71 1.71 1.59 1.62 1.72 2.52 2.07 1.18 1.18

0.09 0.04 0.03 0.04 0.03 0.04 0.04 0.99 1.35 1.76 1.41

0.04 0.02 0.02 0.02 0.02 0.03 0.02 0.28 0.39 0.60 0.54

5.36 7.39 8.04 8.53 7.47 8.28 6.51 7.43 8.33 4.24 14.9 10.4 5.54 5.91 6.07 3.94 3.84 3.69 3.24 3.59

1.34 1.11 0.98 1.12 1.13 1.24 1.04 1.28 1.30 0.95 1.88 1.28 0.95 0.93 0.97 1.00

2.80 4.98 5.55 7.37 5.32 5.86 3.98 5.28 5.97 2.06 11.3 8.51 3.21 3.28 3.80 1.54

0.68 0.82 0.85 0.87 0.82 0.83 0.79 0.81 0.82 0.68 0.86 0.87 0.77 0.78 0.80 0.61

0.89 0.89 1.10

1.36 1.06 1.13

0.60 0.54 0.51

4.81 4.29 3.76 3.95 3.60 3.37 3.44 3.25 3.47 3.45 3.41 3.28 3.01 3.32 3.66

1.22 1.14 1.11 1.25 1.15 1.29 1.22 1.20 1.19 1.15 1.13 1.17

2.53 2.15 1.38 1.62 1.27 1.08 1.08 0.97 1.16 1.21 1.06 1.05

0.67 0.65 0.55 0.57 0.52 0.46 0.47 0.45 0.49 0.51 0.48 0.47

(continued on next page)

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Table 1 (continued) Sample

Deptha (cm)

Al (wt.%)

Transgressive gray – black shale boundary: 289.3-14GR 8789.65 8.97 289.3-10GR 8807.00 9.51 289.3-9GR 8808.20 9.38 289.3-8GR 8809.60 9.67 289.3-7GR 8810.50 10.37 289.3-6GR 8812.20 9.59 289.3-5GR 8813.20 9.49 289.3-4GR 8814.10 8.80 298.3-3GR 8815.05 8.24 289.3-2GR 8816.17 8.04 289.3-1GR 8818.40 7.73

Mn (ppm)

Mob (ppm)

8954.51 cm depthc 585 386 348 365 370 317 330 264 25 229 32 372 14 424 9

Regressive gray – black shale boundary: 8820.12 cm depthc 289.3-1B 8821.15 8.15 562 6 289.3-2B 8822.85 6.82 483 13 289.3-3B 8824.20 6.77 381 16 289.3-4B 8826.00 4.00 497 34 289.3-5B 8827.50 5.89 573 289.3-6B 8828.70 6.18 576 289.3-7B 8829.95 6.83 418 15 289.3-8B 8831.50 6.49 457 289.3-9B 8833.40 5.81 513 31 289.3-12B 8841.68 6.28 410 289.3-13B 8846.93 5.30 361 80 289.3-15B 8857.03 5.85 607 33 294.0-19B 8919.95 7.22 211 142 294.0-17B 8931.25 6.52 387 294.0-16B 8938.00 7.05 344 294.0-15B 8944.15 8.34 348 294.0-14B 8945.70 7.77 312 10 294.0-13B 8947.40 7.50 284 294.0-12B 8949.65 8.19 336 294.0-9B 8951.55 7.42 512 Transgressive gray – black shale 294.0-1GT 8955.30 294.0-2GT 8956.45 294.0-3GT 8957.55 294.0-4GT 8958.60 294.0-5GT 8959.70 294.0-6GT 8960.70 294.0-7GT 8961.55 294.0-8GT 8962.50 294.0-9GT 8963.65 294.0-10GT 8965.00 294.0-11GT 8966.50 294.0-12GT 8968.25 294.0-13GT 8969.90 294.0-14GT 8971.30 294.0-15GT 8972.70 a

boundary: 6.42 6.63 7.26 7.38 7.18 7.63 7.08 7.21 6.76 7.36 7.42 6.84 6.72 6.62 7.24

8954.51 cm depthc 578 541 51 526 747 14 745 751 7 779 5 766 5 716 3 707 4 617 2 673 2 709 600 785 9

V (ppm) 132 148 154 140 152 159 156 154 145 174 223

Nib (ppm) 48

37

70 85 56 55

Zn (ppm) 131 99 110 117 84 118 98 117 263 321 163

398 588 663 449 752 728 719 649 700 767 834 1.05  10 3 1.90  10 3 990 602 248 239 207 190 246

172 220 228 219 195 221 206 167 142 193 107 103 163 207 120 88 72 93 132 147

206 272 321 785 240 256 194 299 222 111 123 270 1.56  10 3 603 836 275 222 180 175 195

274 270 220 191 154 131 126 134 110 103 118 106 117 121 133

159 146 89 110 80 61 80 50 84 67 88 66

878 247 148 140 172 108 90 73 65 61 44 62 57 51 62

Below the surface. b Blank values indicate no measurement was made on that sample. c Based on first/last appearance of laminations in X-radiographs of cores.

Pbb (ppm)

87 67 23 22

65 99 131 242

128 245 347 206 87

48

66 41 29 26 23 24 27 19 23

36

U (ppm) 3.3 2.9 4.1 4.0 4.2 3.8 4.6 4.1 4.2 4.1 4.9

7.0 6.4 7.0 101 8.8 7.9 7.3 8.0 10.1 8.3 6.4 15.1 56.5 51.8 34.4 13.5 11.9 16.3 8.3 7.4

34.6 8.5 7.4 9.1 7.2 6.7 6.6 5.7 5.5 5.1 5.4 5.4 5.8 6.0 6.1

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conversion to sulfur dioxide and analyzed via continuous-flow mass spectrometry. Sulfur isotope data are reported using standard delta notation, where d34S is expressed in permil (x) units, and are calculated as the relative difference in the 34S/32S ratio of the sample, Rs, and the Vienna Canyon Diablo Troilite standard, RVCDT:   Rs  Rvcdt d34 S ¼  1000 ð2Þ Rvcdt Sulfur isotope measurements on replicate samples were reproducible to within F 0.2x. The concentrations of metals (Fe, Ti, Al, K, Mn, V, Ni, Zn, U) were determined for the powdered, whole-rock samples using instrumental neutron activation analysis (INAA) performed at the Missouri University Research Reactor (MURR). Mo and Pb concentrations in the powders were determined via multi-acid digestion and inductively coupled plasma atomic emission spectrometry (ICP-AES) analysis, as described by Lyons et al. (2003).

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4. Results 4.1. Carbon –sulfur– iron concentrations The patterns in C – S –Fe concentrations for these gray – black shale sequences are discussed in detail in Cruse and Lyons (2000) and are reviewed here because of their importance as indicators of watercolumn redox and siliciclastic input. Concentrations of Corg in the black shale facies of C-TW-1 and IRC are greatly enriched relative to the over- and underlying gray shales, with values ranging from 5 to 13 wt.% in C-TW-1 and from 7 to 35 wt.% in IRC (Figs. 3A and 4A; Tables 1 and 2). The transgressive gray shales in C-TW-1 contain between 2 and 6 wt.% Corg, with the concentrations increasing approximately 4 cm above the first appearance of laminations, which marks the transgressive gray – black shale boundary (8955 cm depth; Fig. 3A) and presumably the onset of anoxic bottom-water conditions. In contrast, Corg concentrations decrease abruptly across the regressive gray – black boundary (8820 cm depth; Fig. 3A)

Fig. 4. Section of core from drillhole IRC in southwestern Iowa. Plotted are concentrations of (A) organic carbon (Corg; diamonds) and inorganic carbon (TIC; triangles), (B) pyrite-sulfur (Spy), (C) degree-of-pyritization (DOP) values, and (D) d34Spy versus depth. The stippled area corresponds to the black shale facies containing laminations indicative of deposition under an anoxic/euxinic water column.

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from 6.7 wt.% in the upper laminated interval to 1.8 wt.% in the basal regressive gray shale. In IRC, the transgressive gray shales contain between 0.3 and 2 wt.% Corg, with a sharp increase in Corg concentrations at the transgressive gray – black shale boundary (31,683 cm depth; Fig. 4A). Concentrations of Corg in the IRC regressive gray shale decrease from 9.3 to 0.3 wt.% with distance upcore from the regressive gray – black shale boundary (31,609 cm depth; Fig. 4A). Concentrations of carbonate carbon (total inorganic carbon or TIC) in the black shale facies range between < 0.7 and 0.8– 1.2 wt.% in C-TW-1 and IRC, respectively. TIC concentrations in C-TW-1 are highest in the

transgressive gray shales and are low throughout the black and regressive gray shales (Fig. 3A). The black shale in IRC contains between 0.2 and 1.5 wt.% TIC (1.7 – 12.5 wt.% CaCO3), with concentrations in both gray shales increasing with distance away from the gray –black shale boundaries (Fig. 4A). Note that the two deepest analyzed samples of the transgressive gray shale facies in IRC, 1039.23BT and 1039.2-4BT, are shaley limestones or marlstones. The Spy concentrations in C-TW-1 and IRC show very different patterns with depth (Figs. 3B and 4B; Tables 1 and 2), despite gross similarities in Corg and TIC trends (Figs. 3A and 4A). In core C-TW-1, Spy

Table 2 Geochemical data for core IRC, Iowa Sample

Deptha (cm)

1036.7-7GR 1036.7-6GR 1036.7-5GR 1036.7-4GR 1036.7-3GR 1036.7-2GR 1036.7-1GR

31,589.46 31,593.11 31,596.66 31,599.46 31,602.31 31,605.21 31,608.36

Corg (wt.%) 0.29 0.46 0.58 0.71 1.31 3.79 9.31

TIC (wt.%)

Spy (wt.%)

5.35 1.08 0.60 0.63 0.86 0.88 0.77

1.67 2.09 2.87 3.27 2.20 1.80 1.24

Regressive gray – black shale boundary: 31,609.70 cm depthc 1036.7-2B 31,611.66 11.31 0.76 1036.7-3B 31,615.66 9.19 0.58 1036.7-4B 31,619.86 9.62 0.69 1036.7-5B 31,623.34 14.41 1.46 1036.7-6B 31,626.36 19.24 1.15 1036.7-7B 31,629.89 17.16 0.24 1036.7-8B 31,634.41 15.96 1.03 1036.7-9B 31,638.96 21.95 1.01 1039.2-10B 31,646.43 6.65 1.26 1039.2-9B 31,649.08 11.41 0.77 1039.2-8B 31,652.83 23.37 0.53 1039.2-7B 31,656.68 11.18 1.23 1039.2-6B 31,660.03 24.59 0.57 1039.2-5B 31,664.98 28.29 1.10 1039.2-4B 31,668.08 33.08 0.62 1039.2-3B 31,672.28 24.34 1.24 1039.2-2B 31,676.83 25.37 0.81 1039.2-1B 31,680.93 15.25 0.99

1.28 1.31 1.54 1.71 1.59 1.44 1.95 1.57 1.36 1.40 1.12 1.94 1.52 1.88 1.21 2.08 1.39 1.52

Transgressive gray – black shale boundary: 31,682.90 cm depthc 1039.2-1GT 31,684.18 2.05 3.09 1.65 1039.2-2GT 31,686.43 0.50 7.17 2.70 1039.2-3GT 31,689.33 0.41 10.9 0.67 1039.2-4GT 31,696.63 0.34 10.3 0.44

d34Sb x , CDT  34.1

 41.6  39.1  38.5

 37.8  36.8

 38.0  33.8

 34.4

 34.4  35.1

 34.9  31.3  29.9

Fet (wt.%)

FeHCl (wt.%)

Fepy (wt.%)

DOP

3.35 3.91 4.85 5.33 4.29 4.06 3.70

0.96 0.85 1.09 0.83 1.30 1.28 1.21

1.45 1.82 2.50 2.85 1.91 1.57 1.08

0.60 0.68 0.70 0.77 0.60 0.55 0.47

3.53 3.44 3.53 3.67 3.35 3.35 4.07 3.17 3.25 3.54 3.13 3.99 3.35 3.48 3.01 3.65 3.15 3.57

1.26 1.01 1.01 1.09 0.96 0.87 1.12 0.98 1.08 1.28 0.95 1.17 0.82 0.84 0.66 0.96 0.86 0.95

1.11 1.14 1.34 1.49 1.39 1.25 1.70 1.37 1.19 1.22 0.98 1.69 1.33 1.64 1.05 1.81 1.21 1.32

0.47 0.53 0.57 0.58 0.59 0.59 0.60 0.58 0.52 0.49 0.51 0.59 0.62 0.66 0.61 0.65 0.58 0.58

3.09 3.22 0.72 0.63

0.85 0.55 0.10 0.19

1.44 2.35 0.58 0.38

0.63 0.81 0.85 0.67

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329

Table 2 (continued) V (ppm)

Ni (ppm)

Zn (ppm)

Pbb (ppm)

67 120 125 127 133 190 345

46 71 70 83 109 139 211

34 44 56 56 56 109 149

16

Regressive gray – black shale boundary: 31,609.70 cm depthc 1036.7-2BR 611.66 6.24 197 71 1036.7-3BR 615.66 6.49 181 1036.7-4BR 619.86 6.37 164 90 1036.7-5BR 623.34 5.13 186 1036.7-6BR 626.36 4.73 147 1036.7-7BR 629.89 6.46 144 372 1036.7-8BR 634.41 5.62 193 1036.7-9BR 638.96 4.66 149 306 1039.2-10BT 646.43 6.45 197 1039.2-9BT 649.08 6.03 174 1039.2-8BT 652.83 4.54 134 1039.2-7BT 656.68 5.70 198 206 1039.2-6BT 660.03 5.39 156 1039.2-5BT 664.98 4.25 148 1039.2-4BT 668.08 4.11 129 428 1039.2-3BT 672.28 4.29 158 1039.2-2BT 676.83 4.66 146 364 1039.2-1BT 680.93 6.08 197

619 407 864 1796 2.53  10 3 2.91  10 3 2.04  10 3 2.45  10 3 589 1.00  10 3 1.68  10 3 1.42  10 3 2.22  10 3 2.17  10 3 2.75  10 3 2.70  10 3 2.35  10 3 759

244 186 209 284 349 266 336 311 161 249 336 223 349 398 384 356 409 286

559 452 960 1.77  10 3 2.73  10 3 1.65  10 3 1.94  10 3 2.43  10 3 346 1.01  10 3 3.07  10 3 1.17  10 3 2.53  10 3 2.88  10 3 2.57  10 3 3.32  10 3 2.07  10 3 702

Transgressive gray – black shale boundary: 31,682.90 cm depthc 1039.2-1GT 684.18 5.63 262 15 1039.2-2GT 686.43 2.06 379 19 1039.2-3GT 689.33 0.50 364 6 1039.2-4GT 696.63 0.86 379

91 24 5 10

Sample 1036.7-6GR 1036.7-5GR 1036.7-4GR 1036.7-3GR 1036.7-2GR 1036.7-1GR 1036.7-1BR

Deptha (cm)

Al (wt.%)

589.46 593.11 596.66 599.46 602.31 605.21 608.36

4.27 7.02 7.74 7.39 7.48 7.33 7.45

Mn (ppm) 379 244 269 272 265 261 237

Mob (ppm) 47

58 39 52

70 24

34 12 4 6

13 16 21

26 39

50 43

38

35 34

22 3 2 8

U (ppm) 7.8 8.3 7.6 8.7 12.6 17.4 23.0

42.0 44.2 69.9 79.0 126 134 135 212 60.3 60.6 64.6 95.6 53.5 132 87.1 187 105 29.2

11.7 4.6 2.3 2.2

a

Below the surface. Blank values indicate no measurement was made on that sample. c Based on first/last appearance of laminations in X-radiographs of cores. b

concentrations (1) differ by a factor of 10 between the two gray facies, averaging approximately 1 wt.% in the transgressive gray shale and 0.l wt.% in the regressive gray shale; and (2) are enriched in the black shales relative to the over- and underlying gray shales, with concentrations in the upper black shales generally higher (2.4 – 12.9 wt.%) than the lower shale samples (1.1 –4.4 wt.% Spy; Fig. 3B). In contrast, the black shale in core IRC contains relatively invariant Spy concentrations, averaging 1.5 wt.% (range: 1.2– 2.1 wt.%); these concentrations are generally lower than those found in the transgressive and regressive gray shales (1.7 – 3.3 wt.%; Fig. 4B).

The trends for DOP versus depth are similar to those observed for Spy concentrations in both cores (Figs. 3D and 6D; Tables 1 and 2). DOP values in the transgressive black shale of core C-TW-1 range from 0.51 to 0.78, whereas those in the regressive facies range from 0.68 to 0.86. DOP values in C-TW-1 drop dramatically across the regressive boundary (from black to gray shale), in sharp contrast to the slight change observed across the transgressive boundary. DOP values in core IRC are largely invariant throughout the black shale facies, ranging between 0.47 and 0.66 (Fig. 4C). As expected from the sulfur trends, these values are less than those found in the IRC

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transgressive and regressive gray shales (0.47 – 0.85; Fig. 4C). 4.2. Sulfur isotope ratios The d34S values of pyrite (d34Spy) for cores C-TW1 and IRC are reported in Tables 1 and 2 and are shown as depth profiles in Figs. 3D and 4D, respectively. Values for d34S in both cores are significantly depleted in 34S relative to coeval seawater sulfate values of f + 13xrecorded in Desmoinesian-aged sulfates from Colorado (Claypool et al., 1980). Within the black shale facies of core C-TW-1, d34S values range from  22.5xto  36.3xand are low (34S-

depleted) relative to values in the gray shales. In the transgressive gray shale facies of C-TW-1, d34S values increase downward from  31.0xto  21.3x with increasing depth below the gray –black shale boundary. The regressive gray shale facies in core C-TW-1 is strongly 34S-enriched relative to the underlying shales, with values between  0.34xand 2.4x(Fig. 3D). Within core IRC, the d34S values are in general lighter than those observed in C-TW-1 and exhibit less variability within both the gray and black shale facies. Within the black shale of IRC, d34S values are between  33.8xand  38.1x(Fig. 4D). The values in the transgressive gray shale are enriched relative to those in the regressive gray shale,

Fig. 5. Normalized metal concentrations in a section of core from drillhole C-TW-1, northeastern Oklahoma. Plotted are values for: (A) Mn/Al; (B) Fe/Al; (C) Mo/Al; (D) U/Al. Dashed lines indicate the metal/aluminum concentration ratios for PAAS shale standard (Table 3; Taylor and McClennan, 1985). Ratios are calculated as wt.% metal/wt.% Al for Fe, and ppm metal/wt.% Al for Mn, Mo, and U. The stippled area corresponds to the black shale facies containing laminations indicative of deposition under an anoxic/euxinic water column.

A.M. Cruse, T.W. Lyons / Chemical Geology 206 (2004) 319–345

ranging from  29.9xto  34.9xand  34.1x to  41.6x, respectively (Fig. 4D). 4.3. Transition metal concentrations Bulk transition metal concentrations are mixtures of both detrital and authigenic components, but it is only the concentrations of the authigenic component that vary in response to changes in water-column redox conditions (unless primary phases are lost [remobilized] under changing redox conditions). Transition metal concentrations reported in Tables 1 and 2 are shown as depth profiles in Figs. 5 and 7 for core C-TW-1 and Figs. 6 and 8 for core IRC. To evaluate metal enrichment relative to the detrital fraction, and avoiding the complications associated with dilution by biogenic fluxes—i.e., calcium carbonate and organic matter (Sageman and Lyons, 2003)—the plotted metal data have been normalized to Al. Despite potential complications intrinsic to the normalization process (Van Os et al., 1993), Al can be treated as a conservative tracer of siliciclastic deposition. Although contamination of the samples with Al during crushing is viewed as a potential problem, it would only act to decrease the magnitude of enrich-

331

ment expressed in the metal/Al ratios. In plots of normalized metal concentrations versus depth for each core (Figs. 5 –8), the ratio calculated for the Post-Archean Average Shale (PAAS; Taylor and McClennan, 1985; Table 3) is indicated by a dashed line to provide an approximate baseline for the continental (detrital) mean. In most cases, general agreement with the oxically deposited gray shales validates the use of PAAS as a proxy for the detrital end member in the midcontinent setting and argues against significant Al contamination during sample preparation. In both cores, transition metal concentrations, with the exception of manganese, are generally elevated throughout the black shale facies relative to the overand underlying gray shales. Manganese concentrations are highest within the gray shales rather than the black shales. Bulk metal concentrations within the black shales are also elevated up to an order of magnitude relative to PAAS. Transition metal concentrations, with the exception of Fe and Pb, are significantly higher in core IRC than in core C-TW-1. For example, the maximum Zn concentration in C-TW-1 is 1557 ppm, with other values ranging from f 200 to 700 ppm. In IRC, the maximum Zn concentration is

Fig. 6. Normalized metal concentrations in a section of core from drillhole IRC, southwestern Iowa. Plotted are values for: (A) Mn/Al; (B) Fe/ Al; (C) Mo/Al; (D) U/Al. Dashed lines indicate the metal/aluminum concentration ratios for PAAS shale standard (Table 3; Taylor and McClennan, 1985). Ratios are calculated as wt.% metal/wt.% Al for Fe, and ppm metal/wt.% Al for Mn, Mo, and U. The stippled area corresponds to the black shale facies containing laminations indicative of deposition under an anoxic/euxinic water column.

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Table 3 Metal/Al ratios in the Post-Archean Average Shale standard (PAAS; Taylor and McClennan, 1985) Metal/Al wt.%/wt.% Fe

0.51

ppm/wt.% Mn Mo V Ni Zn Pb U

90 0.10 15 5.5 8.5 2.0 0.31

2880 ppm, with the remaining concentrations generally ranging from f 1000 to 2000 ppm.

were hypothesized to relate to differences between the siliciclastic flux and the organic carbon reservoir in each core. Differences in metal concentrations between gray and black shales of a given core are dominantly controlled by differences in bottom-water redox conditions. However, as for the C –S – Fe systematics (Cruse and Lyons, 2000), variations in the siliciclastic flux, Corg reservoir, and potential hydrothermal inputs to the Pennsylvanian midcontinent sea result in variations in the levels of transition-metal enrichment in the black shales of each core. Although our study is not intended to be an exhaustive study of regional geochemical variability, recognition of how variations in the principal controlling factors impact transition metal enrichment/depletion processes at two sites within the Pennsylvanian sea serves to strengthen regional mechanistic models for the accumulation and preservation of organic-rich sediments throughout the geologic record.

5. Discussion 5.1. Bottom water redox—Mn While the cores in C-TW-1 and IRC are similar in terms of their overall lithologic sequence (gray shale – laminated black shale –gray shale), there are striking differences in the concentrations of Corg, Spy and redox-sensitive transition metals. The relative accumulation versus depletion of metals in organic-rich sediments are strongly influenced by depositional redox conditions, with enrichment/depletion patterns controlled by (1) the solubilities of the reduced versus the oxidized species as either oxide or sulfide minerals (e.g., Mn, Fe, V, Zn, and Pb); (2) the preferential adsorption of reduced or oxidized species onto, and subsequent cycling with, other authigenic minerals such as Fe-oxyhydroxides or pyrite (V, Zn and Ni); and (3) the preferential adsorption of reduced species with particulate Corg (V, Mo, Ni, and U) (Breit and Wanty, 1991; Crusius et al., 1996; Morford and Emerson, 1999; Pedersen and Calvert, 1990; Shaw et al., 1990; Spencer et al., 1972). Using sedimentological and geochemical criteria, Cruse and Lyons (2000) proposed a model for paleoenvironmental conditions during deposition of the Hushpuckney Shale Member and its equivalents in which euxinic conditions were generally persistent throughout the Midcontinent during deposition of the black shales. However, differences in the geochemical proxies of paleoredox conditions (i.e., C –S – Fe concentrations)

Manganese is depleted in sediments in dysoxic to anoxic environments due to the reductive dissolution of manganese oxyhydroxides (Calvert and Pedersen, 1993; Force and Cannon, 1988; Morford and Emerson, 1999). Ratios of Mn-to-Al throughout the black shale facies of core C-TW-1 and IRC are generally depleted relative to the PAAS standard and the overand underlying gray shales (Figs. 5A and 6A). However, several horizons within the upper portion of the black shale of C-TW-1 are characterized by Mn/Al ratios that are equal to or slightly greater than those in PAAS (Fig. 5A). Under the persistently euxinic watercolumn conditions that existed during black shale deposition in both cores, Mn concentrations became depleted relative to average shale because any Mn delivered to the sediments as oxyhydroxide minerals underwent reductive dissolution and was remobilized. Although the transgressive gray shales in both cores, which were deposited underneath an oxic water column, do have Mn/Al ratios equal to that of PAAS in both cores, the regressive gray shales are marked by depletions in Mn/Al relative to PAAS. This indicates that while the water column was oxic, the sediments were likely reducing, with oxygen penetration depths less than 1 cm (Morford and Emerson, 1999). Under these conditions, Mn is remobilized from sediments

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via reductive dissolution of oxyhydroxide minerals, but enrichments due to reprecipitation as an oxide mineral at the sediment –water interface and subsequent cycling (i.e., the manganese pump of Calvert and Pedersen, 1993) are not preserved. Because euxinic redox conditions lead to a depletion in Mn concentrations in black shales, the concentrations are not affected by such factors as siliciclastic flux or the organic carbon reservoir. 5.2. Siliciclastic flux—Fe Iron is enriched in organic-rich sediments and black shales due to its reduction under anoxic conditions and subsequent fixation as pyrite. A recent model for iron distributions in marine basins indicates that regional patterns of Fe accumulation—as manifested in DOP values and ratios of Fe-to-Al and highly reactive (toward sulfide) Fe-to-total Fe—are largely controlled by the relative proportions of iron delivered (1) with detrital sediments, such as Fe oxides and silicates that are part of the siliciclastic flux and (2) as a fraction that is decoupled from the local detrital flux. This second pool of Fe is derived from the scavenging of dissolved Fe in euxinic waters column by syngenetic pyrite formation in the water column (Canfield et al., 1996; Lyons, 1997; Lyons et al., 2003; Poulton and Raiswell, 2002; Raiswell and Canfield, 1998; Raiswell et al., 2001; Sageman and Lyons, 2003). According to this model, sites of rapid euxinic siliciclastic accumulation show low to intermediate DOP values in association with Fe/Al ratios that are similar to that of siliciclastic fluxes in this study, the ratio of PAAS; (Lyons et al., 2003; i.e., Raiswell et al., 2001; Sageman and Lyons, 2003, and references therein). Euxinic settings with slower sediment accumulation rates can exhibit dramatic augmentation of the total iron pool by the iron component that is scavenged from the water column, which leads to both elevated DOP values and Fe/Al ratios. The black shale facies of core C-TW-1 is characterized by maximum Corg concentrations that are only half those observed in core IRC (14% versus 33%; Figs. 3A and 4A). At the same time, core C-TW-1 has Spy concentrations (Fig. 3B) and Fe/Al ratios (Fig. 5B) that are substantially higher than those in core IRC (Figs. 4B and 6B). The DOP values in core C-

333

TW-1 are generally high (0.51 – 0.78; Fig. 4C) and are in the range of those observed in the modern euxinic Black Sea (Lyons, 1997). Values for Corg, Spy, Fe/Al, and DOP increase upcore within the black shale from the lower (transgressive) gray – black boundary, so that the uppermost black shales are enriched in these components relative to the lower samples. Overall, the C – S – Fe trends at C-TW-1 have been interpreted (Cruse and Lyons, 2000) to reflect deposition beneath a euxinic water column, with scavenging of dissolved iron during formation of pyrite within the water column leading to the elevated Fe/Al ratios within the black shale facies. Again, such scavenged Fe is decoupled from the local detrital flux and is manifest in elevated ratios of total Fe-to-Al, and highly reactive Fe-to-total Fe, relative to those of primary siliciclastic inputs (Canfield et al., 1996; Lyons, 1997; Lyons et al., 2003; Poulton and Raiswell, 2002; Raiswell and Canfield, 1998; Sageman and Lyons, 2003). In contrast, core IRC has lower DOP values and enrichments in Fe compared to core C-TW-1 (Figs. 4C and 6B). More specifically, Fe/Al ratios elevated relative to PAAS are not observed in core IRC (Fig. 6B), despite the euxinic conditions that persisted during deposition (Cruse and Lyons, 2000). One possible explanation is that the elevated ratios expected from iron scavenging in the water column were precluded (swamped) by high inputs of unreactive (toward sulfide) iron that was delivered with the detrital flux. Given that the position of IRC within the basin is not typically considered to be located within the marginal environments where elevated (as compared to the basin center) siliciclastic fluxes are present (Heckel, 1980), this relationship may indicate that previous paleogeographic reconstructions have underestimated the potential for regional variation in the quantity of siliciclastic sedimentation and specifically regarding the relatively high inputs at IRC. As demonstrated by Lyons et al. (2003), siliciclastic sedimentation rates do not have to be dramatically increased in order to swamp the scavenged Fe component. 5.3. Organic carbon—Mo, U, V, Ni, Zn, and Pb The redox-sensitive metals Mo, U, V, Ni, Zn, and Pb are enriched in organic-rich shales under anoxic water-column conditions due to direct or indirect

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interactions with organic matter (Mo, U, V, Ni) or due to precipitation as sulfide minerals or with pyrite (Zn and Pb; Anderson et al., 1989; Coveney et al., 1991; Crusius et al., 1996; Lewan and Maynard, 1982; Morford and Emerson, 1999). Within the black shales of both cores, the Al-normalized concentrations of these metals are enriched relative to the gray shales and the PAAS standard (Figs. 5C – D, 6C –D, 7, and 8). However, the metal/Al ratios in the black shales of core IRC are higher than those in the black shales of core C-TW-1, with the exception of Pb. In core CTW-1, there is little or no correlation observed between Corg and normalized metal concentrations (Fig. 9). In contrast, the Al-normalized metal concentra-

tions in IRC are all well correlated with Corg concentrations, including Zn and Pb, despite their ability to directly precipitate as sulfide minerals (Fig. 10). Of the various redox-sensitive trace metals, Mo appears to be the most robust tracer of sediment deposition under conditions where high concentrations of dissolved sulfide are present either in the water column or sedimentary porewaters (e.g., Adelson et al., 2001; Helz et al., 1996; Lyons et al., 2003; Werne et al., 2002; e.g., Yarincik et al., 2000). Although Mo is thought to be directly scavenged by, and complexed with, organic matter (Coveney et al., 1991; Helz et al., 1996), elevated concentrations of HS within the water column or sedimentary pore-

Fig. 7. Normalized metal concentrations (calculated as ppm metal/wt.% Al) in a section of core from drillhole C-TW-1, northeastern Oklahoma. Plotted are values for: (A) V/Al; (B) Ni/Al; (C) Zn/Al; (D) Pb/Al. Dashed lines indicate the metal/aluminum concentration ratios for PAAS shale standard (Table 3; Taylor and McClennan, 1985). The stippled area corresponds to the black shale facies containing laminations indicative of deposition under an anoxic/euxinic water column.

A.M. Cruse, T.W. Lyons / Chemical Geology 206 (2004) 319–345

waters may be required to facilitate a shift in the dominant Mo species from molybdate to particlereactive thiomolybdate, so that the scavenging can occur (Adelson et al., 2001; Crusius et al., 1996; Erickson and Helz, 2000; Helz et al., 1996; Zheng et al., 2000). Recent work in the Cariaco Basin focusing on the transition from sediments deposited under oxygenated waters during the last glacial episode to those deposited under the persistently euxinic conditions of the last f 14.5 ky shows that pronounced Mo enrichments are confined to the anoxically deposited sediments (Dean et al., 1999; Lyons et al., 2003). However, the presence of water-column hydrogen sulfide does not appear to be a universal requirement for Mo enrichment in marine sediments. A strong correlation between Mo/Al ratios and Corg concentrations in the microlaminated sediments indicates that organic matter plays a key role in this process by enhancing HS production via bacterial sulfate reduction, by providing a substrate for the scavenging of molybdenum as a thiomolybdate species, or both (Lyons et al., 2003). The C-TW-1 and IRC black shales both have Mo/ Al ratios that greatly exceed that of PAAS, with the IRC black shales enriched the most (Figs. 5C and 6C). Variations in the Mo/Al ratio in IRC are strongly coupled to fluctuations in the amount of Corg present (compare Figs. 4A and 6C, also Fig. 10A), whereas very little correlation exists between Mo/Al and Corg in core C-TW-1 (compare Figs. 3A and 5C, also Fig. 9A). In core IRC, where euxinic water-column conditions persisted during deposition of the black shale, Mo was likely efficiently scavenged from the water column and/or underlying sedimentary porewaters. The lower enrichments expressed in the Mo/Al ratios in core C-TW-1 (Fig. 5C) may be related to either (1) differences in the type and amount of organic matter present (Coveney et al., 1991) or (2) efficient scavenging of dissolved sulfide by Fe, which limited Mo enrichment mechanisms (Meyers et al., in revision; Sageman and Lyons, 2003). Coveney et al. (1991) suggested that variations in bulk Mo abundance in stratigraphically equivalent Middle Pennsylvanian black shales reflected deposition in nearshore versus offshore environments. The environmental variation yielded differences in both the type of organic matter deposited (terrestrial versus marine), as well as the inferred pH of porewaters

335

(acidic in nearshore environments and higher in offshore settings). Regional differences in the organic geochemistry of the stratigraphically lower Mecca Quarry Shale Member of the Linton Formation (Coveney et al., 1987) indicate that core IRC may have contained a greater proportion of terrestrial organic matter than core C-TW-1. Humic acids, which are characteristic of terrestrial organic matter, have been shown to be efficient scavengers of Mo, with the efficiency increasing when dissolved sulfide was present in the experimental solutions (Helz et al., 1996). The observed differences in Mo/Al ratios between these two cores may indicate that IRC was located closer to the paleoshoreline as compared to core C-TW-1, which is consistent with the lower observed Fe/Al ratios at IRC. Alternatively, the greater enrichment of Mo/Al in core IRC (Fig. 6C) relative to core C-TW-1 (Fig. 5C) could also reflect differences in the scavenging of dissolved sulfide by reactive Fe in the water column and porewaters. In seawater, Mo is present as the molybdate ion, but at a critical value (switchpoint) of aHS, (e.g., 10 –3.6 – 10– 4.3; Helz et al., 1996) the molybdate ion is converted to particle-reactive thiomolybdate. Therefore, when dissolved sulfide is efficiently fixed by Fe through pyrite formation, watercolumn and sediment concentrations may not build up to the switchpoint value, thereby decreasing the particle reactivity of the Mo reservoir (Meyers et al., in revision). This process serves to limit the extent of Mo accumulation in sediments. High Fe/Al ratios, DOP values and Spy concentrations in core C-TW-1 point to the efficient fixation of sulfide by Fe, whereas in core IRC, Fe that was reactive toward sulfide appears to have been less abundant. Which of these two mechanisms is the key control on Mo sequestration in sediments remains a question that requires additional work. Differences in the type and/or amount of organic matter present in these shales may also be a key control on the enrichment of the other transition metals. Aluminum-normalized concentrations of U, V, Ni, and Zn are all higher in the core IRC black shales (Fig. 8) as compared to the black shales in core C-TW-1 (Fig. 7). Also, the normalized concentrations of the transition metals are generally highly correlated with Corg concentrations in IRC, which is not the case in core C-TW-1 (Figs. 9 and 10). The close correlation

336

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Fig. 8. Normalized metal concentrations (calculated as ppm metal/wt.% Al) in a section of core from drillhole IRC, southwestern Iowa. Plotted are: (A) V/Al; (B) Ni/Al; (C) Zn/Al; (D) Pb/Al. Dashed lines indicate the metal/aluminum concentration ratios for PAAS shale standard (Table 3; Taylor and McClennan, 1985). The stippled area corresponds to the black shale facies containing laminations indicative of deposition under an anoxic/euxinic water column.

between normalized transition metal and Corg concentrations in core IRC may also be further evidence of the greater proportion of terrestrial organic matter in IRC versus C-TW-1. Again, terrestrial organic matter is composed of a greater proportion of humic substances, which are strong chelators of transition metals. For example, the comparatively unreactive nature of the iron in core IRC—as manifest in lower DOP values despite the euxinic bottom-water conditions present during deposition (Cruse and Lyons, 2000)— may be related to high Corg concentrations in IRC as well as to variations in the siliciclastic flux as compared to C-TW-1. High concentrations of organic carbon may have reacted with the iron in such as way as to prevent reaction with sulfide. Humic and fulvic substances are known to be strong chelators of iron in modern aquatic systems, and it is possible that the sedimentary organic matter deposited in core IRC may have reacted with Fe in a similar manner. Among the transition metals, Fe3 + has one of the strongest affinities for humic acids (Sholkovitz and Copland, 1981). The extent of adsorption of humic acids to iron oxyhydroxides increases with decreasing pH (Tipping, 1981), and Fe will outcompete other metals

such as Zn and Cu with natural organic matter (Tipping et al., 2002). Using pyrolysis-gas chromatographic analysis (Leventhal, 1981), Coveney et al. (1987) demonstrated a trend of increasing amounts of terrestrially derived organic carbon in the Mecca Quarry Shale Member—a black shale contained in an older cyclothem—along a trend from Kansas to Indiana. A similar trend may also characterize other Pennsylvanian cyclothemic shales, such that a greater proportion of the organic carbon in core IRC could be terrestrially derived and thus, richer in humic material. Indeed, detailed petrographic analyses reported by Hoffman et al. (1998) for Hushpuckney Member shale samples from several cores in eastern Kansas indicate that the dominant maceral types are vitrinite and inertinite. Liptinitic macerals, which would have been derived from marine algal material, were not detected in their analyses. Although this may be an artifact of preservation conditions, it could also indicate a true dominance of terrestrial organic matter in the black shales of the Hushpuckney Shale Member (Hoffman et al., 1998). A dominance of terrestrial versus marine organic carbon in a given black shale will not directly affect observed Fe/Al ratios. However, a correlation

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Fig. 9. Organic carbon concentrations versus normalized metal concentrations (calculated as ppm metal/wt.% Al) for a section of core from drillhole C-TW-1, northeastern Oklahoma. Plotted are values for: (A) Mo/Al; (B) U/Al; (C) V/Al; (D) Ni/Al; (E) Zn/Al; (F) Pb/Al. Sample 289.3-4BR (8826 cm depth) was excluded because it contained a large phosphate nodule. The correlation coefficients (r2) for the normalized metal concentrations and organic carbon for the black shales, excluding sample 289.3-4B, are given on each plot.

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Fig. 10. Organic carbon concentrations versus normalized metal concentrations (calculated as ppm metal/wt.% Al) for a section of core from drillhole IRC, southwestern Iowa. Plotted are values for: (A) Mo/Al; (B) U/Al; (C) V/Al; (D) Ni/Al; (E) Zn/Al; (F) Pb/Al. The correlation coefficients (r2) for the normalized metal concentrations and organic carbon for black shales are indicated on each plot.

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between the proportion of terrestrial organic matter present and detrital Fe/Al ratios may be observed because these components are derived from a common source region. 5.4. Hydrothermal augmentation—Fe and Pb One puzzling question is the cause of the extreme enrichment of Fe in core C-TW-1. In the Black Sea, the modern type euxinic basin, slowly accumulating sediments in the deep euxinic basin contain Fe/Al ratios that are at most twice that of the local detrital value (Lyons and Berner, 1992; Lyons et al., 2003). Some horizons in core C-TW-1 are enriched as much as six times relative to PAAS (Fig. 5B). These enrichments in Fe also correspond to enrichments in Spy, suggesting the iron is present as pyrite. Petrographic observations do not indicate that the pyrite is present as concretions or nodules that would have formed relatively late during diagenesis. Similarly, 34S-depleted isotopic values of sulfide argue against late formation of pyrite. The extreme enrichments of transition metal concentrations in Pennsylvanian black shales have led other workers to postulate additional sources of metals to the shales in addition to overlying seawater (Coveney and Glascock, 1989; Coveney et al., 1987). For example, assuming a water depth of 10 m, and modern seawater concentrations of Mo, Zn, V, and U, Coveney and Glascock (1989) calculated that f 1.7 million years would have been required to fix the metal concentrations observed in the Mecca Quarry Shale Member, a time far in excess time of estimates for the entire cycle ( < 0.4 my; Heckel, 1986). An obvious source for the additional metals would be the syngenetic venting of hydrothermal fluids into the basin during black shale deposition (Coveney and Glascock, 1989). Mineralized tubes that may be relics of ancient hydrothermal vents have been observed in Pennsylvanian limestones from Kansas (Coveney, 1992), indicating that fluids were likely venting into the basin concomitantly with cyclothem deposition. Also, Pb/Al ratios are elevated in core CTW-1 (Fig. 7D) compared to core IRC (Fig. 8D). Given the proximity of C-TW-1 to an economic-grade Pb –Zn deposit (Tri-State District, Fig. 1), which is thought to have formed in Pennsylvanian –Permian times (Brannon et al., 1996; Symons et al., 1996), the

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differences in Pb concentrations in the two cores may also indicate the possible presence of hydrothermal vents in the Pennsylvanian seas. Modern hydrothermal vent fluids are enriched in the transition metals, with concentrations up to several orders of magnitude greater than that of bottom oxygenated seawater (see Von Damm, 1995 for a review). Iron is the most enriched of the transition metals in such fluids—concentrations can approach 20 mmol/kg fluid (Von Damm, 1995). When hydrothermal fluids are vented at the seafloor, mixing with cold, oxidizing seawater causes the rapid precipitation of both sulfide and oxyhydroxide minerals and metal deposition within the chimney structures and in the sediments near the vents. Precipitation of sulfide minerals that form the ‘‘black smoke’’ commonly associated with modern vents occurs only at those vents with measured exit temperatures >300 – 350 jC. Cooler vent fluids, with temperatures between 150 and 250 jC, such as those found at Middle Valley, northern Juan de Fuca Ridge, or Escanaba Trough, Gorda Ridge, do not generate the ‘‘black smoke’’ because metal and sulfide concentrations are lower than those observed at the hottest vents. However, metal concentrations in these cooler vent fluids are still enriched relative to oxic bottom seawater and thus represent additional source of metals to the basin. For example, iron concentrations measured in hydrothermal fluids from Middle Valley and Escanaba Trough are up to 20 and 10 Amol/kg fluid, respectively, as compared to 5 nmol/kg fluid in oxygenated bottom seawater (Butterfield et al., 1994; Campbell et al., 1994). Hydrothermal vents could thus serve to augment the bottom-water concentrations of the transition metals to levels higher than those resulting from redox processes alone. This model indicates that core CTW-1 was likely closer to a venting system than core IRC, leading to the observed differences in Fe/Al ratios, principally through scavenging during watercolumn pyrite formation—as indicated by the typical euxinic d34Spy relationships. However, the fact that core C-TW-1 was not enriched in the other transition metals, as compared core IRC, indicates that the venting hydrothermal fluid was either not as enriched in those metals or that the organic matter present in core C-TW-1 was not as efficient at fixing metals relative to the organic matter in core IRC.

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The hydrothermal metal enrichment was most likely not a post-depositional process, as we observed no petrographic evidence in these cores, such as crosscutting veins of ore minerals, to indicate that the metal enrichments were epigenetic. The geochemistry of these shales is also consistent with the hypothesis that venting was contemporaneous with deposition. For example, Mo is not enriched in the C-TW-1 black shales relative to IRC, indicating that sulfide was probably not present at the ‘‘switchpoint’’ concentrations. The injection of iron into the bottom-waters by hydrothermal vents may have served to buffer sulfide concentrations below the value required for thiomolybdate formation in the waters near C-TW-1. One would not expect the post-depositional addition of metals to the shales to produce relative depletions of Mo in C-TW-1 as compared to IRC, since large amounts of pyrite in C-TW-1 could have served as an efficient sink for Mo from such fluids (Helz et al., 1996; Huerta-Diaz and Morse, 1992). The hypothesized Pennsylvanian hydrothermal vent fluid may have been rich in metals, but based on the d34Spy in the black shales of core C-TW-1, it was not a significant source of the sulfide. Values of d34Spy in the black shales are isotopically depleted, ranging from  22xto  41x(Tables 1 and 2). If the d34S value of sulfate in Pennsylvanian seawater was f + 13x(Claypool et al., 1980), the Spy values are depleted by f 50x—a value characteristic of pyrite produced as a result of bacterial sulfate reduction under relatively open reservoir conditions, such as a euxinic water column (Lyons, 1997). The uniformity of the sulfide isotopic values is also consistent with pyrite formation in a relatively open euxinic environment (Sageman and Lyons, 2003). Assuming that the sulfide preserved in pyrite is precipitated from sulfide dissolved in a single fluid derived from the mixing of two components—lowtemperature bottom waters/porewaters and hydrothermal fluids—an isotopic mass balance can be constructed: d34 SPY ¼

d34 SWBW XBW f þ d34 SHT XHT ð1  f Þ XBW f þ XHT ð1  f Þ

ð3Þ

where X indicates the concentration of H2S in a fluid, f is the fraction of fluid from the sedimentary porewater source, and the subscripts PY, BW, and

HT represent pyrite, bottom waters, and hydrothermal fluid, respectively (see also Chapter 9 in Faure, 1986). We apply this equation specifically to calculate the fraction of low-temperature bottom waters that contributed to the final fluid from which the preserved pyrite was precipitated. This calculation assumes that pyrite is precipitated at low temperatures with little or no fractionation from the aqueous sulfide composition (Price and Sheh, 1979). The porewater fluid was assumed to have a H2S concentration of 0.5 mmol/kg fluid, which is an average value for porewaters in anoxic sediments from the deepest basin of the Black Sea (Lyons, 1997). The isotopic composition of sulfide produced by dissimilatory bacterial sulfate reduction in pure cultures is depleted in 34S relative to the original sulfate by approximately 2 – 46x (Chambers et al., 1975; Detmers et al., 2001; Habicht and Canfield, 2001), a depletion that is less than the values of as much as 60xand more observed in the Phanerozoic sedimentary record (Sageman and Lyons, 2003). Recent models indicate that bacterial disproportionation of sulfur intermediates (e.g., elemental S, thiosulfate, sulfite) serves to exacerbate the 34S depletions observed between biogenic HS and pyrite (Canfield, 2001). Therefore, the isotopic composition of the porewater sulfide was calculated by first assuming an apparent total fractionation factor (a), which includes the effects of bacterial sulfate reduction and disporportionation, between sulfate and sulfide of 1.060. Given that: a¼

ðd34 Ssulfate þ 1000Þ ðd34 Ssulfide þ 1000Þ

ð4Þ

and using a value of + 13xfor the sulfate isotopic composition (Claypool et al., 1980), an isotopic composition of  44xis calculated for the porewater sulfide. Modern hydrothermal fluids from sedimented midocean ridge environments have been collected from Guaymas Basin, East Pacific Rise, Middle Valley, northern Juan de Fuca Ridge, and Escanaba Trough, Gorda Ridge. The concentration of H2S in fluids from these areas ranges from 1.1 to 7.9 mmol/kg fluid, with sulfur isotopic compositions ranging from  2.3xto + 10.4x(Campbell et al., 1988, 1994; Cruse, 2003). In order to maximize the amount of hydrothermal

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fluid that could have contributed to the final mixture, a H2S concentration of 1 mmol/kg fluid with an isotope composition of + 3x was chosen. This H2S concentration also corresponds to that assumed to be present in fluids thought to form MVT deposits (Sicree and Barnes, 1996) during Late Pennsylvanian – Permian times (Brannon et al., 1996; Symons et al., 1996). From this calculation, the fraction of sulfide in the black shales of core C-TW-1 derived from the low-temperature fluid, which could be euxinic porewaters or bottom water, ranges between 79x and 94%. If the higher H2S concentrations observed in modern vent fluids are used in the isotope mass balance calculation, an even greater proportion of the final fluid would have to be derived from sedimentary porewaters in order to produce the observed d34Spy values.

6. Conclusions Pennsylvanian deposits of midcontinent North America are classic cyclic stratigraphic sequences, the origins of which remain much debated. Through the application of geochemical proxies and comparisons with modern oceanic environments, this study is intended to provide a template for developing refined depositional models for the Pennsylvanian midcontinent seaway and for black shale environments in general. While our database is not large, the implications are far-reaching. Within a single Pennsylvanian cycle, geochemical proxies vary in concert with sedimentological/ecological indicators of dynamic paleoceanographic conditions during cyclothem deposition. However, local variations in such factors as siliciclastic fluxes, the nature of the organic carbon reservoir, and potential inputs from hydrothermal vents, are recorded in the levels of transition metal enrichments in individual cores. The classic models for cyclothem deposition of black shales invoke accumulation in relatively deep basinal settings, removed from the ancient shoreline. However, variations in the reactive iron reservoirs in core C-TW-1 and IRC, as manifested in differences in DOP values and Fe/Al ratios, indicate that C-TW1 experienced enhanced inputs of sulfide-reactive Fe relative to the delivery of sulfide-unreactive detrital Fe. Core C-TW-1 was enriched in reactive Fe (as

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pyrite) through inputs of iron-rich hydrothermal fluids to the basin and the subsequent scavenging of this Fe through water-column pyrite formation under potentially lower rates of siliciclastic deposition. The less reactive nature of the iron reservoir in core IRC may reflect differences in siliciclastic inputs, with IRC receiving more detrital flux than core C-TW-1. If so, this hypothesis challenges traditional paleogeographic reconstructions of the Pennsylvanian seaway. The proportion of organic matter derived from terrestrial sources is hypothesized to be greater in core IRC as compared to C-TW-1 and this possibility is supported by independent studies of other Pennsylvanian black shales (Coveney et al., 1987). Thus, it may be that the less-reactive (toward sulfide) Fe reservoir in core IRC is due to differences in the organic matter reservoir whereby reaction of Fe with Corg precluded pyrite formation. Reconstructing the location of ancient peat swamps that may have been the source of this organic matter (Wenger and Baker, 1985) and the final source-tosink input relationships could place constraints on ancient ocean circulation patterns. However, more work on the scavenging of transition metals by a range of organic matter types in general, and specifically the organic geochemistry of these Pennsylvanian shales, is required to test this hypothesis. Variations in Mo/Al ratios observed in core C-TW1 and IRC are consistent with the key role hypothesized for H2S in Mo fixation by organic matter. Enrichments in Mo/Al due to scavenging within the water column or sulfide-rich pore fluids is consistent with the euxinic conditions that likely characterized the deposition of the black shale facies. The lower level of Mo/Al enrichment in core C-TW-1, as compared to core IRC, is consistent with the hypothesized augmentation of bottom-water Fe concentrations by hydrothermal fluids. The addition of such Fe to waters near core C-TW-1 could have buffered sulfide concentrations to levels below the concentration required for Mo sequestration in the sediments. Furthermore, the extents to which Mo and other metals (U, V, Ni, Zn, and Pb) are enriched in organic-rich sediments may also be affected by the concentration and type of organic matter that accumulates in local settings. However, rigorous testing of this hypothesis awaits detailed organic geochemical analyses of these and other black shale samples.

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Importantly, our model indicates that processes other than bottom-water redox conditions must be considered in models for transition metal accumulations in organic-rich sediments. Factors such as the rate of siliciclastic accumulation, the nature of the organic carbon reservoir, and external inputs (e.g., hydrothermal vents), which may vary spatially and temporally within a single basin, can produce variations in the extent of metal sequestration in sediments, regardless of the uniformity of bottom-water redox conditions. Additionally, external inputs of metals, such as from hydrothermal vents, may also add to the complexity of utilizing metal accumulations as redox proxies. However, by integrating sedimentological data and a suite of geochemical indicators, such complexities can be unraveled, leading to more robust models for organic-rich sediment accumulation in ancient and modern basins.

Acknowledgements We thank Mike Formolo and Jon Fong for assistance with the d34S analyses, and Mike Glascock and the MURR for performing the INAA analyses. Thoughtful reviews of the manuscript by Mike Lewan, Joe Hatch, Joe Werne, and an anonymous reviewer for comments were greatly appreciated. Funding was provided by NSF grant EAR-9875961. Phil Heckel provided the initial stimulus for this project and helped with sample acquisition, and Ray Coveney furthered our understanding of the potential roles of hydrothermal inputs. References Adelson, J.M., Helz, G.R., Miller, C.V., 2001. Reconstructing the rise of coastal anoxia; molybdenum in Chesapeake Bay sediments. Geochim. Cosmochim. Acta 65, 237 – 252. Anderson, R.F., Fleisher, M.Q., LeHuray, A.P., 1989. Concentration, oxidation state and particulate flux of uranium in the Black Sea. Geochim. Cosmochim. Acta 53, 2205 – 2213. Arthur, M.A., Dean, W.E., 1998. Organic-matter production and preservation and evolution of anoxia in the Holocene Black Sea. Paleoceanography 13, 395 – 411. Berner, R.A., 1970. Sedimentary pyrite formation. Am. J. Sci. 267, 19 – 42. Brannon, J.C., Podosek, F.A., Cole, S.C., 1996. Radiometric dating of Mississippi Valley-type ore deposits. In: Sangster, D.F. (Ed.),

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