deep sea drilling project initial reports volume 75

21. ORIGIN AND GEOCHEMISTRY OF CRETACEOUS DEEP-SEA BLACK SHALES ANDMULTICOLORED CLAYSTONES, WITH EMPHASIS ON DEEP SEA DRILLING PROJECT

SITE 530, SOUTHERN ANGOLA BASIN1

Walter E. Dean, U.S. Geological Survey, Denver, ColoradoMichael A. Arthur,2 Department of Geology, University of South Carolina, Columbia, South Carolina

andDorrik A. V. Stow, Grant Institute of Geology, University of Edinburgh,

Edinburgh, Scotland, EH9 3JW, United Kingdom

ABSTRACT

Deep-water sedimentary sequences of mid-Cretaceous age, rich in organic carbon, have been recovered at manyDSDP sites in the Atlantic Ocean. Most of these sequences have a marked cyclicity in amount of organic carbon result-ing in interbedded multicolored shale, marlstone, and (or) limestone that have cycle periods of 20,000 to 100,000 yearsand average 40,000 to 50,000 years. These cycles may be related to some climatic control on influx of terrigenous or-ganic matter and sediment, rates of upwelling and sea-surface production of organic matter, and preservation oforganic matter related to deeper-water dissolved oxygen concentration. These variations in supply of organic matterhad pronounced effects on the potential of the sediment for subsequent diagenetic changes and geochemical partition-ing in adjacent beds.

Many trace elements are enriched in organic-carbon-rich lithologies relative to interbedded organic-carbon-poorlithologies. Elements that are most commonly enriched are Cr, Ni, V, Cu, Zn, and Mo. The association of high trace-element concentrations with organic matter may be the result of concentration of these elements by organisms or bychemical sorption and precipitation processes under anoxic conditions. Detailed trace-element profiles from organic-carbon-rich strata at Site 530 suggest that there may be differential mobility of trace elements, with diffusion of someelements over distances of at least tens of meters. The sequence of trace-element mobility, from highest to lowest, is ap-proximately Ba, Mn, Pb, Ni, Co, Cr, Cu, Zn, V, Cd, and Mo. Slowly deposited, oxidized clays directly overlying someblack shale sequences are enriched in some metals, particularly Fe, Mn, Zn, and Cu, relative to normal pelagic clays,and this enrichment may be the result of upward migration of metals in pore waters during compaction or diffusionfrom the underlying black shale.

Most depositional models that have been used to explain the accumulation of the organic-carbon-rich strata implythat reducing conditions in the sediments (and therefore the increased degree of preservation of organic matter) werethe result of anoxic or near-anoxic conditions in oceanic bottom waters, or in a midwater oxygen-minimum zone.Evidence from several DSDP sites in the Atlantic, however, indicate that some of these middle Cretaceous "blackshale" beds may be the result of variations in rate of supply of organic matter that produced anoxia or near-anoxiawithin midwater oxygen-minimum zones and possibly, under extreme conditions, throughout much of the bottom-water mass. Although bottom-water anoxia may have occurred during periods of organic-carbon-rich strata, it was notnecessarily the only cause for accumulation of these strata. The main reason for the accumulation of organic-carbon-rich strata was an increase in the relative amount of organic debris being deposited. Some of this organic debris wasderived from continental-margin areas of increased production, accumulation, and preservation of organic matter frommarine, terrestrial, or mixed sources and transported to slope and basinal sites by turbidity currents.

INTRODUCTION

Deep-water sedimentary sequences of middle Creta-ceous age that contain beds rich in organic carbon (Corg)have been recovered at many Deep Sea Drilling Project(DSDP) sites, particularly in the Atlantic Ocean. Mostof these sequences display a marked cyclicity in amountof Corg resulting in interbedded lighter and darker shale,marlstone, and (or) limestone. Lower Cretaceous (Neo-comian) Corg-rich strata accur within carbonates, butmost of the other Cretaceous strata in the Atlantic arecarbonate-poor and clay-rich, resulting in cyclic inter-bedding of red, green, and black claystone or shale. Inthis chapter we use "black shale" to refer to dark-colored mudstone, but because we are dealing with

1 Hay, W. W.( Sibuet, J.-C, et al., Init. Repts. DSDP, 75: Washington (U.S. Govt.Printing Office).

2 Present address: Graduate School of Oceanography, University of Rhode Island, Nar-ragansett, Rl 02882.

freshly cut core material it is not always possible to tellif these mudstones would develop fissility on outcrop tomake them "shale" in the classical sense.

Accumulation of Corg-rich strata in the Atlanticbegan in the Early Cretaceous and has continued inter-mittently to the present. The interbedding of green andblack, and green and red lithologies that is so commonthroughout the entire sedimentary sequence in the At-lantic Ocean is clearly a manifestation of varying dia-genetic redox conditions within the sediments, but it isnot clear whether fluctuations in amount of organicmatter, that are related to sediment color, were causedby or were the cause of varying redox conditions in bot-tom waters just above the sediment/water interface.Nor is it clear to what extent deoxygenation of bottomwaters overlying the sediments may have caused fluctua-ting redox conditions within the sediments.

Organic matter is preserved in relatively high concen-trations in deep-sea Corg-rich strata either because of

819

W. A. DEAN, M. A. ARTHUR, D. A. V. STOW

high rate of supply of organic matter, enhanced preser-vation under anoxic conditions, enhanced preservationdue to high sedimentation rate, or some combination ofthese factors. Once the organic matter reaches the sedi-ment, some portion of it usually is decomposed, andvarious oxidizing agents are reduced in a series of oxida-tion/reduction reactions that are largely mediated bybacteria. The by-products of these reactions and of thedecomposition of organic matter are various reducedmetal species, ammonia, phosphate, sulfide, and largeamounts of carbon dioxide and bicarbonate which con-tribute to increases in alkalinity in interstitial waters.Changes in the composition of organic matter with pro-gressive diagenesis will not be considered in detail in thischapter, but the nature of the original organic matterand subsequent breakdown of organic compounds is ex-tremely important in determining the ultimate chemis-try and mineralogy of Corg-rich sediments because theamount and reactivity of organic matter in part controlsthe amount of oxidant that is reduced. In addition, or-ganic matter may supply a large portion of trace metalsfound in Corg-rich sediments, and, therefore, there maybe a significant correlation between Corg content andconcentrations of certain trace metals. Finally, the chem-ical nature of preserved organic matter is often used todraw inferences about the depositional environment(e.g., Didyk et al., 1978).

The purposes of this chapter are: (1) to outline the oc-currence of Corg-rich facies of Cretaceous age at Site 530and other DSDP sites in the North and South Atlantic,(2) to discuss the extent to which environment of deposi-tion determines the supply of organic carbon to deeper-water sediments and therefore predetermines the extentto which diagenetic reactions and enrichment of certaintrace elements can occur, and (3) to contrast the chem-istry of Corg-rich facies from different depositional set-tings, and to briefly outline the potential effects of Cre-taceous black-shale deposition and diagenesis on paleo-cean chemistry.

ORGANIC-CARBON-RICH STRATA IN THECRETACEOUS ATLANTIC OCEAN

The oldest occurrences of Corg-rich strata are withinbioturbated white to light-gray limestone of Tithonianto Neocomian age recovered from DSDP Sites 101, 105,367, 387, and 391 (Figs. 1 and 2A; Dean et al., 1978;Jansa et al., 1978 and 1979; Dean and Gardner, 1982).Upper Jurassic (Callovian-Oxfordian) strata containingas much as 3% organic carbon were recently recoveredat DSDP Site 534 in the Blake-Bahama Basin (Sheridan,Gradstein, et al., 1982), but details of these strata arenot yet available. The Neocomian laminated dark marl-stone beds generally contain more than 5% organic car-bon, and one bed at Site 367 in the Cape Verde Basin offnorthwest Africa contains 33% organic carbon (Dean etal., 1978; Dean and Gardner, 1982).

The end of deposition of the Neocomian carbonatesapparently was caused by a sudden rise in the carbonatecompensation depth (CCD) (Arthur, 1979a; Thierstein,1979), and the remainder of the section in the North At-lantic through the Eocene is dominated by multicolored

clay-rich lithologies. At most DSDP sites in the NorthAtlantic, the Neocomian carbonates are overlain by theso-called middle Cretaceous (here defined as Aptian-Albian to Turonian) black-shale facies, but at some sitesthe two are separated by a thin (< 10 m) unit of in-terbedded red and green claystones of late Aptian toearly Albian age (Jansa et al., 1978; Dean and Gardner,1982).

The middle Cretaceous black-shale facies representsthe main period of accumulation of organic matter inthe Atlantic, and at most DSDP sites it consists of inter-bedded green and black clay-rich lithologies (see reviewsby Arthur, 1979a; Tucholke and Vogt, 1979; Thierstein,1979; Arthur and Natland, 1979; Tissot et al., 1979 and1980; Dean and Gardner, 1982; Weissert, 1982). Thisfacies has been recovered at DSDP Sites 4, 5, 101, 105,135, 137, 138, 144, 367, 368, 369, 370, 386, 387, 391,398, 400, 402, 415, 416, 417, 418 and 534 in the NorthAtlantic Ocean, and Sites 327, 328, 330, 356, 357, 361,363, 364, 511, and 530 in the South Atlantic Ocean (Fig.1). Upper Jurassic (Oxfordian) black shale was recov-ered at Site 327 on the Falkland Plateau. This shale wasdeposited in shallow water and contains mostly terres-trial organic matter (McCoy and Zimmerman, 1977; Bar-ker et al., 1977). Other than the Site 327 black shale, themain periods of black-shale deposition were during theAptian-early Albian and sporadically from late Albianthrough Coniacian-Santonian (Arthur and Natland,1979). The black-shale sequences are overlain by UpperCretaceous red claystone. Concentrations of organiccarbon in the black lithologies in the North Atlanticrange from 3 to 37% and usually are more than 5%(Lancelot, Seibold, et al., 1978; Tissot et al., 1979 and1980; Summerhayes, 1981; Dean and Gardner, 1982).Concentrations of organic carbon generally are lower inblack lithologies from western North Atlantic sites thanin those from eastern North Atlantic sites. True blackshale usually amounts to less than half of the so-calledblack-shale facies.

The middle Cretaceous black-shale unit in the CapeVerde Basin (DSDP Site 367) is overlain by at least 160m of interbedded green and brown clay of Late Creta-ceous to late Paleocene/early Eocene age. The greenand brown clay section is overlain by at least 85 m of in-terbedded green and black zeolitic, radiolarian clays ofEocene age; the black clay beds contain up to 4% or-ganic carbon (Dean and Gardner, 1982). The Creta-ceous/Tertiary red and green clays, like the Aptian/Al-bian red and green claystones, represent a period ofaccumulation of more-oxidized, Corg-poor sedimentsandwiched between strata deposited during periods ofaccumulation of less-oxidized, Corg-rich sediment.

The organic detritus in Corg-rich layers may be fromautochthonous marine planktonic production, or fromterrestrial organic debris. According to Tissot et al.(1979 and 1980), and Summerhayes (1981), autochtho-nous marine organic matter (that is, organic matter con-taining type II kerogen of Tissot et al. (1974) in con-tinuous sections of Lower Cretaceous rocks in the NorthAtlantic is found only in the Cape Verde Basin (Site367), on Cape Verde Rise (Site 368), and on Demerara

820

CRETACEOUS CLAYSTONES AND BLACK SHALES

45'

45'

Figure 1. Map showing relative locations of DSDP sites where Lower and middle Cretaceous organic-carbon-rich strata havebeen recovered plotted on a continental reconstruction at 100 ± 10 m.y. ago. (Modified from Sclater et al., 1977.)

Rise (Site 144). Everywhere else in the North Atlantic,terrestrial organic matter with admixtures of highly de-graded, residual marine organic matter dominates Low-er Cretaceous Corg-rich strata. The Cretaceous Corg-richstrata in the southeastern North Atlantic, therefore, ap-pear to have more in common with similar strata in theSouth Atlantic, where most of the organic matter is de-

rived from marine sources, than with equivalent stratain the rest of the North Atlantic (Tissot et al., 1979 and1980; Summerhayes, 1981).

In the South Atlantic sequences (Fig. 2B), organic-carbon concentrations average about 3 to 4% in the Ap-tian-lower Albian at Site 361 in the Cape Basin, withvalues as high as 14% (Arthur and Natland, 1979), and

821

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Figure 2. A. Summary stratigraphic column for Mesozoic strata recovered from DSDP sites in the North Atlantic Ocean. B. Summary stratigraphic columns for oc-currences of Cretaceous Corg-rich strata at South Atlantic DSDP sites. Index map (inset) is modified from Sclater et al. (1977).


5 to 7%, with values as high as 24%, in the lower Albianat Site 364 in the eastern Angola Basin. Average organ-ic-carbon concentrations of 4 to 5% are typical of theHauterivian-lower Aptian strata at Site 511 in theArgentine Basin (von der Dick et al., 1983), and lessthan 1% in the Aptian-Albian. The black-shale beds ofAlbian to Coniacian-Santonian age at Site 530 containas much as 16.59b organic carbon with an overallaverage of 5.4% organic carbon as compared with lessthan 0.5% organic carbon in intercalated beds of redand green claystone (Meyers, Brassell, Hue, this vol-ume). The black-shale beds occur within about 160 m ofsection but most are within 20 m of section in Cores 97and 98 (Fig. 3). Most of the organic matter in blackshales at Site 530 is of marine origin (Brassell et al.,Deroo et al., this volume). The organic matter in blackshales at Sites 364 and 511 also is mostly autochthonousmarine organic matter (Tissot et al., 1980), whereasmuch of the organic matter at Site 361 in the Cape Basinapparently is redeposited terrestrial organic matter. Inspite of these differences in sources of organic matter,the organic-carbon-rich strata in all Atlantic basins areinterbedded with bioturbated, organic-carbon-poqr strata.

The cyclic couplets of organic-carbon-rich and or-ganic-carbon-poor beds are mainly several decimetersthick, but can range from several centimeters to as muchas 80 cm thick. Periodicities of these cycles at most sitesrange from 20,000 to 100,000 years, with averages ofabout 40,000 to 50,000 years (Dean et al., 1978; Arthurand Fischer, 1977; Arthur, 1979a; McCave, 1979a). Thethickness of an individual cycle depends mainly on sedi-mentation rate. For example, the average cycles in theNeocomian-Aptian carbonate facies are thicker thanthose in the more slowly deposited Aptian-Albian black-shale facies (Dean et al., 1978; Arthur, 1979a). Regard-less of the exact mechanism involved in the formation ofthe sedimentary cycles, the pronounced cyclicity of thesedimentary section in the Atlantic, and particularly inthe middle Cretaceous, suggests that there probablywere cyclic variations in the influx of terrigenous organ-ic matter and detrital material, sea-surface productionof organic matter and CaCO3, dissolved oxygen concen-tration in deeper-water masses, and (or) preservation ofCaCO3. The cyclicity of supply and (or) preservation oforganic matter and CaCO3 had a pronounced influenceon subsequent diagenetic events and on geochemicalpartitioning in adjacent beds.

BLACK-SHALE DEPOSITION AT SITE 530

The age of the oldest sediment above basaltic base-ment at Site 530 is late Albian-Cenomanian. The sea-floor at Site 530 is presently at 4645 m, and backtrack-ing of the site using the methods of Sclater et al. (1977)and van Andel et al. (1977) gives an estimated paleo-depth of about 2500 m during the Albian. This estimateis reasonable for the location of Site 530 at or near theridge crest at that time. The site was several hundredkilometers from the African continental margin (Fig. 2),and at that time probably was separated from it by atrough. To the south, however, the site was flanked bythe aseismic Walvis Ridge. There appears to have been a

continuous paleoslope or steep scarp between the Wal-vis Ridge and the part of the Cretaceous mid-oceanridge flank upon which Site 530 was located. There is agreater likelihood that any redeposited sediment at Site530 was derived from the flank of Walvis Ridge thanfrom the more distant African continental margin. Thishas implications for the origin of the Corg-rich intervalsin the mid-Cretaceous of Site 530.

Thin black shale beds are intercalated within greenand red mudstone beds throughout the 165-m interval(1105-940 m sub-bottom) from upper Albian to Conia-cian-Santonian (Fig. 3). Black-shale beds make up 5 to50% of the recovered section and occur as 260 thin beds(< 1 to 62 cm with an average of 4.3 cm). Very fine-grained, thin-bedded turbidites make up between 5 and20% of the section and occur both in the black shalesand in the interbedded green and red mudstones. Or-ganic carbon contents range from 1.4 to nearly 19%over the interval (Meyers, Brassell, and Hue, and Derooet al., this volume) and Rock-Eval pyrolysis results sug-gest that most of the organic matter is marine in origin,at least in the more Corg-rich intervals (Brassell, Derooet al., Meyers, Brassell, and Hue, this volume). There isa good correlation between Corg content and the kerogenhydrogen index (Fig. 4) suggesting that the degree of ox-idation is the main control of hydrogen and oxygen in-dices of kerogen in Site 530 black-shale beds. Allkerogen samples that have a hydrogen index greaterthan 100 mg HC/g Corg have low oxygen indices (lessthan 50 mg O2/g Corg), indicating good preservation oforiginal autochthonous marine organic matter (Fig. 5).Sediment samples containing less than 2 to 3% Corg havehydrogen indices of less than 200 mg HC/g Corg (Fig. 4).This suggests that there was a critical Corg content ofabout 3% at this site below which subsequent early dia-genetic oxidation was deleterious, resulting in poor pres-ervation. The high oxygen and low hydrogen indicestherefore probably reflect mainly oxidation of marineorganic matter rather than an increased flux of terres-trial higher plant debris, although there may be a back-ground residual of terrestrial Corg as in Cretaceous blackshales from sites in the North Atlantic (e.g., Summer-hayes, 1981).

The excellent preservation of organic matter in cer-tain thin beds within the Cretaceous multicolored clay-stone sequence (Unit 8) at Site 530 could have originatedby one or more of the following mechanisms (e.g., De-maison and Moore, 1980; Muller and Suess, 1979): (1)periodic high productivity and high flux of Corg to thedeep-sea floor; (2) enhanced preservation under anoxicconditions; and (3) enhanced preservation under ex-ceptionally high sedimentation rates. All three mech-anisms have been called upon either singly or in combi-nation by other authors to explain Cretaceous deep-water black shales at other DSDP sites. Depending uponpaleolatitude, proximity to continental sources, andrates of circulation and oxygen content of deep-watermasses, a different mechanism could satisfactorily ex-plain enhanced preservation of Corg in each locality(e.g., Arthur and Natland, 1979; Dean et al., 1977;Summerhayes, 1981; Ibach, 1982). There is undoubt-

823

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700 r

0 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20Organic carbon (wt.%)

Figure 4. Scatter diagram of hydrogen index and percent Corg in redand green claystones and black shales in Hole 53OA. Data fromDeroo et al., this volume.

700 r 200r

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600

500

400

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Oxygen Index (mg O2/g Corg)

Figure 5. Scatter diagram of hydrogen and oxygen indices for organicmatter in red and green claystones and black shales in Hole 530A.Data from Deroo et al., this volume.

edly a complex interplay between each of the threemechanisms. At Site 530, however, none of the threeseems to have played a direct role in preservation ofCorg, although they all may have indirectly affected theaccumulation of black clays there.

Evidence for exceptionally high productivity duringdeposition of mid-Cretaceous sediment at Site 530 islacking. Although concentrations of Corg are high, theyoccur intermittently in discrete beds that commonly areseparated by substantial thicknesses of red or greenclay stone. Except for Cores 97 and 98, black beds gen-erally comprise less than 10% of the section (Fig. 3).Other evidence of high surface productivity, such ashigh concentrations of biogenic silica and phosphate,also is largely absent (see Dean and Parduhn, thisvolume). The argument that a high sedimentation ratealso aids in the preservation of Corg does not directly ap-ply to Site 530 black shale. Accumulation rates of Corgare not particularly high (maximum of 46 g/cm2/m.y.;Table 4) and are about one order of magnitude lower

lantic black-shale sequences. Sedimentation rates varyfrom 9 to 14.6 m/m.y. at Site 530, which are relativelylow. The range of accumulation rates of Corg at Site 530(Table 4), with an average of about 10 g/cm2/m.y., aresomewhat low relative to those reported by Muller andSuess (1979) and Ibach (1982). (The accumulation ratesreported by Ibach are 100 times too high because of anerror in her equation for computing total organic car-bon accumulation rates.) At sedimentation rates of 9to 14.6 m/m.y., an average Corg accumulation rate of10 g/cm2/m.y. falls slightly below the typical "blackclay" facies of Ibach (1982) but within her "black clay-siliceous overlay." However, when considered on anindividual basis, the black shale beds contain an aver-age of 5.4% Corg (Meyers, Brassell, and Hue, thisvolume), but make up less than 10% of the total section.Interbedded red and green claystone layers contain nosignificant amount of Corg. The black layers must beconsidered separately because they represent episodicsedimentation events and because they contain a dispro-portionate share of the Corg in the sequence.

We find it difficult to explain the interbedding andsedimentary structures in the multicolored claystone se-quence using models that require periodic anoxia as acause of the interbedding. The frequency of changesfrom oxic to anoxic conditions would have to have beenhighly variable, but with an average period of about60,000 years (Table 4). Abrupt decreases in deeper-wa-ter oxygen content are unlikely, so that the black layerscontaining better preserved Corg should not have sharpbasal contacts with underlying clay layers depleted inCorg. Burrowing organisms should continue to mix sedi-ment until deep-water oxygen contents reach about 0.5ml/1 (e.g., Rhoads and Morse, 1971). Therefore, thereshould be a mottled transition zone between the under-lying red or green clay and the black clay layer. Sharpupper contacts could occur as a result of abrupt oxy-genation events, but subsequent burrowing activitywould tend to disrupt upper contacts of black clay lay-ers. Therefore, if the black clay layers represent epi-sodes of anoxia at the seafloor, one would usually ex-pect gradation in organic content, color, and laminationfrom underlying oxidized intervals into laminated blackclays and to overlying oxidized sediment. About 80% ofthe black-shale beds at Site 530 are at least partlylaminated (Stow and Dean, this volume). Some of thislamination is in the form of very thin silt laminae (Site530 summary chapter; Stow and Dean, this volume).Many of these silt laminae clearly are parts of very thin-bedded, fine-grained turbidites, and these turbidites oc-cur within at least 25% of the black-shale beds. About40% of the black-shale beds are bioturbated to some ex-tent, although the burrows generally are smaller than ininterbedded green mudstones. Much of the evidence forturbidites in the red and green mudstones has beenobliterated by bioturbation. We suggest that the black-shale layers at Site 530 contain material that was re-deposited from the flank of Walvis Ridge and representa distal turbidite facies where the sediment source waspredominantly within an intense oxygen-minimum zone.

than average accumulation rates of Corß in other At- This could explain the high Corg content of individual

825


layers and the excellent preservation of autochthonousmarine organic matter. Surface organic productivitymay have been high in upwelling zones located along thesouthwest African continental margin (e.g., Parrish andCurtis, 1982) which led to an intense and relatively thickmid-water oxygen-minimum zone over much of theAngola Basin. Although productivity was not necessari-ly high along the seaward portions of the Walvis Ridge,organic matter was preserved along the ridge slopewithin the intensified oxygen-minimum zone. Frequentslumps and turbidity currents might have occurred be-cause of high sedimentation rates, steep slopes, periodicearthquakes, and (or) intense gas generation in the Corg-rich slope sediments.

Redeposition of black clay to deeper water does notnecessarily negate the hypothesis of "oceanic anoxicevents" (Schlanger and Jenkyns, 1976). The widespread,nearly synchronous deposition of Corg-rich strata overmuch of the Atlantic and Tethys during, for example,the Aptian-Albian and the Cenomanian-Turonian ar-gues for some fundamental difference between theoceans during these time periods and most of the LateCretaceous-Cenozoic. At Site 530 we can see the im-print of such an oxygen-deficient event in the muchgreater frequency of black-shale interbeds near the Ce-nomanian/Turonian boundary. This record of maxi-mum deposition of black shale occurs in Cores 97 and98 (Fig. 3.) in which the accumulation rate of Corg is 4 to10 times that of the rest of the sequence (Fig. 10; Table4). We suggest that this culmination of black shale de-position reflects an increase in productivity near thesouthwest African margin, an expansion and intensifi-cation of the oxygen-minimum zone that impinged onWalvis Ridge, or both.

GEOCHEMISTRY OF ORGANIC-CARBON-RICH STRATA

Sulfur

Sulfur is incorporated into anoxic, organic-carbon-rich sediments mainly through bacterial sulfate reduc-tion. This process is relatively well known (e.g., Berner,1971, 1974, 1978, 1980; Goldhaber and Kaplan, 1975and 1980; Goldhaber et al., 1977; Jorgensen, 1977;Oremland and Taylor, 1978) and involves the utilizationof pore-water sulfate as an electron donor for the oxida-tion of sedimentary organic matter by bacteria in theabsence of dissolved molecular oxygen. Organic carbonis consumed in the proportion of two moles for everymole of sulfate reduced. Sulfate reduction usually pro-ceeds after all other oxidants are consumed. For ex-ample, Froelich et al. (1979) found that for pelagic sedi-ments in the eastern equatorial Atlantic, oxidants wereconsumed in the following order of decreasing energyproduced per mole of organic carbon oxidized: O2 >Mn-oxides > NO3~ > Fe-oxides > SOJ~. And yet, inmost anoxic sediments, sulfate accounts for most of theoxygen consumption and organic-matter decomposition.For example, Jorgensen (1977) found that sulfate reduc-tion accounted for 53% of the mineralization of organic

matter in anoxic sediments of a fjord in Denmark.Jorgensen also found that only 10% of the sulfide pro-duced by sulfate reduction was precipitated as metalsulfides and the rest diffused to the surface where it wasreoxidized. Berner (1978) suggested than an averageof 0.2 to 0.4 weight percent of the organic carbon islost from sediments during sulfate reduction, but thisamount may vary depending upon the rate of sedimen-tation and the amount of sulfate available by diffusioninto the sediment, as well as on the type and reactivity oforganic matter present (Toth and Lerman, 1977; Ber-ner, 1980; Goldhaber and Kaplan, 1980). The reducedsulfur is free to diffuse upward to the sediment-waterinterface, where it can be oxidized by bacteria, or to dif-fuse to sites of metal-sulfide precipitation, a major sinkfor sulfur in anoxic sediments. Modern anoxic sedi-ments contain concentrations of total sulfur of up toseveral weight percent. The weight ratio of sulfur toorganic carbon in modern anoxic sediments averages0.4, but there are significant variations from this value.Levanthal (1983) suggested, based on data for S:C ra-tios from the Black Sea that S:C ratios greater than 0.4in black shales, with a slope intersect at a positive valueof sulfur at zero carbon, signify deposition under ananoxic water column, rather than having only anoxicsediment pore waters. Sulfur in excess of that diffusinginto sediments from overlying seawater is presumablyadded to sediment by sulfate reduction and formationof metal sulfides in the overlying anoxic water column.

Figure 6 is a plot of organic carbon versus sulfur forthose samples of Cretaceous strata for which both ór-ganic-carbon and sulfur analyses are available. This plotsuggests that a S:C ratio of about 0.4 is a reasonableaverage for these strata, which is similar to the S:C ratioin most modern anoxic sediments. Kendrick (1979; dataincluded in Fig. 6) measured concentrations of totalsulfur of up to 4.3 weight percent in organic-carbon-richstrata from DSDP Sites 386 and 387. As much as 40%of this sulfur apparently is in the organic fraction,and the rest in pyrite. Kendrick interpreted this sulfurenrichment as evidence for deposition and preservationof organic carbon under an anoxic water column. Lowconcentrations of sand organic carbon were interpretedas an indication of oxic bottom-water conditions.

Metal sulfides of varying morphologies are present innearly all of the Cretaceous organic-carbon-rich stratarecovered from DSDP sites in the Atlantic. Pyrite is themost commonly reported sulfide mineral and occurs asframboids, test-fillings of foraminifers, micronodules,replacements or molds of radiolarian tests, and largediscrete nodules and lenses. Localized concentrations ofpyrite and (or) marcasite are commonly observed atcontacts between black shale and underlying green clay-stone beds (Fig. 7). In fact, sulfide sulfur appears to bevery mobile in sequences of interbedded organic-car-bon-rich and organic-carbon-poor strata (Berner, 1969and 1970), and green claystone beds often contain asmuch total sulfur as adjacent beds of black shale.

Sedimentation rates were generally high, and burrow-ing apparently was sufficiently ineffective during de-

826


4 -

cm90 ,—

1 2 3 4 5 6 7 8 9

Organic carbon (%)

4 Atlantic DSDP, Lower Cretaceous: Groups 1,2, and 3of Brumsack (1980)

0 Atlantic DSDP, Upper Cretaceous (1) and Cenozoic (2)of Scharf (1980)

• Atlantic, DSDP, Cretaceous black shales (Sites 386 and 387;Kendrick, 1979)

A Atlantic, DSDP, Upper Cretaceous black shales (Site 367,Core 17: Dean and Gardner, 1982)

Δ Atlantic, DSDP, Upper Cretaceous green claystones(Site 367, Core 17: Dean and Gardner, 1982)

V Atlantic, DSDP, Upper Cretaceous green and red claystones(Site 367, Cores 15 and 16; Dean and Gardner, 1982)

• Atlantic, DSDP, Lower Cretaceous olive marlstone (Site 367,Cores 25 and 26; Dean and Gardner, 1982)

D Atlantic, DSDP, Lower Cretaceous white limestones (Site 367,Cores 25 and 26, Dean and Gardner, 1982)

Figure 6. Plot of concentrations of organic carbon and total sulfur(presumably as sulfide sulfur) in Cretaceous organic-carbon-richstrata from Atlantic DSDP sites. Although there is a large amountof scatter, the slope of a regression line through the data is about0.4, similar to that for plots of organic carbon versus total sulfurin modern anoxic Black Sea sediments (Leventhal, 1983).

position of organic-carbon-rich strata in many Creta-ceous deep-sea environments to inhibit extensive diffu-sion of sulfate into the sediment. Availability of ferrousiron probably was never a limiting factor in metal sul-fide formation in Cretaceous Corg-rich strata.

Phosphate

Atlantic Cretaceous Corg-rich strata usually have alarge range of phosphorus concentrations (0.005 to 1.8weight percent P2O5), but we know of no reported phos-phate-mineral aggregates, such as pellets or nodules,from these strata. Most phosphorus is supplied to sedi-ments in productive areas of the oceans in organic mat-ter and fish debris (Froelich et al., 1982). Phosphate isreleased from organic matter by decomposition both inthe water column during settling of organic matter andfecal pellets to the bottom, and in sediment pore watersduring aerobic and anaerobic decomposition (e.g., Froe-lich et al., 1979). The C:P ratio in organic matter in-creases systematically as organic matter settles throughthe water column and during early diagenesis (Suess,

95

100

105 —

Figure 7. Pyrite nodule at the contact between a bed of black shale andan underlying bed of green claystone: Sample 530A-104-3, 90-105

1980; Froelich et al., 1982). As organic matter decom-poses in the sediment, phosphate is liberated to pore wa-ters; it can diffuse out of the sediment into overlyingseawater and (or) precipitate in sediment as some phos-phatic phase, or by adsorption on Fe-oxyhydroxides atan oxic/anoxic interface (e.g., Berner, 1974; Curtis,1980). In general, dissolved phosphate in pore waters oforganic-carbon-rich sediments increases with depth to amaximum that may be controlled by precipitation of acarbonate-fluorapatite phase (e.g., Baturin et al., 1970;Berner, 1980; Suess, 1981) due to decomposition oforganic matter and possibly through dissolution of fishdebris (Suess, 1981). Fish debris usually is well preservedin organic-carbon-rich strata, suggesting that dissolu-tion is buffered by the increases in phosphate concentra-tions due to decomposition of organic matter. Baturinet al. (1970) found maximum phosphate concentrationsof 9.7 µg-at/1 in pore waters of Walvis Bay muds, overthree times the highest concentration in open oceanic

827


deep waters. Concentrations of phosphate several or-ders of magnitude higher than those in overlying seawa-ter have been reported from sediments in nearshore en-vironments (Baturin, 1972; Nissenbaum et al., 1972;Sholkovitz, 1973; Berner, 1974; Suess, 1976). A limit tothe amount of phosphate that can be buried in sedimentis the initial C:P ratio in the sedimenting organic matter.Deep-water organic-carbon-rich sediments may not be amajor sink for phosphorus because, although much or-ganic matter may be preserved in anoxic strata, much ofthe organic phosphorus may have been liberated todeep-water masses by decay of organic matter settlingto the bottom. Thus, the C:P ratio of sedimented or-ganic matter may already be elevated, and the totalphosphorus content of deeper-water organic-carbon-rich sediments will be relatively low (Glenn and Arthur,1982). The average P2O5 content of modern organic-carbon-rich sediments from the Black Sea is 0.17 weightpercent, which is similar to or lower than concentrationsin oxic sediments where phosphate is adsorbed onto fer-ric oxyhydroxides (Berner, 1974). Analyses of Creta-ceous deep-water organic-carbon-rich strata and inter-bedded organic-carbon-poor strata show that averageP2O5 values are generally below 0.18 weight percent inAtlantic DSDP sites (e.g., Sites 105, 364, 386, 387,417D, and 530), with a range of 0.06 to 1.77 weight per-cent. The few available analyses from DSDP Site 369 onthe northwest African continental slope (Donnelly, 1980)suggest that organic-carbon-rich Cretaceous nannofos-sil marl and limestone at that site have concentrationsof P2O5 that are higher than deeper-water strata ofequivalent age. Phosphate probably occurs mainly inthe form of preserved phosphatic fish debris in mostdeep-sea organic-carbon-rich strata.

The geometric mean concentration of P2O5 in blackshale from Site 530 is 0.18 percent (Dean and Parduhn,this volume). The average concentration of organic car-bon in all samples of black shale reported by Meyers,Brassell, and Hue (this volume) is 5.4%. These twoaverage values yield an average C:P atomic ratio of 168which is relatively high (compare with 106, the averageRedfield ratio in freshly fixed planktonic organic mat-ter; Redfield, 1958). Phosphorus accumulated at ratesbetween 0.013 and 0.58 g/cmVm.y. (average of about0.2) in the black shales at Site 530, which are not excep-tionally high values.

Trace Elements

The Cretaceous Corg-rich strata in the Atlantic pro-vide insight into the general problem of incorporationof trace metals in organic-carbon-rich sediments. Suchmarine sediments and rocks are known to be significant-ly enriched in a variety of redox-sensitive trace elements(e.g., Vine and Tourtelot, 1970) relative to average mar-ine sediment and shale, but the degree of enrichmentvaries greatly. Few systematic studies have been made ofmajor- and trace-element concentrations in Cretaceousdeep-sea organic-carbon-rich lithologies, but such stud-ies are important because these relatively thick and lat-erally extensive lithologic units are potential sinks for

organic carbon, sulfur, and trace elements, particularlyAs, Ba, Cd, Co, Cr, Cu, Mo, Ni, Pb, V, and Zn. Duringthe deposition of organic-carbon-rich sediments, it ispossible that trace elements, organic carbon, and sulfurare extracted from the oceans at greater than steady-staterates, thereby perturbing the chemical and isotopic bal-ances of the ocean. The possible effects of Cretaceousblack-shale deposition on the global carbon cycle hasbeen discussed by Scholle and Arthur (1980) and Arthur(1982). These authors found that average δ13C values ofoceanic total dissolved carbon increased during mid-Cretaceous black-shale deposition, as did the rates ofaccumulation of organic carbon in the deep sea. Aver-age deep-sea accumulation rates of organic carbon wereas much as seven times those of today during the peakof Aptian-Albian black-shale deposition. The reasonsfor the enhanced preservation of organic carbon in theCretaceous deep sea basins, however, are still a matterof debate (see discussion below). Anderson et al. (1982)suggested that increased rates of burial of sulfur as met-al sulfides in Cretaceous black shales caused significantchanges in the oceanic sulfur cycle. Brumsack (1980)also considered the possible effects of Cretaceous black-shale deposition on trace-metal and carbon cycling, buthis estimates of the fluxes of those components to theblack shales probably are too high because he consid-ered the entire thickness of a so-called "black shale"unit to have the composition of his "average" blackshale (see Tables 1 and 2). As we have already pointedout, black-shale-bearing units recovered from AtlanticDSDP sites commonly contain less than 10 to 20% ofbeds with an organic-carbon concentration of morethan 1 %, but some intervals have higher proportions. Inthis section, we contrast the geochemistry of interbedd-ed black, green, and red Cretaceous strata, and comparetheir geochemistry and metal accumulation rates withthose of modern anoxic sediments.

The enrichment of certain trace elements—especiallyCu, Zn, Mo, V, Ni, and Cr—in sediments and rocks en-riched in organic matter has been reported by many in-vestigators (e.g., Wedepohl, 1964; Brongersma-Sanders,1965; Calvert and Price, 1970; Vine and Tourtelot, 1969and 1970; Volkov and Fomina, 1974; Lange et al., 1978;Brongersma-Sanders et al., 1980; Dean and Gardner,1982; Dean and Parduhn, this volume). The associationof high trace-element concentrations with organic mat-ter may be the result of concentration of these elementsby organisms (bi° c o n c e ntration), by chemical sorptionand precipitation with organic detritus, clays, etc. set-tling through the water, or by formation of sulfides un-der anoxic conditions created by a flux of excess organiccarbon. Marine plankton are known to concentrate traceelements, especially Ba, Pb, Ni, Cu, Zn, Mn, and Fe(Vinogradov, 1953; Goldberg, 1957; Boyle and Lynch,1968; Knauer and Martin, 1973; Martin and Knauer,1973; Vinogradova and Kovalskiy, 1962; Bostrom et al.,1974; Chester et al., 1978; Moore and Bostrom, 1978;Leinen and Stakes, 1979), and this suggests that biocon-centration is a potentially important mechanism for in-corporation of certain trace elements in organic-carbon-rich marine sediments. However, the great effectiveness

828


Table 1. Average or representative chemical analyses for modern organic-carbon-rich sediments, oxic deep-sea sediments, and Cretaceous blackshales and interbedded strata.

Component Black Seaa

Oxide or Element (%)

SiO2

A12O3

TiO2

Fe2O3

CaOMgONa2OK2OMnOP2O5SCOrg

32.58.970.294.82

20.92.211.611.720.0660.272.902.25

Element (ppm)

AsBaCdCoCrCuMoNiPbVZn

265

271074526

10810

18876

(51.8)(14.3)

(0.46)(7.69)

(3.53)(2.57)(2.75)(0.11)(0.43)

(423)

(49)(171)(72)(42)

(172)(16)

(316)(121)

BunnefjordNorway"

34.98.770.43

10.47.31.830.932.274.780.614.265.44

71552

601071352645

167276944

(40.4)(10.1)(0.50)

(12.1)

(2.10)(1.08)(2.63)(5.54)(0.71)

(82)(640)

(70)(124)(157)

(30)(52)

(194)(320)

(1094)

Mean of52 anoxicsamples"

32735

271251333355

148181571

Mean of65 oxic

samples"

26768

24113543355

148181571

Walvis Baydia tornclayc

51.22.920.231.627.261.33

0.71

1.591.599.35

198

6853

10812

68

(57.6)(3.28)(0.26)(1.82)

(1.49)

(0.80)

(1.79)

(222)

(76)(60)

(121)(13)

(76)

Baltic Sead

39.014.20.727.72.22.70.34.35.20.232.84.6

750

229078354325

130110

(42.0)(15.3)

(0.77)(8.3)

(2.9)(0.3)(4.6)(5.6)(0.25)

(806)

(24)(97)(84)(38)(46)(27)

(140)(118)

Deep-seapelagicclaye

60.419.10.928.93

3.731.673.560.520.16

132300

0.217490

25027

22580

120165

Cretaceousblack shales

from Atlantic'

2.196.18

5300.92

2720215659

18616

822828

(38)(257)(194)

(84)(263)

(21)(1281)(1207)


Site 5308

56.711.81.118.5

2.81.393.170.070.18

32660

859

18015026

13027

460320

[57.3][11.9][1.12][8.61]

[2.8][1.40][3.19][0.08][0.19]

[48][1111]

[28][69]

[222][171][26]

[181][27]

[692][782]

Cretaceousred claystone,

Site 5308

58.212.61.23

10.0

3.31.463.430.110.17

830

3311273

77

210110

Note: All values in parentheses are recalculated on a carbonate-free basis.a Average of modern Black Sea deep-water sediments; data of Hirst (1974)." Analysis of organic-carbon-rich anoxic sediment in Bunnefjord, Norway; Doff (1969), quoted in Calvert (1976).c Average of diatomaceous clay, Walvis Bay Southwest Africa; data of Calvert (1976).d Analysis of Gotland Basin anoxic sediment; data of Manheim (1961).*; Average of typical oxidized, deep-sea pelagic clay samples; data of Chester and Aston (1976).f Average of 38 samples of Atlantic black shales; data of Brumsack, 1980.g Geometric-mean concentrations of 28 samples of black shale, and 12 samples of red claystone from DSDP Site 530. Values in brackets are arithmetic-mean concentrations for the black-shale

samples. All values are on a carbonate-free basis. Data of Dean and Parduhn (this volume).

Table 2. Selected mean and range of concentration of trace elements in modern anoxicsediments and Atlantic Cretaceous "black shales" and associated strata.

Element

AsBaCdCoCrCuMoNiPbVZn

Black Seaa

Mean

327—26

14338138212

22598

Range

92-1594—

11-461-4131-2141-188

24-2651-33

144-46339-150

Green mudWalvis Bayb

Mean Range

— 1-52440 (198) 1-3660— 1-79— 1-7.7— 1-216— _14 (53) 1-10430 (108) 1-166

_ _— 12-83


DSDP Site 53Oc

Mean

32660

859

18015026

13027

460320

(48)(UIO)

(28)(69)

(222)(171)

(26)(181)

(27)(692)(780)

Range

<20-92270-4400<6-114

13-18057-57075-430

<5-19040-660

<20-100120-220068-5900


North Atlantic11

Mean Maximum

530 —0.92

2720215659

18616

822828

(38) 94(257) 430(194) 635

(84) 606(263) 870

(21) 46(1281) 3627(1207) 7158

Note: All values are in parts per million (ppm) in bulk sample, and not on a carbonate-free basis, unless otherwise noted.

a Average of all Black Sea sediment analyses in 11 cores. Data from Hirst (1974)." Mean concentrations are from Baturin and others (1970 and 1972) and Calvert and

Price (1970) or Price and Calvert (1977; in parentheses). Concentration ranges arefrom Brongersma-Sanders et al. (1980).

c Geometric mean and range on a carbonate-free basis of concentrations of elements in 28samples of black shale (from Dean and Parduhn, this volume). Values in parentheses arearithmetic means for comparison with other mean values in this table.

Mean and maximum concentrations of elements in 38 samples of black shale containingmore than 1.0% organic carbon (Brumsack, 1980). Mean values in parentheses are on acarbonate-free basis.

of absorption of trace metals by clay minerals and or-ganic matter, and coprecipitation of trace metals, par-ticularly as sulfide minerals, suggests that these pro-cesses also may play an important role in the removal oftrace metals from seawater (Tourtelot, 1964 and 1979;Volkov and Fomina, 1974; Holland, 1979).

Tourtelot (1964) studied the minor-element composi-tions of contemporaneous Upper Cretaceous marineand nonmarine shales in the western interior of theUnited States and found that element concentrations

were consistently higher in the marine shales. He con-cluded from this observation that sorption and syngene-tic mineral formation in the marine environment werean important, if not the most important, processes af-fecting the minor-element content of shales. The finalminor-element composition is thus the result of cycles ofsorption-desorption and (or) precipitation-dissolutionas the sediments are subjected to varying chemical con-ditions in depositional and diagenetic environments.The greatest control is probably exerted by redox condi-

829


tions during diagenesis. Vine and Tourtelot (1970) alsoconcluded that sorption processes are important in me-tal enrichment in black shales, but felt that some priorconcentration of metals from seawater (for example, bybioconcentration) was required. They stated that theavailability of metals in the various solutions that theorganic matter comes in contact with throughout its his-tory is probably the most significant factor in determin-ing metal enrichment, and that interstitial pore fluidsare probably the most important of these solutions.

If thin beds of black shale are interstratified withother lithologies, these beds may become sinks for ele-ments mobilized and transported in pore waters duringcompaction (Berner, 1969 and 1975). The importance ofinterbedded more-reduced and less-reduced strata onthe migration and accumulation of trace elements, andthe relative mobility of different elements under reduc-ing conditions was pointed out by Dean and Parduhn(this volume) for Cretaceous black shale interbeddedwith red and green claystone in the Angola Basin atDSDP Site 530. This 160-m-thick interbedded sequenceis Albian through Coniacian-Santonian in age and rep-resents the upper part of mid-Cretaceous black-shale de-position in the Atlantic. The black-shale beds in this se-quence have an average of 5.4 weight percent organiccarbon, compared with less than 0.5% for red and greenclay stones in the sequence (Meyers, Brassell, and Hue,this volume). Geometric-mean concentrations of traceelements from each lithology show that concentrationsof As, Cd, Co, Cr, Cu, Mo, Ni, Pb, V, and Zn are allhigher in black shales than in red and green claystones(Fig. 8). Concentrations of Ba and Fe are generallyhigher in red claystone than in green claystone and blackshale. Concentrations of Mn are generally higher ingreen claystone than red claystone, and much lower inblack shale. Dean and Gardner (1982) found the samegeneral relationships in cycles of black shale and greenand red claystone from DSDP Site 367 off northwestAfrica; Cr, Mo, Ni, Zn, and V were enriched in blackshale, whereas Co and Cu were of about equal concen-trations in all lithologies (Fig. 9).

The maximum concentrations of elements in the Site530 black shales, however, do not all occur at the samedepth (Fig. 10). Concentrations of Cd, Zn, V, Cu, andMo are highest in the part of the section where blackshale beds are most abundant, which suggests that theyhave not migrated far from their organic-matter source.These elements also are among the ones most commonlyreported to be enriched in black shales from other areas.

The pattern of trace-element mobility in the black-shale sequence at Site 530, illustrated in Figure 10, is inaccordance with basic concepts concerning the controlof redox reactions and solubility of various redox-sensi-tive elements in pore waters by a combination of organiccomplexing (e.g., Elderfield, 1981) and sulfide-carbon-ate equilibria (e.g., Berner, 1980; Suess, 1979). Iron andmanganese are examples of redox-sensitive elementsthat behave differently in anoxic sediments, and this isreflected in the black-shale sequence at Site 530. The en-richment pattern for the concentration of iron in black-shale beds is most similar to that of barium in that the

10

Parts per million100 1,000 10,000 100,000

0.001% 0.01% 0.1%

Figure 8. Observed range (length of bar) and geometric mean (symbol)of concentrations of major-element oxides and trace elements insamples of red and green claystone and black shale from DSDPHole 53OA, Cores 89-105.

concentrations are highest in the upper part of the black-shale-bearing section and lowest in the lower part of thesection, the exact opposite of the enrichment pattern formanganese (Fig. 10). In addition, the concentration ofiron is not significantly correlated with concentrationsof any of the trace elements that are closely associatedwith the black shale beds. The concentration of iron isabout the same in black shale and green claystone, andperhaps slightly enriched in red claystone (Fig. 8). Elder-field (1981) demonstrated that up to 80 percent of irondissolved in pore waters of anoxic sediments is com-plexed by dissolved organic molecules; most of these or-ganic complexes have large ionic radii and diffuse slow-ly in fine-grained sediment. The solubility of iron, there-fore, initially is controlled by organic complexing, andiron resides near its initial site of reduction and solubili-zation eventually to be precipitated as iron sulfide or asiron oxyhydroxides in sediments where the Eh is suffi-ciently high. Dissolved iron may migrate as much asseveral centimeters to local sites of sulfide or oxyhy-droxide precipitation, but probably not over distancesof tens of meters. Manganese, to the contrary, is muchmore mobile than iron because it does not form manyorganic complexes (Elderfield, 1981) and does not pre-cipitate as sulfide minerals. In fact, manganese may besolubilized and desolubilized several times as the low-Ehinterface moves upward in the sediment column, where-

830


10

Parts per million

100 1,000 10,000

Parts per million

100 1,000 10,000 100,000

Site 367Core 17

Total-C

Mn

Cu

Cr

Ni

Zr

Mo

Figure 9. A. Observed range (length of bar) and geometric mean (symbol) of concentrations of major-element oxides and trace elements on acarbonate-free basis for samples of interbedded black shale and green daystone from DSDP Hole 367, Core 17. B. Observed range Gength of bar)and geometric mean (symbol) of concentrations of major-element oxides and trace elements on a carbonate-free basis for samples of interbeddedgreen and red daystone from DSDP Hole 367, Cores 15 and 16.

as iron may remain behind depending upon the extent ofthe reduction. Some anoxic sediments may be relativelyenriched in manganese due to local precipitation of Mn-rich carbonate phases (e.g., Suess, 1979 and Pedersonand Price, 1982). That the high manganese concentra-tions in black-shale beds at Site 530 (Fig. 10) probablyare associated with carbonates is indicated by a highpositive correlation between concentrations of Mn andCa (r = 0.85; Table 3). The carbonate content of theblack shales is highest in the lower part of the section,which explains the higher concentration of manganesethere. Any manganese that was not coprecipitated withcarbonate quickly diffused out of the sediment.

Variation in concentrations of Cd, Cu, Mo, V, andZn all are similar as evidenced by similar patterns ofconcentration versus depth (Fig. 10) and high correla-tion coefficients (Table 3). Concentrations of Mo, V,Zn, and Cu in organic matter and (or) biogenic silicahave been reported by many investigators (e.g., Volkovand Fomina, 1974; Presley et al., 1972; Leinen andStakes, 1979), but once solubilized these elements candiffuse until they are scavenged by sulfides. Concentra-tions of these elements in Site 530 black shales are allpositively correlated with the concentration of P2O5(Table 3) and negatively correlated with SiO2, whichsuggests an organic association. There is no correlationbetween SiO2 and A12O3 and the Si:Al ratio is high(average of 4.8) for black, red, and green lithologies,which suggests that much of the SiO2 is biogenic, al-though no recognizable remains of siliceous organismswere observed (Site 530, summary chapter, this vol-ume). This biogenic silica must be a dilutant rather thana source of trace metals, or else the biogenic silica and

trace metals have become efficiently separated duringdiagenesis.

The distributions of concentrations of Pb and Crwith depth are very different from those of most othertrace elements (Fig. 10). Concentrations of Co and Niare highly correlated (r = 0.80; Table 3) and have max-ima in certain black-shale beds (Fig. 10), but not in thesame beds as most of the other trace elements. In the in-terbedded red clay stones, concentrations of Co, Cr, Cu,Ni, V, and Zn are all lower than in the black shales butare all highly correlated with each other (r > 0.90 for allcorrelations) and are all highly correlated with the con-centration of barium (r > 0.8 for all correlations).

Brumsack (1980) analyzed 38 samples of Cretaceousblack shale from the Atlantic having more than 1 % or-ganic carbon (average of 6.12%; Table 2), and com-pared the values obtained from 11 samples of associatedlithologies containing 0.5 to 1.0% organic carbon, and38 samples of associated lithologies containing less than0.5% organic carbon. His results showed enrichment ofCu, Cr, Mo, Ni, and Zn in organic-carbon-rich bedsrelative to associated organic-carbon-poor beds. Fe andMn concentrations were highest in black and red bedsbut not in green beds. Tables 1 and 2 compare Brum-sack^ (1980) data with those from Site 530 (Dean andParduhn, this volume). The mean values for samplesfrom Site 530 given by Dean and Parduhn and plotted inFigure 8 are geometric means, which are more realisticmeasures of central tendency because trace-element con-centrations usually have frequency distributions thatmore closely approximate log-normal than normal dis-tributions. In Tables 1 and 2, however, the Site 530 dataare expressed as arithmetic means for purposes of com-

831


% Black beds C o r g M A R F e 2 ° 3 ~ c f Fe:AI-blk. MnO-cf Ba-cf Co-cf Ni-cf

0 25 50 0 25 50 6 9 12 0.4 0.8 1.2 0.0 0.1 0.2 0 2000 4000 0 100 200 0 300 600

0 25 50 0 25 50 6 9 12 0.4 0.8 1.2 0.0 0.1 0.2 0 2000 4000 0 100 200 0 300 600% Black beds C Q r g MAR Fe 2 O 3 -cf Fe:AI-blk. MnO-cf

Pb-cf Cr-cf Cu-cf Cd-cf Mo-cf V-cf Zn-cf

0 50 100 0 300 600 0 250 500 0 50 100 0 100 200 0 1500 3000 0 3000 6000

-

- * -

‰:

I• i • 1 1 1 "

•

j- -\ -

•

0 50 100

Pb-cf

300 600

Cr-cf

0 250 500

Cu-cf

0 50 100

Cd-cf

0 100 200

Mo-cf

1500 3000 0 3000 6000

V-cf Zn-cf

Figure 10. Plots of carbonate-free (cf) concentrations versus depth for Fe2θ, MnO, Ba, Co, Ni, Pb, Cr,Cu, Cd, Mo, V, and Zn in samples of black shale from DSDP Site 530, southern Angola Basin. Con-centrations of Fβ2θ and MnO are in percent; all other concentrations are in parts per million. TheFe:Al ratio in black-shale samples, the percentage of black-shale beds in each 9.5-m cored interval,and the mass accumulation rate (MAR; g/cm2/m.y.) of organic carbon for each 9.5-m cored intervalalso are plotted. The vertical line through each concentration plot is at the geometric-mean con-centration for that element or oxide. All values greater than the geometric mean are shaded. Thenumbers (87-105) within the plot for percentage of black-shale beds are the numbers of each core. Thetwo horizontal lines drawn through all plots are drawn at the top of Core 97 and the bottom of Core 98and mark the zone of maximum black-shale-bed concentration (Fig. 3).

parison. The average concentrations of Cr, Cu, Ni, andPb in both sets of black-shale samples are comparable,whereas Brumsack's (1980) average concentrations ofMo, V, and Zn are about twice those of the Site 530black shales, and concentrations of Ba, Cd, and Co arehigher in Site 530 black shales. Some of these differ-ences may reflect analytical uncertainties, but it appears

that, in general, Cretaceous black shales from the At-lantic are similarly enriched in trace elements.

In comparison to modern anoxic sediments from avariety of environments, the Site 530 black shales are atleast as enriched or more enriched in Ba, Cd, Co,Cr, Cu, Ni, Pb, V, and Zn (on a carbonate-free basis;Table 1). Of the trace elements analyzed, only Mo is de-

832


Table 3. Correlation coefficients among concentrations of selected trace elements and major-element oxides in 28 samples of black shale from DSDPSite 530 (computed from data of Dean and Parduhn, this volume).

SiO2

A12O3

Fe2O3

MgOCaOK2O

P2O5MnOBaCoCrCuMoNiPbV

A12O3

0.01

Fe2O3

-0.490.08

MgO

0.640.34

-0.43

CaO

0.20-0.48-0.26

0.20

K2O

0.150.710.200.35

-0.39

P2O5

-0.63-0.17

0.10-0.38

0.19-0.09

MnO

0.18-0.41-0.18

0.200.85

-0.220.18

Ba

-0.070.300.22

-0.25-0.22

0.250.05

-0.19

Co

-0.560.010.66

-0.45-0.40

0.100.23

-0.370.51

Cr

-0.350.060.36

-0.33-0.45

0.010.20

-0.560.260.38

Cu

-0.86-0.14

0.31-0.57-0.11-0.35

0.55-0.21

0.400.490.53

Mo

-0.75-0.35

0.06-0.56

0.23-0.46

0.660.16

-0.070.310.030.67

Ni

-0.82-0.17

0.51-0.68-0.21-0.17

0.55-0.20

0.350.800.180.630.66

Pb

0.180.210.040.240.080.00

-0.180.03

-0.07-0.11

0.13-0.11-0.10-0.23

V

-0.87-0.30

0.22-0.68-0.02-0.45

0.61-0.08

0.100.460.150.800.870.82

-0.22

Zn

-0.75-0.39

0.11-0.59

0.17-0.49

0.470.110.000.180.050.760.850.550.170.90

pleted in the Site 530 black shales relative to modernanoxic sediments. In general, the trace element enrich-ments in Site 530 deep-water black shales, as well asmost of those reported by Brumsack (1980) and Deanand Gardner (1982) are higher than those in sedimentsfrom modern anoxic environments from the Black Seaand Bunnefjord (Table 1). These two modern anoxic en-vironments have anoxic water columns above the sedi-ment rather than only anoxic sediment-interstitial wa-ters. The ranges and mean concentrations of most traceelements in Site 530 black shales also are larger thanthose in very organic-carbon-rich green muds depositedwithin an intense oxygen-minimum zone in Walvis Bay,Southwest Africa (Table 2).

Enrichments of Cd, Mo, Ni, Co, and Zn can be re-lated to enhanced supply and preservation of organicmatter (e.g., Bruland et al., 1980), with subsequent lib-eration during decomposition of the organic matter andscavenging by sulfides during early diagenesis in anoxicsediment. The high concentrations of Zn and Co alsomay have been supplied directly to the sediment as metalsulfide particles produced in an anoxic, sulfide-richwater column as in the modern anoxic Black Sea (e.g.,Spencer et al., 1972; Brewer and Spencer, 1974), butthere is other evidence to suggest that the black shales atSite 530, and probably in much of the Cretaceous At-lantic Ocean, were not deposited under an entirely anox-ic water column.

A possible example of trace-element enrichment bypore waters in oxidized sediment overlying reduced sedi-ment was recovered at DSDP Site 105 in the NorthAmerican Basin of the western North Atlantic Ocean.Here, the middle Cretaceous black shale unit is overlainby Upper Cretaceous-lower Tertiary red and green clays(Lancelot et al., 1972). Jansa et al. (1979) interpretedthe oxidized clays as having been deposited in an oxy-genated, pelagic deep-sea environment similar to the en-vironment presently accumulating pelagic clays in thecentral North Pacific. This unit was recovered at eightother sites in the North American Basin (Jansa et al.,1979), but at Site 105 the multicolored clays are enrichedin heavy metals, especially Mn, Zn, Cu, Pb, Cr, Ni, and

V; this led Lancelot et al. (1972) to suggest that the clayswere enriched in metals by hydrothermal volcanic exhal-ations and in this regard were analogous to metal-en-riched sediments associated with Red Sea brine deposits,or metal-enriched basal sediments overlying basementalong the East Pacific Rise.

Arthur (1979b) proposed that the metals in the ox-idized clays at Site 105 were not derived from volcanicexhalations but from chemically reduced metallic ions inthe anoxic pore waters of the underlying black shales.The reduced ions in pore waters diffused upward andwere oxidized in the slowly accumulating Upper Creta-ceous clays. The equivalent clays at DSDP Site 367 inthe Cape Verde Basin in the eastern North Atlantic,however, are not enriched in any metals, nor are thereany significant differences in metal concentrations be-tween interbedded red and green clays (Dean and Gard-ner, 1982).

The relative mobility of trace elements in anoxic porewaters is of interest because we would like to understandthe patterns of trace-element enrichment in black shalesand to use them as indicators of depositional environ-ment and paleocean chemistry. The fact that some ele-ments may migrate over fairly large distances (tens ofmeters or more) in anoxic pore waters may invalidatebulk analyses of black-shale beds if only a few selectedbeds are analyzed. For example, if we had only analyzedblack-shale beds at the bottom of the section at Site 530(below 1050 m, Fig. 10), we would not have found themost significant trace-element enrichments. It is alsoimportant to examine the compositions of interbeddedlithologies because of trace-element migration. Relativeconcentrations (parts per million) of trace elements alsomay be misleading because of dilution effects in high-sedimentation-rate sequences. For example, a compar-ison of trace-element concentrations between blackshales and average red clay (Table 1) suggests that aver-age pelagic red clay is as enriched in Co, Cu, Ni, Pb,and Zn as many black shales. However, the accumula-tion rates of these trace elements probably are higher inmost black shales because the average rates of bulk-sediment accumulation are higher.

833


Table 4. Mass accumulation rates of selected components in red, green, and black lithologies, Site 530.

Core

87888990919293949596979899

100101102103104105

Depth(m)

940-949949-958958-967967-976976-985985-990990-999999-1,008

1,008-1,0171,017-1,026-1,035-1,044-1,053-1,062-1,071-

1,026,035,044,053,062

1,071,080

1,080-1,0851,085-1,0941,094-1,103

Recovery(cm)

703448922446574

0918273656795607501735672995800600790800

Blackbeds(ft)

8635004659

52369

125328

10

Redbeds(ft)

72748785950

9079804500

20253570600

20

Greenbeds(ft)

20201010506

151546486471666025379472

Period(I03 yr.)a

40617061

3930

5237614622103339

12922222221955

Znb inblack beds

(ft)

0.1800.140

—0.1100.00.00.0140.0420.0090.0230.2750.3000.0190.0120.0130.0130.0140.0120.014

Cr b inblack beds

(ft)

0.0290.045

—0.0180.00.00.0250.0250.0920.0330.0400.0160.0340.0200.0420.0270.0080.0100.010

Nib inblack beds

(ft)

0.0360.039

—0.049

—0.00.0300.0150.0650.0080.0060.0300.0090.0100.0120.0200.0110.0100.005

Znb inred beds

(ft)

0.0100.0070.025

—0.0110.0

0.0070.009

———_—

0.0090.006

——

0.009

Crb inred beds

(ft)

0.0110.0110.025

—0.0100.0

0.0080.009

———

—0.0090.006

——

0.006

Nib inred beds

(ft)

0.0040.0040.015

—0.0060.0

—0.0060.007

—————

0.0090.004

——

0.008

Zn b ingreen beds

(ft)

0.0080.013

—0.0130.0150.0

—0.0160.0310.0140.0130.0240.0070.010

—0.0100.0100.0100.008

Crb ingreen beds

(ft)

0.0150.019

—0.0150.0150.0

—0.0120.0120.0080.0100.0070.0070.008

—0.0100.0060.0090.008

* Based on average accumulation rates of 14.6 m/m.y. for Cores 87-94 and 9 m/m.y. for Cores 95-105.b Data of Dean and Parduhn (this volume). Dash indicates no analyses for that core.c MAR for bulk sediment = [(accumulation rate) 100 (1 - porosity/100) grain density] (Thiede and Rea, 1981), where accumulation rate = 14.6 m/m.y. for Cores 87-94 and

9.0 m/m.y. for Cores 95-105; porosity = 33ft for all cores, and grain density is assumed to be 2.7 g/cm3.d MAR for C o r g = (MAR of black beds) (5.4/100), where 5.4 is the grand average concentration of organic carbon in black-shale samples analyzed by Meyers, Brassell,

and Hue (this volume).e MAR for trace elements = (MAR of each colored bed) (<% of trace element in each colored bed/100).

For each 9.5-m cored interval in the black-shale-bear-ing sequence at Site 530, we calculated mass accumula-tion rates (MAR; g/cmVm.y.) for bulk sediment, foreach different-colored lithology, for organic carbon, forthree trace-elements (Zn, Cr, and Ni) in each different-colored lithology, and for the sum of each of the threetrace elements (Table 4). An average black shale was as-sumed to have a concentration of organic carbon of5.4%, which is the average of analyses reported byMeyers, Brassell, and Hue (this volume). The MARswere calculated (Table 4) using (1) estimates of the per-centages of black, red, and green beds in the recoveredpart of each 9.5-m cored interval, (2) biostratigraphic-ally determined accumulation rates and porosity (Site530 summary, this volume), and (3) an average grain den-sity of 2.7 g/cm3. The MAR for organic carbon variesfrom 0 to 46 g/cm2/m.y. As expected, maximum valuesfor organic-carbon and trace-element MARs occur inCores 97 and 98 where black-shale beds are far moreabundant than in any other part of the section (Fig. 10).Maximum MARs for Zn, Cr, and Ni in black-shale bedsare 2.38, 0.35, and 0.18 g/cmVm.y., respectively, inCores 97 and 98. MARs for these elements are all about0.01 to 0.02 g/cm2/m.y. in the lower part of the black-shale-bearing sequence, and 0.01 to 0.4 g/cmVm.y. inthe upper part of the sequence (Table 4). The changes inZn MAR are the most extreme. Maximum values of totalMARs for Zn, Cr, and Ni are 2.48, 0.66, and 0.42 g/cm2/m.y., respectively, and only the maximum for Zn occursin the zone of maximum black-shale concentration inCores 97 and 98. Total MARs for all three elements areuniformly low in the lower part of the black-shale-bear-ing section, and relatively high in the upper part of thesection (Cores 87 to 96). For comparison, Leinen andStakes (1979) calculated average MARs of Zn and Ni of0.17 g/cm2/m.y. and 0.08 g/cm2/m.y. respectively for

sediment on the crest of the East Pacific Rise, and 0.06g/cm2/m.y. and 0.03 g/cm2/m.y. respectively, for pe-lagic sediment in the eastern equatorial Pacific Ocean.The average MARs from Leinen and Stakes (1979) aresomewhat higher than the average MARs of Site 530black shales, but lower than total MARs in all cores ex-cept those in the lower part of the section. As we ex-pected, the percentages of total MAR for each elementthat are contributed by black shale are highest in Cores97 and 98, but black shales also contribute significantproportions of the total MARs for Zn and Ni in Coresabove 96 (as much as 63% of the Zn MAR in Core 87),although black-shale beds comprise less than 8% of thispart of the section (Fig. 3).

Site 530 probably is an example of the lower limit oftrace-element accumulation in Cretaceous black shales.The organic-carbon accumulation rates are low (max-imum of 46 g/cm2/m.y.) compared with other AtlanticCretaceous black-shale sequences. For example, Arthurand Natland (1979) calculated that organic carbon in theAptian-lower Albian black-shale sequence at Site 361 inthe Cape Basin accumulated at an average rate of 258 g/cm2/m.y., and organic carbon in lower to middle Al-bian black, dolomitic sapropels, and marly limestonesaccumulated at rates of about 500 g/cm2/m.y. We cal-culated that organic carbon in Albian-Turonian blackshales at Site 367 in the southeastern North Atlantic offnorthwest Africa accumulated at rates of about 460 g/cmVm.y. using measured organic-carbon, porosity, andsedimentation-rate data of Lancelot, Seibold, et al.,1978. For the Site 367, we calculated Zn and Ni MARsof 90 and 50 g/cm2/m.y. respectively, using averageconcentrations of 330 and 180 ppm respectively (data ofLange et al., 1978). These values clearly are at least twoorders of magnitude higher than those in pelagic claysstudied by Leinen and Stakes (1979). The trace-element

834


Table 4.

Nib ingreen beds

( « )

0.0100.005

—0.0080.0340.0

0.0070.0050.0060.0050.0060.0060.006

0.0050.0050.0080.005

(Continued.)

MARBulk sed.c

(g/cm^/m.y.)

2,6412,6412,6412,6412,641

02,6412,641

,664,664,664,664,664,664,664,664,664,664,664

MARBlack beds

(g/cm^/m.y.)

21115879

13200

10615883

150865599150200835033

13317

MARRed beds

(g/cirr/m.y.)

1,9011,9542,2982,2452,509

02,3772,0861,331

74900

333832582

1,165998—333

MARGreen beds

(g/cmVm.y.)

528528264264132

0158396250765799

1,0651,1811,098

998832616

1,5641,198

MARCorgd

(g/cm Vm.y.)

11847006848

46328

1143271

MAR-Znin blacke

(g/cmVm.y.)

0.4000.2200.2000.1400.00.00.0100.0700.0100.0302.3801.8002.0200.0200.0100.0100.0050.0200.002

MAR-Crin blacke

(g/cm2/m.y.)

0.0600.0700.0500.0200.00.00.0300.0400.0800.0500.3500.1000.0500.0400.0300.0100.0030.0100.002

MAR-Niin blacke

(g/cmVm.y.)

0.0800.0600.0600.0600.00.00.0300.0200.0500.0100.0500.1800.0100.0200.0100.0100.0040.0150.001

MAR-Znin rede

(g/cm2/m.y.)

0.190.140.570.250.280.00.170.150.120.070.00.00.030.070.050.070.070.00.03

MAR-Crin rede

(g/cm2/m.y.)

0.190.210.570.220.230.00.190.170.120.070.00.00.030.070.050.070.060.00.02

MAR values for Site 367 also are much higher thanthose estimated for modern anoxic deep-basin sedi-ments from the Black Sea (Table 5), although the MARsfor organic carbon are about the same as those for Site367.

Cretaceous black shales obviously were an importantsink for trace metals in the Cretaceous Atlantic Ocean.We think that the above concepts and observations de-serve further research, and may ultimately explain muchof the variability in metal enrichment in black shales ingeneral. Studies of major and trace-element geochemis-try of black-shale sequences, in conjunction with studiesof sulfur, organic carbon, and changes in sedimentationrate, also are needed.

MODELS OF DEPOSITION

Bottom-Water Oxygenation

Interbedding of sediments alternately rich and poorin organic matter can be the result of either fluctuatingpreservation of organic matter or fluctuating rate ofsupply of organic matter. Differential preservation canresult if bottom waters at the site of accumulation arealternately oxic and anoxic (or near-anoxic). Therefore,alternating periods of oxygen-deficient and relativelywell-oxygenated bottom waters may be required if thesite of accumulation is on a basin floor. Cyclic variationin amount of organic matter deposited at a continentalmargin site can be explained by cyclic fluctuations in thethickness and intensity of a midwater oxygen-minimumlayer. Both depositional models (basin deoxygenationand expansion and intensification of an oxygen-mini-mum layer) have been proposed to explain the accumu-lation of organic carbon-rich strata (e.g., Schlanger andJenkyns, 1976; Ryan and Cita, 1977; Fischer and Ar-thur, 1977; Thiede and van Andel, 1977; Arthur andSchlanger, 1979). Both models require very low to zeroconcentrations of dissolved oxygen in part of the watercolumn as a result of reduced advection of oxygenatedwater and (or) increased supply of organic matter, and

both models imply that reducing conditions in the sedi-ments (and therefore the increased degree of preserva-tion of organic matter) are the result of anoxic or near-anoxic conditions in the overlying waters.

Black, laminated, Corg-rich Cretaceous strata are sowidespread in the Atlantic that a closed-basin model hasbeen proposed as the most likely analog for the Earlyto middle Cretaceous South and North Atlantic basins.The implication is that there was a stable salinity strati-fication, either because of lower-salinity surface wateroverlying bottom water of more-normal salinity as inthe modern Black Sea, or because of normal-salinitysurface water overlying more-saline bottom water, whichled to increased residence time of deep-water masses andgradual oxygen depletion with or without increased sur-face productivity. A higher-salinity bottom-water masswould have a lower oxygen solubility which, togetherwith an increased residence time, would result in greateroxygen depletion. The salinity-stratification model ismade even more plausible by the fact that isolated At-lantic basins were accumulating great thicknesses ofevaporites in the Late Jurassic and again in the Aptian,and therefore probably were more saline as well as morestagnant in the Early Cretaceous. In contrast, Brass etal. (1982) suggested that sinking of warm, saline bottomwater produced over warm, arid, extensive shelf areascould generate bottom-water turnover rates as high asthose of modern thermqhaline deep circulation. Modelstudies by Southam et al. (1982) suggest that it would bedifficult to create a totally anoxic water column in theopen ocean because of the delicate balance between ox-ygen demand and supply to deeper waters, and nutrientupwelling to support surface productivity. However, theBlack Sea sediment record is proof that rapid variationsin oxygenation and anoxia can occur in large, isolatedbasins (Deuser, 1974; Degens and Stoffers, 1976).

Supply of Organic Carbon to Sediments

We find it difficult, however, to explain all of the in-terbedding of more reduced and less-reduced strata byfluctuating deoxygenation of bottom waters particularly

835


Table 4. (Continued.)

Core

87888990919293949596979899

100101102103104105

MAR-Niin rede

(g/cmVm.y.)

0.080.080.340.130.140.00.140.130.090.060.00.00.030.060.050.030.060.00.03

MAR-Znin greene

(g/cmVm.y.)

0.040.070.030.030.020.00.030.060.080.110.100.260.080.110.100.080.060.160.10

MAR-Crin greene

(g/cnrr/m.y.)

0.080.100.040.040.020.00.020.050.030.060.080.070.080.090.100.080.040.140.10

MAR-Niin greene

(g/cm^/m.y.)

0.050.030.020.020.040.00.010.030.010.050.040.060.070.070.060.040.030.130.06

MAR-Zntotal

(gW/m.y.

0.630.430.800.420.300.00.210.280.210.212.482.060.130.200.160.160.140.180.13

MAR-Crtotal

(g/cmVm.y.)

0.330.380.660.280.250.00.240.260.230.180.430.170.160.200.180.160.100.150.12

MAR-Nitotal

(g/cmVm.y.)

0.210.170.420.210.180.00.180.180.400.120.090.240.110.150.120.080.090.150.09

MAR-Zn inblack as °?oMAR-Zn

total

63512533

004

255

1496871610664

112

MAR-Cr inblack as °IoMAR-CR

total

18188700

1315352881593120176372

MAR-Ni inblack as "laMAR-Ni

total

38351429

00

171125

85675

9138

134

101

Table 5. Organic carbon, phosphorus, and trace-element concentra-tions and accumulation rates in modern sediments from the BlackSea.

COrgPZnNiCuCr

Core no.

Concentration(wt.%)

2.080.10

91 × I 0 " 6

149 × 10~ 6

56 × 10~6

167 × 10~6

1432a

Accumulationrate

(gW/m.y.)178

8.470.771.260.471.41

Core no.

Concentration

67673752

(wt

3.0.

××××

.%)

581710~6

1 0 ~ Λ10 ~ 6

10~ 6

1462b

Accumulationrate

(g/cm2/m.y.)

52524.9

0.980.980.540.54

Note: Data of Hirst (1974); chronology from Ross and Degens (1974); bulk-sediment accumulation rates were computed using a dry-bulk density of1.1 g/cm3 (Keller, 1974).

a Depth, 2248 m; average bulk-sediment accumulation rate, 8470 g/cmr/m.y.

D Depth 2179 m; average bulk-sediment accumulation rate, 14,660 g/cmz/m.y.

in deep-basin settings. Most of the Jurassic and Creta-ceous strata in the Atlantic contain abundant evidencethat they were deposited in well-oxygenated bottom-wa-ter environments with at least some benthic infauna; on-ly the interbedded dark-colored, organic carbon-richlithologies suggest anaerobic depositional conditions be-cause they commonly are laminated and apparently un-bioturbated. Ekdale (1980) has shown that trace-fossildiversity varies within Cretaceous deep-water redoxcycles, decreasing in organic-carbon-rich beds. Chon-drites, a shallow-burrowing, branching ichnofossil type,is typically the last trace fossil found at the transitionbetween a green or red and a black bed, and the first tocolonize the bottom following the deposition of lami-nated black-clay layers. Our data from Site 530 confirmEkdale's observations (Stow and Dean, this volume).This suggests that at least some of the redox cycles maybe the result of a change in bottom-water oxygen con-centration at or above the sediment/water interface. In-terbedding of oxic-anoxic strata in the North Atlanticbegan in the Jurassic and continued at least into the

Eocene. However, only the Lower to middle Cretaceousand Eocene strata contain more than 1% organic car-bon. It is not difficult to visualize abrupt changes inbottom-water circulation in a restricted basin such asthe Early Cretaceous Atlantic Ocean, but it is difficultto imagine total stagnation of the Atlantic during theLate Cretaceous and Tertiary.

We therefore suggest that a third depositional modelis necessary to explain some examples of interbeddedmore-reduced and less-reduced strata. This model in-volves varying diagenetic redox conditions in the sedi-ments that are caused by varying oxygen demand withinthe sediments, largely in response to varying rates of ac-cumulations of marine and (or) terrestrial organic de-tritus. Figure 11 summarizes our concept of how varia-tions in rate of supply of organic matter may result incyclic interbeds. This model does not necessitate majorchanges in bottom-water oxygen content, only that thesupply of organic matter increases, either by periodic in-creases in surface productivity or by periodic rede-position of organic matter in the deep basin from siteson the outer shelf or slope where more organic matter ispreserved in sediments, for example, under an oxygen-minimum zone. Similar environmental interpretationsof cyclic facies of the Liassic of Yorkshire, England,were made by Morris (1979) and Demaison and Moore(1980).

In the environment illustrated in Figure 11 A, the rateof supply of organic detritus is low, the bottom watersare more or less well oxygenated, and the zero-oxygenisopleth is at some distance below the sediment-waterinterface, and red clay accumulates. Iron would be pres-ent mainly as ferric oxides and hydroxides and sulfurwould be present mainly as SO^~. In the environment il-lustrated in Figure HB, the rate of supply of organicdetritus is higher (indicated diagrammatically by moreblack arrows); the bottom waters are still more or lesswell oxygenated but increased oxygen demand in thesediments, mainly from the increased decomposition ofthe organic matter, has resulted in a zero-oxygen iso-pleth at or near the sediment-water interface, and in the

836


Organic detritus

Red clay SO4= , Fe2O3; Fe(OH)3

Green clay SO4~; H2S; HS FeS; Fe""

O , = 0

Black clay FeS

Figure 11. Variations in rate of organic decomposition and position ofthe boundary between oxidizing and reducing conditions (O2 = 0isopleth) with variations in rate of supply of organic detritus todeep-sea sediment. A. A low rate of supply of organic detritus re-sulting in red sediment. B. A moderate rate of supply of organicdetritus resulting in green sediment. C. A rapid rate of supply oforganic detritus resulting in black sediment. See text for discus-sion.

accumulation of green, reduced sediments. Iron wouldbe present mainly as Fe2+ and ferrous sulfide (FeS orFeS2), and sulfur would be present as SO^~, H2S, HS~,and FeS/FeS2 In the environment illustrated in FigureHC, the rate of accumulation of organic detritus is sogreat that the sediments are completely anoxic andreduced, the zero-oxygen isopleth may be some distanceabove the sediment-water interface, and H2S may accu-mulate locally in the bottom waters. The extent of de-oxygenation of the bottom waters would be controlledby the rate of oxygen consumption, rate of H2S pro-duction, rate of oxygen supply to the bottom waters,strength of bottom currents (advection of oxygen), andrate of sedimentation. Sluggish bottom-water circula-tion and (or) extreme drawdown of oxygen by oxidationof organic carbon in the water column would aid in H2Saccumulation and may even be necessary for this to oc-cur, but the primary cause of anoxic conditions wouldbe the high rate of biological oxygen demand within thesediments. Iron would be present as ferrous sulfide (plussome Fe2+), sulfur would be present mainly as SO|~,H2S, HS~, and FeS, and the benthic epi- and infaunawould be eliminated. The black color would result

mainly from organic matter and FeS. The FeS wouldeventually be converted to FeS2 (pyrite or marcasite)which would tend to lighten the color of the sediments(Berner, 1970). However, maturation of the organicmatter would tend to darken the color of the sedimentswith the result that the black muds probably wouldbecome black, or at least dark gray, shales through in-creasing thermal alteration and diagenesis. If the initialdeposition resulted in interbedding of layers more andless enriched in organic matter but otherwise homogene-ous, then migration of SO2,- into layers enriched inorganic matter to replace SO2.- consumed by sulfatereduction may produce localized concentrations of FeSand an intensification of color differences (Berner, 1969).With time, these localized sulfide concentrations wouldproduce concentrations of pyrite at the boundary be-tween interbeds of more- and less-organic-carbon-richmud rocks (Fig. 3).

One problem with this model is how organic-carbonsupply to deep-water sediment can increase withoutcreating an excess of production over consumption oforganic carbon in the water column. In the modernocean, it has been shown that less than 0.1% of theorganic carbon produced survives to a depth of 4000 mor greater (Suess, 1980) because of high rates of con-sumption and oxidation in the water column. It is there-fore unlikely that changes in surface productivity alonecould account for organic-carbon enrichment in theblack shales without concomitant depletion of dissolvedoxygen in the bottom waters. However, enough organiccarbon could be supplied periodically to cause inter-bedding of green and red sediment because such changesin oxidation state are found in Holocene oceanic deep-water sediments (Lynn and Bonatti, 1965; McGeary andDamuth, 1973; Gardner et al., 1982). Presumably theperiodicity in initial organic-carbon-preservation wouldbe aided by periodic increases in sedimentation rate(Heath et al., 1977; Muller and Suess, 1979) which couldalso explain the red and green cycles.

Influences of Turbidity Currents

The Corg-rich strata in the Atlantic Ocean commonlyare interpreted as pelagic sediments that accumulatedunder anoxic conditions. However, there are now manyexamples described from DSDP sites which show thatturbidity currents have been at least partly responsiblefor the deposition of some black shales.

In the western North Atlantic, Sites 101 and 105(Lancelot et al., 1972), 386 and 387 (McCave, 1979a),391 (Benson, Sheridan, et al., 1978), 417 and 418 (Don-nelly et al., 1980), and 534 (Sheridan, Gradstein, et al.,1982) all penetrated mid-Cretaceous black shales. Theblack shales are interbeded with reddish, greenish, andgrayish mudstones and marlstones, and commonly makeup less than 50% of the section. The black shales arevariously laminated, massive, and bioturbated, andcontain variable concentrations of organic carbon ofmixed terrigenous and marine origins (Tissot et al., 1979and 1980). Fine-grained turbidites make up parts (main-ly less than 30%) of the mid-Cretaceous section at eachsite, occurring in both black shales and associated li-

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thologies. There are thin-bedded siltstone and mudstoneturbidites, mostly derived from the North Americancontinental margin, as well as thicker-bedded biogenicturbidites, probably derived locally from intrabasinhighs. With more detailed analyses, however, the per-centage of turbidites may well increase for, as McCave(1979b) found after detailed size analyses of sedimentsfrom Sites 386 and 387, "More units in the black andgreen mudstone sequence may be graded and of possibleturbidite origin than at first presumed."

In the eastern North Atlantic off the Euro-Africancontinental margin, Sites 367, 368, 369 and 370 (Dean etal., 1977; Dean and Gardner, 1982), 397 (Cornford,1979), 398 (Arthur, 1979a), 400 and 402 (de Gracianskyet al., 1979), and 415, and 416 (Lancelot, Winterer etal., 1980) also penetrated mid-Cretaceous black shalesthat show evidence of turbidite deposition. The blackshales again form part (usually less than 50%) of a morefully oxidized section comprising multicolored mud-stones and marlstones. Most of the turbidites recog-nized are fine-grained siltstone or mudstone, althoughat Sites 370, 397, 398, 402, and 416 there are interbedd-ed coarser-grained and thicker-bedded turbidites, de-bris-flow deposits, and slumped units. There is no con-sistent association of turbidites with any one particularfacies. For example, at Site 398 thin-bedded turbiditesgrade upward into biogenic pelagic strata, and at Site400 biogenic calcilutite turbidites grade up into blackshales, parts of which may represent pelagic sedimenta-tion.

It appears that the Eocene black clays in the NorthAtlantic at Sites 367 (Dean et al., 1978; Dean and Gard-ner, 1982) and 386 (McCave, 1979b) are more consis-tently fine-grained turbidites that have been redepositedinto a fully oxic oceanic basin. They are very similar tosome of the Neogene black-shale turbidites in the Medi-terranean and Black seas (e.g., Sites 378 and 380, Hsü,Montadert et al., 1978).

At least three DSDP sites in the South Atlantic thathave recovered mid-Cretaceous black shales show clearevidence of turbidites. At Site 361 (Natland, 1978; Ar-thur and Natland, 1979) a sequence of thick sandstoneand pebbly sandstone turbidites and slumps are as-sociated with green, gray, and black mudstones. Thecoarse-grained turbidites commonly are rich in terrige-nous organic debris and grade upward into thinly lami-nated black shales, some of which contain more marineorganic matter. Both pelagic and turbiditic processesmay have been involved in the deposition of these units.The black shales and associated facies at Sites 364(Bolli, Ryan et al., 1978) and 530 (Stow and Dean, thisvolume) have a more distal or basinal character.

From the above discussion it is clear that turbiditescommonly are associated with black shales in the At-lantic Ocean. However, they rarely form the dominantpart of the section. Some black shales were depositedentirely or partly as thin-bedded siltstone or mudstoneturbidites, whereas others were deposited by pelagicprocesses. Where turbidites do occur in black-shale se-quences they provide: (1) a mechanism for periodicallyincreasing the supply of organic matter to sediments; (2)

an element of randomness to the cyclicity of Corg-richand Corg-poor strata, both regionally and temporally;and (3) a means of transporting oxygen-depleted shelfwaters to ocean basins and (or) causing bottom-watermixing.

Climatic and Oceanographic Controls

Turbidity currents and pelagic deposition, operatingseparately or together, can provide a variable supply oforganic-carbon-rich sediment and result in the inter-bedding of lithologies of varying color and organic con-tent. But why was the supply of organic debris so muchgreater during the middle Cretaceous? Local pockets ofaccumulation of organic matter in continental marginareas today are the result of local circulation patterns(e.g., Walvis Bay, southwest Africa; Santa Barbara Ba-sin, southern California). It is unlikely that such condi-tions would have existed around the entire margin of theEarly to middle Cretaceous Atlantic Ocean as well as onisolated plateaus and seamounts in the Pacific (Schlang-er and Jenkyns, 1976; Dean et al., 1981; Thiede et al.,1982) without the aid of some worldwide conditions ofocean circulation and productivity. Global climate dur-ing the Early to middle Cretaceous was warm, eustaticsea levels were high, spreading rates were fast, pelagicsediments in the world ocean were accumulating rapid-ly, and oceanic surface- and bottom-water temperatureswere high (Douglas and Savin, 1975; Fischer and Ar-thur, 1977; Brass et al., 1982). Increased surface-watertemperatures would have had two main effects: first,thermohaline deep-water circulation, driven today bythe sinking of cold, oxygen-rich surface waters in highlatitudes, would have been more sluggish; and second,the warmer water would contain lower concentrationsof dissolved oxygen. Bottom-water circulation was suf-ficient to supply some oxygen to maintain oxidizingconditions in the deep basins of the Pacific but slowenough to permit depletion of dissolved oxygen atmid-water depths in areas of high productivity of or-ganic matter. Accumulations of organic-carbon-rich sedi-ments at many places in the world ocean at times thatwere not always strictly synchronous undoubtedly arethe results of coincidences of several factors acting toproduce and preserve organic matter. During the middleCretaceous, much of the world ocean may have been sopoised that relatively small changes in the flux of organ-ic matter and (or) circulation at any one place may havecaused anoxia or near-anoxia within midwater oxygen-minimum zones and possibly, under extreme condi-tions, throughout much of the bottom-water mass. Anexpanded and intensified oxygen minimum would ex-plain the excellent preservation of organic carbon, andthe increase in accumulation rate of organic carbon overa much larger area in slope and deep-sea (largely by re-deposition) environments during much of the Early andmiddle Cretaceous. It is not inconceivable that the in-tensified oxygen-minimum zone extended as deep as2500 to 3000 m; this would help to explain the presenceof Albian black shales containing well-preserved organ-ic carbon on or near the crest of the Mid-Atlantic Ridgein the North Atlantic (Sites 386 and 417/418) where it is

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unlikely that continental-margin-derived turbidites couldhave supplied all of the organic matter.

Interrelationship of Variables

From the above discussion, it is apparent that the ori-gin of interbedded more- and less-reduced lithologieswith variable amounts of organic matter and variableamounts of pelagic, hemipelagic, and terrigenous sedi-ment is complex and probably is not the result of anyone simple process. In Figure 12 we have attempted toshow a range of variables that may have operated dur-ing a period of time when less organic matter was ac-cumulating (Fig. 12A; e.g., Albian-Cenomanian), incontrast with a period of time when more organic mat-ter was accumulating (Fig. 12B; Aptian-Albian or Ce-nomanian-Turonian). The main variables are the rateof influx of terrigenous sediment and organic matter,and surface-water productivity. The amounts of organicmatter from these two sources (terrestrial and marine) inturn determine the thickness and intensity of a mid-water oxygen minimum zone. Surface-water produc-tivity is determined, at least in part, by the intensity ofupwelling of nutrient-rich water from the oxygen-mini-mum zone. Other variables are the frequency and mag-nitude of turbidity currents, as well as climatic, Oceano-graphic, and paleogeographic conditions.

In the situation shown diagrammatically in Figure12A, surface-water productivity is fairly high so thatthere is a mid-water oxygen minimum. Shelf sedimentswould consist largely of terrigenous clastic material andorganic matter with some contributions from autoch-thonous marine organic matter and biogenic calcareousand siliceous debris. Sediment that accumulated wherethe continental margin intersected the oxygen minimummight be organic-carbon-rich, black, and laminated ifthe dissolved oxygen concentration was near zero through-out a sufficient thickness of the water column. The com-position 01 any redeposited sediment would depend onlocal sources. Slope sediments along arid continentalmargins that are characterized by coastal upwellingwould be primarily fine grained and rich in carbonate orbiogenic silica (e.g., Sites 398, 400, 402, and 370). Sedi-ment accumulating in well-oxygenated bottom-waterenvironments would be mostly red clay. Occasionally,there would be green clay interbeds that represent peri-ods of accumulation of slightly higher concentrations oforganic matter, high enough to consume all of the dis-solved oxygen within the sediment and to impart a greencolor (mainly from reduced iron) to the sediment, butnot enough to have a marked effect on the benthicfauna. The result would be interbedded red and greenclay units that are both intensely bioturbated.

In the situation shown diagrammatically in Figure12B, there is higher surface-water productivity in re-sponse to a greater supply of nutrients from more in-tense upwelling. In addition, there may be an increase inthe influx of terrestrial organic matter as a result of arise in sea level and (or) an increase in amount of ter-restrial vegetation. This increased supply of organicmatter from marine and/or terrestrial sources would in-crease the oxygen demand in the water column and in

ShelfGreen sediments

(organic carbon 1—2%)

Black sediments(organic carbon >2%)

Dark green toblack sediments(organic carbon >2%)

Green orblack sediment(depending on amountof organic carbon)

Figure 12. Illustrations showing how influx of terrigenous organicmatter, surface-water organic productivity, upwelling, and turbid-ity currents may contribute to accumulation and redeposition ofsediments containing varying amounts of organic carbon and,therefore, ranging in color from red to black. A. Example of a pe-riod of moderate supply of organic detritus from autochthonous-marine and terrigenous sources and moderate upwelling resultingin accumulation of green sediment in shelf areas, sediment withhigh organic-carbon concentrations, perhaps black in color, withinthe O2-minimum zone, and dominantly red clays in the deep basin.Such conditions might be representative of the Late Cretaceous(e.g., Santonian-Coniacian) in the Atlantic Ocean. B. Example ofa period of rapid rate of supply of organic detritus and strongerupwelling resulting in an expanded and more intense O2-minimumzone that may include much of the water mass and allow organic-carbon-rich sediment to accumulate over a much greater area andprovide a greater source area for supply of organic-carbon-richsediment to the deep basin by downslope transport. Such condi-tions might be representative of periods of accumulation of organ-ic-carbon-rich strata in the Atlantic Ocean (Aptian-Albian andCenomanian-Turonian).

the sediments which would cause an expansion and in-tensification of the oxygen minimum zone. As a result,sediment accumulating along the continental marginwould be more reduced and contain higher concentra-

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tions of organic matter, and this sediment would cover amuch greater area. Because there would be more sedi-ment containing higher concentrations of organic mat-ter available for down-slope transport into deeper-waterenvironments, the basin sediment also would be morereduced and contain higher concentrations of organicmatter. In extreme cases, when bottom waters were deli-cately poised at a low oxygen level, oxygen-deficientconditions may have extended into deeper waters, butmost of the time the bottom-water mass was sufficientlyoxic to support at least a benthic epifauna. The resultingsediment facies would consist of dark green to black,organic-carbon-rich sediment deposited in continental-margin settings, and interbedded green and black, or-ganic-carbon-rich basinal sediment. Sediment depositedin either the continental margin or basinal environmentsat times of extreme oxygen deficiency might be lami-nated, indicating complete elimination of the benthicfauna.

We believe that a complex interrelationship of pro-cesses of the sort outlined above is the most appropriatemodel for black-shale accumulation at Site 530 as wellas in many other parts of the Atlantic during the Earlyto middle Cretaceous. The most important process forany one area or bed is difficult if not impossible toresolve, and the extent to which different processes op-erated most likely varied at different times and places.

The diagenetic redox cycles of variable duration(20,000 to 140,000 years on average) and their colormanifestations are superimposed on a longer period cy-clicity that has resulted in three periods of maximumaccumulation of organic carbon in the Cretaceous At-lantic Ocean—the Neocomian, the Aptian-Albian, andthe Cenomanian-Turonian. Unfortunately, we know lit-tle about the periodicity of the main variables involved;these variables are as follows:

1) Variation in bottom-water oxygenation (influencedby climatic, Oceanographic, and paleogeographic fac-tors). Low oxygen levels are inferred for much of theEarly to middle Cretaceous, and periods of lowest ox-ygen may correspond to the three periods of maximumaccumulation of organic carbon.

2) Variation in the extent and intensity of a mid-water oxygen minimum (influenced by climatic and Oce-anographic effects on surface productivity and verticalmixing). An expanded oxygen minimum is inferred forthe periods of maximum accumulation of organic car-bon, but probably there were shorter-period variationsin the extent and intensity of the oxygen minimum aswell.

3) Variation in the supply of organic matter to sedi-ments (influenced by surface productivity, terrestrialsupply, and redeposition processes). The supply of or-ganic matter probably had a complex and irregular peri-odicity, in part following that of surface productivity.

The preservation of organic carbon in the sedimentand the development of black, green, or red colorationdepends therefore on the interrelationships of thesethree main variables in a way similar to that suggestedby Tucholke and Vogt (1979). It is not surprising, there-fore, that the cyclicity, organic-carbon content, and

depositional characteristics of middle Cretaceous blackshales in the Atlantic are so variable.

CONCLUSIONS

1. Most of the Cretaceous section in the AtlanticOcean is characterized by an interbedding of more-re-duced and less-reduced lithologies, most commonly redand green or green and black mudstones and marl-stones. The black lithologies usually contain more than2% organic carbon and commonly contain more than10% organic carbon.

2. Most of the organic matter in organic-carbon-richCretaceous strata in the western and northern North At-lantic Ocean is of terrestrial origin, whereas most of theorganic matter in equivalent strata in the southeasternNorth Atlantic and South Atlantic is of marine origin.Mixtures of terrestrial and marine organic matter, how-ever, are common in both the North and South Atlantic.

3. Periodicities of the redox cycles usually range be-tween 20,000 and 100,000 years/cycle and average about40,000 to 50,000 years/cycle. These cycles may be re-sponses to climatic forcing, similar to Cenozoic climate-induced depositional cycles, but they may also reflect, inpart, periodic sediment redeposition events.

4. Certain trace elements, especially Cd, Cu, Zn,Mo, V, Ni, and Cr, are enriched in organic-carbon-richstrata. This trace-element enrichment may be the resultof concentration by clays and organic matter and copre-cipitation with sulfide minerals under reducing con-ditions. The enrichments may reflect deposition of sedi-ment under anoxic conditions in the deep sea, or may beinherited from original deposition under low-oxygenconditions upslope and subsequent redeposition to ox-ygenated deeper-water environments.

5. Compactional dewatering of some black-shale se-quences resulted in upward diffusion of some reducedmetal ions after deposition of high concentrations oforganic matter and resulting anoxic conditions withinthe sediment. Slowly deposited, oxidized claystonesoverlying some black-shale sequences are enriched insome metals, particularly Fe, Mn, Zn, and Cu, relativeto normal pelagic clays, and this enrichment may be theresult of upward migration of metals from the underly-ing black shale.

6. The origin of interbedded more- and less-reducedlithologies containing varying amounts of organic mat-ter and varying amounts of pelagic, hemipelagic, andterrigenous sediment is complex and probably is not theresult of any one process. The main variables are therate of influx of terrigenous sediment and organic mat-ter, and surface-water productivity. Variations in rateof supply of these two sources of organic matter fromshallow-water areas to deeper, basinal sites of accumu-lation, produced cyclic interbeds of more-reduced andless-reduced lithologies. During several periods of time(e.g., Neocomian, Aptian-Albian, and Cenomanian-Turonian), a coincidence of several climatic and Oce-anographic factors apparently resulted in maximumproduction and preservation of organic matter in deep-sea environments that otherwise were fairly well ox-ygenated and were preserving lesser amounts of organic

840


matter. Bottom water anoxia may have occurred duringthese periods of maximum accumulation of organic-carbon-rich strata, but the commonly observed cyclicinterbeds of green and black strata are in part also theresult of diagenetic fluctuations in redox potential in thesediment in response to fluctuation in amount of organ-ic debris being deposited, regardless of source. At times,much of the organic debris was transported to basinalsites from continental-margin settings within oxygen-minimum zones by turbidity currents; at other times,the organic debris was largely of pelagic origin.

ACKNOWLEDGMENTS

We would like to thank Joseph Hatch, Douglas Kirkland, PhilipMeyers, Harry Tourtelot, and Douglas Waples for helpful reviews ofthe manuscript. Louise Reif was particularly helpful in massagingseveral long drafts of the manuscript through the word processor.

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