Version of Record: https://www.sciencedirect.com/science/article/pii/S0016703720301605
Manuscript_a65d6a875e845ac794b79d1224525a5a
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Beyond redox: control of trace-metal enrichment in anoxic marine facies by watermass
chemistry and sedimentation rate
Jiangsi Liua,b,c,*, Thomas J. Algeob,c,d,*
a
State Key Laboratory of Marine Environmental Science, Xiamen University, Xiamen 361102,
China
b State Key Laboratory of Biogeology and Environmental Geology and School of Earth Sciences,
China University of Geosciences, Wuhan 430074, China
c Department of Geology, University of Cincinnati, Cincinnati, OH 45221-0013, U.S.A.
d State Key Laboratory of Geological Processes and Mineral Resources, China University of
Geosciences, Wuhan 430074, China
*Correspondence: jiangsiliu@xmu.edu.cn (JL); thomas.algeo@uc.edu (TJA)
Abstract
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Although redox conditions are the dominant control on authigenic enrichment of trace
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metals in marine sediments, other factors may be important within environments having
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relatively uniform redox characteristics, such as some anoxic silled basins. Notably, watermass
21
chemistry (specifically, aqueous trace-metal concentrations) and sedimentation rate can also
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influence the authigenic accumulation of redox-sensitive trace metals such as molybdenum (Mo)
23
and uranium (U) in the sediment, although these effects have received less attention than redox
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controls to date. Here, we (1) utilize a diffusion-reaction model to evaluate the effects of
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variations in watermass chemistry and sedimentation rate on authigenic trace-metal enrichment,
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and (2) present case studies of Mo and U enrichment in modern Black Sea sediments and North
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American Devonian-Carboniferous boundary (DCB) black shales that illustrate these influences.
28
In both case studies, redox conditions were assessed using non-trace-metal-based proxies (i.e., C-
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S-Fe, FeT/Al, and Corg:P). Stations 6 and 7 of the modern Black Sea, at water depths of 380 and
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1176 m, respectively, exhibit marked differences in authigenic Mo and U enrichment: median
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Mo/TOC is 13.2 at Station 6 (range 11.5-14.8) versus 5.7 at Station 7 (range 3.7-7.6), and
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median U/TOC is 2.6 at Station 6 (range 1.5-3.0) versus 1.3 at Station 7 (range 0.7-1.9) (note:
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units are ppm/% or 10−4, and ranges are 16th-84th percentiles). Given the nearly identical redox
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conditions and sedimentation rates at these two sites, the most likely cause of the >2×
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enrichment of Mo and U at Station 6 relative to Station 7 is differences in aqueous Mo and U
36
concentrations, which decline steeply through the upper part of the Black Sea water column,
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demonstrating the influence of watermass chemistry on patterns of authigenic trace-metal
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enrichment in the sediment. For the DCB black shales, the median Mo/TOC is 22.2 in the Lower
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and Upper Bakken (range 15.8-27.0), 28.8 in the Sunbury (range 19.2-37.2), and 14.3 in the
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Cleveland (range 9.3-22.5), and the median U/TOC is 6.2 in the Lower and Upper Bakken (range
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3.5-9.6), 2.6 in the Sunbury (range 1.7-3.4), and 1.2 in the Cleveland (range 0.7-1.7). Differences
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in Mo and U concentrations between these formations show no relationship to inferred aqueous
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trace-metal concentrations, Mn-Fe particulate shuttles, or paleoenvironmental redox conditions,
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but they broadly correlate with variation in sedimentation rates, providing evidence that
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sedimentation rates can measurably influence the degree of authigenic trace-metal enrichment of
46
marine sediments.
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Keywords: molybdenum; uranium; Black Sea; Devonian; Bakken Formation; Cleveland Shale
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1. Introduction
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Redox-sensitive trace metals (RSTMs, shortened herein to TMs), especially molybdenum
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(Mo) and uranium (U), have been extensively utilized to reconstruct redox conditions in
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paleomarine systems (Algeo and Maynard, 2004, 2008; Och and Shields-Zhou, 2012; Scholz et
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al., 2017; Costa et al., 2018). Under oxic conditions, molybdenum (Mo(VI)) is present mainly in
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the form of molybdate (MoO42-), and uranium (U(VI)) as uranium carbonate complexes
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(UO2(CO3)34-) (Tribovillard et al., 2006). Under anoxic conditions, molybdate is transformed to
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thiomolybdate ( MoOxS24−−x ), and U(VI) is reduced to U(IV) as UO2+ or uranous oxide complexes.
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Particle-reactive thiomolybdates and insoluble U(IV) are rapidly removed to the sediment,
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leading to potentially large authigenic Mo and U enrichments (e.g., Algeo and Maynard, 2004).
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In addition to first-order control by local redox conditions, sedimentary enrichment of TMs
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is influenced by a number of “second-order” factors. One such factor is watermass chemistry,
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specifically, the aqueous concentrations of trace metals. Most TMs have long residence times in
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seawater (Tribovillard et al., 2006; Algeo and Maynard, 2008), so their concentrations in
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unrestricted marine systems are nearly invariant, except at long time scales. However, in
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restricted anoxic marine basins, limited resupply of TMs can result in their aqueous
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concentrations being strongly modified by authigenic uptake in the sediment (Algeo and Lyons,
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2006; Algeo and Rowe, 2012). Changes in the degree of restriction of such basins (due to
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eustatic or tectonic processes) can lead to changes in TM resupply rates or authigenic uptake
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rates, rapidly modifying the TM chemistry of the watermass, as documented in paleomarine
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systems of Devonian (Algeo et al., 2007), Jurassic (McArthur et al., 2008), and Cretaceous age
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(Hetzel et al., 2011), among others.
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Another potentially important second-order influence on sediment TM concentrations is
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sedimentation rate, which controls the dilution of subordinate sediment fractions such as organic
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matter and the downward diffusion of trace metals and reactants such as sulfate (Sommerfield
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and Nittrouer, 1999; Jobe et al., 2012; Mintz et al., 2017). Generally, higher sedimentation rates
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lead to lower TM concentrations owing to a combination of sediment dilution and limitation of
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diffusion, but this effect should be less strong for U, the enrichment of which is linked to an
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aqueous reactant (sulfate) that is not affected by dilution, compared to Mo, whose enrichment is
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linked to a solid phase (organic matter) that can be diluted. Although sedimentation rate effects
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on TM enrichment have been considered tangentially in a few studies (Morford et al., 2007;
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Hardisty et al., 2018), this issue has not received systematic examination to date.
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In addition to watermass chemistry and sedimentation rate, other potential second-order
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influences on authigenic trace-metal enrichment levels have been identified, including the size of
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the oceanic TM reservoir (Scott et al., 2008; Partin et al., 2013), operation of particulate shuttles
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(Algeo and Tribovillard, 2009; Dellwig et al., 2010), trace-metal concentration through
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watermass recycling processes (Algeo and Herrmann, 2018), biological sequestration (Little et
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al., 2015), and the availability of organic substrates (Jin et al., 2018). We will not consider these
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influences further in the present study, although some of them are potentially amenable to
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investigation using the tools developed herein.
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The present study is designed specifically to test the influences of watermass chemistry and
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sedimentation rate on authigenic enrichment of TMs in anoxic marine facies. We focus on
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molybdenum (Mo) and uranium (U) owing to their general importance in paleoceanographic
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research (Section 2). We then model the effects of aqueous trace-metal concentrations and
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sedimentation rates on authigenic Mo and U enrichment using a diffusion-reaction model, thus
95
defining theoretically expected relationships (Section 3). Finally, we test for these effects based
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on case studies of authigenic Mo and U enrichment in modern Black Sea sediments (Section 4)
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and Devonian-Carboniferous boundary black shales of the North American Seaway (Section 5).
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This study thus provides new insights regarding non-redox controls on the enrichment of TMs in
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anoxic marine facies.
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2. Background
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2.1. Mo and U aqueous chemistry
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Mo and U are particularly useful in paleoceanographic research because they are
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dominantly of authigenic origin in sediments, owing to their low concentrations in upper
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continental crust, river waters, and marine plankton (Algeo and Tribovillard, 2009; Tribovillard
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et al., 2012). They have relatively high concentrations (Mo = 110 nM; U = 13 nM) and long
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residence times (Mo ~440 kyr; U ~500 kyr; Dunk et al., 2002; Miller et al., 2011) in modern
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seawater and, thus, are well-mixed in the global ocean. These factors tend to result in
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enrichments of Mo and U in anoxic marine facies that significantly exceed the low background
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levels present in continental crustal rocks and their weathering products (average ~1.5 ppm Mo,
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~2.8 ppm U; McLennan, 2001).
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Authigenic Mo uptake from aqueous sources is strongly promoted by euxinic conditions.
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Conversion of molybdate (MoO42-) to particle-reactive thiomolybdates ( MoOxS24−−x , x = 0 to 3)
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requires the presence of at least small amounts of H2S (Helz et al., 1996; Erickson and Helz,
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2000). This process may be catalyzed by clay-mineral surfaces under mildly acidic conditions
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(Vorlicek and Helz, 2002) as well as by aqueous intermediate S species (Vorlicek et al., 2004).
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Euxinic micro-environments, e.g., associated with decaying organic matter, can lead to local Mo
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enrichments in Fe-sulfides (Tribovillard et al., 2008), possibly through an intermediate Fe-Mo-S
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colloidal phase (Helz et al., 2011; Vorlicek et al., 2018). This process has been argued to account
120
for strong Mo-TOC correlations, being regarded as the incidental outcome of microbial sulfate
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reduction (MSR)-driven organic matter decay rather than a product of organic-Mo complexation
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(Helz and Vorlicek, 2019). However, the ubiquity of strong Mo-TOC and concurrent weak Mo-S
123
correlations (e.g., Algeo et al., 2007) as well as a strong tendency toward organic-Mo ligand
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formation in both marine sediments (Malcolm, 1985) and soils (Wichard et al., 2009) suggest
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that organic matter is generally the dominant host phase of Mo. However, in depositional
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systems with active redox cycling of Fe and Mn through the water column, adsorption of Mo
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onto sinking Mn-Fe-oxyhydroxide particulates can accelerate its transfer to the sediment, a
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process known as a “particulate shuttle” (Algeo and Tribovillard, 2009; Dellwig et al., 2010).
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Authigenic U uptake is also favored by euxinic conditions (Wanty and Goldhaber, 1992;
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McManus et al., 2005) but its enrichment begins at a higher redox threshold (i.e., at the Fe(III)-
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to-Fe(II) transition within the suboxic zone) than that of Mo (Anderson et al., 1989; Zheng et al.,
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2002). Consequently, enrichment of authigenic U typically begins before that of authigenic Mo
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(Algeo and Tribovillard, 2009). Multiple U uptake pathways exist, although organic-metal
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ligands may be the most important quantitatively. An experimental study showed that U(VI)
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uptake is favored by mildly acidic conditions (pH = 4-6), binding by certain functional groups,
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and the presence of humid acids, which can both complex with U and reduce the competitive
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effects of Ca2+ and other ionic species (Liu et al., 2016). Adsorption onto clay minerals is
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possible, especially those with high cation exchange capacities (CEC) such as montmorillonite,
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although organic-clay complexes (“clay biopolymers”) are more effective at U uptake, partly due
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to their stronger negative surface charges (Olivelli et al., 2013). In soils, U uptake is influenced
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by the types and amounts of organic matter, clay minerals, and oxides, with dead organic matter
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being more effective than either living biomass or abiotic components (Choi and Park, 2005).
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Although TM enrichment levels are frequently expressed as Al-normalized ‘enrichment
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factors’ (EFs) (Tribovillard et al., 2006), this protocol is not optimal for evaluation of controls on
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Mo and U enrichment. The amount of authigenic Mo and U in marine sediments is strongly
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dependent on total organic carbon (TOC) content because organic matter represents the main
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substrate for Mo and U uptake (e.g., Lüning and Kolonic, 2003) and, thus, TOC content
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ultimately limits authigenic TM enrichment (Algeo and Lyons, 2006; Algeo and Rowe, 2012).
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For this reason, we will report Mo and U enrichments on a TOC-normalized basis (i.e., Mo/TOC
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and U/TOC) in case studies of modern Black Sea sediments and Devonian-Carboniferous
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boundary (DCB) black shales (Sections 4 and 5), and we will make comparisons to our
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diffusion-reaction model calculations (Section 3) on the basis of burial fluxes of authigenic Mo
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and U (i.e., in units of nmol cm-2 yr-1).
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2.2. Evaluation of redox conditions
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Prior to analysis of second-order influences on authigenic TM enrichment, a necessary
157
preliminary step is to evaluate the first-order (dominant) influence on their uptake by the
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sediment: redox conditions. In order to recognize second-order effects, one must either (1)
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compare units deposited under nearly identical redox conditions, or (2) if comparing units of
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dissimilar redox conditions, demonstrate that the second-order control of interest has had such a
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large influence that it has overridden primary redox effects. For the purpose of assessing redox
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conditions in case studies of modern Black Sea and DCB sediments (Sections 4 and 5), we will
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make use of three non-trace-metal-based redox proxies: (1) C-S-Fe relationships, (2) FeT/Al, and
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(3) Corg:P ratios. We eschew use of trace-metal-based proxies because they would conflict with
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our goal of evaluating non-redox controls on authigenic TM enrichment.
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Carbon-sulfur-iron (C-S-Fe) relationships have been widely used to assess redox conditions
167
in marine systems (Dean and Arthur, 1989; Arthur and Sageman, 1994; Montero-Serrano et al.,
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2010; Georgiev et al., 2015; Moradi et al., 2016). In C-S-Fe ternary diagrams, a S/TOC ratio of 0.36
169
is indicative of oxic conditions, with higher or lower values indicative of euxinic conditions that are
170
C- and Fe-limited, respectively (Berner and Raiswell, 1983). Additional redox information is
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yielded by the line representing S/Fe ratios, with values of ~0.42, ~0.75, and 1.15 (the last being
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equivalent to stoichiometric pyrite) representing dysoxic, anoxic, and strongly euxinic conditions,
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respectively (Berner, 1984; Dean and Arthur, 1989; Arthur and Sageman, 1994; El-Shafeiy et al.,
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2016).
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FeT/Al ratios (units: wt. %/wt. %) are a useful redox proxy because average upper continental
176
crust (UCC) has a value of 0.44, and higher values are indicative of addition of excess Fe
177
(generally as pyrite) to the sediment under reducing conditions (Raiswell and Canfield, 1998;
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Poulton et al., 2010; Clarkson et al., 2014). Normalization of total Fe to Al is not sensitive to
179
dilution effects, making it a robust proxy over a wide range of sedimentation rates (Raiswell and
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Canfield, 1998; Raiswell et al., 2008). The FeT/Al ratio varies from 0.42 to 0.56 from both oxic
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and euxinic sediments (Lyons et al., 2003), thus higher FeT/Al values possibly represents Fe-
182
enriched anoxic condition (Raiswell et al., 2008).
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Corg:P molar ratios (units: mol/mol) are a useful redox proxy because (1) most organic
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carbon and phosphorus in marine sediments come from algal biomass (e.g., Anderson et al.,
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2001), (2) eukaryotic algae have a relatively uniform initial Corg:P ratio, approximated by the
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‘Redfield ratio’ of 106:1 (Redfield, 1958; Klausmeier et al., 2004), and (3) the fate of
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remineralized organic carbon and phosphorus differs as a function of redox conditions (Ingall et
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al., 1993; Algeo and Ingall, 2007). P-bearing biomolecules (including nucleic acids and lipids)
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tend to break down more rapidly than many other biomolecules, and the organic fraction of the
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sediment becomes progressively more P-depleted through burial decay (Ingall and Jahnke, 1994;
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Joshi et al., 2015). Under anoxic conditions, the remineralized P is lost from the sediment
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through diffusion, leading to a progressive increase in bulk sediment Corg:P ratios with time
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(Slomp et al., 1996; Ingall et al., 2005; Algeo and Ingall, 2007). On the other hand, oxic
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conditions favor retention of remineralized P in the sediment through adsorption onto Fe-
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oxyhydroxides, even as remineralized carbon diffuses out of the sediment. Consequently, Corg:P
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values of approximately <50, 50-100, and >100 in ancient marine sediments are indicative of
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oxic, suboxic, and anoxic depositional conditions, respectively (Algeo and Ingall, 2007).
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2.3. Evaluation of watermass chemistry
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For modern marine systems such as the Black Sea, watermass chemistry is a known
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parameter, with existing observational data permitting evaluation of short-term secular and
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spatial variation in aqueous TM concentrations (e.g., Pohl and Hennings, 2005; Ludwig et al.,
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2009). For ancient marine systems, this is not the case, and an assessment of watermass
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chemistry, specifically in relation to the trace metals of interest, must be undertaken. Deep-time
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oceans may have had seawater TM concentrations that deviated markedly from present-day
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seawater concentrations (e.g., Algeo, 2004; Scott et al., 2008; Hetzel et al., 2011; Partin et al.,
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2013). Furthermore restricted marginal-marine basins, both modern and ancient, are prone to
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development of unique watermass chemistries, as demonstrated by a survey of modern systems
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(Algeo and Maynard, 2008). Thus, both secular and spatial variations in seawater TM
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concentrations are common and have the potential to influence the degree of authigenic TM
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enrichment of sediments.
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Evaluation of aqueous TM concentrations in deep-time systems is a relatively new field
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(e.g., Algeo et al., 2007; McArthur et al., 2008; Scott et al., 2008; Partin et al., 2013). The
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concentrations of Mo and U in ancient seawater can potentially be estimated from authigenic
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enrichment levels in sediments that are normalized to TOC content (Section 2.1). If multiple
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samples are collinear in a TM-vs-TOC crossplot, they can used to define an m(TM/TOC) value
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(where m indicates a slope) that is thought to be proportional to the aqueous concentration of the
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TM of interest (Algeo and Lyons, 2006; Algeo and Rowe, 2012; note: all m values throughout
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this study are in units of ppm/%, or 10‒4). For example, in the modern, m(Mo/TOC) ranges from
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~45 for unrestricted marginal-marine basins, in which aqueous [Mo] is close to that of global
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seawater (e.g., >90% for Saanich Inlet), to ~4.5 for highly restricted basins, in which aqueous
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[Mo] is strongly depleted relative to that of global seawater (~3 % for the Black Sea; Algeo and
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Lyons, 2006). A similar range of m(Mo/TOC) values (~3-65) was reported for a large group of
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North American Upper Devonian-Lower Carboniferous black shales (n = 55; Algeo et al., 2007),
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suggesting that this proxy has general validity in deep-time marine systems. Similar analyses
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have not been undertaken for U or other TMs in modern marine systems to date (although some
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applications to ancient systems have been made, e.g., Hetzel et al., 2011; Partin et al., 2013), but
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the same principles should apply to other TMs as to Mo.
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3. Modeling the influences of watermass chemistry and sedimentation rates on authigenic
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trace-metal enrichment
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The authigenic uptake of trace elements such as Mo and U below the sediment/water
233
interface (SWI) can be modeled as a function of three sets of processes: (1) diffusion and bio-
234
irrigation, which modulate the downward flux of dissolved trace elements into the sediment, (2)
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deposition and compaction, which determine the depth of the site of trace-element uptake and the
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permeability of the overlying sediment column, and (3) reaction processes, which control the
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transfer of trace elements from porewater to the sediment (Fig. 1). (Note that our model
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explicitly does not address trace-metal transfer to the sediment via adsorption on sinking
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particulates.)
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In anoxic marine systems, bio-irrigation (see Morford et al., 2007, for modeling details) is
241
effectively reduced to zero owing to exclusion of benthic animals, so diffusion becomes the
242
dominant control on downward elemental fluxes. Diffusion in a one-dimensional system is given
243
by Fick’s Second Law of Diffusion (Fick, 1855):
∂c
∂ 2c
= Dsed 2
∂t
∂x
244
245
(1)
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where Dsed is the diffusion coefficient, c is the concentration of a given element, t is time, and x
247
is the distance below the SWI. Diffusion coefficients for sediments (Dsed) can be calculated as
248
(McDuff and Ellis, 1979; Morford et al., 2009):
Dsed =
249
DSW
θ2
250
(2)
251
where Dsw is the diffusion coefficient for a trace element in solution, and θ2 is the tortuosity of
252
the diffusion pathway. Tortuosity (θ2) has been empirically linked to sediment porosity as follows
253
(Boudreau, 1996):
θ 2 = 1 − ln(φ 2 )
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(3)
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where ϕ is porosity. Sediment porosity is a function of burial depth, which we modeled based on
257
profiles from two recent Black Sea cores (Cores CH12 and CH18) (Opreanu, 2003; his table 1).
258
The equation of the best-fit curve (Fig. 2) is:
φ = 0.9 × exp( −7.5 × 10 −4 × x )
259
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(4)
261
where x is depth in units of cm.
262
The processes of deposition and compaction can be expressed as:
263
∂c
∂c
= −w
∂t
∂x
264
(5)
265
where the derivative w is defined as the linear sedimentation rate (i.e., a net sediment
266
accumulation rate at depth, w = dx/dt), with w > 0 representing deposition and w < 0 representing
267
compaction. In areas of active sediment accumulation, the net value of w is positive; however, it
268
is a depth-dependent parameter, declining as compaction increases with depth below the SWI.
269
The direct physical effect of sedimentation is a simple vertical displacement of the sediment-
270
water interface. However, sedimentation also exerts an influence on reaction rates, if the main
271
reactant is a solid phase that can be variably diluted by siliciclastic sediment influx (as, for
272
example, organic matter; see Eq. 8b).
The effect of chemical reaction processes with a first-order reaction rate dependent on solute
273
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concentration can be written as (McNaught and Wilkinson, 1997):
∂c
= −kG
∂t
275
276
(6)
277
where k is the reaction rate constant, and G is the concentration of the main reactant driving the
278
reaction. If the concentration of the main reactant changes as a function of depth below the SWI,
279
then the reactant concentration equation can be written as (Berner, 1964):
∂G
∂G
= −w
− kG
∂t
∂x
280
281
(7)
282
For the steady state condition, ∂G/∂t = 0. In this case, reorganization of Eq. (7) yields:
k
G = G 0 exp[ −( ) x ]
w
283
284
(8a)
285
where G0 is the initial concentration of the main reactant. The form of this relationship depends
286
on whether the key reactant is aqueous or a solid-phase component. If the key reactant is an
287
aqueous phase (e.g., sulfate), then its initial concentration (G0) is not affected by sedimentation
288
rate because the reactant is present only in the porewater. However, if the key reactant is a solid
289
phase (e.g., organic matter), then its initial concentration is sedimentation rate-dependent because
290
higher sedimentation rates will lead to its dilution in the sediment. In this case, the preceding
291
equation must be modified to:
G=
292
G0
k
exp[ −( ) x ]
w
w
293
(8b)
294
where the term G0/w accounts for dilution effects associated with higher sedimentation rates.
295
The key reaction process for uptake of aqueous Mo by the sediment is the conversion of
296
unreactive molybdate (MoO42-) to reactive thiomolybdate ( MoOxS24−−x , x = 0 to 3), which
297
requires the presence of free H2S in solution (Helz et al., 1996). An H2S concentration of >11
298
µM is needed to completely convert molybdate to thiomolybdate, which enhances the particle-
299
reactivity of Mo allowing rapid adsorption onto sedimentary particles (Erickson and Helz, 2000).
300
The key reaction process for uptake of aqueous U by the sediment is the reduction of U(VI) to
301
U(IV) (Tribovillard et al., 2006). In its oxidized state, U is present in solution mainly as soluble
302
carbonate complexes ( UO 2 (CO 3 ) 34 − ), whereas in its reduced state, U readily forms solid-phase
303
oxides such as UO2, U3O7, or U3O8. A recent study proposed that “reduction hotspots”, e.g.,
304
aggregates of cells, organic matter, FeS, and aluminosilicates, facilitated the transformation of
305
U(VI) to U(IV) under sulfate-reducing conditions followed by adsorption of U(IV) to organic
306
matter and clay particles (Bone et al., 2017). Thus, the main reactants driving authigenic uptake
307
of these trace metals are H2S (aqueous phase) for Mo and organic matter (solid phase) for U,
308
necessitating use of Eqs. (8a) and (8b), respectively, for calculation of authigenic Mo and U
309
fluxes.
310
311
The preceding equations can be combined to yield a full expression for solute concentration
changes with time (cf., Berner, 1964; Hardisty et al., 2018):
312
∂c
k
∂c
DSW
∂ 2c
=
×
− w − kG0 exp[−( ) x]
2
2
∂t 1 − ln(φ ) ∂x
∂x
w
(9a)
313
∂c
DSW
∂ 2c
∂c
G
k
=
×
− w − k 0 exp[−( ) x]
2
2
∂t 1 − ln(φ ) ∂x
∂x
w
w
(9b)
314
where Eq. (9a) is for TMs linked to aqueous-phase reactants (e.g., Mo), and Eq. (9b) is for TMs
315
linked to sold-phase reactants (e.g., U). For the case of steady state, ∂c/∂t = 0. The general
316
solution of this equation is:
317
c = J exp[(
318
c = J exp[(
w × (1 − ln(φ 2 ))
w 2G 0
k
) x] + (
) exp[ −( ) x] + C1
D
DSW
w
SW
w2 + (
)k
1 − ln(φ 2 )
w × (1 − ln(φ 2 ))
wG 0
k
) x] + (
) exp[ −( ) x ] + C1
DSW
DSW
w
w2 + (
)k
2
1 − ln(φ )
(10a)
(10b)
319
where C1 is a constant, and J is an arbitrary constant of integration. Eq. (10a) is for TMs linked
320
to aqueous-phase reactants (e.g., Mo), and Eq. (10b) is for TMs linked to sold-phase reactants
321
(e.g., U). These equations are solved by setting the boundary conditions c (0 ) = c0 and c(∞) = 0
322
and by setting J to zero (cf., Berner, 1964), yielding:
323
c=
324
c=
w 2G 0
k
exp[ −( ) x] + C1
Dsw
w
w2 +
k
2
1 − ln(φ )
wG 0
k
exp[−( ) x] + C1
Dsw
w
w2 +
k
2
1 − ln(φ )
(11a)
(11b)
325
where Eq. (11a) is for TMs linked to aqueous-phase reactants (e.g., Mo), and Eq. (11b) is for
326
TMs linked to sold-phase reactants (e.g., U). In this manner, the TM concentration of the
327
porewater can be simplified to:
328
k
c = c0 exp[ − ( ) x ]
w
329
To estimate reaction rate constants (k) for Mo and U, we made use of three porewater
330
concentration profiles each for Mo and U from the uppermost 30 cm of the sediment column on
331
the Peru Margin (cores MUC 19, 39, and 53), Long Island Sound, New York (FOAM), and
332
Hingham Bay, Massachusetts (Cores 1 and 2) (Morford et al., 2007; Scholz et al., 2011; Hardisty
333
et al., 2018). Porewater Mo and U concentrations are shown in Figure 3, and the equations of the
334
best-fit curves are given in Table 1. kMo was calculated as 0.12 yr-1, 0.014 yr-1 and 0.0038 yr-1 at
335
sites Core 1, FOAM and MUC19, respectively, and kU was calculated as 0.12 yr-1, 0.0052 yr-1,
336
and 0.0087 yr-1 for sites Core 2, MUC53, and MUC39, respectively. The variability of k is due to
337
its dependence on seawater temperatures, the nature of organic substrates, and the concentrations
338
of reaction catalysts such as SO42‒ (Erickson and Helz, 2000; Helz et al., 2011), which can vary
339
greatly in different marine areas. In our subsequent modeling of modern and ancient marine
340
systems, we varied k in order to determine the best match to our datasets.
341
342
(12)
TM fluxes in the porewater were calculated per Morford et al. (2009) with substitution of
the porosity parameter ( φ ) into Equation (4):
Fp = −φDsed
343
dc 0.9 Dsed kc0
k
=
exp[( − − 7.5 × 10 − 4 ) x ]
dx
w
w
344
(13)
345
Increasing TM fluxes to the sediment are related to decreasing TM concentrations in the
346
porewater (Morford et al., 2007). The mass accumulation rate of authigenic TMs in the sediment
347
(Fs) can thus be calculated from the sink flux of TMs from the porewater (Fp):
Fs =
348
0.9 Dsed kc0 0.9 Dsed kc0
k
−
exp[( − − 7.5 × 10 − 4 ) x ]
w
w
w
349
(14)
350
TM concentrations in the sediment (Cauth) thus correspond to the time-integrated sink flux of
351
TMs from the porewater, which can be calculated as follows (Morford et al., 2007; see Item S1
352
in the Supplemental Materials for details of equation integration):
Cauth =
353
Fs
ρ s (1 − φ ) w
(15)
354
where Cauth is in units of nmol g-1, Fs is the accumulation rate in the sediment (units: nmol cm-2
355
yr-1), w is the sedimentation rate, and ps is the dry bulk-sediment density (here taken as 1.88 g
356
cm-3; Böning et al., 2004). Thus, authigenic TM concentrations ([TM]auth, unit: ppm) are
357
expressed as:
[TM ]auth = C auth M TM = (
358
0.9Dsed kc0
0.9 Dsed kc0
k
−
exp[(− − 7.5 × 10 −4 ) x]) × M TM × 10 −3
2
2
w
ρ s (1 − φ ) w ρ s (1 − φ )w
359
(16)
360
where MTM is the molar mass, i.e., MMo = 95.94 g mol-1 and MU = 238.028 g mol-1.
361
In this diffusion-reaction model, TM concentrations in both porewater (Eq. 12) and
362
sediment (Eq. 16) are thus controlled by two factors: (1) watermass chemistry (i.e., c0); and (2)
363
sedimentation rate (i.e., w). In this context, we will now evaluate the roles of these factors in
364
influencing levels of authigenic Mo and U enrichment in two marine systems, the modern Black
365
Sea and the Devonian-Carboniferous boundary (DCB) North American Seaway.
366
367
4. Case study: Modern Black Sea
368
4.1. Geological background
369
The Black Sea, the largest anoxic marine basin in the modern world, has an area of 423,000
370
km2 with a maximum depth of 2212 m and a sill depth of 33 m in the Bosporus Strait (Gunnerson
371
and Özturgut, 1974; Murray, 1991a; Jørgensen et al., 2004). The depth of its O2-H2S redoxcline
372
varies from ~50 m in the basin center to ~120-150 m around the basin margins, although its position
373
has fluctuated in both the pre-modern and modern periods (Brewer and Spencer, 1974; Murray et
374
al., 1989; Tugrul et al., 1992; Lyons et al., 1993; Anderson et al., 1994; Wilkin and Arthur, 2001).
375
The modern Black Sea exhibits variable but generally low rates of organic carbon (~1-10 g m-2 yr-1)
376
and bulk sediment accumulation (10-200 g m-2 yr-1), both increasing toward the basin margins
377
(Shimkus and Trimonis, 1974; Calvert et al., 1991; Karl and Knauer, 1991; Arthur et al., 1994).
378
Gravity and box cores collected on oceanographic expeditions, e.g., the 1969 Atlantis II (Degens
379
and Ross, 1974) and 1988 R/V Knorr cruises (Murray, 1991b) show that the surface sediments
380
consist of a 50- to 60-cm-thick, laminated, white-brown organic coccolith ooze younger than
381
~2000 yr B.P. (Unit 1), which is underlain by a ~70- to 100-cm-thick, laminated, olive-black
382
marine sapropel dating to ~2000-7160 yr B.P. (Unit 2a) (Jones and Gagnon, 1994; Major et al.,
383
2002). Intercalated within Units 1 and 2a are homogeneous, greenish-gray mud layers up to 20
384
cm thick that have been termed “turbidites”, although the lack of basal erosion draws into
385
question the exact mode of emplacement (Lyons, 1991; Arthur et al., 1994). Unit 3 consists of
386
laminated terrigenous muds that were deposited in a lacustrine setting (Mazzini et al., 2004).
387
Units older than Unit 2a are non-marine and will not be considered in this study.
388
389
4.2. Materials and methods
390
The gravity cores of Station 6 (43.68°N, 30.13°E, 380 m water depth) and Station 7
391
(43.52°N, 30.22 °E, 1176 m water depth) were collected during a research cruise of R/V Petr
392
Kottsov in September, 1977 (Fig. 4; Lüschen, 2004). The cores penetrated to total depths of 850
393
cm (Station 6) and 622 cm (Station 7) and contain Units 1, 2a, 2b, and the upper part of Unit 3.
394
The thicknesses of Units 1 and 2a are 60.0 and 70.0 cm in Station 6 and 60.0 and 85.3 cm in
395
Station 7, respectively. Units 1 and 2a both have relatively high TOC contents (Station 6 = 7.3 ±
396
4.7 %; Station 7 = 16.5 ± 6.8 %). Given the age model above, Units 1 and 2a accumulated at
397
average sedimentation rates of 0.03 and 0.014 cm yr -1, respectively, at Station 6 (mean 0.018 cm
398
yr -1), and 0.03 and 0.017 cm yr -1, respectively, at Station 7 (mean 0.020 cm yr -1).
399
All geochemical data for Stations 6 and 7 are from Lüschen (2004). The methodology used
400
in that study entailed digestion of 50 to 100 mg of sample powder in 0.3-M nitric acid, collection
401
of the supernatant in a Teflon capsule, and further digestion in HF and HClO4 at 180 °C for 6-12
402
h. Element concentrations were analyzed using a Finnigan MAT “Element” ICP-MS, with
403
precisions better than 2% for major elements and 7% for trace elements.
404
405
4.3. Results
406
Total Mo concentrations range from 58 to 97 ppm (median 65 ppm) at Station 6 (n = 35),
407
and from 34 to 110 ppm (median 45 ppm) at Station 7 (n = 50; note: ranges given as 16th-84th
408
percentiles). Total U concentrations range from 11 to 16 ppm (median 14 ppm) at Station 6 (n =
409
35), and from 11 to 17 ppm (median 15 ppm) at Station 7 (n = 50). To remove the influence of
410
terrestrial inputs, authigenic (auth) Mo and U concentrations were calculated as the difference
411
between total concentrations and estimated detrital-fraction (detr) concentrations: Xauth = Xtotal ‒
412
Al × (X/Al)detr, where X is the trace metal of interest, (X/Al)detr is the detrital metal-to-aluminum
413
ratio based on average upper continental crustal composition (McLennan, 2001, his table 5), and
414
all values are weight-based concentrations. Authigenic Mo concentrations range from 58 to 95
415
ppm (median 64 ppm) at Station 6, and from 33 to 109 ppm (median 45 ppm) at Station 7 (Table
416
2); the median values are statistically significantly different, as shown by a Mann-Whitney-
417
Wilcoxon test (p(a) < 0.01). Authigenic U concentrations range from 10 to 14 ppm at Station 6,
418
and from 10 to 16 ppm at Station 7; these ranges of values are not significantly different.
419
420
4.4. Discussion
421
4.4.1. Evaluation of redox conditions and watermass restriction
422
The modern Black Sea exhibits exclusively euxinic conditions below a shallow redoxcline
423
(at a water depth of ~100 m in the study area; Anderson et al., 1994). Below the redoxcline, O2
424
concentrations are zero and H2S concentrations rise continuously with depth. At the water depths
425
of Stations 6 and 7 (i.e., 380 m and 1176 m), mean H2S concentrations are ~140 µM and ~340
426
µM, respectively (Neretin et al., 2001). For comparison, peak H2S concentrations on the Black
427
Sea abyssal plain are typically 300-600 µM (Grasshoff, 1975; Glenn and Arthur, 1985). Although
428
the depth of the redoxcline has ranged from ~50 m to ~200 m over the 7-kyr depositional history
429
of Units 1 and 2a (Lyons et al., 1993; Wilkin and Arthur, 2001), there is no evidence that it has
430
descended as deeply as 380 m during this interval (Sinninghe Damsté et al., 1993).
431
Although the redox history of Units 1 and 2a is well-established from earlier studies (Wilkin
432
and Arthur, 2001; Neretin et al., 2004; Becker et al., 2018), we analyzed the same suite of
433
sedimentary redox proxies (C-S-Fe, FeT/Al, and Corg:P) as for the DCB study units (Section 5)
434
for comparative purposes. Black Sea Stations 6 and 7 exhibit similar median S/TOC ratios (0.18,
435
range 0.16-0.22; and 0.13, range 0.10-0.28, respectively; Fig. 5A) and Corg:P values (172, range
436
128-186; and 175, range 109-238, respectively; Fig. 5C), consistent with uniformly euxinic
437
conditions at these two sites [note: all ranges given as 16th-84th percentiles to avoid the influence
438
of extreme outliers]. Station 6 exhibits slightly lower S/Fe ratios (median 0.43, range 0.36-0.46)
439
than Station 7 (median 0.70, range 0.60-0.84) (Fig. 5A), as well as slightly lower FeT/Al ratios
440
(median 0.54, range 0.52-0.56 vs. median 0.73, range 0.53-0.83; Fig. 5B). The higher S/Fe ratios
441
at Station 7 reflect more intense sulfate reduction rates and enhanced sedimentary Fe-sulfide
442
sequestration than at Station 6 (Jørgensen et al., 2004). The higher FeT/Al ratios at Station 7
443
reflect the activity of an “Fe shuttle” in the upper slope region of the Black Sea (Wijsman et al.,
444
2001; Anderson and Raiswell, 2004; Severmann et al., 2008), which transfers Fe downslope
445
through physical transport processes (Lenstra et al., 2019). In this context, the differences in S
446
and Fe content between the two study sites are not indicative of redox differences. The present
447
redox conditions and post-7-ka redox histories of Stations 6 and 7 are similar, and, thus, redox
448
conditions are unlikely to account for the differences in authigenic Mo and U enrichment
449
between these sites.
450
Mn-Fe particulate shuttles can increase rates of accumulation of authigenic Mo (but not
451
authigenic U) in the sediment (Algeo and Tribovillard, 2009; Dellwig et al., 2010). The Black
452
Sea is characterized by active redox cycling of Mn and Fe in the vicinity of the aqueous
453
chemocline (Lewis and Landing, 1991; Scholz et al., 2013), but because of the strongly reducing
454
conditions of the deeper watermass, Mn-Fe-oxyhydroxides that precipitate at the chemocline are
455
reductively dissolved before reaching the sediment-water interface. For Mn particulates, most of
456
this dissolution takes place within the first 50 m below the chemocline (Lewis and Landing,
457
1991; Scholz et al., 2013), leading to no shuttle-related Mo enrichment of the sediment at the
458
depths of Stations 6 and 7. The relative stability of redox stratification within the Black Sea is
459
generally unfavorable for shuttle-related TM enrichments, whereas strongly fluctuating
460
chemoclines are known to promote such enrichments, as in the modern Saanich Inlet (Berrang
461
and Grill, 1974; Algeo and Tribovillard, 2009).
462
Sedimentation rates are also unlikely to account for differences in authigenic Mo and U
463
enrichment between Stations 6 and 7 given the similar sedimentation rates at these two sites
464
(~0.018 and ~0.020 cm yr
465
sedimentation rates of Unit 1 (~0.03 cm yr -1) relative to Unit 2a (~0.014-0.017 cm yr -1) are
466
linked to higher porosity (~0.74 % vs ~0.47 %), and that the bulk accumulation rates of Units 1
467
and 2a are nearly identical (~1.5 g m-2 yr-1) (Calvert et al., 1987). As illustrated by our diffusion-
468
reaction model (see Section 2), if sedimentation rates are not an important control on authigenic
469
TM enrichment, then differences in watermass chemistry are likely to be the dominant control.
-1,
respectively; see Section 4.2). We note that the higher
470
471
4.4.2. Influence of watermass chemistry on sediment Mo-U enrichment
472
Watermass chemistry (i.e., dissolved TM concentrations) can have a major influence on
473
sediment TM enrichment (Algeo and Lyons, 2006; Algeo and Rowe, 2012). In the Black Sea,
474
aqueous Mo concentrations ([Mo]aq) gradually decrease from 80 nM to 5 nM over the depth
475
range of 0 to 600 m and then are relatively stable at ~5 nM from 600 to 2250 m (Emerson and
476
Huested, 1991; Fig. 6A). Thus, measurable differences in [Mo]aq exist between Station 6 (15.3
477
nM) and Station 7 (4.8 nM). Similarly, aqueous U concentrations ([U]aq) gradually decrease from
478
8.8 nM to 5.5 nM over the depth range of 0 to 600 m and then stabilize at ~5.5 nM from 600 to
479
2250 m (Anderson et al., 1989; Colodner et al., 1995; Fig. 6A). Thus, measurable differences in
480
[U]aq also exist between Station 6 (8.3 nM) and Station 7 (5.5 nM) (Fig. 6A).
481
These depth-dependent differences in watermass chemistry are related to TM source and
482
sink fluxes in the Black Sea. The semi-enclosed Black Sea basin is connected to the
483
Mediterranean Sea via the Marmara Sea and two shallow straits. Whereas the lower-salinity
484
surface layer (~18 psu) flows outward through the Bosporus Strait, the higher-salinity deep layer
485
(~23 psu) is constantly, albeit slowly, recharged by inflow from the Mediterranean Sea (Özsoy et
486
al., 2002; Murray et al., 2007; Soulet et al., 2010). Mo and U from the Mediterranean source flux
487
is mixed more-or-less uniformly into the deeper water column, so the existence of vertical
488
concentration gradients for aqueous Mo and U are due to the depth-dependency of sink fluxes.
489
Aqueous TMs are strongly removed to the sediment by uptake at the sediment-water interface,
490
with most sequestration occurring across the broad abyssal plains of the Black Sea, at water
491
depths >2000 m. Strong water-column stratification limits vertical mixing of aqueous TMs,
492
maintaining pronounced vertical gradients (Algeo and Maynard, 2008) (Fig. 6A).
493
Authigenic TM enrichment of the sediment in Units 1 and 2a of the Black Sea can be
494
evaluated based on TOC-normalized Mo and U concentrations (cf. Algeo and Lyons, 2006;
495
Algeo and Rowe, 2012). ). The median Mo/TOC is 13.2 at Station 6 (range 11.5-14.8) versus 5.7
496
at Station 7 (range 3.7-7.6), and median U/TOC is 2.6 at Station 6 (range 1.5-3.0) versus 1.3 at
497
Station 7 (range 0.7-1.9) (note: Mo/TOC and U/TOC ratios have units of ppm/% or 10−4, and
498
ranges represent 16th-84th percentiles). The median Mo/TOC and U/TOC values of Station 6 are
499
significantly larger than those of Station 7, as demonstrated by a Mann-Whitney-Wilcoxon test
500
(p(a) < 10-4 both for Mo/TOC and U/TOC). Stations 6 and 7 exhibit substantially different
501
degrees of authigenic TM enrichment: at Station 6, m(Mo/TOC) and m(U/TOC) are ~16.3 and
502
~0.8, respectively, whereas at Station 7, m(Mo/TOC) and m(U/TOC) are ~4.7 and ~0.3,
503
respectively (Fig. 6B). Thus, both m(Mo/TOC) and m(U/TOC) are ~3× higher at Station 6 than
504
at Station 7. For Mo, this corresponds quite closely to a 3× difference in [Mo]aq concentrations
505
between the two sites (Fig. 6A), which is consistent with our hypothesis that watermass
506
chemistry is the dominant control on authigenic Mo uptake within a uniformly euxinic
507
environment. For U, the ~2.7× difference in m(U/TOC) between the two sites is larger than the
508
~1.6× difference in [U]aq (i.e., 8.8 nM versus 5.5 nM; Fig. 6A). This discrepancy suggests that
509
some additional factor is enhancing U uptake at Station 6 relative to Station 7. Despite this
510
uncertainty, watermass chemistry appears to be a major influence on spatial variations in
511
authigenic Mo and U enrichment of the sediment within the deep (euxinic) Black Sea watermass.
512
513
4.4.3. Comparison with diffusion-reaction model results
514
To further investigate controls on authigenic Mo and U enrichment of modern Black Sea
515
sediments, we applied the diffusion-reaction model developed above (see Section 3). The
516
specific values of c0, k, Dsed applied to Station 6 and Station 7 sediments in calculating authigenic
517
Mo and U uptake are shown in Table 3. Based on Eq. (16), we tested the sensitivity of authigenic
518
Mo and U enrichment of the sediment to variation in initial aqueous trace-metal concentrations.
519
This test showed that Mo and U concentrations increase linearly as a function of initial aqueous
520
trace-metal concentrations (Fig. 7; Table 3) at a c/c0 ratio that is proportional to
521
φDsed k
M (see Eq. 16). Thus, sediment trace-metal concentrations depend significantly
ρ s (1 − φ ) w2 TM
522
on watermass chemistry, which may vary through space (as a function of local watermass
523
restriction) and time.
524
Considering that the sedimentation rates at Station 6 (0.018 cm yr -1) and Station 7 (~0.020
525
cm yr-1) are similar, we employed a uniform sedimentation rate of ~0.020 cm yr-1 for modeling
526
purposes and simulated the relationship between aqueous trace-metal concentration and
527
authigenic trace-metal concentration of sediments on basis of Eq. (16) in the diffusion-reaction
528
model (Fig 7). Owing to variability in the reaction rate constant (k), we utilized a range of k for
529
Mo (minimum 0.01 yr-1; median 0.03 yr-1; maximum 0.1 yr-1) and U (minimum 0.003 yr-1;
530
median 0.01 yr-1; maximum 0.03 yr-1) (Table 1; Fig. 7). Based on these parameters, the modeled
531
authigenic trace-metal concentrations match closely actual measured values for Mo (Station 6:
532
median 64 ppm, range 58-95 ppm; Station 7: median 45 ppm, range 33-109 ppm) and U (Station
533
6: 12 ppm, range 10-14 ppm; Station 7: median 13 ppm, range 10-16 ppm; Fig. 7). These
534
considerations support the hypothesis that watermass chemistry plays an important role in
535
control of trace-metal enrichment of anoxic marine facies.
536
537
5. Case study: Devonian-Carboniferous boundary North American Seaway
538
5.1. Geological background
539
The interior of North America accumulated black shales widely during the Late Devonian to
540
earliest Carboniferous (Jaminski et al., 1998; Hartwell, 1998; Caplan and Bustin, 1999; Kuhn,
541
1999). For this study, we analyzed four black shale units deposited near the centers of their
542
respective deepwater basins: the Lower and Upper Bakken shales of the Williston Basin, and the
543
Cleveland and Sunbury shales of the Appalachian Basin (Fig. 8). Lithologically, the four units are
544
quite similar, consisting of black to brown, laminated, sub-platy to blocky, non-calcareous and
545
highly carbonaceous shale. Their geochemistry is also similar: the Lower and Upper Bakken shales
546
contain ~10-12 % TOC and ~2-5 % TS (Schmoker and Hester, 1983; Nandy et al., 2014), the
547
Cleveland Shale ~9.3±1.9 % TOC and ~2.5±1.9 % TS, and the Sunbury Shale ~10.1±2.9 % TOC
548
and ~3.5±1.4 % TS (Jaminski et al., 1998; Kuhn, 1999). The thermal maturity levels of these
549
formations are also similar, with Ro values of 1.3-2.0 for the Cleveland and Sunbury shales and
550
1.62-1.73 for the Bakken shales (Price et al., 1984; East et al., 2012).
551
These shale units are well dated: faunal (primarily conodont) evidence assigns the Lower
552
Bakken and Cleveland shales to the expansa-praesulcata Zone of the late Famennian, and the
553
Upper Bakken and Sunbury shales to the Lower Siphonodella crenulata Zone of the early
554
Tournaisian (Gutschick and Sandberg, 1991; Playford and McGregor, 1993). Deposition of the
555
Lower Bakken and Cleveland shales was controlled by a single transgressive-regressive cycle
556
culminating in an end-Devonian glacio-eustatic lowstand, whereas deposition of the Upper
557
Bakken and Sunbury shales in the earliest Carboniferous was tied to a rapid post-glacial eustatic
558
rise (Johnson et al., 1985; Sandberg et al., 2002; Algeo et al., 2007). Owing to control by high-
559
amplitude glacio-eustatic fluctuations, deposition of the lower Bakken and Cleveland shales was
560
fully coeval within dating uncertainties, as was also deposition of the upper Bakken and Sunbury
561
shales. Based on the most recent international time scale (Ogg et al., 2016), the depositional interval
562
of the Lower Bakken-Cleveland shales was ~800±100 kyr, and that of the Upper Bakken-Sunbury
563
shales was ~500±50 kyr. More detailed geological background information can be found in Item
564
S2 of the Supplementary Materials.
565
566
5.2. Materials and methods
567
The Bakken Shale was analyzed in two drillcores: the Lower Bakken Shale in the Sun-
568
Marathon Shobe #1 core (47.895 °N, 102.627 °W) in Mountrail County, North Dakota, and the
569
Upper Bakken Shale in the Texaco Thompson #5-1 core (47.229 °N, 103.263 °W) in Billings
570
County, North Dakota (Hartwell, 1998). The Sun-Marathon Shobe #1 core is located ~30 km east
571
of the basin depocenter, and the Texaco Thompson #5-1 core ~35 km south of the basin
572
depocenter. Core recovery was nearly 100 % for both cores.
573
The Cleveland and Sunbury shales were both analyzed in three drillcores: the KEP-3 core
574
(38.48°N, 83.41°W) in Lewis County, northeastern Kentucky, the OHRS-5 core (39.26°N,
575
83.08°W) in Ross County, south-central Ohio, and the OHDW-1 core (40.20°N, 83.08°W) in
576
Delaware County, central Ohio. All three cores are located in the central Appalachian Basin, south
577
of significant influence by the Catskill Delta and north of the thinned black shale succession (i.e.,
578
Chattanooga Shale) located over the Cumberland Saddle (Schieber, 1994). These cores were
579
originally sampled for the purpose of cm-scale analysis of selected dm-thick (~10-20 cm) cycles
580
that were distributed at intervals through the study units in order to reconstruct paleo-environmental
581
changes at sub-millennial timescales (Jaminski, 1997; Kuhn, 1999; Liu et al., 2019). Although these
582
cores were not sampled continuously, dm-scale cycles were selected from the lower, middle, and
583
upper parts of each formation, providing a representative test of Mo and U concentrations
584
throughout each study unit.
585
For all study units, major and trace element concentrations were analyzed using a
586
wavelength-dispersive Rigaku 3040 X-ray fluorescence spectrometer in the Department of
587
Geology, University of Cincinnati. The results were calibrated using U.S. Geological Survey and
588
internal laboratory standards. The analytical precision was better than ±2% for major and ±5%
589
for trace elements (see Algeo and Maynard, 2004, or Liu et al., 2019, for a full description of XRF
590
techniques).
591
592
5.3. Results
593
Mo concentrations range from 146 to 450 ppm (n = 31) for the Lower Bakken, from 160 to
594
300 ppm (n = 24) for the Upper Bakken, from 58 to 183 ppm (n = 152) for the Cleveland Shale,
595
and from 176 to 412 ppm (n = 87) for the Sunbury Shale (Table 2). U concentrations range from
596
37 to 144 ppm (n = 34) for the Lower Bakken, from 26 to 57 ppm (n = 23) for the Upper
597
Bakken, from 6 to 19 ppm (n = 50) for the Cleveland Shale, and from 14 to 35 ppm (n = 78) for
598
the Sunbury Shale. Authigenic trace-metal concentrations were calculated using the same method
599
as for the Black Sea samples (see Section 4.3). Authigenic Mo concentrations range from 145 to
600
449 ppm (n = 31) for the Lower Bakken, from 160 to 299 ppm (n = 24) for the Upper Bakken,
601
from 56 to 181 ppm (n = 152) for the Cleveland Shale, and from 174 to 410 ppm (n = 87) for the
602
Sunbury Shale (Table 2). Authigenic U concentrations range from 34 to 142 ppm (n = 34) for the
603
Lower Bakken, from 24 to 56 ppm (n = 23) for the Upper Bakken, from 3 to 16 ppm (n = 50) for
604
the Cleveland Shale, and from 10 to 31 ppm (n = 78) for the Sunbury Shale.
605
606
5.4. Discussion
607
5.4.1. Evaluation of depositional environmental conditions
608
Similar benthic redox conditions for the four DCB study units are indicated by a
609
combination of redox proxies (i.e., C-S-Fe systematics, FeT/Al, and Corg:P). The Cleveland and
610
Sunbury shales both yield S/Fe ratios of 0.75 to 1.15, and the Lower and Upper Bakken shales
611
yield S/Fe ratios of ~1.15, consistent with dominantly euxinic conditions but somewhat more
612
intensely reducing conditions for the latter units (Fig. 9A). The Bakken Shale yields higher FeT/Al
613
ratios (Lower Bakken: median 0.52, range 0.35-0.83; n = 35; Upper Bakken (median 0.68, range
614
0.46-1.08; n = 43) relative to the Cleveland (median 0.32, range 0.22-0.44; n = 152) and Sunbury
615
shales (median 0.30, range 0.25-0.40; n = 87) (note: ranges given as 16th-84th percentiles in order
616
to limit the influence of outliers) (Fig. 9B). Whereas the Fe/Al ratios of the Cleveland and
617
Sunbury shales are typical of upper crustal values, the unusually high Fe/Al ratios of the Bakken
618
Shale may reflect additional input of Fe-oxyhydroxides from arid-zone coasts around the
619
margins of the Williston Basin (Witzke and Heckel, 1988; Berwick, 2008). All four of the DCB
620
units exhibit elevated Corg:P ratios: median values are 387 for the Lower Bakken (range 83-789;
621
n = 35), 550 for the Upper Bakken (range 187-674; n = 43), 512 for the Cleveland (range 325 to
622
595; n = 152), and 721 for the Sunbury (range 542 to 847; n = 87) (Fig. 9C), which are consistent
623
with intensely euxinic conditions (cf. Algeo and Ingall, 2007). The observation that the four
624
DCB study units exhibit no consistent sequence of low to high values with regard to these three
625
redox proxies (C-S-Fe, Fe/Al, and Corg:P) suggests that there was no major difference in redox
626
conditions between them (cf. Jaminski et al., 1998; Hartwell, 1998; Kuhn, 1999; Nandy et al.,
627
2014).
628
Devonian-Carboniferous basins of the North American Seaway had somewhat varying
629
degrees of deepwater restriction, as determined from Mo/TOC ratios (Algeo et al., 2007),
630
although this proxy may have been locally influenced by other factors such as sedimentation
631
rates, as discussed herein. The slope of the Mo-TOC regression (i.e., m(Mo-TOC) or ‘m’) can be
632
used to estimate degree of watermass restriction (see Section 2.3). The Lower Bakken Shale
633
yields an m value of 34.9 (Fig. 10A), and the Sunbury and Cleveland shales m values of 34.6 and
634
12.7, respectively (Fig. 10B). The Upper Bakken Shale does not yield a significant Mo-TOC
635
correlation (r = 0.06; p(a) >0.05; n = 43) and, thus, a true m value, although its average Mo and
636
TOC concentrations are close to those of the Lower Bakken Shale (Fig. 10A). These values
637
imply similar degrees of watermass restriction within the Appalachian and Williston basins
638
during the Late Devonian to Early Carboniferous, except for relatively greater restriction in the
639
Late Devonian Appalachian Basin (cf. Algeo et al., 2007). U-TOC relationships are similar to
640
Mo-TOC relationships for the study units, with median values of 6.2, 2.6, and 2.1 for the Lower
641
and Upper Bakken, Sunbury, and Cleveland shales, respectively (Fig. 10C-D), suggesting
642
somewhat greater watermass restriction in the Appalachian Basin than in the Williston Basin. It
643
is possible that these more restricted conditions influenced watermass chemistry and, thus, Mo
644
and U uptake, an issue that will be considered below (see Section 5.4.2).
645
The degree of restriction of an epicontinental sea or marginal-marine basin is commonly
646
reflected in its salinity, with lower salinities associated with greater restriction owing to reduced
647
seawater/freshwater mixing ratios. Paleo-salinities can be estimated in carbonate-free shales on
648
the basis of Sr/Ba ratios, with values <0.2 and >0.5 indicative of freshwater and fully marine
649
conditions, respectively (Wei et al., 2018; Wei and Algeo, 2019). Mean Sr/Ba ratios in the DCB
650
study units are 0.26±0.09 for the Lower Bakken, 0.30±0.07 for the Upper Bakken, 0.24±0.02 for
651
the Cleveland, and 0.23±0.02 for the Sunbury, supporting brackish conditions in both the
652
Williston and Appalachian basins (Table 4). That these values are indicative of brackish rather
653
than marine conditions is demonstrated by the systematically higher Sr/Ba ratios of the
654
Woodford Shale (0.30-0.36) and Chattanooga Shale (0.28-0.45; Table 4), both of which
655
accumulated on the paleo-southern margin of North America and thus experienced relatively
656
greater influence from open-ocean waters to the south. The slightly higher Sr/Ba ratios of the
657
Lower and Upper Bakken shales (0.26-0.30) relative to the Cleveland and Sunbury shales (0.22-
658
0.25) may reflect more limited freshwater runoff into the Williston Basin, which was located in
659
the arid subtropics and did not experience orogenic precipitation effects as in the Appalachian
660
Basin (Witzke and Heckel, 1988; Berwick, 2008). However, the Sr/Ba ratios of the four DCB
661
study units are sufficiently similar that the implied small differences in their depositional
662
watermass salinities are negligible for present purposes. It should be noted nonetheless that
663
similar watermass salinities do not preclude the possibility of differences in aqueous trace-metal
664
concentrations between the two study basins, e.g., as a result of watermass chemical evolution
665
(cf. Algeo and Maynard, 2008).
666
667
5.4.2. Influence of sedimentation rates on Mo-U enrichment
668
For the DCB black shales, the median Mo/TOC is 23.5 in the Lower Bakken (range 14.9-
669
26.6; n = 31), 21.7 in the Upper Bakken (range 15.8-29.3; n = 24), 28.8 in the Sunbury (range
670
19.2-37.2; n = 87), and 14.3 in the Cleveland (range 9.3-22.5; n = 152). The median U/TOC is
671
7.2 in the Lower Bakken (range 5.0-12.5; n = 34), 4.1 in the Upper Bakken (range 2.1-6.6; n =
672
23), 2.6 in the Sunbury (range 1.7-3.4; n = 78), and 1.2 in the Cleveland (range 0.7-1.7; n = 49)
673
(note:
674
Bakken:Sunbury:Cleveland ratios of ~1.5:2:1 for Mo and ~5:2:1 for U.
ranges
given
as
16th-84th
percentiles).
These
median
values
thus
reflect
675
The differences in authigenic Mo and U enrichment levels between the DCB study units can
676
be accounted for mostly through differences in sedimentation rates. In the study cores, the
677
thicknesses of the Lower Bakken and Upper Bakken shales are 4.0 m (n = 1) and 2.5 m (n = 1),
678
and those of the Cleveland and Sunbury shales are 13.4-21.6 m (n = 3) and 5.0-5.1 m (n = 3),
679
respectively (Jaminski, 1997; Kuhn, 1999). Duration estimates of ~800 kyr for the Lower
680
Bakken/Cleveland interval and ~500 kyr for the Upper Bakken/Sunbury interval (see Section 5.1)
681
yield average sedimentation rates of 5.0 m Myr‒1 for both the Lower and Upper Bakken shales,
682
10.0-10.2 m Myr‒1 for the Sunbury Shale, and 17-27 m Myr‒1 for the Cleveland Shale. These
683
values represent sedimentation rate ratios for the Bakken:Sunbury:Cleveland shales of ~1:2:3.5-
684
5.5, which is approximately the inverse of the TM enrichment ratios for authigenic U above.
685
Control of authigenic TM enrichment in black shales by sedimentation rates depends on
686
differential diffusion of TMs within the upper sediment column. Diffusion rates are an integrated
687
function chiefly of sedimentation rate, TM diffusion coefficients, sediment porosity, and pore
688
network tortuosity (Morford et al., 2009; Scholz et al., 2011). Although the diffusion coefficient
689
for a given TM is temperature- and pressure-dependent (Franks, 1972), it would not have varied
690
much under the similar environmental conditions that prevailed in the deep Williston and
691
Appalachian Basins. Furthermore, the Bakken, Sunbury, and Cleveland shales are all fine-
692
grained organic-rich siliciclastic units and probably had similar sediment porosity and tortuosity
693
characteristics, so Dsed (Eq. 2) can be regarded as a constant. However, the DCB shales are
694
characterized by markedly different sedimentation rates, which were ~2× to 5× greater for the
695
Cleveland and Sunbury shales relative to the Lower and Upper Bakken shales (see above). Other
696
factors being equal, higher sedimentation rates should result in the sediment being removed more
697
rapidly from contact with the overlying water column and, thus, less TM enrichment per unit
698
volume of sediment. This inference is fully consistent with the differences in U concentrations
699
among the DCB shales noted above and, thus, with a sedimentation rate control on authigenic U
700
enrichment.
701
Authigenic Mo enrichment also may have been subject to a sedimentation rate control, as
702
suggested by the similar median authigenic Mo concentrations and sedimentation rates for the
703
Lower and Upper Bakken shales, and the 2× difference between the Cleveland and Sunbury
704
shales (with the Sunbury exhibiting higher authigenic Mo concentrations and lower
705
sedimentation rates than the Cleveland). However, differences in sedimentation rates between
706
the Bakken shales on the one hand and the Cleveland-Sunbury shales on the other suggest that
707
the former should exhibit ~3× greater Mo enrichment (or the latter ~3× less enrichment) than
708
actually observed. We infer that an additional factor must have influenced Mo enrichment, e.g.:
709
(1) aqueous Mo concentrations, which might have been lower in the Williston Basin than in the
710
Appalachian Basin (cf. Algeo and Maynard, 2008), (2) operation of a Mn-Fe particulate shuttle
711
in the Appalachian Basin, leading to relatively greater Mo enrichment of the Sunbury and
712
Cleveland shales (cf. Algeo and Tribovillard, 2009), or (3) unrecognized differences in redox
713
conditions. Note that all three of these potential influences are operative mainly at a basinal scale
714
rather than a local scale, and that they are thus consistent with the need for a mechanism to
715
account for differences in Mo enrichment between the Williston and Appalachian basins but not
716
for differences within each basin, for which sedimentation rate variation is a sufficient
717
explanation. We cannot determine with certainty which of these three factors was responsible for
718
inter-basinal differences in Mo enrichment, but we suspect differences in aqueous Mo
719
concentrations because certain considerations weigh against the other two factors. First, there is
720
no evidence to support operation of a Mn-Fe particulate shuttle during Cleveland-Sunbury
721
deposition, although a shuttle is known to have influenced Mo uptake by the Chattanooga Shale
722
(the upper part of which is laterally equivalent to the Cleveland Shale) at the southern end of the
723
Appalachian Basin (Algeo and Tribovillard, 2009). Second, although the Appalachian Basin was
724
probably deeper than the Williston Basin and thus potentially subject to more reducing
725
bottomwater conditions, the multiple redox proxies examined in Section 5.4.1 provide no support
726
for this redox scenario, and depositional models for the Appalachian Basin suggest that it
727
received episodic inflows of oxygenated waters from the Catskill Delta (e.g., Jaminski et al.,
728
1998), whereas freshwater discharge into the arid-zone Williston Basin was almost certainly
729
more limited (Algeo et al., 2007). This considerations suggest that greater authigenic Mo
730
enrichment in the Appalachian Basin was not due to more reducing bottomwater conditions,
731
making a higher aqueous Mo concentration in the Appalachian Basin (i.e., ~3× greater than in
732
the Williston Basin) the most likely reason for relatively greater Mo enrichment of the Sunbury
733
and Cleveland shales.
734
735
5.4.3. Comparison with diffusion-reaction model results
736
We tested the potential influence of sedimentation rates on TM enrichment patterns in the
737
DCB black shales using the diffusion-reaction model developed in Section 3. Initial seawater TM
738
concentrations must be estimated or inferred, and given a lack of age-specific constraints we
739
adopted the following estimates: modern seawater concentrations, i.e., Moaq = 110 nM; Uaq = 13
740
nM (Algeo and Tribovillard, 2009), with a reduction in Williston Basin seawater Mo by a factor
741
of 3× (i.e., to Moaq = 37 nM) to account for interbasinal differences in authigenic Mo enrichment
742
levels, as discussed in Section 5.4.2. We also adopted modern reaction rate constants for Mo and
743
U in sediment porewaters, with specific values of c0, k, Dsed for Mo and U given in Table 5.
744
Based on Eq. (16), we tested the sensitivity of authigenic Mo enrichment as a function of depth
745
below the sediment-water interface (Fig. 11; Table 5). As in similar published models (Morford
746
et al., 2007; Hardisty et al., 2018), authigenic Mo enrichment is quite sensitive to sedimentation
747
rate, with lower sedimentation rates yielding greater enrichment, and vice versa. These results
748
are a robust demonstration that sedimentation rate can exert a strong influence on the level of
749
authigenic enrichment of trace metals in the sediment.
750
Total concentrations of Mo and U in the sediment increase with depth as a result of
751
authigenic uptake, eventually reaching a stable maximum value (Fig. 11). Mo and U
752
concentrations increase more quickly with depth and reach higher maximum values at low
753
sedimentation rates than at high sedimentation rates. Authigenic Mo and U concentrations in the
754
sediment thus exhibit a linear negative relationship to sedimentation rate. The results of our
755
diffusion-reaction model conform well to measured sedimentation rates and authigenic Mo and
756
U concentrations in the four DCB black shales (Fig. 12). The study units exhibit an
757
approximately four-fold range of sedimentation rates, from ~0.0005 cm yr‒1 in the Lower Bakken
758
and Upper Bakken shales to ~0.0010 cm yr‒1 in the Sunbury Shale and ~0.0022 cm yr‒1 in the
759
Cleveland Shale. The median authigenic Mo concentrations are 279 ppm (range 159-369 ppm;
760
16th-84th) for the Lower and Upper Bakken, 241 ppm (range 174-410 ppm; 16th-84th) for the
761
Sunbury, and 104 ppm (range 56-181 ppm; 16th-84th) for the Cleveland. The median authigenic
762
U concentrations are 55 ppm (range 31-115 ppm; 16th-84th) for the Lower and Upper Bakken, 24
763
ppm (range 10-31 ppm; 16th-84th) for the Sunbury, and 9 ppm (range 3-16 ppm; 16th-84th) for the
764
Cleveland. Sedimentation rate variation among the four DCB black shales can thus account fully
765
for differences in authigenic Mo and U enrichment, with the additional requirement that aqueous
766
Mo concentrations were lower in the Williston Basin than in the Appalachian Basin by a factor of
767
~3× (Fig. 12). These results demonstrate that authigenic Mo and U concentrations in marine
768
sediments are highly sensitive to variation in sedimentation rates.
769
770
6. Conclusions
771
Authigenic Mo-U enrichment patterns are associated with specific environmental conditions
772
(e.g., redox state, degree of deepwater restriction) and depositional processes (e.g., sedimentation
773
rate). We utilized a diffusion-reaction model to illustrate the effects of watermass chemistry and
774
sedimentation rates on authigenic trace-metal enrichment, focusing on molybdenum (Mo) and
775
uranium (U). From a theoretical perspective, sedimentary trace-metal concentrations scale
776
directly to their aqueous concentrations, and inversely to sedimentation rate owing to its
777
influence on (i) the diffusional gradient of Mo or U, and (ii) the concentration profile of key
778
reactants (i.e., H2S for Mo, TOC for U). We tested these relationships through case studies of
779
authigenic Mo and U enrichment in modern Black Sea sediments and in Devonian-Carboniferous
780
boundary (DCB) black shales from the Williston and Appalachian basins. In the Black Sea case
781
study, two sites (Stations 6 and 7) record >2× differences in authigenic Mo and U enrichment,
782
despite similar redox conditions and sedimentation rates. The upper part of the Black Sea water
783
column exhibits strong vertical gradients in aqueous trace-metal concentrations, matching almost
784
exactly the observed differences in sedimentary Mo and U enrichment and thus demonstrating
785
the influence of watermass chemistry on authigenic trace-metal uptake. In the DCB case study,
786
the four black shales record >4× differences in authigenic Mo and U enrichment, despite similar
787
redox and salinity conditions. The most likely control on authigenic Mo and U enrichment
788
appears to be sedimentation rates, which were ~2× higher in the Sunbury and ~4-5× higher in the
789
Cleveland relative to the Lower and Upper Bakken. The inferences of both case studies are
790
supported by the results of our diffusion-reaction model, which closely match observed
791
sedimentary Mo and U concentrations based on calculated sedimentation rates and diffusion
792
coefficients from modern marine systems. This study thus demonstrates the potential influences
793
of watermass chemistry and sedimentation rates on trace-metal enrichment in anoxic marine
794
facies.
795
796
Acknowledgments
797
JSL was financially supported by the China Postdoctoral Science Foundation (grand No.
798
K2819003), Outstanding Postdoctoral Scholarship, State Key Laboratory of Marine
799
Environmental Science at Xiamen University (grant No. K04002, K4318012), National Key
800
R&D Project of China (2016YFA0601104), National Natural Science Foundation of China (grant
801
No. 41290260, 41472170, 41772004), the 111 project (grant No. B08030) and the China
802
Scholarship Council Fund (File No. 201706410081). TJA thanks the Sedimentary Geology and
803
Paleobiology program of the U.S. National Science Foundation, and the China University of
804
Geosciences-Wuhan
805
BGL21407).
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(SKL-GPMR
program
GPMR201301,
and
SKL-BGEG
program
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Table 1. Compilation of sedimentation rates, initial concentrations (c0), and reaction rate
constants (k) for Mo and U in selected cores
Core a
Sedimentation
rate (cm yr‒1)
c0 for Mo
Best-fit curve for Mo
k for Mo
(yr-1)
Core 1
0.6
150
y = 150·exp[(-0.12/0.6)·x]
0.12
FOAM
0.2
163
y = 163·exp[(-0.014/0.2)·x]
0.014
MUC19
0.05
110
y = 110·exp[(-0.0038/0.05)·x]
0.0038
Core a
Sedimentation
rate (cm yr‒1)
0.6
0.026
0.058
c0 for U
Best-fit curve for U
13
13
13
y = 13·exp[(-0.12/0.6)·x]
y = 13·exp[(-0.0087/0.026)·x]
y = 13·exp[(-0.0052/0.058)·x]
Core 2
MUC39
MUC53
1224
1225
1226
1227
a
k for U
(yr-1)
0.12
0.0087
0.0052
The Mo and U profiles for Cores 1 and 2 are from Morford et al. (2007), FOAM from Hardisty
et al. (2018), and MUC19, MUC 39, and MUC53 from Scholz et al. (2011).
1228
1229
Table 2. Median values and range data for Station 6, Station 7, Cleveland, Sunbury, and Bakken
shales
Characteristics
64
45
104
241
279
Median authigenic concentration (U) a
12
13
9
24
55
58-95
10-14
33-109
10-16
56-181
3-16
174-410
10-31
159-369
31-115
Range
for authigenic Mo
th
th
Range (16 -84 ) for authigenic U
1231
1232
1233
1234
1235
Devonian case study
Cleveland Sunbury Bakken
Median authigenic concentration (Mo) a
(16th-84th)
1230
Modern case study
Station 6 Station 7
a
Authigenic concentrations were calculated following Algeo and Maynard (2004) as:
[TM ]
[TM ]auth = [TM ]tot − (
)UCC[ Al ]tot , where [TM]tot and [Al]tot are the total TM and Al
[ Al]
[TM ]
)UCC is the average Al-normalized
[ Al ]
concentration of upper continental crust (UCC; McLennan, 2001).
b MARs for authigenic Mo and U were calculated per Eq. (14).
concentrations of the sample, respectively, and (
1236
Table 3. Diffusion-reaction model parameters for modern Black Sea
Characteristics
Initial concentration (c0) for Station 6 (nM) a
Mo
15.3
U
8.3
Initial concentration (c0) for Station 7 (nM) a
4.8
5.5
Reaction rate constant (k, yr-1) b
0.03
0.01
391
134
cm2·yr-1) c
Diffusion coefficient (Dsed,
Sedimentation rate for Station 6 (cm yr‒1)
Sedimentation rate for Station 7 (cm yr‒1)
Porosity (ϕ) d
Density (ps, g·cm-3) e
1237
1238
1239
1240
1241
1242
1243
a
0.018
0.02
0.7-0.9
1.88
TM concentrations at the sediment/water interface based on Figure 6A.
Reaction rate constants for Mo and U based on Figure 3.
c Diffusion coefficients for Mo and U from Malinovsky et al. (2007) and Li and Gregory (1974),
respectively.
d Porosity values from Opreanu (2003), with porosity-depth relationship given by Eq. (7).
e Dry bulk-sediment density from Böning et al. (2004).
b
1244
1245
1246
1247
1248
1249
Table 4. Sr/Ba paleosalinity proxy data for DCB black shales
Mountrail County, ND (47.895 °N, 102.627 °W)
Sr/Ba
(mean ±
1σ)
0.26 ± 0.09
35
Thompson
Billings County, ND (47.229 °N, 103.263 °W)
0.30 ± 0.07
43
Cleveland
KEP-3
Lewis County, KY (38.409°N, 83.437°W)
0.23 ± 0.02
29
Cleveland
Cleveland
Sunbury
Sunbury
Sunbury
OHRS-5
OHDW-1
KEP-3
OHRS-5
OHDW-1
Ross County, OH (39.258°N, 83.076°W)
67
18
47
24
16
Woodford
Woodford
Woodford
Chattanooga
RSP
CLY
AJD
LC
Yoakum County, TX (33.14°N, 102.89°W)
0.24 ± 0.01
0.22 ± 0.01
0.23 ± 0.02
0.23 ± 0.01
0.25 ± 0.01
0.34 ± 0.10
0.36 ± 0.16
0.30 ± 0.16
Leslie County, KY (37.156°N, 83.389°W)
0.28 ± 0.12
Formation
Section/
Corea
Location (latitude-longitude)
Lower Bakken
Shobe
Upper Bakken
Delaware County, OH (40.288°N, 82.784°W)
Lewis County, KY (38.409°N, 83.437°W)
Ross County, OH (39.258°N, 83.076°W)
Delaware County, OH (40.288°N, 82.784°W)
Pontotoc County, OK (34.674°N, 96.641°W)
Murray County, OK (34.46°N, 97.15°W)
n
18
29
7
31
Humphreys County, TN (36.075°N, 87.495°W)
Chattanooga
DGHS
0.45 ± 0.08 26
a All study sections and cores are part of the present study except: RSP = Ryan Shale Pit; CLY =
Classen Lake YMCA; AJD = Amoco A.J. Davis #9; LC = Leslie County; DGHS = Dupont
GHS. Note: for the raw data, see the Supplemental Information.
1250
Table 5. Diffusion-reaction model parameters for DCB black shales
Characteristics
Initial concentration (c0)
Mo
110
U
13
Reaction rate constant (k, yr-1) a
0.03
0.01
391
134
Diffusion coefficient (Dsed, cm2·yr-1) b
‒1
Sedimentation rate for Cleveland Shale (cm yr )
Sedimentation rate for Sunbury Shale (cm yr‒1)
Sedimentation rate for Bakken Shale (cm yr‒1)
Porosity (ϕ) c
1251
1252
1253
1254
1255
1256
1257
0.0017-0.0027
0.0010
0.0005
0.7-0.9
Density (ps, g·cm-3) d
1.88
a
Reaction rate constants for Mo and U based on Figure 3.
b Diffusion coefficients for Mo and U from Malinovsky et al. (2007) and Li and Gregory (1974),
respectively.
c Porosity values from Opreanu (2003), with porosity-depth relationship given by Eq. (7).
d Dry bulk-sediment density from Böning et al. (2004).
1258
1259
1260
1261
1262
1263
1264
1265
1266
1267
1268
1269
1270
1271
1272
1273
1274
1275
1276
1277
1278
1279
1280
1281
1282
1283
1284
1285
1286
1287
1288
1289
1290
1291
1292
1293
1294
1295
1296
1297
1298
1299
1300
1301
1302
1303
Figure captions
Fig. 1. Diffusion-reaction model for authigenic trace-metal uptake by sediments. The principal
reactions for Mo and U are conversion of Mo(VI) from molybdate (MoO42-) to thiomolybdate
(MoS42-) and reduction of U(VI) to U(IV), respectively. SWI = sediment-water interface.
Fig. 2. Porosity versus depth with best-fit curve (red). Data for cores CH12 (blue) and CH18
(green) from the recent Black Sea (Opreanu, 2003; his table 1).
Fig. 3. Mo and U concentrations in cores MUC 19, 39, and 53 of Peru Margin (Scholz et al.,
2011), FOAM core of Long Island Sound, New York (Hardisty et al., 2018), and Cores 1 and 2
of Hingham Bay, Massachusetts (Morford et al., 2007). Best-fit curves (dotted lines) were
calculated for the purpose of determining reaction rate constants for Mo and U uptake (Table 1).
SR = average sedimentation rate.
Fig. 4. (A) Location map and (B) stratigraphic columns for modern Black Sea sediments. Map
modified from Algeo and Lyons (2006). Station 6 and 7 stratigraphy from Lüschen (2004).
Fig. 5. Redox proxies for Stations 6 and 7 of Black Sea. (A) C-S-Fe ternary system, (B) FeT/Al,
and (C) molar Corg:P. Data from Lüschen (2004).
Fig. 6. Black Sea data: (A) Aqueous Mo and U profiles (Anderson et al., 1989; Emerson and
Huested, 1991; Colodner et al., 1995; Algeo and Lyons, 2006). The arrows labeled 110 and 13
represent global-ocean seawater Mo and U concentrations in units of nM. (B) Sediment Mo
versus TOC and (C) sediment U versus TOC for Stations 6 and 7; data from Lüschen (2004).
Fig. 7. Modeled sediment concentrations of Mo (A) and U (B) as a function of initial aqueous
TM concentrations, per Eq. (16). Calculated curves based on three reaction rate constants (k) for
Mo (minimum 0.01 yr-1; median 0.03 yr-1; maximum 0.1 yr-1) and U (minimum 0.003 yr-1;
median 0.01 yr-1; maximum 0.03 yr-1). The shaded area represents observed k values, which
range from 0.0038 to 0.12 yr-1 in A, and from 0.0052 to 0.12 yr-1 in B (see Figure 3). Authigenic
Mo and U enrichment levels are linearly related to [TM]aq and to k. The authigenic Mo and U
data can be found in Table 2.
Fig. 8. (A) Paleogeography, (B) location map, and (C) stratigraphic columns for Late Devonian
Williston and Appalachian basins. The base map in A is from Blakey (2005). Abbreviations:
AOB = Antler Orogenic Belt, WB = Williston Basin, TA = Transcontinental Arch, AB =
Appalachian Basin.
Fig. 9. Redox proxies for North American DCB shales. (A) C-S-Fe ternary system, (B) FeT/Al,
and (C) molar Corg:P. Data sources: Bakken Shale (Hartwell, 1998) and Cleveland-Sunbury
shales (Jaminski, 1997; Kuhn, 1999).
Fig. 10. Total Mo versus TOC for (A) Bakken Shale and (B) Cleveland-Sunbury shales.
Correlations are significant for Lower Bakken (r = +0.96; p(a) <0.05; n = 35), Cleveland (r =
+0.81; p(a) <0.05; n = 52), and Sunbury (r = +0.82; p(a) <0.05; n = 79). Total U versus TOC for
1304
1305
1306
1307
1308
1309
1310
1311
1312
1313
1314
1315
1316
(C) Bakken Shale and (D) Cleveland-Sunbury shales. Correlations are significant for Lower
Bakken (r = +0.75; p(a) <0.05; n = 34), Cleveland (r = +0.49; p(a) <0.05; n = 50), and Sunbury
(r = +0.69; p(a) <0.05; n = 79). Data sources: Bakken Shale (Hartwell, 1998), ClevelandSunbury shales (Jaminski, 1997; Kuhn, 1999).
Fig. 11. Modeled sediment concentrations of Mo (A) and U (B) as a function of depth, for four
different sedimentation rates (SR). Arrows at the top of each panel show average [Mo] and [U]
for the Bakken (BK), Sunbury (SB), and Cleveland (CL) shales. Final authigenic Mo and U
enrichment levels (at depth) are linearly related to SR0.5.
Fig. 12. Authigenic concentrations of (A) Mo and (B) U as a function of sedimentation rate, per
Eq. (16) and using seawater concentrations (c) of Mo and U as shown. The vertical gray bars
represent measured values for the four study units.