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Economic Geology
Vol. 91, 1996, pp. 36-47
Sedimentary Geochemistry of Manganese: Implications for the Environment
of Formation of Manganiferous Black Shales
S. E. CALVERT AND T. F. PEDERSEN
Department of Earth and Ocean Sciences, University of British Columbia, Vancouver, B.C., Canada V6T 1Z4
Abstract
The sedimentary geochemistry of manganese is dominated by the redox control of its speciation, higher
oxidation states (Mn s+ and 4+) occurring as insoluble oxyhydroxides in well-oxygenated environments and the
lower oxidation state (Mn '•+) being much more soluble in oxygen-deficient settings. Its geochemical behavior
is therefore quite different in oxic and anoxic environments, and where oxic and anoxic conditions are
juxtaposed, Mn is recycled between the two environments. In modern marine sediments, Mn is present above
its crustal abundance as an oxyhydroxide in all slowly accumulating (pelagic) sediments of the deep ocean
and in surficial deposits of continental margin environments. Diagenetic recycling of Mn in the latter causes
surficial deposits to have larger Mn enrichments than in many pelagic sediments. Bottom sediments of
permanently anoxic basins show no enrichments and have Mn concentrations that are controlled solely by
the aluminosilicate fraction. Manganese carbonates (kutnohorite and calcic rhodochrosite) are found only in
anoxic sediments accumulating beneath surface oxic horizons (and therefore under oxygenated bottom waters)
in many nearshore environments. Such enrichments are due to delivery of Mn by burial of surface oxyhydrox-
ides into the subsurface anoxic environment where they are dissolved. Pore-water Mn levels can reach
saturation with respect to a mixed Mn-Ca carbonate phase in such sediments. The diagenetic origin for these
phases is shown by their carbon isotope compositions, which typically indicate a carbon source from decompos-
ing organic matter. The presence of Mn carbonates therefore signifies that the host sediment must have
accumulated under oxygenated bottom waters.
On the basis of this information it is proposed that, in contrast to several current explanations for the
formation of Mn carbonates (kutnohorite and rhodochrosite) in ancient organic-rich shales, limestone, and
marl sequences and in many Mn ore deposits, the occurrence of these mineral phases indicates that the
sediments originally accumulated beneath oxygenated bottom waters. By analogy with the present, Mn carbon-
ates could not have formed in the bottom waters of anoxic basins. These diagenetic phases, however, did
form where Mn was supplied at a high rate, namely, by the burial of oxyhydroxide-enriched surface sediments,
to a subsurface anoxic environment. This situation could only have occurred under oxygenated bottom waters.
The presence of Mn carbonates in ancient black shales (and in some carbonate-rich rocks) lends strong
support to the notion that these rocks did not necessarily form in anoxic basins but owe their carbon richness
to a high supply of organic matter to sediments deposited under oxygenated bottom waters, probably in
continental margin settings.
Introduction
THE ENVIRONMENT of formation of organic-rich (black)
shales in the geologic record is receiving increasing attention
because of interest in the mechanisms by which sedimentary
deposits are enriched in organic matter and heavy metals.
The fact that such rocks are carbon rich is commonly ascribed
to the preferential preservation of deposited organic matter
under anoxic (sulfide-bearing) bottom waters (Demaison and
Moore, 1980; Demaison et al., 1984), so that many black
shales are presumed to have formed in anoxic basins (Wool-
nough, 1937) or in intermediate depth oxygen minima
(Schlanger and Jenkyns, 1976; Thiede and Van Andel, 1977)
on continental margins or on the margins of oceanic pedestals
and plateaus. An alternative explanation is that black shales
owe their formation to relatively high settling fluxes of organic
matter brought about by high primary plankton production
in the surface waters of the ocean (Calvert, 1987; Pedersen
and Calvert, 1990; Calvert and Pedersen, 1992). Enrichments
of many metals in such rocks have also been ascribed to
bottom-water anoxia in the basin of sedimentation (Holland,
1979; Brnmsack, 1986a, 1986b, 1991; Dean and Arthur, 1986;
Arthur et al., 1990; Hatch and Leventhal, 1992; Turner,
1992). Hence, if we had a sound knowledge of the marine
geochemistry of deposits which are enriched in minor and
trace elements, then the relative concentrations of these ele-
ments might be used as guides to conditions that led to the
formation of these unique rocks.
Higher manganese contents, in some deposits reaching ore
grade, are associated with some black shales, either as depos-
its of oxides or as carbonates (Roy, 1988, 1992). Because the
behavior of Mn in sediments is known to be strongly affected
by changes in redox potential, such concentrations have been
used to infer deposition of the oxide deposits under highly
oxidizing conditions (Cannon and Force, 1983; Frakes and
Bolton, 1984, 1992; Bolton and Frakes, 1985; Force and Can-
non, 1988) and the carbonate deposits under reducing condi-
tions either (1) in anoxic or oxygen-deficient bottom waters
(Bencini and Turi, 1974; Kim, 1979; Bolton and Frakes, 1985;
Jenkyns, 1988; Jenkyns et al., 1991; Pratt et al., 1991; Frakes
and Bolton, 1992); (2) at the boundaries between oxygenated
and anoxic or dysaerobic waters as a consequence of the
mixing of deeper anoxic waters with more oxidizing waters
at shallower depths (Cannon and Force, 1983; Frakes and
Bolton, 1984; Force and Cannon, 1988; Force and Maynard,
1991; Fan et al., 1992); or (3) in anoxic sediments (Hein and
Kosld, 1987; Force and Cannon, 1988; Oldta et al., 1988;
0361-0128/96/1800/36-1255.00 36
Mn GEOCHEMISTRY: IMPLICATIONS FOR BLACK SHALES 37
Force and Maynard, 1991; Minoura et al., 1991; Polgaff et
al., 1991; Sugisaki etal., 1991; Okita and Shanks, 1992; Pol-
gaff, 1993). Force and Cannon (1988) thought that Mn car-
bonates formed under both oxic and anoxic water columns.
Hence, the redox chemistry of Mn and the possibility that
ocean waters were anoxic in the past are critical factors in
some models for the formation of many of the large oxide
and carbonate Mn orebodies in the geologic record.
The purpose of this paper is to reexamine the environment
of formation of Mn carbonates in black shale settings in light
of new information on the geochemistry of Mn in the modern
ocean. Our interpretation of these data demonstrates that
some of the current models for the formation of Mn carbon-
ate ores which are based on the behavior of Mn in the modern
ocean should be revised and that the presence of Mn carbon-
ates in some black shales supports the hypothesis that such
deposits accumulated under oxygen-beaffng bottom waters.
Oxic and Anoxic Sedimentation
In the discussion that follows, we shall refer to two situa-
tions in which anoxic conditions prevail in bottom sediments.
One, of course, involves a basin of sedimentation in which
the bottom waters are themselves anoxic, that is, they contain
dissolved H2S. Such environments most commonly occur
where a topographic barrier restricts deep-water replen-
ishment in a coastal basin, so that the rate of consumption
of dissolved oxygen exceeds its rate of replenishment by deep-
water renewal. Well-known examples of this category include
the Black Sea, the Caffaco trench, and several Norwegian and
Canadian fjords (see Richards, 1965). The second situation, in
which sedimentary anoxic conditions occur, is found when
the rate of burial of particulate organic matter exceeds the
rate of replenishment of dissolved oxygen from bottom waters
by diffusion or irrigation. Such sediments accumulate along
continental margins under oxygenated bottom waters; the
surficial sediments are therefore oxic, but they become anoxic
(sulfidic) at a variable depth below the sediment-water inter-
face owing to the microbial utilization of oxygen, nitrate, and
Fe and Mn oxyhydroxides during degradation of organic mat-
ter, after which H2S is produced as a by-product of bacterial
sulfate reduction (Froelich etal., 1979). Such sediments are
distinctive in having a vaffably thick reddish to dark brown
surface layer because of the presence of manganese and iron
oxyhydroxides; this layer overlies olive-green reducing sedi-
ments at depth (Gorshkova, 1931, 1957; Lynn and Bonatti,
1965).
Manganese in Seawater
Manganese occurs as the Mn •+, Mn 3+, and Mn 4+ oxidation
states in seawater. Mn •+ and MnC1 + are the dominant species
in oxic seawater, but Mn •+ is thermodynamically unstable in
the presence of oxygen and is sluggishly oxidized at neutral
3+ 4+
pH to insoluble hydrated Mn and Mn oxides, normally
referred to collectively as Mn oxyhydroxides. The vertical
distribution of dissolved Mn in the open ocean (Fig. 1A)
generally shows a near-surface maximum, ranging from 1 to
8 nM (0.05-0.4/•g-1) in the North Pacific to 20 nM in the
North Atlantic, low and relatively constant values of about
0.5 nM in deep water, and occasional maxima in bottom
waters that can reach levels of 8 nM (Bender etal., 1977;
1000
5ooo
Dissolved Mn (nM)
I 2 3 4
OS 08 tg,u•O 0 s 0 ,5,•
ß
ß o
ß o
ß ß
ß
ß
A
! !
50 100
Dissolved 0 2 (glVl)
o
o
• GEOSECS
Station 343
!
150 200 250
Dissolved Mn (nM)
0 5 10 15
Eastern
hi Tropical
oo Pacific
•o o o Galapagos
o o Oo Riff
• oOO •g ©o o
Fxc. 1. A. Vertical distribution of dissolved Mn (filled circles) and oxygen
(open circles) in the northeastern Pacific (16031.6 ' N, 12301.4 ' W) showing
the low Mn concentrations between 1,000 and 3,500 m, higher concentra-
tions at the sea surface and below 3,500 m, and the concentration maximum
coincident with the oxygen minimum (from Klinkhammer and Bender,
1980). B. Vertical distribution of dissolved Mn in the eastern tropical Pacific
(same as A on a different concentration scale) and over the Galapagos rift
(0048.6 ' N, 8609.3 ' W). The rapid increase in Mn with depth over the rift
is due to the flux of hydrothermal Mn from the Galapagos spreading center.
From Khnkhammer (1980b) and Klinkhammer and Bender (1980).
Klinkhammer and Bender, 1980; Landing and Bruland, 1980;
Murray etal., 1983). High concentrations of dissolved Mn,
reaching 3 riM, are also found at intermediate depths where
dissolved oxygen concentrations are below about 100/•M (the
oxygen minimum) (Bender et al., 1977; Klinkhammer and
Bender, 1980) or at locations close to the influence of ridge
crest hydrothermal activity (Fig. lB) (Klinkhammer etal.,
1977; Weiss etal., 1977; Edmond et al., 1979). The surface
enrichments are mainly due to the lateral supply of dissolved
Mn from nearshore regions, either from rivers or by regenera-
tion from sediments (see Manganese in Marine Sediments,
below), with atmospheric supplies being of minor importance.
Low concentrations in deep waters reflect active removal of
dissolved Mn onto particle surfaces by oxidative precipitation
(Goldberg, 1961; Morgan, 1967). Increases in concentration
close to the sea floor imply supply of Mn to the ocean from
bottom sediments (Klinkhammer and Bender, 1980).
The intermediate depth Mn maximum that is often ob-
served coincident with the oxygen minimum (Fig. 1A) has
been ascribed to the reductive dissolution of Mn oxyhydrox-
ides from settling particles or the transport of dissolved Mn
that diffuses out of anoxic shelf and slope sediments (Klink-
hammer and Bender, 1980). Johnson etal. (1992) measured
the flux of Mn from such sediments in the eastern Pacific
and showed that this source is inadequate to account for
the Mn concentration maximum observed within the oxygen
minimum. These authors concluded that processes within
the water column itself, involving reductive dissolution of
manganese oxyhydroxides falling through the water column
(Klinkhammer and Bender, 1980), the release of dissolved
Mn from decomposing organic particles settling through the
water column (Landing and Bruland, 1980), and/or a reduc-
tion in the rate of Mn scavenging in the oxygen minimum,
38 CALVERT AND PEDERSEN
Dissolved Mn (nM)
0 2000 4000 6000 8000 10000
1500
2OOO
Dissolved
Black Sea
Station BS3-2
2 4 6 8 10
Particulate Mn (nM)
FIG. 2. Depth distribution of particulate and dissolved Mn in the Black
Sea (42o50 ' N, 32ø00 ' E. From Lewis and Landing (1991).
must be responsible for this phenomenon. We shall return
to this subject when examining the possible role of the oxygen
minimum in promoting Mn deposition on the adjacent sea
floor.
Dissolved Mn 2+ accumulates in deep sulfidic waters of an-
oxic basins (Spencer and Brewer, 1971). The depth distribu-
tion of dissolved Mn in the Black Sea (Fig. 2) reflects the
reductive dissolution of particulate Mn 4+ oxyhydroxides that
settle from the upper, oxygenated layer into the deeper, sul-
fide-bearing waters. The peak in the concentration of particu-
late Mn lies immediately above the boundary between oxy-
genated and anoxic waters and represents the main locus of
the oxidative precipitation of Mn •+ that diffuses or is mixed
upward from the deeper waters into an oxygenated region of
the water column. Manganese is actively recycled between
oxidized (particulate) and reduced (dissolved) forms across
this redox boundary. This process appears to be catalyzed
microbiologically (Emerson et al., 1982; Tebo et al., 1984;
Tebo and Emerson, 1986; Tebo, 1991).
Manganese in Marine Sediments
A good understanding of the geochemistry of Mn and the
mechanisms controlling the deposition and recycling of Mn
in modem marine sediments stems from the interest in the
mechanisms of formation of ferromanganese nodules that
occur abundantly on the deep-sea floor and on many conti-
nental shelves (see Glasby, 1977; Cronan, 1980). Manganese
occurs at concentrations well above its crustal abundance in
the sediments of the deep ocean (Goldberg and Arrhenius,
1958) and in surface sediments on the continental margin.
Solid-phase MnO2 and MnOOH have been identified as the
solid forms of the fine-grained, poorly ordered material in
such sediments (Grill, 1982; Kalhorn and Emerson, 1984;
Murray et al., 1984). The concentration of the metal in pelagic
sediments (Fig. 3) is due to the slow accumulation rates of
terrestrial and biogenic detritus relative to precipitation from
seawater of authigenic Mn under oxic conditions (Krishnas-
wami, 1976), which is enhanced by the autocatalytic precipita-
tion of oxyhydroxides on preexisting Fe and Mn oxide surfaces
(Morgan, 1967). The very small settling fluxes of organic mat-
ter (Suess, 1980) into the oxygenated bottom waters of the
deep ocean and the low degradability of this material at abys-
sal depths ensure that oxie conditions are maintained to great
depth in the bottom sediments. The predominant form of
Mn in such sediments is therefore Mn 4+ oxyhydroxides.
In continental margin environments and in areas of diver-
genees in the open ocean, where accumulation rates of or-
ganie matter are much higher, oxygen is completely con-
sumed at relatively shallow depths below the sea floor by the
microbial degradation of labile organic matter. Under these
conditions active Mn recycling occurs between surface and
subsurface horizons in a manner analogous to the processes
occurring in the waters of the Black Sea (Fig. 2). The progres-
sive utilization of electron aeeeptors with depth in the sedi-
ment, including Mn oxyhydroxides, during the microbial oxi-
dation of deposited organic matter (Froelieh et al., 1979)
means that anoxie conditions occur at variable depths below
the sea floor (Strakhov, 1972). Burial of surface oxyhydroxides
transports Mn 4+ into an anoxie environment, where it dis-
solves. Mn •+ concentrations are therefore commonly much
higher in the pore waters of such sediments immediately
below the oxyhydroxide horizon (Fig. 4). Dissolved Mn dif-
fuses down into the sediment and up into the overlying oxie
horizon, where it is repreeipitated. This "zone refining" of
Mn (Froelieh et al., 1979) leads to surfieial solid-phase Mn
concentrations in hemipelagie and nearshore sediments that
are often much higher than they are in pelagic sediments,
but sediments that are permanently buried will have much
lower levels of Mn (but see later in this section). Iron is
diagenetieally recycled in a manner similar to Mn, but Fe is
more efficiently removed from anoxie pore waters by forma-
tion of sulfide phases (monosulfide and pyrite). Therefore,
Manganese (Wt %)
0 I 2 3 4 5
5
• l0
ß ßell •
20 ß •1•..•. LL44-GPC3
ß . .%•.' 5705 m
FIG. 3. Distribution of solid-phase Mn in core LLA4-GPC3 from the
North Pacific (30ø1.9.9" N, 157049.9" W). The Cretaceous-Tertiary boundary
in this core lies at a depth of 20.5 m. Manganese is present throughout the
core as an oxyhydroxide and in the upper 10 m is due largely to hydrogenous
(seawater) supply. Increasing hydrothermal influence below this depth has
caused a marked increase in Mn abundance in the middle section of the
core. Suboxic conditions are found below a depth of 21 m, and these may
have caused some Mn remobilization. The changing Mn sources reflect the
movement of the core site northward away from the East Pacific Rise during
the Tertiary. From Kyte et al. (1993).
Mn GEOCHEMISTRY: IMPLICATIONS FOR BLACK SHALES 39
Pore Water Mn (gM)
0 5 10 15 20
10 ß
•, 20
50 Core 14GC1
4170 m
0 1000 2000 3000 4000
Solid Phase Mn (pprn)
FIG.. 4. Distribution of pore-water and solid-phase Mn in a gravity core
from the equatorial Atlantic (0ø00.1 ' S, 12019.3 ' W). The solid-phase Mn
maximum is produced by the precipitation of oxyhydroxides from upward-
diffusing dissolved Mn from the underlying zone of Mn reduction. From
Froelich (1979) and Froelich et al. (1979).
less of the buried oxide Fe is returned to surface horizons,
and this results in surface oxyhydroxides having Mn concen-
trations much larger than those of Fe (Cheney and Vreden-
burg, 1968).
The ease with which Mn oxyhydroxides are solubilized
from surface sediments is demonstrated by the experimental
results of Balzer (1982), who deployed a sealed chamber onto
the sea floor in a nearshore region of the Baltic Sea and
monitored changes in the composition of waters overlying
the sediment. After the bell jar was emplaced, dissolved oxy-
gen levels in the supernatant water decreased steadily to zero
over a 50-day period, and H2S appeared when the 0.2 had
disappeared completely (Fig. 5). Dissolved Mn increased in
concentration as the 0.2 level fell; it first appeared in the
water in measurable amounts after 16 days, when the dis-
solved oxygen level had fallen by only 50 percent, and there
was a sharp increase in concentration before the 0.2 had com-
pletely disappeared. Manganese continued to increase to a
more or less constant value after 80 days at levels that are
only observed in sediment pore waters. Iron did not appear
in the water until all of the 0.2 had been removed and H2S
was present. The source of the Mn and Fe that appeared in
the supernatant water in this bell jar experiment was the
upper 3 cm of bottom sediment, where the two metals were
originally present as oxyhydroxides which had accumulated
there under oxygenated conditions. Solubility calculations
showed that the concentrations of dissolved Mn .2+ and
CO•- in the supernatant waters were supersaturated with
respect to ideal MnCO3, demonstrating that a large supply
of Mn (from surface oxyhydroxides) is required for a carbon-
ate phase to precipitate in seawaters. The difference in the
solubility behavior of Mn and Fe corroborates the widely
accepted model of the separation of these transition metal
neighbors in the sedimentary cycle (Krauskopf, 1957). The
phenomenon observed experimentally by Balzer (1982) also
occurs naturally in areas of restricted circulation where bot-
tom waters have marked seasonal cycles of dissolved oxygen
levels (Kawana et al., 1980).
In spite of the accumulation of dissolved Mn in the bottom
waters of anoxic basins, the concentrations of solid-phase Mn
in the underlying sediments are invariably very low and reflect
exclusively the aluminosilicate fraction (Fig. 6A). This is be-
cause Mn oxyhydroxides are unstable in anoxic conditions
and the concentrations of dissolved Mn in the bottom waters,
even in the presence of relatively high alkalinity, do not reach
levels that would permit precipitation of an Mn s+ solid phase
(i.e., MnCO3). Thus, as well as the Saanich inlet, the modern
sediments of the Black Sea, the Cariaco trench, and Fram-
varen all have very low bulk Mn concentrations that correlate
closely with the content of A1 (Calvert and Pedersen, 1993).
An apparent exception to this generalization is the Baltic Sea,
where Mn carbonates are found in sediments of several anoxic
basins (Debyser, 1961; Manheim, 1961; Hartmann, 1964;
Suess, 1979). This situation is anomalous, however, because
the sediments of these basins are not in steady state, owing
to two factors. First, the basins have become anoxic only
relatively recently because of the increasing isolation of the
Baltic Sea by postglacial isostatic uplift of the strait separating
it from the North Sea and because the intensity of the pycno-
cline that separates the bottom waters from the near-surface,
oxygenated waters has increased owing to a change in the
fresh-water balance of the Baltic as a whole. This has resulted
in the reductive solution of nodular Mn oxyhydroxides that
occur abundantly over large areas of the Baltic proper
• 80
•D 20
o•
ß ,"• 0
Mn/
:t,.., ...... ,
0 20 40 60 80 100
8O
• 300
• 2o0
12 o 0 20 40 60 80 100
Days
FIG. 5. Temporal changes in the concentrations of dissolved Mn, Fe,
02, and H2S in the bell jar experiment of Balzer (1982). The bell jar (3.14
m 2) was deployed on the sediment surface in Kiel Bight (Baltic Sea) at a
depth of 20 m, and the supernatant water was sampled at intervals through
a sampling port.
40 CALVERT AND PEDERSEN
Solid Phase Manganese (%)
0 1 2 3 4 5
0 •
20
30
0
, ,
Saanich Inlet '•
200ra
I I
I 2 3 4 5
Organic Carbon (%)
Solid Phase Manganese (%)
I 2 3 4 5
q• Jetvia Inlet •.•
•.• 300m •B
0 1 2 3 4 5
Organic Carbon (%)
FIG. 6. Distribution of Mn and organic carbon in sediments from (A)
intermittently anoxic Saanich inlet (48040.4 ' N, 123ø30.15 ' W) and (B) neigh-
boring fully oxygenated Jervis inlet (49ø48.1' N, 124002.2 ' W), British Colum-
bia. In the anoxic basin, Mn concentrations are at crustal abundances, be-
cause there is no mechanism for supplying excess Mn to the surface or the
subsu•ace environment. In the oxygenated environment, Mn concentrations
are high in the subsurface anoxic sediment as well as at the core surface,
because Mn is continually supplied to the subsu•ace levels by burial of the
surface oxyhydroxides; this causes solution of the solid Mn phase and the
production of high pore-water concentrations, which, in turn, lead to the
precipitation of Mn carbonate.
(Grippenberg, 1934; Winterhalter, 1966) and has produced
high levels of dissolved Mn in basin waters. Second, the pore
waters of the recent anoxic sediments are significantly influ-
enced by upward-diffusing Ca from underlying fresh-water
sediments formed in the postglacial lake phase of the Baltic
(Suess, 1979; Manheim, 1982), which, together with the high
levels of dissolved Mn in the basin waters, has had the effect
of increasing the saturation state of pore waters with respect
to a mixed Mn-Ca carbonate phase. Suess (1979) pointed out
that the concentration of authigenic phases in the sediments
of one of the Baltic deeps is so high that they cannot be
regarded as the precipitates of a normal quasi-closed pore-
water system but that they formed as a consequence of a
continuous supply of dissolved solutes, mainly from the un-
derlying sediments of a different facies. Hence, formation of
Mn carbonates in the deep basins of the Baltic should not
be regarded as being typical of permanently anoxic basins in
general.
In contrast to the lack of Mn concentrations in sediments
accumulating beneath anoxic waters, Mn '•+ carbonates com-
monly form in subsurface anoxic sediments on continental
margins or in nearshore basins (Calvert and Price, 1970; Ped-
ersen and Price, 1982; Hamilton-Taylor and Price, 1983;
Shimmield and Price, 1986). Such authigenic mineral phases
form because very high concentrations of dissolved Mn are
generated in the pore waters of the anoxic horizons as a
consequence of the dissolution of Mn oxyhydroxides that are
continually supplied from the overlying oxic horizon by burial
(Fig. 4). Consequently, nearshore anoxic sediments that accu-
mulate beneath oxygenated bottom waters have significantly
higher solid-phase Mn concentrations compared with sedi-
ments of anoxic basins (Fig. 6B). Macroscopic Mn carbonate
concretions have occasionally been recovered from such sedi-
ments (Calvert and Price, 1970; Suess, 1979; Pealersen and
Price, 1982). Where the abundance of such authigenie ear-
bonates is below detection limits, using, for example, X-ray
diffraetometry, the solid-phase concentration of Mn is sig-
nificantly higher than its crustal abundance and is positively
correlated with the amount of carbonate carbon present
(Doff, 1969; Hamilton-Taylor and Price, 1983). Moreover,
pore-water profiles from such sediments showy clear evidence
of Mn removal within the sediments, because dissolved Mn
concentrations decrease below the maximum values that lie
immediately below the surface oxie horizon (see Fig. 4; Li et
al., 1969; Calvert and Price, 1972; Pedersen and Price, 1982;
Sawlan and Murray, 1983). Boyle (1983) has also ascribed
the presence of high Mn/Ca ratios of foraminifera in some
sediment horizons which lie deeper than the Mn oxide reduc-
tion zone to the formation of diagenetic Mn overgrowths. A
number of diagenetic models of the distribution of solid- and
aqueous-phase Mn in marine and lacustrine sediments have
been developed to account for these distributions and pro-
cesses (Holdren etal., 1975; Robbins and Callender, 1975;
Burdige and Gieskes, 1983). The carbonate is invariably a
mixed Mn-Ca phase rather than pure rhodochrosite, and
some occurrences have compositions that are close to ideal
kutnohorite (Mn,Ca(CO3)2; Calvert and Price, 1977).
The diagenetic origin of some Mn carbonates in modem
sediments is supported by their carbon isotope composition.
An Mn-Ca carbonate concretion from Loch Fyne, Scotland,
had a 613C value of -5.75 per mil (Calvert and Price, 1970),
whereas a manganoan calcite from one of the Baltic basins
had a value of -13 per mil (Suess, 1979). Such light
values indicate that the carbon in the carbonate lattice was
derived partly from organic matter in addition to normal sea-
water (Irwin et al., 1977). The formation of these phases
therefore occurred within the sediment during degradation
of organic matter at depths where the composition of the pore
waters was heavily influenced by organic matter diagenesis. In
contrast to the results from Loch Fyne and the Baltic Sea,
Pedersen and Price (1982) reported that a mixed Ca-Mn-Mg
carbonate from a Panama basin core had a 5•3C value of 2.59
per mil, not very different from the value of +1 to -1.5
per mil in benthic foraminifera from the same basin; they
concluded that carbon must have been supplied to the grow-
ing Mn carbonate by diffusion from bottom waters or from
dissolving calcite tests.
Liet al. (1969) were the first to show that the pore waters
of a modem marine sediment from the Arctic basin had con-
centrations of dissolved Mn and carbonate that were roughly
in equilibrium with ideal rhodoehrosite. Similar estimates by
Pealersen and Price (1982) demonstrated that pore water-
dissolved Mn concentrations were roughly in equilibrium
with MnCO3 in sediments from the Panama basin •vhere
macroscopic Mn carbonate concretions were recovered. The
pore-water Mn concentrations measured in that study (120-
180 •M) were at least five orders of magnitude higher than
oxie bottom-water values.
Middelburg et al. (1987) reviewed the thermodynamic the-
ory of precipitation of manganese-calcium carbonates and
concluded that there are three stability fields for such phases.
The first corresponds to equilibrium with Mn calcite, where
Mn GEOCHEMISTRY.. IMPLICATIONS FOR BLACK SHALES 41
the dissolved Mn concentration is controlled by chemisorp-
tion and eopreeipitation on calcite surfaces (as discussed by
Franklin and Morse, 1983). The second and third are charac-
terized by equilibria between the dissolved Mff 2+ in pore
waters and either kutnohorite (where the [Mn2+]/[Ca '2+] ratio
is between about 0.0015 and 0.0085) and ealeie rhodoehro-
sites (where [Mn2+]/[Ca '2+] > 0.0085). Thus, in carbonate-
rich deposits, discrete precipitates of kutnohorite are unlikely
because sorption and/or eopreeipitation keep the Mff 2+ con-
centration too low to allow such formation, but in carbonate-
poor sediments kutnohorite or ealeie rhodoehrosite are the
probable phases that control manganese solubility. Further-
more, Middelburg et al. (1987) used the approach of Lipp-
mann (1980) to point out that rhodoehrosite (MnCOa) cannot
exist as a pure phase in modem marine sediments because
the range of aqueous mole fractions observed is not thermo-
dynamically consistent with solutions that would be in equilib-
rium with this phase.
Empirical studies by Mueei (1988) support the conclusion
that a kutnohoritelike phase ("pseudokutnohorite") primarily
controls manganese solubility in pore waters, and this reflects
stoiehiometrie solubility constants that show that rhodoehro-
site (pK = 8.49 at 25øC; Johnson, 1982) is much more soluble
than kutnohorite in seawater (pK = ca. 15.2; A. Mueei,
McGill University, oral eommun.). The compositions of the
Mn-Ca authigenie carbonates found in sediments by a num-
ber of workers are consistent with this predicted thermody-
namic behavior (see Calvert and Price, 1977; Pedersen and
Price, 1982).
The solubility product of disordered (poorly crystalline)
kutnohorite at 25øC is about 6 x 10 -•6, whereas that of its
ordered counterpart is some two orders of magnitude lower
(A. Mueei, oral eommun.). In order to saturate normal sea-
water with respect to the disordered phase, the dissolved
Mn 2+ concentration would need to be on the order of 48
mM (at STP and pH 8), assuming that [Ca •2+] = 10 -.2 M;
that [COl-] = 2 x 10 -4 M; and that the total ion activity
coefficient product, Tgtotal(Mn 2+)Tgtotal(Ca •2+)Tg2t2otal(CO• -) is
2.8 X 10 -5 (see Mucci, 1988). Increasing the alkalinity (via
the reduction of sulfate) would reduce this exceptionally
high estimate, but not enough to approach the range of
manganese levels observed in pore waters or seawaters in
nature. In contrast, saturation with respect to the ordered
phase at the specified conditions would require a manganese
concentration of roughly 500/zM, based on Mucci's empiri-
cal solubility measurements. This is still several-fold higher
than that observed in natural pore waters where the concen-
tration of Mn •+ in apparent equilibrium with a kutnohorite-
like phase has been directly observed (e.g., Pedersen and
Price, 1982). The point of these rather crude estimates is
to show that saturation of seawater with respect to either
an ordered or disordered mixed Mn-Ca carbonate appears
to require a dissolved manganese concentration at least four
orders of magnitude higher than has been observed in the
water columns of anoxic basins such as the Black Sea, where
the highest dissolved Mn concentrations lie between 6 and
8/zM (Landing and Lewis, 1991). Apparently, only in sedi-
ments underlying an oxic water column where manganese
oxides are being actively buried and reduced can conditions
• 10 -3
¸
_• 10 "t OM o
'•O SW o
¸
o•,•
• 104 •
10 © 10 -½ 10-s
XxK t ß CB
• spm•
A •
EA ß •
BS•c• PB
CTo H • •M•.LF
I I I I I
10 '? 10 's 10 4 10"* 10 '3 10 '2
Dissolved Manganese (M)
FIG. 7. Relative degree of saturation of seawaters (open circles) and
sediment pore waters, ,Mth respect to ideal MnCOa (squares = sediments
;vith an Mn carbonate phase; solid circles = sediments without an Mn
carbonate phase). Data points: SW = surface seawater ([Mn] from Klinkham-
ruer and Bender, 1980; alkalinity and pH from Skirrow, 1975); OM = oxygen
minimum ([Mn] from Johnson et al., 1992; alkalinity and pH from Skitrow,
1975); CT = Cariaco trench ([Mn] from Bacon et al., 1980; alkalinity and
pH from Zhang and Millero, 1993); BS = Black Sea ([Mn] from Landing
and Lewis, 1991; alkalinity and pH from Goyet et al., 1991); A = Arctic
basin (data from Li et al., 1969); EA = eastern Atlantic (data from Froelich,
1979); C = Manop C (data from Jahnke et al., 1982); H = Manop H, M =
Manop M (data from Emerson et al., 1980; Klinkhammer, 1980a); PB =
Panama basin, LF = Loch Fyne (data from Pedersen and Price, 1982); jI
= Jervis inlet, B.C. (data from Grill, 1978); CB = Chesapeake Bay (data
from Holdren et al., 1975). Dissolved [CO]-] data were derived from avail-
able alkalinity or total CO2 values using the equations in Skirrow (1975) and
ignoring the small (<3%) contribution of borate to total alkalinity. Activity
coefficients are from Pedersen and Price (1982). The diagonal line marked
"K•p" is the apparent solubility product constant for MnCO3 from Grill
(1978).
of supersaturation with respect to a carbonate phase be
reached.
As an illustration of the overriding control on the solubility
of manganous carbonate imposed by the dissolved manga-
nese content, we have plotted in Figure 7 data that summa-
rize the extent of super- or undersaturation of a range of
seawaters and pore waters. The points include both oxygen-
ated and anoxie seawaters as well as modern sediment pore
waters collected from deposits where an Mn carbonate
phase does not occur, and where such a phase has been
identified or where the solid-phase Mn concentrations in
the anoxie part of the sediment column imply that Mn ear-
bonate is present. We have used the solubility product of
MnCOa for comparison because the factors controlling the
precipitation of kutnohorite in sediments, e.g., precursor
carbonate surfaces (Mueei, 1988), inhibitors, etc., are not
well understood. The data show that anoxie seawaters and
those represented by the intense oxygen minimum of the
eastern Pacific are undersaturated with respect to MnCOa.
On the other hand, the intensely reducing subsurface sedi-
ments accumulating under oxygenated bottom waters in
which Mn carbonate is precipitating are supersaturated with
respect to this phase, whereas other sediments lacking authi-
genie Mn carbonate are, as predicted, undersaturated. Fig-
ure 7 also reinforces the conclusion that the principal control
of Mn carbonate precipitation in marine sediments is the
dissolved Mn 2+ concentration (which ranges over at least six
orders of magnitude) rather than the carbonate ion eoneen-
42 CALVERT AND PEDERSEN
tration, which, as suggested by Pedersen and Price (1982),
plays a relatively minor role.
Oxygen Minima
The oxygen minimum in the ocean is a depth interval,
normally lying between a few hundred to roughly 1,000 m
deep, where dissolved oxygen levels are significantly lower
than those in the waters above and below this horizon. The
minimum zone owes its existence to the high rate of oxygen
consumption by organic-rich particulate detritus that settles
from the surface-productive layers of the sea, coupled with
the fact that the waters at intermediate depths circulate some-
what more sluggishly than those influenced by the wind at
sea surface and those deeper waters that are affected by
the vigorous thermohaline circulation (Wyrtki, 1962). The
intensity of the oxygen minimum (the difference between the
in situ O2 concentrations and those at sea surface) varies
greatly within the ocean depending on the intensity of surface
production and the time since the intermediate waters were
last in contact with the atmosphere. Vanishingly small concen-
trations of oxygen are found, for example, in the oxygen mini-
mum of the eastern tropical Pacific (Brandthorst, 1959) and
the Arabian Sea (Wyrtki, 1971; Sen Gupta and Naqvi, 1984),
whereas the concentrations in the North Atlantic are at least
half those at sea surface (Richards and Redfield, 1954; Miller
and Lohmann, 1982).
Dissolved Mn concentrations are often significantly higher
in some oxygen minima compared with shallower and deeper
waters (Fig. 1). As discussed previously, these high concentra-
tions have been ascribed to the reductive dissolution of Mn
oxyhydroxides from particles settling through the water col-
umn or the transport of dissolved Mn that diffuses out of
anoxic shelf and slope sediments (Klinkhammer and Bender,
1980). The recent observations of Johnson et al. (1992) have
effectively ruled out the second explanation, suggesting that
the excess Mn is either released from settling particles (Klink-
hammer and Bender, 1980) or is not as effectively scavenged
from solution onto particles (Johnson et al., 1992) in this zone
of lower redox potentials.
In accordance with the control of the solid phase Mn con-
tent of surface sediments by bottom-water dissolved oxygen
concentrations described previously, Mn contents of surface
sediments collected through modern oxygen minima are uni-
formly very low and roughly the same as crustal abundances
(Fig. 8). In the Arabian Sea, the concentration of Mn in-
creases in sediments collected deeper than the core of the
oxygen minimum zone, most likely because of the enrichment
of oxyhydroxides in more slowly accumulating pelagic sedi-
ments that are found on the lower part of the continental
slope. These results clearly show that in spite of the higher
dissolved Mn concentrations observed in some oxygen min-
ima, they are still far too low to cause Mn carbonate precipita-
tion in surface sediments bathed by the oxygen-deficient wa-
ters. Moreover, there is little evidence for significant Mn
increase in sediments located above the oxygen minima.
Higher rates of sedimentation of aluminosilicate and biogenic
detritus that are characteristic of the shallower oxygenated
areas of the shelf and slope may dilute any Mn oxyhydroxide
that is precipitated from dissolved Mn derived from the oxy-
gen minima in these regions.
MrdA1
0 0.05 0.1 0.15 0
I_•. Sea
3000 ß
ß A
I I I I
0 50 100 150 200 250
Dissolved Oxygen (gM)
MrdA1
0.05 0.1 0.15
ß '•of
2 California
ß
ß
ß ß B
! ß ! ! !
0 50 lOO 150 2oo •
Dissolved Oxygen (•M)
F•c. 8. Depth distribution of MrdA1 weight ratios (to compensate for
variable carbonate and organic matter dilution) in the surface sediments
collected through the well-developed oxygen minima of the eastern Arabian
Sea and the Gulf of California. In the Gulf of California, the rather abrupt
increase in Mn below a depth of 1,800 m is probably due to the supply of
hydrothermal particulate Mn from the spreading center located in the Guay-
mas basin (Von Datum et al., 1985).
Implications for the Geologic Record
The mechanism discussed so far by which modern marine
sediments may be enriched in Mn strongly suggests that high
sedimentary concentrations of solid-phase Mn, in the form
of both oxyhydroxide and Mn carbonate phases, indicate that
the sediments originally accumulated under oxic bottom-wa-
ter conditions. Where high concentrations of Mn oxyhydrox-
ides are found, as in modern, slowly accumulating pelagic
sediments, the surface and interstitial environments are fully
oxygenated and the higher oxidation states of Mn are stable.
Preservation of such oxyhydroxides is possible, of course, only
where the environment remains oxidizing, because a decrease
in the ambient dissolved oxygen to low or zero levels leads
to reduction and solubilization of the solid phases. In the case
of Mn carbonate precipitates, a mechanism is required for
increasing the concentration of dissolved Mn in the sediment
to levels where the solubility product of the solid phase is
reached. Because the dissolved Mn concentrations in perma-
nently anoxic basins or oxygen minima are far too low for this
to occur, such a mechanism apparently does not operate in
the water column. However, where surface sediments accu-
mulate Mn oxyhydroxides, which are subsequently buried,
very high dissolved Mn levels can build up in subsurface
reducing horizons. Such a process (the "manganese pump")
can occur only beneath oxygenated bottom waters, because
only under these conditions are surface Mn oxyhydroxides
stable. The highest concentrations of surficial oxyhydroxides
appear to be found in continental margin settings where the
rapid accumulation of organic matter leads to the complete
consumption of oxygen at relatively shallow depth in the sedi-
ments. Continuous burial then supports active recycling of
buried solid-phase oxyhydroxides and the secondary enrich-
ment of the surface, oxic horizons in Mn because of the
supply of dissolved Mn by diffusion from deeper layers.
Mn GEOCHEMISTRY: IMPLICATIONS FOR BLACK SHALES 43
This reasoning has been applied to the interpretation of
the high solid-phase Mn concentrations in the upper part of
the Holocene sapropel in the Black Sea, which reach 1 per-
cent by weight (Calvert, 1990). The moderu sediments have
Mn levels at crustal abundances, consistent with the lack of
sedimentary Mn accumulation in fully anoxic basins. The high
Mn concentrations in the sapropel signify the presence of an
Mn carbonate phase, because the sediments are highly anoxic
and any Mn that might have been originally deposited as
an oxide would have been dissolved. The occurrence of Mn
carbonate in these sediments therefore implies that this sedi-
mentary unit formed when the deep basin contained oxygen-
ated waters, and this is supported by the concentration of
iodine, another trace element indicator of oxic sedimentation
(Price and Calvert, 1973; Fran9ois, 1987) in the same facies.
This in turn implies that the formation of the sapropel itself
was not caused by the presence of anoxic bottom waters and
supports the suggestion that it was formed as a consequence
of increased organic matter supply to the sediments during
a period of higher plankton production during the transition
from a lake to a marine basin (Calvert et al., 1987; Calvert,
1990; Middelburg et al., 1991).
Manganese carbonates (rhodochrosite and kutnohorite) oc-
cur in many organic-rich shale sequences (Hein and Koski,
1987; Jenkyns, 1988; Xu et al., 1990; Jenkyns et al., 1991;
Fan et al., 1992) and either are present or constitute the
principal Mn-bearing phase in some of the giant Mn ore
deposits (Strakhov, 1965; Kim, 1979; Cannon and Force,
1983; Bolton and Frakes, 1985; Roy, 1988; Maynard et al.,
1990; Okita and Shanks, 1992). In addition, Mn concentra-
tions in some limestone sequences are also inferred to be
caused by the presence of Mn carbonates (Bencini and Turi,
1974; Stamm and Thein, 1982; Pomerol, 1983; Pratt et al.,
1991). As noted previously, the presence of a manganous
phase has often been used to infer euxinic conditions during
the formation of such deposits (Frakes and Bolton, 1984,
1992). However, many of the horizons of Mn enrichment in
shales and carbonates as well as many of the host rocks of
Mn rhodochrosite ores are bioturbated or contain benthonic
fossils (Strakhov, 1965; Stature and Thein, 1982; Force and
Cannon, 1988; Maynard et al., 1990; Pratt et al., 1991), mak-
ing it highly unlikely that the sea floor was anoxic at the time
of deposition. Moreover, our understanding of the marine
geochemistry of Mn strongly suggests that Mn concentrations
significantly exceeding crustal abundances, whether as oxyhy-
droxides or carbonates, can be produced only in sediments
which accumulate under oxygenated bottom waters-in other
words, where the manganese pump operates.
Further evidence for the immediate locus of precipitation
of Mn carbonate in ancient rocks and ores is furnished by
the carbon isotope composition of stratiform sedimentary
rhodochrosites. Values for 8•aC often show that this phase is
highly depleted in •aC; for example, values as light as -7.8
per rail are found in the Jurassic Molango ore deposit (Okita
et al., 1988; Okita and Shanks, 1992), -31 per rail has been
reported in the Jurassic Urkut ore deposit from Hungary
(Polgari et al., 1991), and a value of -37 per rail was mea-
sured in Triassic and Jurassic rhodochrosites from Japan (Mi-
noura et al., 1991; Sugisaki et al., 1991). These very isotopi-
cally depleted Mn carbonates sometimes occur together with
calcium carbonate that has a 6•aC value very close to seawater
values, much closer to 0 per mil. The isotopic composition
of the Mn phases indicates clearly that the source of the
carbon is CO2 derived from organic matter or methane oxida-
tion (Irwin et al., 1977) rather than from ambient bottom-
water carbonate. Hence, the formation of the Mn carbonates
takes place where organic matter degradation is occurring in
an environment removed from the ocean reservoir, namely,
within sediment pore waters. The co-occuring calcites in
these deposits, on the other hand, must have formed from
normal seawater carbonate, probably much closer to the sedi-
ment-water interface. A very similar conclusion was reached
by Okita et al. (1988), Minoura et al. (1991), Polgari et al.
(1991), Sugisaki et al. (1991), and Okita and Shanks (1992).
Okita et al. (1988) also suggested that the Molango Mn car-
bonates formed as a consequence of the reduction of manga-
nese oxides, a process identical to that proposed in this paper.
The formation of some large Mn oxide ore deposits has
been explained by the supply of dissolved Mn from oxygen
minima or from anoxic waters into shallower depths that were
well oxygenated (Cannon and Force, 1983; Frakes and Bol-
ton, 1984). Thus, the Albian Groote Eylandt deposit of Aus-
tralia (Slee, 1980; Frakes and Bolton, 1984; Bolton and
Frakes, 1988), the Cenomanian-Turonian Imini deposit of
Morocco (Force et al., 1986), and the Oligocene Chiatura
deposit of Georgia (Strakhov, 1965; Bolton and Frakes, 1985)
are thought to have formed in relatively shallow shelf environ-
ments that received dissolved Mn from deeper anoxic or oxy-
gen-deficient waters. In each of these deposits, Mn carbon-
ates also occur in organic-rich beds that formed in deeper
waters farther offshore; this has been used to support the
notion that anoxic or oxygen-deficient waters were juxtaposed
with the oxide deposits. Dickens and Owen (1993, 1994)
recently discovered significant depletions of Mn in Cenoman-
ian-Turonian strata from some slope and pelagic settings that
are temporally coincident with the deposition of some, but
not all, large shallow-water Mn ore deposits. They interpret
those findings as a reflection of a more expanded oxygen
minimum or bottom water anoxia at those times, a widely
accepted explanation for the formation of widespread coeval
black shales (Arthur and Schlanger, 1979; Demaison and
Moore, 1980; Jenkyns, 1980; Arthur et al., 1984, 1987; de
Graciansky et al., 1987). In this view, the presence of the
oxygen-depleted waters led to the trapping of dissolved Mn
derived from the solution of surface oxides or the direct en-
trapment of hydrothermal Mn in oxygen-deficient waters and
its subsequent accumulation in oxic shelf sediments lying
above the oxygen-poor water mass. This explanation is consis-
tent with the models of Cannon and Force (1983), Frakes
and Bolton (1984), Force and Cannon (1988), and Force and
Maynard (1991). If such a hypothesis is valid, there should
be sediments depleted in manganese adjacent to other large
oxide ores. Such Mn depletions have not yet been recorded;
the search for such occurrences would constitute a critical
test of the hypothesis. Based on the geochemical behavior of
Mn in moderu sediments discussed in this paper, the forma-
tion of Mn carbonates within the oxygen-deficient zone as
suggested by Pratt et al. (1991) and Dickens and Owen (1993)
would be unlikely. Indeed, the occurrence of Mn carbonates
in deeper water facies of the Groote Eylandt, Imini, and
44 CALVERT AND PEDERSEN
Chiatura deposits is more plausibly explained by their diage-
netic precipitation in sediments accumulating in offshore en-
vironments that were oxygenated at the sea floor but anoxic
below a variable depth. Such an environment would be con-
ducive to the formation of an organic-rich deposit because
of the higher productivity of continental margin settings (Ber-
ger et al., 1989), the high organic matter flux being ultimately
responsible for the reducing interstitial environment of the
sediments (Calvert, 1987).
Conclusions
Based on the evidence presented in this paper, we suggest
that the presence of Mn carbonate in marine sediments can
be used as an indicator of sedimentation under oxygenated
bottom-water conditions. The association between Mn car-
bonate and organic-rich sediments, both in the modern ocean
(Calvert and Price, 1977) and in the geologic record (Hein
and Koski, 1987; Jenkyns, 1988; Xu et al., 1990; Jenkyns et
al., 1991; Fan et al., 1992), is consistent with this hypothesis,
because a high rate of accumulation of organic matter will
promote the development of subsurface anoxic conditions in
the sediments, which in turn will lead to the active redox
recycling of Mn. This will cause the surface sediment to be
enriched in Mn oxyhydroxides, which, when buried, will pro-
duce high pore-water concentrations of dissolved Mn. This
diagenetic manganese pump produces supersaturation of the
pore fluids with respect to Mn carbonate and consequent
precipitation of Ca-Mn carbonates.
The fact that the occurrence of Mn carbonate in organic-
rich rocks indicates original sedimentation under oxygenated
conditions lends further support to the suggestion of Calvert
(1987), Pedersen and Calvert (1990), and Calvert and Pe-
dersen (1992) that the organic content of marine sediments
is primarily controlled by the supply of organic matter to the
depositional site and not by the oxygen content of the bottom
water. Indeed, organic-rich sediments in the modern ocean
are restricted to the margins of the ocean basins, especially
in nearshore basins and on the upper parts of the continental
slope, palcogeographic settings that have also been proposed
for many of the large Mn carbonate ore bodies (Strakhov,
1965; Maynard, 1983; Bolton and Frakes, 1985; Maynard et
al., 1990).
Acknowledgments
We thank Alfonso Mucci for very informative discussions
on the thermodynamics of Mn carbonate precipitation in sea-
water, Eric Force and Jack Dymond (who suggested the prep-
aration of Fig. 7) for constructive criticism of an early draft
of this paper, and two Economic Geology referees for formal
reviews. Support for this research was provided by the Natu-
ral Sciences and Engineering Research Council of Canada
and the Canada Council.
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