0016-7037 79,0701-1075SO?.M) 0
Early oxidation of organic matter in pelagic sediments of
the eastern equatorial Atlantic: suhoxic diagenesis
P. N. FROELICH,* G. P. KLINKHAMMER,M. L. BENDER, N. A. LUEDTKE, G. R. HEATH,t
DOUG CULLEN and PAUL DAUPHIN
Graduate School of Oceanography,
University of Rhode Island. Kingston. RI 02881, U.S.A.
DOUG HAMMONDand BLAYNE HARTMAN
Department of Geology, University of Southern California, Los Angeles, CA 9ooo7, U.S.A.
and
VAL MAYNARD
Department of Marine Sciences. University
(Received 11 August
of South
Florida,
St. Petersburg,
1978; accepted in reaisedform
FL 33701, U.S.A.
7 March 1978)
Abstract-Pore
water profiles of total-CO,,
pH, PO:-, NO; plus NO;, SO:-, S2-, Fe’+ and Mn*+
have been obtained
in cores from pelagic sediments of the eastern equatorial Atlantic under waters
of moderate to high productivity. These profiles reveal that oxidants are consumed in order of decreasing energy production
per mole of organic carbon oxidized (0, > manganese
oxides 5 nitrate > iron
oxides > sulfate). Total CO1 concentrations
reflect organic regeneration
and calcite dissolution.
Phosphate profiles are consistent
with organic regeneration
and with the effects of release and uptake
during inorganic
reactions.
Nitrate profiles reflect organic regeneration
and nitrate reduction,
while
dissolved iron and manganese
profiles suggest reduction
of the solid oxide phases. upward fluxes of
dissolved metals and subsequent
entrapment
in the sediment column. Sulfate values are constant and
sulfide is absent, reflecting the absence of strongly anoxic conditions.
INTRODUCTION
THE DIAGENESISof marine organic matter in anoxic
basins and sediments has been extensively studied.
largely to gain a better understanding
of reactions
occurring during oxidation
of organics by sulfate,
with emphasis on the consumption
of sulfate and the
generation of CO,, CH4, alkalinity, PO:- and NH:
and several
metals
(RITTENBERG et al., 1955;
RICHARDS and VAC~ARO, 1956; BERNER et al., 1970;
BISCHOFF and Ku, 1971; NIS~ENBAUMet al., 1972;
GAINES and PILSON, 1972; SHOLKOVITZ, 1973; HARTMAN et (I/.. 1973; MARTENS and BERNER. 1974; KAPLAN, 1974; REEBURGH, 1976; BARNES and GOLDBERG,
1976; MARTENS and BERNER, 1977; GOLDHABERrt al.,
1977; and others). In comparison, less effort has been
applied to studying pore water profiles resulting from
the early oxidation of organic matter in the upper
sediment column of pelagic sediments overlain by
oxygenated waters (LYNN and BONATTI, 1965; PRESLEY er al.. 1967; LI et al., 1969; BISCHOFF and Ku,
1970; BENDER et al., 1977). This paper presents pore
water data from pelagic cores collected in the eastern
equatorial Atlantic; we will use these data to elucidate
the chemical processes associated with oxidation of
organic matter in these sediments
and the fates of
* Present address: Dept of Oceanography,
Florida State
University Tallahassee,
FL 32306, U.S.A.
t Present address:
School of Oceanography.
Oregon
State University, Corvalis, OR 97.730. U.S.A.
nutrients, gases and metals released and consumed
during organic diagenesis. In this sediment, water
exchange is sufficiently slow that oxidants are more
depleted than in the overlying water column, but the
organic content is low enough that the various oxidants are consumed over a large depth range. As a
result, the observed concentrations
provide information from which the sequence
inferred.
SEQUENCE
OF
of reactions
may be
OXIDATION
The simplest plausible model of organic diagenesis
is one in which marine organic matter having the
Redfield composition
(FLEMING, 1940; REDFIELD,
1958) [i.e. (CH20)10e (NH3)16 (H,PO,)]
is oxidized
by the oxidant yielding the greatest
free energy
change per mol of organic carbon oxidized (McKn+
NEY and CONWAY, 1957; RICHARDS, 1965; CLAYPOOL
and KAPLAN, 1974; and others). When this oxidant
is depleted, oxidation will proceed utilizing the next
most efficient (i.e. most energy producing) oxidant,
and so on until either all oxidants are consumed or
oxidizable organic matter is depleted. The predicted
sequence of reactions is outlined in Table 1. Note
that the sequence is similar to that of SILL~N (1965),
in which the primary oxidants present in the earth’s
ocean. atmosphere
and crust are titrated with Hz
(also, see STUMMand MORGAN, 1970, and MECHALAS.
1974). In marine sediments,
the primary reduced
1075
I076
P. N. FKOI~LICH t'rtrl.
Table
1.
1. Oxidation
(O1201106 ‘“HjIIb
reactions
of sedimentary
(H3PO4) + 138 02 ___f
organic
matter
106 CO2 + 16 HN03 + H3PO4 + 122 HZ0
AGo' = -3190 kJ/mole of Slucose
2.
(CH20)lo6 (W3)16
(H3PO4) + 236 t&O, + 472 H++
i
236 h2+
+ 106 CO
2
+ B N2 + H3P04 + 366 H20
b,GQ'E -3090 kJ/mole (BIRNESSITE)
-3050 kJ/mole (NSUTITE)
-2920 kJ/mole (PYROLLISITE)
3a.
(M20jlo6
(NH3j16 (HjP04)
l
94.4 HN03 +
106 CO2 + 55.2 N2 + H3P04
+ 177.2 HZ0
AGo' = -3030 kJ/mole
3b.
(CH20)106 (NH3)16 (H3P04) + 84.8 HNO3 4
106 CO2 + 42.4 Y2 + 16 NH
3
+ H3P04 + 148.4 H20
AGo' = -2750 kJ/mole
4.
(o120)106 vH3)16 (H3P04) + 212 Fe,03 (or 424 F&OH)
+ 848 H+ W
1
424 Fe-+ + 106 CO2 + 16 NH3 + HjPOI + 530 H,O (or 74: H,O)
AGo' E -1110 kJ/mole (HEMATITE, FeZOjj
-1330 kJ/mole (LIMINITIC GOETHITE, F&OH)
5.
(CH20)106 (W3)l6
(H3P04) + 53 SOi- 3
106 CO2 + 16 NH; + 53 S2- + HjP04
+ lob H20
AG"' = -380 kJ/mole
6.
(ctizo)106wj)16
(H3P04) +
53 CO, + 53 cH4 + 16 NH3 + HjPOl
AG"' = -350 kJ/mole
Standard
free energies were taken from LATIMER(1952). POURBAIX(1966). WHITE
tv a/. (1968), STUMM and MORGAN (1970). BERNER(1971) and LEHNINGER(1975). ‘CH,O
was taken to represent organic matter having an average oxidation state corresponding
to glucose. Nitrogen
was calculated
as if it all starts as R-NH,,
a primary amine.
Phosphorus
is calculated as if it all starts as glucose-l-phosphate.
The AC”’ are presented as kilojoules per mole of glucose (C,H120,).
Multiplication by the factor 17.67
will convert the? values to kJ per ‘Redfield-molecule’-mole. Gibbs free energy changes
were calculated at standard conditions
and at the biochemical
reference state (pH 7).
species being oxidized is organic carbon. In a closed
system (for example see BENDER et al., 1977), reaction
1 continues until sufficient oxygen has been consumed
to drive the redox potential low enough to favor the
next most efficient oxidant. If nitrate reduction occurs
by reaction 3a (all nitrogen goes to NJ then the free
energy changes of reactions 2 and 3 overlap (depending on the MnO, phase being reduced), so that we
might expect nitrate and MnO, reduction to occur
almost simultaneously. Alternatively, if nitrate reduction occurs by reaction
3b (organic nitrogen
is
released as ammonia rather than being oxidized to
N2), then MnO, is apparently reduced before nitrate.
After consumption
of nitrate and labile Mn02, oxidation continues by iron oxide reduction (reaction 4).
sulfate reduction (reaction 5). and disproportionation
(reaction 6). Reaction 1 occurs in oxic environments,
reactions 2-1 in suboxic environments,
and reactions
5 and 6 in anoxic environments.
In writing the reactions in Table I and applying them to the subsequent
discussion, we have assumed that (1) organic matter
is represented
by the Redfield
formulation;
(2)
ammonia released during reaction 1 is quantitatively
oxidized to nitrate, all organic nitrogen oxidized in
1077
Suboxic diagenesis
reactions 2 and 3 is converted to N, (i.e. ignoring
the possibility of reaction 3b), and all ammonia
released in subsequent reactions remains unoxidized;
(3) 02, NO;, FelO, or FeOOH, MnO,, and SO:are the only electron-acceptors in marine sediments;
and (4) organic matter is the only electron-donor.
Several manganese and iron oxides have been included in Table 1 to demonstrate that the predicted
sequence of reactions is similar for several reasonable
Mn(IV) or Fe(III) mineral phases. However, it is likely
that the redox-active metal phases in marine sediments are the oxyhydroxides, which constitute only
a fraction of the total. These phases are poorly characterized chemically and thermodynamically.
hd+:;
QLM)
po,3-
o
A
3
(pM)
E*
STOICHIOMETRY OF OXIDATION
If we initiate the reactions in Table 1 in a closed
system starting with excess organic carbon, and with
dissolved constituent concentrations equal to North
Atlantic Deep Water values (taken as average bottom
water concentrations at our coring sites), then we can
calculate the resultant metabolite, oxidant and reductant concentrations as a function of metabolic CO2
produced. In doing this we expand the model of
BENDER et al. (1977) to include MnO, and Fe,O,
(or FeOOH) reduction. We will start with concentrations equal to those of bottom water at our coring
sites, given in column A of Table 2 (see Fig 1, point
A). We will assume that the oxidants are limiting,
and that each reaction proceeds to completion before
the next commences. We will further assume that the
system is in equilibrium with sedimentary calcite.
Since [Ca’+] in the pore waters is not expected to
vary appreciably before the onset of extensive sulfate
diagnenesis, we have the constraint that [CO:-] must
remain constant at a value of K,,‘/[Ca”].
Assuming
that the variation of borate alkalinity during early
diagenesis is small (i.e. that pH is fairly constant),
and neglecting changes in the CO2 and H2C03 concentrations, we can write these three equations
(BROECKER,1974):
ATC02 = A[HCO;]
ATA = A[HCO;]
+ A[CO:-1
(7)
+ 2A[CO:-]
(8)
METABOLIC-
CO,
PRODUCTION,
pmoles
Fig. 1. Predicted changes in concentrations of dissolved
constituents in a closed system during oxidation of organic
matter. Line segments A-B, EC, C-D, and D-E correspond to reactions 10, 11, 12 and 13, respectively. Metabob&O2 production corresponds to CO, produced from
CH,O oxidation alone.
A[CO:-]
= ATA - ATCO, = 0.
where TCOz is total-CO,
(9)
and TA is total-alkalinity.
Equation 9 provides us with the conditions approximating calcite saturation during oxidation reactions.
[An analysis of the sensitivity of our assumptions
(neglecting ACa’+, ACOZ, ApH and ARA) over the
TC02 range from 2.185 to 2.700 mM indicated that
a cumulative error of less than 5% is introduced.]
In the following discussion, ATCOI and ATA
values are in units of moles or equivalents per
106 mol of organic carbon oxidized. Temporarily
omitting N, CH20 will be oxidized by O2 q follows
(see reaction 1, Table 1):
(CH20)ie6 + 106 O2 -+ 106 CO2
Table 2. Metabolite and oxidant concentrations during
oxidation reactions in a closed system
r\
rce*> RH
02 I UM
NO; , UH
2.185
s
D
c
E
2.598
2.5i6
2.640
2.609
250
0
0
0
0
22
51
51
0
0
PO:-, UM
1.45
3.26
&x2*, "M
0
0
20
20
20
0
0
0
0
20
3.34
3.83
3.93
;;;I; ',"
23.3
%
.YM
0
A = initial bottom
23.3
0
23.3
0
28.8
0
28.8
0.6
water concentrations, B = after
O1 reduction, C = after MnO, reduction, D = after
NO; reduction, E = after FeOOH reduction.
+ 106 H20; ATCO* = + 106, ATA = 0.
If ATA = ATCOI, then about one mol of CaCO,
must dissolve for every mol of CO2 produced according to this summed reaction:
(CH20)io6 + 106 O2 + 106 CaC03
-+ 106 Ca2+ + 212 HCO;.
This yields ATA = +212, and ATCOI = ACHCO;]
= +212, in agreement with eqn (9).
Now considering
organic nitrogen
oxidation
(BREWER and GOLDMAN, 1976)
(NHs)te
+ 320*+16NO;
+ 16HsO
+ 16H+
excess protons
released
16CaC0,
+ 16H++
so that the net reaction
(NH3hh
reaction
will react with CaCO,
16Ca2+ + 16HCOj
(CH,O),,,
(NH,),,
+ ll.6CaC03-+55.2N2
+ 11.6Ca’” + H,PO,
yields
(H,PO,)
+ 94.4 NO,
+ 117.6HC0,
+ 71.2 H,O
(12)
+ 32 O2 + I6 CaCO,
+l6NO,
+ 16Ca”
+ l6HC0,
ATA = + 117.6: and ATCO? = + 117.6.
+ l6H,O
where ATA = + 16, and ATCOz = AHCO, = + 16.
The net effect of oxygen reduction in the oxidation
of organic matter is thus:
(CH,O),,,
(NH&i6 (H,POJ
+ 138 0,
+ 122 CaCO, -+ 228 HCO;
+ 122 Ca2+ + 16 H,O
+ 16 NO;
+ H,PO,
(IO)
ATA = +228
and
ATCO, = AHCOJ
where
= +228.
In oxic sediments, reaction 10 consumes 250gmol
oxygen, producing 413 prnol of CO,. I .8l pmol of
PO:-, and 29 pmol of NO,. The values of dissolved
constituents after completion of reaction 10 are given
in column B, Table 2 (Fig. 1. point B).
After oxygen has been consumed, MnO, reduction
commences
(reaction 2). The reaction for Mn02
reduction
(assuming
calcite equilibrium)
can be
handled in a similar manner:
(CH,O),,,
(NH3)i6 (H,PO,)
+ 366 Ca2+ + 260 HCO;
ATA = - 260. and ATC02
The net effect
+ 2Mn0,
+ 260 H,O
= AHCO;
is to precipitate
because alkalinity production
during MnO, reduction
+ 236 MnO,
+ 366 CaCO,
+ 236 Mn2+ + 8 N, + H,PO,
CH,O
3a for calcite equilibrium:
(11)
= -260.
CaCO,.
primarily
(i.e. acid consumption)
+ 4H+
-CO*
+ 2Mn”
+ 3Hz0
requires consumption
of TC02 (i.e. calcite precipitation). This reaction may go until Mn2’ in the pore
waters is at saturation with respect to rhodochrosite
(LYNN and BONATTI, 1965; LI et al.. 1969) or it may
cease earlier if the kinetics of Mn(IV) reduction are
too slow to drive the MnZf concentration
to this
point. Rather than try to estimate the Mn2+ concentration (I priori, we adopt the value of 20 PM for the
in our pore waters
asymptotic Mn2’ concentration
as a guide to the amount of Mn2+ dissolved. Production of 20 pmol Mn2+ would then coincide with production of 0.08 icmol of PO:-.
and 31 pmol of
while consuming
31 pmol of Ca2+ and
CaCO,.
22 pmol of TC02 (HCO;). The final concentrations
after completion
are given in column C, Table 2
(Fig. 1, point C). Note that the changes in TCOz and
PO:- are barely detectable, and the CaCO, precipitated is totally insignificant, while the manganese concentration increases dramatically.
At some point during or after MnO, reduction,
nitrate reduction commences (reaction 3). Rewriting
Reducing 5 1 pmol of NO; would dissolve 6 pmol of
CaCO,
and oxidize 57pmoI of CH,O,
releasing
64~cmol of TCO, (HCO;).
6pmol of Ca2+ and
0.54 pmol of PO:
The final concentrations
are given
in column D. Table 2 (Fig. I. point D).
oxidation
of organic
After NO3 consumption.
matter proceeds by Fe III reduction
(reaction 4).
Writing the reaction with calcite equilibrium:
(CH,O),,,
+ 758 Ca’_
+ 16NH(
(NH,),,
(H,POJ
+ 652 HCOj
+ H,P04
+ 424 FeOOH
--+ 758 CaCO,
t 424Fe”
+ 636H,O
(13)
ATA = ATCO? = -652.
Here CaCO, is precipitated,
consuming
TC02 to
balance the TA gain by proton consumption
upon
FeOOH reduction. Based on the data from cores
4GCl and 5GCl. in vvhich we seem to be approaching
we again adopt
the asymptotic Fc2 ’ concentration.
the value of 20 HIM for the asymptotic value. Production of 20 pmol of Fe’ * coincides with production
of 0.05 ilrnol of PO:
precipitation
of 35 pm01 of
CaCO,. and consumption
of 36,umol of Ca”
and
31 pmol of TCO, (HCO, ). As during MnO, reduction, the TC02 decrease and PO:increase are
barely detectable. the CaC03 precipitated
is totally
insignificant. and the Fe 2+ increase is very large and
easily measurable. Final concentrations
after completion of reaction 13 are given in column E, Table 2
(Fig. I. point E).
After Fe(Il1) reduction.
organic diagenesis continues by sulfate reduction (reaction 5). which is not
plotted on Fig. I. On the same scale, starting with
28.8 mM - SOi- ~ the horizontal axis would have to
be extended IO-fold to include SOi- diagenesis.
This discussion points out the effect of calcite equilibrium on TC02 concentrations
during organic diagenesis. It also demonstrates
the influence of metal
diagenesis on the TCO, and TA balances in calcareous marine sediments. In the sediments we studied.
pore water Fe and Mn concentrations
do not exceed
about 50/1M. The corresponding
changes in TCO,
and TA of about 50 60 [cM or lieq I ’ are less than
the scatter of our TCOz data, but it is certainly detectable provided good samples can be collected. In
sediments
where metal diagenesis
is more pronounced. the effect on the CO,-system
should be
greater (compared with simple organic regeneration).
The above discussion is based on a first-order understanding of the reactions occurring in marine sediments. Nevertheless.
it provides us with a starting
point for describing pore fluid chemistry in pelagic
sediments. In principle. marine sediments are not
Suboxic
closed systems, but are open to molecular diffusion.
However, for the first-order approach we take here,
diffusion can be neglected since diffusivities of ions
do not differ from one another by more than about
a factor of two. Gases, however, diffuse much faster,
so that a real system is more open with respect to
O2 than, for example, HCO;. Discrepancies
between
our predicted and observed TC02 values may be
partly due to this diffusivity difference.
If we take the horizontal axis in Fig. 1 as representing depth and age in the sediment, then the primary
changes we expect as we go downcore are (1) an increase in the NO;
concentration
to a maximum
value, (2) an increase in dissolved Mn’+ and a decrease in NO; to zero, (3) an increase in dissolved
Fe* +, and (4) fairly monotonic increases in TC02 and
PO:-. The pore water data we will present here are
consistent with this picture.
1079
diagenesis
UNWASHED FILTERS
CaCOs - RICH SEDIMENT
I-III
0
PROCEDURES
Three types of cores were recovered for pore water analyses. A modified Bent/m gravity corer with 25in. dia.
liners without core catcher, cutter, or pipe barrel was used
to collect gravity cores (labelled GC) and trigger-weight
cores at piston core sites (labelled TW). An 8 in. i.d.
sphincter corer (BURKE, 1968) was used at one site (labelled
24SC). This core was subsampled
immediately
after retrieval by pushing three 2iin. i.d. core liner pipes into
the mud. Two of these were treated as duplicate cores for
pore water analysis (24SCl and 24SC2).
Supernatant
seawater
was gravity siphoned
from the
tops of cores, which were then immediately
placed upright
in a reefer at in situ temperatures.
Extruding,
slicing and
squeezing generally commenced
within 6 hr of collection.
Two types of squeezers were used. Most samples were
squeezed with REEBURGH (1967) type nylon squeezers provided by Kent Fanning.
A USC squeezer
(KALIL and
GOLDHABER, 1973) was used for some gas analyses.
Cores were extruded,
sliced and sealed into squeezers
under a positive-pressure
helium atmosphere
inside two
interlocked glove bags in a walk-in reefer at 4°C. Two-centimeter thick sediment cakes were extruded from the core
liners, sliced. and placed on filters in the squeezers. Nuclepore (0.45 pm) and Whatman 42 filters were supported
by
Nitex screens. Cakes were covered with a polyethylene
disc
and tamped down to exclude as much helium as possible.
Squeezer tops were then screwed on tightly. The squeezers
were transferred
to a thermostated
refrigerator.
Pore
waters were expressed at pressures increasing
from 5 to
80 psig.
Filters for all stations except 16 and 23 were washed
in redistilled HCI and dried. For stations 16 and 23, alternate samples were squeezed using acid-washed
and unwashed filters because it was discovered that washed filters
apparently
contain an acid residue that dissolved CaCO,
and increased TC02 by 1 mM or more.
Samples analyzed at sea were stored refrigerated
until
they were drawn by the analysts.
Trace metal samples
(acidified to pH 2 with redistilled HCI) and sulfate samples
were stored in l-ml plastic vials.
Anulyticd
techniques
A total of ID30 ml of pore water was obtained
from
each 2-cm slice of sediment. This volume was collected
as several 1-3-ml sequential squeezing aliquots which were
labeled with a sequential aliquot number or, where possible, with the cumulative volume expressed. These sequential aliquots
were analyzed
individually
in an effort to
I
4
._
III
6
I
a
III
III
IO
12
14
ACID-WASHED
FILTERS
CaC03- RICH SEDIMENT
0
SAMPLING
2
2
4
CUMULATIVE
6
a
SQUEEZING
12
IO
VOLUME,
ml
14
Fig. 2. Variations
in TCOz concentrations
as a function
of water
volume
expressed
from individual
samples
(sequential squeezing aliquots).
evaluate the integrity of the squeezed samples. Since many
of the analyses were done on-board
within 24 hr, we were
able to continuously
adjust our procedures
to eliminate
apparent
problems, As a result .the data from the later
cores are of higher quality than those from the beginning
of the cruise. We therefore established criteria for questioning and discarding
data based on internal
arguments
within the data set itself. All data have been screened
according
to these criteria. After this process was completed, the remaining
values of sequential squeezings were
averaged and recorded as a single value for each depth.
Total CO2 was measured immediately
by acidifying an
aliquot of sample (5 1 cm3), extracting CO2 with a helium
carrier
in a Swinnerton
stripper,
passing
the effluent
through columns of indicating
Drierite and of silica gel
to separate air and C02, and through a thermistor
detector. An aliquot of Coleman Grade COz (99.9% purity) was
ADJACENT
SAMPLES
LAST SEQUENTIAL
OUT
-._
2.0
3.0
TCO,,
USC
SQUEEZER
4.0
mM
(unwashed
filter)
Fig. 3. Comparison
of TCOz values obtained by squeezing
adjacent samples (presumbably
with identical original pore
water TCOz concentrations)
with two different squeezers.
1080
P. N.
FROELICH er ul.
ACID- WASHED AND
UNWASHED
FILTERS
15
.
.
n
2
I
SEQUENTIAL
3
SQUEEZE
4
5
6
ALIOUOT
Fig. 4. Variations in phosphate concentrations as a function of water volume expressed from individual samples
(sequential squeezing aliquots). The first sequential
squeezed through acid-washed filters has been omitted.
injected through a gas sampling valve after every fifth
sample for calibration. The detector output was monitored
after every fifth sample for calibration. The detector output
was monitored with a strip chart recorder equipped with
an integrator. The area under the CO, peak was monitored for 6 min after the peak appeared and corrected for
drift to obtain the signal for both standards and samples.
Pore waters expressed with Fanning squeezers suffered loss
of both N2 and CO1 due to stripping by He inside the
squeezer. Thus TC02 values determined on water expelled
through “unwashed” filters at the beginning of the squeezing process are suspected of being 10% + 10% low (Fig. 2).
Gas loss in samples squeezed through USC squeezers
was less than in water squeezed through Fannning
squeezers, because there was far less free air (He) in the
USC squeezers. As a result, TCO, concentrations in the
last aliquot of pore water expelled from Fanning squeezers
is about 7% less than the TC02 in the last aliquot expelled
from an adjacent sample with the USC squeezer (Fig. 3).
Acid washed filters contained an acid residue that, upon
squeezing, released H+ in exchange for cations and dis-
r
solved CaCO,, particularly in calcareous sediment (Fig. 2).
This caused TCO, values for acid-washed filters to be too
high by as much as 1 mM on the first few milliliters
squeezed. Where TCO, varied between 2_6mM, values
determined on the last aliquot squeezed through unwashed
filters were accepted as correct, and values determined on
the last aliquot squeezed through washed filters were
accepted as upper limits.
Reactive phosphate was determined within 24 hr after
sample collection with a Technicon AutoAnalyzer (Ind.
method 155-71/w) using standard techniques (MURPHY
and RILEY, 1962). Absorbances were measured in S-cm
flowcells at 885nm. Standards were prepared in surface
seawater, and reagent blanks in deionized water. Reactive
phosphate samples were anomalously high in those
sequentials sampled for TC02, evidently due to contamination from phosphoric acid used in the TCO, analyses.
Therefore all phosphate data from which aliquots were
drawn for TCO, analysis have been discarded. In addition.
all phosphate analyses on the first sequentials squeezed
through acid-washed filters were discarded because acid
released from filters dissolved solid phosphate. The remaining data showed only minor sequential squeezing effects
(Fig. 4). The phosphate data reported here are considered
accurate to about +-5”/”
Nitrate plus nitrlte was determined within 24 hr after
sample collection with a Technicon AutoAnalyzer II (Ind.
Method 158--71/W) using a cadmium-copper
reduction
column (STRICKLAND and PARSONS, 1968). Absorbances
were measured at 540 nm in 2-cm flowcelIs. Standards were
prepared in surface seawater, and the method was blanked
against deionized water. Data from first sequentials
squeezed through acid-washed filters were discarded due
to a small squeezing effect. The remaining acid-washed
data show a small decrease (- 1PM) with sequentials,
while the unwashed data show no squeezing effects
(Fig. 5). Nitrate data reported here are considered accurate
to about + I PM.
Dissolved sulfide concentrations were determined immediately after sample collection by the CLINE (1969)
method. Absorbances were measured at 670 nm in I cm
cells with a Beckman DlJ spectrophotometer. Free dissolved sulfide concentrations in all samples were below
the detection limit of about 3 FM.
Sulfate samples (-I ml) were obtained from the last
sequential squeezing aliquot of each sample, stored in
30
UNWASHED
ACID -WASHED
CUMULATIVE
SQUEEZING
VOLUME,
FILTERS
ml
Fig. 5. Variations in nitrate plus nitrite concentrations as a function of water volume expressed from
individual samples (sequential squeezing aliquots). The first sequential squeezed through acid-washed
filters has been plotted here. but was omitted from the data base.
1081
Suboxic diagenesis
-0
160
-
120
-
,”
::
3
m
._
::
0
0
’
Mn,
’
’
60
pfvl
’
’
’
120
’
I6
0
(unwashed)
Fe, ihl
(unwa:hed)
Fig. 6. Comparison of Mn and Fe values obtained by squeezing adjacent samples (presumably with
identical original pore water Mn and Fe concentrations) with acid-washed and unwashed filters. Many
of these samples were from depth intervals in piston cores that typically showed no Fe and Mn
gradients.
Zdram polyethylene vials in a humid atmosphere, and
returned to the shore-based lab for analysis. One half ml
of sample was added to one half ml of an acidified 130 mM
Ba(NO& solution containing radioactive “‘Ba (10.4yr
half-life, 81 keV and 356 keV y-rays). The resulting BaSO,
precipitated at pH 5 (to minimize BaCO, interference) and
was collected quantitatively on a 0.45 pm Nuclepore filter
and counted with a Ge(Li) detector for 8000 sec. Standards
were prepared from Na,SO, covering the range O-30 mM
SO:-. Linear calibration curves were obtained from plots
of SOi- vs the sum of the two ‘33Ba peak areas. The
sulfate data presented here are considered accurate to
better than *5x.
Iron and manganese samples (1 ml) were acidified to pH
2 with redistilled HCI and stored in acid-washed polyvials
for analysis after return to the laboratory. Both elements
were determined by direct injection into a graphite furnace-atomic
adsorption
unit (Perkin-Elmer
360 or
503 AAS and Model 2100 Graphite Furnace). Iron concentrations in samples expressed through acid-washed filters
were about a factor of two higher than in samples squeezed
from adjacent sediment through unwashed filters (Fig. 6).
‘Unwashed
values are considered accurate to about
& 10%. Operational detection limits (squeezing blanks)
varied from about 0.04 to 0.4pM.
Manganese concentrations were anomalously high (by
25 + 15%) in samples squeezed through acid-washed filters
(Fig. 6). The unwashed data are considered accurate to
+ 5%. Operational detection limits (squeezing blanks)
ranged from 0.02 to 0.2 PM.
pH was determined with punch-in electrodes. pH and
reference electrodes were punched into the top of the
extruded core at in situ temperatures inside the glove bag
during core slicing. A reading was taken when the drift
rate was less than 0.3 mV min- ‘, generally about IO min
after punch-in. Standardization was established at in situ
temperatures using the same drift rate criteria. The electrodes were then extracted and cleaned, while the core was
extruded and sliced. The electrodes were then punched into
this fresh surface for the next reading. pH values are accurate to within 0.1 pH unit. The major problem was due
to slow equilibration of the electrodes at low temperature.
Precision is probably as good as k 0.02 pH units.
Ammonia was determined immediately after collection
by the S~LORZANO (1969) method. Absorbances were
measured at 64Onm in l-cm cells with a Beckman DU
spectrophotometer.
The ammonia data scatter badly
between 1 and about 30pM, with most values being less
than IOpM. The scatter was presumably due to low-level
sample contamination, the source of which was not identified. The data are presented here as a guide to the probable
upper limit of NH: concentrations in these cores.
Coring locations
All cores were collected during RV Gyre cruise G76-5
in May 1976. Coring locations and water depths are presented in Table 3, and in the map in Fig. 7. Productivity
in this area is fairly high due to upwelling along the coast
and at the equatorial divergence (MAHNKEN,1969).
Cores l&23 displayed a distinct change in lithology over
a lo-cm zone centered at about 4Ocm, which consistently
corresponded to the depth at which NO; vanished (see
below). Sediments above 35cm were typically tan, highly
calcareous (70-90x CaCOs), low organic-carbon (0.2-0.5x
C,,,.J post-glacial sediments; they overlie dark-olive-green,
silty, less calcareous ( ~60% CaCO,), moderately organicrich (0.5 to > 1.0% C_), terrigenous glacial sediments.
Porosities in the upper section (measured on duplicate
Table 3. Cores recovered during G76-5 for pore water analysis
core
No.
4GCl
5CCl
24SCl G 2
lOGC1
23GCl
lllnl
12GC2
14GCl
16GCl
Latitude
3O51.5'N
2O51.8'N
2'50.9'N
1'05.1'N
1°06.0'N
0'01.8'N
0°04.1'N
o"oo.l's
0'02.5's
Longitude
5'54.6'W
6'42.7'W
6'41.O'W
S011.6'W
8'12.5'W
9OO3.9'W
10°33.8'W
12019.3'W
16'07.1'W
Depth, m
Length, cm
3612
4563
4572
4956
4901
4980
3880
4170
3310
93
65
27
65
55
89
90
63
48
I’. N.
FKotLIctt 01 u/.
5GC1,
24SCl+
l
140
16’W
Fig. 7. Map of eastern
equatorial
120
Atlantic
showing
locations
GYRE
76-5
DEPTH
IN METERS
8’
IO”
2
6’
of cores squecrcd
for pore water analyses
cores not squeezed
for pore waters) averaged
about
Phosphate
values generally increase from bottom
0.76 k 0.03. Below 40 cm porosities increased. The top and
water values (1.85 ELM)to about 5-6pM at the depth
bottom of this sedimentological
transition zone (35-45 cm)
where nitrate goes to zero. Below this depth, phoshas been “‘C dated at 10,OOQ and 13,000yr B.P., respectphate increases to over 10pM.
ively, in core lOGC1, and probably
was deposited
when
a rising sea-level flooded the continental
shelf, trapping
and dijhsion-reaction
zones in
fluvial clays in estuaries and on the shelf (RICHARDSON. Oxidation-reduction
1974). Sedimentation
rates based on ‘%I dates in cores
marine sediments
lOGC1. SGCl, and 14GCl
are about 4cm 10-3yr-’
Figure 17 is a schematic representation
of trends
above this transition
zone.
RESULTS
Pore water
data for all pelagic cores are shown
in Figs. 8-16. The data plotted here have been
screened according to the criteria established in the
analytical discussion. Acid-washed filter data are plotted with solid symbols to distinguish them from unwashed filter data (open symbols). In general, the scatter is due to squeezing artifacts, contamination
during
handling, or uncertain operational blanks.
The data from all cores other than 4GCl follow
a number of simple trends. NO; concentrations
increase from the ambient bottom water value (22 PM)
to a maximum, then decrease linearly to zero at about
the depth of the lithologic break. Dissolved Mn’+
concentrations
are very low (co.02 PM) at the core
top, but, at some depth between the NO;
maximum and the NO; zero, begin increasing towards
an asymptotic
value of 15-50 PM. Dissolved FeZi
concentrations
are below the detection limit of 0.4 PM
to a depth below the zone where NO; goes to zero,
then begin to increase. Sulfate concentrations
never
detectably differ from bottom water values. Sulfide
concentrations
(not shown) are always less than the
detection limit of 3pM. TCO, concentrations
range
from 2.2mM
to nearly 6mM. TC02 values in
samples squeezed through unwashed filters only (cores
16GCl and 23GCl) are all below 2.8 mM with the
exception of two 23GCl values. pH values consistently fell between 7.6 and 7.8, and show no trends.
in pore water profiles summarized
in the previous
paragraph. The concentration
and depth axes are arbitrary. Zone 1 represents the interval over which
oxygen is being consumed
by organic matter oxidation, releasing ammonia
which is oxidized
to
nitrate. Below the nitrate maximum, nitrate diffuses
downwards to be reduced (presumably by denitrification) near the depth where the nitrate concentration
goes to zero. The remarkable linearity of the downward diffusion gradient suggests that over this interval (zone 2) nitrate is neither produced nor consumed
(BENDERet ~1.. 1977).
Zone 4 represents the interval over which organic
carbon is oxidized by manganese oxides, releasing dissolved Mn2+ to the pore water. The upward convexity of the Mn-profiles indicates production of Mn2+
over this interval. consistent with reduction of solid
phase Mn(IV)O, (ELDERFIELD,1976). MnZ+ then diffuses upwards and is consumed at the top of the gradient (zone 3). We propose that this consumption
is
due to oxidation of Mn 2+ by a small amount of O2
leaking through the base of zone 1. The proposed
reaction is
2Mn”
+ 0,
+ 2H20-+2Mn02
+ 4H’.
(14)
We can check whether this hypothesis is reasonable
by estimating the downward oxygen gradient needed
to oxidize the manganese
diffusing upward. The
required oxygen flux is
Fol = -hFhln~. . where F represents
the diffusive
1083
Suboxic diagenesis
4GCl
TC02,
mM
NO;,
PH
NH,+. .LLM
P’M
0
0
50
100 150
TCO,, mM
pti
24%
24SC2
I 0,.
NO;,
FM
A,.
NH,,+.
pIA
20
E
”
I
+
-40
NO
DATA
i
a
w
0 60
.
E
80
c
.
L
Po,JYpM
Fe,
Mn.
/AM
50 100
, , I,
'.
0
S042-v
p’M
mM
E5
0
l
.
.
IO 20
T
= IO
1
.
%
.
m
.
I
Ei
015
-:
.
t
20
.
L
-
.
t
Fig. 10.
0.
.
.
Fig. 8
lOGC1
TC02,
mM
5GCI
TC02,
mM
PH
NO;. p’M
20
40
0
NH:.$4
IO
6.0
PM
20
0
DR
t
.*.
1
l.
O@
f
.
.
.
0
.
0
0.
0.
PO.+? /.LM
Fe, p’M
0
Or---L
IO
SO,‘,
0
rr
Po43;pM
mM
20
.
.
’
.
.
60-
NH.,+.
40
0
0
20
0
E
0 20 !z
P
; 40 -
NO;.
25
0
Or-----
pH
40
Mn,
0
0
PM
10.
Fe, PM
2
20 0
SO,‘; mM
0
20
E
0
E
20
”
20 =
a
-
," 40-
60 1
cl%
I
.
I
la
‘.
cl%
.
.
; 40
.
.
-I
.
.
Fig. 9.
60 L
/
L
t
Fig. 11.
Figs. 8-16. Plots of pore water data vs depth in core. Values obtained with acid-washed filters are plotted
with solid symbols (A, 0). Values from unwashed filters are plotted with open symbols (0, A).
PM
1084
23GCl
TCOZ.
Pod-,
mM
pLM
llTW1
TC02.
pH
Mn, PM
Fe.
PM
SO,‘-,
mM
PH
I
Fig.
Fig. 12.
12GC2
PO;-,
mM
/AM
2
pH
NO;,
/_A
Fe. $4
0
4
NH4+.
SO:-.
2
O0
.
l..
.
r
20
.
.
E
u
.
.
.
\
-1
.'
a
W
n
60
.
.
c
r
60
:
@A
mM
2460
TC02,
NO;,
t ..
1
\’
i
Fig. 14
;
‘0
_
NO
DATA
mtvl
/.LM
I3
0
2
Suboxic diagenesis
14GCl
TCO,,
mM
NO;.
0
pM
20
40
:I:
:
t
F”
PO;-,
j.LM
5
O0
IO
.
.
.’
.’
.*
I_
Fe. pM
SO:;
0
0
2
mM
20
40
.
E
0
.
.
1
:
I
t-
t
a
w
0
:
40
r
.
r--
- 20
r
l*.
l
.
Fig. 15.
16GCl
TC02. mM
pH
NO;.
0
02&
NH,+.pM
pt.4
20
0
40
IO
‘i%+O
Fe.
pM
2
4
0
SO,‘;
1085
of infinite dilution, or 3/14 (using the values of LI
and GREGORY,1974, and LINGANE,1958). The maximum Mn2+ gradient at the top of zone 3 in our cores
is about 0.1-0.2 x 10T9 mol cmm4. If this oxygen gradient is linear and extends from the bottom of zone 1
to the center of zone 3, and this distance is 20cm
(estimated from our results), then the oxygen concentration at the base of zone 1 is estimated to be about
2-4pM. Our predicted O2 curve is given in Fig. 17
by the dashed line. The predicted low oxygen concentration after termination of oxygen utilization in the
oxidation of organic matter (base of zone 1) is consistent with other observations suggesting that nitrate
reduction commences and oxic diagenesis ceases at
low, but non-zero O2 concentrations (GOERING,1968;
RICHARDS,1971; DEVOL,1975, 1978).
The process of manganese reduction and mobilization in zone 4, followed by upward diffusion and
reoxidation in zone 3, provides a mechanism for stripping manganese from sediments as they accumulate
and redepositing manganese oxides in a discrete layer
(WANGERSKY,1962; RICHARDSON,1974; LYNN and
BONATTI,1965). MnOl deposited on the sediment surface is buried as additional sediment accumulates. As
sediment passes through zone 4 (which is migrating
upwards at the same rate as sediment accumulates,
so as to remain at a given depth below the sedimentwater interface) MnO, is reduced to Mn’+ and diffuses up to be oxidized and redeposited. With further
sedimentation, this new MnO, passes into zone 4 and
is again remobilized. The result is a sedimentary
manganese trap, the redox equivalent of zone refining,
which gleans manganese passing through the accumuCHARACTERISTIC
CONCENTRATION
mM
z
::
.
0
I
I.-<0
23
NO
DATA
.*
I--
4.
a
Since F = -D,,dc/dz,
where
F is the diffusive flux, D,,is the apparent interstitial
=
8
ano2
--z--’
1 D.,Mn2+
d[Mn’+]
Daqo,
I
Fez+
dCO,l
_D
=
This equation may be rearranged as follows:
= -2
d%‘+]<
d2 [NO-]
L
dz2
n
diffusion coefficient, and dc/dz is the vertical concentration gradient, then
d[O,]
dz
>.
Dtffusmn
14
o
11
l-
flux of the components.
dz
: 0
I
W
dCMn'+l
dr’
dz2
I
Fig. 16.
112Da.unz+
d’[NO-]
dz2
S
R
.
IO
dz2
3, d2 h2+]
N
REACTION
d*[NO;]
dz
.
The ratio of apparent diffusion coefficients in the
sediments is taken as equal to the ratios of the values
c
‘i
\
I
6.
‘,
r
Ei
1
I
d2 [Fe’+]
12
>o
>.
?
dz2
d2[Fe2+]<
o
13
dz2
I
Fig. 17. Schematic representation of trends in pore water
profiles. Depth and concentration axes are in arbitrary
units. The zones, characteristic curvature of the gradients,
and reaction numbers are discussed in the text.
IOU6
P. N.
FKCELKH
lating sediment column and concentrates
it into a
layer highly enriched in manganese. Such layers and
Mn-enriched
upper oxidized sediments
have been
observed by many workers, and have been assumed
to be due to such a remobilization
process (e.g. LI
et al., 1969).
We propose that the depth of the manganese spike
is governed by the balance of O2 diffusing downward
and Mn2+ diffusing upward. Both species must be
completely consumed within the spike. If there were
excess 02, it would diffuse downward
to oxidize
Mn’+ deeper in the core. If there were excess Mn*+.
it would diffuse upwards to be oxidized higher in the
sediment section. The system is at steady state only
when the flux of Mnzf up through the top of zone
4 is twice the flux of O2 leaking down through the
base of zone 1. and this condition is satisfied for only
one depth of the solid phase manganese spike. In a
steady-state system, the concentration
of manganese
in this layer will increase until the sedimentary input
of reactive MnO, is balanced by the efficiency of
reduction and remobilization:
i.e. the peak concentration rises until incomplete remobilization
causes
a loss from the bottom of our diagenetic manganese
trap which just balances the sedimentary input.
A steady state system would display a highly concentrated spike of solid Mn near the top inflection
in the dissolved Mn’+ gradient. The MnO, solid
phase concentrations
above and far below the zones
of diagenesis would be equal. Figure 18 is a schematic
of such a steady-state system.
Figure 19 shows plots of pore water and solid
phase Mn data from cores 4GC1, 5GCl. 23GCl.
10GCl. 14GC1, and 16GCl (solid phase Mn was
determined by instrumental neutron activation analysis). As a first approximation,
these data are similar
to our anticipated
steady-state
case. The dissolved
manganese profiles are convex up and there is a solidphase manganese peak near the top of each gradient.
In detail, however, they differ considerably
from our
predictions. The pore water gradient extends above
the major solid phase spike, and the solid phase concentrations above the peak are 225 times higher than
those below the peak.
In at least two cores (5GCl and lOGC1) there
appear to be two peaks, the upper one at the top
of the interstitial gradient. Double peaks may result
from a change in bottom water [O,] or organic
matter burial rate.
If the oxygen concentration
of the deep water in
the Guinea Basin has decreased (due to fluctuations
in bottom-water
renewal), then the oxygen gradient
into the upper sediment must have decreased, permitting the interstitial
Mn*+
gradient
to migrate
upwards. The same effect would also result from an
increase in the organic burial and oxidation rates in
zone 1. The lower solid Mn peak would then be a
relict feature which is slowly being reduced and redeposited at the depth of the upper solid Mn peak. The
characteristic time for a pore water profile to readjust
c’t u/.
MANGANESE
CONCENTRATION
l
F I
Fig. 18. Schematic
phase Mn profiles
representation
of dissolved and solid
in a hypothetical
steady-state
system.
to a 10 cm shift is of the order of only a year, whereas
the characteristic time for a solid phase readjustment
to this new profile is of the order of a thousand years
(see Table 4). Thus a ‘rapid shift in oceanographic
conditions and in the Mn*+ pore water profile will
be followed by a much slower readjustment
of the
location of the solid phase peak concentration.
A second process that may influence the distribution of solid phase Mn is bioturbation
in the upper
several tens of centimeters. Mixing of the sediment
will tend to smear the peak concentration
upward,
since MnO, is stable above zone 3, but unstable
below. Thus the fact that Mn concentrations
above
the peak are greater than below the peak may be
at least partly due to biological mixing.
We can check whether the MnO,-rich
layer is the
result of diffusional entrapment
(rather than simply
burial of a past surficial layer) by demonstrating
that
the burial rate of solid Mn in the spike roughly
balances the upward diffusion of Mn*+. Using the
values in Table 4 and an average sedimentation
rate
of 4cm lO_“yr-‘.
we estimate the solid Mn burial
flux to be about 15 mg Mn cm-’ 10m3 yr-r and the
upward diffusional flux to be about 5 mg Mncm-’
lo- ’ yr- ‘. These values are similar enough (within
the error of our estimation) to demonstrate
that the
diffusional transport is capable of balancing the burial
fll.lX.
The dissolved iron profile drawn in the schematic
representation
of Fig. 17 is not as well documented
from our field data as is the manganese profile. The
curvature is derived by analogy with MnO, reduction. If this presentation
is correct, then zone 7 represents production of dissolved Fe’+ by reduction of
ferric oxides during organic carbon oxidation (reaction 4, Table 1). Dissolved iron then diffuses upwards
to be consumed
near the top of zone 7. Unlike
manganese (almost all of which is remobilized). only
Suboxic
0
I
4
Mns. w/go0
8
I
0
10x7
diagenesis
20
I
I
IO
40
I
,
20
60
I
0
Tr-G-iG
30
0
4
6
12
A
20
E
u
- 40
fr
l-=--l
I
In
2
Y
5GCl
A DISSOLVED
l SOLID
Ma
Mn2+
lOGC1
t
60
I pLM 0
I
Mns.
0
8
1
I
I
16
1
I
24
I
I
0
1
0
I
16
III
0
I
4
Mn2+
I
0
r-
8
I
I
I
16
I
I
4
l
%
.P-7Fc
Fig.
19. Plots
1,
16GCl
14GCl
of pore water (A) and solid phase (0) Mn data in six cores. Acid-washed
are omitted from pore water data for core 16GC1 and 23GCI.
a small portion of the total iron in these sediments
is mobile, presumably mostly surface coatings (?) of
ferric oxyhydroxides
(STUMM
and MORGAN, 1970).
Thus the visible effect of the iron trap is small, and
no iron peak is observed at the top of zone 7. (We
are presently attempting to differentiate mobile iron
and manganese from more immobile mineral phases
(e.g. in clay lattices) by differential
leaching techniques.)
Deposition of dissolved iron in zone 6 may be due
to oxidation, in which case an electron acceptor must
be involved. The most likely oxidants are O2 which
has survived transport through the O2 and MnO,
reduction zones, or, more likely, nitrate. The downward flux of nitrate is more than sufficient to oxidize
the Fe’+ fluxing upwards. Alternatively,
Fe*+ consumption may occur by incorporation
of reduced iron
into solid phases such as mixed carbonates (siderite),
iron-rich smectites and/or glauconites.
TCOz values in cores 16GCl and 23GCl (‘unwashed’ data) are about 2.6mM by the completion
of O2 consumption,
in agreement with our predicted
value of 2598 pm (Figs. 12 and 16). The TCOL decrease during MnO,
reduction
is not detectable
(zone 4). and values at the depth where nitrate
vanishes are about 2.8-3.2 mM, where we would have
predicted 264OpM. This discrepancy may be partly
due to sampling and/or analytical
problems,
but
could also be accounted for by differences in diffusi-
Mn data
vity of HCO; and O2 or NO;, and by COZ production during deep SOireduction and diffusion of
HCO; upward.
Dissolved
phosphate
concentrations
above the
depth where nitrate goes to zero are usually higher
than the predicted values. We suspect this is due in
part to dissolution of any of several potential phosphorus-bearing
solid phases, including calcium carbonates, hydroxyapatites.
and carbonate fluorapatites.
Phosphate concentrations
higher than our predicted
values may also be due partly to preferential release
(fractionation)
of phosphorus relative to carbon during burial and oxidation. In this case the C:P ratio
in the organic matter undergoing oxidation would increase with burial depth (HARTMANN et al., 1973). In
most of these cores that penetrate the iron reduction
zone (zone 7), a dramatic increase in dissolved phosphate is evident that suggests a release of phosphate
to pore waters during mobilization
of iron host
phases (oxyhydroxide coatings). We can eliminate the
possibility that this deep increase in phosphate is due
to anoxic diagenesis alone, since the observed phosphate increase (about 5 PM) would be accompanied
by an increase in NH; of about 80/1M (reaction 5.
Table 1). The observed NH: increase is certainly less
than 30 PM. Until our ongoing studies of solid phase
phosphorus geochemistry of these cores is completed,
further discussion of phosphorus diagenesis is unwarranted.
P. N.
FROELICH et ul.
Table 4. Characteristic
times for readjustment
1.
time required
is given by.
of pore water
layers (2)
The characteristic
shift in conditions
for
profiles
a pore water profile
(I) and solid phase
to adjust
to a 10 cm
t=g
where t = time of diffusion
x = distance
of diffusion
D = diffusion
coefficient
For x=10
2.
cm and I&,, =
1.5 x 10
-6
2
cm set -’
The characteristic
time required for
thick solid phase layer is given by:
where [MIS
=
concentration
is given
IhIs
by:
F?n in solid
sediment
the layer
of Mn in 1 cm slice
density
flux
-9
a 1 cm
(as moles cmS2).
- sediment),
of
B = bulk dry
-2
of Mn (moles cm
se=-+.
Fbh = -DM
%141* averages
2 x 10
to transport
per wet volume (g-sedimenUcm3).
is about 1 cm thick,
d[Mn2+l r
dz
flux
t = -year.
1971))
[Mnls = B(%Mn), where %Mns = concentration
FIti is given by:
cores,
a diffusive
phase peak (moles-Mn/g
and FM = diffusive
For these
(Bender,
about 91 umoles/g,
B averages
Dh = 1.5 x 10-6 c”2 set -’
moles cm-4 (average
maximum dissolved
about 0.75
(Bender,
-3
g cm
19711,
Mn gradient
,
and
in these
cores.
?herefore
t =
(%My) (B)
f
DELd 7,
z
700 years.
“1
As would be expected in the absence of intense
anoxic diagenesis, SOi- values are not detectably different from the overlying sea water and there is no
detectable
free dissolved
sulfide. However,
slight
SOi- reduction (less than that necessary to produce
about 10pM NH:) cannot be discounted
since (1)
we cannot detect sulfate decreases of less than 1 mM,
and (2) the resulting S2- may have been consumed
by excess Fe’ ’ to produce FeS.
SUMMARY
We have presented and discussed pore water data
for nine eastern equatorial Atlantic gravity cores. We
have shown that oxidants are consumed in the predicted sequence (0,. then MnO, and NO;,
then
Fe,O, or FeOOH). Our results confirm earlier conclusions that the manganese rich bands in equatorial
Atlantic sediments are diagenetic features. TCOs concentrations at the base of the 0, reduction zone agree
well with predicted values, but deeper in the core
exceed predicted values for reasons that are not understood. PO:- concentrations
are higher than predicted from oxidation of organic matter with the Redfield stoichiometry,
and suggest that inorganic phosphorus is both mobilized and precipitated within the
sediment column. Sulfate values are not detectably
different from overlying seawater, demonstrating
the
absence of strongly anoxic conditions. Free dissolved
sulfide was undetectable, confirming either absence of
anoxic conditions or complete titration of low sulfide
levels by excess dissolved Fe’+.
Acknowledgements-The
pore water work
was made poss-
ible through the cooperation of Drs. DAVESCXINK(Texas
A & M University) and KENT FANNING(Univ. of South
Florida) and the captain and crew and marine technicians
of R/V Gyre. We are grateful to Drs. MICHAELP~LSON.
HARRY ELDERFIELDand SAVE EMERSONfor discussion.
This work was supported by NSF grant 7642318 (MLB
and GRH).
Suboxic diagenesis
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