8
Sulfur Cycling and Methane Oxidation
294
the sea bed (Brüchert and Pratt 1996). In continental slope and deep sea sediments this provides
a net burial of reducing power over geological
time that contributes to maintain an oxidized
surface of the earth and thus, together with the
burial of organic carbon, to regulate the atmospheric oxygen level (Berner 1982, 1989). Berner
and Raiswell (1983) found that organic carbon and
reduced sulfur were buried in the present ocean
bed at a mean ratio of ca. 2.8 (±1.5) on a weight
basis. This corresponds to a C org :S red burial ratio
of 7.4 on a molar basis. Since the complete
oxidation of organic carbon to CO 2 represents a
change of 4 oxidation steps while the complete
oxidation of reduced sulfur (pyrite) to sulfate
represents 7 oxidation steps, the burial ratio is 4.3
in terms of reducing equivalents. The latter ratio
corresponds to the amount of oxygen that can
potentially accumulate in the atmosphere due to
the burial of organic carbon versus reduced sulfur.
Thus, the reduced sulfur is equivalent to ca. 20%
of the buried reducing equivalents in the sea bed.
In deltaic and shelf sediments, where most
sulfide production takes place, the intensive
sulfide burial during interglacial periods is
interrupted during glaciations that bind ocean
water in polar ice caps and thereby may lower the
sea water level by 100 m. The glacial sea water
low-stand causes erosion of accumulated shelf
sediments and partly reoxidation of the exposed
pyrite, thereby returning much of the accumulated
sulfide into the oceanic sulfate pool. Berner (1982)
estimated the current burial of pyrite in marine
sediments to be 39 Mt S yr
-1
or 0.12 · 10
13
mol S yr
-1
.
This is equivalent to the sulfide production over
5-6 days per year or to 1.6% of the total sulfate
reduction in the global ocean floor (cf. Fig. 8.4).
Thus, 98.4% of the produced sulfide is reoxidized
back to sea water sulfate on a geological time
scale (million years).
The short-term (years to thousand years)
burial of sulfide in the form of pyrite in ocean
margin sediments is more efficient and generally
accounts for 5-20% of the entire sulfide production (Jørgensen et al. 1990; Lin and Morse 1991;
Canfield and Teske 1996). This burial provides a
sink in the dynamic cycling of sulfur that is limited
by the availability of reactive iron to bind the
large excess of sulfide. Raiswell and Canfield
(1998) found that, of the total iron in marine
sediments, on average only 25-28% is highly
reactive, 23-31% is poorly reactive, and 41-42% is
unreactive. In spite of the small fraction of net
pyrite burial, a much larger fraction of the sulfur
cycle, however, does go through pyrite and more
labile iron sulfides, in particular in the surficial
sediment. These metal-bound sulfides are recycled
within the sediment together with the free sulfide
so that 80-95% of the entire sulfide production is
gradually oxidized back to sulfate. The reoxidation
takes place at all depths of the sediment, most
rapidly in the upper oxidized zone but also in the
deeper and sulfidic part where the process is more
difficult to detect.
The oxidation of organic material by sulfate reduction yields only a fraction of the energy that is
available by aerobic respiration of the same organic compounds. Consequently, a large part of the
potential chemical energy is still conserved in the
product, H 2 S, from sulfate reduction and this
energy may become available to other microorganisms, provided a useful oxidant such as O 2
or NO 3
-
is present. In coastal sediments where the
organic deposition, and therefore the sulfate
reduction, is particularly high, the reactive metal
oxides may become completely reduced by sulfide.
In this extreme case, H 2 S may diffuse freely up to
the sediment surface and reach the thin oxic skin
of the surface sediment. A H 2 S-O 2 interface
thereby develops within the uppermost few mm of
the sediment where the gradient-type of colorless
sulfur bacteria may flourish on the chemical
energy from H 2 S. Such hotspots of sulfide oxidation may be recognizable from the dark coloration
of the sediment surface due to black iron sulfide
(“black spots”; Rusch et al. 1998). The sediments
may also develop a distinct coating of filamentous
sulfur bacteria, such as Beggiatoa, that store light
refracting sulfur globules in their cells and thus
provide the sediment with a distinct whitish
appearence. Such white Beggiatoa mats are typical of the sediments around hydrothermal vents
and cold seeps that bring H 2 S from the subsurface
in direct contact with oxygenated sea water
(Jannasch et al. 1989). In extreme cases where the
water column overlying the sediment is anoxic,
e.g. in the permanently stratified Black Sea or
during summer in some eutrophic coastal
environments, sulfide is not retained at the sediment surface but penetrates directly up into the
sea water (e.g., Roden and Tuttle 1992).
In oxic marine sediments, a brown layer rich in
iron and manganese oxides generally separates O 2
and H 2 S and thereby prevents a direct sulfide
oxidation by oxygen (e.g. Thamdrup et al. 1994a).
In this suboxic zone, neither O 2 nor H 2 S is present
Sulfur Cycling and Methane Oxidation
294
the sea bed (Brüchert and Pratt 1996). In continental slope and deep sea sediments this provides
a net burial of reducing power over geological
time that contributes to maintain an oxidized
surface of the earth and thus, together with the
burial of organic carbon, to regulate the atmospheric oxygen level (Berner 1982, 1989). Berner
and Raiswell (1983) found that organic carbon and
reduced sulfur were buried in the present ocean
bed at a mean ratio of ca. 2.8 (±1.5) on a weight
basis. This corresponds to a C org :S red burial ratio
of 7.4 on a molar basis. Since the complete
oxidation of organic carbon to CO 2 represents a
change of 4 oxidation steps while the complete
oxidation of reduced sulfur (pyrite) to sulfate
represents 7 oxidation steps, the burial ratio is 4.3
in terms of reducing equivalents. The latter ratio
corresponds to the amount of oxygen that can
potentially accumulate in the atmosphere due to
the burial of organic carbon versus reduced sulfur.
Thus, the reduced sulfur is equivalent to ca. 20%
of the buried reducing equivalents in the sea bed.
In deltaic and shelf sediments, where most
sulfide production takes place, the intensive
sulfide burial during interglacial periods is
interrupted during glaciations that bind ocean
water in polar ice caps and thereby may lower the
sea water level by 100 m. The glacial sea water
low-stand causes erosion of accumulated shelf
sediments and partly reoxidation of the exposed
pyrite, thereby returning much of the accumulated
sulfide into the oceanic sulfate pool. Berner (1982)
estimated the current burial of pyrite in marine
sediments to be 39 Mt S yr
-1
or 0.12 · 10
13
mol S yr
-1
.
This is equivalent to the sulfide production over
5-6 days per year or to 1.6% of the total sulfate
reduction in the global ocean floor (cf. Fig. 8.4).
Thus, 98.4% of the produced sulfide is reoxidized
back to sea water sulfate on a geological time
scale (million years).
The short-term (years to thousand years)
burial of sulfide in the form of pyrite in ocean
margin sediments is more efficient and generally
accounts for 5-20% of the entire sulfide production (Jørgensen et al. 1990; Lin and Morse 1991;
Canfield and Teske 1996). This burial provides a
sink in the dynamic cycling of sulfur that is limited
by the availability of reactive iron to bind the
large excess of sulfide. Raiswell and Canfield
(1998) found that, of the total iron in marine
sediments, on average only 25-28% is highly
reactive, 23-31% is poorly reactive, and 41-42% is
unreactive. In spite of the small fraction of net
pyrite burial, a much larger fraction of the sulfur
cycle, however, does go through pyrite and more
labile iron sulfides, in particular in the surficial
sediment. These metal-bound sulfides are recycled
within the sediment together with the free sulfide
so that 80-95% of the entire sulfide production is
gradually oxidized back to sulfate. The reoxidation
takes place at all depths of the sediment, most
rapidly in the upper oxidized zone but also in the
deeper and sulfidic part where the process is more
difficult to detect.
The oxidation of organic material by sulfate reduction yields only a fraction of the energy that is
available by aerobic respiration of the same organic compounds. Consequently, a large part of the
potential chemical energy is still conserved in the
product, H 2 S, from sulfate reduction and this
energy may become available to other microorganisms, provided a useful oxidant such as O 2
or NO 3
-
is present. In coastal sediments where the
organic deposition, and therefore the sulfate
reduction, is particularly high, the reactive metal
oxides may become completely reduced by sulfide.
In this extreme case, H 2 S may diffuse freely up to
the sediment surface and reach the thin oxic skin
of the surface sediment. A H 2 S-O 2 interface
thereby develops within the uppermost few mm of
the sediment where the gradient-type of colorless
sulfur bacteria may flourish on the chemical
energy from H 2 S. Such hotspots of sulfide oxidation may be recognizable from the dark coloration
of the sediment surface due to black iron sulfide
(“black spots”; Rusch et al. 1998). The sediments
may also develop a distinct coating of filamentous
sulfur bacteria, such as Beggiatoa, that store light
refracting sulfur globules in their cells and thus
provide the sediment with a distinct whitish
appearence. Such white Beggiatoa mats are typical of the sediments around hydrothermal vents
and cold seeps that bring H 2 S from the subsurface
in direct contact with oxygenated sea water
(Jannasch et al. 1989). In extreme cases where the
water column overlying the sediment is anoxic,
e.g. in the permanently stratified Black Sea or
during summer in some eutrophic coastal
environments, sulfide is not retained at the sediment surface but penetrates directly up into the
sea water (e.g., Roden and Tuttle 1992).
In oxic marine sediments, a brown layer rich in
iron and manganese oxides generally separates O 2
and H 2 S and thereby prevents a direct sulfide
oxidation by oxygen (e.g. Thamdrup et al. 1994a).
In this suboxic zone, neither O 2 nor H 2 S is present