301
1992). Depending on the abundance and reactivity
of the Fe(III) minerals present as well as on the
rate of sulfate reduction, H 2 S may be undetectable
despite its production at high rates. As the H 2 S
accumulates in the pore water, it diffuses up
towards the sediment surface where most of it is
ultimately oxidized back to sulfate. Only a small
fraction, 5-20%, of the produced H 2 S is trapped as
solid phase iron-sulfur minerals or organic sulfur
within the sediment. Over longer periods of thousands to millions of years, most of this trapped
sulfur becomes exposed to oxidants and again
returns to the great sulfate reservoir of the ocean.
Over geological time, ocean sulfate is thus
continuously recycled through sulfate reduction
in the sea bed. During this recycling, the sulfur
atom in sulfate undergoes a major redox transformation of 8 oxidation steps between sulfate (+6)
and sulfide (-2) and is thereby an effective oxidant
for the mineralization of marine organic matter.
Importantly, the sulfur atom in sulfate is separated
from the oxygen atoms during the reduction
process (SO 4
2→ H 2 S), and is subsequently
combined with new oxygen atoms from sediment
pore water upon reoxidation of the sulfide. How
fast is this recombination of the elements and how
long is the residence time of sulfate in sea water?
We have calculated above that the magnitude of
sulfate reduction in the global sea bed is 7.5 · 10
13
mol SO 4
2yr
-1
. The mean sulfate concentration of
sea water is 29 mM and the total volume of ocean
water is 1.37 billion km
3
(Garrison 1997). Thus, the
ocean contains 1.37 · 10
21
liter of sea water with a
total pool of 4.0 · 10
19
mol SO 4
2. The turnover time
of this global ocean sulfate pool through bacterial
sulfate reduction is (4.0 · 10
19
)/(7.5 · 10
13
) = ca. 0.5
million years.
Sulfate is reduced in the sea bed not only
through bacterial catalysis but also through thermochemical catalysis at high temperature. At midoceanic ridges sea water is slowly convected
through the hot magmatic crust where a part of it
is heated to >350°C (see Chapter 13). At such high
temperatures sulfate reacts as an oxidant for
ferrous iron and other reduced minerals and the
sulfate is reduced to H 2 S without the participation
of microorganisms which are excluded by
temperatures above ca. 100°C (Weber and
Jørgensen 2002). The global sea water flux that
undergoes heating to >350°C has been estimated
to be 3-6 · 10
13
liter yr
-1
(Elderfield and Schultz
1996). By a complete reduction of the 29 mM of
sea water sulfate, this corresponds to the
reduction of 0.08-0.16 · 10
13
mol SO 4
2yr
-1
or 1-2%
of the global microbiological sulfate reduction.
The turnover time of the entire volume of ocean
water at >350°C, and thus of the oceanic sulfate
pool through thermochemical reduction, in hot
mid-oceanic crust is thus 20-40 million years (cf.
Elderfield and Schultz 1996).
The hydrothermal systems associated with the
mid-oceanic ridges harbor some of the richest
biological communities on the ocean floor and
include giant clams, mussels, tube worms and
shrimps. These animals are nourished by symbiotic bacteria that oxidize sulfide and methane from
the vent water and use the energy for chemoautotrophic biomass production that in turn feeds
the hosts. It is an interesting perspective that
their energy source is independent of light and
photosynthesis but is instead of geothermal
origin and based on chemical reactions at high
temperature. (It should be remembered, however,
that the oxygen used for the oxidation of sulfide
and methane is derived from photosynthesis in
the surface ocean, so the chemosynthetic processes would still not run without sunlight.)
Although the hydrothermal vent communities are
rich oases of life on the sea floor, their total
contribution to the oceanic carbon cycle is
marginal. Bach and Edwards (2003) estimated that
the global chemosynthetic biomass production
from oxidation of new ocean crust may be up to
10
12
g C yr
-1
. The total potential for chemosynthetic primary production at the deep sea hydrothermal vents is globally estimated to be about
10
13
g biomass per year (McCollom and Shock
1997). This represents only about 0.02% of the
global primary production by photosynthesis in
the oceans. As the chemosynthetic production
takes place in the otherwise nutrient-poor deep
sea, however, it makes a much more important
contribution to the local carbon supply to the
deep sea floor. Based on the oxygen uptake data
presented in Fig. 8.4, the global organic carbon
mineralization by oxygen in the deep sea bed at
>3000 m water depth is only 5 · 10
13
mol yr
-1
,
relative to which the chemoautotrophic production at mid-oceanic ridges corresponds to 20%.
8.7
The Sulfur Cycle
1992). Depending on the abundance and reactivity
of the Fe(III) minerals present as well as on the
rate of sulfate reduction, H 2 S may be undetectable
despite its production at high rates. As the H 2 S
accumulates in the pore water, it diffuses up
towards the sediment surface where most of it is
ultimately oxidized back to sulfate. Only a small
fraction, 5-20%, of the produced H 2 S is trapped as
solid phase iron-sulfur minerals or organic sulfur
within the sediment. Over longer periods of thousands to millions of years, most of this trapped
sulfur becomes exposed to oxidants and again
returns to the great sulfate reservoir of the ocean.
Over geological time, ocean sulfate is thus
continuously recycled through sulfate reduction
in the sea bed. During this recycling, the sulfur
atom in sulfate undergoes a major redox transformation of 8 oxidation steps between sulfate (+6)
and sulfide (-2) and is thereby an effective oxidant
for the mineralization of marine organic matter.
Importantly, the sulfur atom in sulfate is separated
from the oxygen atoms during the reduction
process (SO 4
2→ H 2 S), and is subsequently
combined with new oxygen atoms from sediment
pore water upon reoxidation of the sulfide. How
fast is this recombination of the elements and how
long is the residence time of sulfate in sea water?
We have calculated above that the magnitude of
sulfate reduction in the global sea bed is 7.5 · 10
13
mol SO 4
2yr
-1
. The mean sulfate concentration of
sea water is 29 mM and the total volume of ocean
water is 1.37 billion km
3
(Garrison 1997). Thus, the
ocean contains 1.37 · 10
21
liter of sea water with a
total pool of 4.0 · 10
19
mol SO 4
2. The turnover time
of this global ocean sulfate pool through bacterial
sulfate reduction is (4.0 · 10
19
)/(7.5 · 10
13
) = ca. 0.5
million years.
Sulfate is reduced in the sea bed not only
through bacterial catalysis but also through thermochemical catalysis at high temperature. At midoceanic ridges sea water is slowly convected
through the hot magmatic crust where a part of it
is heated to >350°C (see Chapter 13). At such high
temperatures sulfate reacts as an oxidant for
ferrous iron and other reduced minerals and the
sulfate is reduced to H 2 S without the participation
of microorganisms which are excluded by
temperatures above ca. 100°C (Weber and
Jørgensen 2002). The global sea water flux that
undergoes heating to >350°C has been estimated
to be 3-6 · 10
13
liter yr
-1
(Elderfield and Schultz
1996). By a complete reduction of the 29 mM of
sea water sulfate, this corresponds to the
reduction of 0.08-0.16 · 10
13
mol SO 4
2yr
-1
or 1-2%
of the global microbiological sulfate reduction.
The turnover time of the entire volume of ocean
water at >350°C, and thus of the oceanic sulfate
pool through thermochemical reduction, in hot
mid-oceanic crust is thus 20-40 million years (cf.
Elderfield and Schultz 1996).
The hydrothermal systems associated with the
mid-oceanic ridges harbor some of the richest
biological communities on the ocean floor and
include giant clams, mussels, tube worms and
shrimps. These animals are nourished by symbiotic bacteria that oxidize sulfide and methane from
the vent water and use the energy for chemoautotrophic biomass production that in turn feeds
the hosts. It is an interesting perspective that
their energy source is independent of light and
photosynthesis but is instead of geothermal
origin and based on chemical reactions at high
temperature. (It should be remembered, however,
that the oxygen used for the oxidation of sulfide
and methane is derived from photosynthesis in
the surface ocean, so the chemosynthetic processes would still not run without sunlight.)
Although the hydrothermal vent communities are
rich oases of life on the sea floor, their total
contribution to the oceanic carbon cycle is
marginal. Bach and Edwards (2003) estimated that
the global chemosynthetic biomass production
from oxidation of new ocean crust may be up to
10
12
g C yr
-1
. The total potential for chemosynthetic primary production at the deep sea hydrothermal vents is globally estimated to be about
10
13
g biomass per year (McCollom and Shock
1997). This represents only about 0.02% of the
global primary production by photosynthesis in
the oceans. As the chemosynthetic production
takes place in the otherwise nutrient-poor deep
sea, however, it makes a much more important
contribution to the local carbon supply to the
deep sea floor. Based on the oxygen uptake data
presented in Fig. 8.4, the global organic carbon
mineralization by oxygen in the deep sea bed at
>3000 m water depth is only 5 · 10
13
mol yr
-1
,
relative to which the chemoautotrophic production at mid-oceanic ridges corresponds to 20%.
8.7
The Sulfur Cycle