255
two elements which was attributed to dissimilatory iron reduction and subsequent iron reoxidation by manganese oxides which in turn results in
the production of Mn
2+
aq.
.
Similar to the possibility of concurrent reduction of sulfate and ferric iron by a culture of a
single bacteria (Coleman et al. 1993; see section
7.4.3.4) other iron reducing bacteria were found to
additionally maintain dissimilation with more than
one electron acceptors under suboxic conditions
(Lovley and Phillips 1988) or even under oxic
conditions (Myers and Nealson 1988a). In the
presence of Fe(III) and Mn(IV) strain MR-1 was
found to reduce both but additional manganese
reduction occurred due to the immediate abiotic
reaction with released Fe
2+
(Myers and Nealson
1988b). The interactions of biotic and abiotic
reactions are shown in Fig. 7.17.
7.4.3.3 Iron-bound Phosphorus
In section 7.4.2.1 the theoretical significance of
phosphate adsorption onto iron oxides was
illustrated. Numerous studies on natural sediments suggest that iron oxides control phosphate
pore water and solid phase concentrations, as well
as the overall sedimentary phosphate cycle (Krom
and Berner 1980; Froelich et al. 1982; Sundby et al.
1992; Jensen et al. 1995; Slomp et al. 1996a,b). A
generalized representation of the sedimentary
phosphorus cycle is shown in Fig. 6.11. Apart
from the Fe-bound P, organic P and authigenic
carbonate fluorapatite are the principal carriers of
solid phase P. HPO 4
2is the predominant dissolved
P species under sea water conditions (Kester and
Pytkowicz 1967).
The following information and simple calculation allows the reader to assess and understand
the important role of iron oxyhydroxides and their
interactions with phosphorous: A maximum of 2.5 -
2.8 µmol m
-2
of adsorbed phosphate on iron oxides
were found (Goldberg and Sposito 1984; Pena and
Torrent 1984). In order to approximate a maximum
adsorbed phosphate concentration in the sediment
one can assume 1 cm
3
of sediment with a porosity
of 75 %, a dry weight density of 2.65 g cm
-3
, 50
µmol/g Sediment Fe bound to iron oxides and an iron
oxide specific surface area of 120 m
2
g
-1
(see
section 7.4.2.1). For the wet sediment we can
calculate an iron concentration of 33 µmol cm
-3
which is bound to iron oxides. This fraction has a
specific surface area of ∼ 0.22 m
2
within 1 cm
3
of
wet sediment which may then adsorb up to ∼ 0.57
µmol P (assuming an adsorption capacity of 2.6
µmol P per square meter of iron oxide). For a
comparison of adsorbed and dissolved phosphate
concentration we can furthermore assume a
concentration of 5 µM phosphate within the
interstitial water which is equivalent to 0.0037
µmol cm
-3
of wet sediment. Consequently, the
adsorbed fraction of phosphate can be more than
two orders of magnitude greater than the
dissolved fraction due to the presence of iron
oxides. In reality, the ratio of Fe bound to poorly
crystalline iron oxides and P bound by these
phases has been determined to be ∼ 10Fe : 1P for
coastal and shelf sediments (Slomp et al., 1996a).
This differs significantly from the given
theoretical sample in which we find a ratio of 58
(33 µmol Fe : 0.57 µmol P per 1 cm
3
). Either P is
additionally bound in the crystalline lattice of the
iron oxides (i.e. Torrent et al., 1992) or the
adsorption capacity for P in shelf sediments is
much greater than derived from the calculated
example. In the latter case, one must conclude that
the specific surface area of Fe oxides is higher
than assumed as another quantitatively important
adsorbent of P in sediments is not likely. Instead
of 120 m
2
g
-1
iron oxide one needs to encounter a
specific surface area of 650-700 m
2
g
-1
to justify
such high P adsorption.
7.4.3.4 The Formation of Siderite
In the marine environment siderite (FeCO 3 ) is
hardly found relative to iron sulfides because it is
thermodynamically not stable in the presence of
even low dissolved sulfide activities. Postma
(1982) proved the calculation of the solubility
equilibrium between siderite and ambient pore
water chemistry to be a reliable approach for the
investigation of present-day siderite formation. In
salt marsh sediments where the influence of salt
and fresh water varies temporarily and spatially
the formation of siderite and pyrite are closely
interlinked. Mortimer and Coleman (1997)
demonstrated that siderite precipitation is
microbiologically induced. They could show that
δ
18
O values of siderite precipitated during the
culturing of one specific iron-reducing microorganism, Geobacter metallireducens, were
distinctively lower than expected according to
equilibrium fractionation between siderite and
water (Carothers et al. 1988) which is a contradiction to a pure thermodynamically induced
reaction.
7.4
The Early Diagenesis of Iron in Sediments
two elements which was attributed to dissimilatory iron reduction and subsequent iron reoxidation by manganese oxides which in turn results in
the production of Mn
2+
aq.
.
Similar to the possibility of concurrent reduction of sulfate and ferric iron by a culture of a
single bacteria (Coleman et al. 1993; see section
7.4.3.4) other iron reducing bacteria were found to
additionally maintain dissimilation with more than
one electron acceptors under suboxic conditions
(Lovley and Phillips 1988) or even under oxic
conditions (Myers and Nealson 1988a). In the
presence of Fe(III) and Mn(IV) strain MR-1 was
found to reduce both but additional manganese
reduction occurred due to the immediate abiotic
reaction with released Fe
2+
(Myers and Nealson
1988b). The interactions of biotic and abiotic
reactions are shown in Fig. 7.17.
7.4.3.3 Iron-bound Phosphorus
In section 7.4.2.1 the theoretical significance of
phosphate adsorption onto iron oxides was
illustrated. Numerous studies on natural sediments suggest that iron oxides control phosphate
pore water and solid phase concentrations, as well
as the overall sedimentary phosphate cycle (Krom
and Berner 1980; Froelich et al. 1982; Sundby et al.
1992; Jensen et al. 1995; Slomp et al. 1996a,b). A
generalized representation of the sedimentary
phosphorus cycle is shown in Fig. 6.11. Apart
from the Fe-bound P, organic P and authigenic
carbonate fluorapatite are the principal carriers of
solid phase P. HPO 4
2is the predominant dissolved
P species under sea water conditions (Kester and
Pytkowicz 1967).
The following information and simple calculation allows the reader to assess and understand
the important role of iron oxyhydroxides and their
interactions with phosphorous: A maximum of 2.5 -
2.8 µmol m
-2
of adsorbed phosphate on iron oxides
were found (Goldberg and Sposito 1984; Pena and
Torrent 1984). In order to approximate a maximum
adsorbed phosphate concentration in the sediment
one can assume 1 cm
3
of sediment with a porosity
of 75 %, a dry weight density of 2.65 g cm
-3
, 50
µmol/g Sediment Fe bound to iron oxides and an iron
oxide specific surface area of 120 m
2
g
-1
(see
section 7.4.2.1). For the wet sediment we can
calculate an iron concentration of 33 µmol cm
-3
which is bound to iron oxides. This fraction has a
specific surface area of ∼ 0.22 m
2
within 1 cm
3
of
wet sediment which may then adsorb up to ∼ 0.57
µmol P (assuming an adsorption capacity of 2.6
µmol P per square meter of iron oxide). For a
comparison of adsorbed and dissolved phosphate
concentration we can furthermore assume a
concentration of 5 µM phosphate within the
interstitial water which is equivalent to 0.0037
µmol cm
-3
of wet sediment. Consequently, the
adsorbed fraction of phosphate can be more than
two orders of magnitude greater than the
dissolved fraction due to the presence of iron
oxides. In reality, the ratio of Fe bound to poorly
crystalline iron oxides and P bound by these
phases has been determined to be ∼ 10Fe : 1P for
coastal and shelf sediments (Slomp et al., 1996a).
This differs significantly from the given
theoretical sample in which we find a ratio of 58
(33 µmol Fe : 0.57 µmol P per 1 cm
3
). Either P is
additionally bound in the crystalline lattice of the
iron oxides (i.e. Torrent et al., 1992) or the
adsorption capacity for P in shelf sediments is
much greater than derived from the calculated
example. In the latter case, one must conclude that
the specific surface area of Fe oxides is higher
than assumed as another quantitatively important
adsorbent of P in sediments is not likely. Instead
of 120 m
2
g
-1
iron oxide one needs to encounter a
specific surface area of 650-700 m
2
g
-1
to justify
such high P adsorption.
7.4.3.4 The Formation of Siderite
In the marine environment siderite (FeCO 3 ) is
hardly found relative to iron sulfides because it is
thermodynamically not stable in the presence of
even low dissolved sulfide activities. Postma
(1982) proved the calculation of the solubility
equilibrium between siderite and ambient pore
water chemistry to be a reliable approach for the
investigation of present-day siderite formation. In
salt marsh sediments where the influence of salt
and fresh water varies temporarily and spatially
the formation of siderite and pyrite are closely
interlinked. Mortimer and Coleman (1997)
demonstrated that siderite precipitation is
microbiologically induced. They could show that
δ
18
O values of siderite precipitated during the
culturing of one specific iron-reducing microorganism, Geobacter metallireducens, were
distinctively lower than expected according to
equilibrium fractionation between siderite and
water (Carothers et al. 1988) which is a contradiction to a pure thermodynamically induced
reaction.
7.4
The Early Diagenesis of Iron in Sediments