13.4 Early Diagenesis
The processes mentioned above are described in more
detail, e.g., by Lippmann (1973), Irwin et al. (1977),
Goldhaber and Kaplan (1980), Rudson (1982), Kelts and
McKenzie (1982), Gautier and Claypool (1984), Curtis
(1987), Carpenter et al. (1988), Coleman and Raiswell
(1995), Middleton and Nelson (1996).
The zone of microbial methanogenesis extends deep
below the sea floor, but a specific depth range is difficult to define. In organic matter-rich sediments on the
continental margins of the present oceans, methane occurs in large quantities in the form of gas hydrates
(clathrates; cf. Sect. 14.3.6).
Infresh water environments where pore fluids are
poor in sulfate, the sulfate reduction zone and the precipitation of pyrite are sub ordinate or absent (Fig.
13.23c). As a result, the zone of methanogenesis expands upward. Under anoxic conditions, it may extend
into the water body. The principal authigenic mineral
is siderite, which may occur in large quantities, while
pyrite is rare.
13.4.2 Biogenie Carbonates
(Muds and Grainstones)
Fine-grained sediments rich in carbonate (e.g., hemipelagic oozes and marls), but also grainstones, show
additional early diagenetic features, which result from
the relatively good solubility ofthe primary carbonate
minerals. Aragonitic and high-Mg calcitic skeletal particles, in particular, are easily dissolved and
reprecipitated as carbonate cement in neighboring layers (Figs. 13.l6a,b and 13.24). Some cement types are
illustrated in Fig. 13.27. With increasing burial depth,
pressure solution may become an important mechanism providing carbonate cement (Sect. 13.3.3).
In this way, indistinct primary bedding is enhanced
by diagenetic overprint. Even minor variations in primary composition and pore space are sufficient to
cause significant diagenetic modifications. Typical
examples demonstrating this internal carbonate redistribution process are limestone-marl successions, layered carbonate concretions, and nodular limes tones
(Fig. 13.24). Hardgrounds may result from both the
reaction of surface sediment with the overlying sea
water and internal carbonate redistribution. Marked
differences in compaction between limestone concretions, limestones, and adjacent marls testify to the fact
that carbonate is commonly redistributed at relatively
shallow burial depths (several meters to tens of meters). In contrast, pelagic oozes consisting predominantly of calcite skeletons usually dissolve and
recrystallize at deeper burial depths (cf. Fig. 13.3e,f).
The diagenesis of shallow-water carbonate buildups
and reefs is briefly described in Sect. 3.4.7. It largely
depends on the lithofacies and pore types of these
rocks (mudstones, wackestones, packstones, . grainstones, boundstones). Lagoonallime muds and skeletal
687
shoals may become dolomitized soon after deposition
via direct contact and water exchange with the overlying, possibly hypersaline water body. If such an environment is maintained in a subsiding basin for considerable time, thick dolomites of early diagenetic nature
can be built up. The late diagenetic transformation of
thick limestones into dolomites (pervasive dolomitization) requires a different mechanism (Sect.
13.5.3).
Pressure solution occurs in those layers, which have not been
lithified by previous cementation (e.g., Rird and Tucker
1988; J~s and Bone 1989). The dissolved carbonate is
transported by diffusion over short distances (a few to tens of
centimeters). For further details see, e.g., Ricken (1986),
Ricken and Eder (1991), Bathurst (1987, 1991). A comprehensive and detailed treatment of carbonate diagenesis, inc1uding descriptions ofthe various cement types and numerous references, has been provided by Bathurst (1975) and
McIlreath and Morrow (1990). For abriefsummary see, e.g.,
Jordan and Wilson (1994). Detailed studies of carbonate
cements, inc1uding their response to sea-level changes, have
been carried out by many workers (e.g. Budd et al. 1995;
Koch anel Liedmann 1996; Schroeder et al. 1996; Grötsch
1996; Westphal 1998).
13.4.3 Biosiliceous Sediments
The diagenesis of siliceous, predominantly deep-sea
sediments is described in Sect. 5.3.5. The phenomena
of silica diagenesis in shelf carbonates are in many
ways similar to those observed in pelagic carbonates.
Due to more pronounced undersaturation of shallow
waters with respect to silica, however, opal skeletons,
such as sponge spicules, are frequently dissolved and
replaced by carbonate as long as theyare in diffusive
contact with sea water. Opaline silica preserved within
the sediment usually forms early diagenetic chert nodules which are little affected by compaction. Under
special conditions, calcareous fossil remains become
silicified. This may result, for example, from silicarieh alkaline pore fluids below hypersaline lagoonal
waters. Silicification of carbonate mayaiso occur at
the contact with fresh pore water, when some of the
dissolved silica in saline pore fluids is precipitated due
to lowered pH.
For further information see, e.g., Brachert et al. (1987),
Iijima (1988). Chert in oceanic sediments is usually found in
sections ofMiocene age and older. However, very early chert
formation has been detected in sediment cores from the
southernmost Atlantic ocean where the siliceous ooze is little
diluted by terrigenous material (Bohrrnann et al. 1994).
In highly saline, alkaline, sodium carbonate-bicarbonate
lakes, a special type of early diagenetic chert may be precipitated (Sheppard and Gude 1986). This was first observed in
Upper Pleistocene deposits of Lake Magadi in Kenya, but
since then has been found il1 many lake deposits of Jurassic
to Pleistocene age. This Magadi-type chert occurs as thin,
discontinuous beds or as plates and nodules of irregular
The processes mentioned above are described in more
detail, e.g., by Lippmann (1973), Irwin et al. (1977),
Goldhaber and Kaplan (1980), Rudson (1982), Kelts and
McKenzie (1982), Gautier and Claypool (1984), Curtis
(1987), Carpenter et al. (1988), Coleman and Raiswell
(1995), Middleton and Nelson (1996).
The zone of microbial methanogenesis extends deep
below the sea floor, but a specific depth range is difficult to define. In organic matter-rich sediments on the
continental margins of the present oceans, methane occurs in large quantities in the form of gas hydrates
(clathrates; cf. Sect. 14.3.6).
Infresh water environments where pore fluids are
poor in sulfate, the sulfate reduction zone and the precipitation of pyrite are sub ordinate or absent (Fig.
13.23c). As a result, the zone of methanogenesis expands upward. Under anoxic conditions, it may extend
into the water body. The principal authigenic mineral
is siderite, which may occur in large quantities, while
pyrite is rare.
13.4.2 Biogenie Carbonates
(Muds and Grainstones)
Fine-grained sediments rich in carbonate (e.g., hemipelagic oozes and marls), but also grainstones, show
additional early diagenetic features, which result from
the relatively good solubility ofthe primary carbonate
minerals. Aragonitic and high-Mg calcitic skeletal particles, in particular, are easily dissolved and
reprecipitated as carbonate cement in neighboring layers (Figs. 13.l6a,b and 13.24). Some cement types are
illustrated in Fig. 13.27. With increasing burial depth,
pressure solution may become an important mechanism providing carbonate cement (Sect. 13.3.3).
In this way, indistinct primary bedding is enhanced
by diagenetic overprint. Even minor variations in primary composition and pore space are sufficient to
cause significant diagenetic modifications. Typical
examples demonstrating this internal carbonate redistribution process are limestone-marl successions, layered carbonate concretions, and nodular limes tones
(Fig. 13.24). Hardgrounds may result from both the
reaction of surface sediment with the overlying sea
water and internal carbonate redistribution. Marked
differences in compaction between limestone concretions, limestones, and adjacent marls testify to the fact
that carbonate is commonly redistributed at relatively
shallow burial depths (several meters to tens of meters). In contrast, pelagic oozes consisting predominantly of calcite skeletons usually dissolve and
recrystallize at deeper burial depths (cf. Fig. 13.3e,f).
The diagenesis of shallow-water carbonate buildups
and reefs is briefly described in Sect. 3.4.7. It largely
depends on the lithofacies and pore types of these
rocks (mudstones, wackestones, packstones, . grainstones, boundstones). Lagoonallime muds and skeletal
687
shoals may become dolomitized soon after deposition
via direct contact and water exchange with the overlying, possibly hypersaline water body. If such an environment is maintained in a subsiding basin for considerable time, thick dolomites of early diagenetic nature
can be built up. The late diagenetic transformation of
thick limestones into dolomites (pervasive dolomitization) requires a different mechanism (Sect.
13.5.3).
Pressure solution occurs in those layers, which have not been
lithified by previous cementation (e.g., Rird and Tucker
1988; J~s and Bone 1989). The dissolved carbonate is
transported by diffusion over short distances (a few to tens of
centimeters). For further details see, e.g., Ricken (1986),
Ricken and Eder (1991), Bathurst (1987, 1991). A comprehensive and detailed treatment of carbonate diagenesis, inc1uding descriptions ofthe various cement types and numerous references, has been provided by Bathurst (1975) and
McIlreath and Morrow (1990). For abriefsummary see, e.g.,
Jordan and Wilson (1994). Detailed studies of carbonate
cements, inc1uding their response to sea-level changes, have
been carried out by many workers (e.g. Budd et al. 1995;
Koch anel Liedmann 1996; Schroeder et al. 1996; Grötsch
1996; Westphal 1998).
13.4.3 Biosiliceous Sediments
The diagenesis of siliceous, predominantly deep-sea
sediments is described in Sect. 5.3.5. The phenomena
of silica diagenesis in shelf carbonates are in many
ways similar to those observed in pelagic carbonates.
Due to more pronounced undersaturation of shallow
waters with respect to silica, however, opal skeletons,
such as sponge spicules, are frequently dissolved and
replaced by carbonate as long as theyare in diffusive
contact with sea water. Opaline silica preserved within
the sediment usually forms early diagenetic chert nodules which are little affected by compaction. Under
special conditions, calcareous fossil remains become
silicified. This may result, for example, from silicarieh alkaline pore fluids below hypersaline lagoonal
waters. Silicification of carbonate mayaiso occur at
the contact with fresh pore water, when some of the
dissolved silica in saline pore fluids is precipitated due
to lowered pH.
For further information see, e.g., Brachert et al. (1987),
Iijima (1988). Chert in oceanic sediments is usually found in
sections ofMiocene age and older. However, very early chert
formation has been detected in sediment cores from the
southernmost Atlantic ocean where the siliceous ooze is little
diluted by terrigenous material (Bohrrnann et al. 1994).
In highly saline, alkaline, sodium carbonate-bicarbonate
lakes, a special type of early diagenetic chert may be precipitated (Sheppard and Gude 1986). This was first observed in
Upper Pleistocene deposits of Lake Magadi in Kenya, but
since then has been found il1 many lake deposits of Jurassic
to Pleistocene age. This Magadi-type chert occurs as thin,
discontinuous beds or as plates and nodules of irregular