13.2 Compaction, Flow Mecbanisms
ity, n 2 , at the average depth, Z2' below the top of the
sequence (Fig. 13.6a and b).
The ratio between h} and h 2 is the compaction ratio, CR, which directly indicates the reduction in
thickness between a higher and lower interval in a
sedimentary column. Since h} is usually not known,
the compaction ratio is determined from the porosities, n\ and n 2 (Fig. l3.6a, Eq. I).
In tbis case, nz is measured from rock sampies, and n \ is
assumed from published porosity-depth curves for the corresponding sediment type. One has, however, to take into
account that the value of n] and thus the compaction ratio
significantly depend on the height, h I> chosen for the topmost (initial) layer as compared to that ofthe deeper buried
layer (Fig. 13.6b). A thick top layer renders a lower n]
value than a thin one, and therefore yields a lower compaction ratio for the same sequence and depth range considered.
Under favorable conditions, compaction ratios can be
determined or estimated directly (Fig. 13.6c-e). For
example, horizontal, originally cylindrical, sedimentfilled burrows, generated several to tens of centimeters below the sediment-water interface, are deformed under higher burial load to bodies with elliptical cross sections. Shallow-burial concretions preserve more or less the original thickness of parallel
sediment laminae, while the same laminae become
compacted outside. Sediment-filled body fossils, the
shells of which are dissolved on or slightly below the
sea floor, also register subsequent sediment compaction.
Differential Compaction
The effects of compaction are difficult to detect in
homogeneous, widely extended sediments, except for
the above mentioned changes in mass physical properties. Compaction becomes evident, however, in
cross sections of strata of laterally varying compaction ratios, as weIl as in basins with inclined or irregular basin floors (Fig. 13.7). For example, mechanically resistant reef structures undergo little or no
compaction, particularly if their pore space is cemented early. In contrast, the adjacent interreef sediments displayahigh compaction ratio. Such differences in compaction, i.e., differential compaction of
sediments not far apart, may enhance the primary
relief of a bioherm (Fig. l3.7a), or accentuate and
modify the shape of sand-filled channels within a
shale sequence (Fig. 13.7b). Differential compaction
may generate faults on top of a step-like relief in the
underlying bedrock (Fig. 13.7c) and lead to a synthetic fault system on top of an inclined basin floor
(Fig.13.7d).
Differential compaction is a common feature in
marine deltaic environments. Where fluvial sands of
657
distributary channels, river mouth bars and the delta
front (cf. Sect. 3.5) accumulate on top of thick
prodelta muds, they cause greater compaction than in
adjacent areas less loaded. The resulting slight depressions in the topography are immediately filled up
with more sand, which adds to the already existing
higher overburden press ure at this site. In this way,
so-called growth faults are generated (Fig. 13.8a and
b).
Such faults are known, for example, from the Niger delta
front (Cohen and McClay 1996). The growth faults can
cause a rotation of the more loaded sediment bodies. High
sand input into the basin and a seaward prograding delta
front or shelf edge during sea level lowstand may lead to a
series of basinward prograding minor to medium-sized
faults (Fig. 13.8a), while a balance between subsidence and
deposition tends to form a single or only a few large
growth faults (Fig. 13.8b). Some of these initially steeply
inclined faults may be affected by the lateral sliding of sediment ma~ses, such that they then continue basinward as
gently dipping reflectors (Bruce 1973). As a result, a rather
complicated extensional fault pattern can develop, as observed in seismic records of the Texas coast (Fig. 13.8c).
Cross-sections of the continental margin in the northwestem Gulf of Mexico, which is little affected by salt diapirs,
show Tertiary sands and sandy shales subdivided by
growth faults into aseries of coastal-parallel troughs. These
rest on the irregular top of older, thick shales which form
"ridge-like" structures between the sand-filled troughs. The
growth faults die out near the deepest parts of the troughs.
The underlying shales appear to be more compacted than
the shales forming the ridges, which are underconsolidated
and exhibit excess pore fluid pressures from subsurface
depths of approximately 1500 m downwards.
The true origin of all these structures is revealed by
decompacting the sediments with the aid of the
(downward increasing) compaction ratio. In this way,
the state and thickness of a sediment body prior to
compaction can be reconstructed. As demonstrated in
Figs. 13.6b and l3.7, decompaction can be carried
out in several steps to show the evolution of a specific structure. U sing this method, the compactioninduced structures can also be dated.
13.2.3 Compaction flow
Model Calculations
The reduction of the water-filled pore space of sediments versus depth is accompanied by the expulsion
of considerable volumes of water. This water moves
as compaction flow or as a kind of advective flow
either vertically upward through the entire sedimentary column, or uses inclined, weIl permeable layers
to escape lateraIly.
In the following consideration, the amount of pore
water is expressed in terms of height of a pure water
column. The height or volume of expelled pore water
ity, n 2 , at the average depth, Z2' below the top of the
sequence (Fig. 13.6a and b).
The ratio between h} and h 2 is the compaction ratio, CR, which directly indicates the reduction in
thickness between a higher and lower interval in a
sedimentary column. Since h} is usually not known,
the compaction ratio is determined from the porosities, n\ and n 2 (Fig. l3.6a, Eq. I).
In tbis case, nz is measured from rock sampies, and n \ is
assumed from published porosity-depth curves for the corresponding sediment type. One has, however, to take into
account that the value of n] and thus the compaction ratio
significantly depend on the height, h I> chosen for the topmost (initial) layer as compared to that ofthe deeper buried
layer (Fig. 13.6b). A thick top layer renders a lower n]
value than a thin one, and therefore yields a lower compaction ratio for the same sequence and depth range considered.
Under favorable conditions, compaction ratios can be
determined or estimated directly (Fig. 13.6c-e). For
example, horizontal, originally cylindrical, sedimentfilled burrows, generated several to tens of centimeters below the sediment-water interface, are deformed under higher burial load to bodies with elliptical cross sections. Shallow-burial concretions preserve more or less the original thickness of parallel
sediment laminae, while the same laminae become
compacted outside. Sediment-filled body fossils, the
shells of which are dissolved on or slightly below the
sea floor, also register subsequent sediment compaction.
Differential Compaction
The effects of compaction are difficult to detect in
homogeneous, widely extended sediments, except for
the above mentioned changes in mass physical properties. Compaction becomes evident, however, in
cross sections of strata of laterally varying compaction ratios, as weIl as in basins with inclined or irregular basin floors (Fig. 13.7). For example, mechanically resistant reef structures undergo little or no
compaction, particularly if their pore space is cemented early. In contrast, the adjacent interreef sediments displayahigh compaction ratio. Such differences in compaction, i.e., differential compaction of
sediments not far apart, may enhance the primary
relief of a bioherm (Fig. l3.7a), or accentuate and
modify the shape of sand-filled channels within a
shale sequence (Fig. 13.7b). Differential compaction
may generate faults on top of a step-like relief in the
underlying bedrock (Fig. 13.7c) and lead to a synthetic fault system on top of an inclined basin floor
(Fig.13.7d).
Differential compaction is a common feature in
marine deltaic environments. Where fluvial sands of
657
distributary channels, river mouth bars and the delta
front (cf. Sect. 3.5) accumulate on top of thick
prodelta muds, they cause greater compaction than in
adjacent areas less loaded. The resulting slight depressions in the topography are immediately filled up
with more sand, which adds to the already existing
higher overburden press ure at this site. In this way,
so-called growth faults are generated (Fig. 13.8a and
b).
Such faults are known, for example, from the Niger delta
front (Cohen and McClay 1996). The growth faults can
cause a rotation of the more loaded sediment bodies. High
sand input into the basin and a seaward prograding delta
front or shelf edge during sea level lowstand may lead to a
series of basinward prograding minor to medium-sized
faults (Fig. 13.8a), while a balance between subsidence and
deposition tends to form a single or only a few large
growth faults (Fig. 13.8b). Some of these initially steeply
inclined faults may be affected by the lateral sliding of sediment ma~ses, such that they then continue basinward as
gently dipping reflectors (Bruce 1973). As a result, a rather
complicated extensional fault pattern can develop, as observed in seismic records of the Texas coast (Fig. 13.8c).
Cross-sections of the continental margin in the northwestem Gulf of Mexico, which is little affected by salt diapirs,
show Tertiary sands and sandy shales subdivided by
growth faults into aseries of coastal-parallel troughs. These
rest on the irregular top of older, thick shales which form
"ridge-like" structures between the sand-filled troughs. The
growth faults die out near the deepest parts of the troughs.
The underlying shales appear to be more compacted than
the shales forming the ridges, which are underconsolidated
and exhibit excess pore fluid pressures from subsurface
depths of approximately 1500 m downwards.
The true origin of all these structures is revealed by
decompacting the sediments with the aid of the
(downward increasing) compaction ratio. In this way,
the state and thickness of a sediment body prior to
compaction can be reconstructed. As demonstrated in
Figs. 13.6b and l3.7, decompaction can be carried
out in several steps to show the evolution of a specific structure. U sing this method, the compactioninduced structures can also be dated.
13.2.3 Compaction flow
Model Calculations
The reduction of the water-filled pore space of sediments versus depth is accompanied by the expulsion
of considerable volumes of water. This water moves
as compaction flow or as a kind of advective flow
either vertically upward through the entire sedimentary column, or uses inclined, weIl permeable layers
to escape lateraIly.
In the following consideration, the amount of pore
water is expressed in terms of height of a pure water
column. The height or volume of expelled pore water