13.6 Transition from Diagenesis to Metamorphism
- In sandstones. Clay minerals (e.g. kaolinite) - quartz
- carbonates (including dolomite) - sometimes sulfates
(anhydrite) and chlorides. Late calcite is often found to
replace quartz. In volcaniclastic sands, glass may be
followed by zeolites which later trans form into feldspar or mica. Greywackes and sands rich in feldspar
and lithic components develop a matrix consisting of
altered primary particles and cement which was precipitated later (clay minerals, chlorite, and other silicates).
- In shallow-marine carbonates or mixed siliciclasticcarbonate sediments. High-Mg calcite (in addition to
aragonite) - ankeritic calcite (Fe,Mn<:Mg) - ferroan
calcite -low-Mg calcite. In each ofthese stages, calcite
may be replaced by dolomite.
Recent work has demonstrated that the number of cement generations is often much higher than listed here.
In many instances 10 to 20 generations have been
identified. The main reason for these findings is the
fact that meteoric and basin-wide fluid flow have reactivated the diagenetic processes in the deep-burial
stage. As a result, some cement types formed repeatedly and then are labelled, e.g. calcite I, 11, III, etc.
Some examples of cement minerals frequently observed in bioclastic carbonates and quartz sandstones
are illustrated in Fig. 13.27. Detailed descriptions are
presented in many publications on carbonate and sandstone diagenesis.
Numerous publications contain infonnation on cement mineral sequences, besides observations on burial depths, fracturing and dissolution events of certain rock types in specific
tectonic and environmental settings (e.g. Schmidt and McDonald 1979; Edman and Surdam 1984; Füchtbauer 1988;
Schroeder 1988; Dreyer et al. 1990; Schroeder et al. 1996;
Swennen et al. 1996; Heydari 1997; Wojcik et al. 1997).
More references are listed earlier in this chapter.
13.6 The Transition from Diagenesis
to Metamorphism
At deep burial depth, hydrous mineral phases tend to
trans form into denser, less hydrous phases (Fig. 13.28;
cf. mineral reactions in Fig. 13.19). Smectite, for example, forms illite at temperatures between 60 and
100°C, or kaolinite is converted to illite and quartz
(120 to 150°C). Zeolites, derived from volcanic glass,
and other unstable minerals form chlorite, low-temperature feldspar (preferentially potash feldspar or albite),
and quartz. Oxyhydrates of iron and manganese form
less hydrous compounds or, after reduction, may be
ineorporated into carbonates. Most of these reaetions
release water and dissolved species and thus modify
the ionic strength and composition of the pore fluids.
In conjunction with continued porosity reduetion,
possibly enhanced by lateral tectonic stress, shales
transform into slates. Their porosities drop tovalues
around 3% and less. This transition from deep-burial
697
diagenesis to low-grade or "sub-greensehist facies"
metamorphism is characterized by increasing
crystallinily 01 illitelmuscovite (e.g., Dunoyer de
Segonzac et al. 1968; Frey 1987; Robinson 1987).
A clear boundary between diagenesis and metamorphism cannot be set. Evaporites and organic matter
show significant transformations earlier than the common silicate materials. Many authors hold the opinion
that metamorphism of silicate rocks commences at
temperatures of 150±50°C (Bucher and Frey 1994).
Beyond this limit, a number of specifie minerals appear, such as glaucophane, lawsonite, ete.
Regional metamorphism norrnally oecurs in deeply
subsiding basins accumulating great thicknesses of
sediment::;. Metamorphism at shallower burial depths is
caused by specific faetors, such as abnormally high
temperature gradients and lateral compressional stress,
or along the contact to ascending hot plutonic rocks or
magma.
Assuming a ± normal geothermal gradient of 30° Ikm, a specific rock unit must be buried almost 7 km in order to attain
the temperature needed for the onset of significant low-grade
metamorphism (around 200°C).
Some specific problems related to the smectite-illite transformation in mud rocks are discussed by Awwiller (1993).
13.7 Thermal History of Basin Fills
13.7.1 Subsiding Basins Affected by
Conductive Heat Flow
The temperature distribution in a sedimentary basin is
eontrolled by conductive heat flow from the deeper
ernst and, in particular cases, by advective heat transfer due to meteoric water circulation. In addition,
changes in the thermal conductivity due to variations
in the eomposition and compaction (porosity) of the
basin-filling sediments have to be considered.
Figure 13.29 demonstrates examples of the reconstrnction ofthe thermal history and temperature distribution of basins assuming that meteoric flow was absent or insignificant. Rift basins evolving into passive
eontinental margin basins experience an early phase of
enhanced conductive heat flow (Sect. 8.3), i.e., high
temperature gradients, which is followed by cooling
and decreasing temperature gradients. Such scenarios
have been modeled for a number ofbasins; two exampIes ofthe North American Atlantic margin are shown
in (Fig. 13.29a,b).
In the model for the Georges Bank basin offNewfoundland,
it is assumed that the stretching factor (lithospheric attenuation) was ß=2.5 (cf. Sect. 8.3). During the initial rifting stage,
200 Ma before present, the temperature gradient may have
been about 50°C/km, but later it dropped to approximately
30°C. Because ofthe rapid subsidence and sediment accumulation during the Jurassic, the base of these sediments
- In sandstones. Clay minerals (e.g. kaolinite) - quartz
- carbonates (including dolomite) - sometimes sulfates
(anhydrite) and chlorides. Late calcite is often found to
replace quartz. In volcaniclastic sands, glass may be
followed by zeolites which later trans form into feldspar or mica. Greywackes and sands rich in feldspar
and lithic components develop a matrix consisting of
altered primary particles and cement which was precipitated later (clay minerals, chlorite, and other silicates).
- In shallow-marine carbonates or mixed siliciclasticcarbonate sediments. High-Mg calcite (in addition to
aragonite) - ankeritic calcite (Fe,Mn<:Mg) - ferroan
calcite -low-Mg calcite. In each ofthese stages, calcite
may be replaced by dolomite.
Recent work has demonstrated that the number of cement generations is often much higher than listed here.
In many instances 10 to 20 generations have been
identified. The main reason for these findings is the
fact that meteoric and basin-wide fluid flow have reactivated the diagenetic processes in the deep-burial
stage. As a result, some cement types formed repeatedly and then are labelled, e.g. calcite I, 11, III, etc.
Some examples of cement minerals frequently observed in bioclastic carbonates and quartz sandstones
are illustrated in Fig. 13.27. Detailed descriptions are
presented in many publications on carbonate and sandstone diagenesis.
Numerous publications contain infonnation on cement mineral sequences, besides observations on burial depths, fracturing and dissolution events of certain rock types in specific
tectonic and environmental settings (e.g. Schmidt and McDonald 1979; Edman and Surdam 1984; Füchtbauer 1988;
Schroeder 1988; Dreyer et al. 1990; Schroeder et al. 1996;
Swennen et al. 1996; Heydari 1997; Wojcik et al. 1997).
More references are listed earlier in this chapter.
13.6 The Transition from Diagenesis
to Metamorphism
At deep burial depth, hydrous mineral phases tend to
trans form into denser, less hydrous phases (Fig. 13.28;
cf. mineral reactions in Fig. 13.19). Smectite, for example, forms illite at temperatures between 60 and
100°C, or kaolinite is converted to illite and quartz
(120 to 150°C). Zeolites, derived from volcanic glass,
and other unstable minerals form chlorite, low-temperature feldspar (preferentially potash feldspar or albite),
and quartz. Oxyhydrates of iron and manganese form
less hydrous compounds or, after reduction, may be
ineorporated into carbonates. Most of these reaetions
release water and dissolved species and thus modify
the ionic strength and composition of the pore fluids.
In conjunction with continued porosity reduetion,
possibly enhanced by lateral tectonic stress, shales
transform into slates. Their porosities drop tovalues
around 3% and less. This transition from deep-burial
697
diagenesis to low-grade or "sub-greensehist facies"
metamorphism is characterized by increasing
crystallinily 01 illitelmuscovite (e.g., Dunoyer de
Segonzac et al. 1968; Frey 1987; Robinson 1987).
A clear boundary between diagenesis and metamorphism cannot be set. Evaporites and organic matter
show significant transformations earlier than the common silicate materials. Many authors hold the opinion
that metamorphism of silicate rocks commences at
temperatures of 150±50°C (Bucher and Frey 1994).
Beyond this limit, a number of specifie minerals appear, such as glaucophane, lawsonite, ete.
Regional metamorphism norrnally oecurs in deeply
subsiding basins accumulating great thicknesses of
sediment::;. Metamorphism at shallower burial depths is
caused by specific faetors, such as abnormally high
temperature gradients and lateral compressional stress,
or along the contact to ascending hot plutonic rocks or
magma.
Assuming a ± normal geothermal gradient of 30° Ikm, a specific rock unit must be buried almost 7 km in order to attain
the temperature needed for the onset of significant low-grade
metamorphism (around 200°C).
Some specific problems related to the smectite-illite transformation in mud rocks are discussed by Awwiller (1993).
13.7 Thermal History of Basin Fills
13.7.1 Subsiding Basins Affected by
Conductive Heat Flow
The temperature distribution in a sedimentary basin is
eontrolled by conductive heat flow from the deeper
ernst and, in particular cases, by advective heat transfer due to meteoric water circulation. In addition,
changes in the thermal conductivity due to variations
in the eomposition and compaction (porosity) of the
basin-filling sediments have to be considered.
Figure 13.29 demonstrates examples of the reconstrnction ofthe thermal history and temperature distribution of basins assuming that meteoric flow was absent or insignificant. Rift basins evolving into passive
eontinental margin basins experience an early phase of
enhanced conductive heat flow (Sect. 8.3), i.e., high
temperature gradients, which is followed by cooling
and decreasing temperature gradients. Such scenarios
have been modeled for a number ofbasins; two exampIes ofthe North American Atlantic margin are shown
in (Fig. 13.29a,b).
In the model for the Georges Bank basin offNewfoundland,
it is assumed that the stretching factor (lithospheric attenuation) was ß=2.5 (cf. Sect. 8.3). During the initial rifting stage,
200 Ma before present, the temperature gradient may have
been about 50°C/km, but later it dropped to approximately
30°C. Because ofthe rapid subsidence and sediment accumulation during the Jurassic, the base of these sediments