Clay Minerals (1998) 33, 15-34
Clay mineral diagenesis in sedimentary
basins
a key to the prediction of rock
properties. Examples from the North Sea
Basin
K. BJORLYKKE
Department of Geology, Box 1047 Blindern, University of Oslo, N-0316 Oslo, Norway
(Received 23 September 1996," revised 9 June 1997)
A B S T RA C T : Dissolution of feldspar and mica and precipitation of kaolinite require a through flow
of meteoric water to remove cations such as Na + and K § and silica. Compaction driven pore-water
flow is in most cases too slow to be significant in terms of transport of solids. The very low solubility
of A1 suggests that precipitation of new authigenic clay minerals requires unstable Al-bearing
precursor minerals. Chlorite may form diagenetically from smectite and from kaolinite when a source
of Fe and Mg is present. In the North Sea Basin, the main phase of illite precipitation reducing the
quality of Jurassic reservoirs occurs at depths close to 4 km (130-140~ but the amount of illite
depends on the presence of both kaolinite and K-feldspar. Clay mineral reactions in shales and
sandstones are very important factors determining mechanical and chemical compaction and are thus
critical for realistic basin modelling.
The presence of clay minerals and clastic sheet
silicates strongly influence the physical and
chemical properties of both sandstones and shales.
The primary sediment composition and the early
diagenetic reactions determine the burial diagenetic
reactions and rock properties as a function of depth.
Clay minerals will also, in most cases, reduce their
shear strength and increase the surface area of the
sediments and change chemical properties such as
ion exchange capacity. The primary clastic composition of sedimentary rocks is related to source
rocks, weathering and erosion in the source area,
transport processes and to the depositional environment. Each basin has a different basin subsidence
and depositional history and clay diagenesis is
influenced by many different factors.
Diagenetic reactions are driven towards higher
t h e r m o d y n a m i c stability at a rate which is
controlled by the kinetics of the mineral reactions.
The main principles for clay mineral diagenesis
should therefore be the same for all basins even if
they have very different initial mineralogy and
thermal history. If these principles can be agreed
upon, the main problem is making the right
assumptions about variables such as provenance,
facies, sedimentation rates and geothermal gradients. The same diagenetic reactions that we can
study in sandstones probably also apply to
mudstones, even if the texture and mineralogy
may be different. The North Sea Basin and
Haltenbanken (Mid-Norway) Basin are particularly
good 'laboratories' for studying clay mineral
diagenesis. Both basins are extensively cored and
large amounts of geochemical and mineralogical
data are available on the composition of the
sediments, as well as the pore-water (Egeberg &
Aagaard, 1989; Aagaard et al., 1992; Warren &
Smalley, 1994; Bjorlykke et al., 1995). With the
exception of the marginal parts of the basins, there
has been almost continuous subsidence and
sedimentation through the Cenozoic. An overview
of the regional geology and stratigraphy of the
North Sea Basin is provided by Glennie (1990) and
of Haltenbanken by Koch & Heum (1995).
9 1998 The Mineralogical Society
K. Bj~rlykke
16
The present-day geothermal gradients in the
North Sea vary mostly between 30-40~
(Glennie, 1990; Hermanrud et al., 1991) and there
is no evidence to suggest that the geothermal
gradients were very much higher earlier in
Cenozoic times. Even if the geothermal gradients
should have been considerably higher during the
Mesozoic or Lower Tertiary, it is not likely that the
temperature of a particular rock should have had a
higher absolute temperature because of the Pliocene
and Pleistocene subsidence. This is also supported
by the fact that fluid inclusions in quartz from the
North Sea and Haltenbanken record temperatures up
to the present-day bottom-hole temperature but not
higher (Walderhaug, 1990, 1994; Saigal et al.,
1992). Where there has been no hydrothermal
activity or uplift, the present burial depths and
temperatures can, in most cases, be taken as
maximum values because of the rapid late
Cenozoic subsidence. It is reasonable to assume
that the North Sea Basin has experienced recent
progressive burial diagenetic processes with
increasing temperatures, except in the uplifted
marginal parts of the basins.
The depth ranges of authigenic minerals such as
kaolinite, illite, chlorite and quartz provide very
important constraints on the interpretations derived
from petrographic analyses. This paper is an
attempt to present a summary of the main clay
mineral reactions typical of the North Sea and
Hattenbanken basins and to discuss the principles of
diagenetic processes involving clay minerals.
However, a detailed discussion of the regional
variations in clay mineralogy within these basins is
beyond the scope of this paper. For recent reviews
of elastic diagenesis see Wilson (1994).
CLAY MINERALOGY
AND
SANDSTONE DIAGENESIS
Diagenetic reactions must have a thermodynamic
drive so that the minerals precipitated are more
stable than the minerals which are dissolving. At
shallow depths and low temperatures, hydrous
minerals such as gibsite, kaolinite and smectite
form as a result of weathering or early diagenetic
processes during meteoric water flow. Such early
diagenetic processes may be considered a continuation of the weathering process even if the porewater is reducing. The overall reaction:
rock (feldspar, mica) + water = clay + cations.
These minerals become unstable at greater burial
depth and higher temperatures and this reaction is
often referred to as reversed weathering:
clay (kaolinite, smectite) + cations (K+) =
aluminosilicate (illite) + quartz + water.
The above reactions are modified from Velde
(1995).
In the North Sea and Haltenbanken basins,
diagenetic studies have focused mostly on the
reservoir sandstones of Jurassic age. The main
primary minerals such as feldspar and mica are
unstable when exposed to meteoric water of low
ionic strength near the surface (weathering), but
comprise a stable mineral assemblage during burial
diagenesis at higher temperatures and lower flow
rates. It is well known that arkoses have their
feldspars well preserved after exposure to greenschist facies or higher grades of metamorphism.
Only if kaolin, smectite or other potentially
unstable clay minerals form at shallow depth will
clay mineral reactions such as precipitation of illite
take place at greater burial due to higher
temperatures. In such well-sorted reservoir sandstones, nearly all the clay minerals are authigenic
and the distribution of clay minerals then depends
on the diagenetic processes.
Early (shallow) diagenesis
The term early diagenesis is used here to include
processes near the surface where the diagenesis
may be strongly influenced by meteoric water and
sea-water. Early marine diagenesis is strongly
influenced by the accumulation of biogenic
carbonate and silica on the sea floor and by
interaction with sea-water by diffusion near the
red/ox boundary. These changes in the primary
elastic sediment composition are related to marine
facies and may strongly influence diagenetic
reactions at greater burial. Meteoric water may
flow deep into sedimentary basins and many of the
reservoirs in the northem North Sea have salinities
which are about 50% of that of normal sea-water
(Warren & Smally, 1994 ). However, the flux of
meteoric water is highest in the marginal and
shallow parts of the basin. It also depends on the
climate, topography, water table and on aquifers
and aquitards in the basin. Fluvial and shallow
marine sediments will be flushed by meteoric
water after deposition while more distal shelf
facies and turbidites normally will be subjected
Clay mineral diagenesis in sedimentary basins
17
FIG. 1. Leached feldspar and authigenic kaolinite from the Brent Group (Ness Formation, Huldra Field, depth,
3722 m (Nedkvitne & Bjorlykke, 1992). Porosity is gained by feldspar dissolution (secondary porosity) but lost
through precipitation of kaolinite.
to much less meteoric water flushing. The
dissolution of feldspar and mica and precipitation
of kaolinite (Fig. 1) is a weathering reaction and
this type of early diagenesis may be referred to as
subsurface weathering:
2K(Na)A1Si3Os + 2 W + 9H~O = A12SiOs(OH)4
+ 2HaSiO4 + 2K(Na +)
It is clear that this reaction cannot take place in a
closed system because it requires the supply of
protons and the removal of cations such as Na + and
K + and silica by fluid flow. Most ground-waters are
in the stability field of kaolinite (Garrels & Christ,
1965). Calculations show that a total flow of
103--104 m3/m2 through sandstones is required to
dissolve significant quantities of feldspar and mica
and precipitate a few percentages of kaolinite
(Bjorlykke, 1994). This rate of flow is obtained in
fluvial and shallow marine environments if the
climate is humid. Silica will normally not
precipitate as quartz at such low temperatures and
must be removed along with the alkali ions in order
for the pore-water to remain in the stability field for
kaolinite. Meteoric water may penetrate deeply into
sedimentary basins in some cases, but meteoric
water fluxes, significant in relation to dissolution of
feldspar and mica, probably occur mostly at depths
shallower than I00 m - - in many cases shallower
than 10 m (Bjorkum et al., 1990). The model
indicating mixing of meteoric water and compaction driven flow from opposite direction into the
Brent Gr. (Osborne et al., 1994), is hydrodynamically very problematic. Meteoric water flowing into
sedimentary basins will gradually approach equilibrium with the mineral phases present, including
also K-feldspar and mica as the distance from the
18
K. Bjorlykke
area of recharge increases. High fluxes are required
for the pore-water to remain in the stability field of
kaolinite (low K+/H+), but only a few exchanges of
pore-water are required before the pore-water
composition will be dominated by the isotopic
signature of meteoric water. Osborne et al. (1994)
suggested that authigenic kaolinite formed in the
Brent sandstones because of meteoric water flow
into the basin in the Late Cretaceous to Early
Eocene. At that time, however, the Middle Jurassic
Brent Group was in most places covered by a
1 - 2 km thick sequence of mudstones, which
probably had low permeability, thereby reducing
the potential for large fluxes of meteoric water to
flow into the Jurassic sediments and up to the
surface again. In addition, the land areas adjacent to
the northern North Sea such as the Shetland
platform and western Norway were transgressed
by the sea during Late Cretaceous and Early
Tertiary times (Hancock, 1984; Jordt et al., 1995)
providing little head for meteoric water flow into
the basin. Late Eocene and Oligocene smectitic
mudrocks with very low permeability probably
reduced the potential for fluid into the underlying
Mesozoic sequence as well as back up to the
surface.
In the North Sea Basin it has been shown that the
distribution of kaolinite in sandstones can be related
to facies and climate (Bjorlykke & Aagaard, 1992).
The Permian and Triassic sandstones in the North
Sea Region, which were deposited in a dry climate,
generally contain little kaolinite compared to the
Jurassic fluvial and shallow marine sandstones
deposited in a more humid climate. Practically all
samples of sandstones from the Brent Group which
are not carbonate cemented contain authigenic
kaolin and show evidence of feldspar dissolution
(Morton e t al., 1992). Below ~4 km depth,
however, much of the kaolinite has been dissolved
and replaced by illite (Fig. 2). The Brent Group
represents a deltaic facies where both the fluvial
and shallow marine sediments would have been
flushed by meteoric water shortly after deposition.
The most effective leaching of feldspar and mica
occurs at shallow depths (<10-20 m), even though
meteoric water may extend much deeper into the
basin. The degree of dissolution of feldspar and
mica varies through the Brent Group as a function
of facies (Nedkvitne & Bjorlykke, 1992).
The 8180 values of authigenic kaolinite indicate
rather low formation temperatures. The exact
temperatures cannot be calculated because of the
uncertainty of the isotopic composition of the porewater but it is highly unlikely that it could be
formed at high (>100~ temperatures (Glasmann et
al., 1989a,b).
Authigenic kaolinite is common in sandstones of
the Brent Group in the shallowest reservoirs at
<1700 m in the Emerald Field (Osborne et al.,
1994) and at c. 1.8 km depth in the Gullfaks Field
(Bjorlykke et al., 1992; Giles et al., 1992) and in
Upper Jurassic sandstones of the Troll Field at
1.5-1.6 km indicating that kaolinite has precipitated at shallower depth. Osborne et al. (1994)
suggested that kaolinite precipitated from meteoric
water at temperatures between 25-80~ at a burial
depth between 571-2143 m at a time interval
between 47-86 Ma. They assumed that kaolinite
was precipitated from a mixture of meteoric water
and compactional water with 81So values between
-6.5 to -3.5~
SMOW. The composition of
Jurassic meteoric water is, however, poorly
constrained and may vary locally depending on
various factors such as wind directions during
rainfall. Assuming that the kaolinite precipitated
from meteoric ground water with 8180 values
between - 7 to -9%0 SMOW, most of the kaolinite
could have precipitated at low temperatures
(20-30~
which is also suggested by McAulay
et al. (1990). Seventy six analyses of diagenetic
kaolinite from the Brent Group (Osborne et al.,
1994) showed that 5180 values decrease with
increasing present-day burial depth from an
average of 17.2%o at 1600-1700 m, 16.4%o
between 2-3000 m and 14.2%o at >3300 m. This
could indicate that some degree of re-equilibration
had occurred during burial, but then the pore-water
must have continued to have low 5180 at greater
depth. Possibly, some of the kaolinite now observed
in reservoir sandstones could have been recrystallized from amorphous aluminium phases or from
less crystalline kaolinite thus explaining somewhat
elevated temperatures. Some of the kaolinite in the
deeper reservoirs may be dickite and this may also
change the oxygen isotopic composition. Although
kaolinite is a pore-filling mineral, it is often partly
enclosed in authigenic quartz at greater burial depth
(Fig. 3).
Upper Jurassic sandstones from the North Sea
Basin representing more distal shelf facies and
turbidites contain very little or no authigenic
kaolinite and much less evidence of feldspar
leaching compared to the underlying Brent Group
of delta facies where these features are ubiquitous.
19
Clay mineral diagenesis in sedimentary basins
Biogenic
carbonate
s i Iic a
/
I
Detrital supply
and
Basin f
Carbonate cement
BURIAL
DEPTH Opal A-CT - quartz
Little aut. kaolinite
o-3.s(4)
Met..water
i
l
l
~I
Verdine(Fe)
Facies
Extensive
quartz cementation
Little illite if
kaolinite and
s m e c t i t e are absent
N
Dissolution of feldspar
and mica, precipitation
of authigenic k a o l i n i t e
2KAISi308+2H++ 9H20 =
AI2Si205(OH) 4 +4H4SiO 4 +2K +
km
BURIAL
DEPTH >
3.5(4)
Km
~
Chlorite
Quartz cement,
coatings? Illitization if
Little
kaolinite and
quartz
K-feldsparare present
cement
KAISi308+AI2Si2Os(OH)4 =
KAI3Si3010(0H)2+2Si02 +2H20
FI~. 2. Model for relationships between provenance, facies-related early diagenesis and diagenesis at greater
burial depth. Provenance and early diagenesis in meteoric or marine environments strongly influence the
diagenesis at deeper burial.
In the Claymore Field, a thin turbiditic sandstone
(Ten Foot Sand) within the Kimmeridge Clay
Formation shows no significant authigenic kaolinite
and feldspar dissolution (Spark & Trewin, 1986),
while the Piper Formation, which is a paralic
deposit, contains authigenic kaolinite and intensively leached feldspar and mica. The Fulmar
Formation is an example of a distal shelf and
turbidite facies which contain little diagentic
kaolinite (Stewart, 1986). This may be because it
has received too little meteoric water flushing to
cause leaching of feldspar and mica (Saigal et al.,
1992). If the leaching was related to the generation
of CO2 or organic acids as suggested by several
authors (Schmidt & McDonald, 1979; Surdam et
al., 1984, 1989; Burley et al., 1985; Burley, 1986),
then the sandstones which were close to the source
rock would be expected to show the most leaching.
This is clearly not the case. In Upper Jurassic
sandstones, representing more proximal shallow
marine facies such as in the Piper and Tartan
formations, authigenic kaolinite is observed, probably because they have been more extensively
flushed by meteoric water (Burley, 1986). Near the
top of rotated fault blocks, when they were
submerged as islands, the Brent Group has been
exposed to meteoric water flushing (Bjorlykke &
Brendsdal, 1986). Because of uplift and erosion,
relatively little of the section affected by meteoric
flushing below the unconformity may be preserved
(Bjorkum et al., 1990).
Dissolution and precipitation of minerals due to
fluid flow during deeper burial cannot be expected
to be facies selective as is the case with early
diagenetic reactions. Shallower sandstones are
rarely cored but authigenic kaolinite has been
20
K. Bjorlykke
FIG. 3. Authigenic kaolinite enclosed by quartz overgrowth. Rarmock Formation (Brent Group, Statfjord Field).
The scale bar represents 0.001 mm.
observed in cuttings from shallower sandstones, i.e.
from the Pliocene Nordland Group (36/1-2) at
500 m present burial depth (Singh, 1996). Most
Lower Tertiary sandstones in the North Sea contain
little authigenic kaolinite, probably because they
represent distal marine and turbidite facies which
have not been subjected to extensive flushing by
meteoric water (Bjorlykke & Aagaard, 1992).
Minor amounts of authigenic kaolinite may,
however, be found in Paleocene sandstones
(Stewart et al., 1990). The Permian and Triassic
sandstones, which were deposited in a rather dry
climate, contain little authigenic kaolinite and
mostly smectite and illite, which is typical of
desert environments today (Weaver, 1989). In
Upper Triassic sediments, i.e. the Lunde
Formation at the Snorre Field, the kaolin content
is higher, probably due to a slightly less arid
climate and it is clearly linked to increased ground
waterflow through fluvial sandstone.
The depth and temperature of formation of
kaolinite in the basins like the North Sea has
been the subject of considerable controversy. It was
a widely held view that the dissolution of feldspar
and precipitation of kaolinite occurred in connection with release of acids from generation of oil
from kerogen (Burley, 1986; Surdam et al., 1984,
1989). As shown above, observations from cores
Clay mineral diagenesis in sedimentary basins
show that abundant kaolinite has already precipitated prior to deep (>1.5-2.0 km) burial. This is
also supported by isotopic data suggesting a
relatively low temperature (Glasmann et al.,
1989a). Precipitation from meteoric water at
1 - 2 km depth as suggested by Osborne et aL
(1994) cannot be disproved by data since there are
no cores from depths <1.5 km. The interpretation
that kaolinite is formed by leaching at very shallow
depth is based on indirect reasoning about meteoric
water fluxes required to produce significant
dissolution (weathering).
Authigenic clays change the pore-size distribution and therefore also oil saturation and the
production capability of reservoir sandstones
(Pittman, 1978). Prediction of such rock properties
depends very much on the diagenetic model for
minerals like kaolinite which is important in
shallow reservoirs and even more important as
precursor for illite at greater depth.
Burial diagenesis
Meteoric water flow may reach deep into
sedimentary basins driven by the head of ground
water from nearby land areas. The rate of flow is
highest near the surface and decreases with depth
and distance from land areas, depending on the
distribution of aquifers. At a certain depth,
however, the compaction process in subsiding
basins will build up sufficient over-pressure and
prevent penetration of meteoric water. This depth is
very difficult to estimate but it follows from what
has been stated above, that meteoric water can flow
more deeply into uplifted sediments which are not
undergoing compaction.
Weathering reactions like dissolution of feldspar
and precipitation of kaolinite require that K§ and
silica are removed so that the pore-water can
remain in the stability field of kaolinite (Fig. 4).
The pore-water need not be acidic but must have a
low K+/H+ ratio. In sandstones of the North Sea
Basin, there is nearly always some carbonate
present and the pore-water is in equilibrium with
calcite and the pH is to a large extent determined
by the COz. At greater depth (>3-4 km), clay
mineral reactions have the highest pH buffering
capacity. Organic acids have much lower buffering
capacities than both the silicate and the carbonate
system and therefore do not influence the pH very
much (Hutcheon et al., 1992). In the presence of Kfeldspar, the K+/H+ ratio of the pore-water will be
21
too high to be in the stability field of kaolinite. If
some K-feldspar or mica should dissolve, the K
concentration in the pore-water will increase until
the reaction stops, since there is normally no other
mechanism for removing K.
At burial depths > 2 - 3 km, the kaolin mineral
present is commonly not kaolinite but dickite
(Ehrenberg et al., 1993). Much of what has
previously been described as kaolinite from North
Sea reservoir sandstones should mineralogically be
classified as dickite. The transition of kaolinite to
dickite is still poorly understood. It is not known if
there are factors other than temperature influencing
this transition and to what extent it involves total
dissolution of kaolinite and precipitation of dickite
so that the oxygen isotopic ratios are reset.
Kaolinite in shales is probably mostly clastic
although this is difficult to prove because the
primary textures are usually difficult to observe
because they have been destroyed by compaction.
Mudstones are normally not subjected to much
meteoric water flow due to their low permeability.
At least in the presence of K-feldspar, the porewater in the shales should be expected to be in the
stability field of illite. Authigenic kaolinite,
however, may be observed in coarse-grained
mudstones. It is often not quite clear to what
extent kaolinite in these cases has precipitated due
to leaching of feldspar and mica or has precipitated
from clastic gibbsite minerals or amorphous
aluminous gels (Foscolos, 1984) reacting with
biogenic silica. The stability of smectite is
reduced with higher temperatures and, as the rate
of quartz precipitation increases, the pore-water will
be less supersaturated with respect to quartz.
Illite. Authigenic illite often occurs as a fibrous
pore-filling mineral which strongly reduces the
permeability in reservoir sandstones (Fig. 5). The
percentage of authigenic illite is difficult to
quantify by XRD because of interference from
clastic illite and mica. A strong increase in the
amount of illite relative to kaolinite in the clay
fraction is observed below 3.7-4.0 km, both in the
northern North Sea (Giles et al., 1992; Bjorlykke et
al., 1992) and Haltenbanken (Bjorlykke et al.,
1986; Ehrenberg & Nadeau, 1989). North Sea
Jurassic reservoir sandstones which have been
buried to depths <3.5 km generally show little
pore-filling illite in thin-section or by SEM. Even if
the pore-water composition in the North Sea Basin
in most cases falls in the stability of illite, little
precipitation occurs due to an extremely low kinetic
22
K. Bjorlykke
ol
""...
""..........
60
K-Feldspar
"%
%~
,
80 ~E]~ o ~
""''""" ................................
o•'•lOO
g
,
tll
"\
1~.
o
Illite
D Viking Graben, < 50,000 ppm CI
%o
9 Viking Graben, > 50,000 ppm CI
99
,, Central Trough, < 50,000 ppm CI
140t~-111
I1~
9 Central Trough, > 50,000 ppm CI
~60 1 I.A~
o Haltenbanken, < 50,000 ppm CI
7~T
9 Haltenbanken, > 50,000 ppm CI
l
~
Kaolinite~\
~8o, ,~
0
, ~ , . ,
500
.....
...,,
....
K
(ppm)
, ....
, ....
, ....
I
1000 1500 2000 2500 3000 3500 4000
Equilibrium between kaolinite and illite at 0% salinity
Equilibrium between kaolinite and illite at 10% salinity
........................ Equilibrium between illite and K-feldspar at 0% salinity
......
Metastable equilibrium between kaolinite and K-feldspar
at 0% salinity
FIG. 4. Geochemical composition of formation water from North Sea and reservoirs in relation to the stability
field of kaolinite, illite and K-feldspar. All the pore-water analyses fall in the stability field of illite but
precipitation of illite depends on the available A1 from precursor minerals and on the kinetics which is very slow
below 120-140~ At higher temperatures, the pore-water composition falls close to the boundary between the
stability field of kaolinite and illite as should be expected when kaolinite is replaced by illite. From Bj~rlykke et
al. (1995).
precipitation rate at low temperature (< 120-140~
(Bjgrlykke et al., 1995).
High concentrations of authigenic illite are nearly
always associated with dissolution of an unstable
aluminous mineral phase which in North Sea
Jurassic reservoirs is mostly kaolin. Another
precursor mineral for illite is smectite but analyses
of shallow samples (<2 km) suggest that the
Jurassic sandstones had a low primary smectite
content. Tertiary sandstones, however, may have
had a high smectite content (Bjorlykke et al., 1995).
Authigenic illite may form by different reactions
(Bjorlykke et al., 1995):
smectite + K § = illite + silica (via mixed-layer
minerals) (1)
A12SiOs(OH)4 + KAISi308 =
Kaolinite
K-feldspar
KA13Si3Olo(OH)2 + 2SIO2 + H20 (2a)
Illite
Quartz
3A12Si2Os(OH)4 + 2KA1Si308 + 2Na + =
Kaolinite
K-feldspar
2KA13Si301o(OH)2 + 2NaA1Si308 + 2H + +
3H20 (2b)
Illite
Albite
Clay mineral diagenesis in sedimentary basins
23
FIG. 5. Pore-filling authigenic illite from a Jurassic reservoir, Haltenbanken (4.2 km depth).
The distribution of smectite in the North Sea
sediments suggest that smectite dissolves at
temperatures of ~65-75~
and at 80-100~ the
mixed-layer minerals contain >70% illitic layers
(Dypvik, 1983). This reaction depends, however, on
several other factors including the supply of K and
the time factor (Boles & Franks, 1979).
Reaction 2a is isochemical and does not require
the supply or removal of ions by pore-water flow. It
does require, however, that K +, A13+ and silica are
transported from the surface of the dissolving
K-feldspars and kaolinite to the site of illite
growth. The SEM pictures frequently show that
the authigenic illite growth is closely associated
with or replacing dissolved kaolinite. The ratelimiting process for illite growth will then be the
kinetics of illite precipitation and the transport of
K + from dissolving K-feldspar. At low temperature
and slow precipitation rate, the pore-water will be
highly supersaturated with respect to illite and less
under-saturated with respect to K-feldspar, thus
reducing the dissolution rate of K-feldspar and the
diffusive transport of K. In the second reaction,
illitization is combined with albitization and there is
no excess silica which can be precipitated as quartz.
There are several pieces of evidence suggesting that
illite requires relatively high temperatures to form.
(1) Geochemical analyses of pore-water from the
North Sea show that the pore-waters are mostly
supersaturated with respect to illite (Fig. 4). (2) A
strong increase in the amount of diagenetic illite is
observed in reservoir sandstones at depths close to
3 . 8 - 4 . 0 km c o r r e s p o n d i n g to 1 2 0 - 1 4 0 ~
(Bjorlykke et al., 1986; Ehrenberg, 1990).
(3) Illite may also form from dissolving smectite
at somewhat lower temperatures (80-100~
The stability of smectite is reduced as quartz
starts to precipitate, reducing the supersaturation
with respect to quartz (Aagaard & Helgeson, 1983;
Sass et al., 1987). The rate of illitization of smectite
also depends on the supply of K (Hower et al.,
1976) and in sediments without zeolites and
K-bearing evaporites, this will mainly be K-feldspar.
It is still not known how far K can be transported
by diffusion within sandstones and between
sandstones and shales. Detailed petrographic and
24
K. Bjorlykke
XRD analyses from the Garn Formation at
Haltenbanken show, however, that samples with
relatively high kaolinite content at 4 km depth have
little K-feldspar (Ehrenberg, 1991). Potassium
appears not to have been supplied by diffusion
from K-feldspar 1 0 - 2 0 m away. Mudstones
containing smectite may represent a sink for K
from adjacent sandstones during illitization but if
the mudstones contain K-feldspar there is no
concentration gradient to drive such transport.
Illite datings. A large number of K-Ar dates of
illite from the North Sea and Haltenbanken have
been published. The ages obtained from Jurassic
reservoir sandstones range from 100-30 Ma, often
with a concentration of ages between 40-60 Ma
(Thomas, 1986; Liewig et al., 1987; Jourdan et al.,
1987; Glasmann et al., 1989a,b). Analyses of Jurassic
sandstones from Haltenbanken gave ages from
55-31 Ma, but Ehrenberg & Nadeau (1989)
interpreted these ages to be much too old due to
contamination of old feldspar and illite and
interpreted the illite to have formed in the last few
Ma at temperatures close to 140~ Also from other
basins, like the Paris Basin, the possible detrital
contamination of illite and the validity of K/Ar
datings have been debated (SpiStl et aL, 1996; Clauer
et al., 1996). If these illite dates represent the time of
the main phase of illite growth, this poses several
problems. The problems of such datings have
recently been discussed in detail by Clauer &
Chaudhuri (1995). If the Early Tertiary Jurassic
reservoir sandstones, presently at -4 km depth, were
buried to only about 2 km or less, the illite must then
have formed at rather low temperatures (50-80~
or the geothermal gradients must have been about
twice that of the present day for a long time in the
Late Cretaceous and Early Tertiary. However, the
distribution of authigenic illite in reservoir rocks
seems to be strongly controlled by the present burial
depth. In the North Sea and the Haltenbanken basins,
there is a strong increase in the amount of illite
below 3.7-4.0 km depth. Below this depth authigenic kaolinite can commonly be observed to have
been replaced by illite (Bjorlykke et aL, 1986,, 1992;
Ehrenberg & Nadeau, 1989; Ehrenberg, 1990).
In the Brent Group, a marked decrease in the
K-feldspar content is commonly observed
suggesting that it has been dissolved in the
process of illitization of kaolinite and possibly
also smectite (Bjorlykke et al., 1992). In the Brent
Group, the illite content in sandstones commonly
increases at 3 . 7 - 4 k m (11000-12000 It) depth
(Giles et al., 1992; Scotchman et aL, 1989). This is
the same depth as Haltenbanken despite lower Late
Cenozoic subsidence suggesting that the illitization
is controlled by the present depth.
A lower degree of illitization and more unaltered
kaolinite have been observed in reservoir sandstones like those of the Hild Field where the Kfeldspar content is very low, suggesting that the
supply of K is the limiting factor for illitization
(Lonoy et al., 1986; Bjorlykke et al., 1992;
Thyberg, 1993).
The amount of kaolinite formed at shallow depth
can be the limiting factor for the formation of illite
during deeper burial diagenesis, particularly in the
distal shelf and ~n'bidite facies where meteoric
water flushing is not usually very effective and
contains little authigenic kaolinite, thus reducing
the potential for illitization (Fig. 2).
The illite datings suggest that authigenic illite
formed at 75~ at -2 km or less (Hamilton et al.,
1992). As documented above, a pronounced
increase in the illite content is observed at
-3.7-4.0 km depth, both in the North Sea, although
the thickness of the Pliocene/Pleistocene is much
greater at Haltenbanken (1 km) than in the North
Sea, suggesting that the distribution of illite is
controlled by the present burial depth. Observations
both from the North Sea and Haltenbanken
(Ehrenberg, 1991; Lonoy et al., 1986; Thyberg,
1983) show that kaolinite remains stable to higher
temperatures than 140~ (4 km) when K-feldspar is
not locally available to supply the K. The reduced
illitization and improved reservoir quality can then
be related to the amount of clastic K-feldspar.
Chlorite. Chlorite is common as a clastic mineral
in the Pliocene-Pleistocene sequences of the North
Sea Basin because of limited weathering in this
partly glacially-influenced cold climate. In the
warmer climate of the Lower Tertiary and
Mesozoic, the clastic sediments supplied to the
North Sea probably contained little chlorite. Most
of the chlorite minerals in the Lower Tertiary
sequence were probably formed diagenetically
during early diagenesis on the sea floor or by
burial diagenesis from smectite or volcanic detritus.
Similar smectite-rich mudstones of volcanic origin
have been found in Upper Mesozoic and Lower
Tertiary sequences along the Atlantic margin
(Chamley, 1992). Authigenic chlorite occurs as
grain-coating cement in some sandstones and is
particularly common in the Jurassic Tilje Formation
at Haltenbanken (Ehrenberg, 1993). Grain-coating
Clay mineral diagenesis in sedimentary basins
25
FIG. 6. Authigenic chlorite coating quartz grains retarding the growth of authigenic quartz thus preserving
abnormally high porosity. From Jurassic reservoirs buried to -5 km at Haltenbanken.
chlorite cement (Fig. 6) inh"bits quartz overgrowth,
thus preserving higher porosity than normal at
depths of 4.0-5.5 km (Ehrenberg, 1993). Chlorite
crystals become more coarse grained with
increasing burial depth as a result of grain
coarsening (Jahren & Aagaard, 1989).
Precipitation of chlorite requires a source of Fe
and Mg and possible sources are clastic biotite, basic
rock fragments and volcanic rock fragments or early
diagenetic Fe minerals (verdine and glaucony)
formed in deltaic or estuarine environments by the
supply of Fe from rivers as demonstrated from the
Niger Delta (Odin et al., 1988; Ehrenberg, 1993). It
is possible that such early diagenetic Fe minerals
formed in tidal or estuarine environments could be
important precursors for chlorite coatings forming at
greater depth. In sandstones of the Brent Group, the
North Sea chlorite cement is in most cases rare or
absent (Giles et al., 1992; Bjorlykke et al., 1992),
possibly because it represents mostly a proximal
marine and fluvial facies.
Lower Jurassic Statfjord Formation and Intra
Dunlin sand in the Veslefrikk Field of the North
Sea, however, contain chlorite coatings (Ehrenberg,
1993; Hillier, 1994).
DIAGENESIS
OF M U D S T O N E S
SHALES
AND
Mechanical compaction is the reduction in volume
due to reorientation and breakage of grains, a
function of grain strength and effective stress.
Chemical compaction involves mineral dissolution
and precipitation and is a function of mineral
stability and kinetics of precipitation of cements,
processes that are strongly influenced by temperature. These two types of compaction have very
different driving forces and must be treated
separately. In basin modelling, however, compaction of mudstones is assumed to be a function of
effective stress (Hermanrud et al., 1991; Illiffe &
Dawson, 1996). Hermanrud pointed out the great
variations in shale porosity in the published data,
i.e. from Rieke & Chillingarian (1974). Similar
variations are found in North Sea mudstones but
these are functions of the initial grain size and
26
K. Bjgrlykke
mineralogy and can be predicted. As a general rule,
it can be stated that with increasing burial depth,
the rate of compaction becomes more chemical and
more a function of temperature and less of the
effective stress.
Mechanical compaction o f mudstones
The North Sea and Haltenbanken basins are
characterized by three main types of mudstones and
mixtures of them: (1) Glacial marine mudstones of
Pliocene and Pleistocene age. These are mineralogically immature mudstones with a high feldspar
content, dominantly clastic chlorite and also,
frequently, unstable rock fragments such as
pyroxenes and amphiboles (Thyberg, personal
communication; Rundberg, 1989; Karlsson et al.,
1979). (2) Smectitic mudstones, mostly of Lower
Tertiary age. In particular, Eocene and Oligocene
mudstones representing a distal facies may have a
very high smectite content (>50%) and almost no
quartz or feldspar (Huggett, 1992; Thyberg,
personal communication). These mudstones are
derived from volcanic ash resulting from subaerial
volcanicity during the opening of the NorwegianGreenland Sea. It is also present onshore in
Denmark (Nielsen & Heilman-Clausen, 1988).
Mudstones of more proximal facies contain more
kaolinite and quartz (Rundberg, 1989; Thyberg et
aL, 1998). (3) Mesozoic mudstones and shales
consisting mostly of illite and variable amounts of
kaolinite, smectite and mixed-layer I-S and chlorite.
Most of these sediments have been buried to
>2.0-3.0 km and the smectite and the mixedlayer content may have been higher at shallower
depths of burial. These types of mudstones have
very different properties, particularly during
mechanical compaction (Rieke & Chillingarian,
1974). Mudstones and shales should, therefore, not
be treated as one category during basin modelling.
The difference between these types, with respect to
compaction, is clearly seen on velocity and density
logs from the North Sea (i.e. 34/7-1) (Thyberg,
personal communication).
The Pliocene and Pleistocene glacial marine
mudstones are characterized by high velocities
(2.5-2.7 kin/s) and densities (low porosity),
producing a strong velocity and density inversion
compared to the underlying smectitic mudstones,
which have typical velocities o f - 2 . 0 km/s. The
driving force for mechanical compaction is the
effective stress (~e) transmitted at the grain
contacts. Compaction causes reduction of porespace and can only occur if the fluid (water) is able
to escape. The coarse-grained glacial marine
sediments are relatively permeable, allowing
rapid dewatering, and thus avoiding overpressure.
The smectitic mudstunes have very high surface
area and low permeability and in such fine-grained
sediments the rate of fluid expulsion may be the
rate-limiting process in compaction (Thyberg,
personal communication). Fractures in mudstones
can only remain open if the pore-pressure is equal
to the least horizontal stress (~h), which in such
sediments is 80-90% of the vertical stress (Crv)
(Garenstroom et al., 1993). The effective stress
(~e) to drive compaction would then be very small
and there would be low pressure gradients for
water to flow into the fracture. In addition, the
fracture must connect with high permeability
pathways all the way to the surface. According to
Gaarenstrom (1993) the degree of overpressure in
the Central graben may have been increasing
through the Tertiary reaching fracture pressures
in the Late Tertiary. The effective stress would
then have been reduced and the mechanical
compaction must have stopped, but the chemical
compaction which is not so sensistive to stress has
continued.
Chemical compaction o f mudstones and shales
The onset of chemical compaction depends on the
stability of the mineral phases. As discussed above,
smectite is replaced by mixed-layer minerals and illite
in the temperature range of 70-100~ which, in
basins like the North Sea, corresponds to burial
depths of 2 - 3 km and this causes compaction which
could not have been obtained mechanically.
Experimental data by Chillingarian & Knight (1960)
suggest that even at effective stresses corresponding
to 15000 ft, the porosity of the montmorillonite clay
is only reduced to ~45%, while kaolinite and illite are
compacted much more efficiently at the same
overburden stress (Fig. 7). This suggests that smectitic
mudstones can only compact mechanically to
-30-40%, depending on the smectite content. Both
porosities derived from density logs (Fig. 7) and from
analyses of core samples (Tyridal, 1994; Tyridal,
personal communication) suggest that Lower Tertiary
smectitic mudstones have porosities in this range
down to ~2 km. Further compaction probably
depends on the transformation of smectite to illite
via mixed-layer minerals, which increases the particle
Clay mineral diagenesis in sedimentary basins
27
Gulf C o a s t T e r t i a r y
1
It
~
-
Venezuela Tertiary
2
-- - -- "- -
3 84
.5
Experimental
6
7
.o
Mio
O k l a h o m a - Palaeozoic
4
o
0
o
C o m l S o s i t e - R e c e n t
.....
Gulf C o a s t - T e r t i a r y
.......
JQpo~- Tertiary
Experimental
8
L
n
g
0
--"
10
n
r~
11
to
12
13
o
E
14
1
50
0
"
20
40
t
i
60
80
Porosity ('/.)
F~G. 7. Experimental and observed porosity depth curves from mudstones. Porosity depth data (points) for
Tertiary mudstone from the Central Graben based on density logs and estimates of mineral matrix composition
from six wells (Lauvrak, 1996). The mudstones are fine grained and smectite-rich and show little reduction in
porosity down to nearly 2 km depth. This may be due partly to moderate overpressure reducing the effective
stress. Experimental compaction curves (Chillingarian & Knight, 1960) show that pure clay minerals do not
compact readily even at stresses equivalent to 5 km. The much greater rates of compaction between 2 - 3 km
depth may reflect chemical compaction, mainly dissolution of smectite and precipitation of mixed-layered
minerals and illite.
size and therefore the permeability and the rate of
compaction.
Also in shales, K-feldspar reacts with kaolinite to
form illite, but the reaction rate may be lower due
to slower diffusion of K from K-feldspar in the
shale m a t r i x w h i c h has low p o r o s i t y and
permeability.
Kaolinite in shales, as in sandstones, is subject to
dissolution and precipitation of illite at depths close
to 4 km (130-140~
if K-feldspar is present to
supply K. Authigenic illite is more fine grained
than the dissolving kaolinite and this process may
contribute to the reduction in permeability.
Clastic kaolinite may be a part of the clastic
framework carrying effective stress. When dissolved,
illite may precipitate in the available pore-space and
allow for more efficient compaction. Both the
reaction from smectite to illite and kaolinite to
illite release water, which may contribute to the
build-up of pore-pressure. Dehydration of minerals
28
K. Bjorlykke
involves a partial phase change from solid to fluid tectonic stress (low strain rates) and will therefore
thus increasing the porosity and fluid/solid. reduce the potential to transmit plate tectonic stress
Dehydration of clay minerals can generate a in sedimentary basins during subsidence (Bjorlykke
significant percentage of the total compaction- & Hoeg, 1997).
driven flux. Shales containing 20% kaolinite may
generate water corresponding to ~4% of the rock
DISCUSSION
volume which could contribute to the build-up of
overpressure (Bjorlykke, 1996). Most North Sea Predictions of burial diagenetic reactions depend on
mudstones, however, will probably have had a lower whether or not the chemical composition of the
initial kaolinite content. Compaction and generation sediments can be assumed to be constant during
of petroleum are in most cases important contribu- burial. Changes in the bulk composition during
tors to overpressure (Buhrig, 1989). If a constant burial must be due to transport in pore-water by
porosity is assumed and the permeability/depth curve diffusion or by fluid flow (advection). With
is constant, the modelling of overpressure will increasing temperature, the pore-water approaches
necessarily be a function of the sedimentation rate. equilibrium with the constituent minerals and the
However, since both mechanical and chemical concentration gradients for driving diffusion will be
sediment compaction of mudstones are a function very small. Where there are important differences
of time, high sedimentation rates will imply that the in the primary lithology and mineral assemblage,
porosity and permeability at a certain depth is also there may locally be higher concentration gradients
higher. The higher permeability may partly or totally between pore-waters buffered by different minerals.
compensate for the increased flux due to the higher The presence or absence of minerals like calcite, Kfeldspar and kaolinite may be critical for such
sedimentation rate. The observed porosity/depth
trends of mudstones vary greatly as a function of buffering producing concentration gradients near
grain size and mineralogy (Fig. 7). As shown above, lithological boundaries (Thyne et al., 1996). It is
there is a limit to how much mudstones will compact difficult to estimate how effective diffusion driven
mechanically and dissolution and precipitation of by mineral buffered pore-water is, but it is
minerals make it possible to form a mineral fabric significant probably to the order of a few metres.
the pore-water is
with a much lower porosity (Fig. 8). Mudstones At temperatures >80-100~
often have silt- or sand-sized quartz grains and probably close to quartz saturation and the K
dissolution and reprecipitation of quartz may be an concentration will normally be in the stability field
important factor in producing closer packing of illite and in some cases also chlorite. The K§ and
(Fig. 8). This process has been shown in sandstones Mg++ may be transported by diffusion locally but
to be mainly a function of temperature and textural the North Sea formation water does not have high
relationships and relatively insensitive to variations concentrations of these elements except near
in stress (Bjorkum, 1995; Walderhaug, 1996). The evaporites (Warren & Smalley, 1994). During
same processes probably also apply to mudrocks and early marine diagenesis near the sea floor there
shales but it is more difficult to study textural may be effective diffusion from sea water into the
relationships between quartz, mica and clay minerals sediments. Fluid flow rates are orders of magnitude
in such f i n e - g r a i n e d s e d i m e n t s (Fig. 9). greater during meteoric water flushing than during
Overpressures built up at depths where chemical burial diagenesis (Giles, 1987). The potential for
compaction is d o m i n a n t ( > 2 - 3 km depth, mass transport is also greater during early
70-100~
should not be expected to have diagenesis because at low temperatures the poresignificantly higher porosity than normally pressured water may be highly over saturated or under
rocks. Chemical compaction cannot be modelled saturated with respect to mineral phases. During
based on effective stress because the temperature is burial diagenesis, when the temperature is greater
the most important factor (Fig. 9). The transition and the flow rates are much smaller, the pore-water
from mechanical to dominantly chemical compaction is closer to equilibrium with the minerals present.
is not fixed and will depend on the mineralogy and The volume of minerals dissolved or precipitated
the burial history. Smectite-rich mudstones will due to advective flow can then be calculated (Wood
compact chemically at shallower depth than & Hewett, 1984; Bjorlykke, 1994):
kaolinite-rich mudstones. Both mechanical and
Vc = F t sin[3(0T/~Z) ~v/p
(3)
chemical compaction provide a ductile response to
29
Clay mineral diagenesis in sedimentary basins
Mechanical
IO'v (overburden
, ....... lit _ stre_ ss) "
a function
compactionof effective
X
I
t
i
I
,
I
I
I
__J
OC
) f" . . . . . . . . . .
stress
oe
4
I'+
Effective
stress at grain
contacts G e
-'Gv - P p
Chemical compaction involving dissolution and precipitation
of minerals mainly as a function of temperature
Smectitic mudstones (30-60% porosity)
(partly bound water)
70~
/
I
Via mixedlayer minerals
to
lo0oc , Illite
M ~ : : ~ - , .
~
l L{
{7~
t7_:,:~ _;, - ~ _
#,'::7:"
b__~_:•
~l
I
= Clastic kaolinite
Dissolution of clastic kaolinite
and precipitation of authigenic illite
in the avilable pores will
increase compaction.
FQrther compaction depends on the
dissolution and precipitation of quartz
FIG. 8. Schematic representation of mechanical and chemical compaction of mudstones.
Here, Vc is the volume of cement precipitated, F
is the total flux of pore-water (cm3/cm2 s--l), t =
time (s), ~ is the angle between the direction of
flow and the isotherm, ~T/~Z is the geothermal
gradient, Ctr is the solubility-temperature function
(transfer coefficient) and p is the density of the
mineral. In the case of clay minerals, the mobility
of A1 is particularly critical. Calculations using
K. Bj#rlykke
30
Porosity
IP
/
/
2-3
km
D
E
P
~
H
Effective stress:
~v=.(prgh-P)and
compaction modulus
"
i
T
j
/
compaction
"
4kin
me
.
/ Compact,on
/
t 70-1000C
Chemical
compaction
dominant.
Mainly a function
of temperature
and mineralogy
FIG. 9. Relationships between compaction, effective stress and temperature in mudstones. The depth at which
chemical compaction becomes dominant over mechanical compaction depends on the mineralogy and textural
relationships.
SOLMINEQ 88 show that the solubility of A1 is
<1 ppm at temperatures <140~
and that the
solubility is not increased in the presence of
o r g a n i c a c i d s ( B j o r l y k k e et al., 1995).
Significantly higher A1 concentrations have not
been reported from North Sea reservoirs. Because
of the low mobility of AI, growth of diagenetic clay
minerals like illite and chlorite require a local
precursor aluminous mineral like smectite and
kaolinite and the supply of K and Mg. The
distribution of illite and chlorite at depth must be
linked to provenance and facies and the distribution
of early diagenetic kaolinite due to meteoric water
flushing (Fig. 5). Modelling of compaction driven
flow indicates that flow rates are very low so that
the system is characterized by low Peclet numbers
and thus both heat transport and mass transport are
dominated by diffusion (Bethke, 1985; Ungerer et
al., 1990; Ludviksen et al., 1993; Bjorlykke, 1994).
Diagenesis is still, to a large extent, based on what
was traditionally called 'sedimentary petrology' and
careful mineralogical petrographic data and observations are always valuable. Diagenetic theories
should be tested and calibrated against observations,
but the processes cannot be inferred from petrographic data alone. Changes in sediment composition and mineralogy with burial depth are often
inferred to be due to diagenetic processes. In a
single well there are frequently very distinctive
stratigraphical variations in the lithology. To
observe changes in one stratigraphic interval with
depth, several wells where this unit is cored must
be studied and it must be considered that the
observed trend with depth may also be due to other
factors such as provenance and facies. Because of
lateral variations in facies and provenance, it is
impossible to study the same rock at different burial
depths. Sandstones may vary laterally with respect
to facies and provenance but this is also the case
with mudstones. An increase in illite and K content
with depth may be due to primary enrichment of
illite or smectite in the distal fine-grained facies.
Because there may always be lateral changes in the
primary composition of sandstones and shales,
geochemical changes cannot be inferred from
sampling at different burial depths.
If the sediments undergo diagenesis in a
geochemically closed system, the diagenetic reac-
Clay mineral diagenesis in sedimentary basins
tions and the rock properties of the sediments are
direct functions of the primary sediment composition and the burial history. Predictions of burial
diagenetic reactions must therefore be linked to
facies and provenance. Rock properties like
porosity, permeability, specific surface and
thermal conductivity all depend very much on the
primary clay content and on the burial diagenetic
reactions.
CONCLUSIONS
Clay minerals play a crucial role in both
mechanical and chemical compaction of sandstones
and shales. The amount and type of clay mineral is
a function of the provenance of clastic minerals and
of diagenetic reactions at shallow and greater depth.
Early diagenetic reactions may be relatively open
(from a geochemical point of view), due to
diffusion from seawater or due to meteoric water
flow. Dissolution of feldspar and mica and
precipitation of authigenic kaolinite require that K
and silica are removed by fluid flow as is the case
during weathering. During burial diagenesis, the
pore-water flow is very much smaller and
calculations show that advective transport of
dissolved ions is in most cases relatively insignificant. Diffusion, however, may be significant over
shorter distances depending on differences in
mineral composition controlling the concentration
gradients of the dissolved ions. Clay mineral
reactions are therefore close to being isochemicaI
within the range of effective diffusion transport
(1-10 m?). Sandstones without clay minerals such
as kaolinite and smectite and unstable rock
fragments will remain stable and porosity will
then be reduced mainly by mechanical compaction
and dissolution/precipitation of quartz and carbonate minerals.
Prediction of diagenetic reactions at greater
burial must be linked to provenance facies and
early diagenetic reactions. North Sea formation
water is nearly always in the stability field of illite,
but precipitation of illite requires an aluminous
precursor mineral like smectite or kaolinite and is
mainly a function of temperature and supply of K
from K-feldspar. Mudstones and shales from the
North Sea and Haltenbanken vary greatly in
mineralogical composition and grain size.
Immature, partly glacial Pliocene and Pleistocene
mudstones in which illite and chlorite are the main
clay minerals, compact much faster and have
31
greater velocities than the underlying Eocene and
Oligocene smectitic mudstone and commonly
develop overpressures. Dissolution of smectite and
kaolinite and precipitation of illite releases crystalbound water which adds to the fluid flux and may
potentially contribute to the build-up of overpressure. At burial depths >2.0-3.0 km, porosity
loss by compaction is mainly chemical, involving
dissolution and precipitation of minerals and this
compaction can not be calculated as a function of
effective stress as is assumed in most basin
modelling programmes. Realistic prediction of
rock properties can only be made if the primary
sediment composition and burial diagenetic
processes are known. Mineralogy, temperature and
time are the main factors controlling compaction
both of reservoir sandstones and shales, and the
degree of overpressure plays a minor role at depths
> 2 - 3 km (70-100~
During progressive burial,
sediments and mudstones in particular compact both
mechanically and chemically and show a ductile
response to stress at slow strain rates. This reduces
the transfer of tectonic stress in sedimentary basins.
ACKNOWLEDGMENTS
Financial support from the Norwegian Research
Council (NFR) and from Norwegian oil companies
and useful discussions with Per Aagaard are gratefully
acknowledged. Two anonymous reviewers provided
useful comments to help improve the manuscript.
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K. Bjorlykke
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