Middleton G.V. (Ed.) Encyclopedia of Sediments and Sedimentary Rocks
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MIXED SILICICLASTIC AND CARBONATE SEDIMENTATION
445
Punctuated and facies mixing are the most prevalent, and
both will lead to the lateral mixing of coeval juxtaposed
environments, producing lateral interfingering of elastics and
carbonates. Spatial mixing invokes local autocyelic processes,
and thus occurs along a timeline in an ancient deposit. Spatial
mixing requires no changes in relative sea-level.
Holocene examples that are affected by spatial mixing
include: (a) humid fluvio-deltaic systems proximal to a
carbonate
shelf,
such as the nearshore clastic belt and offshore
barrier reef of the Belize shelf (Wantland and Pusey, 1975),
and the northeastern Australian shelf and offshore Great
Barrier Reef (Maxwell and Swinchatt, 1970; Flood and Orme,
1988);
(b) Arid eolian dunes proximal to carbonate ramp
system, such as along the leeward southeast coast of the Qatar
Peninsula in the Persian Gulf (Shinn, 1973); (c) the deep axial
trough of Persian
Gulf,
where fine-grained elastics (clays and
silts) from the Tigris-Euphrates delta accumulate adjacent to
fine-grained carbonates of the Persian Gulf carbonate ramp
system (Purser, 1973). Ancient examples include Cambrian
ramp-to-deeper shale shelf transitions in the Upper Cambrian
intra-shelf basin of the Virginia Appalachians (Markeilo and
Read, 1981), and the Upper Musehelkalk (Triassic) outer ramp
deposits of the Catalan Basin (Spain; Calvet and Tucker,
1988),
among others (Houlik, 1973; Mount, 1984; Brett and
Baird, 1985; Grotzinger, 1986; Sonnenfeld and Cross, 1993;
Adams and Grotzinger, 1996; Cowan and James, 1996).
Temporal variations in environments or
evolution in sedimentation
Over geologic time, mixtures may occur as a result of relative
sea-level change and/or variations in sediment supply, which
causes vertical variations in the stratigraphic succession of
facies.
The concept of "cyclic and reciprocal sedimentation"
(Van Siclen, 1958; Wilson, 1967; Silver and Todd, 1969;
Meissner, 1972) which invokes changes in relative sea level was
developed to explain the origin of mixed clastic-carbonate
shelf-to-basin cyclic deposits. This model suggests that
carbonate sedimentation dominates during relative sea-level
highstands and rises in sea-level, when elastics are trapped
updip in flooded fluvial valleys and narrow clastic shelves while
the carbonate factory is fully operational. Additionally, during
such highstands, rates of carbonate sedimentation (between
20 cm to
30
cm/1000 years in lagoons to greater than
1
m/1000
years in reefs and shoals) are sufficient to keep pace with
typical background subsidence rates such that the subtidal
factory outeompetes any clastic influx. Shelfal carbonate
deposits are typically thick and the coeval carbonate basin is
starved and only thin carbonate deposits are preserved.
Clastic sedimentation dominates during lowstands when
carbonate shelves are subaerially exposed and thus largely shut
down. During a relative fall in sea-level, widespread braided
fluvial systems, shallow deltas or nearshore stiallow marine
elastics may migrate across the shelf completely and ultimately
source deep-water clastic deposits (Van Siclen, 1958; Wilson,
1969;
Meissner, 1972; Rankey etal., 1999), or elastics may
migrate laterally by longshore prcesses. During true lowstands,
fluvial-deltaic systems may incise the antecedent shelf (e.g., the
Pleistocene of Belize, Choi and Ginsburg, 1982) and funnel
clastic sediments to point-sourced submarine canyons into the
basin in humid climates, or in arid climates, eolian dunes may
migrate across subaerially deflated carbonate shelves, often
truncating older carbonate features (Goldhammer etal., 1991;
Borer and Harris, 1991; Atchley and Loope, 1993). During
such lowstands, the subaerially exposed carbonate shelf may
be karstified and often completely by-passed such that clastic
material is deposited on adjacent slopes and basinal areas as
onlapping sheets, wedges and fans that terminate against the
defunct carbonate slope, for example the Upper Permian
Capitan depositional system of west Texas and New Mexico
(Bebout and Kerans, 1993; Tinker, 1998). Such basinal
deposits may be then subject to reworking by contour-hugging
bottom currents. The cyclic and reciprocal model further
suggests that lowstand shelfal elastics will be thin and coeval
basinal elastics (submarine fans, turbidites) will be much
thicker. However, the height and steepness of the underlying
carbonate slope is an important consideration in determining
the character of any lowstand clastic sand body. Alternatively,
if an updip clastic source is limited, basinal settings may
receive little or no clastic material and only record the
shutdown of the carbonate platform in the form of submarine
diastems (e.g., mineralized, cemented hardgrounds).
The character of the lowstand clastic sediments reflects
both proximity to the source area and depth of deposition.
For example in the late Devonian of the Canning Basin
(western Australia), Southgate etal. (1993) record tiie evol-
ution of several Frasnian-Fammenian "third-order" deposi-
tional sequences whereby the shelfal carbonates are completety
bypassed. Within the basin, however, there exist several
sequences of lowstand elastics each composed of basin-floor
fans,
slope fans, and prograding complexes, which overall
form a fining-upward trend from basal coarse elastics to fine-
grained mudstone and shales vertically within a particular
lowstand. This grain-size trend is interpreted to depict initial
stream rejuvenation upon initial subaerial shutdown of the
carbonate
shelf,
followed by subsequent waning of coarse
proximal sediment supply as renewed transgression of the shelf
occurred.
An unequivocal and detailed example of mixed clastic-
carbonate sedimentation driven by sea-level changes is
provided by the Quaternary reefs in the southernmost Belize
iagoon (Choi and Ginsburg, 1982). During the last glacial
maxima, eustatic sea-level was about
120
m beneath its present
day level, exposing and karstifying much of the Belize
carbonate
shelf.
As coastal rivers advanced seaward onto the
shelf,
they created morphological features including fluvial and
submarine channels, fluvial channel bars, levess and deltaic
lobes.
Subsequently, when eustatic sea-level rose, such clastic
features with positive morphologic relief served as nueleation
sites and nodes for the initial development of Holocene reefs
and other biogenic carbonates. Another example, provided
by Davies et al. (1989) illustrates the cyclic and reciprocal
nature of Cenozoic clastic-carbonate sequences develped in
response to eustasy along the northeastern shelf of Australia
(Figure Ml8).
Conclusions
Mixed siliciclastic and carbonate sedimentation is a common
theme throughout the stratigraphic record. Holocene systems
enable us to examine mixtures due to spatial variability, where
mixing occurs principally by lateral facies mixing of coeval
sedimentary environments. Studies of spatial mixing empha-
size process and provide clues to examing ancient mixed
successions. Where such examples are cyclically inter-
bedded, they are often attributed to temporal evolution in
446
MIXED SILICICLASTIC AND CARBONATE SEDIMENTATION
(A) High Sea Level Stage
Sea
Level
Lagoonai Carbonates
Clastic Dominated
Carbonate Dominated
Slope
Facies
(B) Low Sea Level Stage
Subaeriaiiy-Exposed
Fioodplain Deposits
Reef...,
SheifEdge
Fan Deita
Sea Level
Clastic Dominated
Figure M18 Schematic sections showing high (A) and low (B) sea-level control on the stratigraphic evolution and sedimentary geonnetry of
shelf facies in the central Great Barrier Reef province. Note the predominance of siliciclastic facies in the lowstand. Modified from Davies et al.
(1989).
sedimentation, induced by sea-tevel changes arid/qi- variations
in sediment supply, causing a vertical variatior) in \hp
stratigraphic succession. Many, if not most, ancient mixed
successions can be interpreted as examples of cyclic and
reciprocal sedimentation, controlled by cyclic variations in
relative sea-level.
Robert K. Goldhammer
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Cross-references
Facies Models
Neritie Carbonate Depositional Environments
Sands
and
Sandstones
MIXED-LAYER CLAYS
Clay mineral crystals contain from
a few to few
hundred
silicate layers
(1:1 or
2:1, dioctahedral
or
trioctahedral,
see
figure
I in
Illite Group
Minerals), spread over tens
to
thousands
448
MIXED-LAYER CLAYS
of nm in the a* and b* crystallographic directions and stacked
parallel in the c* direction. Layers are either electrically neutral
or they have a negative charge, resulting from isomorphous
substitutions within the tetrahedral and/or octahedral sheet.
The charged layers are bounded by interlayer cations or sheets
of interlayer hydroxide. If the interlayer cations are hydrated,
the interiayer (smectite or vermiculite type) has swelling
properties (i.e., it may change thickness and accept organic
molecules). Clay crystals may contain non-identical layers and/
or interlayers, that is, be mixed-layered (interstratified).
Mixed-layering is defined as random and characterized by
the ordering parameter (Reichweite) RO if there is no preferred
sequence in stacking of layers (or interlayers) A and B. tf some
sequences are privileged, such mixed-layering is called ordered
(Rl if ABAB..., R2 if AABAAB... etc.; details in Moore and
Reynolds, 1997). If only one type of sequence is present, the
mineral is called regular. Such minerals (only of RI type) exist
in nature; they are given separate names, and their origin is
explained by different mechanisms. Other names of mixed-
layer clays are combinations of the components' names, for
example, illite-smectite.
Mixed-layering is identified and quantified by means of
X-ray diffraction (XRD), supplemented by chemical analysis
and electron microscopy. OOI XRD reflections of mixed-layer
minerals appear between positions of corresponding reflec-
tions of pure components. The nature of layers and swelling
interlayers is revealed by XRD of specimens heated or
treated with different exchange cations or organic solvents.
Computer modeling of such diffraction effects is used to
measure layer ratios and ordering (Moore and Reynolds,
1997).
tf the imaging conditions are properly selected, and a
distinctive and stable under vacuum spacing of the swelling
interlayer achieved, the nature of layers can be identified in
transmission electron microscope by measuring the layer
spacings and obtaining AEM chemical analyzes of the bulk
crystal.
Mixed-layer clays with smectitic interlayers can be separated
in dilute suspensions, by infinite osmotic swelling along these
interlayers, into individual layers or blocks of layers bounded
by stable interlayers (fundamental particles, see figure 1 in
Illite Group
Minerals). When such separation is stabilized in the
solid state, the mixed-layering diffraction effect is destroyed
and such blocks of layers (individual layers do not diffract
coherently) can be studied by XRD, as if they were pure end
member clays (Ebert etal., t998).
Mixed-tayer clay minerals are most often intermediate
products of reactions altering pure end-member clays.
Mixed-layer minerals occur in natural environments ranging
from surface to low-grade metamorphic and hydrothermal
conditions (see review by Srodoti, 1999). The majority of
mixed-layer clays have been synthesized in the laboratory.
Most often, the mixed-layering is 2-component, but more
complicated interstratifications have also been documented.
Mixed-layer clays are either di or trioctahedral; di/trioctahe-
dral interstratifications are very rare. Most mixed-layer clays
contain swelling interlayers of smectite type (typical of
diagenetic alteration) or vermiculite type (typical of weathering
alteration). In general, the weathering reactions producing
mixed layering are reversals of the corresponding high
temperature reactions, but the reaction paths are quite
different. Solid-state transformation and dissolution/crystal-
lization are the two mechanisms responsible for the formation
of different mixed-layer clays.
Dioctahedral mixed-layer clays
Kaolinite-Smectite (K-S)
K-S is known from Recent soils developed on acidic to alkaline
parent rocks and from the corresponding paleosols (Hughes
etal., 1993). If present in sedimentary rocks, K-S is most often
a detrital component, which survives burial temperatures of at
least 100°C. Its further evolution has not been traced. Most
probably, the abundance of K-S is underestimated, because it
is difficult to detect and also can be misidentified by XRD as
halloysite.
K-S is typically very fine-grained, randomly interstratified,
and it always evolves from smectite (origin from kaolinite has
not been reported). The reaction has been reproduced in
laboratory and is promoted by a high supply of Al.
lllite-Smectite (I-S)
I-S is the most abundant, most ubiquitous and the most
studied mixed-layer clay. 1-S can be detected by XRD in the
<0.2 \im fraction of almost every sample of sedimentary rock.
Smectite or recycled I-S is stable in sedimentary environments
and starts to react toward more illitic composition at about
70°C.
This reaction has been originally interpreted as fixation
of potassium in smectitic (swelling) interlayers plus solid-state
transformation of layers, but currently it is explained as
dissolution of smectite layers and simultaneous nueleation of
2 nm illite particles in the swelling interlayers and subsequent
growth of these particles (see Illite Group Minerals).
Evolution of I-S is continuous, always from RO (decreasing
to ca. 45 percent smectite), through R1 (^down to 20 percent S),
to R>1, and finally to illite. Regular 1:1 interstratification (K-
rectorite) is rare, known mostly from hydrothermal environ-
ments. Chemical evolution involves enrichment of
I-S in K (and in some cases NH4) and Al, a decrease in
cation exchange capacity, and liberation of Ca, Mg, Fe, and
Si,
which contribute to quartz and carbonate cementation.
I-S crystals have a lognormal distribution of thickness with
the mean of 5-6 layers/crystal, which increases below
20 percent S.
The percent S in I-S is an indicator of diagenetic grade,
analogous to organic indicators (see Maturation, Organic), and
can be used for evaluation of maximum paleotemperatures
and their K-Ar dating (see Illite Group Minerals). Because of
its abundance and swelling layers, I-S controls mechanical,
electrical and exchange properties of most shales.
Weathering of illite or I-S is not a simple reversal of
diagenetie reaction and is not well understood. Smectite can be
produced directly by dissolution-precipitation (Sucha et al.,
2001) or by transformation via ordered illite-vermiculite and
vermiculite (Wilson and Nadeau, 1985).
Glauconite-Smectite (G-S)
^
Minerals in monomineral glauconite pellets are iron-rich
analogues of 1-S. Glauconitization (see Glaucony and
Verdine)
proceeds at the sediment/water interface, through the same
steps of random and ordered interstratifications as illitization.
Regular interstratification has not been observed. Anomalous
chemical characteristics of some samples may result from
interlayering with berthierine (iron serpentine). Data are
lacking on the evolution of G-S during burial diagenesis.
MIXED-LAYER CLAYS
449
Na-illite-smectite (rectorite)
This mineral is reported only as a regular variety, called
rectorite (allevardite in older literature), and its geological
occurrence is restricted to hydrothermal alteration zones and
low-grade metamorphic rocJcs. In both cases, it is associated
with pyrophyllite and occasionally with cookeite, which
indicates crystallization tempearatures of ca. 280°C. Some-
times,
Ca may prevail over Na, which is expressed by the
presence of margarite-like layers. Rectorite crystals are always
very thick, as is evident from their very sharp diffraction peaJcs
and electron microscope images.
Chlorite-smectite (tosudite)
This clay is known only as a regular interstratification (named
tosudite) of beidellite (Al-smectite) and di-dioctahedral chlor-
ite (donbassite), or di-trioctahedral chlorite (sudoite, Mg-rich;
or cookeite, Li-rich). Likewise, only regular interstratification
was produced in hydrothermal experiments. Tosudite is known
from hydrothermal, pneumatolytic, or low-temperature meta-
morphic alteration zones, in association with dioctahedral
chlorite, serpentine, rectorite, and pyrophyllite. This associa-
tion suggests that tosudite forms during acid hydrothermal
alteration, at temperatures well below 35O°C.
Mica (Illite)-Vermiculite (I-V)
I-V is a product of potassium removal during weathering of
micas and illite. It occurs as thick crystals, with tri-dimensional
order, frequently with almost regular interstratification. The
vermiculitic layer may evolve toward smectite, giving rectorite-
like clay (Wilson and Nadeau, 1985). Partial chloritization of
vermiculitic interlayer is, however, more common. Clays of
this type are called soil vermiculites, soil chlorites, chlorite
intergrades, or swelling chlorites (Barnhisel and Bertsch, 1989).
Formation of soil vermiculites is a massive phenomenon in
temperate climates and this material has been identified in
fresti sediments. Its further evolution during diagenesis
remains obscure.
chloritization at elevated temperatures (commonly during
burial diagenesis). The process is not fully analogous to
illitization of smectite and is not fully understood. Instead of
a continuous process, overlapping succession of three phases
is observed: C-S with <20 percent C, corrensite (regular
interstratification), and C-S > 85 percent C. Diagenetic
corrensite appears in sedimentary rocks at 60°C to 160°C. A
decrease of Si and interlayer exchangeable cations, and an
increase of tetrahedral Al accompanies chloritization.
Weathering of chlorite commonly produces corrensite as
an intermediate phase. Distinct from diagenetic corrensite,
this mineral is characterized as chlorite-vermiculite, that is,
high-charge corrensite. The final product is vermiculite (^. v.),
which may evolve latter into smectite. The process involves
loss of Fe and Mg to such extent that the end member
vermiculite is strictly dioctahedral (Proust etal., 1986).
Mica-Chlorite (M-C)
M-C is typically close to regular interlayering and the mica
component is almost exclusively biotite. All M-C minerals
occur as big crystals producing excellent quality electron
microscope images. The majority of described occurrences are
from hydrothermally altered rocks or the products of low to
medium-grade regional metamorphism of granite, gneiss,
basalt, or pelites. M-C may form both by alteration of biotite
and chlorite.
Talc-Smectite (T-S)
T-S clays are known from saline lake and hot-spring deposits,
weathered ophiolites and dolomites. They are frequently
associated with sepiolite. Full range of interstratification was
reported (Wiewiora et al., 1982), including regular aliettite.
The temperature of formation (hot springs) and solution
chemistry are probably the factors responsible for the variation
in expandability. It remains unclear whether the reaction
proceeds from smectite to talc, or vice versa, or in both
directions.
Trioctahedral mixed-layer clays
Serpentine-Chlorite (Sp-C)
Sp-C is exclusively Fe-rich clay (berthierine-chamosite). It is
known to form in Recent "verdine facies" sediments in very
shallow tropical seas (see Glaucony and Verdine), often as
sandstone grain coatings. Either both layers crystallize
simultaneously, or chloritization starts at the sea bottom.
During burial, the berthierine content decreases and disap-
pears completely at the transition zone between shales and
slates.
Sp:C ratio can be now quantified precisely by XRD
(Moore and Reynolds, 1997). Regularly interstratified Sp-C is
called dozyite. Sp-C interstratification is accompanied by
interstratification of polytypes.
Chlorite-Smectite (vermiculite) (C-S)
C-S is the second most common mixed-layer mineral in
sedimentary rocks, often occurring together with I-S, and is
characteristic of hypersaline facies, volcanoclastic rocks and
graywackes (Reynolds, 1988). Unknown from Recent
sedimentary environments, C-S is a product of smectite
MIca-Vermiculite (M-V)
M-V is a common product of K-removal during weathering of
biotite. The reaction is easily reproducible in laboratory. The
complete compositional range is known (Pozzuoli etal., 1992),
including a regular variety (hydrobiotite). Little is known
about the behavior of M-V at the subsequent stages of the
rock cycle.
Jan Srodori
Bibliography
Barnhisel, R.I., and Bertsch, P.M., 1989. Chlorites and hydroxy-
interlayered vermiculite and smectite. In Dixon, J.B., and Weed,
S.B. (eds.). Minerals in Soil Environments. Soil Science Society of
America, pp. 729-788.
Eberl, D.D., Nuesch, R., Sucha, V., and Tsipursky, S., 1998.
Measurement of fundamental particle thicknesses by X-ray
diffraction using PVP-10 intercalation. Clavs atid Clay Minerals,
46:
89-97.
Hughes, R.E, Moore, D.M., and Reynolds, R.C. Jr., 1993. The
nature, detection, occurrence, and origin of kaolinite/smectite. In
Kaolin Genesis and Utilization. The Clay Minerals Society, Special
Publication, 1, pp. 291-323.
450
MIXING MODELS
Moore, D., and Reynolds, R.C, t997. X-Ray Diffiaction and the Iden-
tification and Anaiysisof Clay Minerals. Oxford University Press.
Pozzuoli, A., Vila, E., Franco, E., Ruiz-Amil, A., and de la Calle, C,
1992.
Weathering of biotite to vermiculite in Quaternary lahars
frotn Monti Ernici, central Italy. Clay Minerals, 11: 175-184.
Proust, D., Eymery, J.P., and Beaufort, D., 1986. Supergene
vermiculitization of a magnesian chlorite: iron and magnesium
removal process. Clays and Clay Minerals, 34: 572-580.
Reynolds, R.C. Jr., 1988. Mixed layer chlorite minerals, tn
Hydrous Phyllosilicates (exelusive of
tnieas),
Mineralogical Society
of America, Reviews in Mineralogy, 19, pp. 601-629.
Srodoii, J., 1999. Nature of mixed-layer clays and mechanisms of their
formation and alteration. Annual Review of Earth and Planetary
Seienees, 27:
19-53.
Sucha, V., Srodori, J., Clauer, N., Elsass, F., Eberl, D.D., Kraus, t.,
and Madejova, J., 2001. Weathering of smectite and illite-smectite
in Central-European temperate climatic conditions. Clay Minerals,
36:
403-419.
Wiewiora, A., Dubiriska, E., and twasiriska, I., 1982. Mixed-layering
in Ni-containing talc-like minerals from Szklary, Lower Silesia,
Poland. In Proeeedings of tlie International Clay Conference, ttaly:
Bologna-Pavia, pp. 111-125.
Wilson, M.J., and Nadeau, P.H., 1985. Interstratified clay minerals
and weathering processes. In Drever, J.I. (ed.). The Chemistry of
Weathering, D. Reidel, pp. 97-118.
Cross-references
Bentonites and Tonsteins
Berthierine
Chlorite in Sediments
Clay Mineralogy
Diagenesis
Glaucony and Verdine
Illite Group Clay Minerals
Kaolin Group Minerals
Mudrocks
Smectite Group
Vermiculite
Weathering, Soils, and Paleosols
principle of mass/chemical species conservation, and their
mathematical forms depend on the particular assumptions
made about the nature of the mixing. For example, mixing due
to the random movement of small infauna, the feeding of
head-down (conveyor belt) species, the infilling of vacated
burrows, and the resuspension of sediment by clams or waves
all generate different types of sediment motions and, in
principle, these differences are ret!ected in different mathema-
tical forms for the appropriate mixing model (Boudreau,
1997).
However, in practice, workers in this field tend to use a
simple array of models, that is, well-mixed models, diffusive
models, and nonlocal models. In addition, these models can be
distinguished further on the basis of the nature of the inputs
they require and outputs they produce. Thus, models can be
either of forward type, where the user inputs all parameter
values and expects a prediction of the distribution of a
sediment component with space and/or time, or they can be of
backward type, where parameters or the input history are
calculated from the observed depth profile of a sediment
property.
Well-Mixed Models. Mixing can be so intense that the
sediment component has uniform distribution (concentration,
C) in the mixed tayer. For example, the distribution of '^C
often exhibits a surficial zone some 4 cm to
10
cm thick of
constant "age", which retlects relatively fast mixing compared
to the half-life of this isotope. This type of mixing will cause
homogenization of input signals to the sediment and add long
tails of the signals when they enter the sediment record. An
example is schematically illustrated in Figure M19.
When mixing is so intense, the concentration of the
homogenized component, C (mass or numbers per unit
volume), in the mixed layer is given by
dC
dt
•-R
(Eq. 1)
MIXING MODELS
where t is time, Fjjf) is the input (flux) at the sediment-water
interface (mass or numbers per unit area and time), w is the
burial velocity relative to the sediment-water interface (length
per unit time), L is the depth of the surticial layer that is
mixed, and R represents any sources or sinks for the component
within the mixed layer, for example, radioactive decay of an
isotopic tracer (mass or numbers per unit volume and time).
Equation 1 balances the temporal change in concentration in
the mixed layer (left-hand side) with the sum of the input at the
In general, sediments are not faithful recordings of the history
of sedimentary input at a given location because they are
subject to a variety of postdepositional processes that can alter
their eomposition, structure, and therefore the apparent
record. Foremost amongst these processes is widespead mixing
of sediments by biological or physical agents. Geologists,
geochemists, biologists, and engineers often find it necessary to
quantify the effects of mixing on the distribution of solid
sedimentary components. For example, the fate of a buried
contaminant can be controlled by its unanticipated release
from the sediments by mixing. In addition, isotopic distribu-
tions can be significantly affected by mixing, such that
estimates of sediment accumulation rates become highly
skewed to unreasonably large values if the intluence of mixing
is not taken into account.
Correct prediction of the effects of mixing can be obtained
through mathematical models; these models are based on the
Concentration
Concentration Concentration
Initial
input spike
Rapid
homogenization
in mixed layer
After
deposition
equal to
several
mixed-layers
Figure M19 Schematic evolution ofa concentration spike deposited at
the interface of a sediment with a well-mixed layer of thickness.
MUDKtK.KS
451
itiferface. the output (bitrial) at the base of the mixed layer,
and the net ptoduction (consumption) by sources and sinks on
the right-hand side.
Diffusion Modets. Mixing is often not strong enough to
homogenize the active layer, and in this case the concentration
of the sediment component will also be a function of space, as
well as time. The most popular model in this case likens the
effects of mixing to those of diffusion (see ChemicalDiffiision).
whereby mixing imparts small random motions to the
sediment particles. The mathematical form of this model is.
at constant porosity,
(Eq. 2)
where. ,v is depth in the sediment. Ds is the tnixini! coeftieient.
analogous to the molecular/ionic diffusion coefficient (length
squared per unit titne), and the other terms are as in equatioti
(1).
Here the temporal change at a given depth (left) is the
result ofa balance (right) between the effects of mixing, burial
and reaction.
Fquation 2 is typically applied to isotopic data in the tnixed
zone to obtain Du. as the accumulation rate ir generally cannot
be extracted unatnbiguously from such data. AssutnJng steady
state {cCldt = 0). constant paratneters. R-/.C (A = decay
constant), and ir" <| 4/.Ds. the solution to equation (2) is
ln\C(.x)] = -(
fEq. 3)
where C\0) is the concentration/aetivity at the sediment-water
interface. An example fit of equation ? to ^'"Pb^.^ data to get
D/i
is iilustntted in Figure M20.
A large marine data base of Dg values, calculated as above,
is available; this data indicates that Df) is correlated to the
sediment accumulation velocity w. i.e..
= l.S.7ir
(Eq. 4)
so that ihis eqttation may be used in situations where D/j
cannot be calculated from data.
Nonlocal Models. Biological organisms commonly displace
sediment in ways that can hardly be thought of as diffusive,
i.e.. stnall scale and/or random. In such cases another type of
Figure M20 Chtinge oi excess ''"'Pb with depth in marine seditnents
from tlu' C\in.idi.in continenlai shelf. When plotted in a semi-
logarithmit manner js in ihis figure, the data closely follow a str<5ight
line,
that is, equation
L3).
with the stated D,, value.
model must be employed, wherein the biological mixing acts
like related sources and sinks of material (see Boudreau. 1997.
for mathematical details). This allows sedimentary compo-
nents to be transported large distances through the alimentary
canal of large infauna and by directed (non-random) motions.
For example, with these models it is possible to describe
situations dominated by head-down deposit feeder, a.k.a.
conveyor-belt feeders, where the organisms ingest sediment at
depth and defecate it at the sediment water interface: the
resulting feeding voids are ftlled-in by eotnpaction. A sitnilar
model can account for effects of burrowing fiddler crabs in
marshes, and a nonlocal model can explain persistent subsur-
face maxitna in radioactive isotopes in some seditnents as a
result of subsurface defecation.
Nonlocal tnodels differ fundamentally from the other two
models in how they operate on abrupt tetnporal change in the
cotnposition or structure (e.g.. latninations) of seditnents. In
both the well-tnixed and diffusion models, such discontinuities
are very rapidly lost, which begs the question how any such
changes can be preserved in biotic sedimenls. Nonlocal
models, in contract, are not as effective in dissipating these
structures, and their chances of survival are much higher. As
the geologic record abounds with many such changes, this
faet argues that real mixing is probably nonlocal in character.
In addition, nonlocal tnixing can actually act to .segregate
sediment components.
Forward and Backward Models. The two groups do not
constitute alternatives to the above models, but reclassifies
them in terms of what the model produces. In forward models,
we solve the equations given above for C. and to do so we
must input values of the parameters, that is. w. Dg. etc.. and
the tlux at the interface, F{t), This type of model describes the
sensitivity studies cotnmon in the geochetnical literature. With
inverse models, we start with observed values of C. and then
run the equations "backwards" to solve for either the value of
a chosen parameter or the input function. Least-squares
regression analysis of isotope data to obtain values of
H'
or
Du in seditnents is an example of a simple inverse model as in
Figure M20. Other more impressive applications have ineluded
objective proof of gasoline-bttsed lead reduction in sediments
and the need for the use of nonlocal models (see the examples
reviewed in Boudreau. 2000).
Bernard P. Boudreau
Bibliography
BerntT, R.A . 1980. Early
Diagenesis:
ATheoretieat Approach, Princeton
University Press.
Boudreau. B.P.. 1997. Diagenetic Models und their Implementation,
Springer-Verlag.
Boudreau. B.P.. 2()()(). The mathematics of early diagenesis: Irotn
worms to waves. Reviews in
Geophysics,
38: 389 416.
MUDROCKS
Introduction
Mudrocks are one of the several names for fine-grained,
argillaceous sedimentary roeks. also broadly called shales and
mudstones, all of which consist mostly of terrigenous clay and
452
MUDROCKS
(A)
CLAY
CARBONATE
ROCKS
SANDSTONE
SILT3T0NE
Figure M21 (A) Terminology for muddy unconsolidated mixtures
{Shepartl,
1954, figure 71; and (B) Ihp three most common lithologic
transitions of mudrocks.
sandstones and it truly cotnprehensive and integrated view of
earth history can be obtained.
Practical uses abound as well. Mudrocks serve as the source
of nearly all oil and gas, have many uses ranging from barriers
to lluid flow (eonfinement of loxic waste, caps for aquifers, and
seals for petroleum Iraps) to bricks, china, tooth paste,
cosmetics, catalysts to decoloring agents and hundreds more.
It is fair to say that without tnud and mudrocks, modern life as
we know it could not exist all this having started from sun
dried bricks some 6.000 years ago. And. mudroeks, because of
their propensity to compact, expand, or fail, also detnand
special attention in modern construction.
The literature of mudroeks is vast and widely scattered,
because so many researchers have difTerenl interests in them -
basin analysis (petroleum and stratiform ore deposits), biology
and paleontology, construction, their role in environmental
protection, organie, and inorganic geochetnistry, industrial
minerals and products, sedimentology, and many more. Four
recent texts include those by Chamley (1989). Weaver (1989).
O"Brian and Slatt (1990) and the two volumes of Schieber ctai.
(1998).
which contain 30 research and review articles and many
references. Much of the earlier literature is foutid in Potter
ctai. (1980).
silt along with sotne carbonate and organic matter. Mudrocks
occur either as thin beds or in thick sections, but some
siltstones. sandstone, or carbonate beds are nearly always
present and rare is the section of mudroeks that does not pass
laterally into interbedded carbonates or sandstones somevv-here
along its limits (Figure M21). Mudrocks thus may form but a
stnall part of a deposit or almost all of it with names sueh as
London Clay Formation, Rhinestreet Shale. Quamby Mud-
stone Group. Cambridge Argillite. ete.
Mudrocks deserve attention for many reasons. They are the
most abundant of all the sedimentary rocks, they occur
throughout the geologic colutnn in all the major sedimentary
environtnents. and are less studied and understood than either
carbonates or sandstones. Thus there is much to learn about
and from them. Unlike carbonates and sandstones, whose
modern deposits have recei\ed much attention and play
AW
itnportant role in interpreting ancient analogs, much more of
our understanding of mudrocks is based on studies of the
ancient.
From a scientific viewpoint, understanding the origin of
mudrocks has the potential to tell us much about ancient
environments of deposition and how they may have changed
throughout geologic time. Thought of in this manner,
mudrocks. and satidstones nicely cotnpletnetit one another in
the study of earth history—sandstones provide information
about source rocks atid their plate tectonic setting and the
paleo-hydraulics of sand transport into a basin, whereas
mudrocks are keys to the paleo oxygen levels of the basin, and
through their tnicrofauna and body fossils, insights to its
paleo-circulation system, environment of deposition and
possible paleoclimates. Exeeptionally. they are also rich
La^cr.stdttcit for fossils such as the famous Cambrian Burgess
Shale Formation of the Canadian Rockies, because in a
muddy basins fossils may be quiekly buried by a slurry, may
die and fall into loxic bottotn waters free of scavangers and.
because the low pcrtneability of tnud tninimi/es dissolution by
migrating aeid solutions, escape much diagenesis. Add to this
information frotn the study of interbedded carbonates and
Texture and nomenclature
By convention, the ftne-grained sediments inelude both silts
and clays and it is. in fact, their relative proportions that are
used to classify them texturally (Table M3).There are several
competing systems of nomenclature for mudrocks. the board
general genetic term for fine-grained argilleceous rocks, but all
depend on the relative proportions of clay and silt. Once this is
understood, it becomes easy to use the different systetns. When
the clay fraction (less than 4 microns) forms more than
67 percent of the rock, the terms claystone or ctayshale are
used: when the clay fraction tbrms .'^.3 66 pereent. the tertns
miulstone or ntudshalc are used; but when silt (>4 mierons)
forms 67 or tnore percent, the term sittstoiw is used. This
important "purity" distinction is easily carried out in the field
as a first approximation by putting a small pieee between the
teeth and determining its grittiness. Fine carbonate is also
comtnon in tnany mudrocks associated with limestones or
dolomites and locally mudrocks tnay also be silieeous or
phosphatic. Color is exceptionally variable ranging from black
to brownish black through many shades of gray (the tnost
comtnon) to red and purple and even whitish gray and thus far
more variable than other seditnentary rocks. And, unlike
sandstones and earbonates, the colors of mudrocks have tnuch
genetic significance.
Processes
Sources of clay and silt
The immediate source of the clay and silt fractions of
mudrocks are soils and preexisting tnudrocks. Weathering in
soils is favored by high rainfall and good drainage, which
together remove cations frotn complex silicates formed at high
tetnperatures to tnake progressively sitnpler clay mineral
structures stable at the earth's surfaee. Where weathering
is incomplete because of either insufficient rainfall or high
steep slopes, elay and silt compositions closely reflect
MUDROCKS
Table
M:1
Dest rihins; mtiflriK ks' after Stow
(1981:
Biisic Terms
Unlit hij ted
Silt
Mud
Clay
Texture
Silly
Sandy >10"/.
Pebbly
Lithified'lion -ji.'isite
Siiistonc
Miidstotie
Claystone
Lithified/fissile
Sillshalc
Mudshiile
Clav-sliale
Composition
Calcareous, siliceous, carbonaceous (all •>!()"'..), pyritiferous (I to 57")
atid lerritiiitiotis. niicaccoiis. and fossilifcroiis
Stratijieation (bedding, if >
1
cm & iamination. if less}
May be parallel, wavy (rippled), ctoss-laininated. graded, continiious-disconiiiuiou.s. bioturbated. or defortned
(iisi;
appropriate cotiibinatioti.s)
Metamorphie
eijuiv(denis
Argillite
Slale
Peltte
Sliglitly
Mciamorpfiosed/fissite
liither ;iri;ilite or slale
Proptirtions &
grainsize
>\ silt-sized''
silt & clay mixture
>^ clay-sized''
silt and
c-lay
mixture
'' More ttian
511"
<i
ti;rn>iL'iiou> ul
i^liiLh
mint' \\r.m >H'.. must be smaller than (i.' micron^: '' 4 63fm!. ' - 4nni
source rocks, but where rainfall is abundant, soil drainage
good, and slopes low. the resulting fine debris mostly eonsists
of kaolinite and quartz (silica, iiluminuni, and iron oxides) and
only poorly reflects source rock composition. Thus, rcmlenci'
titne in the source area is all important for mudrock
composition and is ultimately linked to tectonic stability.
The volume of fines depends on relief and rainfall u
eotnbination of high relief and rainfall (monsoons against a
high Htmalayas) produces large volutnes of debris whereas low
rainfall and low relief (Sahara craton). high rainfall and low
relief (Aniai^on craton) or wide carbonate landscapes produce
but little.
But source rocks also need sotne consideration too when
fines are considered. Advanced weathering of granite and
gneiss yields a kaolinitc-ricli clay fraction plus coarse and fine
quartz silt, whereas volcanic rocks yield lines rich in unstable
glass,
tl'ldspar. and amphiboles that transform rapidly into
/eolttes and stnectites. Low grade tnetamorphic rocks, on the
other hand, especially the greenschist facies, tend to yield
micas,
illites, chlorites. and silt and fine sand-sized quartz and
feldspar. It also appears that most fines travel in jumps as
streams rework their alluvium; that is. most of their time is
spent in storage on the lloodpKiin. This implies that headwater
mineralogy may alter down stream, especially on wartn. wet
stable cratons. Another major source of the silt fraction of
mudrocks is wind erosion of deserts (Sahara, central Asia, and
north China) to deposit loess, whtch tnay have been far tnore
important in pre-Devonian times before extensive plant cover.
And finally, some clay minerals in mudrocks are authigenic
such as the illiles and chlorites formed in deep burial whereas
palygorskite (saline lakes and seits) and glauconite (shallow
shcKes) itre fortned on sea bottoms. In addition, some quartz
silt may be born in place at tempertitures well below 90 C
trom the diageneic alteration of opaline quart/ of biogenic
origin. Although deciphering the proportions of detrital to
authigenic components ofa tnudrock can at times be difficult.
it seems safe to say that most of the clay and silt fraetion of
mudrocks is detrital and that mitch cotnes frotn the erosion of
preexisting tnudrocks and siltstones. Ultitnately, however,
most is derived IVom the weathering of granites, gneisses,
volcanics. and low-rank metamorphic rocks.
Tratisport
Differential transport is responsible for the segregation of the
large volumes of clay and silt that form the world's mudrticks
and, on a tnuch stnaller scale, it is differential transport that
also segregates by particle size and shape the clay minerals of
one clay bed from another—an all important distinction in
the tnining and utilization of clays for their many diverse
industt"iul uses.
In ail of these, virtually all the clay fraction is transported in
suspension by either wind or water as is tnuch of the line silt
fraction, whereas coarser silt and sand, are transported
intertnittently by saltation. Fifieen microns seems to be an
important size for the hydraulie behavior of silt experimen-
tally it has been found that below 15 microns most silt is
deposited as planar lamination whereas above 15 mierons
small scale bedforms develop.
In both wind and water it is the upward cotnponent of
turbulence, r,, that prevents a suspended particle tVotn settling
to the bottom. The tnagnitude of this component is roughly
proportional to the average veloeity of the flow so the stronger
the flow, the greater i,, and the larger the size of the particle
that can be suspended. In detail, the criteria for suspension is
V,,
> vt', where n', is settling or fall velocity which is given to a
good approximation by Stokes Law for stnall particles as
(Eq. I)
where d is particle diatneter. A/J is the density difference
between particle and fluid, fi is dynamic viscosity, and t,' is
gravity.
454 MUDROCKS
As particle size decreases, so too does the downward driving
force acting on it. Consequently, particles a few microns in
diameter have exceedingly small fall velocities, remain
suspended by the weak upward components of turbulence in
gentle currents for weeks and are thus carried great distances.
It is this small fall velocity of elay and tine silt that is
responsible for the great lateral continuity of many mudroeks,
be they thick or thin. Stokes law is also the key control on the
concentration of organic matter that occurs commonly in
mudrocks. bottom oxygen levels permitting. The low density
of most organic matter permits it to preferentially settle to the
bottom in quiet water along with the clay fraction to produce
organic rich shales—the source beds for petroleum. Becatise
clays and organic material settle together in quiet water, the
mudrocks of a topographic low those deposited in a
protected paleo sinkhole, in a gentle sag on the seafloor. or
abandoned channel or eanyon are nearly always thieker,
darker colored and more clay- and organic rich ihan lateral
equivalents outside the low.
Several observations are needed about Stokes law. First, it
only applies to single spherieal partieles smaller than about
180 micrometers settling in quiet, isothermal, nonturbulent
water hardly the conditions of clay flakes in nearly always
turbulent natural flows. Thus predicted fall velocities are
conservative ones. Still another complicating factor is that
many, perhaps most, clays, and fine silts are transported not as
single particles, but as loose, large, aggregates called floccules,
Jiocs for short, which settle differently than single partieles.
Floccules are composed of networks of clay flakes, silt, fine
skeletal debris and organic material all held together by
organic slimes or. for the clay minerals, by ionic bonding.
Floccules occur in streams, but are most eommon where (resh
water mixes with saline. Saline water favors (locculation,
because its positively charged sodium and negatively charged
chlorine ions promote the attraction of elay partieles. The large
planar face surfaces of clay particles are negatively eharged
and thus brought together by sodium ions. Edge to edge
bonding also occurs but. because the edges have both positive
and negative charges, the process differs. Depending on the
charge, there will be either an edge to face or face to face
contact. In low pH environments, say the acid waters of a
swamp, hydrogen ions are bonding agents for clay flakes.
Finally and exceptionally, in .semiarid climates expandable
clays can aggregate to form sand-sized partieles (peds) whieh
locally form clay dunes and a few, rare alluvial sands.
Biological aggregation also oecurs in seas and lakes, is far
more important, and is called peUelizalion the product of
infauna feeding on delicious organic rich mud.
In sum. while both weak and strong currents can transport
clay in suspension long distances, weak ones are necessary for
it to Iinally settle to the bottom flocculated or not. This means
quiet water, either shallow or deep is needed for deposition—a
small lagoon fed by a muddy river, a temporary lake on a flood
plain during spring Hoods, a deep lake in a rift, a deep shelf
whose bottom is rarely stirred by storm waves, or an oceanic
trench only reached by distal, slow moving turbidity eurrents
that deposit thin graded muddy laminations.
It appears that much, perhaps most, transport of clay and
silt is episodic. On land this equates to floods- maximum sheet
wash erosion of
slopes,
gullies, and stream banks provides high
suspended eoneentrations and volumes in streams while in
lakes and seas, episodic storms and turbidity currents play a
comparable role.
Sedimentary structures
Mudrocks and sandstones share the same types of sedimentary
structures—hydraulic, chemical, and biologic. These differ
only in abundanee and scale, all the hydraulic ones in
mudrocks being much smaller, because of weaker currents.
Thus mudrocks mostly have lamination rather than beds
{By convention all stratiflcation thinner than lO.Omm is called
laminalion and that thieker, heddjiig).
Lamination in mudrocks is mostly seen by variations in
proportions of silt and clay, bul is also deiined by thin
carbonates (pelagic rain or algal mats'), b> minor textural
variations within muddy layers and exceptionally by variations
of organic matter (kerogcn). Lamination takes many different
forms from rare, silty planar streaks only a few grains thick in
almost completely muddy sections to almost all siltstone,
where each siltstone is .separated by a thin elay parting with
all possibilities in between. Individual laminations may be
continuous or discontinuous, wavy, or planar, massive, or
graded, inclined (cross laminated) or convoluted, random, or
rhythmic. Rhythmic lamination is typically defined by altera-
tions of clay and silt and less commonly by carbonate. The
presence or absenee of lamination in mudrocks is used to
divide them into two broad classes—those that have lamina-
tion (classically called shale.s) and those that lack lamination
and are massive (often called mutlstoncs). In several black
shale basins individual laminations as well as packets of them
have been convincingly correlated more than 100 km after
cores were X-radiographed, thus demonstrating a virtual total
laek of bottom currents.
Several dilTerent processes are responsible for absence of
lamination in mudrocks. The principal control is the amount
ol"
bottom oxygen that inhibits benthonic organisms from
scavenging on or in muddy bottoms and thus destroying
primary stratification. In ancient basins paleo oxygen levels arc
approximated by noting the degree of bioturbation and how it
varies with the color, pyrite, and organic content of the
mudroek. Lighter mudrock colors, less pyrite and less organic
material typically are found as bioturbation increases and
lamination decreases. A few massive mudrocks also are the
product of muddy slurries that flowed along the bottom, and
exeeptionally, some are formed by deposition from denser than
normal suspensions.
The term hlocky is used to describe the weathering of a
massive, poorly, or non-laminated mudrock. Two other terms
are often used in the description of mudrocks. fissility and
purling. Fissility refers to the fanlike separation of one
lamination from another during weathering. Thus, while
fissility is not a primary sedimentary strueture it reflects one.
For example, well, and uniformly laminated mudroeks may
become paper like (Blaitcrnion) upon weathering as their
individual laminations separate one another by swelling oi
expandable clays. An argillitc. on the other hand, is usually so
well indurated that its laminations do no separate one from the
other upon weathering. The term parting is used for a single
lamination separating more massive lithologies above and
below. X-radiography is needed to see the full details of the
physical, chemical, and biological structures of many muds,
although careful observation of weathered outcrops captures
most details.
Present on the bottoms of some thin silty laminations are
small flutes and load casts plus borings, traces, and impres-
sions made by botlom dwellers. The orientation of flutes and