Middleton G.V. (Ed.) Encyclopedia of Sediments and Sedimentary Rocks
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MATURATION, ORGANIC
425
Wetting inereases bulk weight
and
hence shear stress, while
changing effective stress
as the
groundwater table rises above
the level
of
the potential slip plane. Such concomitant changes
in stresses
may
occur within
a few
hours, explaining why many
slope failures occur during major storms (Selby, 1993),
On
a
longer timescale, slope undercutting
by
rivers
and
glaciers causes slope steepening
and
hence shear stress
increase. This induces slow creep movements within
a
slope,
causing
a
progressive reduction
in
both cohesion
and
frictional
resistance over time
and
hence long-term material fatigue.
Seismic shock often provides
the
final trigger
for
many
catastrophic slope failures
in a
wide range
of
materials,
ranging from hard, jointed rock
to
unconsolidated clays.
Types
of
mass movement
A common failure type
is the
translational slide, involving
a
slab-shaped detachment along
a
well-defined rupture surface.
In rock slopes, rock slides typically follow discontinuities such
as bedding, Joints, foliation
and
faults. Sliding
is
feasible only
where
the dip of the
discontinuity
is
less than
the
slope angle,
allowing
the
critical plane
to
crop
out on a
slope (Hoek
and
Bray, 1981), Large rock slides (20 Mm^
to
200 Mm')
are
often
seismically triggered
and
produce rock avalanches, since total
disintegration
of the
mass allows
Oow to
occur, Entrainment
of saturated valley-floor material provides
a
liquefied base
for
the
avalanche, allowing velocities
of
10
m
per
second
to
30 m
per
second. Exceptionally large liquefied failures, termed
lahars, result from collapse
of
stratovolcanoes (Vallance
and
Scott, 1997),
The
largest lahar deposits approach lO'm^, cover
tens
of km , and
rank
as the
largest landslides
on
Earth, High
mobility
of
lahars
is
caused
by
saturation
by
melting
of
ice
and
snow, drainage
of
volcanic crater lakes,
and
entrainment
of
saturated alluvial material.
On soil
and
debris covered slopes, translational debrisslides
occur along
the
soil-bedrock interface since soil shear strength
is lower than that
of
rock. Debris slides
are
typically
lO^m to
10^
m^
and are
triggered mainly
by
rainstorms. Scour
of
saturated material along
a
stream channel transforms many
debris slides into debris flows (Iverson
et al.,
1997), Debris
flows occur
in
every climatic zone, attaining velocities
of
5
m
per second
to 10m per
second
and
typical volumes
of
10"^m''
to
10^
ml
On slopes
cut
into thick clay
or
shale, deep-seated slumps
are common, involving
a
roughly cylindrical-arc slip surface
oblique
to the
stratification witii backward rotation
of the
landslide mass. Slumping creates
a
steep, debuttressed head-
scarp prone
to
further slumping,
a
process termed retro-
gression. Prolonged retrogression
may
produce
an
elongated
earth flow,
up to
several kilometers
in
length, moving
at
1
ma~' to
5ma~', with
a
total volume exceeding
lO^m , In
low density saturated silts
and
clays, liquefaction
may
occur
at
depth
as the
deposit collapses under shear loads, Frictional
strength declines notably
as
pore pressure rises, allowing
rapid acceleration
as a
quick clay flow-slide (Turner
and
Schuster, 1996), Flow-slide retrogression
may
occur
at
meters
per minute, with flow velocity attaining
1
m per
second
to
5
m per
second.
In permafrost areas, seasonal thaw
of
ice-rich soil produces
a 0,5 m
to LOm
thick water-saturated active layer above
permafrost material, Solifluction involves movement
of the
active layer
at
10~^ma~'
to
10~'
ma~' on
slopes
as low as 2°,
Movement
on
slopes well below
the
theoretical lowest angle
predicted
by
Mohr-Coulomb theory occurs
by a
combination
of frost
creep
and
gelifluction (Harris, 1987), Frost creep
is
soil
movement
by
alternating cycles
of
frost heave
and
thaw-
settlement
in
soils subjected
to ice
segregation, GeliHuction
is a
type
of
soil flow, occurring during seasonal thaw,
in
which
high water pressures develop during melting
of low
density,
ice-rich soil. Overburden pressure
is
largely borne
by the
pore-
water, allowing
low
effective stress
to
develop,
A
transient
phase
of
negligible frictional strength occurs, allowing soil
to
move
on
slopes substantially flatter than possible
in
saturated,
consolidated soils.
On
steeper slopes
in
permafrost areas, skin
flows involve rapid downslope detachment
of the
entire active
layer. Retrogressive thaw slumps
are
triggered
by
river under-
cutting
of
slopes
and
exposure
of
ice-rich sediments. Thaw
slumps
are the
largest slope failures found
in
permafrost areas,
Michael
J,
Bovis
Bibliography
Craig,
R,F,, 1997,
SoilMechanies.
6th edn,,
London: Spon Press,
Harris,
C, 1987,
Mechanisms
of
mass movement
in
periglacial areas,
tn Anderson,
M,G,, and
Richards (eds,). Slope Stability.
J,
Wiley
& Sons,
pp,
531-559,
Hoek,
E,, and
Bray,
J,W,, 1981,
Roek Slope Engineering.
3rd edn,,
London: Institution
of
Mining
and
Metallurgy,
tverson,
R,M,,
Reid,
M,E,, and
LaHusen,
R,G,, 1997,
Debris flow
mobilization from landslides. Annual Review of Earth and Planetary
Sciences,
25:
85-138,
Kenney,
T,C,, 1984,
Properties
and
behaviours
of
soils relevant
to
slope stability,
tn
Brunsden,
D,, and
Prior,
D,B,
(eds,). Slope
Instability. John Witey
&
Sons,
pp,
27-66,
Selby,
M,J,, 1993,
Hillslope Materials
and
Proeesses.
2nd edn,
Oxford
University Press,
Sidte,
R,C,,
Pearee,
A,J,, and
O'Loughhn,
C,L,,
1985, Hillslope Stabi-
lity and Land
Use.
American Geophysical Union,
Turner, A,K,,
and
Schuster,
R,L,
(eds,), 1996, Landslides:Investigation
and Mitigation. Washington,
DC:
National Academy Press,
pp,
585-606,
Vallance,
J,W,, and
Scott,
K,M,, 1997, The
Osceola Mudflow from
Mount Rainier; sedimentology
and
hazard implications
of a
huge
clay-rich debris flow, Geologieal Soeiety
of
Ameriea Bulletin,
109:
143-163,
Cross-references
Atterberg Limits
Avalanche
and
Rock Fall
Bentonites
and
Tonsteins
Debris Flow
Earth Flows
Grain Flow
Gravity-Driven Mass Flows
Liquefaction
and
Fluidization
Mudrocks
Slide
and
Stump Structures
Slurry
MATURATION, ORGANIC
Organic maturation refers
to the
progressive
and
mainly
irreversible transformation
of
organic matter (OM)
in
response
initially
to
biological
and
later
to
thermal energy. Organic
426
MATURATION, ORGANIC
maturation is also referred to as organic diagenesis, organic
metamorphism, and as coalification when referring to coal.
Organic matter is a minor component in most sedimentary
rocks,
however its importance by far outweighs its abundance.
Organic matter comprises the fossil fuel resources petroleum
and coal. The products of OM diagenesis govern many mineral
reactions and are important in the transport and deposition of
many ore minerals. Studies of organic maturation have mainly
been pursued form two separate avenues: organic geochemists
have investigated the chemical reactions and products of
organic maturation particularly from the perspective of
petroleum whereas coal petrologists have studied organic
maturation of coal mainly microscopically.
Introduction
Diagenesis of OM has been considered in terms of three or
four stages by various authors (e,g,, Tissot etal., 1974; Bustin,
1989;
Diessel, 1992), Eogenesis has been used in reference to
biological, physical and chemical changes in OM that occur at
low temperatures where reactions are mainly biologically
mediated, Diagenesis, at higher temperatures normally asso-
ciated with greater depths of burial, is successively referred to
as catagenesis and metagenesis. Diagenetic reactions of OM
due to contact with free oxygen or meteoric water has been
refereed to as telogenesis. The bouridary between different
diagenetic stages is gradational as a result of the heterogeneity
of the OM and biological processes. The different coal ranks
from peat to lignite, subbituminous, high-, medium-, and low-
volatile bituminous coal, semi-anthracite and anthracite, are
names sequentially assigned to successively more carbon-rich
and volatile-poor coals that form in response to progressive
diagenesis.
Substantial research exists on the kinetics of organic
diagenesis because of the importance of predicting the timing
and volume of hydrocarbon generation from the diagenesis of
kerogens. As well, sophisticated numerical models have been
proposed that attempt to predict organic maturation based on
the thermal history of the strata and the activation energy of
various kerogen types. Because many of the products of
organic diagenesis are metastable and rocks are open to
migration of products (hydrocarbons), thermodynamic equili-
brium is rarely ever established,
Eogenesis
Eogenesis refers to low temperature diagenesis of OM, where
reactions are at least in part biochemical resulting from
metabolic processes of organisms, and the effects of tempera-
ture and pressure are subordinate (Figure M7), Due to the
diversity of mierobial processes and organic substrates,
reactions during eogenesis are complicated and poorly under-
stood. Different depositional settings give rise to different
diagenetic processes and sources of OM, which in turn
influence the type and rate of degradation of the OM mainly
by bacteria, actinomyces and fungi. For OM to be preserved at
all to further undergo diagenesis requires rather specific
depositional conditions. It is clear that in subaerial environ-
ments chemical oxidation and aerobic mierobial decomposi-
tion invariable leads to complete decomposition of the OM,
Anaerobic mierobial processes are generally slower and hence
eogenetie reactions are commonly retarded sueh that there is
greater potential for a portion of the OM to be preserved. For
Biogenic
poiymers
Monomers
03
m
z
CO2,
H2O,
CH4
NO3,
NH4+
SO4-,
S-,
HS-
Ro-0,5
Geochemicai fossiis
free iHCs and reiated
compounds
en
CO
z
LU
Low to
med,
MWHCs
High
MW
HCs
" Ro-2,0
eking
-^J
Methane, iight HCs condensates
Methane (dry gas)
Oii
Gas&
METAGENESIS
Graphite, iarge
condensed aromatics
Figure M7 Schematic diagenetic pathway of organic matter in
sediments. Modified after Barnes et al. (1990),
OM to survive eogenesis and to move on to catagenesis, it
must be rapidly removed from oxic conditions. Proteins,
carbohydrates (^sugars, cellulose, and chitin), lipids (fats,
waxes, and steryl esters) and lignins are the main biopolymers
contributed by organisms to the sediment, Biopolymers are
variously depolymerized during eogenesis forming what is
colloquially referred to geopolymers, which are essentially
fulvic and humic acids, humins and kerogens (Bustin, 1999),
Geopolymers are detined on their solubility, their molecular
weight and, because they form by random recombination,
their structures, Geopolymers vary depending on available
monomers and diagenetic conditions. With increasing cross-
linkage and loss of acidic functional groups (carboxyl and
phenolic hydroxyl groups), the humic and fulvic acids lose
their base solubility and form an insoluble humin/kerogen
residue. The term humin is used for base insoluble fraction of
the OM in soils, whereas kerogen refers to the organic fraction
in rocks, which is insoluble in organic solvents. The details of
diagenetic reactions leading to kerogen formation remain
poorly understood. Some kerogen forms through loss of
functional groups and cross-linkage formation early in
diagenesis, but others later during catagenesis and still other
kerogen appears to arise directly from monomers without an
intervening humic stage particularly under anoxic conditions.
The lipid-rich components are particularly resistant to
biochemical processes and unless degradation is particularly
intense, these components survive eogenesis. An important
product of eogenesis is the formation of mainly methane (up to
99 percent), carbon dioxide (0 percent to 8 percent) and minor
MATURATION, ORGANIC
427
heavier gases by fermentation reactions. This biogenic gas may
source significant economic accumulations both in the coal
(coalbed methane) and conventional reservoirs. Carbon
dioxide generated together with organic degradation products
influences the Eh, pH, and ionic composition of pore waters
and thus plays a role in diagenesis of minerals.
The end of eogenesis generally occurs when the microscopic
reflectance of coal maceral vitrinite is about 0.5 percent, which
also corresponds to the boundary between subbituminous
coal and high-volatile bituminous coal. At this stage
base-soluble compounds are usually negligible and thermal
energy is required for further diagenesis.
Catagenesis
Diagenesis beyond the eogenetic stage is in response to thermal
exposure (temperature and time), generally accompanying
burial in a sedimentary basin or high heat flow adjacent
intrusive or extrusive bodies (Figure M7). The rate of organic
diagenesis is exponential with temperature and linear with
time.
Confining pressure has little effect during organic
diagenesis apart from compaction during the peat and lignite
stages and some experiments suggest high confining pressure
retard catagenetic and metagenetic reactions. Catagenetic
reactions take place at temperatures of around 40°C to
150°C and moderate pressures (30-150 MPa), often producing
significant amounts of petroleum, methane and carbon
dioxide. The amount and type of liquid hydrocarbons
generated varies with the type of" kerogens present. Generally
most coals are considered to have little liquid hydrocarbon
potential whereas algal-rich, fine-grained rocks are excellent
liquid hydrocarbon source rocks. During catagenesis coal rank
progressively increases from subbituminous to anthracite and
kerogen increases in carbon content, decreases in volatiles,
aliphatic chains are cleaved and the size of aromatic clusters
grows through condensation and development of ordering.
Coincident with the generation of gas, the surface area of the
kerogen increases such that significant quantities of the
generated gas are adsorbed in the matrix (up to about 30 mV
tonne for coal) giving rise to coalbed methane. Coalbed
methane (and gas shale), is an important energy resource,
however, it is also a potential explosive hazard during coal
mining. The exact level of catagenesis for liquid hydrocarbon
generation (oil birth line) varies with the chemical composition
of the kerogen. The catagenetic level at which liquid
hydrocarbons begin to crack is referred to as the oil death
line.
The zone of catagenesis between the birth and death
"lines"
is referred to as the oil window or oil kitchen
(Figure M7).
Organic matter undergoes noticeable transformations that
are observable microscopically even at the beginning of the
eogenetic stage. The fraction of the OM derived from higher
plants, mainly cellulose and lignin, give rise to the micro-
scopically defined maceral huminite (at low maturation levels)
and later vitrinite. The reflectivity of huminite/vitrinite
progressively increases, as observed microscopically in re-
flected light through diagenesis. This degree of reflectance can
be measured with a specially designed microscope and is
referred to as vitrinite reflectance. Vitrinite reflectance is the
most widely used method for determining the degree of
diagenesis of coals and OM in rocks. During catagenesis
vitrinite reflectance progressively increases from about 0.5
percent to 2.0 percent (Figure M7). Liptinite components
change little until the oil birth line which corresponds to a
vitrinite reflectance of about 0.5 percent and then it become
progressively brighter and merge in brightness with vitrinite at
a reflectance of about 1.4 percent (the death line of oil). The
fraction of the OM rich in lipids is microscopically darker
than vitrinite and is referred to as liptinite. The liptinite
macerals change little with diagenesis until the onset of the oil
window at which time there is a "jump" in diagenesis as HCs
are generated (Figure M7). That microscopically defined
fraction of the OM that is semi charred or charred at the site
of deposition (charcoal) is referred to respectively as semi-
fusinite and inertinite. These materials being variously pre-
coalified undergo little change during diagenesis.
Metagenesis
Metagenesis refers to the level of diagenesis of the OM at
which crystalline ordering of the OM begins (Figure M7).
Aromatic clusters increase in size and C-C bonds are broken
generating additional methane. Any remnant aliphatic mole-
cules and any previously formed heavier hydrocarbons are
cracked to methane. The remaining kerogen becomes further
enriched in carbon with loss of volatiles. Microscopically the
reflectance of vitrinite progressively increases from about
2.0 percent to 10 percent or more and the liptinite component
and lower reflecting inertinite components become visually
indistinct from vitrinite. Early studies argued that graphite is
the end point of coal diagenesis however field and experimental
evidence now suggest that the activation energy for formation
of graphite is so high that it is unlikely that it will form in
nature by thermal exposure alone (Figure M7).
Telogenesis
Organic matter diagenesis is progressive and for the most part
irreversible. Organic matter at all stages of diagenesis is
reduced and subsurface environments are typically anaerobic.
Where, however, OM is uplifted to the surface or comes in
contact with oxygenated meteoric water, it is subjected to
oxidation. Oxidative processes are collectively refereed to as
teleogenesis and result in marked changes in the OM. During
oxidation nonaromatic groups are selectively removed and
acidic functional groups such as carboxyl, carbonyl and
phenolic hydroxyl are formed leading to production of humic
acids.
The oxidative diagenetic processes have a deleterious
effect on coal quality: coking coals loose their ability to form
coke,
thermal coals decline in calorific value and coal particles
loose their hydrophobicity and thus are difficult to separate
from mineral matter during coal processing (Copard et al.,
2002;
Teichmiiller, 1982). If oxidation is intense it may be
microscopically evident by distinctive darkish halo along
boundaries of organic particles. With pooled hydrocarbons
two distinct telogenesis processes occur: (1) water washing
which results in selective removal of lighter hydrocarbons by
groundwater; and (2) bacterial degradation which can lead to
loss of normal alkanes and isoprenoids leaving a heavier oil
residue enriched in cycloalkanes and aromatics. These two
processes are considered responsible for formation of the
heavy oil deposits.
428 MATURATION, ORGANIC
METAGENESIS CATAGENESIS
EOGENESIS
0)
Anthracite
emi-
thra
-< 1-
3
•
^
Medium
volatile
bituminous
A |B| C
Higfi vol. bituminous^
JSub-bitum.
Lignite Peat
10 20 30
40
50 60 70
kJ/kg
36,000 36,000
29,000 23,000
17,000
2.5 5.5 10
25
%H
35
% H2O
dry gas
wet
gas
oil
and gas
early gas
1.51.35
1.2 1.0 0.8 0.6 0.6 0.4
0.3 0.15
no fluorescence
-500 Sporinite
"600 fluorescence
•700 (rnax.)
no fluorescence
black dark dark-brown
brown-black
4.0 3.8 3.7 3.5
red-
orange- yellow-
brown brown orange
1.0
black black very dark
grey brown dark brown
brown to very pale brown pale yellow
1.5
£.3
o-O
485°C 47O''C
450°C
435°C
425°C
based on:
(C25-C33)]
kerogen type I and II
DIAGENETIC LEVEL
Coal rank
Volatile matter
Coal (daf)
Calorific value
Moisture (ash free)
Hydrogen (daf)
HC generation
Vitrinite reflectance
(Rpmax %)
1 Sporinite
1
Fluorescence
Spectral-
quotient
Spore colour
and
TAI
Conodont colour
and alteration
index(CAI)
Atomic
ratio
Pyrolysis (S2)
^ Carbon
-1.4 Preference
Q Index (CPI)
I
Figure M8 Correlation of major organic maturation indices. The correlation shown is based on virtinite reflectance. Modified after
Bustin et al. (1990).
Quantifying organic diagenesis
Both the kerogens and bitumen fraction of the OM have been
used to quantify diagenesis by chemical, physical and
microscopic methods (Bustin etal., 1990; Tissot and Welte,
1978).
In Figure M8, the correlation between the major
different diagenesis indices is provided. A few of the more
common techniques are summarized below. The reflectance of
vitrinite, and its precursor huminite is the most widely
accepted method of quantifying organic diagenesis and the
method to which all others are compared. Vitrinite reflectance
is measured microscopically using a highly polished specific
and a specialized microscope equipped with a photometer.
With increasing levels of diagenesis the reflectivity of vitrinite
increases in a predictable fashion. Characteristics of coal have
often been used as diagenetic indicators. The volatile matter,
moisture, and hydrogen content of coal decrease with
increasing diagenesis and the carbon content and calorific
value increase in a predictable manner (Figure M8). The
chemistry of kerogens has also been found to be useful
indicator. For example the hydrogen to carbon ratio and
pyrolysis yield of kerogens decrease with increasing diagenesis.
The carbon preference index (ratio of odd to even n-alkanes)
also progressively decreases with diagenesis from values
greater than 1 (due to initial organic selectively) to values
approaching one as diagenesis progresses.
The fluorescence of liptinite macerals has been utilized in
some studies. With increasing levels of diagenesis the intensity
of fluorescence decreases and the wavelength of maximum
fluorescence progressively increase. Such measurements re-
quire a microscope equipped with a scanning monochromatic
filter and sensitive photometer.
Kerogen coloration in transmitted light including pollen,
spores and eonodonts has been successively used to quantify
maturation. As with toast, the OM darkens with increases with
degree of thermal exposure and specialists attribute the colors
to various numerical schemes based on visual examination.
The color of kerogens has been referred to as the Thermal
Alteration Index, whereas the color of eonodonts is referred
to the Conodont Alteration Index (CAI) (Figure M8).
R. Marc Bustin and Raphael A.J. Wust
Bibliography
Barnes, M.A., Barnes, W.C, and Bustin, R.M., 1990. Chemistry and
diagenesis of organic matter in sediments and fossil fuels. In
Mcllreath, I.A., and Morrow, D.W. (eds.), Diagenesis. Geological
Association of Canada, pp. 189-204.
Bustin, R.M., 1989. Diagenesis of kerogen. In Hutcheon, I.E. (ed.),
BuriaiDiagenesis. GAC/MAC Short Course Handbook. Montreal,
Quebec: Mineralogical Association of Canada, pp. 1-38.
Bustin, R.M., 1999. Coal: origin and diagenesis. In Marshall, C.P., and
Fairbridge, R.W. (ed.). Encyclopedia of Geochemistry. Kluwer
Academic, pp. 90-92.
Bustin, R.M., Barnes, M.A., and Barnes, W.C, 1990. Determining
levels of organic diagenesis in sediments and fossil fuels. In
Mcllreath, I.A., and Morrow, D.W. (eds.), Diagenesis. Geological
Association of Canada, pp. 205-226.
Copard, Y., Disnar,
J.-R.,
and Becq-Giraudon, J.F., 2002. Erroneous
maturity assessment given by Tmax and HI Rock-Eval parameters
on highly mature weathered coals. Internalionai Journal of Coal
Geology, 49 (1): 57-65.
Diessel, C.F.K., 1992. Coal-bearing Deposilional Systems. Springer-
Verlag.
MATURITY: TEXTURAL AND COMPOSITIONAL
429
Teichmiiller, M., 1982. Origin of the petrographic constituents of coal.
In Stach, E. et at. (eds.), Textbook of Coal Petrology. Gebriider
Borntrager, pp. 219-294.
Tissot, B., Durand, B., Espitalie, J., and Combaz, A., 1974. Influence
of nature and diagenesis of organic matter in formation of
petroleum. American Association of Petroleum Geologists Bulletin,
58 (3): 499-506.
Tissot, B.P., and Welte, D.H., 1978. Petroleum Formation and Occur-
rence. A New Approach
to
Oil and
Gas
Exploration. Springer-Verlag.
Cross-references
Coal Balls
Diagenesis
Kerogen
MATURITY: TEXTURAL AND COMPOSITIONAL
Maturity refers to the degree to which clastic sediment has been
modified by physical and chemical processes at Earth's surface.
Textural maturity refers to the degree to which physical
characteristics of grains and populations of grains approach
the "ultimate end product" (Pettijohn, 1975, p.491). Com-
positional maturity (stability) refers to the degree to which
chemical characteristics approach the "ultimate end product,"
so that grains are more in equilibrium with Earth's surface
conditions.
Textural maturity
As rocks approach Earth's surface during removal of overlying
rocks,
individual fragments of rock are defined by diverse
physical, chemical and biologic weathering processes. Joints
and other structural discontinuities tend to define gravel clasts,
whereas preexisting grain size tends to define sand and silt
clasts.
Intensity of chemical weathering largely controls clay
mineralogy
{q.v.).
Chemical weathering causes the least stable
components in go into solution and/or to combine into more
stable minerals during weathering. The result of these varied
processes is the creation of sedimentary clasts of diverse grain
size,
shape and composition. As the clasts are eroded,
transported and deposited (possibly through several cycles),
they are modified by diverse physical and chemical processes.
Most of these processes produce increased textural maturity
through time.
Various measures of textural maturity have been developed.
Roundness is a measure of the degree to which a grain has
attained a continuously curving surface. Sphericity is a
measure of the degree to which a grain has attained a spherical
shape (equidimensional). These characteristics can be deter-
mined by painstaking measurements of curvature and dimen-
sions of individual grains, or more commonly, by visual
comparison with standard charts. A more sophisticated
method for analyzing large populations of grain shapes utilizes
Fourier analysis (e.g., Ehrlich and Weinberg, 1970).
Grain-size distributions reflect diverse processes at work
during erosion, transportation and deposition. Some deposi-
tional environments tend to concentrate certain grain sizes; in
the case of high-energy beaches, coarser clasts are left at river
mouths and finer clasts are transported offshore, resulting in a
concentration of fine to medium sand along beaches. Such an
environment produces a narrow range of grain sizes, and is,
therefore, texturally mature. In contrast, mass wasting and
flash floods produce texturally immature deposits.
Standard petrographic methods of studying sandstone have
shown a strong correlation between grain size and composition
(e.g., Boggs, 1968). As rock fragments break into their
constituent minerals (especially quartz and feldspar), QFR
compositions change from dominantly R to dominantly Q and
F.
Thus, finer-grained sandstone (texturally more mature;
higher QFR percent Q and F) is produced by the breakage of
coarser grains (texturally less mature; higher QFR percent R).
Interpretation of compositional maturity of sandstone is
complicated by this artifact of method because of the
dependence of composition on grain size, but interpretation
of textural maturity is straightforward. Thus, the use of QFR
petrographic methods is more useful for paleoclimate and
other studies, where textural maturity is diagnostic (e.g.,
Suttner et al., 1981; Suttner and Basu, 1985). Where
reconstruction of original source rocks is the goal, the Gazzi-
Dickinson method (QFL) is recommended (see below).
Compositional maturity (stability)
Most minerals and rocks are unstable at Earth's surface
because they originated at higher pressures and temperatures,
and in different geochemical environments. As a result, surface
conditions are continuously acting on these minerals and rocks
in ways that create more stable assemblages. Metastability is
common at Earth's surface due to slow reaction rates at low
temperatures; nonetheless, given the immensity of geologic
time,
even metastable minerals are gradually altered.
In general, minerals crystallized at higher pressures (P) and
temperatures (T) are less stable at Earth's surface than are
lower-PT minerals. Thus, the inverse of Bowen's reaction
series is Goldich's weathering series, which indicates that, for
example, albite is more stable at Earth's surface than is
anorthite, and biotite is more stable than olivine. Thus, one
measure of compositional maturity is the stability of mineral
assemblages relative to their parent assemblages in their
provenance areas.
The ZTR index (Hubert, 1960) is a commonly used mineral-
stability measure. It is determined by dividing the quantity of
zircon, tourmaline and rutile (super-stable accessory minerals)
by the quantity of total accessory minerals. Thus, super-stable
mineral assemblages have ZTR values close to unity.
Petrographically determined measures of stability include
the proportion of total quartz in the QFL population (using
the Gazzi-Dickinson method, which lessens the dependence
of composition on grain size; lngersoll et al., 1984). Sub-
populations also indicate relative stability; for example, Qm/
(Qm + Qp), Qp/(Qp + Lm + Lv + Ls), Ls/(Lm + Lv -f Ls), Qm/
(Qm
4-
Fk -h Fp) values all tend toward unity in super-stable
sandstone. Polycrystallinity and undulosity tend to weaken
quartz grains, so that ultra-quartzose sandstone tends to have
relatively low proportions of both (Basu, 1985). Super-stable
conglomerate consists predominantly of quartzite clasts,
analogous to super-stable quartz arenite (including ortho-
quartzite).
Stability of mudrock is highly dependent on climate because
stability of clay minerals is determined primarily by chemical
weathering regime. Thus, light leaching produces montmor-
illonite and illite, and related clays, depending on starting
430 MEANDERING CHANNELS
material. With increased leaching, kaolinite
is
favored. With
extreme leaching, aluminum
and
iron oxides
are the
primary
residuum (laterites,
q.v.),
thus representing extreme composi-
tional maturity (super stable).
Ratios
of
major oxides also indicate compositional maturity
because
the
oxides behave differently
at
Earth's surface.
MgO,
CaO,
and
Na20 tend
to
decrease, whereas
K2O,
SiO2, TiO2,
and Fe2O3 tend
to
increase during weathering (Pettijohn,
1975).
Various ratios
of the
latter oxides
to
total oxides
approach unity with increasing stability.
lngersoll,
R.V.,
BuUard,
T.F.,
Ford,
R.L.,
Grimm,
J.P.,
Pickle,
J.D.,
and Sares,
S.W.,
1984.
The
effect
of
grain size
on
detrital modes:
a
test
of the
Gazzi-Dickinson point-counting method. Journal
of
Sedimentary
Petrology,
54:
103-116.
Kay,
M., 1951.
North American Geosynclines. Geological Society
of
America Memoir,
48,
143pp.
Pettijohn,
F.J.,
1975. Sedimentary Rocks. Harper
and Row, 3rd edn.
Suttner,
L.J., and
Basu,
A.,
1985.
The
effect
of
grain size
on
detrital
modes:
a
test
of the
Gazzi-Dickinson point-counting method—
discussion. Journalof Sedimentary Petrology,
55:
616-617.
Suttner,
L.J.,
Basu,
A., and
Mack,
G.H.,
1981. Climate
and the
origin
of quartz arenites. Journalof Sedimentary
Petrology,
51: 1235-1246.
Discussion
Textural maturity
and
compositional maturity
are
distinctly
different measures
of the
degree
to
which sediment
has
been
modified during
the
diverse processes
of
weathering, erosion,
transportation
and
deposition. Distinction between these
two
attributes, however,
has not
always been clear.
The
"gray-
wacke problem"
is an
informative example
of
possible'
confusion.
Geosynclinal theory equated "graywacke" (poorly sorted
sandstone, with abundant "fine-grained matrix") with
"eu-
geosynclinal sedimentation" (e.g.,
Kay,
1951). When most
"graywackes" were recognized
as
turbidites,
it was
said that
turbidites were poorly sorted mixtures
of
sand
and mud. The
presence
of
abundant "matrix"
in
"graywackes"
was
cited
as
evidence. Detailed petrographic study
of
"graywackes"
has
shown that,
in
almost
all
cases, most
of the
"matrix"
is
pseudomatrix, that
is,
physically deformed
and
chemically
altered detrital grains (Dickinson, 1970). Recognition
of
pseudomatrix allows
for the
correct interpretation
of
prove-
nance (based
on
assignment
of
pseudomatrix
to
detrital-grain
categories)
and the
correct interpretation
of
textural maturity
(little original detrital matrix, which means that sand
was
well
sorted prior
to
diagenetic alteration). Recognition
of
pseudo-
matrix
as
detrital grains
is
consistent with
the
observation that
modern turbidites
are
well sorted, that
is,
have less than
5
percent matrix (e.g., Hollister
and
Heezen, 1964). Thus,
the
"graywacke problem" resulted from
the
confusion
of
compo-
sitional maturity (unstable lithic grains
and
micas readily
become pseudomatrix during diagenesis) with textural matur-
ity (pseudomatrix
was
interpreted
as
detrital matrix, which
led
to
the
inference that turbidites were poorly sorted). Careful
petrographic study
is
needed
to
recognize
the
original detrital
texture
and
composition
of
sandstone,
and for the
correct
interpretation
of
both textural
and
compositional maturity.
Raymond
V.
lngersoll
Bibliography
Basu,
A., 1985.
Reading provenance from detrital quartz.
In
Zuffa,
G.G.
(ed.).
Provenance
of Arenites.
D.
Reidel,
pp.
231-247.
Boggs,
S. Jr., 1968.
Experimental study
of
rock fragments. Journalof
Sedimentary
Petrology,
38:
1326-1339.
Dickinson,
W.R.,
1970. Interpreting detrital modes
of
graywacke
and
arkose. Journal of Sedimentary Petrology,
40:
695-707.
Ehrlich,
R., and
Weinberg, 1970.
An
exact method
for
characterization
of grain shape. Journalof Sedimentary
Petrology,
40:
205-212.
Hollister,
CD., and
Heezen,
B.C., 1964.
Modern graywacke-type
sands.
Science, 146: 1573-1574.
Hubert, J.F., 1960. Petrology
of
the Fountain
and
Lyons Formations,
Front Range, Colorado. Colorado School
of
Mines Quarterly,
55:
242pp.
Cross-references
Climatic Control
of
Sedimentation
Feldspars
in
Sedimentary Rocks
Grain Size
and
Shape
Heavy Minerals
Provenance
Quartz, Detrital
Sands
and
Sandstones
Tectonic Controls
of
Sedimentation
Weathering, Soils,
and
Paleosols
MEANDERING CHANNELS
The planform
of
rivers—the channel pattern
as
seen from
an
overflying aircraft—is quite varied
and has
invited many
attempts
at
classification
by
river scientists.
An
important early
planform classification
by
Leopold
and
Wolman (1957)
divided rivers into three categories: braided, meandering
and
straight. Braided channels
are
multiple-thread patterns
and
contrast with
the
single-thread meandering
and
straight
channels; meandering channels
are
distinguished from their
straight counterparts
by
their sinuously winding course.
The
term "meander"
has
become part
of our
everyday language
and derives from
the
apparently aimless wandering habit
of
the
Maiandros
in
Phrygia,
a
winding river named
by the
ancient
Greeks
in
what
is now
modern Turkey.
More recently river scientists have recognized
the
limitations
of this simple tripartite river-planform division
and
have
stressed instead
the
continuum
of
river planforms. Never-
theless,
the
meandering channel remains
an
important
end-
member
in all
modern classifications
of
river planform
and is
one
of the
most common patterns found
in
nature.
It may be
defined
as a
single-thread river channel consisting
of a
continuous series
of
channel bends
of
alternating curvature
together displaying elements
of
both periodic
and
random
geometry.
The planform
and
channel geometry
of
river meanders
One
of the
most intriguing properties
of
meandering channels
is
the
sometimes remarkably regular periodic waveform they
display (see Figure M9).
In its
purest form this regular sinuous
waveform
has a
characteristic shape (sine-generated)
and a
geometry that scales with channel width
so
that meanders tend
to
be
geometrically self-similar regardless
of the
size
of the
river.
For
example,
on
average, meander wavelength typically
MEANDERING CHANNEI
S
431
2CX)m
Figure M9 (A) Aeri.il photofiraph of the freely meandering Pemliina River, Alberta, Canada; iB) aerial photograph afcunlined meandering Beaver
River, Alhert.i, Canada.
is 8-12 times the biinkfiill channel width and meander-bend
curvature Is about 2-3 times the channel width. Because
channel width (WO is related to discharge IQ). meander
wavelength (L) is also related to river diseharge
(where k ranges from 30 to 60. depending on the return period
ol"
Q), a relationship whieh has been exploited in paleo-
hydrologic studies of aneient rivers (see Dury. 1964).
Initially river scientists were preoeeupied with the periodie
character of river-nieaiider traces and only more recently have
acknowledged their non-periodie. inlermiitcnt, or random
beha\ior. In most meandering channels the orderliness of
highly periodic sinuosity is blended with, and disturbed by, a
variety of planform disorder ihat relates lo ehannel migration
processes. Kor example, regular meander-bend development
can be distorted by heterogeneities in the alluvium through
whieh bends are migrating, meander bends can migrate
laterally to forin complex meander loops whieh are mueh
larger than the associated meander bends, and high sinuosity
meander bends can erode back on themselves to form meander
cutotTs. Typical variations in the relative importance of these
two planform elements led Kellerhals ef al. (1976) to
distinguish three eategories of meandering: irregular, strongly
repeating, and tortuous.
Although the shape of channel bends in a freely migrating
meander train can be syiiimetrical and sinusoidal, even slight
river confinement eommonly leads to bend asymmetry as the
train migrates downvalley. In eases of pronounced confine-
ment (e.g., meanders formed in the confines of abandoned
glaeial meltwater channels of the Canadian Prairies), meander
asymmetry is extreme and is an indicator of flou and
migration direction (Figure M9(BK.
Channel cross-sections at or somewhat dounstream of the
axis of meander bends arc distinctly asymmetrical with deeper
flow oeeupying the outer bank /one and the deposition of a
point bar eausing shoaling at the opposite inner bank. At
points of meander infleetion flow typically is shallower over a
432 MEANDERING CHANNELS
ehannel-wide bar ealled a riffle. This "pool and riffle" sequence
is a distinctive morphological characteristic of all meandering
channels.
Flow in meandering channels
Although much debated, there is no generally aecepted
universal explanation for why meanders form in rivers (see
reviews by Callander. 1978 and Knighton, 1998). Whatever the
eause of meandering, however, the general pattern of flow
through the alternating channel bends of meanders is
well-
known and understood. In the absence of secondary flow,
bend flow seeks to conserve angular momentum so that it
tends to conform to that of
a
fi'ee vortex with high velocity at
the smaller radius of the inner bank and lower velocity al the
outer bank where radial acceleration is lower. But secondary
flow redistributes momentum in the bend to achieve a reversal
of this relationship in the zone of fully developed bend flow.
Near the water surface where velocity is highest, secondary
flow is dominated by eentrifugal "forces" and flow moves
toward the outer bank. Near the bed. where veloeity and thus
the eentrifugal effects are lowest, the balance of forces is
dominated by the inward hydraulic gradient of the super-
elevated water surface and secondary flow moves toward the
inner bank. The three-dimensional expression of
these
forces is
helical flow which has an alternating sense through successive
meander bends. One of the important consequenees of helical
flow in meanders is that .sediment eroded from the outside ofa
meander bend tends to be moved to the inner bank or point
bar of the next downstream bend.
The process of meander formation and maintenance
Meander planform geometry is controlled by the process of
lateral migration. Meandering channels erode the outer banks
of channel bends and maintain a eonstant channel width by
achieving a matching rate of deposition on the point bar
forming the inner bank. Rates of migration may be exceed-
ingly slow (a small fraction of channel width/year) or very
rapid (several channel widths/year) depending on the ratio
stream power/bank strength, and on the degree of bend
curvature (Hlekin. 1974; Hiekin and Nanson. 1975. 1984).
Depending on the particular pattern of bend erosion and
deposition,
meanders ean exhibit a variety of
evolution.
Simple
bends can increase in amplitude by extension until decreasing
bend curvature deflects or constrains the direction and rate of
motion.
Phase differences in the flow and bend alignment can
lead to bend rotation, lobing and compound meander loops.
Where confinement consti'ains meander development, bend
motion is mainly restricted to downvalley translation. These
motions are further modified by the formation of gooseneck
and chute cutoffs which short-circuit the meandering channel
path,
sometimes forming ox-bow lakes.
The deposits of meanders: meander floodplains
The nature of sediments deposited to form the floodplains of
most meandering rivers is intimately related to the sediment
available for point bar deposition and to the process of lateral
migration.
The classical point bar model of deposition in sand-
and gravel-bed rivers, originating primarily from the work of
Allen (1963). is illustrated in Figure MIO. Within-ehannel
lateral accretion deposits of gravel overlain by sand arc capped
by overbank fines deposited during flood discharges. Lateral
accretion deposits consist of a basal gravel platform of bed
material on which are deposited upward-flning sands and silts
representing the declining energy of the depositional environ-
ment as the point bar aggrades to less frequently inundated
elevations. Internal sedimentary structures vary from dune to
ripple related cross strata and trough cross-lamination in the
lateral accretion deposits to horizontal strata in the overbank
deposits. This upward-lining sediment column with its
assoeiated upward downsealing of sedimentary structures, is
%::••.•
SCROLL BARS & SWALES
LATERAL ACCRETION
(FINING-UPWARD THAI WFri\v-f
^t'^??'-.
SEQUENCE) --^^r^LWEG
^J^^i^
y: QLDER
•
- ^
CHANNEL
i FILL yjC
s^-Jp CHUTE CUT-OFF
DUNES
Figure MIO The geomorphic setting and sedimentological elements ot ihe upward-fining sequence, the basis of the classical meandering-river
facies model (modified tVom Walker and Cant, 1984|.
MEANDEKINC; CHANNFIS
packaged in units bound by signioidal epsilon ceoss strata
(ECS),
corresponding to distinct phases of channel migration
and point bar deposition. These packages of sediment often
form arcuate ridges or scroll bars, separated one from the
other by troughs in the surface of the floodplain.
More recent investigations, however, have recognized many
variations on this elassieal model of meander deposition and
still others describe no{idplain sedimentology which seems to
be entirely unconventional. For example, in MialFs (1985)
models of meandering river floodplain (gravel meandering,
gravel-sand meandering, sand meandering and Jine-grained
meandering) he recognizes the important control of sediment
type on the details of floodplain sedimentology. These river
channel styles reflect a progressive increase in the abundance of
fines although coarser basal lateral accretion deposits are
common to all models. Floodplains with lateral accretion
deposits consisting of sandy silt with negligible upward fining
liave been described by Nanson (1980). Others have also noted
that meandering river floodplains formed in environments
where silt is abundant and gravel and sand are mueh less
available, particularly in low-gradient tropieal regions, also do
not appear to be adequately described by the elassical Allen
model (Jackson. 1981: Smith, 1987). Sitn'ilarly, the negligible
role played by lateral aeeretion in some steep upland valleys
formed in non-cohesive sediments, as well as in low-gradient
channels in cohesive sediments, is formalized in the genetic
classification of lloodplains by Nanson and Croke (1992).
The floodplains of meanders may also include a signilicani
proportion of counterpoint (or eddy-accretion) sediments of
horizontally bedded silt and fine sand deposited in flow
separation zones at points of abrupt direction change in the
channel. In confined meanders the abrupt angle bends that
occur where the channel impinges on (he valley wall
(Figure M9(B)) form systematically alternating zones of
flow separation downstream at the valley walls. These
alternating gyres Irap eddy accretion sediments whieh form
concave bank benches with a eurvature opposed to that of the
convexly curved scroll bars on the adjoining point bar (Hiekin.
1979.
1986; Nanson and Page. 1983; Burge and Smith. 1999).
River planform facies and architectural element
models of meander deposits
Attempts to characterize meandering river sediments in terms
of bedform-scale sedimentary struetures and assemblages
(planform tacies models) that would distinguish them with
some certainty frotn other river types (especially braided) have
not fulfilled early promise. Although general sedimentological
properties for sand/gravel-bed rivers are well represented by
these faeies models, overlap among planform groups is so
significant that their diagnostic utility is weak and olten
misleading (Jackson, 1978; Brierley and'Hickm. 1991; Hiekin,
199.V Miall, 1996). In recent years some tluvial sedimentolo-
gists have advocated the development of bar-seale or
arehitectural element-scale models to charaeterize fluvial
sediments (Miall, 1985). It has been argued that such planform
element models may be iar more effective diagnostic toots than
are bedform-scale facies models. This development represents
a return to earlier geomorphic concepts sueh as those
developed by Happ etal. (1940) and Fisk (1944, 1947) and
espoused by Leopold. Wolman and Miller (1964) who
recognized that floodplains of rivers sueh as the Mississippi
consist of an assemblage of distinct morphological and
sedimentological elements (such as channel fills, point bars.
meander scrolls, sloughs, levees, baekswamp deposits, and
crevasse splays). But such an element approach to the
depositional reeord requires three-dimensional mapping of
sueh features and the typically poor access to exposures in
the field is rarely adequate for this purpose. There is some
prospect, however, that remote sensing of the subsurfaee
architecture by means such as ground-penetrating radar
(GPR).
shallow seismic and electrical resistivity surveys, may
yield the necessary data to allow this approach to be developed
further (Alexander
(•/a/..
1994; Leclerc and Hiekin, 1997).
Summary
Meandering channels, single-thread river-ehannels consisting
of a continuous series of ehannel bends of alternating
curvature, are an important end-member in all modern
classifications of river planform and is one of the most
common river patterns found in nature.
River meanders tend to have a sinuous waveform of
eharacleristic shape (sine-generated) and a geometry that
scales with channel width so that meanders tend to be geo-
metrically sell-similar regardless ol' the size of the river. The
regularity of this sinuous pattern, also reflected in a eor-
responding "'pool and riffle" bed geometry, can be disturbed
by such effeets as tloodplain heterogeneities, meander cutotTs.
complex loop development, and ehannel confinement.
In ihe classieal point bar model of deposition in sand-bed
and gravel-bed ri\ers within-ehannel lateral accretion deposits
of gravel overlain by sand are capped by overbank fines
deposited during flood discharges. This upward-fining sedi-
ment column with its upward downsealing of sedimentary
structures, is packaged in units bound by sigmoidal epsilon
cross strata (ECS), corresponding to distinct phases of channel
migration and point bar deposition.
More recent investigations have recognized many variations
on this classical model of meander deposition, encompassing
Ihe broader range of sediments aetually found in nature,
including counterpoint deposits common in confined mean-
ders.
Mud-dominated floodplains remain poorly studied and
the subject of current researeh.
Planlbrm facies models at the bedform scale have enjoyed
limited diagnostic success. Bar-scale or architectural element-
seale models of Hoodplain sedimentology hold more promise
but require three-dimensional data dependent on remote
sensing of the stibsurface by techniques such as ground-
penetrating radar.
Edward J. Hiekin
Bibliography
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Gawthorpc. R.L.. 1994, Holocene meander boll cvolulion in an
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Cross-references
Erosion
and
Sediment Yield
Floodplain Sediments
Placers, Fluvial
Rivers
and
Alluvial Fans
MELANGE; MELANGE
Melange (spelled melange in much English-language literature)
describes heterogeneous rock units that contain blocks,
typically of hard or structurally competent lithologies,
surrounded and supported by a fine-grained matrix of shale,
slate, or serpentinite. Considerable controversy has sur-
rounded the term melange since its first use by Greenly
(1919), in the description ofa unit from Anglesey in the British
Isles.
Interest in meianges was revived during the development
of plate tectonics, when melanges were identified as tectonic
products of subduction (e.g., Hsu, 1968). Controversy has
surrounded both the definition of the term melange, and the
interpretation and origin of units so described.
Definitions
Almost all definitions of melange include as critical compo-
nents the presence of blocks of varied sizes and lithologies, not
in contact with each other, and surrounded by a finer-grained
matrix. Most geologists have also agreed that a melange has to
be a unit that is mappable, typically at a scale such as
1:25,000
(Silver and Beutner, 1980) or
1:24,000
(Raymond, 1984).
Authors have differed as to whether the presence of a tectonic
fabric in the matrix is essential to the definition of melange.
Most described examples do have a fabric of some sort; rock
bodies in which blocks are present in a clearly untectonized
matrix are typically labeled using more interpretive terms
such as "olistostrome" or "debris flow"
{q.v.),
which imply
specifically sedimentary processes of origin. Some authors
have suggested that such sedimentary units be excluded from
the definition of melange (Hsu, 1974). Nonetheless, the term
has been applied to many units entirely lacking tectonic fabric,
and the type example described by Greenly (1919) locally lacks
fabric. A purely descriptive definition is therefore preferable;
units of known tectonic origin are commonly distinguished as
"tectonic melanges."
Also in dispute is whether the blocks have to include
"exotic" lithologies, such as the blueschist and ultramafic units
in the well-known Franciscan melanges of California.
Although exotic blocks were regarded as essential in the
definition adopted by Raymond (1984), it is often difficult to
decide what lithologies are truly "exotic", and close textural
and fabric similarities exist between chaotic units with and
without exotics (Brandon, 1989). Thus a definition consistent
with most modern usage is: "an internally disorganized unit
that is mappable at
1:25,000
scale, that is characterized by the