Middleton G.V. (Ed.) Encyclopedia of Sediments and Sedimentary Rocks
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HYDROCARBONS IN SFDIMFNTS
retained within
the
source rock
and to a
lesser extent
on
mineral surfaces along migration pathways.
Occurrences
Although hydrocarbons have been reported almost everywhere
in the earth's crust, they are
far
more abundant
in
sediments
than
in
igneous
or
metamorphic rocks. This reflects
the
biogenie origin
of
most hydrocarbons
in
sedimentary roeks.
Hydrocarbons make up
a
very low proportion
of
the biomass
and henee their abundanee in recent sediments is generally low,
in the 20ppm
to
100 ppm range (Tissot and Welte. 1984. p.95).
In aneient non-reservoir sediments, especially organic-rich,
fine-grained
rocks,
their abundance
is
much higher,
up to
several thousand
ppm. The
large scale accumulations
of
hydrocarbons within commercial
oil
fields
are due to
their
migration out
of
large volumes of source rocks and into more
permeable and porous beds where they have been trapped.
A diverse range of alkanes and aromatic hydrocarbons have
been reported
in
petroleum
and
sediments.
The
average
hydroearbon composition
of a
crude
oil has
been stated
to
be
?>?•
percent acyclic (normal
and
branched alkanes).
32
per
cent cycloalkanes
and 35
percent aromatic hydro-
carbons {Killops and Killops, 1993, p.137). These proportions
vi'ill vary depending on the nature of the organic matter
in
the
oil source rock
and the
level
of
thermal maturity, with
saturaled hydrocarbons becoming more predotninant at higher
tnaturity. Organic geochcmists usually analyze for compounds
with less than (brty carbon atoms, but the recent development
of high temperature gas chromatography
has
shown higher
molecular v\eight hydroearbons, up to
C1211
may be ubiquitous
in oils (Hsieh and Philp, 2001). The simplest and usually
the
most abundant hydroearbons are the straight-chain n-alkanes,
which
can
range from
C]
(methane)
to
C^d,..
Particularly
in
recent sediments,
the
C25-C13 n-alkanes often show
a
pronouneed
odd
carbon number preference. This
is due to
ihe predominance
of odd
carbon number n-alkanes
in
plant waxes. Simple branched alkanes sueh
as
the iso-alkanes
(2-methylalkanes) and anteiso-alkanes (3-mcthylalkanes) and
cyclic alkanes sueh
as the
eyelohexanes
and
cyclopentanes
with
a
similar carbon number range
to
the n-alkanes are also
ubiquitous
in
crude oils
and
ancient sediments. Many
hydrocarbons have carbon skeletons that can be easily related
10
a
precursor originally synthesized
by an
organism
(Figure H7). Such compounds
are
called biomarkers (short
for biological markers)
or
chemical fossil.s (Peters
and
Moldowan, 1993). Many
of
the most commonly found
and
abundant biomarkers are those that are based on the isoprcne
niiit. These include acyclic isoprenoids such
as
pristane (Ci^i)
and phytane (Ci^i)
and
cyclic compounds such
as
steranes.
OH
cholestane
Figure H7 Example
of
biological precursor—biomarker relationship.
Cholesterol is an example of a sterol, compounds found
in
all higher
(euktiryolic) organisms inc luHinj; humnns. During diagenesis,
dehydration and hydrogendtion ol cholesterol occurs to give
cholestane, a sterane that is almost ubiquitous
in
petroleum and
sediments.
hopanes, carotenoids (e.g.. /^carotane)
and
other terpanes
(Figure H8). Sterane and terpane distributions (usually hopane
and tricyclic terpanes)
are
those most commonly used
by
petroleum geochemists to "fingerprint" oils so that they can be
compared to other oils or potential source rocks for oil-oil and
oil source correlation. Aromatic compounds
are
also abun-
dant
in
petroleum
and
extracts from aneient sediments.
Benzene
and
toluene
are the
major tight aromatics while
heavier fractions
of
petroleum contain alkylated benzenes,
naphthalenes, phcnanthrenes, triaromatic steroids and triter-
penoids and other polyeyelic aromatic hydrocarbons (PAHs)
(Figure H8). Alkylated compounds predominate
in
oils
as
opposed
to
anthropogenic combustion sources where parent
PAHs are dominant. The distributions of several saturated and
aromatic hydrocarbons (e.g., steranes
and
methylphenan-
threnes) ean be used as an indieator
of
the thermal stress that
organic matter
has
been subjected
to.
Although
a
range
of
0.02 percent
to 5
10 pereent have been reported tor olefins
in
oils (Curiale and Frolov. 1998), their concentration
is
usually
toward the lower part of this range. Olefins are generally much
more abundant
in
recent sediments where they
are
directly
inherited from cotnpounds originating from living organisms.
Some
of the
olefins found
in
petroleum
are
thought
to be
biogenic compounds directly derived from the souree rock
or
picked up during migration when petroleum eame into eontaet
with lower maturity sediments. Other olefins have
been attributed
an
abiogenic origin, either from radioiytic
dehydrogenation
of
saturated hydrocarbons due
to
exposure
to radioactive elements or lo enhanced thermal activity such as
when
an
igneous intrusion comes into close contact with
a
reservoir (Curiale and Frolov, 1998).
Origin
Although
a
biogenic origin
of
hydrocarbons is easily the most
currently tavored theory, this has not always been the case and
nor is
it
accepted by everybody today. There are three general
classes
of
ideas that have been
put
forward
to
explain
the
phylane
[i-carotane
hopane
alkyl benzene
phenanthrene
tricyclic
triterpane
C,, triaromaticsterane
(cf. choleslanein Fig.H7)
Figure H8 Examples
of
hydrocarbons commonly found
in
sediments.
R indicates alkvi chain.
366
HYDROXIDES AND OXYHYDROXIDE MINERALS
origin of petroleum, and by imphcation, hydrocarbons in
sedimentary rocks: inorganic, biogenic, and duplex origins.
In the nineteenth century, many people did not think that
the formation of petroleum involved the biosphere. Reactions
using inorganic substrates such as the Fischer-Tropsch
reaction, which commercially is used to synthesize hydro-
carbons from carbon monoxide and hydrogen on a metallic
catalyst (Olah and Molnar, 1995, p. 66-75), were suggested as
likely mechanisms. During the twentieth century, most people
came to accept the biogenic theory outlined below although
some workers still support an inorganic origin (e.g., Porfir'ev,
1974;
Gold and Soter, 1982). A vast body of evidence indicates
that this is very unlikely for most sedimentary hydrocarbons.
Small amounts of abiogenic hydrocarbons, mostly methane,
are found in Precambrian Shield and mantle-derived rocks but
not enough to constitute economic deposits (Sherwood-Lollar
etal., 1993; Jenden etal., 1993; Sugisaki and Mimura, 1994).
The duplex theory holds that petroleum is a mixture of
compounds having inorganic and biogenic origins. It was
proposed (Robinson, 1963) at a time when it was more difficult
to relate many of the constituents of petroleum to a biogenic
origin. Many of today's proponents of an inorganic origin for
petroleum tend toward a duplex origin in order to explain
some of the more obvious biogenic compositional character-
istics of oils, such as the occurrence of biomarkers.
Most geologists and geochemists believe that petroleum is
the product of the thermal degradation of insoluble organic
matter derived from dead organisms buried in sedimentary
rocks.
This is the only theory that satisfactorily explains the
composition and geographical and geological distribution of
petroleum. In the past, it was suggested that petroleum was
simply an accumulation of the hydrocarbons derived from
living fiora and fauna that occur in recent sediments. Hence
there was no need for any chemical conversion of organic
matter into petroleum with time or depth of burial. This is no
longer thought to be credible for a number of reasons, most
notably because hydrocarbons are found in very low
abundance in living organisms and recent sediments (these
contain mostly non-hydrocarbon compounds), and petroleum
contains abundant hydrocarbons not found in living organ-
isms,
especially in the lower molecular weight C4-C8 range.
There are three stages in the transformation of biogenic
organic matter into petroleum (see also Maturation, Organic).
Diagenesis is the low temperature (up to 50-60°), largely
mierobial mediated, irreversible alteration of organic debris in
the water column, at the water/sediment interface and in near
surface sediments. The biogenically derived organic matter
that survives bacterial attack at this stage is preserved as
mostly insoluble organic matter called kerogen. During
diagenesis, hydrocarbons are in low abundance. Most of the
compounds that can be extracted from sediments still contain
functional groups and are directly inherited from living
organisms. Biogenic methane due to methanogenie bacteria
activity is the only hydrocarbon generated in significant
amounts during diagenesis. Catagenesis is the second stage of
organic maturation and takes place between about 50°C and
150°C. This is a thermally mediated process in which kerogen
cracks to form oil and gas. The catagenesis stage thus includes
the principle zone of oil formation and later, the initial stages
of gas formation when C2-C4 hydrocarbon gases with an
increasing amount of methane are generated in large amounts
from the cracking of previously generated larger molecules and
kerogen. Biogenically derived molecules are diluted by these
neo-formed hydrocarbons which have a less obvious structural
relationship to their biogenic precursors. Metagenesis is the
final thermal degradation of both residual kerogen and the
bitumen formed during catagenesis to the most thermodyna-
mically stable products, carbon (pyrobitumen or graphite) and
methane. It occurs at temperatures between 175°C and 250°C.
Martin Fowler
Bibliography
Curiale, J.A., and Frolov, E.B., 1998. Occurrence and origin of olefins
in crude oils. A critical review.
Organic
Geochemistry, 29: 397-408.
Gold, T., and Soter, S., 1982. Abiogenic methane and the origin of
petroleum. Energy Exploration and Exploitation, 1: 89-104.
Hsieh, M., and Philp, R.P., 2001. Ubiquitous occurrence of high
molecular weight hydrocarbons in crude oil.
Organic
Geochemistry,
32:955-956.
Hunt, J.M., 1996. Petroleum Geochemistry and
Geology,
2nd edn. New
York: W.H. Freeman.
Jenden, P.D., Hilton, D.R., Kaplan, I.R., and Craig, H., 1993.
Abiogenic hydrocarbons and mantle helium in oil and gas fields. In
Howell, D.G. (ed.). The Future of Energy Gases. US Geological
Survey Professional Paper 1570, pp. 31-56.
Killops, S.D., and Killops, V.J., 1993. An Introduction to Organic Geo-
chemistry. Longman Scientific & Technical.
Olah, G.A., and Molnar, A., 1995. Hydrocarbon Chemistry. John
Wiley & Sons.
Peters, K.E., and Moldowan, J.M., 1993. The Biomarker Guide.
Interpreting Molecular Fossils in Petroleum and Ancient Sediments.
Prentice Hall.
Porfir'ev, V.B., 1974. Inorganic origin of petroleum. American Asso-
ciationof Petroleum Geologists Bulletin, 58: 3-33.
Robinson, R., 1963. Duplex origin of petroleum. Nature, 199:
113-114.
Sherwood Lollar, B., Frape, S.K., Weise, S.M., Fritz, P., Maeko, S.A.,
and Welhan, J.A., 1993. Abiogenic methanogenesis in crystalline
rocks.
Geochimica et Cosmoehimiea Aeta, 57: 5087-5097.
Sugisaki, R., and Mimura, K., 1994. Mantle hydrocarbons: abiotic or
biotic? Geochimicaet Cosmoehimiea Acta, 58: 2527-2542.
Tissot, B.P., and Welte, D.H., 1984. Petroleum Formation and Occur-
rence. 2nd edn., Springer-Verlag.
Cross-references
Diagenesis
Gases in Sediments
Kerogen
Maturation, Organic
Oil Sands
Mutation, Organic
HYDROXIDES AND OXYHYDROXIDE
MINERALS
The group of oxides, hydroxides and oxyhydroxides comprises
several cations such as Al, Fe, Mn or Ti, where the anionic
part is either an oxygen (-0), a hydroxyl group (-0H) or
an oxygen and a hydroxyl (-OOH). This article deals
essentially with the iron minerals (collectively called iron
oxides), because aluminum minerals such as gibbsite (a-
AI(OH)3), diaspore (a-AlOOH) and boehmite (y-AlOOH) are
handled in the article on bauxites
{q.v.),
whereas manganese
oxides and phyilomanganates are essential constituents of
HYDROXIDES AND OXYHYDROXIDE MINERALS
367
iron-manganese nodules (^.
v.).
The TiO2 polymorphs rutile,
anatase and brookite are of minor importance in sediments
and sedimentary rocks.
Minerals
In sediments and sedimentary rocks, oxides (hematite a-Fe2O3,
magnetite Fe^"*"Fe2^O4, maghemite y-Fe2O3, wuestite FeO,
ilmenite FeTiO3) and oxyhydroxides (goethite a-FeOOH,
lepidocrocite y-FeOOH, ferrihydrite ~FeOOH) do occur.
Sehwertmannite, Fe8O8(OH)6SO4, is transitional between
oxyhydroxides and sulfates (for details see Cornell and
Schwertmann, 1996).
Occurrences
The iron contents of sediments are highly variable. On
average, carbonate rocks contain ca. 6mgg~
3,
sand-
2
stones ca. I4mgg~' and elaystones ca. 71mgg"'. The iron
oxides are either of detrital origin or neoformed in the
sediments. Titanomagnetites and ilmenite belong to those
detrital minerals that survive weathering due to high stability
in surface environments, and, hence, may concentrate, for
example, in placers (^.
v.),
up to mining quality. Hematite,
which is the major pigmenting mineral in red beds (^.
v.)
or any
other red-colored rock, may be either of detrital origin or it
formed diagenetically. The banded
iron
formations (BIF) or
itablrites are Precambrian, thin-bedded or laminated chemical
sediments, where hematite and magnetite are interbedded with
chert layers. In other kinds of sedimentary iron ores, hematite
may be associated with goethite and/or magnetite (see Iron-
stones and Iron Formations, and Flichtbauer, 1988). Yellowish
to brown colors indicate the presence of goethite, which is
present in small concentrations ubiquitously but in Paleozoic
and older rocks to a lesser extent than is hematite. Laterites
(^.
V.),
ferricretes and bauxites contain goethite, hematite and
also maghemite. Quaternary bog ores, which are the younger
analogs to ferricretes and bauxites, are dominated by goethite
and ferrihydrite, whereas hematite is absent. Tertiary sedi-
ments may contain lepidocrocite in addition to goethite.
Wuestite has been reported as rims around magnetite in
sandstones.
Properties
The basic structural unit in most of the iron oxides are Fe^"*"
cations coordinated by six anions (O^, and/or OH~). These
units are connected either by sharing one oxygen across a
corner, two oxygens across an edge or three along a face. With
the addition of fourfold coordinated Fe^'*' and/or Fe''''' cations
in the spinel phases, distinct structures with densities between
4gmL~' and 5.2gmL~' can arise by mixing several kinds of
bonds. In cases where one iron oxide is present, the color of
sediments {q.v.) is an indicative property. Variations in crystal
size,
however, or mixtures of two minerals (e.g., goethite and
hematite) not only change the hues, but also might mask other
iron oxides. Identification and characterization is therefore
better done by X-ray diffraction using Co radiation. Depending
on "crystallinity", detection limits may range from a few
mg g~' for micron-sized hematites in sandstones to ~0.2gg~'
for ferrihydrite. Mossbauer spectroscopy offers even lower
detection limits because of being sensitive to solely ^^Fe.
Crystal size and deviation from the ideal chemical composi-
tion reflect the physico-chemical environment in which
authigenic iron oxides have grown. In laterites, ferricretes,
and bauxites, crystal sizes of goethite and hematite range from
tens to hundreds of nanometers, whereas in sandstones
hematites may have grown to several microns in size. The
first mentioned environments with their intensive chemical
weathering yield also goethites and hematites, in which AP'*'
substitutes for Fe^+. Maximum substitutions of
O.33molmor' Al/(A1-|-Fe) have been found for goethites; in
hematites the substitution seems to be limited to
0.16molmol"'. Smaller crystal sizes and Al substitution
distinguishes iron oxides derived from supergene weathering
processes from those found in sandstones, carbonates and
other rocks, where crystal sizes are usually much larger and
substitution of foreign cations is almost absent.
Besides Al, other heavy metals such as V, Cr, Co, or Ni may
substitute in iron oxides, although with much lower absolute
contents. In ferricretes so much phosphate is associated with
iron oxides that it spoils their use as iron ores. Phosphate sorbs
strongly on iron oxide surfaces by forming a bidentate
mononuclear complex, in which two oxygens of the P04~
anion are shared with two surface oxygens (which were OH
groups before adsorption) bound to two Fe^'*' beneath the
surface. Selenite and sulfate adsorb similarly, but less strongly.
This kind of inner-sphere complexation has also been observed
for arsenate and chromate bound to sehwertmannite in
sediments from acid mine drainages.
Unpaired electrons of structural Fe(III) and Fe(II) result in
magnetic moments, which interact with each other and
depending on thermal motion produce magnetic ordering in
all iron oxides. The ordering temperatures of micronsized
crystals of magnetite, hematite and to some extent also
goethite are high enough (850, 956, and 400 K, respectively)
to preserve magnetizations even on a geological timescale.
When such particles settle freely during sediment formation,
they orient along the Earth's magnetic field and build up a
detrital remanent magnetization. Alternatively, a very stable
chemical remanent magnetization is acquired, when these iron
oxides grow authigenically in a sediment (see Dunlop and
Ozdemir, 1997).
Formation
Chemical weathering of primary, Fe(II)-containing silicates
liberates Fe^'"', which oxidizes and hydrolyses to form any of
the above-mentioned phases. Which iron oxide forms is
governed by pH, the rate of oxidation, the presence of
inorganic and organic anions, temperature, activity of Fe^'"'
and other foreign compounds. Near-neutral pH values and
higher temperatures favor formation of hematite over goethite.
With decreasing oxidation rates of Fe^^, ferrihydrite, lepido-
crocite or goethite are formed. In agreement with the
occurrences of less crystalline and less stable iron oxides in
younger sediments, kinetic factors seem to outweigh thermo-
dynamic stabilities, because at conditions prevailing in surficial
environments (i.e., low temperature, activity of water =1,
presence of oxygen), large crystals of goethite (a-FeOOH) are
the only thermodynamically stable phase. The ubiquitous
presence of hematite (a-Fe2O3) can be explained by lower
activities of water due to stronger binding of water in small
pores and/or due to higher temperatures, where hematite
becomes stable despite of a^fi =
1
{cf the formation of
368 HYDROXIDES AND OXYHYDROXIDE MINERALS
hematitic red beds
(^.v.)
in a warmer climate). A further factor,
which influences thermodynamic stability, is crystal size
because of increasing contributions of surface energy.
Transformation
By heating in dry conditions, all OH-containing oxides
transform finally to hematite via solid state reactions, but
depending on the nature of the compound, its crystallinity, the
extent of isomorphous substitution and other chemical
impurities, temperatures between 140°C and 500°C are
required. In the presence of water, dehydroxylation of goethite
to hematite requires hydrothermal conditions, whereas the
reverse reaction (hydroxylation of hematite) has not been
observed in situ because of kinetic hinderance. '^O isotope
measurements of such transformations have shown that a
solution step with transport via water molecules is always
involved (Bao and Koch, 1999).
The solubilities of Fe(ni) oxides at naturally occurring pH
values are so low that the concentration of Fe in solution is
orders of magnitudes lower than the solubility as Fe^'*'.
Although complexing organic compounds increase the total
solubility, the most efficient way to mobilize iron is by redu-
ctive processes mediated by bacteria or reducing compounds
(e.g., HS~). The high mobility of Fe^''' then allow redistribu-
tion of iron from the nanometer scale to landscapes. An
indicator for Fe^'*' as the feeding source is substitution of V^'*'
in lateritic goethites and hematites. Under natural pH and Eh
values, where Fe^"*' is stable, tetravalent VO^'*" coexists, but
not the unstable V^'^. A reduction VO^'^ to V^'^ followed by
incorporation is feasible, however, by a surface-mediated
reduction after sorption of Fe^'*' onto an oxide surface, which
decreases the redox potential of Fe'''''/Fe^''' in solution from
0.77 Volt to a value of ~0.35 Volt low enough to reduce
sorbed V''+ to V^+ (Schwertmann and Pfab, 1996).
Helge Stanjek
Bibliography
Bao,
H., and Koch, P.L., 1999. Oxygen isotope fractionation in ferric
oxide-water systems: low temperature synthesis. Geochimica
et Cosmoehimiea Acta, 63: 599-613.
Cornell, R.M., and Schwertmann, U., 1996. The Iron Oxides.
Weinheim: VCH...
Dunlop, D.J., and Ozdemir, O., 1997. Rock Magnetism. Cambridge
University Press.
Fuchtbauer, H., 1988. Sedimente und Sedimentgesteine. Stuttgart:
Schweizerbart.
Schwertmann, U., and Pfab, G., 1996. Structural vanadium and
chromium in lateritic iron oxides: genetic implications. Geochimica
et Cosmoehimiea Acta, 60: 4279-4283.
Cross-references
Bauxite
Colors of Sedimentary Rocks
Diagenesis
Iron-Manganese Nodules
Ironstones and Iron Formations
Laterites
Magnetic Properties of Sediments
Red Beds
ILLITE GROUP CLAY MINERALS
The term 'illite' usually is used in its petrographic sense as a
name for the K-rich, argillaceous component of sedimentary
rocks,
identified by ca.
1
nm spacing for its 001 X-ray
diffraction (XRD) peak. At 30 percent, illite is the second
most abundant mineral component of sedimentary rocks (after
quartz, 34 percent). Defined that way, it is not a mineral, but is
the "illitic fraction" of a rock, which, in addition to the
mineral illite, may contain admixtures of similar minerals,
most often fine-grained micas, glauconite and mixed-layer
illite-smectite. Used in the mineralogical sense, "illite"
identifies a clay structure, which is of 2:1 type, dioctahedral,
non-expanding, aluminous, and contains nonexchangeable K
as the major interlayer cation. The potassium balances a
negative layer charge, generated by Al for Si substitution in the
tetrahedral sheet and (Mg, Fe^"*") for Al in the octahedral sheet
(Figure II). This definition refers both to illite as a separate
mineral (discrete illite), and as a non-expanding component of
mixed-layer illite-smectite (I-S), known as the
illite
funda-
mental particle (Figure II). Due to its complicated chemical
composition and broad range of possible isomorphous
substitutions, illite becomes an important sink for most major
elements liberated at the Earth surface by interaction between
rocks and the hydrosphere. Thin illite crystals contribute
considerably to a rock's cation exchange capacity and they
reduce its permeability. Illite group minerals are covered in
detail in recent reviews by Srodoii (1999a,b).
Origin and evolution of illite in the rock cycle
Common surface environments are not favorable for the
formation of the mineral illite. Reports of illite formation in
weathering or marine environments often are misleading
because of eolian contamination (Sucha et al., 2001). Illite
may be partially stripped of its potassium by weathering and
may fix it back in sedimentary environment without changing
its silicate layers. Experiments indicate that smectites also may
fix potassium at surface temperatures without changing its
silicate layers, thus forming mixed-layer I-S during the process
of alternate wetting and drying, however, the geological
significance of this process has not been assessed. Illite can
be crystallized at surface temperatures at pH > 10 (Bauer etal.,
2000).
In natural high pH environments (lacustrine, playa),
ferruginous illite, having a composition intermediate between
illite and glauconite, is a common product of low-temperature
diagenesis.
Illitization becomes a major chemical reaction in
sedimentary rocks above approximately 70°C. Three types of
processes have been recognized: (1) alteration of smectite into
illite via mixed-layer I-S; (2) direct alteration of kaolinite and
feldspar into discrete illite; (3) crystalization of filamentous
illite in the pore space. The first process, volumetrically
dominant, is known in the most detail. Illitization of smectite
is a continuous process, producing pure non-expandable illite
at the bottom of the anchizone {ca. 275°C—this zone is the
highest grade of burial diagenesis). If not uplifted and eroded,
it gradually recrystallizes into coarse-grained mica up to 325°C
(biotite isograd, McDowell and Elders, 1980).
In the weathering environment, eroded illite is very stable
and commonly becomes recycled into new sediments either
chemically intact or partially stripped of its potassium (soil
vermiculite).
Mixed-layering, polytypes, and structural defects
Illite fundamental particles, which are only a few layers thick,
are very flexible. They can join each other or smectite
monolayers face-to-face to form multi-particle crystals
(Figure II), called mixed-layer crystals because the particle
interfaces display not illite but smectite characteristics (e.g.,
exchangeable cations, swelling properties). In the course of
diagenesis, smectite particles dissolve and illite particles grow
in thickness (Srodori etal., 2000), thus mixed-layering evolves
along with increasing percent of illite layers from so-called RO
(smectite monolayers plus illite bilayers) into Rl (no smectite
monolayers, bilayers dominant) and then R >
1
(3 nm or
thicker particles dominant).
370
ILLITE CROUP CLAY MINERALS
Tetrahedral sheet
Octahedral sheet
Tetrahedral sheet \ Si,
5-^ O
Exchangeable cations plus water
0
Si,Al
O.OH
Al,
Mg,
Fe
O.OH
AI
2:1 silicate
layer
Fixed cations
Reynolds, 1997) and used for tracing geological processes.
For example, the lM(j/2M| ratio is measured to quantify the
ratio of diagenetic to detrital illite in the diagenetic zone, and
the progress toward metamorphism in the anchizone.
Generally, the density of defects decreases during diagenesis,
with trans-vacant layers becoming dominant, but exceptions to
this rule are known.
\_
~7
Illite
fundamental particle
] ] \ Smectite
/ \ [J fundamental particle
^1
X
L
Discrete illite
crystal
Mixed-layer
illite-smectite
crystal
Illite
fundamental particle
Figure II Schematic structure of the 2 :1 silicate layer. The layer has
two tetrahedral and one octahedral sheet and is dioctahedral, i.e., 2
out of 3 octahedral positions are occupied by cations. The lower part
of the figure shows two types of occurrence of illite layers: as discrete
crystals and as fundamental particles in mixed-layer illite-smectite
crystals.
Mixed-layer crystals produce 001 XRD reflections at
positions intermediate between the positions of end-members.
The illite: smectite ratio is measured from these XRD
positions (Moore and Reynolds, 1997) and used for tracing
the increasing grade of diagenesis. Mixed-layering becomes
undeteetable at approximately 20 layers mean thickness.
Because of the pseudohexagonal symmetry of layer surfaces,
subsequent layers in illite particles can be superimposed not
only in the same direction, but also rotated by 60° or 120°. If
such pattern repeats systematically a polytype is created. If the
pattern is random, the structure contains stacking faults
(structural defects). A second type of structural defect
recognized in illites arises from the location of cation vacancies
in the octahedral sheet. Only two of the three sites are occupied
in dioctahedral minerals and the sites are not equivalent (two
"cis"
and one "trans" site). Illite layers can be either trans or
cis vacant and both types are commonly found in one mineral
(Moore and Reynolds, 1997).
Among common illite polytypes, IM is characteristic of
diagenesis and 2M| appears in the anchizone (intermediate
between diagenesis and metamorphism). While 2M| is
typically devoid of structural defects, 1M most often contains
defects, which broaden and weaken its XRD diffraction peaks
(so-called lMj illite). Polytypes and defects can be quantified
by computer modeling of XRD patterns (Moore and
Crystal size, shape, and "crystallinity" (Kubler index)
Illite crystals and fundamental particles are plates terminated
by 001 planes, either equidimensional (the stiape common for
shales) or highly elongated (the shape typical of "hairy illite"
that clogs sandstone porosity). These shapes are controlled by
crystal growth rates (Bauer etal., 2000).
The thickness of illite fundamental particles, first measured
by Nadeau etal. (1984), was studied in detail for bentonites
(Srodori etal., 2000). Illite nucleates as particles 2 layer (2nm)
thick, which grow during the course of diagenesis, producing
lognormal distributions of increasing mean thickness, broad-
ening and skewness. The growth of fundamental particles is
the major factor responsible for a decrease in the XRD 001
peak breadth ("Kubler index") during anchimetamorphism.
Particle thickness distributions can be now measured directly
byXRD(Eberle?a/., 1998).
Layer charge and fixed cations of illite
The typical structural formula for pure illite from bentonites:
(K, Na, NHJo 5o(Al,
corresponds to infinitely thick illite crystals. The thinner the
crystals, the lower their bulk fixed charge, because the number
of interlayers is always smaller by one than the silicate layers
(Figure il). In the extreme case of 2nm illite crystals, the fixed
charge equals 0.45 and there is a substantial number of
exchangeable cations at crystal surfaces.
The charge of illite layers is satisfied mostly by fixed
potassium. Minor Na substitution is well-known but there is
no complete solid solution, and if sodium is abundant, Na-mica
(paragonite) is formed in the anchizone (Frey, 1987). Solid
solution between K and NH4 is possible and NH4-rich illites
(tobelites) have been reported from deep diagenesis (Schroeder
and McLain, 1998) and anchizone (Lindgreen etal., 2000).
Illite as paleogeothermometer
Illite crystal growth has been traced in the diagenetic zone by
measuring the illite: smectite ratio of mixed-layer 1-S and in
the anchizone by measuring the Kubler index. Students of the
anchizone believed that the Kubler index indicates paleotem-
peratures; and anchizone limits were defined using its values
(;Frey, 1987).
In the diagenetic zone, the effect of bulk rock chemistry on
the illite: smectite ratio has been demonstrated and explained
by a deficiency of K in bentonites. Because of abundant K and
relatively closed system, shales are considered the most
suitable lithology for studying illite response to basin
temperature. The approach, introduced by Hoffman and
Hower (1979), assumes that the illitization reaction in shales
is so fast that the time factor can be neglected and that
maximum paleotemperatures can be evaluated from measured
illite:
smectite ratios. Other investigators take into account the
IMBRICATION AND FLOW-ORIENTED CLASTS
371
time factor and chemistry and use various kinetic equations to
model the experimental illitization profiles.
K-Ar dating of illite
The interpretation of K-Ar dates depends on the type of host
rock. In pure beach or eolian sandstones all illite can be
authigenic (hairy illites) and the radiometric date indicates the
time of hot fluid flow and perhaps oil emplacement (Hamilton
etal., 1992). Shales, even in the finest fractions, usually contain
a mixture of authigenic and detrital illite. Techniques for
obtaining both ages have been proposed. If extracted properly,
the K-Ar date of diagenetic illite from shale represents the
mean time that has passed between the onset of illitization and
the time of maximum paleoternperature. This mean strongly
depends on burial history (Srodori et al., 2002). Many
bentonites are free of detrital contamination, but the K-Ar
date represents an even longer period of time than I-S from
surrounding shales because of slow diffusion of potassium into
the bentonite. This period of time can be evaluated by dating
bentonites that are zoned, that is, less highly illitized in the
center of the bed, and/or fractions of fundamental particles
separated from a bentonite.
Jan Srodori
Bibliography
Bauer, A., Velde, B., and Gaupp, R., 2000. Experimental constraints
on illite crystal morphology. Clay Minerals, 35: 587-597.
Eberl, D.D., Nuesch, R., Sucha, V., and Tsipursky, S., 1998.
Measurement of fundamental particle thicknesses by X-ray
diffraction using PVP-10 intercalation. Clays and Clay Minerals,
46:
89-97.
Frey, M., 1987. LowTemperatureMetamorphism. Blackie.
Hamilton, P.J., Giles, M.R., and Ainsworth, P., 1992. K-Ar dating of
illites in Brent Group reservoirs: a regional perspective. Geological
Society Special Publication, 61, pp. 377-340.
Hoffman, J., and Hower, J., 1979. Clay mineral assemblages as low
grade metamorphic geothermometers: application to the thrust
faulted disturbed belt of Montana, USA. SEPM, Special Publica-
tion, 26, pp. 55-79.
Lindgreen, H., Drits, V.A., Sakharov, B.A., Salyn, A.L., Wrang, P.,
and Dainyak, L.G., 2000. Illite-smectite structural changes during
metamorphism in black Cambrian Alum shales from the Baltic
area. American Mineralogist, 85: 1223-1238.
McDowell, S.D., and Elders, W.A., 1980. Authigenic layer silicate
minerals in borehole Elmore 1, Salton Sea geothermal field,
California, USA. Contributions to Mineralogy and Petrology, 74:
293-310.
Moore, D.M., and Reynolds, R.C., 1997. X-Ray Diffiaction and ihe
Identification and Analysis of Clay Minerals. Oxford University
Press.
Nadeau, P.H., Wilson, M.J., McHardy, W.J., and Tait, J., 1984.
Interstratified clay as fundamental particles. Science, 225: 923-925.
Schroeder, P.A., and McLain, A.A., 1998. Illite-smectites and the
influence of burial diagenesis on the geochemical cycling of
nitrogen. Clay Minerals, 33 : 539-546.
Srodori, J., 1999a. Use of clay minerals in reconstructing geological
processes: current advances and some perspectives. ClavMinerals,
34:
27-37.
Srodori, J., 1999b. Nature of mixed-layer clays and mechanisms of
their formation and alteration. Annual Review of Earth and
Planetary Sciences, 27:
19-53.
Srodori, J., Eberl, D.D., and Drits, V., 2000. Evolution of funda-
mental particle-size during illitization of smectite and
implications for reaction mechanism. Clays and Clay Minerals,
48:
446-458.
Srodori, J., Clauer, N., and Eberl, D.D., 2002. Interpretation of K-Ar
dates of illitic clays from sedimentary rocks aided by modelling.
Ameriean Mineralogist, 87: 1528-1535.
Sucha, V., Srodori, J., Clauer, N., Elsass, F., Eberl, D.D., Kraus, I.,
and Madejova, J., 2001. Weathering of smectite and illite-smectite
in Central-European temperate climatic conditions. Clay Minerals,
36:
403-419.
Cross-references
Bentonites and Tonsteins
Clay Mineralogy
Diagenesis
Fabric, Porosity, and Permeability
Glaucony and Verdine
Mixed-Layer Clays
Mudrocks
IMBRICATION AND FLOW-ORIENTED CLASTS
Imbrication is the overlapping arrangement of similar parts, as
of roof tiles or fish scales. The earliest use of "imbricate" or
related words by geologists quoted in the Oxford English Dic-
tionary are by Dana (1852, 1862) and by A. Geike (1858) to
describe biologic or paleontologic objects. An illustration of
sedimentary imbrication is given as early as Jamieson (1860),
reproduced here in Figure 12. A modified version of this figure
appears in the llth and 12th editions of Lyell's (1872, 1875,
p.
342) Principles, but neither Jamieson nor Lyell use any form
of the word "imbricate" in their discussions of the figures.
Imbrication, which is a special case of flow-oriented clasts, is
a common phenomenon, but observations and theory con-
cerning it are published rarely. The flow-orientation of a single
clast develops in many ways. To understand the flow
orientation of clasts in different environments can improve
interpretation of rocks, and lead to a better understanding of
sediment transport.
Imbricated clasts must have unequal lengths for their a, b, c
axes {a longest, c shortest). Usually, imbricated clasts are
approximately tabular in shape, where the a and b axes
significantly exceed the c axis. Roof tiles, or fish scales, or
tabular rock fragments can be imbricated; basketballs, or
grapes, or spherical sand grains cannot. As suggested by
Figure 12, imbricated clasts must be approximately similar in
absolute size: flat pebbles can form an imbricated group, but a
flat pebble cannot imbricate with a flat boulder.
Jamieson (1860, p. 350) suggests that flow orientation, such
as shown in Figure 12, "is best exemplified when the stones are
pretty large, say, from six to twelve inches in length."
However, flow orientation and imbrication occurs in finer
sediments, including sand (Gibbons, 1972), although com-
pared to pebbles or boulders, sand grains have more random
orientations because their size is smaller relative to the
turbulent eddies, and their submerged weight is relatively less
than the drag exerted by such eddies. On the larger scale,
boulders have been imbricated in debris flows and floods
(Blair, 1987, his figure 7C).
Sediment sorting determines the likelihood of imbrication.
A well-sorted collection of flat pebbles commonly contains an
imbricated array sueh as in Figure 12, but a poorly sorted sand
372 IMI^KICATION AND KLOW-ORIENTED CLASTS
Figure 12 An early illuslralion of imbricated and flow-orienteri clasts (lamjeson, 1860] and photos of the same phenomena trom the
ColoraHo Plateau. These illustrations and photos show regularly oriented clasts, where the d-b planes of the clasts dip down toward the
upstream direction. Full scale is 30cm.
with a gravel component will rarely contain imbricated pebble
arrays.
Regularly-oriented clasts
Isolated tabular clasts are easily oriented by flowing water.
Analysis of the fluid mechanics involved in overturning a
tabular clast suggests that there is a critical overturning Froude
number (V\) above which the clasts overturns,
IF;
= V/v/^ (Eq. I)
where V is the mean flow velocity affecting the clasts, g is the
acceleration due to gravity, and c is the thickness (short axis)
of the tabular clasts (Galvin. 1987). The critical overturning
value of F,, above which the tabular clast is oriented by the
flow, is about 3.0 to 3.5. This critical value decreases as the
suspended mud content of the water increases. The analysis
assumes that flow depth exceeds the a axis of the clast. a
condition where the immersed weight of the clast resisting the
overturning is significantly less than its weight in air.
Usually, the acceleration, g, in a Froude number can be
treated as just another constant necessary to make the units
balance, but Pathfinder photos of clasts on the surface
ol"
Mars
suggest that these clasts may be imbricated or flow-oriented.
The lower Martian value of
,<,'
to be used in equation 1 for
explaining this imbrication or clast orientation implies that a
given size clast overturns in flowing water on Mars when the
velocity is only about 62 percent that needed to overturn it on
Earth.
The expected orientation of an imbricated group
ol"
tabular
clasts or an isolated tabular clast is one in uhich the planes
containing the longer ((/. h) axes dip down toward the source
of the (low (Figtirc 12). For a clast in a stream, (his means that
the (/ h plane usually dips toward the upstream direction.
Clasts having this expected dip toward the upstream direction
are regularly oriented clasts.
Reversely-oriented clasts
On aclive foresct beds, there may be exceptions to this
upstream dip (Figure 13, top). On active foresets, isolated
clasts may dip in the down-slope direction parallel to. or more
steeply than, the slope
itself,
if turbulence shed by the clast
scours a hole at the downstream edge of the clast. and
(possibly) if grain flow exerts an overturning moment on the
upslopc edge of the clasl. Clasts whose a-b planes dip in the
downstream direction are reversely oriented clasts. so-called
because the dip is opposite that expected from intuition
developed by observations such as shown in Figure 12.
IMBRICATION
AND
FLOW-ORIENTED CLASTS 373
Figure 13 Reversely oriented clasts. Above: downstre.im
dip of
larger
pchhies
in ,i
s<indy member
of
the Cretaceous Potomac Eorm,5tion,
LorUin Sldlion, Virginia. The horizonlal
is
indit<3teri
by the
central
bubble
ol the
level. Inferred t'low
is
from right
ttj
left. Notice
the
(ollection
of
smaller pebbles
in ihe
wake
of the
large reversely
pebble.
At .i
higher level
in the
same outcTo|) where
the
is more nearly hori^onial, flow-oriented pebbles have
regularly inclined orienlation. Below:
in
Brushy Basin Member
of
Mori son
formation, reversely oriented blocky sandstone on "popcorn"
munlmorillorite shaie, apparently
the
result
ol
scour
on Ihe
downslope edge
of
the clasts during sheet flow runoff.
In arid regions it is common to tind clasts on eroding slopes
whose a b planes dip downslope more steeply than the slope
on vthich they rest. This commonly occurs when volcanic rocks
capping the uplands break up and slide down slopes of fine-
grained volcanic ash. Such an orientation of clasts (flints) in
surficial gravels were reported toward the end of his life by
Charles Darwin from the south of England. Probably, the
cause is runoff by overland sheet flow during rare, heavy
thunderstorms. For sheet llow, the water depth is typically less
than the <-axis, a condition where the effective weight of the
clast on the bottom is nearly that of its weight in air. The sheet
flow scours on the downslope side of the clasts. and the clast
settles into the scoured depression with the result that the a b
plane of the clast dips more steeply than the slope (Figure 13.
bottom). With respect to the presentation shoun in Figure 12,
the clasts in Figure 13 are reversely oriented because they dip
toward the downstream direction of the flow. Such orienta-
tions of the a b planes steeper than the underlying slope may
superficially resemble the condition produced by waves on
beaches.
Finally, a form of reversely oriented clasts develops when
wedge-shaped or pear-shaped clasls are subject (o sheet flow.
The thick end of the clast e.\erts more weight on the underlying
Figure 14 Imhricaled boulders
al
Monument Cove, Acadia Nalional
Park, Maine. Boulders
are
locally derived and abraded iWaag
and
Ogren,
1984). The long axes
of
boulders
dip
into the waves,
usually more steeply than
the
foreshore slope.
Table
II (
lassification
of
imbricated
and
flow-oriented clasts
Class
Flow C"ondition
Regularly inclined
Reversely inclined
Wiivc-orientcd
V'ertiai))v oricoK'tl
Stream flow (clear water)
Debris flow
Delta flow (i'oreset beds)
Sheet flow (erodible
bed)
Sheet flow (pavemciU)
Waves
Longshore
Sie;id\
tl(*\
Fi>3.nto
3.5
F|<3.0
Scour
at
downstream
end of
clasts wiih lilt
due lo
fluid
flow, gravity,
and
grain flow.
Scour
at
downstream
end of
cliisl
uUh
lilt due
to
gravity.
Pivot about Iliick
end of
clasls. Thick
end
fi.ved
by
friclioii
Onshore offshore tlow dominates. Cktsls
dip
loivard
wiives. more steepK than foreshore.
Signiiicani longshore flow aeeompanies waves. Cliisls
dip
into longshore current.
HiL'h concenlralions
of
plale-like clasts.
374
IMPREGNATION
surface. If the weight produces sufficient friction along the
underlying surface, the thin edge of the clast will be moved by
the current, pivoting like a rudder about its post, to point
downstream. Such orientations might be expected on large
clasts after sudden extreme floods such as the floods that
produced the ehaneled scablands on the Columbia Plateau or
jokulhlaup deposits in Iceland.
Wave-oriented clasts
For clasts on beaches, the orientations of a-b planes have
aspects of both regularly and reversely oriented clasts. For a
clast subject to wave action on a beach foreshore, the a-b
plane dips toward the incoming waves (Figure 14). The
incoming wave is the dominant flow, so this seaward dip
may be considered the expected regular orientation. Relative
to the foreshore on which the clast rests, the clast dips more
steeply so that the lower end of the clast is embedded in the
foreshore. It appears that this embedding of the clast at its
lower end is at least in part due to the currents draining
seaward from the swash, in which perspective, the clast is
reversely oriented. When the waves produce a significant
longshore current, especially at the distal ends of spits, the a-b
plane of the clast will reorient to dip toward (into) the
longshore currents, producing a regular orientation from the
perspective of the longshore current.
On low spits during times when waves actively overtop the
spit, the overtopping due to the waves acts as a uniform stream
which orients flat pebbles with their a-b planes dipping into
that current. This results in clasts on the land-facing shore of
spits having a seaward dip to the a-b plane.
Vertically-oriented clasts
Finally, there is the vertical orientation of a-b planes
introduced when exceptionally thin clasts, such as relatively
flat shells or slate, are subject to strong flows, such as surging
waves (Sanderson and Donovan, 1974). This phenomenon is
relatively rare on a worldwide basis but locally common,
particularly where clams are shucked commercially. The origin
of beds of shells with vertically oriented a-b plane appears to
be similar to the origin of the more common phenomenon of
leaf jams in streams flowing through deciduous woodlands.
There, a heavy concentration of large leaves, such as oak
leaves, in flowing water produces a stable orientation with the
leaf planes vertical.
Summary
The relationship of the various factors to flow direction and
slope are summarized in Table II.
Cyril Galvin
Bibliography
Blair, T.C., 1987. Sedimentary processes, vertical stratification
sequences, and geomorphology of the Roaring River Alluvial
Fan, Rocky Mountain National Park, Colorado. Journai of Sedi-
mentary Researcit, 57: 1-18.
Galvin, C, 1987. Imbrication of tabular boulders, tn Proceedings of
lite
National Conferenceon Hydraulic Engineering. American Society
of Civil Engineering, pp. 279-284.
Gibbons, G.S., 1972. Sandstone imbrication study in planar
sections: dispersion, biasses, and measuring methods. Journai of
Sedimentary
Pelroiogy,
42
:
966-972.
Jamieson, T.F., I860. On the drift and rolled gravel of the north of
Scotland. Quarterly Journal of the Geological Society of London,
16:
347-371.
Sanderson, D.J., and Donovan, R.N., 1974. The vertical packing of
shells and stones on some recent beaches. Journal of Sedimentary
Petrology, 44: 680-688.
Waag, C.J., and Ogren, D.E., 1984. Shape evolution and fabric in a
boulder beach, Monument Cove, Maine. Journal of Sedimentary
Petrology, 54: 98-102.
IMPREGNATION
Impregnation is a technique used to obtain high quality
sections of soft or poorly consolidated sediments for preserva-
tion or petrographic analysis. The hardened sediment samples
are also valuable for elemental and microfabric analyzes. It is
also frequently beneficial to impregnate partially consolidated
samples (e.g., diamicts) prior to facing and thin section
preparation. The primary goal of any sediment impregnation
method is to stabilize the material of interest in order to obtain
undisturbed samples. Due to the range of sample character-
istics and the varying needs of researchers, a large number of
impregnation methods have been developed. Workers in the
fields of sedimentology and soil science have made parallel
developments in impregnation techniques and both fields,
often in isolation, have adopted many common techniques.
However, despite the wide use of impregnation techniques in
sedimentology, geologists have not adopted one or more
standard methods for sample preparation.
The main developments in sediment impregnation have
resulted from the availability of new embedding materials. In
particular, the advent of low viscosity epoxy resins has made
impregnating clay-rich and compacted samples possible. Key
criteria for selecting a particular method and embedding media
depend on the sample cohesiveness, texture, and cementation.
In most cases, the most appropriate method will produce an
imbedded sample that is fully stabilized and has undergone
minimal disturbance, particularly shrinkage. Thorough dis-
cussions of various materials and the comparative advantages
of a large number of different techniques can be found in
Bouma (1969) and Murphy (1986).
Main approaches
One of the earliest approaches used to obtain an impregnated
sample from unconsolidated sediment samples was the
production of a tape or glue peel (see Peeis). This method is
most appropriate for unconsolidated sediments typically found
in sections and from split cores. The common approach is to
saturate the sediment surface with diluted, water-soluble glue
that is allowed to dry over several days. In coarse-grained
samples, the glue typically penetrates to a depth of 0.5 cm to
2 cm. When dry, the glue surface can be separated from the
remaining sediment, yielding a permanent sample for
storage, mounting, and analysis. Although suitable for large
exposures and field use, the chiefdisadvantage of water-soluble
glues is that some degree of shrinkage occurs. The use of a