Petrography of Siliciclastic Rocks

Sedimentology for master
Part III
Petrography of siliciclastic rocks
Adrian Immenhauser, summer 2002
„Sandstones are sediments composed mainly of quartz
or silicate-grains and rock fragments with an average
diameter of 0.063-2 mm. They may be lithified by various types of cement.“
Sandstones and mudrocks and their unconsolidated counterparts form 80 to 90% of the sediments on Earth. All inorganic components of sandstones are derived from crystalline rocks,
which form approximately 95% of the earth’s crust. Therefore, the person studying sandstones
must be versed in igneous and metamorphic petrology to some extent. However, as shown in
chapter 4.14, a small number of minerals form most of the volume of a sandstone. Most important amongst them is quartz (chapter 4.15).
The mineralogy of sediments, unlike that of crystalline rocks, does not represent an equilibrium assemblage. In many cases, mineral associations are found in sandstones which neither
in this composition exist in crystalline nor in metamorphic rocks.
In addition to components from crystalline rocks, clastic sediments usually contain large proportions of grains whose source was older sediment. The distinction between “fresh” and “recycled” components is fundamental for geologic reconstructions. As we will learn during this
course, a worn secondary growth on a quartz grain indicates derivation of that grain from
older sediment. Nevertheless, in most cases such evidence will remain scarce (see chapter
4.12 ‘Maturity’).
Some of the recycled components in sandstones may be carbonate rocks or bioclasts from carbonate rocks. However, fossils may also have lived at the site of sandstone deposition and became part of the sediment after their death. In this case these bioclasts do not indicate reworking of older carbonate rocks but rather reflect a specific (e.g., shallow marine) environment of deposition.
Sandstones are lithified by various types of cement. These cements are an important source of
information (cf. chapter 4.9). In most cases, they are reprecipitation products of dissolved
components within the sandstone in other cases they may precipitate from fluids, which originate from outside the sandstone body.
Most attempts to classify sandstones are based on the mineral composition of the sand-sized
grains. Three mineral species or groups of rock fragments are commonly selected as poles of
an equilateral triangle, which then is subdivided in internal domains. The result is chaotic!
Today there are much more than 50 sandstone classifications! Figure 1 shows four of the more
widely used triangles. However, there seems to be a growing tendency to use quartz plus
chert at one pole, feldspar at the second, and unstable rock fragments at the third. In the
framework of this lecture we follow the majority of workers and use the classification systems
shown in figure 1.
The terms graywacke, arkose, orthoquartzite, and lithic are in common use among sandstone petrologists. However, not all users agree on their definition. The definitions proposed
below are from Blatt et al. (1980).
• Graywacke - A field term. Describes gray and “dirty” sandstones with a high content of
mica both as matrix and as foliated metamorphic rock fragments. The mineral composition
can be specified by using terms such as “feldspatic graywacke” (Fig. 1).
• Arkose - (Oriel 1949) Coarse clastic rock containing more than 25% feldspar. Some
arkoses have a clay matrix others not.
• Orthoquartzite - (Krynine 1948) Detrital sediments composed largely (90% or more) of
quartz and chert. The cementing material is nearly always silica or carbonate.
• Lithic - (Gilbert 1955) Sandstones that contain less than 90 to 95 % quartz plus chert and
more rock fragments than feldspar.
If we study sandstones, the basic questions will be:
• What was the facies of this sediment at the time of deposition? and,
• to which degree has diagenesis altered this sediment?
In order to answer this question, we have to understand all post-depositional events and their
effect on the rock, which changed the sediment into what we see today.
• What of the primary rock at the time of deposition has been dissolved, silicified, carbonated, dolomitized, recrystallized? or,
• how much deviates the present day form/shape/roundness etc. of grains from their
original form/shape/roundness?
In most cases we may recognize that the primary habit of these grains has been changed drastically a long time ago and that we probably never will be able to answer this question.
Fig. 1. Some triangles in common use for the classification of sandstones. From Blatt
et al. (1980; p. 371).
4.1 Internal sedimentary structures
Sedimentary structures are difficult to recognize in thin-sections. This mainly because the
field of observation is so small. Most difficult to recognize are sedimentary structures in very
fine-grained sandstones or siltstones. Among a large number of internal sedimentary structures (see e.g., Demicco and Hardie 1994 for a detailed documentation) the following may be
of importance in this course:
• Parallel lamination - Not all types of lamination have the same mode of origin. Laminae
in mud may be formed by periodic changes in the physical or chemical conditions of deposition. Laminae in sand may be produced rapidly and form under various conditions but is
commonly related to changes in the paleo-hydrodynamic regime.
• Cross bedding - indicates uni-, or bi-directional water flow either due to waves or due to
• Grading - decrease of grain-size commonly indicates decrease in current velocity.
• Bioturbation - burrowing of the sediment by endobiotic organisms, this leads to a blurring
of primary structures. Depending on the degree of bioturbation, the sediment may become
totally homogenized.
• Imbrication - orientation of pebbles in water currents such as shingles or tiles on a roof.
The dip of components is towards the current direction.
• Geopetal markers - partially infilled pore space in sedimentary rocks, may be used as topbottom markers if studied statistically in sufficient number.
• Porosity – term applied to all the openings or interstices within the framework of a sediment. Porosity is normally expressed as percentage of the bulk volume of the material occupied by pore space. Different types of porosity are summarized in figure 2.
4.2 Color
The color of a hand specimen may be rather different from the color of a thin-section. For example, mature quartz sandstones with a calcareous matrix are whitish as hand-specimen but
clear in thin-sections. Immature sandstones appearing reddish in hand specimen (orthoclase)
or greenish (basic volcanics) are indicative for a number of depositional environments such as
tropical or cold climate and are characteristic for rapid erosion, transport and deposition. The
most common colors are due to secondary iron oxides, respectively organic matter.
black - organic matter FeS
yellowish - pyrite, markasite FeS2
yellowish or bluish - sulphates, carbonates and chlorides
yellowish to brownish - limonite, goethite FeO(OH)
reddish - iron oxides FeO
Black, yellowish and greenish-bluish colors are stable under reducing conditions. Black color
in paleosols may be related to poor destruction of organic matter by bacteria, palustrine
(swamps) environments or moderately humid climate. Black color in the marine environment
indicates poor destruction of organic matter due to rapid sedimentation rates or an anoxic environment (black shales). Greenish colors are often the product of decaying brown limonite in
the marine environment. These colors are not stable under subaerial conditions. Dark organic
matter becomes light brown under oxidizing conditions; iron-rich minerals turn into yellowish
or brownish limonite (weathering of surfaces in hand specimen).
Yellowish, brownish or red colors are common under oxidizing conditions (subaerial environment) if iron turns into iron oxides (yellowish to brownish limonite and goethite) and hydrated iron oxides (reddish hematite).
Hematite turns into limonite when transported in rivers. Limonite turns into greenish ironsalts in seawater but may also become hydrated and then turns into red hematite. Hematite is
common in red soils of warm humid climates (red podsolitic soils).
Violet (Mn) colors are indicative for red sandstones, which underwent slight metamorphic
Light greenish colors may be due to glauconite. Dull greenish colors are common in slightly
metamorphic sediments containing chlorite.
4.3 Grain-size
There are four main reasons for grain-size analysis. These are:
• The grain-size is a basic descriptive measure of the sediment,
• grain-size distribution may be characteristic of sediment deposited in certain environments,
• grain-size distributions may yield some information about the physical mechanisms acting
during transport and deposition,
• grain-size may be related to permeability, variations in these properties may be predicted
from grain-size.
The basic concept of determine the size of components in thin-sections is a linear dimension.
The volume of the particle is probably a better approximation to “size” but impossible to obtain in two-dimensions.
The Udden-Wentworth grade scale (Udden 1898; Wentworth 1922; Fig. 3) is commonly
used, but others such as the Atterberg-scale are also found. Krumbein (1934) has transformed
the Udden-Wentworth grade scale into a logarithmic scale named (phi) φ-scale, where d is
the diameter in mm. The φ-equivalents of the Udden-Wentworth grade scale are shown in
figure 4. One millimeter is the equivalent to zero φ.
Φ = -log2 x d (d=diameter)
Fig. 2. Summary diagram of porosity fabrics in sedimentary rocks (based on various
sources, from Tucker (1988, p. 153).
Fig. 3. Various size grade scales in common use. After Blatt et al. (1980; p. 57)
Grain-size in thin-sections - Ancient sedimentary deposits that cannot be readily disintegrated
present major difficulties to the potential size-analysts. In thin-sections only the maximum
size observed corresponds to the true size of the parent population. The probability that a
grain will be cut by a random section increases proportionally with the size of the grain. Size
distributions in thin-sections are thus number frequencies already weighted by the size of the
grains. If grain-sizes are point-counted, the frequencies are already weighted by the volume of
the grains.
The problem to determine the grain sizes in thin-sections is further complicated by the fact
that most particles show anisotropic dimensional fabrics. Therefore, the result of measurements will depend on
• the direction in which the section is cut relative to the grain fabric, and
• the degree of anisotropy of the fabric.
Limiting factors - The thickness of thin-sections is approximately 30 µ (0.03 mm). It was
shown that measuring grains of < 30 µ (fine silt and clay) produces unreliable results (Chayes
1956). Within the framework of this course, we will therefore treat all grains smaller than <
30 µ as “matrix”. The following diagenetic and other effects may markedly change the grainsize distribution of a parent-population:
• Detritic clay may recrystallize to fine-grained mica during diagenesis.
• The recrystallization of unstable phases leads to an apparent increase of the clay content in
fine-grained sandstone. Secondary porosity and ghost structures are indicative.
• Clay matrix between the grain-sized components of a sandstone may recrystallize and
“chlorite” or “sericite” may grow instead. The margins of quartz grains become locally dissolved during this process, leading to a decrease in grain-size. The terms “chlorite” and
“sericite” are purely descriptive in this context. “Chlorite” indicates fine-grained, greenish
mica, whereas “sericite” stands for clear to yellowish micas.
• Carbonate matrix may dissolve the surface of detritic quartz; a quartz-sandstone may become entirely carbonatic during diagenesis. Small fissures in quartz grains may become
filled with authigenic carbonates, which separate the fragments during further growth. It
then may become difficult to recognize the original grain size.
• Pressure solution may decrease the average grain-size distribution drastically.
• Tectonic overprinting in fault zones may result in fracturing of components and, thus, decrease of grain-size. Similarly may components become elongated, flattened etc.
• Authigenic overgrowth, e.g. on quartz grains, may result in an increase of the previous
Grain-size and sedimentary environment - The size of a largest component in a sandstone
provides us with some information about the site of deposition.
• Aeolian (wind-transported) sandstones are commonly built by grains of < 200µ (0.2mm) in
average grain size.
• A denser medium, such as ice or mud-rich water may transport blocks of up to tens of
meters in size.
• In the case of marine, fluvial or lacustrine sandstones it is useful to compare the largest
fraction of detrital components with the largest non-detrital components such as bioclasts, mud-pebbles etc.
Several attempts were made to classify grain-size and to relate the various groups to specific
environmental conditions (Friedman 1961; Passega 1964). None of the methods suggested is
an infallible guide to discrimination of environments. Determination must be based on sieve
analyses of several samples and the size distributions must not have been modified by diagenesis. This approach, however, is not possible in the case of thin-sections.
4.4 Form
The definition of roundness and spericity of a particle is insufficient to define the shape even
for ideal ellipsoids. Four main shape classes were defined by Zingg (1935; Fig. 5):
oblate (tabular or disk shaped)
prolate (rod shaped)
Equant particles have the highest spericity but oblate, bladed and prolate particles may all
have the same spericity values. The different shapes of grains influence their behavior as
components in sediment. For example, prolate particles might roll much easier than oblate or
bladed particles.
4.5 Roundness
The most commonly used method of estimating roundness is the visual comparison of grains
with the standard images of Powers (1953; Fig. 6). The roundness classes are given descriptive names reaching from angular to well-rounded. Folk (1955) suggested a logarithmic of
roundness analogues to what he called (rho) ρ-scale. Rho-values extend from zero (perfectly
angular) to six (perfectly rounded), classes corresponding to Powers (1953; Fig. 6). However,
it has been shown that the accuracy of the visual comparison method is low. Different persons
may estimate different roundness classes that differ by a whole class. More sophisticated
methods involve Fourier analysis and have considerable potential but are not widely applied
Roundness as environmental parameter - Roundness is, after all, the result of the abrasion of
irregularities on the grain surface due to grain-grain or grain-ground collisions. It is rather
unusual that individual grains will brake during such collisions unless they brake along
schistosity plains. A few very general rules are the following:
• Larger components are rounded more rapidly than smaller ones since their weight is higher.
Fig. 4. φ (phi)-scale after Krumbein (1934) and corresponding sizes in mm.
Fig. 5. The four shape classes after Zingg (1935; p. 80).
Fig. 6. Roundness classes, redrawn from Powers (1953).
Comparably soft lithologies (e.g., limestones) become more rapidly rounded than hard ones
(e.g., granites).
• Angular limestone pebbles are well rounded after a transport distance of roughly 15-35 km
within a river; quartzite pebbles take about 70 km to reach the same degree of roundness
(Plumley in Pettijohn 1975). Basalt pebbles along the beach of the recently emerged volcanic island Surtsey (near Iceland) take only few months to become well rounded.
• Abrasion of sand grains in rivers is negligible (Kuenen 1964), somewhat more important at
beaches but mainly found in the aeolian regime.
Attention! The study of etched grains may result in totally misleading directions and should,
therefore, be avoided.
Structural Inversion - The co-occurrence of e.g., well rounded quartz grains in a clay matrix.
Structural inversion is important for the interpretation of a sandstone because it indicates
commingling of two or more levels level of hydro-/aero-dynamic energy or several rock
sources. Well-sorted and well-rounded grains may be blown from a coastal dune in a lagoon
located behind the dunes. In the lagoon, the grains become part of the fine-grained sediments
deposited in this environment. The formation of the well-rounded grains took place in an environment of relatively high aerodynamic energy but deposition occurred under low hydrodynamic conditions. We therefore judge sediments with structural inversion after the environment of lowest energy. Examples of structural inversion are shown in figure 7.
4.6 Surface texture
In most cases, the scanning microscope is used to study the texture of grain-surfaces of components larger than 0.2 mm in size. The shape, orientation and number of pits in grain surfaces may remain somewhat speculative but it was suggested that certain combinations of
textures are characteristic for specific sedimentary environments. For example, v-shaped pits
may be produced by chemical etching or grain-grain-collisions.
• Impacts produce abundant, poorly oriented pits as well as small grooves.
• Solution and re-deposition produce smooth or irregular surfaces.
Unfortunately, in many ancient sandstones diagenesis has destroyed original surface textures.
4.7 Spericity
Spericity is a concept developed by Wadell (1932). It measures the degree to which a particle
approaches a spherical shape. However, the spericity does not exactly express the dynamic
behavior of the particle in a fluid. Therefore, Sneed (1958) defined the maximum projection
spericity as the ratio between the maximum projection area of a sphere with the same volume
as the particle and the maximum projection area of the particle.
Fig. 7. Examples of structural inversion, after Folk (1974).
4.8 Sorting
How to measure the sorting? Pettijohn (1975) suggested a rough approximation of sorting
classes (Fig. 8). These are:
• ≤ 3 φ-classes (see Fig. 4) - well sorted (largest diameter of the largest component is up to 8
times the diameter of the smallest grain)
• 4-6 φ -classes - moderately sorted (largest diameter of the largest component is up to 64
times the diameter of the smallest grain)
• ≥ 7 φ -classes - poorly sorted (largest diameter of the largest component is more than 64
times the largest diameter of the smallest grain)
The meaning of sorting
• The sorting of the source rock/sediments. Glacial deposits, when eroded e.g. along a coast,
provide poorly sorted beach deposits. A river, eroding into well-sorted Tertiary sands will
deposit well-sorted fluvial deposits.
• hydrodynamic regimes - Strong fluctuations in the hydrodynamic regime of a river will
cause relatively poor sorting. A constant current of intermediate velocity will deposit wellsorted sands.
• The type of transport medium (wind, water, ice, etc.)
• The amount of detritus transported to the site of deposition. Poorly sorted sediments are
e.g., typical for beaches near cliffs where rivers and the erosional potential of waves produce very high amounts of sediments per time unit. Coasts with little relief where little
sediment is delivered are typified by well-sorted sandstones.
4.9 Cements in sandstones
Well sorted sandstones with a low clay content have commonly a large porosity and permeability. Groundwater currents circulate and cements can precipitate from the groundwater.
The most common cement phases in sandstones are silica, carbonate and iron oxides.
• Silica cements appear in sandstones in many morphologic and crystalline forms. Quartz
precipitated during diagenesis may occur as microcrystalline blocks (chert), as coarse
crystalline mosaic of sub-equant blocks, as elongate fibbers (chalcedony), as euhedral secondary growths on detrital quartz grains, as micro- or macro-quartz replacement of carbonate minerals (either fossils or matrix), and, less commonly, as replacement of glauconite
and the materials. In thin sections, the boundary between host grain and overgrowth is typically marked by impurities that form an incomplete coating on the detrital grain. If grains
were “scrubbed clean”, e.g., by groundwater, prior to overgrowth precipitation, no hostovergrowth boundary may be visible.
• Carbonate cements are generally chemical precipitates therefore crystalline and transparent. Large calcite crystals may overgrow several quartz grains (poikilitic structure). Small
amounts of iron oxides (limonite) in the carbonate cements are responsible for a slightly
brownish to yellowish color. Carbonate cements may dissolve locally the quartz grains.
Carbonate cements are uncommon in Holocene sandstones. This because silica sands are
deposited in a rather high hydrodynamic regime where soft carbonate that nucleates on
grain surfaces is rapidly abraded. In ancient sandstones, low-magnesium calcite is the
dominant carbonate cement. The calcite occurs as mosaic of anhedral crystals 20 µ or more
in diameter. In most sandstones the volume of carbonate cement does not exceed about
30%, the original porosity of the sediment. In some cases, where the percentage is much
higher, the carbonate replaces the detrital grains. The quartz particles are etched into irregularly shaped fragile grains that could not survive transport and thus are not detrital. It is
commonly observed in thin-sections that young cement phases replace older calcite cement
phases in sandstones.
• Siderite cement. Siderite (iron calcite FeCO3) might form as a replacement of pre-existing
calcite cement. Calcite will be dissolved and replaced by siderite if it is in contact with a
solution containing ferrous iron ions in a concentration more than 1/20 that of calcium
ions. Conversely, there will be desideritization by calcite if the pore solution has more than
20 times as much calcium ions as ferrous ion.
• Hematite cement. Red pigmented matter in sandstones is present as a primary pore filling
precipitated from solution, as microcrystalline particulate iron oxide deposited with the
clastic grains, or as iron ions absorbed on clay mineral surfaces. The sources of the iron atoms in this pigment are the common iron-bearing accessory minerals of igneous or metamorphic rocks, such as hornblende, chlorite, biotite, ilmenite and magnetite. Magnetite
may average iron oxide in silica rocks of more than 95%. Although red coloration can start
to develop penecontemporaneously with sediment deposition, it can continue until all ironbearing detrital minerals have lost their iron, perhaps for tens or hundreds of millions of
4.10 Fabric of sandstones
Fabric means „ the manner of mutual arrangement in space of the components of a rock body
and of the boundary between these components“ (International Tectonics Dictionary). It thus
includes the packing and the orientation of grains. Grain packing strongly affects the porosity and permeability whereas grain orientation affects the permeability.
Packing is the „spacing or density pattern of mineral grains in a rock“ (AGI Glossary). Even
under the assumption of a sediment composed of perfect spheres with an uniform size, six different systematic ways of arranging the spheres are possible. The loosest arrangement is the
‘cubic packing’ with a porosity of 47.6%, the tightest (‘rhomboedral’) packing allows porosity of 26.0%. However, in nature, an infinite number of combination of these six and random packing are developed. The type of contact between grains can also be studied in thin
section. In the ideal case of packed spheres, the only observed contacts would be tangential
ones. However, the four possible types of contacts are (Taylor 1950):
long (straight line in the plane of section)
The frequency of concave-convex and sutured grains relative to the two other types has been
used as a measurement of the intensity of compaction. The four types are illustrated in figure
4.11 Compaction
Sands compact much less than mudrocks. The lesser compaction of sands is due to the following two reasons: First, sandstones are composed largely of quartz grains, quartz grains are
undeformable under most sedimentary conditions.
Second, mudrocks yield initially a high content of water and this water is quickly expelled under pressure. Quartz sands compacted under laboratory conditions have only shown a decrease
of the aggregate thickness of 10 - 15% due to rearrangement of grains and chipping of grain
Compaction in sandstones takes place in two ways:
• In un-cemented sediments at grain-grain contacts - Once the grains within an uncemented
sediment have reached their densest configuration by slippage on grain surfaces, grain reorientation and fracture of radical grains, overburden is transformed through grain-tograin-contacts (Fig. 10a). Continuous stress causes dissolution of the contact area (Fig.
10b). During solution compaction, the area immediately adjacent to the stressed grain to
grain contact is taken into solution and may be flushed away from the immediate area of
pore fluid migration. These fluids are considered the source fluid of cements within the
subsurface, either quartz or calcite cements (Sibley and Blatt 1976). The dissolved quartz
or calcite may be precipitated locally or carried for some distance in solution (e.g., Moore
1985). As cementation proceeds in clean sandstones or grainstones, pressure is transmitted
through the entire rock body and there is no excess of stress at the grain to grain contacts.
Solution compaction between individual grains therefore ceases, but continues over a much
wider area along stylolites, seam and sutures within the sediment (Fig. 10c, d, e).
• In cemented sediments along irregular surfaces (stylolites) - Stylolites and solution seams
are present within many sediment types (e.g. Gillett 1983 and references therein). Both
stylolites and solution seams transect the cemented sediment and develop perpendicular to
the axis of maximum stress such as overburden pressure or tectonic stress. Because stylolites and solution seams are zones of preferential solution within sediment, small fragments
of less soluble minerals tend to accumulate along the solution seam (Fig. 10e).
Fig. 8. Sorting images and classes after Anstey and Chase (1974). Sehr gut = very well
sorted; gut = well sorted; Mässig = moderately sorted; schlecht = poorly sorted.
Fig. 9. Definition of grain contact types and packing proximity.
(After Taylor 1950 and Kahn 1956).
Fig. 10. Solution compaction between individual grains (porosity stippled). After Tucker
& Bathurst (1990). a. Point grain to grain contact; b. stressed grain to grain contact,
leading to dislocation in grain lattice and subsequent dissolution, with later fluid transport of solutes; c. planar grain to grain contact; d. interpenetrating grain to grain contacts;
e. sutural grain to grain contacts.
Criteria for minor compaction (loose packing) in thin sections
• The primary porosity in sandstones (including cement-filled pore-space) is in the order of
25-50 vol.-% (see ‘packing’). Higher porosity values indicate dissolution of components or
enlargement of the primary distances between components due to cement growth (Taylor
• The average number of contacts in a sandstone was found to be in the order of 1.51.6/component. A higher number of contacts indicate compaction a lower number of contacts (combined with the presence of clay) indicate that some of the components may have
decayed into clays (Pettijohn 1975).
• The presence of un-deformed components - such as clay flakes, calcareous fossils or
plant remaining - which are highly sensible to deformation, indicates minor compaction.
Criteria for increased compaction (dense packing) in thin-sections
• Porosity (including cement-filled pore space) of less than 25-50 vol.-%. Sandstones buried
in depths of about 4 to 7 km have a mean porosity of only 10-15 % (Maxwell 1960).
• Mean numbers of grain-grain contacts larger than 1.5-1.5. The number of grain contacts is
about 4 to 5 in 4 to 5 km depth (Pettijohn 1975).
• High numbers of long, concave-convex and sutured grain boundaries. Sutured contacts
form below 1.5 to 2 km burial depth (Fig. 9).
• Presence of deformed, broken or fragmented mica, feldspar etc.
Relative ages of packing and cementation
Relative ages of packing and various cement generations may be deduced by carefully studying grain-grain, grain-cement and cement-cement contacts. Some examples are shown in figures 11 and 12.
4.12 Maturity
The maturity of a sandstones depends on how many cycles of erosion and redeposition its
components have “seen”. Pettijohn (1957) distinguished so called ‘compositional maturity’
and ‘structural’ respectively ‘textural maturity’ (Fig. 13).
A sandstone that consists exclusively of well rounded and well sorted quartz and chert grains
is very mature (e.g., a lithified eolianite), both in terms of composition (only quartz and
chert) and of structure/texture (well rounded and sorted). A sandstone that consists of angular,
Fig. 11. Basic styles of solution compaction (after various sources, in Tucker 1988, p. 138).
Fig. 12. Two examples of relative ages of compression and cement formation.
Modified after Roep (1994).
Fig. 13. Maturity classes (In Scholle 1979, p. 103, after Folk 1974).
poorly sorted quartz and chert grains and a fair amount of fresh feldspars or rock fragments is
Compositional maturity = quartz + chert / feldspars + rock fragments
The compositional maturity increases with decreasing amount of weathering sensitive components (feldspars and rock fragments). In other words the less feldspars and rock fragments a
sandstone comprises the higher is its maturity. Feldspars and rock fragments cannot undergo
many erosion-redeposition cycles because they are relatively unstable, disintegrate or alter
into clays during diagenesis.
Structural maturity = sorting and roundness of sand-sized grains
The structural/textural maturity increases with the increasing degree of roundness and sorting.
The structural/textural maturity of a sandstone increases with every cycle of erosionredeposition. The maturity may therefore be a measurement on how many depositional cycles
components have gone through. Attention: the composition of the source rock is also important and may lead to misinterpretation of the maturity if, e.g., no feldspars or rock fragments
are present.
4.13 Diagenesis, and changes in texture and mineralogy in sandstones with increasing burial depth and temperature
Temperatures at which most chemical reactions occur on Earth have a range of not more than
40°C. In contrast burial temperatures range between 0°C and 200°C and then grade into what
is called “metamorphism”. The precipitation of solid substances - under subsurface conditions
- within the pore space of sands turns loose sand into a more or less hard rock.
Table I shows some of the chemical/physical reactions occurring in sandstones under increasing burial depth and increasing temperature. Table I is based on Blatt et al. (1980, p.
Changes in detrital minerals
Other changes
K-feldspar sericDissolution
Compaction of all sands;
itization and kao- and crystalli- precipitation in pores of
linitization; prezation of
cements opal, aragonite,
cipitation of
calcite, quartz, hematite,
and some clay; also replacement of one cement by
Coarsening of
crystal-size in
pressure solution of
Feldspars -> albitization and dissolution; continued alteration of
Alteration of
mafic volcanic fragments to clay,
zeolites, and
Fabric deformation in sands
containing abundant foliated lithic fragments
Pressure solution
of quartz
Feldspar ->
Mafic fragments largely
destroyed, chloritization of
Growth of
micas and
Development of foliation
and a largely recrystallization fabric
Recrystallization of polycrystalline
Destruction of
potassic feldspars; growth of
abundant albite
4.14 Terrigenous components
The presence of terrigenous components in sandstone depends basically on three factors:
• Availability - A limestone source rock delivers no crystalline fragments; a granite delivers
no chert pebbles; the absence of feldspar in a sandstone may e.g. indicate humid climates in
the hinterland.
• Mechanical durability - Depends on the cleavabilty and the physical hardness of the components
• Chemical stability - Minerals which precipitate late in magma chambers, i.e. under cooler
conditions and the presence of water, are chemically more stable than early precipitates.
The relative stability of minerals being subjected to weathering (Folk 1974) is shown below. Olivine is the least stable whereas quartz is the most stable.
The same accounts for detrital heavy minerals of which some are more stable towards weathering than others (modified after Folk 1974).
Moderately stable
Very unstable
Rutile, zircone, tourmaline, anatase
Apatite, garnet (iron-poor), biotite, magnetite, and others
Epidote, kyanite, garnet (iron-rich), and others
Hornblende, augite, and others
Some of the most common heavy minerals are shown in figure 14.
Fig. 14. Common heavy minerals. Degree of weathering increases for each mineral to the
right. After Grimm (1973).
Frequency of terrigenous components
Terrigenous components, i.e. products of the erosion of crystalline rocks or sediments with
crystalline components, form about 60% of the total rock volume of sandstones. Chemical
precipitates form about 20-40%. The most common rock type in continental lithosphere is
granite, formed predominantly by quartz, feldspar and mica. Oceanic lithosphere consists
mainly of basalt; however, erosion of oceanic lithosphere is less common. The most important
terrigenous components are therefore quartz, feldspar (possibly decayed to clay) and mica.
Other components (e.g., heavy minerals) form commonly below 1 vol.-% of sandstones. Visual estimation of frequencies in thin-sections is made by the use of diagrams such as shown in
figure 15.
Folk (1974, p. 66) has summarized estimated values of the terrigenous and the chemical components in sandstones (table II).
Terrigenous components
Clay minerals, commonly the products of
unstable minerals
Metamorphic minerals and rock-fragments
(phyllites, schist’s, metaquartzites etc.)
0.2-1 %
Chert (microcrystalline qtz)
Heavy minerals
Chemical components
Sulfates, gypsum, anhydrite etc.
Others: feldspar, hematite, limonite,
pyrite, glauconite, tourmaline, zircone
Table III at the end of this handout summarizes minerals of various groups, their chemical
composition, color, cleavage, relief, birefringence as well as other features. Table III is modified after Blatt (1980).
4.15 Quartz
Amongst all components of sandstones, quartz is the most important. The average sandstone
contains 65% quartz, the average shale 30%, the average limestone 5%, and the average present-day pelagic deposits 10% (Blatt et al. 1980). Because of its importance, the various types
of quartz are shortly described below. Nearly all quartz-bearing crystalline rocks are massive
plutonic granitoid rocks, gneisses, or schists (Fig. 16).
• Quartz from plutonites - Granitoid rocks disintegrate to more or less equal amounts of
polycrystalline and monocrystalline quartz grains. Monocrystalline grains are in average 1φ
= 0.5mm. Polycrystalline grains are larger than monocrystalline grains. An average 80 to
90% of the crystals have been deformed and show undulatory extinction.
• Quartz from gneisses - The average gneiss yields 20 to 25% mono- and 75 to 80% polycrystalline quartz. The size of the monocrystalline grains is much smaller than in plutonites
(in average 2.2φ = 0.2mm). As in granites, nearly all crystals show undulatory extinction.
The crystals are frequently elongated. The quartz crystals have a bimodal size distribution
(the smaller ones have grown later).
• Quartz from schists - Quartz from schists have some characteristics of granitic quartz and
some of gneissic quartz. Schists have 40% mono- and 60% polycrystalline quartz. Moncrystalline grains average 2φ = 0.25 mm in size. Such as in gneisses grains are often elongated/flattened.
• Quartz from volcanics and subvolcanics - Quartz from volcanic rocks is usually monocrystalline, has non-undulatory extinction and is clear, i.e. has no inclusions. Some grains
include glass.
Quartz from hydrothermal rocks - Quartz from hydrothermal rocks may be mono- or
polycrystalline and is generally characterized by its milky color due to a large content of
water-filled inclusions.
Fig. 15. Diagrams for visual estimation of percentages of various components in rock
sections. After Shvetsov (1954).
Fig. 16. Summary diagram of quartz grain fabrics for use as provenance indicators (after
Tucker, 1988, p. 124).
Kaolinite Al2Si2O5(OH)4
Orthoclase KAlSi3O8
Color- Planar
KAl2(AlSi3)O10(OH less
Fe2+)3(Al, Brown- Planar
Fe ) Si3O10(OH,F)2 green
Ab: NaAlSi3O8
An: CaAl2Si2O8
Moder- Gray
Fine grained
Clay minerals
Acicular chert (chalcedony),
megaquartz and microquartz;
all types diagenetic
Detrital grains and cement
Form and occurrence
Common as
burial cement
Detrital minerals, alteration
products of silicates and as
Common detrital mineral
Common detrital mineral
Common detr. material
May form over- Present as detrital mineral,
growth cements commonly of cloudy appearance due to alteration in clay
Common detr. material
Parallel extinc- no
Parallel extinc- no
tion, pleochroic
Multiple twins
Simple twins
Overgrowth cement
Chalcedony and
megaquartz occlude voids
Birefringence Other features Cement
Moder- Bright colors
Moder- Bright colors
Ca(Fe, Mg, Mn)
Anhydrite CaSO4
Dolomite CaMg(CO3)2
Very high colors
Very high colors
Very high colors
Rectilinear Moder- Bright colors
Color- Rhombic
Color- Rhombic
h or colorless
CaCO3 (rhomb)
Moder- Gray
Rectilinear Moder- High colors
Color- Planar
Color- Planar
Color- Planar
Aragonite CaCO3 (orth.)
by staining
Commonly replaces pellets
Low sedimentation rates
Commonly replacement mineral in ironstones
Detrital grains, cements, fossils, replacement fabrics in
Common burial Commonly crystalline due to
replacement of evaporite sequences
Many cement
Common as ce- Associated with volcanogenic
ments near vol- sediments
Chamosite (Mg, Fe2+)3
Berthierin Fe3+3(AlSi3)O10
Magnetite FeFe2O4
Hematite Fe2O3
Opaque no
Opaque, no
Green- no
Opaque no
Iron Oxides
Gray, masked by
In ooids of ironstones and as
pore fills
Distinguished in Common authigenic minerals
reflected light
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