7 Siliciclastic sediments II: Mudrocks 7.1 Introduction

Mudrocks are the most abundant type of sedimentary rocks, constituting about 45-55% of the sedimentary sequences. But since they are easily weathered, they are covered with vegetation and are poorly exposed.

Mudrocks can be deposited in all sedimentary environments, but the majority are deposited in river floodplains, lakes, large deltas, the more distal areas of clastic shelves, basin slopes, and deep sea floors. The main constituents of mudrocks are the clay minerals and silt-grade quartz.

In terms of grain size, clay refers to particles less than 4 µm in diameter, whereas silt is between 4 and 62 µm. On the other hand, clay as a mineral is a hydrous aluminosilicate with a specific sheet structure (to be discussed below). The term mud (also lutite) loosely refers to a mixture of clay- and silt-grade material.

The mudstone, the indurated or lithified equivalent of mud, is a blocky, non-fissile rock, whereas shale is usually laminated and fissile (fissility is the ability of rock to split into thin sheets). Argillite is used for more indurated mudrock and slate is the metamorphic mudrock possessing a cleavage. Claystone is the sedimentary rock consisting of a clay- grade material. Siltstone consists of more silt-grade particles than clay. Calcareous mudrocks are called marls. Fig. 7.1 gives a scheme for mixed sand-silt-clay deposits.

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Fig. 7.1: Classification of siliciclastic sediments based on sand, silt and clay content.

In the field, the terms mudstone, shale, claystone and siltstone are best qualified by attributes referring to color, degree of fissility, sedimentary structures, and mineral, organic and fossil content (Tab. 7.1).

Tab. 7.1: Features to note in the description of mudrocks

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7.2 Textures and structures of mudrocks The particle size on unconsolidated muds can be measured using sedimentation chamber or settling tube, whereas lithified mudrocks can be studied by the scanning electron microscope.

The use of grain-size data in interpretation of the depositional environments is complicated due to the fact that clay minerals are deposited as floccules and aggregates, and that sediment feeding organisms may generate pellets of muds. Also bioturbation disrupts the depositional textures of mud. However, the following are the important textures and structures of mudrocks.

Preferred orientation of clay minerals and mica flakes parallel to bedding is a common texture of mudrocks. It is the result of deposition of clay flakes parallel to bedding, and less commonly due to compaction and dewatering.

Related to preferred orientation is fissility of shale (Fig. 7.2).

Fig. 7.2: a- Shale characterized by fissility (left); b- Massive non-laminated mudstone (right).

3 The origin of fissility is mainly due to alignment of clay minerals as a result of compaction, in addition to the presence of lamination. Absence of fissility, as in mudrocks, can be explained by bioturbation, the presence of much quartz silt or calcite, and the flocculation of clays during sedimentation that produces random fabric which could be retained on compaction.

The most common sedimentary structure of mudrocks is lamination (Fig. 7.3).

Fig. 7.3: Lamination in mudrock: rhythmites of glacial verves, where graded silt passes upward into clay-grade material

Lamination is the result of variation of grain size and/or changes in composition. Size- graded laminae may be deposited from low-density turbidity currents followed by deposition from suspension currents in relatively short periods of time (hours or days). Other laminae develop over longer periods of times (months or years) if there is a seasonal or annual fluctuation in sediment supply and/or biological activity. Organic laminae in mudrocks, for example, may be produced by seasonal microbial blooms. Also the varved couplets of glacial lakes are taken to reflect the annual spring melting.

Small-scale current ripples occur in siltstones and give rise to cross-lamination (Fig. 7.4). Symmetrical wave-formed ripples also can form in siltstones.

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Fig. 7.4: Photograph showing ripple cross-lamination in siltstone indicating a flow from right to left; the ripple structure is picked out by alternation of dark clay-rich and pale clay-poor laminae.

In tidal flats, mud and fine-sand to silt are deposited alternatively through fluctuating current regimes and/or sediment supply giving rise to flaser (ripple-shaped fine sand-silt occur in mudstone), and lenticular bedding (ripple-shaped mud occurs within fine sand- silt) (Fig. 7.5).

Fig. 7.5: Flaser-lenticular bedding in Ordovician Dubaydib Formation of Jordan.

5 Small-scale scour and fill structures, and micro-cross-lamiantion may be present in siltstones.

Mudrocks lacking any internal sedimentary structures are called massive. This massive nature is due to deposition from high viscosity currents as mudflows and debris flows; or due to bioturbation that destroys any original structures as lamination, mass sediment movement (sliding), dewatering, soil processes (pedogenesis) and root growth.

The structures in muds and mudrocks can be studied using X-ray radiography (photographing using X-ray) which may reveal lamination or bioturbation in an otherwise massive-looking rock.

Other sedimentary structures that may occur in mudrocks include: erosional structures cut in muds and preserved on soles of overlying sandstones (grooves and flutes); slump structures; desiccation cracks formed through subaerial exposure; rain-spot prints and biogenic structures.

7.2.1 Nodules and concretions Many mudrocks contain nodules, also called concretions (Fig. 7.6).

Fig. 7.6: Calcareous nodules in red mudrock.

These nodules are regular to irregular, spherical, ellipsoidal to flattened bodies, commonly composed of calcite, siderite, pyrite, chert, calcium phosphate, together with some original sediment.

Nodules may grow within the sediment during diagenesis, either just below sediment- water interface, or much deep in the sediment column.

Early diagenetic nodules form in sediments that are still soft and uncompacted. These can be recognized from the presence of uncrushed fossils within the nodule and from the folding of laminae in the mudrock around the nodule, showing that compaction took place after growth of the noudle.

6 Nodules that form after compaction of the host sediment during burial diagenesis are characterized by laminae in the sediment pass unaffected through the nodule.

The growth of nodules arises from the localized precipitation of cement from pore waters within the sediment. The composition, Eh, pH of these pore waters are important in controlling nodule mineralogy and growth rates.

In some case nodules formed around a nucleus, a fossil for example, as a result of local chemical conditions. More commonly nodules are without nucleus, and form along definite horizons or within particular beds, reflecting a level at which supersaturation of pore waters was achieved. Elongate nodules may show a preferred orientation, reflecting the direction of pore water movement.

7.3 The color of Mudrocks The color of mudrocks depends on the mineralogy and geochemistry of the rock. The main controls of color are the organic content, pyrite and oxidation state of iron. With increasing organic matter and pyrite, mudrocks take a darker grey color and eventually become black. Many marine and deltaic mudrocks have a dark grey to black color because of finely disseminated organic matter and pyrite.

Red and purple color results from the presence of ferric oxide, hematite, occurring chiefly as grain coatings and intergrowth with clay particles. Red color develops after deposition, though an aeging processes of a hydrated iron oxide precursor, as the case with red sandstones. The precursor is mainly detrital in origin, and may be in situ solution of metastable mafic mineral grains. Red color of flooding plain mudstones reflects oxidizing nature of depositional and early diagenetic environment.

Green mudrocks contain no hematite, organic matter or iron sulfide, but the color comes from ferrous iron within the lattices of illite and chlorite. The green color may be original or may develop in mudrocks that were red originally but subjected to reduction of hematite by migrating ground water. Green spots and patches in some red mudrocks are sites of iron reduction from local occurrence of organic matter.

Other colors in mudrocks result from mixing of color-producing components. For example, olive and yellow mudrocks may owe this color to a mixing of green minerals and organic matter.

Some mudrocks have a color mottling, where there are different shades of grey that may be the result of bioturbation. In yellows/reds/browns mottles, pedogenic processes resulting by moving water through soil causes an irregular distribution of iron oxide/hydroxide and /or carbonate, and the effect of roots. The color mottling is common in lacustrine and floodplain muds and marls.

7.4 Mineral constituents of mudrocks 7.4.1 Clay minerals

7 Clay minerals are hydrous aluminosilicates with a sheet or layered structures; they are phyllosilicates like the micas. The sheets are of two basic types. One is a layer of silicon- oxygen tetrahedral with three of the oxygen atoms in each tetrahedron shared with adjacent tetrahedral and linked together to form a hexagonal network (Fig. 7.7). The basic unit is Si2O5 but within these silica layers aluminum may replace up to half of the silicon atoms.

Fig. 7.7: Diagrams illustrating the structures of clay minerals.

The second type of layers consists of Al in octahedral coordination with O2- and OH- ions, so that in effect the Al3+ ions are located between two sheets of O/OH ions (Fig. 7.7). In this type of layer, not all the Al (octahedral) positions may be occupied, or Mg2+, Fe and other ions may substitute for Al3+.

Layers of Al-O/OH in a clay mineral are referred to as gibbsite layer because the mineral gibbsite (Al(OH)3) consists entirely of such layers. Similarly, layers of Mg-O/OH are referred to as brucite layers after the mineral brucite (Mg(OH)2) composed solely of this structural unit.

Clay minerals consist of sheets of silica tetrahedral and Al or Mg octahedral linked together by oxygen atoms common to both. The stacking arrangement of the sheets determines the clay mineral type, as does the replacement of Si and Al by other elements.

8 Structurally, there are two basic groups of clay minerals are the kandite group and the smectite group.

Members of the kandite group have two-layered structure consisting of a silica tetrahedral sheet linked to an Al-octahedra (gibbsite) sheet by common O/OH ions (Fig. 7.7). There is no substitution of Al for Si in the kandites so that the structural formula is (OH)4Al2Si2O5.

Members of the kandite group include: kaolinite, the most common member, the dickite and nacrite that have different lattice structure, and the halloysite consisting of kaolinite layers separated by sheets of water. Kaolinite has a basal spacing (distance between on silica layer and the next one) of 7 Ǻ.

Members of the smectite group have a three-layered structure in which an Al-octahedra layer is sandwiched between two layers of Si-tetrahedra layers (Fig. 7.7). The typical basal spacing is 14 Ǻ but smectites have the ability to water molecules and this changes the basal spacing; it may vary from 9.6 Ǻ (with no adsorbed water) to 21.4 Ǻ. This feature of smectites, as a result of which they are called ‘expandable clays’, is utilized in their X-ray identification.

The common smectite is montmoroillonite; it approximates to Al4(Si2O10)2(OH)4.nH2O but substitution of the Al3+ by Fe2+, Mg2+ and Zn2+ can take place. A net negative charge resulting from this substitution is balance by introduction of other cations, especially Ca2+ and Na+ that are contained in interlayer position.

Other members of the smectite group are; nantronite, where Fe3+ replaces Al3+ in the octahedral layer; saponite and stevensite, where Mg2+ substitutes fro Al3+.

Vermiculite has a structure similar to smectite, although it is less expandable, with all the octahedral positions occupied by Mg2+ and Fe2+ and much substitution of Si4+ by Al3+.

Illite is the most common clay mineral in sediments and it is related to the mica muscovite. It has a three-layered structure, like the smectites, but Al3+ substitution for Si4+ in the tetrahedral layer results in a deficit of charge, which is balanced by K ions in interlayer positions (Fig. 7.7). Some hydroxyl OH-, Fe2+ and Mg2+ ions also occur in illite. The basal spacing is about 10 Ǻ.

Other clay minerals are chlorite, like smectites and illite, has three-layered structure but with a brucite layer (Mg-OH) in between (Fig. 7.7). Substitution by Fe2+ occurs in chlorite (imparting green color) and the basal spacing is 14 Ǻ.

Galuconite is related to illite and the micas, but contains Fe3+ substituting for Al3+. Sepiolite and palygorskite are Mg-rich aluminosilicates.

In addition to the above major clay minerals, mixed-layer clays are also common. These consist of stacked sheets of the common clays, in particular illite-montmorillonite and

9 chlorite-montmorillonite. Specific names have been applied where there is a regular mixed-layering: corrensite for a chlorite-montmorillonite mixed layer clay, for example.

During weathering and diagenesis, interlayer cations can be leached out of the clay minerals by percolating waters. Such non-stoichiometric clays are termed degraded and in fact many illites and smectites in modern sediments are degraded, as well as chlorites, biotites and muscovites.

The identification of clay minerals in mudrocks can be carried out employing X-Ray diffraction of the less than 2µm fraction of the sediment, or scanning electron microscope. The polarized-light microscope is usually not used to identify clay minerals because of their small size which is beyond the resolution of this microscope.

7.4.2 Quartz Quartz in mudrocks is mainly of silt-size, although coarser sand-size may occur, especially where mudrocks grade laterally or vertically into sandstones. Quartz in mudstones is mainly detrital, with much less common of diagenetic origin. It is invariably angular, in comparison with more rounded quartz sand.

7.4.3 Other constituents Feldspars may be present in mudrocks but with much less abundance than quartz, because of their lower chemical and mechanical stability relative e to quartz. But because of the lower permeability of mudstones than sandstones, feldspars are better preserved and might be utilized in provenance investigations.

Muscovite is common but biotite is much less abundant. Calcite may occur as skeletal grains as well as digenetic form, in addition to other carbonate minerals such as dolomite and siderite that all can be disseminated or concentrated as nodules.

Pyrite occurs as cubes, framboids and nodules in dark, organic rich mudrocks. Other minerals present locally are glauconite, berthierine, hematite, gypsum, anhydrite and halite. Organic matter is common in mudrocks, particularly black shales.

7.5 The formation and distribution of clay minerals in modern sediments Clay minerals in sediments and sedimentary rocks have three origins: inheritance, neoformation and transformation.

Inherited clay minerals are the detrital clay minerals that have been formed in another area, probably at much earlier tome, transported and finally deposited away from the source area, but they are stable in their present location.

The neoformed clay minerals are the clay minerals that have formed in situ by precipitation from solution or formed from amorphous silicate material.

10 The transformed clay minerals are inherited clay minerals that have been modified by ion exchange (ions in aqueous solution can be adsorbed onto and desorbed from clays) or cation arrangement.

There are three major locations where clay-mineral formation takes place: 1- In the weathering and soil environment. 2- In the depositional environments. 3- During diagenesis and low-grade metamorphism.

The major site of clay-mineral formation is in the weathered mantles and oil profiles developed upon solid bedrock and unconsolidated sediments. Soils develop through physio-chemical and biological processes (pedogenesis) and usually possess distinct horizons (Fig. 7.8).

Fig. 7.8: Schematic diagram of a soil profile.

All clay mineral types can be formed in horizon A of the soil by the various pedogenic processes, particularly chemical weathering of feldspars and micas, and other minerals, in the parent rock.

Illite is formed in soils by limited leaching in temperate areas; chlorite forms in acid soils by intermediate leaching in humid and arid regions; montmorillonite forms in temperate soils by intermediate leaching with good drainage and neutral pH; and kaolinite forms particularly in acid tropical soils by intensive leaching.

These clay minerals, as well as colloidal organic matter and ions in solution, percolate downward from A horizon towards B horizon through a process termed eluviation. The downward percolated clay minerals accumulate in the B horizon through a process called

11 illuviation. In addition to clay minerals accumulating in B horizon, iron oxides/hydroxides, and carbonates are precipitated in this soil horizon.

The clay minerals formed in the soil profiles and in the weathering mantles on provenance region are now subject to erosion then transportation then deposition. Their depositional environments will be discussed below. Besides these detrital clay minerals, others can be formed by other processes. They may precipitate directly from water or pore water in surficial sediments. Also clay mineral may be formed by alteration of volcanic materials. Moreover, clay minerals can be precipitated within coarser siliciclastic sediments as cements during diagenesis.

The distribution of clay minerals in modern sediments is mainly a reflection of the climate and weathering pattern of the source rock area. Kaolinite is dominant in low- latitude areas, particularly off major rivers draining regions of tropical weathering. Illite is more common in ocean muds of higher latitudes. The distribution of smectites, derived mainly from alteration of volcanic material, is related to the active mid-ocean ridge systems and volcanic oceanic islands.

7.6 Mudrocks and their depositional environments There are three major groups of mudrocks in the geologic record: 1- Residual mudrocks formed in situ through contemporaneous processes of weathering and soil formation upon preexisting rocks and sediments. 2- Detrital mudrocks formed through normal sedimentary processes of erosion, transportation and deposition. 3- Volcaniclastic mudrocks formed through in situ weathering and/or later alteration of volcaniclats deposits.

7.6.1 Residual mudrocks and paleosoils Paleosoils or ancient soils are common in the geological record but weathering mantles developed on bedrocks are quite rare. Soils developed upon/within sediments occur as old as the Precambrian, but are much common when plants developed on land since the Silurian Period.

Of these paleosoils are two important types, the calcretes and the seatearths. Calcretes or caliches vary from scattered to densely packed nodules of CaCO3 and are typical semi-arid climatic areas where evaporation exceeds precipitation. They occur in many river floodplain sediments and clay minerals formed in these soils include smectite, sepiolite and playgorskite.

Seatearths or underclays are clayey soils occurring below or beneath coal seams. They are commonly vertisols, formed under a humid tropical climate with seasonal shrink – swell processes. They are typically massive with rootlets and siderite nodules, but they commonly have polygonal and vertical crack systems.

7.6.2 Detrital mudrocks

12 The majority of clay minerals and silt-grade quartz in mudrocks are derived from erosion of continental rocks and soils. These fine terrigenous clastic particles are transported largely in suspension by water, with deposition taking place in quiet, low-energy environments.

Rivers transport vast quantities of silt and clay in suspension to be deposited in floodplains, lakes, deltaic environments and nearshore and offshore marine environments.

Wind can also transport dust up to thousands of kilometers from source areas to the desert where loess (aeloian silt) is deposited. Also dust can be carried by wind to the ocean where it is deposited as hemipelagic sediment.

Mud accumulated on continental shelves and solpes may be resedimented by storms or by gravitational sliding and slumping (turbidity currents) to reach and to be deposited on deeper parts of the ocean.

Also on alluvial fans, glacial-proglacial regimes, and on submarine fans mud can be transported as viscous, sediment-laden, water-poor flows known as mudflows.

Non-marine mudrocks The mudrocks of river floodplains are best identified by their association with fluvial sandstones. They are overbank deposits and may represent upper parts of fining upward sequences. They could be red-colored and contain calcareous pedogenic nodules if they were deposited under arid to semi-arid climate.

Lacustrine (lake deposits) mudrocks vary considerably depending on the chemistry of lake water, organic productivity and climate. However, they could be characterized by presence of millimetric-scale rhythmic laminations resulting from seasonal clastic influx, coupled with phytoplankton growth. Also, glacial deposits could be characterized by varved rhythmic laminations consisting of alternation of coarse and fine laminae. The coarse laminae consisting of silt to fine sand-grade were deposited from low density suspension currents during spring melting. The fine laminae consisting of clay-grade material were deposited from suspension during summer and winter.

Marine mudrocks In the marine environment, mudrocks are deposited in five main locations: muddy coastlines; nearshore and mid-shelf mud belt; open shelf mud blanket; basinal slope; and basin floor.

Muddy coastlines are close to rivers bringing large quantities of suspended mud to the sea, including tidal flats, lagoons behind barrier islands, salt marches and mangrove swamps. Off major deltas, marine shelves may be covered completely in mud to give a shelf mud blanket.

Ancient mudrocks of shoreline environment are identified by restricted fossil assemblages that may suggest brackish water or hypersaline conditions; the presence of

13 rootlets that indicate emergence; presence of mud cracks, rippled lenses of sandstone and association with other channel, beach or barrier sandstone; and dark grey color due to high organic content.

The nearshore mud belt occurs in water depths of 5-20 m, beyond the foreshore- shoreface sand belt of many coastlines. Mud is deposited out of suspension below fairweather wave-base.

The typical mudrocks of the open-shelf, deposited mainly below wave-base, are various shades of grey and in some case rich in fossils. The fossils are both epifaunal (living on the sediment surface) and infaunal (living within the sediment) with some pelagic forms (free-swimming and free-floating species). Bioturbation is common. Thin sharp-based and graded beds of sandstone and limestone within the mudrocks may represent storm deposits (tempestites).

Muds and mudrocks deposited in deeper water largely from suspension are termed hemipelagic. Such muds (hemipelagites) cover the sea floor on the deep, outer parts of continental shelves, on continental slopes and on fast areas of ocean basin. The deep- ocean floors are usually well oxygenated. The reason of this is that cold, dense oxygen- rich waters are produced in Polar Regions that descend and flow above ocean bottom to lower latitudes, thus keeping the ocean floor oxygenated.

Many pelagic mudrocks are grey in color, although red, brown, green and black varieties also occur. They are characterized by pelagic fauna, such as diatoms, planktonic foraminifera and Coccolithophoridae from the Mesozoic to the present day; radiolarian from the Paleozoic; cephalopods in the later Paleozoic and Mesozoic; and graptolite in the early Paleozoic. Hemipelagic mudrocks are commonly interbedded with siliciclastic and carbonate turbidites.

Hemipelagic mudrocks may grade laterally or vertically into pelagic limestone or radiolarites which form in areas and at times of minimum clay and silt sedimentation. Modern hemipelagic sediments accumulating below CCD are red and brown clays that cover the abyssal plains of the central pacific and occur in the Atlantic and Indian Oceans. They consist of detrital clay and silt, clay minerals and zeolites derived from alteration of volcanic ash, and radiolarians, diatoms and sponge spicules.

Organic-rich mudrocks and black shale One particular important group of mudrocks comprises those rich in organic matter including black shales and carbonaceous and bituminous mudrocks which typically contain 3-10% organic carbon. With an increasing organic content, organic-rich mudrocks pass into oil shales, which yield a significant amount of oil on heating.

In many depositional environments organic matter is decomposed and destroyed at the sediment surface but if the rate of organic productivity is high, then organic matter can be preserved. The accumulation of organic matter is favored if the circulation of water is

14 restricted to some extent so that insufficient oxygen reaches the bottom sediments to decompose the organic material (Fig.7.9).

Fig. 7.9: The relationships between mudrock facies and oxicity-anoxicity, fauna and organic matter content.

Locations where this commonly takes place are lakes, fjords, silled basins (such as the Black Sea), sediment-starved basins and deep ocean trenches (such as the Cariaco trench).

As a result of poor circulation and restriction, the water body becomes stratified and the sea or lake floor may become oxygen deficient (dysaerobic) or totally anoxic.

Dysaerobic conditions also occur where the sea floor is within the oxygen-minimum zone, generally in the depth range 100-1000 m. This zone of low O2 results from the bacterial decomposition of organic matter sinking from fertile, oxic, surface waters. Irrespective of the degree of anoxia, however, the major control on organic-carbon accumulation does appear to be the primary production rate.

Where there is oxygen deficiency on the sea floor, organic matter will be preserved, but the surface sediments could still support a benthic epifauna, although of low diversity.

15 Where there are anoxic conditions on the sea floor, there usually is much H2S in the water and benthic organisms are absent (Fig. 7.9). This is the case with the Black Sea and Cariaco Trench at the present time.

Mudrocks deposited in an anoxic environment would contain only pelagic fossils. Pyrite is common in marine organic-rich mudrocks and siderite in non-marine ones.

One feature of organic-rich sediments is that they contain high concentrations of certain trace elements, such as Cu, Pb, Zn, Mo, V, U, and As. The trace elements are adsorbed onto the organic matter and also onto the clay minerals. It is likely that the source of these elements is the sea water and that they are scavenged by the organisms and organic matter.

Loess and loessite Loess is a yellow-to-buff-colored clastic deposit composed mainly of silt-sized quartz grains, generally in the size 20-50 µm. A distinctive feature is the well-sorted nature of the silt, together with a dominantly angular shape to the grains.

Loess is usually unstratified and unconsolidated, but it may contain shells of land snails and concretions formed around roots.

Loess deposited during the late Pleistocene occurs over vast areas of central Europe, Eastern South America and China.

Loess is regarded primarily of an aeolian (wind) deposit, but once deposited can be modified by fluvial reworking and pedogenesis. There are two types of loess: loess of cold, periglacial regimes derived from deflation of glacial-outwash plains (this accounts for most of the late Pleistocene occurrences); and loess derived from hot, arid, desert areas.

7.6.3 Mudrocks of volcaniclastic origin Mudrocks formed from the alteration of volcaniclastic material are known as bentonites (also fuller‘s earth) if montmorillonite is the main clay mineral present and tonstein if kaolinite is dominant. Zeolites are also present.

The volcaniclastic deposits may be subaerial or subaqueous, but because of the metastable nature of volcanic glass, devitrification soon takes place and clay minerals and zeolites form.

7.7 Diagenesis of clay minerals and mudrocks Clay minerals can be modified and altered during early and late diagenesis, and into metamorphism.

The main physical diagenetic process affecting mudrocks is compaction. Compaction in mudrocks expels water, reduces thickness of the deposited sediment and reduces

16 porosity. Upon deposition, muds contain 70-90% water by volume, which could be reduced to 30% at a burial depth of 100m (Fig. 7.10).

Fig. 7.10: Diagram illustrating the stages of water loss from muddy sediments with increasing depth of burial.

Much of this water is not free pore water but is contained in the lattice of the clay minerals and adsorbed onto the clays. Further compaction due to increasing burial depth causes further water loss. At a burial depth of about 2-4 km dehydration and some change in clay mineralogy take place. Final compaction to give a mudrock with only a few per cent water requires a much longer period of overburden pressure with elevated temperatures.

Evidence of compaction in mudrocks is given by the fracture of shells, flattening of burrows, and the bending of laminae around shells and early diagenetic nodules.

Chemical diagenesis, or the change of chemistry and/or mineralogy of clay minerals, takes place mainly through the rise of temperature accompanying increased burial depth. The main change is the alteration of smectites to illite via mixed-layer clays of smectite- illite (Fig. 7.11).

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Fig. 7.11: Diagram illustrating the changes of clay minerals with increasing depth of burial and into metamorphism.

This alteration involves the incorporation of K+ ions into the smectite structure and loss of interlayer water. This process is temperature dependant, where the smectite starts to disappear at about 70-95 ºC, corresponding to a burial depth of about 3-3 km under normal geothermal gradient (30 ºC km-1). At slightly higher temperatures and greater depths, kaolinite is replaced by illite and chlorite.

There is a change in the clay mineralogy of mudrocks through geological time (Fig. 7.12).

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Fig. 7.12: The distribution of clay minerals through time.

Mudrocks of the Upper Paleozoic, Mesozoic and Cenozoic contain a variety of clay minerals, whereas Lower Paleozoic and Precambrian mudrocks are dominated by illite and chlorite. Similar to changes of clay minerals through increasing burial depth, with greater ages more time is allowed for diagenetic reactions.

The irregular pattern of smectite behavior through time may be related to orogenic periods. At these times, volcanism is generally more widespread, leading to the formation of much smectite.

Passing into the realm of incipient metamorphism (catagenesis or anchimetamorphism) clay minerals are further altered and replaced (Fig. 7.11). The phyllosilicate pyrophyllite (related to tlac) and laumontite (a zeolite) may develop at the expense of clay minerals. Although smectite, mixed-layers clays and kaolinite do not survive into metamorphism, illite and chlorite do.

With increasing degree of incipient and low grade metamorphism, the order or crystallinity of the illite lattice increases (measured from XRD by sharpness ratio of the peak).

There are also changes in the chemical composition (an increase in the Al/(Fe + Mg) ratio). Illite is then replaced by sericite, a finely crystalline variety of muscovite. The percentage of smectite-illite mixed layering also decreases with increasing burial.

19 Studies of clay mineralogy, if combined with measurements of the rank of associated coal and vitrinite reflectance, they can give an indication of the temperature to which the formation as a whole has been subjected.

20