A certain amount of modification of the sediment occurs at the sediment–water and sediment–air interfaces: cements formed at this stage are referred to as eogenetic cements and they are essentially synsedimentary, or very soon after deposition. Most chemical changes occur in sediment that is buried and saturated with pore waters, and cements formed at this stage are called mesogenetic. Rarely cement formation occurs during uplift, known as telogenetic cementation. During these diagenetic stages, chemical reactions take place between the grains, the water and ions dissolved in the pore waters: these reactions take place at low temperatures and are generally very slow. They involve dissolution of some mineral grains, the precipitation of new minerals, the recrystallisation of minerals and the replacement of one mineral by another.
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Dissolution
The processes of grain dissolution are determined by the composition of the grain minerals and the chemistry of the pore waters. Carbonate solubility increases with decreasing temperature and increasing acidity (decreasing pH): the presence of carbon dioxide in solution will increase the acidity of pore waters and leaching of compounds from organic matter may also reduce the pH. It is therefore common for calcareous shelly debris within terrigenous clastic sediment to be dissolved, and if this happens before any lithification occurs then all traces of the fossil may be lost. Dissolution of a fossil after cementation may leave the mould of it, which may either remain as a void or may subsequently be filled by cement to create a cast of the fossil. Silica solubility in water is very low compared with calcium carbonate, so large-scale dissolution of quartz is very uncommon. Silica is, however, more soluble in warmer water and under more alkaline (higher pH) conditions, and opaline silica is more soluble than crystalline quartz. Most quartz dissolution occurs at grain boundaries as a pressure dissolution effect, but the silica released is usually precipitated in adjacent pore spaces.
Precipitation of cements
The nucleation and growth of crystals within pore spaces in sediments is the process of cementation. A distinction must be made between matrix, which is fine-grained material deposited with the larger grains, and cements, which are minerals precipitated within pore spaces during diagenesis. A number of different minerals can form cements, the most common being silica, usually as quartz but occasionally as chalcedony, carbonates, typically calcite but aragonite, dolomite and siderite cements are also known, and clay minerals. The type of cement formed in a sediment body depends on the availability of different minerals in pore waters, the temperature and the acidity of the pore waters. Carbonate minerals may precipitate as cements if the temperature rises or the acidity decreases, and silica cementation occurs under increased acidity or cooler conditions. Growth of cement preferentially takes place on a grain of the same composition, so, for example, silica cement more readily forms on a quartz grain than on grains of a different mineral. Where the crystal in the cement grows on an existing grain it creates an overgrowth with the grain and the cement forms a continuous mineral crystal. These are referred to as syntaxial overgrowths. Overgrowths are commonly seen in silica-cemented quartz sands; thin-section examination reveals the shape of a quartz crystal formed around a detrital quartz grain, with the shape of the original grain picked out by a slightly darker rim within the new crystal. In carbonate rocks overgrowths of sparry calcite form over biogenic fragments of organisms such as crinoids and echinoids that are made up of single calcite crystals. Cementation lithifies the sediment into a rock and as it does so it reduces both the porosity and the permeability. The porosity of a rock is the proportion of its volume that is not occupied by solid material but is instead filled with a gas or liquid. Primary porosity is formed at the time of deposition and is made up mainly of the spaces between grains, or interparticle porosity, with some sediments also possessing intraparticle porosity formed by voids within grains, usually within the structures of shelly organisms. Cements form around the edges of grains and grow out into the pore spaces reducing the porosity. Secondary porosity forms after deposition and is a result of diagenetic processes: most commonly this occurs as pore waters selectively dissolve parts of the rock such as shells made of calcium carbonate. Permeability is the ease with which a fluid can pass through a volume of a rock and is only partly related to porosity. It is possible for a rock to have a high porosity but a low permeability if most of the pore spaces are not connected to each other: this can occur in a porous sandstone which develops a partial cement that blocks the 'throats' between interparticle pore spaces, or a limestone that has porosity sealed inside the chambers of shelly fossils. A rock can also have relatively low porosity but be very permeable if it contains large numbers of interconnected cracks. Cement growth tends to block up the gaps between the grains reducing the permeability. Pore spaces can be completely filled by cement resulting in a complete lithification of the sediment and a reduction of the porosity and permeability to zero.
Recrystallisation
The in situ formation of new crystal structures while retaining the basic chemical composition is the process of recrystallisation. This is common in carbonates of biogenic origin because the mineral forms created by an organism, such as aragonite or high magnesium calcite, are not stable under diagenetic conditions and they recrystallise to form grains of low magnesium calcite. The recrystallised grains will commonly have the same external morphology as the original shell or skeletal material, but the internal microstructure may be lost in the process. Recrystallisation occurs in many molluscs, but does not occur under diagenetic conditions in groups such as crinoids, echinoids and most brachiopods, all of which have hard parts composed of low magnesium calcite. Recrystallisation of the siliceous hard parts of organisms such as sponges and radiolaria occurs because the original structures are in the form of amorphous opaline silica, which recystallises to microcrystalline quartz.
Replacement
The replacement of a grain by a different mineral occurs with grains of biogenic origin and also detrital mineral grains. For example, feldspars are common detrital grains and to varying degrees all types of feldspar undergo breakdown during diagenesis. The chemical reactions involve the formation of new clay minerals that may completely replace the volume of the original feldspar grain. Feldspars rich in calcium are the most susceptible to alteration and replacement by clay minerals, whereas sodium-rich, and particularly potassium-rich, feldspars are more resistant. These reactions may take millions of years to complete. Silicification is a replacement process that occurs in carbonate rocks: differences between the mineralogy of a shelly fossil and the surrounding carbonate rock may allow the calcium carbonate of the fossil to be partly or completely replaced by silica if there are silica-rich pore waters present in the rock.
Nodules and concretions
Most sedimentary deposits are heterogeneous, with variations in the concentrations of different gain sizes and grain compositions occurring at all scales. The passage of pore waters through the sediment will be affected by variations in the porosity and permeability due to the distribution of clay particles that inhibit the flow. The presence of the remains of plants and animals creates localised concentrations of organic material that influence biochemical reactions within the sediment. These heterogeneities in the body mean that the processes that cause cementation are unevenly dispersed and hence some parts become cemented more quickly than others. Where the distinction between well-indurated patches of sediment and the surrounding body of material is very marked the cementation forms nodules and concretions. Irregular cemented patches are normally referred to as nodules and more symmetrical, round or discoid features are called concretions. Nodules and concretions can form in any sediment that is porous and permeable. They are commonly seen in sand beds (where large nodules are sometimes referred to as doggers), mudrock and limestone. Sometimes they may be seen to have nucleated around a specific feature, such as the body of a dead animal or plant debris, but in other cases there is no obvious reason for the localised cementation. Concretions formed at particular levels within a succession may coalesce to form bands of well-cemented rock. A variety of different minerals can be the cementing medium, including calcite, siderite, pyrite and silica. In places there is clear evidence that concretions in mudrocks form very soon after deposition: if the layering within the mudstones drapes around the concretion. This is evidence that the cementation occurred locally before the rock as a whole underwent compaction.
Septarian concretions
The interiors of some carbonate concretions in mudstones display an array of cracks that are often filled with sparry calcite. These are known as septarian structures, and they are believed to form during the early stages of burial of the sediment. The precise mechanism of formation of the cracks is unclear, but is believed to be either the result of shrinkage (in a process similar to syneresis), or related to excess pore fluid pressure in the concretion during compaction, or a combination of the two.
Flints and other secondary cherts
Chert can form directly from siliceous ooze deposited on the sea floor: these primary cherts occur in layers associated with other deep-water sediments. Chert may also form in concretions or nodules as a result of the concentration of silica during diagenesis. These secondary cherts are diagenetically formed and are common in sedimentary rocks, particularly limestones. They are generally in the form of nodules that are sometimes coalesced to form layers. The diagenetic origin to these cherts can be seen in replacement fabrics, where the structures of organisms that originally had carbonate hard parts can be seen within the chert nodules. The edges of a chert nodule may also cut across sedimentary layering. The nodules form by the very fine-scale dissolution of the original material and precipitation of silica, often allowing detailed original biogenic structures to be seen. The source of the silica is generally the remains of siliceous organisms deposited with the calcareous sediment. These organisms are sponges, diatoms and radiolarians that originally have silica in a hydrated, opaline form, and in shelf sediments sponge spicules are the most important sources of silica. The opaline silica is relatively soluble and it is transported through pore waters to places where it precipitates, usually around fossils, or burrows as microcrystalline or chalcedonic quartz in the form of a nodule. Flint is the specific name given to nodules of chert formed in the Cretaceous Chalk.
Colour changes during diagenesis
The colour of a sedimentary rock can be very misleading when interpretation of the depositional environment is being attempted. It is very tempting to assume, for example, that all strongly reddened sandstone beds have been deposited in a strongly oxidising environment such as a desert. Although an arid continental setting will result in oxidation of iron oxides in the sediment, changes in the oxidation state of iron minerals, the main contributors to sediment colour, can occur during diagenesis. A body of sediment may be deposited in a reducing environment but if the pore waters passing through the rock long after deposition are oxidising then any iron minerals are likely to be altered to iron oxides. Conversely, reducing pore waters may change the colour of the sediment from red to green. Diagenetic colour changes are obvious where the boundaries between the areas of different colour are not related to primary bedding structures. In fine grained sediments reduction spots may form around particles of organic matter: the breakdown of the organic matter draws oxygen ions from the surrounding material and results in a localised reduction of oxides from a red or purple colour to grey or green. Bands of colour formed by concentrations of iron oxides in irregular layers within a rock are called liesegangen bands. The bands are millimetre-scale and can look very much like sedimentary laminae. They can be distinguished from primary structures as they cut across bedding planes or cross-strata and there is no grain-size variation between the layers of liesegangen bands. They form by precipitation of iron oxides out of pore waters. Other colour changes may result from the formation of minerals such as zeolites, which are much paler than the dark volcanic rocks within which they form.
