sedimentary geochemistry

sedimentary geochemistry Sedimentary geochemistry applies chemical principles and techniques to decipher the history of sedimentary rocks from their formation as sediments through lithification to form rocks, to eventual erosion to form the components for new sediments. Sedimentary geochemists also model this sedimentary cycle to better understand elemental cycles, such as the carbon cycle, and the fate and transport of chemical compounds, including pollutants, in soils, surface waters and groundwater.

During formation, sediments develop primary chemical signatures reflective of their origins. These signatures include the mineralogy, and the isotopic and trace-element compositions of minerals. After deposition, sediments are subjected to a number of changes which can add new minerals, dissolve old minerals, and alter the organization of components. These processes also cause the loose sediment grains to be bound together to form sedimentary rocks. These processes, collectively termed sedimentary diagenesis, occur through interactions between sediment grains and the waters in contact with them. Consequently, a large part of the discipline of sedimentary geochemistry is concerned with study of the chemistry of diagenetic waters.

A number of clues are present in sedimentary rocks that a sedimentary geochemist can use in reconstructing rock history and diagenetic fluid chemistry. Rocks may be analysed to determine their bulk chemistry or mineralogical compositions, or the chemical compositions of individual constituents may be determined separately. In addition, during growth, some minerals incorporate small pockets of liquid, termed fluid inclusions, and these too can be analysed.

Materials making up sedimentary rocks can be divided into three general categories: clastic minerals, produced by weathering of precursor rocks; minerals that are chemically or biochemically precipitated, including carbonates, evaporites, sulphides, oxides, and phosphates; and organic matter, including kerogen, coal, and petroleum.

Clastic sediments

Clastic rocks are composed predominantly of those minerals most resistant to weathering; these are also the igneous minerals which form at the lowest temperatures in Bowen's reaction series: quartz, feldspar, and mica. Classic rock groups include sandstones, siltstones and shales. The more soluble minerals in source rocks dissolve, contributing chemicals such as potassium, calcium, sodium, iron, magnesium, and silicon to surface waters and sediment pore waters. These chemicals often subsequently reprecipitate during diagenesis to form new minerals as cements in sediment pore spaces. Common cements include clay minerals such as kaolinite, montmorillonite, or illite; quartz or chalcedony; iron oxides such as haematite; or calcite. Diagenesis may also lead to dissolution of detrital minerals after the sediment has lithified, creating pockets of porosity which may subsequently be filled by petroleum or water.

The geochemistry of clastic rocks focuses on determining the source rocks of weather-resistant grains and on the geochemical conditions that lead to cement formation and diagenetic dissolution. Bulk-rock geochemistry is often determined and usually correlates well with mineralogy. For sandstones, major-element geochemistry also correlates well with tectonic setting because of differences in igneous and metamorphic source rocks that are present in ocean island arcs and passive or active continental margins.

Precipitated sediments

Carbonates

Carbonate sediments are composed predominantly of fragments of the skeletal hard parts of marine organisms, including corals, clams, and other molluscs, and calcareous algae. The hard parts are composed of calcite and aragonite, two mineral polymorphs of calcium carbonate. The calcite may contain some magnesium substituted for calcium to produce low-magnesium (less than 5 per cent substitution) or high-magnesium (5–20 per cent substitution) calcite. High-magnesium calcite is characteristic of organisms in tropical to subtropical marine environments and hypersaline settings; low-magnesium calcite of organisms in freshwater or temperate marine environments. Aragonite is mostly restricted to tropical and subtropical marine and hypersaline environments. After these grains are deposited as sediments, these same minerals are also commonly precipitated in marine environments as cements between sediment particles. In terrestrial, freshwater settings, inorganically precipitated low-magnesium calcite forms in subsurface aquifers, coastal zones where sea water mixes with groundwater, hydrothermal springs, and cave deposits. Dolomite also forms in very small quantities in some modern sediments, apparently as a product of early diagenetic alteration, but it is abundant in ancient carbonate rocks. Siderite (iron carbonate), magnesite (magnesium carbonate), strontianite (strontium carbonate), rhodochrosite (manganese carbonate), and smithsonite (zinc carbonate) may form as minor secondary phases in some sedimentary rocks.

By comparison with clastic rocks, carbonates are very reactive and commonly undergo substantial mineralogical and textural changes during diagenesis. When shallow marine sediments are exposed to fresh water, high-magnesium calcite and aragonite are readily replaced by low-magnesium calcite, often with the structure of the organism's hard parts still preserved. These minerals may also dissolve without being replaced, resulting in considerable increases in rock porosity.

The nature of the organisms whose hard parts are preserved in carbonate rocks helps geologists to identify the type of environment that was the source of the sediment. After deposition and diagenetic alteration, mineralogy and trace-element and isotopic chemistry are keys to deciphering rock history and the chemistry of diagenetic fluids. Ancient carbonate rocks have mainly been converted to low-magnesium calcite or dolomite, and so one of the first tasks of a carbonate geochemist is to reconstruct the original mineralogy of the constituent cements and grains and the chemistry of the waters causing the mineral alteration. For the skeletal fragments, the mineralogy of similar or equivalent modern organisms is assumed for their ancient counterparts. However, some organisms do not have surviving examples, and their original mineralogy is inferred from studying the geochemistry of their fossil hard parts. One clue for determining the mineralogy of an original cement is the shapes of crystals, which are commonly preserved after mineral alteration: aragonite tends to form long, needle-like crystals, high-magnesium calcites tend to form spiky dog-tooth shapes, and low-magnesium calcites tend to be equant, blocky crystals. In addition, the replacement is usually not complete; small domains of aragonite or high-magnesium calcite can be found by electron microscopy in low-magnesium calcite or dolomite crystals. Also, aragonite tends to have much higher amounts of strontium as a trace element in the calcium carbonate. Low-magnesium calcite will thus have higher strontium values when it has replaced aragonite than when it replaces high-magnesium calcite.

Other precipitated rocks

Evaporites are formed by precipitation of minerals as a body of water evaporates. Evaporite deposits formed from sea water have a characteristic mineral sequence, beginning with calcite, gypsum, and halite, followed by potash salts. In lakes on continental crust, such as the Great Salt Lake, or the hypersaline lakes of the East African Rift Valley, the mineral salts are different, and may include trona, mirabilite, glauberite, epsomite, and natron, as well as gypsum, halite, and carbonate minerals. Evaporites also form from early or late-stage diagenetic processes in the subsurface. Geochemists often have the task of determining whether evaporite minerals are primary or diagenetic. Isotopes, trace elements, and analysis of fluid inclusions are valuable approaches for differentiating between the two.

Chert is another precipitated sedimentary rock. It is composed of various polymorphs of silicon dioxide: chalcedony, opal, and quartz. Chert may derive from shells of microscopic silica-precipitating marine organisms: radiolaria and diatoms, and some sponges. Chert (and flint, which is a distinctive form of chert found in the chalk of north-western Europe) also forms diagenetically as stringers and nodules in limestones, or as a replacement for carbonate mineral fossil fragments. Trace elements are common in chert-forming minerals and these give chert some spectacularly colourful banding, creating the complex patterns found in chert nodule geodes and jewellery stones. Jasper, for example, is red chert, coloured by trace amounts of iron. Trace elements and oxygen isotopes are useful geochemical indicators of the chemistry of the waters from which the chert precipitated.

Iron-rich sedimentary rocks are particularly abundant in Precambrian sequences, but were also formed during the Early Palaeozoic, the Jurassic, and Cretaceous. They can be divided into iron formations, which are commonly banded, consisting of magnetite and haematite interbedded with chert, and are primarily of Precambrian age; and ironstones, which are Palaeozoic and Mesozoic in age, commonly consisting of goethite and haematite associated with carbonate minerals, and usually are massive or oolitic. There are no modern analogues for iron-rich sedimentary rocks, and their origins remain enigmatic. Today, iron is deposited as disseminated sulphides and oxides in carbonate and clastic rocks, and as ferromaganese nodules on the sea floor. The chemistry of iron depends on the amount of dissolved oxygen present and the acidity of solutions; precipitation of iron as oxides or hydroxides occurs rapidly unless acid, anoxic (or reducing) conditions are present. Thus, identifying the geochemical mechanism that kept iron in solution long enough to accumulate it in these enriched ancient deposits remains one of the classic problems in sediment geochemistry; also puzzling is why this mechanism is not operating today.

Phosphorites are sedimentary rocks containing more than 15–20 per cent P₂O₅, present as the mineral apatite, often concentrated in carbonate or chert host rocks. The remains of fish skeletons, shark teeth, nodules, pellets, phosphatized carbonate fossil fragments, and ooids are common constituents. Phosphorites are presumed to originate as marine deposits in zones of upwelling currents, but there is, again, a need to find a mechanism to explain how these high amounts of phosphate (up to 40 per cent in some ancient deposits) become concentrated from the low levels present in sea water. Most geochemists believe that this concentration process occurs diagenetically in the pore waters of ocean-floor sediments. Measurement of oxygen and carbon isotopes, as well as trace-element geochemistry, can be used to constrain the compositions of diagenetic fluids.

Carbonaceous sedimentary rocks contain at least 10 per cent organic matter. Two types of organic matter accumulate in sediments: humus and sapropel. Humus is produced from the decay of plant organic matter in soils. In soils, humus is usually only a few per cent of the total material, but in environments where there is little oxygen, such as bogs or swamps, enough humus is preserved to create peat, which is over 40 per cent organic. Sapropel is organic matter produced predominantly from the decay of phytoplankton and zooplankton in marine, estuarine, or lake environments. With increasing burial and temperature, diagenetic processes change organic matter in sedimentary rocks into kerogen, the precursor of petroleum. Organic matter consists of carbon, with much smaller amounts of hydrogen, oxygen, nitrogen, and sulphur. The relative amounts of these elements in kerogen are related to the nature of the source material. Oxygen, carbon, and hydrogen isotopes can also be useful indicators of the diagenetic history.

Thermodynamics

Thermodynamics enables geochemists to determine whether a mineral, in contact with water of a specified composition, will precipitate or dissolve. This is dependent on the quantity of the mineral's constituent ions that are present in solution. A simple experiment with table salt (sodium chloride, chemically the same as the mineral halite) and a glass of water illustrates the concept. If salt is added to the water, it dissolves initially, contributing sodium and chloride ions to solution. If more and more salt is added, a point is eventually reached when no further dissolution occurs, and salt crystals remain on the bottom of the glass. At this point, the water is said to be exactly saturated or at equilibrium with respect to sodium chloride. Before equilibrium was reached, fewer sodium and chloride ions were present in the water than the amounts needed at equilibrium, and the solution was undersaturated with respect to sodium chloride. Similarly, if more sodium and chloride ions were present, the solution would be supersaturated with respect to sodium chloride. The amounts of sodium and chloride ions in solution at equilibrium can be measured to express a quantity called the solubility constant. The value of the solubility constant is in fact constant only for a specific set of geochemical conditions. If the glass of water at equilibrium with the salt at room temperature were heated, the salt at the bottom of the glass would dissolve further, coming to a new equilibrium concentration at the higher temperature; if the water were cooled, some of the sodium and chloride ions would start to precipitate, forming new salt crystals until a new equilibrium was established. If further experiments were done by putting the salt solution in a pressure cooker, or by adding other compounds to the water, the dependence of the solubility constant on pressure, and on the amounts of other ions present in solution could also be demonstrated.

The situation for the geochemistry of natural waters is akin to having a glass of water with many dissolved ions in it and many minerals in contact with it. To determine whether the water is supersaturated, saturated, or undersaturated with respect to the minerals present, a complete water analysis must be done, and the temperature and pressure of the environment must be known. From extensive databases of solubility constants and their dependence on bulk solution chemistry, geochemists can determine which minerals are dissolving or precipitating. These reactions not only determine the types of alteration occurring in the sediments, but also indicate which minerals are influencing the compositions of the diagenetic fluids. For the sedimentary geochemist, however, the situation is similar to having a suite of minerals in the bottom of the glass, but no water present. The goal is to use the minerals to work out what the water composition was at an earlier stage.

Thermodynamics can also be used to assess the stability of minerals relative to each other, under a given set of geochemical conditions. Whether one mineral is more stable than another is determined by the energy associated with the reaction in which one mineral is converted into another, and by the amount of disorder in the system. Reactions tend to proceed to a state of lowest energy and maximum disorder.

From thermodynamic calculations of mineral solubility and mineral conversion reactions, diagrams can be constructed which show stability fields for various minerals. The axes for these diagrams are various geochemical parameters, such as temperature, ion concentrations in solution, or concentration ratios of solution. A particularly important type of stability diagram plots the acidity of solutions, measured as pH, against the amount of available oxygen in solution, measured as Eh, oxidation–reduction, or redox, potential. Eh–pH diagrams are used extensively in sedimentary geochemistry to predict the behaviour of metals in water, and the formation of metal-rich sedimentary rocks, including iron formations.

Kinetics

Thermodynamics predicts the direction and magnitude of a reaction, but the rate of reaction and the pathway by which it occurs is the subject of kinetics. A geochemist once compared thermodynamics and kinetics by use of a feather allowed to drop to the floor. Thermodynamics described its initial state, resting in his hand, the fact that it would fall downward, not upward, and its final state, resting on the floor. But the complex pathway of fall, with all of the feather's various floating motions, is described by kinetics.

Sedimentary geochemists study the kinetics of mineral reactions in order to understand why thermodynamically predicted states do not occur. A classic example is the case of dolomite. Dolomite, from a thermodynamic standpoint, is more stable than calcite or aragonite, and modern sea water is hundreds of times oversaturated with respect to dolomite. However, dolomite is rarely, if ever, precipitated directly from sea water, nor does it replace calcite or aragonite except in special environments (see dolomite). Kinetic studies provide some answers: dolomite precipitation rates are extremely slow compared with calcite and aragonite growth rates. Thus, calcite and aragonite, competing for the same constituent ions in solutions, end up predominating over dolomite, in spite of its thermodynamically favoured status.

Isotope geochemistry

Ratios of stable isotopes in minerals are important clues for deducing the chemistry of the water and environment in which the minerals have been precipitated, but the interpretation of these isotopic signatures can be extremely complex. Stable isotopes are atoms of an element which have the same number of protons but different numbers of neutrons in their nuclei. These extra neutrons do not, however, make the nuclei unstable, and so they do not decay radioactively. Sedimentary geochemists commonly measure stable isotopic ratios for carbon, oxygen, strontium, hydrogen, and sulphur in a variety of minerals.

Because of differences in their chemical behaviour, isotopes of an element are present in minerals in ratios that are different from those that are found in the waters from which they were precipitated. The degree of difference, termed the isotopic fractionation factor for a mineral, depends on the temperature of precipitation, but is independent of most other environmental conditions. This means that isotopic ratios in minerals can be used to determine the isotopic ratio of the water if the temperature of precipitation is known; conversely, the temperature can be determined if the isotopic ratio of the water is known.

Carbon

About 99 per cent of the carbon atoms on Earth have an atomic weight of 12 (¹²C), their nuclei consisting of six protons and six neutrons; there is also, however, about 1 per cent of carbon atoms that have extra neutrons, giving them atomic weights of 13 (¹³C). Most of the carbon in the world is found in two places (i.e. reservoirs): carbonate minerals and organic matter. Carbonate minerals preferentially incorporate more ¹³C, giving them higher ¹³C/¹²C ratios than the waters from which they are precipitated. Plants selectively take up ¹²C, and so any organic matter in rocks or sediments from plants or phytoplankton (or from any animals that eat plants or phytoplankton) has less ¹³C than would be expected if plants took in both isotopes indiscriminately from their carbon sources. Thus, carbon is present in one reservoir of mineral or inorganic carbon that is relatively enriched in ¹³C, and in one of organic carbon that is relatively depleted. The ¹³C/¹²C ratios of carbon-containing ions in water reflect the amount of interaction of the water with these organic and inorganic reservoirs: if decay of organic matter is occurring (as in swamps and wetlands), waters will have relatively low ¹³C/¹²C ratios; if waters are dissolving carbonate minerals (as in aquifers near recharge areas), they will have relatively high ¹³C/¹²C ratios. Any carbonate minerals that are subsequently precipitated from these waters will also have relatively low and high ¹³C/¹²C ratios, respectively, offset from the water values by their fractionation factors at the temperature of precipitation.

Oxygen

Oxygen isotopes give different information about the source the water from which minerals have formed. Oxygen has two important isotopes: ¹⁶O (99.76 per cent) and ¹⁸O (0.20 per cent). When water, H₂O, evaporates, ¹⁶O is preferentially removed, leaving residual water enriched in ¹⁸O; thus, rainwater has lower ratios than the sea water from which it evaporates; and rain at progressively higher latitudes, formed from multiple evaporation–precipitation cycles, has progressively lower ratios of ¹⁸O/¹⁶O. Thus freshwater carbonates show a relative depletion in ¹⁸O relative to marine carbonates, and there is a progression with latitude in the signatures of freshwater carbonates. However, changes in temperature also produce differences in the oxygen isotope fractionation factors for minerals, making if difficult to separate these effects.

Oxygen isotope ratios can aid in determining sources for detrital particles in clastic sedimentary rocks and can be used to identify diagenetic waters from which clays, silica, and carbonate cements precipitated.

Hydrogen

Hydrogen has two important stable isotopes: ¹H, making up 99.984 per cent and ²H, also known as deuterium or D, making up 0.016 per cent. As with oxygen, when water evaporates, the lighter isotope, ¹H, is preferentially removed, leaving the residual water relatively enriched in the heavier isotope. Hydrogen isotope ratios are measured to aid in determining the origin of diagenetic waters. They can also be measured in clays or other mineral that contain hydroxyl (OH) groups.

Strontium

Strontium ratios (⁸⁷Sr/⁸⁶Sr) can also be used to trace water sources. As noted above, carbonate minerals incorporate strontium in trace amounts. Higher ratios indicate that formation waters have been in contact with clays or other minerals containing radioactive rubidium, which decays to produce stable ⁸⁷Sr. Carbonates forming from deep groundwaters in sedimentary basins may have high ratios. Sea water appears to have a constant ratio today, but differences in the rate of interaction of sea water with basalt at mid-ocean ridges have produced changes in the sea water ratio over geological time. In turn, these changes are reflected in the ⁸⁷Sr/⁸⁶Sr ratios of marine carbonates.

Sulphur

Sulphur has two isotopes of importance to sedimentary geochemistry, ³⁴S and ³²S, representing 4.2 per cent and 95 per cent of total sulphur, respectively. In small percentages, metal sulphide minerals are common in both carbonate and clastic rocks. They form when sediment pore waters become depleted in oxygen, which is used up as bacteria break down organic matter. Some bacteria can use sulphate as a substitute for oxygen, and this process causes the sulphate to be transformed to sulphide. If metal ions are presented in solution, the sulphide ions readily precipitate out, forming minerals such as pyrite within the pore spaces of sedimentary rocks.

Trace-element geochemistry

The presence of trace amounts of some elements can be used to determine the conditions of formation of precipitated minerals or the source rocks of detrital minerals. For carbonate minerals, the amounts of strontium and magnesium in calcite are important indicators for the alteration of precursor aragonite or high-magnesium calcite. The amount of magnesium incorporated in calcite also apparently increases with increasing temperature. Iron and manganese concentrations in carbonate cements can be determined qualitatively by the use of a cathodoluminescence microscope, or quantitatively by a microprobe. Incorporation of manganese and iron indicates that reducing conditions prevailed when the cements were precipitated. In clastic rocks, various trace elements are used to match source rocks with detrital quartz and feldspar grains. Cathdoluminescence of quartz, which measures trace-element concentrations and lattice defects, may also be diagnostic. Trace elements in authigenic clays can also be used to constrain diagenetic water compositions.

Trace elements and rare-earth elements (REE) (elements 57 to 71 in the periodic table) can also be used to better understand Earth surface processes and the operation of the sedimentary cycle. These elements occur in specific concentration ratios in different types of rocks: sediments and sedimentary rocks are depleted in some elements but enriched in others relative to igneous and metamorphic rocks. Also, specific elements are concentrated in different types of sedimentary rocks. This geochemical ‘sorting’ is the result of differences in the chemical properties of elements that govern their behaviour during the processes of weathering, transport, deposition, and diagenesis that make up the sedimentary cycle.

E. Burton

Bibliography

Drever, J. I. (1997) The geochemisty of natural waters. Prentice-Hall, Englewood Cliffs, New Jersey.