The transformation of soft sediment to hammer-ringing rock is driven by physical and chemical diagenesis during burial.
Diagenesis of sediment is the inevitable consequence of its burial. The diagenesis of any mineral, whether it be precipitation, dissolution, or recrystallization does not take place in isolation but in concert with other minerals and aqueous fluids that may be reacting in very different ways. During burial, intensive variables like temperature, pressure, and concentration (i.e. independent of the size of the system under investigation) and extensive variables like volume and entropy, change continually. How these variables change determines whether solid minerals and dissolved aqueous species are thermodynamically stable or involved in water-rock reactions.
Diagenetic systems are physically and chemically complex.
The initial reactants subjected to burial diagenesis are the solid detrital components plus any authigenic minerals that form at the surface or the sediment -water interface (e.g., calcite, aragonite, amorphous silica, glauconite, iron oxides). Water is the dominant fluid phase in all diagenetic reactions. Initial fluid compositions range from relatively fresh to saline water (seawater, closed basin brines). Detrital components that are unstable with respect to the ambient fluid composition will tend to react. Different water-rock interactions will take place as the sedimentary pile thickens, with concomitant increases in temperature and pressure at depth. Detrital and authigenic minerals that were relatively stable before burial, may become unstable as burial proceeds resulting in dissolution, precipitation of new phases, or recrystallize (e.g. neomorphism), with concomitant changes in fluid composition.
The important variables and processes involved in burial diagenesis are outlined in the following notes. Keep in mind that none of these variables acts in isolation – they are interdependent.
The pore pressure – depth curve typical of the Gulf Coast Basin shows the transition from hydrostatic conditions to overpressures approaching lithostatic values at 2.5 – 4 km, coinciding approximately with abundant quartz and clay cementation, and carbonate neomorphism. Modified from Bethke, 1986 – PDF available.
Temperature: Temperatures generally increase with depth in Earth’s crust where values are determined by the geothermal gradient (degrees/kilometre). The gradient in continental crust commonly ranges from 25o–30o/km; oceanic crust gradients are typically in the range 60o-100o C/km. The important consequences of increasing temperature with depth include:
- An increase in chemical reaction rates; a rule-of-thumb states that reaction rates double for every 10o C increase in temperature.
- Most common minerals become increasingly soluble with temperature. This is commonly reflected in the observed increase in total dissolved solids (TDS) with depth. A notable exception is calcite, where the solubility in aqueous solution decreases at temperatures above 25o-30o C.
- Organic matter begins to break down into its hydrocarbon components at about 60o C, becoming more rapid and sustained above 80o C. A typical Oil and Gas maturation temperature window is 80o – 120o C. The production of organic acids during these reactions can have a profound impact on pH, pH buffering capacity, and the stability of minerals, particularly the carbonate suite of minerals.
Pressure
Pressure
The increase in pressure with depth is conveniently expressed in two components: lithostatic pressure and hydrostatic pressure. The important consequences of increasing pressure include:
- Compaction of sediment begins almost immediately following deposition, driven by lithostatic pressure. Compaction continues through the entire thickness of a sedimentary basin fill. The decrease in sediment-rock volume results in the expulsion of fluid – compaction is an important driver of subsurface fluid flow (but not the only driver) and mass transport of dissolved solids.
- Compaction results in textural changes (e.g., grain packing) and importantly, decreased porosity and permeability.
- Pressure solution at grain-to-grain and crystal-to-crystal contacts in carbonate rocks, commonly manifested as stylolites, but can also occur in silicates like quartz. Solution depends on the existence of pressure and concentration gradients across the grain contact, and a mechanism to remove dissolved solids beyond the contact via mass transfer – principally by diffusion.
Porosity-permeability
Porosity and permeability are fundamental textural properties of all sediment. The reduction of both during mechanical compaction is well documented, summarized in the porosity-depth curves above. However, porosity is also reduced by cementation at all stages of burial – this is particularly the case for carbonates where authigenic calcite and aragonite cements can form at the sea floor. The precipitation of authigenic quartz and clay cements is also common in siliciclastic deposits although this kind of porosity occlusion tends to become more pervasive at depths greater than 1-2 km depending on burial temperatures and fluid compositions.
Mudrocks usually have very high initial porosities (commonly >60%) but early compaction at shallow burial depths reduces these values to just a few percent. Initial water loss from mudstone-shale compaction also occurs at much shallower depths than is generally observed for arenites and carbonates. Dehydration and physical compaction of shales decreases significantly with depth because of decreasing permeability. However, diagenetic dehydration of smectites below about 2-3 km burial will release crystal lattice-bound water during the smectite-illite transformation, water that is expelled by diffusive flow. Compaction disequilibrium develops where anomalous fluid pressures arise when expelled water cannot escape because of permeability barriers.
Fracture porosity is important in impermeable rock subjected to tensile stresses that form joints or fractures. Fracture networks that have a high degree of connectivity provide pathways for mass transport of aqueous and non-aqueous fluids.
Fluid flow
Interstitial fluids in buried sediment and rock are rarely static. Two principal types of fluid flow are recognized: advective flow that moves fluids en masse through porous and permeable media, and diffusive flow.
- Topography-driven advective flow is forced by gravitational potential energy. Topographically driven groundwater can reach depths of 3000-4000 m in sedimentary basins that border topographically elevated terrains, such as those associated with foreland basin or accretionary prism deformed belts. Advective flow also applies to the mass transport of fluid and dissolved solids (a diagrammatic representation of meteoric flow systems is shown below).
- Tectonically-driven advective flow: The role of elevated pore pressures in promoting faulting, particularly thrust faulting, is well established (Hubbert and Rubey, 1959). The elastic response to tectonically derived compressional stress results in compaction and fluid loss. If escaping fluids are trapped by permeability barriers, then pore pressures will rise above hydrostatic values. Hydrostatic pressures that approach lithostatic values have been recorded from thrust fault systems, like those generated in accretionary prisms. The faults themselves may act as conduits for fluid flow.
- Fluid flow, or mass transfer at the scale of individual pores, pore throats, and grain boundaries is dominated by diffusion, of which two main types are recognized: Mechanical diffusion where the flow of fluid and dissolved solids occurs as small eddies around grains or through fractures. Molecular diffusion occurs in response to concentration and electrical charge gradients in dissolved ionic species. Both diffusive processes are responsible for the dispersion of dissolved solids beyond the limits of advective flow.
- Smectite dehydration during burial will also provide a source of water via mass transfer through shale (primarily diffusion).
- Rates of diffusive flow (mass transfer) are orders of magnitude lower than for advective flow (mass transport).
- Large-scale temperature- and density-driven convection is an important process in many surface water-bodies but is probably less common in sedimentary basin successions because of multiple permeability barriers (e.g., mudrocks, salt deposits).
Fluid composition
Initial fluid compositions range from fresh water to hypersaline, depending on the basin setting and depositional environment. Ion activities, pH, and REDOX conditions change during burial in response to increasing temperature, pressure, and evolving water-rock interactions. These changes are commonly represented as total dissolved solids (TDS) that measure the summed concentration of aqueous ions and soluble organic compounds; TDS is also an approximation of salinity. Charged ions commonly included in TDS analyses are calcium, magnesium, sodium, potassium, bicarbonate, sulfate, chloride, nitrate, and silica. The characterization of water types is commonly based on TDS. TDS values generally increase with burial depth.
The reactions responsible for these compositional effects typically involve precipitation, dissolution, oxidation and reduction (e.g., Fe3+ to Fe2+, SO42- to S2−), ion exchange (e.g., clays, zeolites), dehydration (e.g., smectite), and transitions from disordered crystal lattices to ordered structures (e.g., dolomite, smectite-illite transformations). Individual ion activities will change according to whether they are consumed or released from a reaction. For example, K+ that evolves from the dissolution of potassium feldspars is available for illite precipitation. The concomitant production of SiO42- is available for any number of silicate reactions involving clays, authigenic albite, and quartz overgrowths. Ca2+ produced by dissolution of gypsum-anhydrite is available for calcite or dolomite precipitation. The transformation of organic matter produces soluble organic acids that can significantly alter carbonate equilibria.
Changing fluid chemistry can also be represented by comparing various ion ratios that respond to reactions involving precipitation and dissolution. The average Mg/Ca ratio in seawater (concentrations expressed in mg/L) is about 3. The decrease in Mg/Ca as salinity increases is a response to reactions like the precipitation of dolomite or the inclusion of Mg2+ in smectite. Likewise, SO42- decreases because of sulphur reduction (to sulphides) but could also be due to gypsum precipitation – the latter would also imply a decrease in Ca2+ which my limit the Mg/Ca trend.
Bulk fluid compositions at depth can also change because of the ingress of topography-driven meteoric groundwater. Groundwater composition, even at shallow depths, will reflect the different kinds of bedrock through which it flows. Long residence periods at increasing depths of meteoric ingress will tend to increase dissolved solid concentrations. A hypothetical example of evolving meteoric water compositions is shown in the Piper plot below. Mixing of relatively fresh meteoric and residual basin fluids can also reduce the average model age for fluids – this is the reported case in Alberta Basin where meteoric waters that infiltrated 2-3 km deep, flushed and diluted older brines (Gupta et al., 2015; OA).
Equilibria and chemical reactions
Chemical reactions are driven by energy transfers from reactants to products. The concept of equilibrium applies to reversible reactions; at equilibrium, the energy transferred in a forward reaction equals that for the reverse reaction. Equilibria are statements of energy balance; they are not concerned with how reactants are transformed to products (i.e. reaction paths). However, they do require knowledge of reactant-product compositions. This is straight forward for relatively ‘simple’ minerals like calcite and quartz. But for complex minerals like clays, and minerals that are part of solid solution series (e.g., feldspars), establishing accurate compositions of mineral and aqueous species at the time of reaction is difficult.
Dissolution or precipitation of a mineral also depends on its solubility at a given temperature, and the concentration (or activity) of ion species in relation to the solubility product. Consider the reaction:
CaCO3(solid) ↔ Ca2+(aq) + CO32-(aq)
The solubility (activity) product at equilibrium is Ksp = (aCa2+).(a CO32-) – the activity of the solid phase is 1. At equilibrium the aqueous solution is saturated with respect to the solid phase. If the activity of the products increases, the solution is supersaturated, and the reaction is pushed to the left – calcite will precipitate until equilibrium is reestablished. If the opposite occurs, the solution becomes undersaturated and calcite will dissolve. Changes in the degree of saturation, and therefore mineral stability during burial are strongly influenced by companion water-rock interactions that alter solution pH or change the concentration (or activity) of Ca2+ and CO32-, such as the dissolution of gypsum and calcic plagioclase, or the precipitation of dolomite.
Reference not linked
Carpenter, A. B. (1978) Origin and chemical evolution of brines in sedimentary basins. In 13th Industrial Minerals Forum (eds. K. S. Johnson and J. R. Russell): Oklahoma Geological Survey Circular, v. 79, p. 60-70.