Geol 135 Sedimentation

J Bret Bennington

Updated 10/99


Diagenesis – postdepositional modification of sedimentary layers by both physical and chemical processes. Diagenesis grades into metamorphism, the boundary between the two defined by a temperature – pressure line. Generally, diagenetic changes are those that occur below 200° C and 25 km of depth.

Early post-depositional structures:

Soft sediment deformation

  1. Dewatering structures – caused by the expulsion of water due to loading or shock. Includes dish and pillar structures and dewatering pipes.
  2. Slumped beds – caused by downslope movement of layers deposited on a shallow slope, such as along the front of a prograding delta. These can be recognized because they are locally folded, but are bounded above and below by horizontal beds.
  3. Load structures – caused by the deposition of relatively dense sand over watery mud, wherein the sand partially sinks into the underlying mud. Mud may also be forced upward into the sand to form a flame structure. Spherical loads of sand sunken into underlying mud are often called ball and pillow structures.
  4. Diapirs – deposition of thick layers of sand over watery muds can cause the mud to rise upward through the sand due to the density contrast. The mud diapir may become detached from its source and rise to the surface of the sands. A similar phenomenon occurs with layers of salt deposited as evaporites. During compaction, the weight of overlying sediments can compress the salt, causing it to concentrate in one area and flow upward, disrupting the overlying sediments.

Lithification – the process of turning unconsolidated sediment into rock.

The timetable for lithification is highly variable and depends on type of sediment, depth of burial, and chemical environment after burial. Some sediments lithify as they are being deposited, others remain unlithified for millions of years after burial (the Cretaceous sediments of the New Jersey coastal plain, for example.)

Compaction – the first stage of burial, caused by overburden pressure.

  1. Clasts rearrange and water is expelled. Sands lose about 10% volume by compaction, muds may compact up to 50%. Early cemented carbonates my compact little, if at all.
  2. Pressure solution – high pressure at grain-grain contacts causes mineral to dissolve and re-precipitate away from contacts. This produces a fabric of progressively more tightly sutured grains.
  3. Grain deformation / pulverization – weaker grains such as lithic fragments and micas are deformed between stronger grains such as quartz sand.

Cementation – ions dissolved in one place due to pressure solution or local ph conditions can migrate with pore waters and be precipitated as cement. In marine environments the most common cement is calcium carbonate, which can be found in both clastic and carbonate marine rocks. Silica also can form cement, but commonly only in the absence of calcium carbonate (more common in terrestrial sediments).

Sparry cement – large crystals which grow out into pore spaces or envelop grains.

Micrite cement – microscopic crystals of cement.

Cements that are the same minerals as the grains commonly form as overgrowths, extending the original grain, but leaving a recognizable outline of the original grain boundary. In marine organisms that form single crystal skeletal elements (eg. Echinoderms) the cement extension of the grain is called a syntaxial overgrowth.

Cementation reduces both porosity and permeability.


Nodules and concretions

Pyrite – indicative of sulphate and organics, often occur in marine sediments

Siderite – indicative of absence of sulphate, presence of organics, associated with freshwater

Carbonate – common in marine shales, silts, and sands

Flint / chert – common in carbonates, source of silica may be sponge spicules, radiolarians, or diatoms.

Concretions often form early in diagenesis, shortly after compaction. Evidence for this is seen in the lack of compaction in layers and fossils preserved within the concretion relative to the layers outside of the concretion.

Siderite and Pyrite

siderite (Fe2CO3) and pyrite (FeS2)

The formation of stable pyrite or siderite requires an anoxic sedimentary environment where reduced iron (Fe2+) can exist in solution. Both pyrite and siderite are unstable in oxidizing environments. In the presence of O2, pyrite and siderite are oxidized to goethite or hematite (the deep red outer coloration of exposed siderite nodules suggests that hematite is the final weathering product).

Dissolved iron is present in both marine and fresh waters and is derived from reduction of fine-grained detrital iron oxide minerals by organic material in the sediment. O2 is rapidly depleted upon deposition in sediments rich in organic carbon, due to exhaustive aerobic decay. Even if the overlying water is oxygenated, aerobic decay within the upper few millimeters of organic-rich sediment will maintain anoxic reducing conditions

Given anoxic conditions and reduced iron, the primary factor that determines whether pyrite or siderite will form is the presence of sulfide (H2S and HS-). Pyrite is the end product of a series of reactions that first combine reduced iron and sulfide to form metastable monosulfide iron minerals, which then subsequently react with aqueous sulfur species to form the disulfide. The overall reaction is given by:

3H2S + S0 + Fe2O3 = 2FeS2 + 3H20 (Berner, 1981)

Sulfides are produced from sulfate (SO4=) during bacterial decomposition of organic matter. Marine waters are enriched in sulfate relative to fresh waters by a factor of 100 (Berner, 1971), so that marine sediments rich in organics produce abundant sulfide, while freshwater sediments are sulfide poor. Organic decay also provides some sulfate to the sedimentary environment, but this source is negligible relative to the quantity of sulfate available from interstitial marine water and from the diffusion into the sediment of sulfate from the essentially infinite source in the overlying water column. Marine sedimentary environments are thus termed "sulfidic environments" in Berner's classification, with pyrite being the characteristic early diagenetic mineral.

Siderite forms through the combined effects of iron reduction and bacterial methanogenesis of organic carbon compounds by the following overall reaction:

7CH2O + 2Fe2O3 = 3CH4 + 4FeCO3 + H2O (Curtis, 1986)

As with pyrite, the source of iron is the reduction of detrital iron oxides in a strongly reducing, organic-rich sedimentary environment. Siderite is inhibited from forming in marine environments because Ca++ reacts preferentially with bicarbonate at normal marine concentrations. The Fe++/Ca++ ratio in normal marine waters is two orders of magnitude too small to permit siderite precipitation (Matsumoto, 1981). In a reducing marine sedimentary environment, Fe++ is prevented from building up to a concentration that would permit siderite formation by its reaction with sulfides to form pyrite. Therefore, only in a "nonsulfidic environment" can Fe++ be precipitated as siderite (Berner, 1981).

The most likely way to achieve a nonsulfidic environment is to exhaustively react all available sulfide in the sediment pore waters. This is most easily accomplished in sulfate-poor, non-marine sedimentary environments. Organic-rich, non-marine sedimentary environments harbor optimal conditions for siderite formation. The scarcity of sulfate (and therefore sulfide) in non-marine water limits sulfate reduction of organic matter and pyrite formation. Reduced iron accumulates in solution as methanogenesis produces abundant bicarbonate ion that combines with Fe++ to produce siderite. Attainment of this "methanic environment" during the early stages of diagenesis suggests that the initial depositional environment was sulfate-poor and, therefore, non-marine (Berner, 1981).


Rock Color

Potter et al. (1980 ) have shown that color in mudrock is influenced almost exclusively by two parameters, the ratio of Fe3+ to Fe2+ and the percent organic carbon in the rock. In the absence of organic carbon, mudrock color varies from red to greenish gray as Fe3+ is progressively reduced to Fe2+. The addition of increasingly greater amounts of organic carbon darkens the mudrock from gray to black. These two parameters are not independent because the presence of any appreciable organic carbon in the sediment virtually assures that all interstitial oxygen will be consumed, leaving undecomposed organic matter that reduces Fe3+ to Fe2+ (Berner, 1981).

Red (oxidized iron) ----------------Green (reduced iron)

1.5% - dark gray

1% - dark gray to medium dark gray

5% - medium dark gray

4% - medium dark gray to medium gray

.3% - medium gray

.2% - medium gray to light gray

.1% - light gray

The amount of organic carbon preserved in sediments is related both to its rate of input into the sediment as well as its survival once deposited. Organic matter is decomposed in the upper layers of sediment, first by aerobic microrganisms, and then by anaerobic sulfate reducing and fermenting bacteria (Drever, 1988). Aerobic decay is inhibited by dysaerobic or anaerobic conditions at or beneath the sediment-water interface. Anaerobic decay of organic matter can be reduced by increasing the rate at which the organic matter is buried to a depth below the near-surface zone of intense bacterial decomposition. Johnson Ibach (1982) has shown that preservation of organic matter first increases with increasing rate of clastic deposition, and then decreases as the effect of diluting the available organic carbon with inorganic sediments overwhelms the enhanced preservation of organics due to rapid burial.