Authigenic Sediments

A substantial number of authigenic minerals are precipitated in situ on the sea floor, but only a few common examples will be discussed. Formation of these minerals depends on local geochemical conditions, including elemental abundances, water characteristics, proximity of hydrothermal sources, and rate of sediment accumulation. Precipitation of minerals on or within the sediments of the sea floor generally results from supersaturation of the element or compound required to form the mineral. Supersaturation may occur as the result of change in oxidation state of an element from a soluble, reduced state to a lower solubility oxidized state, resulting in precipitation of a hydrogenous phase, such as iron and manganese crusts. Because authigenic mineral accumulation rates are often less than 1 mm/1000 years, resulting sediments are common only where terrigenous and biogenic accumulation rates are nearly zero. In many cases, crusts of authigenic minerals form where bottom currents prevent the accumulation of other sediments.


Barite (BaSO4) occurs in crystalline or microcrystalline phases or as replacement material in fecal pellets in deep-sea sediments. Barite concentrations average 1% in deep sea sediments, but can make up as much as 10% by weight of the carbonate-free fraction on the East Pacific Rise, where it is associated with hydrogenous iron oxide. Most (80%) of the elemental barite in the oceans enters through rivers, about 20% comes from hydrothermal vents. A major conduit of barium to ocean sediments is secretion by a group of deep-sea protozoans, the xenophyophorans that produce barite crystals in large quantities. Elemental barite is found in biogenic sediments and has been attributed to production by these organisms or by concentration in organic matter following the death of the organism. Deep-sea sediments tend to be richer in barite than slope-depth deposits. Sediment pore waters in the deep sea are saturated with respect to barite; preservation potential is estimated at 30% in oxidized sediment and much lower in anoxic sediments.

In the Pacific, barite is found in radiolarian oozes beneath the equatorial upwelling zone. In the Atlantic, elevated barite concentrations are found on the mid-ocean ridges in areas of low sedimentation rates and where there is an abundance of ferromanganese or iron oxide from hydrothermal sources.


Glauconite is a well-ordered K- and Fe-rich mica-structure clay mineral. It occurs as flakes or pellets, and may occur as infilling in foraminiferal shells and sponge spicules. It may occur in fissures in feldspars, as crusts on phosphorite nodules, and as replacement mineral in coproliths. The color is usually blue-green, but this depends on the original clay-type and chemical composition. For example, dark-green illitic clays alter to dark-green glauconite, while yellowish smectite clays alter to yellowish glauconite. It is usually associated with organic residues, indicating that organic matter plays a role in formation of the mineral. Bacterial activity may promote glauconite formation by producing micro-reducing conditions in the sediment.

Glauconite deposits occur from 65o N to 80o N, but are most common on lower latitude outer shelves and slopes from 20-700 m water depth. Glauconite forms from micaceous minerals or muds of high iron content where sedimentation rates are relatively low. Associated sediments are mainly calcareous, with a high proportion of fecal pellets.

Marine Phosphates

Phosphate concentrations are typically very low within the euphotic zone of the oceans because phytoplanktons extract phosphate nutrients to photosynthesize organic matter. Vertebrates also concentrate phosphate into apatite, from which their bones are constructed. Vertically migrating fish and invertebrates feed on phytoplankton and zooplankton in surface waters at night and retreat to the shelter of darker subsurface waters during the day. Excretion of wastes in subsurface waters, along with decay of organic matter settling through the water column, concentrates inorganic phosphate ions and compounds below the euphotic zone, especially within thermocline depths. Both organic matter and skeletal remains accumulate on the sea floor, where decay and dissolution return phosphate to solution in bottom waters. Where the seafloor is at thermocline depths, especially beneath upwelling surface waters, that promotes export of organic matter to the bottom and phosphate ions may become sufficiently concentrated to precipitate phosphatic nodules or crusts.

The most important phosphatic mineral is microcrystalline carbonate fluorapatite. Phosphatic nodules and crusts typically form along continental shelves, upper continental slopes and on oceanic plateaus beneath upwelling surface waters and where bottom currents limit accumulation of detrital sediments. Typical areas of phosphatic deposition are the continental margins of Peru, Chile, and southwest and northwest Africa. Phosphorite nodules or crusts average 18% phosphate. Conglomerates of phosphatized limestone pebbles and megafossils in a matrix of glauconite may have up to 15% phosphate.

Marine phosphates and phosphorite deposits are also found associated with anoxic sediments. Phosphorite may form by replacement of carbonate by phosphate. Upwelling occurs in the southern Caribbean in the surface waters above the Cariaco Basin, resulting in export of organic matter to bottom sediments. Phosphate precipitation is occurring along the rim of the basin where anoxic water from the trench mix with oxygenated waters from above. Phosphate may also be adsorbed by hydrous iron minerals, aluminum oxides and clay minerals. This accounts for phosphate concentrations of 1-2% in some iron-rich, clay or zeolite sediments in the deep sea.

Heavy Metals

Iron oxides are an important constituent in slowly accumulating deep-sea clays where they occur as amorphous or poorly crystalline reddish-brown coatings on clays and other minerals and as minute globules in the sediments. Iron-rich basal deposits are found in oxidizing environments on the crests and flanks of actively spreading ocean ridges Figure 8.23 . Here, brownish-stained carbonate oozes may contain up to 14% Fe2O3. Iron-manganese minerals in these sediments are commonly attributed to hydrothermal activity associated with ocean-ridge volcanism. These associations result from penetration of seawater into hot volcanic rock, where the seawater is heated and becomes acidic and reducing, by geochemically reacting with fresh lava. As the hot solution mixes with cold seawater, sulfides precipitate first. With further mixing, iron and manganous oxides precipitate, producing iron-rich basal sediments.

As seawater percolates into hot, volcanic rocks, seawater sulfate reacts with reduced iron. Where the hot solutions are forcibly expelled from the rocks (vents and fumeroles), metal sulfides precipitate as crusts and chimneys up to several meters high ridges Figure 8.24 . Localized accumulation rates can be a meter per year. Deposits rich in Fe, Mn, Cu, and Zn can occur where there is hydrothermal activity on the sea floor. One of the most spectacular examples of ridge-crest metalliferous deposits was discovered in the Red Sea in 1963. Rather than localized vents, metals are concentrated in deep, brine-filled basins. Manganese micronodules (less than 1 cm in diameter), nodules (1-10 cm in diameter) and crusts or coatings form in sediments or on exposed hard surfaces in the deep sea ridges Figure 8.25 . These oxides are brown-black agglomerations of manganese and iron oxides in fine-grained silicates or iron oxide-rich groundmasses in detrital and biogenic grains. Accessory metals include Ni, Cu, K. Ca, and Co. Elemental distribution patterns within nodules are variable and depend both on the environment of deposition and the nature of the mineral phases they contain. Where redox potential is lower, nodules are more iron rich; in well-oxidized deep-sea settings, nodules are richer in Mn.

A nodule commonly forms around a nucleus such as a shark's tooth or volcanic fragment. Nodules grow in concentric layers that may represent changes in seawater composition during growth. Rates of nodule growth are 1-4 mm/106 years. They commonly occur where sedimentation rates are less than 5 mm/1000 years. Apparently, sporadic movement by benthic organisms burrowing through the sediments is sufficient to keep most nodules at the sediment surface, where they can grow. The greatest area of manganese nodule development occurs in the Pacific, where 75% of the equatorial and North Pacific deep sea floor is covered with nodule patches Figure 8.26 . Fields of nodules develop in areas swept clean of fine detrital sediments by bottom currents. Where nodules cover 100% of the sediment surface, the area is called a manganese nodule pavement. In some cases, nodules join to form a solid surface. Such pavements are found on deep plateaus including the Blake Plateau in the western North Atlantic and the Agulas Plateau south of South Africa.

The manganese comes from terrestrial sources by wind and water transport. In the water column, plankton extract manganese from solution, then carry it to the bottom. Manganese is also scavenged from seawater and deposited on the bottom by organic aggregates. Local deep-water sources of manganese may be interstitial waters leaching sediments rich in Mn and Fe near basaltic rocks. Near mid-ocean ridges, nodules may derive their Fe, Mn and accessory minerals from volcanic sources, as noted above.

: Although there is economic interest in both metalliferous sulfide deposits and in manganese nodules, the costs of mining currently exceed the value of the minerals.

Organic-Rich Sediments

Organic material is measured in sediment as total organic carbon (TOC) or particulate organic matter (POC) which in ocean water is primarily living organisms or the remains of dead organisms. Upon the death of an organism, its remains are subjected to chemical and bacterial degradation processes. Detrital POC, which is produced in surface waters by primary production, may sink through the water column as fecal pellets or as marine snow and flocculate into what is called the fluffy layer. Skeletal remains, including coccoliths, diatom frustules, foraminiferal tests and radiolarian skeletons, as well as clay particles and volcanic ash, sink along with the organic matter. Both organic and inorganic particles influence to some degree the water chemistry of the waters they pass through. Within the water column, organic matter provides food for filter-feeding animals, which remove usable compounds and package unusable materials, including inorganic debris, into fecal pellets. The greater size and density of these pellets greatly increases settling rates of this material.

When the organic matter reaches the sea floor, it provides food for benthic filter-feeding and detritus feeding organisms, reducing the concentration of POC accumulating in the sediments relative to what reaches the sea floor. In the Panama Basin, which is an upwelling area, depth-stratified sediment trap studies indicate that approximately 5% of the particulate matter reaching the bottom are POC, yet TOC concentrations in the sediments are less than 2%. Utilizable organic matter is known as labile organic matter. The least degradable materials, which often include terrestrial cellulose brought to the deep ocean in gravity flows, are called refractory solid organic matter. In typical pelagic sediments, TOC concentrations are less than 1%.

Most organic carbon in sediments accumulates under conditions of high primary productivity in surface waters and low oxygen in bottom waters or interstitial pore waters. As a result of coastal upwelling and runoff from land that provide nutrients to phytoplankton communities in surface waters, combined with relatively rapid sedimentation rates in these regions, roughly 50% of all organic carbon burial occurs on continental shelves and margins.

Organic-rich sediments that accumulate where bottom waters are depleted of oxygen (anoxic) are called sapropels. Anoxic conditions develop either because of rapid influx of POC or because of stagnation of bottom waters. Though limited in extent in modern oceans, sapropels occur in a variety of settings, including semi-isolated basins with restricted bottom circulation and portions of continental margins or slopes that lie within the mid-water oxygen minimum zone and below upwelling zones.

Late Quaternary deep-water sediments in the Black Sea provide an example of restricted bottom circulation under which sapropels (ooze or sludge rich in organic matter) formed. From 23,000 to about 9,000 years ago, when sea level was 40 m or more lower than today, the Black Sea was completely isolated from the Mediterranean and was a large, freshwater lake which was aerobic thoughout. As sea level rose following the last glacial advance, seawater began to occasionally spill over the Bosphorus Sill into the Black Sea, filling the deeper parts of the basin with dense seawater. However, river runoff into the Black Sea kept surface waters fresh. Because of higher evaporation rates in the Mediterranean, most of the flow of water through the Bosphorus was freshwater from the Black Sea to the Mediterranean. The seawater filling the basin of the Black Sea was isolated from air beneath a layer of low density fresh water. Primary productivity in the surface waters rained organic matter into the deep waters, depleting all oxygen, so that by 7,000 years ago, anoxic conditions were fully developed. About 3,000 years ago, two-way circulation developed with the Mediterranean, driving turnover of the deep waters of the Black Sea and allowing deep sea marine faunas to become established.

Examples of modern sapropel formation within the oxygen minimum zone beneath upwelling high productivity surface waters can be found on the continental slope of the Arabian Peninsula and in the California borderlands. Upwelling in the northwest Indian Ocean provides sufficient surface productivity to provide an excess of organic matter to sediments on the continental slope of the Arabian Peninsula where the oxygen minimum zone intersects the slope. Off California, the combined effects of sluggish circulation in semi-isolated basins, continental margin depths within the oxygen minimum zone, and high surface water productivity all contribute to accumulation of laminated, organic-rich sediments in the Santa Barbara basin.

Anoxic sediments have been widespread in the past and are of great economic importance as source rocks for hydrocarbon deposits. Expansion and intensification of the oceanic oxygen minimum zone, probably during times of reduced thermohaline circulation, is one mechanism that seems to account for many sapropels. Deep basins connected only by shallow connections, which resulted in restricted bottom circulation, were especially common during early stages of continental rifting that formed the Atlantic basins Figure 8.27.