Figure 6.2.2 shows a cross-section through a typical reef in a tropical environment (normally between 40° N and 40° S). Reefs tend to form in areas with clear water (e.g., not close to the mouths of large rivers), and near the edges of steep drop-offs because the reef organisms thrive on nutrient-rich upwelling currents. As the reef builds up, it is eroded by waves and currents to produce carbonate sediments that are transported into the steep offshore fore-reef area and the shallower inshore back-reef area. These sediments are dominated by reef-type carbonate fragments of all sizes, including mud. In many such areas, carbonate-rich sediments also accumulate in quiet lagoons, where mud and mollusc-shell fragments predominate (Figure 6.2.3a) or in offshore areas with strong currents, where either foraminifera tests accumulate (Figure 6.2.3b) or calcite crystallizes inorganically to form ooids—spheres of calcite that form in shallow tropical ocean water with strong currents (Figure 6.2.3c).

Figure 6.2.2 Schematic cross-section through a typical tropical reef.

Figure 6.2.3 Carbonate rocks and sediments: (a) mollusc-rich limestone formed in a lagoon area at Ambergris, Belize, (b) foraminifera-rich sediment from a submerged carbonate sandbar in Belize (c) ooids from a beach at Joulters Cay, Bahamas.

Limestone also accumulates in deeper water, from the steady rain of the carbonate shells of tiny organisms that lived near the ocean surface. The lower limit for limestone accumulation is around 4,000 meters. Beneath that depth, calcite is soluble so limestone does not accumulate.

Calcite can also form on land in a number of environments. Tufa forms at springs (Figure 6.2.4) and travertine forms at hot springs. Similar material precipitates within limestone caves to form stalactites, stalagmites, and a wide range of other speleothems. Tufa, travertine and speleothems make up only a tiny proportion of all limestone.

Dolomite (CaMg(CO₃)₂) is another carbonate mineral, but dolomite is also the name for a rock composed of the mineral dolomite (although some geologists use the term dolostone to avoid confusion). Dolomite rock is quite common (there’s a whole Italian mountain range named after it), which is surprising since marine organisms don’t make dolomite. All of the dolomite found in ancient rocks has been formed through magnesium replacing some of the calcium in the calcite in carbonate muds and sands. This process is known as dolomitization, and it is thought to take place where magnesium-rich water percolates through the sediments in carbonate tidal flat environments.

Chert

As we’ve seen, not all marine organisms make their hard parts out of calcite; some, like radiolarians and diatoms, use silica, and when they die their tiny shells (or tests) settle slowly to the bottom where they accumulate as chert. In some cases, chert is deposited along with limestone in the moderately deep ocean, but the two tend to remain separate, so chert beds within limestone are quite common (Figure 6.2.5), as are nodules, like the flint nodules of the Cretaceous chalk of southeastern England. In other situations, and especially in very deep water, chert accumulates on its own, commonly in thin beds.

Figure 6.2.5 Chert (brown layers) interbedded with Triassic Quatsino Fm. limestone on Quadra Island, B.C. All of the layers have been folded, and the chert, being insoluble and harder than limestone, stands out.

Banded iron formation

Banded iron formation (BIF) is a deep sea-floor deposit of iron oxide that is a common ore of iron (Figure 6.2.6). BIF forms when iron dissolved in seawater is oxidized, becomes insoluble, and sinks to the bottom in the same way that silica tests do to form chert. BIF is prevalent in rocks dating from 2400 to 1800 Ma, a result off changes in the atmosphere and oceans that took place over that time period. Photosynthetic bacteria (i.e., cyanobacteria, a.k.a. blue-green algae) consume carbon dioxide from the atmosphere and use solar energy to convert it to oxygen. These bacteria first evolved around 3500 Ma, and for the next billion years, almost all of that free oxygen was used up by chemical and biological processes, but by 2400 Ma free oxygen levels started to increase in the atmosphere and the oceans. Over a period of 600 million years, that oxygen gradually converted soluble ferrous iron (Fe²⁺) to insoluble ferric iron (Fe³⁺), which combined with oxygen to form the mineral hematite (Fe₂O₃), leading to the accumulation of BIFs on the sea floor. After 1800 Ma, little dissolved iron was left in the oceans and the formation of BIF essentially stopped.

Figure 6.2.6 An example of a banded iron formation with dark iron oxide layers interspersed with chert stained red by hematite. This rock is 2.1 billion years old.

Evaporites

In arid regions many lakes and inland seas have no stream outlet and the water that flows into them is removed only by evaporation. Under these conditions, the water becomes increasingly concentrated with dissolved salts, and eventually some of these salts reach saturation levels and start to crystallize. A common environment this takes place is at Death Valley, California (Figure 6.2.7.). Although all evaporite deposits are unique because of differences in the chemistry of the water, in most cases minor amounts of carbonates start to precipitate when the solution is reduced to about 50% of its original volume. Gypsum (CaSO₄·H₂O) precipitates at about 20% of the original volume and halite (NaCl) precipitates at 10%. Other important evaporite minerals include sylvite (KCl) and borax (Na₂B₄O₇·10H₂O).

Figure 6.2.7 An aerial view of the Dante’s Salt Flats in Death Valley National Park, California.

Media Attributions

• Figures 6.2.1, 6.2.2, 6.2.3ab, 6.2.5: © Steven Earle. CC BY.
• Figures 6.2.4. & 6.2.7.: Wikimedia Commons
• Figure 6.2.3c: JoultersCayOoids by Wilson44691. Public domain.
• Figure 6.2.6: © Andre Karwath. CC BY-SA.