Box 65 Abiological precipitation of calcium carbonate

Where a skeletal source cannot be identified, calcium carbonate (CaCO3) grains and finegrained muds may be of abiological origin. The most famous occurrences occur in shallow, warm, saline waters of the Bahamas and the Arabian Gulf. In these areas two distinctive morphologies are present, ooids and needle muds (Fig. 1).

Ooids are formed by aggregation of aragonite* crystals around a nucleus, usually a shell fragment or pellet. Successive layers of aragonite precipitation build up a concentric structure, which may vary in size from about 0.2 to 2.0 mm in diameter. Needle muds are also aragonitic; typically each needle is a few micrometres wide and tens of micrometres long.

It has long been thought that the warm,

Salinity Soil Magnetized Waters

Fig. 1 (a) Ooid-rich sediment from the Great Bahama Bank. Individual grains are typically 1 mm in diameter (photograph courtesy of J. Andrews). (b) Scanning electron microscope photograph of aragonitic needles from the Great Bahama Bank. Scale bar = 1 mm (photograph courtesy of I.G. Macintyre, from Macintyre and Reid 1992).

Fig. 1 (a) Ooid-rich sediment from the Great Bahama Bank. Individual grains are typically 1 mm in diameter (photograph courtesy of J. Andrews). (b) Scanning electron microscope photograph of aragonitic needles from the Great Bahama Bank. Scale bar = 1 mm (photograph courtesy of I.G. Macintyre, from Macintyre and Reid 1992).

(continued)

shallow, saline waters where these deposits are found favour increased concentrations of carbonate ions (CO|-), increasing the ion activity product of aCa2+.aCO|- such that precipitation of CaCO3 occurs. The formation of ooids probably requires fairly agitated, wind- or wave-stirred waters, allowing periodic suspension of the grains into the CaCO3-saturated water, whereas aragonitic needles may precipitate as clouds of suspended particles, known as whitings.

There has been, and still is, much debate about the origins of these particles. Firstly, it is difficult to disprove the effects of microbial mediation in their formation. Thus we might regard the grains as non-skeletal, while accepting a possible microbial influence. Secondly, various geochemical and mineralogical studies have produced equivocal results in attempting to demonstrate an abiological origin. Having said this, recent work based on crystal morphology and strontium substitution lends support for inorganic precipitation of needle muds.

Despite the considerable interest these phenomena provoke, we should remember that they are of minor significance to the modern oceanic CaCO3 budget. The relative importance of inorganic CaCO3 in the geological past is more difficult to assess, but may have been more significant before the evolution of shelly organisms about 570 million years ago.

*Aragonite and calcite are known as polymorphs of CaCO3. Both minerals have the formula CaCO3 but they differ slightly in the structural arrangement of atoms.

Clay Distribution Present Day
Fig. 6.9 The present-day distribution of the principal types of marine sediments. After Davies and Gorsline (1976), with kind permission from R. Chester.
Coccosphere Emiliania Huxleyi

Fig. 6.10 (a) Planktonic coccolithophore Emiliania huxleyi, a very common species in the modern oceans. This specimen has a diameter of 8 mm. The skeleton is clearly composed of circular shields packed around a single algal cell. After death the coccosphere breaks down, releasing the shields to form the microscopic particles of deep-sea oozes and chalks (photograph courtesy of D. Harbour). (b) Skeleton of modern planktonic foraminifer Globigerinoides sacculifer, common in tropical oceans. Scale bar = 50 mm (photograph courtesy of B. Funnell).

Fig. 6.10 (a) Planktonic coccolithophore Emiliania huxleyi, a very common species in the modern oceans. This specimen has a diameter of 8 mm. The skeleton is clearly composed of circular shields packed around a single algal cell. After death the coccosphere breaks down, releasing the shields to form the microscopic particles of deep-sea oozes and chalks (photograph courtesy of D. Harbour). (b) Skeleton of modern planktonic foraminifer Globigerinoides sacculifer, common in tropical oceans. Scale bar = 50 mm (photograph courtesy of B. Funnell).

degree of CaCO3 undersaturation in the deep waters (Fig. 6.11). In the Atlantic Ocean the CCD is at about 4.5 km depth; above the CCD, at about 4km depth in the Atlantic, there is a critical depth, known as the lysocline (Fig. 6.11). Here the rate of calcite dissolution increases markedly and all but the most robust particles dissolve rapidly. It is estimated that about 80% of the CaCO3 settling into deep waters is dissolved, either during transit through the water column or on the seabed. As a consequence, pelagic carbonate deposits are most common on the shallower parts of the deep ocean floors (Fig. 6.11) or on topographic highs that project above the CCD.

Planktonic coccolithophores and foraminifera did not evolve until the mid-Mesozoic (about 150 million years ago), whereas shallow-water shelly organisms are known to have existed throughout Phanerozoic time (570 million years to the present day). This means that the locus of carbonate deposition has shifted to the deep oceans only in the last quarter of Phanerozoic time.

The removal of Ca2+ by CaCO3 precipitation can be estimated directly from the abundance of CaCO3-rich ocean sediments and their sedimentation rates (Table 6.2). From equation 6.4, we see that two moles of HCO3 are removed with each mole of Ca2+, a process that releases dissolved CO2 into seawater, ultimately to be returned to the atmosphere. Calcium carbonate also incorporates a small but significant amount of Mg2+ by isomorphous substitution for Ca2+ (see Box 4.6) and this is used to derive the Mg2+ removal in Table 6.2.

Degree of saturation (Q) for calcite

Under-

saturated Supersaturated

Fig. 6.11 Schematic diagram showing depth relationship between degree of saturation for calcite in seawater and rate of CaCO3 dissolution. At 4 km depth, as seawater approaches undersaturation with respect to calcite, rate of dissolution of sinking calcite skeletons increases. The lysocline marks this increased rate of dissolution. Below the lysocline only large grains (foraminifera) survive dissolution if buried in the seabed sediment. Below the calcite compensation depth (CCD; see text) all CaCO3 dissolves, leaving red clays.

Modern surface seawater is demonstrably supersaturated with respect to CaCO3, and the presence of carbonate sediments in rocks of all ages suggests that it has been throughout much of Earth history. Oceanic pH is unlikely to have fallen below 6 as such a shift requires a 1000-fold increase in atmospheric pCO2 over its present value of 10-36atm. An atmospheric pCO2 of this magnitude may have occurred very early in Earth history, but the existence of 3.8 billion-year-old CaCO3 sediments shows that seawater was still close to saturation with calcite. Achieving calcite saturation at these high pCO2 values implies that oceanic Ca2+concentrations were at that time 10 times larger than modern values. Seawater pH has probably never exceeded 9 in the geological past, because at that pH sodium carbonate would be a much more common mineral precipitate than calcite. There is no evidence in ancient marine sediments that sodium carbonate has ever been a common marine precipitate.

6.4.5 Opaline silica

Opaline silica (opal) is a form of biologically produced silicon dioxide (SiO2.nH2O) secreted as skeletal material by pelagic phytoplankton (diatoms) and one group of pelagic zooplankton (radiolarians) (Fig. 6.12). Opaline silica-rich sediments cover about one-third of the seabed, mainly in areas where sedimen-

Under-

saturated Supersaturated

Fig. 6.11 Schematic diagram showing depth relationship between degree of saturation for calcite in seawater and rate of CaCO3 dissolution. At 4 km depth, as seawater approaches undersaturation with respect to calcite, rate of dissolution of sinking calcite skeletons increases. The lysocline marks this increased rate of dissolution. Below the lysocline only large grains (foraminifera) survive dissolution if buried in the seabed sediment. Below the calcite compensation depth (CCD; see text) all CaCO3 dissolves, leaving red clays.

Fig. 6.12 (a) Skeleton of siliceous radiolarian Theocorythium vetulum, early Pleistocene, equatorial Pacific. Scale bar = 75 mm. (b) Siliceous diatom Coscinodiscus radiatus, early Pleistocene, equatorial Pacific. Scale bar = 38 mm. Photographs courtesy of B. Funnell.

tation rates are high, associated with nutrient-rich upwelling waters and polar seas, particularly around Antarctica (Fig. 6.9). Seawater is undersaturated with respect to silica and it is estimated that 95% of opaline silica dissolves as it sinks through the water column or at the sediment/water interface. Thus, the preservation of opaline silica only occurs where it is buried in rapidly accumulating sediment, beneath the sediment/water interface. Subsequent dissolution of opal in the sediment saturates sediment pore waters with silica. The pore water cannot readily exchange with open seawater and saturation prevents further opal dissolution. High sedimentation rates in the oceans can be caused by high mineral supply rates from the continents, but are usually caused by high production rates of biological particles (Section 6.5.4). In high productivity areas, for example parts of the Southern Ocean bordering Antarctica, diatoms are the common phyto-plankton species, and this enhances the importance of these regions as silica sinks. The biological removal of silicon (Si) from seawater is calculated from the opal content of sediments and rates of sedimentation (Table 6.2).

6.4.6 Sulphides

The oxidation of organic matter proceeds by a number of microbially mediated reactions once free oxygen has been used up (see Section 5.5 & Table 4.7). Although small amounts of nitrate (NO-), manganese (Mn) and iron (Fe) are available as electron acceptors in marine sediments, their importance is small in comparison with SO4-, which is abundant in seawater (Table 6.1). At seawater pH around 8, sulphate-reducing bacteria metabolize organic matter according to the following simplified equation.

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