where R is a characteristic size of sediment particles. Using this approach, we can identify the following cases: (1) when (f> < 1 the time scale for diffusion is short compared with that for chemical reaction and reactivity will determine the rate of change of concentration; and (2) for (j> > 1 the time scale for diffusion is long and diffusion will play an important role in determining the reaction rate.
Finally, it must be stressed that diffusion of dissolved species in solutions is a key physico-chemical process for the sea/sediment interaction and energy exchange at the sediment-water interface. The reader is referred to Cussler (1984) for a more comprehensive presentation of diffusion in fluid systems.
Reducing and oxidizing conditions in a sediment determine the chemical stability of the solid compounds and the direction of spontaneous reactions. The redox state can be recognized as a voltage potential measured with a platinum electrode. This voltage potential is usually referred to as E or Eh defined by the Nernst equation, which was introduced in Chapter 5, Section 5.3.1:
nF J [Reactants]
The Nernst equation is applicable only if the redox reaction is reversible. Not all reactions are completely reversible in natural systems; activities of reacting components may be too low or equilibrium may be reached very slowly. In a sediment, the biotic microenvironment may create a redox potential that is different from the surrounding macroenvironment. For this reason, measurements of Eh in natural systems must be cautiously evaluated and not used strictly for calculations of chemical equilibria. Calculations of redox equilibria are in some cases valuable, in the sense that they will give information about the direction of chemical reactions.
If a system is not at equilibrium, which is common for natural systems, each reaction has its own Eh value and the observed electrode potential is a mixed potential depending on the kinetics of several reactions. A redox pair with relatively high ion activity and whose electron exchange process is fast tends to dominate the registered Eh- Thus, measurements in a natural environment may not reveal information about all redox reactions but only from those reactions that are active enough to create a measurable potential difference on the electrode surface.
In marine sediments, usually only the uppermost layer of the sediment exhibits oxidizing conditions while the rest is reduced. The thickness of the oxidized layer and the reducing capacity of the sediment below depend on:
1. concentration of oxygen in the ambient bottom water;
2. the rate of oxygen penetration into the sediment;
3. the accessibility of utilizable organic matter for the bacterial activity.
The depositional rate of organic matter is higher in the coastal areas than in the open sea. Even though the exchange of water is higher in shallow areas of the ocean the deep water is not deficient in oxygen. We should thus expect to find a general trend of thicker oxidized sediments with increasing distance from the shoreline. Lynn and Bonatti (1965), for example, found that the thickness of the oxidized upper sediment in the Pacific Ocean between 15° S and 20° N increased with distance from the continent. Thicknesses were < 1 cm near the shore and 8-15 cm at distances of 800 km offshore.
The inflection point of the redox gradient, constituting the boundary between oxidizing and reducing environments in sediments lies at around +250 mV. This boundary, the redox-cline (Hallberg, 1972), is recognized as comparable to other natural boundaries, such as the halocline and thermocline. The redoxcline is usually situated close to the sediment-water interface resulting in a redox turnover from oxidizing to reducing conditions during early diagenesis. The redox turnover will, in turn, produce disequilibrium within the sediment, causing dissolution of certain minerals and compounds and increased ion exchange across the redoxcline. The change of redox potential is a useful tool for describing the sedimentary environment.
In oceanic sediments macro- (> 1 mm) and microorganisms play an important role in the mixing of surface sediment layers. Burrowing by organisms in marine sediments is so common that it is the preservation of depositional structures that requires explanation, not their destruction (Arrhenius, 1952, p. 86). Bioturbation is the mixing of sediments by in-situ fauna, and is usually attributed to macro- and meso-fauna (0.1-1 mm of sieved sediments). Very little is known about the role of microfauna. The mixing of sediments by the former can easily be observed in coastal areas where their abundance is high and the resuspension and cycling of the annual sediment influx can be as high as 99% (Young, 1971). In anaerobic sediments, however, where macro- and mesofauna are absent due to lack of oxygen, there is still a large community of microorganisms; 107 individuals/ cm3 of sediments is not an uncommon value. Several types are motile and have been observed to swim 3 mm/day (Oppenheimer, 1960). If all bacteria in a cubic centimeter of sediment moved 1 mm/day they would cover a total distance of 10 km/day. These organisms cannot move particles but may have a significant effect on the mixing of the interstitial water, and thus also on the exchange between water and sediments.
8.6 Soils, Weathering, and Global Biogeochemical Cycles
The soil may represent a thin film on the surface of the Earth, but the importance of soils in global biogeochemical cycles arises from their role as the interface between the Earth, its atmosphere, and the biosphere. All terrestrial biological activity is founded upon soil productivity, and the weathering of rocks that helps to maintain atmospheric equilibrium occurs within soils. Soils provide the foundation for key aspects of global biogeochemical cycles.
The soil, an open system in the context of biogeochemical cycles, receives inputs and outputs of C and N, and mineral elements and is the foundation of the primary production of terrestrial ecosystems. The amount of carbon present in the soil is closely connected to the C02 concentration of the atmosphere, but atmospheric C02 is regulated mainly by the ocean rather than by the soil (see Chapters 10 and 11). The amount of N in the soil also does not influence the N in the atmosphere because the atmosphere is a huge reservoir regulated mainly by the ocean (see Chapter 12). Nevertheless, the soil has a tremendous influence on the nitrate load of rivers. Biogeochemical processes that occur in soils and the processes controlling the delivery of the products of these reactions to the oceans exert profound influences on global biogeochemical cycles.
Weathering processes take an active part in the cycling of oxygen and carbon, but does chemical weathering affect these cycles to a significant extent? Consider the following examples.
Oxygen is formed by photosynthesis according to the reaction
The annual primary production of organic carbon through photosynthesis is on the order of 70Pg/yr. The major part of this carbon is decomposed or respired in a process that also involves the biogeochemical transformation of nitrogen, sulfur, and many other elements. Only a small part of the annual primary production of organic carbon escapes decomposition and is buried in marine sediments. On average, sediments and sedimentary rocks contain 3% carbon, which corresponds to a net production of oxygen of about 300 Tg/yr. This production has the potential to change atmospheric oxygen over a time scale of 4 Myr. In contrast, the fossil record indicates that the partial pressure of oxygen in the atmosphere has fluctuated very little, at least during the last 600 Myr (Conway, 1943). Obviously there must be a sink that can accommodate the net annual production of oxygen. This sink is the annual weathering of rocks, which can be estimated from the approximately 18 x 1023 g of sediments formed during the last 600 Myr (Gregor, 1968). During this period, soil formation and sediment formation have fluctuated greatly, but the average weathering rate is 30 x 1014 g/yr. The average continental rock contains ferrous iron and sulfide sulfur (mainly as pyrite) and organic carbon. During the weathering processes these constituents are oxidized and consume oxygen. The amounts of oxygen necessary for each of these reactions are given in Table 8-8.
The total annual consumption of oxygen by weathering can be estimated (Holland, 1978) as follows:
Note that this estimate of the annual 02 loss to weathering processes is approximately equal to the estimated annual production of oxygen estimated above. Hence, the weathering of rocks and burial of organic carbon in sediments during their formation are important processes for the oxygen content of the atmosphere.
Weathering of rocks is also a sink for C02. Garrels and Mackenzie (1971) have estimated that the formation of 1500 g of sedimentary rocks requires 100 g of C02. If we use the same
Table 8-8 Average oxygen consumption during weathering of rocks"
number as above for the annual formation of sedimentary rocks, we have a C02 sink of 200 Tg/yr. About half of that comes from the oxidation of organic carbon in the weathered rock. The rest is from the atmosphere. The burning of fossil fuel increases the partial pressure of C02 in the atmosphere, but increased C02 levels increase the weathering rate. Thus, over geologic time, the weathering rate and sediment formation ultimately control the carbon dioxide content of the atmosphere.
8-1 Discuss how physical weathering operates in each of the following environments: (1) sea shore, (2) hot desert, (3) temperate forest.
8-2 Rewrite the weathering reactions shown in Section 18.104.22.168 using HN03 in place of H2C03.
8-3 Why do we speak in terms of soil horizons and sediment layers?
8-4 Discuss the significance of clay minerals in a description of the solid phase of a sediment.
8-5 Give examples of early diagenetic processes.
8-6 Why do continental margins play a dominant role on the biogeochemical cycling of elements?
8-7 Give examples of bacterial transformations in a sediment that are of special importance for biogeochemical cycles.
8-8 Write a balanced reaction for formation of hematite from Fe2+.
8-9 Explain the difference between a Mollisol and a Spodosol. How would cycling of Ca differ in each? N?
8-10 Track the possible fate of a K+ ion from the moment it is released by weathering in a soil to its burial at sea.
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