UNIT III1. Microbial transformations of nitrogen, phosphorus, sulphur, iron and manganese in soil; 2. Biochemical composition and biodegradation of soil organic matter and crop residues,

humus formation; 3. Cycles of important organic nutrients.

Nitrogen cycle

Schematic representation of the flow of nitrogen through the environment. The importance of bacteria in the cycle is immediately recognized as being a key element in the cycle, providing different forms of nitrogen compounds assimilable by higher organisms.

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The nitrogen cycle is the process by which nitrogen is converted between its various chemical forms. This transformation can be carried out through both biological and physical processes. Important processes in the nitrogen cycle include fixation, mineralization, nitrification, and denitrification. The majority of Earth's atmosphere (approximately 78%) is nitrogen,[1] making it the largest pool of nitrogen. However, atmospheric nitrogen has limited availability for biological use, leading to a scarcity of usable nitrogen in many types of ecosystems. The nitrogen cycle is of particular interest to ecologists because nitrogen availability can affect the rate of key ecosystem processes, including primary production and decomposition. Human activities such as fossil fuel combustion, use of artificial nitrogen fertilizers, and release of nitrogen in wastewater have dramatically altered the global nitrogen cycle.

A 2011 study found that nitrogen from rocks may also be a significant source of nitrogen, that had not previously been included.

Accounting

The first step in understanding the nitrogen cycle is to examine the distribution of N on earth. The Table below gives the distribution of N in x1015 grams. Notice that the largest pool of available N is in the atmosphere.

Rocks and sediments 190,400,120 (deep, unavailable)

Atmosphere 3,900,000Ocean 23,348Soils 460Land plants 14land animals 0.2 In the Atmosphere:N2 3,900,000N2O  1.4NOx 0.0006

(less than 1 billionth %)

Contents

1 Ecological function2 The processes of the nitrogen cycle

2.1 Nitrogen fixation 2.1.1 Conversion of N2

2.2 Assimilation2.3 Ammonification2.4 Nitrification2.5 Denitrification2.6 Anaerobic ammonium oxidation

3 Marine nitrogen cycle 3.1 New vs. regenerated nitrogen

4 Human influences on the nitrogen cycle

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4.1 Wastewater treatment4.2 Environmental impacts

Ecological function

Nitrogen is essential for many processes; it is crucial for any life on Earth. It is a component in all amino acids, as incorporated into proteins, and is present in the bases that make up nucleic acids, such as DNA and RNA. In plants, much of the nitrogen is used in chlorophyll molecules, which are essential for photosynthesis and further growth. Although Earth’s atmosphere is an abundant source of nitrogen, most is relatively unusable by plants. Chemical processing, or natural fixation (through processes such as bacterial conversion), are necessary to convert gaseous nitrogen into forms usable by living organisms, which makes nitrogen a crucial component of food production. The abundance or scarcity of this "fixed" form of nitrogen, (also known as reactive nitrogen), dictates how much food can be grown on a piece of land.

The processes of the nitrogen cycle

Nitrogen is present in the environment in a wide variety of chemical forms including organic nitrogen, ammonium (NH4

+), nitrite (NO2-), nitrate (NO3

-), nitrous oxide (N2O), nitric oxide (NO) or inorganic nitrogen gas (N2). Organic nitrogen may be in the form of a living organism, humus or in the intermediate products of organic matter decomposition. The processes of the nitrogen cycle transforms nitrogen from one form to another. Many of those processes are carried out by microbes, either in their effort to harvest energy or to accumulate nitrogen in a form needed for their growth. The diagram above shows how these processes fit together to form the nitrogen cycle.

Nitrogen fixation

Atmospheric nitrogen must be processed or "fixed" for use by plants. Some fixation occurs in lightning strikes, but most fixation is done by free-living or symbiotic bacteria. These bacteria have the nitrogenase enzyme that combines gaseous nitrogen with hydrogen to produce ammonia, which is then further converted by the bacteria to make their own organic compounds. Most biological nitrogen fixation occurs by the activity of Mo-nitrogenase, found in a wide variety of bacteria and some Archaea (Archaea and bacteria are quite similar in size and shape, but a few archaea have unusual shapes eg, flat and square-shaped cells. Despite the visual similarity to bacteria, archaea possess genes and several metabolic pathways that are more closely related to those of eukaryotes). Mo-nitrogenase is a complex two component enzyme. Some nitrogen fixing bacteria, such as Rhizobium, live in the root nodules of legumes (such as peas or beans). Here they form a mutualistic relationship with the plant, producing ammonia in exchange for carbohydrates. Nutrient-poor soils can be planted with legumes to enrich them with nitrogen. A few other plants can form such symbioses. Today, about 30% of the total fixed nitrogen is manufactured in ammonia chemical plants.

Conversion of N2

The conversion of nitrogen (N2) from the atmosphere into a form readily available to plants and hence to animals is an important step in the nitrogen cycle, which distributes the supply of this essential nutrient. There are four ways to convert N2 (atmospheric nitrogen gas) into more chemically reactive forms:

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1. Biological fixation: some symbiotic bacteria (most often associated with leguminous plants) and some free-living bacteria are able to fix nitrogen as organic nitrogen. An example of mutualistic nitrogen fixing bacteria are the Rhizobium bacteria, which live in legume root nodules. These species are diazotrophs. An example of the free-living bacteria is Azotobacter.

2. Industrial N-fixation: Under great pressure, at a temperature of 600 C, and with the use of an iron catalyst, hydrogen (usually derived from natural gas or petroleum) and atmospheric nitrogen can be combined to form ammonia (NH3) in the Haber-Bosch process which is used to make fertilizer and explosives.

3. Combustion of fossil fuels: automobile engines and thermal power plants, which release various nitrogen oxides (NOx).

4. Other processes: In addition, the formation of NO from N2 and O2 due to photons and especially lightning, can fix nitrogen.

Assimilation

Plants take nitrogen from the soil, by absorption through their roots in the form of either nitrate ions or ammonium ions. All nitrogen obtained by animals can be traced back to the eating of plants at some stage of the food chain.

Plants can absorb nitrate or ammonium ions from the soil via their root hairs. If nitrate is absorbed, it is first reduced to nitrite ions and then ammonium ions for incorporation into amino acids, nucleic acids, and chlorophyll. In plants that have a symbiotic relationship with rhizobia, some nitrogen is assimilated in the form of ammonium ions directly from the nodules. Animals, fungi, and other heterotrophic organisms obtain nitrogen by ingestion of amino acids, nucleotides and other small organic molecules.

Ammonification

When a plant or animal dies, or an animal expels waste, the initial form of nitrogen is organic. Bacteria, or fungi in some cases, convert the organic nitrogen within the remains back into ammonium (NH4

+), a process called ammonification or mineralization. Enzymes Involved:

GS: Gln Synthetase (Cytosolic & PLastid) GOGAT: Glu 2-oxoglutarate aminotransferase (Ferredoxin & NADH dependent) GDH: Glu Dehydrogenase:

o Minor Role in ammonium assimilation.o Important in amino acid catabolism.

Nitrification

The conversion of ammonia to nitrate is performed primarily by soil-living bacteria and other nitrifying bacteria. In the primary stage of nitrification, the oxidation of ammonium (NH4

+) is performed by bacteria such as the Nitrosomonas species, which converts ammonia to nitrites (NO2

-). Other bacterial species, such as the Nitrobacter, are responsible for the oxidation of the nitrites into nitrates (NO3

-). It is important for the ammonia to be converted to nitrates because accumulated nitrites are toxic to plant life.

Due to their very high solubility, and because soils are largely unable to retain anions, nitrates can enter groundwater. Elevated nitrate in groundwater is a concern for drinking water use

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because nitrate can interfere with blood-oxygen levels in infants and cause methemoglobinemia or blue-baby syndrome. Where groundwater recharges stream flow, nitrate-enriched groundwater can contribute to eutrophication, a process that leads to high algal, especially blue-green algal populations. While not directly toxic to fish life, like ammonia, nitrate can have indirect effects on fish if it contributes to this eutrophication. Nitrogen has contributed to severe eutrophication problems in some water bodies. Since 2006, the application of nitrogen fertilizer has been increasingly controlled in Britain and the United States. This is occurring along the same lines as control of phosphorus fertilizer, restriction of which is normally considered essential to the recovery of eutrophied water bodies.

Denitrification

Denitrification is the reduction of nitrates back into the largely inert nitrogen gas (N2), completing the nitrogen cycle. This process is performed by bacterial species such as Pseudomonas and Clostridium in anaerobic conditions. They use the nitrate as an electron acceptor in the place of oxygen during respiration. These facultatively anaerobic bacteria can also live in aerobic conditions.

Anaerobic ammonium oxidation

In this biological process, nitrite and ammonium are converted directly into molecular nitrogen (N2) gas. This process makes up a major proportion of nitrogen conversion in the oceans.

Aquatic nitrogen cycle

A proposed revision to the aquatic/marine nitrogen cycle with direct nitrification-anammox coupling in the suboxic zone. It illustrates the fact that nitrification and anammox, previously considered as clearly oxic and anoxic processes respectively, can actually cooccur under microaerobic conditions. Nitrification in the suboxic zone provides a direct source of nitrite for anammox. It acts as a short circuit channeling regenerated NH4

+ to direct N loss.

PON=Particulate organic nitrogen, DON= Dissolved organic nitrogen, DNRA= Dissimilatory nitrate reductase to ammonium

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Anammox, an abbreviation for ANaerobic AMMonium OXidation, is a globally important microbial process of the nitrogen cycle. In this biological process, nitrite and ammonium are converted directly into dinitrogen gas. This process contributes up to 50% of the dinitrogen gas produced in the oceans. It is thus a major sink for fixed nitrogen and so limits oceanic primary productivity. The overall catabolic reaction is:

NH4+ + NO2

− → N2 + 2H2O.

The bacteria that perform the anammox process belong to the bacterial phylum Planctomycetes, of which Planctomyces and Pirellula are the best known genera. Currently four genera of anammox bacteria have been (provisionally) defined: Brocadia, Kuenenia, Anammoxoglobus, Jettenia (all fresh water species), and Scalindua (marine species).

The nitrogen cycle is an important process in the ocean as well. While the overall cycle is similar, there are different players and modes of transfer for nitrogen in the ocean. Nitrogen enters the water through precipitation, runoff, or as N2 from the atmosphere. Nitrogen cannot be utilized by phytoplankton (Phytoplankton are photosynthesizing microscopic organisms that inhabit the upper sunlit layer of almost all water bodies. It is too small to be individually seen with the unaided eye, but when present in high enough numbers, they may appear as a green discoloration of the water due to the presence of chlorophyll or even other colours) as N2

so it must undergo nitrogen fixation which is performed predominately by cyanobacteria (aquatic, photosynthetic and usually unicellular bacteria often grow in colonies large enough to see, it is a relative of bacteria, chloroplast in eukaryotic algae eg, BGA). Without supplies of fixed nitrogen entering the marine cycle the fixed nitrogen would be used up in about 2000 years. Phytoplankton needs nitrogen in biologically available forms for the initial synthesis of organic matter. Ammonia and urea are released into the water by excretion from plankton. Nitrogen sources are removed from the euphotic zone by the downward movement of the organic matter. This can occur from sinking of phytoplankton, vertical mixing, or sinking of waste of vertical migrators. The sinking results in ammonia being introduced at lower depths below the euphotic

zone. Bacteria are able to convert ammonia to nitrite and nitrate but they are inhibited by light so this must occur below the euphotic zone. Ammonification or Mineralization is performed by bacteria to convert the ammonia to ammonium. Nitrification can then occur to convert the ammonium to nitrite and nitrate. Nitrate can be returned to the euphotic zone by vertical mixing and upwelling where it can be taken up by phytoplankton to continue the cycle. N2 can be returned to the atmosphere through denitrification.

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A schematic representing the Marine Nitrogen Cycle

NH4+ is thought to be the preferred source of fixed nitrogen for phytoplankton because its

assimilation does not involve a redox reaction and therefore requires little energy. However NO3

is more abundant so most phytoplankton have adapted to have the enzymes necessary to undertake this reduction (nitrate reductase). There are a few notable and well-known exceptions that include Prochlorococcus and some Synechococcus. These species can only take up nitrogen as NH4

+.

The nutrients in the ocean are not uniformly distributed. Areas of upwelling provide supplies of nitrogen from below the euphotic zone. Coastal zones provide nitrogen from runoff and upwelling occurs readily along the coast. However, the rate at which nitrogen can be taken up by phytoplankton is decreased in oligotrophic waters all year-round and temperate water in the summer resulting in lower primary production. The distribution of the different forms of nitrogen varies throughout the oceans as well.

Nitrate is depleted in near-surface water except in upwelling regions. Coastal upwelling regions usually have high nitrate and chlorophyll levels as a result of the increased production. However, there are regions of high surface nitrate but low chlorophyll that are referred to as HNLC (high nitrogen, low chlorophyll) regions. As of now the best explanation for HNLC regions relates to iron limitation in the ocean. In recent years iron has become an important player when discussing ocean dynamics and nutrient cycles. The input of iron varies by region and is delivered to the ocean by dust (from dust storms) and is leached out of rocks. Iron is under consideration as the true limiting element in the ocean.

NH4+ and NO2 show a maximum concentration at 50–80 m (lower end of the euphotic zone) with

decreasing concentration below that depth. This distribution can be accounted for by the fact that NO2 and NH4

+ are intermediate species. They are both rapidly produced and consumed through the water column. The amount of NH4

+ in the ocean is about 3 orders of magnitude less than nitrate. Between NH4

+, NO2, and NO3, NO2 has the fastest turnover rate. It can be produced during NO3 assimilation, nitrification, and denitrification; however, it is immediately consumed again.

New vs. regenerated nitrogen

Nitrogen entering the euphotic zone is referred to as new nitrogen because it is newly arrived from outside the productive layer. The new nitrogen can come from below the euphotic zone or from outside sources. Outside sources are considered to be upwelling from deep water or by nitrogen fixation. If the organic matter is eaten, respired, delivered to the water as ammonia, and re-incorporated into organic matter by phytoplankton it is considered recycled/regenerated production.New production is an important component of the marine environment. One reason is that only continual input of new nitrogen can determine the total capacity of the ocean to produce a sustainable fish harvest. Harvesting fish from regenerated nitrogen areas will lead to a decrease in nitrogen and therefore a decrease in primary production. This will have a negative effect on the system. However, if fish are harvested from areas of new nitrogen the nitrogen will be replenished.

Human influences on the nitrogen cycle

As a result of extensive cultivation of legumes (particularly soy, alfalfa, and clover), growing use of the Haber-Bosch process in the creation of chemical fertilizers, and pollution emitted by

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vehicles and industrial plants, human beings have more than doubled the annual transfer of nitrogen into biologically-available forms. In addition, humans have significantly contributed to the transfer of nitrogen trace gases (NOx) from Earth to the atmosphere, and from the land to aquatic systems. Human alterations to the global nitrogen cycle are most intense in developed countries and in Asia, where vehicle emissions and industrial agriculture are highest.

Nitrous oxide (N2O) has risen in the atmosphere as a result of agricultural fertilization, biomass burning, cattle and feedlots, and industrial sources. N2O has deleterious effects in the stratosphere, where it breaks down and acts as a catalyst in the destruction of atmospheric ozone. In the atmosphere nitrous oxide is a greenhouse gas, and is currently the third largest contributor to global warming, after carbon dioxide and methane. While not as abundant in the atmosphere as carbon dioxide, it is for an equivalent mass, nearly 300 times more potent in its ability to warm the planet.

Ammonia (NH3) in the atmosphere has tripled as the result of human activities. It is a reactant in the atmosphere, where it acts as an aerosol, decreasing air quality and clinging to water droplets, eventually resulting in nitric acid (HNO3) that produces acid rain. Atmospheric ammonia and nitric acid damage respiratory systems.

All forms of high-temperature combustion have contributed to a 6 or 7 fold increase in the flux of NOx to the atmosphere. Its production is a function of combustion temperature - the higher the temperature, the more NOx is produced. Fossil fuel combustion is a primary contributor, but so are biofuels and even the burning of hydrogen. The higher combustion temperature of hydrogen produces more NOx than natural gas combustion. The very-high temperature of lightning produces small amounts of NOx, NH3, and HNO3.

Ammonia and nitrous oxides actively alter atmospheric chemistry. They are precursors of tropospheric (lower atmosphere) ozone production, which contributes to smog/pollution, and acid rain, damages plants and increases nitrogen inputs to ecosystems. Ecosystem processes can increase with nitrogen fertilization, but anthropogenic input can also result in nitrogen saturation, which weakens productivity and can damage the health of plants, animals, fish, and humans.

Decreases in biodiversity can also result if higher nitrogen availability increases nitrogen-demanding grasses, causing a degradation of nitrogen-poor species diverse heath lands.

Wastewater treatment

Onsite sewage facilities such as septic tanks and holding tanks release large amounts of nitrogen into the environment by discharging through a drainfield into the ground. Microbial activity consumes the nitrogen and other contaminants in the wastewater.

However, in certain areas, the soil is unsuitable and some or all of the wastewater, with the contaminants, enters the aquifers. These contaminants accumulate and eventually end up in drinking water. One of the contaminants most concerned about is nitrogen in the form of nitrates. A nitrate concentration of 10 ppm (parts per million) or 10 milligrams per liter is the current EPA limit for drinking water and typical household wastewater can produce a range of 20–85 ppm.

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One health risk associated with drinking water (with >10 ppm nitrate) is the development of methemoglobinemia and has been found to cause blue baby syndrome. Several American states have now started programs to introduce advanced wastewater treatment systems to the typical onsite sewage facilities. The result of these systems is an overall reduction of nitrogen, as well as other contaminants in the wastewater.

Environmental impacts

Additional risks posed by increased availability of inorganic nitrogen in aquatic ecosystems include water acidification; eutrophication of fresh and saltwater systems; and toxicity issues for animals, including humans. Eutrophication often leads to lower dissolved oxygen levels in the water column, including hypoxic and anoxic conditions, which can cause death of aquatic fauna. Relatively sessile benthos, or bottom-dwelling creatures, are particularly vulnerable because of their lack of mobility, though large fish kills are not uncommon. Oceanic dead zones near the mouth of the Mississippi in the Gulf of Mexico are a well-known examples of algal bloom-induced hypoxia. The New York Adirondack Lakes, Catskills, Hudson Highlands, Rensselaer Plateau and parts of Long Island display the impact of nitric acid rain deposition, resulting in the killing of fish and many other aquatic species.

Ammonia (NH3) is highly toxic to fish and the level of ammonia discharged from wastewater treatment facilities must be closely monitored. To prevent fish deaths, nitrification via aeration prior to discharge is often desirable. Land application can be an attractive alternative to the aeration.

Phosphorus cycle

The phosphorus cycle

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The phosphorus cycle is the biogeochemical cycle that describes the movement of phosphorus through the lithosphere, hydrosphere, and biosphere. Unlike many other biogeochemical cycles, the atmosphere does not play a significant role in the movement of phosphorus, because phosphorus and phosphorus-based compounds are usually solids at the typical ranges of temperature and pressure found on Earth. The production of phosphine (PH3) gas occurs only in specialized, local conditions.

Low phosphorus (chemical symbol, P) availability slows down microbial growth, which has been shown in studies of soil microbial biomass. Soil microorganisms act as sinks and sources of available P in the biogeochemical cycle. Locally, transformations of PO4 are microbially driven; however, the major transfers in the global cycle of P are not driven by microbial reactions, but by tectonic movements in geologic time. Further studies need to be performed for integrating different processes and factors related to gross phosphorus mineralization and microbial phosphorus turnover in general.

Contents

1 Phosphorus in the environment 1.1 Ecological function1.2 Biological function1.3 Process of the cycle1.4 Phosphatics minerals

2 Human interference

Phosphorus in the environment

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The aquatic phosphorus cycle

Ecological function

Phosphorus is an essential nutrient for plants and animals in the form of ions. Phosphorus is a limiting nutrient for aquatic organisms. Phosphorus forms parts of important life-sustaining molecules that are very common in the biosphere. Phosphorus does not enter the atmosphere, remaining mostly on land and in rock and soil minerals. Eighty percent of the mined phosphorus is used to make fertilizers, and a type of phosphorus such as dilute phosphoric acid is used in soft drinks. Phosphates may be effective in such ways but also causes pollution issues in lakes and streams. Enrichment of phosphate can lead to eutrophication of fresh and inshore marine waters, leading to algae bloom because of the excess nutrients. Bacteria consume the algae and a bacterial bloom ensues. Cellular respiration of bacteria and decomposers use all the oxygen in the water, causing many fish to die.

Phosphorus normally occurs in nature as part of a phosphate ion (PO4)3-, consisting of a phosphorus atom and 4 oxygen atoms, the most abundant form is orthophosphate. Most phosphates are found as salts in ocean sediments or in rocks. Over time, geologic processes can bring ocean sediments to land, and weathering will carry these phosphates to terrestrial habitats. Plants absorb phosphates from the soil, then bind the phosphate into organic compounds. The plants may then be consumed by herbivores who in turn may be consumed by carnivores. After death, the animal or plant decays, and the phosphates are returned to the soil. Runoff may carry them back to the ocean or they may be reincorporated into rock.

Biological function

The primary biological importance of phosphates is as a component of nucleotides, which serve as energy storage within cells (ATP) or when linked together, form the nucleic acids DNA and RNA. The double helix of our DNA is only possible because of the phosphate ester bridge that binds the helix. Besides making biomolecules, phosphorus is also found in bones, whose strength is derived from calcium phosphate, in enamel of mammalian teeth, exoskeleton of insects, and phospholipids (found in all biological membranes). It also functions as buffering agent in maintaining acid base homeostasis in the human body.

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Process of the cycle

Phosphates move quickly through plants and animals; however, the processes that move them through the soil or ocean are very slow, making the phosphorus cycle overall one of the slowest biogeochemical cycles.

Unlike other cycles of matter compounds, phosphorus cannot be found in air as a gas, it only occurs under highly reducing conditions as the gas phosphine PH3. This is because at normal temperature and circumstances, it is a solid in the form of red and white phosphorus. It usually cycles through water, soil and sediments. Phosphorus is typically the limiting nutrient found in streams, lakes and fresh water environments. As rocks and sediments gradually wear down, phosphate is released. In the atmosphere phosphorus is mainly small dust particles.

Initially, phosphate weathers from rocks. The small losses in a terrestrial system caused by leaching through the action of rain are balanced in the gains from weathering rocks. In soil, phosphate is absorbed on clay surfaces and organic matter particles and becomes incorporated (immobilized). Plants dissolve ionized forms of phosphate. Herbivores obtain phosphorus by eating plants, and carnivores by eating herbivores. Herbivores and carnivores excrete phosphorus as a waste product in urine and feces. Phosphorus is released back to the soil when plants or animal matter decomposes and the cycle repeats.

Phosphatic minerals

The availability of phosphorus in ecosystem is restricted by the rate of release of this element during weathering. The release of phosphorus from apatite dissolution is a key control on ecosystem productivity. The primary mineral with significant phosphorus content, apatite [Ca5(PO4)3OH] undergoes carbonation weathering releasing phosphorus contained different forms. This process decreases the total phosphorus in the system due to losses in runoff.

Little of this released phosphorus is taken by biota (organic form) whereas, large proportion reacts with other soil minerals leading to precipitation in unavailable forms. The later stage of weathering and soil development. Available phosphorus is found in a biogeochemical cycle in the upper soil profile, while phosphorus found at lower depths is primarily involved in geochemical reactions with secondary minerals. Plant growth depends on the rapid root uptake of phosphorus released from dead organic matter in the biochemical cycle. Phosphorus is limited in supply for plant growth.

Low-molecular-weight (LMW) organic acids are found in soils. They originate from the activities of various microorganisms in soils or may be exuded from the roots of living plants. Several of those organic acids are capable of forming stable organo-metal complexes with various metal ions found in soil solutions. As a result, these processes may lead to the release of inorganic phosphorus associated with aluminium, iron, and calcium in soil minerals. The production and release of oxalic acid by mycorrhizal fungi explain their importance in maintaining and supplying phosphorus to plant.

The availability of organic phosphorus to support microbial, plant and animal growth depends on the rate of their degradation to generate free phosphate. There are various enzymes such as phosphatases, nucleases and phytase involved for the degradation. Some of the abiotic pathways in the environment are hydrolytic reactions and photolytic reactions. Enzymatic hydrolysis of organic phosphorus is an essential step in the biogeochemical phosphorus cycle,

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including the phosphorus nutrition of plants and microorganisms and the transfer of organic phosphorus from soil to bodies of water. Many organisms rely on the soil derived phosphorus for their phosphorus nutrition.

Human interference

Nutrients are important to the growth and survival of living organisms, and hence, are essential for development and maintenance of healthy ecosystems. However, excessive amounts of nutrients, particularly phosphorus and nitrogen, are detrimental to aquatic ecosystems. Natural eutrophication is a process by which lakes gradually age and become more productive and may take thousands of years to progress. Cultural or anthropogenic eutrophication, however, is water pollution caused by excessive plant nutrients, which results in excessive growth in algae population. Surface and subsurface runoff and erosion from high-P soils may be major contributing factors to fresh water eutrophication. The processes controlling soil P release to surface runoff and to subsurface flow are a complex interaction between the type of P input, soil type and management, and transport processes depending on hydrological conditions.

Repeated application of liquid hog manure in excess to crop needs can have detrimental effects on soil P status. In poorly drained soils or in areas where snowmelt can cause periodical waterlogging, Fe-reducing conditions can be attained in 7–10 days. This causes a sharp increase in P concentration in solution and P can be leached. In addition, reduction of the soil causes a shift in phosphorus from resilient to more labile forms. This could eventually increase the potential for P loss. This is of particular concern for the environmentally sound management of such areas, where disposal of agricultural wastes has already become a problem. It is suggested that the water regime of soils that are to be used for organic wastes disposal is taken into account in the preparation of waste management regulations.

Human interference in the phosphorus cycle occurs by overuse or careless use of phosphorus fertilizers. This results in increased amounts of phosphorus as pollutants in bodies of water resulting in eutrophication. Eutrophication devastates water ecosystems.

Total excess input from 1600 to 3600 AD is 1860 Tg (teragrams) of phosphorus. Given that, in the marine environment, between 106 and 170 units of carbon are buried per unit of phosphorus one can predict that excess phosphorus would effectively bury 76,000 to 123,000 Tg carbon. In essence, this burial removes carbon from the atmosphere through the biological fixation of carbon dioxide during photosynthesis. The present annual rate of anthropogenic carbon addition to the atmosphere is 7900 Tg carbon, so the phosphorus eutrophication effect would only account for about 10–15 years of anthropogenic carbon emissions to the atmosphere over the next 2000 years (i.e. only 0.6% of total projected carbon emissions, if emissions stay constant).

Although the net effect as a carbon sequestration mechanism is minimal, the ecological impact of phosphorus fertilization to the ocean could be extreme. Given the other assaults on marine ecosystems, including warming, and acidification of surface ocean waters from higher carbon dioxide levels, it would be pure speculation to project how P eutrophication would affect ecosystem structure and distribution in the future. However, those who have witnessed local eutrophication in ditches, streams, ponds, and lakes can attest to the ecological devastation that excess nutrients and the proliferation of monocultures can cause in such isolated environments. The eutrophication of coastal and open-marine ecosystems would result in a grim future for ecological diversity.

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Sulfur CycleThe sulfur cycle is the collection of processes by which sulfur moves to and from minerals (including the waterways) and living systems. Such biogeochemical cycles are important in geology because they affect many minerals. Biogeochemical cycles are also important for life because sulfur is an essential element, being a constituent of many proteins and cofactors.

The Sulfur cycle (in general)

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Steps of the sulfur cycle are:

Mineralization of organic sulfur into inorganic forms, such as hydrogen sulfide (H2S), elemental sulfur, as well as sulfide minerals.

Oxidation of hydrogen sulfide, sulfide, and elemental sulfur (S) to sulfate (SO42–).

Reduction of sulfate to sulfide. Incorporation sulfide into organic compounds (including metal-containing derivatives).

Structure of 3'-phosphoadenosine-5'-phosphosulfate, a key intermediate in the sulfur cycle.

These are often termed as follows:

Assimilative sulfate reduction (sulfur assimilation) in which sulfate (SO42–) is reduced by plants,

fungi and various prokaryotes. The oxidation states of sulfur are +6 in sulfate and –2 in R–SH. Desulfurization in which organic molecules containing sulfur can be desulfurized, producing

hydrogen sulfide gas (H2S, oxidation state = –2). An analogous process for organic nitrogen compounds is deamination.

Oxidation of hydrogen sulfide produces elemental sulfur (S8), oxidation state = 0. This reaction occurs in the photosynthetic green and purple sulfur bacteria and some chemolithotrophs. Often the elemental sulfur is stored as polysulfides.

Oxidation of elemental sulfur by sulfur oxidizers produces sulfate. Dissimilative sulfur reduction in which elemental sulfur can be reduced to hydrogen sulfide. Dissimilative sulfate reduction in which sulfate reducers generate hydrogen sulfide from

sulfate.

Contents

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1 Sulfur sources and sinks2 Biologically and thermochemically driven sulfate reduction3 δ34S4 Evolution of the sulfur cycle5 Economic importance6 Human Impact

Sulfur sources and sinks

Sulfur is found in oxidation states ranging from +6 in SO4 to -2 in sulfides. Thus elemental sulfur can either give or receive electrons depending on its environment. Igneous rocks such as pyrite (FeS2) comprised the original pool of sulfur on earth. Owing to the sulfur cycle, the amount of mobile sulfur has been continuously increasing through volcanic activity as well as weathering of the crust in an oxygenated atmosphere. Earth’s main sulfur sink is the oceans as SO2, where it is the major oxidizing agent.

When SO4 is assimilated by organisms, it is reduced and converted to organic sulfur, which is an essential component of proteins. However, the biosphere does not act as a major sink for sulfur, instead the majority of sulfur is found in seawater or sedimentary rocks especially pyrite rich shales and evaporite rocks (anhydrite and baryte). The amount of sulfate in the oceans is controlled by three major processes:

1. input from rivers2. sulfate reduction and sulfide reoxidation on continental shelves and slopes3. burial of anhydrite and pyrite in the oceanic crust.

There is no significant amount of sulfur held in the atmosphere with all of it coming from either sea spray or windblown sulfur rich dust, neither of which is long lived in the atmosphere. In recent times the large annual input of sulfur from the burning of coal and other fossil fuels adds a substantial amount SO2 which acts as an air pollutant. In the geologic past, igneous intrusions into coal measures have caused large scale burning of these measures, and consequential release of sulfur to the atmosphere. This has led to substantial disruption to the climate system, and is one of the proposed causes of the great dying.

Dimethylsulfide [(CH3)2S or DMS] is produced by the decomposition of dimethylsulfoniopropionate DMSP) from dying phytoplankton cells in the shallow levels of the ocean, and is the major biogenic gas emitted from the sea, where it is responsible for the distinctive “smell of the sea” along coastlines. DMS is the largest natural source of sulfur gas, but still only has a residence time of about one day in the atmosphere and a majority of it is redeposited in the oceans rather than making it to land. However, it is a significant factor in the climate system, as it is involved in the formation of clouds.

Biologically and thermochemically driven sulfate reduction

Sulfur can be reduced both biologically and thermochemically. Dissimilatory sulfate reduction has two different definitions:

1. the microbial process that converts sulfate to sulfide for energy gain, and

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2. a set of forward and reverse pathways that progress from the uptake and release of sulfate by the cell to its conversion to various sulfur intermediates, and ultimately to sulfide which is released from the cell.

Sulfide and thiosulfate are the most abundant reduced inorganic sulfur species in the environments and are converted to sulfate, primarily by bacterial action, in the oxidative half of the sulfur cycle. Bacterial sulfate reduction (BSR) can only occur at temperature from 0 up to 60–80 °C because above that temperature almost all sulfate-reducing microbes can no longer metabolize. Few microbes can form H2S at higher temperatures but appear to be very rare and do not metabolize in settings where normal bacterial sulfate reduction is occurring. BSR is geologically instantaneous happening on the order of hundreds to thousands of years. Thermochemical sulfate reduction (TSR) occurs at much higher temperatures (160–180 °C) and over longer time intervals, several tens of thousands to a few million years.

The main difference between these two reactions is obvious, one is organically driven and the other is chemically driven. Therefore the temperature for thermochemical sulfate reduction is much higher due to the activation energy required to reduce sulfate. Bacterial sulfate reductions requires lower temperatures because the sulfur reducing bacteria can only live at relatively low temperature (below 60 °C). BSR also requires a relatively open system; otherwise the bacteria will poison themselves when the sulfate levels rise above 5–10%.

The organic reactants involved in BSR are organic acids which are distinctive from the organic reactants needed for TSR. In both cases sulfate is usually derived from the dissolution of gypsum or taken directly out of the seawater. The factors that control whether BSR or TSR will occur are temperature, which is generally a product of depth, with BSR occurring in shallower levels than TSR. Both can occur within the oil window. Their solid products are similar but can be distinguished from one another petrographically (Petrography: study of rocks for mineral content and the textural relationships within the rock), due to their differing crystal sizes, shapes and reflectivity.

Relative attributes of BSR and TSRParticular BSR TSRSubstrate Sulfur SulfurSource gypsum or directly out of seawater gypsum or directly out of seawaterProcess driven Organically ChemicallyTemperature 0 to 60 oC 160-180 oCSystem Relatively open with SO4

+ level upto 5-10%

Not needed

Depth of sea Shallower More deepOccurrence Within oil window Within oil windowProducts Sulfide, thiosulfate, sulfate and

finally H2SSulfide, thiosulfate, sulfate and finally H2S but distinguished petrographically

δ34S

Although 25 isotopes are known for sulfur, only four are stable and of geochemical importance. Of those four, two (32S, light and 34S, heavy) comprise (99.22%) of S on Earth. The vast majority

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(95.02%) of S occurs as 32S with only 4.21% in 34S. The ratio of these two isotopes is fixed in our solar system and has been since its formation. The bulk Earth sulfur isotopic ratio is thought to be the same as the ratio of 22.22 measured from the Canyon Diablo troilite (CDT), a meteorite. That ratio is accepted as the international standard and is therefore set at δ0.00. Deviation from 0.00 which is also a ratio is expressed as δ34S in per mill (part per thousand, ‰). Positive values correlate to increased levels of 34S, whereas negative values correlate with greater 32S in a sample.

δ34S‰ = (Rsample/Rstandard - 1)1000

where "R" is the ratio of the heavy to light isotope in the sample or standard.

Formation of sulfur minerals through non-biogenic processes does not substantially differentiate between the light and heavy isotopes, therefore sulfur values in gypsum or barite (CaSO4.2H2O) should be the same as the overall ratio in the water column at their time of precipitation. Sulfate reduction through biologic activity strongly differentiates between the two isotopes because of the more rapid enzymic reaction with 32S. Sulfate metabolism results in an isotopic depletion of -18‰, and repeated cycles of oxidation and reduction can result in values up to -50‰. Average present day seawater values of δ34S are on the order of +21‰.

Throughout geologic history the sulfur cycle and the isotopic ratios have co-evolved along with the biosphere becoming overall more negative with the increases in biologically driven sulfate reduction, but also show substantial positive excursion. In general, positive excursions in the sulfur isotopes mean that there is an excess of pyrite deposition rather than oxidation of sulfide minerals exposed on land.

Evolution of the sulfur cycle

The isotopic composition of sedimentary sulfides provides primary information on the evolution of the sulfur cycle.

The total inventory of sulfur compounds on the surface of the Earth (nearly 1022g S) represents the total out gassing of sulfur through geologic time. Rocks analyzed for sulfur content are generally organic-rich shales meaning they are likely controlled by biogenic sulfur reduction. Average seawater curves are generated from evaporites deposited throughout geologic time because again, since they do not discriminate between the heavy and light sulfur isotopes, they should mimic the ocean composition at the time of deposition.

4.6 billion years ago (Ga) the Earth formed and had a theoretical δ34S value of 0. Since there was no biologic activity on early Earth there would be no isotopic fractionation. All sulfur in the atmosphere would be released during volcanic eruptions. When the oceans condensed on Earth, the atmosphere was essentially swept clean of sulfur gases, owing to their high solubility in water. Throughout the majority of the Archean (4.6–2.5 Ga) most systems appeared to be sulfate-limited. Some small Archean evaporite deposits require that at least locally elevated concentrations (possibly due to local volcanic activity) of sulfate existed in order for them to be supersaturated and precipitate out of solution.

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3.8–3.6 Ga marks the beginning of the exposed geologic record because this is the age of the oldest rocks on Earth. Metasedimentary rocks from this time still have an isotopic value of 0 because the biosphere was not developed enough (possibly at all) to fractionate sulfur.

3.5 Ga anoxyogenic photosynthesis is established and provides a weak source of sulfate to the global ocean with sulfate concentrations incredibly low the δ34S is still basically 0. Shortly after, at 3.4 Ga the first evidence for minimal fractionation in evaporitic sulfate in association with magmatically derived sulfides can be seen in the rock record. This fractionation shows possible evidence for anoxygenic phototrophic bacteria.

2.8 Ga marks the first evidence for oxygen production through photosynthesis. This is important because there cannot be sulfur oxidation without oxygen in the atmosphere. This exemplifies the coevolution of the oxygen and sulfur cycles as well as the biosphere.

2.7–2.5 Ga is the age of the oldest sedimentary rocks to have a depleted δ 34S which provide the first compelling evidence for sulfate reduction.

2.3 Ga sulfate increases to more than 1 mM; this increase in sulfate is coincident with the “Great Oxygenation Event", when redox conditions on Earth’s surface are thought by most workers to have shifted fundamentally from reducing to oxidizing. This shift would have led to an incredible increase in sulfate weathering which would have led to an increase in sulfate in the oceans. The large isotopic fractionations that would likely be associated with bacteria reduction are produced for the first time. Although there was a distinct rise in seawater sulfate at this time it was likely still only less than 5–15% of present-day levels.

At 1.8 Ga, Banded iron formations (BIF) are common sedimentary rocks throughout the Archean and Paleoproterozoic; their disappearance marks a distinct shift in the chemistry of ocean water. BIFs have alternating layers of iron oxides and chert. BIFs only form if the water is be allowed to supersaturate in dissolved iron (Fe2+) meaning there cannot be free oxygen or sulfur in the water column because it would form Fe3+ (rust) or pyrite and precipitate out of solution. Following this supersaturation, the water must become oxygenated in order for the ferric rich bands to precipitate it must still be sulfur poor otherwise pyrite would form instead of Fe3+. It has been hypothesized that BIFs formed during to the initial evolution of photosynthetic organisms that had phases of population growth, causing over production of oxygen. Due to this over production they would poison themselves causing a mass die off, which would cut off the source of oxygen and produce a large amount of CO2 through the decomposition of their bodies, allowing for another bacterial bloom. After 1.8 Ga sulfate concentrations were sufficient to increase rates of sulfate reduction to greater than the delivery flux of iron to the oceans.

Along with the disappearance of BIF, the end of the Paleoproterozoic also marks the first large scale sedimentary exhalative deposits showing a link between mineralization and a likely increase in the amount of sulfate in sea water. In the Paleoproterozoic the sulfate in seawater had increased to an amount greater than in the Archean, but was still lower than present day values. The sulfate levels in the Proterozoic also act as proxies for atmospheric oxygen because sulfate is produced mostly through weathering of the continents in the presence of oxygen. The low levels in the Proterozoic simply imply that levels of atmospheric oxygen fell between the abundances of the Phanerozoic and the deficiencies of the Archean.

750 million years ago (Ma) there is a renewed deposition of BIF which marks a significant change in ocean chemistry. This was likely due to snowball earth episodes where the entire

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globe including the oceans was covered in a layer of ice cutting off oxygenation. In the late Neoproterozoic high carbon burial rates increased the atmospheric oxygen level to >10% of its present day value. In the Latest Neoproterozoic another major oxidizing event occurred on Earth’s surface that resulted in an oxic deep ocean and possibly allowed for the appearance of multicellular life.

During the last 600 million years, seawater SO4 has varied between +10 and +30‰ in δ34S, with an average value close to that of today. This coincides with atmospheric O levels reaching something close to modern values around the Precambrian–Cambrian boundary.

Over a shorter time scale (ten million years) changes in the sulfur cycle are easier to observe and can be even better constrained with oxygen isotopes. Oxygen is continually incorporated into the sulfur cycle through sulfate oxidation and then released when that sulfate is reduced once again. Since different sulfate sources within the ocean have distinct oxygen isotopic values it may be possible to use oxygen to trace the sulfur cycle. Biological sulfate reduction preferentially selects lighter oxygen isotopes for the same reason that lighter sulfur isotopes are preferred. By studying oxygen isotopes in ocean sediments over the last 10 million years were able to better constrain the sulfur concentrations in sea water through that same time. They found that the sea level changes due to Pliocene and Pleistocene glacial cycles changed the area of continental shelves which then disrupted the sulfur processing, lowering the concentration of sulfate in the sea water. This was a drastic change as compared to preglacial times before 2 million years ago.

Economic importance

Sulfur is intimately involved in production of fossil fuels and a majority of metal deposits because of its ability to act as an oxidizing or reducing agent. The vast majority of the major mineral deposits on Earth contain a substantial amount of sulfur inclusion to: sedimentary exhalative deposits* (SEDEX), Mississippi Valley-Type (MVT) and copper porphyry deposits*. Iron sulfides, galena and sphalerite will form as by-products of hydrogen sulfide, as long as the respective transition or base metals are present or transported to a sulfate reduction site. If the system runs out of reactive hydrocarbons economically viable elemental sulfur deposits may form. Sulfur also acts as a reducing agent in many natural gas reservoirs.

* Sedimentary exhalative deposits (SedEx deposits) are ore deposits which are interpreted to have been formed by release of ore-bearing hydrothermal fluids into a water reservoir (usually the ocean), resulting in the precipitation of stratiform ore. SedEx deposits are the most important source of lead, zinc and barite. Porphyry copper deposits are copper orebodies from magma to rock typically contain between 0.4 and 1% copper with smaller amounts of other metals such as molybdenum, silver and gold.

Important sources of sulfur in ore deposits are generally deep-seated, but they can also come from local country rocks, sea water, or marine evaporites. The presence or absence of sulfur is one of the limiting factors on both the concentration of precious metals and its precipitation from solution. pH, temperature and especially redox states determine whether sulfides will precipitate. Most sulfide brines will remain in concentration until they reach reducing conditions, a higher pH or lower temperatures.

Ore fluids are generally linked to metal rich waters that have been heated within a sedimentary basin under the elevated thermal conditions typically in extensional tectonic settings. The redox

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conditions of the basin lithologies exert an important control on the redox state of the metal-transporting fluids and deposits can form from both oxidizing and reducing fluids. Metal-rich ore fluids tend to be by necessity comparatively sulfide deficient, so a substantial portion of the sulfide must be supplied from another source at the site of mineralization. Bacterial reduction of seawater sulfate or a euxinic (anoxic and H2S-containing) water column is a necessary source of that sulfide. When present, the δ34S values of barite are generally consistent with a seawater sulfate source, suggesting barite formation by reaction between hydrothermal barium and sulfate in ambient seawater.

Once fossil fuels or precious metals are discovered and either burned or milled, the sulfur become a waste product which must be dealt with properly or it can become a pollutant. There has been a great increase in the amount of sulfur in our present day atmosphere because of the burning of fossil fuels. Sulfur acts as a pollutant and an economic resource at the same time.

Human Impact

Human activities have had a major effect on the global sulfur cycle. The burning of coal, natural gas, and other fossil fuels has greatly increased the amount of S in the atmosphere and ocean and depleted the sedimentary rock sink. Without human impact sulfur would stay tied up in rocks for millions of years until it was uplifted through tectonic events* and then released through erosion and weathering processes. Instead it is being drilled, pumped and burned at a steadily increasing rate. Over the most polluted areas there has been a 30-fold increase in sulfate deposition.

* Tectonic events are the group name for such occurrences as earthquakes, volcanoes and tsunamis. They are called tectonic because they are all associated with earth movements of one kind or another.

Although the sulfur curve shows shifts between net sulfur oxidation and net sulfur reduction in the geologic past, the magnitude of the current human impact is probably unprecedented in the geologic record. Human activities greatly increase the flux of sulfur to the atmosphere, some of which is transported globally. Humans are mining coal and extracting petroleum from the Earth’s crust at a rate that mobilizes 150 x 1012 gS/yr, which is more than double the rate of 100 years ago. The result of human impact on these processes is to increase the pool of oxidized sulfur (SO4) in the global cycle, at the expense of the storage of reduced sulfur in the Earth’s crust. Therefore, human activities do not cause a major change in the global pools of S, but they do produce massive changes in the annual flux of S through the atmosphere.

When SO2 is emitted as an air pollutant, it forms sulfuric acid through reactions with water in the atmosphere. Once the acid is completely dissociated in water the pH can drop to 4.3 or lower causing damage to both man-made and natural systems. According to the EPA, acid rain is a broad term referring to a mixture of wet and dry deposition (deposited material) from the atmosphere containing higher than normal amounts of nitric and sulfuric acids. Distilled water (water without any dissolved constituents), which contains no carbon dioxide, has a neutral pH of 7. Rain naturally has a slightly acidic pH of 5.6, because carbon dioxide and water in the air react together to form carbonic acid, a very weak acid. Roughly 2/3 of all SO2 and 1/4 of all NO3

come from electric power generation that relies on burning fossil fuels, like coal.

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Iron Cycle or Transformation

Iron exists in nature either as ferrous (Fe++) or ferric (Fe+++) ions. Ferrous iron is oxidized spontaneously to ferric state, forming highly insoluble ferric hydroxide. Plants as well as microorganisms require traces of iron, manganese copper, zinc, molybdenum, calcium boron, cobalt etc. Iron is always abundant in terrestrial habitats, and it is oftenly in an unavailable form for utilization by plants and leads to the serious deficiency in] plants.

General aspects of the iron cycleRedox transformations of iron, as well as dissolution and precipitation and thus mobilization and redistribution, are caused by chemical and to a significant extent by microbial processes (Fig. 1). Microorganisms catalyze the oxidation of Fe(ll) under oxic or anoxic conditions as well as the reduction of Fe(Ill) in anoxic habitats. Microbially influenced transformations of iron are often much faster than the respective chemical reactions. They take place in most soils and sediments, both in freshwater and marine environments, and play an important role in other (bio)geochemical cycles, in particular in the carbon cycle. Microbial iron cycling impacts the rate of both organic and inorganic pollutants, including those released from industrial and mining areas (Thamdrup 2000; Straub et al. 2001; Comell and Schwertmann 2003).

Soil microorganisms play important role in the transformations of iron in al number of distinctly different ways such as:

A. Certain bacteria oxidize ferrous iron to ferric state which precipitate as ferric hydroxide around cells

B. Many heterotrophic species attack on in soluble organic iron salts and convert into inorganic salts

C. Oxidation-reduction potential decreases with microbial growth and which leads to the formation of more soluble ferrous iron from highly insoluble ferric ions

D. Number of bacteria and fungi produce acids such as carbonic, nitric, Sulphuric and organic acids which brings iron into solution

E. Under anaerobic conditions, the sulfides formed from sulphate and organic sulphur compounds remove the iron from solution as ferrous sulfide

F. As microbes liberate organic acids and other carbonaceous products of metabolism which results in the formation of soluble organic iron complex. Thus, iron may be precipitated in nature and immobilized by iron oxidizing bacteria under alkaline soil reaction and on the other hand solubilization of iron may occur through acid] formation.

Some bacteria are capable of reducing ferric iron to ferrous which lowers the oxidation-reduction potential of the environment (eg. Bacillus, Clostridium, Klebsiella etc). However, some chemoautotrophic iron and sulphur bacteria such as Thiobacillus ferroxldans andFerrobacitlus ferrooxidans can oxidize ferrous iron to ferric hydroxide which accumulates around the cells.

Most of the aerobic microorganisms live in an environment where iron exists in the oxidized, insoluble ferric hydroxide form. They produce iron-binding compounds in order to take up ferric

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iron. The iron-binding or chelating compounds / ligands produced by microorganisms are called"Siderophores". Bacterial siderophores may act as virulence factors in pathogenic bacteria and thus, bacteria that secrete siderophores are more virulent than non- siderophores producers. Therefore, siderophore producing bacteria can be used as biocontrol agents eg. Fluorescent pseudomonads used to control Pythium, causing damping-off diseases in seedlings. Recently Vascular - Arbusecular – Mycorrhiza (VAM) has been reported to increase uptake of iron.

Solubility and chemical transformation of Fe(II) and Fe(III) mineralsDifferent Fe(ll), Fe(Ill) and mixed Fe(ll)-Fe(Ill) minerals are found in the environment and many are used, produced or transformed by microbial activities (Table 1). Fe(Ill) minerals are characterized by low solubility at circumneutral pH and usually only very low, hardly detectable concentrations in the range of 10-9 M of Fe(III) are present in solution (Fig. 2).

Figure 1. Microbial and chemical iron cycle

However, colloid formation or complexation by organic compounds can lead to elevated concentrations of dissolved Fe(III), even at neutral pH (Comell and Schwertmann 2003; Kraemer 2004). At strongly alkaline or strongly acidic pH, ferric iron oxides can be dissolved because of their amphoteric character. Ferric iron oxides can be reduced chemically by a range of organic and inorganic reductants. However, the environmentally most important reducing agent for Fe(III) is hydrogen sulfide, which is a common end product of microbial sulfur and sulfate reduction (Thamdrup 2000; Comell and Schwertmann 2003).

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The manganese cycle or transformation

Geochemical cycling of manganese

The carbon and nitrogen cycles may get most of the attention in high school science class, but the manganese cycle also has a tremendous impact on the Earth's geochemistry. Manganese is essential for life as a cofactor in enzymes such as Photosystem II and manganese superoxide dismutase and is one of the most abundant redox-active transition metals in the Earth's crust [1]. In Nature, manganese is commonly found in three oxidation states: Mn(II), Mn(III) and Mn(IV) [2]. The reduced form, Mn(II), is generally soluble and stable in the absence of oxygen. Mn(IV), the most oxidized form, is insoluble, forming oxides which are strong oxidants, capable of oxidizing inorganic and organic compounds, for example, Fe(II) and UO2 [3]. The intermediate oxidation state, Mn(III), is unstable as an ion under normal environmental conditions unless complexed with organic or inorganic ligands (see below) [1]. Mn(III) also occurs as insoluble Mn(III) oxy(hydrox)oxide phases or in mixed Mn(III,IV) oxides. Mn(III,IV) oxides also sorb metals and other compounds from the environment; as a result, Mn(II) oxidation can control the distribution of many other elements (e.g. copper, cobalt, nickel, lead, iron, radium, uranium and rare earth elements).A one-electron transfer for the oxidation of the soluble Mn(II) species Mn(H2O)6

2+ by O2 is thermodynamically unfavourable at pH < 9, but a two-electron oxidation is favourable at pH > 3 [6]. Complexed forms of Mn(II) can oxidize at neutral pH, but abiotic Mn(II) oxidation at circumneutral pH is quite slow under normal environmental conditions; bacteria and fungi can greatly increase the rate of this reaction by up to 5 orders of magnitude [1]. Thus it is thought that the bulk of environmental Mn(II) oxidation is carried out by micro-organisms. In the present mini-review, we focus specifically on Mn(II)-oxidizing bacteria.Mn(II)-oxidizing bacteria are widespread and found in diverse environments such as soil, natural waters and sediments. These organisms are phylogenetically diverse, with representatives in the Firmicutes, Actinobacteria and the Alpha-, Beta- and Gamma-proteobacteria [7]. The physiological role Mn(II) oxidation plays in these species is unknown. Because Mn(IV) formation is thermodynamically favourable, the bacteria could derive energy from the reaction; however, this has not been shown conclusively for any organism. The oxidation of Mn(II) may also help to protect the cells from ROS (reactive oxygen species) or other free radicals [8]. Alternatively, the Mn(IV) oxides formed in this reaction are themselves highly reactive and may be used to oxidize refractory organic material that can then be utilized by the micro-organism as a carbon source. Other possible functions for the oxides are as terminal electron acceptors, manganese storage or, since the solids coat the cell, protection from environmental hazards such as UV radiation, predation or phage infection.Mn(III,IV) oxide deposits can be found in many environments [1,10]. These include ferromanganese nodules in the deep sea, lakes and soils and terrestrial ferromanganese crusts, also called rock varnish. Mn(IV) oxides are also found in ore deposits and metalliferous sediments associated with spreading centres and as ferromanganese crusts on seamounts on the ocean floor. Given that the oxides reflect the redox conditions of the local environment, they are unsurprisingly often found associated with oxic–anoxic interfaces in sediments and water. The extent to which microbial activity is responsible for these formations is unclear. However, the ubiquity of Mn(II)-oxidizing organisms and their presence in environments where Mn(IV) oxides are found suggest that their role is substantial.

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Molecular biochemistry of Mn(II) oxidation

Despite the diversity of organisms capable of the reaction, there are common themes in bacterial Mn(II) oxidation [9]. The product is primarily Mn(IV)O2, with oxidation occurring through two sequential one-electron steps [Mn(II)→Mn(III)→Mn(IV)] that require O2 [11,12] (Figure 1). Organisms commonly initiate Mn(II) oxidation at the onset of stationary phase and deposit the oxides on their outer surface [1,9]. However, some variations on these common themes occur. Both fungal Mn(II)-oxidizers and the alphaproteobacterium Erythrobacter sp. SD-21 produce Mn(III), which is stabilized by biogenic ligands [13–16]. Another less direct mechanism of Mn(II) oxidation is through the action of enzymatically produced superoxide, as has been observed with the alphaproteobacterium Roseobacter sp. AzwK-3b [17–19]. The iron-chelating siderophores produced by bacteria under iron-limiting conditions are also able to bind with high affinity to manganese; this interaction itself can promote Mn(II) oxidation [20–23].

Fig. The updated manganese cycle

The enzymes responsible for bacterial Mn(II) oxidation have been identified in several species and fall into two general categories. The alphaproteobacteria Aurantimonas manganoxydansSI85-9A1 and Erythrobacter sp. SD-21 both employ calcium-binding haem peroxidases named MopA to oxidize Mn(II) [16]. Many other species possess MCO (multicopper oxidase) Mn(II) oxidase enzymes (Figure 2). These include MofA in Leptothrix sp. and the twin Mn(II) oxidase MCOs of Pseudomonas putida 

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ORGANIC MATTER

ORGANIC MATTER - Organic matter is defined as a grouping of carbon containing compounds which have originated from living beings and been deposited on or within the earth’s structural components. Soil organic matter includes the remains of all plant and animal bodies which have fallen on the earth’s surface or been purposely applied by man in organically synthesized forms eg, pesticides. A fertile soil should contain from 2-8 percent organic matter; most soils contain less than 2%. When organic matter is burned, there remains a residual ash. The residual ash is composed of the minerals and trace elements required by plants and animals during their normal growth processes. Thus organic matter contains mineral elements required by plants.An accurate measurement of the soil organic matter would be helpful in monitoring soil fertility. Currently the best extractant for removing organic matter from a soil is 0.5N sodium hydroxide (NaOH) (working under N2). The second best extractant is sodium pyrophosphate decahydrate (Na4P2O7.10H2O at pH 9.8). Neither one of these extractants is able to remove all of the organic matter from a soil sample. Obviously since these chemicals are the best known it is impossible to determine the exact amount of organic matter present within a soil. In reality soil organic matter is not a measurable soil component. The organic matter content of a soil sample, reported on soil tests, is only an estimate. The organic carbon content of a soil can be measured and would be a much more valuable indication of the potential humic-chemistry of a soil. The soil’s carbon content would be a desirable part of a soil test report.

The term "humus" dates back to Roman times when the term was commonly used to designate the soil as a whole.  In 1761 Wallerius first defined "humus" in terms of decomposed organic matter.  However, the prevailing ideas concerning the chemical nature of humus and the mechanism of its formation at that time were still very vague.   Most often it was considered as a complex formed in soils, in bogs, or in composts, from plant residues, by a special process of "humification". The famous work of De Saussure, "Recherches Chimiques Sur La Vegetation", devotes considerable attention to humus.   He reasoned that it is not a homogeneous substance, but that it consists of various readily removable complexes.  They are differentiated between "mild humus", formed in the presence of sufficient oxygen, and "acid humus" or peat, formed with limited admission of oxygen. Thus, the term "humus" came into general use at a time when organic chemistry was still in its infancy. We now regard most organic and inorganic compounds as more complex   substances rather than of simple in chemical composition.In leached and acid soils, which are often sandy, substantial portions of the organic matter is in the form of plant debris and fulvic acids (FAs). In neutral and alkaline soils a large percentage of the organic matter is in the form of humic acids (HAs) and humin.

Humification(Microbial Soil Organic Matter Decomposition)Soil microorganisms exist in large numbers in the soil as long as there is a carbon source for energy. A large number of bacteria in the soil exist, but because of their small size, they have a smaller biomass. Actinomycetes are a factor of 10 times smaller in number but are larger in size so they are similar in biomass to bacteria. Fungus population numbers are smaller but they dominate the soil biomass when the soil is not disturbed. Bacteria, actinomycetes, and protozoa are hardy and can tolerate more soil disturbance than fungal populations so they dominate in tilled soils while fungal and nematode populations tend to dominate in untilled or no-till soils.

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There are more microbes in a teaspoon of soil than there are people on the earth. Soils contain about 8 to 15 tons of bacteria, fungi, protozoa, nematodes, earthworms, and arthropods.

Table 1: Relative number and biomass of microbial species at 0–6 inches (0–15 cm) depth of soilMicroorganism Number/g soil Biomass (g/m2)Bacteria 108-109 40-500Actinomycetes 107-108 40-500Fungi 105-106 100-1500Algae 104-105 1-50Protozoa 103-104 VariesNematodes 102-103 Varies

Organic matter decomposition serves two functions for the microorganisms, providing energy for growth and supplying carbon for the formation of new cells. Microbes need regular supplies of active (readily decomposable) SOM in the soil to survive in the soil. Long-term no-tilled soils have significantly greater levels of microbes, more active carbon, more SOM, and more stored carbon than conventional tilled soils. A majority of the microbes in the soil exist under starvation conditions and thus they tend to be in a dormant state, especially in tilled soils.

Decomposition of soil organic matter

Soil organic matter can be broken down into its component parts. One hundred grams (g) or 100 pounds (lbs) of dead plant material yields about 60–80 g (lbs) of carbon dioxide, which is released into the atmosphere. From decomposition of the remaining 20–40 g (lbs), energy and nutrients is released and turned into about 3–8 g (lbs) of microorganisms (the living), 3–8 g (lbs) of non-humic compounds (the dead), and 10–30 g (lbs) of humus (the very dead matter, resistant to decomposition).

Soil Organic Matter NutrientsThe molecular structure of SOM is mainly carbon and oxygen with some hydrogen and nitrogen and small amounts of phosphorus and sulfur. Soil organic matter is a by-product of the carbon and nitrogen cycles.

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Consider the following three scenarios. Soils typically turnover 1 to 3% of nitrogen stored in SOM. Tilled or unhealthy soils release a lower percent of nitrogen due to lower microbial activity. A tilled soil with 2% SOM (2,000 lbs of N) may release 1% N or 20 lbs of N per year. A soil that is more biologically active and has 4% SOM (4,000 lbs N) may release 1.5% N or 60 lbs N while a 6% SOM soil (6,000 lbs N) may release 2% N or 120 lbs of N [(6%/100)x6000 lbsx2%=120 lbs]. In tilled soils, excess nutrients released are often lost and the carbon stores are depleted so that future storage of nutrients is reduced. Farmers often see this occur when they till a virgin soil, an old pasture, or a fence row. For several years, crops on the newly tilled soil will grow better than the surrounding soils, but over time the soil will be depleted of carbon and the newly tilled soil will become less fertile because the carbon is oxidized as carbon dioxide and lost to the atmosphere. Tillage results in the oxidation and destruction of carbon in the soil by increasing the soil oxygen levels, thereby promoting bacteria populations to expand and consume active carbon in the soil.

Factors effecting SOM Climate, Temperature, and pH SOM is affected by climate and temperature. Rate of microbial biochemical reactions doubles with every 10oF change in temperature. In the tropics, the topsoil has very little SOM because high temperatures and moisture quickly decompose SOM. Moving north or south from the equator, SOM increases in the soil. The tundra near the Arctic Circle has a large amount of SOM because of cold temperatures. Moisture, pH, soil depth, and particle size affect SOM decomposition. Hot, humid regions store less organic carbon in the soil than dry, cold regions due to increased microbial decomposition. The rate of SOM decomposition increases when the soil is exposed to cycles of drying and wetting compared to soils that are continuously wet or dry. Other factors being equal, soils that are neutral to slightly alkaline in pH decompose SOM quicker than acid soils; therefore, liming the soil enhances SOM decomposition and carbon dioxide evolution. Decomposition is also greatest near the soil surface where the highest concentration of plant residues occur. At greater depths there is less SOM decomposition, which parallels a drop in organic carbon levels due to less plant residues. Small particle sizes are more readily degraded by soil microbes than large particles because the overall surface area is larger with small particles so that the microbes can attack the residue.Hardwood forests and tree tap roots are high in lignin, and deciduous trees left large amounts of leaf litter on the soil surface. Hardwood tree roots do not turn over quickly so organic matter levels in the subsoil are fairly low. In forest soils, most of the SOM is distributed in the top few inches. In tall grassland prairies, the landscape and topsoil are formed from deep fibrous grass root systems. Fifty percent of a grass root dies and is replaced every year and grass roots are high in sugars and protein (higher active organic matter) and lower in lignin. So soils that formed under tall grass prairies are high in SOM throughout the soil profile. These prime soils are highly productive because they have higher percentage of SOM (especially active carbon), hold more nutrients, contain more microbes, and have better soil structure due to larger fungal populations.

Carbon to Nitrogen RatioThe break down of organic residues by microbes is dependent upon the carbon to nitrogen (C:N) ratio. Microbes in a cow’s rumen, a compost pile, and soil microbes rely on the C:N ratio to break down organic (carbon-based) residues. Consider two separate feed sources, a young tender alfalfa plant and oat or wheat straw. A young alfalfa plant has more crude protein, amino acids, and sugars in the stalk so it is easily digested by microbes whether it is in a cow’s rumen, a compost pile, or in the soil. Young alfalfa has a high nitrogen content from protein (amino

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acids and proteins are high in nitrogen and sulfur), so it has a lower carbon to nitrogen ratio (less carbon, more nitrogen). However, oat and wheat straw (or older mature hay) has more lignin (which is resistant to microbial decomposition), lower crude protein, and less sugars in the stalk and a higher C:N ratio. Straw is decomposed by microbes but it takes additional time and nitrogen to break down this high carbon source.A low nitrogen content or a wide C:N ratio is associated with slow SOM decay. Immature or young plants have a higher nitrogen content, lower C:N ratios and faster SOM decay. For good composting, a C:N ratio less than 20 allows the organic materials to decompose quickly (4 to 8 weeks) while a C:N ratio greater than 20 requires additional N and slows down decomposition. So if a high C based material with low N content is added to the soil, the microbes will tie up soil nitrogen. Eventually, the soil N is released but in the short-term the N is tied up. The conversion factor for converting N to crude protein is 16.7, which relates back to why it is so important to have a C:N ratio of less than 20. The C:N ratio of most soils is around 10:1 indicating that N is available to the plant. The C:N ratio of most plant residues tends to decrease with time as the SOM decays. This results from the gaseous loss of carbon dioxide. Therefore, the percentage of nitrogen in the residual SOM rises as decomposition progresses. The 10:1 C:N ratio of most soils reflects an equilibrium value associated with most soil microbes (Bacteria 3:1 to 10:1, Fungus 10:1 C:N ratio). Bacteria are the first microbes to digest new organic plant and animal residues in the soil. Bacteria typically can reproduce in 30 minutes and have high N content in their cells (3 to 10 carbon atoms to 1 nitrogen atom or 10 to 30% nitrogen). Under the right conditions of heat, moisture, and a food source, they can reproduce very quickly. Bacteria are generally less efficient at converting organic carbon to new cells. Aerobic bacteria assimilate about 5 to 10 percent of the carbon while anaerobic bacteria only assimilate 2 to 5 percent, leaving behind many waste carbon compounds and inefficiently using energy stored in the SOM.

Graph of Relative Available N with Length of Time for Decomposition

Fungus generally release less carbon dioxide into the atmosphere and are more efficient at converting carbon to form new cells. The fungus generally captures more energy from the SOM as they decompose it, assimilating 40 to 55 percent of the carbon. Most fungi consume organic matter higher in cellulose and lignin, which is slower and tougher to decompose. The lignin content of most plant residues may be of greater importance in predicting decomposition velocity than the C:N ratio.Mycorrhizal fungi live in the soil on the surface of or within plant roots. The fungi have a large surface area and help in the transport of mineral nutrients and water to the plants. The fungus life cycle is more complex and longer than bacteria. Fungi are not as hardy as bacteria, requiring a more constant source of food. Fungi population levels tend to decline with

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conventional tillage. Fungi have a higher carbon to nitrogen ratio (10:1 carbon to nitrogen or 10% nitrogen) but are more efficient at converting carbon to soil organic matter. With high C:N organic residues, bacteria and fungus take nitrogen out of the soil (see the graph on net immobilization).Protozoa and nematodes consume other microbes. Protozoa can reproduce in 6–8 hours while nematodes take from 3 days to 3 years with an average of 30 days to reproduce. After the protozoa and nematodes consume the bacteria or other microbes (which are high in nitrogen), they release nitrogen in the form of ammonia (see the graph on net mineralization). Ammonia (NH4

+) and soil nitrates (NO3-) are easily converted back and forth in the soil. Plants absorb

ammonia and soil nitrates for food with the help of the fungi mycorrhizal network.Microorganism populations change rapidly in the soil as SOM products are added, consumed, and recycled. The amount, the type, and availability of the organic matter will determine the microbial population and how it evolves. Each individual organism (bacteria, fungus, protozoa) has certain enzymes and complex chemical reactions that help that organism assimilate carbon. As waste products are generated and the original organic residues are decomposed, new microorganisms may take over, feeding on the waste products, the new flourishing microbial community (generally bacteria), or the more resistant SOM. The early decomposers generally attack the easily digested sugars and proteins followed by microorganisms that attack the more resistant residues.

Decomposition of Cover Crop Residues: Cowpeas with a low C:N ratio (<20) will decompose in 4 to 8 weeks and result in net mineralization or release of N. Sudan grass or cereal rye with a higher C:N ratio (>38) will decompose slowly (3 months to 1 year or more) and will result in net immobilization or will tie up soil N.

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Graph of Cowpeas (C:N<20) being decomposed by bacteria and fungus, the carbon dioxide evolution and protozoa and nematodes consuming the bacteria and fungus and excreting ammonia into the soil for plant growth. NO3

- and NH4

+ are easily converted in the soil.

Cover crops supply food (active carbon like glucose and proteins) to the microbes to feed on. In the soil, there are 1,000 to 2,000 times more microbes associated with roots than are living in bare or tilled soil. The microbes in turn build SOM and store soil nutrients. Building SOM requires soil nutrients like N-P-K-S to be tied up in the soil. Winter cover crops soak up excess soil nutrients and supply food to all the microbes in the soil during the winter months rather than microbes having to use up SOM reserves for nutrients. In a conventional tilled field, soil nutrients are quickly released as SOM is burned up and the microbes and soil organisms habitat are destroyed. In a no-till field, high levels of SOM are reserves of soil nutrients which are slowly released into the soils. Adding a living cover crop to a no-till field increases active organic matter (sugars and proteins) for the soil microbes. Soil microbes have two crops to feed on instead of one crop per year. Microbes thrive under no-till conditions and winter cover crops. Cover crops and manure can be used to feed soil microbes and recycle soil nutrients. As soil microbes decompose organic residues, they slowly release nutrients back into the soil for the winter cover crops or for the preceding crop. Cover crops prevent the nutrients from being lost through soil erosion, leaching, volatilization, or denitrification.

Transformation of organic matter into humusThe process of "humification" can occur naturally in soil, or in the production of compost. The importance of chemically stable humus is thought to be the fertility it provides to soils in both a physical and chemical sense, though some agricultural experts put a greater focus on other features of it, such as its ability to suppress disease. It helps the soil retain moisture by increasing microporosity, and encourages the formation of good soil structure. The incorporation of oxygen into large organic molecular assemblages generates many active, negatively charged sites that bind to positively charged ions (cations) of plant nutrients, making them more available to the plant by way of ion exchange. Humus allows soil organisms to feed and reproduce, and is often described as the "life-force" of the soil. Yet, it is difficult to define humus precisely; it is a highly complex substance, which is still not fully understood. Humus should be differentiated from decomposing organic matter in that the latter is rough-looking material, with the original plant remains still visible, whereas fully humified organic matter is uniform in appearance (a dark, spongy, jelly-like substance) and amorphous in structure, and may remain such for millennia or more. It has no determinate shape, structure or character. However, humified organic matter, when examined under the microscope may reveal tiny plant, animal or microbial remains that have been mechanically, but not chemically, degraded. This suggests a fuzzy boundary between humus and organic matter. In most literature, humus is clearly considered as an integral part of soil organic matter. Plant remains (including those that passed through an animal gut and were excreted as faeces) contain organic compounds: sugars, starches, proteins, carbohydrates, lignins, waxes, resins, and organic acids. The process of organic matter decay in the soil begins with the decomposition of sugars and starches from carbohydrates, which break down easily as detritivores initially invade the dead plant organs, while the remaining cellulose and lignin break down more slowly. Simple proteins, organic acids, starches and sugars break down rapidly, while crude proteins, fats, waxes and resins remain relatively unchanged for longer periods of time. Lignin, which is quickly transformed by white-rot fungi, is one of the main precursors of humus, together with by-products of microbial and animal activity. The end-product of this process, the humus, is thus a mixture of compounds and complex life chemicals of plant,

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animal, or microbial origin that has many functions and benefits in the soil. Earthworm humus (vermicompost) is considered to be the best organic manure there is.

Stability of humusCompost that is readily capable of further decomposition is sometimes referred to as effective or active humus, though scientists would say that, if it is not stable, it is not humus at all. This kind of compost, rich in plant remains and fulvic acids, is an excellent source of plant nutrients, but of little value with respect to long-term soil structure and tilth. Stable (or passive) humus consists of humic acids and humins, which are so highly insoluble, or so tightly bound to clay particles and hydroxides, that they cannot be penetrated by microbes and are greatly resistant to further decomposition. Thus stable humus adds few readily available nutrients to the soil, but plays an essential part in providing its physical structure. Some very stable humus complexes have survived for thousands of years. The most stable humus is that formed from the slow oxidation of black carbon, after the incorporation of finely powdered charcoal into the topsoil. This process is at the origin of the formation of the fertile Amazonian dark earths or Terra preta do Indio.

Benefits of soil organic matter and humus The process that converts raw organic matter into humus feeds the soil population of

microorganisms and other creatures, thus maintains high and healthy levels of soil life. The rate at which raw organic matter is converted into humus controls (promote/fast or

limit/slow) the existence of plants, animals, and microbes in soil. Effective humus and stable humus are further sources of nutrients to microbes, the former

provides a readily available supply, and the latter acts as a longer-term storage reservoir. Decomposition of dead plant material causes complex organic compounds to be slowly

oxidized (lignin-like humus) or to break down into simpler forms (sugars and amino sugars, aliphatic, and phenolic organic acids), which are further transformed into microbial biomass (microbial humus), and further oxidized into humic assemblages (fulvic and humic acids), which bind to clay minerals and metal hydroxides. There has been a long debate about the ability of plants to uptake humic substances from their root systems and to metabolize them. There is now a consensus about how humus plays a hormonal role rather than simply a nutritional role in plant physiology.

Humus is a colloidal substance, and increases the soil's cation exchange capacity, hence its ability to store nutrients by chelation. While these nutrient cations are accessible to plants; they are held coated around the soil particles and safe from being leached by rain or irrigation.

Humus can hold moisture equivalent to 80–90% of its weight, and therefore increases the soil's capacity to withstand drought conditions.

The biochemical structure of humus enables it to moderate – or buffer – excessive acid or alkaline soil conditions.

During the humification process, microbes excrete sticky gum-like mucilages; these contribute to the crumb structure (tilth) of the soil by holding particles together, and allowing greater aeration of the soil. Toxic substances such as heavy metals, as well as excess nutrients, can be chelated (that is, bound to the complex organic molecules of humus) and so prevented from entering the wider ecosystem.

The dark color of humus (usually black or dark brown) helps to warm up cold soils in the spring.

A soil’s ability to retain water and stimulate plant growth depends on the soil organic matter (SOM) and especially on its humic substances (HSs) fraction. Arable soils contain up to 10%

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SOM, and HSs typically account for 80% of the SOM. Figure 1 shows the major precursors of soil organic matter. Dead leaves, a major soil input, consist of 50–60% cellulose, 15–20% lignins, and 15–20% lipids. Soil chemical and microbiological oxidation of dead animals and plants (humification) initially is exothermic, but then becomes slow synthesis and degradation of HSs. Degradation of HSs ultimately leads to coal (mostly aromatic), crude oil (mostly aliphatic), and carbon dioxide. The aromatization of aliphatic soil components or HSs to yield coal is an oxidation (dehydrogenation) process. The CO2 product of HS respiration completes the carbon cycle, prevents the earth from being covered with HS soup, and reminds us that HSs are long lived but eventually transient on geological time scales.

Figure 1. Typical soil inputs to humus

Soil has an important role in the air–soil–water cycle. HAs are among the most active components of soil. Consider just a few of the facts that involve humic materials in that cycle. HSs contain more carbon than all living things. Soil respiration contributes much more to global CO2 levels than fossil fuel combustion for

heating and transportation. Solid HSs act as pH buffers, metal binders, solute sorbents, and redox catalysts, and they

are photosensitizers. No other natural materials have so many functions in so many different places.

HSs are more versatile than any other synthetic or natural material and they are biodegradable and non-allergenic if free from harmful metals, xenobiotics and microorganisms.

The deserts are growing, populations are exploding, and huge amounts of soil and humic substances wash away every year by erosion. HSs lost from soils need to be replenished.

Long-term intensive farming depletes SOM. Solutions to SOM loss include promoting organic farming, replacing incinerators with waste composting, and seeking alternative natural sources of HSs, the pivotal components of the air–soil–water system.

Small wonder that this class of natural substances has been the object of so much study over the last 100 years. Scientific interest in humic substances is continually expanding, despite inherent natural obstacles. For example, it is important to know which components of HSs are responsible for some of the major processes in soils and waters, and the proportions of these components in a humic source. For meaningful structural studies, it is desirable to deal with pure substances. However, HSs are gross mixtures whose separation into discrete substances is a continuing challenge.

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