**1. Introduction (Histosols)**

Organic soils are soils which have diagnostic horizons with more than 20% organic matter and essentially reside in marshes, bogs, and swamps where anaerobic soil conditions support a low rate of organic matter decomposition relative to the rate of organic matter production. Thus, organic soils are observed to have a carbon input rate that is initially greater than the carbon loss rate resulting in an annual carbon accumulation, then with continued soil genesis the rate of carbon input approximately equals the carbon loss rate and a carbon quasi-equilibrium is attained.

These organic soils are frequently associated with extremely wet landscapes, or extremely acidic soils, or soils lacking available nutrients or some combination of these influences. Organic soils (Histosols) as defined in the United States [1] are soils that have an abundance of organic soil materials with additional criteria specifying that they lack sufficient andic properties and lack permafrost plus these soils possess certain thickness, water saturation duration, and decomposition status associated with their fresh and rubbed fiber contents. According to the United States Keys of Soil Taxonomy [1], organic soil horizons have (i) 12% organic carbon (approximately 21% soil organic matter) if the clay content is 0% and (ii) 18% organic carbon if the clay content is 60% or greater. For horizons that have clay contents between 0 and 60% the organic carbon content is a linear relationship to clay content involving the 12% organic carbon if the clay content is 0% and 18% organic carbon if the clay content is 60%.

Histic epipedons are surface organic horizons that are water saturated for at least 30 days in most years (typically an aquic soil moisture regime) are generally 0.2 to 0.4 m thick and have sufficient organic carbon as a function of clay content. Folistic epipedons are surface horizons that are not water saturated for at least 30 days in most years (not artificially drained), typically are more than 0.20 m thick, and are largely composed of 75% or more sphagnum fibers or have a bulk density of less than 0.1 g cm−3. The Keys of Soil Taxonomy [1] partition histic epipedons into fibric, hemic and sapric materials. Fibric materials (Of) are minimally decomposed where three quarters or more of its volume is made up of fibers after rubbing the sample. Sapric materials (Oa) are highly decomposed; less than one-sixth of the volume of sapric material contains fibers after a sample is rubbed. Hemic materials (Oe) are intermediate with respect to decomposition. In general, fibric materials possess a very low bulk density (0.05 to 0.15 Mg m−3), a large total pore space (85%) with a high distribution of large pores spaces, a low bearing capacity, and a hydraulic conductivity ranging from 1.6 to 30 m day−1.

Generally, the Histosol soil order is recognized if more than half of the upper 0.8 m of the soil profile is organic or if organic soil material rests on rock or fragmental material showing interstices filled with organic material. In colloquial terms the Histosol order contains soils formally described as bogs, moors, peatlands, muskegs, fens or are composed of peats and mucks. Histosols make up about 1% of the world's glacier-free land surface (325 to 375 million ha). Suborders of Histosol order are based on the degree of organic material decomposition and the length of water saturation. The Histosol suborders are: Fibrists, Hemists, Saprists and Folists. The World Reference Base for soil resources [2] states that Histosols are soils having a histic or folic horizon either 0.1 m or more thick from the soil surface to a lithic or paralithic contact or 0.4 m or more thick and starting within 0.3 m from the soil surface, and having no andic or vitric horizon starting within 0.3 m of the soil surface.

#### **2. Histosol soil forming processes**

Histosols occur in all latitudes; however, Histosols are particularly common in the boreal zone, a feature Histosols share with Spodisols. The dominant feature of Histosols is the accumulation of organic materials, which may be characterized as:

$$\text{Organic material content} = \text{organic matter input} - \text{organic matter loss} \tag{1}$$

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soil conditions for a sustained time interval to restrict soil organic matter decomposition. Topography influences Histosol formation by directing water flux within the landscape position. Lateral groundwater may create seepage on sideslopes, whereas peatlands may form in poorly-drained basins. Fens occur where surface water inflow or groundwater discharge concentrates nutrient rich water. Pocosins or bogs on coastal plains or interior flatlands are frequently located on slightly raised

The degree of soil organic matter decomposition has a significant influence on soil properties. Buol et al. [3] reviewed literature to describe the soil genesis and classification of Histosols. Key soil properties that are influenced based on the degree of soil organic matter decomposition include: organic carbon, total nitrogen, carbon to nitrogen ratio, cellulose content, pH, cation exchange capacity, bulk density, water contents at field capacity and permanent wilting point, hydraulic conductivity. Upon transition from fibric to sapric soil conditions the following properties typically increase in magnitude: total nitrogen, pH, cation exchange capacity, bulk density, and the water contents at field capacity and permanent wilting point. Most notably the vertical and horizontal hydraulic conductivities decrease on transition from fibric to sapric soil conditions. However, many Histosols exhibit greater soil organic matter decomposition with increasing soil profile depth, thus the corresponding reduced hydraulic conductivity and increased water content at

Buol et al. [3] alluded to two adjacent Histosols in Michigan that differ in nutrient sources. The Napoleon soil series (dysic, mesic Typic Haplohemists) receives nutrients only from precipitation and dry deposition, whereas the Houghton (euic, mesic Typic Haplosaprists) primarily receives nutrients from seepage water that transverses calcareous sandy glacial till. The Napoleon mucky peat has an Oa1-Oa2-Oe1-Oe2 horizon sequence, with all horizons having a pH near 4, whereas the Houghton muck has an Oa1-Oa2-Oa3-Oa4-Oa5-Oa6 horizon sequence with all horizons having a pH near 7. Vegetation associated with the Napoleon mucky peat comprised various maples, swamp white oak, and dogwood, whereas the Houghton muck is vegetated with marshy grasses. Thus, water chemistry dramatically influences the soil's pH and exchangeable cation expression and coupled with hydrology influences vegetation

Aide and Aide (two authors of this manuscript) have unpublished field data of several soil series in northeastern Wisconsin. The Lupton series (Euic, frigid Typic Haplosaprists) are very deep, very poorly-drained organic soils formed in depressions on lake and outwash plains. The horizon sequence is Oa1-Oa2-Oa3-Oa4-Oa5 and has little inorganic material, a very low bulk density, a pH in 0.01 *M* CaCl2 of 5.7 to 6.0 and a cation exchange capacity ranging from 107 to 199 cmol kg−1 across multiple pedons. The dominant surrounding soil consists of pedons of the Padus series (coarse-loamy, mixed, superactive, frigid Alfic Haplorthods). The tupical Padus horizon sequence is A-E-Bs1-Bs2-E/B-B/E-2C. The texture is sandy loam above the lithologic discontinuity and sandy textured at greater depths (2C). These very deep, well-drained and very strongly acidic pedons are moderately deep to stratified sandy outwash with an abundance of clay films in the B material of the E/B and B/E horizons. The organic carbon content of the A horizon is less than 2% and the cation exchange capacity is very low, reflecting the sandy loam texture and diminished quantity of soil organic matter. Water extracts from both soils show an abundance of calcium, reflecting that calcium is the dominant exchange cation. These two soils have very distinctive profiles, whose properties are directly related to the contrast-

ing oxidation–reduction environments imposed by the local hydrology.

Parent materials for Histosols are mostly hydrophytic plants [1]. Sphagnum consists of both living and dead tissue from the genus Sphagnum, with approximately

greater soil profile depth support continuance of the sapric condition.

*DOI: http://dx.doi.org/10.5772/intechopen.94399*

interfluvial positions.

establishment.

The rate of organic matter decomposition in Histosols is usually very slow, a feature attributed to specific conditions of climate, topography and hydrology. In boreal biomes, cool summer temperatures restrict microbial activity, with biologic zero being approximately 4 to 5°C. Low soil temperatures must be further associated with anoxic soil conditions to support Histosol genesis. In tropical climates, warmer temperatures support greater ecosystem productivities; however, the combined effects of precipitation, topography and hydrology may create anoxic

#### *Soil Genesis of Histosols and Gelisols with a Emphasis on Soil Processes Supporting Carbon… DOI: http://dx.doi.org/10.5772/intechopen.94399*

soil conditions for a sustained time interval to restrict soil organic matter decomposition. Topography influences Histosol formation by directing water flux within the landscape position. Lateral groundwater may create seepage on sideslopes, whereas peatlands may form in poorly-drained basins. Fens occur where surface water inflow or groundwater discharge concentrates nutrient rich water. Pocosins or bogs on coastal plains or interior flatlands are frequently located on slightly raised interfluvial positions.

The degree of soil organic matter decomposition has a significant influence on soil properties. Buol et al. [3] reviewed literature to describe the soil genesis and classification of Histosols. Key soil properties that are influenced based on the degree of soil organic matter decomposition include: organic carbon, total nitrogen, carbon to nitrogen ratio, cellulose content, pH, cation exchange capacity, bulk density, water contents at field capacity and permanent wilting point, hydraulic conductivity. Upon transition from fibric to sapric soil conditions the following properties typically increase in magnitude: total nitrogen, pH, cation exchange capacity, bulk density, and the water contents at field capacity and permanent wilting point. Most notably the vertical and horizontal hydraulic conductivities decrease on transition from fibric to sapric soil conditions. However, many Histosols exhibit greater soil organic matter decomposition with increasing soil profile depth, thus the corresponding reduced hydraulic conductivity and increased water content at greater soil profile depth support continuance of the sapric condition.

Buol et al. [3] alluded to two adjacent Histosols in Michigan that differ in nutrient sources. The Napoleon soil series (dysic, mesic Typic Haplohemists) receives nutrients only from precipitation and dry deposition, whereas the Houghton (euic, mesic Typic Haplosaprists) primarily receives nutrients from seepage water that transverses calcareous sandy glacial till. The Napoleon mucky peat has an Oa1-Oa2-Oe1-Oe2 horizon sequence, with all horizons having a pH near 4, whereas the Houghton muck has an Oa1-Oa2-Oa3-Oa4-Oa5-Oa6 horizon sequence with all horizons having a pH near 7. Vegetation associated with the Napoleon mucky peat comprised various maples, swamp white oak, and dogwood, whereas the Houghton muck is vegetated with marshy grasses. Thus, water chemistry dramatically influences the soil's pH and exchangeable cation expression and coupled with hydrology influences vegetation establishment.

Aide and Aide (two authors of this manuscript) have unpublished field data of several soil series in northeastern Wisconsin. The Lupton series (Euic, frigid Typic Haplosaprists) are very deep, very poorly-drained organic soils formed in depressions on lake and outwash plains. The horizon sequence is Oa1-Oa2-Oa3-Oa4-Oa5 and has little inorganic material, a very low bulk density, a pH in 0.01 *M* CaCl2 of 5.7 to 6.0 and a cation exchange capacity ranging from 107 to 199 cmol kg−1 across multiple pedons. The dominant surrounding soil consists of pedons of the Padus series (coarse-loamy, mixed, superactive, frigid Alfic Haplorthods). The tupical Padus horizon sequence is A-E-Bs1-Bs2-E/B-B/E-2C. The texture is sandy loam above the lithologic discontinuity and sandy textured at greater depths (2C). These very deep, well-drained and very strongly acidic pedons are moderately deep to stratified sandy outwash with an abundance of clay films in the B material of the E/B and B/E horizons. The organic carbon content of the A horizon is less than 2% and the cation exchange capacity is very low, reflecting the sandy loam texture and diminished quantity of soil organic matter. Water extracts from both soils show an abundance of calcium, reflecting that calcium is the dominant exchange cation. These two soils have very distinctive profiles, whose properties are directly related to the contrasting oxidation–reduction environments imposed by the local hydrology.

Parent materials for Histosols are mostly hydrophytic plants [1]. Sphagnum consists of both living and dead tissue from the genus Sphagnum, with approximately

380 species. Sphagnum leaf tissue consists of chlorophyllose and hyaline cells, with the former having photosynthetic activity and the latter consisting of larger, clear and non-living cells with a large capacity to hold and store water. The cell walls contain an abundance of phenolic compounds that are resistant to decomposition. Sphagnum also has a substantial uptake capacity for calcium, magnesium and other nutrients, predisposing the underlying mineral soil to an acidic reaction. Typically, Sphagnum is the dominant plant genus in mires, raised bogs and blanket bogs. Other plant species commonly associated with Sphagnum include sedges, various dwarf shrubs, *Betula nama* (Dwarf birch) and *Salix* spp. (Willows).

Paludification or the geologic accumulation of organic materials across a landscape is influenced by soil pH, soil temperature, microbial activity, nutrient availability, oxidation–reduction and vertebrates (example: beavers or *Castor canadensis*). One criterion for paludization is the maintenance of anaerobic soil conditions sufficient to inhibit plant material decomposition. In glacial lake settings or ox-bows in fluvial systems, sediment infusion may occur resulting in lacustrine sediment accumulation. When sediment accumulation is sufficient to permit acceptable light levels to penetrate to the submerged sediment surface and if the water oxygen levels are appropriately anaerobic then plant material preservation prevails. When Histosols evolve because of sediment deposition with subsequent soil organic matter accumulation then this process is termed terrestrialization.

### **3. Gelisols**

In the United States the Keys of Soil Taxonomy support 12 soil orders at the highest level of soil taxonomy [1]. Gelisols (Cryosols in the World Reference Base of Soil Resources [2]) are soils that have permafrost within two meters from the soil surface. Permafrost is a soil climatic condition where soil material has continuous temperatures at or below 0°C. Because of the permafrost requirement, Gelisols occur extensively in boreal, subarctic and arctic environments and comprise approximately 18 km<sup>2</sup> (13.4%) of the ice-free land area [1]. Gelisols having a short period of seasonal thawing have an upper zone that thaws, creating an "active layer" approximately a few cm to 1.5 m thick. This active layer may experience soil forming processes, including sufficient biotic activity to form histic epipedons (suborder histels) [3].

The boundary between the active layer and permafrost is termed the "permafrost table". In moist soil and with the return of winter conditions, soil freezing begins at the permafrost table and also at the soil surface, which subsequently finalizes in the active layer. Thus, the active layer experiences freezing fronts from both the soil surface and from the permafrost table, giving rise to compaction and a loss of any soil structure. In the active layer of many Gelisols, dark streaks of organic matter that are distinguished from the soil matrix colors, suggesting soil material redistribution because of cryoturbation. The permafrost table is frequently impermeable to percolating water and therefore develops an accumulation of soil organic matter.

In very cold and low precipitation areas Gelisols are mostly shallow and relatively featureless soils; however, where temperatures are relatively mild and precipitation is more extensive, Gelisols are deeper and likely have an active layer that exhibits accumulation of soil organic matter. Gelisol vegetation includes lichens, moss, liverwort, sedge, grass and boreal forest species. Soil inhabiting organisms include prokaryotes (most notably N-fixing Azotobacter), fungi, actinomycetes, anthropoids, nematodes, protozoa and algae [1, 3].

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Solifluction may occur on sloping landscapes. Cryopedogenic processes

(Haploidization), soil structure formation, seasonal ice lens formation above the permafrost table, landscape collapse (thermokarst), and the formation of redoximorphic features. Additionally, soil carbon pool sizes, redistribution within the soil profile, and bioavailability are strongly affected by (1) cryoturbation, which is the soil-mixing action of freeze/thaw processes, and (2) by the presence of permafrost itself, which has strong controls over soil temperature and moisture and runoff. Overall, permafrost affected soils represent 16% of all soils on the globe, and contain up to 50% of the global belowground soil carbon pool [4]. Histels are Gelisols consisting of organic materials, with suborder groups listed as: (i) Folistels, (ii) Glacistels [have the upper boundary of a glacic layer (75% or

include cryoturbation causing a reduction in soil profile horizonation

more visible ice)], (iii) Fibristels, (iv) Hemistels, and (v) Sapristels.

northern permafrost region, noting that approximately 3.56 x 10<sup>6</sup>

examples do exist in grassland and deciduous forest biomes.

deeper soil depths because of cryoturbation.

**4. Organic carbon and peatlands**

Tarnocai et al. [4] performed an extensive review of carbon pools in the

this region at peatlands. These authors provided data illustrating that Histels (66.6 kg m−2) and Histosols (69.6 kg m−2) have the highest soil organic carbon contents. Histels alone are estimated to contain 184 Pg C, whereas histosols contribute 94.3 Pg C. Turbels show extensive soil organic carbon incorporation to

Peatland ecosystems are well represented in the majority of the world's biomes. In this manuscript we define a biome as a community of associated ecosystems characterized by their prevailing vegetation and by organism adaptation to that particular environment. Different sources define the types and number of biomes differently; herein, we specify six biomes: (i) tundra, (ii) taiga, (iii) grassland, (iv) deciduous forest, (v) desert, and (vi) tropical rainforest. Tundra, taiga and tropical rainforests are commonly accepted biomes having considerable expanses of peatlands; however,

Peatlands, as defined by the National Working Group (Canada), are wetlands containing more than 0.4 m thickness of peat [5]. Ombrotrophic peatlands or oligotrophic peatlands include soil and vegetation which receive water and nutrients primarily from precipitation, thus they are environments isolated hydrologically from the surrounding landscape. Given that rainfall is acidic because of equilibrium with the partial pressure of CO2 and the rainfall nutrient composition is relatively low, ombrotrophic peatlands are typically considered nutrient deficient and exhibit reduced microbial activity. Frequently the vegetation is dominated by Sphagnum mosses. Minerotrophic peatlands are wetlands whose water availability comes mainly from nutrient-enriched surface waters that have neutral to alkaline pH reactions. Typically, minerotrophic wetlands have a high-water table, low internal drainage and exhibit moderately-well to well-decomposed sedges, brown mosses and related vegetation. Carbon content is variably defined to represent the carbon concentrations on a surface area basis or a soil volume basis. Typically, carbon content defined as the mass of carbon per unit land area (kg carbon m−2) is presented to indicate landscape variability, whereas carbon content on a volume basis (kg carbon m−3) is presented to indicate intra-pedon or inter-pedon differences. Carbon content as expressed as the carbon concentration per volume is a soil or landscape property influenced by bulk density and horizon depth. Carbon accumulation is the net gain or loss of carbon content, typically at century or millennial scales. Peatlands reside on nearly

 km<sup>2</sup> in

*DOI: http://dx.doi.org/10.5772/intechopen.94399*

*Soil Genesis of Histosols and Gelisols with a Emphasis on Soil Processes Supporting Carbon… DOI: http://dx.doi.org/10.5772/intechopen.94399*

Solifluction may occur on sloping landscapes. Cryopedogenic processes include cryoturbation causing a reduction in soil profile horizonation (Haploidization), soil structure formation, seasonal ice lens formation above the permafrost table, landscape collapse (thermokarst), and the formation of redoximorphic features. Additionally, soil carbon pool sizes, redistribution within the soil profile, and bioavailability are strongly affected by (1) cryoturbation, which is the soil-mixing action of freeze/thaw processes, and (2) by the presence of permafrost itself, which has strong controls over soil temperature and moisture and runoff. Overall, permafrost affected soils represent 16% of all soils on the globe, and contain up to 50% of the global belowground soil carbon pool [4]. Histels are Gelisols consisting of organic materials, with suborder groups listed as: (i) Folistels, (ii) Glacistels [have the upper boundary of a glacic layer (75% or more visible ice)], (iii) Fibristels, (iv) Hemistels, and (v) Sapristels.

Tarnocai et al. [4] performed an extensive review of carbon pools in the northern permafrost region, noting that approximately 3.56 x 10<sup>6</sup> km<sup>2</sup> in this region at peatlands. These authors provided data illustrating that Histels (66.6 kg m−2) and Histosols (69.6 kg m−2) have the highest soil organic carbon contents. Histels alone are estimated to contain 184 Pg C, whereas histosols contribute 94.3 Pg C. Turbels show extensive soil organic carbon incorporation to deeper soil depths because of cryoturbation.

#### **4. Organic carbon and peatlands**

Peatland ecosystems are well represented in the majority of the world's biomes. In this manuscript we define a biome as a community of associated ecosystems characterized by their prevailing vegetation and by organism adaptation to that particular environment. Different sources define the types and number of biomes differently; herein, we specify six biomes: (i) tundra, (ii) taiga, (iii) grassland, (iv) deciduous forest, (v) desert, and (vi) tropical rainforest. Tundra, taiga and tropical rainforests are commonly accepted biomes having considerable expanses of peatlands; however, examples do exist in grassland and deciduous forest biomes.

Peatlands, as defined by the National Working Group (Canada), are wetlands containing more than 0.4 m thickness of peat [5]. Ombrotrophic peatlands or oligotrophic peatlands include soil and vegetation which receive water and nutrients primarily from precipitation, thus they are environments isolated hydrologically from the surrounding landscape. Given that rainfall is acidic because of equilibrium with the partial pressure of CO2 and the rainfall nutrient composition is relatively low, ombrotrophic peatlands are typically considered nutrient deficient and exhibit reduced microbial activity. Frequently the vegetation is dominated by Sphagnum mosses. Minerotrophic peatlands are wetlands whose water availability comes mainly from nutrient-enriched surface waters that have neutral to alkaline pH reactions. Typically, minerotrophic wetlands have a high-water table, low internal drainage and exhibit moderately-well to well-decomposed sedges, brown mosses and related vegetation.

Carbon content is variably defined to represent the carbon concentrations on a surface area basis or a soil volume basis. Typically, carbon content defined as the mass of carbon per unit land area (kg carbon m−2) is presented to indicate landscape variability, whereas carbon content on a volume basis (kg carbon m−3) is presented to indicate intra-pedon or inter-pedon differences. Carbon content as expressed as the carbon concentration per volume is a soil or landscape property influenced by bulk density and horizon depth. Carbon accumulation is the net gain or loss of carbon content, typically at century or millennial scales. Peatlands reside on nearly

2.7% of the global land surface, yet peatlands possess a significant portion of the terrestrial soil carbon pool with deep soil organic matter accumulations created over millennia. Estimates suggest that boreal and subarctic peatlands contain 455 Pg C [6] and 462 Pg [7], repectively. Boreal peat deposits tend to be deeper than subarctic peatlands, a feature attributed to long carbon accumulation intervals [8].

Peat-forming systems have been partitioned into acrotelm and catotelm zones [9]. The acrotelm portion of a peat-forming soil system is defined as the relatively more oxygenated (oxic) upper portion of the peat forming soil system, where aerobic decomposition is comparatively greater, the hydraulic conductivity is more rapid and the bulk density typically ranges from 0.1 to 0.4 g cm−3. Conversely the catotelm is the suboxic to anoxic lower portion of the peat-forming soil system that is characterized by a comparatively slower hydraulic conductivity and a bulk density typically ranging from 0.8 to 1.2 g cm−3.

Soils being open thermodynamic systems receive water and particulate soil organic matter and energy at their boundaries, most notably at the soil-atmosphere interface. Matter and energy may also be transferred by lateral flow at the pedonpedon interface or vertical flow at the soil-sediment interface. Water infiltration and percolation within the acrotelm is rapid; however, percolation slows substantially in the catotelm, creating the upper oxic and deeper anoxic oxidation–reduction regimes within the soil profile. As soil organic matter decomposition progresses at the base of the acrotelm, the resulting loss of pore space, attributed to an increase in the bulk density, supports water retention and conversion of the lowermost portion of the acrotelm into that of the catotelm, thus elevating the acrotelm-catotelm boundary with progressive soil development.

The primary vegetation productivity (P [=] g cm−2) is the annual production of particulate organic matter and its subsequent incorporation in the soil's surface horizons. The transformation of particulate matter to humus is predicated on soil temperature, microbial acidity, the soil's oxidation–reduction status, pH and nutrient availability. The rate of organic matter accumulation per unit surface area (x) is the difference between the annual production of particulate organic matter per unit area and the rate of soil organic matter loss per unit area, expressed as a first-order linear ordinary differential equation:

$$\mathbf{dx} \,/\,\mathrm{d}\mathbf{t} = \mathrm{P} - \alpha \mathbf{x},\tag{2}$$

where α is the decay coefficient, and t is time (years). Integration using an integration factor provides a solution:

$$\mathbf{x} = \left(\mathbf{P}/\alpha\right)\left(\mathbf{1} - \mathbf{e}^{-\mathrm{at}}\right). \tag{3}$$

**187**

occurring after 1950.

**Figure 1.**

*respectively.*

**molecular weight carbon species**

*Soil Genesis of Histosols and Gelisols with a Emphasis on Soil Processes Supporting Carbon…*

carbon stocks because of stimulated plant production. In Poland, Sienkiewicz et al. [11] investigated Histosol soil organic carbon and its relationship to total nitrogen, dissolved organic carbon and dissolved organic nitrogen. Carbon and nitrogen loss rates were independent, and soil organic carbon losses were dependent on the soil organic carbon content. The ratio of dissolved organic carbon to soil organic carbon increased with respect to the intensity of soil organic matter decomposition. Turunen et al. [12] investigated wet deposition of nitrogen (0.3 to 0.8 g nitrogen m−2 yr.−1) in ombrotrophic peatlands in eastern Canada, noting that nitrogen

*Illustration of mass accumulation per year (0 to 3500 g m−2 yr.−1) versus time (40 years) using Eq. (2) . The primary vegetation productivity was 150 and 450 g m−2 yr.−1 and the decay coefficients were 0.05 and 0.15 year−1,* 

Qui et al. [13] modeled northern peatland areas and carbon changing aspects during the Holocene. They recognized that the net primary production (NPP) and heterotrophic respiration increased over the past century in response to climate change and increased atmospheric CO2 activity. In their study net primary productivity was a greater influence than heterotrophic respiration, with 11.1 Pg C accumulated carbon storage since 1901, with the majority of the carbon storage increase

**5. Research studies focusing on soil chemistry with emphasis on low** 

The literature is replete with compelling research documenting biologically mediated geochemical pathways that are instrumental in creating vibrant biomes that have substantial accumulations of soil organic matter. Microbial populations secrete extracellular enzymes that are specific for degrading organic functional groups. The effectiveness of these extracellular enzymes is a complex function of (i) peat chemistry and litter quality, (ii) nutrient status, (iii) moisture content, (iv) plant community composition, (v) microbial community representation, and (vi) temperatures [14]. The absence of oxygen may also result in the accumulation of phenolic compounds that impost a negative feedback on microbial activity. Key enzyme activities

additions supported a greater diversity of vascular plants.

*DOI: http://dx.doi.org/10.5772/intechopen.94399*

From Clymo [9] typical decay constant values include α = 0.05 and 0.15 year−1. Also, from Clymo [9] typical annual production of particulate organic matter values includes: 150 and 450 g m−2 yr.−1. Using Eq. 3, The mass accumulation is presented for two scenarios: (i) P = 450 g m−2 yr.−1 and α = 0.15 year−1 (upper line in **Figure 1**) and (ii) P = 150 g m−2 yr.−1 and α = 0.05 year−1 (lower line in **Figure 1**). The scenario (i) P = 450 g m−2 yr.−1 and α = 0.15 year−1 provides a greater annual production of particulate organic matter and a faster rate of decay, such that the ratio P/α is a limit point as t approaches infinity. The asymptotic approach to P/α as a limit point implies that the net annual accumulation of organic matter ultimately becomes constant.

Street et al. [10] in Svalbard considered the influence of phosphorus (P) on the decomposition potential of carbon stocks. Nitrogen additions supported carbon stock reductions because of enhanced soil organic matter decomposition; however, the combination of added nitrogen and phosphorus supported an increase in the

*Soil Genesis of Histosols and Gelisols with a Emphasis on Soil Processes Supporting Carbon… DOI: http://dx.doi.org/10.5772/intechopen.94399*

#### **Figure 1.**

*Illustration of mass accumulation per year (0 to 3500 g m−2 yr.−1) versus time (40 years) using Eq. (2) . The primary vegetation productivity was 150 and 450 g m−2 yr.−1 and the decay coefficients were 0.05 and 0.15 year−1, respectively.*

carbon stocks because of stimulated plant production. In Poland, Sienkiewicz et al. [11] investigated Histosol soil organic carbon and its relationship to total nitrogen, dissolved organic carbon and dissolved organic nitrogen. Carbon and nitrogen loss rates were independent, and soil organic carbon losses were dependent on the soil organic carbon content. The ratio of dissolved organic carbon to soil organic carbon increased with respect to the intensity of soil organic matter decomposition. Turunen et al. [12] investigated wet deposition of nitrogen (0.3 to 0.8 g nitrogen m−2 yr.−1) in ombrotrophic peatlands in eastern Canada, noting that nitrogen additions supported a greater diversity of vascular plants.

Qui et al. [13] modeled northern peatland areas and carbon changing aspects during the Holocene. They recognized that the net primary production (NPP) and heterotrophic respiration increased over the past century in response to climate change and increased atmospheric CO2 activity. In their study net primary productivity was a greater influence than heterotrophic respiration, with 11.1 Pg C accumulated carbon storage since 1901, with the majority of the carbon storage increase occurring after 1950.

## **5. Research studies focusing on soil chemistry with emphasis on low molecular weight carbon species**

The literature is replete with compelling research documenting biologically mediated geochemical pathways that are instrumental in creating vibrant biomes that have substantial accumulations of soil organic matter. Microbial populations secrete extracellular enzymes that are specific for degrading organic functional groups. The effectiveness of these extracellular enzymes is a complex function of (i) peat chemistry and litter quality, (ii) nutrient status, (iii) moisture content, (iv) plant community composition, (v) microbial community representation, and (vi) temperatures [14]. The absence of oxygen may also result in the accumulation of phenolic compounds that impost a negative feedback on microbial activity. Key enzyme activities

important to mineralization include: (i) alpha-glucosidase, (ii) beta-glucosidase, (iii) cellobiohydrolase, (iv) N-acetylglucosaminidase, (v) acid phosphatase, and (vi) leucine aminopeptidase.

Fox [15] reviewed literature involving low-molecular-weight organic acids. Low-molecular weight organic acids are approximately 10% of a typical forest soil's dissolved organic carbon pool, but they may have a disproportionate influence on soil processes, including metal complexation. Common low molecular weight organic acids include: acetic, aconitic, benzoic, cinnamic, citric, formic, fumaric, gallic, lactic, malic, maleic, malonic, p-hydroxybenzoic, phthalic, protocatechuic, oxalic, salicylic, succinic, tartaric, and vanillic. Common functional groups include (i) acidic groups [carboxylic (R-COOH), enolic (R-CH=CH-OH), phenolic (Ar-OH) and quinones (Ar = O)], (ii) neutral groups [alcoholic OH (R-CH2OH), ethers (R-CH2-O-CH2-R), ketones (R-C=O (−R)), aldehydes (R-C=O(-H)) and esters (R-C=O(-OR))] and (iii) neutral nitrogen-bearing amines (R-CH2-NH2) and amides (R-C=O(NH-R)). When considering root extracts oxalic, citric and malic are quite abundant. Sources of low molecular weight organic acids are root respiration, leaching from the litter floor, decomposition of soil organic matter, and rainfall. Herbert and Bertsch [16] further detailed dissolved and colloidal organic matter in the soil solution. Based on their review of literature dissolved organic matter is primarily composed of hydrocarbons, chlorophyll, carotenoids, phospholipids and long-chain fatty acids, tannins, flavonoids and other polyphenols, fulvic and humic acids, aromatic and aliphatic acids, and proteins /amino acids. In most studies the dominant organic materials were humic substances.

Kane et al. [17] measured pore water chemistry associated with an artificiallyinduced warming of a nutrient poor fen. The dissolved organic carbon (DOC) concentration was greater in the warmed fen (73.4 ± 3.2 mg L−1) compared to the untreated check (63.7 ± 2.1 mg L−1). The amount of dissolved organic nitrogen (DON) was greater in the warmed fen; however, the DON/DOC ratio was smaller. The reduced DON/DOC ratio was primarily attributed to a smaller capacity of the microbial community to yield labile nitrogen via the decomposition process and the greater utilization efficiency of the nitrogen by the microbial community. In Manitoba (Canada) Aide and Cwick [18] studied Eluviated Eutric Brunisols having an Of-Bm-C horizon sequence and Orthic Eutric Brunisols having an Oh or Of-Bm-C horizon sequence. Located in the glacial Lake Agassiz these soils formed in fine-graine lacustrine sediments interspersed with organic soils and fens. The surface horizons of the Eluviated Eutric Brunisols possessed organic carbon contents ranging from 19.8 to 29.4% with C/N ratios of 29.5 to 27.4, whereas the surface horizons of the Orthic Eutric Brunisols possessed organic carbon contents ranging from 27.3 to 41.7% with C/N ratios of 39.5 to 25.4. The C/N ratios and associated nitrate-N concentrations suggests that nitrogen limits the rates of soil mineralization. In a near companion manuscript Aide et al. [19] documented that the silty sediments were dominated by hydroxy Al-interlayered vermiculite, smectite, hydrous mica, and kaolinite in the clay separate. The potential for potassium fixation by vermiculite was reduced by Al-interlayering.

Van Cleve and Powers [20] isolated state factors involved in carbon storage in forest soils, noting the role of climate, parent material, topography, vegetation, and soil organisms. The chemistry of soil organic carbon, including root exudates and leachates, strongly influence the microbial processing of detritus, the materials synthesized in this process and the intensity of the roles that low and high molecular weight organic acids have in soil development. Observed effects show that synthesized products are more resistant to further decomposition and possessed smaller nitrogen contents, which over time supports soil organic matter accumulation.

**189**

*Soil Genesis of Histosols and Gelisols with a Emphasis on Soil Processes Supporting Carbon…*

**6. Research studies having a focus on carbon loss as greenhouse gas** 

Peatlands are an important terrestrial carbon sink and any increased microbial activity may result in soil organic matter oxidation, with subsequent CO2 release. Northern peatlands historically have had the benefit of cool to frigid temperatures that limit microbial activity. Low oxygen activity attributed to water saturation further limits mineralization. Climate change may result in warmer soils, with the cavate that the effective length of the increasingly warmer summer interval is also increased. The encroachment of vascular plants will be expected to proceed, leading to a positive feedback on microbial activity. Thus, studies on peatland functioning in higher latitudes and their potential to accelerate climate change are becoming

In Canada, Dieleman et al. [21] established mesocosms, where peat production of dissolved organic carbon was measured. The production of dissolved organic carbon from peat was estimated to be a function of temperature, CO2 concentration and the influence of the water table, wherein increased temperatures increased the dissolved organic carbon contents, lowered water tables increased decomposition rates and reduced pore water dissolved organic carbon concentrations. In the Alaskan arctic Euskirchen et al. [22] established eddy covariance flux towers across various ecosystems for three years to document peak CO2 uptake patterns. Peak CO2 uptake centered from June to August at a mean of 51 to 95 g C m−2 across the various ecosystems. Warmer spring seasons promoted greater CO2 uptake patterns, whereas warmer late seasons supported greater soil respiration rates, reducing the

In Canada, Frolking et al. [23] employed the Holocene Peat Model to simulate the vegetation community composition and the annual net primary productivity. Northern peatlands take up CO2 at rates of 40 to 80 g carbon m−2 yr.−1, with carbon leaching as DOC at rates of 10–20 g DOC m−2 yr.−1. Decomposition was estimated to be 95% of the Net Primary Productivity. Similarly, Frolking et al. [23] observed undisturbed Canadian peatlands and determined that these peatlands were a weak sink for carbon and a moderate source of methane emission. McLoughlin and Webster [24] performed a review of peatland dynamics, primarily within the Hudson Bay Lowlands. Long term carbon accumulation, CO2 sequestration, peat depth and land age were positively correlated. Carbon dioxide sequestration showed the greatest variability, with bogs (−1.7 to 1.5 g carbon m−2 day−1), fens (−4.3 to 1.6 g carbon m−2 day−1), and palsa peat (−0.8 to 1 g carbon m−2 day−1). Methane and evapotranspiration were greater in the wettest ecosystems, with methane emission for bogs (3.3 to 28 mg carbon m−2 day−1), fens (0.1 to 204 mg carbon m−2 day−1), and

On paludified soils Schneider et al. [25] measured methane (CH4) flux for forest and peatland areas. Open peatlands exhibited a methane emission rate of 21.9 ± 1.6 g m−2 yr.−1 in contrast with forested peatland transition zones (7.9 ± 0.5 g m−2 yr.−1). The forested peatland transition zones demonstrated an inflow of less acidic surface water that supported a higher biological diversity and greater plant productivity. These authors noted that methane emission was more influenced by increased temperatures than the water table depths. In Sweden, Sagerfors et al. [26] established eddy covariance measurements across oligotrophic mires. Based on the vertical exchange of CO2 their sites were a net sink for carbon (55 ± 7 g carbon m−2 yr.−1). The non-growing seasons exhibited a carbon loss; however, the growing season sequestration of carbon more than compensated for

*DOI: http://dx.doi.org/10.5772/intechopen.94399*

**emissions**

commonplace [14].

Net Ecosystem Exchange (NEE).

palsa peat (−1.6 to 24 mg carbon m−2 day−1).

the non-growing season carbon loss.

*Soil Genesis of Histosols and Gelisols with a Emphasis on Soil Processes Supporting Carbon… DOI: http://dx.doi.org/10.5772/intechopen.94399*
