use of stable isotopes to identify sources and ...

Sci.Int.(Lahore),28(3),2719-2724,2016

ISSN 1013-5316;CODEN: SINTE 8

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USE OF STABLE ISOTOPES TO IDENTIFY SOURCES AND FORMATION PROCESSES OF GREENHOUSE GASES AT WETLAND – ATMOSPHERE INTERFACE: A REVIEW Tariq M. Munir1,2,*, Ghulam Murtaza Jamro3, 4, Ashfaq A. Rahi5, Hameed Ullah6, Iftikhar Ahmad6, Zia-ul-hassan4 1

Department of Geography, University of Calgary, 2500 University Dr. NW Calgary, Alberta T2N 1N4, Canada Department of Geology, St. Mary’s University, 14500 Bannister Road SE, Calgary, Alberta T2X 1Z4, Canada 3 Department of Renewable Resources, University of Alberta, Edmonton, Alberta T6G 2E3, Canada 4 Department of Soil Science, Sindh Agriculture University, Tandojam, Sindh, Pakistan 5 Soil & Water Testing Laboratory for Research, Agriculture Farm, Old Shujabad Road, Multan 60500, Punjab, Pakistan 6 Mango Research Institute, Agriculture Farm, Old Shujabad Road, Multan 60500, Punjab, Pakistan *Corresponding author email: [email protected] 2

ABSTRACT: Northern wetlands contain a large variety of organic substrates of carbon dioxide (CO2) and methane (CH4) greenhouse gases (GHG). Different substrates exhibit different decomposition and respiration rates which will emit GHGs at different rates to the atmosphere. Therefore, without the knowledge of the exact source of carbon (C) and hydrogen (H) reactants for the CO2 and CH4 formation reaction, it is difficult to predict how the rate of formation processes of these GHGs will respond to climate change. Therefore, our objectives of this review were to (i) concisely review various sources and formation processes of peatland GHG emissions, and (ii) describe the potentials and limitations of using isotopic compositions of these GHGs for identification of their sources and formation processes. The C and oxygen isotopic ratios can be used to estimate the proportions of photosynthesis and respiration components of ecosystem CO2 exchange, whereas C and H isotopic compositions of CH4 can be used to determine the relative contribution of the different formation pathways. This will improve our understanding on how the potential changes in climate and environmental factors will change the proportions of sources and formation processes of these GHGs in wetland ecosystems. Therefore, there is a tremendous potential for use of stable isotope compositions of GHGs for identifying sources and partitioning their contributing processes provided the limitations in our knowledge and instrument requirements are addressed. Keywords: carbon dioxide, climate change, environment, methane, stable isotope, wetland, 1. INTRODUCTION Significantly more carbon (C) is stored in the world soils including wetlands (peatlands) and permafrost, than is present in the atmosphere [1]. Northern peatlands store an estimated 30% (~547 Pg C (1 Pg=1015 g) of the global soil C stocks [2], which is equivalent to approximately 75% of the preindustrial mass of C stored in the atmosphere [3]. If the large reserves of soil C are transmitted to the atmosphere by a climatic-warming-induced accelerated decomposition of organic matter, a positive feedback (loss of carbon dioxide (CO2) and methane (CH4) from peatland) to climate change would occur due to temperature sensitivity of the peat soil organic matter [4]. The releasing major greenhouse gases (GHGs; CO2 and CH4) have been recently estimated to in turn warm the atmosphere linearly by the end of this century [5]. Rising levels of these GHGs will enhance warming leading to larger changes in the global climate system [1] and causing unstoppable changes in the climate for at least another 1000 years in the best case scenarios [6]. The formation rates of the GHGs are driven by environmental controls [7], for e.g., climatic [8], hydrological [9] and nutrient supply [10]; however, investigations to untangle their influence on sources and formation processes of the GHGs, are limited. Atmospheric CO2 is the leading source of C for plants. Plants uptake and convert atmospheric CO2 to organic compounds, transfer solar energy in photosynthetically active radiation (PAR) into energy stored in chemical bonds, and in the meantime release oxygen (O2) as a by-product into the atmosphere. The synthesized organic compounds (e.g., glucose) are assimilated for growth, storage and plant respiration, while a smaller fraction is leached as root exudates and consumed by microorganisms [11]. Death and decay of the plants results in addition of organic matter to the

soil in the form of plant litter which is decomposed and mineralized by microbes through the process of respiration. The sum of autotrophic respiration (plant respiration) and heterotrophic respiration (microbial mineralization) is ecosystem respiration (ER) which is the original source of CO2 to atmosphere. 1.1 Environmental Controls on CO2 and CH4 formations A number of major environmental controls, e.g., atmospheric temperature, precipitation, peatland water-table (WT) level and moisture, availability of soil nitrogen (N) and phosphorus (P) may simultaneously be the controls on the sources and formation processes of CO2 and CH4 [4]. They may also play roles in obscuring the intrinsic sensitivity of the sources and formation processes of these GHGs [4]. The CH4 formation in soils occurs under anaerobic (saturated) condition with availability of suitable substrate, or under aerobic (oxic) environment as a result of biochemical reduction of CO2. The estimation protocols of ER and CH4 emission quantify only the bulk quantities of these GHG emissions from peatland but lack information of the contributing source (substrate) of C and H used by the formation processes. Under changing environmental controls, the substrate may change after vegetation succession [12]. Different substrates exhibit different decomposition and respiration rates which will emit GHGs at different rates to the atmosphere. Therefore, without the knowledge of the exact sources of C and H reactants for the CO2 and CH4 formation reaction, it is difficult to predict how the rate of formation processes of these GHGs will respond to the changing environmental controls as reported by IPCC [1]. Therefore, the objectives of this paper are; 1) to briefly review various sources and formation processes of peatland GHG (CO2 and CH4) emissions, and 2) to describe the

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potentials and limitations of using isotopic compositions of these GHGs for identification of their sources and formation processes. The atmospheric warming and WT lowering are the major drivers of CO2 release to the atmosphere [e.g., 12]. The CO2 emissions significantly increased with increased warming and lowered WT in a boreal treed bog [12]. Increasing soil temperature with elevating WT level significantly increased CH4 production or emission [13, 14], in contrast to increased production or emission with lowering of WT level reported by Strack, Waller [15]. Therefore, CO2 and CH4 production and emission sources and formation processes needs to be investigated. The use of stable isotopic compositions of the GHGs can take up this challenge by isolating or distinguishing a source or formation process from the others. The stable C, H and O2 isotopes have been used to separate individual sources or processes contributing to the balance of GHGs in soil, plant and atmosphere [e.g., 9]. 1.2 Stable Isotopes of CO2 and CH4 Isotopes are atoms of an element with different number of neutrons. The mass differences, due to a difference in the number of neutrons, will result in partial separation of the light isotopes from the heavy isotopes during chemical reactions such as methanogenesis and methanotrophy [16]. This process is called isotope fractionation where one isotope is enriched relative to another in a chemical or physical process occurring in soil, plant or atmosphere. For example, the difference in mass between the two stable isotopes of hydrogen (H), 1H (1 proton, no neutron, also known as protium) and 2H (1 proton, 1 neutron, also known as deuterium) is that the later is almost 100% heavier [16]. Therefore, there will be a significant fractionation between the two isotopes. Heavier isotopes are stable because they form stronger chemical bonds and have lower reaction rates in enzymatic and microbial reactions [17]. The difference in reaction rates results in variation in their abundance in chemical processes. Isotopic composition is mostly expressed in terms of δ values (‰) as the ratio of heavy and light isotope in a compound in question relative to that ratio in a standard: δ= (Rsample/Rstandard – 1) x 103 (1) 13 2 where, δ is C or H, and R is the corresponding ratios of 13 12 C/ C or 2H/1H. For isotopic investigations of organic soil C sources or processes, the elements of interest are C, O and H, where the isotopic standard for C is PeeDee Belemnite (PDB; limestone) and for both oxygen and hydrogen the standard is standard mean ocean water (SMOW). Only C and H are being discussed. Changes in isotopic composition resulting from an enzymatic process can be expressed as a discrimination (Δ) or fractionation (α) factor: Δ=Rsource/Rproduct – 1 = (δsource – δproduct)/(1 + δproduct/1000) (2) α = Rsource/Rproduct = (δsource + 1000)/(δproduct + 1000) (3) 2. LITERATURE REVIEW (CO2) 2.1 Use of Stable Isotope Composition of CO2: Potentials and Limitations Most of the C cycling studies measuring net ecosystem exchange (NEE) as the balance of gross ecosystem

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photosynthesis (GEP) and ER do little separation of these values into components which combine to create the C flux values. Also, empirical derivation of the value of GEP from NEE and ER creates errors [18]. Even if this NEE is resolved into its components of GEP and ER, the individual contribution of vegetation type (C3 or C4) or community to the system’s productivity, and the partial contributions of different types of substrates (e.g., leaf, root or old recalcitrant organic matter) to the ER is lacking. Different vegetation types and substrates have different rates of production and emission; therefore, without an understanding of the magnitude of the smaller scale reactions that contribute to net C fluxes, it is difficult to predict how wetlands will respond to change in environmental controls. Stable isotopes of C and O2 can be used to separate individual reactions contributing to the CO2 balance of wetlands. The stable isotopes can record two types of information: 1) where chemical and physical reactions fractionate stable isotopes, the resulting isotopic ratios reflect process information, and 2) they also record information about origins of reactants (source partitioning) of the process reaction. Due to interdependency between sources, their isotopic ratios and GHG formation processes, the potentials and limitations of using them are being discussed together. 2.2 Net Ecosystem Exchange of CO2: A Source and Formation Processes Photosynthesis and respiration rates are affected differently by environmental factors. Therefore, to gain a clear understanding of ecosystem response to environmental controls, it is useful to split NEE into its components: photosynthesis and respiration. These components influence the C and O2 stable isotopes differently; therefore measuring the isotopic composition of ecosystem CO2 provides an opportunity to differentiate these processes [19] and check which component is more sensitive to a perturbation. A twomember mixing model can be used in which the net flux is the balance of photosynthesis (P) and respiration (R), each with a unique isotopic signature such that: N (net flux) = P + R (4) NδN= PδP + RδR (5) P = [NδN(N-P)δR]/δP (6) R= [NδN – (N-R)δP]/δR (7) where δN, δP and δR are the C or O2 isotopic ratios of CO2 for net flux, photosynthesis and respiration, respectively [20]. Both C and O isotopes in CO2 can be used to differentiate P and R and using both will provide an independent check of findings [19, 21]. For the remainder of this discussion, only C isotopes will be considered; however, similar techniques can be applied using O isotopes. The isotopic composition of atmospheric CO2 is altered by photosynthesis because plants discriminate against 13C. This alteration of isotopic composition occurs due to: 1) differential rates of diffusion of 12CO2 and 13CO2 through stomata, and 2) enzymatic discrimination by different plant species using different photosynthetic pathways (e.g., C3 and C4 plants) [22]. The relative importance of diffusional and enzymatic fractionation is related to the ratio of the concentration of CO2 in the atmosphere and the chloroplast,

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controlled by leaf physiology, photosynthetic capacity and stomatal conductance [19]. Also, the enzyme involved in the fixation of CO2 in the leaves of C3 plants (oxygenase, or Rubisco) has a much larger discrimination for 12CO2 than the enzyme used in C4 vegetation [phosphoenol pyruvate carboxylase; 22]. Therefore, the relative preference between C3 and C4 plants will alter the photosynthetic impact on atmospheric isotopic ratios. However, the limitation is that the isotopic discrimination will, in either case (C3 or C4) lead to enrichment of atmosphere in 13CO2. The discrimination (∆) resulting from photosynthesis can be estimated based on atmospheric CO2 concentrations and stomatal conductance/leaf CO2 concentration, which two concentrations can also be used to calculate ratio of atmospheric to intercellular CO2 concentrations [23]. Also, if the vegetation canopy is considered as a “big leaf”, the concept of isotopic discrimination across the leaf can be applied to the canopy with atmospheric and intercellular CO2 concentrations being replaced with concentration outside and inside the canopy [23]. Discrimination can then be calculated as: Δ = 1000 [Co/(Co – Ci)](δi – δo)/[1000 + δi –[(Co/(Co–Ci)](δi – δo) (8)

where, C is concentration, δ is isotopic ratio and, o and i represent air outside and inside the canopy, respectively [23]. An δ of instantaneous photosynthesis can be obtained by determining the isotopic fraction for recently produced biomass, but this ratio will vary in response to changes in environment [20]. This technique can be applied relatively easily when vegetation present has a homogeneous isotopic discrimination; however, this will not be the case when C4 plants are present or if different vegetation types vary in their discrimination [e.g., 24]. In these cases, δP must be estimated by weighting the isotopic effects of the different vegetation types based on their abundance or productivity, potentially introducing a large degree of uncertainty. Respiration does not appear to cause significant fractionation of C isotopes like photosynthesis [25]; however, since it results from the breakdown of 12C-enriched compounds resulting from photosynthesis, CO2 released by respiration leads to an ultimate dilution of 13CO2 in the atmosphere. Different classes of C molecules may vary in isotopic composition (for e.g., lignin is diluted of 13C relative to whole cellulose [26] and their differential rate of decomposition leads to spatial variability in δ13C values of respired CO2 [19]. Despite this variability, the isotopic ratio of CO2 respired by the ecosystem can be determined using a mixing model [27] such that: δe = Ca(δa – δs)(1/Ce) + δs (9) where, C is the concentration of CO2, δ is the isotopic ratio for CO2 and the subscripts a, e and s denote the atmosphere, ecosystem and ecosystem sources, respectively [20]. If these measurements are made at night when respiration is the only ecosystem source, δs will be equivalent to the isotopic ratio of ecosystem respiration (δR) [20, 28]. 2.3 Methods (CO2) Earlier, the C flux gradient approach determined NEE by using aerodynamic parameters and isotopic concentration of CO2 in the upwind where fast-responding instruments were not required [21]. In contrast, currently, eddy covariance

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requires rapid measurement of wind speed and CO 2 concentrations [29]. Griffis, Baker [30] used tunable diode laser absorption spectroscopy to measure the mixing ratios of 13 CO2 and 12CO2 for isotopic partitioning of CO2 fluxes. Because this method provides more rapid estimation of isotopic ratios and is less labour intensive than the old flask sampling technique, it holds good for future studies of ecosystem CO2 flux. Despite the advances made using these techniques, most studies have been carried out over limited sampling intervals and more research is required to determine whether the isotopic flux partitioning will be successful in long term monitoring studies. More recently, long term stable isotope labeling of newly fixed organic matter (as a result of atmospheric CO2 enrichment) studies has enabled further investigation into the contribution of soil components to ecosystem respiration [31]. Also, in order to study the effects of potential atmospheric CO2 fertilization (expected under changing climatic conditions) on ecosystem function, several studies are being conducted in areas where local CO2 concentrations have been experimentally enriched (free air CO2 enrichment or FACE). Since the CO2 used to create this enrichment is derived from fossil fuels it is depleted in 13C relative to the atmosphere. Therefore, carbon compounds created using this CO2 are depleted in 13C compared to those fixed before the enrichment. These isotopic differences can then be used to track the fate of this newly formed organic matter. A full review of these techniques is given in Pataki, D.S. Ellsworth [31]. Similar techniques could be applied if CO2 enrichment studies are extended to wetland ecosystems. 2.3.1 Global Surface – Atmosphere Carbon Accounting The use of stable isotope composition of CO2 is making substantial contributions to understand the controlling factors of formation and sources of CO2 balance in the atmosphere. The combustion of fossil fuel and oxidation of forest organic matter could be the possible sources of atmospheric CO2 balance. While the history of fossil fuel combustion has been tracked, a little attention is paid to the terrestrial burning and decomposition of organic matter. However, the amount of terrestrial CO2 being shifted to the atmosphere can be estimated through the use of C cycle model parametrized for including the magnitude of changes in 13C content of atmospheric CO2 since pre-industrial time [16]. Addition of large quantities of CO2 both from fossil fuel combustion (δ13C ~-27‰) and oxidation of terrestrial organic matter (δ13C ~-28‰) has lowered the δ13C values of atmospheric CO2 (δ13C ~-7‰) [32]. However, if the time since fossil fuel combustion started and the amount burnt are known, the magnitude of terrestrial biomass C released to atmosphere can be calculated by difference method. But this application would require C cycle model that can occlude the amount of atmospheric C exchanges with ocean and biosphere [33]. The analyses of δ13C from tree rings have inferred an average decrease of 1.5 ‰ in atmospheric 13C concentration since pre-industrial. These findings have been supported by the δ13C values quantified from Antarctic ice cores [34]. The analysis of recent atmospheric δ13C values corroborate the earlier findings. This example illustrates the potential of the isotopic tracer approach for solving mass balance challenges, especially

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where a well-mixed reservoir such as global atmosphere can be sampled over time. Alternate methods of determining the loss of biospheric C through analysis of land use patterns and deforestation records are uncertain due to the lack of accurate historical land use and deforestation records [35]. However, if these records are available, they can be considered along with atmospheric values of δ13C, to trend the global C exchanges over temporal scales [16]. If δ differences between two peatlands are very large versus within the peatlands, the use of isotopic compositions can be a powerful tool for peatland – atmosphere C exchange, and only few samples may provide effective solution of determining sources and processes of these exchanges. For example, it is easy to distinguish between C4 (δ13C ~-13‰) and C3 (δ13C ~-28‰) plant ecosystems using stable C isotopes. On the other hand, it would be very difficult to estimate nitrogen fixation using δ15N technique because soil organic matter δ is about +2‰ and atmospheric δ15N is 0‰. It would be far better to choose an ecosystem where the soil organic matter had δ15N values of +6 to +10‰ which contrast strongly to the atmospheric value. 3. LITERATURE REVIEW (CH4) 3.1 Use of Stable Isotope Composition of CH4: Potentials and Limitations The CH4 flux from a wetland is the result of the balance between CH4 production in the anoxic zone and CH4 oxidation in oxic layers. Once CH4 is produced it can be released to the atmosphere via diffusion, ebullition (bubbling) or plant-mediated transport [36], each of which exposes it to varying degree of oxidation. The CH4 emissions are measured as bulk fluxes without information of the contribution of various substrates to the formation processes. Stable isotopes of C and H can be used to separate the individual contributions. Analyses of both C and H isotopes in CH4 provide information about production pathways, the extent of oxidation and mechanisms of its release to the atmosphere. The δ of CH4 flux from wetlands and other major bodies to the atmosphere provides an important information in mass balance models of the global CH4 cycle [37]; however, a significant challenge remains in assigning of accurate isotopic values to the different production pathways and their sources. 3.2 CH4 Sources and, Production and Emission Processes Wetlands exhibit a wide range of δ13C and δ2H values between emission and the subsurface CH4 [e.g., 38]. Considerable efforts have been made to understand the systematic patterns in the distribution of δ13CCH4 and δ2HCH4 values of source organic matter, production pathways, processes of transport and emission [e.g., 39]. To describe the potentials and limitations of the use of δ of CH4 for identifying sources and formation processes, an understanding of the sources and formation processes is a prerequisite. 3.3 Pathways, Potentials and Limitations Biogenic formation of CH4 is carried out by bacteria called methanogens. Two primary pathways are possible: acetate fermentation and/or CO2 reduction with H2 [40, 41]. Variations in fractionation of C and H between these two production pathways reflect the differences between stable isotopic compositions of substrates and the methanogenic

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pathways. This allows using δ of CH4 for pathway determination. The isotopic fractions of C and H in CH 4 and CO2 have been used by Whiticar, Faber [42] for determining the fractionation factors for fresh water and marine water/sediment zones. They found that the acetate fermentation process and CO2 reduction process usually dominated in freshwater wetlands and marine waters, respectively. They determined that freshwater and marine sediments had mean δ13CCH4 of -59‰ and -68‰, respectively. The CH4 formed in freshwater through acetate fermentation receives three H from the methyl group of acetate and the rest one from the source water, whereas the CH4 produced in marine water (where O2 is limited) via CO2 reduction receives all four H from source water. Regardless of the cause, they determined large differences in hydrogen fractionation factors (αD) between the 2 pathways with values of 1.45 and 1.20 for acetate fermentation and CO 2reduction, respectively. The isotopic composition of precursor compounds affects δHCH4 values via influence of the δD in water [43]. The minerotrophic wetlands which have both higher net ecosystem productivity and higher rates of CH4 flux may contain pore water that has more positive δ13C and more negative δD values than ombrotrophic bogs [44]. As the CH4 production process shifts from acetoclastic to CO 2 reduction pathway, the value of δ13C tends to decrease while the value of δD tends to increase. Recent research limits the use of isotopic ratios of DCH4 to separate pathways of CH4 production in-spite of the large obvious variation in δD. It is because, the δDCH4 is preferentially controlled by the δD of source water than by the production pathway [43], highlighting the probability of H transfer from source water to acetate which causes increase in number of H atoms derived from water in CH4 produced by this mechanism. This approach is contradicted by Hornibrook, Longstaffe [45] who reported a large variation in δD between CH4 pools, although variation within the δD of the source water was very little. Chanton, Crill [39] suggested that changes in the stable isotope composition of CH4 precursors do not appear to influence δ13CCH4 values significantly in northern peatlands. While hydrogen isotopic fractionation during CH4 production deserves more investigation, several studies have used C and H isotopic ratios to determine the relative contribution of each production pathway in a variety of wetlands [e.g., 46, 47]. Evidence from a study of CH4 produced using other substrates, such as methanol and dimethyl sulphide, suggests that δ may also be useful for separating these pathways [48], although they are of minor importance in most wetlands. 4. CONCLUSIONS 1. The soil organic matter is a vitally important source as well as a sink of CO2 and source of CH4 production. The production and emission rates of these GHGs are affected by several environmental controls. The soil organic matter is a leading player in setting the atmospheric budgets of the GHGs; however, the exact sources of these greenhouse formations are not partitioned. 2. The large heterogeneity of soil organic matter within and between the peatland ecosystems requires exact source

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partitioning of components of emissions, as different sources and their components are affected by various environmental factors at different extents. This is possible by using stable isotope compositions of these GHGs during formation and at emission processes. 3. The C and O2 isotopic ratios can be used to estimate the proportions of photosynthesis and respiration components of ecosystem CO2 exchange, whereas C and H isotopic compositions of CH4 can be used to determine the relative contribution of the different formation pathways. This will improve our understanding on how the potential changes in climate and environmental factors will change the proportions of sources and formation processes of these GHGs in wetland ecosystems. Therefore, there is a tremendous potential for use of stable isotope compositions of GHGs for identifying sources and partitioning their contributing processes provided the limitations in our knowledge and instrument requirements are addressed. 5. REFERENCES [1]. Ipcc, Summary for policymakers., in Climate change 2007: The physical science basis. Contribution of Working Group I to the fourth assessment report of the Intergovernmental Panel on Climate Change, S. In: Solomon, Qin, D., Manning, M., Chen, Z., Marquis, M., Averyt, K.B. et al (eds.), Editor. 2007: Cambridge University Press, Cambridge, United Kingdom. [2]. Yu, Z., "Northern peatland carbon stocks and dynamics: A review". Biogeosciences. 9(10): 40714085(2012). [3]. Strack, M., Peatlands and Climate Change. 2008: IPS, International Peat Society. [4]. Davidson, E. A. and Janssens, I. A., "Temperature sensitivity of soil carbon decomposition and feedbacks to climate change". Nature. 440(7081): 165-173(2006). [5]. Peake, S. and Smith, J., Climate Change. Second ed. 2009: Oxford University Press. [6]. Gillett, P. G., Arora, V. K., Zickfeld, K., Marshall, S. J., and Merryfield, W. J., "Ongoing climate change following a complete cessation of CO2 emissions". Nature Geoscience. 4: 83-87(2011). [7]. Moore, P. D., "The future of cool temperate bogs". Environmental Conservation. 29(3-20): 2002). [8]. White, J. R., Shannon, R. D., Weltzin, J. F., Pastor, J., and Bridgham, S. D., "Effects of soil warming and drying on methane cycling in a northern peatland mesocosm study". J. Geophys. Res. 113(G3): G00A06(2008). [9]. Choi, W., Chang, S. X., and Bhatti, J. S., "Drainage affects tree growth and C and N dynamics in a minerotrophic peatland". Ecology. 88(2): 443453(2007). [10]. Bragazza, L. and Gerdol, R., "Are nutrient availability and acidity-alkalinity gradients related in Sphagnumdominated peatlands?". Journal of Vegetation Science. 13(4): 473-482(2002). [11]. Vasander, H. and Kettunen, A., Carbon in Boreal Peatlands, in Boreal Peatland Ecosystems, R.K. Wieder and D.H. Vitt, Editors. 2006, Springer Berlin Heidelberg. p. 165-194.

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