Universit degli Studi di Milano Bicocca

SCUOLA DI SCIENZE

Dipartimento di Scienze dellAmbiente e del Territorio e di Scienze della Terra

Corso di Laurea Magistrale in Scienze e Tecnologie per lAmbiente e il Territorio

The relationships between water

table and redox potential in

peatlands

______________________________________________________________________________

_________

Anno Accademico 2013 /2014

Relatore:

Prof. Roberto Comolli

Tesi di Laurea di:

Iacopo Federico Ferrario

Correlatore:

Prof. Ruurd Van Diggelen

Matricola:

770166

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TABLE OF CONTENTS

1 INTRODUCTION ................................................................................................................... 4

1.1 Peatlands ............................................................................................................................... 4

1.2 Ecohydrology ....................................................................................................................... 6

1.2.1 Height above water table .............................................................................................. 9

1.2.2 Micro-topography ....................................................................................................... 10

1.3 Carbon cycle ....................................................................................................................... 10

1.3.1 Decomposition ........................................................................................................... 11

1.3.2 Spatial process of decomposition ............................................................................... 12

1.4 Microbial community ......................................................................................................... 15

1.5 Vegetation season: drying and rewetting ........................................................................... 15

1.6 Redox potential .................................................................................................................. 17

1.7 Redox potential is a complex indicator .............................................................................. 22

1.7.1 Assembling the picture ............................................................................................... 33

1.8 Aim and research questions ................................................................................................ 35

1.8.1 Hypotheses ................................................................................................................. 35

2 MATERIALS AND METHODS .......................................................................................... 36

2.1 Field work .......................................................................................................................... 36

2.1.1 Site description ........................................................................................................... 36

2.1.2 Experimental design ................................................................................................... 37

2.1.3 Redox measurements .................................................................................................. 41

2.1.4 Water table measurements ......................................................................................... 42

2.1.5 Pore water sampling ................................................................................................... 43

2.1.6 Decomposition ........................................................................................................... 44

2.1.7 Peatsoil sampling ........................................................................................................ 45

2.2 Laboratory work ................................................................................................................. 45

2.2.1 Water chemistry analysis ........................................................................................... 45

2.2.2 Peat analysis ............................................................................................................... 46

2.2.3 Data calculation .......................................................................................................... 53

3 RESULTS AND DISCUSSION ........................................................................................... 54

3.1 Introduction ........................................................................................................................ 54

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3.2 Redox potential and water table ......................................................................................... 56

3.3 pH, Dissolved Inorganic Carbon and Dissolved Organic Carbon ..................................... 68

3.4 Cellulose decomposition and TBI ...................................................................................... 78

3.5 Peat quality and degradability ............................................................................................ 83

4 CONCLUSIONS ................................................................................................................... 93

5 Acknowledges ....................................................................................................................... 98

6 Appendices ............................................................................................................................ 99

7 References ........................................................................................................................... 102

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1 INTRODUCTION

1.1 Peatlands

Peatland ecosystems are terrestrial environments where, over the long term, net primary

production exceeds organic matter decomposition, leading to the accumulation of a deposit of

poorly decomposed organic matter, named peat, thicker than 30 cm (Wieder and Vitt, 2006).

Peat formation is the result of complex interaction among anoxic conditions, low

decomposability of the plant material and hydrology. Peatlands are distinguished by the presence

of water close to the surface, they often have unique soil conditions that differ from the adjacent

uplands, they support vegetation adapted to wet conditions and they are characterised by an

absence of flooding-intolerant vegetation. Climate and geomorphology control at the larger scale

the degree to which peatland can exist, but hydrology underpins its ultimate development,

affecting the physicochemical environment and influencing the biota (Gosselink and Mitsch,

2000). The persistence of peatlands depends from the constant supply of water, while source of

water influences its structure and function. Origin of water has been long employed for

classifying peatlands. When water comes from surrounding or underlying mineral soil, peatlands

are termed minerogenous (or geogenous). If peat has grown thick enough to progressively

become isolated from the mineral soil, precipitation becomes the only source of water and

peatlands are termed ombrogenus. Minerogenous water carries cations, anions and nutrients and

the resulting chemical composition has great effects on influencing flora, vegetation and

ecosystem functions. Minerotrophy is the term describing this condition and forms the ecological

basis for the peatland type named fen. On the contrary, ombrogenous water has very low

dissolved minerals and provides the ombrotrophic condition for the development of the peatland

type named bog.

Two main environmental gradients are responsible for the differentiation of peatland

habitat types. One gradient follows the variation in moisture and aeration as a function of water

table position, which changes in time and space, and of the pore structure of the peat. The

distance between the soil surface and the water table is the height above water table (HWT) and

represents roughly the depth of the aerated zone, which has high biological and ecological

relevance. The other is a complex gradient that combines the variation of pH, base richness and

nutrient availability. The gradient of nutrient availability and increasing productivity is described

using three terms borrowed from limnology, which are respectively oligotrophic, mesotrophic

and eutrophic. This gradient changes independently from the mineral gradient composed of pH,

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base cations and associated anions (Wieder and Vitt, 2006), indeed nitrogen, potassium and

phosphorus have their own chemistry and variation and all assumptions of strong correlation

with pH and base richness are not obvious (Rydin and Jeglum, 2013). The factors governing the

differentiation into the main habitats are the same as those causing vegetation differentiation, so

that a useful system to describe habitat variation in peatlands relies on vegetation pattern.

Therefore, the floristic composition mirrors change in pH and base richness, creating the bog-

poor fen-rich series. Bogs are always oligotrophic while rich fens can be either eutrophic or

oligotrophic. As the peatland becomes more isolated from the groundwater, the pH usually

decreases because Sphagnum mosses have the ability to acidify water, which cannot be

contrasted by the low buffering capacity of rainwater. The pH of bogs is comprised between 3.5

- 4.2, while it increases along the poor-fen rich-fen gradient from 4 to 8 (Rydin and Jeglum,

2013). The complete absence of bicarbonate alkalinity below pH 5.5 is a fundamental dividing

point in the habitat limits of many peatland species. Below 5.5 pH fens are dominated by

oligotrophic species of Sphagnum, while above 5.5 pH Sphagnum abundance decreases and true

mosses mostly dominate (Wieder and Vitt, 2006). Mire ecologists observed that the measure of

the water table from the surface (HWT) was a very strong predictor of vegetational gradients in

peatlands. The water table gradient allows the micro-topography of peatlands to be classified in

the following elements, which were first described in Sjrs (1948; cf. in Rydin and Jeglum,

2013):

Hummocks are mounds of 20-50 cm raised above the lowest surface level. The thick

aerobic peat supports dwarf shrubs and vascular plants.

Lawns are intermediate microforms, 5-20 cm above the water table, where graminoids

are most common. Mosses reach the highest species diversity in lawns. Firm surfaces that

you can walk on.

Carpets are from 5 cm below to 5 cm above the water table. They are characterised by a

soft, green layer of mosses and a sparse cover of cyperaceous plants. Softer surface into

which your feet sinks.

Mud-bottoms are most of the time submerged; they lack vascular plants and are

dominated by mosses or they can also expose the bare peat, usually covered by algae.

Pools are water-filled depressions.

Within the peatland, a vegetational gradient develops largely in parallel to HWT, which is deeper

at the uplands and shallower towards the open mire, passing through the mire margin and the

mire expanse. The mire margin has usually thinner peat and vascular plants can reach the

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mineral layer, so that small, creeping trees and shrubs can grow. pH is lower in the centre and

higher in margin and calcium ion shows the same pattern. In bogs and especially in raised bogs,

where the peat meets the upland soil, usually a narrow fen called lagg develops. The lagg is a

feature of the margin that surrounds the bog and receives water both from the bog and from the

surrounding mineral soils.

The peculiar shape of a peatland, its surface morphology and patterning is the result of

interactions between substrate, climate, hydrology and vegetation. Raised bogs are ombrotrophic

mires raised above the level of surrounding uplands, which are usually at least 500 m in diameter

and with a convex cupola that can be several metres higher than the edges and surrounding

uplands (Rydin and Jeglum, 2013). Water slowly flows from the centre to the edge. Where raised

bogs develop over round-shaped lakes, they grow circular and hummocks and hollows form

concentric patterns, perpendicular to the lateral through-flow.

1.2 Ecohydrology of raised bogs

Raised bogs are ombrotrophic mires that depend only on precipitation for their supply of water,

so that they can be found only in climatic regions where the input of water exceed the loss.

Raised bogs are characterised by vertical oscillation of the surface during drying and rewetting

(the so called mire breath), small temporal fluctuation of the water table, reduction of

evapotranspiration occurring at shallow water level and large storage coefficients (Schaaf, 1999).

In saturated condition, the content of water in undisturbed bog peatsoil can range between 88-97

% (Ivanov, 1981).

Peatlands such as raised bogs are not made of a uniform body, but they are characterised

by a double layer, and mire researchers introduced the terms acrotelm and catotelm to define this

particular feature (figure 1.1).

Figure 1.1 Bogs are formed by a double layer. The acrotelm, thin and biologically active and the catotelm, thick and inactive.

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The acrotelm is the superficial layer where water table fluctuation occurs and where

biogeochemical cycles and biological processes are more active, while under a hydrological

point of view, the high hydraulic conductivity and low degree of humification identifies the

acrotelm as an aquifer (Rydin and Jeglum, 2013; Schaaf, 1999). The acrotelm thickness can vary

from 7-8 cm in herb rich fen to 60-70 cm in moss-rich raised mire (Ivanov, 1981). The catotelm

is the main body of a bog and can be about several metres depth. It is the less active layer where

biological processes are slow, peat is generally well decomposed and hydraulic conductivity is

so low that, hydrologically, it is described as an aquitard. The classic Ingrams theory describes

raised bogs as raised water mounds, with the catotelm as a body that loses water by lateral

outflow and that is compensated for by infiltration by the overlying acrotelm (figure 1.2). The

mound shape is a function of the hydraulic conductivity of the catotelm (assumed low and

spatially constant) and the infiltration rate from the acrotelm (assumed constant).

Figure 1.2 Precipitation is the only source of water in raised bogs. Rainwater flows outwards from the centre via runoff, lateral through flow and groundwater flow.

The theory describes raised bogs as fed only by precipitation and isolated from the mineral

surrounding groundwater. However, recent studies have demonstrated that the isolation from

groundwater may not be always a rule in bogs. Relatively high hydraulic conductivity, ranging

from 10-2 to 10-1 m d-1, was measured even in the catotelm (Chason and Siegel, 1986) so that, if

there is no net hydrological distinction between acrotelm and catotelm, there might be significant

downward and upward exchanges of water (Sirin et al., 1996 cited in Rydin and Jeglum, 2013)

(figure 1.3). Internal hydrological mechanisms, the change in precipitation, evapotranspiration

and the degree by which water table fluctuation affects the hydraulic head in the catotelm govern

the process and can cause the reversal of hydraulic gradient (Fraser et al., 2001; Devito et al.,

1997). Consequently, during droughts, the water table drops and mineral groundwater can move

upward reaching 1-2 m depth from the surface thus affecting pore water chemistry (Gosselink

and Mitsch, 2000).

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Figure 1.3 For raised bog with particular climatic and peat conditions upwelling of mineral water can occur.

Hydraulic conductivity (K) is the property that governs the flow of water and transport of

solutes in a porous medium. In a bog, hydraulic conductivity varies with peat type, degree of

humification and bulk density. Peat type is characterised by different structures influencing

hydraulic conductivity, for instance K is lowest for Sphagnum peat and higher for lignoid peat

(Rydin and Jeglum, 2013). Generally, K correlates negatively to degree of humification. The

degree of humification can be measured with the von Posts method, which assigns an increasing

number to the change of some diagnostic features that correlate with degree of decomposition.

Humification can vary either vertically on a scale of decimetres as horizontally on a scale of few

metres (Schaaf, 1999). Humification, however, does not increase necessarily with depth because

it also depends on different factors: vegetation (Silamikele et al., 2007), peat forming conditions,

different decaying rate in the catotelm, physical conditions, micro-topography and climatic

conditions (Schaaf, 1999). Hydraulic conductivity correlates negatively to bulk density. Bulk

density tends to increase with depth since compaction caused by overlaying peat layers reduces

the volume of pores. All the factors described above vary spatially so that also K is spatially

dependent, despite the Ingrams theory assumed constant hydraulic conductivity throughout the

catotelm. Indeed, more recent works showed that in the catotelm K tends to have a vertical

decreasing trend with depth while it tend to decrease from the centre to the bog margin (Beer et

al., 2008; Schaaf, 1999), invalidating the assumption of the Ingrams theory. The way K varies in

a bog controls the vertical and horizontal flow of water. In bogs, water has a preferential

horizontal flow because vertical hydraulic conductivity can be as three orders of magnitude

lower than horizontal K (Beer et al., 2008). In the acrotelm, compaction and decomposition

cause the hydraulic conductivity to decrease strongly with depth from about 105 m d-1 at the

surface to 110 m d-1 at some decimetres below it (Schaaf, 2004; 1999). Fraser et al. (2001)

measured a great decrease of K that reached 10-3 m d-1 at 45 cm below the surface. Yet, hydraulic

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conductivity can be dynamic. Drying can decrease K compressing the pores while precipitation

can cause peat to swell, increasing its hydraulic conductivity. Given the vertical trend of K and

the peat properties change during drying-rewetting cycles, the rate of lateral subsurface flow in

the acrotelm depends on water table position, so that flow decreases as water table decreases.

This mechanism provides an important feedback that augments the water retention power in bogs

during droughts. In the acrotelm, hydraulic conductivity is usually 4-5 times greater horizontally

than vertically (Beer et al., 2008; Schaaf, 1999). Runoff can occurs if the rate of precipitation

fills the storage capacity of hollows and exceeds the rate of infiltration. In ecosystems like raised

bogs, characterised by a small slope and a well developed hummock-hollow topography, runoff

occurs as sheet flow and channel flow in micro channels. The micro-topography adds complexity

to the flow pattern because depressions can be connected during wetting and can be disconnected

during drying and the flow pattern can reverse due to the different filling and emptying pattern of

depressions while some depression can also not participate to surface flow (van der Ploeg et al.,

2011). The transition between subsurface flow and runoff is controlled by complex threshold

mechanisms, and Frei et al. (2010) found that surface flow is mainly a function of depression

storage and that runoff occurred only during intense rainfall. Subsurface flow is then the

dominant flow during dry and wet periods while surface flow occurs only during intense rain

events because micro-topography acts as inhibitors of surface flow (Frei et al. 2010).

1.2.1 Height above water table

Height above water table (HWT, i.e. the distance from peat surface to the water table) is

an important indicator used to predict a number of important eco-hydrological variables in

peatland hydrology, ecology and biogeochemistry including run-off, saturation, redox potential,

biodiversity, soil structure, methane emission, peat quality and organic matter decomposition

(Waddington et al., 2014). Water table governs the water availability to bog plants and

delimitates the zones of aerobic and anaerobic respiration in peat, so that it is considered a main

factor controlling vegetation distribution (Rydin and Jeglum, 2013), microbial diversity and

niche differentiation (Andersen et al., 2013). Water table fluctuation governs the oxygen

penetration in peat and affects redox potential (Niedermeier and Robinson, 2007; Mansfeldt,

2003; Seybold et al., 2002), thus determining the spatial distribution of aerobic conditions. That

condition enhances decomposition not only because aerobic respiration is more efficient than

anaerobic respiration, but also because it allows the degradation of anti-microbial Sphagnum

phenols in shallow peat (Abbott et al., 2013; Fenner and Freeman, 2011). Water table level is

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also the most important control on the relative importance of methane production in peat and of

the amount of CH4 fluxes out of peatlands (Bridgham et al., 2013). Yet, from a hydrological

point of view, water table level can provoke flow reversal in peatlands (Fraser et al., 2001) and it

is an indicator of water discharge and solute transport out of the peatland into the watershed.

1.2.2 Micro-topography

The surface of raised bogs is characterised by distinct hummocks and hollows micro-topographic

elements. Micro-topography features can alter hydrological, physicochemical and biological

attributes (Courtwright et al., 2011; Baldwin et al., 2007; Haraguchi, 1992), adding spatial and

temporal heterogeneity to the turnover of redox-sensitive solutes in peatlands. Therefore, they

are not just depending on peat quality, peat properties and temperature. The most important

difference between hummocks and hollows is water-table depth (Bragazza et al., 1998).

Hydrologically, in continental bogs, above the water table, hummocks have a dominance of

macropores associated with vascular plant roots, so that they have higher hydraulic conductivity

than hollows and the water flow is predominantly vertical. The difference in peat physical

properties and hydraulic conductivity accounted by microforms become less important with

depth to water table. Branham and Strack (2014) showed that hydraulic conductivity was not

influenced any more by microforms at a depth of 20 cm below the water table. Frei et al. (2012)

showed that micro-topography could induce water flow patterns that eventually determined

biogeochemical hot spots, defined as sites characterised by higher biogeochemical activity than

the surrounding areas. For example, when water table increases, hummocks can be the only place

where oxidation can occur. Indeed, a higher concentration of nitrate is usually found in

hummocks (Frei et al. 2012; Wolf et al., 2011; Bragazza et al., 2005). Through the infiltration

zone inside the hummocks, oxidised nutrients can be transported by precipitation to deeper and

more redox-reducing levels where they can be reduced. In this occasion and at the water table

level, hummocks can become hot spots for reducing reactions (Frei et al. 2012).

1.3 Carbon cycle

The study of biogeochemical cycles is of primary importance to unravel the ecological role of

peatlands at different spatial and temporal scales. Regarding the carbon cycle (figure 1.4),

Northern peatlands store half (approximately 1672 Gt) of the global soil carbon pool though

cover an estimated 3.6x106 km2 (Tarnocai et al., 2009), which is equal to about 3 % of the total

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lands. Peatlands can be source, sink or transformers of carbon compounds, and many efforts

were directed to measure its fluxes and to understand their functioning.

Figure 1.4 Carbon cycle in bogs. Circle boxes represent gas phase; broken lines represent microbial-mediated processes.

1.3.1 Decomposition

Measurement of decomposition rate (k) is important for the biogeochemistry of nutrients and

carbon fluxes (Prescott, 2010). The availability of nutrients is due in large part to the decay

dynamics of the organic matter. The process also supports diversity in the microbial population

by supplying a rich set of intermediate degradation products (Berg and McClaugherty, 2014).

Bogs receive very low amount of nutrients from wet and dry deposition that made up their entire

external sources. For that reason, decomposition and recycling of nutrients become an essential

internal source of nutrients in bogs (Bragazza et al., 2013).

Zhang et al. (2008) reviewed the factors that controlled decomposition rate. These were:

(i) climatic factors (mean annual temperature, MAT; mean annual precipitation, MAP; annual

actual evapotranspiration, AET); (ii) litter quality (nitrogen content; carbon:nitrogen ratio, C:N;

lignin content, LIGN; lignin:N ratio, LIGN:N); (iii) vegetation and litter types; (iv) geographical

variables (latitude, LAT and altitude, ALT). The authors found that C:N ratio and the total

nutrient content of the litter (both litter quality factors) were the two most important factors

influencing decaying rate on a global scale. A threshold-based mechanism dominates the relation

between decomposition rate and its factors. There is not a single factor that is, in every

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conditions, more important than the other ones, so that decomposition can slow down despite all

the factors but one is adequate (Prescott, 2010).

In bogs, the litter quality of Sphagnum overwhelms the importance of other factors.

Sphagnum species forming hummocks have relatively higher content of recalcitrant structural

carbohydrates (lignin-like polymeric phenolics) than hollows species. Thus, despite hummocks

are well-aerated and aerobic decomposition rate is higher than anaerobic decomposition, the

decomposition of Sphagnum litter is lower in hummocks (Bragazza et al., 2013). After litter

quality, moisture (i.e. HWT) is the most important factor for decomposition that can show a

threshold effect. In fact, a lower water table caused by droughts can desiccate the litter

hampering decomposition. Laiho (2006), in a review on decomposition constraint in northern

peatlands, included pH as important factor because it can limit the activity of phenol oxidase.

Finally, several studies have shown that temperature increased decomposition only under non-

limiting moisture condition (Bragazza et al., 2013; Mkiranta et al., 2009).

Hydrologic dynamics, interactions between surface ombrogenous water and deeper

groundwater, influence the distribution and transformation of nutrients in wetlands (Frei et al.

2010) and the decomposition rate in the catotelm (Beer and Blodau, 2007). If in the catotelm

diffusion dominates over advection there will be a lack of solute transport and an accumulation

of product of microbial metabolism, such as CH4, dissolved inorganic carbon (DIC) and

recalcitrant dissolved organic carbon (DOC) (Beer et al., 2008). Beer and Blodau (2007) showed

that these phenomena can lead to a thermodynamic constraint for microorganism metabolism

that slows or even cuts off the decay of peat in anoxic layers. It follows that the groundwater

residence time becomes a fundamental factor in assessing the spatial decomposition rate in bogs

(Morris et al., 2011). Advective flux like upwelling of mineral water or downwelling of

rainwater can break the constraint. The influence of rainwater and upwelling depends on the

profile of hydraulic conductivity. It is reported that in bogs at most the first 100 cm can be

affected by rainwater (Dobrovolskaya, 2013; Morris et al., 2011).

1.3.2 Spatial process of decomposition

Main inputs of carbon in peatland are through photosynthetic production of mosses and then

vascular plants. In peatlands, plant biomass can be divided into above ground, rhizome and

coarse roots, and fine roots (diameter < 0.5 mm). Carbon input below ground takes place through

root exudations and decaying vascular plant materials. Most of the organic matter decomposition

occurs in the acrotelm, where rate of decomposition has been estimated to be one hundred times

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higher than in the catotelm. There is an important biogeochemical connection between the upper

layers and the deep peat, so that about 10 % of litter mass produced in the acrotelm reaches the

catotelm (Frolking et al. 2001 cited in Beer et al. 2008), albeit the occurrence of this

translocation is actually strongly dependent on hydraulic conductivity (Beer et al. 2008).

Products of decompositions are CO2 and CH4 gases and DOC. The relative importance of

aerobic and anaerobic respiration depends on the HWT, which controls the oxygen supply into

the peatsoil. The principal product of aerobic respiration is CO2, while anaerobic respiration

produces DOC, CH4 and CO2 (Fenner and Freeman, 2011). Oxic conditions favour CO2

production and inhibit methanogenesis, while anaerobic respiration favours lower CO2

production but greater CH4 production rate (Estop-Aragons and Blodau, 2012). Respiration rate

is higher at water table level and then decreases (Shoemaker et al., 2012) so that most of the

organic matter is decomposed in the acrotelm (Beer and Blodau, 2007). In anoxic layers,

respiratory pathway depends on Eh conditions, peat quality and relative concentration of electron

acceptors. Microorganisms compete for electron acceptors in presence of nitrate, ferric ion,

sulphate and finally CO2. Since bogs are fed only by precipitations, the common electron

acceptors and nutrients have very low concentrations compared to other minerogenic wetlands

and they even decrease with depth (Beer et al., 2008; Proctor, 2003; Steinmann and Shotyk,

1996; Lundin and Bergquist, 1990). Keller and Bridgham (2007) measured that Fe(III) and

nitrate reduction accounted for less than 1% of anaerobic carbon mineralisation, while sulphate

reduction was responsible for 6 26 %. These results are consistent with the frequent

observation that reactive inorganic sulphur (RIS) pool can be very dynamic in bogs (Wieder and

Lang, 1988). Many studies have stressed that respiration driven by common electron acceptors

accounts for only a minor fraction of the CO2 produced in bogs. The sources for the unexplained

CO2 can be the regeneration of electron acceptors during water table drawdown and/or other

respiratory pathways like bacterial respiration with humic substances (Knorr et al., 2009; Deppe

et al., 2009; Heitmann et al., 2007), with organic sulphur species (Kertesz, 2000), and

fermentative processes in absence of electron acceptors. Fermentation can be very important in

bogs and can lead to high production of DOC (Vile et al., 2003). Fermenting microorganisms are

important in degrading complex polymers yielding simpler products used in methanogenesis.

Fermentation, together with DOC reduction, is addressed as responsible of the high fraction of

unexplained carbon mineralization in anoxic layers (Keller and Bridgham, 2007). There is an

increasing need to assess the use of organic electron acceptors (DOC) by microorganisms in

anoxic environment. Comparing sites with different hydrology (bog and fen) Shoemaker et al.

(2012) found a high production of CO2 in anaerobic layer at the ombrotrophic site and found a

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local correlation with concentration of DOC, suggesting that DOC may have driven anaerobic

respiration. A further complication to the role of DOC is that the chemical composition rather

than its amount distinguishes its role as electron donor or electron acceptor (Keller et al., 2009).

The vertical profile of DIC generally increases with depth (Deppe et al., 2009). Several

authors observed that DIC increased at all depths during summer months, along with

temperature. Despite the temperature drop during autumn, DIC is reported to be still increasing

and Shoemaker et al. (2012) reported that the profile of DIC decreased only after frequent rain

events occurred, since rain dilutes bog water and favours CO2 degassing. The DIC profile

depends also by emission to atmosphere, which can occur by diffusive flux or by non-diffusive

flux via roots of vascular plant or via ebullition (Glaser et al., 2004; Tokida et al., 2007).

Substrate quality decreases with depth and the main substrate for decomposition in

deeper layer becomes DOC. DOC concentration is higher in surface layer and decreases with

depth (Deppe et al., 2009; Beer et al., 2008; Fraser et al., 2001). Pore water concentration of

DOC usually increases in summer due to evapotranspiration and enhanced decomposition

(Waddington and Roulet, 1997).

Supply of high quality substrate and water table depth (Rydin and Jeglum, 2013) are the

main factors controlling the production of methane in peatlands. In general, acetoclastic

methanogenesis is favoured in the upper peat layers where abundant labile organic carbon from

decaying vegetation are found, and H2 and CO2-dependent methanogenesis predominates in the

more recalcitrant deeper peat layers (Beer and Blodau, 2007). The zone of most active CH4

production was measured about 10 cm below the average water table (Sundh et al., 1994), where

there is input of fresh organic matter. This zone is closely related to the source of products from

anaerobic respiration, in fact Shoemaker et al. (2012) found a peak of methane production just

below the respiration rate peak near the water table surface (often between 0-5 cm below the

water table). Regarding the role of micro-topography, the hummocks are reported to have lower

methanogenesis than hollows because there is less quality input of substrate (Bubier et al., 1993).

It is also arguable that, being oxygenated, the hummocks provide more energetically favourable

microbial pathways.

Peak of oxidation of methane is found at the water table level and above it and it is higher

in hummocks than in hollows (Frenzel and Karofeld, 2000). Occasionally methanotrophy can be

found in saturated layer. In this case it is driven by oxygen released by roots and by anaerobic

oxidation of methane.

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1.4 Microbial community

Mires are active systems characterized by organic matter turnover (Biester et al. 2012).

Microorganisms play a key role in wetland soils. They are involved in fundamental

biogeochemical cycles and transformation of elements. The rate and extent of their role in these

processes depend on the variability and change of environmental variables on which their

metabolism and growth depend (Andersen et al., 2012). The activity of microorganism depends

on temperature, pH, hydrological regime, oxygen, peat quality, nutrients availability and redox

potential. These variables are spatially dependent and, in peatlands, different habitats can show

changes in the distribution of microbial groups and in their total biomass. Distribution of

microorganism is also depth dependent. Microorganisms face increasing energetic constraints

with depth, caused by a combination of factors such as the availability of oxygen and

thermodynamic disadvantage that are characteristic of a stratified peat with redox and peat

quality change (Robroek et al., 2013). Fungi, as general rule, are dominant in oxic and acid

habitats, so they are abundant in the aerated surface layer. Microhabitats as hollows and

hummocks have different vegetation, hydroperiod and redox dynamic. In Robroek et al. (2013)

an analysis of microbial community between microhabitats has revealed a change in the

fungal/bacterial ratio toward higher values in hummocks than hollows. Generally fungi decreases

in biomass with depth (Andersen et al., 2012), but recently Jassey et al. (2011) have found larger

number in lower peat layer, a finding described also in Dobrovilskaya (2014). Bacteria instead

dominate in anoxic and neutral environment. The bacterial abundance generally decreases with

depth, but some studies have revealed increasing biomass with depth (Golovchenko et al., 2007),

or peaks at some depth (Grodnitskaya and Trusova (2009) cited in Rydin and Jeglum, 2013).

In a peatland, microorganisms respond to alteration of hydrological regime following the

time and space changes and the intensity of the event. Persistent drought could trigger

vegetational shifts and indirectly affect microorganisms. During drought methanogenic bacteria

can survive in anoxic pores still present in aerobic layers (Kotiaho et al., 2013) and aerobic

bacteria can increase in number following the lowering of water table (Karsisto, 1979).

Methanotrophic bacteria have been found throughout the profile but in greater concentration at

certain depth so that under particular condition the population can become active.

Methanotrophic bacteria, which can tolerate extended periods of anoxia, can resume methane

oxidation within few hours of re-exposure to oxygen.

1.5 Vegetation season: drying and rewetting

16

Drought is a period of scarcity of water that is statistically derived from climatic data. Regarding

the effect on biogeochemical processes in peatlands, droughts can be defined according to their

degree of intensity. Severe droughts can affect microorganisms either directly or indirectly, on a

long-term range, after driving a shift in vegetation. Short droughts, although having minor effect

on vegetation, can also affect directly the microbial community. Droughts influence microbes by

lowering the moisture content of peat (i.e. water table) and enabling the penetration of oxygen in

deeper peat layers so that aerobic respiration can establish. If drought is severe, shortage of water

can lead to microbial mortality (Mettrop et al., 2014). Regarding climate change, the effects of

droughts on the carbon cycle is under high concern since an increase of their severity or

frequency could change the role of peatlands from being sinks of carbon to being sources.

Drought intensity and following rewetting are the most important factors that can

enhance decomposition of peat soil. Water table drawdown allows oxygen to reach previous

anaerobic layers of peat. Usually, although the pattern is very peat type-dependent, CO2

increases during drying until an optimum moisture level is reached and afterward it decreases

(Estop-Aragons and Blondau, 2012), because moisture becomes limiting. The higher

decomposition rate occurring during droughts, allows the release of nutrients in the peat, a

process called eutrophication. During the rewetting phase, these nutrients can be transported

downward and can serve to enhance microbial decomposition in low nutrient layers.

Fenner and Freeman (2011) have proposed a theory, called the enzymatic latch theory,

which aims to explain the mechanism underlying the bogs response to droughts. As mentioned

above, the persistence of bogs is due mainly to waterlogged anoxic condition and recalcitrance of

Sphagnum. Sphagnum spp. have high amount of phenolic compounds, with inhibiting effects on

microbes, that can be degraded only in presence of oxygen by phenol oxidases enzymes. When

water table is high, oxygen cannot penetrate in peat and that can prevent the peat deposit to be

released as CO2. When drought introduces oxygen, phenol oxidase can remove phenolic

inhibitors, enabling hydrolases to resume normal mineralization of organic matter and increase

decomposition.

Consequently, droughts influence microbial activity and lead to a hydrochemical shift in

surface water. As mentioned above, eutrophication releases nutrients, while another important

effect of aeration is the re-oxidation and regeneration of electron acceptors (Deppe et al. 2009).

At the same time, the oxidation of reduced redox couples by aerobic respiration releases

hydrogen ions bringing acidification (Brouns et al., 2014; Mettrop et al., 2014; Clark et al.,

2009), which in turn further limits microbial activity. In Fenner and Freeman (2011), sulphate

and nitrate were released after 23 days of drought while potassium and P after 48 days. Sulphate

17

was released within weeks in Brouns et al. (2014) and Proctor (1994), followed by acidification.

The time scale of response of oxidising and reducing processes depends on process, depth and

peat heterogeneity (Knorr et al., 2009). According to the enzymatic latch theory, if a drought is

not intense enough, it is possible that phenolics will not be degraded and then eutrophication and

increasing of CO2 will not occur (Mettrop et al., 2014). Fenner and Freeman (2011) have

observed a strong increase of CO2 after a severe drought. Despite the increasing production of

CO2, in the zone of aeration, DIC can be very low because of degassing (Deppe et al., 2009). The

production of DOC seems related to drought intensity. During aeration, DOC becomes the

substrate of aerobic respiration (Hughes et al. 1997) and it decreases (Mettrop et al., 2014;

Fenner and Freeman, 2011). However, in a mesocosm experiment, Mettrop et al. (2014)

observed that after strong desiccation DOC greatly increased, which the authors proposed that

this is a consequences of microbial mortality or a change in microbial community composition.

Rewetting in bogs occurs by means of precipitations. The infiltration of rainwater to

lower layer is important after droughts (Deppe et al. 2009) because causes the pH to increase

(Fiedler et al. 2007). Fenner and Freeman (2011) measured increasing flux of CO2 after

rewetting, which the authors explained as an effect of eutrophication and removal of pH

constraints. Deppe et al (2009) measured a rapid increase of DIC and CH4 after flooding,

whereas drying and following rewetting had not strong effects on DOC. Clark et al. (2009)

measured that the drying-rewetting cycles influenced DOC down to 55 cm depth. Proctor (2003)

observed a sharp peak of SO4 after rewetting (Proctor, 2003).

1.6 Redox potential

Redox reactions are chemical reactions that involve the transfer of electrons between two

species, molecules, atoms or ions, so that their oxidation state changes. Oxidation is the loss of

electrons or an increase of oxidation state and reduction is the gain of electrons or a decrease in

oxidation state. Reduction and oxidation must occur simultaneously, since, in a redox reaction,

the reducing agent transfers the electrons to an oxidised agent and free electrons cannot exist in

solution. Conceptually, redox reactions are described as two half-reactions, one releasing

electrons and the other gaining electrons, combined to form a whole reaction. The two related

species that exchange electrons and change their oxidation state in a redox reaction are called

redox pair or redox couple. In analogy with pH, the hypothetical electron activity of a solution is

given by p = - log{e-}. The quantity p measures the relative tendency of a solution (with one or

more redox couples) to accept or transfer electrons and it is a measure of the Gibbs free energy

of the redox reaction. Also, the redox potential (Eh) is a measure of the electron availability in a

18

solution, though it is made with an electrochemical cell, where the potential is expressed in

relation to the standard hydrogen electrode. In making an electrochemical measurement of the

redox intensity, an electro motive force (Volts) is measured. Eh is related to P by

hERT

Fp

3.2 Eq. 1.1

Where Eh is the redox potential [V] of the solution (in relation to reference electrode); F is the

faraday constant (= 96940 Cmol-1 e-); R is the gas constant (= 8.314 J mol-1K-1); T is

temperature in K.

The thermodynamic definition of the redox potential of a solution is given by the

Nernsts equation, which describes the relationship between Eh and the activities of oxidised and

reduced species

d

Ox

nF

RThEEh

Relog3.2 Eq. 1.2

Where Eh is the redox potential [V] of the solution (in relation to reference electrode); Eh is the

redox potential [V] under standard conditions (all activities = 1, pH2= 1atm, [H+] = 1M) and is

related to the free energy change for the cell reaction (G); F = 1 is the faraday constant (=

96940 Cmol-1); n is the number of exchanged electrons; R is the gas constant (= 8.314 J mol-

1K-1); T is the temperature in K; 2.3RT/F = 0.059 V (at 25 C); {Ox}/{Red}is the activity of

oxidised/reduced couples.

From equation 1.2, it follows that Eh depends on the ratio of oxidized and reduced forms

(i.e. the relative activities) and not on their absolute quantities. For that reason, Eh is considered

an intensity factor for the reduction/oxidation (i.e. the greater the ratio the more oxidised the

system). Contrarily, redox capacity, a function of the amount of oxidised and reducing

compounds, measures the buffer effect of the solution. The intensity factor determines the

relative ease of the reduction/oxidation whereas the capacity factor denotes the extent to which

the shift will take place. In aqueous natural environments, the water solvent exerts a levelling

influence on the system and restricts the range of accessible Eh between the intensity of water

reduction (E= - 0.83 V) and the intensity of water oxidation (E= + 1.23 V).

Redox reactions are common in nature and they are primary driver of biogeochemical

processes. The metabolism of living organisms relies on redox reactions for energy and for

provision of building blocks, such as for example respiration. Microbes are an important control

19

of redox reactions in soils. They catalyse the kinetic of redox reactions, transferring electrons

from reduced inorganic or organic matter to inorganic or organic terminal electron acceptors

(TEAs), to produce energy for their metabolism. The kind of redox couple, electron donor and

electron acceptor, determines the amount of energy (i.e. Gibbs energy) gained by the

microorganisms out of the process. In a closed aquatic system with organic matter as energy

source and microbes, it is possible to calculate the sequence of oxidation reaction coupled to

TEAs expected on the base of their thermodynamic possibility. Oxygen is the most abundant and

the strongest oxidised agent that is readily available in the atmosphere so that, in aerobic

condition, microbial respiration uses oxygen as final electron acceptor to oxidise organic matter

and produce energy. As water table rises, water saturates the soil and slows down the rate of

oxygen diffusion. Consequently, microorganisms consume oxygen more rapidly than it is

supplied, and, after oxygen is completely depleted, the oxidation of organic matter is coupled to

other TEAs. The temporal sequence predicted on the basis of decreasing Gibbs free energy is the

following: NO3-, Fe(III), Mn(IV), SO4

2- and finally CO2. To put it differently, nitrate has a better

affinity for electrons than sulphate so that the microbes require a higher electron pressure (lower

Eh) to transfer electrons from organic matter to sulphate.

Each of these reactions occurs within defined values of redox potential, or thresholds,

which can be used to predict the occurrence of the relative redox reactions in soils, though there

are currently arguments about where to draw these boundaries and the matter needs to be verified

more extensively. According to Gosselink and Mitsch (2007), local factors like pH and

temperature, influence the values of these thresholds. Nevertheless, if caution is taken in

interpreting the results they can still be useful in characterising redox condition in soils (Reddy

and DeLaune, 2008). In addition, it has to be borne in mind that such boundaries are rather

smooth stripes than straight lines, therefore they should not be taken as exact borders between

soil redox reactions. According to Gosselink and Mitsch (2007), the microbial oxidation of

organic substrate uses oxygen as terminal electron acceptor at a redox potential of between +400

and +600 mV (Reddy and DeLaune (2008) suggested that below 300 mV oxygen is completely

absent). Below +400 mV nitrate starts to being used as electron acceptor; at about +225 mV

manganese is reduced; between +100 and -100 mV iron is transformed from ferric to ferrous

form; from -100 to -200 mV sulphate is reduced to sulphide; eventually, below -200 mV carbon

dioxide is reduced to methane. There are different thresholds and classifications of redox zones

in literature. Figure 1.5 shows a comparison among other classifications.

20

Figure 1.5. Redox zones defined by different authors, based on Fiedler (2007).

As mentioned before, an electrochemical cell measures the difference in potential

(electron motive force, emf) between an inert indicator electrode in contact with a redox couple

in solution and a reference electrode. The indicator electrode is usually an inert metal like

platinum and the reference electrode can be H+/H2 or Ag/AgCl 1mol/l KCl. The reaction of the

cell where the reference electrode is hydrogen is

Pt, H2 (pH2=1)| H+ (a=1) | Mn+ | M

The overall cell reaction is

Mn+ + nH2 M(s) + 2nH+

If the electrode potential is positive, the above reaction is the spontaneous reaction in the

direction left to right. If the electrode potential is negative, the spontaneous reaction is in

the opposite direction. The voltage of the cell, when the activities of all ions in the cell are

21

unity, when gases are at 1 atm pressure and solids are in their most stable form at 25C, is

calculated as

Ecell = Eh(ox-red) Eref Eq. 1.3

Rearranging eq. 1.3 gives

Eh(ox-red) = Ecell + Eref Eq. 1.4

That is the redox potential for the solution measured.

The measurement of redox potential with platinum electrodes has wide applications. The

assessment of anaerobic condition in soils and sediments is important for many disciplines such

as ecology, ecotoxicology, soil science and agronomy. Redox potential influences the

availability of redox sensitive nutrients, their removal and/or translocation in soil profiles and the

efflux of solutes in percolating water (Chadwick and Chorover, 2001). For example, redox

condition influences the nitrogen cycle in soils, determining the nitrogen emission as N2O, N2 or

NH3, or the plant uptake as NH+

4 + or NO3-, with obvious consequence for agriculture and water

quality. Regarding the carbon cycle and global warming, methane, a strong green house gas, is

produced in wetlands depending on redox condition. Most agricultural systems rely on aerated

soil so that roots respiration can occur. Reducing conditions can produce toxic compounds for

plants (reduced forms of Fe and Mn, cyanogenic compounds, ethanol, lactic acid, acetaldehyde

and aliphatic acids such as formic, acetic, butyric acids, and H2S) impairing their growth and the

crops yield. Again, the fate of toxic compounds in the environment depends on redox potential

since the mobility of heavy metals, and that of non-metal like arsenic as well, is redox-sensitive.

Moreover, the degradation of organic pollutants and pesticides in groundwater are affected by

oxidising/reducing condition.

The U.S. National Academy of Sciences Definition states that the minimum essential

characteristics of a wetland are the recurrent, sustained inundation or saturation at or near the

surface and the presence of physical, chemical, and biological features reflective of recurrent,

sustained inundation or saturation. (Gosselink and Mitsch, 2007). In peatlands, where water

saturation and anaerobic conditions are essential characteristics, it is possible to use an electrode

to measure oxygen penetration, in order to characterise the redox conditions and to relate them to

the dominant redox processes (Pezeshki, 2001). When Eh time series is measured, and the

22

relative frequency of redox potential values taken at certain depths is calculated, is possible to

determinate to which redox zones the peat layer corresponds. For instance, Fiedler and Sommer

(2004) related the frequency of Eh to diagnostic features in hydric soil, finding that the

thresholds for Mn reduction was Eh < 450 mV, for Fe(III) reduction was Eh < 170 mV and for

CH4 oxidation was Eh > 75 mV. Measured Eh must be seen as an integrated operational

parameter, which is influenced, as stated, by the activity of living microorganisms but also by

external factors such as variation of water table, precipitation, source of water and chemistry,

pH, temperature and organic matter (Fielder et al., 2007). So far, a number of problems have

hindered the broader application of Eh direct measurement. First, the employment of platinum

electrode probe in in-situ condition requires rugged electrode device to tackle harsh

environmental conditions like storms, prolonged submersion, extreme temperatures and animal

disturbance. Second, irreversibility of coating reactions at the electrode, slow reaction kinetics

and mixed potentials can complicate the interpretation of redox measurements (Peiffer et al.,

1992; Stumm and Morgan, 1981). Last but not least, redox potential shows high spatial and

temporal variability. To conclude, on one hand unpredictable incidents or technical problems

could prevent the collection of reliable data and, on the other hand, the high number of

interplaying variables might prevent simple interpretation of redox potential (Husson, 2013). To

date these issues have slowed the measurement of redox potential in wetlands, leading to a

scarcity of studies (De Mars and Wassen, 1999). Frequently, when a study on redox potential

was performed, it was designed without permanent continuous measurements or without vertical

profiles measurements (Shoemaker et al., 2013; Fiedler et al., 2007). The common use of single

time measurement with intervals in the order of days or weeks cannot account for the variability

expressed by wetland systems. Indeed, in wetland soil redox potential may have periodical or

occasional fluctuations within time of hours and at different depths (Vorenhout et al., 2004).

Moreover, most of the available studies have focused on minerogenous systems with an even

greater lack of publications on redox measurements in ombrogenous bogs.

In the next section, the factors that control redox potential in peatlands are examined, in

order to clarify, where possible, the nature of the relation, the relative importance and the

potential feedbacks.

1.7 Redox potential is a complex indicator

Redox potential in peatland depends on several factors, namely: peat property and quality,

temperature, pH, microbial community, type of vegetation, hydrology and water chemistry

23

(figure 2.2). Since redox potential is an indicator of several processes, the aim here is to

understand how each factor can influence it. The method is based on a review of previous

researches in order to trace emerging trends, which will be discussed in the light of the theory

discussed above and the peculiar characteristic of bog peatlands. In this section, each factor will

be analysed separately for the sake of clarity, though it must be stressed that, in nature, they are

not isolated but they closely interact.

Eh

Hydrology

Water table

Precipitation

Groundwater

Microbial community

Temperature

pH

Peat properties and

peat quality

Vegetation

Water chemistry

Figure 1.6 Controls of redox potential.

Redox potential has a highly dynamic nature and is spatially dependent. The discussion

starts with an evaluation of spatial and temporal variability in the measurement of redox

potential.

Spatial and temporal variability of Eh

In typical wetland soils, Eh values vary from 300 to 700 mV (Reddy and DeLaune, 2008).

According to the work of Fiedler et al. (2007), the range of temporal variation of redox potential

in wetlands can be identified in: short-term, diurnal, single event, seasonal and annual variations.

Periodic diurnal fluctuations can be explained in term of Nernsts equation and the vant Hoffs

law. In this case, the fluctuations are driven by temperature with a temperature maximum

followed by Eh minimum. Plant living roots have daily metabolic cycles. In saturated anaerobic

soils, if aerenchymatous roots are enough dense, redox potential may fluctuate following the

24

release of oxygen that has its peak during the highest photosynthetic rate (i.e. light intensity).

Periodical redox variation can also be driven by periodic physical phenomenon. For example,

Catallo et al. (1999) studied the effect of a 12-hours tide fluctuation in a salt marsh and found

that it corresponded to redox potential fluctuation of about 40-300 mV. In Fiedler et al. (2007),

drying and rewetting are considered single event changes that may bring redox potential to vary

up to 900 mV. Seasonal changes can be related to gradients in soil temperature that affects

microbial activity. Short-term changes are the most difficult to characterise. They can be caused

by soil chemistry dynamics that may result from the production of carbon dioxide, the release of

hydroxyls in ferric iron reduction, the chemistry of precipitation and from inputs of redox

sensitive species. Measurement of redox potential often showed short-term peaks or daily

variation, though so far, there have not been studies that addressed and explained the underlying

causes.

Redox potential can vary with respect to centimetres or even millimetres in soils. This

variability is induced by partial water saturation of soil structure, absorption on soil minerals and

organic matters, plant roots and microbial activity (Yang et al., 2006). Each of these factors

varies spatially and the soil heterogeneity may cause the electrodes positioned at the same site

and depth to show simultaneously a wide range of Eh values (Urquhart et al, 1972). However,

De Mars and Wassen (1999) investigated the role of heterogeneity in different peatlands and,

placing from 4 to 12 electrode replicates within an area of 10 m2 and at 15 cm below surface,

showed that the variability of Eh among the replicates was low, even at low water level. Yang et

al. (2006) measured that spatial redox variability is higher vertically than horizontally, which

seems reasonable given that peat properties that are reported to vary more vertically than

horizontally (Schaaf, 1999). In conclusion, the degree and scale of heterogeneity will affect the

variability of redox potential, though in permanently saturated soils or in areas with minimal

hydrological fluctuations the variability could be minimal (Reddy and DeLaune, 2008).

In wetlands, redox potential generally follows a depth gradient. The vertical profile is

created by local hydrology and addition of electron acceptors and donors and typically decreases

with depth due to oxygen intrusion from the surface (Reddy and DeLaune, 2008). Nonetheless,

other studies showed quite different redox profile patterns. For instance, Vorenhout et al. (2011)

observed that redox potential at deeper soil layers occasionally increased and exceeded the redox

potential at upper layers and that some sensors located at greater depth showed steadily higher

redox potential than overlying layers. Urquhart et al. (1972) observed in a bog a similar pattern,

where some of the redox potential profiles (0 - 30 cm) showed deeper layers having higher Eh

than overlying layers. It follows that redox profile in peatlands can be dynamic and that a simple

25

decreasing gradient is not the general pattern. Hydrology and water circulation are important

factors that can influence the vertical profile and can disrupt the vertical redox stratification that

might be predicted by oxygen intrusion alone (Thompson et al., 2009). However, as mentioned

before, there have not yet been direct field investigations of these eventual relations in wetlands.

Finally, a bogs surface is characterised by micro-topography that adds complexity and affects

the variability of redox potential (Haraguchi, 1992).

Water table

Water table level controls the diffusion of oxygen in peatlands and several studies have found a

good correlation between oxygen concentration and redox potential at low (6-15%) O2 contents

(Callebaut et al., 1982). Almost all the literature examined in the present work involves research

carried out in wetlands other than bogs, highlighting a general lack of studies on bogs. Generally,

redox potential decreases when water table increases so that they are negatively correlated

(Niedermeier and Robinson, 2007; De Mars and Wassen, 1999) and the change of redox

potential trails shortly the drop of water table (Seybold et al., 2002). In a mineral Calcaric

Gleysol soil, Mansfeldt et al. (2003) showed that the principal variable explaining temporal and

spatial variation of redox potential through a vertical profile was water table (r = - 0.97). De

Mars and Wassen (1999) studied different peatlands and showed that the temporal variation

explained by water table level was 2/3 of the total variance. The authors suggested that soil

conditions like physical properties (capillarity), pH and nutritional status (quality of organic

matter) might have explained the residual variance. Rezanezhad et al. (2014) set up two columns

filled with homogenised riparian soil: in one column they imposed a fluctuating water table

regime whereas in the other the water table was kept stable. Redox potential was measured at

10 cm and - 30 cm every 60 seconds for 75 days and they clearly showed that the imposed

regime controlled the spatial and temporal distribution of the soil redox potential. The authors

observed also short-term spikes during high water table extending for hundreds of mV and

probably caused by gas transport and heterogeneity of water composition. Interestingly, these

spikes did not show up in the column with stable water table, so that soil water circulation could

have been responsible of their generation.

While the degree to which water table level influences redox potential is depth

dependent, the major part of studies on redox potential in wetland soils did not carry out

measurements with depth resolution (Shoemaker et al., 2013; Fiedler et al., 2007). The oxygen

concentration in soil depends on diffusion rate and consumption rate by microbial activity. The

process of diffusion as described by Ficks law is driven by the concentration of oxygen at

26

interfaces, diffusion path length and diffusion coefficient. Water table influences indirectly redox

potential lowering the rate of diffusion of O2 by a factor of 10000 in respect to the gaseous

phase (Gosselink and Mitsch, 2007). Soil properties, like porosity, are also important to develop

the degree of soil aeration; in fact, a strong capillarity can support saturation above the water

table level (Knorr et al., 2009; Thompson et al., 2009; Thompson et al., 2007). Previous studies

have measured that oxygen penetration length in peat can be very short; for example Armstrong

and Boatman (1967) found that oxygen, even in extreme cases, would not diffuse further than 6

centimetres. Benstead and Lloyd (1995) measured oxygen in different hollows in a Sphagnum

bog, when water table was above and below 2 cm from the surface, and showed that, at 2 cm

depth, almost all hollows lacked O2; only in one site, the authors detected oxygen at depth of 5

centimetres. These studies suggest that water table efficiently prevents oxygen to penetrate in

peat, though local peat properties and microbial activity can still be important to determine the

length of the vertical extinction curve of oxygen. It is worth adding that during summer the water

table drop might not entail oxygen penetration to deeper layers if oxygen in the upper layer has

been consumed because of high microbiological activity (Barber et al., 2004).

As mentioned before, many studies have measured decreasing redox potential with

increasing depth. Fiedler et al. (2004) found decreasing redox potential in a wetland mineral soil,

while Thomas et al. (2009) found a decreasing gradient in Florida Everglades wetlands. The

relative importance of sources of water shapes the vertical redox profile: for instance, lateral

water flowing from uplands may be responsible of changes of Eh (Wheeler and Shaw, 1995

cited in Thompson et al., 2009) and upwelling of groundwater (when occurs) was suggested to

be responsible of local increasing of Eh in deeper layers. Considering redox profile again, when

two or more redox probes were employed, the measurements have shown that spatial and

temporal variability of Eh decreased with depth (Thomas et al., 2009; Mansfeldt et al., 2003;

Fiedler et al., 2004). However, Mansfeldt et al. (2003) found that the highest variation of redox

potential was at 60 cm below the surface, where the soil most frequently changed between

saturated and unsaturated condition. Therefore, standard deviation is associated to the extent of

water table fluctuation, because it affects the relative proportion of air and water in the pores. In

conclusion, the effects of water table on redox potential should be limited to the zone of water

table fluctuation, or at most, should extend very shortly above or below the water level.

Contrarily, below that zone, other factors should play a major role.

Precipitation

27

Rainfall is a source of dissolved oxygen and oxidised nutrients, like nitrate and sulphate, which

may temporally increase redox potential in surface peat layers. Rainfall also raises water table

and if rainfall rate exceeds storage capacity, runoff water is produced. The effect of rainfall on

oxidation-reduction cycles may be important in peatland and can be greater during dry periods,

because rain can reach deeper anoxic layers (Deppe et al. 2009). Niedermeier and Robinson

(2006) measured redox potential in a fen during two summer rain events, when water table was

below 20 cm. The authors observed a sharp increase of Eh at 10 cm depth during the first rain

(45 mm d-1) and a broad peak in Eh at 10 cm during the second rain (70 mm d-1). The peaks were

about 400-500 mV. There were no effects at 30 cm depth, below the water table. Precipitation

also raised water table lowering the rate of oxygen diffusion and counteracting the former

oxidising effect of rainwater (Niedermeier and Robinson, 2006). The influence of rain is not

always outstanding, in fact Mitchell et al. (2005) has observed just little Eh increment during

rainfall events in the surface of a peatland, and no influence at layers deeper than 15 centimetres.

Contrarily, in the first 50 cm of peat, Haavisto (1974) has measured an averaged decrease of 47

mV, although the procedure used may have produced unreliable values. Fiedler (1999) observed

that at depths > 30 cm, precipitation influenced the potentials only indirectly by raising the water

table.

Water chemistry

Peatland bog soils are electron acceptors limited, but have plenty of electron donors (i.e. organic

matter). Addition of easy degradable organic matter can enhance the electron pressure and can

reduce redox potential while the input of oxidised species can increase it, as explained by the

Nernsts equation (Reddy and DeLaune, 2008). The redox potential measured at the electrode is

characterised by the dominant redox couple, and it depends on standard rate constant for the

redox couple, concentration of oxidised and reduced species, number of electrons transferred per

molecule and electrode surface (Peiffer et al, 1992). Some redox couples can be present in very

low concentration in oligotrophic bog water. For example, Fe(III)/Fe(II) has a high standard rate

constant, so that is often dominant in mineral soils, but total dissolved Fe is very low in bogs

(Wieder and Vitt, 2006) and the major part is bound to and stabilised by DOC (Steinmann and

Shotyk, 1996), so that its reaction may be insignificant (Keller and Bridgham, 2007). The same

can be said for manganese. Reduced inorganic sulphur (RIS) is much less abundant than

organically bound S in bogs, but it is reported to have a very dynamic turnover (Wieder and

Lang, 1988). The concentration of RIS is higher than iron and it can be more important in a bog,

explaining up to 30 % of anaerobic respiration (Keller and Bridgham, 2007). DOC is not only an

28

electron donor but it can also act as electron acceptor for anaerobic respiration. There is an

increasing need of assessing the use of organic electron acceptors by microorganisms in anoxic

environment. Many studies hypothesised their involvement to account for the great fraction of

unexplained carbon mineralization in low nutrient bogs (Knorr et al., 2009; Deppe et al., 2009).

However, the great importance of DOC has been evaluated only in laboratory and, so far, it has

not been studied directly in the field (Blodau et al., 2009). To complicate the issue, the DOC

chemical composition rather than its amount distinguishes its reducing/oxidising role in

peatlands (Keller et al., 2009).

Microbial community

Microorganisms control oxidation-reduction reactions in soils, modifying their environment by

consuming TEAs and lowering redox potential, which, at the same time, determines the

functional microbial type emergent in the bacterial community (Husson, 2013). Despite the low

nutrient content and pH and the recalcitrance of Sphagnum (Dobrovolskaya, 2014), microbial

community in bogs is frequently reported to be as active as other richer peatlands fens

(Mettrop et al., 2014; Fisk et al., 2003). In bogs, microbial activity is greater in surface layers (0-

15 cm below surface) than subsurface ones (15-30 cm below surface) (Fisk et al. 2003). In

general, the physical-chemical and environmental factors that influence microbial activity will

cause the redox potential to change. Studies that took into account the role of microorganisms in

driving Eh changes in bogs are lacking, albeit an important role for them, as observed in other

ecosystems, might be hypothesized. In a tidal marsh, Catallo et al. (1999) showed that microbial

activity alone produced quick Eh variation to the extent of more than 370 mV in less than 48

hours. Oligotrophy of bog ecosystems represents a constraint to microbial activity and input of

nutrients from deposition or other sources can alter this constraint. Bragazza et al. (2012) have

shown that N deposition could alter microbial communities and favour bacterial growth.

Temperature can influence microbial community composition, growth rate, enzyme synthesis

and response. As a rule of thumb, the decomposition rate of organic matter doubles every 10C

increment, following the Arrhenius equation. However, other environmental constraints affect

the rate of decomposition. These physical and chemical constraints are themselves affected by

temperature and climatic factor like flooding, droughts and freezing, which have the effect to

making the relation complex (Davidson and Janssens, 2006) and giving different Q10 for

different ecosystems (Peng et al., 2009). Nevertheless, regardless of the extent of their effect,

warmer conditions increase microbial activity both in oxic and anoxic environments (Estop-

aragons and Blondau, 2012). Acidity affects the microbial community. Acidic pH means a

29

greater amount of energy spent by bacteria pumping hydrogen ions out of their cells to survive.

Enwall et al. (2007), showed a negative correlation between soil pH and the microbial metabolic

quotient, which indicate a decreased efficiency to convert organic carbon in microbial biomass in

acidic soils. Notwithstanding these constraints, N-mineralisation and P-mineralisation can be

higher in bogs than fens (Mettrop et al. 2014; Verhoeven et al. 1990 cited in Wieder and Vitt,

2006). It is possible that bacterial groups in bogs are not constrained by pH because they evolved

for that specific acid condition (Dobrovolskaya, 2014). Moreover, it has to be kept in mind that

fungi are less affected by low pH than bacteria.

Vegetation

Plants have the potential to change redox stratification in soil through at least three distinct

processes: (1) release of O2 through roots into the rhizosphere, (2) primary production,

increasing the quantity of labile organic carbon and releasing roots exudates and (3) direct

nutrients uptake into roots, rhizomes, stems and leaves. The plant species can account for

differences in electron acceptor renewal in anaerobic soils. It is reported that in bogs this process

is usually less important than in fens because of lack of sedges (Deppe et al., 2009). During the

day, photosynthetic activity and active transport of oxygen create an oxic environment in the

rhizosphere. Nikolaustz et al. (2008) observed that, in an artificial Juncus effusus wetland,

reducing condition during night (-320 mV) and oxic condition during day (+300 mV) were a

function of light intensity and dissolved oxygen in the rhizosphere. Instead, microbial

consumption of electron acceptors could explain the reducing condition measured at night. The

authors suggested that this pattern might be applicable to all aerenchymatous plants in wetlands.

Respiration of roots produces a dynamic trend of carbon dioxide in soil. Benstead and Lloyd

(1996) incubated a solid peat core extracted from a Sphagnum-Eriophorum bog and observed a

diurnal fluctuation of CO2 at the depth of 15 cm and 20 cm. They found a minimum at 18:00 and

a maximum at 7:00. The authors placed the same peat in dark and fluctuations disappeared,

suggesting a key role of vegetation in the process. Shoemaker et al. (2012) build a mesocosmos

with Myriophyllum acquaticum and Leersia oryzoides and found that an increase of temperature

occurred almost simultaneously with a rise of redox potential leading to daily fluctuation of

about 100 mV in the top 10 cm. The authors did not find the same pattern in a non-vegetated

mesocosm, which suggested that the observed fluctuations were produced by vegetation roots

and not by temperature.

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Peat structure

Bulk density is measured as the dry weight of peat per unit volume (g cm-3). Bulk density

increases with degree of humification and increases with depth. Porosity is the fraction of soil

volume that is not filled by the solid phase. In Sphagnum peat near the surface porosity is very

high, ranging between 88 and 97 % (Ivanov, 1981). Porosity decreases with increasing

humification, bulk density and degree of compaction, so it decreases with depth. Porosity

correlates positively to hydraulic conductivity (Schaaf, 1999). Porosity can decrease during

water table drawdown (Schlotzhauer and Price, 1999). Peat of different origins can vary in

porosity and in bulk density influencing the capillarity and the air-filled porosity of the peatsoil.

Sphagnum peat may have low capillarity: indeed, Deppe et al. (2009) showed that air filled

porosity sharply increases at 1 cm above the water table in a bog peatsoil. Fen peat with

dominant sedges instead is denser than Sphagnum peat and has higher capillarity, which can

bring to a disconnection between oxic/anoxic boundary and water table dynamic (Knorr et al.,

2009).

Peat chemistry

Peat is composed of an enormous mixture of organic compounds, including carbohydrates

(cellulose, hemicellulose), nitrogenous compounds (proteins, amino acids), phenolics (including

lignin), lipids (waxes, resins, steroids, terpenes) and humic substances (Rydin and Jeglum,

2013). The amount of organic matter, C quality and nutrient availability influence redox

potential through microbial activity. Low input of C, lower nutrient availability and less labile

organic matter can slow the redox pathway catalysed by microbes and the release of electrons,

diminishing the electron pressure and resulting in higher Eh (De Mars and Wassen, 1999). The

acid soluble components (e.g. cellulose) tend to be metabolised by microorganisms more readily

than acid insoluble components (e.g. lignin), so that the relative fraction can be used as index of

degradability. The C:N ratio is an important indices of degradability or recalcitrance, so that

where it is high the decomposition is low (Biester et al., 2014). The relative C:N and C fraction

in litter depends on vegetation type. In Sphagnum bogs, where vascular plants are presents there

is a higher input of lignin (Biester et al. 2014). C:N ratio also varies with plant species, for

example in Sphagnum spp. it is generally around 50-60 but it can reach value of 300 in S.fuscum

litter (Rydin and Jeglum, 2013). Sphagnum tissue is recalcitrant not only because it has high C:N

ratio, but also because it has high content of sphagnum acid in the cell walls (Rudolph and

Sampland, 1985) and other phenolic compounds that act as anti-septic (Dobrovolskaya, 2014).

Phenolic compounds are degraded by phenolic oxidase only in presence of oxygen. Therefore, in

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anoxic conditions phenolic compounds can inhibit the hydrolases, the main enzyme involved in

peat decomposition (Freeman et al., 2001). New fresh organic matter is added into the acrotelm

from the layer of living Sphagnum spp. and from vascular plants. Different vegetation grows on

hummocks and hollows and they have also different HWT, so that litter chemical precursors and

importance of aerobic respiration differ between the two micro-structures, influencing the quality

and quantity of organic matter that is added to the acrotelm. In the hollows, organic matter is

added directly to the zone of water table fluctuation, while in the hummock it has to pass through

the aerobic body where a great part of the labile carbon is decomposed to CO2. Therefore, by the

time that litter has reached the water fluctuation zone, the amount of labile organic matter added

would be lower below hummocks (Nilsson and quist, 2009). Moreover, the hummocks

S.fuscum and vascular plants (e.g. Calluna vulgaris) should produce litter with higher lignin

fraction and higher C:N ratio than hollows, increasing its recalcitrance. Since the effect of water

table fluctuation on decomposition depends on the intrinsic liability of that zone to sustain

decomposition in terms of nutrients, C quality and microbial community (Artz, 2009), also redox

potential will be affected by these factors by showing different sensitivity to water table

fluctuation. Ombrotrophic plants evolved mechanisms for dealing with a nutrient-deficient

environment. Sphagnum spp. and Eriophorum spp. show strong resorption and recycling of

nutrients during tissue senescence (Wang et al., 2014; Rydin and Jeglum, 2013; Bragazza et al.,

2003), resulting in a nutrient-depleted litter. Microbial activity and decomposition rate will be

controlled by the limiting nutrient, however it is also true that a greater amount of nutrients

sustains a higher microbial activity and decomposition. Nutrients exert an indirect effect on Eh

through microbial activity. Thomas et al. (2009) found that the phosphorous gradient in the

Everglades influenced how sensitive the redox potential was to changing water table only in the

surface layer. Thomas et al. (2009) found that the higher the P content was the lower was the Eh.

So far, to the writers knowledge, there have not been studies that addressed in detail the

difference in nutrients between micro-topographic structures. Regarding the pore-water

chemistry, there is some interesting data in the work of Bragazza et al. (2005), who studied

Ryggmossen bog, and found higher K+ and slightly higher orthophosphate, nitrate and ammonia

in hummocks. However, the authors sampled water using wells rather than rhizons. Other studies

found higher nitrate in hummocks (Frei et al., 2012; Wolf et al., 2011). Bragazza et al. (2005)

also analysed peat samples and found that total P was two times higher in hummocks, total

nitrogen was lower and total K was higher. Along the burial process of peat within Sphagnum

bogs, the variation of litter fraction and C:N depends more on decomposition pattern (climate,

hydrology) than plant species (Biester et al., 2014). In the acrotelm the C:N ratio usually

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decreases with depth (Wang et al., 2014), while in more decomposed and deeper layer it

increases (Silamikele et al., 2007). As decomposition proceeds the organic matter becomes

enriched in lignin and N, with lignin suppressing decomposition rate and raised N-level

suppressing lignin degradation (Berg and Meentemeyer, 2002). As mentioned before, the

absence of signicant external P input and relatively elevated atmospheric N deposition in

ombrotrophic systems requires efcient internal cycling of P so that the usual increase of C:P

and N:P ratio with depth might be due to preference of P uptake (Wang et al., 2014). Another

factor of recalcitrance varies with depth, Beer et al. (2008) observed in a bog that aromatic and

phenolic functional groups of organic matter increased with depth, suggesting that degradability

decreased accordingly. The humic fraction of peat is important in affecting redox potential and

electron transfer to humic substances in anoxic systems is considered to competitively suppress

reduction of other terminal electron acceptors (TEAs), including CO2 under methanogenic

conditions (Klpfel et al., 2014). Gondar et al. (2005) measured that the extractable fraction of

humic acid (HA) and fulvic acid (FA) was 10 % of the TOM in the fibric layer and 2 % in the

sapric. Peat has high cations exchange capacity due to humified organic matters (Rydin and

Jeglum, 2013). Most cations can be bound to organic matter diminishing the availability to

microorganisms. For example, Fe(III) can be stabilised by complexation to DOC and can still be

found in anoxic and reducing environment (Steinmann and Shotyk, 1996).

Temperature

Nernsts equation describes how temperature directly affects redox potential, but that the direct

effect of temperature may not be relevant (Shoemaker et al., 2012). Temperature exerts a more

important indirect effect triggering microbial activity (Fiedler et al., 2007). A more subtle

indirect effect occurs through the influence on ionic activity, though in this case a shift from 3 to

22 C produces only a shift of around 25-30 mV (de Mars and Wissen, 1999). It is possible that

temperature effects are more important in explaining variance of Eh in anaerobic layer than in

surface layer because the lack of oxygen in the former does not influence redox potential.

Urquhart et al. (1972) studying four different bogs found significant negative relation between

Eh and temperature at the depth of 30 cm. The authors argued that it was a consequence of

microbial activity. The higher sensibility to temperature of decomposition rate of deep

recalcitrant peat than surface peat (Hilasvuori et al., 2013) could corroborate the importance of

temperature effect on redox potential in anaerobic peat. The heat wave varies in different soil

ecosystem, and, for example, Barber et al (2004) did not find seasonal variation (i.e.

temperature) of redox at depth of 40 cm in a wet grassland.

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pH

Bogs are acid environments and the main source of H+ stems from humic substances and

Sphagnum phenolics. In surface layers, humic substances mainly supply hydrogen ions while at

deeper layers carbonic acid becomes the main source of hydrogen ions (Steinmann and Shotyk,

1996). Humic acids have a more acidic dissociation constant than carbonic acid and for this

reason pH is reported to increase with depth in bogs (Deppe et al., 2009; Steinmann and Shotyk,

1996; Lundin and Berquist, 1990). The Eh responds inversely to change in pH according to the