POSIVA OY

FI-27160 OLKILUOTO, FINLAND

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Mervi Söderlund

Jukka Lehto

June 2012

Working Report 2012-38

Sorption of Molybdenum,Niobium and Selenium in Soils

June 2012

Working Reports contain information on work in progress

or pending completion.

The conclusions and viewpoints presented in the report

are those of author(s) and do not necessarily

coincide with those of Posiva.

Mervi Söderlund

Jukka Lehto

University of Helsinki

Department of Chemistry

Laboratory of Radiochemistry

Working Report 2012-38

Sorption of Molybdenum,Niobium and Selenium in Soils

Base maps: ©National Land Survey, permission 41/MML/12

ABSTRACT An evaluation of the behaviour of radioactive nuclear waste nuclides in the biosphere is included in the safety assessment of spent nuclear fuel final disposal. One part of this estimate is the evaluation of the sorption of molybdenum, niobium and selenium in soils. This literature survey describes the chemistry of these radionuclides and the factors affecting their sorption in soils.

Mo-93, Nb-94 and Se-79 are long-lived activation products formed from the stable isotopes present in the nuclear fuel, structural components and construction materials used in power reactors. Se-79 is also a fission product. Knowledge of their behaviour in the environment is of great interest because of their long physical half-lives and anionic nature.

The migration and sorption of radionuclides in soils is affected by parameters specific to the element and to the soil. Chemical form, speciation, is the most important elemental factor affecting the sorption and migration properties of the element. Soil redox potential, pH and complex forming ligands are features that have great influence on the speciation. Micro-organisms present in soils can affect the speciation of radionuclides indirectly by changing the prevailing Eh-pH conditions. They can also serve as sorbents. The organic matter content and mineral properties of soils have a noticeable influence on the retention of radionuclides. The sorption of anionic radionuclides such as molybdate, selenite and selenate is pronounced in the presence of organic matter and iron and aluminium oxyhydroxides.

This literature review covers certain elemental and soil specific parameters affecting the sorption of radioactive molybdenum, niobium and selenium in soils. The effect of speciation, soil micro-organisms, organic matter and mineral properties are considered.

Keywords: Sorption, molybdenum, niobium, selenium, soil, speciation, Olkiluoto.

Molybdeenin, niobiumin ja seleenin sorptio maaperään

TIIVISTELMÄ

Käytetyn ydinpolttoaineen loppusijoituksen turvallisuusanalyysiin sisältyy arvio ydin-polttoaineen sisältämien jätenuklidien käyttäytymisestä biosfäärissä. Yksi osa tätä arviota on radioaktiivisen molybdeenin, niobiumin ja seleenin pidättyminen maaperään.

Mo-93, Nb-94 ja Se-79 ovat ydinreaktoreissa käytettyjen rakennekomponenttien ja rakennemateriaalien sekä polttoaineen sisältämien stabiilien isotooppien aktivoitumis-tuotteita. Se-79 on myös fissiotuote. Molybdeenin, niobiumin ja seleenin ympäristö-käyttäytymisen tunteminen on hyvin tärkeää johtuen niiden pitkistä fysikaalisista puoliintumisajoista, anionisesta luonteesta ja suhteellisen suuresta liikkuvuudesta.

Radionuklidien kulkeutumiseen ja pidättymiseen maaperässä vaikuttavat sekä alku-aineelle että maaperälle ominaiset erityispiirteet. Alkuaineen kemiallinen muoto, spesi-aatio, on tärkein alkuaineen pidättymiseen ja kulkeutumiseen vaikuttava alkuaine-kohtainen tekijä. Maaperän hapetus- pelkistysolosuhteet, pH ja komplekseja muodos-tavat ligandit vaikuttavat spesiaatioon. Maaperän mikro-organismit voivat vaikuttaa radionuklidien spesiaatioon epäsuorasti muuttamalla maaperässä vallitsevia hapetus-pelkistysolosuhteita ja pH:ta. Ne voivat myös pidättää radionuklideja. Maaperän orgaanisen aineksen pitoisuudella ja mineraaliaineksen ominaisuuksilla on huomattava vaikutus radionuklidien pidättymiseen. Anionisten radionuklidien, kuten molybdaatin, seleniitin ja selenaatin pidättyminen on korostunutta kun orgaanisen aineksen ja rauta ja alumiini oksohydroksidien pitoisuus on korkea.

Tämä kirjallisuuskatsaus kattaa radioaktiivisen molybdeenin, niobiumin ja seleenin maaperäpidättymiseen vaikuttavat tietyt alkuaine- ja maaperäkohtaiset tekijät. Spesiaation, maaperän mikro-organismien, orgaanisen aineksen ja mineraaliaineksen vaikutus on huomioitu.

Avainsanat: Pidättyminen, (ad)sorptio, molybdeeni, niobium, seleeni, maaperä, spesiaatio, Olkiluoto.

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TABLE OF CONTENTS

ABSTRACT

TIIVISTELMÄ

1 INTRODUCTION ........................................................................................................ 3

2 TYPICAL SOIL PROPERTIES ON OLKILUOTO ISLAND ......................................... 5

2.1 Olkiluoto soils ........................................................................................................ 7

3 SOIL TYPES AND PROFILES ................................................................................ 11

3.1 Grain size classes and textural classes .............................................................. 12

3.2 Soil layers and horizons ...................................................................................... 13

3.3 Soils .................................................................................................................... 13

4 FACTORS AFFECTING THE BEHAVIOUR AND SPECIATION OF RADIONUCLIDES IN SOILS ................................................................................... 17

4.1 Retardation processes ........................................................................................ 17

4.2 Soil solution ........................................................................................................ 18

4.2.1 pH of the soil solution .................................................................................. 19

4.2.2 Dissolved organic carbon in the soil solution .............................................. 19

4.2.3 Cations in the soil solution ........................................................................... 20

4.2.4 Anions in the soil solution ............................................................................ 22

4.2.5 Ionic strength of the soil solution ................................................................. 23

4.3 Redox conditions ................................................................................................. 24

4.4 Sorption ............................................................................................................... 25

4.5 Microbial effects .................................................................................................. 26

5 MOLYBDENUM ........................................................................................................ 29

5.1 The influence of soil redox potential and pH ...................................................... 30

5.2 The influence of micro-organisms ...................................................................... 31

5.3 Sorption on organic matter and organic soils ..................................................... 32

5.4 Sorption on mineral soils .................................................................................... 33

5.4.1 Podzols ......................................................................................................... 36

5.5 Sorption on minerals .......................................................................................... 37

6 NIOBIUM .................................................................................................................. 45

6.1 The influence of soil redox potential and pH ...................................................... 46

6.2 Sorption on soils ................................................................................................. 48

7 SELENIUM ............................................................................................................... 55

7.1 The influence of soil redox potential and pH ...................................................... 56

7.1.1 The speciation of selenium .......................................................................... 56

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7.1.2 Sorption behaviour of selenite and selenate ............................................... 59

7.2 The influence of micro-organisms ...................................................................... 60

7.3 Sorption on organic matter and organic soils ..................................................... 61

7.4 Sorption on mineral soils .................................................................................... 62

7.4.1 Selenium in Finnish and Swedish soils ....................................................... 63

7.5 Sorption on minerals .......................................................................................... 65

7.5.1 Iron and alumium oxides ............................................................................. 65

7.5.2 Pyrite and chalcopyrite ................................................................................ 68

7.5.3 Other minerals ............................................................................................. 68

7.6 Volatilisation of selenium ................................................................................... 69

8 CONCLUSIONS ...................................................................................................... 77

REFERENCES ............................................................................................................. 81

APPENDIX 1 ................................................................................................................. 97

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1 INTRODUCTION

A final disposal repository for spent fuel from the Finnish nuclear power reactors in Olkiluoto and Loviisa is to be built in the bedrock at the Olkiluoto site at the depth of approximately 400 m. The final disposal plan includes a safety assessment of the spent nuclear fuel, where the potential dose contributing nuclear waste nuclides for man are specified. As a part of this assessment, the behaviour of these prioritized radionuclides in the biosphere is evaluated.

The biosphere behaviour evaluation of the radionuclides includes aspects concerning the migration of radionuclides in soils, sorption on soil particles and transfer from soil to plants. Distribution coefficient, Kd, is used as a measure to describe the fraction of the radionuclide adsorbed on soils, sediments and suspended material. Kd is defined as the equilibrium ratio of the radionuclide concentration in the solid phase compared with that in the liquid phase (Kd = Csolid/Cliquid). Kd is used to describe the retention and mobility of radionuclides in the environment: with low Kd values, the fraction of radionuclide sorbed in the solid phase is low, and the fraction in the liquid phase is high. This leads to low retention and high potential mobility in the overburden. On the other hand, as the Kd value becomes bigger the fraction of radionuclide sorbed on the solid phase increases as the fraction in the liquid phase decreases. This leads to higher retention on soil and lower potential mobility. With known Kd values and flow velocity of groundwater, the migration rate of radionuclides in soils can be predicted.

The mineralogical and chemical characteristics of the soil layers determine the sorption of radionuclides in the overburden. Such factors are clay minerals and clay mineral content, particle size distribution, organic matter content, redox potential and pH. Radionuclides which have high tendency of forming anionic species in aqueous solutions, such as molybdenum, niobium and selenium, are poorly sorbed on soil mineral constituents, thus leading to high potential mobility in the overburden. The sink for anionic radionuclides, such as molybdenum and selenium, is thought to be organic matter.

Earlier, Söderlund et al. (2011) reviewed literature on iodine, chlorine, technetium and cesium. The present review focuses on molybdenum, niobium and selenium. Due to their anionic nature and high potential mobility in the environment, Mo-93, Nb-94 and Se-79 might be possible radiation dose sources for humans in a long run after the closure of the final disposal repository. The physical half-lives of Mo-93, Nb-94 and Se-79 are 4000, 2.03x104 and 3.77x105 years, respectively. In the long-term safety assessment of spent nuclear fuel, Mo-93, Nb-93m and Nb-94 are considered as level II (high) priority radionuclides, whereas Se-79 is classified as level III (medium) priority radionuclide (Hjerpe et al. 2010). The classification of radionuclides is done according to their expected relevance to the long-term safety of spent nuclear fuel and proportion of the radiation dose posed to humans in the model calculations. Five priority groups are separated ranging from top priority (level I) to low priority (level V) nuclides. The priority of the radionuclides decreases with decreasing level number. Table 1 summarizes the radionuclides included in the priority groups I-IV.

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Table 1. The priority grouping of radionuclides based on their safety relevance in the long-term safety of spent nuclear fuel (Haapanen et al. 2009; Hjerpe et al. 2010).

Priority I (top) Priority II (high) Priority III (medium)

Priority IV (non-immediate)

C-14 Cl-36 I-129

Mo-93 Nb-93m Nb-94 Cs-135

Ni-59 Se-79 Sr-90 Y-90

Pd-107 Sn-126 Sb-126

The aim of this literature study was to gather information on: (a) the chemistry of molybdenum, niobium and selenium in the environment and (b) their sorption on soil. Information was collected especially for soils and soil types currently present in Olkiluoto Island, and for soil types expected to be formed in the future from present soils. The literature reviewed covered research done in several countries and in different soil regimes. Further information concerning the partition of molybdenum and selenium between different fractions (exchangeable, iron and manganese oxides, carbonates, organic matter) and distribution profile in Olkiluoto soil was gained from three excavator pits OL-KK14, OL-KK15 and OL-KK16.

The behaviour of molybdenum, niobium and selenium in soil types represented presently and in the future on Olkiluoto Island can be assumed to be similar to the behaviour of these elements described in the literature for similar soils. Values of distribution coefficients gathered in this survey for molybdenum, niobium and selenium can be used in calculating and predicting the mobility and migration rates of Mo-93m, Nb-94 and Se-79 in the overburden.

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2 TYPICAL SOIL PROPERTIES ON OLKILUOTO ISLAND

Olkiluoto area was covered with continental ice sheet during the last glacial period (Eronen and Lehtinen 1996). The melting of the ice sheet induced the formation of the four main stages of the Baltic Sea during Late Weichselian and Holocene (Figure 1), namely the Baltic Ice Lake (13 500-10 300 BP), the Yoldia Sea (10 300-9 500 BP), the Ancylus Lake (9 500-7 400 BP) and the Litorina Sea (7 400- 4 000 BP) (Agrell 1976; Björck 1995; Eronen 2005).

Figure 1. The four main phases of the Baltic Sea during Late Weichselian and Holocene (Eronen 2005). During the Baltic Ice Lake-phase Olkiluoto area was covered by approximately 1500 m thick continental ice cover (Fjeldskaar 2000). The area was deglaciated at the end of the Yoldia phase only to be covered by a 200-metre water table (Haapanen et al. 2007b; Haapanen et al. 2009; Tikkanen 1981) In the course of the last freshwater stage, the Ancylus Lake, Olkiluoto area was approximately 100 metres under the sea-level

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(Tikkanen 1981). Around 7400 BP, salty sea waters penetrated into the Baltic basin and the sea environment changed from freshwater to brackish. In the early stages of the Litorina Sea Olkiluoto area was still submerged (Tikkanen 1981), but the ongoing isostatic land uplift caused the area to emerge approximately 2500- 3000 years ago (Haapanen et al. 2007b; Haapanen et al. 2009; Mäkiaho 2005).

Sediment cover on Olkiluoto Island was formed during the last glaciation or after that, and is classified as Quaternary deposits (Huhta 2005, 2007, 2009; Lahdenperä et al. 2005; Lahdenperä 2009; Lintinen et al. 2003; Lintinen and Kahelin 2003). At present the representative elevation on the Olkiluoto Island is +5 meters above the sea level (Lahdenperä et al. 2005), and the average land uplift is 6 mm/y (Eronen et al. 1995). The highest points at Olkiluoto Island are Liiklankallio (+18 m), Selkänummenharju (+13 m) and Ulkopäänniemi (+12 m), otherwise the surface being rather flat (Lahdenperä et al. 2005).

New areas are emerged above the sea level due to the ongoing isostatic land uplift, thus increasing the surface area of Olkiluoto Island constantly (Ekman and Mäkinen, 1996; Mäkiaho 2005). The shoreline is in continuous change as former seabed is exposed, and depending on the type of the sediment, different kind of primary successional terrestrial vegetation species occupy the area (Haapanen 2007b). The successional stages for ground vegetation in Olkiluoto Island are presented in a schematic diagram in Figure 2. Usually, the freshly exposed areas are firstly covered by meadows, which are transformed into alder, pine or spruce forests along the successional line. The soil forming processes begin acting immediately after the emerging of the former seabed sediments.

Bedrock is uncovered on Olkiluoto Island in some places, and stones and boulders can be numerous (Haapanen 2007b; Miettinen and Haapanen 2002; Tamminen et al. 2007). In the treeless areas rocks are covered with lichen and dwarf shrub. Liiklankallio and Selkänummenharju are the largest rock forest found on the Olkiluoto Island.

Olkiluoto Island is covered mostly by forests, which are typically classified as mesic upland forests or herb-rich forests (Miettinen and Haapanen 2002). Typical forest types on Olkiluoto Island are spruce-dominated, mixed deciduous and coniferous forest. Characteristic tree species are Scots pine, Norway spruce, birches and black alder. Only a few mires are found on the Island. Reeds, low and high meadows are commonly found at the shoreline of Olkiluoto Island (Haapanen 2007b; Miettinen and Haapanen 2002). Common ground vegetation species found on Olkiluoto Island include e.g. twig plants blueberry and lingonberry, different kind of mosses, spreading woodfern, meadowsweet, wood millet and wavy hair-grass (Tamminen et al. 2007). Bilberry and lingonberry were the prevailing evergreen ground species found.

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Figure 2. A diagram of the successional stages for ground vegetation taking place in the Olkiluoto area (Haapanen et al. 2009).

2.1 Olkiluoto soils

Soils at Olkiluoto are weakly developed due to the climate and the short time span of less than 3000 years for the slow soil forming processes to take place (Eronen and Lehtinen 1996; Jauhiainen 1973; Starr 1991; Tamminen et al. 2007). Undeveloped Arenosols and Regosols, thin or coarse-grained Leptosols and groundwater soils Gleysols are the typical soil types on Olkiluoto Island (Tamminen et al. 2007). Haplic Arenosol is the soil type found usually. The predominant soil forming process in Finland is podsolization (Eronen and Lehtinen 1996; Starr 1991). It is a slow process, as the formation of a mature podzol profile can take up to 500-1500 years (Starr 1991). Characteristic soil horizons for podzols are not met in the weakly developed Olkiluoto soils (e.g. Lintinen and Kahelin 2003; Lusa et al. 2009). The podzolization processes and the formation of the soil layers characteristic of a podzol profile are discussed elsewhere (Söderlund et al. 2011).

Several investigations concerning the soil properties on Olkiluoto Island have been carried out (e.g. Huhta 2005; 2007; 2008; 2009; 2010; Lintinen and Kahelin 2003; Lintinen et al. 2003; Lahdenperä et al. 2009; Lusa et al. 2009). The most common soil types reported on Olkiluoto Island are fine-textured till (53 %), sandy till (39 %), gravelly till (4 %), peat (3.4 %) and outcrops (0.6 %) (Rautio et al. 2004).

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In spite of the irregular bedrock surface, the surface of the overburden is quite smooth even in places where changes in bedrock topography are abrupt. This was caused by the last glaciation activity: bedrock depressions were filled with a thicker layer of overburden, mainly sandy till and fine-grained till, and the bedrock highpoints, which were not buried by the late glacial deposition, protrude through the modest soil layers. The thickness of the till cover is commonly 2-4 m and it is rich in clay fractions in some sites (Lintinen et al. 2003; Lintinen and Kahelin 2003; Lahdenperä et al. 2005; Lahdenperä 2009).

The Olkiluoto bedrock groundwater system is transport-limited and the overburden supply-limited based on the recharge computations carried out with the surface hydrology model. This indicates that there is more supply from the overburden to the bedrock than the bedrock system can transmit, which is due to the lower hydraulic conductivity of the bedrock compared with the overburden soils. The supply-limited overburden groundwater system implies that more precipitation would result in greater runoff and evapotranspiration (Karvonen 2008).

The proportion of fine (<0.063 mm) and clay (<0.002 mm) fractions of the till vary between 32-41 % and 7-14 %, respectively (Lintinen and Kahelin 2003; Lintinen et al. 2003; Lahdenperä 2009). The typical minerals in order of abundance for gravel, sand and silt fractions (Ø < 2 mm) are quartz, plagioclase, potassium- feldspar, mica, chlorite and hornblende (Lintinen et al. 2003; Lusa et al. 2009). Illite, hornblende and chlorite are the typical minerals in the clay fraction (Ø < 0.002 mm) (Lintinen et al. 2003; Lusa et al. 2009). Specific surface area of soil particles varies between 2.3-15 m2/g and density between 2.8-3.0 g/cm3 (Lintinen and Kahelin 2003; Lintinen et al. 2003).

Lusa et al. (2009) studied the dry matter content, the proportion of organic matter, grain size distribution and mineralogy of the soil samples taken from three experimental excavator pits on Olkiluoto Island, namely OL-KK14, OL-KK15 and OL-KK16. pH was determined from the soil samples and from soil solution, from which also dissolved organic carbon (DOC), main cations and anions were specified. The mineral soil layers MS1-MS3 in the excavator pit OL-KK14 were classified as sandy till, whereas MS4 was composed of clay. The mineral soil layers in excavator pit OL-KK16 consisted of fine to coarse sand and sandy till. The composition of excavator pit OL-KK15 was somewhat more heterogeneous; soil layers were classified as clay, fine sand, sand, fine-grained silty/clayish till and coarse-grained sandy till. In Figure 3 the locations of the excavator pits OL-KK14, OL-KK15 and OL-KK16 on Olkiluoto Island can be seen.

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Figure 3. The locations of the excavator pits OL-KK14, OL-KK15 and OL-KK16 on the Olkiluoto Island (Lusa et al. 2009). Map layout: Jani Helin, Posiva Oy.

Ca2+, Mg2+, K+, Na+, Al3+ and H+ are the main exchangeable cations in the Olkiluoto soils (Lahdenperä 2009; Lintinen et al. 2003; Lusa et al. 2009). Surface soils are naturally acidic and pH increases with soil depth (Lahdenperä et al. 2005; Lahdenperä 2009; Lusa et al. 2009). pH in the humus layer, surface soils and C layer (unaltered soil material) are typically 3.7-4.4, 4.0-8.0 and 6.5-8.0, respectively (Lahdenperä 2009; Lusa et al. 2009). The average groundwater surface usually lies approximately 2 meters below the ground surface (Lahdenperä et al. 2005; Lahdenperä 2009).

The cation exchange capacity (CEC) of the Olkiluoto soils samples taken from three different excavator pits OL-KK14, OL-KK15 and OL-KK16 (Table 2) (Lusa et al. 2009). The CEC was calculated as the sum of potassium, sodium, magnesium and calcium ions extracted from the soil samples with 1M BaCl2 (exchangeable cations). CECs were determined for the grain sizes <0.063, 0.063-0.125, 0.125-0.25, 0.25-0.5, 0.5-1.0 and 1.0-2.0 mm. CEC was found to decrease with increasing grain size (Lusa et al. 2009). Highest CECs were found in the humus layer, because high concentrations of ion exchange groups, especially –COOH groups, are found in the organic material creating CEC capacities as high as 1500-5000 mmol H+/kg dry weight (Birkeland 1984). In Lusa et al. (2009) study the CEC of the uppermost mineral soil layer MS1 was lower than the CEC of the underlaying mineral soil MS2 in excavator pits OL-KK14 and OL-KK15. The result corresponded to a higher amount of clay sized particles in MS2 layers. In underlaying soil layers the CEC decreased with depth. In excavator pit OL-KK16 the average CECs decreased with depth.

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Table 2. The average cation exchange capacities (CECs, mmol/kg of dry weight) for the soil layers of the excavator pits OL-KK14, OL-KK15 and OL-KK16 (Lusa et al. 2009).

Soil layer

OL-KK14 OL-KK15 OL-KK16 Layer depth (cm)

CEC (mmol/kg)

Layer depth (cm)

CEC (mmol/kg)

Layer depth (cm)

CEC (mmol/kg)

humus 0-5 74.7 0-7 232 0-10 398 MS1 5-20 10.7 7-50 11.0 10-30 17.8 MS2 20-60 26.3 50-80 27.7 30-50 17.7 MS3 60-105 15.7 80-110 17.5 50-110 11.8 MS4 105-240 14.5 110-160 20 110-300 11.0 MS5 - - 160-260 11.0 - - - soil samples representing mineral soil layer MS5 were not taken from the

excavator pit

The organic matter content of the till on Olkiluoto Island is usually less than 1.8 % (Lahdenperä et al. 2005; Lahdenperä 2009; Lintinen and Kahelin 2003; Lusa et al. 2009). The dry matter content in mineral soils is typically 78-99 % (Lahdenperä et al. 2005; Lahdenperä 2009; Lintinen et al. 2003; Lusa et al. 2009). The proportion of organic and dry matter in the humus layer varies between 68-82 % and 22-33 %, respectively (Lahdenperä 2009; Lusa et al. 2009; Roivanen 2006). The thickness of the humus layer varies between 2.3 cm and 21.1 cm for mineral soils, and 6-115 cm for peat layers (Lahdenperä 2009; Lintinen et al. 2003; Lusa et al. 2009; Rautio et al. 2004).

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3 SOIL TYPES AND PROFILES

Soils are separated into mineral and organic soils (Aaltonen et al. 1949). Mineral soils are composed of soil types, which consist solely of mineral material, or where the amount of organic material is very low (Aaltonen et al. 1949). The naming and classification of mineral soils is done after the dominant grain size group or groups. Mineral soils are divided into sorted and unsorted soils according to the relative percentages of the dominant grain size group(s) (Aaltonen et al. 1949; Haavisto-Hyvärinen and Kutvonen 2007). Sorted mineral soils usually consist of one dominant grain size group and contain minor amounts of other grain sizes (Haavisto-Hyvärinen and Kutvonen 2007). Based on this dominant grain group, coarse and fine grained sorted soil types are separated. On the other hand, unsorted mineral soils, tills, contain all grain size groups or at least several approximately same-sized grain groups mixed up together (Haavisto-Hyvärinen and Kutvonen 2007).

The mineral properties of the soil can have a great influence on the sorption of radionuclides. For example, molybdenum and selenium are retained in the enrichment layer of podzolic soils due to the increased concentration of aluminium and iron sesquioxides (Choppin et al. 2009; Dhillon and Dhillon 1999; Gustafsson and Johnsson 1992; Harada and Takahashi 2009; Keskinen et al. 2009; Vuori et al. 1989). The texture and clay fraction content of the soil affects the sorption of selenium, as the sorption is favoured in fine-textured soils compared with coarse-grained soils (Ashworth et al. 2008; Keskinen et al. 2009; Vuori et al. 1989; Yläranta 1983). The retention of molybdenum, niobium and selenium in soil increases with increasing organic matter content (Choppin et al. 2009; Echeverria et al. 2005; Gerzabek et al. 1994; Gustafsson and Johsson 1992; Lang and Kaupenjohann 2000; Pezzarossa et al. 1999; Sheppard and Thibault 1990).

The formation of a soil profile depends on topography, parent material, vegetation cover, temperature and humidity, groundwater height and flow, mineral composition and grain size distribution, topsoil water conductivity and especially time (FAO 2006; Brady and Weil 2002; Haavisto-Hyvärinen and Kutvonen 2007). Also, human activity, such as grazing, trenching, burn-beating, ploughing and soil acidifying has an impact on soil formation. Altogether, percolating water, air and soil micro-organisms change the chemical composition of a regolith and create a distinct soil profile (Koljonen 1992a). The main impact is on the vegetation cover because it is the principal source of organic matter and plays a vital role in the nutrient cycle (FAO 2006). The nature of the formed soil profile depends on the climate and the chemical composition of the regolith, and thus in different climate conditions dissimilar soils profiles are formed (Koljonen 1992a). In Finland typical regoliths found are Gleysols, Regosols, Arenosols, Cambisols, Podzols, Leptosols, Fluvisols and Histosols (Yli-Halla et al. 2000). On Olkiluoto Island, due to the short time span of less than 3000 years, Arenosols, Regosols, Leptosols and Gleysols are the soil profiles found (Tamminen et al. 2007).

Different soil type or soil profiles found on Olkiluoto Island and Finland are specified in this Chapter. These regoliths consist of Arenosol, Cambizol, Gleysol, Histosol, Leptosol, Regosol and Podzol. Also, soil layers and sub-layers included in these soil types are described. The concept of grain size distribution is discussed. More detailed discussion of the aforementioned topics can be found elsewhere (Söderlund et al. 2011).

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3.1 Grain size classes and textural classes

Mineral soil layers are comprised of minerals and soil particles of different sizes (Mälkönen 2003). Grain size distribution analysis is achieved via dry sieving for larger particles (>0.06 mm in diameter), and by wet sedimentation (e.g Sedigraph instrument) for smaller particles. Results from sieving and sedigraph analyses are combined to grain-size distribution graphs.

Sorted mineral soils are divided into coarse grained and fine grained soils depending on the soil texture, i.e. the proportions of different sized particles (Brady and Weil 2002; Haavisto-Hyvärinen and Kutvonen 2007). In Finnish classification of soil textures, boulders, stones, gravel, sand and fine sand are coarse grained soil types, whereas finer fine sand, silt and clay are fine grained soils (RT classification; construction technical) (Haavisto-Hyvärinen and Kutvonen 2007; Koljonen 1992a). Soil textural class is understood as the classification of soil textural units based on the proportions of different sized particles (Brady and Weil 2002). In Finland, textural classes follow the outline to coarse and fine grained soil textures and seven classes can be distinguished; boulder, stone, gravel, sand, fine sand, silt and clay soils (Aaltonen et al. 1949).

Till is the most abundant mineral soil sediment present in Finland, covering approximately 50 % of the land surface (Koljonen 1992a). It contains all grain size groups or several approximately same sized grain groups mixed together, and are typically weakly sorted (Aaltonen et al. 1949; Haavisto-Hyvärinen and Kutvonen 2007). Till deposits are formed from bedrock, preglacial sediments and “in situ” weathered bedrock during the Glacial Epoch by the action of large ice masses (Aaltonen et al. 1949; Koljonen 1992a). Surface soils are usually quite loose, whereas subsoils are cemented and very compact (Aaltonen et al. 1949). The acidity of the tills soils ranges from pH 4.5 to 6, and increases with depth. Due to the low permeability of the tills, the groundwater surface typically lies near the surface (Haavisto-Hyvärinen and Kutvonen 2007).

Till soils are divided into gravelly till, sandy till and fine-grained till (Aaltonen et al. 1949; Koljonen 1992a), and the respective abundances are 10 %, 75 % and 15 % for gravelly till, sandy till and fine-grained till (Koljonen 1992a). Gravelly till is poor in nutrients and is unsuitable for agricultural uses (Aaltonen et al. 1949; Haavisto-Hyvärinen and Kutvonen 2007). Sandy tills are suited for forest soils due to their excellent moisture, air and nutrient content, even though the coarsest grained sandy tills are not appropriate for agricultural uses because of their dry nature. Fine-grained tills comprise a larger group, which includes silty till, very fine silty till and clayey till (Koljonen 1992a). Clayey tills are rare in Finland and this soil type is mostly met in northern Savo, the northernmost point of Eastern Finland.

3.2 Soil layers and horizons

Soils are comprised of different master horizons or layers depending on the soil forming processes that have occurred (FAO 2006; Tamminen and Mälkönen 1999). Altogether, ten master horizons or layers are recognized and these are denoted with capital letters. In Finnish soils, all the major horizons but one are present, namely H, O, A, E, B, C, R,

13

L and W. The one not included is I (ice formations). In the master horizons, sub-layers can be separated and these are symbolized with small letters (Mälkönen 2003). Sub-layers include f, h, s, hs, g, w, p and m, which all are encountered in Finnish soils. The descriptions of the master horizons and sub-layers can be found elsewhere (Söderlund et al. 2011).

Diagnostic horizons are horizons, which are used as a criterion for the classification of soils (Tamminen and Mälkönen 1999). It can be a whole horizon, a part of a horizon or combined of two or more horizons. Certain physical and chemical characteristics have to be fulfilled by the diagnostic horizons. Physical characteristics include, for example, thickness of the soil layer, grain size distribution or texture. Chemical characteristics include, for instance, pH and base saturation. Diagnostic horizons are (Tamminen and Mälkönen 1999):

Albic E horizon has lost silicate clay and iron oxides so that the colour of the horizon comes from the original colour of the soil particles.

Cambic B horizon contains clay in <2 mm fraction over 8 % and the average grain size typically ranges from clay (<0.002 mm) to coarse silt (0.02-0.063 mm). The minimum thickness of the layer is 15 cm and the lower margin is no less than in the depth of 25 cm. Clay content is higher than in C horizon.

Histic H is 20 to 40 cm thick peat layer. Ochric A is mineral soil layer containing organic matter (Ah).

3.3 Soils

Finnish soils are relatively young and originating from the last glaciation (Yli-Halla et al. 2000). The average age of the soils is 10 000 years, but in spite of the young age several signs of soil forming processes can be seen. In Olkiluoto, the average age of the soils is about 3000 years, and soils and poorly developed due to the young age (Tamminen et al. 2007). The typical soil profiles encountered in Finland include Arenosols, Cambisols, Fluvisols, Gleysols, Histosols, Leptosols, Regosols and Podzols (Yli-Halla et al. 2000), from which Arenosols, Gleysols, Leptosols and Regosols have been found on Olkiluoto Island (Tamminen et al. 2007). Below, a short description of the soil types met in Finland and Olkiluoto are given according to the Finnish soil classification system. For more detailed information concerning podzols and for the comparison of Finnish soil types to soil types encountered in Soil Taxonomy, see Söderlund et al. (2011).

Arenosols are poorly evolved and immature soil profiles developed in coarse grained soils, for which the grain size ranges from sand (0.2-0.63 mm) to gravel (2-60 mm) (Tamminen and Mälkönen 1999). Arenosols do not usually have any other diagnostic soil horizons except Ochric A or Albic E. Roughly 20% of the Finnish forest soils are Arenosols (Lilja et al. 2006).

Cambisols (brown earth soil) are evolved in fine grained soils, the grain size ranging from clay (<0.002 mm) to coarse silt (0.02-0.063 mm) (Tamminen and Mälkönen 1999). Cambisols contain Ochric A horizon and under it Cambic B horizon.

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Approximately 6 % of the Finnish forest soils are classified as Cambisols (Lilja et al. 2006).

Gleysols are soils which have gleyic-properties, i.e. a soil layer or horizon in which the effect of standing groundwater (anoxic conditions for most of the year) can be seen (Tamminen and Mälkönen 1999). Gley horizon can be separated from other horizons according to its blue-greyish colour. Gleyhorizon is situated in the depths of 0-50 cm. Gleysol is rarely found in Finland, as about 4% of the Finnish forest soils are Gleysols (Lilja et al. 2006).

Histosols are peat or bog soils, for which the minimum thickness is 40 cm (Brady and Weil 2002; Tamminen and Mälkönen 1999). For Spaghnum peat the thickness must exceed 60 cm. The content of organic matter has to be >20 % (Brady and Weil 2002). In Finland, about 20 % of the forest soils are classified as Histosols (Lilja et al. 2006).

Leptosols are shallow soils, in which either the thickness of the mineral soil is less than 30 cm, or the fraction of fine particles is less than 20 % (very coarse grained soils) (Tamminen and Mälkönen 1999). Around 13 % of the Finnish forest soils are Leptosols (Lilja et al. 2006).

Regosols are poorly developed and immature soils found in fine grained soils, the grain size ranging from clay (<0.002 mm) to coarse silt (0.02-0.063 mm) (Tamminen and Mälkönen 1999). Merely Ofh/H horizon and/or A horizon can be found. Regosols are rarely developed in forest soils; in Finland they cover only 2 % (Lilja et al. 2006).

Podzols are well developed soils in which five different soil horizons can be distinguished; organic (O), eluvial (leaching) (E), illuvial (enrichment) (B), transition zone (BC) and the slightly changed soil horizon (C) (Tamminen and Mälkönen 1999). The most important factors affecting the structure of podzol soils are the susceptibility of minerals to weathering, grain size, climate, topography and runoff conditions. Podzols are usually developed in soils for which the grain size ranges from coarse silt (0.02-0.063 mm) to gravel (2-60 mm) and have relatively good water permeability. Podzols are the main soil type in Finland (Mälkönen 2003), as 35 % of forest soils are classified as podzols (Lilja et al. 2006). Typical soil profile of podzol is presented in Figure 4.

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Figure 4. Podzol soil profile (Yli-Halla et al. 2000; picture by M. Yli-Halla). O = organic horizon, E = eluvial (leaching) horizon, B = illuvial (enrichment) horizon and C = unchanged parent soil.

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17

4 FACTORS AFFECTING THE BEHAVIOUR AND SPECIATION OF RADIONUCLIDES IN SOILS

4.1 Retardation processes

A retardation process can be defined as any process that delays and inhibits the migration of the material in the groundwater flow. In Söderlund et al. (2011) the retardation processes of radionuclides taking place in soils and bedrock are described in more detail. Here, only a summary of sorption processes is given.

Sorption is understood as the retention of the material on a solid surface through physical or chemical interaction. Physical sorption takes place when the substance is bound to the surface of the adsorbent with weak physical van der Waals forces (Atkins and de Paula 2002). At least one layer of water molecules situates between the surface sorption sites and the ions. The range of attracting forces for physical sorption in solution is more extensive than for chemical sorption.

When a substance is bonded on the surface of the adsorbent through a chemical, covalent bond the term chemical sorption is used. Water molecule layers are not situated between the surface functional group and the sorbed ion. Chemical sorption is also known as inner-sphere complexation, specific sorption or ligand exchange. pHpzc (pzc; the point of zero charge) is usually shifted to more acidic pH-values upon inner-sphere complexation (Goldberg et al. 1996).

In ion exchange reactions the exchange between ions in solid and liquid phase takes place. Outer-sphere complexation (electrostatic interaction) is a synonym for ion exchange. Equation 1 gives an example of a simple binary exchange process between

ions present in the solid phase, BzB , and in the solution, AzA :

ABBAz

Bz

A

z

Az

B AzBzBzAz [1]

Equation 1. Binary ion exchange reaction. AzA and BzB represent ions present in the

liquid phase, zA

A and BzB represent ions present in the solid phase, zA is the charge on

ion A and zB is the charge of ion B.

Ion exchange is divided into cation exchange and anion exchange depending on the nature of the exchanging ions (Jedináková-K ížová 1998). Cation exchange takes place when exchangeable ions are cations, and anion exchange, when exchangeable ions are anions. Charged species are bonded through the electrostatic attraction of the oppositely charged exchangeable ion and the surface of the adsorbent. Typical adsorbents for cationic species in soils are Fe, Al, and Mn oxides, clay minerals and other minerals present. The main sorbent for anionic species is soil humus matter, iron and aluminium oxides and few clay minerals (halloysite and imogolite) (Koch-Steindl and Pröhl 2001). In surface adsorption, the precipitate is usually formed by iron (iron hydroxide) or other multivalent metal cation, such as manganese (manganese oxide). The surface charge of the formed solid has a net positive or negative value depending on the prevailing pH and pHpzc of the solid phase. Metal ions, including those of radionuclides are attracted from the solution to balance the surface charge by adsorption.

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4.2 Soil solution

Soil solution is in continuous parallel interaction with soil constituents. The surface chemistry of soil constituents changes as a response to the changes in the chemistry of soil water, and vice versa. pH, dissolved organic carbon (DOC), ionic strength and concentrations of dissolved anions and cations are factors affecting the chemistry of soils and soil waters.

The protonation degree of aluminol (-AlOH), silanol (-SiOH), hydroxyl groups attached to other metals (-MOH) and further type of functional groups e.g. in organic matter are affected by the pH of the soil solution (Essington 2004; Viers et al. 2001). Typically the sorption of cationic radionuclides increases with increasing pH, whereas the sorption of anions is preferred as the pH decreases. The protonation and deprotonation reactions of surface hydroxyl groups, and other functional groups, occurs through the following reactions (Kamel and Navratil 2002):

M-OH + H+ M-OH2+ pH < pHpzc [2]

M-OH M-O- + H+ pH > pHpzc [3]

The sorption of anions depends significantly on the pH of the soil solution, as the surface charge of the minerals changes when pH decreases below the pH point of zero charge (pHpzc) of the minerals or increases above it (Viers et al. 2001). When pH = pHpzc, the concentration of protonated (M-OH2

+) and deprotonated (M-O-) sorption sites are equal. As long as M-OH2

+ sites are available, anions can be sorbed on the mineral surfaces due to the positive surface charge (M-OH2

+ > M-O-). Once pH increases and overcomes the pHpzc (pH > pHpzc), cations can be sorbed on mineral surfaces and anions are repelled from the vicinity of surfaces due to the negative surface charge (M-OH2

+ < M-O-).

Radionuclides can be sorbed on dissolved organic carbon (DOC) compounds, such as humic acids or fulvic acids (Bibak and Borggaard 1994; Dhillon et al. 2010; Gustafsson and Johsson 1992). Increased DOC concentration may also inhibit the sorption of molybdenum by means of forming coatings on iron oxide surfaces and blocking the entrance to the interdomains and pores of the solid (Lang and Kaupenjohann 2003).

The competition for sorption sites between anionic radionuclides and stable anions present in the soil solution may increase as the anion concentration in the soil solution increases (e.g. Barrow 1970; Brinton and O’Connor 2003; Dhillon and Dhillon 1999; Fujikawa abd Fukui 1997; Pezzarossa et al. 1999). In the case of molybdenum and selenium, especially phosphate and sulphate concentrations have to be taken into consideration (Barrow 1970, 1972; Karamian and Cox 1978; Neal et al. 1987b). Cl- has virtually no effect on the sorption of Mo even in concentrations as high as 1 mol/l (Goldberg and Forster 1998). 4.2.1 pH of the soil solution

In Lusa et al. (2009) pH of the soil solution increased with depth from the humus layer (pH 4-5) to the undermost mineral soil layers (pH 7.5-8) in excavator pits OL-KK14 and OL-KK15. In OL-KK16 pH in the humus layer was approximately 6.7, decreased

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to 6 in the MS1 mineral soil layer, and increased with depth to 7.5 in the undermost mineral soil layer (MS4). Figure 5 presents the pH of the soil solution as a function of the soil layer.

In forest intensive monitoring plot FIP4 (Scots pine forest) and FIP10 (Norway spruce forest) stands situated on Olkiluoto Island, the pH values of the soil solution increased with depth (Haapanen 2006, 2007a, 2008; Haapanen 2009): pH increased approximately from 4.2 to 5.35 as the depth increased from 5 to 30 cm in FIP4 plot. In FIP10, the pH value in 5 cm was Approximately 4.46 and in 30 cm 5.75. These results are in good agreement with the result from Lusa et al. (2009).

Figure 5. pH of the soil solution samples as a function of the soil layer (depth) in excavator pits OL-KK14, OL-KK15 and OL-KK16 (Lusa et al. 2009).

4.2.2 Dissolved organic carbon (DOC) in the soil solution

In filtered (0.45 μm) soil water samples the largest DOC-values were found in the humus layer samples (Figure 6) (Lusa et al. 2009). Among the excavator pits, the highest value (160 mg/l) was found in OL-KK14, whereas the smallest (80 mg/l) was found in OL-KK16. For OL-KK15 DOC was 120 mg/l in the humus layer. DOC-values decreased rapidly to 5-20 mg/l in the second uppermost mineral soil layer (MS2) and remained practically within this range in all underlying mineral soil layers.

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Figure 6. DOC concentrations of the soil solution samples as a function of the soil layer (depth) for excavator pits OL-KK14, OL-KK15 and OL-KK16 (Lusa et al. 2009). In FIP4 and FIP10 stands the DOC concentration decreased with depth (Haapanen 2006, 2007a, 2008; Haapanen 2009). On the other hand, in FIP10 stand the DOC concentrations were higher than in FIP4 plot; in the depth of 5 cm and 30 cm the DOC concentration were approximately 108.0 mg/l and 28.2 mg/l for FIP4 compared with 127.4 mg/l and 51.5 mg/l for FIP10. 4.2.3 Cations in the soil solution

The main cations measured from the soil solution samples were sodium (Na), potassium (K), magnesium (Mg) and calcium (Ca) (Lusa et al. 2009). Other determined cations were aluminium (Al), chromium (Cr), manganese (Mn), iron (Fe), nickel (Ni), copper (Cu), zinc (Zn), arsenic (As), sulphur (S), cadmium (Cd), cesium (Cs), lead (Pb) and uranium (U). The concentrations of Na, K, Mg and Ca are presented in Figure 6, whereas the concentrations of the trace cations are discussed elsewhere (Lusa et al. 2009; Söderlund et al. 2011)

Among the studied mineral soil malyers in excavator pits OL-KK14, OL-KK15 and OL-KK16, the concentrations of the main cations sodium (Na), magnesium (Mg) and calcium (Ca) in the soil solution showed no dependence on depth in the range of 0-300 cm. However, the concentration of potassium (K) increased in general with depth and reached its highest values in the lowest mineral soils layer (Lusa et al. 2009).

Among the excavator pits examined, the maximum sodium concentration was found in the soil solution of MS5 layer in excavator pit OL-KK15 (16 mg/l), whereas the lowest concentration occurred in MS2 layer in OL-KK14 (1.5 mg/l). The average Na concentrations were 3.9±2.1 (value±standard deviation), 8.3±5.4 and 5.1±2.0 mg/l for the entire excavator pits OL-KK14, OL-KK15 and OL-KK16, respectively.

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The highest and lowest concentrations of magnesium were found in the MS2 layers in OL-KK15 and OL-KK14, the respective values being 12.5 and 0.5 mg/l. The average Mg2+ concentrations were 2.5±1.7, 6.5±3.9 and 5.3±2.4 mg/l for the entire excavator pits OL-KK14, OL-KK15 and OL-KK16.

The highest calcium concentration, 24 mg/l, was found in the soil solution sample taken from the humus and MS3 layers of the excavator pit OL-KK16. The lowest Ca2+ concentration, 2 mg/l, was found in the MS2 and MS4 layers from the excavator pits OL-KK14 and OL-KK15. The average Ca2+ concentrations for the entire excavator pits OL-KK14, OL-KK15 and OL-KK16 were 7.7±5.8, 9.9±6.7 and 17.3±6.5, respectively.

The potassium (K) concentrations in the soil solution were the highest in the humus layers, ranging from 12 to 17 mg/l (Lusa et al. 2009). Among the mineral soil layers, the highest K+ concentration in the soil solution was found in the MS5 layer in OL-KK15 (11 mg/l). The lowest K+ concentration was found in the MS2 layer in OL-KK14 (0.5 mg/l). The average K+ concentrations were 5.6±6.2, 7.5±5.5 and 6.1±4.3 mg/l for the entire excavator pits OL-KK14, OL-KK15 and OL-KK16, respectively. Figure 7 presents the concentrations of Na, Mg, Ca and K in the soil solution samples taken from the soil layers in the excavator pits OL-KK14, OL-KK15 and OL-KK16.

Figure 7. The concentrations of Na (A), Mg (B), Ca (C) and K (D) in the soil solution samples as a function of soil layer (depth) (Lusa et al. 2009). The cation concentrations were measured for the samples taken from the soil layers in excavator pits OL-KK14, OL-KK15 and OL-KK16.

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In FIP4 plot the average concentrations of Na+, Mg2+, Ca2+ and K+ for the depth range of 5-30 cm were 3.3±1.1, 1.0±0.3, 3.7±1.3 and 2.4±1.5 mg/l, respectively (Haapanen 2006, 2007a, 2008; Haapanen 2009). In FIP10 plot the average concentrations of Na+, Mg2+, Ca2+ and K+ were 11.0±5.9, 3.6±1.2, 10.7±1.9 and 3.0±2.2 mg/l in 5-30 cm, respectively.

4.2.4 Anions in the soil solution

Fluoride (F-), chloride (Cl-), nitrate (NO3-) and sulphate (SO4

2-) were the anions measured from the soil solution samples (Lusa et al. 2009). Fluoride concentration had no dependence on soil depth. In soil solution samples taken from the excavator pits OL-KK14 and OL-KK16, fluoride concentrations were in the range of 1 to 3 mg/l. In OL-KK15 F- concentration had the highest concentration in the humus layer, 8 mg/l, and the lowest in the undermost MS5 mineral soil layer, 1.5 mg/l. The average F- concentrations for the entire excavator pits OL-KK14, OL-KK15 and OL-KK16 were 2.6±0.5, 4.1±2.6 and 1.8±1.0 mg/l, respectively. Figure 8 presents the concentrations of F-, Cl-, NO3

- and SO4

2- in the soil solution samples taken from different soil layers in excavator pits OL-KK14, OL-KK15 and OL-KK16.

The maximum chloride concentration in excavator pits OL-KK14, OL-KK15 and OL-KK16 was found in the humus layer, 15, 58 and 20 mg/l, respectively (Lusa et al. 2009). When the concentration of chloride decreased with depth to approximately 10 mg/l in MS4 mineral soil layer in OL-KK14 and OL-KK16, in OL-KK15 the Cl- concentration increased from the lowest value (20 mg/l) in MS1 to 33 mg/l in the MS2-MS5 mineral soil layers. The average Cl- concentrations for the entire excavator pits OL-KK14, OL-KK15 and OL-KK16 were 13.1±5.0, 35.4±12.1 and 13.5±7.8 mg/l, respectively.

Nitrate concentration was practically constant throughout the soil layers in excavator pits OL-K14 and OL-KK15, the value being 0.5-1 mg/l (Lusa et al. 2009). In OL-KK16 the highest NO3

- concentration was found in the humus layer (7.5 mg/l), and decreased with depth to 1.5 mg/l in the uppermost MS1 mineral soil layer and 0.5 mg/l in MS2-MS4. The average NO3

- concentrations for the entire excavator pits OL-KK14, OL-KK15 and OL-KK16 were 0.6±0.3, 0.3±0.2 and 2.1±3.0 mg/l, respectively.

Among the studied anions, sulphate concentrations varied the most (Lusa et al. 2009). In OL-KK14, the SO4

2- concentration was 20-40 mg/l throughout the excavator pit. In OL-KK15, the highest concentration was found in MS2 layer, 205 mg/l, and in OL-KK16 in MS3 layer, 160 mg/l. The lowest concentrations were 40 mg/l and 20 mg/l in MS5 and MS1 layers in OL-KK15 and OL-KK16, respectively. The average SO4

2- concentrations for the entire excavator pits OL-KK14, OL-KK15 and OL-KK16 were 27.1±10.9, 88.8±63.8 and 69.2±53.3 mg/l, respectively.

In FIP4 plot the average concentrations for Cl-, NO3- and SO4

2- for the depth range of 5-30 cm were 4.3±1.8, 0.026±0.006 and 1.6±0.9 mg/l, respectively (Haapanen, R. 2006, 2007a, 2008; Haapanen, A. 2009). In FIP10 plot, the average concentrations for Cl-, NO3

- and SO42- were 13.0±5.6, 0.05±0.03 and 7.7±4.9 mg/l in depth of 5-30 cm,

respectively.

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Figure 8. The concentrations of F- (A), Cl- (B), NO3- (C) and SO4

2- (D) in the soil solution samples as a function of soil layer (depth) (Lusa et al. 2009). The anion concentrations were measured for the samples taken from the soil layers in excavator pits OL-KK14, OL-KK15 and OL-KK16.

4.2.5 Ionic strength of the soil solution

The definition of ionic strength (I) is given in Equation 4. As the ionic strength of the solution is being calculated, all the ionic species present in the solution has to be taken into consideration. Ionic strength for the studied soil solutions was calculated from the concentrations of anions and cations given in Lusa et al. (2009).

I = ½ [I]ZI2 [4]

Equation 4. The definition of the ionic strength I. In Equation 4 [i] = the concentration of ion i and Zi = the charge of the ion i.

The highest average ionic strength was calculated for the entire excavator pit OL-KK15 (376 mmol/l), followed by OL-KK16 (297 mmol/l) and OL-KK14 (116 mmol/l). The ionic strength increased with increasing depth from 113 mmol/l in the humus layer to 171 mmol/l in the undermost mineral soil layer MS4 in the excavator pit OL-KK14 (Lusa et al. 2009). For OL-KK15, the ionic strength increased from 240 mmol/l in the humus and MS1 layers to 865 mmol/l in the MS2 layer, and decreased again with depth to 172 mmol/l in the MS5 layer. In the excavator pit OL-KK16, the highest value was

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found in the mineral soil layer MS3 (675 mmol/l), whereas the lowest value was in the uppermost mineral soil layer MS1 (77 mmol/l).

Table 3 summarizes the ionic strengths calculated for the soil solution samples taken from different soil layers from the excavator pits OL-KK14, OL-KK15 and OL-KK16 after the data given in Lusa et al. (2009). The average ionic strengths are also included.

Table 3. The ionic strengths for the soil solution samples taken from different soil layers from the excavator pits OL-KK14, OL-KK15 and OL-KK16 and the average ionic strengths (calculated from the data given in Lusa et al. 2009).

Excavator pit Soil layer (depth cm) Ionic strength (mmol/l)

OL-KK14 Humus (0-5) 113 MS1 (5-20) 62

MS2 (20-60) 83 MS3 (60-105) 153

MS4 (105-240) 171 AVERAGE 116

OL-KK15 Humus (0-7) 243 MS1 (7-50) 239

MS2 (50-80) 863 MS3 (80-110) 524

MS4 (110-160) 215 MS5 (160-260) 172

AVERAGE 376

OL-KK16 Humus (0-10) 256 MS1 (10-30) 77 MS2 (30-50) 254 MS3 (50-110) 675

MS4 (110-300) 223 AVERAGE 297

4.3 Redox conditions

The main factors governing the speciation of radionuclides in soils are the redox potential and pH (Koch-Steindl and Pröhl 2001). Several factors affect the redox potential of the soil, these including e.g. soil type, distance to the water table, biological activity and existence of very low permeability soil horizons. Redox reactions taking place in soils are controlled by the quantity and diffusion rate of oxygen in soil layers (Koch-Steindl and Pröhl 2001). Redox conditions and soil aeration status have been discussed in detail elsewhere (Söderlund et al. 2011).

Table 4 presents the microbially mediated redox-reactions taking place under normal (well aerated), wet and waterlogged soil conditions, and the redox potential values attached to these in pH 7.

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Table 4. Microbially mediated redox reactions taking place under normal (well aerated), wet and waterlogged soil conditions and the redox potential values attached to these in pH 7 (Koch-Steindl and Pröhl 2001).

Soil aeration status Redox reactions Redox potential (mV)

Well aerated soils NO3- reduction begins

Mn2+ formation begins 450-550 350-450

Wet soils O2 is not detected

NO3- is not detected

Fe2+ formation begins

330 220 150

Waterlogged soils

SO42- reduction and S2-

formation CH4 formation begins SO4

2- is not detected

-50

-120 -180

4.4 Sorption

Soils contain different cation adsorbing components especially in the silt and clay fractions (Koch-Steindl and Pröhl 2001). Important sorbents for radionuclides in soils are, for instance minerals like smectite, illite and chlorite as well as oxides and hydroxides of iron, aluminium and manganese. Significant cation adsorbing minerals are illite, smectite, vermiculite and mica due to their intrinsic negative charge, whereas anion adsorbing clay minerals are halloysite and imogolite. In the clay fraction of the Olkiluoto soils major minerals are quartz, potassium feldspar and plagioclase (Lusa et al. 2009). Minor minerals include illite, hornblende, chlorite, mica, amphibole, hematite and kaolinite (Lintinen et al. 2003; Lusa et al. 2009). On the other hand, the sorption of anions is improved in the presence of organic matter (Lang and Kaupenjohann 2000; Choppin et al. 2009; Pezzarossa et al. 1999).

Clay minerals carry permanent negative charge resulting from the isomorphic substitution within the mineral lattice at the time of the formation. In tetrahedral coordination Si4+ -cations are replaced with Al3+-cations, and Al3+ is replaced by Fe3+, Fe2+, Mg2+, etc. in the octahedral coordination (Essington 2004; Koch-Steindl and Pröhl 2001). This permanent negative structural charge is balanced by exchangeable cations present in the interlayer spaces of clays. The capacity of clays to adsorb cations is specific to the mineral in question, and depends on the permanent negative charge (the degree of isomorphic substitution) of the clays and the saturation degree (the amount of cations sorbed in the interlayer spaces). Clay minerals also carry a variable positive or negative charge depending on the protonation degree of the hydroxyl (-MOH), aluminol (-AlOH) and silanol (-SiOH) functional groups (Koch-Steindl and Pröhl 2001).

The surface reactivity of soil organic matter is attributed to rather few different functional groups in spite of its chemical and structural complexity. Some of these, for example carboxyl and sulfonic groups are easily deprotonated and exhibit negative charge leading to moderate or strong interaction with cations. Functional groups that protonate easily and develop positive charge attracting anions are, for instance, amine, imide, phenolic-OH and sulfhydryl groups. Carbonyl group interacts only weakly with soluble species because it is only polarized, not ionized. The most often encountered

26

functional groups in the soil organic matter include carboxyl, phenolic-OH and carbonyl groups (Essington 2004). The scematic figures of soil organic matter functional groups amine, carboxyl, imide, phenolic-OH, sulfonic and sulfhydryl in acidic and alkaline conditions can be seen in Figure 9.

Figure 9. The scematic figures of soil organic matter functional groups amine, carboxyl, imide, phenolic-OH, sulfonic and sulfhydryl in acidic and alkaline conditions (Essington 2004).

4.5 Microbial effects

Micro-organisms can affect the composition of soil and soil solution via several different mechanisms, and thereupon affect the migration of radionuclides and their and geo- and biochemical cycling (Tamponnet et al. 2008). These mechanisms include:

Alter the soil pH via production of organic acids and bases.

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Alter the soil oxidation-reduction conditions from oxic to anoxic via respiration (i.e. the consumption of oxygen).

Alter the structure of soil via the formation of mineral aggregates.

Radionuclides may become fixed to the cell walls of micro-organism or to the

extracellular polysaccharides of bacteria. Fixed radionuclides remain in the micro-organisms even if the environment (e.g. soil solution) of the micro-organisms changes.

Radionuclides can be actively taken up by micro-organisms (bioaccumulation).

Secretation back to the soil solution can take place if the composition of the soil solution changes.

In soil systems with a large input of organic matter, the soil microbiota (including bacteria, fungi, protozoa, algae and microfauna) is highly active and has an important role in the transformation of organic matter (Tamponnet et al. 2008). Soil micro-organisms can modify both the soil structure and the geochemical cycle of elements. Radionuclides associated with the soil organic matter and soil micro-organisms are of great importance since they remain readily exchangeable and are available for plant uptake.

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29

5 MOLYBDENUM

Molybdenum is a transition metal with the atomic number of 42. It belongs to the sixth group in the periodic table of elements together with chromium and tungsten. The electron configuration of Mo is written as [Kr]4d55s1. The atomic and covalent radius of Mo are 124 pm and 145 pm, respectively (Suresh and Koga 2001).

Molybdenum has seven stable and nearly 20 radioactive isotopes. The stable isotopes and their relative abundances are: Mo-92 (14.84 %), Mo-94 (9.25 %), Mo-95 (15.92 %), Mo-96 (16.68 %), Mo-97 (9.55 %), Mo-98 (24.13 %) and Mo-100 (9.63 %). The most important radioisotope of molybdenum in spent nuclear fuel is Mo-93 due to its long physical half-live of 4000 years (Firestone et al. 1998; Haapanen et al. 2009). In the long-term safety assessment of spent nuclear fuel Mo-93 is classified as high priority radionuclide (Haapanen et al. 2009; Hjerpe et al. 2010). Mo-93 is produced in the neutron activation of stable Mo-92 by the reaction Mo-92(n, )Mo-93. Mo-92 is found as an impurity in the constructive materials in the nuclear reactor systems. Mo-93 decays to Nb-93m (88 %) and Nb-93 (12 %) by electron capture (EC) (Lehto and Hou 2010).

Molybdenum is strongly siderophilic and/or chalcophilic element, and its enriched in iron and sulphide rich phases (Wiberg et al. 2001). Molybdenum can form minerals where it is the main cation, for example molybdenite (MoS2), wulfenite (PbMoO4) and powellite (Ca(Mo,W)O4), but typically it is incorporated in sulphide minerals such as galena (PbS), pyrite (FeS2) and sphalerite ((Zn,Fe)S) as a trace element (Ure and Berrow 1982).

The average Mo content in the continental crust has estimated to be 1.1 mg/kg, and 0.8 mg/kg in the bulk crust (Rudnick and Gao 2004). The molybdenum content in soils is affected by the parent material; the concentration is higher for soils of vulcanic origin than for soils consisting of eroded sandstone (Lang and Kaupenjohann 2000). Also, soils derived from clay till contain more Mo than soils formed from sandy or gravelly till (Bibak et al. 1994). The Europian median value for molybdenum content in aqua regia extracted topsoils and subsoils are 0.62 and 0.52 mg/kg, respectively (Salminen 2007). In Finland, Mo content in the topsoils varies between 0.20-1.40 mg/kg, and 0.15-0.91 mg/kg in subsoils (Salminen 2007). The European median values for molybdenum concentration in aqua regia extracted stream sediments and waters are 0.60 mg/kg and 0.22 μg/l, respectively. The corresponding ranges in Finland are 0.25-1.37 mg/kg and 0.03-0.82 μg/l (Salminen 2007).

High iron and aluminium oxide content, clay fraction and organic matter content and low pH increase the sorption of Mo on soils (Barrow 1970, 1972; Bibak et al. 1994; Bibak and Borggaard 1994; Brinton and O’Connor 2003; Goldberg et al. 1996; Karamian and Cox 1978; Lang and Kaupenjohann 2000; McGrath et al. 2010; Mikkonen and Tummavuori 1993b; Roy et al. 1986; Xie and MacKenzie 1991). The high cation concentrations in the solution may increase the sorption Mo on soils (Barrow 1972). Organic matter is the main sorbent for molybdenum (Lang and Kaupenjohann 2000). Soil organic matter and oxalate extracted iron content can be used in regression models as predictors to forecast the bioavailability and toxicity of Mo to plants, and the toxicity decreases with increasing content of soil organic matter and oxalate extracted iron due to sorption of Mo (McGrath et al. 2010).

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The sorption of molybdenum on soils decrease as pH increases (Barrow 1970, 1972; Bibak and Borggaard 1994; Goldberg et al. 1996; McGrath et al. 2010; McGrath et al. 2010, Mikkonen and Tummavuori 1993b), the concentration of competing anions, such as phosphate and sulphate, increases (Barrow 1970, 1972; Brinton and O’Connor 2003; Karamian and Cox 1978; McGrath et al. 2010; Roy et al. 1986; Xie and MacKenzie 1991) and the content of coarse grained sand fraction increases (Mikkonen and Tummavuori 1993b). The surface area occupied by adsorbed MoO4

2--ion was calculated to be 35 nm2 (Motta and Miranda 1989). 5.1 The influence of soil redox potential and pH

Molybdenum forms compounds in multiple oxidation states depending on the prevailing Eh-pH conditions. In acidic solution, when pH is close to 1, the dominant molybdenum species are Mo(OH)5(H2O)+ and Mo(OH)6 (Cruywagen and De Wet 1988). Protonated forms of molybdate-ion (MoO4

2-), namely H2MoO4 and HMoO4-, can be found in pH

interval 0-4. The pKa-values for H2MoO4 and HMoO4- are 3.66 and 3.81, respectively.

MoO42- is the major species in pH over 5, and dominant form in large variety of Eh-pH

conditions. MoO42- is very mobile in highly basic conditions (Morrison and Spangler

1992). Figure 10 presents the Eh-pH diagram of molybdenum.

Figure 10. The Eh-pH diagram of molybdenum (Takeno 2005). Mo=10-10, 298.15 K and 105 Pa.

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The formation of polynuclear molybdenum species can take place in concentrations over 10 mmol/l (Cruywagen and De Wet 1988). Polynuclear species are, for example HMo2O6

+, HMo2O7-, H2Mo7O24

4- and Mo8O264-. The structure of the polynuclear species

becomes more complex as the Mo concentration increases. The protonation and deprotonation degree of the species depends on pH of the solution: highly protonated and positive forms are found in acidic pH, which are turned into less protonated and negative types with increasing pH. The chemisorption of polynuclear molybdenum species does not take place because Mo has octahedral coordination in its compounds (Cruywagen and De Wet 1988).

Molybdenum can be associated with excess amounts of sulphur to form tetrathiomolybdate (MoS4

2-) under anoxic conditions and the presence of sufficient amounts of H2S (Erickson and Helz 2000; Helz et al. 1996; Luther et al. 1986). The sulfidation reaction of MoO4

2- can be expressed as follows (Erickson and Helz 2000):

[5]

Intermediates in the sulfidation reaction include MoO3S2-, MoO2S22- and MoOS3

2-

(Erickson and Helz 2000).

The effect of pH on the sorption of Mo on soils and minerals is affected by the surface charge of soil particles (see Chapter 4.2 for the explanation), and to a lesser extent by the protonation-deprotonation reactions of Mo species present in soils. pH is the main factor affecting the sorption of molybdenum in soils (Brinton and O’Connor 2003). Sorption is the greatest in acidic pH, below pH 5, and decreases with increasing pH both in the mineral and organic soils (Barrow 1970; Bibak and Borggaard 1994; Brinton and O’Connor 2003; Goldberg et al. 1996; Goldberg and Forster 1998). pH also affects the displacement of Mo from soils’ sorption sites with competing anions, such as phosphate (Barrow 1974). The specific influence of phosphate is the most pronounced in pH neutral conditions. Decrease in the temperature from 40 °C to 10 °C enhanced the effect of PO4

3-. The water content of the soil affects the redox condition and desorption of Mo from soils (Barrow 1974). According to Barrow (1974), the fraction desorbed by phosphate of the initially adsorbed Mo increased from approximately 4-4.5 μg Mo/g soil to 7.1-8.5 μg Mo/g soil as the water conditions in soil changed from waterlogged to dry.

5.2 The influence of micro-organisms

Molybdenum has an important role in enzymatic redox reactions within plants, animals and micro-organisms (Sun 2000). Molybdenum is needed in the Fe-Mo cofactor in the nitrogenase enzyme, which is responsible for nitrogen fixation and nitrate reduction to ammonium (NH4

+). Nitrogen-fixing bacteria are the most active in the surface parts of soils, especially in the litter layer (Wichard et al. 2009).

Certain nitrogen-fixing bacteria, for example siderophores producing Azobacter vinelandii, are capable of competing with the organic matter of free molybdenum oxoanions present in soils. Azobacter vinelandii excretes catechol compounds 2,3-dihydroxybenzoic acid, tris(catechol) protochelin and bis(catechol) azotochelin. The last two mentioned catechol compounds are known to complex molybdenum very efficiently and strongly. Mo-catechol complexes are readily available for Azobacter

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vinelandii, short-term rate constant for protochelin was 8.3 x 10-21 mol Mo/cellxmin. Azobacter vinelandii has a regulated uptake system for Mo-catechol complexes; when cells have sufficient amounts of Mo- catechol complexes within, the concentration of Mo in the solution does not decrease (Bellenger et al. 2008).

In anoxic conditions the formation of H2S is induced by the sulphate reducing bacteria (SRB). Molybdenum can form stable sulphides in the reaction with H2S, or be incorporated into the final reduction product of H2S and Fe(II), pyrite (FeS2). Pyrite can also be formed in anoxic peat layers, but in such instances, molybdenum concentrates in the organic peat matter instead of pyrite (Dellwig et al. 2002).

5.3 Sorption on organic matter and organic soils

Organic matter is the most important sorbent for molybdenum in soils and elevated Mo concentrations are often measured from the soil organic humus layer (Lang and Kaupenjohann 2000). The sorption of molybdenum on organic soils increases as the organic matter content increases (Karamian and Cow 1978). Also, molybdenum has high affinity towards humic substances (HS), as the complexation capacity of HS derived from tropical peat samples has a high value of 13.37 mmol Mo/g TOC (Botero et al. 2010). The respective values for Ca2+, Mg2+, Al3+ and Cu2+ were 4.25, 4.33, 13.59 and 13.75 mmol Mo/g. The sorption of molybdenum on humic acid shows intensive dependence on pH (Bibak and Borggaard 1994). For example, Bibak and Borggaard (1994) found that the sorption of Mo on humic acid decreased 90 % as pH increased from 3.5-4 to 7. The maximum sorption found in pH 3.5-4 was 0.45 mmol Mo/g humic acid.

The sorption mechanism for molybdenum on organic matter is complex formation with carboxyl (-COOH), phenol (C6H5OH) and catechol (C6H4(OH)2) groups (Cruywagen and De Wet 1988; Wichard 2009). Furthermore, organic molecules containing two hydroxyl groups in the ortho-position to each other may complex Mo (Xie and McKenzie 1991). Sorption takes place because the octahedral coordination of Mo is preferred instead of tetrahedral, and is achieved via the expansion of the coordination sphere. In octahedral coordination two of the oxygens in MoO4

2- are in the cis-dioxo conformation (Wichard 2009). As a result, two oxygens are in the axial position and four in equatorial position. These equatorial oxygens are bound to a single carbon atom forming two five-membered rings.

The cycling of molybdenum between plants and litter/organic layer takes place in soils (Wichard 2009): Mo is transported to plants, e. g. trees through roots and delivered to leaves. As leaves fall down and begin to decompose, Mo is freed to soil and bounds to tannin and tannin-like substances created in the decomposition reaction.

In tannin and tannin like substances, the complexing group is catechol (C6H4(OH)2), and complexation is dependent on pH (Wichard 2009). Maximum complexation is found in pH 6.1, decreasing in lower and higher pH-values. The decrease is higher in basic pH-values, when only a half of Mo is in a complex, whereas in pH 4.7 most of the Mo is in complex.

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Figure 11 presents molybdenum content in a German Cambisol profile (Lang and Kaupenjohann 2000). Molybdenum concentration decreases as a function of depth as the amount of organic matter and respective sorption sites decreases. The highest concentration is found in the organic humus layer (H), emphasizing the role of organic matter as the primary and the most important sorbent for Mo in soils (Lang and Kaupenjohann 2000; Wichard 2009). The next highest content is seen in the uppermost mineral soil layer (Ah; Figure 14), in which the organic matter content may still be quite high. Molybdenum concentration decreases in BV1 and BV2 layers (Lang and Kaupenjohann 2000), and so does the retention of Mo on organic matter (Wichard 2009). The lowest Mo concentration was found in the unchanged parent soil (Cv; Figure 14). As the organic matter content decreases in soils with depth, the importance of iron and aluminium oxides as sorbents increases (Wichard 2009).

Figure 11. Molybdenum content in a German Cambisol profile (Lang and Kaupenjohann 2000). H=humus, Ah=uppermost mineral soil layer enriched with organic matter (humus), Bv1 and Bv2=illuvial layers containing plinthite and Cv=unchanged parent soil.

5.4 Sorption on mineral soils

In the mineral soils the main Mo sorbing components are iron and aluminium oxyhydroxides (Barrow 1970; Brinton and O’Connor 2003; Goldberg et al. 1996). The sorption of molybdenum on mineral soils increases as the pH of the soil decreases, and the extractable aluminium and iron content increase (Barrow 1970; Gong and Dohahoe 1997). Sorption of Mo on mineral soils is often reversible (Carroll et al. 2006). Figure 12 presents the proposed sorption mechanism of HMoO4

- and HMo2O7- onto a surface

of solid adsorbent.

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Figure 12. Sorption of (a) HMoO4- and (b) HMo2O7

- to a surface of an adsorbent (Cruywagen and De Wet 1988).

The sorption of Mo is siginificant in acidic soils (Mikkonen and Tummavuori 1993b), and the more acidic soil conditions favour the sorption of MoO4

2- over phosphate and sulphate (Barrow 1970). Phosphate is efficient in displacing sorbed MoO4

2- from soils’ surfaces in neutral pH conditions (Barrow 1974). However, the displacement efficiency decreases when the interaction time between MoO4

2- and soil is increased, or the pH of the solution is decreased. It was proposed that the sorption mechanism of MoO4

2- was changed due to mineralization or formation of a stronger bond between MoO4

- and soil surface with increased contact time. Increase in the temperature was also found to increase the nondisplaced fraction of Mo. Ca2+-ions present in the solution were found to increase the sorption of molybdate on soils (Barrow 1972). The explanation given stated that the surface charge of the soils shifted to more positive values upon the sorption of Ca2+, leading to increased uptake of Mo. The difference in the sorption of Mo on soils in solutions with or without Ca2+ was more pronounced for soils with lower pH. Calcium can also form precipitates with MoO4

2- ions (Essington 1990). The solubility constant (Ksp) for CaMoO4, mineral powellite in aqueous solution is 10-7.93 (Felmy et al. 1992).

Molybdenum is retained more efficiently on fine-textured soils than on coarse-textured soils (Gong and Dohahoe 1997; Martinez et al. 2003; Mikkonen and Tummavuori 1993b; Roy et al. 1986). For example, Mikkonen and Tummavuori (1993b) investigated the sorption of molybdenum on Finnish clay, finer fine sand and coarser fine sand soil samples and the effect of pH. The efficiency of the soils to retain molybdenum decreased in order clay > finer fine sand > coarser fine sand. Sorption was found to be pH-dependent: the maximum sorption took place below pH 4.5, where 60-80 % of the initial amount of Mo was sorbed, and decreased as a function of pH to <20 % in pH 7. Clay soil sorbed considerable amounts of Mo in pH 5 and 7, approximately 73 % and 15 %, respectively. The amounts sorbed on finer and coarser fine sand were lower. Desorption of Mo from soils increased in the order clay > finer fine sand > coarser fine sand. The solution final pH was found to decrease in the respective order, thus explaining the desorption sequence.

Martinez et al. (2003) studied the desorption of Mo from metal contaminated silt loam-silty clay loam soil, and the effect of temperature on the desorption reaction. Firstly, the desorption of Mo increased with time from 0 to 550 days. Secondly, the desorption increased upon temperature from 0 °C to 70 °C, but decreased in 90 °C due to a drop in

35

the pH of the solution. They proposed that upon increase in temperature Mo and other heavy metals can be liberated from complexes with organic matter due to the increased solubility of organic matter at elevated temperatures. However, Goldberg and Forster (1998) found that an increase in the temperature from 10 °C to 40 °C had a minimal effect on the sorption of Mo on coarse-loamy soil.

Xie and MacKenzie (1991) examined the sorption-desorption behaviour of Mo on clay and loam soil and the effect of phosphate additions on the sorption-desorption reactions. Experiments were conducted by adding Mo (MoO4

2-) and P (HPO42-) simultaneously.

Sorption of Mo decreased as the amount of P increased, and the decrease was higher for loam soil than for clay soil. This was explained as the different amount of phosphate needed before any effect on the sorption of Mo was visible; PO4

3- amounts were >20 mmol/kg and >30 mmol/kg for loam and clay soil, respectively. The desorbed fraction of Mo was generally below 10 % of the adsorbed fraction. At P concentrations >30 mmol/kg, desorption from loam soil was higher than from clay soil.

Lusa et al. (2009) examined the content and distribution of molybdenum in organic humus layer and mineral soil layer samples taken from three excavator pits dug on Olkiluto Island (OL-KK14, OL-KK15 and OL-KK16). The composition of excavator pit OL-KK14 was clay/sandy till, OL-KK16 fine to coarse sand and sandy till and OL-KK15 clay, fine sand, sand, fine-grained silty/clayish till and coarse-grained sandy till. Molybdenum was extracted from soil samples using five different extraction solutions, namely 1M BaCl2 (extraction step A; exchangeable cations), 1 M CH3COONH4 in 25 % CH3COOH (B; bound to carbonates), 0.04 M NH2OH*HCl in 25 % CH3COOH (C; bound to Fe and Mn oxides), 0.02 M HNO3 and 30 % H2O2 (D; bound to organic matter) and aqua regia (E; residual fraction). In OL-KK14, the highest molybdenum concentration was typically measured from the exchangeable fraction (1.97-2.15 mg/kg), followed by Mo bound to Fe and Mn oxides (1.27-1.45 mg/kg), whereas in OL-KK15 the order was reverse. Smaller amounts of Mo (0.10-0.56 mg/kg) were usually found in the fractions related to carbonates and organic matter in OL-KK14 and OL-KK15. In excavator pit OL-KK16, molybdenum was found solely in the fraction bound to organic matter. Mo was not found in the residual fraction in none of the studied soil layers. In Appendix 1, the concentrations of molybdenum in extraction solutions, soil layers and excavator pits are given. Also the average Mo concentrations in soil layers are included.

For excavator pits OL-KK15 and OL-KK16, molybdenum concentration in the humus layer (1.6-1.9 mg/kg) was noticeably higher than in the mineral soil layers MS1-MS5 (<0.9 mg/kg) (Figure 13). In OL-KK14 the Mo concentration was approximately 1 mg/kg in every soil layer (Figure 13). Molybdenum concentration among the mineral soil layers decreased in the order OL-KK14 > OL-KK15 > OL-KK16. In the humus layer the respective order was OL-KK16 > OL-KK15 > OL-KK14.

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-0,2 0,0 0,2 0,4 0,6 0,8 1,0 1,2 1,4 1,6 1,8 2,0

MS5

MS4

MS3

MS2

MS1

humusS

oil l

ayer

Mo concentration (mg/kg)

OL-KK14 OL-KK15 OL-KK16

Figure 13. Molybdenum concentration in organic humus layer (humus) and mineral soil layers (MS1-MS5) in excavator pits OL-KK14, OL-KK15 and OL-KK16 in Olkiluoto (Lusa et al. 2009). Data points are the average concentrations calculated from the sum of Mo contents in different extraction solutions, namely exchangeable cations (1M BaCl2), bound to carbonates (1 M CH3COONH4 in 25 % CH3COOH), bound to Fe and Mn oxides (0.04 M NH2OH*HCl in 25 % CH3COOH), bound to organic matter (0.02 M HNO3 and 30 % H2O2) and residual fraction (aqua regia) given in Lusa et al. (2009).

5.4.1 Podzols

Figure 14 presents molybdenum content in a German Podzol profile (Lang and Kaupenjohann 2000). The molybdenum content decreased as a function of depth from topmost mineral soil layer (A) to the leaching layer (E), and increase again in B and C horizons (Bibak et al. 1994; Lang and Kaupenjohann 1999, 2000). The highest Mo concentration is found in the organic humus layer (H), followed by the illuvial (enrichment) layers (Bs, Bh and Bv) (Lang and Kaupenjohann 2000). The lowest concentrations are found in the unchanged parent soil (Cv), and the leaching layer (E). The reason for the ‘abnormally’ low Mo content in E horizon is due to the leaching of Mo-sorbing organic matter and aluminium and iron compounds, which are enriched in the illuvial layers. Also, molybdenum is leached from the eluvial layer more easily than iron. Typically, Mo is not in plant available form in podzols due to sorption onto poorly crystalline Al and Fe oxide surfaces and organic matter (Bibak et al. 1994). The fraction representing the plant-available Mo fraction in soils is the free and the non-specifically (adsorbed) phases (Lang and Kaupenjohann 1999).

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Figure 14. Molybdenum content in a German Podzol profile (Lang and Kaupenjohann 2000). H=humus, E=eluvial layer, Bh=illuvial layer enriched with organic matter (humus), Bs= illuvial layer enriched with iron and aluminium sesquioxides, Bv= illuvial layer containing plinthite and Cv=unchanged parent soil. The influx of Mo from the organic layer to the mineral soil layers was approximated to be from 6 to 60 g/ha/a of Mo in German acid forest soils, Cambisols and Podzols (Lang and Kaupenjohann 2000). These influxes correlated with concentration of Mo in current year Norway spruce needles. The uppermost mineral soil layers are more efficient molybdenum sources for trees and plants than the organic humus layer, due to the lower concentration of Mo in the solid phase. Furthermore, it was found that the Mo flux from the organic layer correlates with the nitrate outflow, indicating that the availability of nitrate influences the cycling of Mo. It was also postulated that nitrate may control the uptake of Mo to plants; Mo is needed in the nitrate reductase, and the plants’ demand of Mo is dictated by the nitrate availability in the soil solution (Lang and Kaupenjohann 2000).

5.5 Sorption on minerals

Goldberg et al. (1996) and Goldberg and Forster (1998) studied the sorption of molybdenum on iron oxides (hematite, goethite, amorphous iron oxide, poorly crystalline goethite) and aluminium oxides (gibbsite ( -Al2O3), -alumina ( -Al2O3), amorphous aluminium oxide). Sorption was found to be higher for poorly crystalline oxides on a weight basis (Goldberg et al. 1996). Sorption mechanism was specific sorption, i.e. ligand exchange with surface hydroxyl groups. The sorption maxima took place in acidic pH, which extended to approximately pH 4-5 (Goldberg et al. 1996; Goldberg and Forster 1998). In the pH-range 5-8 sorption decreased rapidly, and above pH 8 virtually no sorption took place. Surface-area corrected sorption (m2/g) decreased in the order hematite > goethite > poorly crystalline goethite > amorphous iron oxide (Goldberg et al. 1996). For aluminium oxides the order was gibbsite > amorphous

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aluminium oxide > -alumina. With increasing pH the sorption mechanism changed to outer sphere sorption for -alumina, gibbsite and goethite.

Bibak and Borggaard (1994) had similar result for synthetic aluminium oxide, goethite (FeO(OH)) and ferrihydrite (Fe2O3 x 0.5 H2O) in pH range 3.5-8. The maximum sorption also took place in pH 4-5, whereas in pH 8 the sorption decreased to a half of the maximum. The sorption mechanism was determined as specific sorption and proceeding through ligand exchange and leading to the formation of a binuclear surface complex.

The retention mechanisms for iron compounds are different: ferrous sulphate (FeSO4) precipitates solid ferrous MoO4

2- (FeMoO4) in low pH-values, whereas the presence of ferric nitrate (Fe(NO3)3 x 9H2O) or ferric sulphate (Fe2(SO4)3 x nH2O) lead to the formation of solid iron compounds (ferric oxohydroxide, FeO(OH)) which acts as sorbent for Mo. Barium may form low-solubility compound BaMoO4 (Ksp 10-7.45) if both ions are present in sufficient amounts. Molybdenum is not sorbed efficiently on montmorillonite or clinoptilolite. Sorption on TiO2 is dependent of pH, decreasing with increasing pH (Morrison and Spangler 1992).

Figure 15 presents the sorption per centages for molybdenum sorbed on goethite and kaolinite as a function of pH (Goldberg and Forster 1998). As can be seen, the sorption of molybdenum on goethite is considerably higher than on kaolinite throughout the studied pH-scale. For goethite, the sorption per centages remains in approximately 100 % in pH-range 2-8 in spite of the initial Mo concentration used. Above pH 8 the sorption of Mo on goethite remains higher in solutions containing 0.292 mol/m3 of Mo than in solutions with 1.04 mol/m3 of Mo; at pH 10 the sorption per centages are 52 and 20, respectively. For kaolinite the sorption per centages reaches their highest values below pH 5, and drop below 10 % at pH above 7.

Figure 16 presents the distribution coefficients of MoO42- on aluminium oxide ( -Al2O3)

and tin oxide (SnO2) as a function of pH (Bürck et al. 1988). Kd values of Mo on -Al2O3 increase rapidly from <10 ml/g to approximately 8000 ml/g as the pH increases from -0.8 to 1.5, and remain practically constant till pH 5. Sharp drop in Kd values is seen in pH 6-9. Above pH 9, hydroxyl ions (OH-) can replace sorbed MoO4

2-. On SnO2, Kd values are quite constant when pH is 1.5-3 (approximately 5000 ml/g), and decrease when pH decrease below 1.5 or increased above 3.

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Figure 15. The sorption per centages of Mo on goethite and kaolinite as a function of pH with two different initial molybdenum concentrations (Goldberg and Forster 1998). [Mo]i=0.292 or 1.04 mol/m3, suspension density for goethite and kaolinite 1.25 g/l and 200 g/l, respectively.

Figure 16. The distribution coefficients of MoO4

2- on aluminium oxide ( -Al2O3) and tin oxide (SnO2) as a function of pH in HNO3-solutions (Bürck et al. 1988). [MoO42-

]i=2.2x10-3 mol/l, contact time 20 h, V/m 8 ml/0.01-2 g, capacities 0.30 mmol/g of Mo for -Al2O3 and 0.19 mmol/g of Mo for SnO2,the specific surface areas 112.6 m2/g for -Al2O3 and 8.5 m2/g for SnO2. The sorption of Mo on iron oxides seems to take place in two stages: fast initial sorption is followed by slow sorption (Strauss et al. 1997). The slow sorption is proposed to occur due to the diffusion of Mo into the pores and interdomains of iron oxides, which can be called high affinity sorption sites (Lang and Kaupenjohann 2003; Strauss et al. 1997). Diffusion into the pores, and especially into the pores with of diameter <2 nm, is an immobilization process (Lang and Kaupenjohann 2003). In the micropores molybdenum is presumed to bound favourably as a polydentate complex. The fraction sorbed in the slow stage is not desorbed easily (diffusion from pores), unlike the

0

20

40

60

80

100

120

0 1 2 3 4 5 6 7 8 9 10 11 12

%so

rpti

on

pH

goethite, 0.292mol Mo/m3

goethite, 1.04mol Mo/m3

kaolinite, 0.292mol Mo/m3

kaolinite, 1.04mol Mo/m3

1

10

100

1000

10000

2 0 2 4 6 8 10

Kd(m

l/g)

pH

Al2O3

SnO2

40

fraction sorbed in the fast stage (outer surfaces) (Strauss et al. 1997). Thus, the desorption of Mo takes also place in two stages. Organic carbon coating on the surface of iron oxides affects the sorption of Mo by blocking the access to pores and interdomains (Lang and Kaupenjohann 2001, 2003). Lang and Kaupenjohann (2001, 2003) found that the desorbed fraction increased when iron oxides had carbon coating. For example, the desorbed amount of Mo from carbon coated iron oxide at time intervals 12, 24 and 48 hours were 48, 25 and 20 mg/g of Mo compared to 21, 4 and 4 mg/g of Mo for pure iron oxide. This was explained as the difference in the bonding of molybdenum: on carbon coated samples Mo forms complex with organic substances sorbed to the surface of iron oxide, whereas in pure iron oxide samples Mo is bound directly to iron via ligand exchange. Carbon coating depresses the access and sorption of Mo into pores, but increases the sorption onto the surface (Lang and Kaupenjohann 2003). At high Mo concentrations, the loss of sorption sites situated in the pores is not balanced by the increase in surface sorption sites.

Xu et al. (2003) investigated the sorption of molybdate (MoO42-) and tetrathiomolybdate

(MoS42-) onto pyrite (FeS2) and goethite (FeO(OH)) in anoxic conditions, and the effect

of pH and competing anions (sulphate, phosphate and silicate) on the sorption. Goethite had higher affinity towards molybdenum, and the affinity of MoS4

2- towards the solid phase was higher than the affinity of MoO4

2-. MoS42- was preferred over MoO4

2-; the sorption capacities of pyrite and goethite for MoS4

2- were 24.3 μmol/g and 377.7 μmol/g, respectively, while the corresponding values for MoO4

2- were 15.3 and 161.9 μmol/g. As an explanation for the high affinity of MoS4

2- to goethite, it was proposed that MoS4

2- cubane like structure (Mo-Fe-S; Figure 17) may be formed. It seemed that the structural S in the iron mineral adsorbent had smaller effect than S present in the sorbed ion.

Figure 18 presents the sorption of MoS42- and MoO4

2- on pyrite and goethite as a function of pH (Xu et al. 2003). As can be seen from Figure 18, sorption has high dependence on pH: maximum sorption takes place in pH<6, decreases in the range 6-8 and drops to below 30 % of the initial at pH>8. The effect of pH was attributed to the pH-dependent acidity of the surfaces of the sorbents.

Figure 17. The structural representations of MoO42- (a) and MoS4

2- (b) sorbed on FeS2 (Bostick et al. 2003). Black circle ( ) = Mo4+, red circle ( ) = O2-, grey circle ( ) = Fe2+ and turquoise circle ( ) = S2-.

41

Figure 18. The sorption percentages of molybdate (MoO42-) and tetrathiomolybdate

(MoS42-) onto pyrite (A) and goethite (B) as a function of pH (Xu et al. 2006). [Mo]i=50

μmol/l, [Fe(II)]=2 g/l for pyrite, [Fe(III)]=0.3 g/l for goethite, background electrolyte 0.1 mol/l NaCl.

Figure 19 presents the effect of competing anions sulphate, phosphate and silicate on the sorption of MoO4

2- and MoS42- onto goethite and pyrite as a function of pH (Xu et

al. 2003). Phosphate was efficient in competing for sorption sites with MoO42- and

MoS42- in the whole studied pH range (3-10). The effect of phosphate was higher for

MoO42- sorption onto goethite and pyrite than for MoS4

2-, because MoS42- formed

irreversibly retained stable complexes with goethite. The effect of sulphate and silicate on the sorption of the both Mo species was negligible, probably because of different sorption mechanism: outer-sphere sorption for SO4

2- and SiO42- compared with inner-

sphere sorption by Mo.

42

Figure 19. The sorption percentages of molybdate (MoO42-) onto pyrite (A) and

goethite (C), and tetrathiomolybdate (MoS42-) onto pyrite (B) and goethite (D) as a

function of pH in the presence of competing anions silicate, sulphate and phosphate (Xu et al. 2006). [Mo]i=50 μmol/l, [Fe(II)]=2 g/l for pyrite, [Fe(III)]=0.3 g/l for goethite, background electrolyte 0.1 mol/l NaCl, concentration of competing anions 0.5 mmol/l.

The sorption mechanism of MoO4

2- on kaolinite (H2Al2S2O8 x H2O) is specific sorption (Goldberg et al. 1996). Sorption on clay minerals kaolinite (Mikkonen and Tummavuori 1993a), illite, montmorillonite and calcite has maxima at pH 3, and decreases rapidly to zero as pH increases to 7 (Goldberg et al. 1996). On a weight basis, the sorption of molybdenum decreased in the order montmorillonite > illite > poorly crystallized kaolinite > well-crystallized kaolinite. The sorption of Mo on clay minerals is smaller than on aluminium and iron oxides. Arsenate (AsO4

-) and phosphate have only a little effect on the sorption of Mo on kaolinite, illite or montorillonite in concentrations equal or twice to that of Mo (Goldberg and Forster 1998).

Kd values for molybdenum collected from the literature and presented in the text are given in Table 5.

43

Table 5. Kd values for molybdenum presented in the text and collected from the literature.

Material Species Kd (ml/g)

Sample description/ other information

Country/ place

Reference

Organic Mo 23 (18-27) arithmetic mean (range) Vidal et al. 2009

Organic Mo 27 (10-74) arithmetic mean (range) IAEA 1994 Organic Mo 25 geometric mean value Sheppard and

Thibault 1990 Organic soil Mo 18-27 min/max value Gil-García et

al. 2009

Peat Mo 1200 native elements measured with ICP

Forsmark, Sweden

Sheppard et al. 2009

Peat Mo 6800 native elements measured with ICP

Laxemar-Simpevarp, Sweden

Sheppard et al. 2009

Clay Mo 90 geometric mean value Sheppard and Thibault 1990

Clay Mo 90 arithmetic mean Vidal et al. 2009

Clay soil Mo 90 Gil-García et al. 2009

Clay soil Mo 270 / 18 pH 5 / 7 Mikkonen and Tummavuori 1993b

Clay gyttja Mo 1200 native elements measured with ICP

Laxemar-Simpevarp, Sweden

Sheppard et al. 2009

Clay gyttja Mo 3300 native elements measured with ICP

Laxemar-Simpevarp, Sweden

Sheppard et al. 2009

Clayey till Mo 70 native elements measured with ICP

Forsmark, Sweden

Sheppard et al. 2009

Clayey till Mo 240 native elements measured with ICP

Forsmark, Sweden

Sheppard et al. 2009

Loam Mo 125 geometric mean value Sheppard and Thibault 1990

Loam Mo 130 arithmetic mean Vidal et al. 2009

Loam soil Mo 130 Gil-García et al. 2009

Sand Mo 50 (7.4-82) arithmetic mean (range) Vidal et al. 2009

Sand Mo 74 (0.82-67)

arithmetic mean (range) IAEA 1994

Sand Mo 10 geometric mean value Sheppard and Thibault 1990

Sand soil Mo 7-82 min-max value Gil-García et al. 2009

44

Table 5. continued.

Material Species Kd (ml/g)

Sample description/ other information

Country/ place

Reference

Sandy till Mo 410 native elements measured with ICP

Laxemar-Simpevarp, Sweden

Sheppard et al. 2009

Soil MoO42- 120 Millhopper soil; loamy,

siliceous oxalate extracted Fe and Al 5.8 g/kg, pH 5.0

Paleudult, Florida, USA

Brinton and O’Connor 2003

Soil MoO42- 21 Immokalee soil; sandy,

siliceous, extracted Fe and Al 0.12 g/kg, pH 5.4

Alaquods, Florida, USA oxalate

Brinton and O’Connor 2003

Soil + biosolid MoO42- 1.3 Pina loam; fine-silty, soil

pH 8.10 Torrifluvent, Arizona, USA sample from 0-15 cm depth,

Carroll et al. 2006

Red earth MoO42-

2.9x103

82 pH 6 a coastal sandy soil from China

Zuyi et al. 2000

Calcareous soil MoO42- 109 pH 8 cultivated soil

from China, depth 0-20 cm

Zuyi et al. 2000

cetylpyridinium bentonite

MoO4- 4/6.14/19 temperature 27/40/50 °C;

sorption capacity for Mo 1.4 mmol/g; surfactant sorbed 0.83 mmol/g

Atia 2008

Alumina MoO42- Chromatographic

alumina, pH 6 Zuyi et al.

2000

45

6 NIOBIUM

Niobium is a transition metal belonging to the fifth group in the periodic table of elements together with vanadium and tantalum. The atomic number of Nb is 41 and its electron configuration is written as [Kr]4d55s1. The atomic and covalent radius of Nb are 135.3 pm and 137 pm, respectively (Suresh and Koga 2001). The ionic radius of Nb(V+) is 64 pm (Salminen 2007). The geochemical behaviour and chemical and physical properties of Nb are similar to those of Ta due to their identical charge and similar size (Barth et al. 2000).

Niobium has only one stable isotope (Nb-93), but over 30 radioactive isotopes. The most important radioisotopes of niobium in the spent nuclear fuel are Nb-93m and Nb-94, which are classified as high priority radionuclides in the long-term safety assessment of spent nuclear fuel (Haapanen et al. 2009; Hjerpe et al. 2010). The physical half-lives of these nuclides are 16.13 and 2.03x104 years, respectively (Firestone et al. 1998).

Nb-93 m and Nb-94 are fission products formed in the fission of uranium and plutonium, but also activation products formed in the neutron activation of stable Nb-93 by the reactions Nb-93(n,n)Nb-93m and Nb-93(n, )Nb-94. Stable Nb-93 is found as an impurity in the cladding of nuclear fuel and structural components of nuclear reaction pressure vessels. Zircaloy alloy contains niobium, as small amounts of stable Nb-93 is added to metal alloys containing nickel as the main component to improve the strength in high temperatures and to inhibit the intercrystalline corrosion (Andersson et al. 1979).

Besides being a fission and activation product, Nb-93m is also formed in the decay of long-lived Zr-93 and Mo-93. The physical half-life of Zr-93 is 1.53x106 years and it decays by 100 % - -decay to Nb-93m. On the other hand, 88 % of Mo-93 decays by EC to Nb-93m. The activation reaction is the major source for Nb-94 and decay of Zr-93 and Mo-93 for Nb-93m due to the low fission yields of Nb-93m and Nb-94 in the fission of U-235 and Pu-239 (1.8x10-11 and 2.3x10-10 % for Nb-93m and 9.4x10-4 and 2.5x10-7 % for Nb-94, respesctively). Nb-94 is a - emitter, with a maximum - energy of 471.68 keV (Firestone et al. 1998).

The solubility of Nb-94 from spent nuclear fuel in water is expected to be 10-5 mol/m3 (Bengtsson and Widén 1991). The maximum instant release rate of niobium through the bentonite barrier was calculated to be 3.92x10-11 mol/a, which corresponded to a release time of 2.17x105 a. The effective diffusivity, De, of Nb-94 in compacted bentonite was proposed to be 3.2x10-3 m2/a (Brandberg and Skagius 1991).

Niobium is a lithophilic element and thus enriched in the lithosphere (Wiberg et al. 2001). The average Nb content in the upper continental crust has proposed to be 11.5 mg/kg, which gives the average concentration of 8 mg/kg in the whole crust (Barth et al. 2000). In the vicinity of mineralization sites 24 mg/kg (Tyutina et al. 1958). In metal ores, the Nb content may be as high as 47.93 % (Culombite ore) or 52.87 % (Edgarite) (Abbasi 1988; Barkov et al. 2000; Tyutina et al. 1958). Furthermore, certain minerals, such as pyroxenes, amphiboles, micas, ilmenite and magnetite, contain niobium at trace levels (Tyutina et al. 1958). The chemical form on Nb in minerals is proposed to be niobite (Fe(II)Nb2O6) or Nb2O5 based complex oxides (Andersson et al. 1979).

46

The Europien median value for niobium content in aqua regia extracted topsoils and subsoils are 9.68 and 9.76 mg/kg, respectively (Salminen 2007). In Finland, Nb content in the topsoils varies between 4.90-9.68 mg/kg and 5-10 mg/kg in subsoil (Salminen 2007). Niobium content in aqua regia extracted brackish-water and brackish-water sediments in the Archipelago Sea were 0.25-0.53 mg/kg and 1.75-4.20 mg/kg, respectively (Åström et al. 2008). In stream waters the European median Nb consentration is 0.0044 μg/l, whereas in Finland the Nb consentration varies between 0.04 and 1.01 μg/l (Salminen 2007). The respective values for aqua regia extracted stream sediments are 13 mg/kg in Europe and 5-26 mg/kg in Finland (Salminen 2007).

The mobility of niobium in soils is expected to be low due to the low solubility of niobium compounds, the formation of sparingly soluble complexes and salts and high sorption on rocks and clays (Andersson et al. 1979; Baston et al. 1992; Baker et al. 1994; Behrens et al. 1982; Berry et al. 1988; Legoux et al. 1992).

6.1 The influence of soil redox potential and pH

Niobium can exist on oxidation states from +V to -I. The most stable oxidation states in groundwater are +V and +III (Andersson et al. 1979), and the prevailing oxidation state in soils is reported being Nb(V+) (Rhodes 1957). The chemical form of niobium in acidic solutions is niobyl oxocation, NbO3+ (Lehto and Hou 2010). Niobium acid, HNbO3, is the principal form met in somewhat acidic and neutral solutions. In basic solutions, niobium hydrolyses and forms anionic species such as NbO4

3- or NbO3-

(Andersson et al. 1979; Brandberg and Skagius 1991; Lehto and Hou 2010). Niobium forms sparingly soluble compounds with alkali and alkali earth metal cations, such as NaNbO3 and Ca(NbO3)2 (Andersson et al. 1979). The Eh-pH-diagram of niobium is presented in Figure 20.

47

Figure 20. The Eh-pH diagram of niobium (Takeno 2005). Nb=10-10, 298.15 K and 105 Pa.

Åström et al. (2008) examined the niobium consentration in stream water samples taken from 26 countries in Europe, including Finland. Samples were also taken from brackish coastal waters, streams, shallow lakes and overburden groundwater tubes situated in Simpevarp and Forsmark areas in Sweden. In filtered (0.45 μm) European stream waters the dissolved niobium consentration was found to correlate positively with dissolved organic carbon (DOC) and iron. The correlation with DOC was explained as the direct bonding of niobium with humic substances, which have negative charge in the natural pH-values (4.5-6.2) in the studied waters. Furthermore, the correlation with iron was suggested occuring due to association with colloidal iron oxyhydroxides. Colloidal iron oxyhydroxides can be mixed with organic substances. Dissolved niobium was found to be solely in the anionic form. The median Nb consentration in European stream waters was 0.004 ppb (Salminen 2007).

In the Simpevarp and Forsmark water samples, Nb consentration was nearly an order of magnitude higher in the overburden groundwater samples (0.138 ppb) than in samples taken from local coastal waters, streams or lakes (0.011-0.044 ppb) (Åström et al. 2008). Interestingly, in surface and groundwater samples niobium was found to correlate solely with dissolved iron. Furthermore, with the same iron concentration the groundwater and the surface water samples, Nb concentration was always higher in the groundwater.

48

6.2 Sorption on soils

The sorption of niobium on soils has received little attention. Compared to the number of publications published on the retention of molybdenum and selenium on soils, only a few articles concerning niobium have been published. These articles have mainly focused on the soil-to-plant transfer of niobium (Echevarria et al. 2005; Gerzabek et al. 1994), or on the sorption of niobium on clays, sands and gravel (Baston et al. 1992; Baker et al. 1994; Behrens et al. 1982; Berry et al. 1988; Legoux et al. 1992). Distribution coefficients for niobium in soils have also been measured via the concentration of stable Nb-93 in soil and water samples (Sheppard et al. 2009).

Niobium is relatively immobile in soils due to its enhanced sorption on soils mineral particles (Echeverria et al. 2005). High organic matter and clay fraction content increase the sorption of niobium in soils (Baston et al. 1992, Echeverria et al. 2005; Gerzabek et al. 1994; Sheppard and Thibault 1990). Niobium is sorbed effectively on clay, and moderately or strongly on sand and gravel (Baston et al. 1992; Behrens et al. 1982). Also the texture of the soil affects the sorption: the predicted Kd values decrease in the order clay>loam>sand (Sheppard et al. 1990). Sorption of niobium shows dependence on pH in alkaline solutions (pH>8), decreasing with increasing pH (Andersson et al. 1979; Baston et al. 1992).

Behrens et al. (1982) investigated the sorption of Nb-95 on sand samples taken from fluviatile and aeolian sand aquifers in Germany. The materials were composed of clayey sand, sandy and gravelly material. The experiments were conducted with artificial groundwater at pH-range 4 to 7, room temperature and aerobic conditions. The Kd values of Nb ranged from 10 to 3600 ml/g. The retardation factors, Rfs, gained from the column tests ranged from 390 to 9000, indicating the low diffusion of Nb in the soil matrix. The definition of the retardation factor is given in Equation 6.

)(1)( xKxR df [6]

Equation 6. The definition of retardation factor Rf for element x. = the bulk density of the material, = the porosity of the material and Kd = distribution coefficient for element x (Behrens et al. 1982).

Charles and Prime (1983) examined the desorption behavior of niobium from estuarine silts contaminated with radioactive pollution derived from Windscale reprocessing plant in the UK. Approximately 46 %, 75 %, 35 % and 4 % of the initial Nb-95 fraction was desorbed with 0.5 M oxalic acid (C2O4H2), 0.5 M citric acid (C6H8O7), concentrated hydrochloric acid (HCl) and concentrated nitric acid (HNO3), respectively. Desorption with oxalic and citric acids indicated that niobium was present as precipitated complex compounds on the surface of the silt samples. The complex compounds were held strongly on the surfaces, but the complexing force of the large oxalic and citric molecules was stronger. Deionized water (H2O), 1.0 M potassium chloride (KCl) or sodium hydroxide (NaOH) did not desorb any Nb-95 from silt samples.

Andersson et al. (1979) studied the sorption of niobium on Swedish granite samples both in oxic and anoxic conditions. They found that Kd decreased from 10 000 ml/g to 100 ml/g as pH increased from 8.0 to 10.5. In the anoxic conditions, Kd was 1200 ml/g.

49

The sorption of niobium on suspended particles and colloids is significant, which is often seen as a decrease in the activity of the liquid phase upon centrifugation and filtration with 0.45 μm or 0.03 μm (30000 molecular weight cut off filters) (Baker et al. 1994; Baston et al. 1992; Lehto and Hou 2010).

Baker et al. (1994) studied the sorption of Nb-95 in Britain onto Nirex vault backfill material consisting of OPC (ordinary Portland cement), lime and limestone flour grout. At pH 12.2 and saline solutions, Kd of Nb-95 was found to range from 35 000 ml/g to 41 000 ml/g.

Baston et al. (1992) had similar results with Baker et al. (1994) with their studies on the sorption of Nb-95 onto London clay and Dungeness aquifer gravel. In the experiments done with clay, Kd values ranged from 2100 to over 10 000 ml/g when solution was filtered with 0.45 μm and 0.03 μm, respectively. In the tests where Kd was determined without filtering the solution, lower values were noted (240-380 ml/g). The results indicated that Nb was sorbed effectively on small particulate (suspended particles) and colloidal matter. Baston et al. (1992) suggested that the main factor affecting the migration of niobium in soils might be the migration mobility of the fine suspended (clay) particles containing niobium on their surfaces, instead of sorption on to solid, stationary phase. In the case of the gravel samples, Kd values increased as the separation method was changed from centrifugation (60-3800 ml/g) to filtration with 0.45 μm filter (>104 ml/g). The sorption of Nb on gravel samples was found to increase slightly as the sand content of the gravel samples increased.

The soil-to-plant transfer of niobium is low (Echeverria et al. 2005; Gerzabek et al. 1994; Tyutina et al. 1958). The transport of element from soil to plant is usually described with transfer factor (TF) or concentration ratio (CR). The definition of these two terms is identical and it is given in Equation 7.

[7]

Equation 7. The definition of the transfer factor. [Nb]= the concentration of Nb (Bq/ g of dry weight) (Gerzabek et al. 1994).

Gerbazek et al. (1994) found that lower CR values of niobium in bean shoot and greenrape were seen in Austrian soil with higher pH; 0.0014-0.0069 in pH 7.5 and 0.0025-0.013 in pH 5.8, respectively. It was proposed that the chelating ability of organic matter towards niobium is enhanced in higher pH values thus making Nb less bioavailable to plants (Gerzabek et al. 1994). It was also suggested that the role of organic matter in the fixation of Nb in soils may be substantial.

Echeverria et al. (2005) determined the CR values for niobium in roots and aerial parts of rye grass and winter wheat grown in French Cambisols. Relatively high CR-values found for roots (0.32-1.51) indicated the active transport of niobium from soil solution to roots, but pore translocation to the aerial parts of the plants (CR 1.0x10-4 – 4.0x10-4).

Kd values for niobium collected from the literature and presented in the text are given in Table 6.

50

Table 6. Kd values for niobium presented in the text and collected from the literature.

Material Species Kd (ml/g) Sample description/ other information

Country/ place

Reference

Sand Nb 170 (160-190) arithmetic mean (range)

Vidal et al. 2009

Organic Nb 2000 arithmetic mean Vidal et al. 2009

Organic Nb 2000 geometric mean value Sheppard and Thibault 1990

Organic soil Nb 2000 Gil-García et al. 2009

Peat Nb 14000

native element measured with ICP

Laxemar-Simpevarp, Sweden

Sheppard et al. 2009

Eutric Cambisol

Nb(V+) 1978 / 4524 Ap- horizon , equilibrium time 3 /30 days

Samples taken from north-eastern France

Echeverria et al. 2005

Eutric Cambisol

Nb(V+) 3228 / 8370 Ap- horizon , equilibrium time 3 /30 days

Samples taken from north-eastern France

Echeverria et al. 2005

Clay Nb 3000 / 2400 (900-4700)

arithmetic mean / geometric mean (range)

Vidal et al. 2009

Clay Nb 900 IAEA 1994 Clay Nb 900 geometric mean value Sheppard and

Thibault 1990

Clay soil Nb 2400 (900-4729)

geometric mean (min/max) value

Gil-García et al. 2009

London clay Nb 240-380 / >104 / >104

Centrifugation – no filtration / filtration 0.45 μm / filtration 30000 MWCO (0.03 μm), pH 8.7-9.0, initial 95Nb conc. 2x10-11 M, N2- atmosphere, clay water simulant

Baston et al. 1992

London clay Nb 630 (pH 8.2), 163 (pH 8.3), 104 (pH 8.8), 7200 (pH 9.0) / >104 / >104

centrifugation – no filtration / filtration 0.45 μm / filtration 30000 MWCO (0.03 μm), pH 8.2-9.0, initial 95Nb conc. 5x10-11 M, N2- atmosphere, clay water simulant

Baston et al. 1992

London clay Nb 1323 sorptivity ( Kd) >6000

Berry et al. 1988

Clay gyttja Nb 12000 native element measured with ICP

Laxemar-Simpevarp, Sweden

Sheppard et al. 2009

51

Table 6. continued.

Material Species Kd (ml/g) Sample description/ other information

Country/ place

Reference

Clayey till Nb 36000 native element measured with ICP

Forsmark, Sweden

Sheppard et al. 2009

Loam Nb 3600 / 2500 (540-8400)

arithmetic mean / geometric mean (range)

Vidal et al. 2009

Loam Nb 540 IAEA 1994 Loam Nb 550 geometric mean

value Sheppard and

Thibault 1990 Loam soil Nb 2500 (540-

8370) geometric mean (min/max) value

Gil-García et al. 2009

Sand Nb 160 geometric mean value

Sheppard and Thibault 1990

Sand Nb 160 IAEA 1994 Sand soil Nb 160-187 min/max value Gil-García et

al. 2009 Sandy till Nb 1700 native element

measured with ICP Laxemar-Simpevarp, Sweden

Sheppard et al. 2009

Calcareous sandy soil

Nb >1957 Kd measured from a mixture of Zr and Nb

Rhodes 1957

Soil Nb 2600 [Nb]i=10-10 M, pH 5.9, sample depth 7.8-7.9 m, specific surface area 8.6 m2/g

Leqoux et al. 1992

Soil Nb 1500 [Nb]i=10-10 M, pH 6.6, sample depth 9.65-9.85 m, specific surface area 12.2 m2/g

Leqoux et al. 1992

Soil Nb 1700 [Nb]i=10-10 M, pH 8.0, sample depth 11.1-11.25 m, specific surface area 14.7 m2/g

Leqoux et al. 1992

Soil Nb 2100 sample depth 12.0-12.2 m, specific surface area 6.6 m2/g

Leqoux et al. 1992

Luvisol Nb(V+) 2512 / 4729 Ap- horizon , equilibrium time 3 /30 days

Samples taken from north-eastern France

Echeverria et al. 2005

Calcareous soil

Nb(OH)6-,

Nb(OH)5 2.98x104 pH 8 cultivated soil

from China, depth 0-20 cm

Zuyi et al. 2000

Red earth Nb(OH)6-,

Nb(OH)5 187 pH 6 a coastal sandy

soil from China

Zuyi et al. 2000

52

Table 6. continued.

Material Species Kd (ml/g) Sample description/ other information

Country/ place

Reference

Dungeness aquifer gravel

Nb 240-470 / 104

Centrifugation – no filtration / filtration 0.45 μm, aerobic conditions, Grain size <2 mm, sample depth 7.5 m, pH ~8, initial 95Nb conc. 1x10-11 M, rock water simulant

Baston et al. 1992

Dungeness aquifer gravel

Nb 260-280 / 5900- 104

Centrifugation – no filtration / filtration 0.45 μm, aerobic conditions, Grain size <2 mm, sample depth 7.5 m, pH ~8, initial 95Nb conc. 1x10-11 M, rock water simulant

Baston et al. 1992

Dungeness aquifer gravel

Nb 910-2800 / 7700- >104

Centrifugation – no filtration / filtration 0.45 μm, aerobic conditions, grain size <2 mm, sample depth 15 m, pH ~8, initial 95Nb conc. 3x10-11 M, rock water simulant

Baston et al. 1992

Dungeness aquifer gravel

Nb 2600- >104 / >104

Centrifugation – no filtration / filtration 0.45 μm, aerobic conditions, grain size <2 mm, sample depth 15 m, pH ~8, initial 95Nb conc. 3x10-11 M, rock water simulant

Baston et al. 1992

Granite Nb 10000 Initial pH 8.0, equilibrium time 6 days

Stripa granite (Sweden), oxic conditions

Andersson et al. 1979

Granite Nb 100 Initial pH 10.5, equilibrium time 6 days

Stripa granite (Sweden), oxic conditions

Andersson et al. 1979

Granite Nb 4000 Initial pH 8.0, equilibrium time 6 days

Finnsjö granite (Sweden), oxic conditions

Andersson et al. 1979

Granite Nb 100 Initial pH 10.5, equilibrium time 6 days

Finnsjö granite (Sweden), oxic conditions

Andersson et al. 1979

Granite Nb 1200 Initial pH 8.0, equilibrium time 20 hours

Stripa granite (Sweden), reducing conditions

Andersson et al. 1979

53

Table 6. continued.

Material Species Kd (ml/g) Sample description/ other information

Country/ place

Reference

Cementitous backfill material

Nb 35000 / 41000

Saline groundwater simulant, pH ~12, N2- atmosphere , Comparison values for fuel ash/Portland cement mix 500-80000 ml/g

Baker et al. 1994

Bentonite Nb 180- 510 Synthetic groundwater, ionic strength 55 mM, pH 8.3-8.7, liquid/solid ratio 10 ml/g

Brandberg and Skagius 1991

Bentonite Nb 510- 1700 Synthetic groundwater, ionic strength 55 mM, pH 8.3-8.7, liquid/solid ratio 20 ml/g

Brandberg and Skagius 1991

Bentonite Nb 200 Fresh water, oxidizing or reducing conditions

Brandberg and Skagius 1991

Bentonite Nb 200 Saline water, oxidizing or reducing conditions

Brandberg and Skagius 1991

Alumina Nb(OH)6-,

Nb(OH)5 2.98x105 Chromatographic

alumina, pH 6 Zuyi et al.

2000

54

55

7 SELENIUM

Selenium is a non-metal belonging to the 16th group in the periodic table of elements together with oxygen, sulfur and tellurium. The chemistry of selenium resembles to that of sulphur and Se can replace S in its inorganic and organic compounds (Pezzarossa and Petruzzelli 2001). The atomic number of Se is 34 and its electron configuration is written as [Ar]3d104s24p4. The covalent radius of Se is equal to the atomic radius (120 pm), whilst the van der Waals radius of Se is 190 pm. The ionic radius of selenate (SeO4

2-) and selenite (SeO32-) are 235 and 225 pm (Huheey et al. 1993).

Selenium has 5 stable and more than 20 radioactive isotopes. The most long-lived isotope of selenium, Se-82, is usually considered as stable due to its long physical half-life of 1.08 x 1020 a (Firestone et al. 1998). The stable isotopes are Se-74, Se-76, Se-77, Se-78, Se-80 and Se-82, for which the isotopic abundances are 0.87 %, 9.0 %, 7.6 %, 23.5 %, 49.8 % and 9.2 % (Shaw and Ashworth 2010). The most important radioisotope of selenium in the spent nuclear is Se-79. It is an activation product formed in the neutron activation of stable Se-78 by the reaction Se-78(n, )Se-79, but also a fission product. Stable Se-78 is found as an impurity in the reactor construction materials and nuclear fuel (Lehto and Hou 2010). Se-79 has physical half-live of 3.77x105 years (Shaw and Ashworth 2010). It is a - emitter, with a maximum - energy of 151 kev (Firestone et al. 1998). In the long-term safety analysis of spent nuclear fuel, Se-79 is classified as medium priority radionuclide (Haapanen et al. 2009; Hjerpe et al. 2010).

Selenium is a chalcophilic element, which easily replaces sulphur in sulphide minerals, such as pyrite (FeS2), chalcopyrite (CuFeS2), pyrrhotite (Fe1-xS) and sphalerite ((Zn,Fe)S) (Salminen 2007). Other less common selenium minerals are known, these including crookesite ((Cu,Tl,Ag)2Se), berzelianite (Cu2Se) and tiemannite (HgSe).

The average Se content in the continental crust has estimated to be 0.09 mg/kg, and 0.13 mg/kg in the bulk crust (Rudnick and Gao 2004). The average selenium content in topsoils is proposed to be 0.33 mg/kg on a world-wide basis (Kabata-Pendias 2001). In stream waters the European median Se content is 0.34 μg/l, whereas in Finland the Se content varies between 0.05 and 0.34 μg/l (Salminen 2007).

The sorption of selenium in soils increases with increasing organic matter content (Choppin et al. 2009; Dhillon and Dhillon 1999; Dhillon et al. 2010; Gustafsson and Johnsson 1992; Pezzarossa et al. 1999; Pezzarossa and Petruzelli 2001; Sharmasarkar and Vance 2002; Yläranta 1983; Zawilanski and Zavarin 1996), increasing clay fraction content (Ashworth et al. 2008; Keskinen et al. 2009; Vuori et al. 1989; Yläranta 1983), increasing amount of CaCO3 in compacted soils (Pezzarossa and Petruzelli 2001; Wang et al. 1999; Wang and Liu 2005), increased weathering of the soil material (Neal et al. 1987a; Vuori et al. 1989), increasing iron and aluminium oxyhydroxide content (Choppin et al. 2009; Dhillon and Dhillon 1999; Gustafsson and Johnsson 1992; Harada and Takahashi 2009; Keskinen et al. 2009 Vuori et al. 1989), decreasing pH (Dhillon and Dhillon 1999; Frost and Griffin 1977; Goldberg and Glaubig 1988; Nakamaru et al. 2005; Neal et al. 1987a; Pezzarossa et al. 1999; Pezzarossa and Petruzelli 2001; Saeki and Matsumoto 1994) and increasing concentration of Ca2+ in the solution (Neal et al. 1987b).

56

The sorption of selenium decreases as the presence of competing anions, such as phosphate, arsenate, carbonate and sulphate increases (Dhillon and Dhillon 1999; Fujikawa abd Fukui 1997; Goldberg and Glaubig 1988; Nakamaru and Sekine 2008; Neal et al. 1987b; Pezzarossa et al. 1999; Saeki and Matsumoto 1994; Yläranta 1983) and increasing pH (Dhillon and Dhillon 2003; Vuori et al. 1994). Also, in some instances the retention of selenium in soils is reported to decrease with increasing clay fraction and CaCO3 content (Dhillon and Dhillon 1999).

The research conducted regarding the sorption of selenium in soils have to a lesser extend concerned in the safety assessment of spent nuclear fuel, and especially the retention and migration of radio-selenium in soils.

Selenium has notably gained interest in soil science because of its essential nature for humans (trace nutrient), narrow range between indeficient and toxic dose, and low or unusually high concentration in soils. Due to these reasons the sorption of selenium in soils and distribution between different fractions (adsorbed, organically bound, residual, etc.) has grown attention. In Finland, the natural plant-available content of selenium in soils is low (Koljonen 1975), which leads to especially low selenium concentration in domestic agricultural products (Koivistoinen 1980). Selenium-containing plant multinutrient fertilizers have been used since 1984 to improve the selenium concentration in agricultural products and to stabilize the daily intake into sufficient level (Ministry of Agriculture and Forestry 1984). Before selenium additions were started, the average Se intake was 0.02-0.03 mg/day and has now steadied to approximately 0.065 mg/day (Eurola et al. 2008; Mutanen 1984; Varo and Koivistoinen 1980). Only recently selenium has been recognized as an important element in the safety assessment of spent nuclear fuel, and especially the retention and migration of radio-selenium in soils has received notice.

7.1 The influence of soil redox potential and pH

7.1.1 The speciation of selenium

Selenium is a redox-sensitive element and can form inorganic and organic compounds in oxidation states ranging from -II to +VI (Sato and Miyamoto 2004). The speciation of selenium depends on the prevailing Eh-pH conditions. Inorganic forms of Se include selenate (SeO4

2-), selenite (SeO32-), elemental selenium (Se(0)) and selenide (Se(-II)).

The oxidation state of selenium in selenate and selenite are +VI and +IV, respectively. These forms are considered as the most mobile and biogeochemically significant, selenate more mobile than selenite (Masschelyen et al. 1991; Pyrzy ska 1998). In soils and solutions, selenate is the dominating species in highly oxiziding and pH neutral conditions, even though SeO4

2- can be the main species found in reducing conditions (Harada and Takahashi 2009; Shaw and Ashworth 2010). As the redox potential decreases, different forms of selenious acid (H2SeO3) become prevailing: seleniuos acid (H2SeO3) at pH<3, biselenite (HSeO3

-) at 3<pH<7 and selenite (SeO32-) at pH>7. The

transformation between selenate and selenite at pH 7 takes place in redox potential of approximately +440 mV. At pH 7, biselenite converts into elemental selenium (Se(0)) when the redox potential is approximately +100 mV. The most reduced form of selenium, selenide (Se(-II)), is formed by the reduction of elemental selenium when the redox potential drops below -100 mV at pH 7. On the other hand, in highly basic

57

solutions (pH 11-13) selenite can be reduced directly into selenide when the redox potential reaches -250 - -400 mV. Selenide is also a common intermediate in the synthesis of selenoenzymes within plants, or in the transformation into extractable organometabolites (Birringer et al. 2002; Ip 1998; Suzuki 2005). Figure 21 presents the Eh-pH diagram of selenium.

Figure 21. The Eh-pH diagram of selenium (Takeno 2005). Se=10-10, 298.15 K and 105 Pa.

The migration of selenium in soils is greatly affected by the soil redox potential; in anoxic conditions and negative Eh-potentials (<-100 mV) the migration of selenium is inhibited and it is strongly immobilized in the solid phase because reduced forms (Se(0), Se(-II)) are formed (Ashworth and Shaw 2006). The solubility, mobility and bioavailability of the inorganic Se-species decrese with decreasing oxidation state: selenate (+VI) > selenite (+IV) > elemental selenium (0) > selenide (-II). Selenide and elemental selenium are usually found in the solid phase.

Selenium can be present in three or four different oxidation states in relatively shallow depths (0-40 cm) in soils, namely selenide, elemental selenium, SeO3

2- and SeO42-

(Harada and Takahashi 2009; Ryser et al. 2006). The oxidation-reduction reactions of selenium in soils can go through microbiological or abiotoc route. Abiotic reduction of SeO4

2- can occur in the presence of green rust, i.e. an analogy to the natural mineral fougerite ((Fe(II),Mg)2Fe(II)2(OH)18 x 4H2O) (Myneni et al. 1997). At pH>4 reduction

58

can occur as homogenous aqueous phase reactions, or on the surface (adsorption and reduction) or in the interlayer spaces (co-precipitation and reduction) of green rust. Reduction in the interlayer spaces is faster compared with the surface reactions and shows less dependence on solution pH. In interlayer reactions, SeO4

2- is first reduced to SeO3

2-, which follows further reduction to elemental Se. In surface reaction, SeO42- is

reduced directly to Se(0) without formation of measureable amounts of SeO32- as an

intermediate product. Green rust is oxidized to magnetite (FeO x Fe2O3) in the reaction. On the other hand, Mn(IV) and to a lesser extent Fe(III) containing minerals may induce the oxidation of Se(IV+) to Se(VI+) (Fujikawa and Fukui 1997; Harada and Takahashi 2009; Myneni 1997). For example, Fujikawa and Fukui (1997) found that the sorption of SeO3

2- onto rocks containing minerals hematite and magnetite was followed by the increase in SeO4

2- concentration in the solution upon oxidation of the sorbed SeO32-.

The increased SeO42- concentration was explained as the oxidative sorption of selenite:

[8]

The influence of micro-organisms on the reduction and oxidation of selenium is discussed in Chapter 7.2.

The transformations between the reduced and oxidized forms of selenium in soils are usually relatively slow (Zawilanski and Zavarin 1996). For example, in Kesterson Reservoir soils (USA) contaminated with high levels of selenium, the major fraction of the initial reduced selenium was unaffected after conditions had changed from reducing into oxidizing five years earlier. In a 2.5-year incubation experiments, the calculated oxidation rates for organic and refractory selenium ranged between 0.058-0.29 a-1 and 0.11-2.4 a-1, respectively.

The factors affecting the concentration ratio between selenate and selenite in the solution include the presence of acids affecting the pH, suspended material, complexing agents and dissolved gases affecting redox potential (mainly oxygen), sorptive surfaces and mineralogy (Dhillon and Dhillon 1999; Pyrzy ska 1998). In Table 7 is given the percentages of SeO4

2-, SeO32- and selenium bound to humic substances in Finnish soil

infiltration water, groundwater, rain and snow samples (Alfthan et al. 1995). SeO32- was

the main component in rain water, and SeO42- dominated in groundwater and snow.

Selenium bound to organic humic substances was the major species in soil infiltration water due to the retention on organic matter in the humus layer, and further leaching deeper to the ground with percolating soil water (Alfthan et al. 1995).

Table 7. The percentages of SeO42-, SeO3

2- and selenium bound to humic substances in Finnish soil infiltration water, groundwater, rain and snow samples (Alfthan et al. 1995).

Species Soil infiltration water

Groundwater Rain Snow

SeO42- (%) 25.7 66.7 33.8 56.3

SeO32- (%) 15.1 3.3 60.5 21.2

bound to humic substances (%)

36.1 6.1 n.a. n.a.

n.a. = not analyzed

59

7.1.2 Sorption behaviour of selenite and selenite

The dissimilar nature and properties of chemical bonds formed in the sorption of SeO42-

and SeO32- can be caused by the differences in their structure (Cotton et al. 1999;

Sharmasarkar and Vance 2002); selenate has tetrahedral and selenite distorted pyramidal structure (Figure 22). Selenite has one lone pair of electrons (LP; two dots above Se (selenite) in Figure 22) in its valence shell and three bond pairs (BP; solid lines between oxygen and selenium in selenite), whereas selenate has only bond pairs (Cotton et al. 1999). The typical sorption mechanism of SeO3

2- is inner-sphere complexation (specific sorption; ligand exchange), whereas SeO4

2- mainly forms outer-sphere complexes (Harada and Takahashi 2009).

Figure 22. The distorted pyramidal structure of selenite and tetrahedral strucuture of selenate (modified after Cotton et al. 1999).

The proposed specific sorption mechanisms, i.e. ligand exchange reactions of SeO32- are

(Saeki and Matsumoto 1994):

[9]

[10]

[11]

SeO32- is sorbed on minerals, soils and sediments much more efficiently than SeO4

2- (Collins et al. 2006; Duc et al. 2003; Goldberger and Glaubig 1988; Sharmasarkar and Vance 2002). Retention of these two species is dependent on pH and decreases with increasing pH; the maximum sorption of selenium takes place in acidic solutions (Frost and Griffin 1977; Fujikawa and Fukui 1997; Goldberg and Glaubig 1988; Neal et al. 1987a). For example, considerable amounts of SeO3

2- was sorbed in pH 2-9 on a sandy loam sample, while SeO4

2- was sorbed only pH<3 (Ahlrichs and Hossner 1987). The effect of pH on the solid phase has been described in Chapter 3.2.

In Finland, Yläranta (1983) established that no sorption of SeO42- took place on clay,

fine sand and peat samples while 85.7 %, 53 % and 43 % of the initial SeO32- (3x10-5

M) was sorbed on clay soil, fine sand soil and Carex peat, respectively. As the initial concentration of SeO3

2- was increased to 3x10-4 M, sorption percentages decreased to 76.7 %, 34 % and 39 % SeO3

2- for clay soil, fine sand soil and Carex peat.

60

7.2 The influence of micro-organisms

Micro-organisms have an important role in the biogeochemistry, speciation and distribution of selenium in soils (Darcheville et al. 2008). Many bacteria are capable of reducing and/or oxidizing selenium species. Reducers are, for example, Wolinella succinogenes (Tomei et al. 1992), Pseudomonas stutzeri (Lortie et al. 1992) and Shewanella putrefaciens (Kenward et al. 2006), whereas oxidizers include Thiobacillus and Leptothrix (Dowdle and Oremland 1998). It is also suggested that micro-organisms may facilitate the retention of selenium into organic matter (Dhillon et al. 2010). The formation of volatile selenium species dimethyl selenide (DMSe) and dimethyl diselenide (DMDSe) is enhanced by the action of micro-organisms (Darcheville et al. 2008; Dhillon et al. 2010; Guo et al. 2001; Martens and Suarez 1999). Moreover, in sterilized soils the volatilisation of selenium does not occur (Darcheville et al. 2008).

Wolinella succinogenes is capable of reducing SeO32- and SeO4

2- to elemental selenium in cultures after reaching the stationary growth rate (Tomei et al. 1992). Amorphous red elemental selenium formed granules in the cytoplasm of the bacteria. SeO3

2- was found to alter the morphology of the cells; normal cells and cells grown in the presence of SeO4

2- were curvy, whereas cells grown in SeO32- containing solutions were straight.

Wolinella succinogenes cannot use selenium as the final electron acceptor (Tomei et al. 1992), whereas Shewanella putrefaciens is capable of doing so (Kenward et al. 2006). Shewanella putrefaciens acts also as a biosorbent for SeO4

2-. The maximum sorption of selenite onto cell walls took place in pH 3, when approximately 40 % of the initial amount of SeO4

2- (not mentioned) was sorbed. Sorption mechanism was a combination of outer-sphere sorption and reduction on the cell surface.

Micro-organisms may also show dependence on the temperature and the presence of other anions. For instance, the reduction of SeO3

2- and SeO42- in concentrations of 48.1

mM was rapid in cultures of Pseudomonas stutzeri in oxic conditions, 25-35 °C and pH 6.5-9.5 (Lortie et al. 1992). From the initial concentration of 6.3 mM, 79% and 68 % of SeO3

2- and SeO42- was reduced to Se(0) within 24 hours. The growth of the culture and

the reduction of selenium were inhibited in the presence of sulphite or chromate in concentrations of 10-3 M.

The microbial oxidation of selenium is substantially slower reaction than the reduction reaction: apparent reduction turnover rate constants for SeO4

2- ranged between 0.0057-0.1447 a-1, while the apparent oxidation turnover rate constants of elemental selenium were 2.5x10-6-3.2x10-5 a-1 (Dowdle and Oremland 1998; Oremland et al. 1990). Reduction rate constants were calculated from the sum amount of selenite and selenate compared to the amount of elemental selenium in a given time t. Examples of micro-organisms capable of oxidizing selenium in soil are Thiobacillus and Leptothrix (chemoheterotrophs and chemoautotrophs) (Dowdle and Oremland 1998). The final products in the oxidation process are SeO4

2- and SeO32-. Interestingly, the oxidation rate

for chemically reduced Se was slower than for microbically reduced Se.

Micro-organisms increase the retention of selenium in soils (Darcheville et al. 2008; Février et al. 2007), for example in sterilized, non-sterilized and amended (glucose or cellulose) soils the sorption percentages of Se were 51 %, 62 % and >90 %, respectively (Darcheville et al. 2008). Furthermore, a greater fraction of SeO3

2- was found in the

61

residual and organically bound fractions compared with non-amended soils (5.2 % and 12.4 % for residual and organic-Se in amended soils, 3.3 % and 5.2 % for residual and organic-Se in non-amended soils). Février et al. (2007) found that the Kd values of SeO3

2- were higher for raw amended (RS) (nitrate, acetate, acetate+nitrate, glucose, glucose+nitrate) calcareous silty clay samples than for raw sterilized (StS) or raw anoxic soil (RSA) (Figure 23). The most efficient amendment was a combination of glucose and nitrate, which increased the Kd values for raw soils samples from approximately 20 l/kg to about 130 l/kg.

Figure 23. Kd values for short-term (48 hours) sorption experiments with raw soil (RS), anoxic raw soil (RSA) and sterilized soil (StS) samples with or without amendment (nitrate, acetate, acetate+nitrate, glucose, glucose+nitrate) (Février et al. 2007).

7.3 Sorption on organic matter and organic soils

The most important sink for selenium in soils is the organic matter (Choppin et al. 2009; Gustafsson and Johsson 1992; Pezzarossa et al. 1999) and the retention of selenium decreases with decreasing organic matter content (Pezzarossa et al. 1999). The mechanism of selenium association with organic matter is proposed to be mediated by soil micro-organism, as micro-organisms reduce selenium to lower oxidation states and enhance its sorption to low-molecular-weight humic substances (Dhillon et al. 2010). Gustafsson and Johsson (1992) found that within the organic matter, selenium was incorporated into hydrophobic fulvates, and the sorption of SeO3

2- was established to be insignificant on humic acid (Saeki and Matsumoto 1994). Organic matter can also increase the retention of volatile selenium species (DMSe and DMDSe) in soils (Martens and Suarez 1999).

Selenium is primarily retained in the soil organic humus layer, topmost 0-4 cm, where the complex-forming processes occur (Choppin et al. 2009; Gustafsson and Johnsson 1992). For example, Gustafsson and Johsson (1992) found that about 87 % of the initial selenium (610 μg) was retained in the topmost 4 cm in organic horizon of a podzol soil.

62

It is suggested that the downward migration of selenium in a soil profile takes place as complexes with organic substances (Choppin et al. 2009; Gustafsson and Johnsson 1992). In some instances solid particles enriched with iron and containing elevated concentrations of selenium can be found mixed within the organic layer (Choppin et al. 2009). Organic matter is also the main sorbent for SeO3

2- in acidic mineral soils (Dhillon and Dhillon 1999).

In groundwaters rich with humic substances, SeO32- is mainly found in a complexed

form associated with dissolved organic carbon, while SeO42- remains noncomplexed in

the solution (Bruggeman et al. 2007). Weng et al. (2011) found that most of the selenium present in soil solution, 67-86 %, was in colloidal fraction bound to DOM (dissolved organic matter), whereas 13-34 % was as SeO4

2-, SeO32- or bound to small

organic compounds (<1 nm).

The Kd value of selenium in organic soils (organic matter content 20 %) ranges from 230 to 1800 ml/g (Gil-García et al. 2009; IAEA 1994; Sheppard and Thibault 1990; Vidal et al. 2009). In Finnish peat, Kd values were as low as 12.8-15.1 ml/g (Yläranta 1983).

7.4 Sorption on mineral soils

Selenium sorption on mineral soils is found to increase with increasing weathering of soil parent material, the increasing content of aluminium and iron sesquioxides, and decreasing particle size fraction (Ashworth and Shaw 2006; Choppin et al. 2009; Dhillon and Dhillon 1999; Gustafsson and Johnsson 1992; Harada and Takahashi 2009; Keskinen et al. 2009 Neal et al. 1987a; Oram et al. 2011; Vuori et al. 1989). The effect of organic matter and micro-organisms is important in mineral soils (Ashworth and Shaw 2006; Dhillon and Dhillon 1999; Vuori et al. 1994). The self-exchange of selenium isotopes in the oxyanions SeO3

2- and SeO42- is insignificant in short time spans

<120 hours (Collins et al. 2006).

The pH of the soil affects the retention of selenium, as acidic soils are more efficient in retaining selenium than calcareous or alkaline soils (Dillon and Dhillon 1999). The sorption mechanism of SeO3

2- on acidic soils is mainly inner-sphere sorption (ligand exchange), whereas in alkaline soils sorption takes place through outer-sphere sorption. Moreover, the increased concentrations of calcium can enhance the sorption of selenite in alkaline pH (pH>5), by means of sorbing onto the surface of an adsorbent and thus creating positive surface charge (Neal et al. 1987b). Other possibility is the formation of precipitate, CaSeO3x2H2O. On the other hand, phosphate (PO4

3-) can effectively compete with SeO3

2- and remove sorbed selenite from the sorption sites in pH 4.0-9.0 (Neal et al. 1987b).

Selenium is enriched in clay soils compared with sand soils (Keskinen et al. 2009), and the efficiency of fine textured soils to sorb selenium is higher than that of coarse textured soils (Ashworth and Shaw 2006). Nonetheless, the affinity of organic matter towards selenium overcomes to that of fine textured clay soils. For example, the Kd values of selenium on organic soil, clay loam and sandy loam in field capacity water content were 274, 138 and 116 ml/g, respectively. Field capacity water content means the amount of soil moisture retained by the soil when no excess water is present and no downward water movement takes place. When the soil moisture content increased to

63

saturated conditions, the retention of selenium on sandy loam soil was higher than on clay loam soil; Kd for sandy loam and clay loam were 265 and 157 ml/g, respectively. The effect of micro-organisms in mineral soils is also obvious, as the Kd value of sandy loam in field capacity and saturated water conditions dropped from 116 and 265 ml/g to 78 and 206 ml/g after fumigation (Ashworth and Shaw 2006).

The vegetation cover affects the speciation of selenium in mineral soils. It can be present in the reduced forms elemental selenium and selenide in the rhizosphere soil because of the high concentration of organic acids and protons excreted from the plant roots and the presence of micro-organisms (Nascimento and Xing 2006; Oram et al. 2011; Verbruggen et al. 2009). Oram et al. (2011) found that SeO4

2- was the main species found in the vascular tissues of the roots, whereas outside the roots and on outer surfaces reduced forms elemental selenium and selenide dominated. It was postulated that plants are capable of oxidizing the reduced forms of selenium into mobile SeO4

2- and actively transport it to roots. The reduced forms were present in soils as mineral or organic selenides and elemental selenium. On the other hand, SeO4

2- and SeO32- were

associated with iron particles. Also, selenium concentration in the bulk soil was found to be higher than in the rhizosphere soil (130-700 μg/g compared with 160-530 μg/g). Even so, selenium in the rhizosphere soil was in more extractable form (soluble, ligand-exchangeable and/or exchangeable) than in the bulk soil (0.96-4.95 % compared with 0.41-1.55 %), indicating the action of plants.

In calcareous soil, CaCO3 is the predominant sorbent for selenium (Wang and Liu 2005). The Kd value for untreated soil was 34.8 ml/g, whereas the Kd values for calcareous soil treated to remove CaCO3 or organic matter were 16.0 and 19.8 ml/g, respectively. When CaCO3 and organic matter were both removed, the Kd value dropped to 12.5 ml/g. The relative contribution of CaCO3 was 54.0 % and the contribution of organic matter was 43.1 %. Calculated retardation factor, Rf, for selenium in calcareous soil was 64 (Wang et al. 2003).

7.4.1 Selenium in Finnish and Swedish soils

Finnish soils are naturally acidic and retain SeO42- poorly (Koljonen 1992). The soil

selenium concentration resembles that of the parent material, bedrock. Fine-textured soils retain selenium more efficiently than coarse-textured soil (Keskinen et al. 2009; Vuori et al. 1989). The retention also increases with the increased weathering degree of the parent material (Vuori et al. 1989). In Finnish clay and sand soils selenium is mainly related with organic matter and iron and aluminium sesquioxides (Keskinen et al. 2009). Higher concentrations of selenium are found in clay soils compared to sand soils. The association of selenium with organic matter and Al and Fe sesquioxides is higher in clay soils than in sand soils. Furthermore, the selenium concentration in iron sesquioxides is higher than in aluminium sesquioxides. The difference between selenium associated to iron and aluminium sesquioxides is pronounced in clay soils, whereas in sand soils approximately equal amounts of Se is bound to Fe and Al sesquioxides. Vuori et al. (1989) established that the sorbed fraction of SeO4

2- on 18 Finnish mineral soil samples varied from 3.6 % to 24.6 % of the initial amount (6.33 mmol/l).

Lusa et al. (2009) examined the content and distribution of selenium in organic humus layer and mineral soil layers from three excavator pits dug on Olkiluoto Island (OL-

64

KK14, OL-KK15 and OL-KK16 fs-cs, st). The composition of excavator pit OL-KK14 was clay/sandy till, OL-KK16 fine to coarse sand and sandy till and OL-KK15 clay, fine sand, sand, fine-grained silty/clayish till and coarse-grained sandy till. The information concerning the extraction solutions can be found from the caption of Figure 24 and Lusa et al. (2009). The highest selenium concentration in soil samples was typically measured from the exchangeable fraction (24.83-55.67 mg/kg), followed by Se bound to Fe and Mn oxides (21.67-43.17 mg/kg) in excavator pits OL-KK14 and OL-KK15. Interestingly, the highest selenium concentration in the humus layer of OL-KK15 was measured from the fraction associated to carbonates (103.67 mg/kg). Smaller concentrations of Se (0.53-13.33 mg/kg) were usually found in the fractions related to carbonates, organic matter and residual in OL-KK14 and OL-KK15. In excavator pit OL-KK16, selenium concentration decreased in the order exchangeable (11.67-62.50 mg/kg) > bound to organic matter (6.52-34.50 mg/kg) > residual (1.80-3.10 mg/kg) > bound to Fe and Mn oxides (1.83-10.23 mg/kg). Se was not found in the fraction associated with carbonates. In Appendix 1, the concentrations of selenium in extraction solutions, soil layers and excavator pits are given. Also the average Se concentrations in soil layers are included.

For excavator pits OL-KK15 and OL-KK16, selenium concentration in the humus layer (35-50 mg/kg) was noticeably higher than in the mineral soil layers MS1-MS5 (<15 mg/kg), and the selenium concentration decreased as a function of depth (Figure 24). In OL-KK14 Se concentration was approximately 20 mg/kg in every soil layer (Figure 24). The average selenium concentration among the mineral soil layers decreased in the order OL-KK14 > OL-KK15 > OL-KK16. In the humus layers the respective order was OL-KK15 > OL-KK16 > OL-KK14.

In Finnish and Swedish podzol soils selenium was found to be associated mainly with organic matter and/or iron oxyhydroxides (Gustafsson and Johnsson 1992; Koljonen 1992). Selenium was mostly enriched in O-Ah and B soil layers (Koljonen 1992). The recycling of selenium in the organic humus layer takes place as plants uptake selenium and liberates it back to soil upon decomposing (Koljonen 1992). In leaching layer (E) selenium is depleted due to the weathering of minerals, which releases selenium into soil solution. Selenium is transported into enrichment layer (B), where it is sorbed on iron and/or aluminium oxyhydroxides and organic matter. For example, Gustafsson and Johnsson (1992) found particularly strong enrichment of selenium in the organic matter. Selenium was mainly incorporated into hydrophobic fulvates, followed by humates and hydrophilic fulvates. It was hypothesized that selenium leached from above laying soil horizons is sorbed on aluminium and iron sesquioxides in the enrichment horizon and assimilated with organic matter afterwards. The assimilation rate was slower in smaller ratios of organic matter : aluminium and iron sesquioxides.

65

-5 0 5 10 15 20 25 30 35 40 45 50 55

MS5

MS4

MS3

MS2

MS1

humusSo

il la

yer

Se concentration (mg/kg)

OL-KK14 OL-KK15 OL-KK16

Figure 24. Selenium concentration in organic humus layer (humus) and mineral soil layers (MS1-MS5) in excavator pits OL-KK14, OL-KK15 and OL-KK16 in Olkiluoto (Lusa et al. 2009). Data points are the average concentrations calcultated from the sum of Se concentrations in different extraction solutions, namely exchangeable cations (1M BaCl2), bound to carbonates (1 M CH3COONH4 in 25% CH3COOH), bound to Fe and Mn oxides (0.04 M NH2OH*HCl in 25 % CH3COOH), bound to organic matter (0.02 M HNO3 and 30 % H2O2) and residual fraction (aqua regia) given in Lusa et al. (2009).

7.5 Sorption on minerals

7.5.1 Iron and aluminium oxides

The sorption mechanism of SeO32- onto goethite (FeO(OH)), fayalite (Fe2SiO4),

magnetite and ferrihydrate (Fe2O3 x 0.5H2O) is inner-sphere complexation in a bidentate or monodentate fashion (Duc et al. 2003; Harada and Takahashi 2009; Hayes et al. 1987; Martínez et al. 2006; Peak 2006). Unlike SeO3

2-, SeO42- forms mainly weak

outer-sphere complexes with goethite and ferrihydrite (Hayes et al. 1987; Peak and Sparks 2002), but in some studies formation of monodentate or bidentate inner-sphere complexes on the surface of goethite and ferrihydrate in highly acidic solutions (pH<3.5) have been detected (Manceau and Charlet 1994; Peak and Sparks 2002). On hematite SeO4

2- is sorbed as inner-sphere complexes (Peak and Sparks 2002). The formation of inner-sphere complexes between SeO4

2- and iron oxides and oxyhydroxides is dependent on pH; at pH 3.5 inner-sphere complexes are formed, between pH 3.5 and 6.0 SeO4

2- forms inner-sphere and outer-sphere complexes on goethite surface, and at pH 6.0 outer-sphere complexes are seen (Peak and Sparks 2002). On the other hand, SeO4

2- forms outer-sphere complexes on the surface of magnetite (Martínez et al. 2006). SeO3

2- is sorbed more efficiently on iron

66

oxyhydroxides than SeO42- (Duc et al. 2003). Kd values of SeO3

2- and SeO42- on

ferrihydrite may be as high as 106 ml/g and 1584 ml/g, respectively (Harada and Takahashi 2009).

Peak (2006) investigated the sorption mechanism of SeO32- and SeO4

2- onto hydrous aluminium oxide (HAO) and corumdum ( -Al2O3). SeO3

2- was found to form mainly bidentate binuclear inner-sphere surface complexes with HAO over the entire studied pH-range 4.5-8.0. Also some outer-sphere complexation was seen. SeO4

2- formed outer-sphere complexes with HAO over the entire studied pH-range. On corumdum, both outer-sphere and monodentate inner-sphere complexes were detected; mainly outer-sphere complexes formed at pH 3.5, whereas at pH 4.5 and 6.0 inner-sphere complexes was the main form.

Figures 25 and 26 presents the sorption of SeO32- onto hematite (25) and magnetite (26)

as a function of the normality (equivalent weight/solution volume) of the competing anion (CO3

2-, HCO3-, Cl- and SO4

2-) (Fujikawa and Fukui 1997). The sorption of SeO32-

is higher on hematite compared with magnetite. SO42- as a competing anion seems to

have a small effect on both minerals, whereas CO32- effectively decreases the sorption

of selenite. The impact of Cl- is insignificant on hematite, but noticeable on magnetite. The presence of HCO3

- has an impact on hematite, but not on magnetite.

Figure 25. Kd values of SeO32- on hematite as a function of the normality of competing

anions (CO32-, HCO3

-, Cl- and SO42-) (data taken from Fujikawa and Fukui 1997).

0,01

0,1

1

10

100

1000

0,001 0,01 0,1

Kd(m

l/g)

normality of competing anion (mol/l)

CO32

HCO3

Cl

SO42

67

Figure 26. Kd values of SeO32- on magnetite as a function of the normality of competing

anions (CO32-, HCO3

-, Cl- and SO42-) (data taken from Fujikawa and Fukui 1997).

Saeki and Matsumoto (1994) studied the sorption of SeO3

2- onto a variety of oxides, including -alumina ( -Al2O3), hydrous alumina (mainly bayerite, Al(OH)3), hematite (Fe2O3), goethite (FeO(OH)), amorphous TiO2, anatase (TiO2), amorphous MnO2, birnessite ((Na0.3Ca0.1K0.1)(Mn(IV)Mn(III))2O4 x 1.5H2O) and silicon dioxide (SiO2, quartz). Quartz was shown to have no affinity towards selenite (Saeki and Matsumoto 1994). The sorption of SeO3

2- on amorphous materials was higher than on crystalline substances. Table 8 summarizes the Kd values for these oxides excluding SnO2 at approximately pH-values of 3, 5 and 7. The sorption of SeO3

2- for all before mentioned oxides (quartz and SnO2 not included) was found to decrease with increasing pH (Table 8). From the sorbed selenium, approximately 25 %, 35 %, 46.7 % and 63.3 % was desorbed from goethite, hydrous alumina, amorphous TiO2 and amorphous MnO2, respectively. Extraction solutions were 0.1 M NaNO3, 0.05 M Na2SO4 and 0.033 M NaH2PO4, of which the latter desorbed most efficiently Se from the solid phase (15-43 %).

Table 8. Kd values for selenite on hydrous alumina, -alumina, goethite, hematite, amorphous TiO2, anatase, amorphous MnO2 and birnessite in approximately pH 3, 5 and 7 (Saeki and Matsumoto 1994).

Oxide Kd (ml/g)* pH 3 pH 5 pH 7

hydrous alumina 30 493 413 -alumina 45.7 10.7 10.7 goethite 1350 350 100

*values are derived from the data given in Saeki and Matsumoto (1994): sorption percentages approximated from graphics and transformed into Kd values.

0,1

1

10

100

0,001 0,01 0,1

Kd(m

l/g)

normality of competing anion (mol/l)

CO32

HCO3

Cl

SO42

68

Table 8. continued.

Oxide Kd (ml/g)* pH 3 pH 5 pH 7

hematite 30.1 13.5-37.5 13.6 amorphous TiO2 600 750 210

anatase - 111 86.2 amorphous MnO2 235 196 78.3

birnessite - 30 7.9 *values are derived from the data given in Saeki and Matsumoto (1994): sorption percentages approximated from graphics and transformed into Kd values. - sorption percentage corresponding to the desired pH value not given in the data

7.5.2 Pyrite and chalocopyrite

The sorption of selenide on pyrite (FeS2) and chalcopyrite (CuFeS2) is high. Pyrite can sorb nearly all of the added Se(-II) from 10-6 – 5x10-4 M solutions in a five-minute reaction period (Liu et al. 2008). Se(-II) was oxidized to Se(0) on the surface of pyrite, and pyrite surface was reduced into FeS-like species by the following reaction:

[12]

Naveau et al. (2007) concluded that the reduction reaction of selenium occurred with pyritic sulphur on natural pyrite and chalcopyrite. The metals (Fe, Cu) were not involved in the oxidation-reduction reaction. Selenium species present on the surface of pyrite was mainly Se(0) and to a lesser extent Se(II-), whereas on chalcopyrite surface diselenide (Se(II-)2) was detected.

The sorption of Se(II-,IV+) on pyrite and chalcopyrite decreases with increasing pH; pyrite and chalcopyrite can immobilize >20 % of the added Se(II-,IV+) (10-4 M) at pH<5, whereas at pH 9 <10 % is sorbed (Naveau et al. 2007). Natural and synthetic pyrite show differences in their sorption behavior towards selenium, as the sorption of Se(II-,IV+) on natural pyrite is smaller than on synthetized pyrite. At pH<5 synthetic pyrite sorbs nearly 95 % of the initial selenium, while only 20-65 % is sorbed on natural pyrite.

7.5.3 Other minerals

Hydroxyapatite and fluorapatite show considerable sorption affinity towards SeO32-, but

only an insignificant sorption of SeO42- occurs (Duc et al. 2003). The sorption

mechanism of SeO32- is specific sorption due to the liberation of PO4

3- ions into the solution upon the sorption of SeO3

2-. Furthermore, the substitution reaction can take place on outer surfaces and in lattice structure, as SeO3

2- was shown to diffuse into the crystal lattice of apatites.

Figure 27 presents the sorption of SeO32- onto calcite as a function of the normality

(equivalent weight/solution volume) of the competing anion (CO32-, HCO3

-, Cl- and SO4

2-) (Fujikawa and Fukui 1997). The effect of CO32- is the most noteworthy, as the Kd

decreases from 24 ml/g to 1x10-5 ml/g with increasing normality from 0.001 N to 0.1 N. The effect of HCO3

2- is obvious, and Cl- and SO42- seem to have only minor effect on

the sorption of SeO32- on calcite.

69

Figure 27. Kd values of SeO32- on calcite as a function of the normality of the competing

anions (CO32-, HCO3

-, Cl- and SO42-) (modified after Fujikawa and Fukui 1997.)

Ticknor et al. (1989) studied the sorption of selenium onto typical primary and fracture-filling minerals met in the East Bull Lake pluton in Canada. They found that different minerals were responsible for the retention of selenium in oxic and anoxic conditions. In oxic conditions Se was sorbed on chlorite ((Mg,Fe)3(Si,Al)4O10(OH)2 x (Mg,Fe)3(OH)6) and serpentine (Mg3(OH)4Si3O5), whereas olivine ((Mg,Fe)2SiO4) and unaltered plagioclase showed moderate and insignificant sorption. In anoxic conditions, the affinity of olivine and unaltered plagioclase remained unchanged and orthopyroxene showed high retention capacity. Overall, the retention of Se on minerals decreased in the order chlorite serpentine > orthopyroxene > olivine > unaltered plagioclase.

Frost and Griffin (1977) and Bar-Yosef and Meek (1987) have studied the sorption of SeO3

2- and/or SeO42- on kaolinite and montmorillonite. Montmorillonite has shown to

sorb more selenite than kaolinite. Sorption maxima occurred at pH 2-4.5 and decreased at higher pH-values. SeO3

2- is sorbed as an outer-sphere complex on motmorillonite (Charlet et al. 2007). If Fe(II) is sorbed on montmorillonite, SeO3

2- can be reduced to Se(0) forming nanoparticles (Charlet et al. 2007). The sorption of SeO3

2- on kaolinite takes place in two phases: the initial fast sorption process accounted for approximately 95 % of the sorption, and was followed by a slow phase depleting selenium from the solution (Bar-Yosef and Meek 1987).

7.6 Volatilisation of selenium

Dimethyl selenide (DMSe) and dimethyl diselenide (DMDSe) are the volatile species of selenium formed in soils through the action of micro-organisms. Also, plants increase the volatilisation of selenium form soils via the production of DMSe and DMDSe. Most of the DMSe and DMDSe arising from the soil are formed in the biologically active layer, uppermost 4-5 cm (Martens and Suarez 1999). DMSe is the main species detected in the gas phase, while DMDSe is sorbed on soil. The formation of DMSe from SeO4

2- takes place in two phases: first, rapid phase occurs when Se(VI+) is reduced to Se(IV+),

0,00001

0,0001

0,001

0,01

0,1

1

10

100

0,001 0,01 0,1

Kd(m

l/g)

normality of competing anion (mol/l)

CO32

HCO3

Cl

SO42

70

which is seen as a fast increase in the production rate of DMSe (Guo et al. 2001). Second, slower phase arises as Se(IV+) is further reduced to Se(0) or organic selenium. These species are generally rather insoluble and react slowly, which cause the production rate of DMSe to decrease. The vapor pressure and solubility of DMSe are 32.03 kPa and 0.0244 g/g water (Karlson et al. 1994). The vapour pressure of DMDSe is lower than for DMSe, namely 0.38 KPa.

It is proposed that the volatilisation of selenium from soil may be a significant pathway for Se to the atmosphere. The migration of volatilized selenium species in soils is affected by the plant cover, organic matter, clay and moisture content of the soil (Ashworth and Shaw 2006; Guo et al. 1999, 2001; Martens and Suarez 1999; Zhang et al. 1999).

Ashworth and Shaw (2006) studied the migration and volatilisation of selenium in planted and unplanted sandy loam columns in Britain. They demonstrated that the volatilisation rate of Se from unplanted soil columns was considerably lower compared to planted soil columns in a nine-month experiment. The volatilisation rate for unplanted soils was always lower than in planted soils, and no volatilisation could be measured from unplanted soil columns after three months. The reason for the loss of Se volatilisation from unplanted soil columns was not given. On the other hand, the volatilisation from planted soils remained quite high throughout the experiment (5-15.5 Bq/m2/day). The highest volatilisation rate for unplanted soil was 9 Bq/m2/day measured for the second month, while the highest rate for planted soil was seen at third month (15.5 Bq/m2/day).

The migration of volatile selenium species in soils is affected by soil moisture content, clay fraction content and the thickness of the clean soil layer, i.e. a layer containing no selenium (Guo et al. 2001; Zhang et al. 1999). In dry, moist (20% moisture content) and water-saturated soil columns the breakthrough peaks for DMSe through the soil to the surface took place in 20 min, 42-81 h and 120-140 h, respectively (Zhang et al. 1999). The reason for different behaviour was attributed to the changes in the porosity and air space of the soils, which decreased with increasing water content. As the water content increased, DMSe had to dissolve in the water phase before volatilisation. Also, the increase in the clay fraction content of the soil decreased the diffusion rate of DMSe; for clay contents of 17.5 %, 25 % and 32.5 % the nonvolatilized fraction at 20 % soil moisture content were 12 %, 39.1 % and 79.6 %, respectively (Zhang et al. 1999).

Guo et al. (2001) demonstrated that the thickness of a clean soil layer (not containing selenium) on the top of a selenium contaminated layer dictates the transport of the formed volatilized Se species in soils. The calculated volatilisation rate for DMSe in uncovered soil was 0.287 % of the initial (20 μg SeO4

2-/g soil), and DMSe was found within 24 hours after the beginning of the experiment. Peak value was found on day 6 (16.55 %). However, the halflife of DMSe in soil ranges from few minutes to few hours (Guo et al. 1999, 2001). The halflife indicates the time taken for the concentration of DMSe in soil to degradate to half of the original. The short halflife implies that longer residence times due to thicker soil layers diminishes the concentration of DMSe. In fact, for soil column with 2 cm thick clean soil coating the peak value of DMSe was only 3.65 % of the initial (20 μg SeO4

2-/g soil), and for thicker coatings, 8 and 16 cm, practically no DMSe was detected (Guo et al. 2001).

71

Soil organic matter may also inhibit the transport of DMSe and DMDSe in soils (Martens and Suarez 1999). DMDSe is mainly sorbed on organic matter and turned into strongly sorbed SeO3

2-, Se(0) and Se(-II) in the solid phase. Only a small amount of DMSe was sorbed on the solid phase, whereas the fraction of DMSe oxidized to SeO3

2- and SeO4

2- accounted for 74-100 % of the initial amount (170 nM) by the action of micro-organisms.

Besides affecting the diffusion of DMSe in soils, the water content also affects its transformations into non-volatile forms. At the moisture contents of 5 %, 15 %, 35 % and 130 % the transformation of DMSe into non-volatile forms were 1.3 %, 1.8 %, 18.4 % and 71.5 %, respectively (Zhang et al. 1999). DMSe was transformed mainly into water-soluble forms, namely oxized dimethylated selenium compounds DMSeO (dimethyl selenoxide) and DMSeO2 (dimethyl selenone). The ODMSe fraction accounted for 80-96 % and the rest was complexed with humic substances, selenite and selenite. In the presence of MnO2 in 35 % moist soil, the non-volatile fraction increased from 18.4 % to 95 %.

Kd values for selenium collected from the literature and presented in the text are given in Table 9.

Table 9. Kd values for selenium presented in the text and collected from the literature.

Material Species Kd (ml/g)

Sample description/ other information

Country/ place

Reference

Organic Se 1000 (230-1800)

arithmetic mean (range) Vidal et al. 2009

Organic Se 1800 IAEA 1994 Organic Se 1800 geometric mean value Sheppard and

Thibault 1990

Organic soil Se 230-1800 min-max value Gil-García et al. 2009

Carex peat SeO32- 15.1;

12.8

[Se]i=3x10-5 M; 3x10-4 M Finnish soil Yläranta 1983

Peat Se 11 native elements measured with ICP

Forsmark, Sweden

Sheppard et al. 2009

Peat Se 130 native elements measured with ICP

Laxemar-Simpevarp, Sweden

Sheppard et al. 2009

Cambisols SeO32- 37-490 min-max value Japanese soil

samples, 0-20 cm depth

Nakamaru and Sekine 2008

72

Table 9. continued.

Material Species Kd (ml/g)

Sample description/ other information

Country/ place

Reference

Clay Se 400/240 (22-2100)

arithmetic mean / geometric mean (range)

Vidal et al. 2009

Clay Se 740 IAEA 1994 Clay Se 740 geometric mean value Sheppard and

Thibault 1990 Clay soil Se 250 (22-

2130) geometric mean (min-max) value

Gil-García et al. 2009

Clay soil SeO32- 119.9;

65.8 [Se]i=3x10-5 M; 3x10-4 M Finnish soil Yläranta 1983

Clay gyttja Se 56 native elements measured with ICP

Laxemar-Simpevarp, Sweden

Sheppard et al. 2009

Clay gyttja Se 140 native elements measured with ICP

Laxemar-Simpevarp, Sweden

Sheppard et al. 2009

Clayey till Se 10 native elements measured with ICP

Forsmark, Sweden

Sheppard et al. 2009

Clayey till Se 14 native elements measured with ICP

Forsmark, Sweden

Sheppard et al. 2009

Clay loam soil SeO32- 1.7-110.9 soil pH 8.2, organic C

0.35 %, CaCO3 10.15 %, free iron 0.25 %

India Dhillon and Dhillon 1999

Clay loam soil SeO32- 1.0-8.3 soil pH 8.2, organic C

0.89 %, CaCO3 9.51 %, free iron 0.63 %

India Dhillon and Dhillon 1999

Loam Se 350/220 (12-1600)

arithmetic mean / geometric mean (range)

Vidal et al. 2009

Loam Se 490 IAEA 1994 Loam Se 500 geometric mean value Sheppard and

Thibault 1990 Loam soil Se 220 (12-

1606) geometric mean (min-max) value

Gil-García et al. 2009

Loamy soil SeO32- 3.8-363.8 soil pH 6.8, organic C

0.27 %, free iron 0.75 % India Dhillon and

Dhillon 1999 Loamy soil SeO3

2- 5.5-570.2 soil pH 6.0, organic C 0.27 %, free iron 1.87 %

India Dhillon and Dhillon 1999

Silty clay loam SeO32- 5.0-62.4 soil pH 8.0, organic C

0.77 %, CaCO3 1.91 %, free iron 1.35 %

India Dhillon and Dhillon 1999

Sandy clay loam soil

SeO32- 1.3-14.0 soil pH 5.8, organic C

0.22 %, free iron 1.66 % India Dhillon and

Dhillon 1999 Sandy clay loam soil

SeO32-

and SeO4

2-

486/49 SeO32-/ SeO4

2-; 6.5 g/kg organic C, 3.11 g/kg citrate- dithionite Fe, specific surface area 24.3 m2/g soil pH 8.4

the Powder Riwer Basin, Wyoming, USA

Sharmasarkar and Vance 2002

73

Table 9. continued. Material Species Kd

(ml/g) Sample description/ other information

Country/ place

Reference

Sandy loam soil

SeO32-

and SeO4

2-

739/439 SeO32- / SeO4

2- ; 4.9 g/kg organic C, 2.43 g/kg citrate-dithionite Fe, specific surface area 13.8 m2/g soil pH 8.3

the Powder Riwer Basin, Wyoming, USA

Sharmasarkar and Vance 2002

Sandy loam soil

SeO32-

and SeO4

2-

979/378 SeO32- / SeO4

2- ; 5.3 g/kg organic C, 2.50 g/kg citrate-dithionite Fe, specific surface area 13.9 m2/g soil pH 8.4

the Powder Riwer Basin, Wyoming, USA

Sharmasarkar and Vance 2002

Sandy loam soil

SeO32-

and SeO4

2-

1305/1495 SeO32- / SeO4

2- ; 8.1 g/kg organic C, 4.12 g/kg citrate-dithionite Fe, specific surface area 13.7 m2/g soil pH 8.3

the Powder Riwer Basin, Wyoming, USA

Sharmasarkar and Vance 2002

Sandy loam soil

SeO32-

and SeO4

2-

19/ 10 SeO32- / SeO4

2- ; 7.5 g/kg organic C, 2.88 g/kg citrate-dithionite Fe, specific surface area 13.8 m2/g soil pH 8.4

the Powder Riwer Basin, Wyoming, USA

Sharmasarkar and Vance 2002

Sandy loam soil

SeO32-

and SeO4

2-

11/8 SeO32- / SeO4

2- ; 6.9 g/kg organic C, 2.80 g/kg citrate-dithionite Fe, specific surface area 13.8 m2/g, soil pH 8.5

the Powder Riwer Basin, Wyoming, USA

Sharmasarkar and Vance 2002

Sandy loam soil

SeO32-

and SeO4

2-

9/8 SeO32- / SeO4

2- ; 5.8 g/kg organic C, 3.93 g/kg citrate-dithionite Fe, specific surface area 17.2 m2/g soil pH 8.0

the Powder Riwer Basin, Wyoming, USA

Sharmasarkar and Vance 2002

Fine sand soil SeO32- 22.6;

10.3 [Se]i=3x10-5 M; 3x10-4 M Finnish soil Yläranta 1983

Silt soil SeO32- 2.7-28 soil pH 7.9, organic C

0.88%, CaCO3 2.74%, free iron 1.41%

India Dhillon and Dhillon 1999

Silt soil SeO32- 5.2-241.1 soil pH 5.5, organic C

0.19%, free iron 1.24% India Dhillon and

Dhillon 1999 Sand Se 150 geometric mean value Sheppard and

Thibault 1990 Sand Se

220/56 (4-1600)

arithmetic mean / geometric mean (range)

Vidal et al. 2009

Sand Se 150 IAEA 1994 Sand soil Se 56 (4-

1616) geometric mean (min-max) value

Gil-García et al. 2009

Sterilized sandy soil

SeO32- 5.1 soil pH 7.74, organic C

4.38 g/kg Isére, France Darcheville et

al. 2008 Sandy soil SeO3

2- 11 soil pH 7.74, organic C 4.38 g/kg

Isére, France Darcheville et al. 2008

Sandy soil with glucose

SeO32- 98 glucose 5 g/kg soil, soil pH

7.74, organic C 4.38 g/kg Isére, France Darcheville et

al. 2008 Sandy soil with cellulose

SeO32- 47 cellulose 4.5 g/kg soil, soil

pH 7.74, organic C 4.38 g/kg

Isére, France Darcheville et al. 2008

74

Table 9. continued. Material Species Kd

(ml/g) Sample description/ other information

Country/ place

Reference

Sandy till Se 41 native elements measured with ICP

Laxemar-Simpevarp, Sweden

Sheppard et al. 2009

Soil SeO32-

and SeO4

2-

255 / 6.4 SeO32- / SeO4

2- ; soil pH 7.1

soil from the USA

Collins et al. 2006

Soil SeO32-

and SeO4

2-

40 / 2.9 SeO32- / SeO4

2- ; soil pH 8.4

soil from the USA

Collins et al. 2006

Soil SeO32-

and SeO4

2-

8.6 / 5.1 SeO32- / SeO4

2- ; soil pH 8.1

soil from the USA

Collins et al. 2006

Soil SeO32-

and SeO4

2-

66 / 1.5 SeO32- / SeO4

2- ; soil pH 7.8

soil from the USA

Collins et al. 2006

Soil Se 16.8 [Se]i=2x10-7 M, pH 5.9, sample depth 7.8-7.9 m, specific surface area 8.6 m2/g

Leqoux et al. 1992

Soil Se 133 [Se]i=2x10-7 M, pH 6.6, sample depth 9.65-9.85 m, specific surface area 12.2 m2/g

Leqoux et al. 1992

Soil Se 89 [Se]i=2x10-7 M, pH 8.0, sample depth 11.1-11.25 m, specific surface area 14.7 m2/g

Leqoux et al. 1992

Soil Se 2.4 [Se]i=2x10-7 M, pH 8.0, sample depth 12.0-12.2 m, specific surface area 6.6 m2/g

Leqoux et al. 1992

Soil (Typic Torriorthent)

SeO32-

and SeO4

2-

7/6 SeO32- / SeO4

2- ; 6.8 g/kg, organic C, 3.91 g/kg citrate- dithionite Fe, clay loam, surface area 27.4 m2/g soil pH 7.8

the Powder Riwer Basin, Wyoming, USA

Sharmasarkar and Vance 2002

Andosols SeO32- 93-1028 min-max Japanese soil

samples, 0-20 cm depth

Nakamaru and Sekine 2008

Andosol SeO32- ~300 non-cultivated soil, soil

pH (H2O) 6.4, oxalate extracted iron and aluminiun 8.6 and 17.9 g/kg

Japan Nakamaru et al. 2005

Andosol SeO32- ~475 cultivated soil, soil pH

(H2O) 6.4, oxalate extracted iron and aluminiun 8.6 and 17.9 g/kg

japan Nakamaru et al. 2005

Fluvisols SeO32- 68-1034 min-max Japanese soil

samples, 0-20 cm depth

Nakamaru and Sekine 2008

75

Table 9. continued. Material Species Kd

(ml/g) Sample description/ other information

Country/ place

Reference

Calcareous soil SeO32- 294 pH 8 cultivated soil

from China, depth 0-20 cm

Zuyi et al. 2000

Calcareous soil HSeO3-

or SeO32-

38.4 / 21.5 batch/column experiment, pH 7.8

Yuzhong county, Gansu province, China

Wang and Liu 2005

Calcareous soil – organic matter removed

HSeO3-

or SeO32-

19.8 / 14.9 batch/column experiment, pH 7.8

Yuzhong county, Gansu province, China

Wang and Liu 2005

Calcareous soil - CaCO3 removed

HSeO3-

or SeO32-

16.0 / 10.8 batch/column experiment, pH 7.8

Yuzhong county, Gansu province, China

Wang and Liu 2005

Calcareous soil – organic matter CaCO3 and removed

HSeO3-

or SeO32-

12.5 / 10.4 batch/column experiment, pH 7.8

Yuzhong county, Gansu province, China

Wang and Liu 2005

Calcareous soil HSeO3-

or SeO32-

32 column experiment, pH 7.8

Yuzhong county, Gansu province, China

Wang et al. 1999

Calcareous soil – organic matter removed

HSeO3-

or SeO32-

27 column experiment, pH 7.8

Yuzhong county, Gansu province, China

Wang et al. 1999

Calcareous soil - CaCO3 removed

HSeO3-

or SeO32-

29 column experiment, pH 7.8

Yuzhong county, Gansu province, China

Wang et al. 1999

Red earth SeO32-

3.3x10 3

pH 6 a coastal sandy soil from China

Zuyi et al. 2000

Sediment SeO32-

and SeO4

2-

9520 / 1.1 SeO32- / SeO4

2- ; sediment pH 6.4

sediment from France

Collins et al. 2006

Sediment SeO32-

and SeO4

2-

3420 / 9.1 SeO32- / SeO4

2- ; sediment pH 6.8

sediment from France

Collins et al. 2006

Sediment SeO32-

and SeO4

2-

4770 / 7.4 SeO32- / SeO4

2- ; sediment pH 6.7

sediment from France

Collins et al. 2006

Sediment SeO32-

and SeO4

2-

9390 / 14.4

SeO32- / SeO4

2- ; sediment pH 6.6

sediment from France

Collins et al. 2006

Sediment SeO32-

and SeO4

2-

412 / 0.9 SeO32- / SeO4

2- ; sediment pH 6.8

sediment from France

Collins et al. 2006

Alumina SeO32-

143

Chromatographic alumina, pH 6

Zuyi et al. 2000

Quartz sand SeO32- ~1 equilibrium time 55 days,

liquid synthetic seawater Jan et al. 2006

76

Table 9. continued. Material Species Kd

(ml/g) Sample description/ other information

Country/ place

Reference

Natural apatite SeO42- 29.3 final pH 7.26, surface

area 10.0 m2/g Florida phosphate rock

Thomson et al. 2003

Tri-calcium-phosphate

SeO42- 48.6 final pH 7.84, surface

area 67.6 m2/g Thomson et al.

2003 Hydroxyapatite SeO4

2- 650.8 final pH 6.85, surface area 56.9 m2/g

Thomson et al. 2003

Hydroxyapatite SeO42- 1.1 final pH 7.89, surface

area 93.9 m2/g Thomson et al.

2003 Hydroxyapatite SeO4

2- 4.6 final pH 7.61, surface area 51.1 m2/g

Thomson et al. 2003

Hydroxyapatite SeO42- 124.1 final pH 7.57 Thomson et al.

2003 Activated magnetite

SeO42- 17.0 final pH 7.67, surface

area 1.7 m2/g Thomson et al.

2003 Bentonite SeO3

2- ~57 equilibrium time 55 days, liquid synthetic seawater

Jan et al. 2006

Illite SeO32- 150 – 75 pH 3 – pH 8 Missana et al.

2009 Smectite SeO3

2- 500 – 100 pH 3 – pH 8 Missana et al. 2009

77

8 CONCLUSIONS

For the spent nuclear fuel repository to be built in Olkiluoto, Finland, the biosphere (dose) assessment programme is focused to Cl-36, I-129, C-14, Mo-93, Nb-94, Cs-135, Ni-59, Se-79 and Sr-90 (Hjerpe et al. 2010, p. 36-38; Haapanen et al. 2009, p. 22-23). Earlier, Söderlund et al. (2011) reviewed literature for sorption of I, Cl, Tc and Cs in soil, and the present report addresses Mo, Nb and Se.

Mo-93, Nb-94 and Se-79 are long-lived activation products formed from the stable isotopes present in the nuclear fuel, structural components and construction materials used in power reactors. Se-79 is also a fission product. Knowledge of their behaviour in the environment is of great interest because of their long physical half-lives and anionic nature.

The migration and sorption of radionuclides in soils is affected by parameters specific to the element and to the soil. Chemical form, speciation, is the most important elemental factor affecting the sorption and migration properties of the element. Soil redox potential, pH and complex forming ligands are features that have great influence on the speciation. Micro-organisms present in soils can affect the speciation of radionuclides indirectly by changing the prevailing Eh-pH conditions. They can also serve as sorbents. The organic matter content and mineral properties of soils have a noticeable influence on the retention of radionuclides. The sorption of anionic radionuclides such as MoO4

2-, SeO32- and SeO4

2- is pronounced in the presence of organic matter and iron and aluminium oxyhydroxides.

Distribution coefficient, Kd, is used as a measure to describe the fraction of the radionuclide adsorbed on soils, sediments and suspended material. Kd is used to describe the retention and mobility of radionuclides in the environment: with low Kd values, the fraction of radionuclide sorbed in the solid phase is low, and the fraction in the liquid phase is high. This leads to low retention and high potential mobility in the overburden. Radionuclides which have high tendency of forming anionic species in aqueous solutions, such as molybdenum, niobium and selenium, are poorly sorbed on soil mineral constituents, thus leading to high potential mobility in the overburden. The sink for anionic radionuclides, such as molybdenum and selenium, is thought to be organic matter. Radionuclides associated with the soil organic matter and soil micro-organisms are of great importance since they remain readily exchangeable and are available for plant uptake.

Molybdenum

Molybdenum is an important element for plants, animals and micro-organism due to its role in the reduction of nitrate (NO3

-) into ammonium ion (NH4+) as a part of Fe-Co

cofactor. Also, certain nitrogen-fixing bacteria are capable of compiting against organic matter for free Mo present in soils by excreting Mo-complexing catechol compounds.

Molybdenum is a redox-sensitive element and can exist on oxidation states ranging from VI+ to 0. Molybdate anion, MoO4

2-, is the main species found in diversity of redox-conditions. Due to its anionic nature, the retention of MoO4

2- on surfaces of soil particles is relatively low. Sorption is favoured when soil contains higher amounts of

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clay fraction, organic matter and aluminium and iron sesquioxides. Single most important factor is the prevailing pH. Sorption is enhanced in acidic pH values.

The main sorbent for molybdenum in soils is the organic matter. It has been found that the concentration of Mo in soils decreases with decreasing amount of organic matter and increasing depth. Highest concentrations of Mo are found on the soil surfaces, in the organic humus layer. The sorption mechanism of Mo on organic matter is complex formation with carboxyl, phenol and catechol groups. Sorption takes place because octahedral coordination of Mo is favoured instead of tetrahedral coordination.

In the mineral fractions of soil the main Mo sorbing components are iron and aluminium oxohydroxides. The sorption mechanism is specific sorption. Sorption is favoured in acidic pH. In pozdol soils elevated concentrations of Mo are found in addition to the organic humus layer in the illuvial (enrichment) layer.

Acidic soils tend to retain higher amounts of Mo compared to neutral ones as the sorption of MoO4

2- on oxides and minerals is favoured in acidic pH (pH<7). Retention is more efficient on poorly crystalline minerals compared to well-crystallized minerals. Iron oxide goethite has found to be excellent sorbent molybdenum, whereas clay minerals are not so efficient.

Phosphate (PO43-) can effectively compete with MoO4

2- for sorption sites. Competition is weaker with lower pH and longer interaction time between molybdenum and soil.

The sorption of molybdenum on fine-textured soils is enhanced in comparison with coarse-textured soils.

The efficiency of the mineral soils to retain molybdenum decreased in order clay > finer fine sand > coarse fine sand. In the soils of the Olkiluoto site, Mo is typically found in fractions related to exchangeable ions, bound to iron and manganese oxides and bound to organic matter.

Niobium

The sorption of niobium on soils has received little attention. Compared to the number of publications published on the retention of molybdenum and selenium on soils, only a few articles concerning niobium have been published. These articles have mainly focused on the soil-to-plant transfer of niobium, or on the sorption of niobium on clays, sands and gravel. The mobility of niobium in soils is expected to be low due to the low solubility of niobium compounds, the formation of sparingly soluble complexes and salts, and high sorption on rocks and clays. High organic matter and clay fraction content increase the sorption of niobium in soils. Also the texture of the soil affects the sorption: the predicted Kd values decrease in the order clay>loam>sand. The soil-to-plant transfer of niobium is low.

Selenium

Selenium has notably gained interest in soil science because of its essential nature for humans (trace nutrient), narrow range between indeficient and toxic dose, and low or unusually high concentration in soils. Due to these reasons the sorption of selenium in

79

soils and distribution between different fractions (adsorbed, organically bound, residual, etc.) has grown attention.

The sorption of selenium in soils increases with increasing organic matter content, clay fraction content, amount of CaCO3, weathering of the soil material, iron and aluminium oxyhydroxide content and and increasing concentration of Ca2+ in the solution. The sorption of selenium decreases as the presence of competing anions, such as phosphate, arsenate, carbonate and sulphate increases. Micro-organisms have an important role in the biogeochemistry, speciation and distribution of selenium in soils

The association of selenium with organic matter and Al and Fe sesquioxides is higher in clay soils than in sand soils. Furthermore, the selenium concentration in iron sesquioxides is higher than in aluminium sesquioxides. The difference between selenium associated to iron and aluminium sesquioxides is pronounced in clay soils, whereas in sand soils approximately equal amounts of Se is bound to Fe and Al sesquioxides.

The speciation of selenium depends on the prevailing Eh-pH conditions. Inorganic forms of Se include selenate (SeO4

2-), selenite (SeO32-), elemental selenium (Se(0)) and

selenide (Se(-II)). The oxidation state of selenium in selenate and selenite are +VI and +IV, respectively. These forms are considered as the most mobile and biogeochemically significant, selenate more mobile than selenite.

The vegetation cover affects the speciation of selenium in mineral soils. It can be present in the reduced forms elemental selenium and selenide in the rhizosphere soil because of the high concentration of organic acids and protons excreted from the plant roots and the presence of micro-organisms. Selenium concentration in the bulk soil has been found to be higher than in the rhizosphere soil. Even so, selenium in the rhizosphere soil is in more extractable form (soluble, ligand-exchangeable and/or exchangeable) than in the bulk soil, indicating the action of plants.

SeO32- is sorbed on minerals, soils and sediments much more efficiently than SeO4

2.

Retention of these two species is dependent on pH and decreases with increasing pH; the maximum sorption of selenium takes place in acidic solutions. For example, considerable amounts of SeO3

2- has been sorbed in pH 2-9 on a sandy loam sample, while SeO4

2- was sorbed only in pH<3.

Dimethyl selenide (DMSe) and dimethyl diselenide (DMDSe) are the most common volatile species of selenium. The volatilisation of selenium in soils is an important pathway for radioselenium from soils into biosphere and atmosphere. Soil micro-organisms and plants enhance the volatilisation of selenium. Thick soil cover and high moisture content inhibit the transport and formation of DMSe and DMDSe in soils. The half-live of volatile selenium species in soils ranges from a few minutes to few hours.

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Appendix 1. The concentrations of molybdenum and selenium in different soil layers in excavator pits OL-KK14, OL-KK15 and OL-KK16 in Olkiluoto. Samples from MS5 mineral soil layer were not taken from excavator pits OL-KK14 and OL-KK16. Concentrations were calcultated from the contents (mg/kg) in the extraction solutions 1M BaCl2 (extraction step A; exchangeable cations), 1 M CH3COONH4 in 25% CH3COOH (B; bound to carbonates), 0.04 M NH2OH*HCl in 25 % CH3COOH (C; bound to Fe and Mn oxides), 0.02 M HNO3 and 30 % H2O2 (D; bound to organic matter) and aqua regia (E; residual fraction). Extraction

step OL-KK14 OL-KK15 OL-KK16

Mo (mg/kg)

Se (mg/kg)

Mo (mg/kg)

Se (mg/kg)

Mo (mg/kg)

Se (mg/kg)

humusA 2.07 55.67 1.40 50.50 < 62.50 B 0.35 13.17 2.52 103.67 < < C 1.30 21.67 2.62 43.17 < 10.23 D 0.20 3.15 0.10 1.01 1.87 34.50 E < 3.03 < < < <

average 0.98 19.34 1.66 49.59 1.87 35.74 MS1

A 2.15 52.00 0.92 26.67 < 11.83 B 0.44 13.33 0.56 11.17 < < C 1.45 25.67 1.43 22.83 < 1.83 D 0.49 9.95 0.26 1.10 0.20 7.00 E < 6.07 < 5.60 < 2.95

average 1.13 21.40 0.79 13.47 0.20 5.54 MS2

A 2.07 52.67 0.72 25.33 < 11.83 B 0.35 12.50 0.47 11.33 < < C 1.45 24.50 1.35 22.67 < 1.98 D 0.29 9.72 < 1.10 0.19 6.98 E < 7.48 < 5.37 < 1.80

average 1.04 21.37 0.85 13.16 0.19 5.65 MS3

A 1.97 51.00 0.71 25.17 < 11.83 B 0.35 12.83 0.46 11.33 < < C 1.30 24.83 1.32 22.17 < 1.97 D 0.30 9.63 < 0.53 0.18 6.68 E < 6.12 < 3.70 < 3.10

average 0.98 20.88 0.83 12.58 0.18 5.90 MS4

A 2.07 53.00 0.76 25.17 < 11.67 B 0.34 12.50 0.48 11.67 < < C 1.27 24.17 1.33 22.50 < 1.85 D 0.35 9.45 < 0.85 0.17 6.52 E < 6.25 < 4.12 < 2.00

average 1.01 21.07 0.86 12.86 0.17 5.51

< concentration below the detection limit

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Appendix 1. continued. Samples from MS5 mineral soil layer were not taken from excavator pits OL-KK14 and OL-KK16. Extraction

step OL-KK14 OL-KK15 OL-KK16

Mo (mg/kg)

Se (mg/kg)

Mo (mg/kg)

Se (mg/kg)

Mo (mg/kg)

Se (mg/kg)

MS5 A 0.69 24.83 B 0.46 10.50 C 1.32 21.83 D 0.13 0.71 E < 3.10

average 0.65 12.19

< concentration below the detection limit