Hydrogeochemical Interpretation of Baseline Groundwater ... · site, geomicrobial studies, PCA and...

P O S I V A O Y

F I N - 2 7 1 6 0 O L K I L U O T O , F I N L A N D

P h o n e ( 0 2 ) 8 3 7 2 3 1 ( n a t . ) , ( + 3 5 8 - 2 - ) 8 3 7 2 3 1 ( i n t . )

F a x ( 0 2 ) 8 3 7 2 3 7 0 9 ( n a t . ) , ( + 3 5 8 - 2 - ) 8 3 7 2 3 7 0 9 ( i n t . )

Hydrogeochemical Interpretationof Baseline Groundwater Conditions

at the Olkiluoto Site

February 2004

POSIVA 2003 -07

Pet ter i P i tkänenSami Par tamiesAr i Luukkonen

POSIVA 2003 -07

February 2004

Pet te r i P i t känen

Sami Par tamies

Ar i Luukkonen

VTT Bu i l d i ng and T r anspo r t

P O S I V A O Y

F I N - 2 7 1 6 0 O L K I L U O T O , F I N L A N D

P h o n e ( 0 2 ) 8 3 7 2 3 1 ( n a t . ) , ( + 3 5 8 - 2 - ) 8 3 7 2 3 1 ( i n t . )

F a x ( 0 2 ) 8 3 7 2 3 7 0 9 ( n a t . ) , ( + 3 5 8 - 2 - ) 8 3 7 2 3 7 0 9 ( i n t . )

Hydrogeochemical Interpretationof Baseline Groundwater Conditions

at the Olkiluoto Site

Base maps: ©National Land Survey, permission 41/MYY/04

ISBN 951 -652 -121 -5ISSN 1239 -3096

T h e c o n c l u s i o n s a n d v i e w p o i n t s p r e s e n t e d i n t h e r e p o r t a r e

t h o s e o f a u t h o r ( s ) a n d d o n o t n e c e s s a r i l y c o i n c i d e

w i t h t h o s e o f P o s i v a .

Tekijä(t) – Author(s)

Petteri Pitkänen, Sami Partamies, Ari Luukkonen VTT Building and Transport

Toimeksiantaja(t) – Commissioned by Posiva Oy

Nimeke – Title

HYDROGEOCHEMICAL INTERPRETATION OF BASELINE GROUNDWATER CONDITIONS AT THE OLKILUOTO SITE Tiivistelmä – Abstract

Olkiluoto at Eurajoki has been selected as a repository site for final disposal of spent nuclear waste produced on Finland. An understanding of the hydrogeochemical groundwater conditions and evolution is essential in evaluating the long-term safety of the repository. The performance of technical barriers and migration of possibly released radionuclides depend on chemical conditions. A prerequisite for understanding these factors is the ability to specify the water-rock interactions, which control chemical conditions in groundwater. The objective of this study is to interpret the processes and factors, which control the hydrogeochemistry, such as pH and redox conditions. A model of the hydrogeochemical evolution in different parts of the crystalline bedrock at Olkiluoto has been created and the significance of chemical reactions and groundwater mixing along different flow paths calculated. This baseline concept of hydrogeochemistry will perform as a reference in evaluating changes resulting from the ONKALO construction and finally in evaluating hydrogeochemical conditions after the closure of repository This interpretation and modelling are based on water samples obtained from Baltic sea, precipitation, eight groundwater observation tubes in the overburden, seven shallow boreholes and thirteen deep boreholes in the bedrock for which a comprehensive data set on dissolved chemical species and isotopes was available. Since the previous interpretation report (Pitkänen et al. 1999a) 86 new samples have been collected. Analyses of dissolved gases and fracture calcite and their isotopic measurements were also utilised. The data covers the bedrock at Olkiluoto to a depth of 1000 m. The results from groundwater chemistry, isotopes, petrography, hydrogeology of the site, geomicrobial studies, PCA and speciation calculations were used in the evaluation of evolutionary processes at the site. The geochemical interpretation of water-rock interaction, isotope-chemical evolution and mixing of palaeo water types were approached by mass-balance calculations (NETPATH). The PHREEQC program was used in the interpretations of the pH and Eh conditions. The hydrochemical data and previous interpretations (Pitkänen et al. 1999a) have already revealed the complex nature of sources of salinity and chemical evolution at the Olkiluoto site. The new data, since 1999, strengthen the previous concept of hydrogeochemical evolution. Actual deviations are not observed compared to the previous interpretations but, clearly, more detailed hydrogeochemical information has been obtained, particularly from shallow depths (the first 10 m), the deep saline groundwater zone (below 500 m), and dissolved gases. Repeated samplings in the bedrock during the years show only minor changes, thus strengthening the hydrogeochemical conceptual model of the site. Changes in past climate and geological environment have caused great variability in the hydrochemical data. Infiltration and mixing of end-members seem to yield a stratified hydrochemical system at least in the water-conducting fracture system. Water-rock interaction, such as carbon and sulphur cycling and silicate reactions, buffer the pH and redox conditions and stabilise groundwater chemistry. Hydrogeochemical interpretations and chemical and isotopic calculations indicate that pH seems to be dominantly controlled by thermodynamic equilibrium with calcite in fractures and there are indications that it may also occur in the overburden layer. Oxic redox conditions, prevailing in recharging groundwater, change abruptly to sulphidic conditions close to the surface, generally in overburden. Methanogenesis may show increasing importance in saline groundwater, although the major part of methane and other hydrocarbons are thermal in origin. The baseline hydrogeochemistry seems to provide the conditions for long-term geochemical stability, which is a requirement for safe long-term disposal of nuclear waste. The concept of hyrogeochemical evolution forms the basis to evaluate changes resulting from the ONKALO construction and finally in evaluating hydrogeochemical conditions after the closure of repository. Avainsanat - Keywords groundwater chemistry, environmental isotopes, nuclear waste disposal, palaeohydrogeology, water-rock interaction, mixing, geochemical modelling

ISBN ISBN 951-652-121-5

ISSN ISSN 1239-3096

Sivumäärä – Number of pages 159

Kieli – Language English

Posiva-raportti – Posiva Report Posiva Oy FIN-27160 OLKILUOTO, FINLAND Puh. 02-8372 (31) – Int. Tel. +358 2 8372 (31)

Raportin tunnus – Report code

POSIVA 2003-07 Julkaisuaika – Date

February 2004

Tekijä(t) – Author(s)

Petteri Pitkänen, Sami Partamies, Ari Luukkonen VTT Rakennus- ja yhdyskuntatekniikka

Toimeksiantaja(t) – Commissioned by Posiva Oy

Nimeke – Title

GEOKEMIALLINEN POHJAVESIOLOSUHTEIDEN PERUSTILATULKINTA OLKILUODON TUTKIMUSPAIKALLA

Tiivistelmä – Abstract

Eurajoen Olkiluoto on valittu käytetyn ydinpoltoaineen loppusijoitupaikaksi Suomessa. Hallitun loppusijoituksen kannalta tutkimusalueen hydrogeokemiallisen luonteen ja kehityksen ymmärtäminen on eräs keskeisimpiä välineitä arvioitaessa pitkäaikaisturvallisuutta. Hydrokemian tuntemus antaa perusteet loppusijoitustilojen teknisten päästö-esteiden toimintaturvallisuuden optimointiin ja mahdollisten vapautuneiden radionuklidien kulkeutumisen arviointiin. Tutkimuksen tavoitteena on arvioida ja tarkentaa aikaisempaa kuvaa Olkiluodon hydrogeokemiallisista olosuhteista, niihin johtaneista ja ohjaavista tekijöistä kohteen eri osissa. Perustilamalli tulee muodostamaan vertailukohdan arvioi-taessa ONKALOn ja loppusijoitustilojen sulkemisen jälkeisiä olosuhteita

Pohjavesien geokemiallinen tulkinta ja mallinnus perustuu vesinäytteisiin, jotka on otettu merestä, sadannasta, kah-deksasta maaperään asennetusta vesiputkesta, seitsemästä matalasta ja kolmestatoista syvästä kallioreiästä. Vuoden 1999 tulkintaraportin (Pitkänen et al. 1999) jälkeen on tullut 86 uutta vesinäytettä. Näytteistä on ollut käytettävissä laajat liuenneiden kiinteiden aineiden, kaasujen ja isotooppien analyysitulokset. Pohjavesievoluutioon vaikuttaneiden prosessien tarkastelussa on hyödynnetty alueen hydrogeologisia, mineralogisia ja geomikrobiologisia havaintoja sekä hydrogeokemiallisen pääkomponenttianalyysin ja termodynaamisten tasapainolaskujen tuloksia. Pohjavesi-kallio-vuorovaikutus, vesityyppien sekoittuminen ja isotooppikemiallinen kehitys on tulkittu massatasapainomallinnuksen (NETPATH) avulla. Pohjavesinäytteiden pH:n ja Eh:n tulkintaan on käytetty PHREEQC –ohjelmistoa.

Hydrokemian aineisto ja sen tulkinta vahvistavat aikaisempaa käsitystä pohjaveden suolaisuuden alkuperästä ja geokemiallisesta kehityksestä. Ratkaisevia poikkeamia ei ole havaittu. Sen sijaan entistä yksityiskohtaisempi käsitys on muodostunut hydrogeokemiallisesta kehityksestä maanpinnan lähellä, syvästä suolaisen veden vyöhykkeestä ja liuenneista kaasuista. Laaja vaihtelu, joka havaitaan kemiallisessa aineistossa, on seurausta muinaisista muutoksista ilmastossa ja geologisessa ympäristössä. Niiden seurauksena kehittyneet suolaiset vesityypit ovat eri aikoina suo-tautuneet ja osin sekoittuneet sekä tuottaneet nykyään rakoverkostossa havaittavan kerroksellisen pohjavesi-järjestelmän. Hiilen ja rikin kierto sekä silikaattireaktiot puskuroivat ja stabiloivat pohjaveden pH- ja redox-olosuhteita. Kalsiitti kontrolloi tulosten perusteella pH-olosuhteita kalliossa ja mahdollisesti myös maaperässä. Aero-biset olosuhteet, jotka vallitsevat suotautuvassa vedessä muuttuvat pikaisesti anaerobisiksi ja sulfidisiksi lähellä maanpintaa, osin jo maaperässä. Metaaniset olosuhteet, joita metanogeneettiset bakteerit ylläpitävät, vallitsevat syvällä suolaisessa pohjavedessä, vaikka suurin osa havaitusta metaanista ja muista hiilivedyistä onkin termistä alkuperää. Vallitsevat hydrogeokemialliset olosuhteet ja käsitys niitä kontrolloivista tekijöistä näyttävät suosivan ydinjätteiden loppusijoituksen pitkäaikaisturvallisuutta ja luovat edellytykset arvioida jatkossa ONKALON ja loppusijoituksen aiheuttamia hydrogeokemiallisia muutoksia.

Avainsanat - Keywords pohjavesikemia, isotoopit, ydinjätteen loppusijoitus, paleohydrogeologia, vesi-kalliovuorovaikutus, sekoittuminen, geokemiallinen mallinnus ISBN ISBN 951-652-121-5

ISSN ISSN 1239-3096

Sivumäärä – Number of pages 159

Kieli – Language Englanti

Posiva-raportti – Posiva Report Posiva Oy FIN-27160 OLKILUOTO, FINLAND Puh. 02-8372 (31) – Int. Tel. +358 2 8372 (31)

Raportin tunnus – Report code

POSIVA 2003-07 Julkaisuaika – Date

Helmikuu 2004

PREFACE

This study is a part of description of the baseline conditions at Olkiluoto site and belongs to the programme for development of a geologic repository at Olkiluoto, Finland, for spent nuclear fuel conducted by Posiva Oy.

This study was performed in VTT Building and Transport. The contact persons were Mia Mäntynen from POSIVA Oy, Margit Snellman from SROY Oy, Petteri Pitkänen from VTT Building and Transport. The authors wish to thank Margit Snellman and Mia Mäntynen for their comments on the draft.

We are particularly grateful to Dr Mel Gascoyne (Gascoyne GeoProjects Inc.) for his review and comprehensive suggestions to improve the manuscript.

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TABLE OF CONTENTS Abstract Tiivistelmä Preface TABLE OF CONTENTS................................................................................................ 1 1. INTRODUCTION................................................................................................. 3 2. GEOLOGICAL SETTING .................................................................................... 5

2.1 Bedrock lithology........................................................................................ 5 2.1.1 Regional scale................................................................................ 5 2.1.2 Local features................................................................................. 8

2.1.3 Fracture minerals ........................................................................... 8 3. HYDROGEOLOGICAL SETTING ..................................................................... 11

3.1 Hydraulic environment ............................................................................. 11 3.2 Groundwater flow indications ................................................................... 14

4. STRATEGY OF GEOCHEMICAL MODELLING ............................................... 17

5. GROUNDWATER CHEMISTRY ....................................................................... 19

5.1 Water samples ......................................................................................... 19 5.2 Sampling and chemical analyses............................................................. 20 5.3 Representativity of hydrochemical data ................................................... 24 5.4 Hydrochemical characteristics ................................................................. 25

5.4.1 Principal components of groundwater compositions.................... 26 5.4.2 Main variables and components .................................................. 31 5.4.3 Trends of main cations................................................................. 34 5.4.4 Trends of main anions.................................................................. 38 5.4.5 Stable isotopes (δ2H and δ18O) .................................................... 44 5.4.6 Tritium (3H) ................................................................................... 46 5.4.7 Strontium isotopic signatures ....................................................... 47 5.4.8 Chlorine isotopes (36Cl and δ37Cl) ................................................ 48

5.5 Hydrogeochemical constraints related to pH and redox conditions ......... 50 5.5.1 Trace elements ............................................................................ 51 5.5.2 Carbon isotopes and carbonate geochemistry............................. 58 5.5.3 Sulphur and oxygen isotopes of aqueous SO4 (δ34S(SO4) and δ18O(SO4)) .................................................................................... 63 5.5.4 Microbes....................................................................................... 66 5.5.5 Dissolved gasses ......................................................................... 67

5.6 Thermodynamic constrains of the groundwater system........................... 75

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6. HYDROGEOCHEMICAL EVOLUTION ............................................................. 81 6.1 Origin of groundwaters and palaeohydrogeology .................................... 81 6.2 Evolutionary processes ............................................................................ 86

6.2.1 Chemical characteristics in shallow (< 30 m depth) groundwater 86 6.2.2 Brackish and saline groundwater (depths below 30 m) ............... 94

7. MASS-BALANCE MODELS .............................................................................. 97

7.1 Introduction .............................................................................................. 97 7.2 Initial conditions........................................................................................ 99

7.2.1 Phases and constraints used in modelling................................... 99 7.2.2 Isotopic calculations and initial values for carbon and sulphur

isotopes...................................................................................... 100 7.2.3 Initial waters ............................................................................... 102 7.2.4 Flow paths.................................................................................. 106

7.3 Results of mass-balance calculations .................................................... 108 7.3.1 Isotopic evolution and mass transfer in redox processes .......... 113 7.3.2 Mixing and palaeohydrogeological implications ......................... 116

7.4 Discussion.............................................................................................. 119 8. CONCLUSIONS .............................................................................................. 121 9. REFERENCES................................................................................................ 125 APPENDICES:.......................................................................................................... 137

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1. INTRODUCTION

The characterisation, interpretation and understanding of groundwater geochemistry form an essential part of repository performance assessment and safety analysis of radioactive waste disposal. Corrosion of canisters, stability of bentonite buffer and transport of radionuclides in groundwaters may be affected by several adsorption, desorption, dissolution and precipitation processes. In order to understand and characterise these reactions in future, a model describing the water-rock interactions present during the interaction controlling the basic hydrochemistry is needed. The ultimate goal is to create a site-specific model that reliably describes changes in groundwater composition and explains their causes. More importantly, the need for evaluating and understanding the undisturbed baseline conditions is essential before the construction of ONKALO (the Finnish underground rock characterisation facility) is launched. This study discusses the geochemical conditions at the Olkiluoto site in SW Finland, and verifies the interpretations of hydrogeochemical evolution presented in a report by Pitkänen et al. (1999a) using new hydrogeochemical material collected from the Olkiluoto site between years 1999 and 2002. The new groundwater sampling, in contrast to the older data, has been accompanied by considerable data from shallow depths, dissolved gases and salinity distribution from deep groundwaters. These subjects have key roles in this baseline evaluation approach and are likely to be most disturbed during the construction of ONKALO. The relatively long sampling history of the deep groundwater samples also enables evaluation of the groundwater sampling results as a function of time i.e. �long-term monitoring�. This may decrease data uncertainty and give information of groundwater reservoirs, which are also important for the underground construction phase. The description of hydrogeochemical data includes a pragmatic statistical approach (principal component analysis) as well as a more detailed approach using binary X-Y plots. Contour plots have been used as an approach to determining essential parameters from the hydrogeochemical data. Interpretation of geochemical relationships within the data is studied by both mass-balance and thermodynamic solubility geochemical modelling approaches. The objectives of this study are to:

- Evaluate and supplement the processes responsible for evolution of different types of groundwater that control the pH and Eh conditions in the bedrock interpreted in the previous report by Pitkänen et al (1999a).

- Evaluate in detail the hydrogeochemical behaviour of the shallow groundwater samples

- Evaluate in detail the nature, behaviour and origin of the dissolved gasses from the groundwater samples.

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- Observe the changes in salinity level over long-term pumping event.

- Monitor the possible indications of transitions in groundwater composition and water-type storage observed during long-term pumping.

The thermodynamic controls on water composition and trends constraining these processes are interpreted in light of chemical, isotopic, petrographic and hydrological data. The mass-balance model NETPATH (Plummer et al. 1994) is used to determine which reactions can explain changes along the flowpath during the evolution of the groundwater system, and to what extent these reactions take place. The thermodynamic code PHREEQC (Parkhurst & Appelo 1999) is used to interpret pH and redox values. Throughout this report sample "depth" means borehole length, i.e. our notation ignores the effect of borehole inclination and true sample depth is always lower than the noted length indicates. However, because the dip of boreholes is steep (>70°) this simplification does not affect the interpretation of data as the accuracy of depth relationships is to within tens of metres.

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2. GEOLOGICAL SETTING

2.1 Bedrock lithology

2.1.1 Regional scale

Olkiluoto is situated 10 km west of the village of Eurajoki and 13 km north of the town of Rauma (Fig. 2-1). The bedrock surrounding the Olkiluoto site is Precambrian in age (Fig. 2-2). The oldest rocks in the area are schists and gneisses deformed and metamorphosed during the Svecofennian orogeny 1900�1800 Ma ago. Large areas to the east and south-east of Olkiluoto Island are covered with roughly 1570 Ma old rapakivi granites (Vaasjoki 1977). The unmetamorphosed Satakunta sandstone formation north-east of the Olkiluoto area accumulated over 1400�1300 Ma (Simonen 1980) and has been protected from erosion in a graben-like depression of the crust. Obviously related to block movements, the youngest rocks in the surroundings of the Olkiluoto site are olivine diabases. The main rock types of Olkiluoto island are (Fig. 2-3), in order of abundance: mica and veined gneisses, migmatite granite, grey gneisses, and diabase. Minor veins and dikes are quartz feldspar gneisses and amphibolites in composition. A recent summary of the petrographic features of the main Olkiluoto rock types can be found Pitkänen et al. (1999a, p. 7�10) and POSIVA (2003).

Figure 2-1. Location of the Olkiluoto site in south-western Finland. Geological domain division after Korsman et al. 1997. A = accretionary arc complex of southern Finland, B = accretionary arc complex of central and western Finland, C = primitive arc complex of central Finland, D = intracratonic and craton margin sequences and intrusions, E = Archean basement.

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Figure 2-2. Precambrian bedrock in the surroundings of Olkiluoto island (modified after Korsman et al. 1997). 1 = migmatitic mica gneiss, 2 = grey gneisses, 3 = microcline granite, 4 = rapakivi, 5 = Satakunta sandstone, 6 = olivine diabase.

The average chemical compositions of the main rock types at the Olkiluoto site have been summarised earlier in the reports of Pitkänen et al. (1999a, p. 11; 2001, p. 9). The average chemical compositions show significant differences in Fe2O3, S and U contents between the main rock types of the site. The Fe and S contents are highest in mica gneiss and amphibolites due to the higher amount of mafic minerals and sulphides in these rock types than in the others. The U content is higher in migmatite granites than in other rock types. Migmatite granites contain clearly least S. These variations in average compositions between the rocktypes may leave a characteristic fingerprint to the compositions of groundwater samples.

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Figure 2-3. Lithological map of the Olkiluoto research site (POSIVA 2003).

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The certain average compositions of the main mineral in different rock types of the Olkiluoto site have been tabulated in Pitkänen et al. (1999a, p. 12). From these earlier studies, it can be concluded that biotite in mica gneisses paleosomes contains dominantly iron over magnesium, and as gneiss grades into to amphibolite gneiss, biotite becomes more phlogopitic (Mg-rich). Similarly, hornblende in amphibolite gneiss is magnesium rich in composition. Plagioclase compositions are mostly oligoclase in mica gneiss and quartzite gneiss, and andesine in amphibolite gneiss. Although not quantitatively analysed, mica and veined gneiss contain pyrite, graphite, pyrrhotite and traces of other sulphides. In places, sulphides can be found in gneisses as bands and dissemination (Paulamäki & Paananen 1996). In Olkiluoto grey gneisses and migmatite granite, biotite contains clearly more Fe and less Mg similarly to mica gneisses. Plagioclase compositions in grey gneisses are mostly andesine and in migmatite granite oligoclase. 2.1.2 Local features

Until end of 2002, 23 boreholes (KR1�KR23) have been drilled into the research site (Fig. 2-3). Of these, groundwater sampling together with geological data collected from boreholes KR1�KR13 are utilised in this study (Appendix 1). In the future studies, the more detail consideration of hydrogeochemical data together with the geological observations is necessary. The geological map in Figure 2-3 is largely based on outcrop, investigation trenches, drilled core (KR1�KR20), geophysical ground and low altitude airborne geophysical measurements. Veined gneiss gradually turns into mica gneiss. Therefore, although veined gneisses are considered to have formed at higher metamorphic conditions than mica gneisses, postmetamorphic thrusting has probably not brought these rock types side by side, but migmatisation has occurred during complex deformation and veining has favoured suitable pressure (and temperature) anomalies (Paulamäki 1995). From the averagely-fractured Olkiluoto bedrock (away from the fracture zones) POSIVA (2003) have presented fracture frequency statistics. The results from drill cores KR1�KR18 indicate that for borehole sections 0�500 m there are about 2.03 fractures per metre (open: 0.10/m; filled: 1.23/m; tight: 0.72/m). For greater than 500 m depths the total fracture frequency estimate diminishes to 1.65 fractures per metre (open: 0.02/m; filled: 1.13/m; tight: 0.51/m).

2.1.3 Fracture minerals

Mica and veined gneisses contain ubiquitously calcite, sulphides and clay minerals as fracture minerals (Gehör et al. 1996, 1997, 2000, 2001a, b, Kärki & Lahdenperä 2002). Very sulphide-rich interlayers have also been found in KR3 at a hole depth of 250 m, where sulphide content varies in the range 1.1�5.0 wt-% (Paulamäki & Paananen 1996). Table 2-1 presents a summary of average coating surface areas and thicknesses for calcite, pyrite and pyrrhotite based on fractures in 12 drill cores.

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Table 2-1. Average coating area (%) and coating thickness (mm) of fracture minerals in the detected fractures of the 12 studied drill cores (based on Gehör et al. 1996, 1997, 2000, 2001a, b, Kärki & Lahdenperä 2002). The coating calculation identifies the mineral filling observations of each fracture in the upper 500m of a drill core (from Luukkonen et al. 2003). The amount of observations is weighted with the total number of fracture obervations, and then with the average coating areas detected from the drilled fracture sections. The average thickness is the average of mineral coating thicknesses detected from individual fractures. The bottom line indicates the overall averages of the tabulated values.

Drill core Calcite Pyrite Pyrrhotite Area (%) Thickness (mm) Area (%) Thickness (mm) Area (%) Thickness (mm)OL-KR1 21.1 0.23 3.3 0.16 0.0 0.04 OL-KR2 33.5 0.45 4.8 0.18 - - OL-KR3 21.5 0.18 10.3 0.18 0.1 0.10 OL-KR4 16.5 0.22 6.3 0.12 0.4 0.07 OL-KR5 34.5 0.22 5.6 0.13 0.2 0.11 OL-KR6 36.8 0.37 2.7 0.22 0.1 0.22 OL-KR7 23.3 0.33 11.7 0.20 - - OL-KR8 32.7 0.44 9.9 0.18 0.2 2.0 OL-KR9 20.6 0.44 6.0 0.41 0.2 0.80 OL-KR10 27.8 0.36 10.5 0.19 - - OL-KR11 31.8 0.80 5.9 0.31 0.3 0.11 OL-KR12 37.4 0.25 6.8 0.14 0.4 0.10 Tot Average 28.1 0.36 7.0 0.20 0.2 0.39

Calcite is found at all drilled depths. It usually forms multi-phased crystallisations and occurs as thin layered coatings (thickness < 1 mm) on fracture surfaces. According to the summary given in Table 2-1, about 28% of the surface area of fractures in the upper 500m of bedrock is covered with a 0.4 mm thick calcite cover. On petrogenetic grounds, Gehör et al. (2002) identified 9 different classes of calcite at Olkiluoto. In the view of low-temperature groundwater chemistry, the main interest is on the evidence of potential low-temperature deposition of calcite. According to Gehör et al. (2002) the late stage calcites form physically homogeneous, scaly layers, and in a few cases, thin layers composed of idiomorphic crystals. Chemically, these calcites are practically stoichiometric (CaCO3) without significant impurities, and the lack of fluid inclusions is considered as evidence of slow crystallisation rates under cool conditions. Further carbon, oxygen and strontium isotope studies with selected calcite samples from moderate depths (~100�200 m) indicate that the analysed late-stage calcite precipitations have occurred at low temperatutes but in methanic redox conditions (Karhu 1999, 2000). Currently, methanic redox conditions are found at greater depths in SO4

-poor, brackish and saline waters of the Olkiluoto bedrock. Therefore, it seems that the studied calcites have been crystallised before the infiltration of the sulphate-rich water currently dominant at this depth range. The measured radiocarbon ages support this interpretation ages over 43 000 years before present have been found for two selected calcite samples, Gehör et al. 2002, Karhu 2000.

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Dissolved carbonate in current groundwater types seems not to be in isotopic equilibrium with fracture calcites, as Blyth et al. (1998) have also reported using fluid inclusion and δ18O data of calcites and groundwater. It should be noted that young calcite in equilibrium with groundwaters may be, because it is difficult to focus sampling of fracture calcites specifically on the last generation, which may be a very thin layer on top of earlier precipitates and can be easily lost during drilling. However Blomqvist et al. (1992) have not observed any low δ13C values indicating CH4 -derived carbonate (< -30�), therefore the δ13C value in calcites indicates that there has been no significant CH4 oxidation by possible oxidative glacial melt-water. These results also do not suggest any significant isotopic exchange between calcite and groundwater. The reason for this imbalance is likely due to organic -derived carbonate in groundwater, because a mere trace of oxidised methane would decrease the δ13C of DIC significantly because of the low DIC content of brackish and saline groundwater types. The result probably emphasises high rates of microbiologically mediated redox processes compared with potential equilibrium between DIC and calcites by isotope exchange. Iron sulphides are found, as well, in all drilled cores and at all depths. Usually sulphide occurs as small deposits of granular pyrite, which may form thin patches or coatings. The thickness of coatings varies from barely observable to 1 mm. On the average it is estimated that about 7% of the area of fractures in the upper 500m of the bedrock is covered with a 0.2 mm thick pyrite cover (Table 2-1). The less common pyrrhotite is frequently found in association with graphite (Gehör et al. 1996, 1997, 2000, 2001a, b, Kärki & Lahdenperä 2002) which in turn is the usual fracture mineral in slickensided fractures. About 0.2% of the area of fractures in the upper 500m of bedrock is covered with a 0.4 mm thick pyrrhotite cover (Table 2-1). As a rule when graphite and pyrrhotite are detected in a fracture, these minerals are found in the surrounding host rock as well. Clay mineral coatings are among the most frequent infills of the Olkiluoto fractures. The thickest coatings are at the most 3�5 mm but normally coatings are less than 1-2 mm in thickness (Gehör et al. 2001). According to Kärki & Lahdenperä (2002) about 34% of the surface area of fractures detected in drill cores KR2, KR3, KR4, KR5, KR7, KR8 and KR10 is covered with a 0.2 mm thick clay cover. Only kaolinite exhibits fine-grained loose powdery variations. Otherwise, clay minerals are usually more crystalline indicating that they have crystallised in more or less hydrothermal or metamorphic conditions. Corroded cavites, i.e. signs of fracture mineral dissolution, are found occasionally in all drilled cores. Previously, some of corroded cavities have been related to fracture having high hydraulic conductivity (e.g. Gehör et al 1996). However, in more recent interpretations this relationship is rejected (e.g. Gehör et al. 1997, 2000, 2001a, b). High-conductivity fractures contain similar mineral assemblages as low-conductivity fractures. Evidently, mineral dissolution and precipitation have been occuring during the whole evolution history of the fractures. As studies with calcite confirm (Gehör et al. 2002), the dissolution-precipitation processes usually have been incomplete and a fracture may contain mineral generations and corroded cavities throughout its geological history.

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3. HYDROGEOLOGICAL SETTING

3.1 Hydraulic environment

The Olkiluoto site at Eurajoki is a coastal island on the Gulf of Bothnia (Figs. 2-1, 2-2). The area is fairly flat, the surface being mostly less than 5 metres above average sea level (m.a.s.l.) and the highest point, located in the SW part of the island, about 18 m.a.s.l. From a hydrological viewpoint (cf. Fig. 2-3) natural hydraulic gradients within the bedrock that are a function of topography differences, must be gentle. The watershed interpreted to begin from the WNW end and terminate at the SE end of Olkiluoto island roughly divides the island into two parts (e.g. Teollisuuden Voima Oy, 1992). The topography is somewhat more varied in the southern part, where boreholes KR4, KR7, KR8 and KR 23 are located, than in the flatter northern part where boreholes KR2, KR5, KR6, KR13, KR19 and KR20 have been drilled. Based on earlier groundwater geochemical studies (e.g. Pitkänen et al. 1994, 1996, 1999a) it is clear that the Quaternary history of the Baltic Sea has had major effects on groundwater compositions at the Olkiluoto site. It is likely that Olkiluoto was covered with a thick ice sheet during the very cold Saale glaciation, which started about 200 000 years ago (Eronen & Lehtinen 1996). The rapidly warming climate which followed some 130 000 years ago was the start of the Eem interglacial period. The climate was distinctly warmer than at present and the Eem Sea was more saline than the present Baltic Sea. According to interpretations based on indirect evidence, Olkiluoto was above sea level during the late Eem stages (Eronen & Lehtinen 1996). At least 75 000 years ago the site was again ice covered by the Weichselian glaciation. Complete Weichselian deglaciation started about 10 000 years ago, and soon Olkiluoto was released from the ice cover but remained below the surface of the mildly saline Yoldia sea. The Olkiluoto site also remained below water level during the stages of the fresh Anchylus lake (starting some 9 500 years ago) and the saline Litorina Sea (starting around 7 500�7 000 years ago) (Fig. 3-1). As a result of land uplift Olkiluoto island begun to emerge from the Baltic Sea about 3 000�2 500 years ago (Eronen & Lehtinen 1996). Currently the post-glacial land uplift at the study site is about 4�6 mm/yr (Eronen et al. 1995, Kakkuri 1987). From a hydrological viewpoint, natural hydraulic gradients in the bedrock related to salinity (i.e. density) differences of groundwater are potentially important.

12

Figure 3-1. Postglacial shoreline in southern Finland from about 9500 BP until present (after Eronen et al. 1995). Table 3-1 presents roughly estimated compositions of glacial melt-water during Weichselian deglaciation and of Litorina seawater. With the exception of 3H isotope values, the Eem Sea most likely resembled the present ocean water composition. The present Baltic Sea as well as mean global ocean water compositions are estimates for surficial seawaters.

Olkiluoto Island is covered mostly with Quaternary soils and only about 4% of the land area is characterised by bedrock outcrops. The soil consists mainly of washed sandy till with an average thickness of about 2�5 m. In a few bog depressions, the thickness of peat is usually less than 1 m (Anttila et al. 1992). The artificial lake of Korvensuo lies in the middle of the Olkiluoto area (Fig. 3-2).

13

Table 3-1. Estimated Quaternary glacial melt and Litorina seawater compositions with inferred Baltic sea and mean ocean water compositions.

1)Glacial water

2)Litorina sea

3)Baltic sea

4)Ocean water

T (°C) 1.0 10.9 8.5 20.0 O2 (mg/l) 7.2 6.6 7.2 4.3 pH 5.8 7.6 7.7 7.5 Density (g/ml) 1.000 1.008 1.002 1.030 HCO3 (mg/l) 0.16 92.5 78.7 144.2 SO4 (mg/l) 0.05 890 450 2540 PO4 (mg/l) 0.0003 0.06 0.02 0.22 Ntot (mg/l) 0.19 0.27 0.21 0.5 Cl (mg/l) 0.70 6500 3025 19550 F (mg/l) 0.00 0.49 0.27 1.3 Br (mg/l) 0.001 22.2 10.3 67 NO3 (mg/l) 0.07 SiO2 (mg/l) 0.01 1.84 0.58 6.61 Fetot (mg/l) 0.0001 0.002 <0.01 0.002 Al (mg/l) 0.0001 0.002 <0.01 0.002 Na (mg/l) 0.15 3674 1760 10860 K (mg/l) 0.15 134 66 391 Ca (mg/l) 0.13 151 82 412 Mg (mg/l) 0.1 448 219 1310 Mn (mg/l) 0.0 0.0 <0.01 0.0 Sr (mg/l) 0.0001 2.68 1.20 8.24 Li (mg/l) 0.0 0.07 0.04 0.18 Charge Balance (%) 1.03 0.97 2.13 0.27 δ2H (o/oo SMOW) -166.0 -37.8 -60.8 -30.0 δ13C (PDB) -25.0 -1.0 -1.68 -1.0 δ18O (o/oo SMOW) -22.0 -4.7 -7.55 -4.0 3H (TU) 0.0 0.0 15.4 15.4 14C (pM) 28.0 43.0 115.8 100 87Sr/86Sr 0.70940 0.70945 0.7094

1) pH, Na, K, and SO4 values estimated from Taylor et al. (1992). δ18O and δ2H values are discussed in section 6.1.3. 14C value is based on conservative isotopic decay and δ13C value indicates organogenic origin of carbon. 2) Regression between average Baltic sea (Puskakari, Eteläriutta) and global mean ocean water (Fairbridge 1972, Harrison 1992) with assumption that Cl-content in Litorina sea has been about 6500 mg/l (Kankainen 1986). 14C value is based on conservative isotopic decay. 3) Average of �94 samples from Puskakari and Eteläriutta. 4) Global mean ocean water(Fairbridge 1972, Harrison 1992).

14

3.2 Groundwater flow indications

Conceptual flow simulation models for the Olkiluoto site have been developed by Koskinen (1992) and Löfman (1996, 1999, 2000). Furthermore, attempts have been made to measure natural groundwater flow directions and intensities from boreholes KR1�KR4 and KR7�KR10 (Rouhiainen 1997, Pöllänen & Rouhiainen 1997). The long-term measurement of the flow and electric conductivity from the deep boreholes has been performed by Pöllänen & Rouhiainen (1996a, b, 2000, 2001a, b and 2002a, b) and Rouhiainen (2000). The method was used for fast determination of hydraulic conductivity and hydraulic head in fractures or fracture zones. The results reveal several anomalous features from the boreholes suggesting possibly divergent characteristics for the interpreted structures. The flow simulations try to model groundwater flow under natural conditions in the fractured bedrock. The 1996 published model (Löfman 1996) is based on the structural bedrock model by Saksa et al. (1993) while the 1998 model (Löfman 1999, 2000) is based on a modification of the bedrock model (Saksa et al. 1996). Both evaluations attempt also to predict changes in groundwater flow conditions over time. The models utilise the present groundwater table, designed geometries for fracture zones, hydraulic properties of bedrock, impacts of land uplift, and salinity variations of groundwater. Interpreted structures of the bedrock model and salinity variations of groundwater play a clear central role in considerations of flux and groundwater flow in the bedrock of the research area (Löfman 1999, 2000). A summary interpretation of groundwater flow directions in the '96 bedrock model is presented in Pitkänen et al. (1999, p. 19�23). A ground level section and a 3-D illustration of a recent 2001/2 Olkiluoto structural model (Saksa et al. 2002) is presented in Figure 3-1. Naturural flow measurements across a borehole in fractures have beenmade with a flowmeter (Rouhiainen 1977, Pöllänen & Rouhiainen 1977). Unfortunately, the measurements with the flowmeter have not been very successful. The majority of transverse flow measurements have failed, mostly due to inadequate sensitivity of the equipment. The current sensitivity is about 1 ml/h for flow across the hole, which corresponds to a Darcian flux value of about 2·10-9 m/s.

15

Figure 3-2. Ground level horizontal section and the 3-D illustration of the Olkiluoto structural model version 2001/2 (Saksa et al. 2002). The 3-D illustration is illuminated from south.

17

4. STRATEGY OF GEOCHEMICAL MODELLING

The modelling approach in this investigation can be described as follows: Sampling points along a flow path in a groundwater system will show progressive changes in chemical and isotopic properties. These changes may reflect processes such as mixing of groundwater types, mineral dissolution and precipitation, ion-exchange, gas-exchange and redox reactions. The goal in chemical modelling is to interpret the probable processes and the quantity of material transferred that are responsible for the evolution of the groundwater composition, as well as any chemical changes that occur in the groundwater composition as a result of these processes. In this work, the overall strategy in modelling is a three-step procedure:

- The first step is to analyse the trends of hydrochemical and isotopic data (Sections 5.4, 5.5 and 6) that constrain the possible processes behind the trends.

- The second step is to determine the tendency of the groundwater to dissolve or precipitate minerals as reflected by the fracture observations and saturation indices (Section 5.6).

- The third step is to derive reaction models that can explain the changes in water chemistry between any points along a chosen flow path by mass-balance calcu-lations (Chapter 7).

The saturation index (SI) is used to evaluate the potential behaviour of minerals at groundwaters in the second step. The SI of a particular mineral that may be reacting in the system is defined as

SI = Log IAP/KT, (Eq.4-1) where IAP is the ion activity product of the mineral and KT is the thermodynamic equilibrium constant adjusted to the temperature of the given analysis. SI is greater than zero for over-saturation, and less than zero for under-saturation indicating precipitation and dissolution for a particular mineral, respectively. Thermodynamic solubility calculations are also used to test internal consistency between interpreted and measured pH and redox conditions. The PHREEQC (version 2) program (Parkhurst & Appelo 1999) is used in solubility calculations. The NETPATH program version 2.0, (Plummer et al. 1994), is used to interpret net geochemical mass-balance reactions between initial and final water along a hydrologic flow path in the third step. The program utilises defined chemical and isotopic data for groundwater samples and calculates those mixing proportions of initial waters and reaction coefficients of the chemical sinks and sources (minerals and gases) that account for the observed changes between an initial and final water. In addition, the program calculates isotopic evolution according to each mass-balance model to predict isotopic composition at the end-point of the flow-path.

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5. GROUNDWATER CHEMISTRY

5.1 Water samples

The hydrogeochemistry of groundwater samples collected in the Olkiluoto area has been discussed in several studies (e.g. Snellman 1991; Pitkänen et al. 1992; Blomqvist et al. 1992; Lampén & Snellman 1993 and refs. therein; Pitkänen et al. 1993, 1994, 1996a; 1999, Snellman et al. 1995a; Tuominen 1994; 1998; Ruotsalainen & Snellman 1996 and refs. therein). These reports also describe sampling points and methods, preparation of samples, analytical methods in the field and in the laboratory, results of water analyses and evaluation of prevailing hydrogeochemical conditions and processes. The most recent sampling methods and samples used in this study have been described by Karttunen et al. (1999, 2000), Mäntynen (1999), Karttunen & Mäntynen (2001), Paaso & Mäntynen (2002) and Rantanen et al. (2002). The sampling procedures follow qualified field manual for water sampling of POSIVA by Paaso et al. (2003). Pitkänen et al. (1999) already evaluated the data sampled before 1997 and recommended reliable data for further studies in hydrogeochemistry. That data were combined with samples collected after 1997. Altogether 181 samples were available for this study. The data were checked (see also ch 5.2-5.3) and 128 water samples (3 of the Baltic seawater, 3 of precipitation, 51 of shallow groundwater (groundwater observation tubes and shallow boreholes) and 71 of deep groundwater), sampled from 78 different points during 1988-2002 from the Olkiluoto site, were used in the interpretations. The analytical results of the new samples since Pitkänen et al. (1999a) report are given in Appendix 2. Appendix contains only revised data, published since Pitkänen et al. (1999) report. Complete hydrochemical baseline data will be revised and published as a whole during 2004. The ID codes (Table 5-1) in Appendix 2 for the samples give an indication of the sample types and various sampling methods. As an exception from the previous report (Pitkänen et al. 1999), the ID coding of all but the "T" and "S" samples (Table 5-1) has changed. In addition to the geographical name or borehole number, the running number is included in the ID code. Furthermore, the depth of the upper packer is added to the ID codes of deep boreholes. Special attention is paid to samples from KR6, where it is possible to follow the changes of the parameters as function of time due to repeated sampling during years 2001 and 2002 from depths of 98.5-100.5 (KR6/99/1-3) and 135-137 metres (KR6/135/1-3) (Paaso & Mäntynen 2002 and Paaso et al. 2003). In the sampling procedure implemented, the first samples were taken from the packered sections without pumping from the upper part of the borehole. This was followed by pumping from the open borehole to turn the groundwater flows towards it. The sampling of the sections were taken synchronously with pumping from the upper part of the borehole. This was to simulate the change in flow resources the seawater or the saline water intrudes to the hole due to the efficient (20 l/min) pumping. In all cases the prepumping of the packer sections preceded the sampling lasting considerably shorter before the first sampling than the followings.

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Table 5-1. ID codes for water samples from Olkiluoto and surrounding areas.

ID code Sample type Sampling method Only geographical name of sampling point and the running number (e.g. Varvinnokka/1)

The Baltic Sea Ruttner sampler

PVP (e.g. PVP1) Shallow groundwater in the overburden observation tube

Submersible pump

PR (e.g. PR4) Shallow groundwater in the bored well

Submersible pump

PP (e.g. PP7) Shallow groundwater in the cored well

Submersible pump

KR /T (e.g. KR2/T7. The 7th packer section is the uppermost and the 1st one is the deepest. T refers to Finnish equivalent for packer).

Deep cored borehole with multipackers

Slim membrane pump

KR /P (e.g. KR3/243/1. P=PAVE) the upper packer depth and the running number form the ID code (e.g. KR3/243/1)

Deep, open cored borehole

Double packer system with membrane pump and PAVE (see section 5-2)

OL-KR /S (e.g. KR1/S1) (Sampled during preliminary site investigations beyond 1992)

Deep, open cored borehole

Double packer system with membrane pump

Three deep groundwater samples from the borehole KR11 (KR11/277/1, KR11/415/1 and KR11/621/1) are all excluded from the modelling procedure due to representativity problems. The results suggest that these sample sequences have reformed to the reverse order as a function of depth that could be estimated on the basis of the geochemical data from other boreholes. The gas results of the sample KR11/621/1, for instance represent deep, highly saline methanic environment, contradict with salinity results representing brackish sulphidic marine derived environment, which support the uncertainties in representativity. According to the report by Karttunen et al. (2000) there have been certain difficulties during the field sampling period as well. Furthermore, the preliminary salinity results (TDS monitoring) from the resampling of borehole KR11 have revealed the existence of the presumed water types comparable to the regional phenomenon. 5.2 Sampling and chemical analyses

Preliminary site investigations for spent nuclear fuel disposal started in the Olkiluoto area in the late 1980s. Hydrogeochemical site investigations (Lampén & Snellman 1993) produced a general characterisation of the local deep groundwaters, precipitation, surface waters and shallow groundwaters in the surrounding area. Groundwater samples from deep, open boreholes were taken with a double packer - membrane pump instrument (Rouhiainen et al. 1992) (S samples in Table 5-1).

21

Deep groundwater sampling During detailed site investigations in 1994-1995, groundwater samples were taken from five deep boreholes (KR1...KR5; 500-1000 m long). To prevent mixing of different bodies of groundwaters, multipackers were installed in the Olkiluoto boreholes in 1992 (Hinkkanen & Oksa 1992), enabling a hydrological steady-state to be reached before groundwater sampling began in 1994 (T samples in Table 5-1). Groundwater samples were pumped with a slim membrane pump (Öhberg 1991; Ruotsalainen et al. 1994) from the packed-off, sections of the deep boreholes. Seven sampling sections between 6-104 m long were isolated with inflatable rubber packers in each borehole. The hydraulic head in each section was constantly monitored for possible effects of sampling (Niva & Ruotsalainen 1995). The long sampling sections could cause mixing of the different water types (Anttila et al. 1999). However, sufficient hydraulic conductivities in the packer intervals were generally limited to a few short sections in a constricted borehole length contributing the vast majority of groundwater extracted during sampling (App. 2 in Pitkänen et al. 1999a). The quality of the groundwater was monitored by continuous on-line measurements of pH, electrical conductivity, Eh(Pt), dissolved O2 and temperature in a combination of flow-through cells. Since the capacity of the slim membrane pump was quite small (max. 100 ml/stroke) and the natural hydraulic conductivity of groundwater sampling sections was often very low (T generally of the order 10-8-10-4 m/s), actual pumping rates ranged from roughly 3 to 133 ml/min. Thus, the pumping period before the groundwater had reached an expected condition generally took 3-5 weeks, during which time roughly 100�3200 l was pumped from each of the sampling sections. About one week before the planned sampling, the amount of remaining drilling water (based on uranine, an organic dye used as a drill water tracer) and tritium activity was checked to evaluate representativity and possible mixing with younger groundwaters. The criteria for representativity are shown in Table 5-2.

Table 5-2. Criteria for representativity of groundwaters (Ruotsalainen & Snellman, 1996)

Parameter(s) Applied criteria in evaluation of representativity

Void volume in sampling section and tubes

Water changed at least 2-3 times where possible

Remaining drilling water

< 2.5%

Field measurements of electrical conductivity and pH

Stabilised values

Field measurements of Eh and O2

Stabilised, lowest possible values

Tritium If no hydrogeological cause (fracture zone), below the detection limit of the direct activity count method (6-8 TU)

All groundwater samples experience a pressure change as they are pumped up to ground level, and dissolved gases are given off. Development work on groundwater sampling equipment that maintains the in situ pressure was started in 1993 (Ruotsalainen et al.

22

1996a). The PAVE sampling system was developed, tested, and subsequently utilised during the groundwater sampling campaigns after 1997 in the study of boreholes KR2, KR3, KR4, KR6, KR7, KR8, KR9, KR10, KR11, KR12 and KR13. During the pumping period, the chemical quality of groundwater (pH, electrical conductivity, dissolved oxygen, redox potential (Pt, Au) and temperature) is monitored on-line with electrodes installed in flow-through cells at the surface as described above. Groundwater bypasses the pressure vessels to preclude accumulation of microbial biofilms, drilling debris and other fine material on the inner walls. When the representativity of the groundwater is good, sampling for field and laboratory analyses will begin. In the final sampling step, the valves of the pressure vessels are opened to increase the hydraulic pressure in the pressure hose of the packers. Next the argon (Ar) or nitrogen (N2) gas in the pressure compartment of the pressure vessel is compressed, the piston moves downwards, and the sample section is filled with groundwater at constant in situ pressure. Groundwater is pumped through the pressure vessels for several hours to ensure good representativity. The valves are then closed and the PAVE equipment is brought up to ground level by winch. The pressure vessels are sent to the laboratories for groundwater analyses. Shallow groundwater sampling After 1997, 16 new shallow groundwater observation tubes were constructed and sampled (together with the older PVP-, PP- and PR-holes), mainly to get representative data for evaluation of infiltration and groundwaters in the possible discharge and recharge areas (Lehto 2001). The analysis programmes were adapted by omitting the redox-sensitive and gas parameters, due to the risks of atmospheric contamination with this sampling technique (Hatanpää 2002, Backman et al. 2002). Analytical programmes Analytical programmes were adapted according to the representativity of the samples and aims of the study. These are detailed in Table 5-3. Hydrochemical data used in this study is given in Appendix 2. Some parameters are infrequently analysed and particularly the data of lathanides, 34S(S2-),222Rn, 232Th, 228Th, 230Th, 226Ra, 228Ra are rare and insufficient relative to the variability of groundwaters at Olkiluoto

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Table 5-3. Analytical programmes used by POSIVA in the analysis of the groundwaters of the Olkiluoto area.

Parameters Deep groundwaters Springs and shallow groundwaters

Measurements and analyses in the field laboratory (on-site)

pH, electrical conductivity, Eh, O2, T, alkalinity, acidity, S2-

tot, Fe2+, Fetot, uranine (=tracer of the drilling water), Cl, Br, F, SO4, NO2, NO3, PO4 (NH4, since 1997) (anions not analysed in the field in the latest samplings)

pH, electrical conductivity, alkalinity, acidity S2-, Fe2+ and Fetot (in case of PR and PP samples)

Main physico-chemical variables (off-site)

pH, electrical conductivity, density, KMnO4, Fetot, Stot, Ntot, Ptot, Btot, DIC (=Dissolved Inorganic Carbon), DOC (=Dissolved Organic Carbon), SiO2, uranine

Same as groundwaters, with the exception of DIC

Cations (off-site) Na, K, Ca, Mg, Mn, Al, Sr, NH4

Same as groundwaters

Anions (off-site) Cl, Br, F, I, NO3, NO2, SO4, PO4


Trace elements (off-site) Rb, Ba, Cs, Li, lanthanides


Evacuated (surface) and dissolved (IN SITU = PAVE) gases (off-site)

O2, CO, CO2, H2, N2, CH4, C2H2, C2H4, C2H6, C3H6, C3H8, He, Ar

-

Isotopes in dissolved gases (off-site)

18O(CO2), 13C(CO2), 13C(CH4), 13C(C2H6), 13C(C3H8), 2H(CH4)

-

Isotopes in water (off-site)

2H, 3H, 18O, 18O(SO4), 34S(SO4), 34S(S2-), 13C(DIC), 14C(DIC), 36Cl, 37Cl, 222Rn, 232Th, 228Th, 230Th, 226Ra, 228Ra, 87Sr/86Sr, 234U/238U, 238U

2H, 3H, 18O 13C(DIC), 14C(DIC), 18O(SO4), 34S(SO4), 222Rn, 87Sr/86Sr, 234U/238U, 238U

Isotopes in particles (off-site)

234U/238U, 238U 234U/238U, 238U

Microbes Direct counting, cultivation of: autotrophic acetogens, heterotrophic acetogens, sulphate reducers, iron reducers, autotrophic methanogens, heterotrophic methanogens, oligonucleotide probes

-

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5.3 Representativity of hydrochemical data

Due to the very low hydraulic conductivities of many sampling sections, and the low capacity of the slim membrane pump used in the multipacker boreholes, transport times to the surface for some groundwater samples could be several hundred hours during the sampling from multi-packered boreholes (Ruotsalainen & Snellman 1996). The part of the polyamide sampling tube above groundwater level is in direct contact with air and diffusion of atmospheric gases through the walls of the tube (Snellman et al. 1995a, b) could cause the anomalously high Eh measurements that are sometimes observed. This may have affected O2--and CO2 -sensitive parameters, mostly alkalinity, Eh, dissolved O2, acidity, gases, Fe2+, Mn, S2-, 13C and 14C. It is known from many previous studies that the natural amount of dissolved carbonates is extremely low in the deep saline groundwaters at Olkiluoto (e.g. Pitkänen et al. 1992, Lampén & Snellman 1993). Thus, the reliability of the carbon isotopic data obtained must be carefully evaluated. Since 1997, the effect of atmospheric gases was eliminated as far as possible by shielding the sample tube with another polyamide tube (Ruotsalainen et al. 1996b). Also, the stability of field measurements was enhanced by installing the flow-through cells in a small N2-flushed box (Ruotsalainen et al. 1996a, 1998)improving also redox measurements (both Au and Pt �electrodes in use since 1999). The representativity of hydrogeochemical data from the five deep multipackered boreholes at Olkiluoto and from springs and shallow borehole wells, and of data from the preliminary investigations during 1987-1992, was evaluated using a procedure developed by Ruotsalainen & Snellman (1996). This process was updated and has been utilised in sampling since 1997 and the evaluation procedure does not reduce the amount of samples; it merely guides the selection process between parallel field and laboratory results. The PAVE sampling for gases is evidently proven to have considerably better intake than the other sampling methods used. The descriptions and evaluation of the alternative other methods and the sampling history of gas sampling from groundwater are discussed in Pitkänen et al. (1999a). Electric conductivity (EC) measurements The electric conductivity measurements of borehole water and fracture specific water can be used to estimate the distribution of total dissolved salinity, expressed as TDS (Öhberg & Rouhiainen 2000) and therefore in evaluation of the representativity of groundwater samples. The EC measurements are carried out together with the flow measurements (Pöllänen & Rouhiainen 1996a, b, 2000, 2001a, b and 2002a, b, and Rouhiainen 2000). The EC measurements of borehole water are based on pumping of the borehole to achieve flow from the formation towards the borehole. In addition to the EC of borehole water, EC of fracture specific water was measured from certain selected fractures (the last values in time series in App. 1). Pumping volumes vary between EC measurements and groundwater sampling, therefore comparison should be done only on

25

qualitative level. The trends of salinity changes during pumping are important for the interpretation of hydrogeological condition of measurement section. The Borehole logs of KR1-KR13 are presented in Appendix 1. The electrical conductivity results are calculated as TDS values by correction to 25 °C using the SFS-EN 27888 standard. The temperature corrections of the borehole KR13 was used the model by Heikkonen et al. (2001). The correspondence between borehole water and fracture specific water in EC measurements is generally good (note scale variation in TDS logs between the boreholes in App. 1). Deviations are most prominent in the boreholes KR1, KR7, KR8 and KR9 where water in boreholes seems to dilute in certain sections and KR10 (and KR1) in which deeper parts the water concentrates considerably during the pumping period comparing to the borehole water. The dilution of the measurement section indicates the migration of more diluted water from shallower depths into greater depths, for example along vertical fractures. The enrichment indicates opposite or slow recovery of fracture from borehole water contamination. The correspondence of TDS between fracture specific measurements and chemical sampling is fairly good, although the differences are larger than between EC measurements themselves. The results so support generally the representativity of hydrochemical samples (in situ). However, certain deviations are distinct. Boreholes KR1, KR3, KR4, KR8, KR9, KR10 and KR11 show typical behaviour where deeper in the borehole groundwater sampling get a more saline input than the EC measurements. The reason may be slow recovery due to small yield of deep fractures. Groundwater samples taken from KR6 during long term pumping experiment are also more saline than EC measurements indicate. Monitored sections do not seem to have connections to dilute groundwater aquifer at shallow depth. Opposite condition are observed in some occasional sampling sections (KR2/T3, T2 and KR5/T2) and particularly in KR12. Representativity problems were observed earlier in the samples from KR2 and KR5 (Pitkänen et al. 1999a). The EC results of KR12 are on an anomalously high level. This deviates strongly from the results of the other boreholes, which correspond to the TDS results of groundwater sampling from KR12. The sedimentation pools are above KR12. They may cause the exceptional TDS distribution in borehole water if they are in hydraulic connection to the lower part of the borehole.

5.4 Hydrochemical characteristics

After evaluation of the hydrochemical data, a grouping (Table 5-4) was developed based on previous hydrochemical results and conclusions, which were evidently connected to changes in the palaeohydrogeological environment (Table 5-5; Pitkänen et al. 1996a, 1999a). This grouping divides the data according to the origin of main water component in a groundwater sample and is used mainly in figures in chapter 5 and presented for each sample in Appendix 2. The palaeo hydrogeological origin of different groups is discussed more in chapter 6.

26

Table 5-4. Grouping of the samples considered in this study.

Group Water type Samples Overburden Ca-Mg-Na-HCO3-SO4-(Cl) Groundwaters from overburden observation tubes.

Dilute or brackish HCO3 groundwaters

Na-Ca-Cl-HCO3-(SO4) Groundwaters from shallow wellwells and deep boreholes

Brackish SO4 groundwaters

Dominantly Na-Cl with notable SO4

Groundwaters from deep boreholes

Brackish Cl groundwaters

Na-Cl, SO4-poor Groundwaters from deep boreholes

Saline groundwaters

Ca-Na-Cl Groundwaters from deep boreholes

Baltic

Seawater Seawater samples off Olkiluoto island

"Litorina"

Seawater Estimated from description in Table 3-1

"Glacial" Fresh Estimated from description in Table 3-1 Table 5-5. Main water types at Olkiluoto (Pitkänen et al. 1999a). Depth, m Salinity Water type Origin of dominant end-

members Age estimate,

BP 0-150 Fresh-slightly brackish

Dominantly Na-Cl-HCO3

Precipitation Baltic seawater

modern - 2 500

100-300 Brackish Na-Cl-(SO4) Litorina seawater

2 500 - 7 500

100-500 Brackish Na-Cl Pre-Litorina water containing fresh glacial melt-water

7 500 - 10 000

> 500 Saline Ca-Na-Cl Preglacial meteoric water, influenced by hydrothermal saline fluids

>> 10 000

5.4.1 Principal components of groundwater compositions

Studied variables in geological data are frequently variously correlated with each other. Examination of a single variable against another seldom gives a good quick overview of the data and the study requires multiple variable considerations. Generally speaking, multiple variable analysis (carried out as a principal component analysis) aims to find a new set of independent variables (oriented at right angles to each other), reduce statistical noise related to input variables, identify outlier samples, and attach together similar samples and variables in a matrix. As a result of analysis, the first component always contains the most obvious information (and accounts for the largest variance) of the data matrix. Subsequent components describe gradually diminishing features, and finally only the noise of the data. Usually the goal of a principal component analysis is to summarise a multivariate dataset as accurately as possible with as few principal components as possible. A usual demand is that a principal component analysis should

27

explain at least 75%, and as high as 90%, 95% or even 99% of variances of the original variables (Jöreskog et al. 1976, Reyment & Jöreskog 1996). In this study, the choice and processing of geochemical variables is similar to studies done with data from the Äspö Hard Rock Laboratory (Laaksoharju 1995, Laaksoharju et al. 1995, Laaksoharju & Wallin 1997), in which Na, K, Ca, Mg, HCO3, Cl, SO4, 3H, δ2H, δ18O concentrations were used. In addition, Br concentrations are available for the Olkiluoto studies. It is assumed that most of the variability related to groundwater composition can be described with these variables. Our principal component analysis is based on the correlation matrix presented in Table 5-6. Several comments can be made directly from the correlation matrix: Na, Ca, Br and Cl have high mutual correlations, indicating that these elements are likely to be combined together; Mg and SO4 have somewhat more significant correlation with each other than K to Mg and SO4; and δ18O and δ2H correlate with each other. The only variables which lack absolute correlations above 0.5 (i.e. less than 25% of the variable�s behaviour can be explained with another variable) are HCO3

- and 3H.

Table 5-6. Pearson correlation matrix based on 93 samples for chemical variables used for principal component analysis. Absolute correlations above 0.5 are shown in bold type. Only samples with all variable values available were accepted from Appendix 2.

Na K Ca Mg HCO3 SO4 Cl Br 3H δ18O δ2H Na 1.000 K 0.360 1.000 Ca 0.911 0.207 1.000 Mg 0.333 0.752 0.169 1.000 HCO3 -0.415 -0.076 -0.279 -0.130 1.000 SO4 0.009 0.548 -0.151 0.869 0.052 1.000 Cl 0.967 0.279 0.985 0.251 -0.347 -0.079 1.000 Br 0.935 0.228 0.991 0.170 -0.318 -0.173 0.991 1.000 3H -0.455 0.024 -0.282 -0.224 0.083 -0.179 -0.357 -0.309 1.000 δ18O 0.034 0.361 0.066 0.333 -0.160 0.301 0.062 0.055 0.427 1.000 δ2H 0.256 0.427 0.346 0.378 -0.125 0.225 0.326 0.340 0.272 0.853 1.000 The loadings of chemical variables onto principal components are shown in Table 5-7. For present purposes, a coefficient of at least 90% of the original variance was considered adequate. This condition is fulfilled with four principal components, all of them exhibiting significant eigenvalues greater than one (cf. Davis 1973, 1986). Table 5-7 shows that with this explanation ratio all variables are heavily loaded onto one principal component, and therefore give their significant contribution to the components in question. Variable communality, i.e. the part of each variable variance, which is explained in this four component model, is shown in the right column of Table 5-7. As communalities indicate, explanation ratios for K and 3H are the lowest, indicating that these components also cause unexplained noise in the analysis.

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Table 5-7. Principal component loadings of geochemical variables in a varimax rotated principal component model, based on Table 5-6. Absolute loadings above 0.65 are shown in bold type. Component score of Fig. 5-2 is calculated with the aid of tabulated component coefficients. As an example, a score for the principal component P1 is calculated as follows: PC1 = -0.0074 + 1.089E-4×[Na+] + 2.643E-3×[K+] + 1.034E-4×[Ca2+] - 3.686E-4×[Mg2+] + 9.745E-4×[HCO3

-] - 6.074E-4×[SO42-] + 3.279E-5×[Cl-]

+ 4.423E-3×[Br-] - 8.381E-3×[3H] - 1.907E-2×[δ18O] + 8.084E-3×[δ2H].

Varimax rotated principal component loadings

PC1 PC2 PC3 PC4 CommunalityNa 0.938 0.197 -0.084 -0.209 0.969K 0.229 0.759 0.269 0.033 0.702Ca 0.985 -0.008 0.064 -0.043 0.976Mg 0.156 0.957 0.077 -0.079 0.953HCO3 -0.235 -0.011 -0.060 0.967 0.994SO4 -0.176 0.936 0.009 0.012 0.907Cl 0.986 0.085 0.011 -0.118 0.993Br 0.992 -0.006 0.047 -0.083 0.9923H -0.367 -0.275 0.728 0.063 0.744δ18O 0.017 0.291 0.886 -0.128 0.886δ2H 0.314 0.289 0.833 -0.004 0.876

Eigenvalues 4.202 2.657 2.104 1.029

Percent of total variance explained 38.20 24.16 19.13 9.36

Cumulative %: 62.35 81.48 90.84

Component score coefficients for non-standardised variables PC1 PC2 PC3 PC4 m -0.0074 -0.3972 5.9849 -1.0265 Na (mg/l) 1.089E-4 1.884E-5 -3.471E-5 -3.254E-5 K (mg/l) 2.643E-3 2.431E-2 5.194E-3 8.944E-3 Ca (mg/l) 1.034E-4 -2.505E-5 1.472E-5 5.158E-5 Mg (mg/l) -3.686E-4 6.087E-3 -1.071E-3 -6.961E-4 HCO3 (mg/l) 9.745E-4 2.480E-4 1.188E-4 9.135E-3 SO4 (mg/l) -6.074E-4 2.467E-3 -6.011E-4 -3.131E-5 Cl (mg/l) 3.279E-5 -2.835E-6 -3.167E-7 6.460E-6 Br (mg/l) 4.423E-3 -1.068E-3 4.631E-4 1.505E-3 3H (TU) -8.381E-3 -2.047E-2 4.777E-2 3.983E-3 δ18O (oo/o SMOW) -1.907E-2 1.513E-2 2.770E-1 -6.323E-2 δ2H (oo/o SMOW) 8.084E-3 1.352E-3 4.190E-2 1.003E-2

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Figure 5-1. Ratios of rotated principal component loadings in vector plot presentations (cf. Table 5-7).

Loadings of Table 5-7 are shown graphically in Figure 5-1. The first principal component (PC1) contains heavy loadings of Ca, Br, Cl and Na. These variables are relatively tightly bound together and reflect quite considerable salinity variations and common sources in the groundwaters of Olkiluoto site. Table 5-7 shows that HCO3 is slightly and negatively loaded onto PC1. There are also certain other indistinct relations. In general, it seems that as the salinity of groundwater increases, concentrations of SO4 and 3H tend to fall and K and δ2H to increase. Essential loadings building up the second principal component (PC2) originate from Mg and SO4, which are usually related to seawater, as well as K, which has also a substantial loading onto PC2. According to Table 5-7, as the seawater character in samples increases, it may be expected that Na, δ2H and δ18O values tend to increase, and 3H values to decrease indistinctly. The third principal component (PC3) is heavily loaded with δ18O and δ2H, and somewhat less distinctively with 3H. The former two variables are temperature sensitive and reflect palaeo conditions: sea and surficial waters tend to have the highest values, and waters variously modified from these exhibit diminished values. 3H is a residence time indicator. As the present surficial or seawater character increases, the 3H concentration usually rises. The fourth principal component is HCO3 sensitive. The principal sources of HCO3 are generally near surface processes such as decay of organic matter and

30

fracture calcite dissolution. At the same time, all other variables are insignificantly loaded onto PC4. This suggests that causes affecting the HCO3 concentration (i.e. carbonate balance) are mostly independent of variables that were related above to components PC1�PC3. The indistinct exception is PC1, which indicates that in saline waters the HCO3 content is low. To summarise, if a sample scores significantly to PC1 it indicates strong content of old saline water in the sample. Similarly, high scoring to PC2 indicates that the sample contains significant amounts of either recent or relict seawater. High scoring to PC3 indicates low residence time of sea or surficial waters in the bedrock. Finally, if a sample exhibit high scores in respect of PC4 it is likely taken from shallow depths below the surficial overburden.

Overburden Dil.-brack.-CO3 Brackish-SO4 Brackish-Cl Saline Baltic

Figure 5-2. Principal component scores for studied samples presented in 3-D.

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The general view of principal component scores for the samples used for the analysis is shown in Figure 5-2. Samples are viewed in 3D to give an oblique view of the PC1�PC2 plane and demonstrate the effect of PC3 and PC4 on different groups of samples. It is evident that on the PC1�PC2 plane there are two general trends among borehole samples. One limb extends from the neighbourhood of overburden and dilute brackish samples towards high PC1 values and the other towards high PC2 values. According to our earlier conclusions the former is related to high salinity groundwaters and the latter to seawaters. The sample population shows that young overburden waters and Baltic Seawater are significantly scored with PC3. On the other hand, the gap between present seawater and seawater type groundwaters in the bedrock indicates that seawaters related to bedrock samples are distinctly older than present seawater. As expected, mainly dilute brackish, and to lesser extent overburden, groundwaters from shallow depths tend to exhibit high scores in respect of PC4, indicating that while CO2 and CaCO3 are dissolved and H+ ions are consumed, HCO3

- concentrations in waters increase. Comparison of the current principal component analysis to that previously published (Pitkänen et al. 1999a, Figure 5-2, p. 38) shows that it can be concluded that the results are almost exactly the same. The earlier analysis was done with 61 samples and the present analysis is done with 95 samples (data included environmental and monitoring samples but did not include parallel samples). This comparision underlines the possibilites and robustness of principal component analysis as long as the explanation ratio of the analysis is kept high enough and variable matrix contains enough variation (as described in first paragraph of this chapter). The sensitivity of principal component analysis to changes in the amount of principal components and variables has been considered in more detail in Pitkänen et al. (1999a, p. 37). The 32 new samples have not generated any new characteristics in the principal component loadings. This fact supports the conclusion that the essential features of the groundwater chemical trends have been captured for the Olkiluoto bedrock prior to disturbances caused by the tunnel and shaft excavation disturbances. The trend features for the earlier 61 samples have been considered in more detail in Pitkänen et al. (1999a, p. 39�40).

5.4.2 Main variables and components

Chloride is considered a conservative ion for the interpretation of hydrogeochemistry, once it is in the aqueous phase (Edmunds et al. 1985; Nordstrom et al. 1985). Pitkänen et al. (1999a) verified the irreversible dissolution of Cl in the groundwater system at Olkiluoto by using solubility calculations. Plots of ion vs. Cl concentrations are especially convenient for estimating mixing and dilution processes. Therefore, most of the graphs in the following sections present variables respective to Cl as well as to show spatial variation with depth. In the graphical presentations, the samples have been grouped according to the principles described in the beginning of section 5.4.

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TDS The TDS values of the unreported new samples correlate with the reported ones (Pitkänen et al. 1999a) (Fig 5-3a). All the overburden samples are determined to be fresh water (TDS<1000 mg/l) (Davis 1964), correspondingly the rest of the samples are brackish (1000 mg/l < TDS < 10 000 mg/l) or saline (TDS > 10 000 mg/l) (Fig 5-3a). The TDS maximum (appr. 73000 mg/l) is reached in sample KR4/860/2. TDS values have an increasing trend as a function of increasing depth. The enrichment of salinity increases strongly below about 400 m, suggesting the effects of deep, highly saline groundwater, which has apparently been influenced by hydrothermal salts (Pitkänen et al. 1996a). A similar increase of the TDS vs. depth slope has been observed in the Canadian Shield where Gascoyne et al. (1987) have interpreted it to indicate a change of hydrological condition from advection to diffusion ie. to fairly stagnant groundwater conditions. The multi-sampled specimens, from borehole KR6 can be classified as a brackish type of water (App. 2). However, there is considerable increase in the TDS values during the extended sampling period according to the Paaso & Mäntynen (2002). This suggests mixing of the more saline water in to the hole, due to pumping of the open borehole. pH The pH ranges between 5.2 and 8.8 showing clear increasing trend as a function of depth (Fig 5-3b). The pH values are more acid near the surface, due to the influence of atmospheric and biogenic CO2. The scattering is a characteristic phenomenon for most of the water types, especially for the overburden and saline type waters. The pH values seem to be slightly lower in deep saline groundwaters in the new samples than in the previous samplings. The pH values from KR6/99/1-3 descend half a unit as Cl increases (Fig 5-4a), whereas, KR6/135 stays fairly stable. However, the decrease in KR6/99 follows the change between the water types.

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Figure 5-3. Depth distribution of a) TDS values and b) pH values of Olkiluoto groundwater samples augmented with new unreported data. SiO2 The dissolved SiO2 content reaches its maximum values in shallow waters and decreases as a function of increasing Cl suggesting the dissolution of silicates in more acid and oxic environments near the surface (Fig. 5-4c). The values of KR6/99/1-3 show gently increasing trend and the values of KR6/135/1-3 do not show any clear trend as a function of pumping volume (Fig. 5-4d). The values in KR6/135 represent high level in brackish SO4 �rich groundwaters.

34

Figure 5-4. pH as a function of chloride for a) all and b) selected (Cl<7000 mg/l) Olkiluoto groundwater samples. SiO2 as a function of chloride c) all and d) selected (Cl<7000 mg/l) Olkiluoto groundwater samples. Symbols and arrows marked with red colour emphasises transition of samples KR6/99/1-3 and with blue colour transition of samples KR6/135/1-3 5.4.3 Trends of main cations

Na Sodium shows a positive correlation as a function of increasing Cl (Fig 5-5a) suggesting a common hydrogeochemical source for these ions. Mixing is proposed to be the controlling process of Na concentration due to the linear behaviour of the different water types. However, the deep saline water data are depleted considerably in Na compared to the seawater dilution line. The steep growth from KR6/99/1 to KR6/99/2 combines with seawater dilution line but the trend deviates in the results of the KR6/99/3 sample (Fig 5-5b) whereas the samples KR6/135/1-3 show somewhat linear behaviour during the pumping.

35

Figure 5-5. Na as a function of chloride for a) all and b) selected (Cl<7000 mg/l) groundwater samples. Ca as a function of chloride for c) all and d) selected (Cl<7000 mg/l) groundwater samples. Mg as a function of chloride for e) all and f) selected (Cl<7000 mg/l) groundwater samples. Symbols and arrows marked with red indicates the transition of samples KR6/99/1-3 and with blue transition of samples KR6/135/1-3. The solid line presents seawater dilution line.

36

Ca Calcium is highly enriched when compared with the seawater dilution line (Fig 5-5c) suggesting a partial rock origin. Calcium is assumed to dissolve from fracture calcite, but may be released from plagioclase as well (Pitkänen et al. 1992, 1994, 1996a, 1999a. However, the linear growth as function of Cl suggests that mixing mainly controls Ca concentration and these elements have a common source in saline groundwaters, i.e. an ancient brine component. The Ca-concentration of the KR6 samples also increases as function of chloride (Figure 5-5d). Mg Magnesium is slightly enriched compared with the seawater dilution line in fresh groundwaters whereas all the brackish and saline samples show Mg depletion relative to the seawater (Fig 5-5e, f). The early enrichment in fresh samples is propably caused by mineral weathering during infiltration. The brackish, SO4 -rich water increases linearly, and is slightly depleted in Mg relative to seawater. The saline water type is strongly depleted in Mg correspond with seawater, its distinguishing clearly own low-Mg part of the diagram. However, the Mg concentration seems to increase in the most saline waters which may reflect the higher Mg-content in the original saline end-member (Pitkänen et al. 1999). According to Pitkänen et al. (1996a), Mg content is enriched in the brackish SO4-rich groundwater at depths of approximately 100-300 m (Fig. 5-5e) with the strongest indication of ancient, Litorina-stage seawaters. During the long-term pumping in KR6 Mg shows a general increase suggesting enrichment of a marine component in the groundwater samples. K Potassium resembles Mg in its behaviour. It is clearly enriched when compared with seawater dilution line in fresh groundwaters suggesting weathering as a source process. There is a moderate increasing trend of K in the saline groundwater samples. This may reflect the complex geochemical processes between K and clay minerals such as illite but, as well, K may originate from a brine component. Both KR6/99/1-3 and KR6/135/1-3 show increasing K content as a function of increasing chloride concentration (Fig 5-6b). The K content of the most saline samples shows notable variation. Sr The behaviour of Sr is similar to Ca because of their similar geochemical properties - the values increase as function of chloride and are highly enriched relative to seawater (Fig. 5-6c) suggesting the same geochemical source. The Sr values of KR6/99/1-3 and KR6/135/1-3 show linear growth and are steeper in the KR6/99 �samples (Fig 5-6d). More generally, the changes of cations in the pumping sections of KR6 tend to represent extreme values of the water type clusters. Sodium plots on the below the other values whereas Ca, Mg, K and Sr tend to plot above them. However groundwater types, i.e HCO3 -rich and SO4 -rich, have internal variance. Therefore, it is not possible to

37

estimate whether these extreme values are generated by pumping or if they are due to conservative mixing of previous SO4 �rich groundwater component which initially represents, slightly extreme composition among the water type. Saline groundwater cannot explain the salinity increase due to the high values of Mg neither current Baltic seawater due to the high values of Ca.

Figure 5-6. Potassium as a function of chloride for a) all and b) selected (Cl<7000 mg/l) groundwater samples. Sr as a function of chloride for c) all and d) selected (Cl<7000 mg/l) groundwater samples. Symbols and arrows marked with red indicates the transition of samples KR6/99/1-3 and with blue colour transition of samples KR6/135/1-3. The solid line presents seawater dilution line.

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5.4.4 Trends of main anions

Alktot The total alkalinity (mainly consisting of HCO3 and CO3) shows an initial increase at shallow depths and decreases as a function of increasing Cl finally reaching the analytical detection limit (Fig. 5-7a). The influence of atmospheric and biogenic CO2 probably accounts for increase in Alktot maximum values in fresh groundwater samples of the overburden and in shallow, dilute or brackish HCO3-type bedrock groundwater samples. Alkalinity may also include other components than carbon compounds e.g. silicates though their magnitude is minor in such high values. Calcite dissolution is also an evident source for alkalinity. However, any other single dissolved component, e.g. Ca, does not show such an enrichment during recharge, thus several weathering processes should dissociate CO2 to HCO3. This also explains HCO3 loading in its own principal component (PC4). The decreasing part of the trend may be due to precipitation of calcite, which may start at shallow depths at Olkiluoto (Pitkänen et al. 1994; 1996a) but may also be due to mixing of water types. The saline samples with low alkalinity can be strongly influenced even by minor CO2 uptake from the atmosphere (disscussed in ch. 5.5.2) during sampling and analysis.The risk of contamination have been minimised by using protective gas (N2) in the analyses. The Alktot values of KR6/99/1-3 show an overall increase culminating in the maximum value seen in sample KR6/99/2, while KR6/135/1-3 show a linear decreasing trend as a function of Cl (Fig 5-7b). DIC Alkalinity measurements are uncertain to evaluate total dissolved carbonate content in acid groundwaters due to dissociation of dissolved carbonate in carbonic acid and HCO3 ion. The analysis of dissolved inorganic carbon shows directly the concentration of total carbonate species. The DIC values are below 10 mg/l in acid recharge zone (Figs 5-7f, 5-8c) indicating mainly dissolved CO2 in these groundwaters. The DIC values reach similiarly to alkalinity values after initial increase the highest values in samples from overburden and dilute HCO3 �rich water and decreases rapidly below 200 meters depth and in samples of saline waters the concentration remain low. SO4 Sulphate also shows an initial enrichment as Cl increases; (Fig 5-7c, d), however the strongest increase is observed in deep bedrock groundwaters. It is noteworthy that SO4 enrichment in HCO3 groundwaters exceeds the seawater dilution line indicating an additional source other than seawater.

39

The SO4 �rich brackish water has been observed (Pitkänen et al. 1999a) to form a significant groundwater body and indicates an ancient seawater between 100 and 300 m depth (Fig. 5-8a). Pitkänen et al. (1996a, 1999a) have linked the SO4 enrichment to infiltration from the Litorina Sea based on the following: 1) the marine SO4/Cl ratio and higher salinity of SO4-rich samples than in present Baltic samples off Olkiluoto island (Fig. 5-7c), 2) the typically marine Br/Cl signature (Fig. 5-8b), and 3) the low 14C content. The microbiologically catalysed reduction of SO4 and mixing with the pre-Litorina SO4 -poor brackish groundwater results in the dramatic decrease of SO4 concentration below this SO4 -rich groundwater body. The deep saline groundwaters have negligible contents of SO4, as they apparently represent methanic redox conditions in which SO4 is reduced (Pitkänen et al. 1996a, 1999a, Ruotsalainen & Snellman 1996). The SO4 values of the samples KR6/99/1-3 increases somewhat linearly becoming approximately constant at sample KR6/99/3, while samples KR6/135/1-3 peak in SO4 concentration at KR6/135/2 (Fig. 5-7d), suggesting Litorina type SO4 -rich water instead of younger seawaters intruding into the borehole due to the pumping. Sulphide Amounts of aqueous sulphide (S2-

(tot) = S2- + HS- + H2S + polysulphides) show a relatively wide range in the Olkiluoto groundwaters (Fig. 5-7e). Small amounts of dissolved sulphide are observed in shallow fresh groundwaters from the overburden and bedrock, indicating anoxic conditions in these groundwaters. The anomalously high concentration observed (12.4 mg/l) was analysed in sample KR13/362/1. Other high sulphide concentrations (> 1mg/l) were observed (Pitkänen et al. 1999a) in KR2/T3, KR2/T2 and KR5/T2, which all are located quite near to each other. Elevated dissolved S2- concentrations seem to coincide with the transition zone from brackish to saline groundwaters. The S2-

(tot) content of the samples from KR6 are absent or very low. The redox species are discussed in more detail in a later section (ch. 5.5).

40

Figure 5-7. Alktot as a function of chloride for a) all and b) selected (Cl<7000 mg/l) groundwater samples. SO4 as a function of chloride for c) all and d) selected (Cl<7000 mg/l) groundwater samples. e) S2

tot as a function of chloride and f) DIC as a function of chloride for all groundwater samples. Symbols and arrows marked with red indicates the transition of samples KR6/99/1-3 and with blue, the transition of samples KR6/135/1-3 The solid line presents seawater dilution line.

41

Cl Chloride is the most abundant anion and it shows a strong increase in concentration as a function of depth. The maximum value ([Clmax] = 45200 mg/l; KR4/860/2) observed is more than double to current ocean water and apparently representing a deep brine of hydrothermal origin, as earlier suggested by Pitkänen et al. (1999a) and discussed in chapter 6. Br The values of Br grow increase linearly with increasing chloride content (Fig 5-9a). The HCO3-rich and SO4-rich groundwaters show mainly Br concentrations, which correspond to the seawater dilution line. The brackish groundwaters with minor SO4 content and saline groundwaters have relatively higher Br values than seawater, indicating an other salinity source (Pitkänen et al. 1994, 1996a, 1999a). In the series KR6/98/1-3 and KR6/135/1-3 samples the Br -concentration also increases along the seawater dilution line reflecting marine origin linearly as a result of the changed flow conditions for infiltrating water changing the salinity during pumping (Fig 5-9b) The first Br value in deeper section is apparently inaccurate. This and information from cations indicate that increasing amount of Litorina derived groundwater affects the observed salinity enrichment during pumping the monitored borehole sections in KR6. Figure 5-8b shows the depth distribution of (molar) Br/Cl ratios. HCO3-rich and SO4-rich samples have a coherent ratio similar to the molar Br/Cl value calculated from the mean composition of ocean water (e.g. Drever 1997) or with the present Baltic Sea, indicating a marine origin for these groundwater types. The saline groundwater group has similar Br/Cl values, which are clearly higher than the seawater value, further indicating the non-marine origin of saline groundwater. The SO4-poor brackish groundwater below the SO4-rich layer shows a similar ratio to that of saline groundwater, suggesting a common salinity source for these deep groundwater types. F Fluoride mainly shows an increasing trend as a function of chloride concentration in fresh and brackish groundwaters (Fig 5-9c). Most of the samples are highly enriched to the seawater suggesting some other source than marine origin F. Edmunds et al. (1987) suggest, that F is released together with Cl during the hydrolysis of biotite (replacing OH in the lattice), thus the initial enrichment may have developed during weathering.

Brackish Cl-type and saline groundwater types seem to show similarly higher F concentrations than HCO3- and SO4-rich types as they have a higher Br/Cl ratio suggesting a common source for all halides (Fig 5-8b). The F concentration does not increase in Ca-rich saline groundwater, which may indicate controll of F solubility by fluorite. However, no observations of fluorite have been reported in mineralogical studies (ch. 2.1.2).

42

PO4 Phosphate analyses (Fig 5-9d) have the same range of values as those reported previously (Pitkänen et al. 1999a). The great majority (52 out of 84) of the PO4 analyse, except for those of HCO3 -type water, are below detection limit. Enrichment in shallow groundwaters is probably caused by the decay of organic matter in recharge zone.

Fig 5-8. Depth distributions of a) SO4, b) Br/Cl and c) DIC. The new samples are marked by red in the diagram.

43

Fig 5-9. Br as a function of chloride for a) all and b) selected (Cl<7000 mg/l) groundwater samples. c) F as a function of chloride and d) PO4 as a function of chloride for all groundwater samples. Symbols and arrows marked with red indicates the transition of samples KR6/99/1-3 and with blue, the transition of samples KR6/135/1-3 The solid line presents seawater dilution line.

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5.4.5 Stable isotopes (δ2H and δ18O)

The contents of stable isotopes generally point to a meteoric origin for the Olkiluoto groundwater samples that are on or in the immediate vicinity of the Global Meteoric Water Line (Fig 5-10). Samples from overburden plot mainly on the line and δ18O ranges between -12 to -11 �. This was slightly lower than expected from results in Pitkänen et al. (1999a) for meteoric recharge (-10.5 to -11�). Most of the HCO3-rich groundwaters form a mixing line from overburden samples towards the Baltic seawater samples, which corresponds to increasing marine input in the water type. However, some of the HCO3 �rich samples show clearly depleted δ2H values. These may be due to analytical uncertainties, because slight differences were observed between analyses from the laboratories used by Hatanpää (2002) and Backman et al. (2002). The examination of possible analytical errors is continuing. Sulphate �rich samples form the heavy end of the mixing line towards seawater composition. Deeper in the bedrock (Fig 5-10b) the stable isotopic composition of brackish groundwaters tends to shift lighter, particularly below the SO4 �rich groundwater. The light stable isotopes in the groundwaters suggest colder infiltration temperatures than today. These light values could also indicate a rather long residence time suggesting infiltration of cold melting waters of the last (Weichselian) glaciation (Pitkänen et al. 1996a, 1999a). The lightest values have been observed in two samples from borehole KR3 (KR3/P1 and KR3/T5). Other isotopic results of these two samples (3H values below 0.8 TU, and low 14C values of 4 and 10.5 pM) are in good agreement with this the glacial hypothesis. According to the database used in the present study, five of the deep saline groundwater samples plot slightly above the GMWL (KR4/860/1, KR4/861/1, KR1/602/2, KR1/497/1 and KR10/498/1). The largest enrichment of deuterium (-49.5�) has been observed in the most saline samples. Similar observations of deep, highly saline groundwaters with stable isotopic data plotting above the GMWL have been interpreted as suggesting hydrothermal, magmatic or metamorphic origins (e.g. Frape et al. 1984, Sheppard 1986). Such groundwaters in the crystalline rocks have been observed in the Fennoscandian (e.g. Blomqvist 1990) and Canadian Shields (Frape et al. 1984). Saline samples also show some low δ2H values, which seems strange because they don�t show any other marine signature. Figure 5-10b shows the variation of δ18O values with depth. Precipitation samples at Olkiluoto (not shown in the figure) have been analysed for δ2H and δ18O contents since 1990, and 29 samples have a large range from -17 to -8�. This range is not reflected in the shallow groundwaters of the overburden, particularly not the light (cold) signatures, because the shallow HCO3 groundwaters show a fairly coherent range of -11�. This may be caused by mixing of different waters during infiltration. Deeper in the bedrock, SO4-rich samples have higher values than shallow groundwaters but also lower as well. Brackish Cl-type groundwater show mainly lower values. The buffering role of shallow HCO3 waters indicates that these shifts in both directions cannot result from current variability of precipitation; i.e. they may reflect ancient changes in infiltration conditions. High isotopic values were previously interpreted (Pitkänen et al. 1996a, 1999a) to result from mixing of Litorina Seawater (estimated value -4.7�, Table 3-1) in

45

the groundwater system, and low values from glacial melt-water infiltration (estimated value -22�, Table 3-1). Saline groundwaters show a linear shift to the right with depth suggesting δ18O value about from �9 to �10� for brine end-member. Figure 5-10c shows the variation of δ18O with Cl. The brackish SO4-rich groundwater shows a trend towards the estimated Litorina composition, supporting the interpretation above. On the other hand, the lowest δ18O values indicate mixing of a melt-water component with the SO4-rich water. Because mixing of heavy Litorina-derived water and light melt-water compensate each other to partions of these end-members may be higher in groundwater than the fairly narrow variation (4 �) of δ18O indicates. Saline groundwaters show an increasing trend with Cl. However, the reason for the scatter at higher Cl-values is unclear, particularly the exceptionally low value.

Figure 5-10. a) Contents of stable isotopes as δ2H and δ18O, b) δ18O vs. depth and c) δ18O vs. Cl in water samples from Olkiluoto. The reference line is the GMWL (δ2H = 8*δ18O + 10; Craig 1961). (The estimated �Litorina� seawater composition (δ18O=-4.7 �, δ2H=-37.8 �) is also shown in figures a and c).

46

5.4.6 Tritium (3H)

Natural low (< 10 TU) levels of tritium (3H) increased by several orders of magnitude since atmospheric thermonuclear experiments began in 1952, with the strongest increases in the 1960s. The half-life of 3H is approximately 12.4 a, therefore no 3H should be detectable in groundwaters that infiltrated prior to 1952. The chronological development of 3H activity in precipitation at different locations is described by Ruotsalainen & Snellman (1996) and in Pitkänen et al. (1999a). 3H contents in precipitation have decreased significantly in Finland since the 1970s and nowadays are almost at the 1950s levels. The complex bacground of the 3H data excludes the use of 3H as an exact dating tool for young waters, but it is an unambiguous indicator of recent water contribution if it is detected in a groundwater sample. Amounts of 3H in Olkiluoto samples decrease as Cl increases (Fig. 5-11a), suggesting minimal mixing with recent groundwaters in deeper parts of the bedrock. The current 3H content of overburden groundwaters and seawater is largely between 10 and 15 TU. The much higher seawater value is already from the 1980�s. In general, groundwater samples reflecting minor recent water input are those with the longest mean recidence times, based on activities of 14C and 3H (Fig 5-11b). As shown in figure 3H decreases to the detection limit (0.8 TU) in HCO3 �rich groundwater and most of the brackish and saline groundwater samples are below or near this value. Some elevated radiocarbon values in brackish and saline samples indicate contamination problems in DIC that is discussed in chapter 5.5.2.

Figure 5-11. 3H activity levels vs. a) Cl and b) 14C in Olkiluoto water samples.

47

5.4.7 Strontium isotopic signatures

Use of the 87Sr/86Sr signature in groundwater studies has been discussed in several papers, e.g. by McNutt et al. (1990) and Bullen et al. (1996). Although 87Sr is radiogenic (a daughter product of 87Rb decay) the extremely long half-life (5 x 1010 a) of the decay process makes the 87Sr/86Sr signature of Sr sources essentially stable on the time scale of groundwater evolution. During water-mineral interaction, the mineral�s isotopic value is reflected in the water�s isotopic value, and is dependent on the amount of Rb in the mineral or, more realistically, minerals. Rubidium is concentrated in K-rich phases such as K-feldspar and micas. Strontium is enriched in Ca-rich phases such as plagioclase and calcite due to the similar geochemical properties of Sr and Ca. Fractionation of Sr isotopes as a result of mineral precipitation is considered negligible. Generally the 87Sr/86Sr ratio in the groundwater at Olkiluoto varies between 0.7167 and 0.75264 and is much higher than in the ocean (0.7092) or modern Baltic Sea (0.70945), reflecting a radiogenic Sr source. The Sr isotope ratio shows an initial increase in diluted shallow groundwaters (Figure 5-12a), suggesting weathering of potassic silicates during recharge near the surface. Actually, according to the results achieved by McNutt et al. (1990), Wallin & Peterman (1994), Blum & Erel (1997) and Bullen et al. (1996, 1997), biotite and muscovite seem to be the most probable sources to dissolve and enrich the radiogenic Sr signature to such a high level as observed in overburden groundwaters at Olkiluoto. The Sr concentration in micas is generally one to two orders of magnitude lower than in feldspars, hence the dissolution of micas should dominate over plagioclase in early weathering during recharge in order to reach the high signature.

Figure 5-12 a) Sr isotope ratio of dissolved Sr against Cl concentration, b) Sr isotope ratio versus Sr concentration in Olkiluoto groundwater samples.

48

In the groundwater flowpath, the portion of radiogenic Sr decreases steeply in HCO3 �rich waters, and then rises a little in brackish groundwaters. Karhu (2000) suggested that calcites have Sr isotopic compositions in the range of 0.716 and 0.719. The steep decrease may so result from calcite dissolution in dilute groundwaters though from plagioclase as well. Finally, the most saline sample (KR4/861/1) shows a higher isotope ratio than all other saline samples. According to Figure 5-12b Sr isotope results seem to have three possible end members: Baltic seawater, deep saline groundwater and overburden water. Seawater mixing may cause a similar effect depletion on the Sr isotopic signature as the initial dissolution of calcite and later of plagioclase. The depletion of 87Sr may result from seawater mixing and/or dissolution of Rb-poor, Ca- and Sr-rich minerals. However, the curved decrease of the 87Sr/86Sr ratio towards brackish groundwaters (Fig.5-12b) is probably a result of dissolution of calcite and mixing of brackish groundwaters suggesting that mixing with Baltic Seawater is only minor. The strontium isotopic signature is stable in more saline groundwater samples despite an obvious increasing input of a saline end-member. This may result from a tendency to equilibrate with fracture calcites or it is controlled by plagioclase dissolution, which is promoted by calcite precipitation. On the other hand, the equilibration theory is not in accordance with the higher signature of the most saline groundwater, which in particular should have a similar Sr isotopic signature to that of fracture calcites. slightly increased value may originate from hydrothermal alteration of primary silicates. Current data do not give an unambiguous solution to the problem, in addition to which the Sr isotopic trend may emerge from both hypotheses

5.4.8 Chlorine isotopes (36Cl and δ37Cl)

The groundwater data from Olkiluoto also contain several analyses of stable chlorine isotopes. The isotopic ratio 37Cl/35Cl is a seldom-used method, but the variations in the ratio, expressed as δ37Cl (� standard mean ocean chloride, SMOC, 0�) may, according to Frape et al. (1996), reveal information about the sources of dissolved chloride. The modern Baltic and possibly palaeo-Baltic seas have a negative δ37Cl signature of about -0.21�, which is considered to be related to salt leachate from extensive Palaeozoic salt deposits south of the Baltic Sea (δ37Cl ranging from -0.58 to 0�). Chlorine bound to rock minerals can show quite a heavy signature, up to +4�, which could be released by water-rock dissolution reactions. Significant enrichment has also been observed in oceanic aerosols, up to +2.53�. The resuts of δ37Cl analyses are presented in Fig 5-13. Three out of 11 samples have duplicate samples when the average values are used. The uncertainty of the analyses is relatively high (±0.2�) given the range of analyses, but the results are considered to show an indication of potential chlorine sources at Olkiluoto. The results suggest three different end-members for Cl in Olkiluoto groundwaters. The brackish SO4-rich groundwater has the only negative values. This water type was interpreted to infiltrate from the Litorina Sea (Pitkänen et al. 1996a, 1999) and the result is consistent with the postulated (Frape et al. 1996) negative value of palaeo-Baltic seas. The chlorine signature in the most saline sample shows a seawater (ocean) origin (Fig. 5-13a). In contrast, in other brackish and saline groundwater samples δ37Cl seems to be

49

enriched, which suggests that palaeo-Baltic waters are not responsible for dilution of the brine end-member. Rather, the results indicate a third Cl source in the groundwater system. The third source should completely dominate in Na-Cl type brackish and saline groundwater (Cl<15 000 mg/l) or its original signature should be significantly higher than observed in groundwater samples. The reason for enrichment of δ37Cl could be rock derived Cl. Alternatively the enrichment could be derived from a glacial process (e.g. freezing during periglacial times), because the sample with the highest δ37Cl values has also the lowest δ18O level in the groundwater data of Olkiluoto (Fig. 5-13b). However, the Cl isotope results have to be regarded as preliminary before more extensive data are available and more knowledge can be gained from Cl fractionation.

Figure 5-13. Stable chlorine (δ37Cl) isotope results of dissolved chloride in groundwater samples at Olkiluoto as function of a) chloride and b) δ18O isotopes.

Chlorine-36 is radioactive isotope (half-life 300,000 years) produced by cosmic radiation in the upper atmosphere and in subsurface by in situ neutron flux. The 36Cl data have been utilized as a tracer for identifying mixing of groundwaters and determining the residence time of groundwater and characteristics of water-rock interaction. The 36Cl data consists of eight analyse results (Figure 5-14) reported by Gascoyne (2001). According to the report, 36Cl/Cl ratios suggests that the two SO4-rich samples (KR4/302/1 and KR8/302/1) were derived from the Litorina Sea and correspond to the low value of Baltic marine water. The 36Cl/Cl ratios of dilute, shallow groundwaters approach the values of surface and atmospheric ratios. The values of deep saline waters correspond to those calculated from in situ neutron production rates and possibly originate from the period preceding the Quaternary age (ie. prior to 1.8 Ma). This agrees with the 36Cl observation by Louvat et al. (1999) from Äspö where the deep saline waters were considered to have penetrated the host rock more than 1.5 Ma ago.

50

Figure 5-14. Variation of 36Cl/Cl relation as a function of 1/Cl.

5.5 Hydrogeochemical constraints related to pH and redox conditions

The measurement of Eh in natural waters is known to be technically difficult to achieve. General aspects of redox reactions in groundwaters and of difficulties obtaining representative Eh measurements have been discussed in several textbooks, articles and reports on hydrogeochemical characterisation (e.g. Stum & Morgan 1981, Lindberg & Runnells 1984, Melchior & Bassett 1990, Appelo & Postma 1993, Snellman et al. 1995b; Ruotsalainen & Snellman 1996) and geochemical modelling (Pitkänen et al. 1994, 1996, 1999a) in the Finnish site characterisation programme. Eh �measurements are normally made with inert Pt-electrodes. Unfortunately, the Pt-electrode responds satisfactorily to very few of the redox pairs important in natural waters. Exceptions include the ferric/ferrous iron couple if ion activities are 10-6 M (0.06 mg/l) or greater. Redox measurements may be disturbed, for example by the lack of equilibrium between different redox couples causing mixing potentials, or by electrode poisoning by precipitating solids. Similar technical problems were encountered in redox measurements of the Olkiluoto groundwater samples performed on-site and reliable redox measurements could not be obtained for all sampling sections. A rough feature of Eh(Pt) measurements at Olkiluoto is that they on average seem to decrease with increasing pH (Fig. 5-15). In part, the Eh results are very high which cannot correspond, for example with the generally observed dissolved sulphide levels in groundwater samples. Therefore, it is important to approach Eh as a quantity that is calculated from the activities of a redox pair (e.g. Drever 1997). At first it is important to study redox sensitive species, isotopes, phases and microbes mediating electron transfer in groundwaters at Olkiluoto to get a general idea of prevailing processes.

51

Significant changes in redox system may affect pH and pH buffering capacity in groundwater. For example, the content of DIC depends partly on prevailing redox processes and may significantly affect the reliability of pH measurements as already noticed earlier in hydrochemical studies at Olkiluoto (e.g. Pitkänen et al. 1992, 1999a). Many of the redox sensitive constraints are also pH sensitive and they are examined together to understand processes controlling pH and redox state in different levels of the groundwater system. Finally, in chapter 5.6, measured pH and Eh levels are evaluated using solubility calculations and values corresponding to interpreted groundwater conditions are presented.

Figure 5-15. Measured Eh values in groundwater samples vs. pH. The lower limit of water stability is set by the half-reaction H2 !" 2H+ 2e-. The lowest possible value is defined by a hydrogen pressure at 1 atm.

5.5.1 Trace elements

Iron and sulphur species

Total iron (Fetot) (Fig. 5-16a) contents are high in the very shallow samples from bedrock and the overburden. This may reflect post-oxic conditions immediately after recharge. Part of these samples also has low pH values, increasing the solubility of ferric iron, and so one should be cautious when making apparently straightforward conclusions about the redox conditions. Many shallow groundwater samples from bedrock contain dissolved sulphide (Fig. 5-16b) further indicationg that anoxic conditions have been reached already at shallow depths. Recently dissolved sulphide

52

has also been observed (Backman et al. 2002) in most of the overburden sampling points, particularly from those with pH over 7.

Ferrous iron (Fe2+) (Fig. 5-16c) has two concentration peaks with depth and it seems that the solubility of iron is limited in the 250 - 450m depth range (note that Fe2+ was not analysed in PR samples before 2001 and in PVP samples before 2003). An apparent sink for aqueous iron may be the precipitation of iron sulphides, e.g. pyrite, as a final precipitate that is poorly soluble in near-neutral conditions. Pyrite has frequently been observed in Olkiluoto drill core samples (e.g. Blomqvist et al. 1992, Gehör et al. 1996a, 1997). Thin pyrite coatings observed on the surfaces of the most recent fracture calcite layer indicate late precipitation of pyrite. Ferrous iron and dissolved sulphide show an inverse distribution in abundance with depth (Fig 5-16). Pitkänen et al. (1999a) suggested that this is caused by changes in dominating constituent buffering redox system in different groundwater layers. Ferrous iron is enriched in groundwaters with significant marine component (brackish HCO3 and SO4-rich types) and in deep, S-poor saline groundwater. High SO4 contents in them are able to buffer the redox system and minimise the content of organic carbon that is necessary for bacteria to reduce SO4 (see Ch. 5.5.5). Sulphide is enriched in the mixing zone of SO4-rich and -poor brackish groundwaters, where bacteria may use CH4 (Ch. 5.5.6) and iron is depleted due to ongoing sulphide production. A few samples at Olkiluoto show anomalously high sulphide contents which are normally restricted by dissolved iron or solid ferric iron phases. The iron-reducing bacteria (Haveman et al. 1998, 2000) discussed in section 5.5.4, for example, can reduce iron oxyhydroxides and release ferrous iron. However, no observations of iron oxyhydroxides have been made from fractures in deep boreholes (Gehör et al. 1996a, 1997). The unusual sulphide enrichment in groundwater may be due to poorly-reactive silicate iron phases (biotite etc.) (Raiswell & Canfield 1996) that control availability of iron for precipitation of insoluble iron sulphides. The observations of dissolved sulphide in shallow groundwater samples indicate a shift from aerobic to prevailing sulphidic conditions mostly in the overburden layer. The samples have high tritium contents and shallow groundwater samples show even slight seasonal variation (Hatanpää 2002, Backman et al. 2002) suggesting a very short lifetime for oxygen during and after recharge. Sulphidic conditions probably dominate down to 450 m although sulphide concentrations in the middle of SO4-rich groundwaters are small due to immediate precipitation with dissolved ferrous iron. Deeper, in the CH4-dominated system, conditions do not favour the occurrence of trace amounts of SO4 and hence also sulphide, and so ferrous iron concentration is not limited by its low solubility with dissolved sulphide.

53

Figure 5-16. a) Fetot in shallow groundwater samples as a function of pH, b) depth distribution of S2- and c) depth distribution of Fe2+

.

Ammonium

Ammonium (NH4+) is in general an important redox species in shallow groundwater

systems particularly in agricultural lands. It also has significance in nuclear safety, because ammonia in higher concentrations might increase the risk for stress corrosion of copper. According to Arilahti et al. (2000) there were no observations of stress corrosive effects in synthetic Olkiluoto groundwater simulations ([NH4+] = 1 - 100 mg/l). The detected NH4 concentrations are low at Olkiluoto(Fig. 5-17). Ammonium shows the highest values at depths of 100-200m in groundwater samples with high SO4 content, i.e. having a strong marine signature. Therefore, the main source of ammonium

54

seems to be related to seawater input in the groundwater system. Seawater has a small NH4 content, but apparently infiltration through marine sediments containing organic nitrogen compounds (amines in plant debris, animal tissue and faecas) and their catalytic hydrolysis (bacterial enzymes) result in NH4 production (Stumm & Morgan 1981, Canter 1997). Bacterial nitrification of ammonium seems to require oxygen, and therefore ammonium is likely to be quite stable in deep groundwater. Below the SO4-rich layer, ammonium contents are occasionally above detection limit but low (<0.3 mg/l). Cation exhange may cause significant depletion of NH4

+ in lower part of SO4-rich groundwater. Cation exchange may fix NH4

+ irreversibly, because the ions tend to induce collapse in highly charged clay minerals when it enters the interlayer space (Appelo & Postma 1993). Such a collapsed structure can only exchange through solid state diffusion, which is 3 to 4 orders of magnitude slower than diffusion in solution. The origin of deep NH4

+ may be different than in marine-derived groundwater types, because the marine component is negligible, especially in deep saline groundwater (Pitkänen et al. 1999a). The slightly higher values in deep saline groundwaters may be related to high hydrocarbon contents (e.g. thermal degradation of organic N-compounds, see Ch. 5.5.5) and to slow ion-exchange in groundwater conditions with high ionic strength.

Figure 5-17. Ammonium (NH4) as a function of Cl-.

55

Uranium (238U) and daughters Uranium has two oxidation states U6+ and U4+ in natural environments. Oxidation state U6+ is much more soluble (even mg/l level) than U4+, which is considered to be in-soluble (ppb level). Therefore, U released in the near-surface oxidising groundwater zone shows mainly increased concentrations in shallow groundwater samples. Deep groundwaters are generally from a reducing environment where U is poorly soluble and typically precipitates. Uranium release is very complicated in such conditions. It depends on such factors as the distribution of U within the rock matrix and on fracture surfaces, the type of U-phases, the chemical state of U in the solid phase and the geochemical properties of the groundwater, such as the availability of reducing and complexation agents The results from Olkiluoto groundwaters (Fig. 5-18a) correspond to general theory. The highest concentrations are observed in near-surface groundwaters. Distinctly high Utot values (>5ppb) are observed in samples from KR5/T7 and PR2. This suggests the dominance of U dissolution in the dilute or brackish HCO3-type samples for a bedrock consisting predominantly of granites and pegmatites (c.f. Fig 2-3). These rocks have clearly more U compared to the other rock types of the site (cf. Ch. 2). The rock type control over the U concentrations in groundwater could probably be used as fingerprint for determining the origin of the U at least in rechargezone. The concentration of Utot decreases with increasing depth in brackish and saline groundwater types (Fig. 5-18a) and thus with longer residence time, as with increasing Cl concentration (Fig. 5-18c). The decreasing trend can be considered as indirect indicators of prevailing reducing redox conditions (e.g. Pearson et al. 1991) in deep groundwaters at Olkiluoto. The greatest variation in Utot values was found in the brackish SO4-type groundwater. This may be due to spatial or lithological variations in the bedrock or mixing between different groundwater types. Evidently the extremely low contents of U in deep groundwaters result from the reducing nature of the host rock, which contains graphite and sulphide minerals, and high methane and hydrogen concentrations (possible reducing agents) (Pitkänen et al. 1999a, 2001). In addition, the low DIC contents in saline groundwaters do not favour U complexation and its maintenance in solution. These factors have probably maintained very reducing conditions throughout the geological history. Additional information on U behaviour is available from the 238U decay series isotope activities. The 234U/238U activity ratio (AR) is frequently used for groundwater because of its link to redox evolution and because the long half-lifes of these nuclides allow dating of groundwaters tens to hundreds of thousands of years old (e.g. Clark & Fritz 1997). The rock mass of Olkiluoto is very old and therefore the decay of 238U and its daughters have achieved equilibrum whereby all activity ratios (AR) are equal to 1.0. As U dissolves in groundwater the AR of U isotopes is generally higher than 1.0 due to the higher solubility of 234U, which is situated in crystal lattices damaged by decay of 238U and auto-oxidation of 234U to U6+during decay (Clark & Fritz 1997). Oxidized groundwater in the recharge zone is aggressive taking most of U into solution. Fractionation between U isotopes is low and the AR value increases only slightly above

56

1.0. In an anoxic environment, 238U is insoluble and the fractionation processes of 238U decay series are able to increase relatively the amount of 234U in solution and so increase the AR value clearly higher. The AR values with U concentration in groundwaters are shown in Figure 5-18b. The samples with high U content (> 5ppb), three of them from PR2, show a low AR indicating oxidising groundwater conditions whereas the sample from KR5/T7 has a high AR. This may indicate retarded reduction of U6+ in a U �rich environment due to lack of electron donors and therefore prolonged residence in groundwater. The other HCO3-rich samples from deep boreholes also show elevated AR values indicating anoxic groundwater conditions. One of the shallow samples with U concentration about 4 ppb (PP7) has also increased AR value (Figs 5-18a & b), which may indicate discharging character of the sampling point. The brackish SO4-rich groundwater shows slightly higher Utot and AR values to the other deep groundwaters. This may reflect infiltration of ancient Litorina Seawater through marine sediments. According to the observation by Ku (1965) a significant proportion of the excess of 234U in seawater may be derived by diffusion from sediments. The slightly elevated U contents in SO4 �rich groundwater may result from a lack of electron donors buffered by the high SO4 concentration (c.f. ch. 5.5.3). The decreasing AR trend deep at Olkiluoto may reflect a long residence time during which the ratio tends to approach equilibrium. That corresponds to million year scale. Radon-222, also a daughter of 238U series, is inert in the hydrogeochemical system and behave similarly to U in terms of the mean values for deep and shallow groundwaters. The activity of Rn, can be a useful qualitative indication of the short-term history of groundwater in the vicinity of the sampling point. Radon is highly soluble in water and is not adsorbed by any solid, thus its activity is generally greater than that of U by a factor of 103 to 105. High Rn contents are typical of waters in porous aquifers with high contents of finely dispersed uranium in contact with pore water, and also for fast transfer of such pore water to groundwater at the sampling point (Pearson et al. 1991). In the bedrock of Olkiluoto, activities of Rn in the groundwaters show a decreasing trend with increasing Cl concentration (Fig. 5-18d). Since 222Rn is a daughter product of the radioactive disintegration of U, Figure 5-18d suggests that uranium located deep in the local bedrock is tightly bound in unaltered minerals, and that the matrix permeability is lower at depth, as indicated by the hydrological investigations (Ahokas et al. 1996). The local lithogeochemical differences in the amounts of U in bedrock minerals cause scatter in the 222Rn data. The sampling equipment may also affect the Rn activities. As Rn is a gas, sampling depths with a lower hydraulic conductivity may result in lower activities of 222Rn, due to out-gassing during the pumping.

57

Figure 5-18 a) depth distribution of Utot, b) Utot vs. 234U/238U, c) Utot vs. Cl, d) concentration of 222Rn vs. Cl in the Olkiluoto groundwaters.

DOC (dissolved organic carbon)

Dissolved organic carbon is suggested to occur at rather low concentrations (< 2 mg/l) in deep granitic groundwater by Pettersson et al. (1990). The DOC concentrations at Olkiluoto were, however, much higher and they were interpreted unreliable due to organic contamination of the sampling equipment (Pitkänen et al. 1999a). New samples show similar level of DOC, although the results from SO4�rich waters occur slightly lower level than the ones reported by Pitkänen et al. (1999a). The data from deep saline groundwaters have high concentrations. The results still seem to contain contamination from the sampling equipments and therefore they are not used in further interpretation. Research continues to minimise DOC contamination during sampling and to evaluate real DOC concetrations.

58

5.5.2 Carbon isotopes and carbonate geochemistry

14C, residence times and uncertainties

The carbon isotopic composition of DIC is a direct reflection of the potential geochemical reactions affecting the history of groundwater (e.g. Fritz et al. 1989; Plummer et al. 1990, Pearson et al. 1991, Clark & Fritz 1997). Unfortunately, carbon isotopic measurements at Olkiluoto are hampered in saline groundwaters by low DIC content and low partial pressure of CO2 (Pitkänen et al. 1999a). Groundwaters are susceptible to CO2 contamination during sampling because the partial pressure is generally lower than in the atmosphere. Deep groundwaters also have an enormous gas content, which may easily be lost from solution due to decreasing pressure near the surface conditions. The minor CO2 may evacuate with other gases, raising the pH of the water sample. Solubility calculations indicated fairly reliable pH measurements (Pitkänen et al. 1999a and details in Ch. 5.6), thus CO2 degassing may not play a significant role in the current groundwater data. Radiocarbon results indicate instead that CO2 uptake may be evident in part of the data, but it may also counteract degassing problems in saline groundwaters. Figure 5-19 shows variation between DIC, tritium and 14C(DIC). Samples from overburden and few dilute groundwaters from shallow wells have fairly low DIC content and modern 14C(DIC) value, which is slightly enriched by nuclear bomb-pulse. In other HCO3-rich groundwaters DIC seems to increase with decreasing 14C content down to 60 pM level. All these samples have a high tritium content indicating that a dead carbon source dilutes the 14C in DIC and the samples represent young groundwater (< 50 years). Dissolution of calcite is probably the most significant dead carbon source in these samples, which have mostly been taken from shallow wells. Calcite also approaches saturation in these samples (details in Ch. 5.6) The DIC content begins to decrease steeply as radiocarbon decreases below 60 pM. Changes in radiocarbon content may dominantly result from mixing of HCO3-rich and SO4-rich brackish groundwaters as shown by calculated mixing lines between reference values of the both groups. However, radioactive decay may have decreased 14C content in some of the HCO3-rich groundwaters, which are plotted left of the mixing lines. These samples have generally low tritium contents (1.1 - 3.4 TU) and represent the most saline lower part of HCO3-rich groundwater layer. Dissolution of fracture calcites is not anymore probable to decrease 14C beyond the mixing band formed by the lines. The 14C depletion in these groundwaters suggests notably longer mean residence time than for the young, shallow part of HCO3-rich groundwaters. The 14C difference between the samples and the mixing band (range 5 to 15 pM units) corresponds roughly from 500 to 3000 years as calculated from 14C decay (half life 5730 years), which is appropriate to the time span of meteoric water infiltration started at Olkiluoto

59

Figure 5.-19. a) Dissolved inorganic carbon and b) tritium versus radiocarbon. The lines represent mixing dependence of 14C(DIC) and DIC between HCO3-rich and SO4-rich waters. The lines are calculated for SO4 and HCO3 groundwater end-member compositions: (20 pM, 8-18 mg/l) and (55-65 pM, 80 mg/l), respectively. The 14C value of samples having the clearest Litorina Sea signature varies between 20 and 30 pM. The difference between these samples and previous HCO3-rich samples (their range 35 to 45 pM) corresponds as calculated 14C decay age to a range of 1300-7000 years. This time span suggests total residence time from 4000 to 7500 years, which corresponds with the estimate in Pitkänen et al. (1999a) and belongs to the Litorina period. These rough estimates do not consider possible mixing effects of older

60

groundwater types or dead carbonate input other than influence during recharge. However, the age estimates are consistent with the concept of palaehydrogeology of the site.

Brackish Cl and saline groundwaters below the SO4-rich layer show more depleted 14C values, some are even below detection limit (3 pM), suggesting still longer residence times. This is also indicated by the 36Cl results for the most saline groundwater samples at Olkiluoto (Gascoyne 2002). However, part of the brackish and saline samples, particularly those having very low DIC contents, plot below the high values of HCO3-rich groundwaters. This must result from contamination by young DIC. Mixing with HCO3-rich groundwater does not seem substantial because DIC contents in these samples are low as are most their tritium contents. Therefore, CO2 ingassing seems to be the most likely factor causing increased radiocarbon values. Diffusion may have occurred through polyamide tubes during sampling. The CO2 partial pressure (see Ch. 5.6) difference between the inner and outer surface of the sampling tube is highest in the very upper part of groundwater column, where diffusive ingassing is possible, but also above the groundwater table in which case CO2 can be either atmospheric or biogenic in origin.

Samples suffering slightly or notably from modern CO2 contamination are probably all brackish Cl-type and saline groundwaters showing higher 14C values than 20 pM. Unfortunately, this also causes reliability problems on δ13C results of corresponding samples and hampers the interpretation of carbon cycling in the deepest part of the site. Carbon-13 has hardly enough time to equilibrate between the sample and environment during sampling. The contamination process can be considered to be a closed system and ingassing of CO2 will generate a similar δ13C signature as in source CO2, i.e. about -23� in biogenic CO2 and �7� in atmospheric CO2 according to Clark & Fritz (1997).

Carbonate evolution in groundwater

Carbon isotopic evolutionary trends are shown in Figure 5-20. Evolution begins in the recharge environment and continues in the subsurface where calcite-water interaction seems to dominate. The samples with modern radiocarbon activity (> 100 pM) form a coherent group showing δ13C value around �25� PDB. This corresponds well with values presented for soil-derived CO2 (e.g. Pearson et al. 1991, Clark & Fritz 1997) produced by plant root respiration and the decay of plant debris in high latitude regions. Infiltrating waters dissolve soil CO2 and the subsequent dissolution of calcite below groundwater table mainly cause, increase in δ13C. Simultaneous increase in pH and DIC (Figs 5-4, 5-7), and decrease in PCO2 (Fig. 5-25b) and 14C are essential reasons to consider that calcite dissolution instead of silicate weathering is the main process to neutralise groundwater in bedrock and that the dissolution occurs mainly under a closed system for soil CO2. If silicate dissolution increased pH, DIC would not neccesarily increase and 14C activity remains stable. An open system would retain a high PCO2 and modern 14C activity, instead of the observed depletion in groundwaters.

61

Figure 5-20. Carbon isotopic evolutionary (DIC) trends (Pitkänen et al. 1999a). Depleted 14C values with increased δ13C values in samples from shallow boreholes show that calcite dissolution starts at shallow depth in bedrock. Carbon isotopic data from observation tubes at overburden are too few to determine any possible indications of calcite dissolution. The recharge zone in the overburden may be an open system for CO2, because PCO2 remains stable. The calcite dissolution trend from carbon isotopes seems to continue as far as 14C decreases to a 60 pM level in HCO3-rich groundwaters. Most of the samples forming this path are from shallow bedrock. The trend corresponds in calcites with δ13C values around �5�, which corresponds to the fracture calcites data in the upper part of the bedrock (cf. Blomqvist et al. 1992, Blyth et al. 1998, 2000, Karhu 2000, Gehör et al. 2002). The rest of HCO3-rich samples below the 60 pM level show a fairly stable δ13C level indicating that calcite dissolution is limited to shallow depths (< 30m) in bedrock. As mentioned above, the further depletion of 14C may result from radioactive decay, but mixing of deeper groundwater types could have a significant role. Part of the SO4-rich samples shows higher δ13C value than HCO3 -rich groundwaters. Calcite dissolution to elevated values is not considered possible in these depths. The probable reason resides in the marine origin of the SO4-rich groundwater type (Pitkänen et al. 1999a). The initial value of δ13C in DIC is different in recharging seawater (about �1 �) than in meteoric derived HCO3 groundwaters. Baltic seawater is already quite near calcite saturation and Litorina seawater, which was more saline, may have been even closer, therefore calcite dissolution may have not been an important process during seawater infiltration. Recharge through sea-bottom sediments represents a totally different path for carbonate evolution than meteoric recharge. Hence a chemical steady-state evolution is interrupted between HCO3- and SO4-rich groundwaters. Smooth,

62

continuous lines trends observable in many chemical parameters are caused by later mixing in bedrock between the groundwater types from different sources.

Chemical conditions in sea-bottom sediments are reducing. Sediments are rich in organic matter that is used by anaerobic bacteria to reduce marine SO4; in particular:

SO42- + 2 CH2O → H2S + 2 HCO3

-. (Eq. 5-1)

Sulphate reduction is a typical process in marine sediments (e.g. Appelo & Postma 1993, Raiswell & Canfield 1996, Whiticar 1999) and if a source of iron is available iron sulphides will be precipitated, giving pyrite as a final phase:

8 SO42- + 4 FeOOH + 15 CH2O → 4 FeS2 + 15 HCO3

- + 9 H2O + OH-. (Eq. 5-2)

Iron sulphides are frequent accessory minerals in marine sediments and are commonly observed in Litorina clays, particularly in western Finland along the coastal area of the Gulf of Bothnia (e.g. Palko 1994). Therefore, significant SO4 reduction may have taken place at Olkiluoto during Litorina seawater infiltration. The SO4-rich samples with high δ13C may represent the final point of evolution, which has mainly been caused by the anaerobic oxidation of organic matter in the sea-bottom sediments. Microbial reduction of SO4 has probably been minor in the bedrock due to the low concentration of organic subtrates. Pitkänen et al (1999) estimated according to this assumption that SO4 may have reduced about 1.2 mmol/l during Litorina seawater infiltration through the sea bed. The corresponding HCO3 input in groundwater is 2.4 mmol/l (about 30 mg/l DIC), which is sufficient to cause calcite saturation. The level of δ13C may have decreased further by anaerobic consumption of CH4 in bacterial SO4 reduction:

CH4 + SO42- → HCO3

- + HS- + H2O. (Eq. 5-3)

This process is particularly active in the border zone of the SO4-bearing groundwater system and only trace amounts could migrate to the inner part of the SO4-rich system (e.g. Niewöhner et al. 1998, Whiticar 1999). Mineralisation of methane, which has a generally low δ13C level (Ch. 5.5.5) may cause very negative δ13C values in DIC. Such a low value has been observed only once in the isotopic data of Olkiluoto (Pitkänen et al. 1999a), in a sample from KR2/T3 representing the lower part of the SO4-rich zone. Low values (-30.2�) has also been recently measured in CO2 gas evacuated from sulphide-rich samples of KR13/362. This supports a CH4 origin of carbonate occurring at Olkiluoto. Bacterial (methanogens) formation of methane fractionates δ13C strongly between CH4 and DIC and isotopic difference (δ13CCO2-δ13CCH4) is about 75�, Clark & Fritz (1997). The process may have two pathways, either by acetate fermentation producing methane and CO2 or carbonate reduction (see Ch. 5.5.5) and both processes act substantially in a SO4-poor groundwater system (e.g. Plummer et al. 1990, Clark & Fritz 1997, Whiticar 1999). Methanogens favour lighter isotopes, thus increasing δ13C in released or residual CO2 depending on methanogenetic pathway.

63

An elevated δ13C signature can be observed in a few groundwater samples representing brackish or saline groundwater below the SO4-rich groundwater layer at Olkiluoto. The high positive value of the saline sample is especially typical to the carbonate reduction pathway (e.g. Clark & Fritz 1997, Whiticar 1999). The two other high values in brackish groundwaters have fairly high 14C contents suggesting contamination by young carbonate, which has been able to decrease original δ13C values.

5.5.3 Sulphur and oxygen isotopes of aqueous SO4 (δ34S(SO4) and δ18O(SO4))

The stable isotopes (34S and 18O) in dissolved SO4 can be used in evaluating the origin of the sulphur, redox processes and, indirectly, residence times of groundwater (Fontes 1994). Sulphate in groundwaters at Olkiluoto may have two significant sources: mineral sulphides and current or ancient seawaters, and several smaller sources such as emissions from fossil fuels, sea spray, and biological sulphur, which may be significant in recharge waters. Sulphide oxidation, as a significant SO4 source, is possible at the surface. The 34S composition of SO4 from oxidation of mineral sulphides is only marginally depleted from the value of original sulphides (Clark & Fritz 1997), which is normally around 0� CDT (Canyon Diablo Troilite standard) or slightly positive in crystalline rocks. The local Baltic seawater value for δ34S, about +20�, is quite similar to the ocean-water value (e.g. +21� in Clark & Fritz 1997), which has been stable during the Quaternary. The δ34S values in SO4 released from fossil fuel or the biosphere, are similar to rock sulphides (possibly slightly higher to about +10�) and sea spray corresponds to the seawater value. The δ18O(SO4) value of atmospheric and seawater sulphates is +10� SMOW (Standard Mean Ocean Water) or more whereas the δ18O value for SO4 oxidised from mineral sulphides may be near 0� at Olkiluoto. Fontes et al. (1989) presented the last value for SO4 in groundwater formed by the oxidation of sulphides at the Stripa site in Sweden and was also expected to be typical for central Canada (Clark & Fritz 1997). The climates in these areas are comparable with Finland. Therefore, δ18O in source oxygen (atmosphere and water) forming SO4 is similar in these areas and δ18O value in SO4 obtained through oxidation of sulphides is also similar. The reduction of SO4 at low temperature is possible in the presence of SO4-reducing bacteria. If accessible organic carbon is available, bacteria will use sulphate as an electron acceptor and produce sulphide and mineralised carbon (as shown in Eq. 5-1 and 5-3.). Fractionation may be 40-50� between SO4 and HS- (Plummer et al. 1990). Sulphide is depleted in 34S, which gradually accumulates in the residual SO4. An enrichment in 18O of the residual sulphate also accompanies reduction of SO4 (Clark & Fritz 1997). The enrichment factor between SO4 and water is about 30� as the water-SO4 system approaches equilibrium at normal temperatures. The δ34S value of the fresh groundwater in the overburden and shallow bedrock at Olkiluoto (about +3.5� CDT) diverges from the other data (Fig 5-21), indicating oxidation of rock sulphides or meteoric fallout as a source for SO4. The δ18O(SO4) value in the single overburden sample suggests solely sulphide oxidation as a source process at the surface. The other δ34S data, as well as δ18O(SO4) data, support seawater (current or Litorina) as the major source for SO4 with some mixing between the end-members (mixing lines are presented in Fig. 5-21a). However, one HCO3-rich sample with low

64

δ18O value, marine level δ34S and low SO4 content (PR1/1) represents a SO4 reduction path more likely starting from overburden composition. The other HCO3-rich sample with marine δ34S and low SO4 content (PR2/1) represents instead a pure marine SO4 origin on the basis of its δ18O (+12�) in SO4. This shallow well is located in microcline granite, which is poor in sulphide minerals (note contrary to U- results). The results of shallow groundwaters show that SO4 has both marine- and sulphide-based initial sources during recharge. The sample, which shows δ34S value at about +10� (PP8) may represent reduction of once-oxidised sulphide at the surface, but also mixing between surface and marine SO4. Oxygen isotopic data could resolve this question, but it is unfortunately missing. Observations of hydrogen sulphide in shallow groundwaters (Hatanpää 2002, Backman et al. 2002) support the isotopic results and they both indicate strongly reducing conditions at very shallow depths in the bedrock at Olkiluoto. The input of marine SO4 in shallow HCO3-rich groundwaters may effectively account for the enrichment of 34S in SO4 due to microbial reduction of SO4. Therefore, more 18O data and also 34S data from dissolved and solid sulphide phases are desirable. The brackish SO4-rich groundwater with Litorina Sea input has clearly higher δ34S values than the present Baltic Seawater indicating that SO4 reduction has occurred. However, the fairly stable level of δ34S, even at lower SO4 concentrations, suggests, that the process is limited and the elevated value may originate from some earlier episode. Pitkänen et al. (1999a) have interpreted that the main reduction episode occurred during Litorina seawater infiltration through organic-rich, sea-bottom sediments. Sulphate reduction in SO4-rich groundwater may be at present limited due to the lack of nutrients (electron donors) for SO4 reducing bacteria. The deep SO4-poor groundwater types usually contain such small amounts of dissolved SO4 that the isotopic analyses were seldom successful. The observed elevated values suggest bacterial reduction of dissolved SO4.

65

Figure 5-21. a) δ34S(SO4) values vs. SO4, b) δ18O (SO4) vs. δ34S(SO4) values in Olkiluoto groundwater samples. The curves represent calculated mixing lines between sulphide derived from SO4 and SO4 in seawater and altered Litorina end-members.

66

5.5.4 Microbes

Microbial populations in Finnish deep bedrock groundwaters have been studied since the late 1990�s. Two sampling campaigns have been carried out at Finnish candidate sites for a nuclear waste repository (Haveman et al. 1998, 1999, 2000). Altogether, 7 samples have been collected from Olkiluoto (Table 5-7) representing brackish and saline groundwater types. As well as the limited number of samples, microbial studies have their own limitations. For example, the methods do not give information about the full diversity of populations and normally only 1 � 2% of total cells can be cultured to determine the most probable number (MPN) of physiological groups; therefore, the results should be considered as indicative only. The samples were studied by Haveman et al. (1998, 2000) for total cell numbers and 6 physiological groups, including iron (IRB) and sulphur (SRB) reducing bacteria, heterotrophic and autotrophic acetogens and methanogens. Heterotrophic acetogens and methanogens (HA, HM) use organic carbon sources to produce acetate and methane, whereas their autotrophic responses (AA, AM) use inorganic carbon sources such as carbonate together with hydrogen. SRB and IRB metabolise simple organic compounds (e.g. acetate) in reducing sulphur and iron. Competition of organic substrates is severe in deep groundwater conditions where the contents of organic carbon are generally small, which may limit the activity of microbes, at least the heterotrophs. Direct cell countings showed that the total number of microbial cells found in subsurface microbial studies was in all samples (105 � 106 cells/ml, which is typical in deep groundwaters in crystalline rock (Haveman et al. 1999). Table 5-7 Total number of cells (cells/ml) and the most probable number of metabolic groups of microorganisms in groundwaters from Olkiluoto sampled for microbial studies (Haveman et. al 1998, 2000). Water type, SO4 (mg/l) and CH4 (ml/l) contents of groundwater samples are shown.

Sample KR8/302/1 KR3/243/1 KR10/324/1 KR3/438/1 KR9/470/1 KR9/563/1 KR4/860/1

Water type Brack. SO4 Brack. Cl Brack. Cl Brack. Cl Saline

Na-Ca-Cl

Saline

Na-Ca-Cl

Saline

Ca-Na-Cl

SO4/CH4 470/0.06 1.4/28.6 8.3/81.5 11.7/51.9 17.1/234 1.3/258 1.2/768

Total±SD*105 2.8±1.0 5.1±1.4 6.5±1.7 7.0±1.2 1.5±0.5 6.2±1.5 1.7±0.5

Autotr. acetogenes - 7.8 2.2*101 - - - -

Heterotr. acetogenes - 3.3*102 9.2*103 9.3*102 1.1*102 - 4.9

Autotr. methanogens - - 4.5*10-1 - - - 7.0

Heterotr. methanogens - - 4.5*10-1 - - - 2.3

SRB 1.6*104 >1.6*104 >1.6*104 4.2*102 9.2*101 1.7 -

IRB not tested 1.5*103 7.0 4.6*102 3.3*101 not tested -

% of tot. cells cult. 5.7 3.5 3.9 0.26 0.16 0.00027 0.0084

- = below detection limit (< 0.2 cells/ml)

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The limited study of the microbiological content of Olkiluoto groundwaters has shown that sulphate reducing bacteria (SRB) are the most abundant species and tend to be particularly associated with groundwaters at an intermediate depth range (~ 250-330 m). The SRB levels and SO4 concentrations appear to be unrelated, however. The deeper, saline groundwaters contain very low amounts of SRB and iron reducing bacteria (IRB). The population of SRB and IRB seems to be strong in the transition zone between SO4 -rich and -poor groundwaters, where redox conditions change from sulphidic to methanic and both methane and dissolved sulphide contents increase. The transition zone may favour microbial activity. Several studies indicate that anaerobic bacterial consumption of CH4 is active at a base of the SO4 reduction zone (e.g. Niewöhner et al. 1998, Whiticar 1999). Above this zone in SO4-rich layer, reduction of SO4 may be curtailed due to deficiency of organic nutrients developed in the system that is buffered by high concentrations of SO4. Deeper in the saline groundwater SO4 may even be lacking. The presented trace values (Table 5-7) may be due to contamination. For example in the first sample taken from KR4/861, SO4 was under the detection limit. Lack of SO4 may therefore decrease the activity of SRB. Haveman et al. (2000) suggested, however, that in deep undisturbed groundwater conditions (at least from the beginning of Weichselian glaciation) SRB have not been able to survive such a long period without the input of new nutrients. Interpretation of MPN results in saline groundwater samples is significantly hampered by the low percentage of total cells which were cultured (Table 5-7) and the appearance of acetogens and methanogens is indistinct (Haveman et al. 2000). Methanogens and acetogens seem to increase in SO4-poor, brackish and saline waters, particularly in highly saline groundwater (KR4/860) where large concentrations of CH4 and H2 are present (cf. next section). This sample most likely represents the hydrogen-driven biosphere proposed by Pedersen (2000). Autotrophic microbes use hydrogen as an energy source in reducing carbonate to produce organic carbon, which in turn may be consumed by heterotrophs. The population of autotrophic methanogens is at least active in Olkiluoto. The better understanding of deep biosphere needs further studies at the site as well as in formation on the distribution of the autotrophic ecosystem.

5.5.5 Dissolved gases

Results of the earlier hydrochemical research programmes at Olkiluoto have already indicated that the saline groundwater has massive amounts of dissolved gases. Unfortunately, the results were quantitatively poor before sampling started with PAVE equipment in 1997. After the last reports (e.g. Anttila et al. 1999, Pitkänen et al. 1999a) the number of gas samples taken with this new system has increased significantly. Altogether 24 samples, between 65 and 1000 m depth, are currently available and 14 different gas species have been measured/observed. Some gas phases such as CO and a few hydrocarbon phases were almost systematically below detection limit. The samples have been taken either using argon or nitrogen as the back pressure gas and flush gas in the pressure vessel. The results shown later are based only on Ar-filled pressure vessels (the N2 results are considered to be feasible, as well). Therefore, Ar is not discussed here.

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Total dissolved gases (measured at NTP) show fairly a coherent increasing trend with depth indicating relatively good reliability of current gas sampling system (Fig. 5-22a). Large variations are also observable in single samples, for example in the results of the deep sample from KR4/860 m (900 � 1900 ml/l) reflecting, however, uncertainty in quantitative results. This is caused by variation in the contents of the main gas phases such as N2, CH4 and H2 in parallel samples. Gascoyne (2002) has discussed the uncertainty question in detail and made suggestions of further development on gas sampling and measurement technologies. He considered that the main reason for uncertainty is due to the variable amount of water recovered during sampling. The amount of dissolved gases deep in the bedrock of Olkiluoto are high compared with the results found at the other sites in Finland, for example, at Hästholmen, with similar elemental compositions, sampled with same equipment and corresponding depths (Pitkänen et al. 2001). Gas contents are about the same in SO4-rich brackish groundwaters at Olkiluoto and Hästholmen but at greater depths, the contents at Olkiluoto are many times higher than at Hästholmen. The higher contents result primarily from large amounts of reactive gases, i.e. hydrocarbons and hydrogen. Of the other atmospheric gases, O2 shows a large scatter with depth (Fig 5-22a) indicating that its appearance is due to contamination from the atmosphere during sampling or sample treatment. Helium and N2 have, instead, relatively coherent increasing trends with depth suggesting that they are relatively reliable data. Helium is considered to originate in the bedrock either by radioactive decay or crustal degassing and N2 by crustal degassing. Significant amounts of hydrocarbons (HC) and hydrogen are, in principle, unstable in the presence of O2 (e.g. Whiticar 1999) particularly if suitable microbes exist (methanotrophes). This supports atmospheric contamination as an origin for O2. Higher HC (C atom number > 1 in molecules, C2H6, C2H4, C3H8 etc.) than CH4 are summed together in Figure 5-22b and they mainly consist of ethane and propane. Hydrocarbons and hydrogen show generally increasing trends with depth, whereas CO2 decreases (Fig 5-22b). However, the trends of both methane and higher HC are bimodal in the upper 400 m. The decreasing trend belongs to brackish SO4-rich groundwater samples in which the signature of Litorina Seawater is strongest. The increasing trends in the upper part of the bedrock are of brackish SO4-rich groundwater samples with smaller marine signature and SO4 content, representing either groundwater mixtures with HCO3-rich waters or the transition zone with SO4-depleted brackish groundwater. These trends correspond well with the concept that hydrocarbons and SO4 are unstable in a common system due to bacterial reduction of SO4 with anaerobic reduction of organics (e.g Niewöhner et al. 1998, Whiticar 1999) and that the simultaneous presence of high contents of CH4 and SO4 indicates a mixing of different waters (Plummer et al. 1994). The decreasing trend of CO2 with increasing depth is commensurate with the decreasing content of DIC in groundwaters. The origin of CO2 and H2 is clear. Carbon dioxide is a dissociation product of dissolved carbonate and H2 is probably a crustal degassing product, but may also have been derived during methane polymerisation. The origin of hydrocarbons is much more complicated and it is worth noting that CH4

69

solubility are near saturation (Gascoyne 2000) in the sample (KR4/861) with the highest measured CH4 content at Olkiluoto.

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Figure 5-22. The content of dissolved gases (ml gas/l water at NTP) with depth in Olkiluoto groundwater samples (PAVE samples), a) total dissolved gases, N2, He and O2, and b) CO2, H2, CH4 and higher hydrocarbons (CnHn). Indicative trend lines are represented by the coloured paths.

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Origin of hydrocarbons

Natural hydrocarbons are largely formed by the thermal decomposition of organic matter (thermogenesis) or by microbial processes (bacteriogenesis) in low temperature. Thermogenesis occurs in deeply buried organic sediments in sedimentary basins. Bacterial methanogenesis may proceed either with fermentation of methylated substrates such as acetate (Eq. 5-4) or through a carbonate reduction pathway using hydrogen gas to reduce CO2 (Eq. 5-5). Formation of higher HC is limited to trace levels.

CH3COOH = CH4 + CO2 (Eq. 5-4)

CO2 + 4H2 = CH4 + 2H2O (Eq. 5-5) During the last decades, an abiogenic source for HC has also been discovered (e.g, Sherwood Lollar 1993, 2002). The abiogenic HC are formed, for example, in hydrothermal systems during water-rock interactions involving the Fischer-Tropsch synthesis reaction (Eq. 5-5) or graphite and hydrogen (Eq. 5-6). Higher hydrocarbons are formed by polymerisation of methane precursors.

C + 2H2 = CH4 (Eq. 5-6) The fermentation pathway is normally restricted in anaerobic conditions to shallow depths, features rich in organic substrates such as in bogs and lake bottom sediments. In sediment pore waters with abundant SO4, SRB outcompete methanogens and methnogenesis is severely curtailed (Whiticar 1999). Therefore, CH4 produced by fermenting bacteria in Olkiluoto should be limited to shallow depths, in goundwater infiltrated in bedrock through ground depressions or through sea-bottom sediments depleted in SO4 due to SO4 reduction. Methane belonging to this group may have been observed from HCO3-rich brackish groundwater in sample KR12/65, which also shows a typical δ13C(CH4) value (-61.5�) for the fermentation δ13C(CH4) pathway (Whiticar 1999). The contents of this type of CH4 in groundwater are small, however, at Olkiluoto (< 1ml/l in KR12/65). Carbonate reduction appears instead to be a more realistic bacterial formation process for CH4 in deep groundwater at Olkiluoto. Evident lack of organic substrates, such as acetate, that is probably consumed by SRB and IRB in SO4-rich groundwater layer, and accumulation of a hydrogen pool in SO4-free deep groundwater zone, support a carbonate reduction path way. This is also indicated by decreasing CO2 (Fig. 5-22b) and low DIC contents (Fig. 5-9c) at corresponding depths, as CH4 and higher HC increase. Few high δ13C(DIC) values below the SO4-rich groundwater zone indicate methanogenetic conditions. Microbial analysis by Haveman et al. (1998, 2000) showing autotrophic methanogens indicate on-going methanogenesis based on carbonate reduction below the SO4-rich zone. On the other hand, Gehör et al. (2002) showed fracture calcite data from the upper 100 � 300m with high δ13C signatures (up to +12� PDB) suggesting that methanogenic conditions have also prevailed in the SO4-rich zone at earlier times.

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Formation of higher HC in bacterial methanogenesis is much more restricted than in thermal processes (Sherwood Lollar et al. 1993, 2002, Clark and Fritz 1997, Whiticar 1999). The molecular ratio C1/(C2+C3) is generally very high on bacterial methanogenesis (more than 103) whereas it is much lower for thermogenic and abiogenic HC (around 101). The ratio varies from 56 to 290 in gas samples from Olkiluoto (Fig. 5-23) suggesting that HC are mixtures of biogenic and thermally formed HC. Samples with higher molecular ratio values (>250) also have lower, bacterial δ13C values. The scatter of δ13C results is quite large, but they seem to show a trend starting from values between �40 to �30�, a range that may represent δ13C values of primary CH4 end-members formed in thermal processes. The carbon-13 value for the bacterial end-member of CH4 may clearly be less than �80�, a level typical of carbonate reduction as a dominant bacterial pathway for CH4 formation (e.g. Whiticar 1999). The molecular ratio and δ13C have a hyperbolic dependence with enrichment of the bacterial end-member in the system, therefore it is not possible to estimate the amount of different end-members. The Figure 5-23 indicates, however, that the main part of CH4 has a thermal origin, particularly in those samples with the highest CH4 contents. The trend presented in Fig. (5-23a) does not shift left enough with increasing molecular number suggesting that the system has also other end-members/processes than the one thermal and microbial HC values shown. Oxidation of methane shifts samples roughly along the presented trend, but if higher HC is also oxidised (i.e. the number maintains stable) the shift is to the right as CH4 oxidizes due to fractionation of δ13C (methanotrophs prefer lighter isotopes and so the residual CH4 becomes heavier). These processes may have diluted the isotopic signature of biogenic CH4. In part, an abiogenic origin of thermal HC may be one explanation that confuses the trend in Fig (5-23a).The deuterium signature of CH4 in groundwater at Olkiluoto is typically slightly low to thermogenic CH4 (Fig 5-24). Abiogenic, hydrothermal methane has a relatively lower deuterium content than thermogenic according to Sherwood-Lollar et al. (1993, 2002) and Whiticar (1999). On the other hand, abiogenic HC has a decreasing pattern of δ13C values among the C1-C4 alkanes, particularly between C1 and higher HC that is not observed in gas samples from Olkiluoto (App. 2). However, bacterial CH4 formation may have lightened the δ13C signature of CH4 at Olkiluoto.

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Figure 5-23. a) δ13C(CH4) and b) CH4 contents vs. C1/C2+C3 ratios for hydrocarbon gasses at Olkiluoto. Evoluotionary trend based on increasing bacterial end-member is shown. Curve is calculated mixing line between for possible bacterial (CO2 reduction) and thermogenic mixtures with end-member compositions of (-100�, 1000) and (-40�, 40).

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Figure 5-24. 13C-2H variation in CH4 at Olkiluoto. The fields classification of bacterial abiogenic and thermogenic origin of CH4 are drawn according to Sherwood-Lollar (1993) and Whiticar (1999). The origin of HC at Olkiluoto is a very complex question. All three sources; bacterial, thermogenic and abiogenic are identified based on isotopic and chemical data. Bacterial formation of CH4 seems to be active currently. Favourable conditions for thermal HC may have prevailed during Precambrian and early Phanerozoic times, which may connect thermal HC to the formation of the brine end-member at Olkiluoto. However, abiogenic HC may have formed in situ, but thermogenic HC may have inflowed from an ancient sedimentary basin, the remnants of which still occur outside in the bottom of Gulf of Bothnia. Abiogenic HC may naturally have been formed at a much greater depth and has migrated to shallower depths more recently. Such slow migration may cause the near saturation state of CH4 in the deep samples (Gascoyne 2000). Bacterial methanogenesis may also increase CH4 content to the near saturation state. On the other hand, why CH4 is not saturated elsewhere in deep, methanic groundwater zone if CH4 is constantly yielded in the groundwater system either by migration or bacterial formation. Hydrogen gas contents are significantly lower above 600m, which limits bacterial carbonate reduction and so indirectly support its significance in creating the near saturation state in great depths.

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5.6 Thermodynamic constraints of the groundwater system

The results of thermodynamic solubility calculations are used to test potential solubility controls and uncertainties in measured pH and Eh values. These values are evaluated by calculating values corresponding the chemical equilibrium of calcite and the activities of available redox pairs. Thermodynamic data generally contain uncertainties, particularly solubility data of solid phases. In PHREEQC (Parkhurst & Appelo 1999) calculations phreeqc.dat are used for aqueous species, gases and mineral phases that are consistent with the aqueous model of WATEQ4F (Ball & Nordstrom 1991) and the compilation of Nordstrom et al. (1990). The latter compilation has been thoroughly revised and has reliable and consistent thermodynamic data for water-mineral reactions that contain relatively few minerals, which are considered to show reversible solubility behaviour necessary to attain chemical equilibrium in natural aquatic systems. However, the data for calcite is reliable and reliable estimates are possible for quartz and its varieties, chalcedony and amorphous silica, by the degree of crystallinity. Solubility calculation results also may contain uncertainties derived from chemical input data. Calcite is one of the most common fracture minerals. Previous studies have already shown its importance in buffering and controlling the pH in groundwater at Olkiluoto. Calculated saturation indices of calcite for groundwater samples show clearly how strongly undersaturated, shallow, fresh groundwaters rapidly attain calcite equilibrium at shallow depths (Fig. 5-25a). Calcite is at or close to equilibrium in most deep groundwater samples although we use ± 0.2 as a strict uncertainty limit. This strongly supports the role of fracture calcites, controlling groundwater pH at Olkiluoto. Oversaturation is most distinct in the lower part of the HCO3-rich groundwaters. These samples may be oversaturated due to retarded precipitation of calcite. Deeper oversaturation is much more limited except for one saline sample, (KR1/S1), which has been sampled during the 1980�s and known to have uncertainties in pH value (e.g. Pitkänen et al. 1999a). High partial pressures of CO2 (expressed here as logPCO2) in overburden groundwaters (Fig. 5-25b and c) are usually derived from the soil produced by vegetation as indicated carbon isotopes. The evolution of PCO2 depends on the �openess� of the system (Clark & Fritz 1997). If the system is open (i.e. continual exchange of CO2 between DIC and the soil atmosphere) PCO2 remains fairly stable, but in closed systems it decreases with increasing DIC. The enrichment of DIC depends on whether recharging groundwater dissolves silicates or carbonates (the latter system has an additional source of carbon to increase the DIC). The fairly constant values in overburden samples that have strongly increasing DIC indicate open system dissolution of calcite. Calcite has not been observed in the overburden at Olkiluoto (bulk analyses in Lintinen et al. 2003), but the indication from high DIC values is rather strong (detailled mineralogical analyses are needed from overburden). Calcite may occur in till (crushed bedrock) or in land depressions where previous sea bottom sediments are preserved. These are a potential calcite pool due ancient anaerobic processes (see ch. 5.5.2-5.5.3).

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Figure 5-25. The variation of a) calculated calcite saturation indices and b) partial pressure of CO2, with depth, and c) the variation DIC with PCO2 in Olkiluoto groundwater samples. Calcite equilibrium tolerance (±0.2), partial pressure (pCO2 = -3.5) for atmostpheric concentrations and shallow wells in HCO3-rich groundwater data are also specified in the figures a ,b and c, respectively. The depletion of PCO2 in the shallow bedrock wells may primarily result from organic derived carbonic-acid promoted dissolution of fracture calcite as carbon isotopes also indicated. This reaction probably dominates to neutralise pH, dissociate CO2 to HCO3, increase DIC and decrease PCO2.The decreased values PCO2 with the highest DIC values in the shallow wells indicate that the system achieves a closed state in the shallow

77

bedrock. At greater depth, where calcite is also saturated, the decrease in PCO2 values may derive from precipitation of fracture calcite and dissolution of silicates and mixing of groundwater types.

The partial pressure of CO2 is also an important measure to evaluate possible uncertainties in measured pH developed during sampling. Samples in the upper part of the bedrock have clearly higher CO2 partial pressures than the atmosphere (-3.5). Therefore some degassing of CO2 is possible during sampling and measurements, particularly as hydrostatic pressure at sampling depth increases relative to atmospheric pressure. If CO2 is released from the sample, pH and the calculated SI of calcite will increase. The group of samples with calcite oversaturation in the lower part of the HCO3-rich groundwaters has a high pCO2 and some degassing during sampling may also explain elevated SI values instead of kinetic constraints.

Deep groundwaters show clearly that they have a lower pCO2 than in the atmosphere and this may cause a CO2 ingassing problem near and at the surface, during sampling. The variation in the results are propably due to contamination problems These waters also have an enormous total gas content, which may easily be degassed due to decreasing pressure in near-surface conditions. The minor CO2 may also evacuate with other gases thus raising the pH of the water sample. Gas samples taken with the PAVE sampler are considered much more reliable and they show that CO2 content is very small in saline groundwaters. These uncertainties hamper the carbon isotope studies of DIC in deep saline groundwaters as shown in ch. 5.5.3.

The results of solubility calculations indicate that fracture calcites are able to control the pH. The fast kinetics of calcite precipitation is considered (e.g. Plummer et al. 1978) to favour attainment of chemical equilibrium in groundwaters at Olkiluoto for the long residence times estimated for deep brackish and saline groundwater types. Therefore, the pH of samples can be calculated assuming calcite equilibrium (SIcalcite = 0). The results are shown in Fig (5-26). Generally, calculated pH values are sligthly smaller than measured values in deep groundwaters. The mean deviation is about 0.3 unit in HCO3 �rich groundwaters and smaller within other groundwater types, indicating good reliability of pH data.

The pH calculations also suggested degassing effect on HCO3-rich samples. Although partial pressures increased up to 0.6 log units relative to calculated values based on measured pH, test calculations showed that the effect on DIC was only few tenths of mmol/l, i.e. few mg/l, thus degassing may have an error that is only a few percent on measured DIC values. The pH calculations also suggested CO2 degassing for some saline samples. The potential error is also a few percent on DIC values. The problem centres mainly on understanding carbon circulation in saline groundwater, if both degassing and ingassing occurs during sampling. Carbon isotope studies should be done in future from pressurised samples taken with the PAVE-sampler.

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Figure 5-26. Calculated (red labels) and measured pH (black labels) versus depth at Olkiluoto groundwaters.

The interpretations of redox conditions given in chapter 5.5 showed that redox conditions at Olkiluoto groundwaters are anoxic (except for a very narrow surface layer) and thus Eh values should clearly be below 0 for the observed pH range. The results also show that prevailing redox pairs that seem to control the redox state at Olkiluoto are SO4

2-/HS- and CO2/CH4, which either do not respond to platinum electrode or respond partially as some species in the whole SO4 reduction series (Macalady et al. 1990, Appelo & Postma 1993). Ferrous iron is also an important indicator of redox conditions at Olkiluoto, but its redox pair, either Fe3+ or FeO(OH) precipitate, seems to be lacking in the deep groundwater system, or the concentrations are too small for observation.

The Eh is calculated from the activities of SO42-/HS- and CO2/CH4 redox pairs which

correspond to pH values equilibrated with calcite (Fig. 5-27). The calculated Eh-values for the sulphur pair vary from �280 to -170 mV and for the carbon pair from � 330 to �270 mV. These levels have been reached in part of Eh measurements (cf. Fig. 5-15), maybe least in saline groundwaters, in which dissolved sulphide is observed most rarely suggesting that dissolved sulphide has, however, the best capacity to maintain low Eh during sampling. Table 5-8 shows statistical parameters of the calculated Eh values for different groundwater types.

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Figure 5-27. Calculated Eh for SO42-/HS- (S) and CO2/CH4 (C) redox pairs versus pH

corresponding calcite equilibrium (c.f. Fig. 5-15).

Table 5-8. Statistics of calculated Eh for SO42-/HS- (S) and CO2/CH4 (C) redox pairs

and pH corresponding to calcite equilibrium for different groundwater types. Dil.-brack. HCO3 Brack. SO4 Brack. Cl Saline Average pH ± std. 7.5 ± 0.25 7.6 ± 0.25 8.1 ± 0.24 7.8 ± 0.33 Average Eh ± std. (SO4

2-/HS) -190 ± 11 -200 ± 20 -250 ± 23 -250 ± 22

Average Eh ± std. (CO2/CH4)

-300± 25 -290 ± 19

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6. HYDROGEOCHEMICAL EVOLUTION

6.1 Origin of groundwaters and palaeohydrogeology

The hydrochemical data and previous interpretations (Pitkänen et al. 1996, 1999a) have already revealed the complex nature of sources of salinity and chemical evolution at the Olkiluoto site. The new data, since 1999, strengthen the previous concept of hydrogeochemical evolution and augment it in particular with data for shallow groundwaters, salinity distribution and information on gases. Changes in climate and geological environment have had a significant effect on local palaeohydrogeological conditions. Chemical and isotopic signatures have left clear imprints on the current groundwater compositions and caused great variability in the hydrochemical data (Fig.6-1) notably in salinity, water types and relative contents of conservative tracers such as Cl, Br, δ2H and δ18O (Fig. 6-2). Hydrochemical data have also revealed the extensive mixing phenomena of different end-member waters from each of the palaeo sources. It is worth to note that the mixing phenomena may include advective mixing of large volumetric groundwater bodies as well as uptake of small, hypersaline pore fluids to cause salinity increase.

Figure 6-1. Vertical variation of the main hydrochemical parameters, microbes and pH and redox buffers in fractures at Olkiluoto (updated since Pitkänen et al. 1999a). Vertical lines represent steady-state or smooth changes. SRB, IRB are sulphate and iron reducing bacteria, respectively. (Note: microbes have been sampled and analysed for the depth range 200-900 m; there is no data for the upper part of the bedrock).

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The interpretation of chemical and isotopic data indicates that there are at least five end-member water types influencing current groundwater compositions at the site. They originate from different epochs, ranging from modern times, through former Baltic stages to preglacial times: modern - meteoric water infiltrated during terrestrial recharge and - seawater from the Gulf of Bothnia (0 � 2 500 BP)

relic - Litorina Seawater (2 500 � 7 500 BP) - fresh water prior to the Litorina Sea stage containing glacial melt water

(7 500 � 10 000 BP) and - saline water (brine) intruded and/or formed under the influence of

hydrothermal activity (pre-Quaternary, probably early Phanerozoic to Precambrian in age).

Groundwater representing modern water types occurs in the upper 100 � 150 m (Fig. 6-1). The latest stage of groundwater formation is dominated by dilute HCO3-rich infiltration, with stable isotopic signatures of water (Figs 5-10, 6-2), corresponding to current, dilute meteoric recharge in shallow depths in soil and bedrock. Seawater input is indicated by increasing salinity, seawater signatures of Br/Cl (Fig. 6-2) and Na/Ca ratios, and increasing SO4 with a marine 34S signature and higher Mg contents (Figs 5-5, 5-7, 5-12). The HCO3-rich groundwater type has been formed since Olkiluoto Island rose above sea level about 2500 years ago, but the general observation of tritium and moderate radiocarbon values (50 - 60 pM or higher) in shallow groundwater, indicate mostly recent recharge, in the last few decades. Therefore, the upper part of the bedrock, characterised by HCO3-rich groundwater, seems to be hydraulically dynamic. Although the HCO3-rich groundwater is mostly recently recharged, the deeper samples of this layer show depletion in 14C content relative to young groundwaters (Fig-5-19). Rough calculations give a range from 500 to 3000 years according to 14C half-life and previous mass-balance calculations (Pitkänen et al. 1999a).

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Figure 6-2. Relationship between δ18O and a) δ2H and b) Br-Cl ratio in Olkiluoto groundwater samples. Arrows depict roughly the composition change with depth. The red arrow indicates the change from overburden and dilute/brackish HCO3 waters towards brackish SO4 waters (highest SO4 concentrations), the blue shows the trend from brackish SO4 waters to brackish Cl waters, and the black shows brackish Cl waters to saline waters. (Note: overburden samples were omitted from part b) for clarity.)

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The distinctly SO4-rich brackish water below the HCO3-rich groundwater (Fig. 6-1) at 100 � 300 m depth, most probably infiltrated during the Litorina Sea stage (Pitkänen et al. 1996), when the island was still below sea level. The marine origin of the water is indicated by the Br/Cl ratio which is typical of seawater (Fig. 6-2), the lower pH and alkalinity, much higher SO4 and Mg than in HCO3-rich groundwater, and the shift of stable isotope composition of water towards a Baltic composition (Fig. 6-2). The Cl content is higher than modern Baltic water (3 200 mg/l) and this is consistent with the estimates presented for the Litorina Sea (Table 3-1). During the main part of the Litorina period, between about 7000 BP and 4000 BP, TDS in the seawater was about 4� higher than in the modern Baltic Sea in Finnish coastal areas (Donner et al. 1999). Neither the Br/Cl ratio, stable isotopes of water, high SO4 and Mg concentrations, Cl isotopic signatures or the principal components (Fig 5-2) support mixing of the brine component as the cause of Cl enrichment. The 14C content of dissolved carbonate in most SO4-rich samples (20 � 30 pM) corresponds to the period 7500 to 4000 years in relation to more recently recharged HCO3-rich groundwater (cf. Ch. 5.5.2) that is appropriate to the Litorina Sea stage. Other evidence for a Litorina Sea intrusion was the observation by Gascoyne (2001) that these groundwaters all had low 36Cl/Cl ratios, comparable to modern Baltic seawater, in contrast to freshwater recharge and underlying saline waters. The interpretation of glacial melt water input is based on the decreasing stable isotopic values of groundwater (Fig. 6-2) in the lower part of the brackish groundwater layer. Salinity shows a slight decrease (Fig. 6-1), and radiocarbon (4 � 25 pM) indicates mean ages corresponding to deglaciation �or the last glacial period. The chemical composition loses its seawater signatures: SO4 and Mg decrease to a minor level (Figs 5-5, 5-7), the Br/Cl ratio increases (Fig 6-2) and Cl isotopes deviate from a marine origin (5-13, 5-14), suggesting a source of dissolved solids other than seawater. The initial source is considered to be same as in saline water, i.e. ancient brine. Groundwater salinity starts to increase significantly at 400 � 500 m depth without showing any indication of levelling off (Fig. 5-3) or showing any discontinuity in the main chemical variables (Pitkänen et al. 1999a). Calcium becomes the dominant cation, the Br/Cl ratio (Fig. 6-2) reaches twice the level of seawater, and the stable isotopes of water, especially 2H, show enrichment above the Global Meteoric Water Line. In particular, highly saline deep groundwater (below 700 m) seems to have been preserved undisturbed during the last glaciation in Quaternary time. Radiocarbon, in spite of its uncertainties, shows values below detection limit (KR4/860). Chemical calculations based on hydrogeochemistry (Pitkänen et al. 1999a), in situ 36Cl/Cl in equilibrium with host production rates (Gascoyne 2001), U-series activity ratios approaching equilibrium and the shift above GMWL, all combine to suggest a very long residence time, million year scale or probably longer. A similar chemical tendency to that of saline groundwater (Ca-Na-Cl) is observed in highly saline (up to 30 wt. %) fluid inclusions in fracture calcites, which are formed at moderate to high temperature (45 � 240 ûC according to Blyth et al. 1998, Gehör et al. 2002). Abundant hydrocarbons in highly saline groundwater also show mainly thermal formation connecting the origin of brine component to hydrothermal conditions. Indications from calcites, fluid inclusions, dissolved gases and isotopic signatures in saline groundwater, suggest an extremely early origin for the original brine end-

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member, which is impossible to connect to any well-known event during the Quaternary. Probably, the formation of the brine end-member which has occurred during the early Phanerozoic or late Precambrian, when this area may have been covered by thick sediments, the remnants of which still occur nearby in the bottom of Gulf of Bothnia. Recently, a theory has been published (Starinsky & Katz 2003) that Fennoscandian brines are residual fluids, which have formed by seawater (i.e. ocean) freezing during Quaternary glaciations (mainly the Weichselian) along the subarctic continental margins, around ice sheets. The theory is significantly short on information that indicates the major ion compositions (Na, Ca, Cl, Br) in Canadian and Fennoscandian Shields are similar and that they trace the evolutionary line of marine brines developed during freezing of seawater, as presented by Herut et al. (1990). This naturally requires long-range migration of the cryogenic brines and infiltration into their current sites under a highly dynamic flow environment. Brine formation by the freezing process has potential and groundwaters may have been evolved under permafrost conditions as suggested by Smellie et al. (2002) giving a possible origin for uncommon Na-SO4 type groundwater at the Palmottu site. The information above from the brine component at Olkiluoto does not support a cryogenic origin, compared with one of elevated temperatures. In addition, the cryogenic theory is not able to explain plausibly why the major ion composition of brine components in Finland changes significantly in a different host rock environment or why the stable isotope composition of water tends to shift clearly above global meteoric water line in saline groundwater. Blomqvist (1999) has shown that major ion relationships are clearly dependent on host rock type and composition deviates strongly from the seawater freezing line (Fig. 6-3). A notable difference is observed even between felsic host rocks. The Mg concentrations of saline groundwaters at Hästholmen, is over 500 mg/l (Pitkänen et al. 2001) whereas it approaches 100 mg/l at Olkiluoto and is less than 10 mg/l at Äspö (Laaksoharju et al. 1999). Mineralogical evidence suggested that the high concentration of Mg in Hästholmen was originally released in groundwater by hydrothermal activity. Freezing does not shift δ2H � δ18O values above the GMWL (Herut et al. 1990). The shift is generally considered to result from strong water-rock interaction where δ18O is fractionated during mineral alteration to hydrated silicates. However, this must take considerably longer time than since the Weichselian glaciation otherwise the shift should be observable almost in all water types at Olkiluoto. Fractionation of 18O needs significant O-isotope exchange between water and wall rock minerals to be measurable, which suggests elevated temperature are required, for example late stage hydrothermal fluids, initially > 120 °C as suggested by Sheppard (1986). The isotopic data for fracture calcites and fluid inclusion temperatures have been used to calculate fluid isotopic compositions that precipitated calcites at Olkiluoto (Blyth et al. 1998). The results show δ18O-values from -2 to +12�, which signifies a remarkable shift compared with present-day groundwater samples at the site (-13 to -9�) or Baltic Seawater (-7�). This shift implies that the fluids were hydrothermal in nature and could have been from several sources such as magmatic, seawater or basinal brines (Blyth et al. 1998). However, measured δ37Cl-values of about 0� from present deep saline groundwater samples, support a seawater origin for chloride.

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Figure 6-3. The Na/Cl-Br/Cl relationship for selected brines from various rock types in Finland. The lines represent freezing and evaporation paths of modern seawater according to Herut et al. (1990). The data from Juuka (serpentinite), Ylivieska (gabbro) and Pori (arcosic sandstone) are from Blomqvist (1999), and Olkiluoto (mica gneiss, KR4/860). Concentration of Cl varies between 52 to 100 g/l at the other sites.

6.2 Evolutionary processes

The interpretation of hydrogeochemical data indicates that mixing controls the wide salinity variation in groundwaters at Olkiluoto. Water-rock interaction, such as carbon and sulphur cycling and silicate reactions, buffer the pH and redox conditions (Fig 6-1) and stabilise groundwater chemistry. However, the extensive difference between the compositions of groundwater types also causes significant deviations in water-rock interaction processes, particularly in redox chemistry. This is observed as discontinuations in data of certain variables such as constraints sensitive to redox conditions, δ13C(DIC) and δ34S(SO4) and emphasises the changes in chemical conditions particularly in transition zones between groundwater types, which must be considered in evaluating previous groundwater evolution and conditions in the future.

6.2.1 Chemical characteristics in shallow (< 30 m depth) groundwater Pitkänen et al. (1999a) presented a comprehensive interpretation of the evolutionary processes behind the hydrochemical changes observed at Olkiluoto. The data have been since extended, particularly with results from shallow depths, the overburden (data from 12 observation tubes) and the first few tens of metres of upper bedrock (data from 10 wells with max. depth 23.8m). Shallow groundwater characteristics are important for evaluating chemical processes during recharge. Changes in shallow hydrogeochemistry are an intial step in understunding groundwater evolution both for natural evolution and

87

in evaluating geochemical evolution during the operation of nuclear waste repository. Infiltration at shallow depths and related evolutionary processes at Olkiluoto, are fairly rapid events, because considerable tritium contents in groundwater samples show short residence times, a few dozen years at most. Weathering processes induced by dissolved gases, i.e. CO2 and O2, dominate typically in shallow groundwater circulating in soil and bedrock. Oxygen is derived from the atmosphere as well as traces of CO2 that already buffer the pH of rainwater to about 5. However, the main resource for the latter gas is the organic soil layer, where the decay of plant debris and plant root respiration produce the majority of CO2 (Eq 6-1) that dissolves in recharging water as indicated by carbon isotopes in DIC (Fig. 5-20). Infiltrating oxygen is partly consumed in these processes and partly oxidises other reduced substances such as ferrous ion and sulphides (Eq. 6-2) according to the isotopic results of SO4. Acid waters that are produced by aerobic oxidising reactions are normally neutralised relatively quickly in natural conditions. Weathering of silicates (Eq. 6-3) is even kinetically sufficient to attain neutral or slightly alkaline pH conditions in groundwater in overburden. In particular, if carbonate minerals occur in the overburden dissolution of them neutralises (Eq. 6-4) infiltrating water quickly (Appelo & Postma 1993). CH2O + O2 → CO2 + H2O ↔ H2CO3 ↔ H++ HCO3

- (Eq. 6-1)

FeS2 + 3.75 O2 + 3.5 H2O → Fe(OH)3 + 2SO42- + 4 H+

pyrite (Eq. 6-2)

2 KMg1.5 Fe1.5AlSi3O10(OH)2 + 0.75O2 + 8 CO2 + 16.5 H2O → biotite

2 K+ + 3 Mg2+ + 8 HCO3- + 3 Fe(OH)3 + Al2Si2O5(OH)4 + 4 H4SiO4

(Eq. 6-3)

CaCO3 + CO2 + H2O ↔ Ca2+ + 2HCO3

- (Eq. 6-4) Mass transfer in weathering processes is small due to low concentration of protons available from dissociation of carbonic acid (Eq 6-1) or oxidation of ferrous minerals. The maximum amount of dissolved CO2 in soil groundwaters is a few millimoles per litre (e.g. Pitkänen et al. 1996, 1998, 1999a, 2001) and the solubility of O2 is few tenths of millimoles (about 10 mg/l, i.e 0.3 mmol/l, Matthess 1982) in temperate climate conditions. Consequently, dissolved solids released in groundwater by weathering processes are also limited to a few millimoles (i.e. few hundreds of milligrams per litre of total dissolved solids depending on available elements). Therefore, pH is probably the most sensitive variable to register the initial chemical interactions in recharging groundwaters. Alkalinity is another variable showing direct changes due to pH variation, but enrichment of alkalinity may also reflect carbonate dissolution in the system. Dissolved iron and/or sulphide may instead indicate an important change from oxic to anoxic redox conditions. Mixing of former water-types in addition to water-rock interaction may significantly affect the concentrations of other main chemical variables

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(Cl, SO4, Na, Ca, K, Mg). Examination of 13C and 14C of DIC and 34S(SO4) and Sr isotopes gives further information of the prevailing processes in groundwater. The plots of pH and alkalinity against Cl (Fig. 6-4a) for shallow groundwaters reveal two clear features. Firstly, the distribution (low or high values) of pH and alkalinity (Fig 6-4b) in shallow groundwaters does not generally depend on sampled from soil or shallow bedrock. This suggests that single sampling points greatly differ with their hydrogeological positions. Secondly, water-rock interaction starts and neutralises infiltrating meteoric water without any significant mixing with older groundwater types, because Cl stays nearly constant during the initial increase of alkalinity and pH.

Figure 6-4. Variation of a) pH and b) alkalinity with Cl in shallow groundwaters at Olkiluoto (PVP = groundwater observation tube in overburden, PR = percussion drilled well in bedrock and PP = shallow borehole in bedrock).

Low pH and alkalinity values belong to sampling points located in the middle of the drilling area of deep boreholes (Fig. 6-5a) and high values occur in shallow groundwaters at the margins of the site. However, the distribution does not correspond with topography at the site. For example, PVP5, PP9 and PR3 are also located at the higher central part of the island, but these groundwater samples are more mature hydrochemically. Boreholes PR1, PR2 and PR4 were sampled for the first time in the late 1990�s, when alkalinity (Fig. 6-5b) in PR2 and PR4 were over 4 and 2 meq/l, respectively. One explanation may be that pumping activities performed at the drilling site have accelerated groundwater infiltration and mixing at shallow depths. Mixing of older, more saline groundwater types in infiltrating meteoric water seems to be more common in the southern part of the site. Particularly high salinity is found in PP7, where Cl (Fig. 6-5c) reaches almost 600 mg/l. The borehole is located in the vicinity of deep borehole KR3 in which brackish palaeogroundwater types were

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observed to occur at shallower depths than in the other parts of the site (Pitkänen et al. 1999a). The borehole PP7 has been drilled just in an area where subvertical fracture zone R10 (Fig. 3-1) has been interpreted (Saksa et al. 2002) to be exposured at the surface and a significant depression (Hagros 1999) in bedrock topography occurs here. These geological features and the water chemistry suggest that PP7 is located in a discharge area. The other shallow holes with increased salinity (PVP3, PVP4, PP2 and PVP10) also locate in the exposure zones of interpreted fractured structures, R18 and R7, and partly in bedrock depressions. The variation of cations, SO4 and DIC can give information on processes underlying the input of dissolved solids in groundwater during infiltration into the bedrock. In addition to alkalinity, DIC (Fig. 6-6a) also increases with Cl and pH in shallow groundwaters indicating a carbonate source in the system. This could be either carbonates, predominantly calcite that should occur in the overburden as well as in fractures and/or oxidation of organic carbon. Unfortunately, carbon isotope data are very limited for overburden groundwaters. Calcite is a likely source, at least in the upper bedrock, according to carbon isotopes (Fig. 5-20). The enrichment of calcium concentration in overburden samples (Fig. 6-6b) corresponds with DIC and alkalinity indicating that calcite dissolves. Calculated constant partial pressures of CO2 with strongly increasing DIC indicated calcite dissolution in the overburden (Fig. 5-25b) in an open system (CO2 exchange with soil) whereas the results show a closed system for calcite dissolution in shallow bedrock. The controlling role of calcite seems evident for the evolution of pH, alkalinity and DIC in groundwaters directly from the recharge at shallow depth. However, the relative enrichment of DIC is higher than can be expected from pure mass-balances in the calcite dissolution reaction (Eq. 6-4), i.e. carbonate addition is double the initial value. Therefore, oxidation of organic carbon in overburden samples with elevated DIC values should be larger than in the low DIC, pH and alkalinity samples (about 7 mg/l) at the central area of the site. The source process may be partly aerobic and partly anaerobic and result from longer interaction in the organic-rich soil layer during recharge in margin areas than in the central area. Mass-balance calculations will further show a more detailed picture of DIC sources, how much is derived from calcite and how much from organic carbon oxidation.

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a)

b)

c)

Figure 6-5. Aerial distribution of a) pH, b) alkalinity (meq/l) and c) Cl (mg/l) in shallow groundwaters sampled from observation tubes (PVP) and shallow wells in bedrock (PP, PR) at Olkiluoto.

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Figure 6-6. Variation of a) DIC and b) Ca against Cl in shallow groundwaters at Olkiluoto (PVP = groundwater observation tube, PR = percussion drilled well and PP = shallow borehole).

The main cations and sulphate all increase from low values observed in shallow sampling points at the central area (Fig.6-5a-c). The enrichment is clearly stronger than could be attained by seawater mixing indicating that other soluble phases than calcite are available. Dissolution of silicates such as plagioclase (Eq. 6-5) was suggested earlier (Pitkänen et al. 1999a) as a potential source in a low pH environment. This is supported by elevated dissolved silica concentrations (Fig. 5-4) even in acid groundwaters, thereby emphasising silicate-water interaction in early weathering reactions. Elevated strontium isotope ratios (Fig. 5-12a, b) in shallow waters suggest the dissolution of K-rich phases such as K-feldspar and biotite (Pitkänen et al. 1999a). Dissolution of biotite (Eq. 6-3) is also supported by initial enrichment of K, Mg and Fe. Biotite hydrolysis may play a role in consuming oxygen during recharge in glacial deposits (c.f. Blum & Erel 1997, Bullen et. al 1997, Pitkänen et al. 1999a) and changing the infiltrating meteoric water into a post-oxic state.

Na0.8Ca0.2Al1.2Si2.8O8 + 1.2 H+ + 0.6 H2O → plagioclase

0.8 Na+ + 0.2Ca2+ + 0.6Al2Si2O5(OH)4 + 1.6 SiO2

(Eq. 6-5)

Sodium and calcium show relatively large concentration ranges with higher Cl contents, indicating different type of evolutionary paths in the shallow groundwater chemistry. For example, PVP4 and PVP5 have the highest Ca contents, but fairly low Na, whereas PVP3 and PVP10 represent the opposite trend (App. 2). However, Na (Fig 6-7a) shows the clearest increasing trend of major cations with Cl. These observations probably

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reflect changes in chemical processes with evolutionary progress. Dissolution of minerals no longer dominates mass changes. Mixing with older groundwater components and cation exchange probably control cation concentrations. Cation exchange is definitely an important process in an aquifer which has been exposed to sequential meteoric and seawater intrusions (Nordstrom 1986, Appelo & Postma 1993). Cation exchange probably retards Ca (Fig. 6-6b) and Mg (Fig. 6-7c) enrichment in the groundwater, particularly as Cl increases and Na is released into the groundwater.

Figure 6-7. Variation of a) Na, b) K, c) Mg and c) SO4 with Cl in shallow groundwaters at Olkiluoto (PVP = groundwater observation tube, PR = percussion drilled well and PP = shallow borehole).

Sulphate concentration in low pH groundwaters is about 5 mg/l increasing to 30 mg/l without any Cl increase and so saline groundwater mixing (Fig 6-7d). Beyond this initial enrichment, SO4 seems to increase along a mixing line towards deep HCO3-rich and SO4-rich groundwaters. The initial enrichment of SO4 is higher than can be expected by oxidation of sulphide minerals below the groundwater table. Although

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dissolved oxygen that is theoretically available would be totally consumed in sulphide oxidation, SO4 production is limited to about 10 mg/l. Isotopic data of SO4 in shallow groundwaters are fairly few at present (Fig 5-21), but the values suggest that in addition to oxidised ferro sulphides, marine SO4 is one important component affecting in the initial enrichment of SO4 in shallow groundwaters. Pitkänen et al. (1999a) suggested that the excess amount of SO4 is released by competitive adsorption between anions according to a theory proposed in the study of a redox experiment at Äspö by Bruton & Viani (1997) and Banwart et al. (1999). This predicts that increasing bicarbonate concentrations at near-neutral pH may displace adsorbed sulphate by anion exchange. Initially, marine SO4 would have been adsorbed onto iron oxyhydroxide surfaces in the soil or near-surface fractures in the bedrock during seawater infiltration e.g. close to the shore-line. The Fe contents in part of shallow groundwaters are high at neutral pH suggesting anoxic redox conditions. Observations of hydrogen sulphide were common during sampling of shallow groundwaters (Hatanpää 2002, Backman et al. 2002) and it also frequently exceeded the detection limit in chemical analyses. The results indicate that groundwater generally approaches sulphidic redox conditions already in the overburden and shallow bedrock, which also means that oxygen is consumed rapidly and in a short flow path during recharge. A few shallow wells (PR1, PR2 and PR4) drilled mainly on the outcrops in the central part of the site are the only exceptions. However, the data from these wells show differences in initial oxidation processes, which reflect differences in sulphide occurrences. The well PR2 is bored in microcline granite, where sulphides are very rare. Isotopic results indicate a pure marine origin for SO4 from this well whereas they indicate oxidative dissolution of high temperature sulphides for SO4 at PR1 and PR4, which represents migmatitic mica gneisses, relatively rich in sulphides. Uranium concentrations at shallow depth in bedrock show a large range. Five of the analyses (samples from PR1, PP8) show low contents below 1 ppb, and activity ratios (AR) that deviate only slightly from 1.00. These values are considered typical of U dissolution in a dynamic oxidising environment (e.g. Pearson et al. 1991). Uranium concentrations in PR2 are the highest observed values (max. 18.2 ppb) at Olkiluoto. The AR value is still near unity reflecting the higher U-content in microcline granite and the importance of the host rock in water-rock interaction processes. The groundwater in well PP7 has a moderate U content (about 4 ppb) and fairly high AR (4.7), which may represent U dissolution in a reducing environment and a prolonged residence time. However, this sample also contains tritium a though a smaller amount (7 TU) compared with other shallow samples (12 - 17 TU). This suggests that the sample is a mixture of an older and younger component as is also indicated by an increased salinity content. Increased AR may originate from a more saline, older water component, which also supports the discharging nature of the sample site. Overall, the isotopic results of SO4 and U show many details in geological and hydrogeological conditions at shallow depth. Therefore, we propose a detailed sampling and analyses programme for shallow groundwater conditions, because it is evident that the results are able to improve significantly the understanding of the hydrogeological system near the surface.

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6.2.2 Brackish and saline groundwater (depths below 30 m)

The ongoing extended sampling programme has also increased the number of samples from deep bedrock. The objective has been to verify previous data, improve the sampling network in the bedrock under study and augment the understanding of baseline hydrogeochemistry achieved, for example, in Pitkänen et al. (1999a). The programme has particularly emphasised to get more information on saline groundwater and gases and their distribution, and possible chemical changes due to long-term pumping. Although the detailed information of processes has increased significantly (e.g. gases, Ch.5.5.6) since Pitkänen et al. (1999a), the basic concept of hydrogeochemical evolutionary processes deep in the bedrock at Olkiluoto has remained unchanged. Deep in the bedrock, dissolution of minerals no longer dominates mass change in HCO3-rich groundwater. Major ion plots in Ch 5 indicate that cation exchange and particularly mixing with older groundwaters that have a dominant marine signature controls the main mass changes. Dissolved inorganic carbon approaches its highest value at a depth of a few tens of meters where calcite dissolution ends due to attaining saturation. Mixing has mainly caused the decrease in DIC content, but minor precipitation of calcite is possible due to a slight input of HCO3 if bacterial SO4 reduction occurs as microbial studies indicate (Haveman et al. 2000). The Litorina seawater character is currently strongest in the upper part of the SO4-rich groundwater layer and the change from HCO3-rich groundwater (at 100-250 m depth depending on location) is fairly abrupt suggesting limited mixing between the groundwater types. This probably resulted from the density difference between the waters. The increased Ca content (likely released by Na to Ca exchange during seawater infiltration) and calcite equilibrium in SO4-rich groundwater have likely fixed pH slightly above 7.5 and adjusted alkalinity to quite a low level. Sulphate-rich groundwater seems to be stable showing significant changes in different constraints (e.g. 13C and 34S) only in the marginal areas of water body, particularly at the base, where different water types tends to mix. The mixing evidently induces anaerobic redox reactions. Due to problems with interpretation of Eh measurements of natural groundwaters (Ch. 5.5 and 5.6), Berner (1981) proposed a classification of redox conditions in terms of indicative redox species or dominant electron accepting processes which usually form a reduction sequence with residence time and flow distance from the ground surface (Figs 6-1 and 6-8). The redox conditions are first divided into oxic, which is limited at Olkiluoto to a very shallow depth (few metres), and anoxic environments depending on whether they contain measurable amounts of dissolved oxygen (> 10-6 M). The anoxic environments are subdivided into post-oxic, sulphidic and methanic zones. The post-oxic zone is often characterised by ferric iron reduction and an elevated Fe2+ content, sulphate reduction occurs in the sulphidic zone (H2S+HS- > 10-6 M), and elevated CH4 is typical of the methanic zone. Eh can be estimated according to dominating redox processes from redox-pH diagrams or calculated with a thermodynamic code as was done in Table 5-8. Evolution to a certain redox zone results normally from domination of certain dissolved species or phases, which are able to buffer the redox system.

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The SO4-bearing groundwaters (both brackish HCO3 and SO4 type) form a Fe-SO4-characterized redox zone (Figs 6-8). Dissolved organic carbon seems to be buffered to very low levels already on the margins of the groundwater body due to a high concentration of reductive SO4 and, consequently, dissolved sulphide remains low and ferrous iron does not have a strict solubility limit, i.e. by pyrite. Deeper in the bedrock, SO4 may be microbially reduced by CH4 in the transition zone from Litorina-derived groundwater to a Na-Cl-type, brackish-saline groundwater at 250 � 450m depth (Fig. 6-8). This process has led locally to very low δ13C(DIC) values (Fig. 5-20). The redox chemistry at this interface seems to be in a metastable state (Pitkänen et al. 1994), as can be observed from the combined occurrence of SO4 and CH4 and occasionally anomalously high sulphide concentrations (the maximum is 12 mg/l). Sulphide production may have been even higher as some sulphide may have been removed due to pyrite precipitation. That is possible to estimate coupling sulphur mass transfer with sulphur isotopic results in mass-balance calculations (Ch. 7). Previous calculations (Pitkänen et al. 1999a) predicted maximum formation of 27 mg/l sulphide in a flow path in bedrock. The higher salinity of Litorina-derived groundwater compared to the SO4-poor layer may have caused the mixing of SO4- and CH4-bearing groundwaters. The unusual sulphide enrichment in groundwater may be due to a poorly reactive silicate iron fraction (Raiswell & Canfield 1996) that control availability of iron for precipitation of insoluble iron sulphides. Methane concentrations increase strongly in this transition zone and below 450m groundwater redox conditions are clearly methanic. The dissolution of iron is not constrained by the solubility of pyrite. The pH is slightly higher, about 8, that is allowed by calcite equilibrium due to smaller amounts of Ca and DIC than in SO4-rich groundwater. Active methanogenesis should be based on available H2 deep in the bedrock where other substrate pools are exhausted (e.g. Whiticar 1999). Hydrogen concentrations in saline groundwater also increase (Fig 5-22b) and this provides a hydrogen-driven biosphere proposed by Pedersen (2000). Low δ13C-values (<-60�) in methane and elevated molecular ratios (Fig. 5-23a) indicating bacterial origin, are observed in low CH4 groundwaters below the SO4-rich layer as well as methanogens and acetogens (Fig. 6-1). The importance of bacterial methanogenesis may increase in very deep, saline groundwater at 800-1000m depth as δ13C-values tend to decrease from �40� to �50� level. This may lead to the possibility of a gas phase as was also indicated by CH4 solubility calculations (Gascoyne 2000). The saline groundwater seems to be chemically stable. The pH and methanic redox conditions are strongly buffered. The high Ca concentration limits the dissolved carbonate to a very low level due to calcite saturation. Calcite equilibrium together with the alteration of silicates (dissolution of primary silicates and secondary precipitation) and complexation of dissolved Ca, Mg and Fe, have caused a slightly lower pH (7.5) than in the overlying brackish groundwater. High CH4 and elevated Fe concentrations (Fig 6-8) buffer sulphur/(sulphates and sulphides) species to very low levels, and even a small input of these is considered unstable and pyrite will be precipitated.

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Figure 6-8. Sequence of redox sensitive species and phases and Berner�s (1981) classification of redox environments with depth at Olkiluoto. Concentration axis is dimensionless and shows only relative changes of a single species or phase. Water types and characteristic redox species in groundwater observed at Olkiluoto are on the right.

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7. MASS-BALANCE MODELS

7.1 Introduction

Geochemical modelling is a particularly important exercise for a system with a complex hydrogeological history such as Olkiluoto, to verify the interpreted concept of hydrogeochemical evolution at the site. Mass-balance reaction models are used to determine the importance of interpreted evolutionary processes based on data from hydrogeochemistry, isotopes, mineralogy, and speciation calculations. The modelling tests the interpretation of hydrogeochemical evolution by giving information on plausible reactions and their extent, on mixing in the system, and on uncertainties in the concept and data. Successful modelling performance may also increase confidence in the ability to predict hydrogeochemical conditions in the geosphere following waste disposal, to identify potential environmental changes in the future and set initial and boundary conditions for radionuclide migration calculations. The modelling tests the reaction and mixing hypotheses (described in Chapter 6) by constructing mass-balance models, which describe the changes in chemical and isotopic composition between recharged water and down-gradient water samples. The model derived between any points along a chosen flow path by mass-balance calculations is of the form:

Initial water(s) + "Reactant phases" → Final water + "Product phases", where reactant phases enter the initial water and products leave it to produce the composition of the final water. The models computed with the NETPATH program (Plummer et al. 1994) define the mixing proportions of initial waters and net geochemical reactions of minerals and gases that may account for the observed composition of a final water. The inclusion of isotopic data in reaction modelling provides additional criteria for testing a reaction hypothesis (Plummer et al. 1994). The modelling approach that follows is similar to that of Plummer et al. (1983) and Plummer (1984) and applied by Nordstrom et al. (1990), Plummer et al. (1990), Busby et al. (1991), Waber & Nordstrom (1992), in Finnish site investigations for nuclear waste disposal by Pitkänen et al. (1994, 1996, 1998, 1999a, 2001), in the Äspö hard rock laboratory (Pitkänen et al. 1999b, Luukkonen 2001) and in the Palmottu natural analogue site by Blomqvist et al. (2000). Calculated mass-balance models presented in this chapter are extensions to previous models of Pitkänen et al. (1999a) and they increase the data on end-member water mixing proportions used in characterisation of 3-D hydrogeochemical conditions at Olkiluoto (Luukkonen et al. 2003) A mass-balance model is defined as the masses of a set of plausible phases that must enter or leave the initial solution or mixture of initial solutions in order to define exactly a set of selected chemical data (elemental, isotopic and electron balance) observed in a final (evolutionary) water. A phase is any mineral or gas that can enter or leave the groundwater along the evolutionary path. Ion exchange and organic matter are included as phases. The validity of the mass-balance models significantly depends on selecting

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appropriate phases in the model (i.e. only phases that occur in the system and have been observed to show relevant water-mineral interaction features should be considered in modelling). The treatment of mass-balance equations, redox state and isotopic calculations with fractionation factors in the NETPATH code is discussed in greater detail in the report of Plummer et al. (1994). The mass-balance modelling of NETPATH applies strictly to the case of a chemical steady-state along the flow path (Plummer et al. 1994). The system may be in a dynamic state, whereby water entering the aquifer today differs in chemical and isotopic composition from the recharge water that has evolved chemically to the currently observed final water. The steady-state assumption may also be violated by elements which can both dissolve and precipitate along a flow path, for example calcite. Therefore the system should be divided into sufficiently short paths in order to avoid potential changes of hydrochemical state . Mixing of different water types causes problems because mass-balance calculation with NETPATH is not capable of determining where along the overall flow path mixing actually takes place, and it assumes that all mixing occurs at the initial condition, followed by subsequent mineral-water reaction. Although the net mineral-water reactions and mixing proportions are unaffected, isotopic calculation with fractionation (Rayleigh distillation) is sensitive to the extent of reaction/mixing progress, if mixing and Rayleigh distillation calculations are considered in the same problem. Isotopic results are distorted if mixing occurs at some point in the reaction progress other than at the initial condition. However, such problems can also be reduced by adding intermediate waters along the flow path, i.e. shorter flow paths (Plummer et al. 1994). Overall, the mass-balance modelling is a useful tool to test hydrogeochemical conceptual models. However, mass-balance models are sensitive to uncertainties, which may be caused by inappropriate selections of initial waters and phases, flow paths and inaccurate chemical and isotopic data. Therefore it is important first to define strictly the initial conditions of mass-balance models, as follows:

• Phases and their composition in water-rock interaction constrained by elemental and isotopic changes along flowpaths (Section 7.2.1)

• Initial values of isotopic compositions and their fractionation in water-rock interaction (Section 7.2.2)

• Define initial hydrogeochemical conditions and flow paths for calculations in order to obey the steady-state requirement of mass-balance models (Sections 7.2.3 and 7.2.4)

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The results of mass-balance models may increase understanding of hydrogeochemical evolution at the site in many ways and calculations are also useful in evaluating data uncertainties:

• Mass transfer in water-rock interaction processes and their temporary and regional occurrence at the site (Sections 7.3.1 and 7.3.2)

• The significance of mixing and palaeohydrogeology in chemical changes (Section 7.3.3)

• Sensitivity calculations may expose data and conceptual uncertainties (Sections 7.3.1 ands 7.3.2), and the needs of further studies (Section 7.4)

Therefore, succesful modelling performance helps to understand pH and redox conditions, buffering processes, palaeohydrogeology, and temporal and regional variation in an integrated way, so they are also able to increase confidence in future predictions important for the assessment of repository safety.

7.2 Initial conditions

7.2.1 Phases and constraints used in modelling

Table 7-1 shows the most obvious mineral phases to be included in mass-balance modelling, their chemical behaviour in reactions based on the discussions above, and the chemical compositions used. The constraints generally used in modelling are the masses of Na, Ca, Mg, K, Al, Fe, Si, C, S, Cl, and redox state (electron balance). In modelled flow steps representing modern recharge conditions, calcite is assumed only to dissolve and no carbon isotope fractionation is expected (closed system, no CO2 escape). Carbon-13 could be used as an additional constraint and carbon isotopes are calculated as an isotope mass-balance problem. The formula CH2O is used only to denote a carbon valence of zero in the oxidation of dissolved organic carbon. Carbon dioxide is allowed to in-gas to the system to reflect the variable activity of vegetation in different parts of the aerobic recharge zone and to degas, for instance, in samples KR6/135/1-3, in which degassing is potentially due to long-term pumping and significant movements of groundwater. Methane gas is also used as a source in SO4 reduction modelling, but in addition, as a sink in carbonate reduction (hydrogen-based methanogenesis). Goethite is a potential sink in the recharge zone and, at depth, cation exchange and biotite are used to provide potential iron phases. Anion exchange acts as a source of SO4 in the recharge zone, i.e. potential competitive surface complexation (c.f. ch. 6.1). Rock-forming silicates dissolve incongruently during weathering by hydrolysis reactions and produce clay minerals and silica. The composition of plagioclase and biotite is based on mineralogical studies (e.g. Gehör et al. 1996, 1997, 2000, 2001b) of mica gneiss, which is the dominating rock type at the study site. Sodium sulphate balances large variations in sulphur contents in intermediate depth samples, which cannot be otherwise solved by SO4 reduction and pyrite precipitation.

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Table 7-1. Selected phases for mass-balance modelling and composition used in modelling. �Source� and �sink� describe allowed chemical behaviour of a phase along modelled flow paths.

Phase Source(+)/Sink(-) Composition

Calcite + / - CaCO3

Organic matter + CH2O

Carbon dioxide +/- CO2

Methane + / - CH4

Hydrogen + H2

Pyrite + / - FeS2

Goethite - FeOOH

Cation exchange + / - (Ca-Na2)X, (Mg-Na2)X, (K-Na)X, (Fe-Na2)X

Anion exchange + (2HCO3-SO4)X

Plagioclase An19 + Na0.79Ca0.19K0.02Al1.23Si2.78O8

Biotite Fe/(Fe+Mg)=0.57 + K0.95Na0.04Mg0.98Fe1.29Al1.85Si2.79O10(OH,O)2

Kaolinite + / - Al2Si2O5(OH)4

Quartz, chalcedony + / - SiO2

Sodium sulphate + Na2SO4

7.2.2 Isotopic calculations and initial values for carbon and sulphur isotopes

NETPATH considers two types of isotopic calculations: isotope mass-balance and Rayleigh calculations (Plummer et al. 1994). The isotope mass-balance calculation corresponds to the mass-balance equation constrained by conservation of a chemical element and electrons and is in general applicable for processes involving the constraining isotope as a source, such as mixing of waters, mineral dissolution or ingassing. When there is both a source and sink for a particular isotope in the reaction, the problem must be treated as a Rayleigh distillation problem, which takes into account isotope fractionation. After each mass-balance model is calculated, NETPATH computes the δ13C(DIC), 14C(DIC), δ34S(SO4

2-, S2-/HS-) and 87Sr/86Sr of the final water as a Rayleigh distillation problem using the equations of Wigley et al. (1978, 1979) for the modelled mass transfer. If isotopes are selected as constraints, the isotopic composition of the final water calculated by the Rayleigh model can be compared with the observed data to examine sensitivity between the fractionation of isotopic evolution and the mass-balance result. Carbon and sulphur isotopic calculations are applied to Olkiluoto data. Equilibrium fractionation factors, identified as Mook factors (Mook 1980), are used for the 13C (DIC) system whereas 13C-kinetic microbial fractionation, between DIC and CH4, is fixed at 75� according to Clark & Fritz (1997). The modelled carbon isotope composition at the end point of the reaction is a function of the initial and final total molalities of: - dissolved inorganic carbon - the initial value of carbon isotopes in the initial waters - the fractionation between precipitating calcite and DIC

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- the fractionation between outgassing CO2 and DIC - the fractionation between CH4 and DIC in methanogenesis - the mass transfer between carbon phases - the average isotopic composition of carbon sources Unfortunately deep groundwaters in the bedrock are mixtures of waters with completely different sources of dissolved carbonate. This impedes calculation of 14C residence times, because it may be impossible to estimate initial 14 C values of DIC or organic carbon in mixing waters. It is possible to test and vary expected initial values and explore the sensitivity of selections and plausible age (Pitkänen 1994), but it is a very demanding task and the results may still suffer from significant uncertainties. As with the carbonate phases, the fractionation factor for sulphur-bearing phases is defined according to the average isotopic composition of sulphur in the solution, both dissolved sulphate and sulphide. Therefore the �observed� values of δ34S in the final waters (Table 7-2) are calculated and deviate from measured values (App. 2), which are of δ34S in SO4. The isotope fractionation factor for the sulphur system applies to precipitation of sulphide phases from solution, and is specifically intended to describe kinetic, microbial fractionation of sulphur accompanying sulphate reduction and precipitation of iron sulphide phases. It is initially assumed that the sulphur isotopic composition of sulphide phases is that of the dissolved hydrogen sulphide in solution. As the sulphur isotopic composition of dissolved hydrogen sulphide is not available in the Olkiluoto data, the relationship introduced by Plummer et al. (1990) is used to estimate δ34S (H2S) based on the observed sulphur isotopic composition of dissolved sulphate and water temperature: δ34S(H2S) = δ34S(SO4) - 54 + 0.40t, (Eq. 7-1) where t is the water temperature in °C. Plummer et al. (1990) defined the equation from the sulphur isotopic composition of waters from two limestone aquifers in the US, probably resulting from kinetic fractionation during biologically mediated sulphate reduction. Clark & Fritz (1997) have emphasised the variation of bacterial fractionation depending, for example, on environmental conditions and bacterial communities, and suggested bacterial enrichment factors between 25 and 40�. In contrast, equation 7-1 produces factors about 50� at current temperatures in Olkiluoto groundwaters. Undoubtedly, sulphur isotopic calculations in future also need δ34S values of dissolved sulphide and mineral phases to make sensitivity calculations by varying fractionation factors. Isotope calculations require some assumptions to be made for phases because the isotopic data of carbon and sulphur species concern dissolved inorganic carbon, gas samples of dissolved hydrocarbons and CO2, determination of SO4 from water samples (App. 2), and fracture calcites (Blomqvist et al. 1992, Karhu 2000). No data are available from biogenic organic carbon. The δ13C value of organic carbon is assumed to be -25� PDB (both CO2 and CH2O in models), which is a general estimation (e.g. Plummer et al. 1990, 1994 and Pearson et al. 1991). Clark & Fritz (1997) suggest �23� for DIC derived from vegetation in a temperate climate. The influence of this difference

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is, however, marginal (observable in second or third decimal number) in mass-balance results compared to other uncertainties in calculations. The δ13C value used for dissolving calcite is -5� corresponding to the uppermost fracture calcites in borehole KR1, the average value of the latest calcites in fractures reported by Blomqvist et al. (1992) and Karhu (2000), and the calcite dissolution trend in Figure 5-20. Application of the 34S calculation requires an estimate of δ34S for the SO4 enrichment observed in dilute groundwaters from shallow depths at Olkiluoto. The δ34S value is assumed to be similar to the seawater value and +20 � is used (see ch 6.2.1). However, in the shallow groundwater system (i.e. modelled by PP8) +10 � is used, because SO4 in an anion exchange site may also be derived from oxidation of mineral sulphides.

7.2.3 Initial waters

Geochemical interpretation and modelling needs evaluation of the initial state of groundwater compositions as a starting point of mass-balance calculations, in order to calculate and understand the mixing and water-rock interaction history of brackish and saline groundwaters. It is impossible to apply the steady-state condition to the whole evolutionary path of hydrogeochemistry observed at Olkiluoto. If end-member compositions were used directly the results of calculations may be a sum of total processes, part of which may have occurred over millions of years (e.g. dilution of brine), part from sea bottom sediments (Litorina infiltration) and part during the postglacial period in bedrock. It is difficult to be aware of changes in initial conditions during the whole Weichselian glacial period and much less is known of the earlier history. Therefore, the initial point for mass-balance calculations so far cannot be fixed. The time since the Weichselian glaciation can be estimated for the mass-balance method (Pitkänen et al. 1999a), although the steady-state evolution has most likely been disturbed at least three times after the ice sheet retreated from the Olkiluoto island. The first stage was during deglaciation, some 10 000 years ago, when dilute melt-water with a high hydraulic gradient dominated the recharge. Secondly, about 7 000 years ago, infiltration from the Litorina Sea began. Finally, once the island had risen above sea level, meteoric recharge started some 2 500 years ago. Pitkänen et al. (1999a) showed that the groundwater data revealed certain extreme groundwater compositions, which enabled initial compositions linked to the above transient states to be interpreted. Conservative ions and isotopes (Cl, Br, δ18O, δ2H) are valuable for identifying possible initial state waters in the groundwater system. These species have also been suggested by principal component analysis (Fig. 5-1). Unfortunately, it is fairly uncertain to determine fully the initial groundwater compositions as these no longer remain, and the compositions in current extreme samples have been affected by reactions whose identity may not be recoverable. However, it has been possible to estimate conservative parameters of the initial groundwaters at the transient stages of the post-glacial period and also to estimate subglacial groundwater composition in the upper part of the bedrock.

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Current groundwater samples are mixtures of ancient end-member water compositions, which represent formation waters from certain extreme conditions such as Litorina seawater, glacial melt and original brine. These pure waters can no longer be found. Current data contain, however, extreme compositions that are derivatives of those end-member waters. These are defined as reference groundwaters. There have also been certain intermediate stage �end-members� (called initial groundwater members), which occurred in the upper part of the bedrock when hydrogeological conditions changed during deglaciation and postglacial times. For example, it is unlikely that marine infiltration from the Litorina Sea mixed with pure melt-water in the bedrock. This pre-Litorina groundwater (one initial groundwater) was probably a mixture of infiltrated glacial melt-water and previous subglacial groundwater (another initial water, modified during the glacial period) which, according to the salinity of the most extreme melt-water reference samples (KR3/243, TDS = 4 600 mg/l, Cl = 2 760 mg/l and no marine signature), had to be brackish to slightly saline in the upper part of the bedrock. In addition, the groundwater before Litorina infiltration probably had only a trace of SO4, if any, as indicated by Cl-dominated brackish and saline groundwater types seen today. Although SO4 does not conform completely to the requirement for conservatism, the salinity (or Cl content) of pre-Litorina and subglacial groundwater in the upper part of the Olkiluoto bedrock was estimated with the help of Cl, δ18O and SO4 data from Pitkänen et al. (1999a). Comparison between δ18O and Cl (Fig. 7-1) indicates four distinct reference groundwaters governing the groundwater compositions by mixing. The reference waters are meteoric (the meteoric end-member), Litorina reference, glacial reference and brine reference by which with the addition of Baltic seawater (basically diluted Litorina) the mixing traces of the other samples can be determined. Because Litorina and glacial reference waters cannot represent the original end-members, the initial groundwater mixed with them should be estimated in order to find out mixing portions of original end-member waters in reference groundwaters and further, using the whole data, to get information of hydrogeological conditions after deglaciation. The assessment of initial groundwater compositions follows mainly the evaluation presented in Pitkänen et al. (1999a), but it is important to repeat it here. The SO4-rich groundwater samples show a linear trend in Fig. 7-2, suggesting that the infiltrated Litorina seawater mixed with a SO4-poor initial groundwater member, which had a δ18O value of about -16�. The estimate assumes no significant mass transfer by reduction of sulphate in SO4-rich groundwater in bedrock. This seems probable based on the discussion of redox conditions (ch. 6.2). The proportional sink of reduced SO4 is considered minor, however, compared to the SO4 concentrations. Based on data from groundwater salinity, climate, temperature and subfossil shells, Kankainen (1986) proposed that δ18O and Cl values in the Litorina Sea may have reached -4.7� and 6500 mg/l, respectively. In this case, a linear estimate for the SO4 concentration of Litorina seawater would be about 900 mg/l (Table 3-1). The trends in Figs 7-1 and 7-2 also support this estimated composition for Litorina seawater.

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Figure 7-1. δ18O versus Cl content for groundwater samples at Olkiluoto.

Figure 7-2. δ18O versus SO4 concentrations in Olkiluoto groundwater samples. The arrow indicates the δ18O value of potential pre-Litorina groundwater, in which infiltration from the Litorina Sea mixed with.

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Using Cl and δ18O values of assumed Litorina composition and -16� for δ18O in pre-Litorina groundwater, the most SO4-rich groundwater sample has about 60% Litorina seawater. Hence the Cl content of pre-Litorina groundwater in the upper part of the bedrock has been about 1700-1800 mg/l. The estimate is sensitive to variation in the proportion of the Litorina end-member, e.g. ± 5% in the Litorina end-member causes a Cl concentration change of ± 600 mg/l in pre-Litorina water. On the other hand the uncertainty cannot be much greater, because if the portion were higher, e.g. 70%, SO4-rich groundwater samples would represent pure two end-member mixing between Litorina seawater and fresh melt-water. If it were much lower, the formation of SO4-rich brackish groundwater requires an additional significant source for SO4 that is not supported either by mineralogy or isotopic results of SO4. Pre-Litorina groundwater was a mixture of infiltrated dilute glacial melt-water and initial groundwater beneath the ice sheet (called subglacial) that had to be more saline and heavier in stable isotopes than the pre-Litorina initial groundwater in the upper part of the bedrock. This is supported by glacial reference groundwaters which show higher Cl and δ18O contents than the estimated pre-Litorina initial groundwater above. Estimation of Cl content in the subglacial groundwater needs an approximation for δ18O in glacial recharge (Cl was 0 mg/l for calculation purposes) and δ18O in subglacial groundwaters. Smellie & Frape (1997) noted a mismatch between reported values which varied from -30� to -18�, but, for instance, one glacier in Iceland shows δ18O values between -22� to-20�. Laaksoharju & Wallin (1997) have proposed -21� based on calcite surface deposits, fracture fillings (Tullborg & Larson 1984) and test calculations. Kankainen (1986) calculated from Hästholmen data that the δ18O value of glacial melt-water was between -26� to -18�. The data for saline waters from Olkiluoto suggest that the δ18O value in subglacial initial groundwater was about -11 to -10� and, assuming a δ18O value of infiltrated melt-water of -22 to -20�, roughly half the pre-Litorina groundwater would have been water derived from the Weichselian ice sheet. Hence the Cl concentration of subglacial groundwater in the upper part of the bedrock would have been about 3500 mg/l before melt-water intrusion. Recently, a sample from KR12/365 has found SO4-poor brackish groundwater with 5100 mg/l Cl and �10� δ18O. These values support the estimate for subglacial groundwater, although salinity is slightly higher but it likely increased with depth also at that time. Mass-balance prediction of this sample needs an estimate for subglacial initial groundwater. Glacial reference sample (KR3/241) and its ionic ratios, together with the above assessed Cl content, calcite saturation and electrical neutrality requirement have been used to determine an approximate composition for the subglacial initialwater: pH 8, Na 1570 mg/l, Ca 600 mg/l, Mg 31 mg/l, K 6 mg/l, DIC 4 mg/l.

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7.2.4 Flow paths

Groundwater samples selected for the calculations had a complete data set by the end of 2002. In addition, for the calculations, the preference was to use samples, which also had isotopic data; therefore, the number of shallow samples is small. Figure 7-3 shows the flow paths used in the post Weichselian hydrogeology at Olkiluoto. The flow paths were essentially based on hydrogeochemistry and apparent age of groundwater samples. Information was also supplemented by the hydrogeological model presented in chapter 3. The upper part of Figure 7-3 describes the environmental and hydrogeological changes at Olkiluoto since deglaciation with salinity estimates of different end-member/initial member waters (orange colour). Infiltration of these waters has created reference water compositions (red) (c.f. previous section) which are used as initial waters to calculate mass-balance models for the rest of the data. Brine and saline reference samples are also paralleled with initial waters, because the brine reference and its dilution to deep saline groundwaters (> 700m) may have occurred beyond the start of mass balance modelling time frame. However, the use of the brine reference sample as an initial water in calculations (KR12/741), may result in a model in which the time and duration of predicted mass transfer along a flow path is impossible to determine. Modelled groundwater samples (black; greens were calculated in Pitkänen et al. 1999a) are grouped in columns (Fig. 7-3) according to their relative age. The column on the left represents groundwaters, which are believed to have been modified during or before the Weichselian glaciation. The evolution of the other samples started after glacial melt infiltration. Flow paths in Fig. 7-3 represent mixing of initial waters and have no age connections; rather, they represent the relative change in average age of the final sample due to mixing in relation to initial waters. Infiltration from the fresh Ancylus Lake is considered improbable, because there was no driving force (hydraulic or salinity gradient) for groundwater recharge at that time

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Figure 7-3. Schematic representation of recharge conditions at Olkiluoto since glacial period and mixing hypothesis for mass-balance modelling. Further explanation see text.

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Single flow paths often had to be tested by trial and error to find plausible solutions for the mass-balance model; therefore, similar final waters are not systematically solved from the same initial waters. Serious contradictions were not observed between hydrogeochemical selections and hydraulic connections. However, certain generalisations had to be made, such as assuming that subhorizontal fracture zones R17, R19, R20 and R26 formed a common hydraulic unit as well as R9 and R10. In addition, reference waters (KR2/T4, KR4/T5 for Litorina seawater and KR3/243 for both glacial melt and subglacial groundwater in the upper part of the bedrock) are not always selected appropriately according to hydraulic connections, because only a few samples are available to represent the most extreme compositions of the external inputs. Previous samples from KR2 and KR10 are frequently used as initial waters because of their location in the central area of the site and the samples used are taken from the R20 fracture zone, which has been interpreted to cut most of the boreholes at intermediate depth (Vaittinen et al. 2001, Saksa et al. 2002). Overburden groundwater from PVP2 is mostly used as recharging groundwater for bicarbonate waters (Fig. 7-3) corresponding to former calculations (Pitkänen et al. 1999a). Recent samplings (Hatanpää 2002) from the overburden, support the earlier values used for carbonate content (analysed as DIC nowadays), which were based on acidity measurements. The chemistry and calculated models suggest a descending trend of groundwater flow from the highest hydraulic head and topographic area between boreholes KR1, KR2, KR12 and KR4, and deeper towards the SO4-rich groundwater layer in subhorizontal fracture zones at 100 to 300m depth. Mixing with former Litorina-derived brackish groundwater characterises this upper bicarbonate zone in the bedrock. The SO4-rich layer that is represented by the Litorina reference in Figure 7-3, divides the flow and seems to prevent mixing with the lower, SO4-poor brackish saline groundwater system. The SO4-rich zone seems to be at a shallower depth near the shoreline in boreholes KR3 and KR6 indicating more upward flow in the upper part of bedrock and possible discharge areas. Slightly elevated salinity in shallow well PP7 also supports discharge near the borehole KR3. Unusual compositions in KR13/112 and KR12/365, compared to other samples from similar depths, may reflect the isolated location of these sampling sections. The first one contains a clear glacial signal with a δ18O value reflecting a glacial melt component that is usually below the highest SO4 contents at Olkiluoto. The second sample seems to be undisturbed by deglaciation and postglacial waters; therefore the sample may represent a certain initial composition in the modelled groundwater system. Exceptionally high dissolved sulphide values have been observed at intermediate depths in boreholes KR2/T3 and T2, KR5/T2 and T1 and KR13/362 (App. 2). The fracture zone model does not indicate any common hydraulic connection for these sampling sections, but borehole logs show that tonalite-granodiorite sections (App.1) connect these sampling sections.

7.3 Results of mass-balance calculations

The mass transfer of selected chemical and isotopic parameters between initial and final waters in the single mass-balance model of Olkiluoto is partly solved by mixing of

109

initial waters and partly by dissolving and ingassing, or precipitating and outgassing, mineral and gas phases, or by using ion exchange (results are given in Table 7-2). The behaviour of Cl (also δ18O) is assumed to be conservative, i.e. no reactive phases are expected to produce or consume Cl along flow paths. Therefore the contents of Cl (or any conservatively manipulated constraint, see footnotes in Table 7-2) in final waters are solved purely by mixing, which adjusts the mixing proportions of initial waters in each mass-balance model. Calculated mixing proportions are presented on the left side of Table 7-2. The mass-balance predictions of active parameters are finally calculated with the code after mixing of initial waters, by constraining the dissolution or precipitation of plausible phases (the right side of Table 7-2) to produce the concentrations in the final water. Some of the initial waters (printed in green in Fig. 7-3) represent data, which are interpreted in a previous modelling report by Pitkänen et al. (1999a). The mass-balance results (Table 7-2) mostly agree well with interpretations of δ13C and δ34S data (similar observed and calculated values) and so support the evolutionary hypothesis used in the models. However, they have a varying degree of uncertainty. In particular, this includes uncertainties in the analytical data, which are linked to uncertainties in determining the flow paths and the mixing of waters. Mixing dominates the mass transfer of Na and Ca along flow paths in the calculations, except in dilute bicarbonate waters. The calculated mass transfers in water-rock interaction operate within analytical uncertainties. Calcium was finally omitted from constraints due to the mass-balance difficulties in some models with higher salinity (Table 7-2), yielding preliminary implausible reaction coefficients for some phases, in particular for plagioclase. This most probably resulted from the inflexibility of the NETPATH code to solve mass-balance equations, i.e. it calculates exact solutions. Nowadays, the PHREEQC code (Parkhurst & Appelo 1999) has capability for inverse geochemical calculations that take place within specified compositional uncertainty limits. On the other hand NETPATH has already been prepared for calculations of C and S isotopic fractionation which was considered important and affected code selection for this study.

110

Table 7-2. Summary of mass transfer results. The left side shows mixing fractions of initial waters for each final water. All mineral and gas mass transfers are in mmol/kg of water. A negative value indicates precipitation/outgassing/sodium uptake and a posive value dissolution/ingassing/sodium release. Prediction of C and S isotopic values and Ca concentrations are on the right.

Initial waters

Final Water

Wat

er ty

pe

Bal

tic se

awat

er

Met

eori

c w

ater

(PV

P1)

Met

eori

c w

ater

(PV

P2)

Mix

ed m

eteo

ric

& L

itori

na (K

R6/

58/1

)

Mix

ed M

eteo

ric

& L

itori

na (K

R12

/65)

Mix

ed M

eteo

ric

& L

itori

na (K

R2/

T7)

Lito

rina

Sea

([C

l- ] = 6

500

mg/

l)

Lito

rina

Alte

red

(KR

2/T4

)

Lito

rina

alte

red

(KR

4/T5

)

Lito

rina

Alte

red

(KR

6/12

5/1)

Mix

ed L

itori

na &

Gla

cial

(KR

2/T3

)

Gla

cial

mel

t ([C

l- ] = 0

mg/

l)

Gla

cial

Alte

red

(KR

3/24

3)

Gla

cial

Alte

red

(KR

10/3

24)

Subg

laci

al g

roun

dwat

er ([

Cl- ] =

350

0 m

g/l)

Salin

e (K

R1/

T3)

Salin

e (K

R9/

563)

Salin

e (K

R1/

S1)

Salin

e (K

R4/

860)

Cal

cite

Pyri

te

"CH

2O"

+CO

2

Ca-

Na 2

X -e

xcha

nge

Mg-

Na

-exc

hang

e

K-N

a -e

xcha

nge

Fe(I

I) -N

a

Ani

on -e

xcha

nge

Na 2

SO4

Goe

thite

Bio

tite

Plag

iocl

ase

Kao

linite

SiO

2

CH

4

H2

δ13C

Obs

erve

d �

(PD

B)

δ13C

Cal

cula

ted

�(P

DB

)

δ34S

Obs

erve

d �

(CD

T)

δ34S

Cal

cula

ted

�(C

DT

)

Ca

Obs

erve

d (N

a in

KR

12/7

41)

Ca

Cal

cula

ted

(Na

in K

R12

/741

)

PP8 0.001 0.999 0.51 1.64 -0.17 -0.21 0.01 0.17 0.08 1.57 -1.07 -2.5 -21.1 -21.1 8.2 9.4 PP7 0.873 0.127 1.77 3.71 2.03 0.00 0.52 -0.21 0.15 0.52 -0.48 -0.99 -17.9 -17.9 19.7 KR6/58 0.921 0.079 1.25 2.68 1.11 -0.41 0.82 -0.12 0.09 2.17 -1.45 -3.50 -16.4 -16.4 21.5 21.6 KR6/99/1 0.827 0.173 1.04 2.93 1.96 -0.28 1.12 -0.02 0.02 2.83 -1.78 -4.42 -15.9 -15.9 22.8 22.7 KR6/99/2 0.551 0.449 1.46 3.45 3.37 0.08 0.83 -0.06 0.06 4.90 -3.08 -7.64 -16.8 -16.8 24.4 24.8 KR6/99/3 0.458 0.542 1.04 2.31 0.62 0.53 -0.19 0.15 0.93 -0.73 -1.59 -15.9 25.7 KR9/149 0.832 0.168 2.08 3.58 3.91 0.17 0.82 -0.23 0.17 2.19 -1.52 -3.51 -15.9 -15.9 22.4 23.1 KR11/125 0.598 0.402 1.66 2.04 4.17 0.36 0.38 -0.23 0.16 4.75 -3.09 -7.46 -15.1 -15.1 24.3 25.5 KR12/65

Dilu

te-b

rack

ish-

carb

onat

e

0.749 0.251 1.68 -0.04 2.33 2.29 -0.76 0.22 0.02 1.63 -1.04 -2.52 -16.7 -16.7 28.2 27.7 KR7/282 0.448 0.552 -1.13 -0.02 0.08 -0.02 0.10 -0.02 0.00 1.72 -1.06 -2.64 -18.2 -17.9 24.7 24.4 KR6/135/1 0.426 0.574 0.22 -0.07 0.24 0.03 -0.64 -0.07 -0.07 0.00 0.64 -0.40 -0.99 -16.3 -16.2 22.6 25.8 KR6/135/2 0.280 0.720 0.18 -0.00 -0.03 -0.53 -1.94 -0.04 -0.01 0.07 2.91 -1.79 -4.52 -15.4 24.2 24 KR6/135/3 0.173 0.827 0.31 -0.00 -0.40 -0.48 -1.51 -0.07 -0.00 -0.10 0.02 -0.01 -0.03 -16.6 24.2 KR13/214 0.375 0.625 -0.66 0.88 0.09 -0.00 0.53 1.72 -1.06 -2.70 0.01 -16.4 -16.7 23.7 26.2 11.8 12.1 KR6/125* 0.680 0.218 0.101 -1.10 -0.32 1.31 - 5.22 1.89 -0.33 0.11 -0.07 0.00 -16.6 -17.5 24.3 25.6 23.1 22.3 KR13/112** 0.408 0.398 0.194 -1.05 -0.45 0.43 0.15 0.00 5.09 -3.13 -7.91 -0.00 -17.1 -17.5 26.2 23.1 KR13/362

Bra

ckis

h-

Chl

orid

e

0.403 0.597 -0.23 -0.00 -0.30 -0.24 0.02 0.74 2.15 -1.32 -3.30 0.37 -40.2 25.4

KR3/443 0.953 0.047 0.03 1.97 0.70 -0.09 0.00 0.09 -0.00 -0.00 -0.06 0.18 -4.3 0.4 5.3 6.2

KR12/365 Bra

ckis

h-ch

lori

de

0.858 0.142 -0.22 -0.04 0.06 1.07 0.05 -0.04 0.14 3.80 -2.34 -5.81 0.07 -16.2 -13.8

KR9/470 0.556 0.444 0.03 -3.18 0.83 0.04 0.00 0.03 0.09 -0.05 -0.00 -16.1 -16.2 24.2 22.6 49.3 56.1 KR6/525 0.250 0.750 -0.15 -0.01 8.14 2.00 0.27 0.06 0.03 -0.02 0 0.03 -25.3 39.4 63.6 75 KR10/498 0.117 0.883 -0.00 3.73 0.73 0.39 0.01 12.46 -7.67 - -0.03 0.14 6.1 KR12/741 0.554 0.446 0.50 1.91 -0.03 0.02 -0.00 0.01 -0.52 2.09 68.4 379 359 KR12/741

Salin

e

0.554 0.446 -0.5 1.91 -0.03 0.02 -0.00 0.01 -0.02 0.08 54.3 379 359

* mixing fixed by Cl and δ18O** mixing fixed by Cl and S

111

The omission of Ca in certain mass-balance models (Table 7-2) predicts Ca concentrations in final waters, which in fact reduces the electrical imbalance compared to the original water analyses. For example, electrical imbalance in the sample KR6/525 is about �4% but it is balanced if the calculated Ca content (75 mmol/l is based on model that is not constrained with Ca) instead of the measured value is used. In many of Ca-omitted models, Ca mass transfer in the reactions represents only a few percent of the total concentration in the final water. Therefore chemical data were changed in part of the lower saline final groundwaters (Table 7-3) with high electrical imbalance (cf. App. 2) and with greatly implausible preliminary mass-balance models. The data manipulation was made on the concentration of the ion with the largest deviation from general trends. Table 7-3. Changes made in the chemical data of groundwater samples for mass-balance models

Sample Parameter Measured value Used value KR6/99/1 Ca 1000 mg/l 800 mg/l KR6/125 Ca 929 mg/l 813 mg/l KR6/135/1 Na 1050 mg/l 1220 mg/l KR6/135/3 Ca

SO4 663 mg/l 392 mg/l

720 mg/l 460 mg/l

KR13/362 Ca 1190 mg/l 1050 mg/l KR2/T4 δ34S(SO4) - 26.75 � CDT* KR2/T3 δ34S(SO4) - 30 � CDT KR10/324 δ34S(SO4) - 30 � CDT * average of two other Litorina reference samples (KR4/T5, KR8/302)

Important δ34S(SO4) isotopic data were partly missing in initial waters which were needed in predicting isotopic compositions of final groundwater samples. Values used in calculations were also estimated according to the hydrochemical data of Olkiluoto. Care must be taken in including mass transfer by Ca-Na ion exchange and plagioclase dissolution (with kaolinite and chalcedony precipitation) particularly in brackish and saline groundwater, and these reactions must be considered as part of other, more reliable reactions. Mass transfer of other major elements (C, S, Mg, K) by reactions, is not as sensitive to mixing calculations because: 1) their concentrations are much smaller, 2) they do not vary as much between different waters as do Na and Ca concentrations, 3) their chemical character is relatively active and 4) mass transfer by reactions is of the same magnitude as the analysed concentrations. However, mixing of a high portion of marine-derived water may cause mass-balance problems due to its anomalous high SO4, Mg and K concentrations compared with waters from other potential sources at Olkiluoto. The amount of precipitated calcite was used in evaluating the reliability of plagioclase dissolution in a previous study by Pitkänen et al. (1999a). Calcite precipitation (Eq. 7-2) should essentially show the level of silicate dissolution (e.g., plagioclase in Eq. 7-3) which is mainly dependent on protons (from carbonic acid) released by calcite precipitation in a nearly neutral and constant pH region.

112

Ca2+ + HCO3- ↔ CaCO3 + H+ (Eq. 7-2)

Na0.8Ca0.2Al1.2Si2.8O8 + 1.2 H+ + 0.6 H2O → Plagioclase

0.8 Na+ + 0.2Ca2+ + 0.6Al2Si2O5(OH)4 + 1.6 SiO2

(Eq. 7-3)

Calcite and organic carbon input to DIC are the dominant processes along the flow paths to give HCO3-rich groundwaters (Table 7-2). Predicted carbonate derivation from organic carbon (several mmol) is much greater than is indicated by DIC contents in low pH (mainly CO2), shallow groundwaters in which they are clearly less than 1 mmol/l. Fracture calcite is the main pH buffer in the hydrogeological system and concern has been raised whether it will be lost and dissolved by CO2-rich recharge waters during the construction and operational phase of the repository. Luukkonen et al. (2003) used 0.7 mmol/l in their calculations examining the effects of CO2 dominated recharge water (pH 5.4) on fracture mineral buffers. However, the chemical data (App. 2 and Figs 6-4, 6-6) show that high DIC values quickly develop at shallow depths, in bedrock and overburden pH values increase above 7 indicating that HCO3 dominates due to carbonic acid dissociation in water. Therefore carbonic acid (dissolved CO2) seems to be lost mainly before meteoric water infiltrates in the bedrock. Solubility calculations also show that HCO3 species dominates and CO2 concentrations are generally below 0.5 mmol/l in shallow groundwaters, either from overburden or bedrock; thus the value used in calculations by Luukkonen et al. (2003) was sufficient. The extensive dissociation of carbonic acid to bicarbonate, in overburden layer, further suggests that a significant part of the calculated mass transfers (Table 7-2) for HCO3-rich groundwaters also occurs near the recharge area. The last sample taken during long-term pumping of KR6/99 does not follow the same systematics as previous samples in the sequence. In the calculations, another Litorina-derived reference water was used to obtain a plausible mass-balance model suggesting that there had been an inflow of a different water type in the borehole. According to current data this could not have been seawater, however. Calcite precipitates at greater depths in brackish and saline groundwaters. However, the mass-balance results suggest slight dissolution of the samples with simultaneous methanogenesis and samples from the long-term pumping test of KR6/135. Methane formation reduces carbonate, thus calcite is the source for DIC. This, as well as the chemical situation during the pumping test, should be further examined by forward calculations (see Plummer et al. 1983) with thermodynamic code, but mass-balance models suggest calcite dissolution with a little CO2 outgassing as the most plausible prediction. The outgassing may be possible due to mixing of different water bodies from different depths during pumping. Although the model determining Litorina seawater infiltration (KR6/125) into bedrock lacks confidence from the steady-state assumption, its predictions are congruent with previous calculations (Pitkänen et al. 1999a), i.e. the anomalous high Na and Mg sorption releasing Ca and Fe. This exchange was interpreted to occur in sea bottom sediments during infiltration.

113

Contradictory results are generally observed in proton balance when comparing the mass transfer of calcite precipitation and plagioclase dissolution and this contradiction is considered to reflect uncertainties in models and data. These problems can also reflect on Ca-Na ion exchange. However, Na is logically released near the surface in HCO3-rich groundwaters, where fresh water flushes the earlier seawater type aquifer. Basically, Ca for Na exchange should dominate in a system that is diluting Ca-rich brine, and a generally higher charged ion is preferred on the exchange site if total solute concentrations decrease in the case of exchange with heterovalent ions (Appelo & Postma 1993). Finally it is concluded that the progress of reactions is plausible, but the obvious uncertainties decrease the reliability of quantitative reactions of phases, in particular the mass transfer of silicates and ion exchange, which depends strongly on salinity variation between initial and final waters. The amount of silicate dissolution in deep groundwater conditions could be approached by forward modelling. Potential dissolution of plagioclase may be evaluated by utilising reaction coefficients calculated by mass-balance models for mixing, calcite formation and the redox system. Adjusting the pH on an observed (or thermodynamically corrected, see ch. 5.6) level may describe the mass transfer of plagioclase, kaolinite and SiO2 in the groundwater system more correctly than mass-balance models can do.

7.3.1 Isotopic evolution and mass transfer in redox processes

The evolution of 13C(DIC) and 34Stot relates closely to pH and redox related processes in groundwater. The δ13C is calculated as an isotopic mass-balance problem (see ch. 7.2.2) in dilute and brackish HCO3-rich groundwaters, i.e. δ13C is one of the constraints adjusting calcite dissolution and organic carbon oxidation, and so calculated δ13C results equal with measured values (Table 7-2). The result of the last sample of KR6/99 is the only value predicted by the mass balance model. All the HCO3-rich groundwaters show reduced δ34S values relative to seawater, but only in one of them is precipitation of pyrite needed to increase δ34S value. In others, mixing of older SO4-rich groundwater is sufficient to increase δ34S to the observed level. Sulphate is also reduced in these models. The formation of minor dissolved sulphide observed in the groundwaters does not, however, affect the δ34Stot value used in mass-balance calculations and presented in Table 7-2, although it does increase δ34S(SO4) value. Brackish, SO4-rich groundwaters require anaerobic oxidation of organic carbon or methane and reduction of sulphate to change their marine-originated 34S and 13C signatures. The required amount of organic carbon remains plausible, the highest value (0.24 mmol/kg) corresponds to the upper limit presented by Pettersson et al. (1990) for deep groundwaters in a granitic environment. The content used in the model of KR6/125 is much higher but organic carbon respiration in this case is assumed to occur in sea bottom sediments that are rich in organic matter. Sulphate reduction and pyrite precipitation are similar to the mass-balance models of other Litorina reference samples (Fig. 7-2) calculated in a previous study (Pitkänen et al. 1999a). Those models indicated pyrite precipitation of 0.38 mmol/kg of water. However, mixing decreases the reliability

114

of the results of Litorina reference samples, because SO4 reduction may have mainly occurred in sea-bottom sediments before mixing and dilution of SO4 concentration started in the bedrock, (probably the opposite of the calculated situation). Therefore, less mass transfer between sulphur phases is sufficient to increase δ34S in the calculated model than occured during the infiltration of Litorina seawater. Sulphate reduction is also simulated with the assumption of no mixing and the measured δ34S value of KR6/125 is the final value after infiltration into the bedrock (Table 7-4). The infiltration model represents the same constraints and phases as the corresponding mixing model in Table 7-2, except that the SO4 content of the final water is increased to such a level (815 mg/l) that no Na2SO4 mass transfer is needed. The results indicate that SO4 reduction during the infiltration of Litorina seawater has been between 0.64 to 0.74 mmol/kg of water (61 � 71 mg/l) in this case. The difference is small, but isotopic predictions are more reliable in the infiltration model. The upper limit is less than predicted in the previous study (1.15 mmol/kg or 110 mg/l) by Pitkänen et al. (1999a), but the δ34S value in KR6/125 (24.3�) is also less than the other Litorina reference samples (about 26.8�). The common results indicate about 1 mmol/l or 100 mg/l reduction of SO4 during infiltration of seawater at Olkiluoto yielded the pyrite and calcite in bottom sediments. Table 7-4. Test calculations of Litorina Seawater infiltration using estimated Litorina composition as as initial water and Litorina reference sample KR6/125 as final water. The mixing model is as used in Table 7-2; the infiltration model assumes similar constraints and phases but no mixing. Signs are as in Table 7-2.

Phase Mixing (Table 7-3) Infiltration Pyrite (FeS2) -0.32 -0.37 CH2O 1.31 1.49 Fe(II)-Na exchange -0.33 -0.37 Calcite -1.10 -1.51 δ34S obs./calc. (� CDT) 24.3/25.6 24.3/24.3 δ13C obs./calc. (� PDB) -16.6/-17.5 -16.6/-16.1

Methane is oxidised in the lower part of the SO4-rich zone to reduce SO4. The last sample of this group contains the highest sulphide content (12.4 mg/l) observed in the groundwater data of Olkiluoto. Pyrite precipitation is minor, but the oxidised methane decreases δ13C to a very low level. Unfortunately, both carbon and sulphur isotopic measurements of this sample have been unsuccessful. The sulphur isotopic value stays constant in the mass-balance model, but the δ34S(SO4) would be clearly higher. The sample locates very near KR2/T3 (used as initial water) which also has a slightly elevated sulphide value (1.3 mg/l) but particularly low δ13C value (-36.3�) that supports the model prediction. The whole flow path (mass-balance models via KR2/T3 to KR13/362) represents the highest hydrogen sulphide production interpreted at Olkiluoto. Previous mass-balance calculations suggested that sulphide formation was along the first step to KR2/T3 (Pitkänen et al. 1999a) 0.84 mmol/kg (27 mg/l) from which 0.8 mmol may have precipitated as pyrite. The flow path to KR13/362 increases total sulphide formation to 1.18 mmol/kg of water (39 mg/l) and this corresponds to about 110 mg/l reduced SO4.

115

Methane is also an obvious sink of reduced carbonate in deeper parts and mass-balance models often predict its formation below SO4-rich groundwaters (brackish-Cl and saline groundwaters in Table 7-2). This interpretation diverges from previous modelling results (Pitkänen et al. 1999a) in which minor SO4 reduction was considered plausible in deep saline groundwaters. Current hydrogeochemical and microbial data indicate a methanogenic environment and occasional sulphur species in saline groundwaters is contamination rather than a sign of sulphidic conditions. The mass-balance results show clearly the significant effect methanogenesis has on δ13C(DIC) value, which is emphasised because of the low DIC contents in deep groundwaters. Even a small volume formed (0.045 mmol corresponds to 1 ml CH4) may elevate 13C dramatically. Unfortunately, carbon isotopic measurements were often unsuccesful and calculated and measured values cannot be compared. However, the mass-balance model of KR3/443 supports the calculated level of methanogenesis as well as the high δ13C measured value from KR4/860/1 (16.8�). Young carbonate contamination was observed in the sample KR3/433 (ch. 5.5.2), which has probably reduced the observed δ13C(DIC) value, too. On the other hand, calculated δ13C values for KR12/741 are anomalously high, possibly beyond all observed limits. The mass-balance results suggest that the amount of microbial methane in deep groundwaters is very small (at a level of ml/l) compared to the observed amounts (hundreds of ml/l). However, the models in Table 7-5 show some uncertainty in the level of formation due to lack of isotopic information. The δ13C(DIC) value does not increase drastically anymore although microbial methanogenesis would be much more abundant. The increase is reduced due to δ13C released in the carbonate pool by dissolution of calcite, the amount of which is similar to CH4 formation (no calcite source problem, cf. Table 2-1). The significance of methanogenesis on CH4 gas cannot be estimated from δ13C(CH4) values either because they also increase due to high δ13C(DIC) value. This may partly explain the observed trend of hydrocarbons in Figure 5-23a, because if δ13C(DIC) is clearly positive, δ13C(CH4) would not be as low as theoretically expected for bacterial reduction of carbonate. For example, δ13C(DIC) is 16.8 � in KR4/860/1, then δ13C in CH4 gas formed first will be about �60 �. Possibly the most critical value for abundant CH4 formation is hydrogen consumption that is larger in the second model of KR10/498 (650 ml/l) than ever observed at Olkiluoto. Finally, in addition to obtaining reliable δ13C(DIC) data, it is important in evaluating bacterial methanogenesis at Olkiluoto to obtain more information from fracture calcites in deep groundwater conditions to see if they show any dissolution features.

116

Table 7-5. Predictions of δ13C values in DIC and CH4 with varying formation of methane. Mixing models as in Table 7-3. δ13C in dissolving calcite is �15 �, and corresponds to values in deep bedrock (Blomqvist et al. 1992, Blyth et al. 1999).

Sample CH4, mmol/kg of water δ13C(CH4), �PDB δ13C(DIC), �PDB KR10/498 0.03 -80.2 6.1 KR10/498 7.4 -6.5 81.9 KR12/741 0.02 -48.4 54.3 KR12/741 0.52 -16.9 68.4

7.3.2 Mixing and palaeohydrogeological implications

Table 7-6 shows the back-calculation of the mixing proportions presented in Table 7-3 to assumed original end-members at Olkiluoto. The end-members are based on current observations of meteoric recharge, seawater, and saline water or interpreted as Litorina seawater and glacial melt (Table 3-1), and the average composition of groundwater in the upper part of the bedrock during glaciation, i.e. subglacial groundwater. The table contains estimated δ18O values of end-member waters according to which δ18O has been calculated for groundwater samples to test the modelling results against the measured values. The value for meteoric water is an estimate from groundwater observation tubes and shallow wells, all with high tritium waters (App. 7). The value for saline water is an estimate according to KR1/S1 and KR4/861. Analytical results of stable isotopes of water are currently under consideration due to fluctuations observed, for example, in shallow groundwater data. Therefore the meteoric parameter of δ18O is decreased to �11 � from �10.6 � used in the previous report. The end-member parameters will be evaluated once again after the comparison study is completed.

117

Table 7-6. Calculated mixing proportions (%) of interpreted original end-members of modelled groundwater samples at Olkiluoto. Calculated δ18O compositions of groundwater samples based on estimated values of end-members are also shown with measured ones.

Sample

Meteoric

Baltic seawater

Litorina seawater

Glacial melt

Subglacial groundwater

Saline

δ18O calculated

δ18O observed

PP8 99.9 0.1 -11.0 -11.1 PP7 87 7.3 2.3 3 -10.7 -10.7 KR9/149 83.2 9.7 3.1 4 -10.7 KR6/58 92.1 4.6 1.4 1.9 -10.9 -10.5 KR12/65 75.1 14.5 4.6 6 -10.6 -11 KR11/125 59.8 23.2 7.3 9.6 -10.23 KR6/98.1 82.8 10 3.1 4.1 -10.7 -10.5 KR6/98.2 55.1 26 8.2 10.7 -10.2 KR6/98.3 45.8 31 11.1 12.1 -10.2

KR6/125 68 24 8 -9.3 -9.3 KR6/135.1 39.2 41 14.4 5.4 -10.0 -9.9 KR6/135.2 25.8 50.2 17.7 6.3 -9.8 KR6/135.3 15.9 57 20.1 6.9 -9.6 KR13/112 39.2 0 24.2 11.6 25 -10.6 -12.7 KR7/282 43 0 33 10.4 13.6 -10.0 -9.9 KR13/214 28.2 41.6 13.1 17.2 -9.8 -9.8 KR13/362 13.6 18.4 59.1 8.9 -11.7 -11.2

KR3/443 0 20.9 78.5 0.6 -12.9 -12.3 KR12/365 0.1 1.2 90.1 8.6 -10.6 -10

KR9/470 0.6 15.3 57.2 26.7 -11.9 -11.4 KR6/525 1 10.8 39.6 48.6 -11.2 -10.7 KR10/498 1.2 9.4 34.5 54.9 -11.0 -11.2 KR12/741 100 -9.5 -11.6

The consistency seems quite good between calculated and measured δ18O- in the samples, supporting the modelling concept and estimation of the composition of end-member waters. Generally measured values are slightly higher than calculated, but in two samples, calculated values deviate instead strongly upwards. Sample KR13/112 has a higher portion of dilute melt water than the model shows and the use of dilute sample KR2/T7 seems not to be a good selection in the model. The measured value of δ18O in saline sample KR12/741 deviates from the saline sample suggesting uncertainties in the analytical result. Calculated δ18O values are slightly lower for groundwaters below the SO4-rich zone than measured values, reflecting too-high a melt water proportion. The reason may be also an unidentified diluting water component affecting the saline groundwater system. Figure 7-4 shows the vertical distribution of the proportional occurrence of different end-member waters in groundwater at Olkiluoto. A significant flow of meteoric water is limited to the upper 150m according to mass-balance calculations. The two major deep occurrences of meteoric water at 200m and 300m (samples KR13/214 and KR7/282) have shown significant decreasing trends in EC-TDS measurements (App. 1) suggesting possible drawdown of meteoric groundwater due to pumping at these sections. The

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monitored sections in KR6 instead behave in an opposite manner and the meteoric water proportion decreases to half during long-term pumping. Litorina-originated water dominates the mixture between 100m and 300m whereas subglacial brackish groundwater below that changes gradually to a predomination of saline groundwater, undisturbed by melt or post-glacial waters around 500m.The current mixing proportion of glacial melt water stays at around the 20% level from 100m down to 400m, decreasing at greater depths to below 10% in the upper layer of saline groundwater. The last estimate may comprise some overestimation as suggested by δ18O results and saline groundwater dilution is probably much more complex than has been able to interpret.

Figure 7-4. Calculated proportions (%) of end-member waters in Olkiluoto groundwater samples. Trend lines show depth distributions of end-members in the Olkiluoto area.

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7.4 Discussion

The results of mass-balance calculations support the interpreted concept of geochemical evolution and baseline conditions at Olkiluoto, and all results are supported by isotopic calculations. The results evaluate the mass transfer of calcite and organic carbon, and indicate SO4 reduction, CH4 formation and ion-exchange. In addition, calculations indicate that the main mass transfer occurs during infiltration (at the level of a few mmols/l), in overburden at the present time and in sea bottom sediments in the past. Compositional variations in bedrock groundwaters are mainly caused by end-member waters and mixing between them and mass transfer between phases is very small (tenths of mmols/l). However, in the transition zone from SO4-rich to CH4-rich groundwaters, the mixing of groundwaters of different redox states initiates the reaction between SO4 and CH4. Although the reactions involve only low concentrations of reactants they are significant in buffering pH and redox conditions in groundwater, as solubility calculations have shown. However, uncertainties in calculations are evident and are caused by analytical errors, insufficient data and selections of flow paths. These uncertainties are reflected in mass transfer between phases such as silicate dissolution, ion-exchange and CH4 formation, or in end-member composition. Although thorough work has already been done to build up a reliable hydrochemical dataset for Olkiluoto, the data should be checked once again to create a final baseline dataset for evaluation and modelling purposes in future. The uncertainties can be decreased significantly by systematic revision of hydrochemical data. The data should be examined as a whole by classifying and sorting it in a similar way as used in figures in this report and this will reveal deviations in the data. Elemental and isotopic results should be evaluated utilising charge imbalances and parallel samples and measurements, taking into account information from samplings and the hydrogeology of sampling sections. New groundwater samples will also be obtained before the disturbance due to ONKALO construction begins. These supplement the data particularly if sufficient isotopic analyses will be available. Completing the baseline dataset may improve the evaluation of flow paths in mass-balance models. Forward modeling should be used in complementing mass-balance models and in evaluating silicate reactions and cation exchange. If Sr isotopic data are available from mineral phases they should be added to the mass-balance calculations, thus achieving independent mass transfer information of mineral sources.

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8. CONCLUSIONS

The interpretation of hydrogeochemical evolution at Olkiluoto is based on the chemical, isotopic and microbiological data of water samples (App. 2) and minerals, and geochemical modelling in conjunction with a knowledge of the hydrogeology and other geological features of the site. The interpretation of baseline conditions at Olkiluoto and evolutionary processes beyond them, supplements the geochemical concept of the site presented by Pitkänen et al. (1999a). The interpretation is based on the data presented in the previous report and subsequently collected until the end of 2002. Actual deviations are not observed compared to the previous interpretations (see e.g. PCA results in ch. 5.4.1) but, clearly, more detailed hydrogeochemical information has been obtained, particularly from shallow depths (the first 10 m) and the deep saline groundwater zone (below 500m), and from dissolved gases. Repeated samplings in the bedrock during the years show only minor changes, thus strengthening the hydrogeochemical concept model of the site. No dramatic changes were observed during the long-term pumping experiment at borehole KR6, in which the, Litorina-derived groundwater signature only became clearer. The interpretation of hydrogeochemical data indicates that mixing of end-member waters controls the wide salinity variation in groundwaters at Olkiluoto. Changes in past climate and geological environment have left distinct chemical and isotopic signatures, and caused great variability in the hydrochemical data. Recognised end-member water types are current, fresh meteoric recharge (infiltration during the last 3000 years), brackish Baltic seawater, brackish Litorina-stage seawater (infiltration 7500 � 4000 years ago), fresh melt water from the Weichselian glaciation and extremely old brine (at least millions of years) that shows a hydrothermal origin. In addition, groundwater data indicate that one further groundwater type, subglacial brackish groundwater with no seawater indications was located in the upper part of the bedrock during glaciation. It is probably diluted from brine over time. Salinity increases evenly with depth and no significant deviations from the general trend have been observed (Fig. 5-3). Mixing calculations indicate meteoric water domination down to 150m depth (Fig. 7-4), following with a Litorina seawater derived groundwater type to 300m depth. Only a few observations of a young meteoric water component occur deeper, found at 200-300m. However, TDS logging suggests that they result from the drawndown effect caused by sample pumping. Below 300m, groundwaters consist of glacial or older water end-members. The subglacial type dominates to 500m depth and the saline type below that. The glacial melt component is calculated to be about 20% from 100m to 450m depth, decreasing with depth and disappearing below 600m. Infiltration and mixing of end-members seems to yield a stratified hydrochemical system at least in the water-conducting fracture system (Fig 6-1, 8-1). The first 100-150m is filled with HCO3-rich fresh or brackish groundwater (meteoric origin) changing sharply to a SO4-rich groundwater that is slightly more saline (marine origin). Sulphate contents decrease with depth and brackish groundwater loses its marine signatures at the depth of 300m. Brackish groundwater changes relatively sharply into saline groundwater between 400 and 500m. Near the shoreline, the interface of HCO3-rich and

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SO4-rich groundwaters seems to climb up to shallower depth (KR3 in Fig. 8-1) and the upper surface of the saline groundwater is deeper. Water-rock interaction, such as carbon and sulphur cycling and silicate reactions, buffer the pH and redox conditions (Figure 9-1) and stabilise groundwater chemistry. However, the extensive difference between the compositions of groundwater types also causes significant deviations in water-rock interaction processes, particularly in redox chemistry. Hydrogeochemical interpretations and chemical and isotopic calculations indicate that pH seems to be dominantly controlled by thermodynamic equilibrium with calcite in fractures. Calcite is the most common fracture-filling mineral in the bedrock and there are indications that it may also occur in the overburden layer. Modelling results suggest that the pH range varies from 7.5 in HCO3- and SO4-rich layers, increasing to 8 in the lower part of brackish groundwaters, decreasing again to 7.5 in saline groundwater. These are slightly less than measured values from groundwater samples.

Figure 8-1. Illustrated west-east cross-section of hydrogeochemical and hydrogeological conditions in the bedrock of Olkiluoto based on interpretation of hydrogeochemistry. Changes in colour describe alteration in water type. Blue arrows represent flow directions. Rounded rectangles contain the main sources with estimated δ13C data, and sinks affecting pH and redox conditions. Rectangles show measured/calculated δ13C(DIC) of selected groundwater samples. Generalised fracture zones (coded by R) are combined on the basis of bedrock models by Vaittinen et al. (2001) and Saksa et al. (2002). Boreholes KR2 and KR4 are combined due to their similar hydrogeochemical character.

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Oxic redox conditions, prevailing in recharging groundwater, change abrubtly to sulphidic conditions close to the surface, within a short residence time. The anaerobic reduction of sulphate with the oxidation of organic carbon seems to be the most important process governing the redox conditions in dilute and brackish HCO3-rich and SO4-rich groundwaters. Methanogenesis may show increasing importance in saline groundwater and occasionally in brackish groundwater, below the SO4-rich layer. Haveman et al. (1998, 2000) have observed microbes that are fundamental for these redox processes to occur in the groundwater at Olkiluoto. Results of equilibrium calculations indicate that the redox level decreases with depth from �200 mV to �250 mV in sulphidic redox conditions in brackish groundwaters, and further, almost to �300 mV in methanic systems, in the lower part of the brackish layer and in saline groundwaters. Mass-balance calculations suggest that the main mass transfer in reactions occurs at the millimole scale during infiltration either in overburden or very shallow bedrock currently or in sea-bottom sediments during the Litorina Sea stage. Deeper in the bedrock mass, transfer may be smaller, only some tenths of mmol per kg of water. Organic carbon oxidation and calcite dissolution are the main reactions in the meteoric recharge zone, whereas bacterial SO4 reduction with oxidation of organic carbon, causing pyrite and calcite precipitation, have dominated during the infiltration of Litorina Seawater. In addition, ion exchange and silicate dissolution are probable, but uncertainties in the data limit mass transfer calculations in these processes and their contribution to pH buffering. Mass transfer in redox reactions seems to increase in the transition zone from SO4-rich groundwaters to SO4-poor containing substantial CH4 content at 300 to 400m depth. The prediction of total SO4 reduction may be about 110 mg/l corresponding to sulphide formation of 40 mg/l, from which a significant part has precipitated as pyrite. The highest observed dissolved sulphide content is 12.4 mg/l. Methane concetration is several hundreds of ml/l in deep saline groundwater at Olkiluoto. Bacterial methane formation is evident deep in the bedrock, but insufficient isotopic data of DIC and hydrocarbons impede detailed evaluation of its magnitude and effect on the carbonate system. The calculations suggest a level of few ml/l for bacterial CH4 production. The data for hydrocarbons indicate that the principal source of CH4 and other hydrocarbons is thermal processes. However, it is unclear whether they are formed by thermal decomposition of organic matter or by hydrothermal reactions between carbonate or graphite with H2. The baseline hydrogeochemistry at Olkiluoto seems to provide the conditions for long-term geochemical stability, which is a requirement for safe, long-term disposal of nuclear waste. Current chemical conditions such as slightly alkaline pH, anoxic redox conditions, low alkalinity and dissolved sulphide, tolerable Cl, and saturation of calcite at about 500m depth are generally favourable for long canister lifetimes, bentonite buffer stability and insolubility of radionuclides and fracture calcites. Pricipal component analysis support the conclusion that the essential features of the groundwater chemical trends have been captured prior to disturbances caused by the excavation of ONKALO.The data and interpretations, however, contain uncertainties, which require further study. Therefore, hydrochemical data collected before ONKALO construction

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should be checked to create a final data set for evaluation and answer open questions in hydrogeochemistry, to confirm the baseline conditions and evolutionary processes behind them. This updated hydrogeochemical model will perform as a reference in evaluating changes resulting from the ONKALO construction and finally in evaluating hydrogeochemical conditions after the closure of repository.

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APPENDICES:

Appendix 1. Lithology, fracture frequency, hydraulic conductivity, EC (electric conductivity) results from fracture water calculated as TDS, interpreted fracture zones and groundwater sampling sections. Lithology and and fracture frequency interpretations are based on studies by Saksa et al. 2001. The EC measurements are made by Pöllänen and Rouhiainen (1996a,b, 2000, 2001a,b and 2002a,b) and Rouhiainen (2000). It is notable, that the scales in TDS logs ranges between the different boreholes.

Appendix 2. Hydregeochemical data collected since the previous study (Pitkänen et al. 1999a). Digits shown in italics and bold = concentration below determination limit, digits with bold type = uncertain result, underlined digits = the underlined result is the most representative of the alternatives, NR = no results due to analytical or sampling problems, empty = not analysed.

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KR1 Appendix 1 (1/13)

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151

Sampling Sec.up. Sec.low. T O2 EhPt ECf Density TDSSample Type date m m

oC mg/l mV mS/m pHf g/ml mg/lPVP1/2 Overburden 071101 6.5 7.8 6 0.9982 54PVP1/3 Overburden 071101PVP3A/1 Overburden 201101 6.4 48.6 7.4 0.9964 368PVP3A/2 Overburden 201101PVP3B/1 Overburden 201101 6.9 40 7.3 0.9961 310PVP3B/2 Overburden 201101PVP4A/1 Overburden 191101 7.0 70.9 7.4 0.9989 528PVP4A/2 Overburden 191101PVP4B/1 Overburden 191101 6.7 70.6 7.4 0.9993 526PVP4B/2 Overburden 191101PVP5A/1 Overburden 201101 6.3 58.8 7.6 0.9962 514PVP5A/2 Overburden 201101PVP8A/1 Overburden 201101 7.1 35.1 7.3 0.9975 293PVP8A/2 Overburden 201101PVP9A/1 Overburden 151101 23.1 7.3 0.9988 194PVP9A/2 Overburden 151101PVP9B/1 Overburden 071101 27.9 7.3 0.9994 245PVP9B/2 Overburden 071101PVP10A/1 Overburden 071101 74 7.5 0.9989 542PVP10A/2 Overburden 071101PVP10B/1 Overburden 071101 68.8 7.5 0.9997 528PVP10B/2 Overburden 071101PP2/1 Dil./Brack. HCO3 121101 6.0 77 7.7 0.9998 665PP2/2 Dil./Brack. HCO3 121101PP5/1 Dil./Brack. HCO3 081101 9 5.7 0.9985 61PP5/2 Dil./Brack. HCO3 081101PP7/1 Dil./Brack. HCO3 131101 6.4 254 8 0.9997 1592PP7/2 Dil./Brack. HCO3 131101PP8/1 Dil./Brack. HCO3 131101 5.6 31.3 7.1 0.9988 269PP8/2 Dil./Brack. HCO3 131101PP9/1 Dil./Brack. HCO3 141101 6.5 35.2 7.2 0.9991 305PP9/2 Dil./Brack. HCO3 141101PR1/2 Dil./Brack. HCO3 061101 6.5 5.8 0.998 49PR1/3 Dil./Brack. HCO3 061101PR2/2 Dil./Brack. HCO3 081101 14.4 6.2 0.9986 102PR2/3 Dil./Brack. HCO3 081101PR4/2 Dil./Brack. HCO3 071101 11.3 6.1 0.9982 89PR4/3 Dil./Brack. HCO3 071101KR3/443/1 Brackish Cl 271198 443.0 448.0 26.0 0.005 173 856 8.3 1.0016 4750KR4/860/1 Saline 200798 860.0 865.0 12.0 0.004 36 10070 7.5 1.0500 67959KR4/860/2 Saline 140502 860.0 866.0 8480 7.0 1.0510 73421KR6/58/1 Dil./Brack. HCO3 280301 58.0 60.0 226 1027KR6/99/1 Dil./Brack. HCO3 260301 98.5 100.5 441 1982KR6/99/2 Dil./Brack. HCO3 180701 98.5 100.5 13.0 0.000 -240 733 7.7 4094KR6/99/3 Brackish SO4 070802 98.5 100.5 788 7.5 4623KR6/125/1 Brackish SO4 090102 125.0 130.0 1360 7.6 8371KR6/135/1 Brackish SO4 240401 135.0 137.0 933 7.7 5088KR6/135/2 Brackish SO4 030701 135.0 137.0 10.0 0.000 -250 1084 7.8 6376KR6/135/3 Brackish SO4 220302 135.0 137.0 1220 7.7 1.0040 6944KR6/135/4 Brackish SO4 081002 135.0 137.0 1223 7.0 1.0030KR6/525/1 Saline 080101 525.0 528.0 14.0 0.000 -120 3410 8.7 1.0130 19446KR6/525/2 Saline 080101 525.0 528.0 14.0 0.000 -120 3410 8.7 1.0130 19497KR7/282/1 Brackish SO4 101298 282.0 289.0 13.0 0.003 -28 760 7.6 1.0016 4761KR9/149/1 Dil./Brack. HCO3 090501 149.0 150.0 13.0 -0.008 -330 325 8.0 0.9996 1979KR9/149/2 Dil./Brack. HCO3 090501 149.0 150.0 13.0 -0.008 -330 310 2041KR9/470/1 Saline 250299 470.0 475.0 10.0 0.002 68 2230 7.8 1.0077 12957KR10/498/1 Saline 270599 498.0 503.0 11.0 -0.009 -152 3550 8.0 1.0147 22099KR11/125/1 Dil./Brack. HCO3 070601 125.0 126.5 7.0 0.000 -290 650 7.9 1.0004 3744KR11/125/2 Dil./Brack. HCO3 070601 125.0 126.5 7.0 -0.001 -290 650 7.9 3695KR11/952/1 Saline 190100 952.4 957.4 8950 7.0 65200KR12/65/1 Dil./Brack. HCO3 120301 65.0 67.0 7.0 0.000 120 410 7.9 1.0000 2460KR12/65/2 Dil./Brack. HCO3 120301 65.0 67.0 7.0 0.000 120 410 7.9 2080KR12/365/1 Brackish Cl 080101 365.0 368.0 12.0 0.000 -390 1420 8.1 1.0044 8389KR12/365/2 Brackish Cl 080101 365.0 368.0 12.0 0.000 -390 1420 8.1 8552KR12/664/1 Saline 140502 664.0 666.0 5730 7.6 1.0270 36839KR12/737/1 Saline 200302 736.5 740.0 6940 7.9 1.0330 46293KR12/741/1 Saline 131100 741.0 751.0 4.0 0.000 50 6650 8.3 1.0360 49484KR12/741/2 Saline 131100 741.0 751.0 4.0 0.000 50 6650 8.3 50651KR13/112/1 Brackish SO4 090102 112.0 116.0 750 8.0 1.0010 4480KR13/112/2 Brackish SO4 090102 112.0 116.0 4080KR13/214/1 Brackish SO4 271101 214.0 220.0 1060 7.8 1.0030 6168KR13/214/2 Brackish SO4 271101 214.0 220.0 5932KR13/362/1 Brackish SO4 130901 362.0 365.0 1430 7.8 0.9890 8402KR13/362/2 Brackish SO4 130901 362.0 365.0 8056

Appendix 2 (1/9)

152

Alk(m) Alktot HCO3 AlkCO3 AlkCO3 Acid. CO2 (field DIC DOCSample Type CB. % meq/l meq/l mg/l mg/l meq/l meq/l analyses) mg/l mgC/l mgC/l

PVP1/2 Overburden 39.28 0.22 0.22 13.4 0.00 0.00 0.4 17.6 9.4 44.0PVP1/3 Overburden 7.1 56.1PVP3A/1 Overburden 6.43 3.03 3.03 185.0 0.00 0.00 0.18 7.9 32.0 16.0PVP3A/2 Overburden 41.2 16.8PVP3B/1 Overburden 4.26 2.59 2.59 158.0 0.00 0.00 0.196 8.6 30.0 17.0PVP3B/2 Overburden 38.2 20.2PVP4A/1 Overburden 0.01 4.30 4.30 262.0 0.00 0.00 0.32 14.1 44.0 3.2PVP4A/2 Overburden 57.7 1.0PVP4B/1 Overburden 0.02 4.27 4.27 260.0 0.00 0.00 0.2 8.8 43.0 3.3PVP4B/2 Overburden 58.6 1.0PVP5A/1 Overburden 2.22 5.40 5.40 329.0 0.00 0.00 0.26 11.4 50.0 15.0PVP5A/2 Overburden 71.7 17.5PVP8A/1 Overburden 2.50 2.84 2.84 173.0 0.00 0.00 0.226 9.9 33.0 13.0PVP8A/2 Overburden 38.3 13.1PVP9A/1 Overburden 16.25 1.72 1.72 105.0 0.00 0.00 0.2 8.8 28.0 35.0PVP9A/2 Overburden 30.6 43.9PVP9B/1 Overburden 11.86 2.41 2.41 147.0 0.00 0.00 0.27 11.9 29.0 40.0PVP9B/2 Overburden 34.7 55.6PVP10A/1 Overburden 1.52 4.32 4.32 264.0 0.00 0.00 0.21 9.2 42.0 7.4PVP10A/2 Overburden 59.0 6.6PVP10B/1 Overburden -1.32 4.40 4.40 268.0 0.00 0.00 0.32 14.1 49.0 8.1PVP10B/2 Overburden 61.7 11.3PP2/1 Dil./Brack. HCO3 -7.41 5.84 5.84 356.0 0.00 0.00 0.24 10.6 51 4.5PP2/2 Dil./Brack. HCO3 42.9 4.0PP5/1 Dil./Brack. HCO3 31.40 0.26 0.26 16.0 0.00 0.00 1.1 48.4 25 27.0PP5/2 Dil./Brack. HCO3 8.8 26.7PP7/1 Dil./Brack. HCO3 -5.38 5.88 5.88 359.0 0.00 0.00 0.15 6.6 47 6.8PP7/2 Dil./Brack. HCO3 62.8 7.2PP8/1 Dil./Brack. HCO3 15.93 2.28 2.28 139.0 0.00 0.00 0.29 12.8 22 14.0PP8/2 Dil./Brack. HCO3 28.8 16.4PP9/1 Dil./Brack. HCO3 1.98 2.92 2.92 178.0 0.00 0.00 0.26 11.4 31 6.0PP9/2 Dil./Brack. HCO3 34.4 5.5PR1/2 Dil./Brack. HCO3 8.39 0.18 0.18 11.0 0.00 0.00 0.71 31.2 20.0 8.7PR1/3 Dil./Brack. HCO3 7.1 12.8PR2/2 Dil./Brack. HCO3 15.95 0.80 0.80 48.8 0.00 0.00 1.2 52.8 27.0 25.0PR2/3 Dil./Brack. HCO3 18.7 30.6PR4/2 Dil./Brack. HCO3 20.03 0.64 0.64 39.0 0.00 0.00 1.11 48.9 17.0 29.0PR4/3 Dil./Brack. HCO3 6.0 29.5KR3/443/1 Brackish Cl -0.58 0.24 0.24 13.1 0.78 0.03 0.00 0.0 2.80 10.8KR4/860/1 Saline -2.12 0.15 0.15 9.2 0.00 0.00 0.12 5.3 1.30 15.8KR4/860/2 Saline 1.92 0.13 0.13 8.2 3.00 0.10 0.19 8.4 1.60 13.2KR6/58/1 Dil./Brack. HCO3 2.17 170.0 32.70 5.4KR6/99/1 Dil./Brack. HCO3 -8.64 140.0 28.00 4.5KR6/99/2 Dil./Brack. HCO3 0.84 3.29 3.29 201.0 0.00 0.00 0.00 0.0 45.90 5.9KR6/99/3 Brackish SO4 0.00 3.09 3.09 188.5 0.19 5.8KR6/125/1 Brackish SO4 1.69 1.48 1.48 0.13 2.3KR6/135/1 Brackish SO4 -4.23 2.43 2.43 148.5 0.00 0.00 0.00 0.0 30.00 8.6KR6/135/2 Brackish SO4 1.00 2.12 2.12 129.0 0.00 0.00 0.00 0.0 24.50 3.5KR6/135/3 Brackish SO4 -1.10 1.91 1.91 116.0 0.14 19.70 3.1KR6/135/4 Brackish SO4 1.81 1.81 110.0 0.10KR6/525/1 Saline -3.93 0.12 2.8 2.22 0.07 0.02 0.9 0.74 6.5KR6/525/2 Saline -6.95 0.00 0.00 0.0 0.67 5.8KR7/282/1 Brackish SO4 -1.73 1.56 1.56 95.2 0.00 0.00 0.07 3.1 18.90 5.2KR9/149/1 Dil./Brack. HCO3 -1.74 5.66 5.66 345.0 0.00 0.00 0.12 5.3 55.00 6.5KR9/149/2 Dil./Brack. HCO3 3.53 5.66 5.66 345.0 0.00 0.00 0.12 5.3 67.20 7.0KR9/470/1 Saline -1.13 0.29 0.29 17.7 0.00 0.00 0.04 1.8 2.30 4.8KR10/498/1 Saline 1.36 0.11 0.11 6.7 0.00 0.00 0.02 0.9 1.30 14.3KR11/125/1 Dil./Brack. HCO3 1.31 4.26 4.26 260.0 0.00 0.00 0.10 4.4 43.70 6.8KR11/125/2 Dil./Brack. HCO3 -1.43 4.26 4.26 260.0 0.00 0.00 0.10 4.4 51.70 8.0KR11/952/1 Saline -1.61KR12/65/1 Dil./Brack. HCO3 -1.96 5.00 305.0 1.20 0.04 0.14 6.2 50.00 14.0KR12/65/2 Dil./Brack. HCO3 6.14 62.80 13.9KR12/365/1 Brackish Cl 1.02 0.15 0.15 9.4 1.20 0.04 0.02 0.7 1.80 3.4KR12/365/2 Brackish Cl -8.57 1.61 5.0KR12/664/1 Saline -5.73 0.10 0.10 6.3 3.00 0.10 0.07 3.1 1.5 7.8KR12/737/1 Saline -2.00 0.17 0.17 10.3 3.00 0.10 0.05 2.2 1.50 10.6KR12/741/1 Saline 0.57 0.13 0.13 7.2 1.20 0.04 0.00 0.0 0.50 4.5KR12/741/2 Saline -7.60KR13/112/1 Brackish SO4 1.47 1.44 1.44 81.8 3.00 0.10 0.15 6.6 16.20 3.5KR13/112/2 Brackish SO4 -11.97 1.44 1.44 81.8 3.00 0.10 0.15 6.6 16.20KR13/214/1 Brackish SO4 -7.66 1.56 1.56 89.1 3.00 0.10 0.09 4.0 17.90 3.4KR13/214/2 Brackish SO4 -2.02 1.56 1.56 89.1 3.00 0.10 13.30 1.4KR13/362/1 Brackish SO4 2.16 0.78 0.78 41.5 3.00 0.10 0.05 2.2 6.50 2.4KR13/362/2 Brackish SO4 -1.63 0.78 0.78 41.5 3.00 0.10 2.60 1.3

Appendix 2 (2/9)

153

Stot S2-tot SO4 Ptot PO4 NO2 NO3 Cl F BrSample Type mg/l mg/l mg/l mg/l mg/l mg/l mg/l mg/l mg/l mg/l

PVP1/2 Overburden 3.2 6 0.041 0.20 2.7 0.11 0.04PVP1/3 Overburden 10 0.02 4.3 0.4 0.50PVP3A/1 Overburden 9.6 25 0.071 0.20 35.0 0.32 0.20PVP3A/2 Overburden 27 0.02 35.0 0.34PVP3B/1 Overburden 8.7 24 0.057 0.20 26.0 0.35 0.26PVP3B/2 Overburden 25 0.02 21.9 0.33PVP4A/1 Overburden 18.0 48 0.071 0.20 60.0 0.5 0.13PVP4A/2 Overburden 50 0.02 48.6 0.5 0.50PVP4B/1 Overburden 18.0 50 0.037 0.20 58.0 0.48 0.11PVP4B/2 Overburden 52 0.02 47.7 0.40 0.50PVP5A/1 Overburden 12.0 31 0.17 0.20 4.3 0.79 0.05PVP5A/2 Overburden 35 0.02 3.6 0.9 0.50PVP8A/1 Overburden 9.4 27 0.037 0.20 3.7 0.38 0.07PVP8A/2 Overburden 29 0.02 3.1 0.42 0.50PVP9A/1 Overburden 4.9 10 0.022 0.20 5.5 0.18 0.23PVP9A/2 Overburden 12 0.02 6.0 0.4 0.50PVP9B/1 Overburden 4.6 8 0.1 0.20 5.2 0.17 0.16PVP9B/2 Overburden 13 0.02 7.2 0.40 0.50PVP10A/1 Overburden 15.0 45 0.35 0.20 64.0 0.59 0.28PVP10A/2 Overburden 46 0.02 66.0 0.5 0.50PVP10B/1 Overburden 16.0 46 0.014 0.20 58.0 0.61 0.27PVP10B/2 Overburden 47 0.11 47.9 0.60 0.50PP2/1 Dil./Brack. HCO3 17.00 0.05 48 0.011 0.2 83.0 0.56 0.22PP2/2 Dil./Brack. HCO3 50 0.02 84.5 0.5 0.50PP5/1 Dil./Brack. HCO3 3.6 0.13 7 0.11 0.20 2.1 0.10 0.12PP5/2 Dil./Brack. HCO3 12 0.02 3.3 0.50 0.50PP7/1 Dil./Brack. HCO3 54.0 0.28 140 0.006 0.20 590 0.10 1.80PP7/2 Dil./Brack. HCO3 141 574.8 0.30 1.40PP8/1 Dil./Brack. HCO3 8.1 0.18 22 0.06 0.20 8.1 0.21 0.15PP8/2 Dil./Brack. HCO3 27 0.02 8.6 0.3 0.50PP9/1 Dil./Brack. HCO3 11.0 0.01 31 0.009 0.20 5.9 0.23 0.09PP9/2 Dil./Brack. HCO3 32 0.10 5.1 0.3 0.50PR1/2 Dil./Brack. HCO3 4.0 0.01 14 0.025 0.20 1.5 0.10 0.04PR1/3 Dil./Brack. HCO3 15 0.04 1.1 0.30 0.50PR2/2 Dil./Brack. HCO3 2.8 0.04 6 0.55 0.20 3.9 0.10 0.08PR2/3 Dil./Brack. HCO3 7 0.02 4.7 0.30 0.50PR4/2 Dil./Brack. HCO3 5.6 0.03 11 0.057 0.20 1.4 0.10 0.18PR4/3 Dil./Brack. HCO3 19 0.09 2.0 0.30 0.50KR3/443/1 Brackish Cl 3.8 0.01 11.7 0.01 0.01 2880 2.30 15.00KR4/860/1 Saline 2.0 0.01 1.2 0.10 0.10 43000 1.70 350.00KR4/860/2 Saline 1.3 0.01 8.4 0.08 0.05 45200 1.50 300.00KR6/58/1 Dil./Brack. HCO3 0.02 126 0.50 0.25 0.25 369 0.61 1.49KR6/99/1 Dil./Brack. HCO3 0.01 203 0.50 0.25 0.25 1000 0.38 2.45KR6/99/2 Dil./Brack. HCO3 0.01 317 0.02 0.02 0.02 2065 0.36 7.70KR6/99/3 Brackish SO4 0.01 330 0.03 0.04 0.01 0.02 2440 0.40 7.20KR6/125/1 Brackish SO4 0.13 541 0.02 0.03 0.02 0.02 4680 0.60 12.00KR6/135/1 Brackish SO4 0.01 352 0.01 0.02 0.02 2850 0.34 4.10KR6/135/2 Brackish SO4 0.03 432 0.02 0.02 0.02 3480 0.32 11.50KR6/135/3 Brackish SO4 0.01 392 0.03 0.03 0.02 0.02 3940 0.40 13.00KR6/135/4 Brackish SO4 0.01 4000 0.50KR6/525/1 Saline 2.4 0.26 5.3 0.04 0.03 0.02 12300 1.50 88.00KR6/525/2 Saline 4.7 0.50 0.25 0.25 12580 1.70 90.00KR7/282/1 Brackish SO4 113.0 0.16 320.0 0.01 0.02 2630 0.41 9.20KR9/149/1 Dil./Brack. HCO3 54.0 0.01 172 0.15 0.37 0.05 780 1.10 2.60KR9/149/2 Dil./Brack. HCO3 50.0 0.01 170 0.11 0.12 0.02 780 1.10 3.20KR9/470/1 Saline 6.0 0.01 17.1 0.01 0.02 8030 1.40 56.00KR10/498/1 Saline 0.4 0.01 1.4 0.02 0.01 13500 1.86 95.00KR11/125/1 Dil./Brack. HCO3 86.0 0.02 250 0.01 0.03 0.05 1850 0.68 7.00KR11/125/2 Dil./Brack. HCO3 84.0 0.02 270 0.02 0.02 0.02 1850 0.68 6.40KR11/952/1 Saline 41300KR12/65/1 Dil./Brack. HCO3 50.0 0.01 150 0.03 0.02 0.02 1160 0.06 3.80KR12/65/2 Dil./Brack. HCO3 48.0 0.10 177 0.50 0.25 0.25 1075 0.49 3.41KR12/365/1 Brackish Cl 2.4 0.01 6.3 0.02 0.03 0.02 5100 1.80 36.00KR12/365/2 Brackish Cl 5.0 0.50 0.25 0.25 5590 1.90 36.70KR12/664/1 Saline 1.3 0.01 5.4 0.03 0.05 23800 1.40 150.00KR12/737/1 Saline 3.1 0.05 5.5 0.03 0.20 29200 1.30 200.00KR12/741/1 Saline 2.5 0.04 5.1 0.25 0.02 0.35 30600 1.20 204.00KR12/741/2 Saline 0.02 3.5 0.50 0.25 0.25 33400 1.30KR13/112/1 Brackish SO4 78.0 0.02 241.0 0.02 0.02 2450 0.90 7.40KR13/112/2 Brackish SO4 78.0 0.02 241.0 0.02 0.02 2450 0.90 7.00KR13/214/1 Brackish SO4 140.0 0.41 431.0 0.02 0.02 3580 0.60 11.00KR13/214/2 Brackish SO4 150.0 0.41 420.0 0.01 0.02 3300 0.60KR13/362/1 Brackish SO4 50.0 12.40 130.0 0.02 0.05 4940 1.30 23.00KR13/362/2 Brackish SO4 38.0 12.30 120.0 0.02 0.06 4870 1.30 24.00

Appendix 2 (3/9)

154

SiO2 Ntot NH4 Fetot (ferrozine FetotFAAS Fe2+ Fe3+Al Na K

Sample Type mg/l mg/l mg/l , field) mg/l mg/l mg/l mg/l mg/l mg/l mg/lPVP1/2 Overburden 15.0 0.78 0.05 2.20 2.20 3.8 1.1PVP1/3 OverburdenPVP3A/1 Overburden 16.0 0.47 0.17 3.50 0.14 45 7.7PVP3A/2 OverburdenPVP3B/1 Overburden 17.0 0.48 0.15 3.60 0.25 33 5.8PVP3B/2 OverburdenPVP4A/1 Overburden 19.0 0.15 0.16 5.20 0.00 27 5.8PVP4A/2 OverburdenPVP4B/1 Overburden 21.0 0.17 0.15 4.50 0.01 25 5.9PVP4B/2 OverburdenPVP5A/1 Overburden 24.0 0.40 0.18 9.00 0.01 19 8.3PVP5A/2 OverburdenPVP8A/1 Overburden 16.0 0.31 0.05 8.90 0.16 5.8 5.7PVP8A/2 OverburdenPVP9A/1 Overburden 17.0 0.94 0.09 4.50 0.82 6.1 4.7PVP9A/2 OverburdenPVP9B/1 Overburden 17.0 0.88 0.05 5.40 0.76 6.7 5.3PVP9B/2 OverburdenPVP10A/1 Overburden 16.0 0.38 0.11 2.10 0.01 80 11.0PVP10A/2 OverburdenPVP10B/1 Overburden 16.0 0.43 0.05 0.05 0.02 63 10.0PVP10B/2 OverburdenPP2/1 Dil./Brack. HCO3 18.0 0.25 0.23 4.80 3.80 3.70 0.10 0.00 53 6.3PP2/2 Dil./Brack. HCO3PP5/1 Dil./Brack. HCO3 17.0 0.40 0.05 4.00 4.50 3.50 1.00 1.30 3.4 2.2PP5/2 Dil./Brack. HCO3PP7/1 Dil./Brack. HCO3 10.0 0.45 0.31 1.70 0.28 0.25 0.03 0.00 350 10.0PP7/2 Dil./Brack. HCO3PP8/1 Dil./Brack. HCO3 15.0 0.29 0.08 8.60 6.90 7.40 -0.50 0.12 16 5.6PP8/2 Dil./Brack. HCO3PP9/1 Dil./Brack. HCO3 13.0 0.37 0.07 4.90 1.90 1.40 0.50 0.03 22 6.0PP9/2 Dil./Brack. HCO3PR1/2 Dil./Brack. HCO3 9.8 0.27 0.05 1.90 2.10 1.80 0.30 0.40 2.1 1.4PR1/3 Dil./Brack. HCO3PR2/2 Dil./Brack. HCO3 13.0 0.42 0.05 2.60 2.60 2.40 0.20 0.89 16 2.2PR2/3 Dil./Brack. HCO3PR4/2 Dil./Brack. HCO3 11.0 0.75 0.05 3.00 2.00 2.50 -0.50 1.20 4.7 2.7PR4/3 Dil./Brack. HCO3KR3/443/1 Brackish Cl 6.8 1.40 0.05 0.01 0.01 -0.01 0.02 1590 7.6KR4/860/1 Saline 4.2 0.67 0.20 2.50 2.40 0.10 0.06 9200 18.0KR4/860/2 Saline 5.6 0.17 3.33 2.90 3.33 -0.43 0.01 9540 28.0KR6/58/1 Dil./Brack. HCO3 11.5 0.38 1.32 1.20 1.24 -0.04 225 7.9KR6/99/1 Dil./Brack. HCO3 10.0 0.25 1.36 1.25 1.12 0.13 460 7.0KR6/99/2 Dil./Brack. HCO3 11.0 0.54 1.43 1.50 0.91 0.59 1100 14.0KR6/99/3 Brackish SO4 12.0 0.52 0.49 1.32 1.40 1.35 0.05 1060 18.0KR6/125/1 Brackish SO4 12.0 0.06 0.17 0.32 0.43 0.35 0.08 1990 18.0KR6/135/1 Brackish SO4 12.4 0.29 0.73 0.64 0.43 0.21 1050 17.0KR6/135/2 Brackish SO4 10.7 0.45 0.85 1.10 0.78 0.32 1440 19.0KR6/135/3 Brackish SO4 12.0 0.32 0.26 0.51 0.57 0.53 0.04 1590 19.0KR6/135/4 Brackish SO4 0.51 0.50 0.01KR6/525/1 Saline 7.0 0.57 0.12 0.02 0.07 0.01 0.06 0.03 4500 7.2KR6/525/2 Saline 7.1 0.02 0.04 4700 7.2KR7/282/1 Brackish SO4 12.0 0.29 0.08 0.12 0.15 -0.03 0.03 1200 8.4KR9/149/1 Dil./Brack. HCO3 13.5 0.57 0.40 0.10 0.10 0.10 0.00 0.01 550 12.0KR9/149/2 Dil./Brack. HCO3 10.0 0.40 0.10 0.13 0.10 0.03 0.01 590 14.0KR9/470/1 Saline 15.0 0.40 0.06 0.15 0.16 -0.01 0.01 2810 9.6KR10/498/1 Saline 8.0 0.83 0.03 0.11 0.11 0.00 0.01 4830 14.0KR11/125/1 Dil./Brack. HCO3 13.9 0.28 0.11 0.04 0.03 0.03 0.00 0.002 1050 3.2KR11/125/2 Dil./Brack. HCO3 12.5 0.11 0.04 0.03 0.03 0.00 0.009 1000 4.1KR11/952/1 Saline 10100KR12/65/1 Dil./Brack. HCO3 12.3 0.29 0.10 0.17 0.18 0.16 0.02 0.10 580 7.2KR12/65/2 Dil./Brack. HCO3 12.0 0.12 0.13 600 7.0KR12/365/1 Brackish Cl 5.4 0.26 0.06 0.16 0.19 0.08 0.11 0.01 2100 9.3KR12/365/2 Brackish Cl 3.0 0.15 0.02 1900 7.3KR12/664/1 Saline 5.8 0.02 0.05 0.05 0.05 0.00 0.01 6330 19.0KR12/737/1 Saline 6.6 0.02 0.07 0.07 0.06 0.00 0.01 7120 19.0KR12/741/1 Saline 5.8 0.33 0.15 0.02 0.20 0.02 0.18 0.00 8300 19.0KR12/741/2 Saline 0.06 8100 17.0KR13/112/1 Brackish SO4 8.4 0.30 0.03 0.11 0.14 0.11 0.03 0.01 1240 7.4KR13/112/2 Brackish SO4 7.6 0.10 0.12 0.00 0.00 990 5.3KR13/214/1 Brackish SO4 7.2 0.13 0.07 0.40 0.40 0.40 0.00 0.02 1400 14.0KR13/214/2 Brackish SO4 6.7 0.14 0.44 0.36 0.08 0.06 1500 16.0KR13/362/1 Brackish SO4 9.7 0.11 0.31 0.05 0.02 0.03 -0.01 0.01 1980 8.5KR13/362/2 Brackish SO4 9.3 0.07 0.31 0.05 0.02 0.03 -0.02 0.02 2000 8.2

Appendix 2 (4/9)

155

Ca Mg Mn Rb Sr B Ni Cu Se SnSample Type mg/l mg/l mg/l mg/l mg/l mg/l mg/l mg/l mg/l mg/l

PVP1/2 Overburden 6 3.1 0.03 0.00PVP1/3 OverburdenPVP3A/1 Overburden 40 12.0 0.73 0.00PVP3A/2 OverburdenPVP3B/1 Overburden 35 9.7 0.70 0.01PVP3B/2 OverburdenPVP4A/1 Overburden 90 14.0 1.20 0.00PVP4A/2 OverburdenPVP4B/1 Overburden 89 15.0 1.30 0.00PVP4B/2 OverburdenPVP5A/1 Overburden 76 19.0 2.20 0.00PVP5A/2 OverburdenPVP8A/1 Overburden 51 8.5 1.20 0.00PVP8A/2 OverburdenPVP9A/1 Overburden 36 7.4 0.59 0.00PVP9A/2 OverburdenPVP9B/1 Overburden 44 9.0 0.77 0.00PVP9B/2 OverburdenPVP10A/1 Overburden 43 17.0 0.36 0.00PVP10A/2 OverburdenPVP10B/1 Overburden 47 18.0 0.37 0.00PVP10B/2 OverburdenPP2/1 Dil./Brack. HCO3 77 17.0 0.82 0.00PP2/2 Dil./Brack. HCO3PP5/1 Dil./Brack. HCO3 3.6 3.0 0.09 0.00PP5/2 Dil./Brack. HCO3PP7/1 Dil./Brack. HCO3 100 28.0 0.26 0.01PP7/2 Dil./Brack. HCO3PP8/1 Dil./Brack. HCO3 44 8.7 0.77 0.00PP8/2 Dil./Brack. HCO3PP9/1 Dil./Brack. HCO3 31 12.0 0.19 0.00PP9/2 Dil./Brack. HCO3PR1/2 Dil./Brack. HCO3 4.1 2.0 0.17 0.00PR1/3 Dil./Brack. HCO3PR2/2 Dil./Brack. HCO3 5.3 2.6 0.10 0.00PR2/3 Dil./Brack. HCO3PR4/2 Dil./Brack. HCO3 10 4.2 0.29 0.00PR4/3 Dil./Brack. HCO3KR3/443/1 Brackish Cl 210 9.6 0.04 0.01 1.8 1.50KR4/860/1 Saline 15100 108.0 2.30 0.12 160.1 0.50KR4/860/2 Saline 18000 130.0 2.30 190.0KR6/58/1 Dil./Brack. HCO3 86 26.0 0.5 0.30KR6/99/1 Dil./Brack. HCO3 118 36.0 1.8 0.40KR6/99/2 Dil./Brack. HCO3 300 72.0 3.6 0.50KR6/99/3 Brackish SO4 430 130.0 4.7KR6/125/1 Brackish SO4 929 180.0 8.1KR6/135/1 Brackish SO4 517 130.0 5.1 0.50KR6/135/2 Brackish SO4 661 184.0 6.3 0.70KR6/135/3 Brackish SO4 663 190.0 7.3 0.66KR6/135/4 Brackish SO4

KR6/525/1 Saline 2500 7.8 0.07 0.19 22.0 0.70 0.01 0.13 0.13 0.03KR6/525/2 Saline 2100 5.6 0.12KR7/282/1 Brackish SO4 390 92.0 0.32 0.01 2.4 0.80KR9/149/1 Dil./Brack. HCO3 74 28.0KR9/149/2 Dil./Brack. HCO3 90 37.0KR9/470/1 Saline 1950 29.0 0.36 0.03 19.0 0.80KR10/498/1 Saline 3570 38.0 0.40 0.03 32.1 0.96KR11/125/1 Dil./Brack. HCO3 240 69.0KR11/125/2 Dil./Brack. HCO3 220 71.0KR11/952/1 Saline 13800KR12/65/1 Dil./Brack. HCO3 180 60.0 0.20 0.00 0.01 0.03 0.03KR12/65/2 Dil./Brack. HCO3 140 64.0 0.23KR12/365/1 Brackish Cl 1100 8.8 0.09 0.02 8.9 1.00 0.01 0.03 0.03 0.03KR12/365/2 Brackish Cl 1000 7.3 0.10KR12/664/1 Saline 6400 44.0 0.41 74.0KR12/737/1 Saline 9600 35.0 0.34 92.0KR12/741/1 Saline 10200 36.0 0.51 0.66 100.0 0.70 0.02 0.25 0.50 0.03KR12/741/2 Saline 9100 28.0 0.35KR13/112/1 Brackish SO4 375 61.0 0.41 0.01 3.3KR13/112/2 Brackish SO4 240 50.0 0.31 0.01 2.5 0.57KR13/214/1 Brackish SO4 526 99.0 0.52 0.02 5.2KR13/214/2 Brackish SO4 470 120.0 0.26 0.02 4.7 0.90KR13/362/1 Brackish SO4 1190 53.0 0.20 0.02 8.8KR13/362/2 Brackish SO4 900 55.0 0.17 0.02 9.2 1.30

Appendix 2 (5/9)

156

Zr3H 238U(H2O) 234U/238U 238UP 234U/238U 18O 2H 222Rn

Sample Type mg/l TU ppb H20 ppb Partic. o/ooSMOW o/ooSMOW Bq/lPVP1/2 OverburdenPVP1/3 OverburdenPVP3A/1 OverburdenPVP3A/2 OverburdenPVP3B/1 OverburdenPVP3B/2 OverburdenPVP4A/1 OverburdenPVP4A/2 OverburdenPVP4B/1 OverburdenPVP4B/2 OverburdenPVP5A/1 OverburdenPVP5A/2 OverburdenPVP8A/1 OverburdenPVP8A/2 OverburdenPVP9A/1 OverburdenPVP9A/2 OverburdenPVP9B/1 OverburdenPVP9B/2 OverburdenPVP10A/1 OverburdenPVP10A/2 OverburdenPVP10B/1 OverburdenPVP10B/2 OverburdenPP2/1 Dil./Brack. HCO3 63.00PP2/2 Dil./Brack. HCO3PP5/1 Dil./Brack. HCO3 130.00PP5/2 Dil./Brack. HCO3PP7/1 Dil./Brack. HCO3 7.30 3.73 4.69 0.0987 1.93 -10.7 -81.0 290.00PP7/2 Dil./Brack. HCO3PP8/1 Dil./Brack. HCO3 16.60 0.69 1.30 0.02 1.16 -11.1 -81.8 22.00PP8/2 Dil./Brack. HCO3PP9/1 Dil./Brack. HCO3 120.00PP9/2 Dil./Brack. HCO3PR1/2 Dil./Brack. HCO3 12.10 0.53 1.24 0.14 1.53 -10.3 -77.3 34.00PR1/3 Dil./Brack. HCO3 14.20PR2/2 Dil./Brack. HCO3 13.30 18.20 1.26 1.47 1.35 -10.7 -79.7 890.00PR2/3 Dil./Brack. HCO3PR4/2 Dil./Brack. HCO3 58.00PR4/3 Dil./Brack. HCO3KR3/443/1 Brackish Cl 1.10 0.08 1.54 0.00 NR -12.3 -92.5 22.00KR4/860/1 Saline 0.80 0.03 0.00 0.00 NR -9.3 -49.5 76.00KR4/860/2 SalineKR6/58/1 Dil./Brack. HCO3 7.10 -10.5 -85.6KR6/99/1 Dil./Brack. HCO3 5.70 -10.5 -85.3KR6/99/2 Dil./Brack. HCO3 3.40KR6/99/3 Brackish SO4KR6/125/1 Brackish SO4 -9.3 -74.6KR6/135/1 Brackish SO4 4.10 -9.9 -80.0KR6/135/2 Brackish SO4 3.90KR6/135/3 Brackish SO4KR6/135/4 Brackish SO4

KR6/525/1 Saline 0.02 2.60 0.07 3.76 0.01 NR -10.7 -83.3 110.00KR6/525/2 SalineKR7/282/1 Brackish SO4 1.60 0.47 4.72 0.00 NR -9.9 -75.5 130.00KR9/149/1 Dil./Brack. HCO3 3.10KR9/149/2 Dil./Brack. HCO3

KR9/470/1 Saline 0.90 0.13 3.58 0.00 NR -11.4 -85.2 41.00KR10/498/1 Saline 0.80 0.02 2.52 0.00 NR -11.2 -77.0 18.00KR11/125/1 Dil./Brack. HCO3 2.40KR11/125/2 Dil./Brack. HCO3

KR11/952/1 SalineKR12/65/1 Dil./Brack. HCO3 1.10 3.41 6.42 NR NR -11.0 -89.2 410.00KR12/65/2 Dil./Brack. HCO3

KR12/365/1 Brackish Cl 0.02 2.60 0.07 3.76 0.01 NR -10.0 -69.5 86.00KR12/365/2 Brackish ClKR12/664/1 SalineKR12/737/1 Saline 0.56 3.30 0.01 58.00KR12/741/1 Saline 0.02 1.00 0.04 3.42 0.01 NR -11.6 -90.4 74.00KR12/741/2 SalineKR13/112/1 Brackish SO4 0.00 1.10 0.67 7.45 0.03 0.85 -12.7 -96.9 650.00KR13/112/2 Brackish SO4 0.00KR13/214/1 Brackish SO4 0.00 0.80 0.62 4.36 0.01 NR -9.8 -75.7 95.00KR13/214/2 Brackish SO4 0.00KR13/362/1 Brackish SO4 0.00 0.80 0.20 2.37 0.01 0.01 -11.2 -83.5 105.00KR13/362/2 Brackish SO4 0.00

Appendix 2 (6/9)

157

14C 14C 13C 34S(SO4)18O(SO4)

37Cl 87Sr/86Sr d87Sr 87Sr/86Sr N2

Sample Type PM BP o/ooPDB o/ooCDT o/ooPDB o/ooSMOC GTK ml/lPVP1/2 OverburdenPVP1/3 OverburdenPVP3A/1 OverburdenPVP3A/2 OverburdenPVP3B/1 OverburdenPVP3B/2 OverburdenPVP4A/1 OverburdenPVP4A/2 OverburdenPVP4B/1 OverburdenPVP4B/2 OverburdenPVP5A/1 OverburdenPVP5A/2 OverburdenPVP8A/1 OverburdenPVP8A/2 OverburdenPVP9A/1 OverburdenPVP9A/2 OverburdenPVP9B/1 OverburdenPVP9B/2 OverburdenPVP10A/1 OverburdenPVP10A/2 OverburdenPVP10B/1 OverburdenPVP10B/2 OverburdenPP2/1 Dil./Brack. HCO3PP2/2 Dil./Brack. HCO3PP5/1 Dil./Brack. HCO3PP5/2 Dil./Brack. HCO3PP7/1 Dil./Brack. HCO3 67.2 3150 -12.7 0.7187 13.38PP7/2 Dil./Brack. HCO3PP8/1 Dil./Brack. HCO3 80.3 1715 -21.1 9.48 0.7212 16.93PP8/2 Dil./Brack. HCO3PP9/1 Dil./Brack. HCO3PP9/2 Dil./Brack. HCO3PR1/2 Dil./Brack. HCO3 106.7 NR -25.5 0.7383 40.96PR1/3 Dil./Brack. HCO3PR2/2 Dil./Brack. HCO3 100.6 NR -23.0 0.7211 16.84PR2/3 Dil./Brack. HCO3PR4/2 Dil./Brack. HCO3PR4/3 Dil./Brack. HCO3KR3/443/1 Brackish Cl 45.0 6365 -4.3 NR NR 0.44 0.7175 11.97KR4/860/1 Saline 3.7 26450 16.8 NR NR 0.09 0.7203 15.70 167KR4/860/2 SalineKR6/58/1 Dil./Brack. HCO3 57.0 4465 -16.4 21.52 11.30KR6/99/1 Dil./Brack. HCO3 49.6 5585 -15.9 22.81 10.94KR6/99/2 Dil./Brack. HCO3 42.0 6915 -16.8 24.45 11.02KR6/99/3 Brackish SO4KR6/125/1 Brackish SO4 34.6 8470 -16.6 24.31 0.7170 11.05 0.7170KR6/135/1 Brackish SO4 46.4 6120 -16.3 22.59 11.60KR6/135/2 Brackish SO4 40.7 7165 -15.4 24.22 11.07KR6/135/3 Brackish SO4KR6/135/4 Brackish SO4

KR6/525/1 Saline NR NR NRKR6/525/2 SalineKR7/282/1 Brackish SO4 33.9 8650 -18.2 24.79 14.51 -0.07 0.7190 13.80 96KR9/149/1 Dil./Brack. HCO3 53.1 5030 -15.9 22.40 12.40KR9/149/2 Dil./Brack. HCO3

KR9/470/1 Saline 40.6 7190 -16.1 22.58 13.12 0.40 0.7194 14.40KR10/498/1 Saline NR NR NR NR NR 0.46 0.7194 14.30KR11/125/1 Dil./Brack. HCO3 50.8 5390 -15.1 24.30 13.40KR11/125/2 Dil./Brack. HCO3

KR11/952/1 SalineKR12/65/1 Dil./Brack. HCO3 43.9 6565 -16.7 28.26 12.86 44KR12/65/2 Dil./Brack. HCO3

KR12/365/1 Brackish Cl 29.4 9780 -16.2KR12/365/2 Brackish ClKR12/664/1 Saline 40.2KR12/737/1 SalineKR12/741/1 Saline NR NR NR 0.20 111KR12/741/2 SalineKR13/112/1 Brackish SO4 32.7 8935 -17.1 26.21 7.44 0.7173 11.38 0.7172 91KR13/112/2 Brackish SO4KR13/214/1 Brackish SO4 31.4 9260 -16.4 23.83 14.45 0.7179 12.31 0.7179 54KR13/214/2 Brackish SO4KR13/362/1 Brackish SO4 NR NR NR NR NR 0.7188 13.48 0.7187 42

Appendix 2 (7/9)

158

O2 CO CO2 H2 Ar He CH4 C2H2 C2H4 C2H6

Sample Type µl/l µl/l ml/l µl/l ml/l ml/l ml/l µl/l µl/l µl/l

PVP1/2 OverburdenPVP1/3 OverburdenPVP3A/1 OverburdenPVP3A/2 OverburdenPVP3B/1 OverburdenPVP3B/2 OverburdenPVP4A/1 OverburdenPVP4A/2 OverburdenPVP4B/1 OverburdenPVP4B/2 OverburdenPVP5A/1 OverburdenPVP5A/2 OverburdenPVP8A/1 OverburdenPVP8A/2 OverburdenPVP9A/1 OverburdenPVP9A/2 OverburdenPVP9B/1 OverburdenPVP9B/2 OverburdenPVP10A/1 OverburdenPVP10A/2 OverburdenPVP10B/1 OverburdenPVP10B/2 OverburdenPP2/1 Dil./Brack. HCO3PP2/2 Dil./Brack. HCO3PP5/1 Dil./Brack. HCO3PP5/2 Dil./Brack. HCO3PP7/1 Dil./Brack. HCO3PP7/2 Dil./Brack. HCO3PP8/1 Dil./Brack. HCO3PP8/2 Dil./Brack. HCO3PP9/1 Dil./Brack. HCO3PP9/2 Dil./Brack. HCO3PR1/2 Dil./Brack. HCO3PR1/3 Dil./Brack. HCO3PR2/2 Dil./Brack. HCO3PR2/3 Dil./Brack. HCO3PR4/2 Dil./Brack. HCO3PR4/3 Dil./Brack. HCO3KR3/443/1 Brackish Cl 1300 12 0.46 20 4.4 51.90 0.1 0.1 200.0KR4/860/1 Saline 34900 58 0.23 7000 4.2 16.3 768.00 0.6 0.6 10300.0KR4/860/2 Saline 270 46 0.03 194 14.7 750 0.46 1.4 13010.0KR6/58/1 Dil./Brack. HCO3KR6/99/1 Dil./Brack. HCO3KR6/99/2 Dil./Brack. HCO3KR6/99/3 Brackish SO4KR6/125/1 Brackish SO4KR6/135/1 Brackish SO4KR6/135/2 Brackish SO4KR6/135/3 Brackish SO4KR6/135/4 Brackish SO4

KR6/525/1 Saline 0.22 22 0.04 13 13.3 340.80 0.2 0.35 2250.0KR6/525/2 SalineKR7/282/1 Brackish SO4 31000 13 2.42 8 1.5 0.3 0.14 0.1 0.1 0.7KR9/149/1 Dil./Brack. HCO3KR9/149/2 Dil./Brack. HCO3

KR9/470/1 Saline 84 17 0.17 33 9.7 234.50 0.2 0.2 1400.0KR10/498/1 Saline 230 23 0.03 14 15.0 348.20 0.2 0.2 2440.0KR11/125/1 Dil./Brack. HCO3KR11/125/2 Dil./Brack. HCO3

KR11/952/1 SalineKR12/65/1 Dil./Brack. HCO3 39 3 4.62 2 0.7 0.1 0.51 0.0 0.0 1.7KR12/65/2 Dil./Brack. HCO3

KR12/365/1 Brackish Cl 960 10 0.12 6 8.6 120.50 0.1 0.1 802.0KR12/365/2 Brackish ClKR12/664/1 Saline 850 17 0.01 9.9 2.1 5.3 275 0.17 0.1 4110.0KR12/737/1 Saline 160 26 0.02 16 12.7 442 0.26 0.4 5800.0KR12/741/1 Saline 9560 36 0.04 64 2.0 13.9 518.20 0.6 2.0 5280.0KR12/741/2 SalineKR13/112/1 Brackish SO4 1537 1.22 3 1.5 0.5 3.51 0.1 0.1 22.6KR13/112/2 Brackish SO4KR13/214/1 Brackish SO4 2318 1.77 2 1.9 0.3 0.61 0.1 0.2 4.7KR13/214/2 Brackish SO4KR13/362/1 Brackish SO4 9907 0.40 3 0.9 4.4 15.24 0.0 0.0 0.1KR13/362/2 Brackish SO4

Appendix 2 (8/9)

159

C3H6 C3H813C(CH4) 13C(C2H6) 13C(C3H8) 13C(CO2) 18O(CO2) 2H(CH4)

Sample Type µl/l µl/l o/ooPDB o/ooPDB o/ooPDB o/ooPDB (o/ooPDB) o/ooSMOWPVP1/2 OverburdenPVP1/3 OverburdenPVP3A/1 OverburdenPVP3A/2 OverburdenPVP3B/1 OverburdenPVP3B/2 OverburdenPVP4A/1 OverburdenPVP4A/2 OverburdenPVP4B/1 OverburdenPVP4B/2 OverburdenPVP5A/1 OverburdenPVP5A/2 OverburdenPVP8A/1 OverburdenPVP8A/2 OverburdenPVP9A/1 OverburdenPVP9A/2 OverburdenPVP9B/1 OverburdenPVP9B/2 OverburdenPVP10A/1 OverburdenPVP10A/2 OverburdenPVP10B/1 OverburdenPVP10B/2 OverburdenPP2/1 Dil./Brack. HCO3PP2/2 Dil./Brack. HCO3PP5/1 Dil./Brack. HCO3PP5/2 Dil./Brack. HCO3PP7/1 Dil./Brack. HCO3PP7/2 Dil./Brack. HCO3PP8/1 Dil./Brack. HCO3PP8/2 Dil./Brack. HCO3PP9/1 Dil./Brack. HCO3PP9/2 Dil./Brack. HCO3PR1/2 Dil./Brack. HCO3PR1/3 Dil./Brack. HCO3PR2/2 Dil./Brack. HCO3PR2/3 Dil./Brack. HCO3PR4/2 Dil./Brack. HCO3PR4/3 Dil./Brack. HCO3 -58.4 -18.9 4.9 -274.0KR3/443/1 Brackish Cl 0.2 0.2KR4/860/1 Saline 1.2 350.0 -63.5 NR NR -25.4 NR NRKR4/860/2 Saline 0.91 390.0KR6/58/1 Dil./Brack. HCO3KR6/99/1 Dil./Brack. HCO3KR6/99/2 Dil./Brack. HCO3KR6/99/3 Brackish SO4KR6/125/1 Brackish SO4KR6/135/1 Brackish SO4KR6/135/2 Brackish SO4KR6/135/3 Brackish SO4KR6/135/4 Brackish SO4 -41.3 -38.4 NR NR NR -257.0KR6/525/1 Saline 0.44 75.0KR6/525/2 Saline -56.5 NR NR -28.2 NR NRKR7/282/1 Brackish SO4 0.3 0.3KR9/149/1 Dil./Brack. HCO3 -43.2 -35.0 -32.9 -17.3 NR -309.0KR9/149/2 Dil./Brack. HCO3 -46.5 -39.2 NR NR NR -277.0KR9/470/1 Saline 0.3 10.0KR10/498/1 Saline 0.5 14.0 -61.5 NR NR -22.9 NR NRKR11/125/1 Dil./Brack. HCO3 -46.6 -36.3 -34.5 -20.0 NR -236.0KR11/125/2 Dil./Brack. HCO3

KR11/952/1 SalineKR12/65/1 Dil./Brack. HCO3 0.1 0.1 -44.8 -27.0 -37.3KR12/65/2 Dil./Brack. HCO3

KR12/365/1 Brackish Cl 0.2 14.0 -44.2 -25.0 -37.2KR12/365/2 Brackish ClKR12/664/1 Saline 0.33 78.0 -48.3 -35.7 -30.2KR12/737/1 Saline 0.53 112.0KR12/741/1 Saline 0.5 120.0KR12/741/2 SalineKR13/112/1 Brackish SO4 0.1 0.1KR13/112/2 Brackish SO4 -49.7 -40.6 -38.7 -15.0 -272.0KR13/214/1 Brackish SO4 0.2 0.1KR13/214/2 Brackish SO4 -22.4 -15.4 5.0KR13/362/1 Brackish SO4 0.0 0.0 -43.2 -35.0 -32.9 -17.3 NR -309.0KR13/362/2 Brackish SO4 -42.9 -21.5 -0.6 -307.0

Appendix 2 (9/9)

LIST OF REPORTS 1 (2) POSIVA-reports 2003 POSIVA 2003-01 Vertical and Horizontal Seismic Profiling Investigations at Olkiluoto,

2001 Calin Cosma, Nicoleta Enescu, Erick Adam, Lucian Balu

Vibrometric Oy March 2003 ISBN 951-652-115-0

115 p. POSIVA 2003-02 Baseline Conditions at Olkiluoto Posiva Oy September 2003 ISBN 951-652-116-9 218 p. POSIVA 2003-03 ONKALO Underground Characterisation and Research Programme

(UCRP) Posiva Oy September 2003 ISBN 951-652-117-7 142 p. POSIVA 2003-04 Thermal Analyses of Spent Nuclear Fuel Repository Kari Ikonen, VTT Processes June 2003 ISBN 951-652-118-5 62 p. POSIVA 2003-05 Programme of Monitoring at Olkiluoto During Construction and

Operation of the ONKALO Posiva Oy December 2003 ISBN 951-652-119-3 92 p. POSIVA 2003-06 Assessment of disturbances caused by construction and operation of

ONKALO Timo Vieno, Jarmo Lehikoinen, Jari Löfman, Henrik Nordman VTT Processes Ferenc Mészáros The Relief Laboratory ISBN 951-652-120-7 92 p.

LIST OF REPORTS 2 (2) POSIVA 2003-07 Hydrochemical interpretation of baseline groundwater conditions at

the Olkiluoto site. Petteri Pitkänen, Ari Luukkonen, Sami Partamies VTT Building and Transport ISBN 951-652-121-5 159 p.

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