POSIVA OY

POSIVA 98-07

Geochemical modelling of groundwater evolution and

residence time at the Kivetty site

Petteri Pitkanen Ari Luukkonen

VTT Communities and Infrastructure

Paula Ruotsalainen Fintact Oy

Hilkka Leino-Forsman Ulla Vuorinen

VTT Chemical Technology

December 1998

Mikonkatu 15 A, FIN-001 00 HELSINKI, FINLAND

Phone (09) 2280 30 (nat.). (+358-9-) 2280 30 (int.)

Fax (09) 2280 3719 (nat.). (+358-9-) 2280 3719 (int.l

POSIVA OY

POSIVA 98-07

Geochemical modelling of groundwater evolution and

residence time at the Kivetty site

Petteri Pitkanen Ari Luukkonen

VTT Communities and Infrastructure

Paula Ruotsalainen Fintact Oy

Hilkka Leino-Forsman Ulla Vuorinen

VTT Chemical Technology

December 1998

Mikonkatu 15 A, FIN-001 00 HELSINKI, FINLAND

Phone (09) 2280 30 (nat.}, (+358-9-) 2280 30 (int.)

Fax (09) 2280 3719 (nat.}, (+358-9-) 2280 3719 (int.)

TOIMEKSIANTO GEOKEMIALLISEST A MALLINT AMISEST A PAIKKATUTKIMUKSISSA: Kivetyn hydrogeokemiallinen mallinnus

TEKIJAORGANISAATIOT:

TILAAJA:

TILAAJAN YHDYSHENKILO:

TILAUSNUMEROT:

TEKIJAORGANISAATIOIDEN YHDYSHENKILOT:

'}~·~ Petteri Pi tkanen

VTT Yhdyskuntatekniikka PL 1900 02044 VTT

VTT Kemiantakniikka PL 1400 02044 VTT

Fintact Oy Hopeatie 1 B 00440 HELSINKI

Posiva Oy Mikonkatu 15 A 00100 HELSINKI

Margit Snellman, Posiva Oy

YKI: 9268/95/MiN, 9569/96/MMK, 9644/97 /MVS KET:9267 /95/MiN, 9661/97 /MVS Fintact: 9576/97/TIMO, 9801198/AJH, 9550/98/ AJH, 9559/98/ AJH

VTT Yhdyskuntatekniikka

Hilkka Leino-Forsman VTT Kemiantekniikka

.,... -~.,:~[A /~-bU-r/J--.(/Gv W'--.-

Paula Ruotsalainen Fintact Oy

TARKASTAJA:

Kai Front VTT Yhdyskuntatekniikka

ISBN 951-652-045-6 ISSN 1239-3096

The conclusions and vievvpoints presented in the report are

those of author(s) and do not necessarily coincide

vvith those of Posiva.

Posiva-raportti - Posiva Report

Posiva Oy Mikonkatu 15 A, FIN-001 00 HELSINKI, FINLAND Puh. (09) 2280 30- lnt. Tel. +358 9 2280 30

Tekija(t) - Author(s) Toimeksiantaja(t)- Commissioned by

Petteri Pi tkanen *, Ari Luukkonen *, Paula Ruotsalainen * * Hilkka Leino-Forsman***, Ulla Vuorinen*** * VTT Communities and Infrastructure Posiva Oy ** Fintact Oy, *** VTT Chemical Technology

Nimeke -Title

Raportin tunnus - Report code

POSIV A 98-07

Julkaisuaika- Date

December 1998

GEOCHEMICAL MODELLING OF GROUNDWATER EVOLUTION AND RESIDENCE TIME AT THE KIVETTY SITE

Tiivistelma - Abstract

An understanding of the geochemical evolution of groundwater is an essential part of the performance assessment and safety analysis of the final disposal of radioactive waste into the bedrock. 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 progress in different parts of the bedrock at Kivetty has been created and the significance of chemical reactions along different flowpaths calculated. Long term hydrodynamics have also been evaluated.

The interpretation and modelling are based on groundwater samples (38 altogether) obtained from the soil layer, shallow wells in the bedrock, and five deep multi-packered boreholes (KR1-KR5) in the bedrock for which a comprehensive data set on dissolved chemical species and isotopes was available. Some analyses of dissolved gases and their isotopic measurements were also utilised. The data covers the bedrock at Kivetty to a depth of 850m. The results from groundwater chemistry, isotopes, petrography, hydrogeology of the site, geomicrobial studies, and 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 C-14 age calculations of groundwater was given a mass-balance approach (NETPATH). Reaction-path calculations (EQ3/6) were used to verify the thermodynamic feasibility of the reaction models obtained.

The hydrogeochemistry of Kivetty is characterised by evolution from low-saline-carbonate-rich recharge water towards Na-Ca-Cl-type water. The salinity remains low. The most important changes in the chemistry of the groundwater are due to carbonate reactions: oxidising of organic carbon, and dissolution and precipitation of calcite. The carbonate reactions and slight hydrolysis of silicates stabilise the pH value at 8-9. In addition to aerobic oxidation of organic matter, oxidative dissolution of biotite seems to be an important oxygen consumer at shallow depth during recharge. The most important process controlling the redox state deeper in the bedrock was interpreted to be the microbially mediated sulphate reduction with simultaneous anaerobic respiration of organic carbon. This process buffers the redox level of about -200 ... -300 mV depending on the pH. Even though the salinities of the groundwater samples and mass-transfer along flow paths remain low, the geochemical evolution has fully developed and has reached quite a stable thermodynamic state. The residence times of the groundwater samples cover the time span back to glaciation. Young ages seem to be limited to the upper part of bedrock, and any really dynamic natural flowpath with deep observed recently recharged water cannot be demonstrated. Deglacial or subglacial ages (over 9,700 years old at Kivetty) are typical below the 150-300m level in the bedrock. Subglacial waters are interpreted to derive from mixing of preglacial water and meltwater, the input of which is estimated to be about 20% at the most. Indications of elevated oxygen intrusion cannot be observed in groundwater having glacial signals.

Avainsanat - Keywords

groundwater chemistry, environmental isotopes, Proterozoic granitoids, nuclear waste disposal, water-rock interaction, geochemical modelling, palaeohydrogeology

ISBN ISSN ISBN 951-652-045-6 ISSN 1239-3096

Sivumaara- Number of pages Kieli - Language 139 English

Posiva-raportti - Posiva Report

Posiva Oy Mikonkatu 15 A, FIN-00100 HELSINKI, FINLAND Puh. (09) 2280 30- lnt. Tel. +358 9 2280 30

Tekija(t) - Author(s) Toimeksiantaja(t)- Commissioned by

Petteri Pitkanen*, Ari Luukkonen*, Paula Ruotsalainen** Hilkka Leino-Forsman***, Ulla Vuorinen*** * VTT Yhdyskuntatekniikka Posiva Oy ** Fintact Oy, *** VTT Kemiantekniikka

Nimeke- Title

Raportin tunnus - Report code

POSIV A 98-07

Julkaisuaika- Date

Joulukuu 1998

GEOKEMIALLINEN MALLI POHJAVEDEN OLOSUHTEISTA JA KEHITYKSESTA KIVETYN TUTKIMUSPAIKALLA

Tiivistelma -Abstract

Pohjaveden geokemiallisen kehityksen ymmartaminen on oleellinen osa ydinjatteiden kalliosijoituksen turvallisuutta ja toimintakykya. Loppusijoitussysteemin teknisten paastoesteiden toiminta seka mahdollisesti vapautuvien radionuklidien kulkeutuminen riippuvat kemiallisista olosuhteista. Naiden seikkojen ymmartamisen edellytyksena on pystya kuvaamaan niita kallio-vesivuorovaikutuksia, jotka saatavat pohjaveden kemiallisia olosuhteita. Tutkimuksen tavoitteena on tulkita pohjaveden koostumukseen vaikuttaneet prosessit ja ne tekijat, jotka ohjaavat hydrogeokemiaa kuten pH- ja redox-olosuhteita. Tutkimuksessa luodaan malli Kivetyn hydrogeokemiallisesta kehityksesta kallioperan eri osissa seka maaritetaan laskennallisesti kemiallisten reaktioiden merkitys pitkin eri virtausreitteja. Lisaksi arvioidaan pohjaveden pitkaaikaisia liikkeita pohjavesikemian avulla.

Tulkinta ja mallinnus perustuu pohjavesinaytteisiin (yht. 38 kpl), joita on otettu maaperasta, porakaivoista ja viidesta syvasta, monitulpatuista kairausrei'ista (KRI - KR5). Naytteista on ollut kaytettavissa laaja liuenneiden kemiallisten aineiden ja isotooppien aineisto. Lisaksi on ollut kaytettavissa kaasuanalyysituloksia ja niiden isotooppimaarityksia. Pohjavesikemian aineisto kattaa Kivetyn kallioperan noin 850 m:n syvyydelle. Lisaksi on pohjaveden kehitys­historiaan vaikuttaneiden prosessien arvioinnissa hyodynnetty alueella tehtyja mineralogisia, geomikrobi- ja hydrogeologisia tutkimuksia seka pohjavesikemiallisen aineiston paakomponenttianalyysin ja termodynaamisten tasapainolaskujen tuloksia. Pohjavesi-kalliovuorovaikutus ja isotooppikemiallinen kehitys seka C-14 viipymat on tulkittu massatasapainomallinnuksen (NETPATH) avulla. Reaktiopolkumallinnuksen avulla (EQ3/6) on tarkastettu tulkinnan termodynaaminen kelpoisuus.

Kivetyn hydrogeokemiaa luonnehtii kallioon suotautuvan vahasuolaisen, karbonaattipitoisen, meteorisen veden kehittyminen Na-Ca-Cl -tyypin suuntaan suolaisuuden pysyessa kuitenkin vahaisena. Merkittavimmat pohjaveden koostumukseen liittyvat muutokset aiheutuvat karbonaattireaktioista: orgaanisen hiilen hapettuminen, kalsiitin liukeneminen ja saostuminen. Karbonaattireaktiot ja vahainen silikaattien hydrolyysi stabiloivat pohjaveden pH:n tasolle 8-9. Aerobisissa olosuhteissa tapahtuva biotiitin hydrolyysi (rauta hapettuu) yhdessa organisen aineksen hapettumisen kanssa vaikuttavat merkittavimmilta happeakuluttavilta reaktioilta maan pinnan laheisyydessa suotaumisvaiheessa. Syvemmalla kalliossa, anaerobisissa olosuhteissa tapahtuva mikrobien katalysoima sulfaatin pelkistyminen orgaanisen hiilen samalla hapettuessa on tulkittu olevan tarkein redox-tilaa saata.va prosessi, joka puskuroi pohjaveden Eh:n pitkalla aikavalilla -200 ... -300 mV:n tasolle riippuen pH:sta. Vaikka pohjavesien suolapitoisuudet ja aineensiirtymat virtauspoluilla pysyvat pienina, geokemiallisesti naytteet ovat kuitenkin kehittyneet pitkalle ja saavuttaneet varsin stabiilin termodynaamisen tilan. Pohjavesinaytteiden keskimaaraiset iat yltavat aina jaakauden aikaisiin saakka. Nuoret iat rajoittuvat aivan kallion ylaosiin eika syvalle kallioon nayttaisi johtavan nopeita, luonnollisia virtausreitteja. Jaakauden tai sen paattymisvaiheen (n. 9700 vuotta sitten) aikaiset iat ovat tyypillisia 150-300 m:n alapuolella paikasta riippuen. Viipymansa perusteella glasiaalisiksi maaritetyt pohjavesinaytteet on tulkittu olevan preglasiaalisen ja sulamisveden sekoituksia, jossa sulamisveden osuus on suurimmillaan n. 20 %. Naissa naytteissa ei ole kuitenkaan havaittavissa merkkeja huomattavasta hapen tunkeutumisesta kallioon.

Avainsanat - Keywords

pohjavesikemia, isotoopit, proterotsooiset granitoidit, ydinjatteen loppusijoitus, vesi-kalliovuorovaikutus, geokemiallinen mallinnus, paleohydrogeologia

ISBN ISSN ISBN 951-652-045-6 ISSN 1239-3096

Sivumaara - Number of pages Kieli - Language 139 Englanti

7

TABLE OF CONTENTS:

1. INTRODUCTION ............................................................................................................ 11 2. GEOLOGICAL SETTING ............................................................................................... 13

2.1 Bedrock lithology ....................................................................................................... 13 2.1.1 Regional scale ...................................................................................................... 13 2.1.2 Local features ...................................................................................................... 16

2.2 Structural features ....................................................................................................... 17 2.3 Fracture mineralogy .................................................................................................... 17

3. HYDROGEOLOGICAL SETTING ................................................................................. 21 3.1 Hydraulic environment ............................................................................................... 21 3 .2Groundwater flow based on a conceptual model. ........................................................ 22 3.3 Groundwater flow indications based on transverse flow measurements .................... 28

4. STRATEGY OF GEOCHEMICAL MODELLING ......................................................... 33 4.1 Consistency checking between hydrogeochemistry and hydrogeology. ..................... 34

5. GROUNDWATER CHEMISTRY ................................................................................... 35 5.1 General ....................................................................................................................... 35 5.2 Sampling and chemical analyses ................................................................................ 35 5.3 Representativeness of hydrogeochemical data ........................................................... 37 5.4 Hydrochemical features .............................................................................................. 38

5.4.1 Principal components of groundwater compositions .......................................... 38 5.4.2 Main variables and components ......................................................................... .42 5.4.3 Trends of main cations ........................................................................................ 45 5.4.4 Trends of main anions ......................................................................................... 46 5.4.5 Redox conditions at Kivetty ................................................................................ 48

5.5 Isotopes ....................................................................................................................... 52 5.5.1 Stable isotopes (8H-2 and 80-18) of groundwater .............................................. 52 5.5.2 Tritium (H-3) ....................................................................................................... 53 5.5.3 Sulphur of aqueous S04 (8S-34 in S04) .............................................................. 54 5.5.4 Uranium isotopes and radon ................................................................................ 55 5.5.5 Carbon isotopes (8C-13, C-14 in DIC) ................................................................ 57

5.5.5.1 Radiocarbon and evidence of uncertainty in carbonate data ........................ 59 5.5.5.2 Carbon isotopes and carbonate evolution ..................................................... 61 5.5.5.3 Implications for palaeohydrogeology ........................................................... 63

5.5.6 Strontium isotopes ............................................................................................... 64 5.6 Thermodynamic controls of the groundwater system ................................................ 66

6. DISCUSSION OF THE GEOCHEMICAL SYSTEM ..................................................... 69 6.1 Hydrogeochemical evolution ...................................................................................... 69

6.1.1 Recharge and carbonate evolution ....................................................................... 69 6.1.2 Salinity changes ................................................................................................... 74 6.1.3 Redox related processes ...................................................................................... 76

6.2 Hydrogeological implications ..................................................................................... 77 7. RESULTS OF MASS-TRANSFER MODELLING ........................................................ 81

7.1 Mass-balance reaction models .................................................................................... 81

8

7.1.1 General ................................................................................................................ 81 7 .1.2 Restrictions and selection of flowpaths ............................................................... 82 7 .1.3 Phases and constraints used in modelling ........................................................... 84 7 .1.4 Isotopic calculations and initial values for carbon isotopes ................................ 86 7 .1.5 Results of mass-balance modelling ..................................................................... 89

7 .1.5 .1 Mass-transfer ................................................................................................ 91 7.1.5.2 Carbon-14 age of groundwater and palaeohydrogeological implications .... 94

7.2 Modelling of reaction paths with EQ6 ....................................................................... 97 7.2.1 Modelling ............................................................................................................ 97 7 .2.2 Results ................................................................................................................. 98 7.2.3 Discussion ......................................................................................................... 100

8. SUMMARY AND IMPLICATIONS FOR SAFETY ASSESSMENT ......................... 109 9. REFERENCES ............................................................................................................... 113

Appendix 1. Lithology, open and filled fracture frequency, hydraulic conductivities, interpreted fracture zones, and groundwater sampling sections with measured hydraulic heads for boreholes KR1-KR5 ........................... 122

Appendix 2. Packed-off intervals used for geochemical groundwater sampling, together with hydraulically conductive depth intervals, inferred locations of conductive fractures, and fracture minerals considered to be in contact with sampled groundwater ............................................ 127

Appendix 3. Hydraulic head field in R9 and R11 based on the study by Taivassalo & Meszaros (1994) with intersection lines of other fracture zones and intersections of boreholes KR1, KR4 and KR5 ....................................... 130

Appendix 4. Hydraulic head field in R12 and R15 based on the study by Taivassalo & Meszaros (1994) with intersection lines of other fracture zones and intersections ofboreholes KR1, KR2, KR4 and KR5 .............................. 131

Appendix 5. Hydraulic head field in R22 and R23 based on the study by Taivassalo & Meszaros (1994) with intersection lines of other fracture zones and intersections of boreholes KR1, KR4 and KR5 ....................................... 132

Appendix 6. Hydrogeochemical data used in the present modelling study ........................ 133

Appendix 7. Evaluation of the hydrogeochemical data of Kivetty .................................... 138

Appendix 8. Samples included in the modelling study in spite of failing the quality classification ................................................................................................. 139

9

PREFACE

This study is a part of the research programme for spent nuclear fuel disposal in Finnish bedrock. The programme is conducted by Posiva Oy.

This study was a collaboration between VTT Communities and Infrastructure, VTT Chemical Technology and Fintact Oy. The contact persons were Margit Snellman from Posiva Oy, Petteri Pitkanen and Hilkka Leino-Forsman from VTT and Paula Ruotsalainen from Fin tact Oy.

The authors wish to thank Margit Snellman for her comments on the draft, Eliisa Hatanpaa and Virpi Karttunen (Imatran Voima Oy) for their expertise in hydrochemical data, Kurt Meling (VTT Energy) for assisting in computing (FEFTRA program) hydraulic head fields, and Adelaide Lonnberg for checking the language.

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

10

11

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, a model describing the water-rock interaction controlling the basic hydrochemistry is needed. The ultimate goal is to create a site specific model that describes reliably the changes in groundwater composition, and explains causes for them.

This study discusses the geochemical conditions at the Aanekoski, Kivetty site, central Finland (Fig. 1-1). The interpretation is based on the data obtained during the preliminary (1987-1992, Teollisuuden Voima Oy 1992; Pitkanen et al. 1992; Lampen & Snellman 1993) and detailed (1993-1996, Posiva 1996; Ruotsalainen & Snellman 1996) site investigations for the final disposal of spent nuclear fuel. Hydrogeochemical studies during the detailed investigation phase have produced a lot of data, both regionally and as a function of depth on the different chemical parameters. Isotopic data, especially, has given more insight into the evaluation of the different geochemical processes occurring in the groundwater system at Kivetty.

Gulf of Finland

Figure 1-1. The Kivetty site in central Finland within the Kymijoki river catchment system.

12

This study continues the specific site modelling studies being a part of the site characterisation process of Posi va. Therefore the results of Ki vetty are often compared to those from the other sites: Romuvaara (Pitkanen et al. 1996a) and Olkiluoto (Pitkanen et al. 1996b). Usage of different tools and methods for interpretation are also described in detail in the Romuvaara modelling study (Pitkanen et al. 1996a).

The objectives of this study are to:

- identify the processes responsible for the evolution of ground water controlling the pH and Eh conditions in the bedrock,

-quantify those reactions that produce the observed groundwater compositions, - create a geochemical conceptualisation of the flow system, - test the consistency of the hydrogeological model and chemical evolutionary

paths and - test the extent of ground water contamination by drilling activities or

groundwater sampling.

Considering chemical, isotopic, petrographic and hydrological data, ion plots and speciation calculations with EQ3NR (Wolery 1992), the thermodynamic controls on the water composition and trends constraining the processes are interpreted. In order to determine the reactions which can explain the changes along the flowpath during the evolution of the groundwater system, and to determine to what extent these reactions take place, the mass-balance model NETPATH (Plummer et al. 1994) is used. Isotopic calculations have been included for the modelling of the carbon isotope evolution of dissolved carbonate in the groundwater. This inclusion of isotopic data provides additional criteria for testing the reaction hypothesis. The reaction path calculation with the EQ6 (W olery 1992) code is used to test the thermodynamic feasibility of the model derived and the chemical conditions of equilibrium solutions.

Throughout this report "depth" means borehole length, i.e. our notation discards the effect of borehole bending and the true sample depth is always smaller than the noted length indicates. However, the dip of boreholes is steep (> 70°) hence this simplification has no effect to interpretations due to used low accuracy (tens of metres) in depth dependencies.

13

2. GEOLOGICAL SETTING

2.1 Bedrock lithology

2.1.1 Regional scale

The Kivetty site is located in central Finland, 80 km north of Jyvaskyla in the town of Aanekoski (Fig. 2-1). The bedrock is Palaeoproterozoic in age, almost exclusively plutonic in origin, and part of the Svecokarelian orogenic complex. Rock types at the research site and its surroundings are, in rough order of abundance, porphyritic granodiorite, porphyritic granite, granodiorite, granite, gabbro and dikes. Coarse porphyritic varieties of porphyritic rocks are common. Gabbro intrusions occur as small bodies and in places exhibit magmatic layering. Mafic and felsic dikes show textures varying from porphyritic to aplitic. According to Anttila et al. (1992) the chronological order of the rock types from the oldest to the youngest is as follows: gabbro, porphyritic granodiorite, porphyritic granite, granodiorite, and granite. In places there are also some schist and gneiss inclusions, which have been interpreted mainly as felsic metavolcanite xenoliths. Gneiss and mafic inclusions in porphyritic granodiorite seem to be distinctly more common than in porphyritic granite.

Figure 2-1. Location of the Kivetty site in central 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, ~ = Archean basement.

AANEKOSKI KIVETTY TUTKIMUSALUEEN KALLIOPERA GEOLOGY OF THE STUDY SITE

MERKKIEN SEUTYS LEGEND

D

• D D

• / ,..

Kvartal-maaailpigneissl Quartz-feldspar gneiss

Gabro Gabbro

Porfyyrlnen granodlorlittl Porphyritlc granodiorite

Granodlorllttl Granodlorite

Granlittllporfyyrlnen graniitti Granlta/porphyritic granite

Mafisla juonla Maflc dykes

Kalranreiki Borehole

LAATINEET COMPILED BY

Seppo Paularnikl Markku Paananen Geologlan tutklmuskeskus Geological Survey of Finland

km

Pohjakartan © Maanmittauslaitos Base map © National Survey of Finland

c:\kartatlklvetty\ldvegeol.wor

Figure 2-2. Litho logical map of the Kivetty area (Front et al. 1998) showing the topography of the area and locations of bore holes.

~

~

15

Average compositions of the granitic and granodioritic rock varieties forming the vast majority of rock bulk at Kivetty are shown in Table 2-1. For comparison, average compositions of dominating rock types at Romuvaara are also included. Although the silica contents of granite and leucotonalite gneiss, and of granodiorite and tonalite gneiss are respectively similar, there are several significant differences with respect to other elements. In general, the rocks of Kivetty contain higher amounts of Fe, Mn, K, Rb, Y, Zr, Nb, Cs, U and F, and lower amounts of Na, Th and Sr than at Romuvaara. Mineralogically this indicates that the former are richer, for example, in potassium feldspar and fluorite. Moreover, indications of metamorphism and alteration, like muscovite/sericite, carbonate and epidote, are more frequent in the rocks of Romuvaara than Kivetty. In respect of dark micas, biotite at Kivetty is clearly more Fe-rich than biotite at Romuvaara.

Table 2-1. Average granitic and granodioritic whole-rock compositions of the Kivetty site compared with average main rock type compositions of the Romuvaara research site (modified after Gehor et al. 1995 & 1996). Main components and sulphur are expressed in wt.%, other elements in ppm.

Kivetty Romuvaara Granite Granodiorite Leucotonalite Gneiss Tonalite Gneiss

n 10 5 3 7 Average Stdev Average Stdev Average Stdev Average Stdev

Si02 73.18 2.44 65.26 0.58 72.77 0.68 70.29 2.01 Ti02 0.30 0.12 0.68 0.06 0.23 0.03 0.30 0.12 Al203 13.35 0.56 15.26 0.09 14.50 0.30 15.19 0.69 Fe203 2.64 0.98 5.57 0.42 1.53 0.09 2.10 0.87 M nO 0.04 0.02 0.09 0.01 0.02 0.01 0.02 0.02 M gO 0.37 0.16 1.03 0.11 0.76 0.09 0.94 0.42 CaO 1.39 0.46 3.04 0.19 3.00 0.85 2.65 0.72 Na20 2.96 0.27 3.44 0.16 4.48 0.08 4.32 0.69 K20 5.16 0.54 4.31 0.14 1.62 0.84 2.80 1.57 P20s 0.06 0.04 0.19 0.02 0.05 0.02 0.11 0.05 Sum 99.44 98.87 98.96 98.71 Th 6.3 3.6 6.5 3.5 14.9 15.8 11.8 17.8 Rb 176 26 135 25 54 20 78 22 y 30 8 41 16 13 6 12 4

Ba 632 302 1052 89 353 230 1087 773 Zr 220 73 355 21 128 11 172 143 Nb 12 2 19 8 <10 <10 Sr 150 49 258 13 378 13 518 239 Cs 2 1 4 2 <1.0 1.0 0.5 u 2.8 0.9 3.3 1.4 0.7 0.3 0.3 0.7 F 431 255 639 579 209 19 324 114 Cl 229 50 313 47 174 23 215 74 s 0.02 0.02 0.03 0.01 0.01 0.01 0.03 0.01

16

Biotite, hornblende and plagioclase compos1t1ons in gran1t1c and granodioritic rock varieties of the Kivetty site are shown in Table 2-2. In general, biotite and hornblende in granites tend to contain more iron than in granodiorites, and plagioclases tend to be more albitic in granites than in granodiorites. Especially in the case of dark minerals, it can be concluded that iron rich end-members are clearly more susceptible to surface weathering via iron oxidation and subsequent transportation than the magnesium rich end-members (e.g. Acker & Bricker 1992, Malmstrom et al. 1995, Bullen et al. 1997). Plagioclases rich in anorthite are more prone to weathering processes than albitic plagioclases. Theoretically biotite and hornblende in granites are more susceptible, and plagioclase more resistant, to weathering than in granodiorites.

Table 2-2. Average composition of biotite, hornblende and plagioclase in granitic and granodioritic rocks (modified after Gehor et al. 1995). Tabulated compositions in the footnotes have been calculated to conform to the following theoretical formulas: biotite K(Mg, Fe) 3(AlSi30 IO)(OH)2, hornblende NaCa2(Mg,Fe,Al) s(Si,Al)s022(0H) 2, and plagioclase (Ca,Na)(Al,Si)AlSi208.

GRANITE GRANODIORITE Biotite (I Hornblende <2 Plagioclase Biotite <4 Hornblende <5 Plagioclase <6

n 9 11 10 5 4 5 Si02 33.39 39.24 61.58 34.02 40.94 60.01 Ti02 3.30 1.12 4.09 1.53 Al203 13.75 9.94 23.63 13.49 9.11 24.59 Cr203 0.00 0.01 0.02 0.02 FeO 29.56 27.23 0.03 27.27 25.06 0.06 MnO 0.40 0.61 0.26 0.53 M gO 3.83 3.05 5.49 4.63 CaO 0.02 10.68 4.95 0.01 10.89 6.02 Na20 0.04 1.51 8.78 0.08 3.31 8.11 K20 8.43 1.48 0.22 8.93 1.21 0.30 NiO 0.06 0.03 0.05 0.09 ZnO 0.11 0.07 0.11 0.09 Tot 92.89 94.96 99.19 93.81 97.39 99.11

~) Annite Fe2+/(Fe2+ +Mg) = 0.81: (Ko.93Nao.01 )(Fe2\.1sMg0.soAl0.31 Tio.22Mnom)(AluoSi2.9oOw)(O,OH)2 ) Ferro-pargasitic hornblende Fe2+/( Fe2+ +Mg) = 0.82:

3) (~<lo.4I Ko.3I )(Ca1.ssNao.o7Mno.o4Fe

2+ o.02)(Fe~+ 3.42Mg~.74Alo.36Fe 3+ o.3o Tio.I4Mno.os)(Si6.43Ali.s7 )022( 0, OHhoi

Ohgoclase An24: (Nao.76Cao.24Ko.oi)(Alo.2sSto.7s)AlSt20s ~) Annite Fe2+/(Fe2+ +Mg) = 0.74: (Ko.97N<io.oi)(Fe2+!.94Mgo.7oTio.26Alo.2sMno.D2)(Al!.l1Si2.890IO)(O,OH)2 ) Ferro-edenitic hornblende Fe2+/( Fe2++M~) = 0.75:

G) (~ <lo.9sKo.2s)(Cai.ssNao.07Mno.o4Fe2+ o.02)(Fe. + 3.23Mg ~ .1tAlo.3I Tio.IsFe

3+ o.12Mno.o4)(Si6.s9Al1.4I)022( O,OH)2.2s

0 hgoclase An29: (N ao.71 Cao.29Ko.o2)(A10.30Sto.7o)A1St20s

2.1.2 Local features

To date, a total of 13 boreholes (KR1-KR13 of length 300-1,000m) have been drilled at the study site (Fig. 2-2). Of these, groundwater samples and geological data from boreholes KR1-KR5 are used in this study (Figs. 2-2, 2-3). The distribution of rock types in drill cores is presented in Appendix 1. Currently, additional detailed geological data are available also from other boreholes, e.g. KR6B, KR8-KR13, but at the outset of this study results of groundwater sampling were missing or were disqualified from our studies. Mutual proportions of rock types in the subsurface, as well as the vertical

I7

distribution of rock types, seem to resemble those at ground level. Gradual transition between porphyritic granodiorite and granite has been observed from boreholes KRI, KR2, KR4, KR8 and KR9 (Paulamaki & Paananen I995). Relatively narrow mafic dikes are found in boreholes KR2, KR4, KR5 and KR9. Among the thickest dikes found at the study area is the mafic dike in KR4 at hole distance I80-I88m. As part of mafic rock investigations (Teollisuuden Voima Oy I993b) borehole KR6B was drilled into a gabbroic intrusion (Fig. 2-2). Studies of the drill core and hole showed that the gabbroic rock is quite heterogeneous, altered and fractured (Lindberg & Paananen I992, Front & Hassinen I993 ). Rock type variation in the drill core also indicates that the ground level section of the intrusion is more discontinuous than Fig. 2-2 assumes.

2.2 Structural features

The general structural model of the Kivetty site (Saksa et al. I993) is based on geological, geophysical and hydrological surface and borehole observations performed during preliminary site investigations. The research site and its surroundings are dominated by NW -SE and NE-SW trending fracture zones. Especially the regional NW­SE orientated lineaments indicate this direction to be the predominating fracture zone orientation at the research area. This conclusion is also supported by the fact that the mean orientation of filled, tight and slickensided fractures measured from drilled cores coincides with the NW -SE orientation. The updated structural model (Saksa et al. I996) used in this study is presented in Figure 2-3. This model consists of 29 interpreted fracture zones, I6 of them based on direct petrographical, geophysical and hydrogeological observations (R4, R5, R8-RI3, RI5, RI6, R22, R23, R25-R28), IO on several indirect indications (RI-R3, R7, RI4, RI7, RI8, R2I, R24, R29), and three on a single indirect indication only (R6, RI9, R20). These zones are highly variable in character and some have been considered significant in respect of groundwater hydraulics, while others are significant in respect of rock engineering only.

The current study concentrates on fracture zones R9, RIO, RII, RI2, RI5, RI6, R22 and R23 (Fig. 2-4). According to the previous model (Saksa et al. I993), RII and RI6 have been considered significant in terms of both bedrock engineering and hydraulic conductivity. Structures R9 and RIO are possibly more important in respect of rock engineering and structure RI2 only in respect of hydraulic conductivity. According to the updated model by Saksa et al. (1996), fracture zones R15 and R23 are now considered hydraulically significant, while R22 is regarded as insignificant, although there are uncertainties concerning its hydraulic significance.

2.3 Fracture mineralogy

Recent detailed investigations (Gehor et al. I995) indicate that general features in respect of low temperature fracture minerals in drill core samples from KRI-KR5, KR8 and KR9 are roughly similar. Fracture mineral assemblages are composed mostly of frequent calcite, iron sulphides and iron oxyhydroxides, and less common iron oxides, clay minerals and quartz. There are certain general depth dependencies among fracture minerals.

R28

'lkikr03l

/

Figure 2-3. Structural model (1:10 000- horizontal section Z = 150m) of the Kivetty area with locations of boreholes KR1- KR5 and KR7- KR1 0 (after Saksa et al. 1996). Rock types as in Figure 2-2.

......... 00

19

Figure 2-4. Orientation of fracture zones R9, RI 0, Rll, R12, Rl5, R16, R22 and R23, and boreholes KRJ, KR2, KR4, KR5 and KR9 viewed from SW based on the bedrock model published by Saksa et al. (1993).

Calcite is a frequent filling mineral in fractures. Normally it is missing at depths close to ground level. Usually the first calcite observations are from 50-lOOm below ground level, and thereafter occur widely down to the deepest drilled levels. The one exception is drill core KR9, which seems to lack calcite along its whole drilled length. This may indicate extensive acid surface water recharge in the area around KR9.

Iron sulphides are found from all drilled cores and their behaviour is somewhat comparable to calcite occurrence, i.e. sulphides tend to be missing from the uppermost depth levels. In places small sulphide grains are nucleated and grown onto calcite grains, while in other places iron sulphides occur as fracture seams and coatings. Sulphide observations are more frequent in drill cores KR5 and KR9 than in others. Frequent sulphide observations begin at 540m in KR5, while in KR9 sulphides are found practically along the whole drilled length. KR3 exhibits somewhat peculiar behaviour, lacking iron sulphides in the interval160-590m, not above or below it.

Normally iron oxyhydroxides tend to be confined to the uppermost depth levels, and occur mostly down to 130-170m below ground level (yellowish hydroxides). In KR4 and KR5, however, this oxyhydroxide zone continues down to 400m and 300m, respectively (mostly yellowish hydroxides). In KRl and KR5 iron oxyhydroxides are also found together with iron oxides at intervals 720-820m and 660-850m, respectively (reddish hydroxides). In KR9 oxyhydroxides together with sporadic iron oxides (see below) are found irregularly down to depths of 400m (mostly yellowish hydroxides) and are replaced almost entirely by an iron sulphide assemblage at 160-300m. Colour differences do not unambiguously define iron oxyhydroxides, but normally goethite is more yellowish-brown to red than lepidocrocite, which is brownish to red (Deer et al. 1992). Lepidocrocite is dimorphous with goethite and both minerals commonly intermingle, also with other minerals like hematite, siderite, pyrite, and magnetite.

20

Cryptocrystalline varieties of these admixtures are often called limonite. Both goethite and lepidocrocite can be artificially prepared at room temperature, the former by hydrolysis of ferric and the latter by oxidation of ferrous salts, thus differences in their occurrences are merely chemical in nature.

As previously mentioned, iron oxides are found only at deep levels of boreholes KRl and KR5, and at moderate depths in KR9 (- 380m). These hematite occurrences are likely of hydrothermal origin, although there are exceptional environments such as gossans where hematite crystallisation is possibly simply a result of weathering (Craig & Vaughan 1994 ). Hematite is also found from certain Phanerozoic sedimentary rocks deposited in shallow marine environments, but in these cases hematite has likely recrystallised from iron oxyhydroxides during the compaction, consolidation and lithification processes of a loose sediment.

Clay minerals are frequently obtained at all depths of drilled cores. There is a loose correlation between clay mineral abundance and fracture density. The major clay minerals found are kaolinite and montmorillonite.

Fractures containing silicified infillings (as quartz) are mainly observed in the granites. Silicification seems to be absent in granodiorite. Silicified fracture observations are most frequent in drill cores KR4 and KR9. In KR9 quartz observations correlate with iron oxyhydroxide findings, due to biotite weathering, and both minerals have almost vanished in the interval 160-300m. Distribution of both minerals may indicate that hydraulically conductive zones are cross-cut by borehole KR9. For example R9, intersecting KR9 approximately at 66m, and R16 intersecting at 309m, are potential candidates for such zones, and may partially explain quartz and oxyhydroxide precipitation at hole distance intervals of 0-160m and 300-400m.

21

3. HYDROGEOLOGICAL SETTING

3.1 Hydraulic environment

The Kivetty area is located in the northern part of the Kymijoki river catchment area (Fig. 1-1). Surface waters flow from the study area into Lake Keitele and thereafter, via Paijanne and Kymijoki, to the Gulf of Finland. Gradients in topography are gentle, except in the eastern part of the study area (Fig. 2-2), where altitude differences reach about 15m over a lOOm distance. The highest point in the area is Kilpismaki, at about +219 metres above average sea level (m.a.s.l.), and the lowest at around+ 154 m.a.s.l. in the westernmost part of the study area. Essential parts of the study area for site investigations are located between + 160 and + 180 m.a.s.l. The Kilpismaki hill acts locally as a watershed (Saksa et al. 1993); thus in the study area surface waters tend to run, following local topographic depressions, towards SE and SW. Given the late Quaternary history of the Baltic Sea, it seems that saline sea water may have had only minor effects on groundwaters at the Kivetty site. Before the latest glaciation (Weichsel) the research site was above the water level of the interglacial Eem Sea (Eronen & Lehtinen 1996), surrendering to the Weichsel glacial ice sheet at least 75,000 years ago. Glaciation continued throughout the Baltic Ice Lake and the Yoldia Sea stages (Eronen 1988). About 9,700 years ago, during the late Yoldia and early Ancylus Lake stages central parts of Finland were released from the ice cover and the research site or parts of it may have been below water level for a relatively short period, but following a rapid land uplift soon emerged as an island. According to Eronen et al. (1995) the research area was above the shoreline during the saline Litorina Sea stage. Currently the post­glacial land uplift at the study site is about 6 mm/a (Kakkuri 1987).

The bedrock of the study area is covered mostly with till and relatively abundant peatlands, especially in SW parts of the study area (Anttila et al. 1992). Bedrock exposures cover only about 2% of the area. Peat bog depressions run mostly in a NW­SE direction. There are also several ponds and small lakes in the study area and its surroundings. The thickness of till cover varies from 1.5m to 23m, but most frequently is in the range 3-6m. In the bog depressions, a l-2m section of peat is underlain by a till layer. The small lakes, ponds and peat bogs indicate that the groundwater table is close to ground level.

The lithology, hydraulic conductivities, fracture zones, hydraulic heads and groundwater sampling sections (coded as Tl-T8 and BT, i.e. multi packed-off samplings during '93-'94 and a double packed-off sampling in '94, respectively) for boreholes KR1-KR5 are shown in Appendix 1 (cf. also A pp. 6, Table 3-1 p. 27). Section lengths for groundwater samples are long, with a median value of 50m. The deep interval of KR5 is especially long, being clearly over lOOm in length. Evidently the long packed-off interval is prone to mixed samples of different types of groundwater and to uncertainties of the source depth of extracted samples. Fracturing of the drill cores has been compared with the results of hydraulic conductivity tests to identify water-conductive fractures (Kuusela­Lahtinen & Front 1990, Vahanne & Front 1990a, 1990b, Melamed & Front 1995). The latest hydraulic tests involved difference flow measurements with 2m intervals

23

3-5. The westemmost borehole KR2 is located in a topographic depression with respect to KRl and KR9, and there is only one modelled structure connecting KRl and KR2, namely Rl5 with significant hydraulic importance (Saksa et al. 1996). The hydraulic head field in structure Rl5 is shown in Appendix 5. The easternmost KR3 has been drilled towards NNE from a local topographic elevation maximum and therefore is hydrogeologically isolated from the other boreholes. According to hydraulic interference tests borehole KR8 is likely connected to KRl via structure R26 (Saksa et al. 1996). This structure was not identified at the time of evaluation of the conceptual flow model. Therefore, in the following, KR8 is dealt with as a separate case. Simple geometrical considerations confirm, however, that groundwater flow from the near ground level parts of KRl to the deep parts of KR8 is impossible, although the reverse is probable.

Modelled groundwater flow in the structures R9, RIO, Rll, Rl2, Rl6, R22 and R23, shown with flow direction arrows, are illustrated together with boreholes KR1, KR4, KR5 and KR9 in Figs. 3-1 and 3-2. Flow directions exhibit moderate or low vertical gradients in all structures except in R9 where steep gradients from surface to depth exist in the NE part of the fracture zone. Appendices 3-5 show that also structures R11, R12, Rl5, R22 and R23 exhibit significant gradients although there is no remarkable vertical component in the variation of hydraulic head values at depths of present interest.

According to the modelled flow directions (Fig. 3-1) the deep packed-off intervals (Tl­T3) of KR1 receive their waters mainly from fracture zone R9, and it is apparent that these waters should evolve from water compositions similar to shallow groundwaters around KR9. The compositional changes between water types should occur mostly because of infiltration of the recharged water along R9. Similarly, the packed-off intervals T4 and T5 of KR1 seem to receive their waters via routes R9-R11 and R9-Rl0, respectively (Figs. 3-1 and 3-2), and no significant compositional differences are expected compared with samples from Tl-T3. The packed-off intervals T6 and T7 of KRl are already relatively near ground level, so the effect of surface fracturing and increased matrix flow cannot be omitted. The interval T6 is quite near the fracture zone R22, and T7 is cross-cut by this zone. Therefore waters from these intervals should be quite similar.

The deepest representative water samples from borehole KR4 are from packed-off intervals T1 and T2. There are, however, no structures in the bedrock model which could be used to interpret the sample evolution. The packed-off interval T3 is cross-cut by the fracture zone R23 and according to modellings it is possible that these waters may be from KR1/T2-T3 (Fig 3-1). As in the case of Tl and T2, the packed-off intervals T4 and T5 of KR4 do not intersect any fracture zones described in the current bedrock model, though water samples are available from these intervals. There are, however, certain indications that structures R23 and Rll are actually parts of the same fracture swarm (Luukkonen et al. 1996), and intervals T4 and T5 may be expected to have received their waters from a region roughly represented by KR1/T4. The packed­off intervals KR4/T6 (Fig 3-2) and KR4/T7 (Fig. 3-1) are cross-cut by structures R12 and R11, respectively. Both intervals are already near ground level and effects of surface fracturing cannot be discarded. Otherwise, according to flow modellings these intervals receive their waters from higher potential regions via infiltration along the fracture zones.

24

Figure 3-1. Modelled flow directions of groundwater in fracture zones R9, Rll, R22 and R23 based on the study by Taivassalo & Meszaros (1994).

25

Figure 3-2. Modelled flow directions of groundwater in fracture zones R9, RIO, Rll, R12 and R23 based on the study by Taivassalo & Meszaros (1994).

26

Figure 3-3. Modelled flow directions of groundwater in fracture zones R15 and R22 based on the study by Taivassalo & Meszaros (1994).

The long packed-off interval (735-853m, BT) in the bottom of KR5 possibly receives groundwaters from many sources. The most likely candidates for flow routes from a standpoint of hydraulic conductivities run via R9-R11 and R9-R22 (Fig. 3-1). In the former case the bottom interval BT should receive its waters from KR1/T4, and in the latter the waters in BT have evolved from shallow groundwaters somewhat similar to KR1/T7. Less likely possibilities for water sources of interval BT are routes R9-R23 (Fig. 3-1) and R9-R10-R22 (Fig. 3-2). In these cases the waters of BT should have developed from waters similar to KR1/T1-T3 or KR1/T4. The packed-off intervals T1-T5 of KR5 have not been assigned reliably to any modelled fracture zones, although there is an uncertain assignment of T4 to the structure R24 (Fig. 2-3, Appendix 1). The near surface packed-off intervals T6 and T7 of KR5 probably receive the major part of their waters from structures R8 and R24, and via incoherent subhorizontal surface fracturing.

It is evident from above and from Appendix 1 that many fractured and hydraulically conductive parts in KR4 and KR5 are without structural interpretation. A striking feature of these boreholes is that hydraulic head values are the same in both boreholes and remain similar as a function of depth (Table 3-1). It is therefore possible that fracture zone R24 (Fig. 2-3) or a fracture zone swarm related to R24 is responsible for most structural features encountered in both boreholes. If this is the case, water samples from KR4 and KR5 should relate more closely to each other than to samples from KR1.

The deepest parts (T1-T3) of the westemmost borehole KR2 fall on relatively intact bedrock, thus no water samples are available from the packed-off intervals T2 and T3. The sample from interval T1 should be a relatively good example of matrix effects on water composition in bedrock. The packed-off intervals T4 and T5 may be related to structure R13 (App. 1), although strictly speaking only T4 falls on this structure. In view of the hydraulic gradients, water flow in R13 should be relatively slow and

27

comparatively long infiltration times may thus be expected. The packed-off intervals T6 and T7 of KR2 are relatively near ground level, therefore the effect of surface fracturing may be important. In any case, intervals T6 and T7 are also related to the fracture zone R15 (Fig. 3-3) and according to flow modellings possibly receive part of their waters from KR1/T6. The surface topography (Fig. 2-2) and groundwater table also suggest some relation between the intervals T6 and T7 and the shallow groundwaters around KR9.

Table 3-1. Interpreted groundwater flow between packed-off intervals of boreholes KR1-KR5 and KR9 based on the conceptual model by Taivassalo & Meszaros (1993, 1994). Acronyms follow the convention: borehole id /interval id/head value. ({No interpretation '' indicates flow from unidentified structures and bedrock matrix. In multi packed-off boreholes, packed-off intervals are labelled T1-T7 (cf App. 1). Other intervals are labelled TP near ground level (KR9) and BT (KR5) in the deep parts of boreholes, both being double packed-off intervals from open boreholes. Hydraulic heads are expressed in metres (after Saksa et al. 1993, Rouhiainen 1996a). NoH indicates that no head value is available. Packed-off intervals with no available hydro­geochemical results are shaded.

KRl

KR2

KRJ

KR4 Surface -> Rll -> KR4mll68.4 Surface -> R12 -> KR4ff6/168.4

KR1ff41169.3 -> R?? -> KR4ff51168 .2 KRlff4!169.3 -> R?? -> KR4ff41168.4 KR1ff3!169.7 -> R23 -> KR4ff3!167.7

No interpretation KR4ff21168.2 No interpretation KR4ffl/168.5

KRS Surface -> R24 -> R8 -> KR5ff71169.4 Surface -> R24 -> R8 -> KR5ff6/168 .5

No interpretation KR5ff5!168.8 No interpretation KR5ff41168.8 No interpretation KR5ff3/168.8 No interpretation KR5ff2/168 .7 No interpretation KRS!fl/168 .8

KR1ff4/169.3 -> Rll -> KRS/BT/167.6

28

The deepest samplings (T1 and T2) from the isolated borehole KR3 are, according to the updated bedrock model (Saksa et al. 1996), percolated from the fracture zones R27 and R28 (Appendix 1). Unfortunately, as the hydraulic properties and connections of these structures are very poorly known, the study in respect of groundwater flow is merely speculative. In the previous bedrock model (Saksa et al. 1993) the packed-off intervals T3-T7 of KR3 are not connected to any known structures. As regards hydraulic conductivities (Appendix 3) it seems that waters in the intervals T3-T5 are mostly the result of matrix flow in the bedrock. It is also worth noting that sampling from T4 was unsuccessful, presumably because of very low water yield. The packed-off intervals T6 and T7 of KR3 are relatively near ground level and probably represent bedrock matrix flow partially related to surface fracturing.

A simplification of groundwater flow interpretations based on the conceptual model of Taivassalo & Meszaros (1993, 1994) is shown in Table 3-1. Some of the features discussed above are not included for the sake of clarity. For example, in the case of sample KR5/BT five possible routes were considered but only one is tabulated. All unidentified structures and matrix flow possibilities discussed and interpreted above are merely given the indication: "no interpretation". As a gross generalisation of the conceptual flow model interpreted above, our study concentrates on three jlowpath hypotheses. In the vicinity of boreholes KR1, KR4 and KR5 there are several potentially conductive fracture zones. The topography in this area indicates that water flow should be from borehole KR1 to holes KR4 and KR5 (jlowpath I- cf. Fig. 3-4). Borehole KR2 is located away from other boreholes in a moderate topography and is therefore considered as a separate case (jlowpath 11- cf. Fig. 3-4). Finally, KR3 is isolated by Kilpismaki into its own case (jlowpath Ill).

3.3 Groundwater flow indications based on transverse flow measurements

At Ki vetty actual ground water flow directions and intensities have been measured in boreholes KR1, KR3, KR5 and KR8 with a flowmeter in two metre sections (Rouhiainen 1996b). The results are extracted in Table 3-2 where classification into fracture zones and normally fractured bedrock is based mainly on the updated bedrock model of Saksa et al. (1996) and its preliminary memo (Saksa 1994). Measured matrix flows in normally fractured bedrock are in the range 0-10 ml/h. According to Rouhiainen (1996b) the sensitivity of the flowmeter is better than 1 ml/h for a cross-hole flow corresponding to a flux of about 2·10-9 m/s. This is, compared to studies of Taivassalo & Meszaros (1994), a relatively high value for a matrix flow if deep levels of bedrock are considered. Measured flow values in the structures vary from almost no­flow ( <1 ml/h) to strikingly high flow conditions (500 ml/h).

Groundwater flow (82 ml/h) in fracture zone R8 measured from borehole KR5 (Fig. 2-3) is northward. The flow direction in R10 measured from KR5 is similarly towards NNW, and its magnitude is of the order of 500 ml/h. Flow measurements from R11 are from KR1 (105 ml/h) and KR5 (130 and 180 ml/h), and according to these groundwater flows should be towards NNW in the structure. In the fracture zone R12 the flow (240 ml/h), measured from KR5, is northward following the structure. Based on

29

measurements in KR1, moderately (50°) northward-dipping fracture zone R15 also conducts ground water northward at about 7 5 ml/h.

Measured groundwater flow directions in the NW-SE oriented structures R8 and R10-12 (Fig. 3-5), and in the E-W R15 all conflict with the conceptual groundwater flow model ofTaivassalo & Meszaros (1993, 1994). Furthermore, as Fig. 2-2 shows, all these measurements indicate an "uphill" flow at depth if compared with the regional ground level topography. When measurements from fracture zones at shallow depths (Table 3-2) are compared with Fig. 2-2, a more logical connection appears between

- topography and flow directions. If direction measurements at depth cannot be attributed to as yet undiscovered deficiencies in the test arrangement, these measured flow directions indicate a hydraulic sink north of the measured locations. This still unidentified structure should guide waters from structures R8 and R10-R12 to the surface over an area with a relatively low altitude. SW parts of the study area in particular are good candidates for such a discharge area.

Figure 3-4. Generalised 3-D illustration of the conceptual hydrogeological model with three potential flowpaths. Flow directions in the fracture zones and bedrock matrix are shown with red and blue arrows, respectively.

30

Table 3-2. Directions and intensifies of groundwater flow in fracture zones R8, RIO­R12, Rl5, R22 and R26-R28, and in normally fractured blocks of bedrock (matrix flow) at the Kivetty site measured with a transverse flowmeter in two metre sections from boreholes KRJ, KR3, KR5 and KR8 (Rouhiainen 1996b). Interpretations of structural relations are based mainly on Saksa et al. (1996) and Saksa (1994).

Matrix flow NNW 538.73 KRl Matrix flow 2 NW 106.00 KR3 R?? 55 KRl Matrix flow no flow 438.55 KR3 R27 <1 KRl Matrix flow <1 482.62 KR3 R28 52 E KRl Matrix flow 4 SW 112.36 KR5 R?? 250 SW Matrix flow 7 SE 252.60 KR5 R8 82 N Matrix flow <1 507.01 KR5 Rl2 240 Matrix flow 577.11 KR5 RIO 500 741.36 KR5 R11 130 N 755.38 KR5 R11 180 276.65 KR8 R26 22 SW KR8 Matrix flow 2 SE 276.65 KR8 R26 10 SW 114.37 KR8 Matrix flow 2 SE

120.38 KR8 Matrix flow <1 156.45 KR8 Matrix flow <1 160.46 KR8 Matrix flow 9 ESE 160.46 KR8 Matrix flow 7 SE 220.56 KR8 Matrix flow 7 WNW 248.61 KR8 Matrix flow 2-6 NE 252.61 KR8 Matrix flow 2 NNW 258.62 KR8 Matrix flow 4 NNW 258.62 KR8 Matrix flow 2 NW

KR8 Matrix flow 2 N KR8 Matrix flow 2 N KR8 Matrix flow 3 E KR8 Matrix flow 3 s KR8 Matrix flow <1

If groundwater flow measurements are considered reliable, several adjustments should be made to Table 3-1. This concept would revoke also our jlowpath generalisation between KRl, KR4 and KR5 (jlowpath I, cf. p. 28). Packed-off intervals at deep levels (T5-T2) in borehole KRl should not get their waters from along structure R9 but mainly along a regionally dominating NW-SE fracture zone swarm (e.g. R8, R10-R12 and R23). Waters from KR4/T7-T6, and even from KR5/BT would be good precursors of waters at KRl/ T5-T2. Furthermore, a modified flowpath interpretation for samples from KR4/T7-T6 and KR5/BT gives no explicit solution. Similarly, if the measured northern flow direction in R15 is extended to a general feature of the fracture zone, the waters from KR1/T6 and KR2/T7 cannot be related to each other. In this case it is likely that the packed-off interval T7 in borehole KR2 inherits most of its chemical character form surface waters surrounding the borehole.

31

Figure 3-5. Measured flow directions of groundwater in fracture zones Rl 0, Rll and R12 (cf Table 3-2).

32

Groundwater flow directions in the normally fractured bedrock in borehole KR1 are, at hole distances greater than 200m, towards N-NW and towards SW and SE, although groundwater flow simulations (Taivassalo & Meszaros 1993, 1994) assume that matrix flow around KR1 is towards SW. According to flow simulations matrix flow around holes KR5 and KR8 should also be mainly towards SW. Flow measurements show, however, that at hole distances greater than 200m matrix flow directions are mostly towards E, NNW and NE. The NW flow direction measured from KR3 also conflicts with the modelled hydraulic head field. Especially eastern and northeastern matrix flow directions in holes KR1, KR5 and KR8 are in disagreement with modelled groundwater flow. These results indicate either clear deficiencies in the test arrangement, or that causes for the flow directions are unidentified structures and predominating fracturing orientations, where the general head field in the area is revoked.

In all, however, simulation results are based on a considerable amount of background data compared with relatively few flow direction measurements. We therefore consider the working hypothesis of the three jlowpaths presented on page 28 viable, at least to start with.

33

4. STRATEGY OF GEOCHEMICAL MODELLING

The modelling approach in this investigation can be described as follows: Sampling points along a flowpath in a groundwater system will show progressive changes in chemical and isotopic properties. These changes reflect geochemical processes such as mineral dissolution and precipitation, ion exchange, gas exchange and redox reactions. The goal in chemical modelling is to identify the probable processes and to predict the quantity of material transferred that is responsible for the evolution of the groundwater, 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 four step procedure:

- The first step is to analyse the trends of hydrochemical and isotopic data (Sections 5.4 - 5.6) that constrain the possible processes behind the trends.

-The second step is to find the tendency of the groundwater to dissolve or precipitate minerals as reflected by the fracture observations and saturation indices.

- The third step is to derive reaction models that can explain the changes in water chemistry between any points along a chosen flowpath by mass-balance calcu­lations.

- The fourth step is to test the thermodynamic feasibility of reaction models derived from the mass-balance calculations.

Proceeding to the second step it is necessary to know the thermodynamic factors that control the composition of the sampled waters (App. 8). The speciation-solubility code EQ3NR in the software package for EQ3/6 version 7.0 (Wolery 1992) with its supporting thermodynamic data base DATAO. COM was used for the equilibrium-specia­tion calculations.

The saturation index (SI) of a particular mineral that may be reacting in the system is defined as

SI =Log IAP/KT ,

where lAP 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. The KT values are also needed for constructing activity diagrams (for further detail cf. Nordstrom and Munoz 1986). The diagrams give important information on potentially stable silicate phases after incongruent weathering sequences of primary silicate minerals in certain groundwater solutions.

The NETPATH program version 2.0 (Plummer et al. 1994) is used to interpret net geochemical mass-balance reactions between an initial and final water along a hydrologic flowpath in the third step. The program utilizes defined chemical and isotopic data for groundwater samples and, calculates those mixing proportions of initial waters and reaction coeficients of chemical sinks and sources (minerals and gases) that account for the observed changes between an initial and final water. In addition the

34

program calculates isotopic evolution according to each mass-balance model to predict isotopic composition at the end point of flowpath.

In the final and fourth step thermodynamic mass-transfer calculations are used to test the thermodynamic feasibility of reaction models derived from the mass-balance calculations. Mass-transfer calculations can better describe where equilibrium is not reached and kinetic information is needed. The forward method, using i.e. the EQ3/6 program (Wolery 1992), is valuable in predicting reaction paths and the composition of equilibrium solutions for hypothetical reactions and conditions (Plummer 1984, Nordstrom et al. 1990).

4.1 Consistency checking between hydrogeochemistry and hydrogeology

Successful modelling performance should give information on long term hydrogeological progress (flowpaths and residence time) at the site (e.g. Plummer et al. 1990, Pitkanen et al. 1996b, Blomqvist et al 1998). The results of geochemical modelling are used to assess the consistency between hydrogeochemistry and hydrodynamic models as part of the integration process to conceptualise the movement of groundwater through the hydrogeological system. The schematic flow chart below (Fig. 4-1) describes the strategy for testing the hydrodynamic model by site scale geochemical information. It is in principal based on the geochemical modelling strategy developed by Plummer et al. (1983).

Potential flowpaths (upper box) are based on the conceptual flow model (Chapter 3). Geochemical modelling (lower box) is uniform with the strategy described in the previous chapter but is divided in steps clarifying hydrogeological information.

Potential flowpaths from conceptual flow model

Comparison of the consistency of hydraulic and chemical

flowpaths, flow velocities and C-14 ages

t Interpretation of chemical processes during the groundwater evolution

Definition of initial and final waters for the single flow step of reaction modelling along

chemical flowpaths

Using carbon isotope data as additional criteria for testing reaction models and adjusting C-14 ages of flow paths

Figure 4-1. Schematic flow chart describing the process for inferring flow information from geochemical data and the consistency between hydrogeochemistry and the hydrogeological model (Korkealaakso & Pitkanen 1997).

35

5. GROUNDWATER CHEMISTRY

5.1 General

The hydrogeochemistry of groundwater samples collected in the Kivetty area has been discussed in several studies (Pitkanen et al. 1992, Lampen & Snellman 1993, Snellman et al. 1995a, Tuominen 1995, Ruotsalainen & Snellman 1996). 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.

Groundwater data used in this study are given in both mgll and mmol/l in Appendix 6. The sampling methods, analytical procedures and evaluation of representativeness of the samples are discussed in the two following chapters (5.2-5.3). The hydrogeochemical features are evaluated in chapters 5.4-5.6.

As chloride is usually considered a conservative hydrogeochemical parameter, most of the graphs present variations of other parameters relative to it. In the graphical presentations the boreholes are arranged according to the jlowpaths defined in chapter 3. The environmental water samples from springs are coded as "Spring" and those from domestic borehole wells or wells at the Kivetty investigation site for flushing water production during the drilling activities as "Well".

5.2 Sampling and chemical analyses

The preliminary site investigations for spent nuclear fuel disposal started in the Kivetty area in the late 1980s. The hydrogeochemical site investigations (Lampen & Snellman 1993) produced a general characterisation of the local deep groundwaters, precipitation, surface waters and shallow groundwaters in the surrounding area. Groundwater samples from the deep, open boreholes were taken with a double packer, bladder pump instrument (Rouhiainen et al. 1992).

During the detailed site investigations in 1993-1994 groundwater samples were taken from five deep boreholes (KR1-KR5; inclined, 500-1,000m long). In order to prevent mixing of different bodies of groundwaters, multipackers were installed in the Kivetty boreholes during 1992 (Hinkkanen & Oksa 1992). Thus a hydrological steady-state could be reached before the groundwater sampling began in 1993.

Groundwater samples were pumped with a slim membrane pump (Ohberg 1991, Ruotsalainen et al. 1994) from the packed-off, 20-102m long sections of the deep boreholes. Seven sampling sections (T1-T7) were isolated with inflatable rubber packers in each borehole. The hydraulic head in each section was constantly monitored for possible effects of sampling (Niva et al. 1994).

Groundwater flowed from the conductive fractures in the sampling section through polyamide tubes to the surface, where the quality of the groundwater was monitored by continuous on-line measurements of pH, electrical conductivity, Eh(Pt), dissolved 0 2

and temperature in a flow-through cell. Since the capacity of the membrane pump was

36

quite small (max. 100 ml/stroke) and the natural hydrological conductivity of sampling sections was often very low (usually 10-8-10-6 m/s), the actual pumping rates ranged between app. 5 and 50 ml/min. Thus the pre-pumping period, before the groundwater had reached an expected representativity, took generally 3-5 weeks, during which time roughly 100-3,500 l was pumped from each of the sampling sections. About 1 week before the planned sampling the amount of remaining drilling water (based on uranine, an organic dye used as a tracer) and tritium activity were checked to evaluate representativity and possible mixing with younger groundwaters. The criteria for representativeness are given in Table 5-1.

Table 5-1. Criteria for representativeness of groundwaters (Ruotsalainen & Snellman 1996)

Parameter(s) Applied criteria in evaluation of representativeness Void volume in sampling section and Water changed at least 2-3 times where possible tubes Remainin_g drilling_ water <2.5% Field measurements of electrical Stabilised values conductivity and pH Field measurements of Eh and 0 2 Stabilised and as low as possible values Tritium If no hydrogeological cause (fracture zone), below the

detection limit of the direct activity count (6-8 TU)

Groundwater samples were shielded as effectively as possible from atmospheric contamination. Most of the samples were collected in vessels enclosed in plastic bags filled with nitrogen gas, and preparations (filtering (0.45 J.tm), complexing, pH adjustment etc.) were made in an N2 filled glove-box in the field laboratory (Ruotsalainen et al. 1994).

Parameters that suffer most from the effects of atmospheric gases (ferrous iron, sulphide, alkalinity and acidity) or transport delays (anions, uranine) were analysed in the field laboratory. Some anions were also analysed in an off-site laboratory. Other samples were preserved where necessary and sent to different laboratories in Finland or abroad.

Reference samples from springs, domestic borehole wells and borehole wells for flushing water production at the site were sampled either with a Ruttner sampler or directly from the water line. Contact with atmospheric gases could not be avoided. The preparation and analytical methods applied for these environmental reference samples have been reported by Lampen & Snellman (1993), Ruotsalainen et al. (1994), Tuominen (1994), and Ruotsalainen & Snellman (1996).

Analytical programmes were adapted according to the representativeness of the samples and aims of the study. These are detailed in Table 5-2.

37

Table 5-2. Analytical programmes

Parameters Deep groundwaters Springs and shallow e;roundwaters

Measurements and pH, electrical conductivity, Eh, 0 2, pH, electrical conductivity, analyses in the field temperature, alkalinity, acidity, uranine alkalinity, acidity laboratory (on-site) (=tracer of the drilling water), Cl, Br, F,

S2-tot• S04, NOz, N03, P04, Fetot• Fe2+

Main physico-chemical pH, electrical conductivity, density, Same as groundwaters variables (off-site) COD, KMn04, Fetot• Stot. Ntot. Ptot• Btot•

DIC (=Dissolved Inorganic Carbon), DOC (=Dissolved Organic Carbon), Si02, uranine

Cations (off-site) Na, K, Ca, Mg, Mn, Al, Sr, NH4 Same as groundwaters Anions (off-site) Cl, Br, F, I, N03, NOz, S04, P04 Same as groundwaters Trace elements (off-site) Rb, Ba, Cs, Li Same as groundwaters Evacuated and dissolved Oz, CO, COz, Hz, Nz, CH4, CzHz, CzH6, -gases (off-site) He,Ar Isotopes in water (off- H-2(H20), H-3(H20), 0-18(Hz0), H-2(H20), H-3(H20), 0-site) 0-18(S04), S-34(S04), 18(H20)

S-34(S2-), C-13(DIC), C-14(DIC),

0-18(C02), 0-18(S04), S-34(S04), C-13(DIC), C-14(DIC), Rn-222, C-13(C02 and CH4), Sr-87/Sr-86, U-234/U-238, Rn-222, Th-232, Th-228, Th-230, Ra- U-238 226, Ra-228, Sr-87/Sr-86, U-234/U-238, U-238

Isotopes in particles (off- U-234/U-238, U-238 U-234/U-238, U-238 site)

5.3 Representativeness of hydrogeochemical data

The representativeness of hydrogeochemical data from the five deep, multipackered boreholes at Ki vetty and from springs and shallow bore hole wells, and of data from the preliminary investigations during 1987-1992, was evaluated using the procedure developed by Ruotsalainen & Snellman (1996). Of a total of 118 water samples taken during 1988-1992 from the Kivetty site or its vicinity, 46 samples were from deep boreholes, 38 from shallow sampling points (springs, wells) and 34 from precipitation (rain, snow). Twenty-nine samples passed the criteria of the evaluation procedure (Appendix 7) and are those primarily considered in this modelling study. Nine other samples were also included; these are shown in Appendix 8 with the reasons for their failing the representativity criteria. All nine samples had charge balances deviating from the set limit of ±5% but all were still within the ±8% range. Samples from the lowest level of KR2 (KR2/T1) had roughly 5% remaining drilling water, compared with the set limit of 2.5%. These minor quality defects were not considered to endanger general hydrogeochemical evaluations of the Kivetty data. The Kivetty groundwaters are very dilute, giving analytical errors an emphasised effect in charge balance calculations.

The results of hydrological pressure monitoring for samples KR2/T1 and KR3/T5 indicated some leakage of the multipacker instruments (Niva et al. 1994). The lengths of these sampling sections were 50m and 65m respectively, thus only very prominent

38

leakages from neighbouring sections would cause severe m1x1ng. Sample KR2/T1 reflects only slight mixing with younger groundwaters (H-3 4 TU and C-14 10.2 pmC). KR3/T5 shows clear indications of recent waters (H-3 24.9 TU and C-14 63.1 pmC) but has been included in the data base to give some idea of near-surface infiltration conditions.

Due to the very low hydraulic conductivities of many sampling sections, and the low capacity of the slim membrane pump, transport times for some groundwater samples could be several hundred hours (Ruotsalainen & Snellman 1996). The part of the polyamide sampling tube above groundwater level was in direct contact with air. Diffusion of atmospheric gases through the walls of the sampling tube (Snellman et al. 1995a) is an obvious reason for the observed anomalously high Eh values measured. This may also have affected oxygen and C02 sensitive parameters, mostly alkalinity, acidity, evacuated gases, Fe2

+, Mn, S2-, C-13 and C-14. Already in the previous

hydrogeochemical investigations (Lampen & Snellman 1993) the Kivetty groundwaters were characterised, using different sampling equipment, to have moderate pH and alkalinity. Thus no strong effect of diffusion on C02 sensitive parameters is expected from the present data either. However, the results of gas analyses (Appendix 6) may have notable uncertainties due to pressure changes and loss of especially light gases (e.g. H2, He) during pumping to the surface. The dissolved 0 2 results also often conflict with the trends of field measurements of Eh and 0 2 . This is apparently due to technical problems in sampling or leakages of the glass ampoule sample vessels during laboratory analyses. Only Eh measurements considered most reliable have been included into this study (Ruotsalainen & Snellman 1996).

5.4 Hydrochemical features

5.4.1 Principal components of groundwater compositions

Recently, principal component analyses have been tested for explaining variations in groundwater compositions (e.g. Laaksoharju & Wallin 1997). Basically, these analyses should be used to find a new set of independent variables oriented at right angles to each other, reduce noise among variables, identify outlier samples, and find and group similar samples in a study matrix. If the purpose of a principal component analysis is to present general regularities hidden in the data, the study matrix should not contain outliers. If the outlier samples are not removed, part of the principal components are potentially weighted in directions caused by these individual samples. 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 explain at least 75%, 90%, 95% or even 99% of the variances of the original variables (Joreskog et al. 1976, Reyment & Joreskog 1996). This explanation ratio increases as the number of principal components considered significant increases, and as the number of variables considered relevant diminishes. Depending on the case, the number of significant principal components may be fixed first, and thereafter the explanation ratio can be increased by dropping insignificantly

39

loaded variables from the study. If all variables are considered important for the study, the explanation ratio is increased by taking more principal components into the analysis.

In this study the choice and processing of hydrochemical variables is similar to studies done with material from the Aspo Hard Rock Laboratory (Laaksoharju 1995, Laaksoharju et al. 1995, Laaksoharju & Wallin 1997), i.e. Na, K, Ca, Mg, HC03, Cl, S04, H-3, 8H-2, 80-18 concentrations are used. It is assumed that most of the variability related to groundwater composition can be described with these variables. In the case of the Aspo studies, extreme groundwater compositions found with principal component analyses are termed brine, glacial, glacial melt water, meteoric, 1960s precipitation, altered marine, sediment pore water, and Baltic and Litorina Sea end-members. Due to the small variability of groundwater compositions at the Kivetty site this approach is inappropriate. For the same reason, principal component analyses of Kivetty samples run together with Aspo end-members turn out infertile because all Kivetty samples focus on one spot in the near vicinity of the meteoric water and glacial end-member compositions of Aspo. Furthermore, it may be questioned whether the analysed end­member compositions of the Aspo site are representative of the named reference waters throughout the Fennoscandian Shield area, e.g. environmental isotopes depend on the geographical location. Therefore the following principal component analysis is carried out with the Kivetty samples only.

The loadings of chemical variables onto principal components are shown in Table 5-3. For present purposes an explanation ratio of at least 75% of the original variance was considered adequate, and this condition is fulfilled with three principal components (77.7%, cf. Table 5-3), all of them exhibiting significant eigenvalues clearly greater than one (cf. Davis 1986). Table 5-3 shows that with this explanation ratio all variables are heavily loaded onto one of the three principal components, and therefore give their significant contribution to the components in question. Variable communality, i.e. the part of each variable variance in common with the other variables, which is explained in this three component model, is also shown.

The first principal component PI describes mostly the salinity vanatlons of groundwaters, and possibly mixing between meteoric and saline waters, but it also contains a climatic signal: with increasing salinities isotope ratios 8H-2 and 80-18 tend to decrease (cf. Table 5-3). Therefore it is also possible that PI describes mixing between meteoric and glacial waters and a sub- or preglacial saline source. Concentration increments of S04 are related positively to salinity variations. The second principal component P2 describes mainly alkalinity changes in groundwater samples, and also compositional variations of variables K, Ca and Mg are positively related to this component. It seems that weathering and the effect of groundwater hosts are partially reflected within this principal component. The division of cations into different principal components suggests different sources for them, e.g. sodium is connected to salinity enrichment and the other cations to water-rock interaction. Tritium is the only variable significantly loaded onto principal component P3. The positive correlation of potassium with tritium may indicate the short residence of potassium in groundwater particularly after recharge, i.e. the weathering zone.

40

Table 5-3. Principal component loadings of geochemical variables in a varimax rotated principal component model, based on samples in Appendix 6. Absolute loadings above 0. 65 are shown in bold type. Component scores of Fig. 5-J are calculated with the aid of tabulated component coefficients. As an example, a score for the principal component PI is calculated as follows (coefficients from the last part of this Table): PI= -9.5694 + 0.0175x[Na+] + 0.1812x[K] + 0.0068x[Ca2+] - 0.0592x[Mg2+] - 0.0030x[HC03-] + 0.0387x[Ct] + 0.19Jlx[SO/-] + 0.0118x[H-3]- 0.0569x[8H-2] -0.2606x[80-18].

Varimax rotated principal component loadings Pr p2 p3 Communality

Na+ 0.656 -0.030 -0.578 0.766 K+ 0.224 0.779 0.405 0.821 Ca2+ 0.288 0.812 -0.298 0.832 Mg2+ -0.220 0.900 0.070 0.863 HC03- 0.007 0.814 -0.462 0.877 er 0.838 0.018 -0.049 0.705 SO/ 0.794 -0.158 0.244 0.716 H-3 -0.110 -0.078 0.884 0.799 8H-2 -0.780 -0.152 0.208 0.675

80-18 -0.738 -0.272 0.319 0.720

Eigenvalues 3.111 2.869 1.794

Percent of total variance explained 31.11 28.69 17.94

Cumulative% 59.79 77.73

Component coefficients for non-standardised variables Pr p2 p3

m -9.5694 -3.5883 0.7998 Na\mg/1) 0.0175 -0.0086 -0.0314 K+ (mg/1) 0.1812 0.4947 0.5587 Ca2+ (mg/1) 0.0068 0.0483 -0.0155 Mg2+ (mg/1) -0.0592 0.2099 0.0523 HC03- (mg/1) -0.0030 0.0091 -0.0080 er (mg/1) 0.0387 -0.0030 0.0115 sol- (mg/1) 0.1911 -0.0367 0.1495 H-3 (TU) 0.0118 0.0061 0.0736 8H-2 (%o SMOW) -0.0569 -0.0036 0.0026

80-18 (%o SMOW) -0.2606 -0.0640 0.0958

Principal component scores for samples studied are shown in Fig. 5-1 and coefficients for linear regression are tabulated in Table 5-3. The plots show some extreme compositions from the data. The bottom sample from KR5 (5/BT in Fig 5-1) stands out to the right with two other Cl rich samples, but these do not identify any coherent end­member composition. Surface waters are clearly distinguished. The upper points in the left plot (Fig 5-la) represent the most carbonate rich samples. In other words they may represent the culmination point of carbonate evolution where calcite saturation is reached, i.e. the most extreme sample of the end-member representing carbonate-water­rock interaction. The lack of coherent end-member compositions is considered to result from low concentration differences in the data. In this case occasional reactions along a single flowpath may exceed the significance of conservative mixing of the saline source.

a) 2.5

1.5

0.5

r£

-0.5

-1 .5

-2.5 -3.0

1m .;. .~.1m

Liimata i1 el).

~rr~1rr6 .;. fT6

4 .;. .;. • KA1

~ KA1 4fT3 .~.Jm 2!T1

.;. 5!T6 ~~ tn

~I .. ltt41 ~r~ 4{ 6 .;. tMJ

~!T1.;. ~~ 3fT5 A5ffi• ~~ 3{ 1

KA2• .;.4 In .;. ~......... ... 1m

1fT4.._ .. '' • J/I.:J

.t2{! 5

2fT6 2 lr4 -'- 2{ 4 .. .;.

m -'-

uorlm~k~ + L emmetyln~n

• YUI -Klvetty Pukola1~pl

• • eln~j~rvt

-2.0 -1.0 0.0 1.0

p1

2.0

b) 3.0

• Surface Waters • Wells .;. Boreholes

2.0

5/BT 1.0

.;.5fT4 0:

0.0

.. 1fT~

-1.0

-2.0

3.0 4.0 5.0 -3.0

1 ~ KA1 • Surface Waters

• Wells • KA

1 ._ Boreholes

.;. 3!T5

Lemn etylnen Helnaf rvl .i. Jffi• . -• ··~·

Yla- l'lvetty • Paskol!lmpl._ 37T ~ 5/BT

... 2m .;. 5!T3 ... srr liimatair en. KA2 .;. 3fT6

~fllA2~?ft Jm

-2.0

1!Tr'l 1Wf1 2fT1 5fT1

~ff6 4/" 1 .;. 31 1 .... .;.~/"6

4n~ ~~ !T1 .;. J!TJ 1fT5 411' '.;.A t.rrlt· fT6 .;.

-1.0

4m T L.IIJ

.;. 1/JL -'- 1ff4 .;. .i. 2{ 4

~fT4 .1. 1/T

0.0 1.0

p1

2.0 3.0 4.0 5.0

Figure 5-1. Principal component scores for studied samples (evaluated data, cf App. 6, have been extended with surface samples from Ylii-Kivetty, Paskolampi and Heiniijiirvi, and borehole samples 2/T7, 2/T6, 3/T7, 4/T7, 5/T7 and 5/T3 - Tuominen 1995, Snellman et al. 1995a) in a varimax rotated three component model. Samples from each bore hole are marked with an acronym denoting the bore hole (KR omitted for the sake of clarity) and packed-off interval. Acronyms are coloured in accordance with borehole numbers.

..j:::.. ~

42

As a conclusion, principal component analysis with Kivetty data alone gives no clear indications of which samples in the data should be referred to as "end-members", i.e. as samples which present the most extreme character, for example, of glacial, meteoric or saline waters. Furthermore, analysis with Aspo end-members gives only an indication of strong affinity for meteoric and glacial waters for all Kivetty data. It seems, therefore, that groundwater samplings might not have gained enough depth to reach "old, likely saline" groundwaters. As regards meteoric or glacial water affinities among samples, the variance among critical variables (e.g. H-3, 80-18, 8H-2) is not yet sufficient for this classification to be made with principal components.

The two first principal components (Fig. 5-1a) seem to group the bulk of groundwater data in clusters formed by the boreholes: KR4 on the left side, KR3 on the right, KR2 at the bottom, and KR1 and KR5 together between the others. The distribution of clusters may indicate separate hydrochemical evolution among the boreholes, further isolated jlowpaths and hydrology, except for KR1 and KR5.

5.4.2 Main variables and components

All of the Ki vetty ground water samples are fresh; the local median value of Total Dissolved Solids (TDS) is 137 mg/1, being well below 1 g/1 (the limit of brackish water, Davis 1964 ). There is a general trend of increasing TDS and chloride with depth (Fig. 5-2), though there is a large amount of variation. Nearly 90% of the Kivetty water samples used in this geochemical study (Appendix 6) had chloride concentrations below 10 mg/l, thus local water-rock interactions may have a major impact on spatial hydrogeochemical vanance.

The geochemically most evolved groundwaters (Clmax = 48 mg/l or 1.35 mmol/1, Snellman et al. 1995a) have been found in the bottom part of borehole KR5 at a sampling depth of 753-853m (cf. Fig 5-2). The extensive hematised breccia structure in KR5 (819-828m) apparently shows signs of an ancient hydrothermal event with a fracture mineral assemblage of calcite-pyrite-iron oxyhydroxide-iron oxide-clay minerals (Gehor et al. 1995).

Two groundwater samples from borehole KR1 do not show very good correlation of Cl vs. depth. The hydrogeochemistry of these samples, T3 (720-795m) and T2 (815-855m), does not offer any reasonable explanation for the low chlorides. There is a negligible amount of remaining flushing water (a few per cent), no measurable tritium (H-3 < 0.8 TU), and the radiocarbon contents do not either indicate young waters (25-34 pmC). Nor have any technical defects (leakages of the multipacker system etc.) been reported (Niva et al. 1994). One explanation is offered by hydrogeological effects of the modelled fracture zones R23 and R16 (Saksa et al. 1993, 1996) and their possible diluting effects on local hydrogeochemical conditions. On the other hand, while the borehole was open there was a downflow along the hole, which intruded into deeply occurring fracture zones (Saksa et al. 1993). Pitkanen et al. (1992) reported that water samples taken from the borehole section T2 were strongly contaminated by groundwater from the upper part of the hole, particularly with similar water sampled from the depths of section T7.

0

0

100 • --(()

200 I~ 0 0

0.2

X ~<

D

A 300 0-

5 400 .c Q. ~ 500

600

700

800

900

~ D

l:J.

<>

D

<> ~

l:J.

43

Cl, mmol/1

0.4 0.6 0.8 1.2

... -l:J.

X

l:J.

I• Spring • Well l:J. KRl o KR4 o KR5 x KR2 <> KR31

1.4

D

Figure 5-2. Molar chloride concentrations of Kivetty shallow and deep groundwater samples vs. sampling depth.

At the Kivetty site the most diluted groundwaters have been observed in borehole KR4 (Cl 1.0-2.1 mg/1, median value 1.4 mg/1). This homogeneity could be explained by the frequent fracture zones (Saksa et al. 1993, 1996), which are intersected by KR4 (c.f. p. 23). This fracturing may connect neighbouring borehole sections.

In general, a trend of increasing pH with greater salinities and Cl concentrations (Fig. 5-3a) along all three flowpaths can be observed, albeit with some variation in the data. Naturally the lowest pH values have been observed in precipitation (pH 4.1-6.5), surface waters (6.4-6.8) and shallow groundwaters (6.1-7.4) due to atmospheric C(h.

The pH values of borehole groundwaters at Kivetty have quite a narrow range from 7.8 to 9.0 (Fig. 5-3a), being slightly to fairly alkaline. The lowest pH value has been observed in KR1, at 300-345m depth, where structure R15 is intersected. Apparently recent waters via this structure could cause the low pH value, although the tritium ( < 0.8 TU) and C-14 ages (30.1 pmC) do not strongly encourage this hypothesis. The highest pH value was measured in KR4, at 170-195m depth, where the borehole intersects a gabbro vein. High pH values were also measured at Romuvaara from sampling sections which penetrate mafic rocks (Pitkanen et al. 1996a).

a)

9.0

8.5

8.0

=§.7.5

7.0

6.5

6.0

5.5

c)

0.5

0.4

~ s 5 o.3 ~ u

0.2

0.1

0.0

e)

s 0

s s ~

0.04

0.00

In

X X C

(])

li~<> ~6. ~-

1-u~

• •• • I •

0

~

~

a>6.

D

I~

0.2 0.4 0.6 0.8 Cl, mmol/1

X

- V

6. > 6. Oil c ~~ X -..

CO •

• •

0

'f::.l

D

~IZl <ll <D

()

CD .6. <> ~

6.

0.2 0.4 0.6 0.8 Cl, mmol/1

c

X

6. IX6.

<> 'i«

0 0.2 0.4 0.6 0.8 Cl, mmol/1

44

D

1.2 1.4

D

1.2 1.4

D

b)

1.6

Sl.2 0 s s ~ Z0.8

0.4

0.0

d)

0.3

~0.2 s ell ~

0.1

0.0

f)

0.002

0.001

0.000

6.

6.

.c.

0 X X c

(J <>""~

D~ c9 ~D ~

~·

0

•

6 A

011"6. g_o

DD <> .. B 6.

0 <>

~6. ~

0

6.

6.

6.

X

0.2 0.4 0.6 0.8 Cl, mmol/1

X

> 6. <> c 6. X

0.2 0.4 0.6 0.8 Cl, mmol/1

c 6./j)(<> 6.

X X -------;-e 6.

to<>

oD D

-1.2 1.4 0 0.2 0.4 0.6 0.8

Cl, mmol/1

I• Spring • Well 1:J.. KRl o KR4 c KR5 x KR2 ~ KR31

D

1.2 1.4

D

1.2 1.4

D

1.2 1.4

Figure 5-3. pH values and trends of main cations vs. chloride in Kivetty groundwater samples.

45

Groundwaters exhibit an evolution (Fig. 5-4) from Ca-Na-Mg-HC03 waters (classification according to Davis & De Wiest 1967) in the upper part of the bedrock to Na-Ca-HC03 type deeper down. Water composition changes abruptly from lake and spring waters to high HC03 waters during the recharge process. The proportion of Na tends to grow as Cl concentration increases. In the bottom part of borehole KR5, the most mature water type is Na-Ca-Cl-HC03 type, obviously due to more extensive water­rock interaction. This trend is similar to Romuvaara"sjlowpath I (Pitkanen et al. 1996a), reflecting moderate interaction between granitoids and groundwaters at low temperatures.

ea Na + K HC03 + C03 20 40 60

CATIONS % meq/1 ANIONS

l( Lake • Spring • Well /:; KR1 ° KR4 D KR5 X KR2 o KR3

80 Cl

Figure 5-4. Piper plot of chemical composition of water samples at Kivetty. Evolutionary trends have been drawn according to increasing Cl concentration.

5.4.3 Trends of main cations

Trends of the main cations (sodium, calcium, magnesium, potassium and strontium) with increasing chloride are shown in Figures 5-3a .. .f. All the cations and pH show an initial increase starting from very dilute spring waters, which reflects weathering of

46

minerals as water descends in the bedrock. Sodium (Na) shows the strongest increase in the samples from boreholes KR1 and KR5 in Figure 5-3b, but the former shows a much higher concentration with higher Cl than the latter. Similar deviation in concentration levels between KR1 and KR5 is observable in most of main cations and pH, suggesting a separate hydrochemical evolution for these boreholes.

In shallow groundwaters (from springs to wells), calcium (Ca) concentrations increase relatively more strongly than Na, possibly due to the start of calcite dissolution after surface water recharge (Fig. 5-3c ). Ground waters from the upper part of KR1 (T7 and T6) have the highest Ca contents. After the initial peak, calcium contents tend to decrease with increasing Cl, due to e.g. precipitation of calcite on fracture surfaces. The most saline sample shows calcium enrichment, reflecting some Ca source deep in the bedrock.

Magnesium (Mg) shows quite similar behaviour with Cl (Fig. 5-3d) to that of calcium. There is an initial rapid increase, but a maximum of Mg is closely reached in shallow groundwaters. The highest Mg contents were observed in borehole sections KR1/T7 (169-181m) and KR2/T1 (450-500m). After an initial increase magnesium shows a decreasing trend with Cl.

Potassium (K) values in the groundwater samples are low (Fig. 5-3e). At any rate the behaviour of K seems to resemble that of Mg, i.e. strong enrichment in shallow wells and depletion in deep borehole groundwaters. The behaviour of Mg and K suggests a common source at shallow depths and a common sink deeper in the bedrock, e.g. dissolution of biotite and ion exchange, respectively. However, the highest K content was observed in sample KR5/BT, indicating some input of K together with Cl.

Strontium (Sr) in Fig. 5-3f shows a similar trend to that of Ca. Groundwaters from KR1 (169-181m) show the maximum values of Sr.

5.4.4 Trends of main anions

The main ions (HC03 + C03) comprising carbonate alkalinity in the groundwaters are shown in Figure 5-5a. There is a general trend of an initial rapid increase of carbonates vs. chloride. The local maximum value is reached in the upper part of borehole KR1 (section T7, 169-180m, at Cl content 0.1 mmol/l). After the maximum, contents of the carbonate species decrease slowly with increasing Cl. The trend of carbonate alkalinity resembles the behaviour of Ca with Cl (cf. Fig. 5-3c) which may reflect the dissolution­precipitation reactions of calcite on fracture surfaces.

A increasing trend of sulphate (S04} is evident in the deep borehole groundwaters (Fig. 5-5b ), but there also seems to be a decreasing trend in the data from spring waters to shallow and very low saline borehole groundwaters (Cl<0.1 mmol/1). This behaviour is different from that of the groundwaters at the Romuvaara site, where S04 decreases with increasing chloride (Snellman et al. 1995b, Ruotsalainen & Snellman 1996), suggesting a sink for S04 with longer residence time (Pitkanen et al. 1996a). At Kivetty, S04 reaches its highest values in borehole sections KR1/T5 (370-415m) and KR5/BT (735-853m), suggesting a common source for S04 and Cl.

47

Amounts of aqueous sulphide (S2-<tot)) are low in the Kivetty groundwaters (Appendix

6). The maximum value in KR1 is only 0.006 mmol/1. Because of the near detection limit values, variation in the data is relatively large. Redox species are discussed in greater detail in section 5.4.5.

Bromide (Br) values are all low (max. 0.0055 mmol/l in the bottom part of KR5), and thus only the highest values have been considered reliable. The highest Br values show a linear correlation with increasing values of Cl (Fig. 5-5c ), suggesting a similar geochemical origin and history. The groundwater sample from KR5/BT has a Br/Cl ratio of 0.00408 (calculated from mmolll), which is much higher than that for the ocean (0.0016, Drever 1982) or the Baltic Sea (0.00157, Pitkanen et al. 1996b). This does not support a marine origin for the salinity in the groundwaters at Kivetty.

a)

2.0

-... Q s s 1.5

0 u + 0 1.0 u ==

0.5

0.0

c)

0.004 -... Q s s ;; = 0.002

0.000

~ ~

b.

b.

b. I no. ~ 0 io ~!>

[ b. <>I ~ P< X -•

~

0 0.2 0.4 0.6 0.8 Cl, mmol/1

X C COD ~b. :i5o m b.<>

0 0.2 0.4 0.6 0.8 Cl, mmol/1

D

1.2 1.4

0

1.2 1.4

b)

0.07

0.06

::::::0.05 Q s 50.04 0 ... CJ:l0.03

0.02

0.01

0.00

d)

0.2

;;;:; Q

s s ~

0.1

0.0

lA

• ) c

<> <> X

<> I"" X

1• 0 x

1&. tY< b.

l~g • A

0 0.2 0.4 0.6 0.8 Cl, mmol/1

b.

b.

X 0 X<>>

~~~ ~ X

• ocP P<

<> [

~

• 0 0.2 0.4 0.6 0.8

Cl, mmol/1

I• Spring • Well A KRl o KR4 c KR5 x KR2 ~ KR31

Figure 5-5. Trends of main anions vs. chloride in Kivetty groundwater samples.

D

1.2 1.4

0

1.2 1.4

48

There is a general increasing trend of fluoride (F) with increasing chloride (Fig. 5-5d). The highest fluoride values are from borehole KR1 (0.27 mmol/1), apparently suggesting more extensive water-rock interaction in samples with longer residence times. Fluoride concentrations are generally slightly higher in the Kivetty than in the Romuvaara data (<0.1 mmol/1 according to Pitkanen et al. (1996a)). This is considered to be due to the higher amounts of F-bearing minerals in the bedrock of Kivetty (Gehor et al. 1995), indicated also by the higher contents of fluorine in whole rock samples than in those from Romuvaara (cf. Table 2-1). Of anions, fluoride shows the most distinct deviation with higher Cl concentration between the KR1 and KR5 samples.

Aqueous silica (Si022 has the highest concentrations in groundwater samples from shallow wells corresponding to Mg and K (Fig. 5-6). This emphasises silicate weathering as a main source of dissolved species in shallow depths. Deeper in borehole samples silica decreases, indicating silica precipitation which is consistent with fracture mineral observations (Gehor et al. 1995).

........ 0 E

0.6

0.5

0.4

E 0.3

0 00

0.2

0.1

0.0

•

.. •

!

C[] 11 0

~X lv

0

• 6

~ ...,..., <>

<>

0 0.2

6

X Cl A

0.4 0.6 0.8 1.2 Cl, rnrnol/1

I• Spring • Well t:.. KRl o KR4 o KR5 x KR2 <> KR3j

Figure 5-6. Silica vs. chloride in Kivetty groundwater samples.

5.4.5 Redox conditions at Kivetty

D

1.4

General aspects of redox reactions in groundwaters and difficulties of obtaining representative Eh measurements are discussed in the report on geochemical modelling of the Romuvaara site (Pitkanen et al. 1996a) and in the summary report of hydrochemical baseline characterisation (Ruotsalainen & Snellman 1996). Similar technical problems (Section 5.3) as at Romuvaara were also faced in on-site redox measurements at Kivetty. In particular, very low amounts of aqueous, redox-active

49

species in local dilute groundwaters hampered the achievement of representative Eh results for all samples. The median value of Fe( tot) in the Kivetty groundwaters is only 1.54·10-6 mol/l comparable with the theoretical lower limit (10-6 mol/1, Grenthe et al. 1992) of iron for successful Eh measurements, if Fe2+/Fe3+ is the redox controlling couple.

There is quite a large scatter in the amounts of redox active species in the dilute groundwaters of Kivetty (Fig. 5-7). The importance of ferrous iron as a redox controlling species seems to decrease with increasing Cl.

Log p(C02) values (calculated with EQ3/6, Fig. 5-8a) that exceeded the atmospheric value of log p(C02) (-3.5) in the most dilute groundwaters especially in boreholes KR1, KR3 and KR5 indicate biochemical oxidation of organic carbon as a source of carbonate during the recharge of groundwaters. The calculated values of log p(C02) and alkalinity are fairly high, therefore the local groundwaters are well buffered against atmospheric C02 during sampling. For this reason atmospheric C02 is not considered to cause any significant error in the analytical results of carbon species.

Dissolved organic carbon (DOC) is usually present in deep granitic groundwaters at levels below 0.17 mmol/1 (Petterson et al. 1990). At Kivetty, organic carbon has been found at all sampling levels, in all boreholes, ranging from 0.06 to 1.4 mmol/1 (median value 0.5 mmol/1). There are indications of organic leachates from polyamide tubes of the sampling equipment (Snellman et al. 1995a), thus these figures have a high degree of uncertainty.

--0 s s

0.006

0.005

0.004

t--~0.003 ~ ;.. 0

~ 0.002

0.001

0.000

• D

D

IU

D

•

• • [J

D • D D Db D D D

JL

~·

' D D D D • D~

0.0 0.2

• •

D

D D

0.4 0.6 0.8 1.0 1.2 1.4 Cl, mmol/1

I • S2- o Fe2+ I

Figure 5-7. Redox active reduced S2- and Fe2

+ in the Kivetty groundwaters.

a)

-2.5

8-3.5 c.

-4.0

-4.5

• • .. • ~·6

6

6 L>

>

t~ 6

101 X

x6

a>

X X

0

0 0.2 0.4 0.6 0.8 Cl, mmol/1

50

D

1.2 1.4

b)

0.015

~ e e 0.010

= ::;

0.005

0.000

•

a. 0

D 0 X .,.LP :M Cl [

?1!. ~0.)( 6 6

0 0.2 0.4 0.6 0.8 Cl, mmol/1

I• Spring • Well A KR1 o KR4 [] KR5 x KR2 <> KR31

D

1.2 1.4

Figure 5-8. Partial pressure of C02 (log p(C02)) and Mn vs. Cl. In diagram a) log p(C02) for atmospheric carbon dioxide (-3. 5) is shown as a bold line.

Manganese is a noteworthy redox species in shallow groundwaters. At Kivetty, the highest Mn values were observed in some borehole wells and boreholes KR4 and KR5. In general, Mn shows a decreasing trend with increasing chloride (Figure 5-8b) similar to iron.

As mentioned above (Fig. 5-7), iron contents are generally very low in the Kivetty groundwaters. Only some borehole wells with near neutral pH values show elevated Fe levels. Iron shows a clear decreasing trend with Cl.

In studies on fracture minerals Gehor et al. (1995) frequently observed iron oxyhydroxides (limonite and goethite) in the upper parts (usually from the surface to depth 130 ... 170m) of nearly all drill cores. In particular, core samples from boreholes KR4 and KR5 had very long sections where fracture mineral assemblages included iron oxyhydroxides. Boreholes KR5 and KR1 also had sections with iron oxides (hematite), apparently due to very old hydrothermal events.

Another sink for aqueous iron is apparently the precipitation of pyrite, which has frequently been observed in core samples (Gehor et al. 1995). Core samples from boreholes KR9 and KR5 have pyrite all along their length. If the deep ground waters host anaerobic microbes which reduce sulphate to sulphide while using organic compounds as energy sources, pyrite will be formed in the presence of any dissolved iron or ferrioxyhydroxides (e.g. Appelo & Postma 1993). According to Figure 5-9a the measured Eh values from Kivetty are lower than the theoretical ones calculated with EQ3/6 using Fe2

+ /Fe3+ as the dominating redox couple. This contradiction refers

apparently to oxidation and precipitation of aqueous iron during the slow sampling procedure. Dark precipitates on inner walls of the sampling tubes reported by Snellman et al. (1995c) are in agreement with this suggestion.

51

Observed sulphide values (S2-<tot) = H2S + HS-+ S2

- + polysulphides) are generally rather low, with a maximum of 6 x 10-3 mmol/1. The trend with Cl (Fig. 5-7) is not very clear and the data show fairly wide variation, especially in dilute near-surface groundwaters. Figure 5-9b suggests a local decreasing trend of Eh with increasing S2

-<tot)· The calculated redox values based on S2-!Sol- ratio are much lower than the measured ones (Fig. 5-9a). This suggests that there has been problems in the field Eh and pH measurements. However, the frequent observations of sulphide in groundwater samples and S04 decrease in low saline ground waters (Fig. 5-5b) suggest reducing conditions and microbially mediated reduction of sulphate. Pyrite on fracture calcites (Gehor et al. 1995) also point to the late stage reduction of S04. This has been recently confirmed with observations of sulphate and iron reducing bacteria in the Kivetty groundwaters (Haveman et al. 1998).

The few observations of evacuated methane show low contents (maximum 18 ~1/1). The uncertainty of this data is extensive, since pressure changes during lift-up and low pumping rates from borehole sections with often very low hydraulic conductivities can cause severe outgassing and atmospheric contamination of samples. The highest C~ content has been observed in very dilute groundwater in borehole KRl (300-345m). The low 8C-13(C~) value (-58%o) of the most reliable sample (KR1/T6) with enough C~ points to a biogenic origin of methane.

a)

• 100 a .. • •• ~ -

-200

16Eh(S) I 6 6

•Eh(Fe) 6 6

-300

-300 -200 -100 0 Measured Eh, m V

•

100 200

b)

100

50

0

> -50

~ -100 .c

~ -150

-200

-250

-300

-350

0

0 ~

0

f-~

0

6KR1

oKR4

o KRS XKR2

<> KR3

~ <>

6

e

0.001 0.002 0.003 0.004 0.005 0.006 s2-(tot), mmol/1

Figure 5-9a. Measured and calculated Eh values for groundwaters at Kivetty. S = S2-

(tot)ISO/- based Eh values calculated with E03/6. Fe =Eh values based on Fe2+/Fe3+.

Figure 5-9 b. Measured Eh values vs. S2-(tatJ in Kivetty groundwaters_

52

5.5 Isotopes

5.5.1 Stable isotopes (8H-2 and 80-18) of groundwater

According to Figure 5-10a, the contents of stable isotopes generally point to a meteoric origin for the Kivetty groundwater samples. The data show a bimodal characteristic. Shallow groundwaters from springs and wells plot under the Global Meteoric Water Line (8H-2 = 8·80-18 + 10, Craig 1961). The depleted deuterium values of shallower groundwaters may reflect surface evaporation (Fritz & Pontes 1980). Most of the deep groundwater samples plot in the same region, as do the shallow groundwaters, suggesting similar infiltration conditions.

There are some groundwater samples (Fig. 5-10a) from nearly all boreholes (KR5, KR2, KR3 and KRl) in which the lighter isotopic composition suggests cooler infiltration conditions, e.g. some glacial melt water input. The lightest values have been observed in sample KR5/T4. This could also indicate a rather long mean residence time of more than 104 a. The other isotopic results of this sample (negligible H-3 and quite low C-14 (19.8 pmC)) are in good agreement with this hypothesis.

Both 80-18 (Figure 5-lOb) and 8H-2 show a decreasing trend with increasing Cl, suggesting longer residence times for geochemically more mature groundwaters.

a)

-90

~ 0 -95 ~ rJ1 0 ~ 0

M'-100 :I:

-105

-110

/ V.

V

/ /'9.

A/, ,QIIo ~

/ 00~-

•

k{'o :

-15.0 -14.0 -13.0 -12.0 0-18, o/oo SMOW

-11.0

b)

-12.5

g ~ -13.0 rJ1 0 0 0 c:IO -13.5 ,.... 6

-14.0

-14.5

l!l[h -~ [J

Cll!J a. ~ ....

A XI:

oP< 0 A

X

0 0.2 0.4 0.6 0.8 Cl, mmol/1

I• Spring • Well 11 KRl o KR4 c KR5 x KR2 ~ KR31

D

1.2 1.4

Figure 5-JOa. Contents of stable isotopes (H-2 and 0-18) in water samples from Kivetty plotted as 8H-2 vs. 80-18. Figure 5-JOb. Values of80-18 vs. Cl in water samples from Kivetty.

53

5.5.2 Tritium (H-3)

The chronological development of tritium activity in precipitation at several locations is shown in Figure 5-11 (Ruotsalainen & Snellman 1996). Atmospheric thermonuclear experiments, which drastically elevated the tritium level in the atmosphere, began in 1952. The half-life of H-3 is approximately 12.5 a, therefore no tritium should be detectable in groundwaters infiltrated prior to 1952. Tritium contents have decreased since 1988 when the site investigations started at Kivetty. Some shallow groundwater samples taken in 1993 at Kivetty (Snellman et al. 1995a) already showed prebomb tritium levels, emphasising the strong dilution of tritium in groundwaters from high atmospheric values in the 1960s.

There is a noteworthy difference in the H-3 activities of the Kivetty groundwaters compared with those at Romuvaara, with a general, below detection limit level of tritium values at Kivetty. At Romuvaara only 8.3% of observations were below the detection limit (0.8 TU, Ruotsalainen & Snellman 1996), whereas at Kivetty the corresponding value was 25.6%.

Amounts of H-3 in Kivetty water samples decrease with Cl (Fig. 5-12a), reflecting minor mixing effects with recent groundwaters in deeper parts of the bedrock. The most obvious reasons for observations of tritium in the deep groundwaters is contamination by open-hole effects or borehole activities in field tests and sampling.

10000

• 1000

• •

- -~

+ 0 • 10 •

~ • +

+

0.1

1950 1955 1960 1965 1970 1975 1980 1985 1990 1995

Time, years

• Europe/N.hemisp./rain • Finland, rain .6. Kivetty, rain + Kivetty, surface/shallow gw.

<> Europe/N.hemisp./snow • Laukaa, rain I:J. Kivetty, snow

Figure 5-11. Tritium in precipitation (and surface waters or shallow groundwaters) in Europe, Finland, Laukaa (50km south from Kivetty), and Kivetty. Data compiled from various sources (Roether 1967, !AEA 1983, STUK 1977, Saxen et al. 1994, Saxen 1995, Keinonen et al. 1992, Tuominen 1995) by Snellman et al. (1995b).

a)

30

25

10

5

0

• •

• • <> <>Q c

~ V V X oo ~~ Xt:. t:.

0 0.2 0.4 0.6 0.8 Cl, mmol/1

54

0

1.2 1.4

b)

u e

lOO

80

':: 60 ~ u

40

20

0

•

0

I~ 'boo a

A

0

• •

• t:. <>

0 0 <>

~

t:. c

5 10 15 20 H-3, TU

I• Spring • Well A. KRl o KR4 c KR5 x KR2 ~ KR31

<

25 30

Figure 5-12. H-3 activities vs. Cl and activities of C-14 vs. H-3 in Kivetty water samples.

In general, groundwaters reflecting minor mixing effects are those with the longest mean residence times, also according to activities of C-14 and H-3 (Fig. 5-12b ). There is, however, great variation in the data.

5.5.3 Sulphur of aqueous S04 (85-34 in S04)

Stable isotopes (8S-34 and 80-18) in aqueous sulphur species (mainly so/-, H2S, HS­and S2

-) can be used in evaluating the origin of the sulphur and determining possible redox processes and, indirectly, residence times (Fontes 1994 ). Reduction of S04 at low temperature is possible in the presence of S04 reducing bacteria. The reduced S04 is depleted significantly in 8S-34, by 40-50%o units in this bacterial process (Claypool et al. 1980), therefore the remaining portion of S04 is enriched in 8S-34.

The 8S-34(S04) value of the shallow groundwater in a spring at Kivetty (Fig. 5-13) is of the same level as those reported for the Swedish Stripa investigations (Fontes et al. 1990), representing precipitation of a non-industrial region (Fritz & Fontes 1980). There is a clear increasing trend of 8S-34(S04) values of Kivetty groundwater samples with increasing Cl, reflecting the reduction of S04 by the catalysing influence of sulphate reducing microbes which Haveman et al. (1998) recently observed in the local groundwaters. Input of marine S04 (with some reduction in the bedrock) would also increase 8S-34(S04), but the mixing of sea-water into the system is not favoured by the observed Br/Cl ratio.

35

30

5

0

55

c

6

o!!>X

0

•

0 0.2 0.4 0.6 0.8 Cl, mmol/1

Cl

1.2 1.4

I• Spring • Well A KRl o KR4 c KR5 x KR2 ~ KR31

Figure 5-13. 8S-34(S04) values vs. Cl in Kivetty water samples.

5.5.4 Uranium isotopes and radon

Uranium isotopes U-234 and U-238 together with their activity ratio, U-234/U-238 or AR, were analysed for nearly all water samples both in the filtrate and in the particle fraction. The aqueous total uranium (Utot=U-234+U-238, calculated by U-238 and AR) reaches its highest values in samples from boreholes KR1 and KR4 (Fig. 5-14a).

The AR is frequently used in groundwater studies as it is not influenced by fractionation during chemical reactions or sorption. Generally groundwaters have activity ratios greater than one. In particular, granitic deep groundwaters seem to have an increasing trend of Utot and U-234/U-238 with residence time and flow (Cherdyntsev 1971). This seems also to be true of the Kivetty data (Fig. 5-14a). An interesting trend of AR vs. C-14 is shown in Figure 5-14b. The increase of AR with decreasing C-14 indicates a positive correlation between Utot and AR with increasing residence time in Kivetty ground waters.

Pearson et al. (1991) have also observed similarly high U, and high AR groundwaters in the crystalline bedrock of Sackingen and Rheinfelden in northern Switzerland, reflecting complex patterns of recharge areas and ages of local groundwaters. Excess of U-234 is correlated with time of water-rock interaction, concentration, chemical state and distribution of uranium in the local rock matrix, and with the contact surface and chemical properties of reacting ground waters. Pearson et al. ( 1991) proposed that high AR with high U-contents indicate long-lasting water-rock interaction with U-bearing rock. Low AR (close to one) in U-rich water is typical of young groundwaters in recharge zones. The groundwater samples from deep boreholes (KR) at Kivetty have high AR and high U-concentration deviating from recent, shallow groundwaters (springs

56

and wells), which indicate long term water-rock interaction and residence time for deep ground water.

There seems to be an increasing trend of Utot down to depths of 150-300m, followed by a decrease (Figure 5-14c). This is apparently due to a combination of contact time, rather rapid hydrological circulation in the active weathering zone, and changing redox conditions. Although Utot is highest in groundwaters from boreholes KR1 and KR4, the lithogeochemical studies of boreholes KR1-KR5 (Lindberg & Paananen 1989, 1990) do not show higher U contents for the core samples from boreholes KR1 and KR4 compared with the other boreholes.

There is a general trend of decreasing activity of Utot with increasing Cl (Fig. 5-14d) after an initial increase of U101 with Cl values above 0.1 mmol/1. This is an indirect indication of prevailing reducing conditions in deep groundwaters at Kivetty (e.g. according to Pearson et al. 1991). In the upper part of the bedrock, oxic conditions of shallow groundwaters allow dissolution of uranium from the rock matrix. Deeper below 300m in geochemically more mature groundwaters, redox conditions are reducing, sorbing U4

+ into the rock matrix during longer residence time. However, reducing conditions may also dominate in shallower depths, because groundwater samples from KR1 with the highest U101 contents contain hydrogen sulphide. Therefore the increasing trend with depth in Figure 5-14b may partly represent heterogeneity between flowpaths and distribution of available U along flowpaths in the oxidising recharge zone.

Thus the Kivetty uranium data suggest that there are reducing conditions in the deeper part of the bedrock (below 150-300m), keeping uranium in the rock matrix. In the uppermost part of the bedrock redox conditions are complex, locally more oxidising, dissolving U6

+ varyingly into the aqueous phase.

The activity of radon (Rn-222) can be a useful qualitative clue to the short-term history of groundwater in the vicinity of the sampling point. Radon is highly soluble in water and is not absorbed by any solid, thus its activity is generally greater than that of uranium or Ra-226 by a factor of 103 to 105

. High Rn-222 contents are typical for waters in porous aquifers with high contents of finely dispersed uranium (or Ra-226) in contact with pore water, and also for a fast transfer of such pore water to groundwater at the sampling point (Pearson et al. 1991).

In the crystalline granitic bedrock of Kivetty, activities of radon in the groundwaters show a clear decreasing trend with increasing Cl (Fig. 5-15). The low amounts of radon in the deep groundwaters are apparently due to the low permeability of bedrock at greater depths, and, consequently, the slower release of Rn into groundwater. Also outgassing during the slow pumping should not be ruled out as affecting the observed trend. The high radon activities in shallow groundwaters in Figure 5-15 indicate more rapid movement of groundwaters in the upper, more fractured part of the bedrock.

a)

800

~600 =:l e c 0

;§'400

200

0

c)

800

~600 =:l e c 0

;§'400

200

0

0.0

I

~ -0

57

A

A 0

0

0 -cal A <>

...,, <> V

<> ~ X X ~A

I [J

.... ~ .... o -2.0 4.0 6.0 8.0

U-234/U-238

A

A 0

0

0 ,...

AP A 0

A .... <>

~~ o<> ~

0 X

[J

hi;! [J

200 400 600 800 1000 Depth, m

b)

6.0

5.0

2.0

1.0

0.0

d)

800

~600 =:l e c 0

;§' 400

200

0

<>

<> <>

4, [ A A

n --~ 0

<> .c!J [J

X ,... X

[J

- ,Q .~ •

0 20 40 60 80 C-14, pmC

A

A 0

0

0

-<> A

OA A

IA...n.<> IV ~

~ <>X X X I

l..cn [

0 0.2 0.4 0.6 0.8 Cl, mmol/1

I• Spring • Well 1l. KRl o KR4 c KR5 x KR2 ~ KR31

• •

100 120

[J

1.2 1.4

Figure. 5-14. U(tot) vs. U-234/U-238, U101 vs. sampling depth, U101 vs. Cl and activity ratio U-234/U-238 vs. C-14 of the Kivetty groundwaters

5.5.5 Carbon isotopes (8C-13, C-14 in DIC)

The naturally occurring radioisotope of carbon, C-14 (half-life 5730 years), can provide information about flow rates and residence times of groundwater in systems with straightforward patterns of flow and geochemical evolution (e.g. Plummer et al. 1990, Pearson et al. 1991). Determination of C-14 contents in groundwater is a common tool for the evaluation of age relations and, less often, absolute ages within the aquifer. To determine isotopic ground water ages using the C-14 of dissolved inorganic carbon (DIC) approach requires detailed information on the C-14 sources and sinks of dissolved carbonate in unsaturated and saturated zones. Ratios of stable carbon 8C-13 and oxygen 80-18 isotopes reflect the origin of minerals and of dissolved carbonate species and may also provide insight into the sources of the groundwater itself (Pearson et al. 1991).

58

1400~----~----~----~~~--~----~----~----~

1200+-----+-----+-----+-----+-----+-----+---~

1000~=~--+-----+-----+-----+-----+-----+---~

~ 800+-----+-----+-----+-----+-----+-----+---~ ~ . N' N N 1:J.

= 600+-----+-----+-----+-----+-----+-----+---~ ~

400+1~~---+-----+-----+-----+----~----~----~

D 0

•• 200+-A~~6-+-----+-----+-----+-----r-----r----~

X X fY X

• I:J. 0 ~o rno

0 0.2

X I:J. c

0.4 0.6 0.8 1.2 Cl, mmol/1

I • Spring • Well t::. KRl o KR4 o KR5 x KR2 o KR31

Figure 5-15. Activities of Rn-222 vs. Cl in Kivetty groundwaters.

D

1.4

The carbon isotopic composition of the aqueous carbon is a direct reflection of the geochemical history of the groundwater. This evolution begins in the recharge environment and continues in the subsurface where mineral-water interaction dominates. In most recharge environments the uptake of isotopically depleted soil C02

produced by plant root respiration and the decay of plant debris (basically reaction 5-1), buffers 8C-13 values in the aqueous carbon by about -25%o PDB (e.g. Pearson et al. 1991, Pitkanen et al. 1996a). In addition to modern vegetation, other organic sources are organic deposits such as peatbogs and ancient dissolved organic carbon in the underlying bedrock. In the first case the organic source would have measurable C-14, whereas in the latter the organic C source could be "dead".

(5-1)

The subsequent dissolution of carbonate minerals by carbonic acid (reaction 5-2) usually causes enrichment in 8C-13, and values as high as O%o can be reached if the incongruent dissolution of dolomites occurs (Plummer et al. 1990, Pearson et al. 1991). Fracture calcites from Canadian and Fennoscandian shields also have 8C-13 values higher than organics; typically bulk values are over -15%a according to the compilation of Frape et al. (1992). Higher values (-10 to O%o) prevail in the upper part of the bedrock, e.g. at Olkiluoto (Blomqvist et al. 1992) where calcite dissolution is most expected. Dissolving old carbonates that are free of C-14 causes a decrease in the C-14 activity of DIC, i.e. isotopic dilution.

59

(5-2)

Biological processes can further modify the isotopic composition of the dissolved carbonate and generate 8C-13 values as high as +20%o (IAEA 1983) if methane production occurs, or very negative values may develop where organic compounds are oxidised (mineralised) and added to DIC (Fritz & Pontes 1980, IAEA 1983, Nordstrom et al. 1985). For 8C-13, bacterial methanogenesis displays 8C-13 ratios close to or above zero. This production of CH4 (strongly depleted C-13) either releases C02 (i.e. fermentation, e.g. reaction 5-3) enriched with 8C-13 as a respiration product or, if C02 reduction is the mechanism, residual DIC is enriched in 8C-13. In contrast, low 8C-13 contents in the DIC indicate either "oxidation" of organic matter or C~, the latter possibly having signatures below -60%o (Fritz et al 1987). Such mineralisation without isotope effects, i.e. the preservation of a typical negative 8C-13 signal of biogenic origin, is to be expected where S04 reduction is strong and where the C02 formed from bacterial respiration is not fractionated with respect to the organic precursor (Drimmie et al. 1991).

(5-3)

Nearly 50% of the area of Finland is covered by bogs, also typical of the bedrock depressions at Kivetty. In such an environment biological processes in the groundwaters can be expected and might be very important locally for dissolved carbonate in the recharging groundwater.

5. 5. 5. 1 Radiocarbon and evidence of uncertainty in carbonate data

Radiocarbon in dissolved carbonate shows a decreasing trend with chloride enrichment (Fig. 5-16a). The highest values belong to spring waters and are clearly over 100 pmC, indicating formation since the start of the nuclear era (i.e after the 1960s), which is confirmed by tritium contents. C-14 depletes steeply with increasing Cl content, reducing to 10-20 pmC in higher salinities. However, the lowest values are found in the samples from KR2 with only moderate salinities. Tritium data indicate that this discrepancy may be partly attributable to young water mixing/contamination in other samples.

In Figure 5-16b, the measured tritium contents of the groundwater samples are plotted against their C-14 contents (DIC). The sample with the highest H-3 content (KR3/T5), which in fact represents KR3/T8 (see App. 1) according to Niva et al. (1994), is considered to be 100% young water, which presumes isotopic dilution of C-14, e.g. by calcite dissolution and/or respiration of old organic carbon. The sample is the best example of the Kivetty data to represent infiltration during the 1960s. The sample is one of the points forming the line of the highest H-3/C-14 ratios of the Kivetty data. The ratio is similar to the highest ratio observed in the Romuvaara data (Pitkanen et al. 1996a). The other sample showing a high tritium level along the line is from KR2/T7. Although this was discarded in the evaluation process (eh. 5.3) it is used in this diagram (Fig. 5-16b) for the sake of evaluation (cf. section 5.5.5.3).

a)

u E

100

80

': 60 ! u

40

20

0

~ •

• OD

[J

•o t-

1~.~ t:. t:. 1.1'> ..

0 )t X

0 0.2 0.4 0.6 0.8 Cl, mmol/1

[J

1.2 1.4

60

b)

20

;:;;! E-~

= [J

10 0 [J

0

0 [J

xO[J 0 0

0 40 C-14, pmC

I• Spring • Well A KRl o KR4 c KR5 x KR2 ~ KR31

• •

80 120

Figure 5-16. Radiocarbon (DIC) vs. Cl and tritium vs. radiocarbon (DJC) in the Kivetty groundwater s.

Other samples comprising the line (KR2/T1; KR3/T3, T2, T1) should also contain a recent water component from the 1960s, although their lower H-3 and C-14 values indicate a lower proportion of young water. The water samples are composed of an old groundwater component possibly contaminated by young groundwater due to borehole activities from the near-surface part of boreholes KR2 and KR3.

In Figure 5-16b several points plot between the line and the samples with no detected H-3 (presented as 0 TU). Compositions in this range may be either mixtures of the young water with older non-tritium, but moderate C-14 containing water, or they could be unmixed young waters recharged after the 1960s when the H-3 in surface water decreased, as the spring waters indicate. Samples containing less than 2-3 TU (KR2/T5; KR4/T4,T2,T1; KR5/T6, BT) belong to the former group and are considered to contain such a small young water component that their carbonate values are pretty reliable, because the dissolved carbonate contents of the Kivetty data do not vary significantly between single samples. Their tritium contents are interpreted as minor contamination of young water from shallow depths, most probably during field activities.

Samples KR3/T6, KR5/T2 and T1 are within the range of modem tritium concentrations (about 10 TU). The radiocarbon level in the samples from KR5 represent KR3/T5 and are interpreted as 100% young water. Sample KR3/T6 is interpreted as a mixture of young and old water due to the relatively low C-14 content. Although the KR5 samples are from great depth, the finding can be explained by hydraulic measurements and tests which show connections between the packer intervals (Ni va et al. 1994, Rouhiainen 1996) and thus a possible subparallel fracture zone near the borehole which short­circuits sampling intervals. This zone (e.g. R8) could have fed near-surface water to sampling depths during technical operations.

61

The moderately high C-14 values of the two most saline samples, KR1/T5 (24.6 pmC) and KR5/BT (25.3 pmC), cannot be explained by their tritium contents. Before borehole KR1 was plugged in the early 1990s, groundwater moved downwards from section T7 and recharged deeper parts of the borehole (Pitkanen et al. 1992, Saksa et al 1993). The uppermost sample from KR1 does not contain detectable tritium; thus it is possible that water from the upper levels mixed with deeper groundwater samples without any tritium contamination. This evidently explains the dilute nature of deep sections in KR1, particularly groundwater from section T2 (cf. section 5.4.2) and its relatively high C-14 content (34 pmC) deep in the bedrock. There are uncertainties with the deepest sample from KR5/BT due to the long borehole interval and immediate sampling after drilling. Therefore the sample represents a mixture from a wide depth interval and may also be disturbed by ground water from the upper part of the borehole.

The carbonate chemistry of the ground water samples can be considered representative in most cases and can be recommended for interpretation and modelling purposes. The samples KR2/T1; KR3/T6, T3, T2, T1 are mixtures of young and old water components and probably do not represent natural groundwater compositions.

5. 5. 5. 2 Carbon isotopes and carbonate evolution

All 8C-13 values of dissolved carbonate are below -15.0%o PDB, emphasising the significance of biogenic carbon in the system (Fig. 5-17a). The lowest values belong to spring and shallow groundwaters and to some samples from borehole KR5 (T1 - T4). 8C-13 values around -23%o PDB in low pH spring waters correspond to the value interpreted as open system equilibration of biogenic soil-C02 at Stripa (Fritz et al. 1989). These waters are thought to represent initial recharge conditions for currently forming groundwaters at the site. Significant enrichment of 8C-13 is not seen in shallow well waters. This indicates that there is no or only trace calcite dissolution occurring in shallow depths and so coincides with the lack of calcite infillings in fractures (Gehor et al. 1995). The low 8C-13 values observed in some samples from KR5 may originate from peat-water interaction. The borehole has been drilled under a peatbog, and relatively diluted C-14 signatures (Fig. 5-16a) in the samples together with high tritium support at least a relatively old organic carbon source. Till under the bog may also contain old organic carbon derived from as early as the preglacial period.

Most of the groundwater samples from the boreholes show a heavier 8C-13 signature than shallow waters (Fig. 5-17a). This indicates the dissolution of calcite in a closed system along the flow during evolution. After initial enrichment 8C-13 seems to deplete, possibly as a result of e.g. the fractionation effect of calcite precipitation on 8C-13 and/or anaerobic oxidation of dissolved organic carbon deep in the bedrock.

The most chloride rich sample (KR5/BT) has exceptionally high 8C-13. Several evolutionary paths may explain this: First, the amount of calcite dissolution after recharge may have been greater than in the flowpaths to other samples. Second, the quantitative evolution of dissolved carbonate could have been similar to that of other samples, but due to the long residence time the groundwater has had time to equilibrate with fracture calcite, i.e. through isotope exchange. The third possibility could be some

62

fermentation reaction giving a low 8C-13 signature to methane and a high one to carbon dioxide. However, the measured methane contents in groundwater samples are so low (App. 6) that fermentation is not considered capable of changing sufficiently the 8C-13 value of the most saline sample. Inmixing of atmospheric C02 cannot be excluded (cf. previous section).

a) -15

-17

~ ~ ~ 0-19 0 c

!"')

0-21

-23

-25

c)

2.0

a-e 1.5

.6 :5 ] 1.0

< 0.5

0.0

0

"'~ CA ~

A ~<>A <> Ax A

ID

• r

• D

D

0 0.2 0.4 0.6 0.8 Cl, mmol/1

0 40 C-14, pmC

u

1.2 1.4

•• 80 120

b) -15

-17

~ ~ ~ -19 dE. !"') -u -21

-23

C-13 = -IO%o in -t-----....--------F-:--,, dissolving calcite -----t

X<>

<> • D •

Mineralisation of D peat derived carbon

-25 -+-------+-----t------1 0 40 80 120

C-14, pmC

d)

<>

~-13 ~ 0

~ X rJl

-; l X A<> ;§ l/ / 0-14

/ D

;j: VI

D

-15

0 40 80 120 C-14, pmC

I• Spring • Well A KRl o KR4 o KR5 x KR2 ~ KR31

Figure 5-17a. 8C-13 vs. Cl in the Kivetty groundwaters. Figure 5-17b. Carbon isotope compositions of dissolved carbonate. The effect of single carbonate sources is indicated by dotted arrows, potential evolutionary trends by solid arrows. Inside the circle is plotted a group of samples discussed in the text. Figure 5-17c. Alkalinity variations vs. radiocarbon content (DJC). The illustrated trends and circle refer to the same samples as in diagram 5-17b. Figure 5-17 d. Oxygen isotopic composition of groundwater samples vs. radiocarbon content (DIC). The illustrated trends (solid lines) and circle refer to the same samples as in diagram 5-17b. Dashed arrows show "uncontaminated" positions of samples presented in Table 5-4.

63

Combined carbon isotope examination together with alkalinity (Fig. 5-17b and 5-17c) reveals more details of potential carbonate reactions prevailing during the evolution of groundwaters. This information can be utilised and tested in mass-balance modelling (Chapter 7). The data do not show the dominating calcite dissolution trend (Fig. 5-17b) expected from the deep occurrence of calcite in open fractures. Organic derived carbonic acid in recharge water seems to be largely consumed in silicate weathering. Second, tritium rich samples from KR5 show a trend of carbonate derived from pure organic carbon (cf. Fig. 5-16b), in which case calcite dissolution cannot dilute their C-14 value. Organic carbon should also have quite a low C-14 signature in these samples, and a peat deposit origin or old soluble organic carbon in till seem more plausible.

The hydrogeological model predicts flowpaths from KR1 towards KR4, but the carbonate chemistry reflects certain differences between the evolution of groundwaters sampled from these boreholes. The samples from KR1 show strong initial enrichment of alkalinity with C-14 depletion (except KR1/T4), whereas alkalinity values of KR4 stay considerably lower (Fig. 5-17c ). However, 8C-13 values stay on a lower level in KR1 (Fig. 5-17b), indicating more extensive organic carbon mineralisation both totally and relatively (organic derived carbonate/calcite dissolution) along flowpaths leading to the upper part of KR1 than during initial alkalinity enrichment in the flow system leading to sampling sections in KR4 (samples inside the circle). The latter seem to be relatively more affected by calcite dissolution. Sample KR1/T4 (also inside the circle) behaves similarly to the data of KR4. Carbonate evolution of the samples from KR5 also seems to deviate from the other two boreholes in jlowpath I (Fig. 3-4), and it is virtually impossible to adjust these samples to give a similar dissolved carbonate evolution to that of samples from KR1 or KR4.

The results for KR2 also indicate the calcite dominated evolution of dissolved carbonate in the Kivetty data as that leading to 8C-13 enrichment in KR4. Therefore the extreme carbonate evolutionary trend of KR1 (locates upstream from KR2) could not explain the evolution in KR2. Dissolved carbonate and carbon isotopes suggest a hydrogeochemical analogy in the south-western part of the study site and a possible hydrogeological connection between boreholes KR2 and KR4 (and KR1/T4).

The carbonate chemistry of the groundwater samples from KR3 is difficult to interpret due to strong contamination by young water. Minor calcite dissolution seems evident.

5. 5. 5. 3 Implications for pa/aeohydrogeology

Radiocarbon and 80-18 levels both give palaeohydrogeological information. Figure 5-17d shows quite a stable 80-18 level with depleting C-14 down to about 30 pmC. Further depletion leads to lighter oxygen composition in groundwater, indicating colder climate precipitation, e.g. input of a glacial meltwater component in the groundwater. The 80-18 of the ice sheet is estimated to have been about -20%o (e.g. Kankainen 1986, Pitkanen et al. 1996b ). Tritium contents suggest C-14 dilution to the 50 - 60 pmC level without radioactive decay (Fig. 5-16b). The meltwater signature begins to increase in samples with 20 - 25 pmC. When the half-life of C-14 (5,730 years) is considered, a rough estimate gives average ages of 8,000- 6,000 years. This could be a consequence

64

of young water contamination depicted for samples from KR2, KR3 and KR5 in section 5.5.5.1. The lowest C-14 values at Kivetty, observed in KR2/T4, seem to represent the retreat stage of the ice sheet, but do not have the lowest 80-18 value.

The original C-14 value can be estimated for the old component in samples representing the highest H-3/C-14 ratio in Figure 5-16b by assuming that the contaminating samples are those with the highest H-3 content in each borehole (KR3/T5 and KR2/T7). The mixing proportion is calculated simply from the H-3 contents of samples based on the presumption that H-3 of the old component is zero (Table 5-4). In calculating C-14 values for old components the dissolved carbonate contents of each mixed sample have been taken into account.

Table 5-4. Calculated C-14 (pmC) and 80-18 (%o SMOW) values for the old component of mixed groundwater samples representing the highest H-3 (TU)/C-14 ratio at Kivetty, and measured H-3, C-14 and 80-18 values for assumed contaminating young waters, with degree of contamination.

Sample Measured Contami- Measured Contami- Old component nating sample nation °/o calculated values

H-3 C-14 0-18 H-3 C-14 0-18 C-14 0-18 KR2/T1 4.0 10.2 -14 KR2/T7 17.0 43.6 -12.7 23.5 6.7 -14.4 KR3/T3 4.7 14.0 -13.6 KR3/T5 24.9 63.1 -12.8 18.9 4.4 -13.8 KR3/T2 9.6 26.4 -13.0 , , , ,

38.5 9.3 -13.1 KR3/T1 11.3 28.4 -13.7 , , , ,

45.4 2.4 -14.2

All the calculated values are less than 10 pmC, representing the same magnitude as tritium deficient samples from KR2/T4 which have the lowest measured C-14 in dissolved carbonate. "Uncontaminated" values for 80-18 are also predicted in Table 5-4 and the corrected positions for the old component are shown in Figure 5-17d with dashed arrows. The new positions of old components correspond fairly well to the position of KR2/T4. A corresponding procedure to correct 80-18 in samples KR1/T5 and KR5/BT is not possible. Regardless of the assessment, 80-18 does not decrease below the value measured from KR5/T4. If we rely on the procedure and new results, the age of 10,000 years when the ice sheet retreated from Kivetty corresponds to C-14 (DIC) values below 10 pmC due to carbon isotope dilution and radioactive decay. These preliminary age estimations are given closer attention once the carbon isotope evolution has been adjusted by reactions using mass-balance calculations (Chapter 7).

5.5.6 Strontium isotopes

Use of the Sr-87 /Sr-86 signature in ground water studies has been discussed in several papers, e.g. by McNutt et al. (1990) and Bullen et al. (1996). Although Sr-87 is radiogenic, a daughter product of Rb-87 decay, the extremely long half life (5·1010

years) of the decay process makes the Sr-87/Sr-86 signature of Sr sources essentially stable on the time scale of groundwater evolution. During water-mineral interaction the mineral's isotopic value will be reflected in the water's isotopic value, which is dependent on the amount of Rb in the mineral or, more realistically, minerals. Rubidium

65

is concentrated in potassium rich phases. Fractionation of Sr isotopes as a result of mineral precipitation is negligible.

Generally the Sr-87/Sr-86 ratio in the groundwater at Kivetty (varying from 0.72218 to 0.78554) is much higher than in the ocean (0.7092), reflecting a radiogenic Sr source such as dissolution of potassic, Rb-rich minerals like biotite and potash feldspar. Radiogenic Sr shows an initial increase with Cl content (Fig. 5-18), suggesting weathering of potassic silicates during recharge near the surface, especially biotite, which often has an extremely high radiogenic Sr signature (e.g. McNutt et al. 1990, Wallin and Peterman 1994, Blum & Erel 1997, Bullen et al. 1997). The highest Sr isotope ratio has been observed in a shallow well (Liimatainen) sampled above the depth of general occurrence of fracture calcite, which could easily release less radiogenic Sr. During the progress of groundwater evolution, the relative enrichment of Sr-86 with chloride refers to Sr release from the dissolution of low rubidium-bearing minerals such as calcite or plagioclase.

Unfortunately no Sr isotopic ratios from fracture calcite are yet available, but the tendency of calcite to dissolve much faster than silicates suggests that the decrease of the Sr isotope ratio has resulted primarily from calcite dissolution. However, the level remains relatively high compared with e.g. Romuvaara, where Sr-87/Sr-86 decreases to the 0.720 level (Pitkanen et al. 1996a) or slightly lower, which is usual during long residence times due to equilibration with plagioclase (McNutt et al. 1990). Radiogenic input of Sr during immediate weathering may be so strong that later plagioclase dissolution deep in the bedrock cannot decrease the Sr isotope signature closer to the level typically controlled by plagioclase. No differences could be discerned between the flowpaths or boreholes.

0.79

0.78

0.77

~

'Z 0.76 rJ).

[::::

~0.75 rJ).

0.74

0.73

0.72

•

0

A

~

<>

~ c

X

0 0.2 0.4 0.6 0.8 Cl, mmol/1

D

1.2 1.4

I• Spring • Well A KRl o KR4 c KR5 x KR2 ~ KR31

Figure 5-18. Strontium isotopic ratio versus Cl.

66

5.6 Thermodynamic controls of the groundwater system

The results of equilibrium speciation calculations are used to investigate thermodynamic controls of the water compositions at the Kivetty site. These results must be taken as indicative and not quantitative, since both the thermodynamic and analytical data contain uncertainties and the minerals considered in the following discussion can represent a range of crystallinity in water-rock interaction. Due to problems in redox measurements the saturation states of minerals calculated from redox sensitive parameters such as pyrite and iron oxyhydroxides are uncertain. They are therefore not presented here, although the minerals occur in the fractures and are important in the interpretation of water-rock interaction. The approach concerns minerals considered capable of reaching thermodynamic equilibrium with groundwater. Therefore clearly undersaturated minerals (e.g. fluorite) or incongruently dissolving minerals, such as most rock-forming minerals, are not included.

The results of speciation calculations show that groundwaters from springs and shallow wells are clearly undersaturated with respect to calcite (Fig. 5-19a). The saturation state increases with depth and all borehole groundwaters are almost saturated. Calcite is interpreted to be in equilibrium with groundwater if the calculated saturation index (SI) is ±0.5, which is the case at most sampling depths in the boreholes (i.e. calcite occurs frequently on fracture surfaces without any dissolution textures in the depths of sampled borehole groundwaters - Gehor et al. 1995). The pH measurements are also highly sensitive to external disturbances when the internal buffer in the solution is relatively weak due to low carbonate contents in the samples. Therefore quite large uncertainty limits have been agreed upon for calcite saturation.

Clearly undersaturated groundwaters have been sampled from KR5 at depth 400-500m. KR5 waters also show other immature, young water features such as quite high tritium (about 10 TU) and pure organic derived 8C-13 (about -24%o PDB). The samples probably do not represent in situ conditions, reflecting instead quite shallow groundwater levels. The descent would have been possible during pumping, since pressure measurements (Ni va et al. 1994) show that the sampling level is connected to the uppermost part of the borehole.

The calculated partial pressure for C02 is typically high in near-surface groundwaters and decreases to around atmospheric values (-3.5) in the borehole samples (Fig. 5-19b). High partial pressures in shallow groundwaters are derived from the soil where high concentrations of C02 are produced by plant-root respiration, and the decay of plant debris corresponds to low 8C-13 values. The depletion of p(C02) deeper in the bedrock may result from carbonic-acid promoted dissolution of fracture calcite (corresponding to increasing calcium and bicarbonate contents, reaction 5-2 under heading 5.5.5) and silicates (corresponding to general cation enrichment). KRl samples with high pressure values also have anomalously high dissolved carbonate concentrations compared with other borehole groundwater samples.

Weathering of rock-forming minerals is a source of dissolved silica. The solubility of different Si02 phases decreases with increasing crystallinity of the solid SiOz phase. The kinetics of precipitation decreases in the same order and is generally low compared with

67

carbonate reactions. Therefore dissolved silica contents in the groundwater are normally highest in the weathering zone and tend to decrease with longer residence time. Samples from shallow wells consistently show the highest saturation indices of chalcedony and also the highest silica concentrations at Kivetty (Fig. 5-19c). Bedrock groundwater samples have lower values, but do not show any clear trend. Groundwater samples are generally oversaturated in respect of amorphous silica. The poorly crystalline chalcedony phase seems to be in equilibrium with silica activity.

a) -5.0 Slcalcite

-3.0 -2.0 -1.0 0.0 1.0 b) -5.0 -4.0 0 0+-----~--~--~~~~----~

• • ••

200

600

<> <> ~6'\0

D 0 D

0

66

D <> X

200

600

D 800~------------------~--~~ 800 --1....---------....1..---------------___J

c)

0.8

0.6

0.2

0.0

-0.2

X

~

~

c. !.:. ••• • D • X

• • CtJ <>

<>

<>

0 0.2 0.4 0.6 0.8 Cl, mmol/1

• 1.2 1.4

I• Spring • Well A KRl o KR4 c KR5 x KR2 * KR31

Figure 5-19a. Calcite saturation index with depth. Waters within boundaries± 0.5 are considered saturated in respect of calcite. Figure 5-19b. Calculated logaritmic partial pressure of C02 with depth (-3.5 = atmospheric pressure). Figure 5-19c. Saturation index of chalcedony as a function of Cl.

68

During weathering of rock-forming minerals, aluminium is practically insoluble in normal pH conditions and is conserved in clay minerals or more probably in some amorphous aluminosilicate analogue due to the slow kinetics of formation of crystalline aluminosilicate clays. The interrelationship of clay minerals and their stability has been studied using activity diagrams (Fig. 5-20) based on the data of Helgeson (1969) and extrapolated to 7°C by polynomial regression. The constant value of logarithmic silica activity was taken to be -3.7 in calculating the diagrams.

The calculated activities plot in both the kaolinite and montmorillonite fields, whereas illite and sodium-montmorillonite are not favoured. Shallow groundwaters (springs and wells) are clearly in the kaolinite field whereas the sorption of Mg and Ca in clay minerals, and further the stabilising of montmorillonite, is suggested by the chemical conditions of the borehole ground waters. Sodium is released in solution according to the diagrams. The disfavour of illite is in accordance with clay mineral observations from fractures dominated by montmorillonite and kaolinite. In that case potassium should be sorbed in the montmorillonite phase.

a)

7.0

+,._ 6.0

= ~ ,g

+,..!..,5.0 = z ~ 0

~ 4.0

3.0

2.0

6.0

Na-Mont I b)

15.0

14.0

,._ 13.0 D +

= M ~12.0 430 6 ~X ~

Kaolinite M

~ i t.A11.0

6~ eJI

610.0 <> eJI

0 Ca-Mont ~ 9.0

•• 8.0

7.0

• • 6.0

10.0 14.0 18.0 Log( Ca2)-2Log(HJ

Kaolinite

• • Ca-Mont

• •

6.0 10.0 14.0 18.0 Log(Ca2)-2Log(HJ

c)

15.0

14.0

,._ 13.0 +

= ~12.0 ~ M

t.A11.0 eJI

610.0 <>

eJI 0 ~ 9.0

8.0 • • 7.0

6.0 -1-------r---.------,--M-----i

0.0 2.0 4.0 6.0 8.0 10.0 Log(K)-Log(H+)

I• Spring • Well J:J. KRl o KR4 c KR5 x KR2 ~ KR31

Figure 5-20. Aqueous activity diagrams for some clay minerals at 7°C, 1 bar and -3. 7 as the logarithmic silica activity. The field boundaries are calculated from the data of Helgeson (1969).

69

6. DISCUSSION OF THE GEOCHEMICAL SYSTEM

6.1 Hydrogeochemical evolution

Hydrogeochemical evolution along a flowpath is the sum of m1x1ng of different groundwater types and water-rock interaction. Considerable mixing usually requires changes in hydrogeological conditions, or transients caused by glaciations and sea-level movements, as is the case at Olkiluoto and Aspo (e.g. Pitkanen et al. 1994, 1996b and Laaksoharju & Wallin 1997). In a stable hydrogeological system physico-chemical conditions vary along the flowpath and water-rock interaction plays a major role in hydrogeochemical evolution. Kivetty is assumed to represent a more or less stable hydrogeological system. Weichselian glaciation and postglacial uplift may have caused mixing of some older water type with descending recharge water, but e.g. postglacial sea water intrusion into the bedrock is impossible in the Kivetty area, as this was already an island when the ice sheet retreated (e.g. Eronen & Lehtinen 1996). The pre-Weichselian Quaternary geology of the area, such as the relation of the area to the saline sea during the Eem interglacial period, is poorly known.

Changes in predominant ions (e.g. Fig. 5-4) and principal component analysis (Fig. 5-1) show the entire evolutionary trend, which is dominated by the geochemical circulation of carbon, mixing of modem climate meteoric water with some glacial water, and a trace of Na-Ca-Cl type saline water. The proportion of the saline end-member is so low that principal component analysis could not recognise it uniformly due to the small variance of groundwater compositions within the data matrix (i.e. total compositional variations) of the Ki vetty site. However, even the slight enrichment of saline elements is difficult to explain without leaching of some old, possibly brine-based saline end­member water in a low temperature shield area. This is discussed later in the chapter.

Flowpath I (cf. p. 28) seems to cover the entire evolutionary trend, but detailed study of the chemical data shows differences in chemical behaviour between boreholes. This is valuable information when considering the geochemical implications for hydrogeology (Section 6.2). In general, the evolutionary path closely resembles that of jlowpath I at Romuvaara (Pitkanen et al. 1996a), indicating that the same major processes have prevailed at both sites. The most distinct differences are in Sr and 8C-13 isotope data, reflecting differences in mineralogical compositions between the sites. The Sr-87 /Sr-86 ratio generally shows higher values at Kivetty, decreasing to roughly the 0.730 level with increasing residence time (at Romuvaara the lowest level is 0.720). The 8C-13 level is concentrated below -20%o PDB at Kivetty and is somewhat over -20%o PDB at Romuvaara. Unfortunately no isotope analyses from minerals or whole rock are yet available, but several probable reactions can be predicted from the water chemistry, specified isotope composition, lithogeochemistry, mineralogy and the description and distribution of fracture minerals.

6.1.1 Recharge and carbonate evolution

The initial enrichment of main cations (Na, Ca, Mg, K) in the recharge zone is considered to result from silicate weathering. At Kivetty a high strontium isotope ratio is

70

indicative of water-rock interaction. The local granitoids are potassic and rubidium-rich compared with the tonalitic gneisses at Romuvaara. The steep rise of the strontium isotope signature in low chloride waters is interpreted to result from hydrolysis of rubidium and potassium rich silicates, particularly biotite that is annitic in composition (Gehor et al. 1995) (reaction 6-1) and potash feldspar in shallow, weathering depths, where dissolution is induced by high concentrations of carbonic acid from organic respiration and oxygen. In addition, easily soluble calcite does not consume carbonic acid since it is missing at shallow depths. The interpretation is consistent with 1) high potassium, magnesium and silica contents, 2) a high Sr-isotope ratio (Liimatainen well) in shallow groundwater, 3) a low remaining 8C-13(DIC) value (alkalinity production only with trace calcite dissolution), and 4) occurrence of silicification and ferric oxyhydroxide coatings on fracture surfaces in the upper parts of boreholes. These reactions probably start in the overburden peat and till as reflected by 8C-13. Unfortunately the overburden is inadequately studied, but generally in Finland the chemical weathering of till is slight due to the cool climate. The till contains primary rock minerals such as biotite, the cover being formed during glaciation under the ice sheet from preglacial overburden and rock debris disintegrated by the ice (Hirvas and Nenonen 1987). Therefore it may also contain old organic matter free from radiocarbon.

Plagioclase shows a typically low Sr-isotopic signature due to its low Rb content. Plagioclase hydrolysis (reaction 6-2) accompanying potassium silicate weathering cannot be excluded and is in fact presumable, according to e.g. the results of Lasaga (1984), whereas annite, a biotite with a high Fe/Mg ratio, shows a relatively high degree of reactivity in a slightly acid, oxygen-bearing weathering environment (e.g. Acker and Bricker 1992, Malmstrom et al. 1995, Blum & Erel 1997, Bullen et al. 1997). Therefore the radiogenic strontium production from biotite exceeds the low Sr signature released from plagioclase dissolution. In addition, the load of potassium in the third principal component with tritium, and their positive correlation, connect the sources of potassium ( + Rb and its daughters) to quick, surface related processes:

2KMgo.7sFe2.2sA1ShOIO(OHh + 1.12502 + 5C02 + 9.25H20 --7

annitic biotite organic derived carbonic acid (6-1)

2K+ + 1.5Mg2+ + 5HC03-+ 4.5Fe(OH)3 + AhSh0s(OH)4 + 4Si02 ferric kaolinite silica

oxyhydroxides

Biotite is more phlogopitic (Mg!Fe ::::: 111) in the bedrock of Romuvaara (Gehor et al. 1996) than at Kivetty, and therefore more resistant to weathering and a possible reason for smaller radiogenic Sr-isotopic signatures in groundwater.

The depletion of radiogenic strontium in the upper part of the bedrock is thought to result mainly from fracture calcite dissolution, which masks the effect of minor contemporaneous silicate hydrolysis. Biotite reactivity is also reduced after oxygen consumption, compared with plagioclase. Calcite dissolution is supported by the stabilised strontium isotopic ratio when calcite saturates (Fig. 6-1 ). The stabilised still elevated Sr isotopic signature with increasing chloride in deeper parts of the bedrock reflects the small amount of plagioclase dissolution during long residence time. It may also reflect the signature of a saline source contributing to the groundwater. This may at

71

any rate reflect the Sr isotope signature of the granitoids of the site, if the saline end­member water has been derived autochthonously in hydrothermal conditions by water­rock interaction.

Several studies (e.g. Nordstrom et al. 1985, Fritz et al. 1989, Banwart et al. 1994, Pontes 1994, Pitkanen et al. 1994 and 1996a) based on 8C-13 in dissolved inorganic carbon have shown the importance of organic derived carbon as a source of dissolved carbonate in addition to calcite dissolution. This would include plant respiration and decay of plant debris in the recharge zone, respiration of dissolved organic compounds (reaction 5-1, p. 58) and possible fermentation (reaction 5-3), all increasing acidity and starting carbon dioxide promoted hydrolysis (e.g. reactions 6-1 and 6-2). Silicate weathering in the recharge zone consumes protons to some extent and produces alkalinity above the zone where descending groundwater infiltrates calcite bearing fractures.

Nao.75Cao.25Alt.25Sh.750s + 1.25C02 + 1.875H20 ~ plagioclase (An25) (6-2)

0.75Na+ + 0.25Ca2+ + 1.25HC03- + 0.625AhSh05(0H)4 + 1.5Si02 kaolinite silica

Calcite is a common fracture mineral over most sampling intervals in deep boreholes. The importance of calcite at Kivetty as a source of dissolved carbonate is not as clear as in the groundwaters at Romuvaara. Dissolved carbonate increases from well waters only slightly before calcite reaches saturation (Fig. 6-2), supporting the interpretation of the importance of silicate weathering in bicarbonate production. However, the increase of 8C-13 from organic derived values (about -23 ... -25%o PDB) indicates some calcite dissolution (Fig 5-17a). The process does not seem to be as strong as at Romuvaara, but mass-balance calculations will determine the quantity. The 8C-13 signature of fracture calcite may also be lower than at Romuvaara.

0.79 ...,..---------------.,-----,

0.78

0.77

">&:>

~ 0.76 'JJ [:::

~ 0.75 'JJ

0.74

0.73 •

•

0

t:.

X

Q

t:.D D

X 0.72 +----,....-----.------r-----,-----+-----1

-5.0 -4.0 -3.0 -2.0 -1.0 0.0 1.0 Slcalcite

I• Spring • Well A K.Rl o KR4 c KR5 x KR2 ~ KR31

Figure 6-1. Strontium isotopic ratio vs. saturation index of calcite.

a)

2.0

s 0 5 5 1.5

6 u + 0 1.0 u ::r::

0.5

0.0

c)

9.0

8.5

8.0

:a. 7.5

7.0

6.5

6.0

5.5

~ ~

A A

A I"" Ill!. D tSo <>>

[ A <>1 ~X X -•

-0 0.2 0.4 0.6 0.8

Cl, mmol/1

1-X

X C (])

li~ ~A I>'

I -cl~ I~

• •• • I •

0 0.2 0.4 0.6 0.8 Cl, mmol/1

72

D

1.2 1.4

D

1.2 1.4

b)

0.5

0.4

--0 5 5 0.3 ~ u

0.2

0.1

0.0

d)

0.0

~

~-2.0 00

-4.0

-6.0

~

a X

<»A

D ~

A i> A a. [

~r:r<x X

• IQ A

•

~

•

0 0.2 0.4 0.6 0.8 Cl, mmol/1

0 X X C

IJf»<> IX a

<>D A

I

• • • ~

0 0.2 0.4 0.6 0.8 Cl, mmol/1

I• Spring • Well A KRl o KR4 c KR5 x KR2 ~ KR31

D

1.2 1.4

D

1.2 1.4

Figure 6-2. Figures showing trends of dissolved carbonate, Ca, pH and saturation state of calcite vs. Cl.

Combined examination of dissolved carbonate, Ca, pH and the saturation state of calcite (Fig. 6-2) gives a rather confused picture of groundwater evolution at Kivetty. Variable behaviour can even be observed inside flowpath I, as mentioned in the discussion on carbon isotopes. The data from borehole KR1 show typical carbonate-calcium enrichment-depletion curves with increasing Cl (Fig. 5-16a). Precisely at the culmination, calcite reaches its saturation and increasing pH stabilises. This is generally interpreted in terms of the buffering character of fracture calcite with respect to the carbonate chemistry in the groundwaters (e.g. Nordstrom et al. 1989, Pitkanen et al. 1996a). The 8C-13 signature of dissolved carbonate also increases with the initial steep rise of carbonate content (cf. Fig. 5-17). This indicates calcite dissolution, because fracture calcites frequently contain heavier carbon than organics (e.g. Frape et al. 1992) which puts their signature on the recharging water. Higher values (> -10%a PDB in calcite) are typical of the upper part of the bedrock and of low-temperature calcites, e.g.

73

at Olkiluoto (Blomqvist et al. 1992), this part normally coming into contact with water undersaturated in calcite.

The highest carbonate values of groundwater were measured from the sampling intervals of KR1, reflecting the most intense carbon cycling at the site. Carbon isotopes assume both biogenic carbon and calcite derived carbonate production (Figs. 5-17a, 5-17b and 5-17c). Biogenic carbon seems to dominate in relation to calcite. The evident precipitation of calcite during later evolution decreases the content of carbonate, calcium and slightly also 8C-13(DIC) in KRl, but anaerobic oxidation of organic carbon is thought to be a more effective process to decrease 8C-13.

The corresponding results for borehole KR4 seem to show only the depletion parts of the carbonate evolution curve (Fig. 6-2). Calcite saturates at higher pH (about 8.3) and at a lower carbonate content than in the data of KRl. Carbonate shows slight depletion with increasing pH to the level of 8.5-9, suggesting minor calcite precipitation with silicate hydrolysis. The whole carbonate evolution takes place over a much smaller Cl enrichment than in KRl. This indicates only a trace input of the saline component and possibly an extra calcium contribution to the groundwater. Thus the pH rises to quite a high level before calcite saturation is reached. Later during the evolution calcite equilibrium is disturbed only weakly and calcite precipitation (i.e. the sink of carbonate) is therefore minimal. The chemical evolution differs from KR1, reflecting mainly a hydraulically separate situation for the boreholes. Only the sample from KR1/T4 (point inside the circle in Fig. 5-17c) shows fairly similar chemistry to KR4, suggesting a common flowpath.

The representativity problems in samples from KR5 make it difficult to interpret the results as a whole. The recent results of flowmeter measurements (Rouhiainen 1996b) suggest an occurrence of parallel fractures/fracture zones in the immediate vicinity of the borehole, which balances hydraulic head values between the packed-off sections. In this case a hydrologically very disturbed situation may develop where water circulates effectively in the fracture zone and borehole before installation of the multipacker system. Thus most of the samples are mixtures of shallow young groundwaters and old groundwaters as suggested by tritium and Cl contents. Five of the samples were accepted as representative enough for geochemical interpretation and modelling as trend-setting groundwaters. Two of them, KR5/T1 and T2, are interpreted from their tritium, 80-18 and Cl contents to be 100% young water, although their original location may be totally different from the sampling interval. The uppermost sample (KR5/T6) represents fairly young water, still undersaturated with regard to calcite (Fig. 6-2) although Cl is somewhat enriched. Furthermore, a small tritium content (2.4 TU) and quite low C-14 (47 pmC) suggest contamination by young shallow groundwater.

The two remaining samples, KR5/T4 and KR5/BT, with the highest Cl content in borehole KR5, correspond to the carbonate evolution curve of KRl (Fig. 6-2). However, the 8C-13 values deviate (Fig. 5-17a) and it seems that KR5/T4 probably has too low and KR5/BT too high a 8C-13 value to adjust via calcite and organic carbon reactions to the same evolution model as the samples from KRl. The samples represent a long residence time, judging from the stable isotopes of water (low 8H-2 and 80-18 indicate some meltwater in-mixing from the ice sheet) and C-14, and the salinity reflects the most matured evolution state at the site. The bottom sample of KR5 has both the highest

74

Cl and C-13 contents at the site. Highest Cl should also indicate high age, but C-14 is only moderate (25 pmC). However, judging from its small tritium content (1.65 TU) the sample is slightly contaminated by young water which has also elevated the C-14 content of dissolved carbonate. Therefore the groundwater in the bottom part of KR5 most likely represents a high age. This is a prerequisite for carbon isotope exchange between calcite and groundwater, i.e. calcite recrystallisation would also be possible, as is suggested by the high 8C-13 of the sample and correspondingly (Pitkanen et al1996a) by the enriched 8C-13 values in the case of Romuvaara.

The carbonate evolution of samples from borehole KR2, i.e. jlowpath 11, is expected to follow calcite dissolution-precipitation behaviour, but the samples represent the saturation state of calcite as do those from KR4. The elevated 8C-13 values in relation to recharge values, as well as the congruent behaviour of carbonate and calcium, support the controlling role of calcite in carbonate hydrochemistry (Fig. 6-2). Low C-14 and depleted 80-18 values are also typical for this borehole, indicating long residence times and far advanced equilibrium with the bedrock. The high level of 8C-13 of KR2 links the samples with KR4, suggesting a similar environment for the evolution of carbonate, and small TDS enrichment causes relatively high pH before calcite equilibrates in both boreholes.

The groundwater samples from KR3, i.e. jlowpath Ill, represent either the enriched or the depleted part of the carbonate evolution curve. Although calcite is slightly undersaturated in the most chloride-rich samples it is presumed to reach saturation. The samples have quite high tritium contents in spite of the low C-14 contents in dissolved carbonate (cf. section 5.5.5.1). Young groundwater contamination is obvious and may decrease pH values enough to make calcite undersaturated in the speciation calculations.

From the examination of carbonate related groundwater chemistry it can be concluded that calcite is the main controlling phase of dissolved carbonate and pH deep in the bedrock. By contrast, in recharging zones and at shallow depths in the bedrock, alkalinity and pH evolution are thought to be connected to organic respiration and silicate weathering.

6.1.2 Salinity changes

The initial enrichment of major cations at the early stage of evolution may well be explained by silicate weathering. Sodium increases regularly with chloride as calcium continues its rise to the carbonate maximum, but K and Mg turn and start to decrease. Calcite precipitation may act as a sink for calcium during carbonate consumption, which releases protons (reaction 6-3) for silicate hydrolysis.

(6-3)

Plagioclase probably dominates the silicate dissolution deep in the bedrock (e.g. McNutt et al. 1990 and the discussion at the start of this chapter). The dissolution acts as a source for sodium (reaction 6-2). Calcite may to some extent also be a sink for magnesium, but a significant sink of all three ions should be ion exchange or precipitation of secondary silicates during hydrolysis reactions, e.g. montmorillonites or some poorly crystalline alumino-silicate phases.

75

The reaction system can be quite complex. It may consist partly of pure cation exchange, especially in the zone of calcite dissolution, and partly of primary silicates weathering to clay minerals, contemporary release of sodium, and sorption of other cations, which can also be considered ion exchange in a wider sense, although this is irreversible. Other cations than Na are depleted after initial enrichment (Fig. 5-3), supporting ion exchange in clay minerals (reactions 6-4, 6-5 and 6-6), which is also favoured by activity diagrams (Fig. 5-20).

(6-4)

(6-5)

(6-6)

Cation exchange could contribute additional carbonate alongjlowpaths. In this case the uptake of Ca and release of Na from exchange sites on clay minerals causes dissolution of fracture calcite. This may also partly explain late stage enrichment of 8C-13 (KR5/BT) and be part of the isotope exchange process. The adsorption of potassium may be irreversible, thus the uptake of potassium might better be termed "illitisation" (Appelo & Postma 1993). The importance of each process will be examined more closely with mass-balance reaction calculations in conjunction with carbon isotopic data to determine the extents of mass transfer of each reaction.

The thermodynamic stability of clay minerals and the behaviour of cations are also important for performance assessment, because groundwater evolution favours alteration of Na-montmorillonite, the major mineral of the bentonite buffer, to (Ca,Mg,K)-montmorillonite. Nevertheless the total mass of the elements is so small that ion exchange is thought to remain at quite a low level, taking into account potential water and bentonite volume (e.g. Melamed et al. 1992, Muurinen et al. 1996). However, accurate amounts should be checked in detail.

The enrichment of salinity and chloride in the groundwaters of Kivetty are quite small (max. TDS 200 mgll with 1.2 mmol/1 Cl). Nonetheless the complex question of the origin of chloride and the enrichment of sulphate, calcium and potassium with higher Cl needs addressing. Chloride is not the main constituent of silicates, but the rocks contain it in concentrations of about 300 ppm (Gehor et al. 1995). No kinds of salts are observed, but e.g. secondary fluid inclusions are frequent in quartz grains (Pitkanen et al. 1992).

The carbonate evolution curves up to the point at which carbonate content and pH stabilise (Fig. 6-2), showing the potential chloride content to be dissolved from the rock by hydrolysis of silicates. This is about 0.3 mmol/1 and part of it could be from fluid inclusions. Chloride dissolution from silicates is not a convincing process once the dissolved carbonate is fixed (no proton production), because mass transfer by hydrolysis of silicates is strongly retarded along the jlowpath. Sea-water is not a probable source according to the Br/Cl ratio and isotope results. A more plausible explanation would be some leaching process of rock salts, fluid inclusions or grain boundary salts as suggested for the groundwaters of Stripa (Nordstrom et al. 1989), Lac Du Bonnet (Gascoyne et al.

76

1992), Olkiluoto (Pitkanen et al. 1994) and Romuvaara (Pitkanen et al 1996a). Especially fluid inclusions in fracture calcite form a potential source at Olkiluoto, according to the results of Frape et al. (1992) and Blyth et al. (1998). The origin of these salts at Kivetty should be ancient hydrothermal activity. Strong hydrothermal activity has occurred e.g. in the bottom part of borehole KR5, and remnants of the same activity have also been observed in other boreholes (Gehor et al. 1995). The alteration has been strongly oxidising, as a result of which hematite has displaced biotite around fractures and feldspars have altered to clay minerals. Calcite and quartz have also been identified as secondary minerals from this zone and sulphides are missing. The description of the alteration would be very suitable as a primary source process for releasing chloride, sodium, calcium and sulphate and also potassium, all of which are to some extent enriched in most saline samples. These elements need an earlier released salt host for leaching into the groundwater, as modem-type low-temperature interaction in reducing conditions cannot explain their enrichment in silicate rock.

6.1.3 Redox related processes

The water chemistry and mineralogy contain imprints of progressive redox reactions down-gradient from the recharge zone. The most obvious observations of these are limonite coatings in fractures in the upper part of the bedrock and traces of hydrogen sulphide in groundwater samples. The redox measurements show even high positive values, but the occurrence of hydrogen sulphide does not support positive values over the measured pH range, indicating instead slightly negative values. Similarly, iron oxyhydroxide formation, i.e. indications of iron-rich biotite weathering, and low 8C-13 values in recharging groundwater suggest strong oxygen consumption in shallow depths and reducing rather than oxygenated or post-oxic conditions at groundwater sampling depths in boreholes. On the other hand, low pumping rates used in sampling require strongly reduced species in order to buffer oxygen diffusion through plastic sampling tubes in the atmosphere and to conserve reducing conditions in the water extracted before redox measurement (Snellman et al.1995b). Unfortunately hydrogen sulphide or iron(II) contents in groundwater samples seem to be generally too low to buffer in situ redox conditions during sampling. However, one indication of a prevailing low redox level is measurements made during preliminary site investigations (Teollisuuden Voima Oy 1992). The sampling technique allowed a much higher pumping rate and -300 ... -350 mV was reached in the upper part of KRl (sample KR1/T7 in App. 6) from the most conductive section of the borehole (Pitkanen et al. 1992).

Biotite (annite) weathering and organic carbon respiration are considered to be the main consumers of oxygen in recharging groundwater in the soil layer and at shallow depths in fractures. Dissolution of iron sulphides as pyrite is also possible, but not necessarily favoured, because the 8S-34(S04) values for recharge waters show a general signature of precipitation of non-industrial regions (according to Fritz and Pontes 1980) and the value also remains quite stable after a short residence time in low chloride samples from the upper sections of the boreholes.

Instead of pyrite oxidation, sulphate reduction should be considered after groundwater infiltration of the bedrock. Occurrence of small grained to cryptocrystalline or platings of pyrite on fracture calcite (Gehor et al.1995) support low temperature sulphide

77

production and precipitation from groundwater. Fundamental S04 reducing bacteria were recently observed by Haveman et al. (1998) as well as iron reducers which are also needed to reduce a potential iron source, i.e. ferric oxyhydroxides, and to release ferrous iron for pyrite precipitation (reaction 6-7). Sulphate decreases along the flow from springs to shallow-intermediate depth (cf. Fig. 5-5b, 0 < Cl < 0.15mmol/l). Simultaneously 8S-34(S04) (Fig. 5-13) increases and hydrogen sulphide appears in groundwater samples, indicating microbially catalysed reduction of sulphate coupled to anaerobic respiration of dissolved organic carbon. The analysed sulphide contents are not necessarily the total occurring amount of sulphate reduction (reaction 6-7), because if a source of iron is available iron sulphides will be precipitated (Drever 1982, Plummer et al. 1990, Busby et al. 1991 Appelo & Postma 1993).

(6-7)

Ferric oxyhydroxides on fracture walls form potential reserves for taking part in anaerobic sulphate reduction by organic carbon oxidation. If sufficient sources of iron are missing or the iron release is kinetically hindered, hydrogen sulphide is formed:

(6-8)

The processes are also favoured by 8C-13 depletion observed in samples with saturated calcite in the carbonate evolution trend (cf. Figs 5-17 and 6-2). Occurrence of sulphate and hydrogen sulphide in the same water samples reflects a very negative redox level. Speciation calculations suggest -200 mV ... -300 mV at pH 7.8 to 8.6 and even lower redox values were measured during preliminary site investigations (Pitkanen et al. 1992, Lampen & Snellman 1993). Uranium data also suggest a reducing condition in the deeper parts of the bedrock at Kivetty, as well as the occurence of hydrogen and methane gases.

6.2 Hydrogeological implications

Hydrogeochemical data and the concept of hydrogeochemical evolution together comprise an important tool for evaluating site-specific long term hydrogeological conditions and ground water flow (e.g. Smellie & Laaksoharju 1992, Pitkanen et al. 1996a). The evolution of geochemical parameters should support postulated hydrogeological flowpaths, or at least groundwater types on the same interpreted hydrogeological path should be logically evaluated from each other. If not, both geochemists and hydrogeologists should review their interpretations. In the best case, geochemical reaction models describe the initial and final water along a hydrological flowpath, provided geochemical variables can be used in the flow modelling. In such a case geochemical and flow modelling could together comprise an iterative process for developing a consistent hydrogeological model.

Dissolved Cl and C-14 (DIC) are regarded as apparent age parameters, and both show logical trends of increasing "age" with decreasing hydraulic conductivity (Fig. 6-3). This supports the use of chemical variables in conceptualising the long term groundwater flow at the site. Generally, the heterogeneity of groundwater flow from recharge to different sampling depths is obvious from hydrogeochemical data. Chloride usually

78

increases with depth but the data (Fig. 5-2) contain several exceptions, pointing out the importance of well-conducting fractures/fracture zones for groundwater flow and residence time. On the other hand the highest "ages" of Cl and C-14 are not congruent; on the contrary, the lowest C-14 was measured from a sample containing less than 0.2 mmol/1 Cl. Also uranium data hint to heterogeneities between single flowpaths. This may also indicate heterogeneities in the hydrogeochemical environment, which may help interpret geochemical flowpaths. The inconsistency between Cl and C-14 could also partly result from contamination by young water from upper reaches in the most saline samples, as shown in Table 5-4.

Boreholes KR1, KR4 and KR5 in flowpath I are, according to the present hydrogeological model, closely related (Saksa et al. 1993, Taivassalo & Meszaros 1993, 1994, Saksa et al. 1996). Boreholes KR4 and KR5 locate downstream in relation to KRI and several sampling sections should form a potential flowpath (Table 3-1 ). The head difference between these connections is small, indicating slow flux, but a geochemical evolution should be observable. However, hydraulic flowpaths from KRI to KR4 seem improbable according to geochemical data. Cl contents alone are adequate to show the contradiction, because Cl contents are smaller in KR4/T5 ... T3 than in the potential initial sections in KRI. In addition, C-14 values suggest a shorter residence time in KR4/T3 than for hydrological initial waters in KRI (cf. Table 3-1). Principal component analyses suggest analogously separate hydrogeochemical evolution although differences are small between the samples.

Although chemical parameters do not indicate flow from KRI to KR4, samples from the upper part of KR4 (T6 ... T3) resemble in many of their chemical variables (pH, Cl, S, alkalinity, H-3, 80-18, 8C-13, C-14, Slcalcite) the sample from KR1/T4 which itself differs quite a lot from its environment in KR1 (e.g. Fig. 5-17), representing less matured chemical conditions than the neighbouring sampling sections. The interpretation of a well-conducting structure assembly R10-Rll-R12-R22-R23 (Saksa et al 1993, 1996), one of the most distinct lineament directions of the drilling area, connects these sampling sections. The chemical similarity supports the interpreted hydrogeological connection and indicates quite equal residence time for each sample regardless of the marked vertical difference between samples.

Results relating to most cations, fluoride, carbonate chemistry and carbon isotopes do not favour the idea that the samples of KR5 could be final waters evolved downstream from KRI. The hydrogeological conditions (Saksa et al. 1993 and Saksa et al. 1996) favour the formation of qualitative flow lines from several packed-off sections from KRI to the sampling sections of KR5, e.g. KR1/T5 via RIO and R22 to KR5/BT. The water compositions of these sections also have certain significant similarities such as enriched Cl, depleted 0-18 and moderately low C-14. The same characteristics are observable in KR1/T3 and KR5/T4 suggesting a deeper, more saline, older aquifer having a glacial component below 300m. However, samples from KRI have a higher N a/Ca ratio and lower K content than samples from KR5, reflecting local features in groundwater evolution, not a flow relationship between the boreholes. Principal component analyses also suggest a different geochemical environment for the deep groundwater samples and therefore a separate flow system.

a)

1.E-5

"' 5 ..Jl.E-7 = 0 C.J

~ "0 ~ ::::

l.E-9

l.E-11

••

' . .,. • . ~ •• • • • • • .

0 0.2 0.4 0.6 0.8 Cl, mmol/1

79

-b)

1.E-5

"' 5 ] l.E-7 0 C.J

~ "0 ~

:::: l.E-9

l.E-11

• •

•

- • • • • • -~· • •

·' • • • •

1.2 1.4 0 10 20 30 40 50 60 70 C-14, pmC

Figure 6-3. Cl (a) and C-14 (DIC) (b) contents in groundwaters vs. hydraulic conductivity at Kivetty.

Between boreholes KR5 and KR4 is a local water divider. This is consistent with the difference in groundwater quality and suggests that the boreholes do not have flow connections even at greater depths. All these observations from flow path I indicate flow barriers in hydrogeological structures between the boreholes. The boreholes form their own sub-aquifers where the recharge locates in the near vicinity of each borehole and water has descended steeply to the depths of different sampling sections. This does not necessarily prevent partial evolutionary similarities between the boreholes, as is implied from the similar chemistry in certain hydrogeological structures which may imprint the groundwater with their own signature during evolution.

The water compositions of KR2, flow path 11 do not suggest a direct flow line from the upper part of KR1 according to the hydrogeological model, but suggest chemical similarities between KR2 and KR4. The samples from KR2 also seem to represent the older, deeper aquifer observed below 300m in KR1 and KR5 but it occurs at a shallower depth in the vicinity of KR2 (already at 160m). The groundwater near KR2 shows lower Cl and C-14 values and higher 80-18 values than around KR1 and KR5, but apparently this discrepancy may be partly due to contamination by shallow groundwaters as shown in Table 5-4 and Figure 5-17.

Flow path Ill is interpreted to be isolated from the other two jlowpaths, except that all of them have a recharge area between boreholes KR1, KR3 and KR5. High tritium contents disturb interpretation, but the chemistry corresponds mainly to jlowpath I except for sulphate, which is invariably high in the groundwater samples of KR3 by companson with other boreholes, emphasising the importance of local water-rock interaction.

Water samples from KR3/T3 and T 1 also show signs of old, cold water existing below 350m around KR3. This water type occurs at varying depth in four out of five boreholes, indicating regional formation (Fig. 6-4 ). Weak signs are also observable in water samples from the lower part of KR4, although they are slightly contaminated by young

80

water judging from the tritium contents. The upper surface of the old water in each borehole seems to be dependent on the hydraulic head (App. 2); the higher the value the deeper the occurrence.

Chemical evidence suggests limited hydrogeological subsystems around each borehole. This indicates flow barriers between boreholes at the Kivetty site, although the latest update of the hydrogeological model (Paulamaki et al. 1996, Saksa et al. 1996) and conceptual flow model (Taivassalo & Meszaros 1993, 1994) suggests continuous, subhorizontal flowpaths between boreholes. The barriers can be discontinuities in single structures or between different structures, or an important missing hydrogeological unit which forms a short-circuit between the boreholes. The observation of an old, deep groundwater system may indicate a mtsstng subhorizontal more dynamic hydrogeological unit which determines the upper boundary of the old aquifer, this varying from depth 160m in KR2 via 300m in KR1 and KR5 to 350m in KR3. Alternatively this horizon may represent the lower boundary of the hydraulically more dynamic upper part of the bedrock A subhorizontal zone might roughly follow the contact between porphyritic granite and granodiorite, intersecting the surface along R13. The carbonate chemistry and related isotopes suggest a hydrogeochemical analogy in the south-western part of the site, which may support a hydrogeological connection between KR2 and KR4 (also KR1/T4) by structure assembly R10-R11-R12-R22-R23. The flow barriers suggest a limited extent of each sub-vertical fractured pattern rather than large, continuous fracture zones presented in the hydrogeological models.

s

-15.0

0

100

200

300

.s 400 c.. Q,j

Q

500

600

700

800

-14.0

X

D

60

X 0

D

80-18, %o SMOW -13.0

• ~

Op ~ l:J.l:J.

X

0 D

p l:J.

0

oo

l:J.

l:J.

-12.0

I• Spring • Well A KRl o KR4 c KR5 x KR2 ~ KR31

-11.0

Figure 6-4. Oxygen isotopic composition of groundwater with sampling depth. Vertical lines define the variation of recent groundwaters at Kivetty.

81

7. RESULTS OF MASS-TRANSFER MODELLING

The chemical model computations are performed using the approach initially developed and described by Plummer et al. (1983). The approach is divided into inverse and forward methods. The inverse problem uses observed chemical, isotopic, petrographical and hydrological information at the initial and final points in the system to define mass­balance reaction models that are consistent with the data. The NETP ATH program (Plummer et al. 1994) is an example of the geochemical mass-balance modelling code for solving an inverse problem. In the forward problem an assumed reaction model is applied to an initial condition to predict the chemical composition of water and rock as a function of the reaction progress. EQ3/6 (Wolery 1992) and PHREEQE (Parkhurst et al. 1980) are widely used examples of forward geochemical modelling codes for calculating the mass-transfer.

An advantage of the inverse problem is that any mass-balance model found to be con­sistent with the observed data predicts the final composition of the system. The inverse problem, however, is not constrained by thermodynamic considerations which must at least be checked through thermodynamic speciation calculations at the initial and final points in the system. In modelling the chemical evolution in the regional groundwater system, for which appropriate hydrogeological data are available, most of the reaction information is gained through the combined use of mass-balance and speciation calculations (Plummer 1984).

The forward problem is useful in predicting details of thermodynamically valid reaction paths between initial and final points, provided each path is constrained by the net mass­transfer derived from the inverse problem. The forward method has its greatest advantage in predicting the previously unknown mass-transfer and final water com­position of equilibrium solutions via hypothetical reaction models resulting from changes in temperature, pressure, mixing of waters etc. (Plummer 1984).

7.1 Mass-balance reaction models

7.1.1 General

Mass-balance reaction models are used to determine the importance of interpreted evolutionary processes based on data of hydrogeochemistry, isotopes, mineralogy and speciation calculations. The modelling is an attempt to test the above reaction and mixing hypotheses (Chapter 6) by constructing mass-balance models which describe the changes in chemical and isotopic composition between recharged water and downgradient water samples. The model derived between any points along a chosen flowpath by mass-balance calculations is of the form:

Initial water(s) + "Reactant phases"~ Final water+ "Product phases".

The models computed with the NETPATH program (Plummer et al. 1994) define net geochemical reactions of minerals and gases that can 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

82

modelling approach that follows is similar to that of Plummer et al. (1983) and Plummer (1984) and that applied by Plummer et al. (1990), Busby et al. (1991), Waber & Nordstrom (1992) and in Finnish site investigations by Pitkanen et al. (1994), (1996a) for the data of Olkiluoto and Romuvaara, respectively.

A mass-balance model is defined as the masses of a set of plausible phases that must enter or leave the initial solution in order to define exactly a set of selected constraints observed in a final (evolutionary) water. A constraint is typically the concentration of a chemical element or may also be the conservation of electrons (redox state) or a particular isotope of an element. 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 in phases. The treatment of mass-balance equations, redox state and isotopic calculations with fractionation factors in the NETP ATH code is discussed in greater detail in the report of Plummer et al. (1994).

Plummer et al. (1983) and Plummer (1984) have pointed out that geochemical modelling rarely leads to unique solutions due to the number of assumed mineral and gas phases normally exceeding the number of constraints in the models. The modelling process is best suited for eliminating reaction models from further consideration. The validity of the mass-balance models significantly depends on selecting appropriate phases in the model, and i.e. only phases that occur in the system and that have been observed to show water-mineral interaction features should be considered in modelling. A mass-balance model can be eliminated if it requires the net precipitation of a phase that is known to be undersaturated in the system. Similarly, a model can be eliminated if the predicted isotopic composition of the final water differs significantly from the observed one. The elimination process could produce a best-fit model which is most compatible with observed changes in water chemistry and with rock minerals.

7 .1.2 Restrictions and selection of flow paths

The user needs to evaluate the appropriateness of the steady-state assumption implicit in mass-balance modelling. The system may be in a dynamic state, where 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. This would lead to misleading results. The mass-balance modelling of NETPATH applies strictly to the case of a chemical steady state along the flowpath (Plummer et al. 1994).

Interaction in a low temperature crystalline rock-water system is vital carbonate chemistry, which is sensitive to external changes in the environment. The steady-state may have been significantly disturbed by events at the end of the last glaciation some 10,000 years ago. During the retreat of the ice sheet, melt water probably dominated the recharge. Compared with modem recharge, differences in salinity are possibly so slight that they could only play a minor role in mass-balance calculations due to the low salinities in both recharges, but differences in isotopic compositions are evident and the amount of dissolved C02 may be crucial. Meltwater was more depleted in the heavier isotopes (e.g. 0-18 and H-2) than modem recharge. Differences in carbonate chemistry between deglacial and current recharge water in bedrock may be quite small because infiltration occurs in principal through the same type of till with organic substances.

83

However, the amount of C02 in meltwater may have been lower because organic activity might also have been lower, especially if the waters have recharged under the ice sheet. The carbon-13 level is assumed to have corresponded to the current value, because organic matter in the till layer probably has marked dissolved carbonate. The original carbon-14 level during deglaciation is difficult to estimate. Partly, the recharging C02 may have been derived from organic matter in subglacial till, which was depleted of radioactive carbon ("dead"), and partly C02 may have been in contact with the atmosphere and therefore shown a modem signature during deglaciation.

The steady state assumption may also be violated by elements which can both dissolve and precipitate along a flowpath. Crossing a culmination point in a chemical trend by a single modelling step must be avoided; the step must be split around the culmination point into separate segments in the calculations in order to predict total mass-transfer. The problem concerns especially the reactions of calcite, because bicarbonate and calcium represent a considerable fraction of total dissolved solids in the Kivetty groundwaters. In addition, if the total mass-transfer is not calculated, the model may lead to a serious error in predicting C-isotope composition in discharge water, as the model may not have taken into account the much stronger potential isotopic effect of calcite dissolution compared to precipitation.

However, it is also important to realise that every sample in a general flowpath has not evolved quantitatively to the same extent even though the system is chemically in a steady state. For example, the enrichment of carbon during recharge and later bicarbonate formation depends on the duration of recharge and type of recharge zone, whether directly into the bedrock or via prolonged interaction in the organic soil layer. The latter system can produce more dissolved carbonate and can be isotopically more strongly buffered against calcite dissolution than the former. This kind of situation can explain the highest alkalinity contents in samples from KR1, but not the highest 8C-13(DIC) values in the Kivetty data. Therefore it is important to be able to estimate the flowpath reliably, which may lead to different flow conceptualisation than that provided by hydrogeological data (hydraulic + structural). This may help in further evaluation of the hydrogeological system and/or in interpretation of the palaeohydrogeology.

Table 7-2 (results of mass-transfer on p. 90) shows the most plausible single flow and evolutionary steps used in mass-balance reaction modelling. The selection of initial and final waters in each flow step is based on the discussion in section 6.2 and testing the plausibility of the mass-balance problem by varying the initial water (trial and error method). Vuorimaki spring water has been used as a representative of a typical recharging initial water if bicarbonate increases in the modelled flow step and calcite is approaching equilibrium in the final water. Although the spring represents discharging water, its chemical composition represents a short residence time and only minor interaction with soil. The interaction with organic matter seems minor compared to recharge through less permeable soils or peat layers common in depressions between outcrops. Therefore the recharging in calculations is interpreted to occur in an open system (unsaturated zone) for oxygen gas and organic carbon in order to ensure sufficient carbonate production along flow steps.

84

Only a few cases seem to require taking into account the culmination point in the carbonate evolution due to calcite dissolution turning into precipitation. This is most obvious in the data of KR1. The selected initial water for a sample evolved over the culmination point must be in equilibrium with calcite. In addition the interpretation of hydrogeochemical evolution (especially carbon isotopes) and hydrogeological position have been crucial in the selection of initial waters. Thus the table does not contain any modelled flowpaths between different boreholes. Test calculations have been done between the most potential initial and final waters selected according to the hydrogeological model (chapter 3) and chemical data e.g. KR1/T7 ---7 KR4/T6, T2 and T1, or KR1/T5 and T4 ---7 KR5/BT. These models were not usually compatible with carbon isotope data, or in order to be so they would need unreliable amounts of additional sulphate or calcite recrystallisation (KR5/BT).

7.1.3 Phases and constraints used in modelling

Table 7-1 lists the most obvious phases to be included in mass-balance modelling, their chemical behaviour in reactions based on the discussions above, and the used chemical compositions. The constraints generally used in modelling are Na, Ca, Mg, K, AI, Fe, Si, C, S, Cl, and the redox state. In most of the modelled flow steps calcite is assumed only to dissolve and no carbon isotope fractionation is expected (no C02 escape). Carbon-13 is used as an additional constraint and carbon isotopes are calculated as an isotope mass-balance problem. Sulphur-34 is also selected once as a constraint (KR1/T6 ---7 KR1/T3).

Table 7-1. Selected phases for mass-balance modelling, their chemical behaviour, and composition used in modelling.

Phase Source(+ )/Sink(-) Com_2osition

Calcite +I- CaC03 Organic matter + CH20 Oxygen gas + 02 Carbon dioxide + C02 Pyrite +I- FeS2 Goethite +I- FeOOH Plagioclase An25 + Nao.7sCao.2sAl1.2sSh7sOs Biotite, annitic + KMgo.7s Fe2.2sAlSi30w(OHh Kaolinite +I- AhSi20s(OH)4 Quartz, chalcedony +I- Si02 K-montmorillonite - Ko.33Al2.33Si3.670Jo(0Hh Ion exchange +I- (Ca-Na2)X, (Mg-Na2)X Chloride salt + Na2CaCl4 Sulphate salt + Na2S04

Organic matter was included in the model as a source of carbon dioxide in the recharge zone and as the most likely electron donor for bacterially mediated sulphate reduction deep in the bedrock. The formula, CH20, is used only to denote carbon valence zero. Sulphate had to be added in some of the reductive steps of the anaerobic respiration of

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organic matter as the sulphate content in the initial water was not adequate. The sulphate contents are very low in general, only a few mg/1 with maximum 0.07 mmol/1 (7 mg/1). The accuracy of the analyses becomes significant for the calculations and an exact flowpath. Variation of the input of sulphate during the recharge stage in oxic and post­oxic zones may be relatively large, although the absolute content varies only by some mg/1. In that case the relation of water flow and the heterogeneity of mineral coatings of single fractures, i.e. the distribution of sulphides and flow channels, possible sulphate sources (sulphides and organic matter) in soil or even variation of S04 in rain water, play an important role in sulphate production. Including sulphate salt as a source in the calculations is not considered to damage plausibility too much if the need is limited to a low level not exceeding the highest observed sulphate contents. It should also be noted that the absolute content of sulphur species in groundwater does not necessarily explain the extent of sulphate reduction because of the negligible solubility of iron sulphides such as pyrite. The mass-transfer of sulphur can be estimated by involving sulphur­isotope compositions in mass-balance calculations (Plummer et al. 1990, 1994).

In models in which final water represents ages reaching the time of glaciation, C02 is independently included in the calculations. Organic carbon is reserved in these models for anaerobic respiration reactions. Recharge conditions during melting of the ice sheet differed from those prevailing today, requiring different modelling treatment than recharge during postglaciation (see next section).

Weathering dissolves incongruently rock-forming silicates by hydrolysis reactions producing clay minerals and silica. For biotite and plagioclase, the used formulae are average compositions of the results of mineral analyses (Gehor et al 1995) which are very coherent for both minerals. Biotite is considered to play a significant part in the mass-transfer process, especially near the surface in oxygen-bearing groundwater conditions. Biotite generally contains trace amounts of chloride replacing OH groups in the mineral lattice. Cl content is estimated to be about 0.3 wt.% in biotite according to lithogeochemical and petrographical data (Gehor et al. 1995). The content corresponds to 0.05 mol in the chemical formula of biotite.

Test calculation requires dissolution of 26 mmol biotite and 15 mmol oxygen in 1 kg of water to produce the observed Cl content (1.35 mmol/1) from KR5/BT. The dissolution is considered far too great as it is not at all comparable with the range of dissolved elements other than Cl, and it leads to mass-balance problems. The concentration of dissolved oxygen exceeds the theoretical solubility about fifty-fold, and biotite dissolution needs protons (or carbon acid), about 65 mmol/kg of water, and produces huge amounts of cations (K, Mg, Fe) which should be precipitated. Thus biotite can only be a trace source for chloride in modem conditions. Chloride enrichment in the models is treated by dissolving salt of Na2CaC14 composition. This corresponds to the main salinity difference between the evolutionary stage following bicarbonate depletion and salinity enrichment in the KR5/BT sample seen in Figures 5-5a and 6-2a, and is therefore assumed to be e.g. conservative dissolution of fluid inclusions (cf. section 6.1.2).

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Ion exchange is considered to concern N a, Ca and Mg ions, but K released from silicate weathering seems to be retarded effectively and quickly, according to Fig 5-4 and PCA results, into montmorillonite according to the interpretation of activity diagrams.

7.1.4 Isotopic calculations and initial values for carbon isotopes

NETPATH considers two types of isotopic calculations: isotope mass-balance and Rayleigh calculations. The isotope mass-balance calculation corresponds to the mass­balance problem constrained by conservation of a chemical element and electrons and is in general applicable for processes involving (reactions) the constraining isotope as a source, such as mixing of waters, mineral dissolution or ingassing. When there is both a source and a 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 8C-13(DIC), C-14(DIC), 8S-34(Sol-, S2-/HS-) and Sr-87 /Sr-86 values 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 value to examine differences between the fractionating differential problem of isotopic evolution and the mass-balance result. For special valid cases where isotopic data are correctly treated as isotope mass-balance problems, the final modelled isotopic composition will always equal the final observed value. For general cases of isotope evolution the calculations usually involve comparison of sensitivity of isotopic and compositional data of selected phases.

Carbon and sulphur isotopic calculations are applied to the Kivetty data. Equilibrium fractionation factors, identified as Mook factors (Mook 1980), are used for the C-13 system, and C-14 fractionation factors are initially defined as twice those for C-13. 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 water

- the fractionation between precipitating calcite and water

- the mass-transfer of carbon

- the average isotopic composition of carbon sources.

In radiocarbon dating this procedure adjusts the initial C-14 composition, A 0, for the reactions in calculating the C-14 composition of the final water, And, but not the radioactive decay. Radiocarbon dating is then applied to the final water on the flowpath using And and the measured value, A, according to the equation:

5730 (A ) J).t(years) = --ln ____!]!!_ • ln2 A

(7-1)

87

The adjusted C-14 age is the travel time, in years, between the initial and the final water sample. If the initial water is recharge water, the age represents the time of the water being groundwater, isolated from the unsaturated zone. To find the actual age of the final water, the calculated age must be added to the residence time of the water in the unsaturated recharge zone, but in the case of Kivetty and generally in areas of thin overburden this time is negligible compared to travel times in bedrock, as is shown by high tritium contents in soil groundwaters.

The isotope fractionation factor for the sulphur system applies to prec1p1tation 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 Kivetty data, the correlation introduced by Plummer et al. (1990) is used to estimate 8S-34(H2S) based on the observed sulphur isotopic composition of dissolved sulphate and water temperature:

8S-34(H2S) = 8S-34(S04)- 54+ 0.40t, (7-2)

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.

Carbon isotope calculations require assumptions because isotope data concern dissolved inorganic carbon and some gas samples (Cfu and C02). No data are available from fracture calcite or biogenic organic carbon, but estimated values are essential for mass­balance reaction calculations. The 8C-13 value of organic carbon derived from the biosphere is assumed to be -25%o PDB, which is a general estimation (e.g. Plummer et al. 1990, Pearson et al. 1991). Organic carbon derived carbonate seems wholly to dominate shallow groundwaters from soil and bedrock. The data from shallow depths from bedrock and deep in borehole KR5 (T2, Tl) also indicate lower values ( < -23%o PDB) than analysed for dissolved carbonate from springs in soil (-22 ... -23%o PDB, corresponds to open system equilibration of soil C02 according to Fritz et al. 1989). This supports the assumption that bedrock groundwaters represent a closed system and C02 escape and fractionation of 8C-13 is considered unlikely, which could enrich the 8C-13 value of dissolved carbonate. The C-14 value for dissolving calcite is assumed to be 0 pmC.

The main uncertainties in carbon isotope calculations are the reliability of the 8C-13 content of dissolving calcite, the initial C-14 fraction of organic (biogenic) carbon and initial waters (A0) in each case. The assessment was performed using certain final waters in test calculations, which represent the most extreme biogenic-carbon derived carbonate chemistry and young flow age, i.e. high tritium content and short advanced evolutionary stage (Fig. 5-16b and 5-17b ). The discussion of carbon isotopic composition (section 5.5.5) suggests that by using samples from KR3/T5, KR5/T2 and T1 it would be possible to determine the mean C-14 of organic carbon and mean 8C-13 of calcite. All the samples represent a young age of less than 50 years, which in practice

88

is less than the accuracy of the C-14 method in groundwater studies. This offers a possibility to determine the mean C-14 content of organic carbon by mass-balance modelling on a level which leads to modem ages for the flowpaths from recharge to the above mentioned samples in Table 7-3 (p. 91 ).

An average estimate for the 8C-13 value of calcite makes it possible to couple in the same mass-balance problems, because calcite is purely dissolving along these flow steps. The A0 value of the initial water is assumed to correspond to the measured value of Vuorimaki spring (110.2 pmC) in these cases of young flow age. The three uppermost mass-transfer models in Table 7-2 are used in "trial and error" determination of required carbon isotope estimates for dissolving calcite and dissolved organic matter presented in Table 7-3. The iterated 8C-13 value for calcite is -5.7%o, which well satisfies the mass­balance models and is used systematically in other models. The value is appropriate with respect to data collected from crystalline bedrock sites (Fritz et al. 1989, Blomqvist et al. 1992, Frape et al. 1992, W allin and Peterman 1994) and comparable to the estimate used for the Romuvaara site (-6.7%o, Pitkanen et al. 1996a).

The iterated value for radiogenic carbon in dissolved organic matter is not single-valued. The sample from KR3/T5 assumes 50 pmC whereas the samples from KR5 assume about 35 pmC for organic carbon dissolved in the present day recharge. The difference between estimates may result from the type of recharging area of both boreholes. KR3 has been drilled into the bedrock under a till-covered hillside and KR5 under a bog depression, which can feed more organic carbon derived from older peat into recharging water than water can leach from till soil. The potential recharging area of borehole KR1 represents a similar overburden to that of KR3, whereas KR2 and KR4 correspond to KR5. Therefore 50 pmC of dissolved organic matter is, in addition to KR3, also applied to the reaction models of borehole KR1, and 35 pmC to the models of KR2, KR4 and KR5.

Because organic carbon seems to buffer the carbon isotope signature during recharge without any significant fractionation of C-13, the average age of organic matter is crucial in C-14 age predictions of flowpaths. In addition, radiocarbon dating is not disturbed although evolution begins in the recharge zone, which is open to oxygen, because the recharge stage in overburden is very short according to tritium results. The determined mean fraction of C-14 in organic matter in the recharge zone can be used only if the final water is also assumed to be young (dynamic system). Probably the mean values for organic carbon were higher during earlier recharge. Thus the amount of radiogenic carbon in the organic matter in the overburden increases with the age of the final water if the initial water is originally recharge water, i.e. the age of the final water corresponds to the age of organic matter, calculated backwards from the determined modem C-14 of the dissolved organic carbon in the initial water (AMDoc = 35 or 50 pmC at Kivetty). The C-14 content of organic carbon during recharge (ARDoc) is calculated according to the adjusted C-14 age of the final water using as many iterative mass­balance computations as necessary to converge the adjusted C-14 age and C-14 content of organic carbon according to the equation:

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1 5730 A Adjusted 4 C age = --ln RDoc

ln2 AMDoc

(7-3)

However, if the age of the final water exceeds the corresponding average age of modem dissolved organic carbon in the recharge zone (AMDoc), i.e. ARDoc > 100 pmC, 100 pmC is used for organic carbon during recharge.

If the initial water is not recharge water sampled deep in the bedrock, the C-14 content of organic carbon at the start of the flow step (14Corg) decreases with the age of the initial water, but the overall age of the flowpath affects the original C-14 content of organic carbon in the recharge water. In this case the above Equation 7-3 is first applied using the adjusted age of the whole flowpath from the recharge. Next the C-14 content of organic carbon is obtained by reducing the radioactive decay corresponding to the C-14 age of the initial water from the C-14 content of organic carbon during the original recharge (ARDoc) of the whole flow/reaction path:

Adjusted 14C age of initial water= 5730

ln ~RDoc ln 2 corg.

(7-4)

The iterative mass-balance calculations are continued as far as ARDoc corresponds to the total age of the whole flowpath. If the total age of the final water exceeds the corresponding average age of modem dissolved organic carbon in the recharge zone (AMDoc), 100 pmC is used for ARDoc.

For the C-14 content of dissolved carbonate of initial waters (A0), either the measured value of the initial sample is used, or a value of 100 pmC in flow steps beginning from recharge and assumed to be older than 50 years (e.g. according to tritium contents or other features indicating long residence time).

7 .1.5 Results of mass-balance modelling

The mass-transfer of selected phases is shown in Table 7-2 and carbon isotope results are listed in Table 7-3. Carbon isotopes are in most cases treated as an isotope mass­balance problem, thus calculated 8C-13 values do not test the validity of the reaction hypothesis although they can verify the model and fix the input of each carbon source along the flow steps. These cases all have recharge water as the initial water and the calculated final 8C-13 equals the final measured value (Table 7-3). The three uppermost models also start from the recharge water, but are used for calibrating 8C-13 of dissolving calcite and C-14 of organic carbon during recharge (cf. section 7.1.3). If calcite is precipitating along the flow steps, carbon isotopes are modelled by solving the Rayleigh distillation equations (Wigley et al. 1978, 1979) for the computed mass­transfer (Plummer et al. 1994 ). This approach is applied to flow steps which initiate deep in the bedrock, and calcite has already been saturated along the previous flow steps of the whole flowpath. Calculated mass-transfer models predict reasonably well 8C-13 values of dissolved carbonate in final waters.

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Table 7-2. Modelled mass-transfer results of flow steps. All mineral and gas mass­transfers are in mmol/kg of water. A negative value indicates precipitation and a positive value dissolution/ingassing. Rech = Vuorimaki spring (App. 6)

Samples Calcite Organic Pyrite Goethite COz gas Oz gas Additional initiaUfinal matter sulphate water NazS04

Rech/KR3T5 0.19 0.74 0.0 -0.15 - 0.79 Rech/KR5T2 0.11 0.86 0.0 -0.34 - 0.94 -Rech/KR5T1 0.02 0.85 -0.01 -0.27 - 0.87 -

Rech!Liimat. 0.03 1.08 -0.0 -0.81 - 1.27 -

Rech/KI-KA2 0.22 0.71 -0.0 -0.13 0.72 -

Rech/KR1T7 0.51 1.01 -0.02 -0.55 - 1.08 -Rech/KR1T6 0.47 1.21 -0.0 -0.47 - 1.31 -KR1T6!T5 -0.82 0.01 -0.0 0.0 - 0.04 Rech/KR1T4 0.39 0.55 -0.0 -0.17 - 0.55 KR1T6!T3 -0.53 0.02 -0.01 0.0 - - 0.02 Rech/KR4T6 0.27 0.60 -0.01 -0.30 0.64 -

Rech/KR4T5 0.20 0.73 -0.01 -0.17 - 0.74 Rech/KR4T4 0.40 0.70 -0.00 -0.20 - 0.72 Rech/KR4T3 0.49 0.57 -0.00 -0.31 - 0.62 -KR4T6!T2 -0.15 0.14 -0.03 0.03 - - 0.08 KR4T6!T1 -0.09 0.16 -0.04 0.03 - - 0.09 Rech/KR5T6 0.20 0.66 0.0 -0.42 - 0.76 -Rech/KR5T4 0.09 0.26 -0.07 0.0 0.42 0.18

Rech/KR2T5 0.34 0.12 -0.03 0.0 0.29 0.07 Rech/KR2T4 0.24 0.34 -0.09 0.0 0.02 0.18 Rech/KR2Tl 0.21 0.25 -0.07 0.0 0.22 0.15

Rech/KR3T6 0.25 0.58 0.0 -0.44 0.70 Rech/KR3T3 0.26 0.31 -0.08 0.0 0.34 0.18 KR3T6!T2 -0.03 0.04 -0.01 0.01 0.02

Samples Plagioclase Biotite Kaolinite SiOz K-montmo- Ca!Mg ... Naz Additional initial/final Anzs rillonite Exchange chloride water NazCaCl4 Rech/KR3T5 0.21 0.06 0.0 -0.39 -0.14 0.0/0.06 ... -0.12 Rech/KR5T2 0.70 0.16 0.0 -0.74 -0.44 0.05/0.0 ... -0.1 0 0.0 Rech/KR5Tl 0.58 0.13 0.0 -0.59 -0.37 0.010.011-0.02 0.0 Rech/Liimat. 0.41 0.36 0.69 0.20 -0.97 0.08/0.0 ... -0.16 0.0 Rech/KI-KA2 0.27 0.07 -0.01 0.0 -0.16 -0.02/0.04 ... -0.04 0.0 Rech/KR1T7 0.16 0.25 0.52 0.03 -0.64 -0.08/0.0 ... 0.16 0.0 Rech/KR1T6 0.43 0.21 0.31 -0.28 -0.59 -0.07/0.0 ... 0.14 0.01 KR1T6!T5 0.97 0.0 -0.54 -1.38 -0.04 0.09/-0.12 ... 0.06 0.13 Rech/KR1T4 0.18 0.08 0.12 0.0 -0.24 -0.20/-0.06 ... 0.52 0.0 KR1T6!T3 0.79 0.0 -0.41 -1.06 -0.07 0.06/-0.11...0.10 0.04 Rech/KR4T6 0.20 0.14 0.26 0.0 -0.40 -0.07 /0.04 ... 0.06 0.0 Rech/KR4T5 0.33 0.08 0.0 -0.35 -0.21 -0.06/0.08 ... -0.04 0.0 Rech/KR4T4 0.11 0.09 0.18 0.0 -0.25 -0.15/0.09 ... 0.12 0.0 Rech/KR4T3 0.10 0.14 0.33 0.0 -0.40 -0.12/0.03 ... 0.18 0.0 KR4T6!T2 0.28 0.0 -0.16 -0.55 -0.01 0.0/-0.09 ... 0.18 0.0 KR4T6!T1 0.10 0.0 -0.06 -0.23 0.0 0.0/-0.07 ... 0.14 0.0 Rech/KR5T6 0.23 0.19 0.35 -0.05 -0.50 -0.02/0.0 ... 0.04 0.01 Rech/KR5T4 0.31 0.03 -0.19 -0.51 -0.01 0.010.0 1...-0.02 0.08

Rech/KR2T5 0.09 0.01 0.0 -0.09 -0.05 -0.17 /0.01...0.32 0.04 Rech/KR2T4 0.27 0.04 0.0 -0.27 -0.16 -0.07/-0.03 ... 0.20 0.02 Rech/KR2T1 0.10 0.03 0.0 -0.13 -0.07 0.13-/0.03 ... -0.20 0.06

Rech/KR3T6 0.25 0.19 0.41 0.0 -0.57 0.07/-0.03 ... -0.08 0.0 Rech/KR3T3 0.23 0.04 0.0 -0.35 -0.14 -0.03/0.01 ... 0.04 0.03 KR3T6!T2 0.11 0.0 -0.07 -0.20 0.0 0.01-0.03 ... 0.06 0.01

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13 J;pble 7-3. Model param~~ers and carbon isotope results ~ Ccalcite = -5. 7%o P DB,

{;calcite = 0% modern, 8 Corganic carbon, C02 = -25%o p DB, Cco2 = 100% modern and 8 Crechargewater = -23.2%o PDB). Samples initiaVfinal 8 13C 14c 14c 14c Adjusted age of 8180 3H in final water water calc./measured org./initial ea le., measured step/ whole flow final water (TU)

(%o PDB) water, Ao (pmC) And (pmC) (pm C) path (years) (%o SMOW) Rech/KR3T5 -21.80 I -21.81 50/110.2 63.6 63.1 71 I moderrn -12.8 24.9 Rech!KR5T2 -23.11 I -23.19 32 I 110.2 55.2 55.2 -1 -12.7 11.8 Rech!KR5Tl -24.23 I -24.22 38 I 110.2 62.9 62.8 14/ modem -12.6 8.1 Rech/Liimat. -23.36 I -23.36 100? I 100 97.8 74.5 2250? -12.5 <0.8 Rech/KI-KA2 -21.30 I -21.30 37 I 100 52.2 49.4 450 -12.6 < 5.5 Rech!KR1T7 -19.60/-19.60 63 I 100 55.6 44.4 1850 -12.5 < 6.1 Rech/KR1T6 -20.40 I -20.40* 100 I 100 78.2 29.8* 8030 -12.7 <0.8 KR 1 T6/KR 1 T5 -20.60 I -20.55 37 I 29.8* 29.8 24.6 1530 I 9560 -13.7 < 0.8 Rech!KR1T4 -19.07 I -19.07 100 I 100 75.8 37.9 5730 -12.6 <0.8 KR 1 T6/KR 1 T3 -20.54 I -20.61 37 I 29.8* 29.8 25.3 1320 I 9350 -13.2 <0.8 Rech!KR4T6 -20.51 I -20.51 55 I 100 59.9 38.6 3630 -12.8 <0.8 Rech!KR4T5 -21.63 I -21.63 65 I 100 67.5 36.0 5190 -12.8 < 0.8 Rech!KR4T4 -19.52 I -19.52 58 I 100 55.8 33.8 4140 -12.9 1.4 Rech!KR4T3 -18.29 I -18.29 58 I 100 52.5 31.7 4170 -12.8 <0.8 KR4T6/KR4T2 -20.98 I -21.56 64 I 38.6 41.1 21.7 5300 I 8930 -13.0 1.2 KR4T6/KR4Tl -21.01 I -21.67 54 I 38.6 40.8 24.1 4360 I 7980 -12.7 2.8 Rech!KR5T6 -21.46 I -21.46 43 I 100 56.8 46.8 1600 -12.4 2.4 Rech!KR5T4 -22.85 I -22.85 0 I 100 71.4 19.8 10600 -14.5 < 8.1

Rech!KR2T5 -18.95 I -18.95 0 I 100 62.4 18.8 9915 -13.6 0.8 Rech!KR2T4 -19.88 I -19.88 0/100 45.8 8.7 13730 -13.3 < 0.8 Rech/KR2Tl -20.74 I -20.74 01100 60.2 10.2 14680 -13.9 4

Rech!KR3T6 -20.58 I -20.58 91 I 100 76.6 42.0 4960 -12.9 8.5 Rech/KR3T3 -20.75 I -20.75 01100 58.8 14.0 11870 -13.6 4.7 KR3T6/KR3T2 -20.75 I -20.80 55 I 42 42.2 26.4 3870 I 8830 -13 9.6

*Average value from two analyses

7. 1. 5. 1 Mass-transfer

Total mass-transfer is generally at a low level and comparable to the results of the Romuvaara site (Pitkanen et al. 1996a). Minor reactions are expected for water­crystalline bedrock interaction at low temperatures in a cool climate environment, and salinity stays at a low level when reactions dominate the mass-transfer and there is no mixing with saline palaeofluid. The results suggest that the dissolution of primary silicates is more significant whereas the dissolution of calcite after recharge is smaller than indicated by the Romuvaara modelling. The result is consistent with the small and quite deep occurrence of calcite infills in fractures. Silicate weathering is calculated as plagioclase and biotite dissolution, and fracture mineral observation of iron released from biotite and precipitated as goethite (or limonite) supports the models. Strontium isotope signatures also support biotite activity. The poor fit of Sr in the primary biotite lattice encourages its rapid loss early in the weathering history during the oxidation process (Bullen et al. 1997). The highest Sr isotope ratios belong to final waters, which need relatively more biotite dissolution (Fig. 7-1) compared with other Sr sources (plagioclase, calcite) along their flowpath (KR1/T6, KR4/T1, Liimatainen: Sr-87/Sr-86 > 0.75) than those with a low Sr signature (KR1/T3, KR5/T4, KR2/T4, KR3/T3: Sr-87/Sr-86 < 0.74).

1.0

0.0

0.70

li X 0

0.72

92

•

li 0

<>

0.74 0.76 0.78 0.80 Sr-87 /Sr-86

I• Spring • Well 11 KRl o KR4 c KR5 x KR2 ~ KR31

Figure 7-1. Calculated dissolution of biotite in relation to the sum of dissolved calcite and plagioclase along the flow path as a function of 87 Sri6Sr of final water. The ratio contains total mass-transfer of each soluble phase from recharge to final water.

The dissolution of pyrite or more generally iron sulphides is not needed, hence goethite formation is derived from silicate weathering and confirms the importance of biotite in oxygen consumption during recharge. Probably a more important user of oxygen is the respiration of organic carbon in overburden. The consumption of oxygen exceeds oxygen solubility in cold water (about 0.3 mmol/1), thus part of the calculated organic carbon oxidation (plant root respiration and decay of plant debris) has to occur in the unsaturated zone above the groundwater table. This means that the concentration of carbon acid in the default recharge water (Vuorimaki spring) is not sufficient

Pyrite precipitation is a minor process, but an important end-product in reaction steps containing anaerobic respiration reactions. Some of these steps need quite a lot of additional sulphate to dissolve in the system (about 0.2 mmol/1), clearly exceeding the highest measured values at the site (0.07 mmol/1). It is impossible to evaluate whether this is realistic. Sulphur isotope calculations would give more information, as use of the sulphur isotope data would show the total amount of sulphate reduction and pyrite precipitation required for an S-34 increase in the system (Plummer et al. 1990), but there is a lack of data. Sulphur isotope calculation was only possible to apply in one flowstep, KR1/T6 --7 KR1/T3. The sulphur isotopic data is treated as a mass-balance problem, even though there is a fractionating output of sulphur in precipitated pyrite in that particular flow step. The purpose is to compare the isotopic composition of the final water calculated by the Rayleigh model with the observed one to get information on accurate sulphate reduction.

93

Table 7-4. Calculated adjusted sulphur isotopic results of the reaction model for KRJIT6 to KRJIT3. Average sulphur isotopic composition in solutions is calculated from measured 834S(S04) and estimated 834S(H2S) using the Equation 7-2.

Samples Average 834S in Modelled I observed Calculated average initial/final water initial water I additional average 834S in final 834S in precipitating

S04 (%o CDT) water (%o CDT) pyrite (%o CDT) KR 1 (T6)/KR 1 (T3) 20.74 I 5.00 25.32 I 26.66 -25.48

In this case the sulphur isotope balance does not require unusual sulphate addition to the ground water along the flow step (Table 7-4 ). Twice as much additional sulphate is needed for pyrite precipitation as in a pure element mass-balance problem, but still the requirement remains very low (0.025 mmol/1, Table 7-2). The need depends on the 8S-34 value of additional sulphate such that a higher value decreases input and the compatibility of calculated and measured 8S-34 improves, but correspondingly 8C-13 values get poorer and vice versa. Modelling results are based on 8S-34(S04) = 5%o CDT corresponding to modem recharge.

The consumption of carbon dioxide (Table 7-2) may describe the subglacial aerobic respiration source, whereas the consumption of organic carbon in those models describes the anaerobic respiration source of dissolved carbonate. Aerobic respiration seems to be smaller during glacial (or de glacial) recharge than during interglacial according to the calculated models.

Potassium montmorillonite acts as a sink for potassium released from the dissolution of biotite. Presumably other cations also take part in the formation of secondary montmorillonites, as the thermodynamics favour samples deep in the bedrock. The irreversible nature of potassium in ion-exchange processes (Appelo & Postma 1993) compared with Na, Ca and Mg also affected the decision of what phases to use. Formation of secondary montmorillonites would not in fact be perfect and some semi­crystalline variety is more probable due to kinetic constraints (Paces 1973). The uptake of potassium may be partly an ion exchange process, partly crystallisation. The proportion of potassium in montmorillonite (beidellite end-member in mass-balance models) is small compared with aluminium, and silica is very small and significantly smaller than in dissolving biotite. Therefore kaolinite and silica dissolve in some models (e.g. Rech. ~ KR1/T7), because aluminium and silica are needed in the precipitation of perfect montmorillonite. However, silica precipitates in most models as fracture infillings suggest (Gehor et al. 1995).

The results of ion exchange are complex. Calcium is mostly adsorbed after recharge and sodium released, as is assumed in systems in which calcite is dissolving in water from fracture surfaces. Exceptions are fast flow paths dominated by silicate weathering. This may be one reason for conserving a high strontium isotope signature after recharge, because strontium behaves similarly to calcium in the geochemical system.

94

7.1.5.2 Carbon-14 age of groundwater and pa/aeohydrogeo/ogical implications

Adjusted C-14 ages for groundwater samples have been calculated on the basis of the reaction models (Tables 7-1. .. 7-3). Most of the ages seem reliable. The isotopic dilution effect of fast reactions on C-14 without radioactive decay seems to be approximately 50% (Fig. 7-2a), i.e. one half-life should be subtracted if uncorrected C-14 (DIC) ages are examined. The dilution effect is comparable to estimates from Romuvaara (Pitkanen et al. 1996a), Hastholmen (Kankainen 1986) and Olkiluoto (Pitkanen et al. 1996b). The modelled ages are slightly higher than those qualitatively estimated according to the isotopic results in section 5.5.5. The shift is caused by the dynamic model of organic­den ved C-14 during recharge that is used in modelling.

Predicted ages are analogous with oxygen-18 data of groundwater samples (Fig. 7-2b). The 80-18 depletes steeply as the calculated age approaches the end of the glaciation period 10,000 years ago. Higher ages are interpreted to be mixtures of glacial melt water and older possibly preglacial water. Tritium contents are also consistent with adjusted ages in most cases (Table 7 -3). Young ages, indicated by a considerable tritium content, are clearly violated (high calculated age) in the four last models. The tritium concentration in the final waters of these models also conflicts with other variables (e.g. 80-18, 8H-2, Cl), indicating a long mean residence time of the final waters. The discrepancy is considered to result from shallow groundwater contamination during field activities. The calculated C-14 ages should represent a lower age than the original groundwater at sampling depth, as shown in Table 5-4.

a) •

70

p

. 0 t.

10

0

0

<> 0 ot.

5 ( <> tr.

0 X'-'

<:

4000 8000 12000 Adjusted C-14 age, years BP

X

16000

b)

~ -13.0 0 ~ 00

-;_

-15.0

0

0 .oil

• t.

0~ <>0 G

-t.

X <: t.

0

4000 8000 12000 Adjusted C-14 age, years BP

I• Spring • Well 1:1. KRl o KR4 c KR5 x KR2 ~ KR31

X

X

16000

Figure 7-2. Measured C-14(DIC) contents and oxygen isotopic composition of groundwater samples vs. reaction adjusted C-14 age.

95

The main uncertainty in the C-14 dating analysis is included in the reliability of the estimated C-14 content of recharging organic carbon ARDoc and the accuracy of the derived mass-transfer in adjusting the C-14 content of dissolved carbonate for reaction effects along the flowpath to final water defining And· The estimation of Ao does not play an important role in the models, because C-14 depleted organic matter (peat-derived, 35-100 pmC) determines the radiocarbon content in dissolved carbonate during a short recharge period in the developed model. If Ao was 120 pmC for young groundwaters, which would be possible since the start of the nuclear era (e.g. Pearson et al. 1991) the value would not affect the calibration of radiocarbon content in organic carbon in cases where KR5/T2 and KR5/T1 are final waters, but in the case of KR3/T5 the radiocarbon content in organic carbon would be 40 pm C. This value only affects KR1/T7, which would be young groundwater. The result is contradictory to other characteristics of the sample, indicating a lower value for Ao than is used in the test, and supports the radiocarbon level measured from groundwaters in overburden (110 pmC).

The uncertainty included in modelling the radiocarbon input from organic matter during recharge is difficult to estimate. The used modem range (35-50 pmC) seems reliable, if bogginess in depressions started quite soon (within 1000 years) after the ice retreat, which is probable. If the used range is too low, flowpaths should also contain dead organic carbon (which is easily possible; especially preglacial organic carbon can exist in a till layer) in order to conserve the low ages of young waters, i.e. the mean C-14 content of reacting organic carbon should remain at the estimated level. In addition, some of the intermediate ages of final waters would increase too much compared with information from other chemical variables. For example, the age of KR4/T2 and T1 would reach the glaciation period, although the stable isotopes do not show a cold climate signature. Especially the calculated ages of KR1 support the method used to estimate the input of organic carbon. The ages of samples with depleted 80-18 correspond to the time interval from glaciation at the site, and the age of KR1/T4 is comparable to the age predictions of KR4/T4 and T3 which are chemically similar and possibly from the same sub-aquifer.

Probably the primary input value of organic carbon (35 and 50 pmC) in adjusting ARDOC

varies more than is used in calculations, as is suggested by the flow step into the Liimatainen well. If ARDoc = 50 pmC is used, the age of the water will be strongly negative. The well does not situate at the site, but the result emphasises local variation in the character of organic input.

The ages of borehole KR4 seem relatively old compared with their chemical character, which is more suggestive of an initial stage of water-rock interaction than matured middle age as suggested by pH and C-14. The mass-transfer results and hydrochemistry of the samples from KR4 suggest that water-rock interaction is small: Calcite saturation is barely reached, silicate hydrolysis plays an important role in bicarbonate production also according to the quite high pH, and without any extra input of calcium into the system calcite precipitation stays a minor process. The slight interaction is consistent with the few fracture infillings observed. The small concentrations prevent significant mass-transfer in mass-balance calculations and limit reactions adjusting the C-14 content in the final water. Therefore the assumed carbon isotope values for carbon input in the models are most critical for the results. Although organic carbon is assumed to be

96

totally dead, which seems impossible, water samples would be several hundred years old and the two deepest ones several thousand. If the 8C-13 value of calcite is lower than assumed, calcite dissolution will be higher but, for instance, the adjusted age of KR4/T5 decreases by only 300 years when the 8C-13 of calcite is -10%o PDB. Carbon isotope exchange with calcite would be one possible process for decreasing the C-14 content in dissolved carbonate without radioactive decay, but simultaneously it strongly increases 8C-13 in the groundwater, which cannot be observed. One potential factor that could decrease the ages of groundwaters from KR4 is that the chemical data lack part of the carbonate evolution, i.e. the extreme peaks in bicarbonate and calcium contents when calcite reaches equilibrium in Fig. 6-2. In that case calcite would dissolve more than is assumed in the models, and it would also precipitate before reaching sampling points. However, the precipitation and dissolution omitted from the models should be slight. The missing dissolution stage cannot significantly decrease the ages because calcite precipitation without silicate dissolution (as salinity does not increase) will acidify the solution and that is not observed.

As a summary of the adjusted ages of water samples from KR4, the modelled mean ages are considered plausible. The results contain some degree of uncertainty, but it is impossible to determine exact limit values for ages, and 35 pmC for organic derived C-14 during modem recharge seems the minimum reliable value. Therefore the mean C-14 ages cannot decrease significantly, at least by no more than 1000 years.

Four final waters (KR5/T4, KR2/T4, KR2/T1 and KR3/T3) clearly precede the time when the ice sheet retreated from the area 9,700 years ago (Eronen & Lehtinen 1996). The subglacial carbon-14 ages are considered unrealistic (Fontes 1994). Test calculations with dead carbon sources, including C02, still show ages over 11,000 years for KR2/T4 and KR2/T2. These observations indicate mixing of old dead carbonate­containing groundwater with younger radiocarbon-containing input. Young contamination could explain part of the mixing, but the results in Table 5-4 also suggest a water component of intermediate age from a dissolved carbonate source in the system. The old water component with no radiocarbon may be preglacial groundwater representing 80-18 values at least as high as modem recharge. In this case the intermediate age water component would be melt water from the retreating ice sheet.

If the deep system contains only these two end-members, it is possible to estimate the mixing ratio with a simple mixing calculation. From Table 5-4 one can approximate the following mean composition: 80-18 = -14%o SMOW and C-14 = 5 pmC, representing the old deeper aquifer at Kivetty (below 300m). If we assume for the old component that 80-18 = -12.5%o SMOW and for melt water that 80-18 = -20%o SMOW (Kankainen 1986), the deeper aquifer contains a melt water component of about 20%. If 80-18 of the melt water had been lower, e.g. -22%o SMOW as estimated by Tullborg and Larson (1984), the mixing proportion of melt water would be 15%, whereas if 80-18 of the old component were 1 %o higher than modem recharge, i.e. -11.5%o SMOW, the melt water proportion would be nearly 30%. An increase of 1%o in 80-18 corresponds, according to temperature coefficients for precipitation (Kankainen 1986), to a climate roughly 2°C warmer than today, which would correspond to the Eemian interglacial period (Eronen & Lehtinen 1996). If both extreme values of 80-18 are provided the mixing proportion of melt water would again be about 20%.

97

7.2 Modelling of reaction paths with EQ6

7.2.1 Modelling

The EQ6 reaction-path code in the EQ3/6 software package (Wolery 1992) was used to test the thermodynamic feasibility of reaction models derived from mass-balance calculations. The objective was especially to test the thermodynamic confidence of the considered reactions and the extent of these processes. The amounts of reaction products give the order of magnitude of the process, and should not be seen as strictly reflecting actual quantities.

The thermodynamic database used for the EQ6 calculations was the same ( dataO.comR2) as for the EQ3 calculations. Of the proposed mass-transfer models in the Kivetty area (Table 7-2) some steps were modelled with EQ6, Table 7-4.

Table 7-4. Reaction steps modelled with EQ6 for the Kivetty site.

Reaction step KRI Rech. ==> KR1/T7

Rech. ==> KR1/T6 KR1/T6 ==> KR1/T5

KR3 Rech. ==> KR3/T6

KR3/T6 ==> KR3/T2

KR4 Rech. ==> KR4/T6 KR4/T6 ==> KR4/T2

KR4/T6 ==> KR4/Tl

In the modelling, the reactions were presumed to occur in a closed system due to the deep confined groundwater system. The used code version does not have an option for ion exchange and hence the ion exchange processes considered were simulated by simply reducing or increasing the amounts of ions in the system. In order to include the aluminium-containing minerals in the modelled systems, a small amount of AI (10-20mmol/l) had to be added to the recharge water, as the analysed result was below the detection limit.

Measured values of Eh are available only for groundwater samples, not for well or spring samples. However, due to difficulties in measuring the Eh (cf. section 5.4.4) the measured values were not representative enough to be used in modelling. Based on the redox interpretations in sections 5.4.5 and 6.1.3 it was considered appropriate to determine the redox state with EQ3 calculations, based on the analysed s2-/SO/­contents of the groundwater samples. Reaction paths which had recharge water (Vuorimaki spring) as initial water in the first reaction step caused some difficulties in model calculations because of 0 2 . Generally, in nature oxygen distributions are controlled by closed system conditions, therefore below the water table gas exchange with the atmosphere ceases, and oxygen is gradually consumed along the flowpath by reactions with reduced substances in the aquifer (Appelo & Postma 1993), in this case

98

organic matter. In some reaction steps where the initial run included oxygen consumption by organic matter producing C02-bearing groundwater, the starting point for the EQ6 modelling was that all organic matter had been consumed by reduction, leaving an equivalent amount of C02 to dissolve in the water. The estimated Eh value at this point was chosen to be -100 m V according to sequences of redox reactions (Appelo & Postma 1993). The temperature of the groundwaters was kept constantly at 7°C during modelling.

The composition of Kivetty plagioclase is An25 (albite high 75%, anorthite 25% ). In modelling the end-members albite and anorthite were selected from the mineral database to represent plagioclase, and amounts of dissolving end-members were given according to the composition. The Kivetty biotite is Fe-rich. The end-members phlogopite and annite were used in the modelling with proportions of 25% and 75%, respectively.

In EQ6 modelling the dissolution and precipitation reactions were set to obey arbitrary kinetics, the rate usually being the same for all reacting minerals. An exception to this was the ion exchange simulation, where ions with + 1 charge were set to "react" at double the rates of ions having a charge of +2. This practice was adopted to maintain the electrical balance during simulation in the modelled groundwater.

When comparing the results of the mass-balance modelling (NETP A TH) and the thermodynamic modelling (EQ6), slight differences are expected in the amounts of precipitated phases because the models use somewhat differing mineral formulae forK­montmorillonite:

EQ3/6 database: NETPATH database:

Ko.33Mgo.33All.67Si40Io(OH)2 Ko.33Ah.33Sb.670Io(0Hh

K-montmorillonite used in the EQ6 modelling contains magnesium, which is not included in the formula for K-montmorillonite in the NETPATH mass-balance modelling. The proportions of aluminium and silica differ as well.

In EQ6 modelling each reaction step (Table 7-2) was divided into a set of EQ6 runs. The consecutive EQ6 runs tried to simulate the true order of processes active in nature. Processes occurring during the initial stage were considered very rapid and for this reason usually no mineral precipitation was allowed during the first run.

7 .2.2 Results

The results of EQ6 reaction path modelling are given in Tables 7-5 ... 7-12, including the processes considered in the final mass-balance modelling, as well as the results for the initial and final waters obtained by EQ3.

The reaction step to KR1/T7 included infiltration of recharge water (Vuorimaki spring) into the bedrock (Table 7 -5). As pointed out above, the consumption of organic matter was assumed to have occurred before the first EQ6 run started. Altogether 11 EQ6 runs were needed to complete this reaction path. The resulting final water reached a pH value of 7.6 and an Eh value of -210 m V. These values agree quite well with the measured pH

99

value and the calculated Eh value for the chosen redox pair, sulphide/sulphate. Somewhat less goethite (0.48 mmol instead of 0.55 mmol), K-montmorillonite (0.61 mmol instead of 0.64 mmol) and pyrite (0.014 mmol instead of 0.02 mmol) had precipitated than predicted from mass-balance calculations. Precipitation of more pyrite was prevented because the final water had not reached the sulphide domain; the sulphur speciation was entirely sulphate. No more K-montmorillonite was precipitated because it had reached equilibrium in the final water as well as goethite. Kaolinite had reached a state of supersaturation.

Also the second reaction step to KR1/T6 included infiltration of recharge water. In the first run, dissolution of 0 2(g), organic matter and calcite was completed (Table 7 -6). Continuing with the other presumed reactions the obtained final water reached a pH value of 7 .5, quite compatible with the measured value. Also the obtained Eh value of-190 m V is in good agreement with the calculated value for the sulphate/sulphide pair. Precipitating phases are in agreement with expectations.

The reaction step in borehole KR1 from T6 to T5 needed three EQ6 runs to completion (Table 7 -7). In this step the final result does not correspond very well with the values of KR1/T5, especially in terms of the pH and Eh values obtained. The EQ6 calculation produced quite a high pH value of 9.2 compared with the measured value of 8.2, and the Eh value obtained was clearly lower ( -320 m V) than the calculated one ( -245 m V) for the sulphide/sulphate pair in KR1/T5. Precipitating phases correspond to expectations.

In borehole KR3 the initial process is again infiltration of recharge water (Vuorimaki spring) into the bedrock (KR3/T6, Table 7 -8). In the final EQ6 run, after all other reactions had reached completion the remaining undissolved kaolinite and small amount of pyrite (0.002 mmol) were allowed to dissolve. The final pH value obtained was 8.6, which agrees quite well with the measured value of 8.4, but the Eh remained high (720 m V) due to high oxygen pressure in the system. The precipitated goethite amount coincided well with the proposed amount, but it was not enough to consume all the oxygen. The system proved extremely sensitive to the dissolution of pyrite; doubling the pyrite amount to 0.004 mmol produced an Eh of -280 m V.

Reaction step in borehole KR3 from T6 to T2 was completed with eight EQ6 runs (Table 7-9). The final pH reached a value of 8.4, which is somewhat higher than the measured one (8.2). The obtained Eh value of -250 m V was clearly lower than the measured one of +20 m V and the one calculated from the sulphide/sulphate pair (-160 m V). The precipitated amounts of minerals agreed well with expectations (NETPATH).

The first reaction step in borehole KR4 includes infiltration of the Vuorimaki spring water (recharge) into the bedrock to KR4/T4 (Table 7-10). The reaction step was completed with four EQ6 runs. After completion of the proposed model the pH value was 7.4 and Eh -185 m V. The obtained pH value was clearly lower than the measured value of 8.1, but the Eh value agreed quite well with the measured value.

In the second reaction step in borehole KR4 from T6 to T2 (Table 7-11) the final pH of 8.8 is in good agreement with the measured value of 8.6, but the obtained Eh value of -340 m V is clearly lower than the calculated one (for the sulphide/sulphate pair) in the

100

final water of KR4/T2. Predicted thermodynamic precipitation is in good agreement with the results of inverse modelling.

The reaction step from KR4/T6 to KR4/T1 was modelled in three EQ6 runs (Table 7-12). The final result is in quite good agreement with the proposed mass-balance model. The final pH value obtained was 8.7, compared with a measured value of 8.6 and the obtained Eh value was -330 m V while that calculated for the sulphide/sulphate pair was -275 m V in KR4/Tl.

7.2.3 Discussion

In most cases the reaction path modelling agrees fairly well with the results of mass­balance modelling, supporting the proposed interpretation of the evolution of ground water at the Ki vetty site. In some modelling cases EQ6 modelling may not be completed because the system does not contain enough material to precipitate the desired amount of phases. This may partly be due to the expected slight deviation in the mineral formula of K-montmorillonite between mass-balance modelling (NETPATH) and thermodynamic modelling (EQ6), and partly to the analytical uncertainties. However, generally when comparing EQ6 modelling results with mass-balance calculations they agree well, especially when the evolution of the deep groundwater is considered, and support the pH level measured in groundwaters.

In EQ6 modelling, the redox states obtained for the final waters were clearly more reducing compared with the measured values. However, the evolution of groundwater proposed on the basis of mass-balance models with thermodynamic reliability gives strong evidence of reducing conditions deep in the bedrock and verifies well-known problems with field measurements of the redox state of deep ground water.

101

Table 7-5. Results of EQ3/6 modelling of flowpath KRJ reaction step 1.

FLOWPATH KRl I reaction step: Rech. ::::::} KR1/T7

Final mass balance calculation results from NETPATH EQ6 I phase mmol/kg mmol/1

Precipitating: pyrite 0.02 0.014 goethite 0.55 0.48 K-montmorillonite 0.64 0.61

Dissolving: 02 (g) 0.07 0.07 C02 (g) 1.01 1.01 calcite 0.51 0.51

Plagioclase* (25%) anorthite 0.16* 0.04 (75%) albite high 0.12

Biotite** (25%) phlogopite 0.25** 0.0625 (75%) annite 0.1875

chalcedony 0.03 0.03 kaolinite 0.52 0.52

Ion exchange: Ca2+ 0.08 out 0.08 out Na+ 0.16 in 0.16 in

EQ3 calculation EQ6 result for results for Final Initial Final Vuorimaki KR1/T7

pH 6.3 7.9 7.6 Eh HS-/S04

2- m V -100*** -260 -210 Alkalinity meq 0.18 2.0 1.8

Ctot mmol/kg 0.46 2.1 2.0 Si02 mmol/1 0.20 0.13 0.008 Fetot mmol/kg 1e-4 0.001 0.07 Allot

11 1e-20 4e-4 0.46 Na+ mmol/1 0.11 0.41 0.39 K+ 11 0.02 0.06 0.07 Ca2+ 11 0.08 0.55 0.55 Mg2+ 11 0.03 0.22 0.02 Mn2+ 11 2e-4 0.003 2e-4

Stot mmol/kg 0.03 0.02 0.003 SO/ mmol/1 0.03 0.001 0.003 HS- 11 -er 11 0.03 0.04 0.03 F 11 - 0.12 -log f(C02) -2.3 -3.0 -2.6 log f(02) -70.7 -75.4 -72.1

* **

NETP ATH mass-balance calculated for indicated plagioclase solid solution. NETP ATH mass-balance calculated for indicated biotite solid solution.

*** set value.

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Table 7-6. Results of EQ3/6 modelling of flow path KRJ reaction step 2.

FLOWPATH KRl I reaction step: Rech. ==> KR1/T6

Final mass balance calculation results from NETPATH EQ6 mhase mmol/kg mmol/1

Precipitating: pyrite 0.003 0.004 goethite 0.47 0.42 chalcedony 0.28 0.36 K-montmorillonite 0.59 0.55

Dissolving: 02(g) 1.31 1.31 or g. 1.21 1.21 calcite 0.47 0.47

plagioclase* (25%) anorthite 0.43* 0.11 (75%) albite high 0.32

biotite** (25%) phlogopite 0.21 ** 0.05 (75%) annite 0.16

kaolinite 0.31 0.16 Na2CaC14 0.01 0.01

Ion exchange: Na 0.14 in 0.14 in Ca 0.07 out 0.07 out

EQ3 calculation EQ6 result for results for Final Initial Final Vuorimaki KR1/T6

pH 6.3 7.9 7.5 Eh measured m V -100*** +80

Hs-;so42- " -205 -190 Alkalinity meq 0.18 1.90 2.02 Ctot mmol/kg 0.46 1.99 2.15 Si02 mmol/1 0.20 0.20 0.07 Fetot mmol/kg 1e-4 0.002 0.05 Altot mmol/kg 1e-20 1e-3 0.45 Na+ mmol/1 0.11 0.61 0.59 K+ " 0.02 0.04 0.05 Ca2+ " 0.08 0.58 0.59 Mg2+ " 0.03 0.18 4e-13 Mn2+ " 2e-4 - 2e-4 Stot mmol/kg 0.03 0.03 0.02 sol- mmol/1 0.03 0.03 -er " 0.03 0.10 0.07 log f(C02) -2.3 -2.9 -2.6 log f(02) -70.7 -71.8 -72.3

* **

NETP ATH mass-balance calculated for indicated plagioclase solid solution. NETP ATH mass-balance calculated for indicated biotite solid solution.

*** set value.

103

Table 7-7. Results of EQ3/6 modelling ofjlowpath KRJ reaction step 3.

FLOWP ATH KRl I reaction step: KR1/T6 ::::::} KR1/T5

Final mass balance calculation results from NETPATH EQ6 I phase mmol/kg mmol/1

Precipitating: calcite 0.82 0.94 pyrite 0.002 0.002 kaolinite 0.54 0.51 chalcedony 1.38 1.29 K-montmorillonite 0.04 0.11

Dissolving: or g. 0.01 0.01

plagioclase* (25%) anorthite 0.96* 0.24 (75%) albite high 0.72

Na2S04 0.04 0.04 Na2CaCl4 0.13 0.13

Ion exchange: Na 0.06 in 0.06 in Ca 0.09 in 0.09 in Mg 0.12 out 0.12 out

EQ3 calculation EQ6 result for results for Final Initial Final KR1/T6 KR1/T5

pH 7.9 8.2 9.2 Eh measured m V +80 -30

HS-/SO/ " -205 -245 -320 Alkalinity meq 1.90 1.37 1.01 Ctot m mol/kg 1.99 1.40 1.07 Si02 mmol/1 0.20 0.20 0.08 Fetot mmol/kg 0.002 7e-4 9e-9 Altot mmol/kg 1e-3 1e-4 4e-5 Na+ mmol/1 0.61 1.7 1.7 K+ " 0.04 0.02 6e-4 Ca2+ " 0.58 0.14 0.11 Mg2+ " 0.18 0.06 0.03 Mn2+ " - 6e-4 -

Stot mmol/kg 0.03 0.07 0.07 sol- mmol/1 0.03 0.07 0.06 er " 0.1 0.63 0.62 F - - 0.13 0.27 0.13 log f(C02) -2.9 -3.4 -4.5 log f(02) -71.8 -72.5 -74.9

* NETP ATH mass-balance calculated for indicated plagioclase solid solution.

104

Table 7-8. Results of EQ3/6 modelling of jlowpath KR3 reaction step 1.

FLOWPATH KR3 I reaction step: Rech. :=:::} KR3/T6

Final mass balance calculation results from NETPATH EQ6 _illhase mmol/kg mmol/1

Precipitating: goethite 0.44 0.43 K-montmorillonite 0.57 0.44

Dissolving: 02(g) 0.70 0.70 or g. 0.58 0.58 calcite 0.25 0.25 pyrite 0.002 0.002

plagioclase* (25%) anorthite 0.25* 0.0625 (75%) albite high 0.1875

biotite** (25%) phlogopite 0.19** 0.0475 (75%) annite 0.1425

kaolinite 0.41 0.41 Ion exchange: Mg 0.03 out 0.03 out

Ca 0.07 in 0.07 in Na 0.08 out 0.08 out

EQ3 calculation EQ6 result for results for Final Initial Final Vuorimaki KR3/T6

pH 6.3 8.4 8.6 Eh measured m V -100*** +70

Hs-;so4z- " -175 +720 Alkalinity meq 0.18 1.47 1.27 Ctot mmol/kg 0.46 1.49 1.29 Si02 mmol/1 0.20 0.20 0.54 Fetot mmol/kg 1e-4 4e-4 2e-18 Altot mmol/kg 1e-20 3e-4 0.60 Na+ mmol/1 0.11 0.23 0.22 K+ " 0.02 0.03 0.07 Ca2+ " 0.08 0.46 0.41 Mg2+ " 0.03 0.14 3e-30 Mn2+ " 2e-4 3e-4 1e-4 Stot mmol/kg 0.03 0.04 0.03 S042

- mmol/1 0.03 0.04 0.03 er " 0.03 0.02 0.04 log f(COz) -2.3 -3.6 -3.8 log f(02) -70.7 -67.8 -2.5

* ** ***

NETP ATH mass-balance calculated for indicated plagioclase solid solution. NETP ATH mass-balance calculated for indicated biotite solid solution. set value.

105

Table 7-9. Results of EQ3/6 modelling of jlowpath KR3 reaction step 2.

FLOWPATH KR3 I reaction step: KR3/T6 ===> KR3/T2

Final mass balance calculation results from NETPATH EQ6 I phase mmoVkg mmol/1

Precipitating: calcite 0.03 0.07 pyrite 0.01 0.01 kaolinite 0.07 0.07 chalcedony 0.20 0.30

Dissolving: or g. 0.04 0.04 goethite 0.01 0.01

plagioclase* (25%) anorthite 0.11 * 0.0275 (75%) albite high 0.0825

Na2S04 0.02 0.02 Na2CaC14 0.01 0.01

Ion exchange: Na 0.06 in 0.06 in Mg 0.03 out 0.03 out EQ3 calculation EQ6 result for results for Final Initial Final KR3/T6 KR3/T2

pH 8.4 8.2 8.4 Eh measured m V +70 +20

HS"/SO/- " -175 -165 -250 Alkalinity meq 1.47 1.05 1.45 Ctot mmol/kg 1.49 1.08 1.46 SiOz mmol/1 0.20 0.19 0.08 Fetot mmol/kg 4e-4 1e-3 4e-3 Altot mmol/kg 3e-3 1e-3 1e-5 Na+ mmol/1 0.23 0.45 0.44 K+ " 0.03 0.03 0.03 Ca2+ " 0.46 0.46 0.43 Mgz+ " 0.14 0.11 0.08 Mn2+ " 3e-4 7e-4 2e-4 Stot mmol/kg 0.04 0.04 0.05 so4z_ mmol/1 0.04 0.04 0.05 er " 0.02 0.05 0.06 F " 0.05 0.10 0.05 log f(COz) -3.6 -3.5 -3.6 log f(02) -67.8 -67 -73

* NETP ATH mass-balance calculated for indicated plagioclase solid solution.

106

Table 7-10. Results of EQ3/6 modelling offlowpath KR4 reaction step 1.

FLOWP ATH KR4 I reaction step: Rech. :::::} KR4/T6

Final mass balance calculation results from NETPATH EQ6 !phase mmol/kg mmol/1

Precipitating: pyrite 0.01 0.01 goethite 0.30 0.26 K-montmorillonite 0.40 0.43

Dissolving: 02 (g) 0.64 0.64 or g. 0.60 0.60 calcite 0.27 0.27

plagioclase* (25%) anorthite 0.20* 0.05 (75%) albite high 0.15

biotite** (25%) phlogopite 0.14** 0.035 (75%) annite 0.105

kaolinite 0.26 0.26 Ion exchange: Ca 0.07 out 0.07 out

Mg 0.04 in 0.04 in Na 0.06 in 0.06 in

EQ3 calculation EQ6 result for results for Final Initial Final Vuorimaki KR4/T6

pH 6.3 8.1 7.4 Eh measured m V -100*** +110

Hs-;so/ " -160 -185 Alkalinity meq 0.18 1.29 1.17 Ctot mmol/kg 0.46 1.33 1.33 Si02 mmol/1 0.20 0.25 0.009 Fetot mmol/kg 1e-4 0.006 0.006 Altot mmol/kg 1e-20 4e-4 0.21 Na+ mmol/1 0.11 0.32 0.32 K+ " 0.02 0.03 0.03 Ca2+ " 0.08 0.32 0.33 Mg2+ " 0.03 0.18 0.05 Mn2+ " 2e-4 0.007 2e-4 Stot mmol/kg 0.03 0.02 0.01 SO/ mmol/1 0.03 0.02 0.01 er " 0.03 0.04 0.04 F " - - -log f(C02) -2.3 -3.3 -2.7 log f(02) -70.7 -75.0 -72.2

* ** ***

NETP ATH mass-balance calculated for indicated plagioclase solid solution. NETP ATH mass-balance calculated for indicated biotite solid solution. set value.

107

Table 7-11. Results ofEQ3/6 modelling ofjlowpath KR4 reaction step 2.

FLOWP A TH KR4 I reaction step: KR4/T6 =::} KR4/T2

Final mass balance calculation results from NETPATH EQ6 I phase mmol/kg mmol/1

Precipitating: calcite 0.15 0.18 chalcedony 0.55 0.41 kaolinite 0.16 0.11 pyrite 0.03 0.03 K-montmorillonite 0.01 0.08

Dissolving: or g. 0.14 0.14

plagioclase* (25%) anorthite 0.28* 0.07 (75%) albite high 0.21

goethite 0.03 0.03 Na2S04 0.08 0.08

Ion exchange: Mg 0.09 out 0.10 out Na 0.18 in 0.18 in

EQ3 calculation EQ6 result for results for Final Initial Final KR4/T6 KR4/T2

pH 8.1 8.6 8.8 Eh Hs-;so/ m V -160 -160 -340 Alkalinity meq 1.29 1.37 1.24 Ctot mmol/kg 1.33 1.37 1.25 Si02 mmol/1 0.25 0.10 0.07 Fetot mmol/kg 0.006 0.002 0.008 Altot mmol/kg 4e-4 0.001 0.008 Na+ mmol/1 0.32 0.91 0.79 K+ " 0.03 0.03 0.004 Ca2+ " 0.32 0.24 0.20 Mg2+ " 0.18 0.08 0.05 Mn2+ " 0.007 5e-4 0.004 Stot mmol/kg 0.02 0.02 3e-6 so42- mmol/1 0.01 0.02 -

er " 0.04 0.06 0.04 F " - 0.15 0.12 log f(C02) -3.3 -3.8 -4.1 log f(02) -75.0 -59.3 -77.6

* NETP ATH mass-balance calculated for indicated plagioclase solid solution.

108

Table 7-12. Results of EQ3/6 modelling offlowpath KR4 reaction step 3.

FLOWPATH KR4 I reaction step: KR4/T6 =::} KR4/Tl

Final mass balance calculation results from NETPATH EQ6 I phase mmoVkg mmol/1

Precipitating: calcite 0.09 0.06 pyrite 0.04 0.03 kaolinite 0.06 0.06 chalcedony 0.23 0.32

Dissolving: or g. 0.16 0.16 goethite 0.03 0.03

Plagioclase* (25%) anorthite 0.10* 0.025 (75%) albite high 0.075

Na2S04 0.04 0.04 Ion exchange: Mg 0.07 out 0.07 out

Na 0.14 in 0.14 in

EQ3 calculation EQ6 result for results for Final Initial Final KR4/T6 KR4/Tl

pH 8.1 8.6 8.7 Eh measured m V +110 -70

Hs-;so42- " -160 -275 -330 Alkalinity meq 1.29 1.34 1.31 Ctot mmol/kg 1.33 1.34 1.39 Si02 mmol/1 0.25 0.16 0.07 Fetot mmol/kg 0.006 0.005 0.007 Altot mmoVkg 4e-4 2e-3 1e-5 Na+ mmol/1 0.32 0.74 0.62 K+ " 0.03 0.03 0.03 Ca2+ " 0.32 0.25 0.28 Mg2+ " 0.18 0.10 0.11 Mn2+ " 0.007 0.001 0.005 Stot mmol/kg 0.02 0.02 3e-6 so42- mmol/1 0.01 0.02 -er " 0.04 0.05 0.04 F " - 0.15 0.12 log f(C02) -3.3 -3.8 -3.9 log f(02) -75.0 -74.1 -77.7

* NETP ATH mass-balance calculated for indicated plagioclase solid solution.

109

8. SUMMARY AND IMPLICATIONS FOR SAFETY ASSESSMENT

The interpretation and modelling of hydrogeochemical evolution at Kivetty was created by extensive use of carbon isotope data in conjunction with water chemistry and thermodynamic speciation calculations. Information on processes of groundwater evolution was also interpreted from fracture mineralogy, isotope hydrochemistry (8H-2, 80-18, 8S-34(S04), Sr-87/Sr-86, U-234/U238) and gas analyses. Carbon-13(DIC) data were especially important because these provided additional criteria linking water chemistry with mineral mass transfer.

Radioactive isotopes C-14 and H-3 were found useful in evaluating uncertainties connected to the groundwater samples, their locality and sampling. A group of samples (KR2/T1, KR3/T6, T5, T3, T2, T1, KR5/T2, T1 in App. 6) contained a considerable young water component at great depths. This contamination is assumed to result from technical activities and problems in boreholes, as there appears to be no natural cause such as a fast flowpath for the deep occurence of young water. However, three of the samples (KR3/T5, KR5/T2, T1) are contaminated so heavily that they actually represent young, recently recharged groundwater and are considered to be chemically representative for further use as reference samples. Other samples from KR3 were also used in interpretations and modelling to get some pieces of information about the evolution around the borehole, but the results must be considered with great reserve. Several samples (KR2/T5, KR4/T4, T2, T1, KR5/T6, BT) show slight indication (H-3 < 2.8 TU) of recent water inmixing. The contamination is considered insignificant, and the carbonate chemistry essential to understanding the hydrogeochemistry at Kivetty is strongly reliable since dissolved carbonate concentrations in deep waters are higher than in shallow ones. The C-14 content in these samples is certainly disturbed and the calculated mean residence times are probably slightly low.

The use of 8C-13 data in mass-balance calculations, either as a constraint or computed as a Rayleigh distillation problem for the modelled mass transfer, permits the estimation of site-specific values for calcite and dissolved organic matter, calculation of the quantity of calcite and organic-carbon-derived carbonate in groundwater, and the detailed separation of flowpaths. In the recharge state a semi -open system evolution of organic-carbon-derived dissolved carbonate is applied. Oxidation of organic matter (peat and plant debris) by oxygen gas during recharge buffers carbon isotope composition of the C02, which later reacts with fracture calcite and silicates along the flowpath. This permits estimation of average C-14 content of dissolved organic matter in peat bogs and till at the site. A dynamic model to calculate the C-14 of decomposing organic matter during recharge was developed. The model takes into account the change of C-14 content with time in organic carbon dissolved e.g. from peat dominated soil which began to form after glaciation. The mass transfer and carbon isotope data for the groundwater indicate a range for modem input of C-14 from the soil. 35 pmC is used for recharge infiltrating through bogs and 50 pmC for infiltration through smaller depressions.

Schematic conceptualisation of the Kivetty study site (Fig. 8-1) summarises generally occurring hydrogeochemical reactions, carbon isotope evolution and hydrogeological features realised during interpretation of recent geochemical and hydrogeological data.

110

All the modelled results, compiled in Tables 7-2 and 7-3, account for observed changes in the water chemistry. They are supported by forward thermodynamic reaction calculations (EQ6), as well as by measured the stable isotopic content of dissolved inorganic carbon, the S-34 content of dissolved sulphate, the Sr-87 /Sr-86 ratio of groundwaters, and also by the fracture mineralogy. The presented conclusions also have implications for performance assessment and safety analysis such as the interpretation of evolution of pH and redox conditions, the consistency between geochemistry and the hydrogeological model, and flow ages adjusted by C-14 data. The main conclusions are:

1. The dominant groundwater processes during recharge in overburden and bedrock are organic carbon oxidation and silicate weathering. Coupled oxidation and hydrolysis of biotite seems important according to Sr isotopic data. Ferrous biotite (annite) decomposition has produced fairly abundant limonite coatings on fracture walls in the upper part of the bedrock. These processes limit modem infiltration of oxygen to shallow levels. This is also inferred by hydrogen sulphide observations already in uppermost sampling sections in the bedrock and uranium concentrations and isotopic systematics. Hydrogen sulphide is probably produced by microbially mediated S04

reduction using dissolved organic carbon as an electron donor. The sulphate reduction is supported by sulphur isotopic results and partial precipitation by fracture mineral observations of late stage pyrite infillings on calcite precipitates deeper in the bedrock (Gehor et al. 1995). Also recent observations of sulphate and iron reducing bacteria (Haveman et al. 1998) in groundwaters augment the interpreted reduction and precipitation processes. Sulphur isotopic test calculation does not at any rate indicate significantly larger sulphur circulation during groundwater evolution than the present concentrations assume for dissolved sulphate and sulphide. Although hydrogen sulphide concentrations are small the reaction paths (EQ3/6 calculations) indicate a thermodynamic redox level of about -200mV ... -300mV depending on pH in deep groundwater. Ferrous biotite is susceptible in aerobic conditions (Acker and Bricker 1992, Malmstrom et al. 1995, Blum & Erel 1997, Bullen et al. 1997) to oxygen and can be important redox buffer material both after closure of waste storage or against possible oxygen-bearing glacial melt water infiltration.

2. Following aerobic processes during recharge, dissolution of fracture calcite and silicates, most likely plagioclase, are active in C02 rich groundwater conditions increasing alkalinity, pH and 8C-13 (DIC), and decreasing the partial pressure of C02

in groundwater. Cation exchange with certain primitive montmorillonite precipitation may control K and Mg concentrations on a low level. Alkalinity enrichment varies between the boreholes and only the data from KR1 show clear initial enrichment -later a depletion curve which is generally interpreted (Nordstrom et al. 1989) to be coupled with calcite saturation and precipitation during evolution of groundwater. Carbonate concentration in the other borehole groundwaters is typically quite low, allowing pH to increase to a fairly high level (about 8.3 - 8.5) in most samples before calcite reaches the saturation pH, whereas the level in KR1 is about 7.8 - 8. Later, probably minor calcite precipitation promotes silicate hydrolysis increasing the pH to 8.5 - 9 during long residence time.

Salinity remains at a low level even in samples showing long residence time, which may explain the minor depletion of dissolved carbonate, because calcite equilibrium

111

is not violated by a significant input of Ca into the groundwater system. The slight enrichment of Cl is difficult to explain without mixing of some old, possibly brine­based saline end-member, which would have been mixing into the system by some leaching process of matrix fluids, fluid inclusions or grain boundary salts as suggested by Nordstrom et al. (1989), Gascoyne et al. (1992) and Pitk:anen et al. (1996a) for various goundwaters on the Fennoscandian and Canadian shields. Groundwater composition is quite sensitive to disturbance of the geochemical system because of its slight buffering capacity of low ionic strength, which could particularly affect pH. On the other hand, alkalinity is higher than in typical saline ground waters in the Fennoscandian shield (e.g. Olkiluoto) but those environments contain greater amounts of pH-buffering calcite in fractures. The buffering capacity in disturbed conditions should be modelled separately as hypothetical cases.

Figure 8-1. 3-D illustration of the conceptual hydrogeological model with hydrogeochemical and isotopic evolution at the Kivetty site. Major fracture zones (coded by R) are based on bedrock models (Paulamiiki et al. 1996, Saksa et a/.1996). Blue arrows represent flow directions based on a hydrogeochemical view. Rectangles contain measured/modelled '6C-13(DJC) in %o PDB, adjusted C-14(DIC) age in years and measured '60-18 in %o SMOW Rounded rectangles contain the main sources with estimated carbon isotopic data and sinks during hydrogeochemical evolution.

112

3. The carbon isotope data coupled with mass balance calculations indicate variation in the input of dissolved carbonate from different sources, which is reflected in the 8C-13 values of samples. Groundwater residence times adjusted for reaction models cover virtually the time span since the site was released from under the ice sheet about 9, 700 years ago. Some sub glacial ages up to 15,000 years were determined. They are considered mixed ages derived from the mixing of preglacial water and melt water from the retreated ice sheet. The melt water contribution is estimated to be about 20% using a simple two end-member mixing model. If melt water intrusion had contained significant oxygen input deep into the bedrock this should be observed. The most likely oxygen consumer is ferrous iron oxidation producing low­temperature ferric precipitates as limonite. The process also increases salinity, because iron oxyhydroxide precipitation aids hydrolysis by acidifying groundwater. For instance 1 mmol/1 dissolved 0 2 produces, by biotite weathering, about 4.5 meq!l of both cations and anions, which clearly exceeds observed dissolved concentrations. Limonite could have dissolved later during sulphate reduction and anaerobic oxidation of organic carbon, but the salinity levels should have been conserved. Anaerobic processes should have caused significant shifts on S and C isotopic data. In addition the mass transfer involved in this anaerobic process can be predicted (Plummer et al. 1990) by mass balance calculations which employ isotopic data. The trends in the hydrochemical and isotopic data from the Kivetty site, the low remaining salinity, total mass transfer in the reaction models or fracture mineral occurrences do not suggest any strengthened chemical processes or oxygen intrusion during the evolution of water samples showing melt water signatures.

4. The interpretation of hydrogeochemical evolution is not in complete accordance with the available hydrogeological models (e.g. Taivassalo & Meszaros 1993, 1994, Saksa et al. 1996, Saksa et al. 1998, Kattilakoski & Meszaros 1999). Congruently flow velocities seem to be slow and any really dynamic natural flowpath with deep observed young groundwater cannot be shown according to C-14 age calculations or groundwater chemistry. However, the hydrogeological system around KR5 requires detailed revision to understand the natural system. Deep in the bedrock is a slightly saline, older subaquifer characterised by subglacial or deglacial residence times has been observed. The observation may indicate a missing subhorizontal, more dynamic hydrogeological unit or generally less permeable rock body which determines the upper boundary of the old aquifer as varying from depth 160m in KR2 via 300m in KR1 and KR5 to 350m in KR3.

The hydrochemistry indicates restricted hydrogeological systems around each borehole deep in the rock where flow barriers may prevent subhorizontal flow connecting boreholes along interpreted assembly of fracture zones. The flow barriers between the boreholes suggest that the assembly consists of hydraulically limited pieces of intense fracturing rather than extensive, uniform fracture zones presented in the hydrogeological models. The interpretation of hydrogeochemical and hydrogeo­logical systems are mainly based on different data and both models consist of hypothesis and uncertainties. Thus integrating the results of hydrogeochemistry more closely in an iterative process to revise the hydrogeological model in future (e.g. Bath & Lalieux 1998, Pitkanen et al. 1998) may improve the consistency of the models and decrease the uncertainties of them.

113

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APPENDICES:

Appendix 1. Lithology, open and filled fracture frequency, hydraulic conductivities, interpreted fracture zones, and groundwater sampling sections with measured hydraulic heads for boreholes KR1-KR5. Lithology and fracture frequency interpretations are based on the studies by Front & Okko (1995 and refs. therein). Hydraulic conductivity measurements of KR1-KR3 and KR5 are after Rouhiainen (1996a), and measurements of KR4 after Kuusela-Lahtinen & Front (1990). Interpreted fracture zones are based on Saksa et al. (1996). Packed-off sampling sections and hydraulic heads are mostly presented in Saksa et al. (1993). Details for KR5/BT are from Snellman et al. (1995a) and Rouhiainen (1996a). PGR = porphyritic granite, PGRDR = porphyritic granodiorite, GR = granite, GRDR = granodiorite, QMZDR = quartz monzodiorite, GB = gabbro, AFB = amphibolite, MY = mylonite, MGN = mica gneiss, MS = mafic schist.

Appendix 2. Packed-off intervals used for geochemical groundwater sampling, together with hydraulically conductive depth intervals, inferred locations of conductive fractures, and fracture minerals considered to be in contact with sampled groundwater. Hydraulic conductivities in bold type are 2-metre interval measurements with the difference flow equipment (Rouhiainen 1996a). Other conductivities are 2-metre estimates from 7-, 10-, and 31-metre interval conductivity measurements. Certain adjacent measurements have been combined, and the shown conductivity is the arithmetic mean.

Appendix 3. Hydraulic head field in R9 and R11 based on the study by Taivassalo & Meszaros (1994) with intersection lines of other fracture zones and intersections of boreholes KR1, KR4 and KR5.

Appendix 4. Hydraulic head field in R12 and R15 based on the study by Taivassalo & Meszaros (1994) with intersection lines of other fracture zones and intersections of boreholes KR1, KR2, KR4 and KR5.

Appendix 5. Hydraulic head field in R22 and R23 based on the study by Taivassalo & Meszaros (1994) with intersection lines of other fracture zones and intersections of boreholes KR1, KR4 and KR5.

Appendix 6. Hydrogeochemical data used in the present modelling study. Digits shown in italics = concentration below detection limit, digits shown with bold type = uncertain concentration determination. KI-KR1/T7* is from preliminary investigations (e.g. Pitkanen et al. 1992) and represents shorter packer interval than entire length of KR1/T7 in App. 1.

Appendix 7. Evaluation of the hydrogeochemical data of Kivetty.

Appendix 8. Samples included in the modelling study in spite of failing the quality classification.

0

E

Kivetty Borehole KR1

Uthology

D PGA

D GR

@ PGRDR

~ GRDR

f?& MY

~ MGN

Fracture Frequency (1/m) 0 15

122

Hydraulic Interpreted Conductivity {m/s) Fracture Zones E-1 0 E-8 E-6

R25 R22

R15

R11

Appendix 1 (115)

Sampling Sections and Hydraulic Head (m) 160 175 190

T8

T7

TG

TS

z 500 ~--~~~~=---r----+-+-+~~~~--~--+-+---------~ Et: w Cl

SCALE 1:5 0 0 0

T4

T2

T1

3D-ROCK VTT

0

250

E

Kivetty Borehole KR2

Uthology

0 PGA

~ PGRDR

f2j GRDR

I MS

123

Fracture Hydraulic Frequency (1/m) Conductivity (m/s) 0 1 5 E-10 E-8 E-6

Appendix 1 (2/ 5)

Interpreted Sampling Sections Fracture Zones and Hydraulic Head {m)

1 60 175 190

R23 TB T7

R15 T6

TS

R17 T4

T3

T2

T1 z 500 r---~~~~--~~---r-+-+-+~-1----~--~+=~------~ tl: w 0

750 r---------1-----r----+-+-+~_,-4r---------+---------~

1000 ----------------~----~~~~~~------------------~

SCALE 1:5000 3D-ROCK VTT

E -:r:.· ~ LU a

0

250

500

Kivetty Borehole KR3

Uthology

PGRDR

0 GR

0 PGA

Fracture Frequency (1/m) 0 15

Appendix 1 (3/5)

124

Hydraulic Interpreted Sampling Sections Conductivity {m/s) Fracture Zones and Hydraulic Head (m) E-1 0 E-8 E-6 160 175 190

T4

T3

R27 T2

R28 T1

1000 ~--------~----~--~~~~~~~--------~--------~

SCALE 1:5000 3D-ROCK VTT

0

250

E

Kivetty Borehole KR4

Uthology

D PGA

-GB

D GR

~ PGRDR

8 QMZDR

fL1

Fracture Frequency (1/m) 0 15

Appendix 1 ( 4/5)

125

Hydraulic Interpreted Sampling Sections Conductivity (m/s} Fracture Zones and Hydraulic Head (m)

E-1 0 E-8 E-6 160 175 190

TB

T7

T6

TS

T4

T3

T2

T1

€ 500 r---~~~~r---+---~~--r-r-+-+---~---r-+~~-----~ UJ a

750 r----------~-----+-----;---r-r-+-+-~----------+------------~

1000 ----------~----~----~~~~~-----------------------

SCALE 1:5000 3D-ROCK VTT

0

Kivetty Borehole KR5

Uthology

D PGA

~ PGA OR

D GR

B GRDR

8 QMZDR

~ AFB

Fracture Frequency (1/m) 0 15

Appendix 1 (5/5)

126

Hydraulic Interpreted Sampling Sections Conductivity (m/s) Fracture Zones and Hydraulic Head ~) E-1 0 E-6 E-6 160 175 190

11111 1111m 1111m 111111 11111

- f ~ TB

T7

-

R8 ~ .......

T4

T3 =

T2

1000 ----------------~------~~~~~---------------------

SCALE 1:5 000 3D-ROCK VTT

Packed-off intervals used for geochemical groundwater sampling, together with hydraulically conductive depth intervals, inferred locations of conductive fractures, and fracture minerals considered to be in contact with sampled groundwater. Hydraulic conductivities in bold type are 2-metre interval measurements with the difference flow equipment (Rouhiainen 1996a). Other conductivities are 2-metre estimates from 7-, 10-, and 31-metre interval conductivity measurements. Certain adjacent measurements have been combined, and the shown conductivity is the arithmetic mean. Pgr = porphyritic granite, Pgrdr = porphyritic granodiorite, Gr =granite, Grdr = granodiorite, Qmzdr =quartz monzodiorite, Gb = gabbro, Amph = amphibolite, My= mylonite. Sec. up= upper end of packer interval in borehole length metres, Sec. low = lower end of packer interval in borehole length metres.

KRl SamJ!Ie Host Sec. up Sec. low Length Hydr.cond. Sec. up Sec. low Length Conductive fractures l) Significant fracture minerals 2)

KRI!f6 Pgr 300.0 345.0 45.0 7.6E-8 3I3.4 3I5.4 2.0 3I4.7, 3I4.5, 3I5.I Cc, Chl 9.2E-7 335.4 337.4 2.0 336.2 Chl

KRI!f5 Pgrdr 370.0 4I5.0 45.0 6.3E-8 381.5 383.5 2.0 382.8 Cc, Qtz, Clay J.OE-8 395.5 399.5 4.0 396.6, 396.7, 397.9-399.5 Chl, Cc, Qtz, Py, Smec

KRI!f4 Pgrdr/My 5I5.0 555.0 40.0 5.4E-7 529.7 531.7 2.0 529.8, 530.3-531.0, 531.2, 531.5 Chl, Cc, Py, Kaol 4.9E-7 537.7 539.7 2.0 537.8-539.4 Chl, Qtz, Kaol, Py, Po, Sul, FeH, Hem

KRI!f3 Pgrdr/My 720.0 795.0 75.0 5.3E-9 742.0 746.0 4.0 742.2-742.6, 743.8, 744.I, 745.I-745.6 Chl, FeH, Py, Kaol l.2E-9 750.0 752.0 2.0 750.44-752.0 Chl, Cc, Kaol l.SE-9 768.1 770.I 2.0 768.6, 769.8 Chl, Cc, Kaol

KRI!f2 Pgrdr/My 8I5.0 855.0 40.0 l.4E-6 820.2 826.2 6.0 820.5-820.9, 821.8-821.9, 823.2-823.9, 824.9-826.I Chl, Cc, FeH

KR2 Sample Host Sec. up Sec. low Length Hydr.cond. Sec. up Sec. low Length Conductive fractures l) Significant fracture minerals 2)

KR2ff7 Pgr 45.0 75.0 30.0 5.8E-8 46.6 48.6 2.0 46.6-48.4 FeH, Kaol 5.7E-8 52.6 56.6 4.0 52.8-53.3, 55.I-56.4 FeH, Hem, Ep, Clay 4.2E-8 62.7 64.7 2.0 64.3 FeH l.4E-8 66.7 70.7 4.0 67.6-67.7, 69.6 FeH, Chl, Kaol, Cc

KR2!T6 Pgr/Pgrdr 75.0 I55.0 80.0 4.9E-8 82.7 84.7 2.0 82.7, 84.0, 84.I, 84.3, 84.5, 84.7 FeH, Qtz, Chl, Sul 2.0E-6 86.7 88.7 2.0 87.4, 88.7 Chl, Qtz, FeH 6.8E-8 90.7 92.7 2.0 91.8, 91.9, 92.0, 92.I Chi, Qtz 5.6E-8 94.7 98.7 4.0 94.8-96.7, 96.9, 97.2, 98.0, 98.6 FeH, Chl, Cc S.SE-8 II4.8 I16.8 2.0 II5.0, II5.7, I16.0 Chl, FeH

KR2!T5 Pgrdr/Grdr I60.0 I90.0 30.0 l.OE-II I60.9 I68.9 7.9 l.OE-II I74.8 I76.8 2.0

KR2!f4 Grdr/Pgrdr 190.0 250.0 60.0 J.SE-8 I92.9 194.9 2.0 193.8, I94.2, I94.3 Chl, Sul 2.2E-8 249.0 251.0 2.0 25l.I, 251.2 Chl

KR2!f1 Pgrdr/Grdr 450.0 500.0 50.0 9.1E-10 467.4 469.4 2.0 469.0-469.I Cc, Sui, Mont, Kaoi

......... N .._J

> "0 "0

('D

::l 0.. ><" N ~

.........

......... VJ '-'

KR3 Sample Host Sec. up Sec. low Length Hydr.cond. Sec. up Sec. low Length KR3ff7 Pgrdr/Gr 85.0 130.0 45.0 3.9E-8 99.0 101.0 2.0

5.4E-7 105.0 109.0 4.0 2.8E-7 119.0 121.0 2.0 1.4E-8 123.0 125.0 2.0 4.2E-8 127.0 129.0 2.0

KR3ff6 Pgrdr 135.0 185.0 50.0 4.2E-8 145.1 149.1 4.0 8.5E-9 153.1 157.1 4.0 1.3E-8 159.1 163.1 4.0 6.5E-9 173.1 175.1 2.0

KR3ff5 Pgrdr 190.0 255.0 65.0 1.3E-8 189.2 191.2 2.0 2.0E-9 199.2 201.2 2.0 8.5E-9 227.2 229.2 2.0

KR3ff3 Pgr 340.0 410.0 70.0 5.7E-9 373.5 375.5 2.0 l.OE-9 389.5 391.5 2.0 S.lE-9 397.5 399.5 2.0 2.4E-9 401.5 403.5 2.0

KR3ff2 Pgr 410.0 460.0 50.0 1.8E-6 437.6 439.6 2.0 KR3ff1 Pgr/Pgrdr 460.0 500.0 40.0 2.1E-7 479.6 483.6 4.0

KR4 Sample Host Sec. up Sec. low Length Hydr.cond. Sec. up Sec. low Length KR4ff7 Pgr 90.0 130.0 40.0 9.9E-6 102.0 125.4 23.4 KR4ff6 Pgr 135.0 155.0 20.0 6.5E-6 141.5 150.0 8.4

6.5E-6 156.5 162.8 6.3 KR4ff5 Pgr/Gb 170.0 195.0 25.0 3.1E-6 174.2 180.5 6.3 KR4ff4 Pgr/Gb 225.00 260.00 35.00 3.2E-8 235.8 237.8 2.0

3.2E-8 240.1 244.3 4.1 1.0E-8 250.0 252.0 2.0

KR4ff3 Pgr/Gr 260.0 310.0 50.0 4.0E-7 265.8 272.1 6.3 1.2E-6 291.8 298.1 6.3 7.0E-8 301.8 308.1 6.3

KR4ff2 Pgr/Pgrdr 320.0 410.0 90.0 4.8E-9 321.5 325.7 4.2 3.1E-8 388.0 396.4 8.4

KR4ff1 Pgrdr/Qmzdr 430.0 500.0 70.0 3.6E-8 439.4 443.5 4.1 3.6E-8 452.3 454.3 2.0

Conductive fractures t) 99.3, 99.5, 99.8, 99.9, 100.0, 100.1, 100.5 106.4, 106.7, 106.9, 107.0-107.9, 108.4, 108.6, 108.7 120.3, 120.5 123.3, 123.4, 125.0 129.0 145.7, 146.3, 146.4, 147.6, 147.7, 147.9, 148.0, 148.6, 149.0 153.9, 156.1, 156.6 159.6, 160.9

191.0

228.3 374.5 391.4 399.4

438.6, 438.8 480.6, 483.2

Conductive fractures t)

103.0-123.0, 123.7, 123.9, 124.8 140.0-150.0 157.0-159.6, 160.1, 162.3, 162.4, 162.5, 162.6, 162.8 175.6-179.0 236.6 240.4, 244.3 250.4, 250.5, 250.6, 250.7, 251.1, 251.4, 251.8 251.9 267 .0, 272.0 293.6, 294.3, 294.4, 295.2, 296.5, 297.3, 297.5, 297.6 303.5, 306.9 322.9, 323.3, 323.4, 323.6, 325.0 388.3, 390.5, 392.3, 393.6, 394.2, 396.2 439.7, 441.6, 442.8, 443.1, 443.2, 443.3 453.3, 453.5

Significant fracture minerals 2)

Chi, FeH,Ep Cc, Chi, Qtz, ill Cc, Chi, Py Chi, Cc, Py Cc,Chl Cc, Chi, Py, Clay FeH, Clay FeH, Kaol Chi, Qtz, ill, Smec, Py Cc, Py, Smec, Chi, Fsp, Qtz Clay, Py Bt, Smec, Qtz, Cc, Py Chi, Cc, Qtz Chi, Cc, Qtz, Sui Cc, Chi, Sui Clay, Cc Cc Chi

Significant fracture minerals 2)

FeH, Qtz, Smec FeH, Ep, Qtz FeH, Ep, Bt, K-fsp, Qtz, Ab, Clay FeH, K-fsp, Qtz, Ep Chi, Cc, Qtz FeH, Clay Chi, Clay FeH, Qtz, Chi, Sui, K-fsp, Kaol Chi, FeH, Cc, Kaol, Fsp, Bt, Smec Chi, Cc, FeH, Fsp FeH,Chl, Cc Chi, Po, Cc, Fsp Chi, Sui, Qtz Cc, Chi

........ N 00

~ '"d '"d (1)

::s 0.. ~-

N ~

N ........... w ,_..,

KRS Sample Host Sec. up Sec. low Length Hydr.cond. Sec. up Sec. low Length Conductive fractures t) Significant fracture minerals 2)

KR5ff7 Pgr/Pgrdr 60.0 162.0 102.0 2.4E-7 61.3 63.3 2.0 61.4, 62.3 FeH, Chi 1.7E-7 77.3 79.3 2.0 78.0-79.1 FeH, Chi, Kaol 1.8E-7 85.3 87.3 2.0 86.9 FeH 2.1E-6 91.3 93.3 2.0 91.3, 92.2, 92.3, 92.5, 92.8, 92.9 FeH,Chl 1.6E-6 111.4 113.4 2.0 112.5, 112.6, 112.9 FeH, Chi 2.7E-7 115.4 117.4 2.0 115.4, 115.6, 116.4, 117.0 FeH, Chi, Kaol

KR5!f6 Pgrdr/Gr 215.0 255.0 40.0 8.6E-7 233.6 235.6 2.0 235.3 Chi, FeH, Cc 6.1E-6 251.6 253.6 2.0 252.4 FeH

KR5ff5 Pgrdr 265.0 300.0 35.0 l.lE-8 267.6 269.6 2.0 268.2 Cc, FeH, Chi, Qtz KR5ff4 Pgrdr 300.0 350.0 50.0 1.0E-9 307.7 309.7 2.0

1.3E-9 312.0 317.6 5.6 314.1-315.1, 315.2, 315.8, 316.0, 316.3, 316.9, 317.0 1.8E-9 323.7 325.7 2.0 2.5E-9 343.5 345.5 2.0 2.5E-9 347.8 349.8 2.0

KR5ff3 Pgrdr 350.0 405.0 55.0 1.5E-9 355.6 357.6 2.0 6.8E-10 363.5 369.8 6.3 1.8E-9 390.0 392.0 2.0 392.0 1.3E-9 393.8 395.8 2.0 395.3

KR5ff2 Pgrdr 415.0 455.0 40.0 4.0E-9 425.9 427.9 2.0 427.5 Cc 2.9E-8 429.9 431.9 2.0 430.1, 430.7, 431.3 Cc, Chi, Qtz 4.1E-9 433.9 437.9 4.0 434.4-435.5, 436.6-437.3 Cc, Chi, Qtz 3.2E-8 443.9 445.9 2.0 445.2 Chi, Cc, Sui, Kaol 1.9E-8 447.9 449.9 2.0 449.1, 449.7 Cc,Chl

KR5ff1 Pgrdr/Qmzdr 455.0 500.0 45.0 4.4E-9 455.9 457.9 2.0 457.7, 457.9 Cc,Chl l.lE-8 461.9 465.9 4.0 462.8-463.7, 464.8, 464.9 Cc, Chi 1.3E-8 488.0 490.0 2.0 489.8 Cc, Chi, Ep 1.3E-9 492.0 494.0 2.0 Cc, Chi, Ep

KR5/BT Pgr/Amph 735.0 853.0 118.0 5.8E-6 740.4 742.4 2.0 740.9, 741.0, 741.3, 741.5 FeH, Chi, Cc, Clay l.OE-6 744.4 750.4 6.0 745.4-746.8, 747.1, 747.6-749.5, 749.9, 750.0, 750.1, 750.2 FeH, Chi, Cc, Clay 7.4E-6 754.4 756.4 2.0 755.1, 755.5, 755.9 FeH, Chi, Qtz, Cc, Clay 3.9E-6 760.4 764.4 4.0 760.9, 761.5-762.4, 763.3, 763.6, 763.8, 764.2 FeH, Chi, Qtz, Cc, Clay 7.1E-6 772.4 774.4 2.0 773.7, 773.8, 774.0-774.4 FeH, Chi, Clay l.SE-6 784.4 786.4 2.0 785.2, 785.7, 785.8, 785.9, 786.0, 786.1, 786.3 FeH, Clay 8.0E-6 798.4 800.4 2.0 799.1, 799.5, 800.2 FeH, Clay l.lE-6 808.5 810.5 2.0 808.9, 809.0, 809.1, 809.3, 809.5, 810.3, 810.4 FeH, Qtz, Cc, Sui, Chi l.SE-6 818.5 820.5 2.0 818.7, 818.9, 819.2-819.3, 819.8, 820.0 FeH,Ep,Cc l.OE-5 820.5 822.5 2.0 820.7, 821.0, 821.4, 821.5 FeH,Hem,Ep 2.2E-6 826.5 828.5 2.0 826.5-827.8, 828.1, 828.2 FeH, Qtz

I) After Vahanne & Front (1990a, 1990b), Kuusela-Lahtinen & Front 1990 and Melamed & Front (1995). 2) After Gehor et al. (1995). Bt = biotite, Cc= calcite, Chi= chlorite, Ep = epidote, FeH =iron hydroxides, Fsp =feldspar, Hem= hematite, ill= illite, Kaol =

kaolinite, K-fsp =potassium feldspar, Mont= montmorillonite, Po = pyrrhotite, Py =pyrite, Qtz =Quartz, Smec = smectite, Sui= unidentified sulphides.

......... N \0

> ""d ""d

(1) :::1 &. X N

,.-._ w ........ w '-"

130 Appendix 3

Hydraulic head field in R9 and Rll with intersection lines of other fracture zones and intersections of boreholes KRl, KR4 and KR5.

KIVETTY R9 Head [m]

187

Ea at

Head [m] 187

KIVETTY R11

18

181

166

16

North 16

131 Appendix 4

Hydraulic head field in R12 and Rl5 with intersection lines of other fracture zones and intersections of boreholes KRl, KR2, KR4 and KR5.

KIVETTY R12

Head [m] 187

184

181

178

175

172

16

Up

166

163 East

160

KIVETTY R15

.a MAX

ON GRI Head [m] 187

18

181

North

132 Appendix 5

Hydraulic head field in R22 and R23 with intersection lines of other fracture zones and intersections of boreholes KRl, KR4 and KR5.

KIVETTY R22

Head [m] ------. , 187

BI!Il ' ' ' ' ' •

KIVETIY R23

18

181

Head [m]

No. Type I Sample

I 1 spring 1

Lemetyinen 2 [spring IVuorimaki 3 shallowKI-KAI 4 ;shallow

1KI-KAI

si shallow[ KI-KA2 6 : shallow1 KI-KA2 7 . shallowi KI-KA2 S I shallow] Liimatainen 9 !FP! Kl-KRlff7*

10 FP! i'Kt-KRlff7* IIIFPI KI-KRlff6 12 FP! lKI-KRlff6

I3jFPI IKI-KRlff5 14, FPI KI-KR lff4 15.FPI -

1

.Kl-KRlff3 16[ FPI KI-KRlff2 17'FPI KI-KR4ff6 ISIFPI IKI-KR4ff5 19tFPI [KI-KR4ff4 20. FPI . KI-KR4ff3

211FPI 11KI-KR4ff2

22:FPI IKI-KR4ffl 23/Pl KI-KR4ffl 24[ FPI I KI-KR5ff6 25 FPI . KI-KR5ff4 261FPI KI-KR5ff2 27: FPI I KI-KR5ff I 2SIFPI 1KI-KR5/BT 29!FP2 I·KI-KR2ff5 30 FP2 KI-KR2ff4 31 I FP2 KI-KR2ff4 32 FP2 j KI-KR2ffl 33 I FP3 . KI-KR3ff6 34:FP3 [Kt-KR3ff5 35[ FP3 1 KI-KR3ff3 36,FP3 iKI-KR3ff2 37[FP3 KI-KR3ffl 3S FP3 IKI-KR3ffl

Hydrogeochemical data of Kivetty

Date Sec.up~ec.lov. EhPt: 0 2 Densityi EC[ pH[ Alkm. Alk"[ Al~01 ~ HCO~~· HCO~IAlkco\f eo~: eo~~· Acid;C021,ccl'co2,,cciKMn041

1DOC E.N.IDrill H20! TDs[ Si02

1 Si02

ddmmyy m ; m I mvl mg/1[ g/ml. mS/m I 1111:411] me4/ll 1111:4/l mg/1 mmnl/11 mc411] mg/1. mmol/1, mval/1, mg/1 mmol/1[ mg/1 mg/1 <7,. 91 I mg/( mg/1. mmnl/1

170S94 ! 4.4!0.99731 4.4 6.1 0.20 0.00 0.20, 12.20[ 0.200 0.00[ 0.001 0.00 0.21 i 9.241 0.210 3.2! 0.95 7.64

060994 12.5 0.9973 1 6.3 6.3 O.IS .. 0.00 .. O.ISI' 10.9S: O.ISOj 0.00[ 0.001 0.001

0.2Sil2.32: 0.2SO 1.5 1.60 4.9S, 0507SS , 12.3 7.4 1.09. 0.00

1,_09 66.51

1

1.091

0.00. 0.00 I 0.00

1

. 0.15 6.60. 0.150

1

1.3, 1.20 -2.31

1

.

I4078S 1

I 10.4. 7.4 1.12

1

. o.oo!_ 1.12' 6S.34 1.121

o.oo1

o.oo;. o.oo o.1oj 4.40

1

, 0.100 2.2: o.so o.s3 2S07SS I I 10.7_ 7.1 1.15 0.00. 1.15 70.17 1.15! 0.00 0.00 0.00, 0.13 5.72 0.130 2.1 1.00 -1.10! 0305S9

1 jio.5l 6.4

1

o.96[. o.oo1

o.96i 5s.ss: o.96o 1 o.oolo.oo, o.oo, o.m, J.os 0.010 2.411.80,-Ho 2205S9

1

, II.OI 6.7[· I.oo. o.oo[~ I.ooi6I.02! 1.001 o.oo_ o.ool o.oo. o.07I' 3.os 1 o.07o 1.9 o.90. -2.21: ISOS94 I 0.5 0.9977 13.6 7.2

1

1.341 0.001

1.34 SJ.76 1.34 0.00:0.001

0.00 0.19 S.36 0.190 2.2, 3.SO 5.29

o712ss I69.oiiso.sj-350 I 1

22.0

1

s.2: 1.901 o.oo1

'_·_9_o, 115.91. 1.90 o .. oojo.oo, o.ool_ .. o.o6 1 2.641 o.or,ol 3.516.30 -0.12.1 1312SS 169.0 IS0.5 -300 I 22.0 S.O • 2.00 0.00 2.00 122.0' 2.00 0.00 0.00 I 0.00 0.06, 2.64 0.060 3.S 5.SO -2.62

1 ~o 194 3oo.o I 345.0 I j o.o1(' o.99S6 ,9.5 .. 7.sl1 ~.1s 1·. o.oo [ 2.Isjl3ul

1

_ ~-'51. o.oo 1 o._oo_ .I o.oo.

1

._ o.o 'I o.53 ·. o.o 12

1

• 2.6 · ~ .. sol' -9 .. 941'_ 0~0594 300.0 345.01 0.00 .. 1.6 7.9 ~.101 0.00 2.10, 12S.l ~.10 0.0010.00 0.00 0.00 0.00: 0.000 : ... so -3.56

090394 i 3700 415.0. -3010.00 0.99S51 25.S[ S.O 1.34: 0.1011.34. 69 .. 56: 1.14.: 0.20 6.001 0.10: 0.00: 0.0·011 o.ooo, ~-517.00: 0.42: 171193 515.0 555.0 0 04 0.99SO I 20.0 J S.l 1.41 , 0.00 1.41 S6.03 1.41 0.00 0.00 0.00 0.07 3.0S 0.070 j ... 2! 3.00 -3.27

050494172o.o 795.0!1

-40_ o.ooi_0.99R2 22-', s.411.691, 0.011 1.6919;-'"II.ssj o.I4, 4_.2oj o.o71 o.ool .. o.o_o. o.ooo.l 2.91

17.0 3.~sj 131293. Sl5.0 S55.0 I 0.01;0.99.S6I 25_,

1

S.I, 2.05. 0.27

1

. 2.05 9 ... 141

1.51_ 0.54!16.2; 0.27 o_ .00 0.00 0.000 3.5

1

10.3_ -2 . ..4_ 0703941135.0 155.0 110 0.0010.99S4 15.1 S.lt 1.341' 0.00 1.341Sl.761 1.341. 0.00 0.001 0.00 0.00 0.001 0.000 1.3 1.95.-4.41 050494 • 1~o.oi 195.01 , o.OI0.9966j: 15.2, 9.0

1 1.51 o.oo. 1.51_ 9~.14 1.51 o.oo; o.oo o.ool

1. o.o5j 2.20 o.oso·l 1.6 uo -7.~7~

221193

1

2 ... 5.0

1

260.0 I 0.01 0.99S3 15.6

1

S.4 1.60 0.12

1

,_60 s ... 9S 1.36. 0.24

1

7.20. 0.12 0.00 0.00 0.000 1.9

1

2.SO -7 ... S

050S93 ·, 260.o: 3Io.o· o.12\·. 1.002s: I4.S1

s.211.54l·,~ o.071. 1.541

1

S5.42

1

I.4o 1 o.I4, 4.2ol' o.071

o.oo: o.ooj' o.ooo: 1.6, 5.5o, o.59' 100294 320.0. 410.01 -70 0.01 o.99S4Il6.2 s.6 1.37 0.10 1.37 71.39 1.17l 0.20 6.00 o.IOI o.ool o.oo o.ooo 3.01 9.30[ 0.09[ I~I093I. 43o.o

1lsoo.o, -7oj' o.oi,0.99S7 17.31 s.6: 1.46 o.o6!' 1.46: SI.76, 1.34 o.I213.6o; o.o6. o.oo_ o.oo 1 o.oool 2.9: -5.4S

0 .. 0594 '430.0 500.0, -35 0.001 : 15.21 S.31 1.49[ 0.00 1.49190.92 1.49, o.oo, 0.001 o.oo. 0.00_ 0.00, o.ooo. 16.10. -2.63_ 120194 215.0, 255.0 1

1

, 0.411_o.9.9S5I 9.0

1 S.31 1.34 1_ o.oo. 1.341. 81.76

1 1.34l 0.00; o.oo, o.oo, 0.01l· 0.57[ 0.013 1.6[ 3.IOI-4.10i

221193 1 30o.o j350.o. o 0.01 :0.99S2 I7.9l S.71 uo, o.o9[ uo_ 6S.34, 1.12 o.1s1 5.40 o.o91 o.oo o.ooj o.oool 6.91

-3.231 060993

1

4I5.o_455.ol :O.o41o.99S2 13.6

1 s.3[1.461 o.oo

1 I.46[8S.72 1.45 o.oi[o.Isj o.oo o.oo o.~o 1 o.ooo

1

3.51

5.90_ 0.63,

0~04941455.01.500.0_ 301.0.00:0.99771. 14.S, s.o, 1.32, o.oo,. 1.32_ S0.541 1.321 0.0010.00. 0.00._ 0.051 2 .... 01 o.oso, 2.2! 4.SOI-1.511 2 .. 0S94. 735.0, S53.0. -75 0.99S6 30.9[ s.sl_ 1.30 0.15[_ 1.30 61.02[ 1.00: 0.301. 9.00, .0.151 0 .. 00 0.00 0.0001 1.3;. 9.70 0.4S 050494 160.0 i 190.01-102, 0.0010.99S2, 16.1: S.3: 1.261 0.05 1.26170.7SI. 1.161. 0.10: 3.001 0.051 0.00 0.001 0.000 ~-'1 ') ! -3.001 osos93 I9o.o 125o.o 1_ o.Iso.99911 13.4

1 s.3

1

uo o.I4

1

. 1.3o. 62.24··. 1.02 o.2sls.4o o.I4I o.oo[ o.oo

1

. o.ooo, ... s .... 9o

1

-2.34l 020694 190.0 ' 250.0 I 0.221 • 14.9. S.S 1.141 0.00 1.14169.56' 1.14: 0.00 0.00 I 0.00. 0.00 0.00 0.000 I I !.SO 6.95 osii93 45o.ol5oo.o [0.330.99SOII5.5. S.7 1.2~: 0.11 1.21 60.4II0.99oj 0.22.6.60 0.11• o.oo, o.oo: o.ooo: 3.s,s.2o· 7.04, 020594 135.0 IS5.o 0.01 ,0.99S3_ 13.61 SAl u ..

1

o.osl I.32j74.44, 1.22. o.Io[3.oo: o.o51 o.ooj o.ooj o.ooo, 3.9. 6.501 o.I4[ 020694 I90.o [255.0 -2oj o.oo[o.99S41 10.1; s.o, 1.10 o.oo, 1.10 67.12 1.10l o.oo: o.ooJ o.oo. o.o2: 0.75 1 o.ol7j 2.sl 5.90, -5.7S

1

100294 340.0 410.0 -105, 0.0010.99S3 17.6[ S.3[ 1.39. 0.06( 1.39, 77.491 1.27i 0.12

13.60

1 0.06

1

1 0.0011 0.001 0.000. 4.4 12.6:-1.73

050494 4IO.o[460.o 1 o.oo1o.9973j 15.3 s.2 u31 o.oo 1.331 SI.I5 1.331 o.oo o.ool o.oo, o.o5 2.20 o.o50[ u 5.sol 2.951

o5os93 46o.o soo.o . o.o9 o.99S6 13.sf s.2 1 u7, o.oo I t.I7, 7 u91 !.17[ o.oo I o.oo: o.oo. o.o2

1

. o.66l o.o 15 2.s 15.3 4.79' 150694 460.0 I 500.0 I 0.00. I 16.5 S.2 1 1.30 i 0.00 1.30 I 79.32 1 1.30 0.00. 0.00 I 0.00 I 0.02 O.SS 0.0201 5.601-2.43!

0.00150.01[ 12.01 0.200 0.00. 50.03112.0 0.200 o.oo, 110.3 9.0: 0.150 o.ool: 122.4, 19.5 · 0.325 0.00 142.7' 32.0: 0.533

0.00·. 113 .. 711

25.51' 0.424 0.00. 120.3 26.5 0.441 o.oo· 150.3 26.9 0.44S

o.6oi_'73.S1

_s.o'l; o .. _I-33 0.20 179.31 S.O 0.133 1.00 ~ 193.5 11.7 0.195 I.oo·

1

.I95.7_ II.s 1 o.I96 1.60 178.9: 12.2' 0.203

0.32 .. ···i.4S.3119.01 0.316 0.20: IS2.1 13.5 0.225 0.62 r 1so.21 14.3 o.23S

0.401. 130.7115.01 0.250 0.60. 13S.2 13.6 0.226 0.2011137.1 •. 12.6:0.210

0.40 135.41 6.010.100 1.16. 126.4 6.0 0.100 1.20: 135.5 I 0.1 ! 0.16S

0.401143.71: 12.51' 0.20S 0.20 13l.S 11.7 0.195 1.20 14S.O 12.5 0.20S 0.40·

1

. 147.3 .

1

• 14.o'l 0.233 0.20 130.6 14.9 0.24S 0.66 199.9 4.6, 0.077

0.601133.61. 12.2\0.203 1.40 125.3 6.7 0.112

0.5S11I31.51 12.1:0 .. 201

s.7s I3s.o'I'I.9[o.I9S 0.20 126.7 12.5. 0.20S 0.20 101.7' 4.21' 0.070 1.14 13S.6 5.5 0.092 1.00 137.S 11.3' O.ISS 0.60 127.9 6.210.103 0.60 133.9 '

1/5

.......... (J.)

V.J

> '"0 '"0

(1J

::s 0.. ;;:r 0\ ~

.......... ....._ Vl '-'

No. Type Sample

I I spring I Lemetyinen 2 . spring . Vuorimaki 3 , shallo~· Kl-KA I

411 shallo~·~~KI-KAI 5 shallo~1 KI-KA2 6 , shallow KI-KA2 7

1:shallo~1Kt-KA2

8 shallo~l Liimatainen 9 i FPI , KI-KRlff7* 10 FPI i'KI-KRlff7* IIIFPI KI-KRlfr6 12

1

FPI .KI-KRlff6

13i'FPI IKI-KRlff5 14 FPI KI-KRlff4

15

1

FPI liKJ-KRlff.3 16 FPI KI-KRlff2

17 1 FPI I Kl-KR4ff6 18\FPI ,KI-KR4ff5 191FPI :KI-KR4ff4

2011

FPI IKI-KR4ft3 21 FPI KI-KR4ff2 22 FPI ! KI-KR4ffl n;FPI jKI-KR4ffl 241FPI (KI-KR5ff6 25 FPI 'IKI~KR5ff4 26 FPI KI-KR5ff2 271FPI tKI-KR5ffl 28tFPI i'KI-KR5/BT 29·FP2 KI-KR2ff5

30! FP2 11

KI-KR2ff4 3 I .

1

. FP2 KI-KR2ff4 32 FP2 KI-KR2ffl 33. FP3 I KI-KR3ff6 341 FP3 .. KI-KR3ff5 35, FP3 I KI-KR3ff3 361FP3 _KI~KR3ff2 37.FP3 KI-KR3ffl 38(FP3 1Kt-KR3ffl

Hydrogeochemical data of Kivetty

Ba Fe'"'l Ft;"1 11

Fe2+ Fe

2+1 Na Nai K K[ Ca Cal Mgl Mg Mni Mnl

1

AI All Rb Rbi Sr Sri Lii Li Ba mg/1 mmol/1 mg/1; mmol/1, mg/1. mmol/1[ mg/1 mmol/11 mgll] mmol/1. mg/1 mmol/11 mg/1' mmol/1, mg/1[ mmol/1. mg/1 mrnol/1 mg/1 rnrnol/1 rng/1, mmol/1• rngll] mmol/1

0.010 I 1.79E-4 2.40. 0.104 0.731' 1.87E-2; 4.51 0.112. 0.85 0.035 0.01 i 1.82E-4 0.010. 3.7fE-4 -b.05115.82E-4! . 0.010 1.79E-4

1

2 5010.109, 0.86 2.20E-2 1 3.2 0.080, 0.76t 0.031 0.01 1.82E-4 0.010 3.71E-4 i , i

0 .. 04. 0'. 7.16E-41 13.62 0.15712.54 :_'_ 6.50E-21._ 11.6. 0.289 3.9R 0.164 0.4518.19-·E·. -3~ 0.030 '. 1.11 E-3 0.050: 5.85E-4[ 0.050 ' .. 5.71 E-4 0.0314.32E-3' .0 .. 05 3.64E-4 0.120,2.15E-3 363 0.158 2.54 6.50E-2 12.5. 0.312 4.73 0.195 0.44f8.01E-3 0.010 3.71E-4 0.050

1

5.85E-410.050 5.71E-4 0.03 4.32E-3 0.05 3.64E-4

0.9401. 1.68E-~ '.. . I 6.231 0.271 1.50 '1, 3 ... 84E-2 t 14.1 '• 0.352. 3.23 ' .. 0.133 0.10; ~ .... 82E-3 '. 0.130 '4.82E-.3. 0.050' 5.85E-4, 0.050 i 5 ... 7l.E-. 4. 0.03. : 4.32E-.. 3: 0.05 3.64E-4 0.720 1.29E-... 5.96 0259 1.29 3.30E-2, 9.2 0.228 3.08, 0.127 0.22.4.01E-3, 0.090 3.34E-3 , , 0.07 4.88E-4

0.2401

4.3-0E-31 : 15971

0.260ii 1.44 3.68E-2 11.3._0.282 3.091

0.1271

. 0.26 4.73E-3····0· .047 1.74E-3:_0.0/0. 1.17E-4; 0.039 4.45E-4_: 0.02;2.31E-31

0.110 1.97E-3 5.90 0.257 2.50. 6.39E-2 12.0 I 0.299 7.40 0.304 1.00 1.82E-2 0.010 3.71 E-4' I 0.085 9.70E-4

0.19011

3.40·E·-3: 0.09. 0 1 .. 1.61 E-3 9.30 I 0.4.05 j 2.3516 .. 0. IE-2122.5·_1' _0.5.61.1_·_5.501 __ 0_._·.2 .. 261 0.19i·3· .46_E~31 0.·1·9·0····f7.04E-... 3_.1·. 0.0}0: .2 .. 34E-4 •. 0.180 f 2.0 .. 5. E~3' 0 .. 0_ .. 21_2.88E-3.1 0.06j14 .... 1.7E-4 0.080 1.43E-3

1

. 0.080. 1.43E-319.50 0.413. 2.44 6.24E-2I' 22.2 0.554 5.40 0.222 0.1813.28E-3 0.150~5.56E-3 0.020 2.34E-4 0.140 1.60E-3 0.02 2.88E-3 0.17 1.24E-3

0.099. 1.77E-3 ... 0 .. 104 i 1.8·6· E. -.3 13.0! .0.5651' 1.50! 3.84E-2 ~ 9.1 I 0.477: 4 .. 90. i Q.202.. 0 .. 21 ' 3.82E-3_: .0 .. 04711 ... 7 .. 4 .. E-\. 0. 0. 05 5.85E-5. 0.093. 1.06E-.3, 0.01 : 1..44E-3: 0.72 5.24E.-3 0.120

1

2.15E-3 0.07011.25E-3

1

14.0

1

0.609 1.5013.84E-2 ...4.0, 0.599

1

4.6010.189. 10.022 8.15E-4 ,0.093 1.06E-3, 0.15 1.09E-3 0.027 4.83E-4i 0.022,3.94E-4 40.0 1.74

1 0.89 2.28E-21 10.0 0:250 1.60 0.066 -0.04

17.28E-4 0.034

1I.26E-3 1 0.002: 2.34E-5 0.068, 7.76E-4 O.OI 2.02E-3 o./5 1.09E-3

0.022 i 3.94E-41

0.076!. 1.36E-3, 18.01 0.783

1

_0.82. 2 ... l.OE .. -.2~127~-- 0.317. ·. ··0 .. 6.5: 0.02_71 0 .. Oll·l·-··8·· .2E-4 j 0.03···8ll . .41E-.3.1.0 .. 0.05

1

5.85E-5 •. 0 .. 097

1

1. .. l .. IE. -310.0 .. 1. 1.44 __ E .. -3. ···0 ... 5 .. 5 14 .. 0 .. 1 E-.3

0.08611.54E-31

0.030 5.37E-4

1

33.0, 1.44 0.69

1

1.76E-2 15.01

0.374; 1.901 0.0781

0.06 1.09E-31 0.029 1.07E-3 0.002 2.34E-5: 0.099 1.12E-3,0.~/,1.44E-3.0./5

11.09E-3

0.100 1.79E-3. 0.115 2.06E-3 27.0 1.17. 1.40 3.58E-2. 15.0 0.374 2.80 0.115 0.21 3.82E-3 0.062 2.30E-3 0.005 5.85E-5, 0.112 1.28E-3 0.01 1.44E-3 0.36 2.62E-3

0.308: 5.52-E.-31 0.203

1

J.6JE-J' 7.501 0.32_6

1

; 1.201

3.07E-21·. 13.0. 0.324 4.40' 0.1811 0.48

1

· 8.74E-3 ·. 0.010:3.71 E-4~. 0.001 i 1.17E-51 0.048 i 5.48E-41 0 .. 02j2.16E-31 0./5 l.09E-3 0.32015.73E-3 0.080 1.43E-3 7.7010.335 1.40. 3.58E-2 12.0 0.299. 4.301 0.177 0.46 8.37E-3. 0.01415.19E-4 0.002

1

2.34E-5 0.04915.59E-4 0.01 1.44E-3 0.15 1.09E-3

0.058. 1.04E-3; 0.057. 1.02E. -31. 7.50 ·. 0.326: 1.10 I_ 2.81 E. -2: 14.311

0 .. 357 .14.70 '. 0.193 · .. 0 .. 391 7 .I OE-31 0.018 6.67E-4! 0.005 . 5.85E-5: 0.0?6. 6.39E-4: 0.0 I ' I ... 44. E-3 ;.0 ... 3.9 t 2 .. 84E-3 0.090il.61E-31 0.069tl.24E-3: 8.60:0.374

1

1.30,3.32E-2119.0 0.474 4.10, 0.16910.32 5.83E-3 0.027 l.OOE-3.0.008 9.36E-510.044,5.02E-410.0/11.44E-310.20 1.46E-3 o.097 1.74E-3 o.oo5j8.95E-5, 21.0. o.9t3 1.00, 2.56E-2 to.o o.250 2.00 o.082 o.o5 8.I9E-4 o.m 1 , t.t5E-3 0.002 2.34E-5 o.o7o, 7.99E-4 o.o1 t.44E-3 o./5 t.09E-3

o.m2[ __ 5.73E-4; o.ot9 1 3 .. 40E-4Jt7.o j o.739, uo j3.32E-2 tOA; 0.259: 2.601 0.101: o. JOIJ.R2E-JI o.o5o -~t.-85E-3. o.oo5j5.85E-5 '. o.o5716.5I E-4, o.o1, t.44E=} o_ .161

t.t7E-3 0.350; 6.27E-31 0.260 i 4.66E-3 I 9.60 I 0.418 I 0.99! 2.53E-2' 16.0 0.399 4.40 0.18lj 0.27; 4.92E-3 0.083 3.08E-3; . 0.044 5.02E-4: ; : 0.15 ' 1.09E-3

O.p01

4.83E-3! .0.0951._1-7. O.E-3!'_ 8 .. 20

1

, __ 0.1571 1. .. 7 .. 0

1

4.35E-.2.1. 13.5·1. 0.3371. 4.20 0.1731. 0.·2·4['_ 4.37E. -.3.1 0.017.: 6.30E-4~t-0.0. 05 )_ 5 .. 85E-51 0.0 •. 4.9, 5.59 .. E---41_0.0/11.44E-3 ..

1

.. 0 ... 5. 54 .. 0 .. 1 E-.3 0.015 2.69E-4 0.037. 6.63E-4 20.0 0.870 2.00 5.12E-2 13.8 0.344 1.60 0.066' 0.10 1.77E-3 I 0.0 ))i4.08E-4 0.005 5.85E-5 0.11011.25E-3 0.0 I 1.44E-3 0.44 3.20E-3 0.299,_5.35E-3 0.026.4.66E-4t 12.0 0.522. 1.40 3.58E~2 1 17.010.424

1

3.70 0.15:21 0.25 4.55E-3 0.010 3.71E-4,0.005 5.85E-5 0.031 3.48E-4 0.01 .1.44E-3 0.20 1.46E-3

0.670 'it.20E-21_0.300 I_ 5.37E-31. 12.0 1·_, 0.52~ I_ 1.201. 3._0IE-21' I 0.01 0.25011

.

1 3.50.1 0.1441 0 ... 3. 6·1:. 6 .. 55E-J I 0.0/0 I' .. 7.1 E-·41: 0 .. 0011_ 1..1. 7E-5 '_, 0.02. 7. 3.0.2 .. E .. -41·0.0·/·1'.· 1.44E-3.; .. 0 .. 1.·5. 1.09. E. -.3

0.0366.45E-4, 0.017t3.04E-4 35.0 1.5 .... 3.00 7.67E-2 22.0, 0.549 2.10 0.086, 0.03 5.46E-4, 0.670.2.48E-2 0.002[ 2.34E-51 0.135 1.54E-3 0.01 1.44E-3[ 0.50. 3.64E-3 0.033

1

5.91E-4j 0.03015.37E-4. 17.00.739!' 0.72 1 1.84E-2. 13.0

1

0.324, 1.4010.058

1

1 0.021

3.46E-41 O.Ot816.67E-4 0.002

1

. 2.34E-5 0.088\l.OOE-3 0.01 1.44E-3f0/5 1.09E-3

0.008 1.43E-4: 0.005 i 8.95E-5120.0 1. 0.870. 0.411

1 1.05E-2113.0 0.324. 0.79 0.033:. 0.01

1

. 1.46E-4! 0.0281t.04E-3. 0.00. 4 4.68E-5: 0.079. 9.02E-4, 0.0/l' 1.44E-3) 0.2···0· _

1

1.46E-3 0.003: 5.37E-5, 0.010 i 1.79E-4. 22.0. 0.957 I 0.40' 1.02E-2: 13.8: 0.3441 0.69 0.028: 0.0 I 1.46E-4. 0.033' 1.22E-31 • 0.093. 1.06E-3 0.20 1.46E-3

0.11211

2.0 I E-31 0.00518.9. 5E-51

10.0 t' 0.435] 1.20, 3.07E-2121.011

0.524 1 5.50.0.2261_0.30 1 5.46E-3l. 0.02218.15E-4: 0.00515.85E-51 0.09011.03E-31 0.02

1 2.45E-3 • 0.44; 3.20E-3

0.025 4.48E-4 0.010. 1.79E-4: 5.40 0.235 1.10 )2.81 E-2 19.0 0.4741 3.40. 0.140: 0.0213.82E-4 ~ 0.007; 2.59E-410.002 I 2.34E-5j 0.059. 6.73E-41 0.01 1.44E-3 0.15 . 1.09E-3

0.01_5_

1

2.69E-41_· 0.220

1

3.94E-3 :. 3.50

1

0.1521. 1.501

3.84E-2, 13.0

1

0.324 3.40

1

0.140

1

. 0.05. 8.37E. -4~· 0.036

1

1.33E-3 '.0 .. 002 t 2.34E-51_ 0.025l'. 2.85E-4.0.0/ 1.4·4· E-3~ 0.20 1 ... 1..46E-3 0.017 3.04E-4 0.011 1.97E-4jt8.0 0.783 0.55

1

1.41E-2116.0 0.39911.70 0.070 0.03 5.64E-4 0.013 4.82E-4 0.003

1

3.51E-5;0.098 l.IIE-3;0.0/ l~E-3 0.15 1.09E-3 0.040 7.16E-4f 0.050 8.95E-4 10.40.452

1

1.203.07E-2 19.0, 0.474 2.70, 0.111: 0.0518.74E-4. 0.026;9.64E-4)0.003: 3.51E-5

1

0.067 7.65E-410.0/ 1.44E-3, 0./5.l.09E-3 0.05419.67E-4l 0.079ft.41 E-31 16.0 I 0.696 0.8712.23E-2116.0 I 0.~991 2.00 I 0.0821· 0.05 ~ 9.10E-41 0.038f1.41 E-3 0.008 i 9.36E-5 0.05316.05E-4 0.01 1.44E-31 0.2011.46E-3 0.028 5.01E-4 0.006 1.07E-4 18.0 0.7831 0.80. 2.05E-2 15.0. 0.374 1.70! 0.070. 0.03f4.55E-4 0.054]2.00E-3: 1 · , 0.20, 1.46E-3

2/5

......... (.;.) ..j:::..

~ '0 '0

(!)

::s 0... ;;t 0\ ,-_ N -Vl .._,

No! Type I Sample

I spring Lemetyinen 2 [spring Yuorimaki 3[shallow[KI-KAI 4 shallow. KI-KA I 51' shallow. KI-KA2 6, shallow] KI-KA2 7 ·shallow, KI-KA2 81 shallow Liimatainen 9 IFPI 'IKI-KRiff7* IOt'FPI KI-KRiff7* 11 FPI IKI-KRiff6 12 FPI KI-KRiff6 13fFPI ,KI-KRiff5 14 FPI I KI-KRiff4 15

1

FPI jKI-KRiff3 16 FPI KI-KRiff2 Ii FP I I KI-KR4ff6 181FPI ,KI-KR4ff5 19'FPI IKI-KR4ff4 201FPI .KI-KR4ff3 21 FPI 1KI-KR4ff2 22 FPI KI-KR4ffl

231FPI lKI-KR4ff-l 24 FPI KI-KR5ff6 251'FPI ·KI-KR5ff4 26 FPI IKI-KR5ff2 21:FPI ,KJ-KR5ffl 281FPI ~·KI-KR5/BT 29 FP2 KI-KR2ff5 30fFP2 :fKI-KR2ff4 31 FP2 KI-KR2ff4

321 FP2 33, FP3 34

1

FP3 35 FP3 36

1

FP3 37 FP3 38jFP3

1

: KI-KR2ffl KI-KR3ff6

IKI-KR3ff5 , KI-KR3ff3 KI-KR3ff2 KI-KR3ffl KI-KR3ffl

Cs Cs'

mg/1 mmol/1

0.050. 3.76E-4

0.050 1: 3.76E-41

0.050 3.76E-4

1

,

0.050 3.76E-4 0.010. 7.52E-5 1

0.020 11.50E-41 0.020 · l.SOE-4, o.oo5[3.76E-s;

~ I 0.034! 2.56E-4: 0.00513.76E-5 i

0.001: 7.52E-6f 0.00513.76E-5. 0.015 1.13E-4

1

:

0.0011 7.52E-6

0.00513.76~-5' 0.006: 4.51 E-5~

0.0 171' 1.28E-4 ', 0.005 3.76E-5

. I 0.00513.76E-51 0.0 17. 1.28E-4 1 o.oo5: 3.76E-sl o.ooii7.52E-6: 0.002: I.SOE-51 0.00211.50E-5: 0.003

1 2.26E-5

1

o.oo513.76E-s: o.oo2 , 1.soE-s1 0.002jt.50ES 0.016 1.20E-4f 0.001 I7.52E-6 0.009: 6.77E-5f

I

Hydrogeochemical data of Kivetty

I :'- :'- I ' I B I B s,ll,: sill' I s s so~ 1 so~ P,ll, P,ll, PO~· PO~ N,ll,. N,ll, mg/1! mmol/li mgll]mmol/1· mg/1 mnml/1 rng/1~ mrnol/1 mg/11 mmol/1 mg/11 mrnol/1 rng/1· mmol/1,

2.2~ o.o69 ---, 6.2o 6.45E~~ o.o 1l 1.61 E-4~ o-:-o 1s, t.58E-4 o.l8 I o.o 13 1.2 0.03711 I 2.8012.91E-2 0.01 1.61E-4 0.015 1.58E-4 0.621 0.044

0.009

1

8 32E-4 1.80 1.87E-2 0.02 6.46E-4 0.10 l.OSE-3 I I I

00098.32E-4 I

1

1.78 185E-2 1 0.0216.46E-41

0.10 1.05E-3, 0.00918.32E-41 I 2.03 2.11E-21 0.02 6.46E-4 0.10 l.OSE-3

I I 0.68 7.08E-3 0.09 2.91E-3 0.10 1.05E-3 I I

0.005 4 62E-4 I 84 1.92E-2 0.02 6.46E-4 0.10 1.05E-3

1

1

0.30 ~1 0 009:

1

, ~1 1.00 ll.(l4E-, 0.10 3.07E-31 0.29 3.05E-31 0.201

0.0 ~·I 0.020 1.85E-3 0.17 0.005 0.190 5.93E-3 0.11 l.ISE-3 0.0216.46E-4~ 0./011.05E-3 0016 1.48E-3 019 0.006 0.070 2.18E-3 0.05 5 20E-4 0.02 6.46E-4 0.10 l.OSE-3

0.04814.44E-311.021 0.03210.019,5 93E-4 243]2.53E-21 0.01. 3.23E-41 0.10 I l.OSE-31 0.491 0.035' I I 0.03019.36E-412.80 2.91E-21 I 0.50 5.26E-3 I

0.240 2.22E-2 1 2.70 0.0841' 0.030 9.36E-4! 6.8017.08E-2

1

0.01 3.23E-4 0.10 1.05E-3 0.47! 0.034 0.051 4.72E-3 049 0.015 0.010 3.12E-4 1.15fl.20E-2 0.01 3.23E-4 0.10 1.05E-3 0.27 0.019

0.09118.42E-311.301 0.041 0.07012.18E-312.50 2.60E-2 0.01, 3.23E-41 0.10 l.OSE-312.101 0.150

1

0.069 638E-3 0.52 001610.0/0 3.12E-4 1.50 1.56E-2 00113.23E-4 0.01 1.05E-4 1.00 0.071 0.010

19 25E-4

1

0.47

1

.. 0.015 0.010: 3.12E-4

1

1 5111.57E-21 0.01 3.23E-41

0.10 l.OSE-3

1

0 14 0.010 0.014 1.29E-3 049 0.015 0.010 13.12E-4 1.10 l.ISE-2 0.01 3.23E-41 0.01 l.OSE-4 0.13 0.009 0.013 I 1.20E-31 0.59 0.01810.010' 3.12E-4 2.1012 19E-2 0.01 I 3.23E-4 0.50 5.26E-3 0.131 0.0091 0.033.13.05E-310.601 0.019:0.050

1·1.56E-3

11.48 1.1.54~-2~ 0.0/13.23E-4 0.50 5.26E-3: 0.40:0.029

0.037 3.42E-3 0.65 0.02010.0/0 3.12E-4 1.90 1.98E-2 0.01 I 3.23E-41 0.01 I.OSE-41 0.99 0.071 0.036

13.33E-3

1

! 0.69. 0.022 0.1001 3.12E-3 1.66·

1

. 1.73E-2, 0.101

3.23E-3 0.50, 5.26E-3 4.90

1

1 0.350:

I 0.931 0.029: 0.030 i 9.36E-4 i 2.30 2.39E-21 I ! 0.50 )5.26E-3: I 0.010 !9.25E-4[ 1.1-0/ 0.03_4[.· 0.022f6.86E-4f. 2.40! 2.50E-2 0.01 3.2.-lE-4

1

, 0.10 ·. 1.05-E-3

1

. 0.14 O.OIOf 0.120

1 t.IIE-212.2o~ o.069 o.ow I 3.12E-4 s.6o s.83E-2: o.o4

1 t.29E-3 1.10 t.I6E-2 1.001 o.soo;

O.o3~~.2.77E-3 1 0.561 0.017:0.0101. 3.12E-4

1

1.s8ji.64E-2

1

0.01

1

3.23E-4; 0.50 15.26E-3: 0.551 ... 0.0391 0.03~12.96E-31 0.48: 0.0151,0.010 3.12E-4' 1.20 1.25E-2. 0.01 I 3.23E-4., 0.00 O.OOE+O. 0.421 0.030 0.061

1

5.64E-3 2.50

1

0.078i0./00. 3.12E-3 6.90 7.18~-2 0.01. 3.23E-4r 1

0.29

1

0.021, 0.097!8.97E-3. 1.42 0.044

1

0.010 13.12E-414.00 4.16E-21 0.0511.68E-3. 0.01 I.OSE-4, 5.40 0.3861 0.087

18.05E-311.10. 0.034.0.050: 1.56E-3

1 2.70 2.81E--2' 0.01. 3.23E-4

1

0.50 5.26E-3

1

0.13. 0.009: 0.83 I 0.026. 0.050 11.56E-312.50 2.60E-2 0.50 5.26E-3 I '

0.061 : 5.64E-311.61: 0.050 I 0.050 1.56E-3 4.50 4.68E-2 11 0.01 13.23E-4: 0.01 l.OSE-4 0.68 0.0491.

0.01019.25E-4\ 1.43

1

0.045:0.0/0,3.12E-4 3.90 4.06E-2, 0.01 3.23E-41 0.501

5.26E-31

0.43 0.0311

0.010 9.25E-4

1

1.97-1

0.061_1.0.050ft.56E-3 5.10 5.31~-2~ 0.0/

1

3.23E-4. 0.5015.26E-3I 0.8110.058 0.073 6.75E-3

1

1.441

0.045:0.0/0,3.12E-4 4.30 4.48E-2 0.01 3.23E-41

0.01 l.OSE-4 1.101

0.0791 0.030 2.77E-3·

11.53. 0.048

1

q.OJO 13.12E-4 3.90 4.06E-2 0.01. 3.55E-4, 0.01 11.05E-4 0.66. 0.047[ 0.045 4.16E-3 1.65

1 0.051_0.050 1.56E-3 4.65 4.84E-21 0.0/13.23E-4f 0.50, 5.26E-3 2.5010.178

1

1

I I 0.060 I 1.87E-3 5.50 5.73E-2 ' 0.50 I 5.26E-3

~ I I I I NH~; NH~ NOo NOo NO~~ NOl Cll Cl mg/1 mmol/1 mg;ll mmol;l! mg/1 nunol/11 mg/1 mmol/1

0.021 9.98E-41. 0.01 1·. 2.17E-4. 0.51[8.23E-31. 0.6410.018 0.01 5.54E-4 0.01 2.17E-4 2.70 1 4.35E-2' 1.20! 0.034

I . . I

0.05 2.77E-3 ·. 0.10: 2.17E-3. 0.10: 1.61 E-3: 2.22, 0.063 0.02 l.IIE-3: O.I0,2.17E-3 0.10 1.61E-3 2.24

10.063

0.07 3.88E-3' 0.10 I' 2.17E-3 0.22 3.55E-3 2.20 0.062 0.03· 1.66E-3: 0.10 2.17E-3i 0.10 1.61E-3 1.77 0.050 0.06i J.33E-31 0.10 2.17E-3: 0.10

1 1.61E-3 1.77 0.050-

0.0814.21E-3 0.01 2.17E-41 0./0:1.61E-3 0.97 0.027

0.03. 1.66E-3_! 0.10 ·1:· 2.17E-3 0.10 I' 1.61 E-3: 1.48 i .0 .. 042 0.03 1.66E-3 0.10 2.17E-3 0.10 1.61E-3 1.46 0.041

0.03 l 1.6-6E-3. 0.02. 4.35E-4: 0.02. 3.23E-4. 3.381 0.095 0.0311.66E-3 0.02, 4.35E-4

1

0.06,9.68E-4 3.50 0.099

0.08 4.44E-3! 0.03 ~.:-6.52E-4·. 0.03

1,4.84E-4~ 22.2! 0._6_ 26

0.03 1.66E-3 0.01

1

2.17E-4 0.01 1.61E-4, 2.05

1

0.058

O.os: 2.77E-3. 0.031

6.52E-4; 0.02,3.23E-4_f_9.70 0.2.74 0.03 1.66E-3 0.01. 2.17E-4 0.01 1.61E-4 4.50 0.127

0.071, 3.88E-3 I 0.031' 6.52E-41 0.031' 4.84E-4; 1.38: 0.039 0.05 2.77E-31· 0.05 1.09E-3 I 0.02 3.23E-4: 1.00 I 0.028

0.03 .. 1.66E-3. 0.20 14.35E-3· .. 0.20 !3.23E-3j' 1.101. 0.031 0.03 1.83E-3

1 0.01, 2.17E-41 0.01 .1.61E-4 1.14

10.032

0.08 4.44E-3: 0.01 12.17E-4 0.03 .4.84E-4! 2.10. 0.059 0.04. 2.22E-3 0.01 2.17E-4 0.01 1.61E-4

1

1.68 1 0.047

o __ .o311.66E-31 _o.o2 1 4.35~--4- i o.1s, 2.42E-3. 1.40 [ o.039

0.03, 1.66E-3! 0.0214.35- E-4_1_0 ... 0. 213.23-E-4·, 3.53-[ 0.10--0 0.03. 1.66E-3, O.I0[2.17E-3 0.10 1.61E-3. 13.8 0.389

0.03 ·~· 1.~6E-3 •. 0.01 )2.17E-4: 0.01 .1.61E-4. 4.3010.121 0.04 2.~2E-3 0.04 8.69E-4 0.02. 3.23E-4 2.10 0.059

0.05 ' 2.77E-3' O.Q214.35E-41 0.01 I' 1.61 E-4 48.0: 1.354

0.0814.44E-31 0.08: 1.74E-3 ·. 0.14 I 2.26E-31 7.80 I 0.220 0.03 1.83E-), 0.01

1

2.17E-4. 0.01 .1.61E-415.00. 0.141 I 0. 05 1.09E-3 0.0213.23E-4 4.60 0.130

0.0311.66E-31 0.01 I 2.17E-4 0.01 1.61 E-4' 11.0' 0.31 I 0.03: 1.66E-3; 0.0214.35E-41 0.04 6.45E-41 0.881 0.025 0.03

1

. 1.66E-31 0.05 1 1.09E-3- I 0.0314.84E-41. 0.84 10.024 0.06 3.33E-3 0.01 I 2.17E-4j 0.06 9.68E-4 6.20, 0.175 0.05' 2.77E-3

1

0.0511.09E-\ 0.03:4.84E-411.901 0.054 0.0311.88E-31 0.01 i 2.17:-4; 0.01 11.61 E-4 I 4.82' 0.136

· 0.0 I i 2.17E-41 0.08 1.29E-3 7.30 0.206

375

........ (.;.)

Vl

>-­"0 '"d (D

~ 0.. ~·

0\ ,..-...., (.;.) -Vl '-'

Noj Type I Sample

I spring -Lemetyinen

2

1

spring ~--Yuorimaki 3 shallow KI-KA I 4 shallow KI-KA I 5 ·shallow' Kl-KA2

611

shallow[ KI-KA2 7 ~shallow Kl-KA2

8 i shall owl Liimatainen 91FPI KI-KRI!f7* 10 FP! Kl-KRI!f7* 11 FPI < KI-KRI!f6

121FPI IKI-KRI!f6 13 FPI ,KJ-KRI!f5

14\FPI ~·Kl-KRI!f4 15 FP! Kl-KRI!f3 16'FPI KI-KRI!f2

171FPI i'Kt-KR4ff6 18!FPI KI-KR4ff5

19/FPI KI-KR4ff4 20 FPI I Kl-KR4ff3

211FPI .Kl-KR4ff2 22FPI ,KI-KR4ffl

23jFPI 1Kt-KR4ffl 241 FPI I KI-KR5ff6 25;FPI ·

1

iKI-KR5ff4 26

1

1 FPI KI-KR5ff2 27 FPI KI-KR5ffl

281FPI I Kl-KR5/BT 29 FP2 I KI-KR2ff5 30~FP2 KI-KR2ff4 3IIFP2 KI-KR2ff4 32

1

FP2 KI-KR2ffl 33[ FP3 Kl-KR3ff6 34 FP3 KI-KR3ff5

351FP3 KI-KR3ff3 36 FP3 I Kl-KR3ff2 371FP3 'KI-KR3ffl 38 FP3 I KI-KR3ff I

F 1 F Br 1 Br I 11 1H1 rso 2H ! . I ! ' I

mg:/1 mmol!ll mg/1 rnrnolll1 mg:/1 rnmol/1 TU'/{,SMOWio/c,SMOWI

0./0

1

•0.005j0.0111.25E-40.01 7.9E-5 10.8, -12.81 -95.8 I ' I I

0.10 0.005• 0.01 1 L25E-4 0.01 7.9E-5113.01 -12.5 -94.6

0.3510.018~ 0.0.5\6.26E-41'0.0/ 7.9E-5. 31.91

. -12.4: -94.9

0.35,0.0181_ 0.0516.26E-4,0.0/ 7.9E-\ 27.1. -12.41 -94.0 2.0710.109 0.05 6.26E-4 0.01 7.9E-5

1

. 5.5

1

. -12.8 -94.1

1.8.00.095 o.o5,6.26E-4j'o.oJ 7.9E-5 6.0 -12.6, -96.4. 1.95 0.103

1

! 0.0516.26E-4 0.01 7.9E-5, 6.0 -12.7' -<J7.91 0.35 0.018 0.01 1.38E-4; 0.01 7.9E-5: 0.8: -12.5. -95.3

2.491 0.13L 0.0516.26E.-4I OAI, 3.2E-31 6.1 -12.51 -9. 6.0,

2.32~0.1221 0.0516.26E-4 0.4113.2E-3 6.1 -12.41

-92.7~

2..491_ 0.13·1'. 0.07.8.7.6E-4 0.04. 3.4E-4 0.8. -12._7, -92.41 2.4010.126 0.03 3.75E-41 I 0.8 -12.71 -95.3 5 .. 20

1~ 0.274.

1

0.16

1

:2.00E-3. 0.21 .• 1.7E.-3· 0.8 -13.71

-100.0:

2.21, 0. I. 1.6. 0.01 1.25E-4; .0.03 2.7E-4-·, 0.8 -12.6 -90 ... 41 3.50 0.184: 0.07 8.76E-41 0.02 1.5E-4l 0.8 -13.21 -94.9

2.10.10.111l0.0/11.25E-4_ 0.0312 .. ·3·E ... -4' 0.8 -12.4 -9-3.2' 2.30 0.121 0.07 8.76E-4 0.03 2.4E-4

1

0.8 -12.8. -92.41

o.4o1

o.o21i o.o/j'l.25E-4[ o.o.4.J2.9.E-4·. o.8·j -12_.8 .. 1 -92.2, 0.4010.0211 0.07 8.76E-4 0.01 7.9E-5. 1.4 -12.9 -95.0 1.46 0.077 0.03 3.75E-41

1 0.03'2.4E-4·~~ 0.81

-12.8. -95.81

2.8oj:o.l47[' o.oJj.L25E-4 o.o6.[4.7E-4, u[ -13.01 -95.01 2.61 0.137 0.07 8.76E-4 0.10 7.9E-4. 2.8, -12.9 -94.51

2.50;0.132: 0.03i3.75E-41 ' I 4.61 -12.71 -96.51

1.5.510.0821 0.0.718.76E-4 1 0.0119_.5E-5j 2.41 -12.41, -92.2! 0.60 0.032 0.10 1.25E-3j 0.10 7.9E-4 8.1 -14.5 -108.0

1.6. I, 0.085

1

! 0.031:3.75E-4f_ 0.031' 2.4E-41 I; -12.71, -94.41. 0.40 I 0.021 0.0 I 1.25E-4 0.02 1.7E-4 -12.6 -90.6

2.401

0 .. 126.: 0.44'5.51E-31·, 0 .. 02

1

1.6.E-4! 1.7 -13.9: -1-01.5·1~ 1.2ojo.o63j omj8.76E-4, o.o8 6.1E-4f o.8j -13.6! -99.3_ 3.17 0.167, 0.03 3.75E~4~ 0.12 9.5E-4, 0.8: -13.3 -100.2.

3.0010.1581 0.0415.01E-4, I I 0.81 -13.3 -102.01 2.100.111

1 0.09

1Ll3E-3

1

-0.70:5.5E-3, 4.0 -14.01 -102.0

1.00 0.0~3~ 0.0313.75E-4! 0.01 17.9E-51 8.5

1

-12.9: -96.41

0.4310.0~3. 0.03i3.7·5·E·-.4~0.0/ 1 7 ... 9.E-5, 24 .. 9 .. -12.81 ~91.9 2.90 0.1531 0.0/

1

1.25E-4 0.0917.1E-4 4.7

1

-13.6 -99.3

1 2.00)0.105, 0.01.1.25E-4. 0.07,5.8E-41 9.6. -13.0: -95.3

1.93 0.1021 o.o3;3.75E-4[ o.03j2.4E-4111.3: -13.7[ -102.21 2.8ol o.I47 o.o6 7.51E-4 1 . , 10.01

Hydrogeochem1cal data of K1vetty

222 Rn/ 1"C/ 11C uCcll4: 11Ccm! rsOcm 14Sso}- qSso4~ rs0so4

I3LJII1 rmd 'Xri'DI3 '.l,rl'lm' o/cri'DI3. o/crl'l>l3 '/!,cnTI '.!,,CnT: o/cri'IH3,

s7 Sr/x6Sri uCcrr4

I '/!ri'DI3

1l 11 s I 2lS ! 2lS Ccm Ocm; Umol U112o

69, 116.21 -23.241

471' 110.21-22.161

270 . !

7701: 1000

280,

45011

650 190

14001' 210

50 1

s7j1

510

3901· 490 520

1201' 17

I

171 35, 17!

3301

1

64

150!: 140 93\

951 370

361' 330

1101 42,

I

49.4: -21.301

74.51·23.36

44.4'

1

-19.601' 'i· :1 30.1 -20.18 . . 29.6, -20.61' -58.001 -9.60,

24.6

1

-20.55. I

37.9 -19.07

25.3' -20.61

34.0)-19:4oj·

~!:~I ~~~:~~I I 33.81-19.52'1 ' 31.7. -18.29 I ~1 21.71-21..56, I'

24.1 -21.671 30.71-2. 1.52 -15.00], -16.60,1'

46 .. 81-21.461 ! 19.8 -22.851 55.2: -23. I 9

62.81-24.22

25.3 -. 15.151' 18.8[' -18.95 9.3 -19.11

8.7,-19.881 10.21-20.74 42.0 ~- -20.58 ~

1 63.1 J -21.81

14.0[-20.7511

26.41-20.80 28.4: -19.76 i

I '

8.10

'/!d'DI31 o/cri'DI3 ml3lJ/I pph

6.701

I

22.60 60i 22.31

I 128.501

. I I ·. . I

1

1.871

i -12.501

I

I

0.729401

I 0.78554

I 0.75187' -3.01

0.730181

I mol I o 72880 j '

123.70

1

, 27.78. 10.501

0.727521 • 1 I

I j i ' I i I 22.701 24.08' 9.961 0.722181 ' -16.41 -6.2

I I I . I • 2000

1

1

1

! 0.73793 I .

I l

1.3/ 0.11 0.5! 0.04

19.2 17.5

2.91

1

1.4 1.1 4.8,

31.6'

23.·51

1

139.0 169.0 51.0: 55.oi

0.11 0.09 0.39

11.30 13.70 4.14

40.01 3.21 60.0 4.86

165.0! 13.40

60.011

4.88 81.7

113.0 31.0 - 2.51

27.9'1 109.0 42.0 13.4.

8.84 3.44

11.81' 15.50 8.1 24.00 8.9 0.72

27.01: 25.0

30.91 30.0

1.0

31.o'[ 43.0

36.0J

2.21

2.44 0.08 2.50 3.52

4/5

.......... w 0\

;:t;> "'d "'d (D

:::s p_. ~·

0\ ,--..,

.,J:::.. -VI '-._/

No: Type I Sample

I spring i Lemetyinen

21spring !Yuon.·maki

3 • shallow~ KI-KA. I 4 shallow KI-KAI 5 shallow KI-KA2

61.shallow1 KI-KA2

7 shallow• KI-KA2

8 , shallowiLiim.ataine.n 9 1 FPI KI-KRiff7* IO:FPI KI-KRiff7*

tt[FPI ·~· KI-KRiff6 12! FPI KI-KRiff6 13):p1 1 KI-KRiff5

141FPI 'KI-KRiff4

15 FPI ·IKI-.KRiff3 16 1 FPI KI-KRiff2

tiiFPI KI-KR4ff6

18 FPI I'Kt-K. R. 4ff5 19

1

FPI KI-KR4ff4 20 FPI KI-KR4ff3 2t~

1,FPt !KI-KR4rr2

22 FPI ] KI-KR4ffl

23lPI IKI-KR4ff.l 24 FPI KI-KR5ff6

25i'FPI 1 KI-KR5ff4

26 FPI 11

KI-KR5ff2 27! FPI KI-KR5ffl 281FPI KI-KR5/BT 291FP2 .KI-KR2ff5

30.FP2 IKI-KR2ff4 3t'I1 FP2 KI-KR2ff4 32 FP2 KI-KR2ff I

33' FP3 I KI-KR3ff6 341' FP3 KI-KR3ff5 35 FP3 KI-KR3ff3

361FP3

11KI-KR3ff2

37[FP3 _ KI-KR3ffl 38FP3 iKI-KR3ffl

2\4 !2\X 12.\X 2.\4 2 \X I U// Ul'aruc·l. Ul'artir: U/. U 23xU 1 m13LJ!l· pph Panic.

1.31 0.20 0.02: 1.26: 0.10 2.40 2.40 2.00

3.30[ 0.7[ 0.10 1.70 1 0.20

0.361 0.029

4.40

1.26 I

Hydrogeochemical data of Kivetty

mThll 2\XU/ 2\.+U: 212Th 2\IITh. 22XTh/2.lii.Th/, 22.xTh/ nxRa.~ nxRal 226Ral22oRa j I · ' '14 ' op · 1

•

pph dpm/kg•dpm/kg1 dpm/kg dpm/kg dpm/kg • U · "Th mi3Lj/kg dpm/kg1 mi3Lj/kg,dpm/kg

N2l 02, co2j H2j1

ul/d ul/1 ul/1 ul/1 -

!:~1: 4.88

4.8 4.44'

0.271 0.02 0.10 0.01

1.521 •• ! I I o.ots 8.38 42.7 o.oo4~ o.otl o.tt~ 0

i 31 too; 6.01[ 2u

i I i 1.28121700 I 820 690 I

3.58 3.99 3.98

3.6511

3.51 3.87 4.66: 3.94i

0.19 0.02i

0.15 ~- I 0.181 0.02!

::11 0.01 li

0.16 0.01

1.491 I

i I

u{ 1 I 1.16~~ I

2.03

1.38 '

9900 I 3600 I 185

I 14000.2000. 245

He I CH4 1 C2H2 1 C2H41 C2Ho ul/1[ ul/1 ul/1 ul/1 ul/1

2t.tlt7.8 t

I I

1.0 5.00, I

I ;

0.0081 0.060

0.0031 0.002

I I I

0.718.10 0.0510.100 0.170

I 4.101 3.01

0.12 0.27 0.14 o.otj

0.021

1.70~ I . 2.541 0.0 I I 6.5 J 19.8 0.003 0.0 I , 0.03 1

1

12.291 100 6f 4.99] 0.31 17500 (2500

I ! 0.9[ 2.20' 0.0061 0.020

3.82) 4.33.

1.981 2.41

4.001. 3.17

2.8011

5.56 2.13, 6.28;

5.761 3.48'

0.241 1.23

0.191 0.17:

0./511 0.12

0.26.1 0.12

0.141

0.15.

I 0.021

I

0.01

0.021 0.01

2.27 I . 1.21 :

: I I I

I , I ,

0.831 1.86 I 0.002. 0.0 I I 0.03 0.007 1.41 I

2.501' I ' : I ' 0.004 1.071 0.00 I I 0: 0.021 0.00 I

~:~~~~ 1

1

I I' I 2.54 . . '

I I I I I

l

18.81

I

201

I I

I 26.78

30 1.8 4.451

I .

11100 jl' 16001

1. 23400 435

li 20300. 3300 •.

I I 0.27 17100

1

1780 I I

:1880019901 111800! 1160

I 15900:2100,

122100 I 470 I

I

3IOI I 54 270.

550

861 I

2001 530,

400i

1301

1.1. 1.70 59.0 13.0

0.5 1.20,

2.0: 1.50.

1.7 0.40

o "1 0.30,

4.2 uo:

15.6115.61

'0.005! 0.013 o.o 1 o.8oo 1~ uo

0.150 0.012

0.008f 0.016

I

:o.03o' o.03o

'i: 6.020 i 0.006

0.020 I 0.060

: o.oto' o.04o

5/E

........... \..;..)

-......J

> '"0 '"0 (!)

~ 0... ~·

0\ ,--, Vl -Vl "-"

138 Appendix 7

Evaluation of the hydrogeochemical data of Kivetty.

Parameter(s) Representative results Notes Amount of remaining Drilling water< 2.5% 4 samples having more than drilling water 2.5% drilling water were

discarded Electrical conductivity Results of laboratory measurements Frequent technical problems in

data storage of field measurements

pH, alkalinity, acidity, Results of field measurements Transportation delay C02(free)

Fetot Results of laboratory AAS analyses Field analysis with FerroSpectral is not perfectly quantitative

2- Results of field measurements, if Transportation delay S tot

larger than the lab value so4 Field results Transportation delay

P04 Field results Transportation delay

NH4 Results of field measurements, if Transportation delay larger than the lab value

N02 Results of field measurements, if Better detection limits than in smaller than the lab value the laboratory IC-analysis

N03 Results of field analyses, if smaller Transportation delay than the lab value

Cl Field values F Results of field measurements Better detection limits than in

the laboratory analysis Br Field values Transportation delay

Fetot Fetot < 20 mg/1 (median value of all Exceeding values discarded Posiva/s sites is 0.06 mg/1)

DOC (Dissolved Organic DOC < 50 mg/1 (median value of all Exceeding values discarded Carbon) Posiva/s sites is 6.40 mg/1) Charge balance Charge balance within ± 5% 18 samples discarded

acceptable Obvious leakage of Indications from hydrological No samples discarded multipacker systems, effects measurements or anomalous of borehole stabilisation with hydrogeochemical results cement The most representative Oz 0.00-0.01 mg/1, 5 samples show very good samples Eh<OmV, representativity according to

0.8 < H-3 < 2.5 TU these criteria

139 Appendix 8

Samples included in the modelling study in spite of failing the quality classification.

Sample Date Reasons for failure in Quality classification Liimatainen 18/08/94 Charge balance> ±5% (+5.28%) KI-KR4/T5 05/04/94 Charge balance> ±5% (-7.23%) KI-KR4/T4 22111193 Charge balance> ±5% (-7.28%) KI-KR4/T1 11110/93 Charge balance> ±5% (-5.48%) KI-KR2/T4 02/06/94 Charge balance > ±5% ( + 7.41%) KI-KR2/T1 08111193 Charge balance> ±5% (+7.81 %);

Drill Water> 2.5% (5.78%); Leakage of multi packer system

KI-KR3/T5 02/06/94 Charge balance> ±5% ( -5.33% ); Leakage of multi packer system

KI-KR3/T2 05/04/94 Charge balance> ±5% (+5.68%) KI-KR3/T1 05/08/93 Charge balance> ±5% (+7.54%)

LIST OF REPORTS 1 (3)

POSIV A REPORTS 1998, situation 12/98

POSIV A 98-01

POSIV A 98-02

POSIV A 98-03

POSIV A 98-04

POSIV A 98-05

Bentonite swelling pressure in strong NaCl solutions - Correlation of model calculations to experimentally determined data OlaKamland Clay Technology, Lund, Sweden January 1998 ISBN 951-652-039-1

A working groups conclusions on site specific flow and transport modelling Johan Andersson Golder Associates AB, Sweden Henry Ahokas Fintact Oy Lasse Koskinen, Antti Poteri VTTEnergy Auli Niemi Royal Institute of Technology, Hydraulic Engineering, Sweden (permanent affiliation: VTT Communities and Infrastructure, Finland) Aimo Hautojiirvi Posiva Oy March 1998 ISBN 951-652-040-5

EB-welding of the copper canister for the nuclear waste disposal­Final report of the development programme 1994-1997 HarriAalto Outokumpu Poricopper Oy October 1998 ISBN 951-652-041-3

An isotopic and fluid inclusion study of fracture calcite from borehole OL-KR1 at the Olkiluoto site, Finland Alexander Blyth, Shaun Frape University of Waterloo, Waterloo, Ontario, Canada Runar Blomqvist, Pasi Nissinen Geological Survey of Finland Robert McNutt McMaster University, Hamilton, Ontario, Canada April1998 ISBN 951-652-042-1

Sorption of iodine on rocks from Posiva investigation sites Seija Kulmala, Martti Hakanen Laboratory of Radiochemistry Department of Chemistry University of Helsinki Antero Lindberg Geological Survey of Finland May 1998 ISBN 951-652-043-X

POSIV A 98-06

POSIV A 98-07

POSIV A 98-08

POSIV A 98-09

POSIV A 98-10

POSIV A 98-11

LIST OF REPORTS

Dissolution of unirradiated U02 fuel in synthetic ground water­Progress report '97 Kaija Ollila VTT Chemical Technology June 1998 ISBN 951-652-044-8

2 (3)

Geochemical modelling of groundwater evolution and residence time at the Kivetty site Petteri Pitkiinen, Ari Luukkonen VTT Communities and Infrastructure Paula Ruotsalainen Fintact Oy Hilkka Leino-Forsman, Ulla Vuorinen VTT Chemical Technology December 1998 ISBN 951-652-045-6

Modelling gas migration in compacted bentonite - A report produced for the GAMBIT Club P.J. Nash, B. T. Swift, M. Goodfield, W.R. Rodwell AEA Technology plc, Dorchester, United Kingdom August 1998 ISBN 951-652-046-4

Geomicrobial investigations of groundwaters from Olkiluoto, Hastholmen, Kivetty and Romuvaara, Finland Shelley A. Haveman, Karsten Pedersen Goteborg University, Sweden Paula Ruotsalainen Fintact Oy August 1998 ISBN 951-652-047-2

Geochemical modeling of groundwater evolution and residence time at the Olkiluoto site Petteri Pitkiinen, Ari Luukkonen VTT Communities and Infrastructure September 1998 (to be published) ISBN 951-652-048-0

Sorption of cesium on Olkiluoto mica gneiss, granodiorite and granite Tuula Huitti, Martti Hakanen Laboratory of Radiochemistry, Department of Chemistry, University of Helsinki Antero Lindberg Geological Survey of Finland September 1998 ISBN 951-652-049-9

POSIVA 98-12

POSIV A 98-13

POSIV A 98-14

POSIVA 98-15

POSIV A 98-16

POSIV A 98-17

LIST OF REPORTS

Sorption of plutonium on rocks in ground waters from Posiva investigation sites Seija Kulmala, Martti Hakanen Laboratory of Radiochemistry, Department of Chemistry, University of Helsinki Antero Lindberg Geological Survey of Finland December 1998 ISBN 951-652-050-2

Solubilities of uranium for TILA-99 Kaija Ollila VTT Chemical Technology Lasse Ahonen Geological Survey of Finland November 1998 ISBN 951-652-051-0

Solubility database for TILA-99 Ulla Vuorinen VTT Chemical Technology Seija Kulmala, Martti Hakanen University of Helsinki, Laboratory of Radiochemistry Lasse Ahonen Geological Survey of Finland Torbjorn Carlsson VTT Chemical Technology November 1998 ISBN 951-652-052-9

Normal evolution of a spent fuel repository at the candidate sites in Finland M.B. Crawford, R.D. Wilmot Galson Sciences Ltd, United Kingdom December 1998 ISBN 951-652-053-7

3 (3)

The social impacts of the final disposal of spent nuclear fuel from the point of view of the inhabitants - Interview research Tytti Viinikainen Centre for Urban and Regional Studies Helsinki University of Technology December 1998 (in Finnish) (to be published) ISBN 951-652-054-5

Possible effects of a proposed high-level nuclear waste repository on consumer demand for goods and services produced in the host community - an overview of the Finnish study Ilpo Koskinen, Mari Niva, Piiivi Timonen Kuluttaj atutkimuskeskus December 1998 (in Finnish) (to be published) ISBN 951-652-055-3